Essentials of Meteorology: An Invitation to the Atmosphere [8 ed.] 9781337515399, 1337515396, 9781305628458, 9781337276108 - DOKUMEN.PUB (2022)

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EIGHTH EDITION

ESSENTIALS OF METEOROLOGY

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Copyright 2018 Cengage Learning. All Rights Reserved. May not be copied, scanned, or duplicated, in whole or in part. WCN 02-200-202

EIGHTH EDITION

Essentials of Meteorology AN INVITATION TO THE ATMOSPHERE

C. Donald Ahrens

Emeritus, Modesto Junior College

Robert Henson

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Essentials of Meteorology An Invitation to the Atmosphere Eighth Edition C. Donald Ahrens and Robert Henson Product Director: Dawn Giovanniello Senior Content Developer: Lauren Oliveira Product Assistant: Marina Starkey Marketing Manager: Ana Albinson Content Project Manager: Hal Humphrey Senior Designer: Michael Cook

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CONTENTS IN BRIEF CHAPTER 1

Earth’s Atmosphere 2

CHAPTER 2

Warming and Cooling Earth and its Atmosphere 26

CHAPTER 3

Air Temperature 54

CHAPTER 4

Humidity, Condensation, and Clouds 78

CHAPTER 5

Cloud Development and Precipitation 114

CHAPTER 6

Air Pressure and Winds 146

CHAPTER 7

Atmospheric Circulations 172

CHAPTER 8

Air Masses, Fronts, and Middle-Latitude Cyclones 208

CHAPTER 9

Weather Forecasting 242

CHAPTER 10

Thunderstorms and Tornadoes 272

CHAPTER 11

Hurricanes 318

CHAPTER 12

Global Climate 350

CHAPTER 13

Earth’s Changing Climate 380

CHAPTER 14

Air Pollution 414

CHAPTER 15

Light, Color, and Atmospheric Optics 442

APPENDICES

A

Units, Conversions, Abbreviations, and Equations 463

B

Equations and Constants 466

C

Weather Symbols and the Station Model 468

D

Average Annual Global Precipitation 470

E

Köppen’s Climatic Classification System 472

F

Humidity and Dew-point Tables (Psychrometric Tables) 473

G

Standard Atmosphere 477

H

Beaufort Wind Scale (Over Land)

478

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CONTENTS Preface xv

CHAPTER 2

CHAPTER 1

Warming and Cooling Earth and Its Atmosphere 26

Earth’s Atmosphere

2

Temperature and Heat Transfer 28 Temperature Scales 28 Latent Heat—The Hidden Warmth 29 Conduction 31 Convection 31

The Atmosphere and the Scientific Method 4 Weather, Climate. and Meteorology 5 Meteorology—The Study of the Atmosphere 5 A Glimpse at a Weather Map 7 Weather and Climate in Our Lives 9

FOCUS ON A SPECIAL TOPIC 2.1

FOCUS ON A SPECIAL TOPIC 1.1

Rising Air Cools and Sinking Air Warms 32

What Is a Meteorologist? 14

Components of Earth’s Atmosphere 15 The Early Atmosphere 15 Composition of Today’s Atmosphere 15 Vertical Structure of the Atmosphere 19 A Brief Look at Air Pressure and Air Density Layers of the Atmosphere 21

Radiant Energy 33 FOCUS ON AN ENVIRONMENTAL ISSUE 2.2

Sunburning and UV Rays 35

Radiation—Absorption, Emission, and Equilibrium 36 Selective Absorbers and the Atmospheric Greenhouse Effect 37 Enhancement of the Greenhouse Effect 38 Warming the Air from Below 40 Shortwave Radiation Streaming from the Sun 41 Earth’s Annual Energy Balance 42 Why Earth Has Seasons 44

19

FOCUS ON AN OBSERVAT RV ION 1.2 RVAT

The Radiosonde 22

Summary 24 Key Terms 24 Questions for Review 24 Questions for Thought and Exploration 25

FOCUS ON A SPECIAL TOPIC 2.3

Space Weather and Its Impact on Earth 44

Seasons in the Northern Hemisphere 46 FOCUS ON A SPECIAL TOPIC 2.4

Is December 21 Really the First Day of Winter? 49

Seasons in the Southern Hemisphere 50 Local Seasonal Variations 51 Summary 52 Key Terms 52 Questions for Review 52 Questions for Thought and Exploration 53

CHAPTER 3

© C. Donald Ahrens

Air Temperature

54

Warming and Cooling Air Near the Surface 56 Daytime Warming 56 Extreme High Temperatures 57 Nighttime Cooling 59 Cold Air Near the Surface 59 Protecting Crops from the Cold Night Air 60 Extreme Low Temperatures 62

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Daily Temperature Variations 63 Regional Temperature Variations 64

Key Terms 111 Questions for Review 111 Questions for Thought and Exploration 112

FOCUS ON A SPECIAL TOPIC 3.1

When It Comes to Temperature, What’s Normal? 65

CHAPTER 5

Applications of Air Temperature Data 67 Air Temperature and Human Comfort 70

Cloud Development and Precipitation 114

FOCUS ON A SPECIAL TOPIC 3.2

A Thousand Degrees and Freezing to Death 71

Measuring Air Temperature 72

Atmospheric Stability 116 Determining Stability 117 A Stable Atmosphere 117 An Unstable Atmosphere 118 A Conditionally Unstable Atmosphere 120 Cloud Development and Stability 121 Convection and Clouds 121 Topography and Clouds 123

FOCUS ON AN OBSERVAT RV ION 3.3 RVAT

Why Thermometers Must Be Read in the Shade 74

Summary 76 Key Terms 76 Questions for Review 76 Questions for Thought and Exploration 77

CHAPTER 4

FOCUS ON A SPECIAL TOPIC 5.1

Atmospheric Stability and Windy Afternoons— Hold On to Your Hat 124

Humidity, Condensation, and Clouds 78

Precipitation Processes 126 Collision and Coalescence Process 126 Ice-Crystal Process 127 Cloud Seeding and Precipitation 129 Precipitation in Clouds 131

Circulation of Water in the Atmosphere 80 Evaporation, Condensation, and Saturation 81 Humidity 82 Vapor Pressure 82 Relative Humidity 83 Relative Humidity and Dew Point 85 Relative Humidity and Human Discomfort 88 FOCUS ON A SPECIAL TOPIC 4.1

Humid Air and Dry Air Do Not Weigh the Same 90

Measuring Humidity 90 Dew and Frost 91 Fog 93 Foggy Weather 96 FOCUS ON AN ENVIRONMENTAL ISSUE 4.2

Clouds 98 Classification of Clouds 98 High Clouds 98 Middle Clouds 99 Low Clouds 100 Clouds with Vertical Development 102 Some Unusual Clouds 103 Clouds and Satellite Imagery 106 Summary 111

© C. Donald Ahrens

Fog Dispersal 97

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vii

CHAPTER 6

FOCUS ON AN ENVIRONMENTAL ISSUE 5.2

Does Cloud Seeding Enhance Precipitation? 131

Air Pressure and Winds

Precipitation Types 132 Rain 132 Snow 133

146

Atmospheric Pressure 148 Horizontal Pressure Variations—a Tale of Two Cities 148 Measuring Air Pressure 150

FOCUS ON A SPECIAL TOPIC 5.3

Are Raindrops Tear-Shaped? 134

Sleet and Freezing Rain 136

FOCUS ON A SPECIAL TOPIC 6.1

FOCUS ON AN OBSERVAT RV ION 5.4 RVAT

The Atmosphere Obeys the Gas Law 150

Aircraft Icing 138

Pressure Readings 152 Surface and Upper-Air Charts 153 Why the Wind Blows 154

Snow Grains and Snow Pellets 138 Hail 138 Measuring Precipitation 140 Instruments 140 Doppler Radar and Precipitation 141 Summary 144 Key Terms 144 Questions for Review 144 Questions for Thought and Exploration 145

FOCUS ON A SPECIAL TOPIC 6.2

Isobaric Maps 155

Newton’s Laws of Motion 155 Forces That Influence the Wind 156 Straight-Line Flow Aloft 159 FOCUS ON AN OBSERVAT RV ION 6.3 RVAT

Estimating Wind Direction and Pressure Patterns Aloft by Watching Clouds 160

Curved Winds Around Lows and Highs Aloft 161 Winds On Upper-Level Charts 162 Surface Winds 163 FOCUS ON AN OBSERVAT RV ION 6.4 RVAT

Winds Aloft in the Southern Hemisphere 164

Winds and Vertical Air Motions 164 Determining Wind Direction and Speed 165 The Influence of Prevailing Winds 166 Wind Instruments 167 FOCUS ON A SPECIAL TOPIC 6.5

Wind Energy 169 Summary 170 Key Terms 170 Questions for Review 170 Questions for Thought and Exploration 171

CHAPTER 7

© C. Donald Ahrens

Atmospheric Circulations

viii

172

Scales of Atmospheric Motion 174 Eddies—Big and Small 175 Local Wind Systems 176 Thermal Circulations 176 FOCUS ON AN OBSERVAT RV ION 7.1 RVAT

Eddies and “Air Pockets” 177

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Sea and Land Breezes 178 Mountain and Valley Breezes 180 Katabatic Winds 180 Chinook (Foehn) Winds 181

Mid-Latitude Cyclonic Storms 229 Polar Front Theory 229 Where Do Mid-Latitude Cyclones Tend to Form? 231 Developing Mid-Latitude Cyclones and Anticyclones 232

FOCUS ON A SPECIAL TOPIC 7.2

Snow Eaters and Rapid Temperature Changes 182

FOCUS ON A SPECIAL TOPIC 8.3

Santa Ana Winds 182 Desert Winds 184 Seasonally Changing Winds—The Monsoon 186 Global Winds 188 General Circulation of the Atmosphere 188 Single-Cell Model 188 Three-Cell Model 189 Average Surface Winds and Pressure: The Real World 192 The General Circulation and Precipitation Patterns 192 Westerly Winds and the Jet Stream 194 Atmosphere-Ocean Interactions 197 Global Wind Patterns and Surface Ocean Currents 197 Winds and Upwelling 198 El Niño , La Niña, and the Southern Oscillation 199 Other Atmosphere-Ocean Interactions 203 Summary 206 Key Terms 206 Questions for Review 206 Questions for Thought and Exploration 207

Nor’easters 233 FOCUS ON A SPECIAL TOPIC 8.4

A Closer Look at Convergence and Divergence 234

Summary 239 Key Terms 239 Questions for Review 239 Questions for Thought and Exploration

240

CHAPTER 9

Weather Forecasting

242

Weather Observations 244 Acquisition of Weather Information 244 Weather Forecasting Tools 245 FOCUS ON AN OBSERVAT RV ION 9.1 RVAT

TV Weathercasters—How Do They Do It? 246

Weather Forecasting Methods 247

CHAPTER 8

Air Masses, Fronts, and Middle-Latitude Cyclones

208

Air Masses 210 Source Regions 210 Classification 211 Air Masses of North America 212 FOCUS ON A SPECIAL TOPIC 8.1

Lake-Effect (Enhanced) Snows 213 FOCUS ON A SPECIAL TOPIC 8.2

Fronts 220 Stationary Fronts 221 Cold Fronts 222 Warm Fronts 224 Drylines 226 Occluded Fronts 227

NASA

The Return of the Siberian Express 215

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ix

Squall-Line Thunderstorms 280 Supercell Thuderstorms 282 Thunderstorms and Flooding 286 Distribution of Thunderstorms 288

The Computer and Weather Forecasting: Numerical Weather Prediction 247 Why Computer-Based Forecasts Can Go Awry and Steps to Improve Them 249 Other Forecasting Techniques 252 Time Range of Forecasts 255

FOCUS ON A SPECIAL TOPIC 10.1

The Terrifying Flash Flood in the Big Thompson Canyon 287

FOCUS ON A SPECIAL TOPIC 9.2

Distribution of Thunderstorms 288

Weather Prediction and The Marketplace 257

FOCUS ON AN OBSERVAT RV ION 10.2 RVAT

Accuracy and Skill in Weather Forecasting 257 Weather Forecasting Using Surface Charts 259 Determing the Movement of Weather Systems 259 A Forecast for Six Cities 259 Using Forecasting Tools to Predict the Weather 264 Help from the 500-mb Chart 264 The Models Provide Assistance 266 A Valid Forecast 267 Satellite and Upper-Air Assistance 267 A Day of Rain and Wind 268 Summary 270 Key Terms 270 Questions for Review 270 Questions for Thought and Exploration 271

Thunderstorms and the Dryline 290

Lightning and Thunder 290 Tornadoes 296 FOCUS ON A SPECIAL TOPIC 10.3

Don’t Sit Under the Apple Tree 297

Tornado Life Cycle 297 Tornado Occurrence and Distribution 298 Tornado Winds 300 FOCUS ON A SPECIAL TOPIC 10.4

The Weird World of Tornado Damage 302 FOCUS ON A SPECIAL TOPIC 10.3

The Evolution of Tornado Watches and Warnings 304

Tornado Outbreaks 305 Tornado Formation 305 Supercell Tornadoes 305 Nonsupercell Tornadoes 309 Waterspouts 310 Observing Tornadoes and Severe Weather 311 Storm Chasing and Mobile Radar 314 Summary 315 Key Terms 315 Questions for Review 315 Questions for Thought and Exploration 316

CHAPTER 10

Thunderstorms and Tornadoes

272

Thunderstorms 274 Ordinary Cell Thunderstorms 275 Multicell Thunderstorms 277

CHAPTER 11

© Robert Henson

Hurricanes

x

318

Tropical Weather 320 Anatomy of a Hurricane 320 Hurricane Formation and Dissipation 323 The Right Environment 324 The Developing Storm 324 The Storm Dies Out 325 Hurricane Stages of Development 325 Investigating the Storm 326 Hurricane Movement 327

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Moist Subtropical Mid-Latitude Climates (Group C) 366

FOCUS ON A SPECIAL TOPIC 11.1

How Do Hurricanes Compare with Middle-Latitude Cyclones? 328

FOCUS ON AN OBSERVAT RV ION 12.2 RVAT

Naming Hurricanes and Tropical Storms 331 Devastating Winds, the Storm Surge, and Flooding 331

A Desert with Clouds and Drizzle 367

Moist Continental Climates (Group D) 370

FOCUS ON A SPECIAL TOPIC 11.2

FOCUS ON A SPECIAL TOPIC 12.3

Devastation from a Tropical Storm—The —The Case of Allison 334

When Does a Dry Spell Become a Drought? 372

Polar Climates (Group E) 374 Highland Climates (Group H) 375

Classifying Hurricane Strength 334 Hurricane-Spawned Tornadoes 336 Hurricane Fatalities 336 Some Notable Hurricanes 337 Galveston, 1900 337 New England, 1938 337 Camille, 1969 337 Hugo, 1989 337 Andrew, 1992 338 Katrina and Rita, 2005 339

FOCUS ON AN ENVIRONMENTAL ISSUE 12.4

Are Plant Hardiness Zones Shifting Northward? 376

Summary 378 Key Terms 378 Questions for Review 378 Questions for Thought and Exploration 379

CHAPTER 13

FOCUS ON AN OBSERVAT RV ION 11.3 RVAT

The Record-Setting Atlantic Hurricane Seasons of 2004 and 2005 340

Earth’s Changing Climate

Sandy, 2012 341 Devastating Tropical Cyclones Around the World 342 Hurricane Watches and Warnings 343

380

Reconstructing Past Climates 382 Climate Throughout the Ages 384 FOCUS ON A SPECIAL TOPIC 13.1

The Ocean’s Influence on Rapid Climate Change 386

FOCUS ON AN ENVIRONMENTAL ISSUE 11.4

Hurricanes in a Warmer World 344 Hurricane Forecasting Techniques 345 Modifying Hurricanes 347 Summary 348 Key Terms 348 Questions for Review 348 Questions for Thought and Exploration 349

CHAPTER 12

Global Climate

350

A World with Many Climates 352 Global Temperatures 352 Global Precipitation 354 FOCUS ON A SPECIAL TOPIC 12.1

Climatic Classification—The Köppen System 357 The Global Pattern of Climate 359 Tropical Moist Climates (Group A) 359 Dry Climates (Group B) 362

© C. Conald Ahrens

Extreme Wet and Dry Regions 356

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xi

Temperature Trends During the Past 1000 Years 385 Temperature Trends During the Past 100-Plus Years 386 Climate Change Caused by Natural Events 388 Climate Change: Feedback Mechanisms 389 Climate Change: Plate Tectonics and Mountain Building 390 Climate Change: Variations in the Earth’s Orbit 391 Climate Change: Variations in Solar Output 394 Climate Change: Atmospheric Particles 394 Climate Change Caused by Human (Anthropogenic) Activities 397 Climate Change: Greenhouse Gases 397

FOCUS ON AN ENVIRONMENTAL ISSUE 13.5

The Impact of Ozone on the Greenhouse Effect and Climate Change 406

Consequences of Climate Change: The Possibilties 406 Climate Change: A Final Note 410 Summary 411 Key Terms 411 Questions for Review 411 Questions for Thought and Exploration 412

CHAPTER 14

Air Pollution

414

A Brief History of Air Pollution 416 Types and Sources of Air Pollutants 418

FOCUS ON AN ENVIRONMENTAL ISSUE 13.2

Nuclear Winter—Climate Change Induced by Nuclear War 398

FOCUS ON AN ENVIRONMENTAL ISSUE 14.1

Climate Change: Land Use Changes 398 Climate Change: Global Warming 399 Recent Global Warming: Perspective 399

Indoor Air Pollution 418

Principal Air Pollutants 419 Ozone in the Troposphere 422 Ozone in the Stratosphere 422

FOCUS ON AN ENVIRONMENTAL ISSUE 13.3

Climate Change and Extreme Weather 400

FOCUS ON A SPECIAL TOPIC 14.2

Future Climate Change: Projections 401

The Formation of Ground-Level Ozone in Polluted Air 423

FOCUS ON A SPECIAL TOPIC 13.4

FOCUS ON AN ENVIRONMENTAL ISSUE 14.3

Climate Models—A Quick Look 404

The Ozone Hole 425

Air Pollution: Trends and Patterns 426 Factors That Affect Air Pollution 429 The Role of the Wind 429 The Role of Stability and Inversions 429 The Role of Topography 431 FOCUS ON AN OBSERVAT RV ION 14.4 RVAT

Smokestack Plumes 432

Severe Air Pollution Potential 433 Air Pollution and the Urban Environment 433 FOCUS ON AN OBSERVAT RV ION 14.5 RVAT

Five Days in Donora—An Air Pollution Episode 434

Acid Deposition 436 FOCUS ON AN ENVIRONMENTAL ISSUE 14.6

© Robert Henson

Heat Waves and Air Pollution: A Deadly Team 437

xii

Summary 439 Key Terms 439 Questions for Review 439 Questions for Thought and Exploration 440

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CHAPTER 15

APPENDIXES

Light, Color, and Atmospheric Optics 442

A Units, Conversions, Abbreviations, and Equations 463 B Equations and Constants 466 C Weather Symbols and the Station Model 468 D Average Annual Global Precipitation 470 E Köppen’s Climatic Classification System 472 F Humidity and Dew-point Tables (Psychrometric Tables) 473 G Standard Atmosphere 477 H Beaufort Wind Scale (Over Land) 478

White and Colors 444 Clouds and Scattered Light 444 Blue Skies and Hazy Days 445 Red Suns and Blue Moons 447 Twinkling, Twilight, and the Green Flash 449 The Mirage: Seeing Is Not Believing 451 FOCUS ON AN OBSERVAT RV ION 15.1 RVAT

The Fata Morgana 453

Halos, Sundogs, and Sun Pillars 453 Rainbows 456 Coronas and Cloud Iridescence 459

Additional Reading Material 479 Glossary 481 Index 496

FOCUS ON AN OBSERVAT RV ION 15.2 RVAT

Glories and the Heiligenschein 460

Summary 461 Key Terms 461 Questions for Review 461 Questions for Thought and Exploration 462

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xiii

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PREFACE The world is an ever-changing picture of naturally occurring events. From drought and famine to devastating floods, some of the greatest challenges we face come in the form of natural disasters created by weather. Yet, dealing with weather and climate is an inevitable part of our lives. Sometimes it is as small as deciding what to wear for the day or how to plan a vacation. But it can also have life-shattering consequences, especially for those who are victims of a hurricane or a tornado. In recent years, weather and climate have become front page news, from record-setting extreme weather events to environmental issues such as global warming and ozone depletion. The dynamic nature of the atmosphere seems to demand our attention and understanding more these days than ever before. Almost daily, there are newspaper articles describing some weather event or impending climate change. For this reason, and the fact that weather influences our daily lives in so many ways, interest in meteorology (the study of the atmosphere) has been growing. This rapidly developing and popular science is giving us more information about the workings of the atmosphere than ever before. One of the reasons that meteorology is such an engaging science to study is that the atmosphere is a universally accessible laboratory for everyone. Although the atmosphere will always provide challenges for us, as research and technology advance, our ability to understand our atmosphere improves as well. The information available to you in this book, therefore, is intended to aid in your own personal understanding and appreciation of our Earth’s dynamic atmosphere.

About This Book Essentials of Meteorology is written for students taking an introductory course on the atmospheric environment. The main purpose of the text is to convey meteorological concepts in a visual, practical, and nonmathematical manner. In addition, the intent of the book is to stimulate curiosity in the reader and to answer questions about weather and climate that arise in our day-to-day lives. Although introductory in nature, this eighth edition maintains scientific integrity and includes up-todate information on large-scale topics, such as global warming, ozone depletion, and El Niño, as well as discussion of recent high-profile weather events. As in previous editions, no special prerequisites are necessary for understanding.

Written expressly for the student, this book emphasizes the understanding and application of meteorological principles. The text encourages watching the weather so that it becomes “alive,” allowing readers to immediately apply textbook material to the world around them. To assist with this endeavor, a color Cloud Chart appears at the back of the text. The Cloud Chart can be separated from the book and used as a learning tool at any place one chooses to observe the sky. To strengthen points and clarify concepts, illustrations are rendered in full color throughout. Color photographs were carefully selected to illustrate features, stimulate interest, and show how exciting the study of weather can be. To enhance the value of the book, several appendices that were only available online in the seventh edition have been reincorporated. Organized into fifteen chapters, Essentials of Meteorology is designed to provide maximum flexibility to instructors of weather and climate courses. Thus, chapters can be covered in any desired order. For example, Chapter 15, “Light, Color, and Atmospheric Optics,” is self-contained and can be covered earlier if so desired. Instructors, then, are able to tailor this text to their particular needs. This book basically follows a traditional approach. After an introductory chapter on the origin, composition, and structure of the atmosphere, it covers solar energy, air temperature, humidity, clouds, precipitation, and winds. Then comes a chapter on air masses, fronts, and middle-latitude cyclonic storms. Weather prediction and severe storms are next. A chapter on hurricanes is followed by a chapter on global climate. A chapter on climate change is next. A chapter on air pollution precedes the final chapter on atmospheric optics. Each chapter contains at least two Focus sections, which either expand on material in the main text or explore a subject closely related to what is being discussed. Focus sections fall into one of three distinct categories: Observations, Special Topics, and Environmental Issues. Some include material that is not always found in introductory meteorology textbooks—subjects such as space weather, the scientific method, and wind energy. Others help to bridge theory and practice. This edition contains several new or rewritten Focus sections, including an updated discussion of nor’easters in Chapter 8 and a new Focus section on tornado damage patterns in Chapter 10. Set apart as “Did You Know?” features in each chapter is weather information that may not be commonly xv

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known, yet pertains to the topic under discussion. Designed to bring the reader into the text, most of these weather highlights relate to some interesting weather fact or astonishing event. Each chapter incorporates other effective learning aids: ● ●

● ●

A major topic outline begins each chapter. Interesting introductory pieces draw the reader naturally into the main text. Important terms are boldfaced, with their definitions appearing in the glossary or in the text. Key phrases are italicized. English equivalents of metric units are immediately provided in parentheses. A brief review of the main points is placed toward the middle of most chapters. Each chapter ends with a summary of the main ideas. A list of key terms with page references follows each chapter, allowing students to review and reinforce their knowledge of key concepts. Questions for Review act to check how well students assimilate the material. Questions for Thought and Exploration encourage students to synthesize learned concepts for deeper understanding. References to 19 concept animations are compiled on pp. xx-xxi. These animations convey an immediate appreciation of how a process works and help students visualize the more difficult concepts in meteorology. Animations can be found on the Meteorology CourseMate, accessed through Cengagebrain.com. At the end of each chapter are questions that relate to articles found on the Global Geoscience Watch website, available on its own or via the Meteorology CourseMate.

Eight appendices conclude the book. In addition, at the end of the book, a compilation of supplementary reading material is presented, as is an extensive glossary. On the endsheet at the back of the book is a geophysical map of North America. The map serves as a quick reference for locating states, provinces, and geographical features, such as mountain ranges and large bodies of water.

xvi

Supplemental Material and Technology Support TECHNOLOGY FOR THE INSTRUCTOR Instructor Companion Website Everything you need for your course in one place! This collection of bookspecific lecture and class tools is available online via www.cengage.com/login. Access and download PowerPoint presentations, images, instructor’s manual, videos, and more. Cognero Test Bank Cengage Learning Testing Powered by Cognero is a flexible, online system that allows you to: ●

● ●

author, edit, and manage test bank content from multiple Cengage Learning solutions create multiple test versions in an instant deliver tests from your LMS, your classroom, or wherever you want

Global Geoscience Watch Updated several times a day, the Global Geoscience Watch is a focused portal into GREENR—our Global Reference on the Environment, Energy, and Natural Resources—an ideal onestop site for classroom discussion and research projects for all things geoscience! Broken into the four key course areas (Geography, Geology, Meteorology, and Oceanography), you can easily get to the most relevant content available for your course. You and your students will have access to the latest information from trusted academic journals, news outlets, and magazines. You also will receive access to statistics, primary sources, case studies, podcasts, and much more! TECHNOLOGY FOR THE STUDENT Earth Science MindTap for Essentials of Meteorology MindTap is well beyond an eBook, a homework solution or digital supplement, a resource center website, a course delivery platform, or a Learning Management System. More than 70 percent of students surveyed said that it was unlike anything they have ever seen before. MindTap is a new personal learning experience that combines all of your digital assets—readings, multimedia, activities, study tools, and assessments—into a singular learning path to improve student outcomes.

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Lab Manual Developed by the Oklahoma Climatological Survey (OCS) research and service facility, in concert with the University of Oklahoma, Explorations in Meteorology places a strong emphasis on helping students understand weather and climate by using real meteorological data. The activities in this lab manual require that students tap into the OCS archives of meteorological data in order to complete meteorological exercises. Full-color pictures and data graphs help students visually understand weather and severe weather topics. The lab manual also challenges students by providing optional questions intended for honors students, making this lab manual appropriate for both introductory and honors meteorology courses.

Eighth Edition Changes This edition of Essentials of Meteorology includes a coauthor—meteorologist and science journalist Robert Henson (Weather Underground). For more than 20 years, Henson produced publications and websites for the University Corporation for Atmospheric Research, which manages the National Center for Atmospheric Research. He is an expert on severe weather, including tornadoes, thunderstorms, and hurricanes. He has also analyzed how television weathercasters cover major storms and report on climate change. Henson is the author of four trade books on meteorology, including The Thinking Person’s Guide to Climate Change (previously The Rough Guide to Climate Change, the first edition of which was shortlisted for the United Kingdom’s Royal Society Prize for Science Books). The authors have carried out extensive updates and revisions to this eighth edition of Essentials of Meteorology, reflecting the ever-changing nature of the field and the atmosphere itself. Dozens of new photos and new or revised color illustrations help students visualize the excitement of the atmosphere. ●

Chapter 1, “Earth’s Atmosphere,” continues to serve as a broad overview of the atmosphere. The text now begins with a discussion of the scientific method and its importance. To help draw students into the material, the introduction to meteorology and the summary of extreme weather types has been placed earlier in the chapter, followed by discussion of the chemistry and vertical structure of Earth’s atmosphere. Among

recent events now included are the severe flooding over the Southern Plains and Southeast in 2015 and the Houston flash flood of April 2016. Chapter 2, “Warming and Cooling Earth and Its Atmosphere,” contains up-to-date statistics and background on greenhouse gases and climate change, topics covered in more detail later in the book. Discussion of the potential impact of clouds on future global warming has been updated. In Chapter 3, “Air Temperature,” several figures and tables have been updated so that they refer to normals drawn from the most recent reference period (1981–2010). Chapter 4, “Humidity, Condensation, and Clouds,” includes updated material on satellite observations, including new background and artwork from the Global Precipitation Mission satellite. Also included is a “Did You Know?” box on the high-impact Atlanta snowstorm of January 2014. Chapter 5, “Cloud Development and Precipitation,” includes a new graphic. New satellite observing techniques are also noted and illustrated in this chapter. Chapter 6, “Air Pressure and Winds,” includes a substantially enhanced description and revised illustrations of the interplay between the pressure gradient and Coriolis forces in cyclonic and anticyclonic flow. Several other illustrations have been revised for clarity, and the discussion of scatterometers has been updated. A box on wind energy features the most recent data on wind energy adoption. Chapter 7, “Atmospheric Circulations,” features a major restructuring, update, and expansion of sections dealing with the El Niño/Southern Oscillation, Pacific Decadal Oscillation, North Atlantic Oscillation, and Arctic Oscillation, including several new and updated images. The opening section, which introduces scales of atmospheric motion, has also been revised for clarity. Chapter 8, “Air Masses, Fronts, and Middle-Latitude Cyclones,” now includes discussion of atmospheric rivers and an illustration of their impacts. In line with recent research, the section on occluded fronts stresses the prevalence of warm-type over cold-type occluded fronts. The discussion of drylines has been expanded, and the Focus box on nor’easters has been reworked to spotlight the record-setting East Coast snowstorm of January 2016. PREFACE

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“Weather Forecasting” (Chapter 9) has undergone substantial revision, with several updated graphics. Three major types of satellite imagery are introduced near the beginning of the chapter. Explanations of watches, warnings, and forecasts of various durations (including seasonal outlooks) are now incorporated in a new section, “Time Range of Forecasts.” The concept of the forecast funnel is also introduced. Chapter 10, “Thunderstorms and Tornadoes,” includes several new and updated illustrations, depicting low- and high-precipitation supercells, a roll cloud, and a shelf cloud. An expanded section covers both flash flooding and river flooding and their connection to thunderstorms, including examples from Colorado (2013) and Texas and Oklahoma (2015). A new Focus box explores the baffling damage patterns that tornadoes can produce. The effort to accommodate new ways of estimating and reporting tornado wind speed (such as mobile radar reports) is also noted. The chapter on “Hurricanes” (Chapter 11) includes a new opening section that introduces students to the terrible impacts of Hurricane Katrina. Several graphics that use satellite imagery to explain basic concepts have been updated with recent tropical cyclones. Charts on hurricane climatology have been brought up to date, and an expanded range of both historical and recent examples are discussed, including the Galveston hurricane of 1900, the New England hurricane of 1938, and Typhoon Haiyan from 2013. Chapter 12, “Global Climate,” includes a number of updates to climatological charts and discussion, drawing on the most recent set of United States climate normals (1981–2010). Chapter 13, “Earth’s Changing Climate,” has been revised throughout to reflect increasing confidence on a variety of climate change indicators and impacts. Also incorporated are graphics, conclusions, and emission pathways from the Fifth Assessment Report (2013-14) of the Intergovernmental Panel on Climate Change. The extremely quiet solar cycle of the late 2000s and early 2010s is noted, along with a variety of weather extremes from recent years that are relevant to climate change. The Paris Accord is discussed in the context of the Kyoto Protocol

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that preceded it. A number of diagrams have been updated. The chapter on “Air Pollution” (Chapter 14) has been revised to include the latest air pollution trends across the United States as well as the latest information on ozone depletion in the Arctic and Antarctic. The U.S. Clean Power Plan is introduced, and the devastating impacts of both indoor and outdoor air pollution are discussed, including the effects of tiny particulates on cardiovascular health. Several of the photos in Chapter 15, “Light, Color, and Atmospheric Optics,” have been replaced with spectacular new examples (e.g., anticrepuscular rays and double rainbows).

Acknowledgments Many people have contributed to the eighth edition of Essentials of Meteorology. A very special and most grateful thank-you goes to Lita Ahrens who proofread each chapter. Special thank you to Charles Preppernau for rendering the beautiful art and to Janet Hansen for careful proofreading. We are indebted to Janet Alleyn, who not only designed the book but, once again, took the art, photos, and manuscript and turned them into a beautiful book. Thanks goes to Judith Chaffin for her careful and conscientious editing. Special thanks to all the people at Cengage Learning who worked on this edition, including Lauren Oliveira, Morgan Carney, Hal Humphrey, and Dawn Giovanniello. Thanks to our friends who provided photos and to those reviewers who offered comments and suggestions for this edition, including: Fidel González Rouco Universidad Complutense de Madrid Redina Herman Western Illinois University Bette Otto-Bliesner National Center for Atmospheric Research David Schultz University of Manchester Alex Huang University of North Carolina at Asheville

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Anthony Santorelli Anne Arundel Community College Dan Ferandez Anne Arundel Community College Dean G Butzow Western Michigan University Douglas K. Miller Purdue University Edward J. Perantoni Lindenwood University Ronald A Dowey Harrisburg Area Community College Shaunna L. Donaher University of Miami Troy Kimmel University of Texas

To the Student Learning about the atmosphere can be a fascinating and enjoyable experience. This book is intended to give you some insight into the workings of the atmosphere. However, for a real appreciation of your atmospheric environment, you must go outside and observe. Although mountains take millions of years to form, a cumulus cloud can develop into a raging thunderstorm in less than an hour. The atmosphere is always producing something new for us to behold. To help with your observations, a color Cloud Chart is at the back of the book for easy reference. Remove it and keep it with you. And, remember, all of the concepts and ideas in this book are manifested out there for you to discover and enjoy. Please take the time to look. Donald Ahrens and Robert Henson

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EXP XPLORE LORE THE CONCEP CONCEPT ANIMATIONS These animations have been carefully created to bring to life key points in the chapters. They are also the perfect tool to help refresh students’ memories of previous concepts, so they can keep building on knowledge already

acquired. Concept Animations are accessed through the MindTap platform, which can be acquired separately or together with print or looseleaf versions of this book. Some examples of Concept Animations are shown here.

Doppler radar images are used extensively throughout this book. To better understand Doppler radar images, watch all 4 parts of this Doppler Radar animation (Chapters 1, 5, and 10).

For a visual interpretation of the energy emitted by the earth without and with a greenhouse effect, watch the Greenhouse animation (Chapter 2).

Learn about how air rises above an area of low atmospheric pressure and sinks above an area of high atmospheric pressure. Converging and Diverging Air (Chapters 1 and 8).

Additional Animations: ● Ice Crystals (Bergeron) Process (Chapter 5) ● General Circulation of the Atmosphere (Chapter 7) ● Geostrophic Wind (Chapter 6) ● Temperature versus Molecular Movement (Chapter 2) ● Condensation (Chapter 4) ● Air Temperature, Dew Point, and Relative Humidity (Chapter 4) ● Daily Temperature Changes Above the Surface (Chapter 3)

Seasons provides a complete picture of Earth revolving around the sun while it is tilted on its axis. While viewing this animation, look closely at how the sun is viewed by a mid-latitude observer at various times of the year (Chapter 2).

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To view air rising over a mountain and the formation of a rain shadow desert, watch Air Rising Up and Over a Mountain (Chapters 5 and 7).

For a visual presentation of the Coriolis force, watch Coriolis Force (Chapter 6).

The concept of atmospheric stability can be a bit confusing, especially when comparing the temperature inside a rising air parcel to that of its surroundings. Watch Stable Atmosphere (Chapter 5) and the two animations Unstable Atmosphere and Conditionally Unstable Atmosphere (Chapters 5 and 10).

For a visualization of a cold front moving across the landscape, watch Cold Front in Winter (Chapters 8 and 9), To see a warm front actually move across the surface, watch Warm Front in Winter (Chapters 8 and 9).

For a visualization of the stages that a wave cyclone goes through from birth to decay, watch the animation entitled Cyclogenesis (Chapter 8).

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CHAPTER

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Earth’s Atmosphere Contents The Atmosphere and the Scientific Method Weather, Climate, and Meteorology

I

remember well a brilliant red balloon which kept me comcom pletely happy for a whole afternoon, until, while I was

playing, a clumsy movement allowed it to escape. Spellbound, I gazed after it as it drifted silently away, gently swaying, growing smaller and smaller until it was only a red point in a blue sky. Att that moment I realized, for the first time, the

Components of Earth’s Atmosphere

vastness above us: a huge space without visible limits. It was

Vertical Structure of the Atmosphere

power over all the Earth’s inhabitants. I believe that many

an apparent void, full of secrets, exerting an inexplicable people, consciously or unconsciously, have been filled with awe by the immensity of the atmosphere. All ll our knowledge about the air, gathered over hundreds of years, has not diminished this feeling. Theo Loebsack, Our Atmosphere

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O

ur atmosphere is a delicate life-giving blanket of air that surrounds the fragile Earth. In one way or another, it influences everything we see and hear—it is intimately connected to our lives. Air is with us from birth, and we cannot detach ourselves from its presence. In the open air, we can travel for many thousands of kilometers in any horizontal direction, but should we move a mere eight kilometers above the surface, we would suffocate. We may be able to survive without food for a few weeks, or without water for a few days, but, without our atmosphere, we would not survive more than a few minutes. Just as fish are confined to an environment of water, so we are confined to an ocean of air. Anywhere we go, air must go with us. Earth without an atmosphere would have no lakes or oceans. There would be no sounds, no clouds, no red sunsets. The beautiful pageantry of the sky would be absent. It would be unimaginably cold at night and unbearably hot during the day. All things on Earth would be at the mercy of an intense sun beating down upon a planet utterly parched. Living on the surface of Earth, we have adapted so completely to our environment of air that we sometimes forget how truly remarkable this substance is. Even though air is tasteless, odorless, and (most of the time) invisible, it protects us from the scorching rays of the sun and provides us with a mixture of gases that allows life to flourish. Because we cannot see, smell, or taste air, it may seem surprising that between your eyes and these words are trillions of air molecules. Some of these may have been in a cloud only yesterday, or over another continent last week, or perhaps part of the life-giving breath of a person who lived hundreds of years ago. Warmth for our planet is provided primarily by the sun’s energy. At an average distance from the sun of nearly 150 million kilometers (km), or 93 million miles (mi), Earth intercepts only a very small fraction of the sun’s total energy output. However, it is this radiant energy*that drives the atmosphere into the patterns of everyday wind and weather, and allows life to flourish. At its surface, Earth maintains an average temperature of about 15°C (59°F).** Although this temperature is mild, Earth experiences a wide range of temperatures, as readings can drop below 85°C (121°F) during a frigid Antarctic night and climb during the day to above 50°C (122°F) on the oppressively hot, subtropical desert. Not everyone will experience such extremes where they live, but all of us have a chance to observe the day-to-day changes in the atmosphere that we refer to as weather. In this chapter, we will examine a number of important concepts and ideas about Earth’s atmosphere, many of which will be expanded in subsequent chapters.

*Radiant energy, or radiation, is energy transferred in the form of waves that have electrical and magnetic properties. The light that we see is radiation, as is ultraviolet light. More on this important topic is given in Chapter 2. **The abbreviation °C is used when measuring temperature in degrees Celsius, and °F is the abbreviation for degrees Fahrenheit. More information about temperature scales is given in Appendix A and in Chapter 2.

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These concepts and ideas are part of the foundation for understanding the atmosphere and how it produces weather. They are built on knowledge acquired and applied through the scientific method, which allows us to make informed predictions about how the natural world will behave.

The Atmosphere and the Scientific Method For hundreds of years, the scientific method has served as the backbone for advances in medicine, biology, engineering, and many other fields. In the field of atmospheric science, the scientific method has paved the way for the production of weather forecasts that have steadily improved over time. Investigators use the scientific method by posing a question, putting forth a hypothesis,* predicting what the hypothesis would imply if it were true, and carrying out tests to see if the prediction is accurate. Many common sayings about the weather, such as “red sky at morning, sailor take warning; red sky at night, sailor’s delight” (see Fig. 1.1) are rooted in careful observation, and there are grains of truth in some of them. However, they are not considered to be products of the scientific method because they are not tested and verified in a standard rigorous way. To be accepted, a hypothesis has to be shown to be correct through a series of quantitative tests. In many areas of science, such testing is carried out in a laboratory, where it can be replicated again and again. Studying the atmosphere, however, is somewhat different, because our Earth has only one atmosphere. Despite this limitation, scientists have made vast progress by studying the physics and chemistry of air in the laboratory (for instance, studying the way in which molecules absorb energy) and by extending those understandings to the atmosphere as a whole. Observations using weather instruments allow us to quantify how the atmosphere behaves and to determine whether a prediction is accurate. If a particular kind of weather is being studied, such as hurricanes or snowstorms, a field study can gather additional observations to test specific hypotheses. Over the last fifty years, computers have given atmospheric scientists a tremendous boost. The physical laws that control atmospheric behavior can be represented in software packages known as numerical models. Forecasts can be made and tested many times over. The atmosphere described by a model can be used to depict weather conditions from the past and to project them into the future. When a model can accurately simulate past weather conditions and provide confidence in its portrayal of tomorrow’s weather, the model can provide valuable information about the weather and climate we may expect decades from now. *A hypothesis is an assertion subject to verification or proof.

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© University Corporation for Atmospheric Research, Photo by Carlye Calvin

FIGURE 1.1 Observing the natural world FI is a critical part of the scientific method. Here a vibrant red sky is visible at sunset. One might use scientific method to verify the old proverb: “Red sky at morning, sailor take warning; red sky at night, sailor’s delight.”

Weather, Climate, and Meteorology When we talk about the weather, we are talking about the condition of the atmosphere at any particular time and place. Weather—which is always changing—is comprised of the elements of: . air temperature—the degree of hotness or coldness of the air . air pressure—the force of the air above an area . humidity—a measure of the amount of water vapor in the air . clouds—visible masses of tiny water droplets and ice crystals or both that are above Earth’s surface . precipitation—any form of water, either liquid or solid (rain or snow), that falls from clouds and reaches the ground . visibility—the greatest distance one can see . wind—the horizontal movement of air If we measure and observe these weather elements over a specified interval of time, say, for many years, we would obtain the “average weather,” or the climate, of a particular region. Climate, therefore, represents the accumulation of daily and seasonal weather events (the average range of weather) over a long period of time. The concept of climate is much more than this, however, for it also includes the extremes of weather—the heat waves of summer and the cold spells of winter— that occur in a particular region. The frequency of these extremes is what helps us distinguish among climates that have similar averages. If we were able to watch Earth for many thousands of years, even the climate would change. We would see rivers of ice moving down stream-cut valleys and huge glaciers—sheets of moving snow and ice—spreading their

icy fingers over large portions of North America. Advancing slowly from Canada, a single glacier might extend as far south as Kansas and Illinois, with ice several thousands of meters thick covering the region now occupied by Chicago. Over an interval of two million years or so, we would see the ice advance and retreat many times. Of course, for this phenomenon to happen, the average temperature of North America would have to decrease and then rise in a cyclic manner. Suppose we could photograph Earth once every thousand years for many hundreds of millions of years. In time-lapse film sequence, these photos would show that not only is the climate altering, but the whole Earth itself is changing as well: Mountains would rise up only to be torn down by erosion; isolated puffs of smoke and steam would appear as volcanoes spew hot gases and fine dust into the atmosphere; and the entire surface of Earth would undergo a gradual transformation as some ocean basins widen and some of them shrink.* In summary, Earth and its atmosphere are dynamic systems that are constantly changing. While major transformations of Earth’s surface are completed only after long spans of time, the state of the atmosphere can change in a matter of minutes. Hence, a watchful eye turned skyward will be able to observe many of these changes. Up to this point, we have looked at the concepts of weather and climate without discussing the word meteorology. What does this word actually mean, and where did it originate? METEOROLOGY—THE STUDY OF THE ATMOSPHERE Meteorology is the study of the atmosphere and its phenomena. The term itself goes back to the Greek *The movement of the ocean floor and continents is explained in the theory of plate tectonics. EARTH’S ATMOSPHERE

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6

NOAA

FIGURE 1.2 Doppler radar image showing precipitation over FI portions of Indiana. The areas shaded light green indicate lighter rain, whereas yellow indicates heavier rain. The dark red shaded areas represent the heaviest rain and the possibility of hail and intense thunderstorms.

severe thunderstorm and unveil its wind, as well as show precipitation intensity (see Fig. 1.2). Recent upgrades to these Doppler radars make it possible for them to distindistin guish raindrops, snowflakes, and hailstones. In 1960, the first weather satellite, Tiros 1, was launched, ushering in space-age meteorology. Subsequent satellites provided a wide range of useful information, ranging from day and night time-lapse images of clouds and storms to images that depict swirling ribbons of water vapor flowing around the globe, as shown in Fig. 1.3. Over the last several decades, even more sophisticated satsat ellites have been developed. These satellites are supplying computers with a far greater network of data so that more accurate forecasts—perhaps extending up to two weeks or more—will be available in the future.

NOAA

philosopher Aristotle, who, about 340 b.c., wrote a book on natural philosophy titled Meteorologica. This work represented the sum of knowledge on weather and climate at that time, as well as material on astronomy, geography, and chemistry. Some of the topics covered included clouds, rain, snow, wind, hail, thunder, and hurricanes. In those days, all substances that fell from the sky, and anything seen in the air, were called meteors, hence the term meteorology, which actually comes from the Greek word meteoros, meaning “high in the air.” Today, we differentiate between those meteors that come from extraterrestrial sources outside our atmosphere (meteoroids) and particles of water and ice observed in the atmosphere (hydrometeors). In Meteorologica, Aristotle attempted to explain atmospheric phenomena in a philosophical and speculative manner. Even though many of his ideas were found to be erroneous, Aristotle’s work remained a dominant influence in the field of meteorology for almost two thousand years. In fact, the birth of meteorology as a genuine natural science did not take place until the invention of weather instruments, such as the hygrometer in the mid-1400s, the thermometer in the late 1500s, and the barometer (for measuring air pressure) in the mid-1600s. With the newly available observations from instruments, attempts were then made to explain certain weather phenomena employing scientific experimentation and the physical laws that were being developed at the time. As more and better instruments were developed, in the 1800s, the science of meteorology progressed. The invention of the telegraph in 1843 allowed for the transmission of routine weather observations. The understanding of the concepts of wind flow and storm movement became clearer, and in 1869 crude weather maps with isobars (lines of equal pressure) were drawn. Around 1920 in Norway, the concepts of air masses and weather fronts were formulated. By the 1940s, daily upper-air balloon observations of temperature, humidity, and pressure gave a three-dimensional view of the atmosphere, and high-flying military aircraft discovered the existence of jet streams. Meteorology took another step forward in the 1950s, when scientists converted the mathematical equations that describe the behavior of the atmosphere into software called numerical models that could be run on new highspeed computers. These calculations were the beginning of numerical weather prediction. Today, computers plot the observations, draw the lines on the map, and forecast the state of the atmosphere for some desired time in the future. Meteorologists evaluate the results from different numerical models and use them to issue public forecasts. After World War II, surplus military radars became available, and many were transformed into precipitation-measuring tools. In the mid-1990s, the National Weather Service replaced these conventional radars with the more sophisticated Doppler radars, which have the ability to peer into a

FIGURE 1.3 This satellite image shows the dynamic nature of FI the atmosphere as ribbons of water vapor (gray regions) swirl counterclockwise about huge storms over the North Pacific Ocean.

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A Satellite’s View of the Weather A good view of the weather can be obtained from a weather satellite. Figure 1.4 is a satellite image showing a portion of the Pacific Ocean and the North American continent. The image was obtained from a geostationary satellite situated about 36,000 km (22,300 mi) above Earth. At this elevation, the satellite travels at the same rate as Earth spins, which allows it to remain positioned above the same spot so it can continuously monitor what is taking place beneath it. The thin, solid black lines running from north-tosouth on the satellite image are called meridians, or lines of longitude. Since the zero meridian (or prime meridian) runs through Greenwich, England, the longitude of any place on Earth is simply how far it is in degrees from the prime meridian up to 180° east or west. North America is west of Great Britain and most of the United States lies between 75°W and 125°W longitude. The thin, solid black lines that parallel the equator are called parallels of latitude. The latitude of any place is how far north or south, in degrees, it is from the equator. The latitude of the equator is 0°, whereas the latitude of the North Pole is 90°N and that of the South Pole is 90°S. Most of the United States is located between latitude 30°N and 50°N, a region commonly referred to as the middle latitudes. Storms of All Sizes Probably the most prominent feature in Fig. 1.4 is the whirling cloud masses of all shapes and sizes. The clouds appear white because sunlight is rere flected back to space from their tops. The largest of the organized cloud masses are the sprawling storms. One such storm shows as an extensive band of clouds, over 2000 km long, west of the Great Lakes. Superimposed on the satellite image is the storm’s center (indicated by the large red L) and its adjoining weather fronts in red, blue, and purple. This middle-latitude cyclonic storm system (or extratropical cyclone) forms outside the tropics and, in the Northern Hemisphere, has winds spinning counterclockwise about its center, which is presently over Minnesota. A slightly smaller but more vigorous storm is located over the Pacific Ocean near latitude 12°N and longitude 116°W. This tropical storm system, with its swirling band of rotating clouds and sustained surface winds of 65 knots* (74 mi/hr) or more, is known as a hurricane. The diameter of the hurricane, as measured by the presence of winds of at least 34 knots (39 mi/hr), is about 800 km (500 mi). The tiny dot at its center is called the eye. Near the surface, in the eye, winds are light, skies are generally clear, and the atmospheric pressure is lowest. Around the eye, however, is an extensive region where heavy rain and high surface winds are reaching peak gusts of 100 knots. Smaller storms are seen as bright spots over the Gulf of Mexico. These spots represent clusters of towering *A knot is a nautical mile per hour where one knot equals 1.15 miles per hour (mi/hr) or 1.9 kilometers per hour (km/hr).

DID YOU KNOW? When it rains, it rains pennies from heaven—sometimes. On July 17, 1940, a tornado reportedly picked up a treasure of over 1000 sixteenth-century silver coins, carried them into a thunderstorm, then dropped them on the village of Merchery in the Gorki region of Russia.

cumulus clouds that have grown into thunderstorms, that is, tall churning clouds accompanied by lightning, thunder, strong gusty winds, and heavy rain. If you look closely at Fig. 1.4, you will see similar cloud forms in many regions. There were probably more than a thousand thunderstorms occurring throughout the world at that very moment. Although they cannot be seen individually, there are even some thunderstorms embedded in the cloud mass west of the Great Lakes. Later in the day on which this image was taken, a few of these storms spawned the most violent disturbance in the atmosphere, tornadoes. A tornado is an intense rotating column of air that usually extends downward from the base of a thunderstorm with a circulation reaching the ground. Sometimes called twisters, or cyclones, they may appear as ropes or as a large cylinder. They can be more than 2 km (1.2 mi) in diameter, although most are less than a football field wide. Most tornadoes have sustained winds below 100 knots, but some can pack winds exceeding 200 knots (230 mi/hr). Sometimes a visibly rotating funnel cloud dips part of the way down from a thunder thunderstorm, then rises without ever forming a tornado. A GLIMPSE AT A WEATHER MAP We can obtain a better picture of the middle-latitude storm system by examining a simplified surface weather map for the same day that the satellite image was taken. The weight of the air above different regions varies and, hence, so does the atmospheric pressure. In Fig. 1.5, the red letter L on the map indicates a region of low atmospheric pressure, often called a low, which marks the center of the middlelatitude cyclonic storm. (Compare the center of the storm in Fig. 1.5 with that in Fig. 1.4.) The two large blue letters H on the map represent regions of high atmospheric pressure, called highs, or anticyclones. The circles on the map represent other individual weather stations or cities where observations are taken. The wind is the horizontal movement of air. The wind direction—the direction from which the wind is blowing*—is given by wind barbs, lines that parallel the wind and extend outward from the center of the station. The wind speed—the rate at which the air is moving past a stationary observer—is indicated by flags, the short lines that extend off each wind barb. Notice how the wind blows around the highs and the lows. The horizontal pressure differences create a force *If you are facing north and the wind is blowing in your face, the wind would be called a “north wind.” EARTH’S ATMOSPHERE

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NOAA/National Weather Service

FIGURE 1.4 This FI satellite image (taken in visible reflected light) shows a variety of cloud patterns and storms in Earth’s atmosphere.

that starts the air moving from higher pressure toward lower pressure. Because of Earth’s rotation, the winds are deflected from their path toward the right in the Northern Hemisphere.* This deflection causes the winds in the Northern Hemisphere to blow clockwise and outward from the center of the highs, and counterclockwise and inward toward the center of the low. As the surface air spins into the low, it flows together and is forced upward, like toothpaste squeezed out of an upward-pointing tube. The rising air cools, and the water vapor in the air condenses into clouds. Notice on the weather map that the area of precipitation (the shaded green area) in the vicinity of the low corresponds to an extensive cloudy region in the satellite image of Fig. 1.4. Also notice by comparing Figs. 1.4 and 1.5 that, in the regions of high pressure, skies are generally clear. As the surface air flows outward away from the center of a high, air sinking from above must replace the laterally spreading surface air. Since sinking air does not usually produce *This deflecting force, known as the Coriolis force, is discussed more completely in Chapter 6, as are the winds.

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clouds, we find generally clear skies and fair weather associated with the regions of high atmospheric pressure. Areas of high and low pressure, and the swirling air around them, are the major weather producers for the middle latitudes. Look at the middle-latitude storm and the surface temperatures in Fig. 1.5 and notice that, to the southeast of the storm, southerly winds from the Gulf of Mexico are bringing warm, humid air northward over much of the southeastern portion of the nation. On the storm’s western side, cool dry northerly breezes combine with sinking air to create generally clear weather over the Rocky Mountains. The boundary that separates the warm and cool air appears as a heavy, colored line on the map—a front, across which there is a sharp change in temperature, humidity, and wind direction. Where the cool air from Canada replaces the warmer air from the Gulf of Mexico, a cold front is drawn in blue, with arrowheads showing the front’s general direction of movement. Where the warm Gulf air is replacing cooler air to the north, a warm front is drawn in red, with half circles showing its general direction of movement. Where the cold front has caught up to the warm front and cold air is now replacing cool air, an occluded front is drawn

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© Cengage Learning®.

FIGURE 1.5 Simplified surface weather map that roughly correlates with the satellite image shown in Fig. 1.4. Because FI this map shows conditions several hours after those in Fig. 1.4, the frontal system across the Midwest is farther east. The shaded green area represents precipitation. The numbers on the map represent air temperatures in °F.

in purple, with alternating arrowheads and half circles to show how it is moving. Along each of the fronts, warm air is rising, producing clouds and precipitation. In the satellite image (Fig. 1.4), the occluded front and the cold front appear as an elongated, curling cloud band that stretches from the low pressure area over Minnesota into the northern part of Texas. Notice in Fig. 1.5 that the frontal system is to the west of Chicago. As the westerly winds aloft push the front eastward, a person on the outskirts of Chicago might observe the approaching front as a line of towering thunderstorms similar to those shown in Fig. 1.6. On a Doppler radar image, the advancing thunderstorms might appear simisimi lar to those shown in Fig. 1.7. In a few hours, Chicago should experience heavy showers with thunder, lightning, and gusty winds as the front passes. All of this weather, however, should give way to clearing skies and surface winds from the west or northwest after the front has moved on by. Observing storm systems, we see that not only do they move but they also constantly change. Steered by the upper-level westerly winds, the middle-latitude storm in Fig. 1.5 gradually weakens and moves eastward, carrying

its clouds and weather with it. In advance of this system, a sunny day in Ohio will gradually cloud over and yield heavy showers and thunderstorms by nightfall. Behind the storm, cool dry northerly winds rushing into eastern Colorado cause an overcast sky to give way to clearing conditions. Farther south, the thunderstorms presently over the Gulf of Mexico in the satellite image (Fig. 1.4) expand a little, then dissipate as new storms appear over water and land areas. To the west, the hurricane over the Pacific Ocean drifts northwestward and encounters cooler water. Here, away from its warm energy source, it loses its punch; winds taper off, and the storm soon turns into an unorganized mass of clouds and tropical moisture. WEATHER AND CLIMATE IN OUR LIVES Weather and climate play a major role in our lives. Weather, for example, often dictates the type of clothing we wear, while climate influences the type of clothing we buy. Climate determines when to plant crops as well as what types of crops can be planted. Weather determines if these same crops will grow to maturity. Although weather and climate affect our lives in many ways, perhaps their most immediate effect is on our comfort. In order to survive EARTH’S ATMOSPHERE

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9

NOAA/National Weather Service

the cold of winter and heat of summer, we build homes, heat them, air condition them, insulate them—only to find that when we leave our shelter, we are at the mercy of the weather elements. Even when we are dressed correctly for the weather, wind, humidity, and precipitation can change our perception of how cold or warm it feels. On a cold, windy day the effects of wind chill tell us that it feels much colder than it really is, and, if we are not correctly dressed, we run the risk of frostbite or even hypothermia (the rapid, progressive mental and physical collapse that accompanies the lowering of human body temperature). On a hot, humid day we normally feel uncomfortably warm and blame it on the humidity. If we become too warm, our bodies overheat and heat exhaustion or heatstroke may result. Those most likely to suffer these maladies are the elderly with impaired circulatory systems and infants, whose heat regulatory mechanisms are not yet fully developed. Weather affects how we feel in other ways, too, not all of them well understood. Arthritic pain is most likely to occur when rising humidity is accompanied by falling pressures. The incidence of heart attacks shows a statistical peak after the passage of warm fronts, when rain and wind are common, and after the passage of cold fronts, when an abrupt change takes place as showery precipitation is accompanied by cold gusty winds. Headaches are common

on days when we are forced to squint, often because of hazy skies or a thin, bright overcast layer of high clouds. Some people who live near mountainous regions become irritable or depressed when there is a warm, dry wind blowing downslope (a chinook wind). Hot, dry downslope Santa Ana winds in Southern California can turn dry vegetation into a raging firestorm. When the weather turns much colder or warmer than normal, it impacts directly on the lives and pocketbooks of many people. For example, the exceptional record warmth observed from January to March 2012 over the United States saved people millions of dollars in heating costs. On the other side of the coin, the colder-than-normal winters of 2013–2014 and 2014–2015 over much of the northeastern United States sent heating costs soaring as demand for heating fuel escalated. Major cold spells accompanied by heavy snow and ice can play havoc by snarling commuter traffic, curtailing airport services, closing schools, and downing power lines, thereby cutting off electricity to thousands of customers (see Fig. 1.8). For example, a huge ice storm during January 1998 in northern New England and Canada left millions of people without power and caused over a billion dollars in damages; and a devastating snow storm during March 1993 buried parts of the East Coast with 14-foot snow drifts and left Syracuse, New York, paralyzed with a

© Robert Henson

FIGURE 1.6 Thunderstorms developing and advancing along FI an approaching cold front.

FIGURE 1.7 The elongated frontal system in Fig. 1.5 shows up FI on Doppler radar as a line of different colors. In this composite image, the areas shaded green and blue indicate where lightto-moderate rain is falling. Yellow indicates heavier rainfall. The red-shaded area represents the heaviest rainfall and the possibility of intense thunderstorms. Notice that a thunderstorm is approaching Chicago from the west.

10

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John Patriquin/Portland Press Herald via Getty Images

FIGURE 1.8 Utility workers in Maine clear off broken tree FI branches from power lines during a major ice storm on December 12, 2008.

snow depth of 36 inches. When the frigid air settles into the Deep South, many millions of dollars worth of temperature-sensitive fruits and vegetables may be ruined, the eventual consequence being higher produce prices for consumers. Prolonged drought, especially when accompanied by high temperatures, can lead to a shortage of food and, in some places, widespread starvation. Parts of Africa, for example, have periodically suffered through major droughts and famine. During the summer of 2012, much of the United States experienced a severe drought with searing summer temperatures and wilting crops, causing billions of dollars in crop losses. California experienced a destructive multi-year drought beginning in 2011. When the climate turns hot and dry, animals suffer too. In 1986, over 500,000 chickens perished in Georgia during a two-day period at the peak of a summer heat wave. Severe drought also has an effect on water reserves, often forcing communities to ration water and restrict its use. During periods of extended drought, vegetation often becomes tinder-dry and,

sparked by lightning or a careless human, such a dried-up region can quickly become a raging inferno. During the winter of 2005–2006, hundreds of thousands of acres in drought-stricken Oklahoma and northern Texas were ravaged by wildfires. Every summer, scorching heat waves take many lives. During the past twenty years, an annual average of more than 300 deaths in the United States were attributed to excessive heat exposure. In one particularly devastating heat wave that hit Chicago, Illinois, during July 1995, high temperatures coupled with high humidity claimed the lives of more than 700 people. In California, during July 2006, more than 100 people died during a two-week period as air temperatures climbed to over 46°C (115°F). Heat waves have been especially devastating in recent years across Europe, where many cities and buildings are not designed for intense heat. In the summer of 2003, tens of thousands died across Europe, including 14,000 in France alone. A record-breaking heat wave in Russia in 2010 killed nearly 11,000 people in Moscow. Each year, the violent side of weather influences the lives of millions. Those who live along the Gulf and Atlantic coastlines keep a close watch for hurricanes during the late summer and early autumn. These large tropical systems can be among the nation’s most destructive weather events. More than 250,000 people lost their homes when Hurricane Andrew struck the Miami area in 1992, and nearly 2000 people along the central Gulf Coast were killed by Hurricane Katrina in 2005. A late-season hurricane called Sandy took a rare path in October 2012, striking the mid-Atlantic coast from the southeast. Although it was no longer classified as a hurricane when it came ashore, Sandy produced wind gusts topping 80 miles per hour in some areas. Because of its vast size and unusual path, Sandy pushed a huge amount of water into the coast, resulting in catastrophic storm-surge flooding and more than 100 deaths across parts of New Jersey, New York, and New England (see Fig. 1.9).

Andrea Booher/FEMA

FIGURE 1.9 A resident of Long Beach, FI New York, digs sand out around his car after Hurricane Sandy, during October 2012, pushed water and sand far inland, destroying homes, commercial businesses, and approximately 10,000 cars in this area alone.

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11

AP Photo/Mark Schiefelbein

AP Photo/Mark Schiefelbein

FIGURE 1.10 Lightning flashes inside a violent tornado that FI tore through Joplin, Missouri, on May 22, 2011. The tornado ripped through a hospital and destroyed entire neighborhoods. (See tornado damage in Fig. 1.11.)

FIGURE 1.11 Emergency personnel walk through a neighborFI hood in Joplin, Missouri, damaged by a violent tornado, with winds exceeding 174 knots (200 mi/hr), on May 22, 2011. The tornado caused hundreds of millions of dollars in damage and took 159 lives, making this single tornado the deadliest in the United States since 1947.

AP Photo/David J Phillip

FIGURE 1.12 Residents evacuFI ate an apartment complex in the Houston area on April 18, 2016, as torrential rains from severe thunderstorms produced severe flash flooding. More than 1000 highwater rescues took place.

It is amazing how many people whose family roots are in the Midwest know the story of someone who was severely injured or killed by a tornado. Tornadoes have not only taken many lives, but annually they cause damage to buildings and property totaling in the hundreds of millions of dollars, as a single large tornado can level an entire section of a town (see Fig. 1.10 and Fig. 1.11). Although the gentle rains of a typical summer thunthun derstorm are welcome over much of North America, the heavy downpours, high winds, and large hail of severe thunderstorms are not. Cloudbursts from slowly moving, intense thunderstorms can provide too much rain too quickly, creating flash floods as small streams become raging rivers composed of mud and sand entangled with uprooted plants and trees. Thunderstorms dumped up to 20 inches of rain in just a few hours over parts of the Houston area in April 2016, leading to severe flash flooding (see Fig. 1.12). If heavy rain covers a large area, devastating river floods can result. Record rainfall produced both types of flooding over the Southern Plains 12

in May 2015 and across South Carolina in October 2015. On the average, more people die in the United States from river floods and flash floods than from either lightning strikes or tornadoes. Strong downdrafts originating inside an intense thunderstorm (a downburst downburst) create turbulent winds that are capable of destroying crops and inflicting damage upon surface structures. Hundreds of people were killed in airline crashes attributed to turbulent wind shear (a rapid change in wind speed, wind direction, or both) from downbursts, before a safety system was implemented in the 1990s, which also has improved since that time. Annually, hail damages crops worth millions of dollars,

DID YOU KNOW? The folks of Elgin, Manitoba, literally had their “goose cooked” during April 1932, when a lightning bolt killed 52 geese that were flying overhead in formation. As the birds fell to the ground, they were reportedly gathered up and distributed to the townspeople for dinner.

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TBD

FIGURE 1.13 Estimates are that lightning strikes Earth about FI 40 to 50 times every second. More than 20 million lightning strikes hit the United States in a typical year. Here, lightning strikes the ground near wind turbines in Texas.

and lightning takes the lives of several dozen people in the United States and starts fires that destroy many thousands of acres of valuable timber (see Fig. 1.13). Up to this point, we have considered the more violent side of weather and its impact on humanity. Weather- and climate-related events can have enormous economic consequences. On average, tens of billions of dollars in property damage occur each year in the United States alone (see Fig. 1.14). Yet even the quiet side of weather has its influence. When winds die down and humid air becomes more tranquil, fog may form. Heavy fog can restrict visibility at airports, causing flight delays and cancellations. Every winter, deadly fog-related auto accidents occur along our busy highways and turnpikes. But fog has a positive side, too, especially during a dry spell, as fog moisture collects on tree branches and drips to the ground, where it provides water for the tree’s root system.

Weather and climate have become so much a part of our lives that the first thing many of us do in the morning is to listen to the local weather forecast or look it up on our smartphones. For this reason, many radio and most television newscasts have their own “weatherperson” to present weather information and give daily forecasts. More and more of these people are professionally trained in meteorology, and many stations require that the weathercaster be certified by the American Meteorological Society (AMS) or hold a seal of approval from the National Weather Association (NWA). To make their weather presentations as up-to-the-minute as possible, weathercasters draw upon time-lapse satellite images, Doppler radar displays, and other ways of illustrating current weather. Many stations work with private firms that create graphics and customized forecasts, largely based on observations and computer models from the National Weather Service (NWS). Since 1982, a staff of trained professionals at The Weather Channel have provided weather information twenty-four hours a day on cable television. (Many viewers believe the weatherperson they see on TV is a meteorologist and that all meteorologists forecast the weather. If you are interested in learning what a meteorologist or atmospheric scientist is and what he or she might do for a living other than forecast the weather, read Focus section 1.1.) The National Oceanic and Atmospheric Administration (NOAA), in cooperation with the National Weather Service, sponsors weather radio broadcasts at selected locations across the United States. Known as NOAA Weather Radio (and transmitted at VHF-FM frequencies), this service provides continuous weather information and regional forecasts (as well as special weather advisories, including watches and warnings) for over 90 percent of the United States. Although millions of people rely on weather broadcasts, many use forecasts obtained on their smartphones or on the internet. Websites operated by the NWS and private forecasting companies provide a wealth of local, national, and global data and forecasts. Smartphone applications

National Centers for Environmental Information (NCEI)

FIGURE 1.14 The number of billion-dollar FI weather and climate events in the United States from 1980 to 2014 (red bar). The total cost (green bar) of the 178 events during this period exceeded $1 trillon. This total is adjusted to the 2014 consumer price index and includes insured and uninsured losses.

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13

FOCUS

ON A SP SPECIAL TOPIC 1.1

Most people associate the term “meteorologist” with the weatherperson they see on television or hear on the radio. Many television and radio weathercasters are in fact professional meteorologists, but some are not. A professional meteorologist is usually considered to be a person who has completed the requirements for a college degree in meteorology or atmospheric science. This individual has strong, fundamental knowledge concerning how the atmosphere behaves, along with a substantial background of coursework in mathematics, physics, and chemistry. A meteorologist uses scientific principles to explain and to forecast atmospheric phenomena. About half of the approximately 9000 meteorologists and atmospheric scientists in the United States work doing weather forecasting for the National Weather Service, the military, or for a television or radio station. The other half work mainly in research, teach atmospheric science courses in colleges and universities, or do meteorological consulting work. Scientists who do atmospheric research may be investigating how the climate is changing, how snowflakes form, or how pollution impacts temperature patterns. Aided by supercomputers, much of the work of a research meteorologist involves simulating the atmosphere to see how it behaves (see Fig. 1). Researchers often work closely with such scientists as chemists, physicists, oceanographers, mathematicians, and environmental scientists to determine how the

NCAR/UCAR/NSF

What Is a Meteorologist?

FIGURE 1 A model that simulates a three-dimensional view of FI the atmosphere. This computer model predicts how winds and clouds over the United States will change with time.

atmosphere interacts with the entire ecosystem. Scientists doing work in physical meteorology may well study how radiant energy warms the atmosphere; those at work in the field of dynamic meteorology might be using the mathematical equations that describe airflow to learn more about jet streams. Scientists working in operational meteorology might be preparing a weather forecast by analyzing upper-air information over North America. A climatologist, or climate scientist scientist, might be studying the interaction of the atmosphere and ocean to see what influence such interchange might have on planet Earth many years from now.

can be tailored to provide conditions and forecasts for your hometown or wherever you may be traveling.

BRIEF REVIEW

On any given day, a wide variety of storms exist on Earth, ranging in size from the very large middle latitude cyclonic storm to the much smaller tornado.

Wind direction is the direction from which the wind is blowing.

In the Northern Hemisphere, winds around an area of surface low pressure blow counter clockwise and inward; around an area of a surface high pressure, they blow clockwise and outward.

Weather and climate impact our lives in many ways. Droughts, floods, heat and cold waves, and violent weather events can cause much suffering and inflict billions of dollars in damage.

In the last section, we looked at many ways in which weather and climate impact our lives. A few of the main points described up to now are: ●

Our understanding of weather and climate is built on knowledge acquired and applied through the scientific method, which allows us to make informed predictions about the natural world.

Weather, the state of the atmosphere at any particular time and place, is composed of the weather elements—air temperature, air pressure, humidity, clouds, precipitation, visibility, and wind.

Climate represents the accumulation of daily and seasonal weather and its extremes over an extended period of time.

Meteorology is the study of the atmosphere and its phenomena.

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Meteorologists also provide a variety of services not only to the general public in the form of weather forecasts but also to city planners, contractors, farmers, and large corporations. Meteorologists working for private weather firms create the forecasts and graphics that are found in newspapers, on television, and on the internet. Overall, there are many exciting jobs that fall under the heading of “meteorologist”—too many to mention here. However, for more information on this topic, visit this website: http://www. ametsoc.org/ and click on “Students.”

Having looked at the many ways that weather can affect our lives, we will now turn to the atmosphere that produces all of the weather we experience and examine its content and structure more closely. Many of the concepts discussed here will be covered in more detail in later chapters.

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Components of Earth’s Atmosphere The Earth’s atmosphere is a relatively thin, gaseous envelope comprised mostly of nitrogen (N2) and oxygen (O2), with small amounts of other gases, such as water vapor (H2O) and carbon dioxide (CO2). Nestled in the atmosphere are clouds of liquid water and ice crystals. Although our atmosphere extends upward for many hundreds of kilometers (km), it gets progressively thinner with increasing altitude. Almost 99 percent of the atmosphere lies within a mere 30 km (about 19 mi) of Earth’s surface (see Fig. 1.15). In fact, if Earth were to shrink to the size of a large beach ball, its inhabitable atmosphere would be thinner than a piece of paper. This thin blanket of air constantly shields the surface and its inhabitants from the sun’s dangerous ultraviolet radiant energy, as well as from the onslaught of material from interplanetary space. There is no definite upper limit to the atmosphere; rather, it becomes thinner and thinner, eventually merging with empty space, which surrounds all the planets. THE EARLY ATMOSPHERE The atmosphere that originally surrounded Earth was probably much different from the air we breathe today. The Earth’s first atmosphere (some 4.6 billion years ago) was most likely hydrogen and helium—the two most abundant gases found in the universe—as well as hydrogen compounds, such as methane (CH4) and ammonia (NH3). Most scientists believe that this early atmosphere escaped into space from Earth’s hot surface. A second, more-dense atmosphere, however, gradually enveloped Earth as gases from molten rock within its hot interior escaped through volcanoes and steam vents. We assume that volcanoes spewed out the same gases then as they do today: mostly water vapor (about 80 percent), carbon dioxide (about 10 percent), and up to a few percent nitrogen. These gases probably created Earth’s second atmosphere.

As millions of years passed, the constant outpouring of gases from the hot interior—known as outgassing— provided a rich supply of water vapor. Moreover, when Earth was very young, some of its water may have originated from numerous collisions with small meteors that pounded Earth, as well as from disintegrating comets. The water vapor condensed into clouds and rain fell upon Earth for many thousands of years, forming the rivers, lakes, and oceans of the world. During this time, large amounts of carbon dioxide (CO2) were dissolved in the oceans. Through chemical and biological processes, much of the CO2 became locked up in carbonate sedimentary rocks, such as limestone. With much of the water vapor already condensed and the concentration of CO2 dwindling, the atmosphere gradually became dominated by molecular nitrogen (N2), which is usually not chemically active. It appears that molecular oxygen (O2), the second most abundant gas in today’s atmosphere, probably began an extremely slow increase in concentration as energetic rays from the sun split water vapor (H2O) into hydrogen and oxygen. The hydrogen, being lighter, probably rose and escaped into space, while the oxygen remained in the atmosphere. We are uncertain whether this slow increase in oxygen supported the evolution of primitive plants, perhaps two to three billion years ago, or if the plants evolved in an almost oxygen-free (anaerobic) environment. At any rate, plant growth greatly enriched our atmosphere with oxygen. The reason for this enrichment is that, during the process of photosynthesis, plants, in the presence of sunlight, combine carbon dioxide and water to produce sugar and oxygen. Hence, after plants evolved, the atmospheric oxygen content increased more rapidly, probably reaching its present composition about several hundred million years ago. COMPOSITION OF TODAY’S ATMOSPHERE ▼Table 1.1 shows the various gases present in a volume of air near Earth’s surface. Notice that molecular nitrogen (N2) occupies about 78 percent and molecular oxygen (O2) about

NASA/JSC

FIGURE 1.15 The Earth’s FI atmosphere as viewed from space. The atmosphere is the thin bluishwhite region along the edge of Earth. The photo was taken from the International Space Station on April 12, 2011, over western South America.

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15

▼ Table 1.1

Composition of the Atmosphere near the Earth’s Surface PERMANENT GASES

Gas

Symbol

VARIABLE GASES

Percent (by Volume) Dry Air

Gas (and Particles)

Symbol

Percent (by Volume)

Parts per Million (ppm)*

Nitrogen

N2

78.08

Water vapor

H2O

0 to 4

Oxygen

O2

20.95

Carbon dioxide

CO2

0.0405

405*

Argon

Ar

0.93

Methane

CH4

0.00018

1.8

Neon

Ne

0.0018

Nitrous oxide

N2O

0.00003

0.3

Helium

He

0.0005

Ozone

O3

0.000004

0.04**

Hydrogen

H2

0.00006

Particles (dust, soot, etc.)

0.000001

0.01–0.15

Xenon

Xe

0.000009

Chlorofluorocarbons (CFCs) and hydrofluorocarbons (HCFCs)

0.00000001

0.0001

*For CO2, 405 parts per million means that out of every million air molecules, 405 are CO2 molecules. **Stratospheric values at altitudes between 11 km and 50 km are about 5 to 12 ppm.

© UCAR

21 percent of the total volume of dry air. If all the other gases are removed, these percentages for nitrogen and oxygen hold fairly constant up to an elevation of about 80 km (or 50 mi). At the surface, there is a balance between destruction (output) and production (input) of these gases. For example, nitrogen is removed from the atmosphere primarily by biological processes that involve soil bacteria. Nitrogen is also taken from the air by tiny ocean-dwelling plankton that convert it into nutrients that help fortify the

FIGURE 1.16 Earth’s atmosphere is a rich mixture of many FI gases, with clouds of condensed water vapor and ice crystals. Here, water evaporates from the ocean’s surface. Rising air currents then transform the invisible water vapor into many billions of tiny liquid droplets that appear as puffy cumulus clouds. If the rising air in the cloud should extend to greater heights, where air temperatures are quite low, some of the liquid droplets would freeze into minute ice crystals.

16

ocean’s food chain. It is returned to the atmosphere mainly through the decaying of plant and animal matter. Oxygen, on the other hand, is removed from the atmosphere when organic matter decays and when oxygen combines with other substances, producing oxides. It is also taken from the atmosphere during breathing, as the lungs take in oxygen and release carbon dioxide. The addition of oxygen to the atmosphere occurs during photosynthesis.The concentration of the invisible gas water vapor, however, varies greatly from place to place, and from time to time. Close to the surface in warm, steamy, tropical locations, water vapor may account for up to 4 percent of the atmospheric gases, whereas in colder arctic areas, its concentration may dwindle to a mere fraction of a percent. Water vapor molecules are, of course, invisible. They become visible only when they transform into larger liquid or solid particles, such as cloud droplets and ice crystals, which may grow in size and eventually fall to Earth as rain or snow. The changing of water vapor into liquid water is called condensation, whereas the process of liquid water becoming water vapor is called evaporation. In the lower atmosphere, water is everywhere. It is the only substance that exists as a gas, a liquid, and a solid at those temperatures and pressures normally found near Earth’s surface (see Fig. 1.16). Water vapor is an extremely important gas in our atmosphere. Not only does it form into both liquid and solid cloud particles that grow in size and fall to Earth as precipitation, but it also releases large amounts of heat— called latent heat—when it changes from vapor into liquid water or ice. Latent heat is an important source of atmospheric energy, especially for storms, such as thunderstorms and hurricanes. Moreover, water vapor is a potent greenhouse gas because it strongly absorbs a portion of Earth’s outgoing radiant energy (somewhat like the glass of a greenhouse prevents the heat inside from escaping

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© Cengage Learning®.

*A more detailed look at the greenhouse effect is presented in Chapter 2.

© Cengage Learning®.

and mixing with the outside air). This trapping of heat energy close to Earth’s surface, called the greenhouse effect, keeps the average air temperature near the surface much warmer than it would be otherwise. Thus, water vapor plays a significant role in Earth’s heat-energy balance.* Carbon dioxide (CO2), a natural component of the atmosphere, occupies a small but important percent of a volume of air, about 0.04 percent. Carbon dioxide enters the atmosphere mainly from the decay of vegetation, but it also comes from volcanic eruptions, the exhalations of animal life, the burning of fossil fuels (such as coal, gasoline, and natural gas), and deforestation. The removal of CO2 from the atmosphere takes place during photosynthesis, as plants consume CO2 to produce green matter. The CO2 is then stored in roots, branches, and leaves. Rain and snow can react with silicate minerals in rocks and remove CO2from the atmosphere through a process known as chemical weathering weathering. The oceans act as a huge reservoir for CO2, as phytoplankton (tiny drifting plants) in surface water fix CO2 into organic tissues. Carbon dioxide that dissolves directly into surface water mixes downward and circulates through greater depths. Estimates are that the oceans hold more than 50 times the im total atmospheric CO2 content. Figure 1.17 illustrates important ways carbon dioxide enters and leaves the atmosphere. Figure 1.18 reveals that the atmospheric concenconcen tration of CO2 has risen by almost 30 percent since 1958, when it was first measured at Mauna Loa Observatory in Hawaii. This increase means that CO2 is entering the atmosphere at a greater rate than it is being removed. The

FIGURE 1.17 The main components of the atmospheric carbon FI dioxide cycle. The gray lines show processes that put carbon dioxide into the atmosphere, whereas the red lines show processes that remove carbon dioxide from the atmosphere.

increase appears to be owing mainly to the burning of fossil fuels, such as coal and oil; about half of the CO2 emitted from fossil fuels remains in the atmosphere, whereas the rest enters the ocean, soil, and plants. Deforestation also plays a role, as cut timber, burned or left to rot, releases CO2 directly into the air, and the trees can no longer remove CO2 from the atmosphere. Deforestation accounts for perhaps 10 to 15 percent of the observed CO2 increase FIGURE 1.18 (a) The FI solid blue line shows the average yearly measurements of CO2 in parts per million (ppm) at Mauna Loa Observatory, Hawaii, from 1958 through 2015. The jagged dark line illustrates how higher readings occur in winter when plants die and release CO2 to the atmosphere, and how lower readings occur in summer when more abundant vegetation absorbs CO2 from the atmosphere. (b) The insert shows CO2 values in ppm during the past 1000 years from ice cores in Antarctica (orange line) and from Mauna Loa Observatory (blue line). (Mauna Loa data NOAA; Ice Core data courtesy of Carbon Dioxide Information Analysis Center, Oak Ridge National Laboratory) EARTH’S ATMOSPHERE

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17

in recent years. Measurements of CO2 also come from ice cores. In Greenland and Antarctica, for example, tiny bubbles of air trapped within the ice sheets reveal that before the industrial revolution, CO2 levels were stable at about 280 parts per million (ppm). (See the insert in Fig. 1.18.) Since the early 1800s, however, CO2 levels have increased by more than 40 percent. With CO2 levels presently increasing by approximately 0.5 percent annually (around 2.0 ppm/year), scientists now estimate that the concentration of CO2 will likely increase from its current value of about 405 ppm to a value perhaps exceeding 750 ppm by the end of this century, assuming that fossil fuel emissions continue at or beyond current levels. Like water vapor, carbon dioxide is an important greenhouse gas that traps a portion of Earth’s outgoing energy. Consequently, with everything else being equal, as the atmospheric concentration of CO2 increases, so should the average global surface air temperature. In fact, over the last 110 years or so, Earth’s average surface temperature has warmed by about 1.0°C (1.8°F). Mathematical climate models, which predict future atmospheric conditions, estimate that if concentrations of CO2 (and other greenhouse gases) continue to increase at or beyond their present rates, Earth’s surface could warm by an additional 3°C (5.4°F) or more by the end of this century. As we will see in Chapter 13, the consequences of this type of climate change (such as rising sea levels and the rapid melting of polar ice) will be felt worldwide. Carbon dioxide and water vapor are not the only greenhouse gases. Others include methane (CH4), nitrous oxide (N2O), and chlorofluorocarbons (CFCs). On average, methane concentrations have risen about one-half of one percent per year since the 1990s, but the pace has been uneven for reasons now being studied. Most methane appears to derive from the breakdown of plant material by certain bacteria in rice paddies, wet oxygen-poor soil, the biological activity of termites, and biochemical reactions in the stomachs of cows, although some methane is also leaked into the atmosphere by natural gas operations. Levels of nitrous oxide— commonly known as laughing gas—have been rising annually at the rate of about one-quarter of a percent. As well as being an industrial by-product, nitrous oxide forms in the soil through a chemical process involving bacteria and certain microbes; it is also produced by fossil fuel burning and other activities. Ultraviolet light from the sun destroys it. Chlorofluorocarbons represent a group of greenhouse gases that, up until the mid-1990s, had been increasing in concentration. At one time, they were the most widely used propellants in spray cans. More recently, they have been used as refrigerants, as propellants for the blowing of plastic-foam insulation, and as solvents for cleaning electronic microcircuits. Although their average concentration in a volume of air is quite small (see Table 1.1, p. 16), they have an important effect on our atmosphere: 18

They not only trap heat as greenhouse gases but also play a part in destroying the gas ozone in the stratosphere, a region in the atmosphere located between about 11 km and 50 km above Earth’s surface. By international law, chlorofluorocarbons are gradually being replaced by other compounds, such as hydrochlorofluorocarbons, which are far less harmful to the ozone layer even though they are still greenhouse gases. On Earth’s surface, ozone (O3) is the primary ingredient of photochemical smog,* which irritates the eyes and throat and damages vegetation. But the majority of atmospheric ozone (about 97 percent) is found in the upper atmosphere—called the stratosphere—where it is formed naturally, as oxygen atoms combine with oxygen molecules. Here, the concentration of ozone averages less than 0.002 percent by volume. This small quantity is important, however, because it shields plants, animals, and humans from the sun’s harmful ultraviolet rays. It is ironic that ozone, which damages plant life in a polluted environment, provides a natural protective shield in the upper atmosphere so that plants on the surface may survive. When CFCs enter the stratosphere, ultraviolet rays break them apart, and the CFCs release ozone-destroying chlorine. Because of this effect, ozone concentration in the stratosphere has decreased over parts of the Northern and Southern Hemispheres in recent decades. Stratospheric ozone levels over springtime Antarctica plummet each year during September and October, to the point where so little ozone is observed that a seasonal ozone hole forms situa (see Fig. 1.19). (We will examine the ozone hole situation, as well as photochemical ozone, in Chapter 14.) Impurities from both natural and human sources are also present in the atmosphere: Wind picks up dust and soil from Earth’s surface and carries it aloft; small saltwater drops from ocean waves are swept into the air (upon evaporating, these drops leave microscopic salt particles suspended in the atmosphere); smoke from forest fires is often carried high above Earth; and volcanoes spew many tons of fine ash particles and gases into the air (see Fig. 1.20). Collectively, these tiny solid or liquid particles of various composition, suspended in the air, are called aerosols. Some natural impurities found in the atmosphere are quite beneficial. Small, floating particles, for instance, act as surfaces on which water vapor condenses to form clouds. However, most human-made impurities (and some natural ones) are a nuisance as well as a health hazard. These we call pollutants. For example, many automobile engines (especially older ones) emit copious amounts of nitrogen dioxide (NO2), carbon monoxide (CO), and hydrocarbons. In sunlight, nitrogen dioxide *Originally the word smog meant the combining of smoke and fog. Today, however, the word usually refers to the type of smog that forms in large cities, such as Los Angeles, California. Because this type of smog forms when chemical reactions take place in the presence of sunlight, it is termed photochemical smog smog.

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DID YOU KNOW? The air density in the mile-high city of Denver, Colorado, is normally about 15 percent less than the air density at sea level. As the air density decreases, the drag force on a baseball in flight also decreases. Because of this fact, a baseball hit at Denver’s Coors Field will travel farther than one hit near sea level.

droplets that are small enough to remain suspended in the air. These particles can obscure visibility and cause respiratory and cardiovascular problems. (More information on these and other pollutants is given in Chapter 14.)

BRIEF REVIEW

NASA

FIGURE 1.19 The darkest color represents the area of FI lowest ozone concentration, or ozone hole, over the Southern Hemisphere on October 2, 2015. Notice that the hole is larger than the continent of Antarctica. A Dobson Unit (DU) is the physical thickness of the ozone layer if it were brought to Earth’s surface, where 500 DU equals 5 millimeters.

reacts with hydrocarbons and other gases to produce surface ozone. Carbon monoxide is a major pollutant of city air. Colorless and odorless, this poisonous gas forms during the incomplete combustion of carbon-containing fuel. Hence, more than half of carbon monoxide in urban areas comes from road vehicles. The burning of sulfur-containing fuels (such as coal and oil) releases sulfur gases into the air. When the atmosphere is sufficiently moist, these gases may transform into tiny dilute drops of sulfuric acid. Rain containing sulfuric acid corrodes metals and painted surfaces and turns freshwater lakes acidic. Acid rain (thoroughly discussed in Chapter 14) is a major environmental problem, especially downwind from major industrial areas. Even the tiniest pollutants are a major concern. Particulate matter refers to solid particles and liquid

Before going on to the next several sections, here is a review of some of the important concepts presented so far: ●

The Earth’s atmosphere is a mixture of many gases. In a volume of dry air near the surface, nitrogen (N2) occupies about 78 percent and oxygen (O2) about 21 percent.

Water vapor, which normally occupies less than 4 percent of a volume of air near the surface, can condense into liquid cloud droplets or transform into delicate ice crystals. Water is the only substance in our atmosphere that is found naturally as a gas (water vapor), as a liquid (water), and as a solid (ice).

The majority of water on our planet is believed to have come from its hot interior through outgassing, although some of Earth’s water may have come from collisions with meteors and comets.

Both water vapor and carbon dioxide (CO2) are important greenhouse gases.

Ozone (O3) in the stratosphere protects life from harmful ultraviolet (UV) radiation. At the surface, ozone is the main ingredient of photochemical smog.

Vertical Structure of the Atmosphere

© David Weintraub/Photo Researchers

When we examine the atmosphere in the vertical, we see that it can be divided into a series of layers. Each layer can be defined in a number of ways: by the manner in which the air temperature varies through it, by the gases that comprise it, or even by its electrical properties. At any rate, before we examine these various atmospheric layers, we need to look at the vertical profile of two important atmospheric variables: air pressure and air density.

FIGURE 1.20 Erupting volcanoes can send tons of particles FI into the atmosphere, along with vast amounts of water vapor, carbon dioxide, and sulfur dioxide.

A BRIEF LOOK AT AIR PRESSURE AND AIR DENSITY Air molecules (as well as everything else) are held near Earth by gravity. This strong, invisible force pulling down on the air squeezes (compresses) air molecules closer together, which causes their number in a given volume to increase. The more air above a level, the greater the squeezing effect or compression. Since air density is the EARTH’S ATMOSPHERE

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it becomes more dense, the air weighs more, and the surface pressure goes up. On the other hand, when fewer molecules are in the column, the air weighs less, and the surface pressure goes down. A change in air density can bring about a change in air pressure. Pounds per square inch is, of course, just one way to express air pressure. Presently, the most common unit for air pressure found on surface weather maps is the millibar (mb), although the metric equivalent, the hectopascal* (hPa), is gradually replacing the millibar as the preferred unit of pressure on surface maps. A more traditional unit of pressure is inches of mercury (Hg), which is commonly used both in the field of aviation and in weather reports on television, radio, smartphones, and the Internet. At sea level, the standard value for atmospheric pressure is: 1013.25 mb = 1013.25 hPa = 29.92 in. Hg.

FIGURE 1.21 Both air pressure and air density decrease with FI increasing altitude. The weight of all the air molecules above Earth’s surface produces an average pressure near 14.7 lb/in2.

Weather conditions may cause the atmospheric pressure to vary from the standard value by 30 millibars or more at a given station. Figure 1.22 illustrates how rapidly air pressure dede creases with height. Near sea level, atmospheric pressure decreases rapidly, whereas at high levels it decreases more slowly. With a sea-level pressure near 1000 mb, we can see in Fig. 1.22 that, at an altitude of only 5.5 km (or 3.5 mi), the air pressure is about 500 mb, or half of the sea-level pressure. This situation means that, if you were at a mere 5.5 km (which is about 18,000 feet) above the surface, you would be above one-half of all the molecules in the atmosphere. At an elevation approaching the summit of Mount Everest (about 9 km, or 29,000 ft), the air pressure would

number of air molecules in a given space (volume), it follows that air density is greatest at the surface and decreases as we move up into the atmosphere.* Notice in Fig. 1.21 that, owing to the fact that the air near the surface is compressed, air density normally decreases rapidly at first, then more slowly as we move farther away from the surface. Air molecules have weight.** In fact, air is surprisingly heavy. The weight of all the air around Earth is a staggering 5600 trillion tons. The weight of the air molecules acts as a force upon Earth. The amount of force exerted over an area of surface is called atmospheric pressure or, simply, air pressure.† The pressure at any level in the atmosphere can be measured in terms of the total mass of the air above any point. As we climb in elevation, fewer air molecules are above us; hence, atmospheric pressure always decreases with increasing height. Like air density, air pressure decreases rapidly at first, then more slowly at higher levels, as illustrated in Fig. 1.21. If in Fig. 1.21 we weigh a column of air one square inch wide, extending from the average height of the ocean surface (sea level) to the “top” of the atmosphere, it would weigh very nearly 14.7 pounds. Thus, normal atmospheric pressure near sea level is close to 14.7 pounds per square inch (lb/in2). If more molecules are packed into the column,

*One hectopascal equals 1 millibar.

**The weight of an object, including air, is the force acting on the object due to gravity. In fact, weight is defined as the mass of an object times the acceleration of gravity. An object’s mass is the quantity of matter in the object. Consequently, the mass of air in a rigid container is the same everywhere in the universe. However, if you were to instantly travel to the moon, where the acceleration of gravity is much less than that of Earth, the mass of air in the container would be the same, but its weight would decrease. †Because air pressure is measured with an instrument called a barometer, atmospheric pressure is often referred to as barometric pressure.

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*Density is defined as the mass of air in a given volume of air. Density = mass/ volume.

FIGURE 1.22 Atmospheric pressure decreases rapidly with height. FI Climbing to an altitude of only 5.5 km, where the pressure is 500 mb, would put you above one-half of the atmosphere’s molecules.

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FIGURE 1.23 Layers of the atmosphere as FI related to the average profile of air tempera temperature above Earth’s surface. The heavy line illustrates how the average temperature varies in each layer.

be about 300 mb. The summit is above nearly 70 percent of all the molecules in the atmosphere. At an altitude of about 50 km (160,000 feet), the air pressure is about 1 mb, which means that 99.9 percent of all the air molecules are below this level. Yet the atmosphere extends upwards for many hundreds of kilometers, gradually becoming thinner and thinner until it ultimately merges with outer space. LAYERS OF THE ATMOSPHERE Up to this point, we’ve looked at how both air pressure and air density decrease with height above Earth—rapidly at first, then more slowly. Air temperature, however, has a more complicated vertical profile.* tem Look closely at Fig. 1.23 and notice that air temperature normally decreases from Earth’s surface up to an altitude of about 11 km, which is nearly 36,000 ft, or 7 mi. This decrease in air temperature with increasing height is due primarily to the fact that sunlight warms Earth’s surface, and the surface, in turn, warms the air above it (investigated further in Chapter 2). The rate at which the air temperature decreases with height is called the *Air temperature is the degree of hotness or coldness of the air and, as we will see in Chapter 2, it is also a measure of the average speed of the air molecules.

temperature lapse rate. The average (or standard) lapse rate in this region of the lower atmosphere is about 6.5 degrees Celsius (°C) for every 1000 meters (m) or about 3.6 degrees Fahrenheit (°F) for every 1000-ft increase in altitude (see Fig. 1.24). Keep in mind that these values are only averages. On some days, the air becomes colder more quickly as we move upward, which would increase or steepen the lapse rate. On other days, the air temperature would decrease more slowly with height, and the lapse rate would be less. Occasionally, the air temperature may actually increase with height, producing a condition known as a temperature inversion. Thus, the lapse rate fluctuates, varying from day to day, season to season, and place to place. The instrument that measures the vertical profile of air temperature in the atmosphere up to an altitude sometimes exceeding 30 km (100,000 ft) is the radiosonde. More information on this instrument is given in Focus section 1.2. Fig. 1.23 shows the region of the atmosphere from the surface up to about 11 km, which contains all of the weather we are familiar with on Earth. Also, this region is kept well stirred by rising and descending air currents, and it is common for air molecules to circulate through a depth of more EARTH’S ATMOSPHERE

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FOCUS

ON AN O OBSERVATION 1.2

The vertical distribution of temperature, pressure, and humidity up to an altitude of about 30 km (about 19 mi) can be obob tained with an instrument called a radiosonde.* The radiosonde is a small, lightsonde weight box equipped with weather instruments and a radio transmitter. It is attached to a cord that has a parachute and a gas-filled balloon tied tightly at the end (see Fig. 2). As the balloon rises, the attached radiosonde measures air tempertemper ature with a small electrical thermometer— a thermistor— located just outside the box. The radiosonde measures humidity electrically by sending an electric current across a carbon-coated plate. Air pressure is obtained by a small barometer located inside the box. All of this information is transmitted to the surface by radio. Here, a computer rapidly reconverts the various frequencies into values of temperature, pressure, and moisture. *A radiosonde that is dropped by parachute from an aircraft is called a dropsonde.

Special tracking equipment at the surface may be used to provide a vertical profile of winds. Radiosondes are equipped with Global Positioning System (GPS) receivers that lead to highly accurate wind computations. (When winds are added, the observation is called a rawinsonde.) When plotted on a graph, the vertical distribution of temperature, humidity, and wind is called a sounding. Eventually, the balloon bursts and the radiosonde returns to Earth, its descent being slowed by its parachute. At most sites, radiosondes are released twice a day, usually at the time that corresponds to midnight and noon in Greenwich, England. Releasing radiosondes is an expensive operation because many of the instruments are never retrieved, and often many of those that are retrieved are often in poor working condition. To complement the radiosonde, modern satellites (using instruments that measure radiant energy) are providing scientists with vertical temperature profiles in inaccessible regions.

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than 10 km in just a few days. This region of circulating air extending upward from Earth’s surface to where the air stops becoming colder with height is called the troposphere— from the Greek tropein, meaning “to turn,” or “to change.” Notice also in Fig. 1.23 that just above 11 km the air temperature normally stops decreasing with height.

FIGURE 1.24 Near Earth’s surface the air temperature lapse FI rate is often close to 3.5° F per 1000 ft. If this temperature lapse rate is present and the air temperature at the surface (0 ft) is 46° F, the air temperature about 4000 ft above the surface would be at freezing, and snow and ice might be on the ground.

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NOAA

The Radiosonde

FIGURE 2 A radiosonde with parachute FI and balloon.

Here, the lapse rate is zero. This region, where, on average, the air temperature remains constant with height, is referred to as an isothermal (equal temperature) zone.* The bottom of this zone marks the top of the troposphere and the beginning of another layer, the stratosphere. The boundary separating the troposphere from the stratosphere is called the tropopause. The height of the tropopause varies. It is normally found at higher elevations over equatorial regions, and it decreases in elevation as we travel poleward. Generally, the tropopause is higher in summer and lower in winter at all latitudes. In some regions, the tropopause “breaks” and is difficult to locate, and here scientists have observed tropospheric air mixing with stratospheric air and vice versa. These breaks mark the position of jet streams—high winds that meander in a narrow channel like an old river, often at speeds exceeding 100 knots.** From Fig. 1.23 we can see that in the stratosphere the air temperature begins to increase with height, producing a temperature inversion. The inversion region, along with the lower isothermal layer, tends to keep the vertical currents *In many instances, the isothermal layer is not present and the air temperature begins to increase with increasing height. **Recall from p. 7 that one knot equals 1.15 mi/hr.

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of the troposphere from spreading into the stratosphere. The inversion also tends to reduce the amount of vertical motion in the stratosphere itself; hence, it is a stratified layer (thus, its name). Even though the air temperature is increasing with height, the air at an altitude of 30 km is extremely cold, averaging less than 46°C (51°F). The reason for the inversion in the stratosphere is that the gas ozone plays a major part in heating the air at this altitude. Recall that ozone is important because it absorbs ultraviolet (UV) solar energy. Some of this absorbed energy warms the stratosphere from below, which explains why there is an inversion. If ozone were not present, the air probably would become colder with height, as it does in the troposphere. Above the stratosphere is the mesosphere (middle sphere). The air here is extremely thin and the atmospheric pressure is quite low (again, refer back to Fig. 1.23). Even though the percentage of nitrogen and oxygen in the mesosphere is about the same as it is at Earth’s surface, a breath of mesospheric air contains far fewer oxygen molecules than a breath of tropospheric air. At this level, without proper oxygen-breathing equipment, the brain would soon become oxygen-starved—a condition known as hypoxia—and suffocation would result. With an average temperature of 90°C (130°F), the top of the mesosphere represents the coldest part of Earth’s atmosphere. The “hot layer” above the mesosphere is the thermosphere. Here, oxygen molecules (O2) absorb energetic solar rays, warming the air. In the thermosphere, there are relatively few atoms and molecules. Consequently, the absorption of a small amount of energetic solar energy can cause a large increase in air temperature that may exceed 500°C, or 900°F (see Fig. 1.25). Moreover, it is in the thermosphere where charged particles from the sun interact with air molecules to produce dazzling aurora displays, which are described in more detail in Chapter 2. Even though the temperature in the thermosphere is exceedingly high, a person shielded from the sun would not necessarily feel hot. This situation arises because there are too few molecules in this region of the atmosphere to bump against something (exposed skin, for example) and transfer enough heat to it to make it feel warm. The low density of the thermosphere also means that an air molecule will move an average distance of over one kilometer before colliding with another molecule. A similar air molecule at Earth’s surface will move an average distance of less than one-millionth of a centimeter before it collides with another molecule. At the top of the thermosphere, about 500 km (300 mi) above Earth’s surface, many of the lighter, faster-moving molecules traveling in the right direction actually escape Earth’s gravitational pull. The region where atoms and molecules shoot off into space is sometimes referred to as the exosphere, which represents the upper limit of our atmosphere. Up to this point, we have examined the atmospheric layers based on the vertical profile of temperature. The

FIGURE 1.25 Layers of the atmosphere based on temperature FI (red line), composition (green line), and electrical properties (blue line). (An active sun is associated with large numbers of solar eruptions.)

atmosphere, however, can also be divided into layers based on its composition. For example, the composition of the atmosphere begins to slowly change in the lower part of the thermosphere. Below the thermosphere, the composition of air remains fairly uniform (78 percent nitrogen, 21 percent oxygen) by turbulent mixing. This lower, wellmixed region is known as the homosphere (see Fig. 1.25). In the thermosphere, collisions between atoms and molecules are infrequent, and the air is unable to keep itself stirred. As a result, diffusion takes over as heavier atoms and molecules (such as oxygen and nitrogen) tend to settle to the bottom of the layer, while lighter gases (such as hydrogen and helium) float to the top. The region from about the base of the thermosphere to the top of the atmosphere is often called the heterosphere. The ionosphere is not really a layer, but rather an electrified region within the upper atmosphere where fairly large concentrations of ions and free electrons exist. Ions are atoms and molecules that have lost (or gained) one or more electrons. Atoms lose electrons and become positively charged when they cannot absorb all of the energy transferred to them by a colliding energetic particle or the sun’s energy. Notice in Fig. 1.25 that the lower region of the ionosphere is usually about 60 km above Earth’s surface. From here (60 km), the ionosphere extends upward to the top of the atmosphere. Hence, the bulk of the ionosphere is in the thermosphere. Although the ionosphere allows TV and FM radio waves to pass on through, at night it reflects standard AM radio waves back to Earth. This situation allows AM radio waves to bounce repeatedly off the lower ionosphere and travel great distances. EARTH’S ATMOSPHERE

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SUMMARY This chapter provides an overview of Earth’s atmosphere and the many ways weather and climate influence our lives. We looked briefly at the weather map and a satellite image and observed that dispersed throughout the atmosphere are storms and clouds of all sizes and shapes. The movement, intensification, and weakening of these systems, as well as the dynamic nature of air itself, produce a variety of weather events that we described in terms of weather elements. The sum total of weather and its extremes over a long period of time is what we call climate. Although sudden changes in weather may occur in a moment, climatic change takes place gradually over many years. The study of the atmosphere and all of its related phenomena is called meteorology, a term whose origin dates back to the days of Aristotle. Weather and climate influence the clothes we wear, the food we eat, and many other parts of our lives. Extreme weather can cause severe damage and major disruption to society. We learned that our atmosphere is one rich in nitrogen and oxygen as well as smaller amounts of other gases, such as water vapor, carbon dioxide, and other greenhouse gases whose increasing levels may result in additional global warming and climate change. We examined Earth’s early atmosphere and found it to be much different from the air we breathe today. We investigated the various layers of the atmosphere: the troposphere (the lowest layer), where almost all weather events occur, and the stratosphere, where ozone protects us from a portion of the sun’s harmful rays. Above the stratosphere lies the mesosphere, where the air temperature drops dramatically with height. Above the mesosphere lies the warmest part of the atmosphere, the thermosphere. At the top of the thermosphere is the exosphere, where collisions between gas molecules and atoms are so infrequent that fast-moving lighter molecules can actually escape Earth’s gravitational pull, and shoot off into space. Finally, we looked at the ionosphere, that portion of the upper atmosphere where large numbers of ions and free electrons exist.

KEY TERMS The following terms are listed (with corresponding page numbers) in the order they appear in the text. Define each. Doing so will aid you in reviewing the material covered in this chapter. atmosphere, 4 weather, 5 weather elements, 5 climate, 5 24

meteorology, 5 middle latitudes, 7 middle-latitude cyclonic storm, 7

hurricane, 7 thunderstorms, 7 tornadoes, 7 wind, 7 wind direction, 7 wind speed, 7 front, 8 outgassing, 15 nitrogen, 15 oxygen, 15 water vapor, 16 carbon dioxide, 17 ozone, 18

aerosols, 18 pollutants, 18 air density, 19 air pressure, 20 lapse rate, 21 temperature inversion, 21 radiosonde, 21 troposphere, 22 stratosphere, 22 tropopause, 22 mesosphere, 23 thermosphere, 23 ionosphere, 23

QUESTIONS FOR REVIEW . What is the primary source of energy for Earth’s atmosphere? . How could a meteorologist use the scientific method in predicting the weather? . List seven common weather elements. . How does weather differ from climate? . Define meteorology and discuss the origin of this word. . Rank the following storms in size from largest to smallest: hurricane, tornado, middle-latitude cyclonic storm, thunderstorm. . When someone says that “the wind direction today is south,” does this mean that the wind is blowing toward the south or from the south? . Weather in the middle latitudes tends to move in what general direction? . Describe at least six features observed on a surface weather map. . Explain how the wind generally blows around areas of low and high pressure in the Northern Hemisphere. . Describe at least six ways weather and climate can influence people’s lives. . How has Earth’s atmosphere changed over time? . List the four most abundant gases in today’s atmosphere. . Of the four most abundant gases in our atmosphere, which one shows the greatest variation from place to place at Earth’s surface? . Explain how the atmosphere “protects” inhabitants at Earth’s surface. . What are some of the important roles that water plays in our atmosphere? . Briefly explain the production and natural destruction of carbon dioxide near Earth’s surface. Give two reasons

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. . .

. . . .

. .

.

for the increase of carbon dioxide over the past 100 plus years. What are some of the aerosols in the atmosphere? What are the two most abundant greenhouse gases in Earth’s atmosphere? (a) Explain the concept of air pressure in terms of weight of air above some level. (b) Why does air pressure always decrease with increasing height above the surface? What is standard atmospheric pressure at sea level in (a) inches of mercury, (b) millibars, and (c) hectopascals? On the basis of temperature, list the layers of the atmosphere from the lowest layer to the highest. Briefly describe how the air temperature changes from Earth’s surface to the lower thermosphere. (a) What atmospheric layer contains all of our weather? (b) In what atmospheric layer do we find the highest concentration of ozone? The highest average air temperature? Above what region of the world would you find the ozone hole? Even though the actual concentration of oxygen is close to 21 percent (by volume) in the upper stratosphere, explain why, without proper breathing apparatus, you would not be able to survive there. What is the ionosphere and where is it located?

QUESTIONS FOR THOUGHT AND EXPLORATION .

Explain how you considered both weather and climate in your choice of the clothing you chose to wear today. . Compare a newspaper weather map with a professional weather map obtained from the Internet.

.

.

.

.

Discuss any differences in the two maps. Look at both maps and see if you can identify a warm front, a cold front, and a middle-latitude cyclonic storm. Which of the following statements relate more to weather and which relate more to climate? (a) The summers here are warm and humid. (b) Cumulus clouds presently cover the entire sky. (c) Our lowest temperature last winter was 29°C (18°F). (d) The air temperature outside is 22°C (72°F). (e) December is our foggiest month. (f) The highest temperature ever recorded in Phoenixville, Pennsylvania, was 44°C (111°F) on July 10, 1936. (g) Snow is falling at the rate of 5 cm (2 in.) per hour. (h) The average temperature for the month of January in Chicago, Illinois, is 3°C (26°F). Suppose a friend poses a question about how weather systems generally move in the middle latitudes. He puts forth a hypothesis that these systems generally move from east to west. How would you use the scientific method to prove his hypothesis to be incorrect? Keep track of the weather. On an outline map of North America, mark the daily position of fronts and pressure systems for a period of several weeks or more. (This information can be obtained from newspapers, the TV news, or from the Internet.) Plot the general upper-level flow pattern on the map. Observe how the surface systems move. Compose a one-week journal, including daily newspaper, weather maps and weather forecasts from a newspaper or from the Internet. Provide a commentary for each day regarding the coincidence of actual and predicted weather.

Go to the News section of the Meteorology portal. Use the search box at left to bring up articles related to the term “storm.” Of the first 25 articles, what are the various kinds of weather phenomena mentioned (dust, tornadoes, etc.)? Which ones are mentioned most often? What terms precede the word “storm”?

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CHAPTER

2

Warming and Cooling Earth and Its Atmosphere Contents Temperature and Heat Transfer Radiant Energy

T

he sun doesn’t rise or fall: it doesn’t move, it just sits there, and we rotate in front of it. Dawn awn means that we are rotatrotat

ing around into sight of it, while dusk means we have turned another 180 degrees and are being carried into the shadow zone. The he sun never “goes away from the sky.” It’s still there

Radiation—A adiation— bsorption, adiation—A Emission, and Equilibrium

sharing the same sky with us; it’s simply that there is a chunk

Why Earth Has Seasons

longer do I drive down a highway and wish the blinding sun

of opaque earth between us and the sun which prevents our seeing it. Everyone knows that, but I really see it now. No would set; instead I wish we could speed up our rotation a bit and swing around into the shadows more quickly. Michael Collins, Carrying the Fire

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A

s you sit quietly reading this book, you are part of a moving experience. Planet Earth is speeding around the sun at thousands of miles per hour while, at the same time, spinning on its axis. When we look down upon the North Pole, we see that the direction of spin is counterclockwise, meaning that we are moving toward the east at hundreds of miles per hour. We normally don’t think of it in that way, but, of course, this is what causes the sun, moon, and stars to rise in the east and set in the west. In fact, it is these motions coupled with energy from the sun, striking a tilted planet, that cause our seasons. But, as we will see later, the sun’s energy is not distributed evenly over Earth. Tropical regions receive more energy than polar regions. It is this energy imbalance that drives our atmosphere into the dynamic patterns we experience as wind and weather. We will begin this chapter by examining the concept of energy and heat transfer. Then we will see how our atmosphere warms and cools. Finally, we will examine how Earth’s motions and the sun’s energy work together to produce the seasons.

Temperature and Heat Transfer Temperature is the measurement that tells us how hot or cold something is relative to some set standard value. But we can look at temperature in another way. We know that air is a mixture of countless billions of atoms and molecules. If they could be seen, they would appear to be moving about in all directions, freely darting, twisting, spinning, and colliding with one another like an angry swarm of bees. Close to Earth’s surface, each individual molecule would travel about a thousand times its diameter before colliding with another molecule. Moreover, we would see that all the atoms and molecules are not moving at the same speed, as some are moving faster than others. This is kinetic energy, the energy of motion. The temperature of the air (or any substance) is a measure of its average kinetic energy. Simply stated, temperature is a

measure of the average speed (average motion) of the atoms and molecules, where higher temperatures correspond to faster average speeds. Suppose we examine a volume of surface air about the size of a large flexible balloon as shown in Fig. 2.1. If we warm the air inside, the molecules will move faster, but they also will move slightly farther apart—the air becomes less dense, as illustrated in Fig. 2.1b. Conversely, if we cool the air back to its original temperature, the molecules would slow down, crowd closer together, and the air will become more dense. This molecular behavior is why, in many places throughout the book, we refer to surface air as either warm, less-dense air or as cold, moredense air. Suppose we continue to slowly cool the air. Its atoms and molecules will move more and more slowly until the air reaches a temperature of 273°C (459°F), which is the lowest temperature possible. At this temperature, called absolute zero, the atoms and molecules will possess a minimum amount of energy and theoretically no thermal motion. Along with temperature, we can also measure internal energy, which is the total energy stored in a group of molecules. Heat, on the other hand, is energy in the process of being transferred from one object to another because of the temperature difference between them. After heat is transferred, it is stored as internal energy. In the atmosphere, heat is transferred by conduction, convection, and radiation. We will examine these mechanisms of energy transfer after we look at temperature scales and the important concept of latent heat. TEMPERATURE SCALES Recall that, theoretically, at a temperature of absolute zero there is no thermal motion. Absolute zero is the starting point for a temperature scale called the absolute scale, or Kelvin scale, after Lord Kelvin (1824–1907), the British scientist who first introduced it. Since the Kelvin scale contains no negative

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FIGURE 2.1 Air temperature is a measure FIGUR of the average speed (motion) of the mol molecules. In the cold volume of air the molecules move more slowly and crowd closer together. In the warm volume, they move faster and farther apart.

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8C 5 5/918F 322 On the Kelvin scale, degrees Kelvin are called kelvins (abbreviated K). Each degree on the Kelvin scale is exactly the same size as a degree Celsius, and a temperature of 0 K is equal to 273°C. Converting from °C to K can be done by simply adding 273 to the Celsius temperature, as K 5 8C 1 273 Figure 2.2 compares the Kelvin, Celsius, and FahrFahr enheit scales. Converting a temperature from one scale to another can be done by simply reading the corresponding temperature from the adjacent scale. Thus, 303 K on the Kelvin scale is the equivalent of 30°C and 86°F.* In most of the world, temperature readings are taken in °C and public weather forecasts use the Celsius scale. In the United States, however, temperatures above the surface are taken in °C, while temperatures at the surface are typically read and reported in °F. Likewise, temperatures on upper-level maps are plotted in °C, while, on surface weather maps, they are in °F. Since both scales are in use, temperature readings in this book will, in most cases, be given in °C followed by their equivalent in °F. LATENT HEAT—THE HIDDEN WARMTH We know from Chapter 1 that water vapor is an invisible gas that becomes visible when it changes into larger liquid or solid

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numbers, it is quite convenient for scientific calculations. Two other temperature scales commonly used today are the Fahrenheit and the Celsius (formerly centigrade). The Fahrenheit scale was developed in the early eighteenth century by the physicist G. Daniel Fahrenheit (1686–1736), who assigned the number 32 to the temperature at which water freezes, and the number 212 to the temperature at which water boils. The zero point was simply the lowest temperature that he obtained with a mixture of ice, water, and salt. Between the freezing and boiling points are 180 equal divisions, each of which is called a degree. A thermometer calibrated with this scale is referred to as a Fahrenheit thermometer, for it measures an object’s temperature in degrees Fahrenheit (°F). The Celsius scale, named after Swedish astronomer Anders Celsius (1701–1744), was introduced later in the eighteenth century. The number 0 (zero) on this scale is assigned to the temperature at which pure water freezes, and the number 100 to the temperature at which pure water boils at sea level. The space between freezing and boiling is divided into 100 equal degrees. Therefore, each Celsius degree (°C) is 180/100 or 1.8 times bigger than a Fahrenheit degree. Put another way, an increase in temperature of 1°C equals an increase of 1.8°F. A formula for converting °F to °C is

FIGURE 2.2 Comparison of Kelvin, Celsius, and Fahrenheit scales, along with some world temperature extremes.

(ice) particles. This process of transformation is known as a change of state or, simply, a phase change. The heat energy required to change a substance, such as water, from one state to another is called latent heat. But why is this heat referred to as “latent”? To answer this question, we will begin with something familiar to most of us—the cooling produced by evaporating water. Suppose we microscopically examine a small drop of pure water. At the drop’s surface, molecules are constantly escaping (evaporating). Because the more energetic, faster-moving molecules escape most easily, the average motion of all the molecules left behind decreases as each additional molecule evaporates. Since temperature is a measure of average molecular motion, the slower motion

DID YOU KNOW? We usually think of average human body temperature as being 37°C (98.6°F). This value, established in European studies more than a century ago, may not be quite right. The original research used thermometers that were less accurate than we now have, and the scientists appear to have rounded their data to the nearest degree Celsius (37°C). More recent studies have found that the actual average is closer to 36.8°C (98.2°F). Body temperature varies widely from person to person and from morning to evening.

*A more complete table of conversions is given in Appendix A. WARMING AND COOLING EARTH AND ITS ATMOSPHERE Copyright 2018 Cengage Learning. All Rights Reserved. May not be copied, scanned, or duplicated, in whole or in part. WCN 02-200-202

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FIGURE 2.3 Heat energy FIGUR absorbed and released.

*By definition, a calorie is the amount of heat required to raise the temperature of 1 gram of water from 14.5°C to 15.5°C. In the International System (Système International, SI), the unit of energy is the joule (J), where 1 calorie = 4.186 J. (For pronunciation: joule rhymes with pool.)

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the change of state is from right to left, heat energy is given up by the substance and added to the environment. The processes of freezing, condensation, and deposition (vapor to ice) all warm their surroundings. Latent heat is an important source of atmospheric energy. Once vapor molecules become separated from Earth’s surface, they are swept away by the wind, like dust before a broom. Rising to high altitudes where the air is cold, the vapor changes into liquid and ice cloud particles. During these processes, a tremendous amount of heat energy is released into the environment (see Fig. 2.4). Water vapor evaporated from warm, tropical water can be carried into polar regions, where it condenses

Robert Henson

suggests a lower water temperature. Evaporation is, therefore, a cooling process. Stated another way, evaporation is a cooling process because the energy needed to evaporate the water—that is, to change its phase from a liquid to a gas—may come from the water or other sources, including the air. The energy lost by liquid water during evaporation can be thought of as carried away by, and “locked up” within, the water vapor molecule. The energy is thus in a “stored” or “hidden” condition and is, therefore, called latent heat. It is latent (hidden) in that the temperature of the substance changing from liquid to vapor is still the same. However, the heat energy will reappear as sensible heat (the heat we can feel and measure with a thermometer) when the vapor condenses back into liquid water. Therefore, condensation (the opposite of evaporation) is a warming process. The heat energy released when water vapor condenses to form liquid droplets is called latent heat of condensation. Conversely, the heat energy used to change liquid into vapor at the same temperature is called latent heat of evaporation (vaporization). Nearly 600 calories* are required to evaporate a single gram of water at room temperature. With many hundreds of grams of water evaporating from the body, it is no wonder that after a shower we feel cold ex before drying off. Figure 2.3 summarizes the concepts examined so far. When the change of state is from left to right, heat is absorbed by the substance and taken away from the environment. The processes of melting, evaporation, and sublimation (ice to vapor) all cool the environment. When

FIGURE 2.4 Every time a cloud forms, it warms the atmosphere. Inside this developing thunderstorm, a vast amount of stored heat energy (latent heat) is given up to the air, as invisible water vapor becomes countless billions of water droplets and ice crystals. In fact, for the duration of this storm alone, more heat energy is released inside this cloud than is unleashed by a small nuclear bomb.

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▼ Table 2.1

FIGURE 2.5 The transfer of heat from the hot end of FIGUR the metal pin to the cool end by molecular contact is called conduction.

and gives up its heat energy. As we will see, evaporationtransportation-condensation is an extremely important mechanism for the relocation of heat energy (as well as water) in the atmosphere. We are now ready to look at other mechanisms of heat transfer in the atmosphere. CONDUCTION The transfer of heat from molecule to molecule within a substance is called conduction. Hold one end of a metal straight pin between your fingers and place a flaming candle under the other end (see Fig. 2.5). Because of the energy they absorb from the flame, the molmol ecules in the pin vibrate faster. The faster-vibrating molecules cause adjoining molecules to vibrate faster. These, in turn, pass vibrational energy on to their neighboring molecules, and so on, until the molecules at the finger-held end of the pin begin to vibrate rapidly. These fast-moving molecules eventually cause the molecules of your finger to vibrate more quickly. Heat is now being transferred from the pin to your finger, and both the pin and your finger feel hot. If enough heat is transferred, your finger will become painful, and you will drop the pin. The transmission of heat from one end of the pin to the other, and from the pin to your finger, occurs by conduction. Heat transferred in this fashion always flows from warmer to colder regions. Generally, the greater the temperature difference, the more rapid the heat transfer. When materials can easily pass energy from one molecule to another, they are considered to be good conductors of heat. How well they conduct heat depends upon how their molecules are structurally bonded together. ▼ Table 2.1 shows that solids, such as metals, are good heat conductors. It is often difficult, therefore, to judge the temperature of metal objects. For example, if you

Heat Conductivity* of Various Substances

SUBSTANCE

HEAT CONDUCTIVITY (WATTS** PER METER PER °C)

Still air

0.023 (at 20°C)

Wood

0.08

Dry soil

0.25

Water

0.60 (at 20°C)

Snow

0.63

Wet soil

2.1

Ice

2.1

Sandstone

2.6

Granite

2.7

Iron

80

Silver

427

*Heat (thermal) conductivity describes a substance’s ability to conduct heat as a consequence of molecular motion. **A watt (W) is a unit of power where one watt equals one joule (J) per second (J/s). One joule equals 0.24 calories.

grab a metal pipe at room temperature, it will seem to be much colder than it actually is because the metal conducts heat away from the hand quite rapidly. Conversely, air is an extremely poor conductor of heat, which is why most insulating materials have a large number of air spaces trapped within them. Air is such a poor heat conductor that, in calm weather, the hot ground only warms a shallow layer of air a few centimeters thick by conduction. Yet, air can carry this energy rapidly from one region to another. How, then, does this phenomenon happen? CONVECTION The transfer of heat by the mass movement of a fluid (such as water and air) is called convection. This type of heat transfer takes place in liquids and gases because they can move freely, and it is possible to set up currents within them. Convection happens naturally in the atmosphere. On a warm, sunny day, certain areas of Earth’s surface absorb more heat from the sun than others; as a result, the air near Earth’s surface is heated somewhat unevenly. Air molecules adjacent to these hot surfaces bounce against them, thereby gaining some extra energy by conduction. The heated air expands and becomes less dense than the surrounding cooler air. The expanded warm air is buoyed upward and rises. In this manner, large bubbles of warm air rise and transfer heat energy upward. Cooler, heavier air flows toward the surface to replace the rising air. This cooler air becomes heated in turn, rises, and the cycle is repeated. In meteorology, this vertical exchange of heat WARMING AND COOLING EARTH AND ITS ATMOSPHERE

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FOCUS

ON A SPECIAL TOPIC 2.1 O

Rising Air Cools and Sinking Air Warms parcel. Recall from Chapter 1 that air pressure always decreases as we move up into the atmosphere. Consequently, as the parcel rises, it enters a region where the surrounding air pressure is lower. To equalize the pressure, the parcel molecules inside push the parcel walls outward, expanding it. Because there is no other energy source, the air molecules inside use some of their own energy to expand the parcel. This energy loss shows up as slower molecular speeds, which represent a lower parcel temperature. Hence, any air that rises always expands and cools.

If the parcel is lowered to Earth (as shown in Fig. 1), it returns to a region where the air pressure is higher. The higher outside pressure squeezes (compresses) the parcel back to its original (smaller) size. Because air molecules have a faster rebound velocity after striking the sides of a collapsing parcel, the average speed of the molecules inside goes up. (A Ping-Pong ball moves faster after striking a paddle that is moving toward it.) This increase in molecular speed represents a warmer parcel temperature. Therefore, any air that sinks (subsides) warms by compression.

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To understand why rising air cools and sinking air warms, we need to examine some air. Suppose we place air in an imaginary thin, elastic wrap about the size of a large balloon (see Fig. 1). This invisible balloonlike “blob” is called an air parcel. The air parcel can expand and conparcel tract freely, but neither external air nor heat is able to mix with the air inside. By the same token, as the parcel moves, it does not break apart, but remains as a single unit. At Earth’s surface, the parcel has the same temperature and pressure as the air surrounding it. Suppose we lift the

FIGURE 1 Rising air expands and cools; sinking air is compressed and warms. FIGUR

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is called convection, and the rising air bubbles are known as thermals (see Fig. 2.6). The rising air expands and gradually spreads outward. It then slowly begins to sink. Near the surface, it moves back

FIGURE 2.6 The development of a thermal. A thermal is a risFIGUR ing bubble of air that carries heat energy upward by convection.

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into the heated region, replacing the rising air. In this way, a convective circulation, or thermal “cell,” is produced in the atmosphere. In a convective circulation the warm, rising air cools. In our atmosphere, any air that rises will expand and cool, and any air that sinks is compressed and warms. To learn more about this important concept, read Focus section 2.1. Although the entire process of heated air rising, spreading out, sinking, and finally flowing back toward its original location is known as a convective circulation, meteorologists usually restrict the term convection to the process of the rising and sinking part of the circulation (see Fig 2.7). The horizontally moving part of the circulation (called wind) carries properties of the air in that particular area with it. The transfer of these properties by horizontally moving air is called advection. For example, wind blowing across a body of water will “pick up” water vapor from the evaporating surface and transport it elsewhere in the atmosphere. If the air cools, the water vapor may condense into cloud droplets and release latent heat. In a sense, then,

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Radiant Energy

FIGURE 2.7 The rising of hot air and the sinking of cool air FIGUR sets up a convective circulation. Normally, the vertical part of the circulation is called convection, whereas the horizontal part is called wind. Near the surface the wind is advecting smoke from one region to another.

heat is advected (carried) by the water vapor as it is swept along with the wind. Earlier, we saw that this is an important way to redistribute heat energy in the atmosphere.

BRIEF REVIEW Before moving on to the next section, here is a summary of some of the important concepts and facts we have covered: ●

The temperature of a substance is a measure of the average kinetic energy (average motion) of its atoms and molecules.

Evaporation (the transformation of liquid into vapor) is a cooling process that can cool the air, whereas condensation (the transformation of vapor into liquid) is a warming process that can warm the air.

Heat is energy in the process of being transferred from one object to another because of the temperature difference between them.

In conduction, which is the transfer of heat by molecule-to-molecule contact, heat always flows from warmer to colder regions.

Air is a poor conductor of heat.

Convection is an important mechanism of heat transfer, as it represents the vertical movement of warmer air upward and cooler air downward.

The horizontal transfer of any atmospheric property by the wind (including smoke and warm or cold air) is called advection.

There is yet another mechanism for the transfer of energy—radiation, or radiant energy, which is what we receive from the sun. In this method, energy can be transferred from one object to another without the space between them necessarily being heated.

DID YOU KNOW? Some birds are weather savvy. Hawks, for example, seek out rising thermals and ride them up into the air as they scan the landscape for prey. In doing so, these birds conserve a great deal of energy by not having to flap their wings as they circle higher and higher inside the rising air current.

On a bright winter day, you may have noticed how warm your face feels as you stand facing the sun. Sunlight travels through the surrounding air with little effect upon the air itself. Your face, however, absorbs this energy and converts it to thermal energy. Thus, sunlight warms your face without actually warming the air. The energy transferred from the sun to your face is called radiant energy, or radiation. It travels in the form of waves that release energy when they are absorbed by an object. Because these waves have magnetic and electrical properties, we call them electromagnetic waves. Electromagnetic waves do not need molecules to propagate them. In a vacuum, they travel at a constant speed of nearly 300,000 km (186,000 mi) per second—the speed of light. Figure 2.8 shows some of the different wavelengths of radiation. Notice that the wavelengt wavelength (which is often expressed by the Greek letter lambda, ) is the distance measured along a wave from one crest to another. Also notice that some of the waves have exceedingly short lengths. For example, radiation that we can see (visible light) has an average wavelength of less than one-millionth of a meter— a distance nearly one-hundredth the diameter of a human hair. To help describe these short lengths, we introduce a new unit of measurement called a micrometer (abbreviated m), which is equal to one-millionth of a meter (m); thus 1 micrometer (m) = 0.000001 m = 10-6 m. In Fig. 2.8, we can see that the average wavelength of visible light is about 0.0000005 meters (or 5 x 10-7), which is the same as 0.5 m. To give you a common object for comparison, the average height of a letter on this page is about 2000 m, or 2 millimeters (2 mm). We can also see in Fig. 2.8 that the longer waves carry less energy than do the shorter waves. When comparing the energy carried by various waves, it is useful to give electromagnetic radiation characteristics of particles in order to explain some of the wave’s behavior. We can actually think of radiation as streams of particles, or photons, that are discrete packets of energy.* As shown in Figure 2.8, an ultraviolet (UV) photon carries more energy than a photon of visible light. In fact, certain ultraviolet photons have enough energy to produce sunburns and penetrate skin tissue, sometimes causing skin cancer. (To learn more about radiant energy and its effect on humans, read Focus section 2.2.) Here are a few important ideas and facts to remember about the concept of radiation: *Packets of photons make up waves, and groups of waves make up a beam of radiation. WARMING AND COOLING EARTH AND ITS ATMOSPHERE

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FIGURE 2.8 Radiation characterized FIGUR according to wavelength. As the wave wavelength decreases, the energy carried per wave increases.

. All things (with a temperature above absolute zero), no matter how big or small, emit radiation. The air, your body, flowers, trees, Earth, and the stars are all radiating a wide range of electromagnetic waves. The energy originates from rapidly vibrating electrons, billions of which exist in every object. . The wavelengths of radiation that an object emits depend primarily on the object’s temperature. The higher the object’s temperature, the shorter are the wavelengths of emitted radiation. By the same token, as an object’s temperature increases, its peak emission of radiation shifts toward shorter wavelengths. This relationship between temperature and wavelength is called Wien’s law* (or Wien’s displacement law) after the German physicist Wilhelm Wien (pronounced Ween, 1864–1928) who discovered it. . Objects that have a high temperature emit radiation at a greater rate or intensity than objects with a lower temperature. Thus, as the temperature of an object increases, more total radiation (over a given surface area) is emitted each second. This relationship between temperature and emitted radiation is known as the Stefan-Boltzmann law** after Josef Stefan (1835–1893) and Ludwig Boltzmann (1844–1906), who devised it. *Wien’s law: lmax 5 constant T where max is the wavelength at which maximum radiation emission occurs, T is the object’s temperature in kelvins (K) and the constant is 2897 m K. More information on Wien’s law is given in Appendix B. **Stefan-Boltzmann law: E 5 sT 4 where E is the maximum rate of radiation emitted by each square meter of surface of an object, (the Greek letter sigma) is a constant, and T is the object’s surface temperature in kelvins (K). Additional information on the Stefan-Boltzmann law is given in Appendix B.

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Objects at a high temperature (above about 500°C) radiate waves with many lengths, some of which are short enough to stimulate the sensation of color. We actually see these objects glow red. Objects cooler than this radiate at wavelengths that are too long for us to see. The page of this book, for example, is radiating electromagnetic waves. But because its temperature is only about 20°C (68°F), the waves emitted are much too long to stimulate vision. We are able to see the page, however, because light waves from other sources (such as light bulbs or the sun) are being reflected (bounced) off the paper. If this book were carried into a completely dark room, it would continue to radiate, but the pages would appear black because there are no visible light waves in the room to reflect off the page. The sun emits radiation at almost all wavelengths, but because its surface is extremely hot—6000 K (10,500°F)— it radiates the majority of its energy at relatively short wavelengths. If we look at the amount of radiation given off by the sun at each wavelength, we obtain the sun’s electromagnetic spectrum. A portion of this spectrum is shown in Fig. 2.9. Notice that the sun emits a maximum amount of radiation at wavelengths near 0.5 m. Since our eyes are sensitive to radiation between 0.4 and 0.7 m, these waves reach the eye and stimulate the sensation of color. This portion of the spectrum is therefore referred to as the visible region, and the radiant energy that reaches our eye is called visible light. The color violet is the shortest wavelength of visible light. Wavelengths shorter than violet (0.4 m) are ultraviolet (UV). The longest wavelengths of visible light correspond to the color red. A rainbow spans the spectrum of visible light from violet to red. Wavelengths longer than red (0.7 m) are called infrared (IR).

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FOCUS

ON AN ENVIRONMENTAL ISSUE 2.2

FIGURE 2 The UV index. FIGUR

Additional protection can come from sunscreens that block UV rays from ever reaching the skin. Some contain chemicals (such as zinc oxide) that reflect UV radiation. (These are the white pastes once seen on the noses of lifeguards.) Others consist of a mixture of chemicals that actually absorb ultraviolet radiation. The Sun Protection Factor (SPF) number on every container of sunscreen dictates how effective the product is in protecting from UVB—the higher the number, the better the protection. Many “broad-spectrum” sunscreens protect against both UVA and UVB, although only the amount of UVA protection is considered in the SPF rating. Protecting oneself from excessive exposure to the sun’s energetic UV rays is certainly wise. Estimates are that, in a single year, over 70,000 Americans will be diagnosed with malignant melanoma, the most deadly form of skin cancer. And in areas where the protective stratospheric ozone shield has weakened, there is an increased risk of problems associated with UVB. Using a good sunscreen and proper clothing can certainly help. The best way to protect yourself from too much sun, however, is to limit your time in direct sunlight, especially between the hours of 10 a.m. and 4 p.m. daylight saving time when the sun is highest in the sky and its rays are most direct. Each day the National Weather Service predicts UV radiation levels for selected cities throughout the United

States. The forecast, known as the UV Index, gives the UV level at its peak, around noon standard time or 1 p.m. daylight saving time. The index corresponds to five exposure categories set by the Environmental Protection Agency (EPA). An index value of 2 or less is considered “low,” whereas a value of 11 or greater is deemed “extreme” (see Fig. 2). Depending on skin type, a UV index of 10 means that in direct sunlight (without sunscreen protection) a person’s skin will likely begin to burn in about 6 to 30 minutes (see Fig. 3).

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Earlier, we learned that shorter waves of radiation carry much more energy than longer waves, and that a photon of ultraviolet light carries more energy than a photon of visible light. In fact, ultraviolet (UV) wavelengths in the range of 0.20 and 0.29 m (known as UVC radiation) are harmful to living things, as certain waves can cause chromosome mutations, kill single-celled organisms, and damage the cornea of the eye. Fortunately, virtually all the ultraviolet radiation at wavelengths in the UVC range is absorbed by ozone in the stratosphere. Ultraviolet wavelengths between about 0.29 and 0.32 m (known as UVB radiation) reach Earth in small amounts. Photons in this wavelength range have enough energy to produce sunburns and penetrate skin tissues, sometimes causing skin cancer. About 90 percent of all skin cancers are linked to sun exposure and UVB radiation. Oddly enough, these same wavelengths activate provitamin D in the skin and convert it into vitamin D, which is essential to health. Longer ultraviolet waves with lengths of about 0.32 to 0.40 m (called UVA radiation) are less energetic, but they are the main ones that produce skin tanning. Although UVB is the primary wavelength responsible for burning the skin, UVA can cause skin redness. It can also interfere with the skin’s immune system and cause long-term skin damage that shows up years later as accelerated aging and skin wrinkling. Moreover, recent studies indicate that the longer UVA exposures needed to create a tan pose about the same cancer risk as a UVB tanning dose. Upon striking the human body, ultraviolet radiation is absorbed beneath the outer layer of skin. To protect the skin from these harmful rays, the body’s defense mechanism kicks in. Certain cells (when exposed to UV radiation) produce a dark pigment (melanin) that begins to absorb some of the UV radiation. (It is the production of melanin that produces a tan.) Consequently, a body that produces little melanin—one with pale skin—has little natural protection from UVB.

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Sunburning and UV Rays

FIGURE 3 If this photo was taken FIGUR around 1 p.m. on a day when the UV index was 10 almost everyone on this beach without sunscreen would experience some degree of sunburning within 30 minutes.

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FIGURE 2.9 The sun’s FIGUR electromagnetic spec spectrum and some of the descriptive names of each region. The numbers underneath the curve approximate the percent of energy the sun radiates in various regions.

Whereas the hot sun emits only a part of its energy in the infrared portion of the spectrum, the relatively cool earth emits almost all of its energy at infrared wavelengths. In fact, notice in Fig. 2.10 that Earth, with an average surface temperature near 288 K (15°C, or 59°F), radiates nearly all its energy between 5 and 20 m, with a peak intensity (max) in the infrared region near 10 m. The sun, with a much higher surface temperature, radiates with a peak emission near 0.5 m. Since the sun radiates the majority of its energy at much shorter wavelengths than does Earth, solar radiation is often called shortwave radiation, whereas Earth’s radiation is referred to as longwave (or terrestrial) radiation.

infrared energy and, by early morning, it may be cooler than surrounding surfaces. Any object that is a perfect absorber (that is, absorbs all the radiation that strikes it) and a perfect emitter (emits the maximum radiation possible at its given temperature) is called a blackbody. Blackbodies do not have to be colored black; they simply must absorb and emit all possible radiation. Since Earth’s surface and the sun absorb and radiate with nearly 100 percent efficiency for their respective temperatures, they both behave as blackbodies. When we look at Earth from space, we see that half of it is in sunlight, the other half is in darkness. The outpouring of solar energy constantly bathes Earth with radiation, while Earth, in turn, constantly emits infrared radiation. If we assume that there is no other method of transferring heat, then, when the rate of absorption of solar radiation equals the rate of emission of infrared Earth radiation, a state of radiative equilibrium is achieved.

If Earth and all things on it are continually radiating energy, why doesn’t everything get progressively colder? The answer is that all objects not only radiate energy, they absorb it as well. If an object radiates more energy than it absorbs, it becomes colder; if it absorbs more energy than it emits, it becomes warmer. On a sunny day, Earth’s surface warms by absorbing more energy from the sun and the atmosphere than it radiates, whereas at night Earth cools by radiating more energy than it absorbs from its surroundings. When an object emits and absorbs energy at equal rates, its temperature remains constant. The rate at which something radiates and absorbs energy depends strongly on its surface characteristics, such as color, texture, and moisture, as well as temperature. For example, a black object in direct sunlight is a good absorber of solar radiation. It converts energy from the sun into internal energy, and its temperature ordinarily increases. You need only walk barefoot on a black asphalt road on a summer afternoon to experience this. At night, the blacktop road will cool quickly by emitting 36

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Radiation—Absorption, Emission, and Equilibrium

FIGURE 2.10 The hotter sun not only radiates more energy FIGUR than that of the cooler Earth (the area under the curve), but it also radiates the majority of its energy at much shorter wavelengths. (The area under the curves is equal to the total energy emitted, and the scales for the two curves differ by a factor of roughly 1 million.)

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SELECTIVE ABSORBERS AND THE ATMOSPHERIC GREENHOUSE EFFECT There are many selective absorbers in our environment. Snow, for example, is a good absorber of infrared radiation but a poor absorber of sunlight. Objects that selectively absorb radiation usually selectively emit radiation at the same wavelength. Snow is therefore a good emitter of infrared energy. At night, a snow surface usually emits much more infrared energy than it absorbs from its surroundings. This large loss of infrared radiation (coupled with the insulating qualities of snow) enables the air above a snow surface on a clear, calm winter night to become extremely cold. Figure 2.11 shows some of the most important selectively absorbing gases in our atmosphere (the purple shaded area represents the absorption characteristics of each gas at various wavelengths). Notice that both water vapor (H2O) and carbon dioxide (CO2) are strong absorbers of infrared radiation and poor absorbers of visible solar radiation. Other, less important, selective absorbers include nitrous oxide (N2O), methane (CH4), and ozone (O3), which is most abundant in the stratosphere. As these gases absorb infrared radiation emitted from Earth’s surface, they gain kinetic energy (energy of motion). The gas molecules share this energy by colliding with neighboring air molecules, such as oxygen and nitrogen (both of which are poor absorbers of infrared energy). These collisions increase the average kinetic energy of the air, which results in an increase in air temperature. Thus, most of the infrared energy emitted from Earth’s surface keeps the lower atmosphere warm. Besides being selective absorbers, water vapor and CO2 selectively emit radiation at infrared wavelengths.* This radiation travels away from these gases in all directions. *Nitrous oxide, methane, and ozone also emit infrared radiation, but their concentration in the atmosphere is much smaller than water vapor and carbon dioxide (see Table 1.1, p. 16).

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The average temperature at which this occurs is called the radiative equilibrium temperature. At this temperature, Earth (behaving as a blackbody) is absorbing solar radiation and emitting infrared radiation at equal rates, and its average temperature does not change. As Earth is about 150 million km (93 million mi) from the sun, Earth’s radiative equilibrium temperature is about 255 K (18°C, 0°F). But this temperature is much lower than Earth’s observed average surface temperature of 288 K (15°C, 59°F). Why is there such a large difference? The answer lies in the fact that Earth’s atmosphere absorbs and emits infrared radiation. Unlike Earth, the atmosphere does not behave like a blackbody, as it absorbs some wavelengths of radiation and is transparent to others. Objects that selectively absorb and emit radiation, such as gases in our atmosphere, are known as selective absorbers.

FIGURE 2.11 Absorption of radiation by gases in the atmoFIGUR sphere. The purple-shaded area represents the percent of radia radiation absorbed by each gas. The strongest absorbers of infrared radiation are water vapor and carbon dioxide. The bottom figure represents the percent of radiation absorbed by all of the atmospheric gases.

A portion of this energy is radiated toward Earth’s surface and absorbed, thus heating the ground. Earth, in turn, radiates infrared energy upward, where it is absorbed and warms the lower atmosphere. In this way, water vapor and CO2 absorb and radiate infrared energy WARMING AND COOLING EARTH AND ITS ATMOSPHERE

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DID YOU KNOW? In a typical year, more than thirty infants and many more pets are killed in the United States by heatstroke when they are left inside a vehicle in direct sunlight with windows rolled up. Just like a florist’s greenhouse, the interior of a car is warmed by the sun’s radiant energy. The trapped heat inside can have deadly consequences, as the temperature inside the vehicle can climb rapidly. Even if outside temperatures are as low as 25°C (76°F), sunshine can cause the air temperature inside a closed vehicle to rise above 43°C (110°F) in less than an hour. On a hot day, readings can top 55°C (131°F).

and act as an insulating layer around Earth, keeping part of Earth’s infrared radiation from escaping rapidly into space. Consequently, Earth’s surface and the lower atmosphere are much warmer than they would be if these selectively absorbing gases were not present. In fact, as we saw earlier, Earth’s mean radiative equilibrium temperature without CO2 and water vapor would be around 18°C (0°F), or about 33°C (59°F), lower than at present. The absorption characteristics of water vapor, CO2, and other gases (such as methane and nitrous oxide depicted in Fig. 2.11) were at one time thought to be similar to the glass of a florist’s greenhouse. In a greenhouse, the glass allows visible radiation to come in, but inhibits to some degree the passage of outgoing infrared radiation. For this reason, the absorption of infrared radiation from Earth by water vapor and CO2 is popularly called the greenhouse effect. However, studies have shown that the warm air inside a greenhouse is probably caused more by the air’s inability to circulate and mix with the cooler outside air than by the entrapment of infrared energy. Because of these findings, some scientists suggest that the greenhouse effect should be called the atmosphere effect. To accommodate everyone, we will usually use the term atmospheric greenhouse effect when describing the role that water vapor, CO2, and other greenhouse gases* play in keeping Earth’s mean surface temperature higher than it otherwise would be. Look again at Fig. 2.11 and observe that, in the bottom diagram, there is a region between about 8 and 11 m where neither water vapor nor CO2 readily absorbs infrared radiation. Because these wavelengths of emitted energy pass upward through the atmosphere and out into space, the wavelength range (between 8 and 11 m) is known as the atmospheric window. Clouds can enhance the atmospheric greenhouse effect. Tiny liquid cloud droplets are selective absorbers in that they are good absorbers of infrared radiation but poor absorbers of visible solar radiation. Clouds even absorb the wavelengths between 8 and 11 m, which are otherwise “passed up” by water *The term “greenhouse gases” derives from the standard use of “greenhouse effect.” Greenhouse gases include, among others, water vapor, carbon dioxide, methane, nitrous oxide, and ozone.

38

vapor and CO2. Thus, they have the effect of enhancing the atmospheric greenhouse effect by closing the atmospheric window. Clouds—especially low, thick ones—are excellent emitters of infrared radiation. Their tops radiate infrared energy upward and their bases radiate energy back to Earth’s surface where it is absorbed and, in a sense, radiated back to the clouds. This process keeps calm, cloudy nights warmer than calm, clear ones. If the clouds remain into the next day, they prevent much of the sunlight from reaching the ground by reflecting it back to space. Since the ground does not heat up as much as it would in full sunshine, cloudy, calm days are normally cooler than clear, calm days. Hence, the presence of clouds tends to keep nighttime temperatures higher and daytime temperatures lower. In summary, the atmospheric greenhouse effect occurs because water vapor, CO2, and other greenhouse gases are selective absorbers. They allow most of the sun’s radiation to reach the surface, but they absorb a good portion of Earth’s outgoing infrared radiation, preventing it from escaping into space. It is the atmospheric greenhouse effect, then, that keeps the temperature of our planet at a level where life can survive. The greenhouse effect is not just a “good thing”—it is essential to life on Earth, for without it, air at the surface would be extremely cold (see Fig. 2.12). ENHANCEMENT OF THE GREENHOUSE EFFECT Although temperature is not measured perfectly at every spot around the globe, more than enough data exist to allow scientists to estimate the global average. Observations indicate that during the past 110 years or so, Earth’s surface air temperature has warmed about 1.0°C (about 1.8°F). Scientific computer models that mathematically simulate the physical processes of the atmosphere, oceans, and ice, predict that should global warming continue unabated, we would be irrevocably committed to major effects from climate change, such as a continuing rise in sea level and a shift in global precipitation patterns. The main cause of this climate change is the greenhouse gas CO2, whose concentration has been increasing primarily due to the burning of fossil fuels and to deforestation (look back at Fig. 1.18, p. 17). However, increasing concentration of other greenhouse gases, such as methane (CH4), nitrous oxide (N2O), and chlorofluorocarbons (CFCs),* has collectively been shown to have an effect approaching that of CO2. In addition, as temperatures warm, more water vapor is added to the air from the world’s oceans. Overall, water vapor accounts for about 60 percent of the atmospheric greenhouse effect, CO2 accounts for about 26 percent, methane about 7 percent, and the remaining greenhouse gases about 7 percent. *To refresh your memory, recall from Chapter 1 that CFCs were once the most widely used propellant in spray cans.

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FIGURE 2.12 (a) Near the surface in an atmosphere with little or no greenhouse gases, Earth’s surface would constantly emit FIGUR infrared (IR) radiation upward, both during the day and at night. Incoming energy from the sun would equal outgoing energy from the surface, but the surface would receive virtually no IR radiation from its lower atmosphere. (i.e., there would be no atmospheric greenhouse effect.) Earth’s surface air temperature would be quite low, and small amounts of water found on the planet would be in the form of ice. (b) In an atmosphere with greenhouse gases, Earth’s surface not only receives energy from the sun but also infrared energy from the atmosphere. Incoming energy still equals outgoing energy, but the added IR energy from the greenhouse gases raises Earth’s average surface temperature to a more habitable level.

Presently, the concentration of CO2 in a volume of air near the surface is just over 0.04 percent, and it is increasing each year. Climate models predict that a continuing increase of CO2 and other greenhouse gases will cause Earth’s current average surface temperature to rise an additional 1°C to 3°C (1.8°F to 5.4°F) or more by the end of this century. How can increasing such a small quantity of CO2 and adding miniscule amounts of other greenhouse gases bring about such a large temperature increase? Mathematical climate models predict that rising ocean temperatures will cause an increase in evaporation rates. The added water vapor—the primary greenhouse gas—will enhance the atmospheric greenhouse effect and roughly double the temperature rise in what is known as a positive feedback. But there are other feedbacks to consider.* The two potentially largest and least understood feedbacks in the climate system are the clouds and the oceans. Clouds can change area, depth, and radiation properties simultaneously with climatic changes. The net effect of all these changes is not totally clear. Oceans, on the other hand, cover 70 percent of the planet. The response of ocean circulations, ocean temperatures, and sea ice to global warming will determine the global *A feedback is a process whereby an initial change in a process will tend to either reinforce the process (positive feedback) or weaken the process (negative feedback). The water vapor–greenhouse feedback is a positive feedback because the initial increase in temperature is reinforced by the addition of more water vapor, which absorbs more of Earth’s infrared energy, thus strengthening the greenhouse effect and enhancing the warming.

pattern and speed of climate change. Unfortunately, it is not now known how quickly each of these feedbacks will respond. Satellite data and computer simulations suggest that clouds overall appear to cool Earth’s climate, as they reflect and radiate away more energy than they retain. (Earth would be about 5°C [9°F] warmer if no clouds were present.) An increase in global cloudiness (if it were to occur) might therefore offset some of the global warming brought on by an enhanced atmospheric greenhouse effect. If clouds were to act on the climate system in this manner, they would provide a negative feedback on climate. The actual result would depend on what types of clouds were present, because some clouds are more reflective and have a stronger cooling effect than others. The most recent models tend to show that changes in clouds as a whole would most likely allow for more heat to be retained, thus providing a small positive feedback on the climate system. Uncertainties unquestionably exist about exactly how much the increasing levels of CO2 and other greenhouse gases will enhance the atmospheric greenhouse effect. Nonetheless, the most recent studies strongly agree that climate change is presently occurring worldwide owing primarily to increasing levels of greenhouse gases. Evidence for this conclusion comes from increases in global average air and ocean temperatures as well as from widespread melting of snow and ice, rising sea levels, and other conditions. These changes are consistent with the effects one would expect from the increase in greenhouse gases. (We will examine the important topic of climate change in more detail in Chapter 13.) WARMING AND COOLING EARTH AND ITS ATMOSPHERE

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39

BRIEF REVIEW

DID YOU KNOW?

In the last several sections, we have explored examples of some of the ways radiation is absorbed and emitted by various objects. Before we continue, here are a few important facts and principles: ●

All objects with a temperature above absolute zero emit radiation.

The higher an object’s temperature, the greater the amount of radiation emitted per unit surface area and the shorter the wavelength of maximum emission.

Earth absorbs solar radiation only during the daylight hours; however, it emits infrared radiation continuously, both during the day and at night.

Earth’s surface behaves as a blackbody, making it a much better absorber and emitter of radiation than the atmosphere.

Water vapor and carbon dioxide are important greenhouse gases that selectively absorb and emit infrared radiation, thereby keeping Earth’s average surface temperature warmer than it otherwise would be.

Cloudy, calm nights are often warmer than clear, calm nights because clouds strongly emit infrared radiation back to Earth’s surface.

It is not the greenhouse effect itself that is of concern, but the enhancement of it due to increasing levels of greenhouse gases.

As greenhouse gases continue to increase in concentration, the average surface air temperature is projected to rise substantially by the end of this century.

With these concepts in mind, we will first examine how the air near the ground warms; then we will consider how Earth and its atmosphere maintain a yearly energy balance. WARMING THE AIR FROM BELOW If you look back at Fig. 2.11 on p. 37, you’ll notice that the atmosphere does not readily absorb radiation with wavelengths between 0.3 and 1.0 m, the region where the sun emits most of its

The average surface temperature on Earth was about 0.2°C (0.36°F) warmer during the first decade of this century (2001–2010) than during the previous decade (1991–2000). If this warming trend were to continue at this rate, the twentyfirst century would warm by 2°C (3.6°F) compared to a warming of around 0.6°C (1.1°F) during the last century.

energy. Consequently, on a clear day, solar energy passes through the lower atmosphere with little effect upon the air. Ultimately it reaches the surface, warming it (see Fig. 2.13). Air molecules in contact with the heated surface bounce against it, gain energy by conduction, then shoot upward like freshly popped kernels of corn, carrying their energy with them. Because the air near the ground is very dense, these molecules only travel a short distance before they collide with other molecules. During the collision, these more rapidly moving molecules share their energy with less energetic molecules, raising the average temperature of the air. But air is such a poor heat conductor that this process is only important within a few centimeters of the ground. As the surface air warms, it actually becomes less dense than the air directly above it. The warmer air rises and the cooler air sinks, setting up thermals, or free convection cells, that transfer heat upward and distribute it through a deeper layer of air. The rising air expands and cools, and, if sufficiently moist, the water vapor condenses into cloud droplets, releasing latent heat that warms the air. Meanwhile, Earth constantly emits infrared energy. Some of this energy is absorbed by greenhouse gases (such as water vapor and carbon dioxide) that emit infrared energy upward and downward, back to the surface. Since the concentration of water vapor decreases rapidly above Earth, most of the absorption occurs in a layer near the surface. Hence, the lower atmosphere is mainly heated from the ground upward.

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FIGURE 2.13 Air in the lower FIGUR atmosphere is heated from the ground upward. Sunlight warms the ground, and the air above is warmed by conduction, convection, and infrared radiation. Further warming occurs during condensation as latent heat is given up to the air inside the cloud.

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FIGURE 2.15 At noon, the sun usually appears a bright white. FIGUR At sunrise and at sunset, sunlight must pass through a thick portion of the atmosphere. Much of the blue light is scattered out of the beam (as illustrated by arrows), causing the sun to appear more red.

SHORTWAVE RADIATION STREAMING FROM THE SUN As the sun’s radiant energy travels through space, essentially nothing interferes with it until it reaches the atmosphere. At the top of the atmosphere, solar energy received on a surface perpendicular to the sun’s rays appears to remain fairly constant at nearly two calories on each square centimeter each minute, or 1361 W/m2— a value called the solar constant.* When solar radiation enters the atmosphere, a number of interactions take place. For example, some of the energy is absorbed by gases, such as ozone, in the upper atmosphere. Moreover, when sunlight strikes very small objects, such as air molecules and dust particles, the light itself is deflected in all directions—forward, sideways, and backward. The distribution of light in this manner is called scattering. (Scattered light is also called diffuse light.) Because air molecules are much smaller than the wavelengths of visible light, they are more effective scatterers of the shorter (blue) wavelengths than the longer (red) wavelengths (see Fig. 2.14). Hence, when we look away from the direct beam of sunlight, blue light strikes our eyes from all directions, turning the daytime sky blue. At midday, all the wavelengths of visible light from the sun strike our eyes, and the sun is perceived as white (see Fig. 2.15). At sunrise and sunset, when the white beam of sunlight must pass through a thick portion of the atmosphere, scattering by air molecules removes the blue light, leaving the longer wavelengths of red, orange, and yellow to pass on through, creating the image of a ruddy or yellowish sun (see Fig. 2.16).

Sunlight can be reflected from objects. Generally, reflection differs from scattering in that during the process of reflection more light is sent backward. Albedo is the percent of radiation returning from a given surface compared to the amount of radiation initially striking that surface. Albedo, then, represents the reflectivity of the surface. In ▼ Table 2.2, notice that thick clouds have a higher albedo than thin clouds. On the average, the albedo of clouds is near 60 percent. When solar energy strikes a surface covered with snow, up to 95 percent of the sunlight may be reflected. Most of this energy is in the visible and ultraviolet wavelengths. Consequently, reflected radiation, coupled with direct sunlight, can produce severe sunburns on the exposed skin of unwary snow skiers, and unprotected eyes can suffer the agony of snow blindness.

*By definition, the solar constant (which, in actuality, is not “constant”) is the rate at which radiant energy from the sun is received on a surface at the outer edge of the atmosphere perpendicular to the sun’s rays when Earth is at an average distance from the sun. Satellite measurements suggest the solar constant varies slightly as the sun’s radiant output varies. The latest measurements and laboratory tests indicate that the solar constant is around 1361 W/m2. It rises by about 1 W/m2 during peaks of solar activity, which occur about every 11 years.

© C. Donald Ahrens

FIGURE 2.14 The scattering of light by air molecules. Air molFIGUR ecules tend to selectively scatter the shorter (violet, green, and blue) wavelengths of visible white light more effectively than the longer (orange, yellow, and red) wavelengths.

FIGURE 2.16 A red sunset produced by the process of scattering. FIGUR WARMING AND COOLING EARTH AND ITS ATMOSPHERE

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41

Typical Albedo of Various Surfaces

SURFACE

ALBEDO (PERCENT)

Fresh snow

75 to 95

Clouds (thick)

60 to 90

Clouds (thin)

30 to 50

Venus

78

Ice

30 to 40

Sand

15 to 45

Earth and atmosphere

30

Mars

17

Grassy field

10 to 30

Dry, plowed field

5 to 20

Water

10*

Forest

3 to 10

Moon

7

*Daily average.

Water surfaces, on the other hand, reflect only a small amount of solar energy. For an entire day, a smooth water surface will have an average albedo of about 10 percent. Averaged for an entire year, Earth and its atmosphere (including its clouds) will redirect about 30 percent of the sun’s incoming radiation back to space, which gives Earth and its atmosphere a combined albedo of 30 percent (see Fig. 2.17). EARTH’S ANNUAL ENERGY BALANCE Although the average temperature at any one place may vary considerably from year to year, Earth’s overall average equilibrium temperature changes only slightly from one year to the next. Each year, then, Earth and its atmosphere combined must send off into space just as much energy as they receive from the sun. The same type of energy balance must exist between Earth’s surface and the atmosphere; that is, each year, Earth’s surface must return to the atmosphere the same amount of energy that it absorbs. If this did not occur, Earth’s average surface temperature would change. How do Earth and its atmosphere maintain this yearly energy balance? Suppose 100 units of solar energy reach the top of Earth’s atmosphere. We already know from Fig. 2.17 that, on the average, clouds, Earth, and the atmosphere reflect and scatter 30 units back to space, and that the atmosphere and clouds together absorb 19 units, which leaves 51 units of direct and indirect (diffuse) solar radiation to be absorbed at Earth’s surface. Figure 2.18 shows approximately what happens to the solar radiation that is absorbed by the surface and the 42

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▼ Table 2.2

FIGURE 2.17 On the average, of all the solar energy that FIGUR reaches Earth’s atmosphere annually, about 30 percent (30⁄100) is reflected and scattered back to space, giving Earth and its atmosphere an albedo of 30 percent. Of the remaining solar energy, about 19 percent is absorbed by the atmosphere and clouds, and about 51 percent is absorbed at the surface.

atmosphere. Out of 51 units reaching the surface, a large amount (23 units) is used to evaporate water, and about 7 units are lost through conduction and convection, which leaves 21 units to be radiated away as infrared energy. Look closely at Fig. 2.18 and notice that Earth’s surface actually radiates upward a whopping 117 units. It does so because, although it receives solar radiation only during the day, it constantly emits infrared energy both during the day and at night. Additionally, the atmosphere above only allows a small fraction of this energy (6 units) to pass through into space. The majority of it (111 units) is absorbed mainly by the greenhouse gases water vapor and CO2, and by clouds. Much of this energy (96 units) is then radiated back to Earth, producing the atmospheric greenhouse effect. Earth’s surface receives nearly twice as much longwave infrared energy from the atmosphere as it does shortwave radiation from the sun. In all these exchanges, notice that the energy lost at Earth’s surface (147 units) is exactly balanced by the energy gained there (147 units). A similar balance exists between Earth’s surface and its atmosphere. Again, observe in Fig. 2.18 that the energy gained by the atmosphere (160 units) balances the energy lost. Moreover, averaged for an entire year, the solar energy received at Earth’s surface (51 units) and that absorbed by Earth’s atmosphere (19 units) balances the infrared energy lost to space by Earth’s surface (6 units) and its atmosphere (64 units). We can see the effect that conduction, convection, and latent heat have in the warming of the atmosphere if we look at the energy balance only in radiative terms. Earth’s surface receives 147 units of radiant energy from the sun

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FIGURE 2.18 FIGUR

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Earth-atmosphere energy balance. Numbers represent approximations based on surface observations and satellite data. While the actual value of each process may vary by several percent, it is the relative size of the numbers that is important.

of energy received each year balance the amount lost. We might conclude that polar regions are growing colder each year, while tropical regions are becoming warmer. But they are not. To compensate for these gains and losses of energy, winds in the atmosphere and currents in the oceans circulate warm air and water toward the poles, and cold air and water toward the equator. The transfer of heat energy by atmospheric and oceanic circulations prevents low latitudes from steadily becoming warmer and high latitudes from steadily growing colder. These circulations are extremely important to weather and climate, and will be treated more completely in Chapter 7.

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and its own atmosphere, while it radiates away 117 units, producing a surplus of 30 units. The atmosphere, on the other hand, receives 130 units (19 units from the sun and 111 from Earth), while it loses 160 units, producing a deficit of 30 units. The balance (30 units) is the warming of the atmosphere produced by the heat transfer processes of conduction and convection (7 units) and by the release of latent heat (23 units). So, Earth and the atmosphere absorb energy from the sun as well as from each other. In all of the energy exchanges, a delicate balance is maintained. Essentially, there is no yearly gain or loss of total energy, and the average temperature of Earth and the atmosphere remains fairly constant from one year to the next. This equilibrium does not imply that Earth’s average temperature does not change, but rather that the changes are small from year to year (usually less than one-tenth of a degree Celsius) and become significant only when measured over many years. One example is the temperature increase of 0.6°C (1.1°F) during the last century produced by the addition of greenhouse gases to the atmosphere. The added gases led to a radiative imbalance as more infrared radiation is trapped by the atmosphere. However, as Earth warms, the balance is restored. Even though Earth and the atmosphere together maintain an annual energy balance, such a balance is not maintained at each latitude. High latitudes tend to lose more energy to space each year than they receive from the sun, while low latitudes tend to gain more energy during the course of a year than they lose. From Fig. 2.19 we can see that only at middle latitudes near 38° does the amount

FIGURE 2.19 The average annual incoming solar radiation FIGUR (yellow arrows) absorbed by Earth and the atmosphere along with the average annual infrared radiation (red arrows) emitted by Earth and the atmosphere. WARMING AND COOLING EARTH AND ITS ATMOSPHERE

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43

FOCUS

ON A SPECIAL TOPIC 2.3 O

The sun is our nearest star, located some 150 million kilometers (93 million miles) from earth. Even at that distance, it provides enough energy for life to flourish. The sun is a giant celestial furnace with a core temperature estimated to be near 15 million degrees Celsius. Here, thermonuclear processes generate an enormous amount of energy, which gradually works its way outward. At its surface, the sun emits vast quantities of radiant energy. It also generates very large amounts of magnetism, and a constant stream of charged particles that stream outward in all directions from its surface. Sometimes these affect our atmosphere, producing a variety of effects known as space weather weather. The eerie yet beautiful light show called the aurora is one of the most commonly observed aspects of space weather. At high latitudes after darkness has fallen, a faint, white glow may appear in the sky. Lasting from a few minutes to a few hours, the light may move across the sky as a yellow-green arc much wider than a rainbow; or, it may faintly decorate the sky with flickering draperies of blue, green and purple light that constantly changes in form and location, as if blown by a gentle breeze. This impressive light show is the aurora (see Fig. 4). The aurora is caused by charged partiparti cles from the sun interacting with our

© Lindsey P. Martin Photography

Space Weather and Its Impact on Earth

FIGURE 4 The aurora borealis is a phenomenon that forms as energetic particles FIGUR from the sun interact with Earth’s atmosphere.

atmosphere, almost always in the thermosphere, more than 80 kilometers (50 miles) above Earth’s surface. From the sun and its tenuous atmosphere comes a continuous discharge of particles. This discharge happens because, at extremely high temperatures, gases become stripped of electrons by violent collisions and acquire enough speed to escape the gravitational pull of the sun. As these charged particles (ions and electrons) travel through space, they are known as the solar wind. When the solar

We now turn our attention to how solar energy produces Earth’s seasons. Before doing so, you may wish to read Focus section 2.3 to learn how the sun’s energy produces space weather and a dazzling light show known as the aurora.

Why Earth Has Seasons Earth revolves completely around the sun in an elliptical path (not quite a circle) in about 365 days and six hours (one year, plus a Leap Day every four years in February). As Earth revolves around the sun, it spins on its own axis, completing one spin in 24 hours (one day). The average distance from Earth to the sun is 150 million km (93 million mi). Because Earth’s orbit is an ellipse instead of a circle, and is slightly off-center from the sun, the 44

wind moves close enough to Earth, it interacts with Earth’s magnetic field, disturbing it (see Fig. 5). This disturbance causes energetic solar wind particles to enter the upper atmosphere, where they collide with atmospheric gases. These gases then become excited and emit visible radiation (light), which causes the sky to glow like a neon light, thus producing the aurora. In the Northern Hemisphere, the aurora is called the aurora borealis, or northern lights; its counterpart in the Southern

actual distance from Earth to the sun varies during the year. Earth comes closer to the sun in January (147 million km) than it does in July (152 million km).* (See Fig. 2.20.) We might conclude from this fact that our warmest weather should occur in January and our coldest weather in July. But, in the Northern Hemisphere, we normally experience cold weather in January when we are closer to the sun and warm weather in July when we are farther away. If nearness to the sun were the primary cause of the seasons then, indeed, January would be warmer than July. However, nearness to the sun is only a small part of the story. Our seasons are regulated by the amount of solar energy received at Earth’s surface. This amount is determined *The time around January 3, when Earth is closest to the sun, is called perihelion (from the Greek peri, meaning “near” and helios, meaning “sun”). The time when Earth is farthest from the sun (around July 4) is called aphelion (from the Greek ap, meaning “away from”).

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FIGURE 5 The stream FIGUR of charged particles from the sun—called the solar wind wind—distorts Earth’s magnetic field. These particles, which spiral in along magnetic field lines, interact with atmospheric gases, and produce the aurora.

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Hemisphere is the aurora australis, or southern lights. The aurora is most frequently seen in the polar regions, where Earth’s magnetic field lines emerge from Earth. Auroral displays are most frequent and intense during the active part of the sun’s cyclic activity, which peaks about every 11 years. During active solar periods, there are numerous sunspots (huge cooler regions on the sun’s surface) and giant flares (solar eruptions) that send large quantities of solar wind particles traveling away from the sun at high speeds (hundreds of kilometers a second). When one of these eruptions is pointed toward Earth, the energetic particles may be able to penetrate unusually deep into Earth’s magnetic field, creating conditions referred to as solar storms. During these conditions in North America, we see the aurora much farther south than usual. Along with the beauty of the aurora, solar storms can produce many negative effects. In 1859, a huge solar storm disrupted telegraph operations around the world. Another solar storm in March 1989 caused millions of people across Quebec to lose electrical power for several hours. Airlines sometimes reroute planes away from polar regions during the most intense solar storms, which reduces the risk of interference with radio communications.

Solar storms can also affect satellite operations, which is a growing concern because of the many types of satellites now used for telecommunications, navigation, and other purposes. A $630 million research satellite from Japan failed during an intense solar storm in October 2003. Because solar storms heat the outer atmosphere, they can increase the drag on satellites and reduce their lifespan in orbit. Space weather was on the quiet side during the solar minimum of 2008–2009, when the number of sunspots observed was at its lowest level in nearly a century.

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primarily by the angle at which sunlight strikes the surface and by how long the sun shines on any latitude (daylight hours). Let’s look more closely at these factors. Solar energy that strikes Earth’s surface perpendicularly (directly) is much more intense than solar energy that strikes the same surface at an angle. Think of shining

FIGURE 2.20 The elliptical path (highly exaggerated) of Earth FIGUR about the sun brings Earth slightly closer to the sun in January than in July.

Activity was also on the low side during the solar maximum that peaked in 2013–2014, which was the weakest maximum in more than a century. Scientists cannot say how long this tendency toward lessened activity during both maximum and minimum might continue, because techniques for predicting the strength of solar cycles are still being researched and tested.

a flashlight straight at a wall—you get a small, circular spot of light (see Fig. 2.21). Now tip the flashlight and notice how the spot of light spreads over a larger area. The same principle holds for sunlight. Sunlight striking Earth at an angle spreads out and must heat a larger region than sunlight impinging directly on Earth. Everything else being equal, an area experiencing more direct solar rays will receive more heat than the same size area being struck by sunlight at an angle. In addition, the more the sun’s rays are slanted from the perpendicular, the more atmosphere they must penetrate. And the more atmosphere they penetrate, the more they can be scattered and absorbed. As a consequence, when the sun is high in the sky, it can heat the ground to a much higher temperature than when it is low on the horizon. The second important factor determining how warm Earth’s surface becomes is the length of time the sun shines WARMING AND COOLING EARTH AND ITS ATMOSPHERE

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FIGURE 2.21 Sunlight that strikes FIGUR a surface at an angle is spread over a larger area than sunlight that strikes the surface directly. Oblique sun rays deliver less energy (are less intense) to a surface than direct sun rays.

SEASONS IN THE NORTHERN HEMISPHERE Notice in Fig. 2.22 that on June 21, the northern half of the world is directed toward the sun. At noon on this day, solar rays beat down upon the Northern Hemisphere more directly than during any other time of year. The sun is at its highest position in the noonday sky, directly above 231⁄2° north (N) latitude (Tropic of Cancer). If you were standing at this latitude on June 21, the sun at noon would be directly overhead. This day, called the summer solstice, is the astronomical first day of summer in the Northern Hemisphere.* Study Fig. 2.22 closely and notice that, as Earth spins on its axis, the side facing the sun is in sunshine and the other side is in darkness. Thus, half of the globe is always illuminated. If Earth’s axis were not tilted, the noonday sun would always be directly overhead at the equator, *As we will see later in this chapter, the seasons are reversed in the Southern Hemisphere. Hence, in the Southern Hemisphere, this same day is the winter solstice, or the astronomical first day of winter.

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each day, the number of daylight hours. Longer daylight hours, of course, mean that more energy is available from sunlight. In a given location, more solar energy reaches Earth’s surface on a clear, long day than on a day that is clear but much shorter. The longer the day, the more surface heating takes place. From a casual observation, we know that summer days have more daylight hours than winter days. Also, the noontime summer sun is higher in the sky than is the noontime winter sun. Both of these events occur because our spinning planet is inclined on its axis (tilted) as it revolves around the sun. As Fig. 2.22 illustrates, the angle of tilt is 231⁄2° from the perpendicular drawn to the plane of Earth’s orbit. Earth’s axis points to the same direction in space all year long; thus, on one side of Earth’s orbit the Northern Hemisphere is tilted toward the sun in summer (June), and on the other side of Earth’s orbit it is tilted away from the sun in winter (December).

FIGURE 2.22 As Earth revolves about the sun, it is tilted on its axis by an angle of 231⁄2°. Earth’s axis always points to the same area in space FIGUR

(as viewed from a distant star). Thus, in June, when the Northern Hemisphere is tipped toward the sun, more direct sunlight and long hours of daylight cause warmer weather than in December, when the Northern Hemisphere is tipped away from the sun. (Diagram, of course, is not to scale. The dates of each solstice and equinox may be a day earlier or later than shown here, depending on the year and on your time zone.)

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FIGURE 2.23 The number of hours of daylight (hours and FIGUR minutes) for different latitudes on June 21. Notice that daylight hours range from 24 at 66 1⁄2°N to zero at 66 1⁄2°S. Also notice that sunlight that reaches Earth’s surface in far northern latitudes has passed through a thicker layer of absorbing, scattering, and reflecting atmosphere than sunlight that reaches Earth’s surface farther south. Sunlight is lost through both the thickness of the pure atmosphere and by impurities in the atmosphere. As the sun’s rays become more oblique, these effects become more pronounced.

and there would be 12 hours of daylight and 12 hours of darkness at each latitude every day of the year. However, Earth is tilted. Since the Northern Hemisphere faces toward the sun on June 21, each latitude in the Northern Hemisphere will have more than 12 hours of daylight. The farther north we go, the longer are the daylight hours. When we reach the Arctic Circle (661⁄2°N), daylight lasts for 24 hours, as the sun does not set. Notice in Fig. 2.22 and Fig. 2.23 how the region above 661⁄2°N never gets into the “shadow” zone as Earth spins. At the North Pole,

the sun actually rises above the horizon on March 20 and has six months until it sets on September 22. No wonder this region is called the “Land of the Midnight Sun”! (See Fig. 2.24.) Even though in the far north the sun is above the horizon for many hours during the summer (see ▼ Table 2.3, p. 48), the surface air there is not warmer than the air farther south, where days are appreciably shorter. Why this is so is shown in Fig. 2.23. When incoming solar radiation (called insolation) enters the atmosphere, fine dust, air molecules, and clouds reflect and scatter it, and some of it is absorbed by atmospheric gases. Generally, the greater the thickness of atmosphere that sunlight must penetrate, the greater the chances that it will be either reflected or absorbed by the atmosphere. During the summer in far northern latitudes, the sun is never very high above the horizon, so its radiant energy must pass through a thick portion of atmosphere before it reaches Earth’s surface. Some of the solar energy that does reach the surface melts frozen soil or is reflected by snow or ice. That which is absorbed is spread over a large area. So, even though northern cities may experience long hours of sunlight, they are cooler than cities farther south, because overall, they receive less radiation at the surface. What radiation they do receive does not heat the surface as effectively. Look at Fig. 2.22 again and notice that, by September 22, Earth will have moved so that the sun is directly above the equator. Except at the poles, the days and nights throughout the world are of equal length. This day is called the autumnal (fall) equinox, and it marks the astronomical beginning of fall in the Northern Hemisphere. At the North Pole, the sun appears on the horizon for 24 hours, due to the bending of light by the atmosphere. The following day (or at least within several days), the sun disappears from view, not to rise again for a long, cold six months. Throughout the northern half of the world on each successive day, there are fewer hours of daylight, and the noon sun is slightly lower in the sky. Less direct

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FIGURE 2.24 Land of the MidFIGUR night Sun. A series of exposures of the sun taken before, during, and after midnight in northern Alaska during July.

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▼ Table 2.3

Length of Time from Sunrise to Sunset for Various Latitudes on Different Dates in the Northern Hemisphere

LATITUDE

MARCH 20

JUNE 21

SEPT. 22

DEC. 21

12 hr

12.0 hr

12 hr

12.0 hr

10°

12 hr

12.6 hr

12 hr

11.4 hr

20°

12 hr

13.2 hr

12 hr

10.8 hr

30°

12 hr

13.9 hr

12 hr

10.1 hr

40°

12 hr

14.9 hr

12 hr

9.1 hr

50°

12 hr

16.3 hr

12 hr

7.7 hr

60°

12 hr

18.4 hr

12 hr

5.6 hr

70°

12 hr

2 months

12 hr

0 hr

80°

12 hr

4 months

12 hr

0 hr

90°

12 hr

6 months

12 hr

0 hr

sunlight and shorter hours of daylight spell cooler weather for the Northern Hemisphere. Reduced sunlight, lower air temperatures, and cooling breezes stimulate the beautiful pageantry of fall colors (see Fig. 2.25). In some years around the middle of autumn, there is an unseasonably warm spell, especially in the eastern twothirds of the United States. This warm period, referred to as Indian summer,* may last from several days up to a week or more. It usually occurs when a large high-pressure area stalls near the southeast coast. The clockwise flow of air around this system moves warm air from the Gulf of

FIGURE 2.25 The pageantry FIGUR of fall colors in New England. The weather most suitable for an impressive display of fall colors is warm, sunny days followed by clear, cool nights with temperatures dropping below 7°C (45°F), but remaining above freezing. Contrary to popular belief, it is not the first frost that causes the leaves of deciduous trees to change color. The yellow and orange colors, which are actually in the leaves, typically begin to show through several weeks before the first frost, as shorter days and cooler nights cause a decrease in the production of the green pigment chlorophyll.

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*The origin of the term is uncertain, but it dates back to the eighteenth century. It may have originally referred to the good weather that allowed Native Americans time to harvest their crops. Normally, a period of cool autumn weather must precede the warm weather period for the latter to be called Indian summer.

Mexico into the central or eastern half of the nation. The warm, gentle breezes and smoke from a variety of sources respectively make for mild, hazy days. The warm weather ends abruptly when an outbreak of polar air reminds us that winter is not far away. On December 21 (three months after the autumnal equinox), the Northern Hemisphere is tilted as far away from the sun as it will be all year (see Fig. 2.22, p. 46). Nights are long and days are short. Notice in Table 2.3 that daylight decreases from 12 hours at the equator to 0 (zero) at latitudes above 661⁄2°N. This is the shortest day of the year, called the winter solstice—the astronomical beginning of winter in the northern world. On this day, the sun shines directly above latitude 231⁄2°S (Tropic of Capricorn). In the northern half of the world, the sun is at its lowest position in the noon sky. Its rays pass through a

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FOCUS

ON A SPECIAL TOPIC 2.4 O

On December 21 (or 22, depending on the year) after nearly a month of cold weather, and perhaps a snowstorm or two (see Fig. 6), someone on the radio or television has the audacity to proclaim that “today is the first official day of winter.” If during the last several weeks it was not winter, then what season was it? Actually, December 21 marks the astronomical first day of winter in the Northern Hemisphere (NH), just as June 21 marks the astronomical first day of summer (NH). Earth is tilted on its axis by 231⁄2° as it revolves around the sun. This fact causes the sun (as we view it from Earth) to move in the sky from a point where it is directly above 231⁄2° South latitude on December 21 to a point where it is directly above 231⁄2° North latitude on June 21. The astronomical first day of spring (NH) occurs around March 20 as the sun crosses the equator moving northward and, likewise, the astronomical first day of autumn (NH) occurs around September 22 as the sun crosses the equator moving southward. Therefore the “official” beginning of any season is simply the day on which the sun passes over a particular latitude, and has nothing to do with how cold or warm the following day will be. In fact, a period of colder or warmer than normal weather before or after a

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Is December 21 Really the First Day of Winter?

FIGURE 6 Snow covers Central Park in New York City on December 17, 2013. Since FIGUR the snowstorm occurred before the winter solstice, is this a late fall storm or an early winter storm?

solstice or equinox is caused mainly by the upper-level winds directing cold or warm air into a region. In the middle latitudes, summer is defined as the warmest season and winter the coldest season. If the year is divided into four seasons with each season consisting of three months, then the meteorological (or climatological) definition of summer over much of the Northern Hemisphere would be the three warmest months of June, July, and August. Winter would be the three coldest months of December, January, and

thick section of atmosphere and spread over a large area on the surface. With so little incident sunlight, Earth’s surface cools quickly. A blanket of clean snow covering the ground aids in the cooling. In northern Canada and Alaska, arctic air rapidly becomes extremely cold as it lies poised, ready to do battle with the milder air to the south. Periodically, this cold arctic air pushes down into the northern United States, producing a rapid drop in temperature called a cold wave, which occasionally reaches far into the south. Sometimes these cold spells arrive well before the winter solstice—the “official” first day of winter—bringing with them heavy snow and blustery winds. (To learn more about this “official” first day of winter, read Focus section 2.4.) Three months past the winter solstice marks the astronomical arrival of spring, which is called the vernal (spring) equinox. The date is March 20 and, once again,

February. Autumn would be September, October, and November—the transition between summer and winter. And spring would be March, April, and May—the transition between winter and summer. So, the next time you hear someone remark on December 21 that “winter officially begins today,” remember that this is the astronomical definition of the first day of winter. According to the climatological definition, winter has been around for several weeks.

the noonday sun is shining directly on the equator, days and nights throughout the world are of equal length, and, at the North Pole, the sun rises above the horizon after a long six-month absence. At this point it is interesting to note that although sunlight is most intense in the Northern Hemisphere on June 21, the warmest weather in middle latitudes normally occurs weeks later, usually in July or August. This situation (called the lag in seasonal temperature) arises because although incoming energy from the sun is greatest in June, it takes time for oceans and landmasses to release the large amounts of incoming energy they have absorbed. As a result, incoming energy still exceeds outgoing energy from Earth for a period of at least several weeks. Once the incoming solar energy and outgoing earth energy are in balance, the highest average temperature is attained. When outgoing energy exceeds incoming energy, the average temperature drops. As in WARMING AND COOLING EARTH AND ITS ATMOSPHERE

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DID YOU KNOW? Seasonal changes can affect how we feel. For example, some people face each winter with a sense of foreboding, especially at high latitudes where days are short and nights are long and cold. If the depression is lasting and disabling, the problem is called seasonal affective disorder (SAD). People with SAD tend to sleep longer, overeat, and feel tired and drowsy during the day. The treatment is usually extra doses of bright light.

SEASONS IN THE SOUTHERN HEMISPHERE On June 21, the Southern Hemisphere is adjusting to an entirely different season. Again, look back at Fig. 2.22 (p. 46), and notice that this part of the world is now tilted away from the sun. Nights are long, days are short, and solar rays come in at a low angle (see Fig. 2.26f). All of these factors keep air temperatures fairly low. The June solstice marks the astronomical beginning of winter in the Southern Hemisphere. In this part of the world, summer will not “officially” begin until the sun is over *Calculating the noon angle of the sun for any latitude is easy. First, determine the number of degrees between your latitude and the latitude where the sun is directly overhead. Then subtract this number from 90°. The result gives you the elevation of the sun above the southern horizon at noon at your latitude.

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the summer, there is also a seasonal temperature lag in winter. Because outgoing Earth energy exceeds incoming solar energy well past the winter solstice (December 21), we normally find our coldest weather occurring in January or February. Up to now, we have seen that the seasons are controlled by the amount of solar energy striking our tilted planet as it makes its annual voyage around the sun. And we know that the tilt of Earth causes a seasonal variation in both the length of daylight and the intensity of sunlight that reaches the surface. These facts are summarized in Fig. 2.26, which shows how the sun would appear in the sky to an observer at various latitudes at different times of the year. Earlier we learned that at the North Pole the sun rises above the horizon in March and stays above the horizon for six months, until September. Notice in Fig. 2.26a that at the North Pole even when the sun is at its highest point in June, it is low in the sky—only 231⁄2° above the horizon. Farther south, at the Arctic Circle (Fig. 2.26b), the sun is always fairly low in the sky, even in June, when the sun stays above the horizon for 24 hours.

In the middle latitudes (Fig. 2.26c), notice that in December the sun rises in the southeast, reaches its highest point at noon (only about 26° above the southern horizon), and sets in the southwest.* This apparent path produces little intense sunlight and short daylight hours. On the other hand, in June, the sun rises in the northeast, reaches a much higher position in the sky at noon (about 74° above the southern horizon) and sets in the northwest. This apparent path across the sky produces more intense solar heating, longer daylight hours, and, of course, warmer weather. Figure 2.26d illustrates how the tilt of Earth influences the sun’s apparent path across the sky at the Tropic of Cancer (231⁄2°). Figure 2.26e gives the same information for an observer at the equator.

FIGURE 2.26 The apparent path of the sun across the sky as observed at different latitudes on the June solstice (June 21), FIGUR the December solstice (December 21), and the equinoxes (March 20 and September 22).

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LOCAL SEASONAL VARIATIONS Look at Fig. 2.26c and observe that in the middle latitudes of the Northern Hemisphere objects facing south will receive more sunlight during a year than those facing north. This fact becomes strikingly apparent in hilly or mountainous country. Hills that face south receive more sunshine and, hence, become warmer than the partially shielded north-facing hills. Higher temperatures usually mean greater rates of evaporation and slightly drier soil conditions. Thus, southfacing hillsides are usually warmer and drier as compared to north-facing slopes at the same elevation. In many areas of the far western United States, only sparse vegetation grows on south-facing slopes, while, on the same hill, dense vegetation grows on the cool, moist slopes that face north (see Fig. 2.27). In the mountains, snow usually lingers on the ground for a longer time on north slopes than on the warmer south slopes. For this reason, ski runs are built facing *For a comparison of January and July temperatures see Figs. 3.14 and 3.15, p. 66.

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the Tropic of Capricorn (231⁄2°S)—remember that this occurs on December 21. So, when it is winter and June in the Southern Hemisphere, it is summer and June in the Northern Hemisphere. Conversely, when it is summer and December in the Southern Hemisphere, it is winter and December in the Northern Hemisphere. If you are tired of the cold, December weather in your Northern Hemisphere city, travel to the summer half of the world and enjoy the warmer weather. The tilt of Earth as it revolves around the sun makes all this possible. We know Earth comes nearer to the sun in January than in July. Even though this difference in distance amounts to only about 3 percent, the energy that strikes the top of Earth’s atmosphere is almost 7 percent greater on January 3 than on July 4. These numbers might lead us to believe that summer should be warmer in the Southern Hemisphere than in the Northern Hemisphere. However, this is not so. A close examination of the Southern Hemisphere reveals that nearly 81 percent of the surface is water compared to 61 percent in the Northern Hemisphere. The added solar energy due to the closeness of the sun is absorbed by large bodies of water, becoming well mixed and circulated within them. This process keeps the average summer (January) temperatures in the Southern Hemisphere cooler than the average summer (July) temperatures in the Northern Hemisphere. Because of water’s large heat capacity, it also tends to keep winters in the Southern Hemisphere warmer than we might expect.*

FIGURE 2.27 In the middle latitudes of the Northern HemiFIGUR sphere, where small temperature changes can cause major changes in soil moisture, sparse vegetation on the south-facing slopes will often contrast with lush vegetation on the northfacing slopes.

north wherever possible. Also, homes and cabins built on the north side of a hill usually have a steep pitched roof, as well as a reinforced deck to withstand the added weight of snow from successive winter storms. The seasonal change in the sun’s position during the year can have an effect on the vegetation around the home. In winter, a large two-story home can shade its own north side, keeping it much cooler than its south side. Trees that require warm, sunny weather should be planted on the south side, where sunlight reflected from the house can even add to the warmth. The design of a home can be important in reducing heating and cooling costs. Large windows should face south, allowing sunshine to penetrate the home in winter. To block out excess sunlight during the summer, a small eave or overhang should be built. A kitchen with windows facing east will let in enough warm morning sunlight to help heat this area. Because the west side warms rapidly in the afternoon, rooms having small windows (such as garages) can be placed here to act as a thermal buffer. Deciduous trees planted on the west side of a home provide shade in the summer. In winter, they drop their leaves, allowing the winter sunshine to warm the house. If you like the bedroom slightly cooler than the rest of the home, face it toward the north. Let nature help with the heating and air conditioning. Proper house design, orientation, and landscaping can help cut the demand for electricity, as well as for natural gas and fossil fuels, which are rapidly being depleted.

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SUMMARY In this chapter, we looked at the concepts of heat and temperature and learned that latent heat is an important source of atmospheric heat energy. We also learned that the transfer of heat can take place by conduction, convection, and radiation—the transfer of energy by means of electromagnetic waves. The hot sun emits most of its radiation as shortwave radiation. A portion of this energy heats Earth, and Earth, in turn, warms the air above. The cool Earth emits most of its radiation as longwave infrared energy. Selective absorbing greenhouse gases in the atmosphere, such as water vapor and carbon dioxide, absorb some of Earth’s infrared radiation and radiate a portion of it back to the surface, where it warms the surface, producing the atmospheric greenhouse effect. The average equilibrium temperature of Earth and the atmosphere remains fairly constant from one year to the next because the amount of energy they absorb each year is equal to the amount of energy they lose. We examined the seasons and found that Earth has seasons because it is tilted on its axis as it revolves around the sun. The tilt of Earth causes a seasonal variation in both the length of daylight and the intensity of sunlight that reaches the surface. Finally, on a more local setting, we saw that Earth’s inclination influences the amount of solar energy received on the north and south side of a hill, as well as around a home.

KEY TERMS The following terms are listed (with corresponding page numbers) in the order they appear in the text. Define each. Doing so will aid you in reviewing the material covered in this chapter. kinetic energy, 28 temperature, 28 absolute zero, 28 heat, 28 Kelvin scale, 28 Fahrenheit scale, 29 Celsius scale, 29 latent heat, 29 sensible heat, 30 conduction, 31 convection, 31 thermals, 32 advection, 32 52

radiant energy (radiation), 33 electromagnetic waves, 33 micrometer, 33 photons, 33 visible region, 34 ultraviolet radiation (UV), 34 infrared radiation (IR), 34 shortwave radiation, 36 longwave (terrestrial) radiation, 36 blackbody, 36 selective absorbers, 37

radiative equilibrium temperature, 37 greenhouse effect, 38 greenhouse gases, 38 atmospheric window, 38 solar constant, 41 scattering, 41

reflected (light), 41 albedo, 41 summer solstice, 46 autumnal equinox, 47 Indian summer, 48 winter solstice, 48 vernal equinox, 49

QUESTIONS FOR REVIEW . Distinguish between temperature and heat. . How does the average speed (motion) of air molecules relate to the air temperature? . Explain how heat is transferred in our atmosphere by: (a) conduction (b) convection (c) radiation. . What is latent heat? How is latent heat an important source of atmospheric energy? . How does the Kelvin temperature scale differ from the Celsius scale? . How does the amount of radiation emitted by Earth differ from that emitted by the sun? . How does the temperature of an object influence the radiation it emits? . How do the wavelengths of most of the radiation emitted by the sun differ from those emitted by the surface of Earth? . When a body reaches a radiative equilibrium temperature, what is taking place? . Why are carbon dioxide and water vapor called selective absorbing greenhouse gases? . List four important greenhouse gases in Earth’s atmosphere. . Explain how Earth’s atmospheric greenhouse effect works. . What greenhouse gases appear to be responsible for the enhancement of Earth’s greenhouse effect? . Why does the albedo of Earth and its atmosphere average about 30 percent? . How is the lower atmosphere warmed from the surface upward? . Explain how Earth and its atmosphere balance incoming energy with outgoing energy. . In the Northern Hemisphere, why are summers warmer than winters even though Earth is actually closer to the sun in January? . What are the main factors that determine seasonal temperature variations?

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. If it is winter and January in New York City, what is the season and month in Sydney, Australia? . During the Northern Hemisphere’s summer, the daylight hours in northern latitudes are longer than in middle latitudes. Explain why northern latitudes are not warmer. . During July, daylight hours in far northern latitudes of the Northern Hemisphere are longer than daylight hours in the middle latitudes. Explain why far northern latitudes are not warmer than middle latitudes. . Explain why the vegetation on the north-facing side of a hill is frequently different from the vegetation on the south-facing side of the same hill.

QUESTIONS FOR THOUGHT AND EXPLORATION

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. Explain why the bridge in Fig. 2.28 is the first to become icy.

. If the surface of a puddle freezes, is heat energy released to or taken from the air above the puddle? Explain. . In houses and apartments with forced-air furnaces, heat registers are usually placed near the floor rather than near the ceiling. Explain why. . How is heat transferred away from the surface of the moon? (Hint: The moon has no atmosphere.) . Which do you feel would have the greatest effect on Earth’s greenhouse effect: removing all of the CO2 from the atmosphere or removing all of the water vapor? Explain your answer. . How would the seasons be affected where you live if the tilt of Earth’s axis increased from 231⁄2° to 40°? . Explain why an increase in cloud cover surrounding Earth would increase Earth’s albedo, yet not necessarily lead to a lower Earth surface temperature. . Why does the surface temperature often increase on a clear, calm night as a low cloud moves overhead? . Would you expect Earth’s surface temperature to continue to rise if CO2 levels continue to increase but levels of atmospheric water vapor begin to decrease? Explain your reasoning. . Explain (with the aid of a diagram) why the morning sun in the Northern Hemisphere shines brightly through a south-facing bedroom window in December, but not in June. . In New York City, the intensity of sunlight and the number of daylight hours are almost identical on October 21 and February 21. Why, then, in New York City is it normally much colder on February 21?

FIGURE 2.28 FIGUR

Go to the Arctic Circle portal. Under Websites and Blogs, access “Arctic Change: A Near-Realtime Arctic Change Indicator Website.” On the left-hand index, click on “Clouds.” What is the trend in springtime cloudiness (March, April, May) over the Arctic? Would this trend act to reduce or increase the amount of solar radiation reaching the Arctic? What if the same trend were observed in December?

ONLINE RESOURCES Visit www.cengagebrain.com to view additional resources, including video exercises, practice quizzes, an interactive eBook, and more.

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CHAPTER

3

Air Temperature Contents Daily Warming and Cooling Air Near the Surface Applications of Air Temperature Data Air Temperature and Human Comfort Measuring Air Temperature

W

e threw a dish of water high into the air, just to see what would happen. Before it hit the ground, it made a

hissing noise, froze, and fell as tiny round pellets of ice the size of wheat kernels. Ice ce became so hard the ax rebounded from it. At such temperatures, metal snapped, wood became petrified, and rubber was just like cement. The dog’s leather harness could not bend or it would break . . . Becoming lost was of no concern. As an observer walked along, each breath remained as a tiny motionless mist behind him at head level. These patches of human breath fog remained in the still air for three or four minutes before fading away. One person even found such a trail still marking his path when he returned 15 minutes later. Itt was easy to freeze your nose without even knowing it. David Phillips, Blame it on the Weather

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Daily Warming and Cooling Air Near the Surface In Chapter 2, we learned how the sun’s energy coupled with the motions of Earth produce the seasons. In a way, each sunny day is like a tiny season as the air goes through a daily cycle of warming and cooling. The air warms during the morning hours as the sun gradually rises higher in the sky, spreading a blanket of heat energy over the ground. The sun reaches its highest point around noon, after which it begins its slow journey toward the western horizon. It is around noon when Earth’s surface receives the most intense solar rays. However, somewhat surprisingly, noontime is usually not the warmest part of the day. Rather, the air continues to be heated, often reaching a maximum temperature later in the afternoon. To find out why this lag in temperature occurs, we need to examine a shallow layer of air in contact with the ground. DAYTIME WARMING As the sun rises in the morning, sunlight warms the ground, and the ground warms the air in contact with it by conduction. However, air is such a poor heat conductor that this process only takes place within a few centimeters of the ground. As the sun 56

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ur opening vignette is an actual account made by two weather observers on February 3, 1947, at Snag Airport in the Yukon of Canada when the air temperature fell to –63°C (–81°F), the lowest temperature ever measured in North America. It illustrates the profound effect air temperature can have on a variety of things, especially when it drops to extremely low readings. Air temperature is a critical weather element. The careful recording and application of temperature data are tremendously important to us all. Without accurate information of this type, the work of farmers, weather analysts, power company engineers, and many others would be a great deal more difficult, not to mention knowing how to dress for the day. Therefore, we begin this chapter by examining the daily variation in air temperature. We will answer such questions as why the warmest time of the day is normally in the afternoon, why the coldest is usually in the early morning, and why calm, clear nights are usually colder than windy, clear nights. After we examine the factors that cause temperatures to vary from one place to another, we will look at daily, monthly, and yearly temperature averages and ranges with an eye toward practical applications for everyday living. Near the end of the chapter, we will see how air temperature is measured and how the wind can change our perception of air temperature.

FIGURE 3.1 On a sunny, calm day, the air near the surface can be substantially warmer than the air a meter or so above the surface, where thermometers are typically located.

rises higher in the sky, the air in contact with the ground becomes even warmer, and, on a windless day, a substantial temperature difference usually exists between the air at ground level and the air directly above it. This explains why runners on a clear, windless, hot summer afternoon may experience air temperatures of over 50°C (122°F) at their feet and only 35°C (95°F) at their waists (see Fig. 3.1). Near the surface, convection begins, and rising air bubbles (thermals) help to redistribute heat. In calm weather, these thermals are small and do not effectively mix the air near the surface. Thus, large vertical temperature differences can exist. On windy days, however, turbulent eddies can mix hot, surface air with the cooler air above. This form of mechanical stirring, sometimes called forced convection, helps the thermals to transfer heat away from the surface more efficiently. Therefore, on sunny, windy days the temperature difference between the surface air and the air directly above is not as great as it is on sunny, calm days. We can now see why the warmest part of the day is usually in the afternoon. Around noon, the sun’s rays are most intense. Incoming solar radiation decreases in intensity after noon, but for a time it still exceeds outgoing heat energy from the surface. This situation yields an energy surplus for two to four hours after noon and substantially contributes to a lag between the time of maximum solar heating and the time of maximum air temperature. The warmest part of the day several feet above the surface occurs when incoming energy from the sun is balanced by outgoing energy from Earth’s surface (see Fig. 3.2). The exact time of the highest temperature readread ing varies somewhat. Where the summer sky remains cloud-free all afternoon, the maximum temperature may

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DID YOU KNOW? During the summer, Death Valley, California, can sizzle at any hour of the day, even at night. On July 12, 2012, the minimum temperature dipped no lower than 107°F, which tied the record for the warmest minimum temperature ever recorded on Earth.

occur sometime between 3 p.m. and 5 p.m. standard time. Where there is afternoon cloudiness or haze, the temperature maximum usually occurs an hour or two earlier. If clouds persist throughout the day, the overall daytime temperatures are usually lower, because clouds reflect a great deal of incoming sunlight. Adjacent to large bodies of water, cool air moving inland can modify the rhythm of temperature change such that the warmest part of the day occurs at noon or before. In winter, atmospheric storms circulating warm air northward can even cause the highest temperature to occur at night. Just how warm the air becomes depends on such factors as the type of soil, its moisture content, and vegetation cover. When the soil is a poor heat conductor (as loosely packed sand is), heat energy does not readily transfer into

EXTREME HIGH TEMPERATURES Most people are aware of the extreme heat that exists during the summer in the Desert Southwest of the United States. But how hot does it get there? On July 10, 1913, Greenland Ranch in Death Valley, California, reported the highest temperature ever observed in the world: 57°C (134°F) (see Fig. 3.3). Here, air temperatures are persistently hot throughout the summer, with the average maximum for July being 47°C (116°F). During the summer of 1917, there was an incredible period of 43 consecutive days when the maximum temperature reached 120°F or higher. One of the hottest urban areas in the United States is Palm Springs, California, where the average high temperature during July is 108°F. Another hot city is Yuma, Arizona. Located along the California–Arizona border, Yuma’s high temperature during July also averages 108°F.

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FIGURE 3.2 The daily variation in air temperature is controlled by incoming energy (primarily from the sun) and outgo outgoing energy from Earth’s surface. Where incoming energy exceeds outgoing energy (orange shade), the air temperature rises. Where outgoing energy exceeds incoming energy (gray shade), the air temperature falls.

the ground. This allows the surface layer to reach a higher temperature, permitting more energy to warm the air above. On the other hand, if the soil is moist or covered with vegetation, much of the available energy evaporates water, leaving less to heat the air. As you might expect, the highest summer temperatures usually occur over desert regions, where clear skies coupled with low humidities and meager vegetation permit the surface and the air above to warm up rapidly. Where the air is humid, haze and cloudiness lower the maximum temperature by preventing some of the sun’s rays from reaching the ground. In humid Atlanta, Georgia, the average maximum temperature for July is 31.7°C (89°F). In contrast, Phoenix, Arizona—in the desert southwest at the same latitude as Atlanta—experiences an average July maximum of 41.1°C (106°F).

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FIGURE 3.3 Death Valley, California, where the air temperature reached a world-record high of 57°C (134°F) during July 1913.

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FIGURE 3.4 Record high temperatures throughout the world.

In 1937, the high in Yuma reached 100°F or more for 101 consecutive days. In a more humid climate, the maximum temperature rarely climbs above 41°C (106°F). However, during the record heat wave of 1936, the air temperature reached 121°F near Alton, Kansas. And during the heat wave of 1983, which destroyed about $7 billion in crops and increased the nation’s air-conditioning bill by an estimated $1 billion, Fayetteville reported North Carolina’s all-time record high temperature when the mercury hit 110°F. Although Death Valley holds the record for the highest officially measured air temperature in the world, it is not the hottest place on Earth. This distinction may belong to Dallol, Ethiopia. Dallol is located near latitude 12°N, in the prospect hot, dry Danakil Depression (see Fig. 3.4). A prospecting company kept weather records at Dallol from 1960 to 1966. During this time, the average daily maximum temperature exceeded 38°C (100°F) every month of the year, except during December and January, when the average maximum lowered to 98°F and 97°F, respectively. On many days, the air temperature exceeded 120°F. The average 58

annual temperature for the six years at Dallol was 34°C (94°F). In comparison, the average annual temperature in Yuma is 23°C (74°F) and at Death Valley, 24°C (76°F). On September 23, 1922, the temperature was reported to have reached a scorching 58°C (136°F) in Africa, northwest of Dallol at El Azizia, Libya (32°N). Until recently, this was considered to be the highest temperature recorded on Earth using standard measurement techniques. However, the reading was declared invalid in 2012 after an investigation by a panel of experts sponsored by the World Meteorological Organization (WMO). The panel found several major concerns with the El Azizia reading, including problematic instrumentation, an observer who was likely inexperienced, and asphalt-like material beneath the observing site that did not represent the native desert soil. Because of these and other factors, the actual temperature in El Azizia may have been 7°C (11°F) cooler than the reported record. Consequently, the 1913 reading in Death Valley was declared to be the world’s hottest officially measured temperature, returning it to the position it had held almost a century before.

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COLD AIR NEAR THE SURFACE A strong radiation inversion occurs when the air near the ground is much colder than the air higher up. Ideal conditions for a strong inversion and, hence, very low nighttime temperatures exist when the air is calm, the night is long, and the air is fairly dry and cloud-free. Let’s examine these conditions one by one. A windless night is essential for a strong radiation inversion because a stiff breeze tends to mix the colder air at the surface with the warmer air above. This mixing, along with the cooling of the warmer air as it comes in contact with the cold ground, causes a vertical temperature profile that is almost isothermal (a constant temperature) in a layer several feet thick. In the absence of wind, the cooler, more-dense surface air does not readily mix with the warmer, less-dense air above, and the inversion is more strongly developed as illustrated in Fig. 3.5. A long night also contributes to a strong inversion. Generally, the longer the night, the longer the time of radiational cooling and the better are the chances that the air near the ground will be much colder than the air above. Consequently, winter nights provide the best conditions for a strong radiation inversion, other factors being equal. Finally, radiation inversions are more likely with a clear sky and dry air. Under these conditions, the ground *Radiation (nocturnal) inversions are also called surface inversions.

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NIGHTTIME COOLING We know that nights are typically much cooler than days. Nights are cooler because, as the afternoon sun lowers, its energy is spread over a larger area, which reduces the heat available to warm the ground. Look back at Fig. 3.2 on p. 57 and observe that sometime in late afternoon or early evening, Earth’s surface and air above begin to lose more energy than they receive; hence, they start to cool. Both the ground and air above cool by radiating infrared energy, a process called radiational cooling. The ground, being a much better radiator than air, is able to cool more quickly. Consequently, shortly after sunset, Earth’s surface is slightly cooler than the air directly above it. The surface air transfers some energy to the ground by conduction, which the ground, in turn, quickly radiates away. As the night progresses, the ground and the air in contact with it continue to cool more rapidly than the air a few meters higher. The warmer upper air does transfer some heat downward, a process that is slow due to the air’s poor thermal conductivity. However, by late night or early morning, the coldest air is next to the ground, with slightly warmer air above (see Fig. 3.5). This measured increase in air temperature just above the ground is known as a radiation inversion because it forms mainly through radiational cooling of the surface. Because radiation inversions occur on most clear, calm nights, they are also called nocturnal inversions.*

FIGURE 3.5 On a clear, calm night, the air near the surface can be much colder than the air above. The increase in air tempera temperature with increasing height above the surface is called a radiation temperature inversion.

is able to radiate its energy to outer space and thereby cool rapidly. With cloudy weather and moist air, much of the outgoing infrared energy is absorbed and radiated back to the surface, retarding the rate of cooling. Also, on humid nights, condensation in the form of fog or dew will release latent heat, which warms the air. So, radiation inversions can occur on any night. But during long winter nights, when the air is still, cloud-free, and relatively dry, these inversions can become strong and deep. On a cold, dry winter night, then, it is common to experience belowfreezing temperatures near the ground, and air more than 10°F warmer at your waist. This process explains why ice or frost can appear on Earth’s surface even if the official low temperature (measured a few feet above) never dips to 0°C (32°F). It should now be apparent that how cold the night air becomes depends primarily on the length of the night, the moisture content of the air, cloudiness, and the wind. Even though wind may initially bring cold air into a region, the coldest nights usually occur when the air is clear and relatively calm. Look back at Fig. 3.2 (p. 57) and observe that the lowest temperature on any given day is usually observed around sunrise. However, the cooling of the ground and surface air may even continue beyond sunrise for a half hour or so, as outgoing energy can exceed incoming energy, because light from the early morning sun passes through a thick section of atmosphere and strikes the ground at a low angle. Consequently, the sun’s energy does not effectively warm the surface. Surface heating can be reduced further when the ground is moist and available energy is used for evaporation. Hence, the lowest temperature can occur shortly after the sun has risen. AIR TEMPERATURE URE

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FIGURE 3.6 On cold, clear nights, the settling of cold air into valleys makes them colder than surrounding hillsides. The region along the side of the hill where the air temperature is above freezing is known as a thermal belt.

On cold nights, plants and certain crops can be damaged by the low temperatures. If the cold occurs over a widespread area for a long enough time to damage certain crops, the extreme cold is called a freeze.* A single freeze in California, Texas, or Florida can cause crop losses in the millions or even billions of dollars. In fact, citrus crop losses in Florida during the hard freeze of January 1977 exceeded $2 billion. In California, several freezes during the spring of 2001 caused millions of dollars in damages to California’s north coast vineyards, which resulted in higher wine prices. And after extremely warm weather in March 2012 led to early blooming of fruit trees, freezing temperatures in April destroyed nearly half of the apple crop in New York and nearly 90 percent in Michigan. Another widespread freeze across the East and Midwest in April 2007 inflicted more than $2 billion in damage to fruit and field crops. The coldest air and lowest temperatures are frequently found in low-lying areas, because cold, heavy surface air slowly drains downhill during the night and eventually settles in low-lying basins and valleys. In middle latitudes, the warmer hillsides, called thermal belts, are less likely to experience freezing temperatures than the valley below (see Fig. 3.6). This phenomenon encourages farmers to plant *A freeze occurs over a widespread area when the surface air temperature remains below freezing for a long enough time to damage certain agricultural crops. The terms frost and freeze are often used interchangeably by various segments of society. However, to the grower of perennial crops (such as apples and citrus) who has to protect the crop against damaging low temperatures, it makes no difference if visible “frost” is present or not. The concern is whether or not the plant tissue has been exposed to temperatures equal to or below 32°F. The actual freezing point of the plant, however, can vary because perennial plants can develop hardiness in the fall that usually lasts through the winter, then wears off gradually in the spring.

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on hillsides those trees and sensitive crops that are unable to survive the valley’s low temperatures. Moreover, on the valley floor, the cold, dense air is unable to rise, so smoke and other pollutants trapped in this heavy air can restrict visibility. Therefore, valley bottoms are not only colder, but also are more frequently polluted than nearby hillsides. PROTECTING CROPS FROM THE COLD NIGHT AIR On very cold nights, protect small plants or shrubs by covering them with straw, cloth, or plastic sheeting. This prevents ground heat from being radiated away to the colder surroundings. If you are a household gardener concerned about outside flowers and plants during cold weather, simply wrap them in plastic or cover each with a paper cup. Fruit trees are particularly vulnerable to cold weather in the spring when they are blossoming. The protection of such trees presents a serious problem to the farmer. Since the lowest temperatures on a clear, still night occur near the surface, the lower branches of a tree are the most susceptible to damage. Therefore, increasing the air temperature close to the ground may prevent damage. A traditional technique for doing this is the use of orchard heaters, which warm the air around the trees by setting up convection currents close to the ground. In addition, heat energy radiated from oil- or gas-fired orchard heaters is intercepted by the buds of the trees, which raises their temperature. Orchard heaters that generate smoke, known as “smudge pots,” were used for many decades but are now prohibited in most areas due to their effects on local air quality. Another way to protect trees is to mix the cold air at the ground with the warmer air above, thus raising the

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AP Photo/John Raoux

© C. Donald Ahrens

FIGURE 3.8 Ice covers citrus trees in Clermont, Florida, that were sprayed with water during the early morning to protect them from damaging low temperatures that dipped into the 20s (°F) on December 15, 2010.

FIGURE 3.7 Wind machines mix cooler surface air with warmer air above.

fairly humid. They do not work well when the air is dry, as a good deal of the water can be lost through evaporation. So far, we have looked at how and why the air temperature near the ground changes during the course of a 24-hour day. We saw that during the day the air near Earth’s surface can become quite warm, whereas at night it can cool off dramatically. Figure 3.9 summarizes these observations by illustrating how the average air temperatempera ture above the ground can change over a span of 24 hours. Notice in the figure that although the air several feet above the surface both cools and warms, it does so at a slower rate than air at the surface.

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temperature of the air next to the ground. Such mixing can be accomplished by using wind machines (see Fig. 3.7), which are power-driven fans that resemble airplane propellers. One significant benefit of wind machines is that they can be thermostatically controlled to turn off and on at prescribed temperatures. Farmers without their own wind machines can rent air mixers in the form of helicopters. Although helicopters are effective in mixing the air, they are expensive to operate. If sufficient water is available, trees can be protected by irrigation. On potentially cold nights, farmers might flood the orchard. Because water has a high heat capacity, it cools more slowly than dry soil, and so the surface does not become as cold as it would if it were dry. Furthermore, wet soil has a higher thermal conductivity than dry soil. In wet soil heat is conducted upward from subsurface soil more rapidly, which helps to keep the surface warmer. If the air temperature both at the surface and above fall below freezing, farmers are left with a difficult situation. Wind machines won’t help because they would only mix cold air at the surface with the colder air above. Orchard heaters and irrigation are of little value as they would only protect the branches just above the ground. Fortunately, there is one form of protection that does work: An orchard’s sprinkling system can be turned on so that it emits a fine spray of water. In the cold air, the water freezes around the branches and buds, coating them with a thin veneer of ice (see Fig. 3.8). As long as the spraying continues, the latent heat—given off as the water changes into ice—keeps the ice temperature at 0°C (32°F). The ice acts as a protective coating against the subfreezing air by keeping the buds (or fruit) at a temperature higher than their damaging point. Care must be taken since too much ice can cause the branches to break. The fruit can be saved from the cold air, but the tree itself may be damaged by too much protection. Sprinklers work well when the air is

FIGURE 3.9 An idealized distribution of air temperature above the ground during a 24-hour day. The temperature curves represent the variations in average air temperature above a grassy surface for a mid-latitude city during the summer under clear, calm conditions. AIR TEMPERATURE URE

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EXTREME LOW TEMPERATURES One city in the United States that experiences very low temperatures is International Falls, Minnesota, where the average temperature for January is –15°C (4°F). Located several hundred miles to the south of International Falls, Minneapolis– St. Paul, with an average January temperature of –9°C (16°F), is the coldest major urban area in the nation. For duration of extreme cold, Minneapolis reported 186 consecutive hours of temperatures below 0°F during the winter of 1911–1912. Within the forty-eight adjacent states, however, the record for the longest duration of severe cold belongs to Langdon, North Dakota, where the thermometer remained below 0°F for 41 consecutive days (January 11 to February 20, 1936). The most extensive cold wave in the United States occurred in February 1899. Temperatures during this cold spell fell below 0°F in every existing state, including Florida. This extreme cold event was the first and only of its kind in recorded history. Record temperatures set during this extremely cold outbreak still stand today in many cities of the United States. The official record for the lowest temperature in the forty-eight adjacent states, however, belongs to Rogers Pass, Montana, where on the morning of January 20, 1954, the mercury dropped to –57°C (–70°F). The lowest official temperature for Alaska, –62°C (–80°F), occurred at Prospect Creek on January 23, 1971. The coldest areas in North America are found in the Yukon and Northwest Territories of Canada. Resolute, Canada (latitude 75°N), has an average temperature of –32°C (–26°F) for the month of January. The lowest temperatures and coldest winters in the Northern Hemisphere are found in the interior of Siberia and Greenland. For example, the average January temperature in Yakutsk, Siberia (latitude 62°N), is –39°C (–38°F). There, the mean temperature for the entire year is a bitter cold –9°C (16°F). At Eismitte, Greenland, the average temperature for February (the coldest month) is –47°C (–53°F), with the mean annual temperature being a frigid –30°C (–22°F). Even though these temperatures are

extremely low, they do not come close to the coldest area of the world: the Antarctic (see Fig. 3.10). At the geographical South Pole, over 9000 feet above sea level, where the Amundsen–Scott scientific station has been keeping records for more than fifty years, the average temperature for the month of July (winter) is –56°C (–69°F) and the mean annual temperature is –46°C (–51°F). The lowest temperature ever recorded there (–83°C or –117°F) occurred under clear skies with a light wind on the morning of June 23, 1983. Cold as it was, it was not the record low for the world. That belongs to the Russian station at Vostok, Antarctica (latitude 78°S), where the temperature plummeted to –89°C (–129°F) on July 21, 1983. ( Figure 3.11 provides more information on record low temperatures throughout the world.)

BRIEF REVIEW Up to this point we have examined daily temperature variations. Before going on, here is a review of some of the important concepts and facts we have covered: ●

During the day, Earth’s surface and the air above will continue to warm as long as incoming energy (mainly sunlight) exceeds outgoing energy from the surface.

At night, Earth’s surface cools, mainly by giving up more infrared radiation than it receives—a process called radiational cooling.

The coldest nights of winter normally occur when the air is calm, fairly dry (a low water-vapor content), and cloud-free.

The highest temperatures during the day and the lowest temperatures at night are normally observed at Earth’s surface.

Radiation inversions usually exist at night when the air near the ground is colder than the air above.

The coldest air and lowest nighttime temperatures are normally found in low-lying areas. Surrounding hillsides are usually much warmer than the valley bottoms.

Farmers use a variety of techniques to protect crops or fruit from damaging low temperatures, including heating the air, mixing the air, irrigating, and spraying water onto trees in below-freezing weather.

NASA/James Yungel

FIGURE 3.10 Antarctica, the coldest continent on Earth, where air temperatures often drop below -100°F.

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DAILY TEMPERATURE VARIATIONS The greatest variation in daily temperature occurs at Earth’s surface. The difference between the daily maximum and minimum temperature—called the daily (diurnal) range of temperature—is greatest next to the ground and becomes progressively smaller as we move away from the surface (see Fig. 3.12). This daily variation in temperature is also much larger on clear days than on cloudy ones. The largest diurnal range of temperature occurs on high deserts, where the air is fairly dry, often cloud-free, and there is little water vapor to radiate much infrared energy back to the surface. By day, clear summer skies allow the sun’s energy to quickly warm the ground which, in turn, warms the air above to a temperature often exceeding 38°C (100°F). At night, the ground cools rapidly by radiating infrared energy to space, and the minimum temperature in these regions occasionally dips below 7°C (45°F), thus giving an extremely high daily temperature range of more than 31°C (55°F). Clouds can have a large effect on the daily range in temperature. As we saw in Chapter 2, clouds (especially

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FIGURE 3.11 Record low temperatures throughout the world.

FIGURE 3.12 The daily range of temperature decreases as we climb away from Earth’s surface. Hence, there is less day-to-night variation in air temperature near the top of a high-rise apartment complex than at the ground level. AIR TEMPERATURE URE

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FIGURE 3.13 (a) Clouds tend to keep daytime temperatures lower and nighttime temperatures higher, producing a small daily range in temperature. (b) In the absence of clouds, days tend to be warmer and nights cooler, producing a larger daily range in temperature.

low, thick ones) are good reflectors of incoming solar radiation, and so they prevent much of the sun’s energy from reaching the surface. This effect tends to lower daytime temperatures (see Fig. 3.13a). If the clouds persist into the night, they tend to keep nighttime temperatures higher, as clouds are excellent absorbers and emitters of infrared radiation—the clouds actually emit a great deal of infrared energy back to the surface. Clouds, therefore, have the effect of lowering the daily range of temperature. In clear weather (Fig. 3.13b), daytime air temperatures tend to be higher as the sun’s rays impinge directly upon the surface, while nighttime temperatures are usually lower due to rapid radiational cooling. Therefore, clear days and clear nights combine to promote a large daily range in temperature. Humidity can also have an effect on diurnal temperature ranges. For example, in humid regions, the diurnal temperature range is usually small. Here, haze and clouds lower the maximum temperature by preventing some of the sun’s energy from reaching the surface. At night, the moist air keeps the minimum temperature high by absorbing Earth’s infrared radiation and radiating a portion of it to the ground. An example of a humid city with a small summer diurnal temperature range is Charleston, South Carolina, where the average July maximum temperature is 33°C (91°F), the average minimum is 23°C (73°F), and the diurnal range is only 10°C (18°F). Cities near large bodies of water typically have smaller diurnal temperature ranges than cities farther inland. This phenomenon is caused in part by the additional water vapor in the air and by the fact that water warms and cools much more slowly than land. Moreover, cities whose temperature readings are obtained at airports often have larger diurnal temperature ranges than those whose readings are obtained in downtown areas, because nighttime temperatures in cities tend 64

to be warmer than those in outlying rural areas. This nighttime city warmth—called the urban heat island—forms as the sun’s energy is absorbed by urban structures and concrete; then, during the night, this heat energy is slowly released into the city air. The average of the highest and lowest temperature observed in a given 24-hour period (typically from midnight to midnight) is known as the mean (average) daily temperature. Sometimes weather websites, TV weathercasts, or daily newspapers will provide the mean daily temperature along with the highest and lowest temperatures for the preceding day. The average of the mean daily temperatures for a particular date across a 30-year period gives the average (or “normal”) temperatures for that date. Currently, the National Weather Service uses the period 1981–2010 to calculate average temperatures. If the average high temperature in a particular city on a certain date is 68°F, does this mean that the high temperature on this date should be 68°F? If you are unsure of the answer, read Focus section 3.1. REGIONAL TEMPERATURE VARIATIONS The main factors that cause variations in temperature from one place to another are called the controls of temperature. In the previous chapter, we saw that the greatest factor in determining temperature is the amount of solar radiation that reaches the surface. This amount is determined by the length of daylight hours and the intensity of incoming solar radiation. Both of these factors are a function of latitude; hence, latitude is considered an important control of temperature. The main controls are: . latitude . land and water distribution . ocean currents . elevation

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FOCUS

ON A SPECIAL TOPIC 3.1 O

When It Comes to Temperature, What’s Normal? When the weathercaster reports that “the normal high temperature for today is 68°F,” does this mean that the high temperature on this day is usually 68°F? Or does it mean that we should expect a high temperature near 68°F? Actually, we should expect neither one. Remember that the word normal normal, refers to weather data averaged over a period of 30 years. For example, Fig. 1 shows the high temperatures measured for 30 years (1981 to 2010) in Salt Lake City, Utah, on May 6. The average (mean) high temperature for this period is 68°F; hence, the normal high temperature for this date is 68°F (dashed line). Notice, however, that only on two days during this 30-year period did the high temperature actually measure 68°F (large red dots). In fact, the most common high temperature (called the mode) was 58°F, which occurred on three days (blue dots). So what would be considered a typical high temperature for this date? Actually, any high temperature that lies between about 46°F and 90°F (two standard deviations* on either side of 68°F) would be considered *A standard deviation is a statistical measure of the spread of the data. Two standard deviations for this set of data mean that 95 percent of the time the high temperature occurs between 47°F and 89°F.

FIGURE 1 High temperatures measured on May 6 each year from 1981 to 2010 in Salt Lake City, Utah. The dashed line represents the normal (average) temperature for the 30-year period.

in the range of typical high temperatures for this date. It would be truly noteworthy, and “abnormal,” if the high temperature happened to be 68°F on May 6 every year, even though 68°F is the normal (average) high temperature for this 30-year period. While a high temperature of 86°F may be quite warm

and a high temperature of 51° may be on the cool side, they are both no more uncommon (unusual) in this period than a high temperature of 68°F. This same type of reasoning applies to normal rainfall rainfall, as the actual amount of precipitation will likely be greater or less than the 30-year average.

We can obtain a better picture of these controls by examining Fig. 3.14 and Fig. 3.15, which show the average monthly temperatures throughout the world for January and July. (The average temperature for each month is the average of the daily mean temperatures for that month.) The lines on the map are isotherms—lines connecting places that have the same temperature. Because air temperature normally decreases with height, cities at very high elevations are much colder than their sea-level counterparts. Consequently, the isotherms in Fig. 3.14 and Fig. 3.15 are corrected to read at the same horizontal level (sea level) by adding to each station above sea level the equivalent average temperature change with height.* Figures 3.14 and 3.15 show the importance of latitude on temperature. Notice that on both maps and in both

hemispheres the isotherms are oriented east-west, indicating that locations at the same latitude receive nearly the same amount of solar energy. In addition, the annual solar heat that each latitude receives decreases from low-to-high latitudes; hence, average temperatures in January and July tend to decrease from lower to higher latitudes. But there is a greater variation in solar radiation between low and high latitudes in winter than in summer. Thus, the isotherms in January (during the Northern Hemisphere winter) are closer together (a tighter gradient)* than they are in July. If you travel from New Orleans to Detroit in January, you are

*The amount of change is usually less than the standard temperature lapse rate of 3.6°F per 1000 feet (6.5°C per 1000 meters). The reason is that the standard lapse rate is computed for altitudes above Earth’s surface in the “free” atmosphere. In the less-dense air at high elevations, the absorption of solar radiation by the ground causes an overall slightly higher temperature than that of the free atmosphere at the same level.

One of the greatest daily temperature ranges ever recorded in North America (100°F), occurred at Browning, Montana, where the air temperature plummeted from a high of 44°F to a low of –56°F in less than 24 hours on January 23, 1916.

*Gradient represents the rate of change of some quantity (in this case, temperature) over a given distance.

DID YOU KNOW?

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FIGURE 3.14 Average air temperature near sea level in January (°F). Temperatures in Central Antarctica are not visible on this map.

FIGURE 3.15 Average air temperature near sea level in July (°F). Temperatures in Central Antarctica are not visible on this map.

more likely to experience greater temperature variations than if you make the same trip in July. Even though average temperatures tend to decrease from low latitudes toward high latitudes, notice on the July map (Fig. 3.15) that the highest average temperatures do not occur in the tropics, but rather in the subtropical deserts of the Northern Hemisphere. Here, sinking air associated with high-pressure areas generally produces clear skies and low humidity. These conditions, along with a 66

high sun beating down upon a relatively barren landscape, produce scorching heat. The most extreme cold over land areas in January is across the interior of Siberia (Fig. 3.14). Even colder readings, on average, occur in Antarctica during the dark winter months, as relatively dry air, high elevations, and snow-covered surfaces allow for rapid radiational cooling. Although not shown in Fig. 3.15, the average temperature for the coldest month at the South Pole is below –70°F.

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And for absolute cold, the lowest average temperature for any month (–100°F) was recorded at the Plateau Station during July 1968. So far we’ve seen that January temperatures in the Northern Hemisphere are much lower in the middle of continents than they are at the same latitude near the oceans. Notice on the July map that the reverse is true. One reason for these temperature differences is the unequal heating and cooling properties of land and water (as discussed in Chapter 2). For one thing, solar energy reaching land is absorbed in a thin layer of soil; reaching water, it penetrates deeply. Because water circulates, it distributes its heat through a much deeper layer. In addition, some of the solar energy striking the water evaporates it rather than heats it. Another reason for the sharp temperature difference between oceans and interior locations is that it takes a great deal more heat to raise the temperature of a given amount of water by one degree than it does to raise the temperature of the same amount of land by one degree.* Water not only heats more slowly than land, it cools more slowly as well, allowing the oceans to act like huge heat reservoirs. Thus, mid-ocean surface temperatures change relatively little from summer to winter compared to the much larger annual temperature changes over the middle of continents. As a result of the warming and cooling properties of water, even large lakes can modify the temperature around them. In summer, for example, the Great Lakes remain cooler than the land and refreshing breezes blow inland, bringing relief from the sometimes sweltering heat. As winter approaches, the water cools more slowly than the land. The first blast of cold air from Canada is modified as it crosses the lakes, and so the first freeze is delayed on the eastern shores of Lake Michigan. Look closely at Figs. 3.14 and 3.15 and notice that in many places the isotherms on both maps tend to bend when they approach an ocean-continent boundary. Such bending of the isotherms along the margin of continents is due in part to the unequal heating and cooling properties of land and water, and in part to ocean currents. For example, along the eastern margins of continents warm ocean currents transport warm water toward the poles, whereas, along the western margins, they transport cold water toward the equator. As we will see in Chapter 7, some coastal areas also experience upwelling, which brings cold water from below to the surface. At any location, the difference in average temperature between the warmest month (often July in the Northern Hemisphere) and coldest month (often January) is called the annual range of temperature. As we would expect, annual temperature ranges are largest over interior *The amount of heat needed to raise the temperature of one gram of a substance by one degree Celsius is called specific heat. Water has a higher specific heat than does land.

continental landmasses and much smaller over larger bodies of water (see Fig. 3.16). Moreover, inland cities have larger annual temperature ranges than do coastal cities. Near the equator (because daylight length varies little and the sun is always high in the noon sky), annual temperature ranges are small, usually less than 3°C (5°F). Quito, Ecuador—on the equator at an elevation of 2850 m (9350 ft)—experiences an annual range of less than 1°C. In middle and high latitudes, annual ranges are large, especially in the middle of a continent. Yakutsk, in northeastern Siberia near the Arctic Circle, has an extremely large annual temperature range of 58°C (104°F). The average temperature of any station for the entire year is the mean (average) annual temperature, which represents the average of the twelve monthly average temperatures.* When two cities have the same mean annual temperature, it might at first seem that their temperatures throughout the year are quite similar. However, often this is not the case. For example, San Francisco, California, and Richmond, Virginia, are situated at nearly the same latitude (38°N). Both have similar hours of daylight during the year; both have a mean annual temperature near 15°C (59°F). But here the similarities end. The temperature differences between the two cities are apparent to anyone who has traveled to San Francisco during the summer with a suitcase full of clothes suitable for summer weather in Richmond. Figure 3.17 summarizes the average temperatures for San Francisco and Richmond. Notice that the coldcold est month for both cities is January. Even though January in Richmond averages only about 8°C (14°F) colder than January in San Francisco, people in Richmond awaken to an average January minimum temperature of –3°C (28°F), which is the lowest temperature ever recorded in San Francisco. Trees that thrive in San Francisco’s weather would find it difficult to survive a winter in Richmond. So, even though San Francisco and Richmond have the same mean annual temperature, the behavior and range of their temperatures differ greatly.

Applications of Air Temperature Data There are a variety of applications for the mean daily temperature. An application developed by heating engineers in estimating energy needs is the heating degree day. The heating degree day is based on the assumption that people will begin to use their furnaces when the mean daily temperature drops below 65°F. Therefore, heating degree days are determined by subtracting the mean temperature for *The mean annual temperature can be obtained by multiplying each of the 12 monthly means by the number of days in that month, adding the 12 numbers, and dividing that total by 12; or by obtaining the sum of the daily means and dividing that total by 365. AIR TEMPERATURE URE

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FIGURE 3.16 Monthly temperature data and annual temperature range for (a) St. Louis, Missouri, a city located near the middle of a continent and (b) Ponta Delgada, a city located in the Azores in the Atlantic Ocean. Notice that the annual temperature range iis much higher in St. Louis, even though both cities are at the same latitude.

FIGURE 3.17 Temperature data for (a) San Francisco, California (38°N) and (b) Richmond, Virginia (38°N)—two cities with very similar mean temperatures.

68

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FIGURE 3.18 Mean annual total heating degree days across the United States (base 65°F).

FIGURE 3.19 Mean annual total cooling degree days across the United States (base 65°F).

the day from 65°F. Thus, if the mean temperature for a day is 64°F, the heating degree day for this day is 1.* On days when the mean temperature is above 65°F, there are no heating degree days. Hence, the lower the average daily temperature, the more heating degree days and the greater the predicted consumption of fuel. When the number of heating degree days for a whole year is calculated, the heating fuel requirements for any location can be estimated. Figure 3.18 shows the yearly average number of heating degree days in various locations throughout the United States. As the mean daily temperature climbs above 65°F, people begin to cool their indoor environment. Consequently, an index called the cooling degree day is used during warm weather to estimate the energy needed to cool indoor air to a comfortable level. The forecast of mean daily temperature is converted to cooling degree *In the United States, the National Weather Service and the Department of Agriculture use degrees Fahrenheit in their computations.

days by subtracting 65°F from the mean. The remaining value is the number of cooling degree days for that day. For example, a day with a mean temperature of 70°F would correspond to (70 – 65), or 5 cooling degree days. High values indicate warm weather and high power production for cooling (see Fig. 3.19). Knowledge of the number of cooling degree days in an area allows a builder to plan the size and type of equipment that should be installed to provide adequate air conditioning. Also, the forecasting of cooling degree days during the summer gives power companies a way of predicting the energy demand during peak energy periods. A composite of heating plus cooling degree days gives a practical indication of the energy requirements over the year. Farmers use an index called growing degree days as a guide to planting and for determining the approximate dates when a crop will be ready for harvesting. A growing degree day for a particular crop is defined as a day on which the mean daily temperature is one degree above the AIR TEMPERATURE URE

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69

base temperature (also known as the zero temperature): the minimum temperature required for growth of that crop. For sweet corn, the base temperature is 50°F and, for peas, it is 40°F. On a summer day in Iowa, the mean temperature might be 80°F. From ▼ Table 3.1, we can see that, on this day, sweet corn would accumulate (80 – 50), or 30 growing degree days. Theoretically, sweet corn can be harvested when it accumulates a total of 2200 growing degree days. For example, if sweet corn is planted in early April and each day thereafter averages about 20 growing degree days, the corn would be ready for harvest about 110 days later, or around the middle of July.* At one time, corn varieties were rated in terms of “days to maturity.” This rating system was unsuccessful because, in actual practice, corn took considerably longer in some areas than in others. This discrepancy was the reason for defining “growing degree days.” In humid Iowa, for example, where summer nighttime temperatures are high, growing degree days accumulate much faster. Consequently, the corn matures in considerably fewer days than in the drier west, where summer nighttime temperatures are lower, and each day accumulates fewer growing degree days. Although moisture and other conditions are not taken into account, growing degree days nevertheless serve as a useful guide in forecasting approximate dates of crop maturity.

Estimated Growing Degree Days for Certain Naturally Grown Agricultural Crops to Reach Maturity

▼ Table 3.1

BASE TEMPERATURE (°F)

GROWING DEGREE DAYS TO MATURITY

Beans (Snap/South Carolina)

50

1200–1300

Corn (Sweet/Indiana)

50

2200–2800

Cotton (Delta Smooth Leaf/Arkansas)

60

1900–2500

Peas (Early/Indiana)

40

1100–1200

Rice (Vegold/Arkansas)

60

1700–2100

Wheat (Indiana)

40

2100–2400

CROP (VARIETY, LOCATION)

Probably everyone realizes that the same air temperature can feel different on different occasions. For example, a temperature of 20°C (68°F) on a clear, windless March afternoon in New York City can feel almost balmy after a long, hard winter. Yet this same temperature can feel uncomfortably cool on a summer afternoon in a stiff breeze. The human body’s perception of temperature— called sensible temperature—obviously changes with varying atmospheric conditions. The reason for these changes is related to how we exchange heat energy with our environment. The body stabilizes its temperature primarily by converting food into heat (metabolism). To maintain a constant temperature, the heat produced and absorbed by the body must be equal to the heat it loses to its surroundings. There is, therefore, a constant exchange of heat—especially at the surface of the skin—between the body and the environment.

One way the body loses heat is by emitting infrared energy. But we not only emit radiant energy, we absorb it as well. Another way the body loses and gains heat is by conduction and convection, which transfer heat to and from the body by air motions. On a cold day, a thin layer of warm air molecules forms close to the skin, protecting it from the surrounding cooler air and from the rapid transfer of heat. In cold weather, then, when the air is calm, the temperature we perceive (the sensible temperature) is often higher than a thermometer might indicate. (Could the opposite effect occur where the air temperature is very high and a person might feel exceptionally cold? If you are not sure how to answer this question, read Focus section 3.2). Once the wind starts to blow, the insulating layer of warm air is swept away, and heat is rapidly removed from the skin by the constant bombardment of cold air. When all other factors are the same, the faster the wind blows, the greater the heat loss, and the colder we feel. How cold the wind makes us feel is usually expressed as a wind-chill index (WCI). The modern wind-chill index (see ▼ Table 3.2, p. 71) was formulated in 2001 by a joint action group of the National Weather Service and other agencies. The index takes into account the wind speed at about 1.5 m (5 ft) above the ground (close to where an adult’s upper body would be) instead of the 10 m (33 ft) where official wind readings are usually taken. In addition, the index translates the capacity of the air to take heat away from a person’s face (the air’s cooling power) into a wind-chill equivalent temperature.* For example, notice in Table 3.2 that an air temperature of 10°F with a wind speed of 10 mi/hr produces a

*As a point of interest, when the air temperature climbs above 86°F in the Corn Belt of the Midwest, the hot air puts added stress on the growth of the corn. Consequently, the corn grows more slowly. Because of this fact, any maximum temperature over 86°F is reduced to 86°F when computing the mean air temperature for growing degree days.

*The wind-chill equivalent temperature formulas are as follows: Wind chill (°F) = 35.74 + 0.6215T T – 35.75 (V 0.16) + 0.4275T T (V 0.16), where T is the air temperature in °F and V is the wind speed in mi/hr. Wind chill (°C) = 13.12 + 0.6215T T – 11.37 (V 0.16) + 0.3965T T (V 0.16), where T is the air temperature in °C, and V is the wind speed in km/hr.

Air Temperature and Human Comfort

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ON A SPECIAL TOPIC 3.2 O

FOCUS

Is there somewhere in our atmosphere where the air temperature can be exceedingly high (say above 500°C or 900°F) yet a person might feel extremely cold? There is such a region, but it’s not at Earth’s surface. You may recall from Chapter 1 (see Fig. 1.25, p. 23), that in the upper reaches of our atmosphere (in the middle and upper thermosphere), air temperatures can exceed 500°C. However, a thermometer shielded from the sun in this region of the atmosphere would indicate an extremely low temperature. This apparent discrepancy lies in the meaning of air temperature and how we measure it. In Chapter 2, we learned that the air temperature is directly related to the average speed at which the air molecules are moving—faster speeds correspond to higher temperatures. In the middle and upper thermosphere at altitudes approaching

300 km (200 mi), air molecules are zipping about at speeds corresponding to extremely high temperatures. However, in order to transfer enough energy to heat something up by conduction (exposed skin or a thermometer bulb), an extremely large number of molecules must collide with the object. In the “thin” air of the upper atmosphere, air molecules are moving extraordinarily fast, but there are simply not enough of them bouncing against the thermometer bulb for it to register a high temperature. In fact, when properly shielded from the sun, the thermometer bulb loses far more energy than it receives and indicates a temperature near absolute zero. This explains why an astronaut, when space walking, will not only survive temperatures exceeding 500°C, but will also feel a profound coldness when shielded from the sun’s radiant energy. At these high altitudes, the

wind-chill equivalent temperature of –4°F. In other words, the skin of a person’s exposed face would lose as much heat in one minute in air with a temperature of 10°F and a wind speed of 10 mi/hr as it would in calm air with a

NASA

A Thousand Degrees and Freezing to Death

FIGURE 2 How can an astronaut survive when the “air” temperature is 1000°C?

traditional meaning of air temperature (that is, regarding how “hot” or “cold” something feels) is no longer applicable.

temperature of –4°F. Of course, how cold we feel actually depends on a number of factors, including the fit and type of clothing we wear, the amount of sunshine striking the body, and the actual amount of exposed skin.

Wind-Chill Equivalent Temperature (°F). A 20-mi/hr Wind Combined with an Air Temperature of 20°F Produces a Wind-Chill Equivalent Temperature of 4°F.*

▼ Table 3.2

WIND SPEED (MI/HR)

AIR TEMPERATURE (°F)

Calm

40

35

30

25

20

15

10

5

–5

–10

–15

–20

–25

–30

–35

–40

5

36

31

25

19

13

7

1

–5

–11

–16

–22

–28

–34

–40

–46

–52

–57

10

34

27

21

15

9

3

–4

–10

–16

–22

–28

–35

–41

–47

–53

–59

–66

15

32

25

19

13

6

–7

–13

–19

–26

–32

–39

–45

–51

–58

–64

–71

20

30

24

17

11

4

–2

–9

–15

–22

–29

–35

–42

–48

–55

–61

–68

–74

25

29

23

16

9

3

–4

–11

–17

–24

–31

–37

–44

–51

–58

–64

–71

–78

30

28

22

15

8

1

–5

–12

–19

–26

–33

–39

–46

–53

–60

–67

–73

–80

35

28

21

14

7

–7

–14

–21

–27

–34

–41

–48

–55

–62

–69

–76

–82

40

27

20

13

6

–1

–8

–15

–22

–29

–36

–43

–50

–57

–64

–71

–78

–84

45

26

19

12

5

–2

–9

–16

–23

–30

–37

–44

–51

–58

–65

–72

–79

–86

50

26

19

12

4

–3

–10

–17

–24

–31

–38

–45

–52

–60

–67

–74

–81

–88

55

25

18

11

4

–3

–11

–18

–25

–32

–39

–46

–54

–61

–68

–75

–82

–89

60

25

17

10

3

–4

–11

–19

–26

–33

–40

–48

–55

–62

–69

–76

–84

–91

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High winds in below-freezing air can remove heat from exposed skin so quickly that the skin may actually freeze and discolor. The freezing of skin, called frostbite, usually occurs on the body extremities first because they are the greatest distance from the source of body heat. In cold weather, wet skin can be a factor in how cold we feel. A cold, rainy day (drizzly or even foggy) often feels colder than a “dry” one because water on exposed skin conducts heat away from the body better than air does. In fact, in cold, wet, and windy weather a person may actually lose body heat faster than the body can produce it. This can even occur in relatively mild weather with air temperatures as high as 10°C (50°F). The rapid loss of body heat can lower the body temperature below its normal level and bring on a condition known as hypothermia—the rapid, progressive mental and physical collapse that accompanies the lowering of human body temperature. The first symptom of hypothermia is exhaustion. If exposure continues, judgment and reasoning power begin to disappear. Prolonged exposure, especially at temperatures near or below freezing, produces stupor, collapse, and death when the internal body temperature drops to about 26°C (79°F). In cold weather, heat is more easily dissipated through the skin. To counteract this rapid heat loss, the peripheral blood vessels of the body constrict, cutting off the flow of blood to the outer layers of the skin. In hot weather, the blood vessels enlarge, allowing a greater loss of heat energy to the surroundings. Perspiration is also a factor. As evaporation occurs, the skin cools. When the air contains a great deal of water vapor (is very humid) and it is close to being saturated, perspiration does not readily evaporate from the skin. Less evaporational cooling causes most people to feel hotter than it “really” is, and a number of people start to complain about the “heat and humidity.” A closer look at how we feel in hot, humid weather will be given in Chapter 4, after we examine the concepts of relative humidity and wet-bulb temperature.

DID YOU KNOW? Possibly the lowest wind chill ever measured was in Antarctica on August 25, 2005, when the Russian Antarctic Station of Vostok recorded an air temperature of –99°F and a wind speed of 113 mi/hr, resulting in a wind-chill equivalent temperature well below –100°F. Under these extreme conditions, any exposed skin would freeze in a few seconds.

opening, or bore, extends from the bulb to the end of the tube. A liquid in the bulb (usually mercury or red-colored alcohol) is free to move from the bulb up through the bore and into the tube. The length of the liquid in the tube represents the air temperature. When the air temperature increases, the liquid in the bulb expands, and rises up the tube. When the air temperature decreases, the liquid contracts, and moves down the tube. Because the bore is very narrow, a small temperature change shows up as a relatively large change in the length of the liquid column. Maximum and minimum thermometers are liquid-in-glass thermometers used for determining daily maximum and minimum temperatures. The maximum thermometer looks like any other liquid-in-glass thermometer with one exception: It has a small constriction within the bore just above the bulb (see Fig. 3.20). As the air temperature increases, the mercury expands and freely moves past the constriction up the tube, until the maximum temperature occurs. However, as the air temperature begins to drop, the small constriction prevents the mercury from flowing back into the bulb. Thus, the end of the stationary mercury column indicates the maximum temperature for the day. The mercury will stay at this position until either the air warms to a higher reading or the thermometer is reset by whirling it on a special holder and pivot. Usually, the whirling is sufficient to push the mercury back into the bulb past the constriction until the end of the column indicates the present air temperature.* A minimum thermometer measures the lowest temperature reached during a given period. Most minimum thermometers use alcohol as a liquid, since it freezes at a temperature of –130°C compared to –39°C

Measuring Air Temperature

72

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Thermometers were developed to measure air temperature. Each thermometer has a definite scale and is calibrated so that a thermometer reading 0°C will denote the same temperature whether it is in Vermont or North Dakota. If a particular temperature reading were to represent different degrees of hot or cold, depending on location, thermometers would be useless. A very common thermometer for measuring surface air temperature is the liquid-in-glass thermometer. This type of thermometer has a glass bulb attached to a sealed, graduated tube about 25 cm (10 in.) long. A very small

*Liquid-in-glass thermometers that measure body temperature are maximum thermometers, which is why they are shaken both before and after you take your temperature.

FIGURE 3.20 A section of a maximum thermometer.

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© Cengage Learning®.

FIGURE 3.21 A section of a minimum thermometer showing both the current air temperature and the minimum temperature in °F.

© Jan Null

for mercury. The minimum thermometer is similar to other liquid-in-glass thermometers except that it contains a small barbell-shaped index marker in the bore (see Fig. 3.21). The small index marker is free to slide back and forth within the liquid. It cannot move out of the liquid because the surface tension at the end of the liquid column (the meniscus) holds it in. A minimum thermometer is mounted horizontally. As the air temperature drops, the contracting liquid moves back into the bulb and brings the index marker down the bore with it. When the air temperature stops decreasing, the liquid and the index marker stop moving down the bore. As the air warms, the alcohol expands and moves freely up the tube past the stationary index marker. Because the index marker does not move as the air warms, the minimum temperature is read by observing the upper end of the marker. To reset a minimum thermometer, simply tip it upside down. This allows the index marker to slide to the upper end of the alcohol column, which is indicating the current air temperature. The thermometer is then remounted horizontally, so that the marker will move toward the bulb as the air temperature decreases. Highly accurate temperature measurements can be made with electrical thermometers. One type of electrical thermometer is the electrical resistance thermometer. This does not actually measure air temperature; rather, it measures the resistance of a wire, usually platinum or nickel, whose resistance increases as the temperature increases. An electrical meter measures the resistance, and is calibrated to represent air temperature. Another type of electrical thermometer is the thermistor. Made of ceramic material, its electrical resistance changes as the air temperature changes. A thermistor is the temperature-measuring device used in the radiosonde, the instrument that measures air temperature from the surface up to an altitude near 30 km. (For additional information on the radiosonde, read Focus section 1.2 in Chapter 1, p. 22.) Electrical resistance thermometers are the type used in the measurement of air temperature at the over 900 fully

FIGURE 3.22 The instruments that comprise the ASOS system. The max-min temperature shelter is the middle box.

automated surface weather stations (known as ASOS for Automated Surface Observing System) that exist at airports and military facilities throughout the United States. (See Fig. 3.22.) They have replaced many of the liquidin-glass thermometers formerly in use. At this point it should be noted that the replacement of liquid-in-glass thermometers with electrical thermometers has raised concern among climatologists. For one thing, the response of the electrical thermometers to temperature change is faster. Thus, electrical thermometers might reach a brief extreme reading that could have been missed by the slower-responding liquid-in-glass thermometer. In addition, many temperature readings that were previously taken at airport weather offices are now taken at ASOS locations situated near or between runways at the airport. This change in instrumentation and relocation of the measurement site can sometimes introduce a small, but significant, temperature change at the reporting station. To reduce the impact of such temperature changes, the United States has created a Climate Reference Network of about 100 weather stations. These stations are carefully placed, calibrated, and maintained to produce a consistent, accurate long-term reading of air temperature. Air temperature can also be obtained with instruments called infrared sensors, or radiometers. Radiometers do not measure temperature directly; rather, they measure emitted radiation (usually infrared). By measuring both the intensity of radiant energy and the wavelength of AIR TEMPERATURE URE

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FOCUS

ON AN OBSERVATION 3.3 O

Why Thermometers Must Be Read in the Shade rate than it can radiate it away, and the liquid keeps expanding and rising until there is equilibrium between incoming and outgoing energy. Because of the direct absorption of solar energy, the level of the liquid in the thermometer indicates a temperature much higher than the actual air temperature. Thus, a statement that says, “Today the air temperature measured 100 degrees in the sun,” has no meaning; a thermometer must be kept in a shady place to measure the temperature of the air accurately. FIGURE 3 Instrument shelters such as the one shown here serve as a shady place for thermometers. Thermometers inside shelters measure the temperature of the air; whereas thermometers held in direct sunlight do not.

© Cengage Learning®.

maximum emission of a particular gas (either water vapor or carbon dioxide), radiometers in orbiting satellites are now able to obtain temperature measurements at selected levels in the atmosphere. A bimetallic thermometer consists of two different pieces of metal (usually brass and iron) welded together to form a single strip. As the temperature changes, the brass expands more than the iron, causing the strip to bend. The small amount of bending is amplified through a system of levers to a pointer on a calibrated scale. The bimetallic thermometer is usually the temperature-sensing part of the thermograph, an instrument that measures and records temperature (see Fig. 3.23).

FIGURE 3.23 The thermograph with a bimetallic thermometer.

74

© Ross DePaola

When we measure air temperature with a common liquid thermometer, an incredible number of air molecules bombard the bulb, transferring energy either to or away from it. When the air is warmer than the thermometer, the liquid gains energy, exex pands, and rises up the tube; the opposite will happen when the air is colder than the thermometer. The liquid stops rising (or falling) when equilibrium between incomincom ing and outgoing energy is established. At this point, we can read the temperature by observing the height of the liquid in the tube. It is impossible to measure air temperature accurately in direct sunlight because the thermometer absorbs radiant energy from the sun in addition to enen ergy from the air molecules. The therther mometer gains energy at a much faster

Thermographs are gradually being replaced with data loggers. These small instruments have a thermistor connected to a circuit board inside the logger. A computer programs the interval at which readings are taken. The loggers are not only more responsive to air temperature than are thermographs, they also are less expensive. Chances are, you may have heard someone exclaim something like, “Today the thermometer measured 100 degrees in the shade!” Does this mean that the air temperature is sometimes measured in the sun? If you are unsure of the answer, read Focus section 3.3 before reading the next section on instrument shelters. Thermometers and other instruments are usually housed in an instrument shelter. The shelter completely encloses the instruments, protecting them from rain, snow, and the sun’s direct rays. It is painted white to reflect sunlight, faces north to avoid direct exposure to sunlight, and has louvered sides, so that air is free to flow through it. This construction helps to keep the air inside the shelter at the same temperature as the air outside. The thermometers inside a standard shelter are mounted about 1.5 to 2 m (5 to 6 ft) above the ground. As we saw in an earlier section, on a clear, calm night the air at ground level may be much colder than the air at the level of the shelter. As a result, on clear winter mornings

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it is possible to see ice or frost on the ground even though the minimum thermometer in the shelter did not reach the freezing point. Many older instrument shelters (such as the one shown in Focus Fig. 3, p. 74) have been replaced by the Max-Min Temperature Shelter of the ASOS system (the middle white box in Fig. 3.22, p. 73). The shelter is mounted on a pipe, and wires from the electrical temperature sensor inside are run to a building. A readout inside the building displays the current air temperature and stores the maximum and minimum temperatures for later retrieval.

Because air temperatures vary considerably above different types of surfaces, shelters are placed over grass where possible, to ensure that the air temperature is measured at the same elevation over the same type of surface. Unfortunately, some shelters are placed on asphalt, others sit on concrete, while others are located on the tops of tall buildings, making it difficult to compare air temperature measurements from different locations. In fact, if either the maximum or minimum air temperature in your area seems suspiciously different from those of nearby towns, find out where the instrument shelter is situated.

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SUMMARY The daily variation in air temperature near Earth’s surface is controlled mainly by the input of energy from the sun and the output of energy from the surface. On a clear, calm day, the surface air warms, as long as heat input (mainly sunlight) exceeds heat output (mainly convection and radiated infrared energy). The surface air cools at night, as long as heat output exceeds input. Because the ground at night cools more quickly than the air above, the coldest air is normally found at the surface where a radiation inversion usually forms. When the air temperature in agricultural areas drops to dangerously low readings, fruit trees and grape vineyards can be protected from the cold by a variety of means, from mixing the air to spraying the trees and vines with water. The greatest daily variation in air temperature occurs at Earth’s surface. Both the diurnal and annual range of temperature are greater in dry climates than in humid ones. Even though two cities may have similar average annual temperatures, the range and extreme of their temperatures can differ greatly. Temperature information influences our lives in many ways, from deciding what clothes to take on a trip to providing critical information for energy-use predictions and agricultural planning. We reviewed some of the many types of thermometers in use: maximum, minimum, bimetallic, electrical, radiometer. Those designed to measure air temperatures near the surface are housed in instrument shelters to protect them from direct sunlight and precipitation.

KEY TERMS The following terms are listed (with corresponding page numbers) in the order they appear in the text. Define each. Doing so will aid you in reviewing the material covered in this chapter. radiational cooling, 59 radiation inversion, 59 freeze, 60 thermal belt, 60 orchard heater, 60 wind machine, 61 daily (diurnal) range of temperature, 63 mean (average) daily temperature, 64 controls of temperature, 64 isotherm, 65 76

annual range of temperature, 67 mean (average) annual temperature, 67 heating degree day, 67 cooling degree day, 69 growing degree day, 69 sensible temperature, 70 wind-chill index, 70 frostbite, 72 hypothermia, 72

liquid-in-glass thermometer, 72 maximum thermometer, 72 minimum thermometer, 72 electrical thermometer, 73 radiometer, 73

bimetallic thermometer, 74 thermograph, 74 instrument shelter, 74

QUESTIONS FOR REVIEW . Explain why the warmest time of the day is usually in the afternoon, even though the sun’s rays are most direct at noon. . On a calm, sunny day, why is the air next to the ground normally much warmer than the air several feet above? . Explain how incoming energy and outgoing energy regulate the daily variation in air temperature. . Draw a vertical profile of air temperature from the ground to an elevation of 3 m (10 ft) on (a) a clear, windless afternoon and (b) an early morning just before sunrise. Explain why the temperature curves are different. . Explain how radiational cooling at night produces a radiation temperature inversion. . What weather conditions are best suited for the formation of a cold night and a strong radiation inversion? . Explain why thermal belts are found along hillsides at night. . List four measures farmers use to protect their crops against the cold. Explain the physical principle behind each method. . Why are the lower branches of trees most susceptible to damage from low temperatures? . Describe each of the controls of temperature. . Look at Fig. 3.14, p. 66 (temperature map for January) and explain why the isotherms dip southward (equatorward) over the Northern Hemisphere continents. . During the winter, white frost can form on the ground when the minimum thermometer in an instrument shelter indicates a low temperature above freezing. Explain. . Why do the first freeze in autumn and the last freeze in spring occur in bottomlands? . Explain why the daily range of temperature is normally greater (a) in drier regions than in humid regions and (b) on clear days than on cloudy days. . Why are the largest annual ranges of temperatures normally observed over continents away from large bodies of water?

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. Two cities have the same mean annual temperature. Explain why this fact does not mean that their temperatures throughout the year are similar. . What is a heating degree day? A cooling degree day? How are these units calculated? . During a cold, calm, sunny day, why do we usually feel warmer than a thermometer indicates? . (a) Assume the wind is blowing at 30 mi/hr and the air temperature is 5°F. Determine the wind-chill equivalent temperature using Table 3.2, p. 71. (b) Under the conditions listed in (a) above, explain why an ordinary thermometer would measure a temperature of 5°F, and not a much lower temperature. . What atmospheric conditions can bring on hypothermia? . Someone says, “Today, the air temperature measured 99°F in the sun.” Why does this statement have no meaning? . Explain why the minimum thermometer is the one with a small barbell-shaped index marker in the bore. . Briefly describe how the following thermometers measure air temperature: (a) liquid-in-glass (b) bimetallic (c) radiometer (d) electrical

QUESTIONS FOR THOUGHT AND EXPLORATION .

How do you think a thick layer of low clouds would affect the lag in daily temperature?

. Which location is most likely to have the greater daily temperature range: a tropical rain forest near the equator or a desert site in Nevada? Explain. . Explain why putting on a heavy winter jacket would be effective in keeping you warm, even if the jacket had been outside in sub-freezing temperatures for several hours. . Why is the air temperature displayed on a bank or building marquee usually inaccurate? . If you were forced to place a meteorological instrument shelter over asphalt rather than over grass, what modification(s) would you have to make so that the temperature measurements inside the shelter were more representative of the actual air temperature? . The average temperature in San Francisco, California, for December, January, and February is 52°F. During the same three-month period the average temperature in Richmond, Virginia, is 40°F. Yet, San Francisco and Richmond have nearly the same yearly total of heating degree days. Explain why. (Hint: See Fig. 3.17, p. 68.) . How would the lag in daily temperature experienced over land compare to the daily temperature lag over water? . In Pennsylvania and New York, wine grapes are planted on the sides of hills rather than in valleys. Explain why this practice is so common in these areas. . Suppose peas are planted in Indiana on May 1. If the peas need 1200 growing degree days before they can be picked, and if the average maximum temperature for May and June is 80°F and the average minimum is 60°F, on about what date will the peas be ready to pick? (Assume a base temperature of 55°F.)

Go to the Document section of the Solar Energy portal. Access the table “Solar Input Statistics for Various Countries.” Using a world map, compare the latitudes of each country to the annual amount of solar energy each country receives at midsummer, midwinter, and annually. Does the amount of energy received vary only with latitude, or are there other differences and similarities among countries that would need to be explained in other ways?

ONLINE RESOURCES Visit www.cengagebrain.com to view additional resources, including video exercises, practice quizzes, an interactive eBook, and more. AIR TEMPERATURE URE Copyright 2018 Cengage Learning. All Rights Reserved. May not be copied, scanned, or duplicated, in whole or in part. WCN 02-200-202

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CHAPTER

4

Humidity, Condensation, and Clouds Contents Circulation of Water in the Atmosphere Evaporation, Condensation, and Saturation Humidity

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t’s 9 a.m. on April pril 26, 2005, in Bangkok, Thailand, one of the hottest and most humid major cities in the world. The streets

are clogged with traffic and on this hot, muggy morning perper spiration streams down the faces of anxious people struggling to get to work. What makes this day so eventful is that a rare weather event is occurring: Presently, resently, the air temperature is 91°F, the relative humidity is 94 percent, and the heat index, which tells us how hot it really feels, is a staggering 130°F.

Dew and Frost Fog Foggy Weather Clouds

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e know from Chapter 1 that, in our atmosphere, the concentration of the invisible gas water vapor is normally less than a few percent of all the atmospheric molecules. Yet water vapor is exceedingly important, for it transforms into cloud droplets and ice crystals—particles that grow in size and fall to Earth as precipitation. The term humidity can describe the amount of water vapor in the air. To most of us, a moist day suggests high humidity. However, there is usually more water vapor in the hot, “dry” air of the Sahara Desert than in the cold, “damp” winter air of New England, which raises an interesting question: Does the desert air have a higher humidity? As we will see later in this chapter, the answer to this question is both yes and no, depending on the type of humidity we mean. So that we may better understand the concept of humidity, we will begin this chapter by examining the circulation of water in the atmosphere. Then we will look at different ways to express humidity. Near the end of the chapter, we will investigate various forms of condensation, including dew, fog, and clouds.

Circulation of Water in the Atmosphere

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Within the atmosphere, there is an unending circulation of water. Since the oceans occupy over 70 percent of Earth’s surface, we can think of this circulation as beginning over the ocean. Here, the sun’s energy transforms enormous quantities of liquid water into water vapor in

a process called evaporation. Winds then transport the moist air to other regions, where the water vapor changes back into liquid (or ice), forming clouds, in a process called condensation. Under certain conditions, the liquid cloud particles (or solid ice crystals) may grow in size and fall to the surface as precipitation—rain, snow, or hail. If the precipitation falls into an ocean, the water begins its cycle again. If, on the other hand, the precipitation falls on a continent, a great deal of the water returns to the ocean only after a complex journey. This cycle of moving and transforming water molecules from liquid to vapor and back to liquid again is called the hydrologic (water) cycle. In the most simplistic form of this cycle, water molecules travel from ocean to atmosphere to land and then back to the ocean. Figure 4.1 illustrates the complexities of the hydrohydro logic cycle. For example, before falling rain ever reaches the ground, a portion of it evaporates back into the air. Some of the precipitation may be intercepted by vegetation, where it evaporates or drips to the ground long after a storm has ended. Once on the surface, a portion of the water soaks into the ground by percolating downward through small openings in the soil and rock, forming groundwater that can be tapped by wells. What does not soak into the ground collects in puddles of standing water or runs off into streams and rivers, which find their way back to the ocean. Even the underground water moves slowly and eventually surfaces, only to evaporate or be carried seaward by rivers. Over land, a considerable amount of water vapor is added to the atmosphere through evaporation from the

FIGURE 4.1 The hydrologic cycle.

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Evaporation, Condensation, and Saturation To obtain a slightly different picture of water in the atmosphere, suppose we examine water in a beaker simimag lar to the one shown in Fig. 4.2a. If we were able to magnify the surface water about a billion times, we would see water molecules fairly close together, jiggling, bouncing, and moving about. We would also see that the molecules are not all moving at the same speed—some are moving much faster than others. Recall from Chapter 2 that the temperature of the water is a measure of the average motion of its molecules. At the surface, molecules with enough speed (and traveling in the right direction) will occasionally break away from the liquid surface and enter into the air above. These molecules, changing from the liquid state into the vapor state, are evaporating evaporating. While some water molecules are leaving the liquid, others are returning. Those returning are condensing, as they are changing from a vapor state to a liquid state. When a cover is placed over the dish (Fig. 4.2b), after a while the total number of molecules escaping from the liquid (evaporating) is balanced by the number returning (condensing). When this condition exists, the air is said to be saturated with water vapor. Under saturated conditions, for every molecule that evaporates, one must condense, and no net loss of liquid or vapor molecules results. If we remove the cover and blow across the top of the water, some of the vapor molecules already in the air above will be blown away, creating a difference between the actual number of vapor molecules and the total number required for saturation. This helps prevent saturation from

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soil, lakes, and streams. Even plants give up moisture by a process called transpiration. The water absorbed by a plant’s root system moves upward through the stem and emerges from the plant through numerous small openings on the underside of the leaf. In all, evaporation and transpiration from continental areas amount to only about 15 percent of the nearly 1.5 quintillion (1018) gallons of water vapor that annually evaporate into the atmosphere; the remaining 85 percent evaporates from the oceans. If all of this water vapor were to suddenly condense and fall as rain, it would be enough to cover the entire globe with about 2.5 centimeters (or 1 inch) of water. The total mass of water vapor stored in the atmosphere at any moment adds up to only a little over a week’s supply of the world’s precipitation. Since this amount varies only slightly from day to day, the hydrologic cycle is exceedingly efficient in circulating water in the atmosphere.

FIGURE 4.2 (a) Water molecules at the surface of the water are evaporating (changing from liquid into vapor) and condens condensing (changing from vapor into liquid). Since more molecules are evaporating than condensing, net evaporation is occurring. (b) When the number of water molecules escaping from the liquid (evaporating) balances those returning (condensing), the air above the liquid is saturated with water vapor. (For clarity, only water molecules are illustrated.)

occurring and allows for a greater amount of evaporation, just as if wind were blowing atop the water surface. Wind, therefore, enhances evaporation. The temperature of the water also influences evaporation. All else being equal, warm water will evaporate more readily than cool water, the reason being that when the water molecules are heated, they will speed up. At higher temperatures, a greater fraction of the molecules have suf sufficient speed to break through the surface tension of the water and zip off into the air above. In other words, the warmer the water, the greater the rate of evaporation. If we could examine the air above the water in Fig. 4.2b, we would observe the water vapor molecules freely darting about and bumping into each other as well as into neighboring molecules of oxygen and nitrogen. We would also observe that mixed in with all of the air molecules are microscopic bits of dust, smoke, and salt from ocean spray. Since many of these serve as surfaces on which water vapor may condense, they are called condensation nuclei. In the warm air above the water, fast-moving vapor molecules strike the nuclei with such impact that they simply bounce away (see Fig. 4.3a). However, if the air is chilled (Fig. 4.3b), the molecules move more slowly and are more apt to stick and condense to the nuclei. When many billions of these water vapor molecules condense onto the nuclei, tiny liquid cloud droplets form. We can see, then, that condensation is more likely to happen as the air cools and the speed of the water vapor molecules decreases. As the air temperature increases, condensation is less likely because most of the water vapor molecules have sufficient speed (sufficient energy) HUMIDITY, CONDENSATION, AND CLOUDS

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capacity for water vapor than does cold air.” At this point, it is important to realize that although these statements are correct, the use of such words as “hold” and “capacity” are misleading when describing water vapor content, as air does not really “hold” water vapor in the sense of making “room” for it.

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Humidity

FIGURE 4.3 Condensation is more likely to occur as the air cools. (a) In the warm air, fast-moving H2O vapor molecules tend to bounce away after colliding with nuclei. (b) In the cool air, slow-moving vapor molecules are more likely to join together on nuclei. The condensing of many billions of water molecules produces tiny liquid water droplets.

to remain as a vapor. As we will see in this and other chapters, condensation occurs primarily when the air is cooled. Even though condensation is more likely to occur when the air cools, it is important to note that no matter how cold the air becomes, there will always be a few water vapor molecules with sufficient speed (sufficient energy) to remain as a vapor. It should be apparent, then, that with the same number of water vapor molecules in the air, saturation is more likely to occur in cool air than in warm air. This fact often leads to the statement that “warm air can hold more water vapor molecules before becoming saturated than can cold air” or, simply, “warm air has a greater

Humidity refers to any one of a number of ways of specifying the amount of water vapor in the air. Most people have heard of relative humidity, which we will examine later, but there are several other ways to express atmospheric water vapor content. Imagine, for example, that we enclose a volume of air (about the size of a large balloon) in a thin elastic container—a parcel—as illustrated in Fig. 4.4. If we extract the water vapor from the parcel, we can specify the humidity in the following ways: . We can compare the weight (mass) of the water vapor with the volume of air in the parcel and obtain the water vapor density, or absolute humidity. . We can compare the weight (mass) of the water vapor in the parcel with the total weight (mass) of all the air in the parcel (including vapor) and obtain the specific humidity. . Or, we can compare the weight (mass) of the water vapor in the parcel with the weight (mass) of the remaining dry air and obtain the mixing ratio. Absolute humidity is normally expressed as grams of water vapor per cubic meter of air (g/m3), whereas both specific humidity and mixing ratio are expressed as grams of water vapor per kilogram of air (g/kg). Look at Fig. 4.4 and notice that we could also express the humidity of the air in terms of water vapor pressure— the push (force) that the water vapor molecules are exerting against the inside walls of the parcel.

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VAPOR PRESSURE Suppose the air parcel in Fig. 4.4 is near sea level and the air pressure inside the parcel is 1000 millibars (mb).* The total air pressure inside the parcel is due to the collision of all the molecules against the walls of the parcel. In other words, the total pressure inside the parcel is equal to the sum of the pressures of the individual gases. Since the total pressure inside the parcel is 1000 millibars, and the gases inside include nitrogen (78 percent), oxygen (21 percent), and water vapor (1 percent), the partial pressure exerted by nitrogen would then be 780 mb and, by oxygen, 210 mb. The partial FIGURE 4.4 The water vapor content (humidity) inside this air parcel can be expressed in a number of ways.

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*You may recall from Chapter 1 that the millibar is the unit of pressure most commonly found on surface weather maps, and that it expresses atmospheric pressure as a force over a given area.

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RELATIVE HUMIDITY While relative humidity is the most commonly used way of describing atmospheric moisture, it is also, unfortunately, the most misunderstood. The concept of relative humidity may at first seem confusing because it does not indicate the actual amount of water vapor in the air. Instead, it tells us how close the air is to being saturated. The relative humidity (RH) is the ratio of the amount of water vapor actually in the air to *When we use the percentages of various gases in a volume of air, these percentages only give us an approximation of the actual vapor pressure. The point here is that, near Earth’s surface, the actual vapor pressure is often close to 10 mb. **When the air is saturated, the amount of water vapor is the maximum possible at the existing temperature and pressure.

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pressure of water vapor, called the actual vapor pressure, would be only 10 mb (one percent of 1000).* It is evident, then, that because the number of water vapor molecules in any volume of air is small compared to the total number of air molecules in the volume, the actual vapor pressure is normally a small fraction of the total air pressure. Everything else being equal, the more air molecules in a parcel, the greater the total air pressure. When you blow up a balloon, you increase its pressure by putting in more air. Similarly, an increase in the number of water vapor molecules will increase the total vapor pressure. Hence, the actual vapor pressure is a fairly good measure of the total amount of water vapor in the air: High actual vapor pressure indicates large numbers of water vapor molecules, whereas low actual vapor pressure indicates comparatively small numbers of vapor molecules. Actual vapor pressure indicates the air’s total water vapor content, whereas saturation vapor pressure describes how much water vapor is necessary to make the air saturated at any given temperature.** Put another way, saturation vapor pressure is the pressure that the water vapor molecules would exert if the air were saturated with vapor at a given temperature. We can obtain a better picture of the concept of saturation vapor pressure by imagining molecules evaporating from a water surface. Look back at Fig. 4.2b and recall that when the air is saturated, the number of molecules escaping from the water’s surface equals the number returning. Since the number of “fast-moving” molecules increases as the temperature increases, the number of water molecules escaping per second increases also. In order to maintain equilibrium, this situation causes an increase in the number of water vapor molecules in the air above the liquid. Consequently, at higher air temperatures, it takes more water vapor to saturate the air. And more vapor molecules exert a greater pressure. Saturation vapor pressure, then, depends primarily on the air temperature. From the graph in Fig. 4.5, we can see that at 10°C, the saturation vapor pressure is about 12 mb, whereas at 30°C, it is about 42 mb.

FIGURE 4.5 Saturation vapor pressure increases with increasing temperature. At a temperature of 10°C, the satura saturation vapor pressure is about 12 mb, whereas at 30°C, it is about 42 mb. The insert illustrates that the saturation vapor pressure over water is greater than the saturation vapor pressure over ice.

the maximum amount of water vapor required for saturation at that particular temperature (and pressure). It is the ratio of the air’s water vapor content to its capacity; thus RH 5

water vapor content . water vapor capacity

We can think of the actual vapor pressure as a measure of the air’s actual water vapor content, and the saturation vapor pressure as a measure of air’s total capacity for water vapor. Hence, the relative humidity can be expressed as RH 5

actual vapor pressure 100 percent. saturation vapor pressure

Relative humidity is given as a percent. Air with a 50 percent relative humidity contains one-half the amount of water vapor required for saturation. Air with a 100 percent relative humidity is said to be saturated because it is filled to capacity with water vapor. Air with a relative humidity greater than 100 percent is said to be HUMIDITY, CONDENSATION, AND CLOUDS

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summary, with no change in air temperature, adding water vapor to the air increases the relative humidity; removing water vapor from the air lowers the relative humidity. Figure 4.6b illustrates that as the air temperature increases (with no change in water vapor content), the relative humidity decreases. This decrease in relative humidity occurs because in the warmer air the water vapor molecules are zipping about at such high speeds they are unlikely to join together and condense. The higher the temperature, the faster the molecular speed, the less likely saturation will occur, and the lower the relative humidity.* As the air temperature lowers, the water vapor molecules move more slowly. Condensation becomes more likely as the air approaches saturation, and the relative humidity increases. In summary, with no change in water vapor content, an increase in air temperature lowers the relative humidity, while a decrease in air temperature raises the relative humidity. In many places, the air’s total vapor content varies only slightly during an entire day, and so it is the changing air temperature that primarily regulates the daily variation dur in relative humidity (see Fig. 4.7). As the air cools during the night, the relative humidity increases. Normally, the highest relative humidity occurs in the early morning, during the coolest part of the day. As the air warms during the day, the relative humidity decreases, with the lowest values usually occurring during the warmest part of the afternoon. These changes in relative humidity are important in determining the amount of evaporation from vegetation and wet surfaces. If you water your lawn on a hot afternoon, when the relative humidity is low, much of the water will evaporate quickly from the lawn, instead of soaking into the ground. Watering the same lawn in the evening,

FIGURE 4.6 (a) At the same air temperature, an increase in the water vapor content of the air increases the relative humidity as the air approaches saturation. (b) With the same water vapor content, an increase in air temperature causes a decrease in relative humidity as the air moves farther away from being saturated.

supersaturated, a condition that does not tend to occur often or last long. A change in relative humidity can be brought about in two primary ways: . by changing the air’s water vapor content . by changing the air temperature

*We can also see in Fig. 4.6a that as the total number of vapor molecules increases (at a constant temperature), the actual vapor pressure increases and approaches the saturation vapor pressure at 20°C. As the actual vapor pressure approaches the saturation vapor pressure, the air approaches saturation, and the relative humidity rises.

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In Fig. 4.6a, we can see that an increase in the water vapor content of the air (with no change in air temperatempera ture) increases the air’s relative humidity. As more water vapor molecules are added to the air, there is a greater likelihood that some of the vapor molecules will stick together and condense. Condensation takes place in saturated air. So, as more and more water vapor molecules are added to the air, the air gradually approaches saturation, and the relative humidity of the air increases.* Conversely, removing water vapor from the air decreases the likelihood of saturation, which lowers the air’s relative humidity. In

*Another way to look at this concept is to realize that, as the air temperature increases, the air’s saturation vapor pressure also increases. As the saturation vapor pressure increases, with no change in water vapor content, the air moves farther away from saturation, and the relative humidity decreases.

FIGURE 4.7 When the air is cool (morning), the relative humidity is high. When the air is warm (afternoon), the relative humidity is low. These conditions exist in clear weather when the air is calm or of constant wind speed.

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or during the early morning when the relative humidity is higher, will cut down the evaporation and increase the effectiveness of the watering. RELATIVE HUMIDITY AND DEW POINT Suppose it is early morning and the outside air is saturated. The air temperature is 10°C (50°F) and the relative humidity is 100 percent. We know from the previous section that relative humidity can be expressed as RH 5

actual vapor pressure 100 percent. saturation vapor pressure

Looking back at Fig. 4.5, p. 83, we can see that air with a temperature of 10°C has a saturation vapor pressure of about 12 mb. Since the air is saturated and the relative humidity is 100 percent, the actual vapor pressure must be the same as the saturation vapor pressure (12 mb), since RH 5

12 mb 100 % 5 100 percent. 12 mb

Suppose during the day the air warms to 30°C (86°F), with no change in water vapor content (or air pressure). Because there is no change in water vapor content, the actual vapor pressure must be the same (12 mb) as it was in the early morning when the air was saturated. The saturation vapor pressure, however, has increased because the air temperature has increased. From Fig. 4.5, note that air with a temperature of 30°C has a saturation vapor pressure of about 42 mb. The relative humidity of this unsaturated, warmer air is now much lower, as RH 5

12 mb 100 % 5 29 percent. 42 mb

To what temperature must the outside air, with a temperature of 30°C, be cooled so that it is once again saturated? The answer, of course, is 10°C. For this amount of water vapor in the air, 10°C is called the dew-point temperature or, simply, the dew point. It represents the temperature to which air would have to be cooled (with no change in air pressure or moisture content) for saturation to occur. Since atmospheric pressure varies only slightly at Earth’s surface, the dew point is a good indicator of the air’s actual water vapor content. High dew points indicate high water vapor content; low dew points, low water vapor content. Adding water vapor to the air increases the dew point; removing water vapor lowers it. Figure 4.8 shows the average dew-point temperatempera tures across the United States and southern Canada for January. Notice that the dew points are highest (the greatest amount of water vapor in the air) over the Gulf Coast states and lowest over the interior. Compare New Orleans with Fargo. Cold, dry winds from northern Canada flow relentlessly into the Central Plains during the winter, keeping

DID YOU KNOW? The highest dew point ever measured and confirmed as accurate in the United States (90°F) occurred at Appleton, Wisconsin, on July 13, 1995. Even though the Midwest is far from the warm Gulf of Mexico, dew points can rise well above 80°F here in midsummer as crops and soils release water vapor into the atmosphere, especially after heavy rain.

this area dry. But warm, moist air from the Gulf of Mexico helps maintain a higher dew-point temperature in the southern states. Figure 4.9 is a similar diagram showing the average dew-point temperatures for July. Again, the highest dew points are observed along the Gulf Coast, with some areas experiencing average dew-point temperatures near 75°F. In fact, most people consider it to be “humid” when the dew-point temperature exceeds 65°F, and “oppressive” when it equals or exceeds 75°F. Note, too, that the dew points over the eastern and central portion of the United States are much higher in July, meaning that the July air contains between 3 and 6 times more water vapor than the January air. The reason for the high dew points is that, in summertime, this region is almost constantly receiving humid air from the warm Gulf of Mexico. The lowest dew point, and hence the driest air, is found in the West, with the lowest values observed in Nevada—a region surrounded by mountains that effectively shield it from significant amounts of moisture moving in from the southwest and northwest. The difference between air temperature and dew point can indicate whether the relative humidity is low or high. When the air temperature and dew point are far apart, the relative humidity is low; when they are close to the same value, the relative humidity is high. When the air temperature and dew point are equal, the air is saturated and the relative humidity is 100 percent.* Under certain conditions, the air can be considered “dry” even though the relative humidity may be be 100 percent. Observe, for example, in Fig. 4.10a that because the air temperature and dew point are the same in the polar air, the air is saturated and the relative humidity is 100 percent. On the other hand, the desert air (Fig. 4.10b), with a large separation between air temperature and dew point, has a much lower relative humidity, 21 percent.** However, since dew point is a measure of the amount of water vapor in the air, the desert air (with a higher dew *As a general rule of thumb, the difference between air temperature and dew point is roughly 10°F when the relative humidity is around 70%, and roughly 20°F when the relative humidity is around 50%. **The relative humidity can be computed from Fig. 4.5 (p. 83). The desert air with an air temperature of 35°C has a saturation vapor pressure of about 56 mb. A dew-point temperature of 10°C gives the desert air an actual vapor pressure of about 12 mb. These values produce a relative humidity of 12/56 100, or 21 percent. HUMIDITY, CONDENSATION, AND CLOUDS

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FIGURE 4.8 Average surface dew-point tempera temperatures (ºF) across the United States and Canada for January.

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FIGURE 4.9 Average surface dew-point tempera temperatures across the United States and Canada (°F) for July.

point) must contain more water vapor. So even though the polar air has a higher relative humidity, the desert air that contains more water vapor has a higher water vapor density, or absolute humidity. (The specific humidity and mixing ratio are also higher in the desert air.) 86

Now we can see why polar air is often described as being “dry” when the relative humidity is high (often close to 100 percent). In cold, polar air, the dew point and air temperature are normally close together. But the low dewpoint temperature means that there is little water vapor

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© C. Donald Ahrens

© C. Donald Ahrens

FIGURE 4.10 In this example, the polar air has the higher relative humidity, whereas the desert air, with the higher dew point, contains more water vapor.

in the air. Consequently, the air is “dry” even though the relative humidity is high. There is a misconception that if it is raining (or snowing), the outside relative humidity must be 100 percent. Look at Fig. 4.11 and observe that inside the cloud the relative humidity is 100 percent, but at the ground the

relative humidity is much less than 100 percent. As the rain falls into the drier air near the surface, some of the drops evaporate, a process that chills the air and increases the air’s water vapor content. The lowering air temperature and rising dew point cause the relative humidity to rise. If the falling rain persists, the air at the surface may become saturated and the relative humidity may reach 100 percent.

BRIEF REVIEW

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Up to this point we have looked at the different ways of describing humidity. Before going on, here is a review of some of the important concepts and facts we have covered: ●

Relative humidity tells us how close the air is to being saturated.

Relative humidity can change when the air’s water-vapor content changes, or when the air temperature changes.

With a constant amount of water vapor, cooling the air raises the relative humidity and warming the air lowers it.

The dew-point temperature is a good indicator of the air’s watervapor content: High dew points indicate high water-vapor content; and low dew points, low water-vapor content.

Where the air temperature and dew point are close together, the relative humidity is high; when they are far apart, the relative humidity is low.

Dry air can have a high relative humidity when the air is very cold and the air temperature and dew point are close together.

FIGURE 4.11 Inside the cloud the air temperature (T) T) and T dew point ((T Td) are the same, the air is saturated, and the relative humidity (RH) is 100 percent. However, at the surface where the air temperature and dew point are not the same, the air is not saturated (even though it is raining), and the relative humidity is considerably less than 100 percent.

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RELATIVE HUMIDITY AND HUMAN DISCOMFORT On a hot, muggy day when the relative humidity is high, it is common to hear someone exclaim (often in exasperation), “It’s not the heat, it’s the humidity!” Actually, this statement has validity. In warm weather, the main source of body cooling is through evaporation of perspiration. Recall from Chapter 2 that evaporation is a cooling process, so when the air temperature is high and the relative humidity low, perspiration on the skin evaporates quickly, often making us feel that the air temperature is lower than it really is. However, when both the air temperature and relative humidity are high and the air is nearly saturated with water vapor, body moisture does not readily evaporate; instead, it collects on the skin as beads of perspiration. Less evaporation means less cooling, and so we usually feel warmer than we did with a similar air temperature but a lower relative humidity. A good measure of how cool the skin can become is the wet-bulb temperature—the lowest temperature that can be reached by evaporating water into the air.* On a hot day when the wet-bulb temperature is low, rapid evaporation (and, hence, cooling) takes place at the skin’s surface. As the wet-bulb temperature approaches the air temperature, less cooling occurs, and the skin temperature may begin to rise. When the wetbulb temperature exceeds the skin’s temperature, no net evaporation occurs, and the body temperature can rise quite rapidly. Fortunately, the wet-bulb temperature is almost always considerably below the temperature of the skin. When the weather is hot and muggy, a number of heat-related problems can occur. For example, in hot weather when the human body temperature rises, the hypothalamus (a gland in the brain that regulates body temperature) activates the body’s heat-regulating mechanism, and more than ten million sweat glands wet the body with as much as two liters of liquid per hour. As this perspiration evaporates, rapid loss of water and salt can result in a chemical imbalance that may lead to painful heat cramps. Excessive water loss through perspiring coupled with increasing body temperature can result in heat exhaustion—fatigue, headache, nausea, and even fainting. If one’s body temperature rises above about 41°C (106°F), heat stroke can occur, resulting in complete failure of the circulatory functions. If the body temperature continues to rise, death may result. In fact, each year across North America, hundreds of people die from heat-related maladies. Even strong, healthy individuals can succumb to heat stroke, as did the Minnesota Vikings’ all-pro offensive lineman, Korey Stringer, who collapsed after practice in Mankato, Minnesota, on July 31, 2001, and died 15 hours later. Before Stringer fainted, temperatures on the practice *Notice that the wet-bulb temperature and the dew-point temperature are different. The wet-bulb temperature is attained by evaporating water into the air, whereas the dew-point temperature is reached by cooling the air.

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field were in the 90s (°F) with the relative humidity above 55 percent. In an effort to draw attention to this serious weatherrelated health hazard, an index called the heat index (HI) is used by the National Weather Service. The index combines air temperature with relative humidity to determine an apparent temperature—what the air temperature “feels like” to the average person for various combinations of air temperature and relative humidity. For example, in Fig. 4.12, an air temperature of 102°F and a relative humidity of 60 percent produce an apparent temperature of 137°F. Heat stroke is very likely when the index reaches this level. However, as we can see from the preceding paragraph, deaths related to heat stroke can occur when the heat index value is considerably lower than 137°F. Also, the values shown in Fig. 4.12 are calculated assuming shade. Under full sunshine, the effective heat index can climb as much as 15°F higher than indicated. Tragically, hundreds and even thousands of people can be killed by a single heat wave. One example is the great Chicago heat wave of July 1995. Dew-point temperatures were extremely high, which led to several days of unusually warm overnight readings that kept residents from gaining relief from the heat. On July 13, the afternoon air temperature reached 106°F at Midway Airport, followed by an overnight low temperature of only 84°F. Many residents either had no air conditioning or could not afford to use it. More than 700 deaths occurred over five days. Since the time of that disaster, Chicago and many other cities in the United States have added neighborhood “cooling centers” and taken other steps to address the danger of heat waves. In a closed vehicle, temperatures can soar far above outdoor readings in a matter of minutes. Since 1998, in the United States, more than 600 children have died after being left in parked vehicles. When sunshine is strong, temperatures need not be blistering outside to cause such a tragedy. On an 80°F day in full sun, the air temperature inside a closed car can rise to 99°F in just ten minutes and to 114°F in half an hour. On a 100°F day, interior temperatures can top 140° within an hour. Unfortunately, “cracking” the windows does little to reduce the buildup of heat. At this point it is important to dispel a common myth about hot, humid weather. Often people will recall a particularly sultry day as having been “90 degrees with 90 percent humidity” or even “95 degrees with 95 percent humidity.” We see in Fig. 4.12 that a temperature of 90°F

DID YOU KNOW? The highest dew point ever measured in the world was 95˚F at Dhahran, Saudi Arabia, near the shore of the Persian Gulf at 2 p.m. on July 8, 2003. At the time of the record, the air temperature was 108˚F, the relative humidity was 67 percent, and the heat index was an incredible 176˚F!

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FIGURE 4.12 Air temperature (°F) and relative humidity are combined to determine an apparent temperature or heat index (HI). An air temperature of 96°F with a relative humidity of 55 percent produces an apparent tempera temperature (HI) of 112°F.

in summer. People suffer, too, when the relative humidity is quite low. The rapid evaporation of moisture from exposed flesh causes skin to crack, dry, flake, or itch. These low humidities also irritate the mucous membranes in the nose and throat, producing an “itchy” or “scratchy” throat.

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with 90 percent relative humidity would produce a heat index of 122°F. Although this weather situation is remotely possible, it is extremely unlikely, as a temperature of 90°F and a relative humidity of 90 percent can occur only if the dew-point temperature is incredibly high (nearly 87°F), and a dew point this high rarely occurs anywhere in the United States, even on the muggiest of days. Similarly, in hot muggy weather, there are people who will remark about how “heavy” or how dense the air feels. Is hot, humid air really more dense than hot, dry air? If you are interested in the answer, read Focus section 4.1. Up to this point we’ve only looked at the discomfort brought on by high humidity. Can a very low relative humidity have an adverse effect on humans, too? During the winter, the relative humidity inside a home can drop to an extremely low value and the inhabitants are usually unaware of it. When cold polar air is brought indoors and heated, its relative humidity decreases dramatically. Notice in Fig 4.13 that when outside air with a temperature and dew point of 5°F is brought indoors and heated to 68°F, the relative humidity of the heated air drops to 8 percent—a value lower than what you would normally experience in a desert during the hottest time of the day. Very low relative humidities in a house can have an adverse effect on living things inside. For example, house plants have a difficult time surviving because the moisture from their leaves and the soil evaporates rapidly. Thus, they usually need watering more frequently in winter than

FIGURE 4.13 When outside air with an air temperature and a dew point of 5°F is brought indoors and heated to a temperature of 68°F (without adding water vapor to the air), the relative humidity drops to 8 percent, placing stress on plants, animals, and humans living inside. (T T represents temperature; Td, dew point; and RH, relative humidity.) HUMIDITY, CONDENSATION, AND CLOUDS

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FOCUS

ON A SSPECIAL TOPIC 4.1

FIGURE 1 On this summer afternoon in Maryland, lighter (lessdense) hot, humid air rises and condenses into towering cumulus clouds.

that changes into liquid cloud droplets and ice crystals, which, in turn, grow large enough to fall to Earth as precipitation. Of lesser importance to weather but of greater importance to sports is the fact that a baseball will “carry” farther in less-dense air. Consequently, without the influence of wind, a ball

Similarly, dry nasal passages permit inhaled bacteria to incubate, causing persistent infections. The remedy for most of these problems is simply to increase the relative humidity. Inside the home, the relative humidity can be increased simply by heating water and allowing it to evaporate into the air. The added water vapor raises the relative humidity to a more comfortable level. In modern homes, a humidifier, installed near the furnace, adds moisture to the air at a rate of about one gallon per room per day. This air, with its increased water vapor, is circulated throughout the home by a forced-air heating system. In this way, all rooms get their fair share of moisture, not just the room where the vapor is added. So, if your throat begins to feel scratchy while you are inside during a cold winter day, boil a pot of water and see if increasing the relative humidity of the air can help soothe your irritated throat. MEASURING HUMIDITY Humidity is most often measured today with automated instruments (though some measurements are still taken manually at some observing sites). One common instrument used to obtain dew point 90

will travel slightly farther on a hot, humid day than it will on a hot, dry day. So when the sports announcer proclaims that “the air today is heavy because of the high humidity,” remember that this statement is not true and, in fact, a 404-foot home run on this humid day might simply be a 400-foot out on a very dry day.

and relative humidity is the psychrometer, which consists of two liquid-in-glass thermometers mounted side by side and attached to a piece of metal that has either a handle or chain at one end (see Fig. 4.14). The thermometers are exactly alike except that one has a piece of cloth (wick) covering the bulb. The wick-covered thermometer—called

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Does a volume of hot, humid air really weigh more than a similar-size volume of hot, dry air? The answer is no! At the same temperature and at the same level in the atmosphere, hot, humid air is lighter (less dense) than hot, dry air. This is because a molecule of water vapor (H2O) weighs appreciably less than a molecule of either nitrogen (N2) or oxygen (O2). (Keep in mind that we are referring strictly to water vapor—a gas—and not suspended liquid droplets.) Consequently, in a given volume of air, as lighter water vapor molecules replace either nitrogen or oxygen molecules one for one, the number of molecules in the volume does not change, but the total weight of the air becomes slightly less. Since air density is the mass of air in a volume, the more humid air must be lighter than the drier air. Hence, hot, humid air at the surface is lighter (less dense) than hot, dry air. This fact can have an important influence in the weather. The lighter the air becomes, the more likely it is to rise. All other factors being equal, hot, humid (less-dense) air will rise more readily than hot, dry (more-dense) air (see Fig. 1). It is, of course, the water vapor in the rising air

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Humid Air and Dry Air Do Not Weigh the Same

FIGURE 4.14 The sling psychrometer consists of two thermo thermometers, one of which has a wick covering the bulb, the wet bulb thermometer.

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the wet bulb—is dipped in clean water, whereas the other thermometer is kept dry. Both thermometers are ventilated for a few minutes, either by being whirled (sling psychrometer) or by having air drawn past it with an electric fan (aspirated psychrometer). Water evaporates from the wick and that thermometer cools. The drier the air, the greater the amount of evaporation and cooling. After a few minutes, the wick-covered thermometer will cool to the lowest value possible. Recall from an earlier section that this is the wet-bulb temperature—the lowest temperature that can be attained by evaporating water into the air. The dry thermometer (commonly called the dry bulb) gives the current air temperature, or dry-bulb temperature. The temperature difference between the dry bulb and the wet bulb is known as the wet-bulb depression. A large depression indicates that a great deal of water can evaporate into the air and that the relative humidity is low. A small depression indicates that little evaporation of water vapor is possible, so the air is close to saturation and the relative humidity is high. If there is no depression, the dry bulb, the wet bulb, and the dew point are the same; the air is saturated and the relative humidity is 100 percent. Instruments that measure humidity are commonly called hygrometers. One type—called the hair hygrometer—uses human or horse hair or a synthetic fiber to measure relative humidity. It is constructed on the principle that, as the relative humidity increases, the length of hair increases and, as the relative humidity decreases, so does the hair length. A number of strands of hair (with oils removed) are attached to a system of levers. A small change in hair length is magnified by a linkage system and rela transmitted to a dial ( Fig. 4.15) calibrated to show relative humidity, which can then be read directly or recorded on a chart. (Often, the chart is attached to a clock-driven rotating drum that gives a continuous record of relative humidity.) Because the hair hygrometer is not as accurate as the psychrometer (especially at very high and very low

FIGURE 4.15 The hair hygrometer measures relative humidity by amplifying and measuring changes in the length of human (or horse) hair or a synthetic fiber.

relative humidities), it requires frequent calibration, principally in areas that experience large daily variations in relative humidity. An automated instrument that measures humidity is the electrical hygrometer. It consists of a flat plate coated with a film of carbon. An electric current is sent across the plate. As water vapor is absorbed, the electrical resistance of the carbon coating changes. These changes are translated into relative humidity. This instrument is commonly used in the radiosonde, which gathers atmospheric data at various levels above Earth. The dew-point hygrometer measures the dew-point temperature by cooling the surface of a mirror until condensation (dew) forms. This sensor is the type that measures dew-point temperature in the hundreds of fully automated weather stations— Automated Surface Observing System (ASOS)—that exist throughout the United States. (A picture of ASOS is shown in Fig. 3.22, p. 73.) Over the last several sections we have seen that, as the air cools, the air temperature approaches the dewpoint temperature and the relative humidity increases. When the air temperature reaches the dew point, the air is saturated with water vapor and the relative humidity is 100 percent. Continued cooling, however, causes some of the water vapor to condense into liquid water. The cooling may take place in a thick portion of the atmosphere, or it may occur near Earth’s surface. In the next section, we will examine condensation that forms near the ground.

Dew and Frost On clear, calm nights, objects near Earth’s surface cool rapidly by emitting infrared radiation. The ground and objects on it often become much colder than the surrounding air. Air that comes in contact with these cold surfaces cools by conduction. Eventually, the air cools to the dew point. As surfaces (such as twigs, leaves, and blades of grass) cool below this temperature, water vapor begins to condense upon them, forming tiny visible specks tempera of water called dew (see Fig. 4.16). If the air temperature should drop to freezing or below, the dew will freeze, becoming tiny beads of ice called frozen dew. Because the coolest air is usually at ground level, dew is more likely to form on blades of grass than on objects several feet above the surface. This thin coating of dew, of course, dampens bare feet, but, more importantly, it is a valuable source of moisture for many plants during periods of low rainfall. Dew is more likely to form on nights that are clear and calm than on nights that are cloudy and windy. Clear nights allow objects near the ground to cool rapidly, and calm winds mean that the coldest air will be located at ground level. These atmospheric conditions are usually associated with large fair-weather, high-pressure systems. HUMIDITY, CONDENSATION, AND CLOUDS

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© Ross DePaola

© C. Donald Ahrens

FIGURE 4.16 Dew forms on clear nights when objects on the surface cool to a temperature below the dew point. If these beads of water should freeze, they would become frozen dew.

On the other hand, the cloudy, windy weather that inhibits rapid cooling near the ground and the forming of dew often signifies the approach of a rain-producing storm system. These observations inspired the following folk rhyme: When the dew is on the grass, rain will never come to pass. When grass is dry at morning light, look for rain before the night!

Visible white frost forms on cold, clear, calm mornings when the dew-point temperature is at or below freezing. When the air temperature cools to the dew point (now called the frost point) and further cooling occurs, water vapor can change directly to ice without becoming a liquid first—a process called deposition.* The delicate, white crystals of ice that form in this manner are called hoarfrost, white frost, or simply frost. Frost has a treelike branching *When the ice changes back into vapor without melting, the process is called sublimation.

FIGURE 4.17 These are the delicate ice-crystal patterns that frost exhibits on a window during a cold winter morning.

pattern that easily distinguishes it from the nearly spherical beads of frozen dew (see Fig. 4.17). In very dry weather, the air may become quite cold and drop below freezing without ever reaching the frost point, and no visible frost forms. Freeze and black frost are words denoting this situation, one that can severely damage certain crops (see Chapter 3, p. 60). As a deep layer of air cools during the night, its relative humidity increases. When the air’s relative humidity reaches about 75 percent, some of its water vapor may begin to condense onto tiny floating particles of sea salt and other substances—condensation nuclei—that are hygroscopic (“water-seeking”) in that they allow water vapor to condense onto them when the relative humidity is considerably below 100 percent. As water collects onto these nuclei, their size increases and the particles, although still small, are now large enough to scatter visible light in all directions, becoming haze—a layer of particles dispersed through a portion of the atmosphere (see Fig. 4.18).

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FIGURE 4.18 The high relative humidity of the cold air above the lake is causing a layer of haze to form on a still winter morning.

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As the relative humidity gradually approaches 100 percent, the haze particles grow larger, and condensation begins on the less-active nuclei. Now water is condensing onto a large fraction of the available nuclei, causing the droplets to grow even bigger, until eventually they become visible to the naked eye. The increasing size and concentration of droplets further restrict visibility. When the visibility lowers to less than 1 km (or 0.62 mi) and the air is wet with millions of tiny floating water droplets, the haze becomes a cloud resting near the ground, which we call fog.*

Fog Fog, like any cloud, usually forms in one of two ways: . by cooling—air is cooled below its saturation point (dew point); and . by evaporation and mixing—water vapor is added to the air by evaporation, and the moist air mixes with relatively dry air. Once fog forms it is maintained by new fog droplets, which constantly form on available nuclei. In other words, the air must maintain its degree of saturation either by continual cooling or by evaporation and mixing of vapor into the air. Fog produced by Earth’s radiational cooling is called radiation fog, or ground fog. It forms best on clear nights when a shallow layer of moist air near the ground is overlain by drier air. Under these conditions, the ground cools rapidly since the shallow, moist layer does not absorb much of Earth’s outgoing infrared radiation. As the ground cools, *This is the official international definition of fog. The United States National Weather Service reports fog as a restriction to visibility when fog restricts the visibility to 6 miles or less and the spread between the air temperature and dew point is 5°F or less. When the visibility is less than one-quarter of a mile, the fog is considered dense.

so does the air directly above it, and a surface inversion forms, with colder air at the surface and warmer air above. The moist, lower layer (chilled rapidly by the cold ground) quickly becomes saturated, and fog forms. The longer the night, the longer the time of cooling and the greater the likelihood of fog. Hence, radiation fogs are most common over land in late fall and winter. Another factor promoting the formation of radiation fog is a light breeze of less than five knots. Although radiation fog may form in calm air, slight air movement brings more of the moist air in direct contact with the cold ground and the transfer of heat occurs more rapidly. A strong breeze tends to prevent a radiation fog from forming by mixing the air near the surface with the drier air above. The ingredients of clear skies and light winds are associated with large high-pressure areas (anticyclones). Consequently, during the winter, when a high becomes stagnant over an area, radiation fog may form on consecutive days. Because cold, heavy air drains downhill and collects in valley bottoms, we normally see radiation fog forming in low-lying areas. Hence, radiation fog is frequently called valley fog. The cold air and high moisture content in river valleys make them susceptible to radiation fog. Since radiation fog normally forms in lowlands, hills may be clear all day long, while adjacent valleys are fogged in (see Fig. 4.19). Radiation fogs are normally deepest around sunrise. Usually, however, a shallow fog layer will dissipate or burn off by afternoon. Of course, the fog does not “burn”; rather, sunlight penetrates the fog and warms the ground, causing the temperature of the air in contact with the ground to increase. The warm air rises and mixes with the foggy air above, which increases the temperature of the foggy air. In the slightly warmer air, some of the fog droplets evaporate, allowing more sunlight to reach the ground, which produces more heating, and soon the fog completely evaporates and disappears. If the fog layer is quite thick, it may

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FIGURE 4.19 Radiation fog nestled in a valley in central Oregon.

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Herbert Spichtinger/Bridge/Corbis

FIGURE 4.20 Advection fog rolling in past the Golden Gate Bridge in San Francisco. As fog moves inland, the air warms and the fog lifts above the surface. Eventually, the air becomes warm enough to totally evaporate the fog.

not completely dissipate and a layer of low clouds (called stratus) covers the region. This type of fog is sometimes called high fog. When warm, moist air moves over a sufficiently colder surface, the moist air may cool to its saturation point, forming advection fog. A good example of advection fog can be observed along the Pacific Coast during summer. The main reason fog forms in this region is that the surface water near the coast is much colder than the surface water farther offshore. Warm, moist air from the Pacific Ocean is carried (advected) by westerly winds over the cold coastal waters. Chilled from below, the air temperature drops to the dew point, and fog forms. Advection fog, unlike radiation fog, always involves the movement of air, so when there is a stiff summer breeze in San Francisco, it’s common to watch advection fog roll in past the Golden Gate Bridge (see Fig. 4.20). As summer winds carry the fog inland over warmer land, the fog near the ground dissipates, leaving a sheet of low-lying gray clouds that block out the sun. Farther inland, the air is sufficiently warm, so that even these low clouds evaporate and disappear. Because they provide moisture to the coastal redwood trees, advection fogs are important to the scenic beauty of the Pacific Coast. The needles and branches of the redwoods absorb moisture from the fog, allowing the trees to grow very tall without having to draw moisture from their roots far below. Additional moisture drips to the ground (fog drip), where it is utilized by the tree’s shallow root system. Without the summer fog, the coast’s redwood trees would have trouble surviving the dry California summers. Advection fogs also prevail where two ocean currents with different temperatures flow next to one another. Such is the case in the Atlantic Ocean off the coast of Newfoundland, where the cold, southward-flowing Labrador Current 94

lies almost parallel to the warm northward-flowing Gulf Stream. Warm southerly air moving over the cold water produces fog in that region—so frequently that fog occurs on about two out of three days during summer. Advection fog also forms over land. In winter, warm, moist air from the Gulf of Mexico moves northward over progressively colder and slightly elevated land. As the air cools to its saturation point, fog will form in the southern or central United States. Because the cold ground is often the result of radiational cooling, fog that forms in this manner is sometimes called advection-radiation fog. During this same time of year, air moving across the warm Gulf Stream encounters the colder land of the British Isles and produces the thick fogs of England. Similarly, fog forms as marine air moves over an ice or snow surface. In extremely cold arctic air, ice crystals form instead of water droplets, producing an ice fog. Keep in mind that advection fog forms when wind blows moist air over a cooler surface, whereas radiation fog forms under relatively calm conditions. Figure 4.21 visually summarizes the formation of these two types of fog. Fog that forms as moist air flows up along an elevated plain, hill, or mountain is called upslope fog. Typically, upslope fog forms during the winter and spring on the eastern side of the Rockies, where the eastward-sloping plains are nearly a kilometer higher than the land farther east. Occasionally, cold air moves from the lower eastern plains westward. The air gradually rises, expands, becomes cooler, and—if sufficiently moist—a fog forms (see Fig. 4.22). Upslope fogs that form over an extensive area can last for days. So far, we have seen how the cooling of air produces fog. But remember that fog can also form from the mixing of two unsaturated masses of air. Fog that forms in this

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FIGURE 4.22 Upslope fog forms as moist air slowly rises, cools, and condenses over elevated terrain.

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manner is usually called evaporation fog because evaporation initially enriches the air with water vapor. Probably a more appropriate name for the fog is evaporation (mixing) fog. On a cold day, you may have unknowingly produced evaporation (mixing) fog. When moist air from your mouth or nose meets the cold air and mixes with it, the air becomes saturated, and a tiny cloud forms with each exhaled breath. A common form of evaporation-mixing fog is the steam fog, which forms when cold air moves over warm water. This type of fog forms above a heated outside swimming pool in winter. As long as the water is warmer than the unsaturated air above, water will evaporate from the pool into the air. The increase in water vapor raises the dew point, and, if mixing is sufficient, the air above becomes saturated. The colder air directly above the water is heated from below and becomes warmer than the air directly above it. This warmer air rises and, from a distance, the rising condensing vapor appears as “steam.” It is common to see steam fog forming over lakes on autumn mornings, as cold air settles over water still warm from the long summer. On occasion, over the Great Lakes and other warm bodies of water, columns of condensed vapor rise from the fog layer, forming whirling steam devils, which appear similar to the dust devils observed on land. If you travel to Yellowstone National Park, you will see steam fog forming above thermal ponds all year long (see Fig. 4.23). Over the ocean in polar regions, steam fog is referred to as arctic sea smoke. Steam fog may form above a wet surface on a sunny day. This type of fog is commonly observed after a rain shower as sunlight shines on a wet road, heats the asphalt, and quickly evaporates the water. This added vapor mixes with the air above, producing steam fog. Fog that forms in this manner is short-lived and disappears as the road surface dries. A warm rain falling through a layer of cold, moist air can produce fog. As a warm raindrop falls into a cold layer of air, some of the water evaporates from the raindrop into the air. This process may saturate the air,

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FIGURE 4.21 (a) Radiation fog tends to form on clear, relatively calm nights when cool, moist surface air is overlain by drier air and rapid radiational cooling occurs. (b) Advection fog forms when the wind moves moist air over a cold surface and the moist air cools to its dew point.

FIGURE 4.23 Even in summer, warm air rising above thermal pools in Yellowstone National Park condenses into a type of steam fog.

and if mixing occurs, fog forms. Fog of this type is often associated with warm air riding up and over a mass of colder surface air. The fog usually develops in the shallow layer of cold air just ahead of an approaching warm front or behind a cold front, which is why this type of evaporation fog is also known as precipitation fog, or frontal fog. HUMIDITY, CONDENSATION, AND CLOUDS

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DID YOU KNOW?

Foggy Weather

Ever hear of caribou fog? No, it’s not the fog that forms in Caribou, Maine, but the fog that forms around herds of caribou. In very cold weather, just a little water vapor added to the air will saturate it. Consequently, the perspiration and breath from large herds of caribou add enough water vapor to the air to create a blanket of fog that hovers around the herd.

back to the driver’s eyes from the fog droplets makes it dif difficult to see very far down the road. Along a gently sloping highway, the elevated sections may have excellent visibility, while in lower regions—only a few miles away—fog can cause poor visibility. Driving from a clear area into fog on a major freeway can be extremely dangerous. In fact, every winter many people are involved in fog-related auto accidents. These usually occur when a driver enters fog and, because of the reduced visibility, puts on the brakes to slow down. The car behind then slams into the slowed vehicle, causing a chain-reaction accident with many cars involved. Airports suspend flight operations when fog causes visibility to drop below a prescribed minimum. The

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The foggiest regions in the United States are shown in Fig. 4.24. Notice that dense fog is more prevalent in coastal margins (especially those regions lapped by cold ocean currents and near the Great Lakes and Appalachian Mountains) than in the center of the continent. In fact, the foggiest spot near sea level in the United States is Cape Disappointment, Washington. Located at the mouth of the Columbia River, it averages 2556 hours or the equivalent of 106.5 twenty-four hour days of dense fog each year. Anyone hoping to enjoy the sun during August and September by traveling to this spot would find its name appropriate indeed. Notice in Fig. 4.24 that the coast of Maine is also foggy. In fact, Moose Peak Lighthouse on Mistake Island averages 1580 hours (66 equivalent days) of dense fog. To the south, Nantucket Island has on average 2040 hours (85 equivalent days) of fog each year. Extremely limited visibility exists while driving at night in thick fog with the high-beam lights on. The light scattered

FIGURE 4.24 Average annual number of days with dense fog (visibility less than 0.25 miles) across North America. (Dense fog observed in small mountain valleys and on mountaintops is not shown.)

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FOCUS

ON AN ENV ENVIRONMENTAL ISSUE 4.2

In any airport fog-clearing operation, the problem is to improve visibility so that aircraft can take off and land. Experts have tried various methods, which can be grouped into four categories: (1) increase the size of the fog droplets, so that they become heavy and settle to the ground as a light drizzle; (2) seed cold fog with dry ice (solid carbon dioxide), so that fog droplets are converted into ice crystals; (3) heat the air, so that the fog evaporates; and (4) mix the cooler saturated air near the surface with the warmer unsaturated air above. To date, only one of these methods has been reasonably successful—the seeding of cold fog. Cold fog forms when the air temperature is below freezing, and most of the fog droplets remain as liquid water. (Liquid fog in belowfreezing air is also called supercooled fog.) The fog can be cleared by injecting several hundred pounds of dry ice into it. As the tiny pieces of cold (–78°C) dry ice descend, they freeze some of the supercooled fog droplets in their path, producing ice crystals. As we will see in Chapter 5, these crystals then grow larger at the expense of the remaining liquid fog droplets. Hence, the fog droplets evaporate and the larger ice crystals fall to the ground, which leaves

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Fog Dispersal

FIGURE 2 Helicopters hovering above an area of shallow fog can produce a clear area by mixing the drier air into the foggy air below.

a “hole” in the fog for aircraft takeoffs and landings. Unfortunately, most of the fogs that close airports in the United States are warm fogs that form when the air temperature is above freezing. Since dry ice seeding does not work in warm fog, other techniques must be tried. One method involves injecting hygroscopic (water-absorbing) particles into the fog. Large salt particles and other chemicals absorb the tiny fog droplets and form into larger drops. A mix of more large drops and fewer small drops improves the visibility; plus, the larger drops are more likely to fall as a light drizzle. However,

resulting delays and cancellations become costly to the airline industry and irritate passengers. With fog-caused problems such as these, it is no wonder that scientists have been seeking ways to disperse or, at least, “thin,” fog. (For more information on fog-thinning techniques, read Focus section 4.2 entitled “Fog Dispersal.”) Up to this point, we have looked at the different forms of condensation that occur on or near Earth’s surface. In particular, we learned that fog is simply many billions of tiny liquid droplets (or ice crystals) that form near the ground. In the following sections, we will see how these same particles, forming well above the ground, are classified and identified as clouds.

BRIEF REVIEW Before we go on to the section on clouds, here is a brief review of some of the important concepts and facts we have covered so far:

since the chemicals are expensive and the fog clears for only a short time, this method of fog dispersal is not economically feasible. Another technique for fog dispersal is to warm the air enough so that the fog droplets evaporate and visibility improves. Tested at Los Angeles International Airport in the early 1950s, this technique was abandoned because it was smoky, expensive, and not very effective. In fact, the burning of hundreds of dollars worth of fuel only cleared the runway for a short time. And the smoke particles, released during the burning of the fuel, provided abundant nuclei for the fog to recondense upon. A final method of warm fog dispersal uses helicopters to mix the air. The chopper flies across the fog layer, and the turbulent downwash created by the rotor blades brings drier air above the fog into contact with the moist fog layer (see Fig. 2). The aim, of course, is to evaporate the fog. Experiments show that this method works well, as long as the fog is a shallow radiation fog with a relatively low liquid water content. But many fogs are thick, have a high liquid water content, and form by other means. An inexpensive and practical method of dispersing warm fog has yet to be discovered.

Dew, frost, and frozen dew generally form on clear nights when the temperature of objects on the surface cools below the air’s dewpoint temperature.

Visible white frost forms in saturated air when the air temperature is at or below freezing. Under these conditions, water vapor can change directly to ice, in a process called deposition.

Condensation nuclei act as surfaces on which water vapor condenses. Those nuclei that have an affinity for water vapor are called hygroscopic.

Fog is a cloud resting on the ground. It can be composed of water droplets, ice crystals, or a combination of both.

Radiation fog, advection fog, and upslope fog all form as the air cools. The cooling for radiation fog is mainly radiational cooling at Earth’s surface; for advection fog, the cooling is mainly warm air moving over a colder surface; for upslope fog, the cooling occurs as moist air gradually rises and expands along sloping terrain. Evaporation (mixing) fog, such as steam fog and frontal fog, forms as water evaporates and mixes with drier air.

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1. High clouds Cirrostratus (Cs) Cirrus (Ci) Cirrocumulus (Cc)

3. Low clouds Stratus (St) Stratocumulus (Sc) Nimbostratus (Ns)

describe clouds as they appear to a ground observer. He named a sheetlike cloud stratus (Latin for “layer”); a puffy cloud cumulus (“heap”); a wispy cloud cirrus (“curl of hair”); and a rain cloud nimbus (“violent rain”). In Howard’s system, these were the four basic cloud forms. Other clouds could be described by combining the basic types. For example, nimbostratus is a rain cloud that shows layering, whereas cumulonimbus is a rain cloud having pronounced vertical development. In 1887, Ralph Abercromby and Hugo Hildebrandsson expanded Howard’s original system and published a classification system that, with only slight modification, is still in use today. Ten principal cloud forms are divided into four primary cloud groups. Each group is identified by the height of the cloud’s base above the surface: high clouds, middle clouds, and low clouds. The fourth group contains clouds showing more vertical than horizontal development. Within each group, cloud types are identified by their appearance. ▼Table 4.1 lists these four groups and their cloud types. The approximate base height of each cloud group is given in ▼Table 4.2. Note that the altitude separating the high and middle cloud groups overlaps and varies with latitude. Large temperature changes cause most of this latitudinal variation. For example, high cirriform clouds are composed almost entirely of ice crystals. In subtropical regions, air temperatures low enough to freeze all liquid water usually occur only above about 20,000 feet. In polar regions, however, these same temperatures may be found at altitudes as low as 10,000 feet. Although you may observe cirrus clouds at 12,000 feet over northern Alaska, you will not see them at that elevation above southern Florida. Clouds cannot be accurately identified strictly on the basis of elevation. Other visual clues are necessary. Some of these are explained in the following section.

2. Middle clouds Altostratus (As) Altocumulus (Ac)

4. Clouds with vertical development Cumulus (Cu) Cumulonimbus (Cb)

HIGH CLOUDS High clouds in middle and low latitudes generally form above 16,000 ft (5000 m). Because the air at these elevations is quite cold and dry, high clouds are composed almost exclusively of ice crystals and are also rather

Clouds Clouds are aesthetically appealing and add excitement to the atmosphere. Without them, there would be no rain or snow, thunder or lightning, rainbows or halos. How monotonous it would be if there were only a clear blue sky to look at. A cloud is a visible aggregate of tiny water droplets or ice crystals suspended in the air. Some are found only at high elevations, whereas others nearly touch the ground (or are classified as fog if they do touch the ground). Clouds can be thick or thin, big or little—they exist in a seemingly endless variety of forms. To impose order on this variety, we divide clouds into ten basic types. With a careful and practiced eye, you can become reasonably proficient in correctly identifying them. CLASSIFICATION OF CLOUDS Although ancient astronomers named the major stellar constellations about 2000 years ago, clouds were not formally identified and classified until the early nineteenth century. The French naturalist Jean-Baptiste Lamarck (1744–1829) proposed the first system for classifying clouds in 1802; however, his work did not receive wide acclaim. One year later, Luke Howard, an English naturalist, developed a cloud classification system that found general acceptance. In essence, Howard’s innovative system employed Latin words to ▼ Table 4.1

▼ Table 4.2

The Four Major Cloud Groups and Their Types

Approximate Height* of Cloud Bases Above the Surface for Various V Locations

CLOUD GROUP

TROPICAL REGION

MID-LATITUDE REGION

POLAR REGION

High Ci, Cs, Cc

20,000 to 60,000 ft (6000 to 18,000 m)

16,000 to 43,000 ft (5000 to 13,000 m)

10,000 to 26,000 ft (3000 to 8000 m)

Middle As, Ac

6500 to 26,000 ft (2000 to 8000 m)

6500 to 23,000 ft (2000 to 7000 m)

6500 to 13,000 ft (2000 to 4000 m)

Low St, Sc, Ns

surface to 6500 ft (0 to 2000 m)

surface to 6500 ft (0 to 2000 m)

surface to 6500 ft (0 to 2000 m)

*Note that the height of a cloud base in each region varies by season.

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FIGURE 4.26 Cirrocumulus clouds.

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The thin, sheetlike, high clouds that often cover the entire sky are cirrostratus (Cs), which are so thin that the sun and moon can be clearly seen through them (see Fig. 4.27). The ice crystals in these clouds bend the light passing through them and will often produce a halo— a ring of light that encircles the sun or moon. In fact, the veil of cirrostratus may be so thin that a halo is the only clue to its presence. Thick cirrostratus clouds give the sky a glary white appearance and frequently form ahead of an advancing mid-latitude cyclonic storm; hence, they can be used to predict rain or snow within twelve to twenty-four hours, especially if they are followed by middle-type clouds.

FIGURE 4.25 Cirrus clouds. Notice the silky “mare’s tail” appearance.

*Small quantities of liquid water in cirrus clouds at temperatures as low as –36°C (–33°F) were discovered during research conducted above Boulder, Colorado.

MIDDLE CLOUDS The middle clouds have bases between about 6500 and 23,000 ft (2000 and 7000 m) in the middle latitudes. These clouds are composed of water droplets and—when the temperature becomes low enough—some ice crystals. Precipitation can form in middle clouds if they become thick enough. Altocumulus (Ac) clouds are middle clouds that appear as gray, puffy masses, sometimes rolled out in parallel waves or bands (see Fig. 4.28). Usually, one part of each

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thin.* High clouds usually appear white, except near sunrise and sunset, when the unscattered (red, orange, and yellow) components of sunlight are reflected from the underside of the clouds. The most common high clouds are cirrus (Ci), which are thin, wispy clouds blown by high winds into long streamers called mares’ tails. Notice in Fig. 4.25 that they can look like a white, feathery patch with a faint wisp of a tail at one end. Cirrus clouds usually move across the sky from west to east, indicating the prevailing winds at their elevation, and they generally occur during periods of fair, pleasant weather. Cirrocumulus (Cc) clouds, seen less frequently than cirrus, appear as small, rounded, white puffs that may occur individually, or in long rows (see Fig. 4.26). When in rows, the cirrocumulus cloud has a rippling appearance that distinguishes it from the silky look of cirrus and the sheetlike cirrostratus. Cirrocumulus seldom cover more than a small portion of the sky. The dappled cloud elements that reflect the red or yellow light of a setting sun make this one of the most beautiful of all clouds. The small ripples in the cirrocumulus strongly resemble the scales of a fish; hence, the expression “mackerel mackerel sky” commonly describes a sky full of cirrocumulus clouds.

FIGURE 4.27 Cirrostratus clouds. Notice the faint halo encircling the sun. The sun is the bright white area in the center of the circle. HUMIDITY, CONDENSATION, AND CLOUDS

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FIGURE 4.29 Altostratus clouds. The appearance of a dimly visible “watery sun” through a deck of light gray clouds is usually a good indication that the clouds are altostratus.

cloud element is darker than another, which helps to distinguish it from the higher cirrocumulus. Also, the individual puffs of the altocumulus appear larger than those of the cirrocumulus. A layer of altocumulus can sometimes be confused with altostratus; in case of doubt, clouds are called altocumulus if there are rounded masses or rolls present. Altocumulus clouds that look like “little castles” (castellanus) in the sky indicate the presence of rising air at cloud level. The appearance of these clouds on a warm, humid summer morning often portends thunderstorms by late afternoon. Altostratus (As) are gray or blue-gray clouds that often cover the entire sky over an area that extends over many hundreds of square kilometers. In the thinner section of the cloud, the sun (or moon) may be dimly

visible as a round disk, which is sometimes referred to as a “watery sun” (see Fig. 4.29). Thick cirrostratus clouds are occasionally confused with thin altostratus clouds. The gray color, height, and dimness of the sun are good clues to identifying an altostratus. The fact that halos only occur with cirriform clouds also helps one to distinguish them. Another way to separate the two is to look at the ground for shadows. If there are none, it is a good bet that the cloud is altostratus because cirrostratus are usually transparent enough to produce shadows. Altostratus clouds often form ahead of mid-latitude cyclonic storms having widespread and relatively continuous precipitation. If precipitation falls from altostratus, the cloud base usually lowers, and the precipitation is steady and not showery as found with cumuliform clouds. If the precipitation reaches the ground, the cloud is then classified as nimbostratus.

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FIGURE 4.28 Altocumulus clouds. Notice the dark-tolight contrasting patterns that distinguish these clouds from cirrocumulus clouds.

FIGURE 4.30 The nimbostratus is the sheetlike cloud from which light rain is falling. The ragged-appearing clouds beneath the nimbostratus are stratus fractus, or scud.

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LOW CLOUDS Low clouds, with their bases lying below 6500 ft (or 2000 m) are almost always composed of water droplets, although in cold weather, they may contain ice particles. Nimbostratus (Ns) are dark-gray, “wet”-looking cloud layers associated with more or less continuously falling rain or snow (see Fig. 4.30). The intensity of this precipitation is usually light or moderate; it is never of the heavy, showery variety, unless well-developed cumuliform clouds are embedded within the nimbostratus cloud. Precipitation often makes the base of the nimbostratus cloud impossible to identify clearly. The distance from the cloud’s base to the top can be over 3 km (10,000 feet). Nimbostratus is easily confused with altostratus. Thin nimbostratus is usually darker gray than thick altostratus, and you normally cannot see the sun or moon through a layer of nimbostratus. Visibility below a nimbostratus cloud deck

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FIGURE 4.31 Stratocumulus clouds forming along the south coast of Flor Florida. Notice that the rounded masses are larger than those of the altocumulus.

is usually quite poor because rain will evaporate and mix with the air in this region. If this air becomes saturated, a lower layer of clouds or fog may form beneath the original cloud base, as seen in Fig. 4.30. Since these lower clouds drift rapidly with the wind, they form irregular shreds with a ragged appearance called stratus fractus, or scud. Stratocumulus (Sc) are low, lumpy clouds that appear in rows, in patches, or as rounded masses with blue sky visible between individual cloud elements (see Fig. 4.31). Often they appear near sunset as the spreading remains of a much larger cumulus cloud. Occasionally, the sun will shine through the cloud breaks, producing bands of light (called crepuscular rays) that appear to reach down to the ground. The color of stratocumulus ranges from light to dark gray. This cloud type differs from altocumulus in that it has a lower base and larger individual clouds. (Compare

Fig. 4.28 with Fig. 4.31.) To distinguish between the two, hold your hand at arm’s length and point toward one of these clouds. Altocumulus cloud elements will generally be about the size of your thumbnail, whereas stratocumulus will usually be about the size of your fist. Although precipitation rarely falls from stratocumulus, light rainshowers or winter snow flurries can occur if the cloud develops vertically into a much thicker cloud with a top colder than about –5°C (23°F). Stratus (St) is a uniform grayish cloud that often covers the entire sky. It resembles a fog that does not reach the ground (see Fig. 4.32). Actually, when a thick fog “lifts,” the resulting cloud is a deck of low stratus. Normally, no precipitation falls from stratus, but sometimes it is accompanied by a light mist or drizzle. This cloud commonly occurs over Pacific and Atlantic coastal waters in summer.

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FIGURE 4.32 A layer of lowlying stratus clouds hides the mountains in Iceland.

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FIGURE 4.33 Cumulus clouds. Small cumulus clouds such as these are sometimes called fair weather cumulus, or cumulus humilis.

A thick layer of stratus might be confused with nimbostratus, but the distinction between them can be made by observing the base of the cloud. Often, stratus has a more uniform base than does nimbostratus. Also, a deck of stratus may be confused with a layer of altostratus. However, if you remember that stratus clouds are lower and darker gray, the distinction can be made. CLOUDS WITH VERTICAL DEVELOPMENT Familiar to almost everyone, the puffy cumulus (Cu) cloud takes on a variety of shapes, but most often it looks like a piece of floating cotton with sharp outlines and a flat base (see Fig. 4.33). The base appears white to light gray, and, on a humid day, may be only a few thousand feet above the ground and half a mile or so wide. The top of the cloud— often in the form of rounded towers—denotes the limit of rising air and is usually not very high. These clouds can be distinguished from stratocumulus by the fact that cumulus clouds are detached (usually with a great deal

of blue sky between each cloud) whereas stratocumulus usually occur in groups or patches. Also, the cumulus has a dome- or tower-shaped top as opposed to the generally flatter tops of the stratocumulus. Cumulus clouds that show only slight vertical growth (cumulus humilis) are associated with fair weather; therefore, we call these clouds “fair-weather cumulus.” If the cumulus clouds are small and appear as broken fragments of a cloud with ragged edges, they are called cumulus fractus. Harmless-looking cumulus often develop on warm summer mornings and, by afternoon, become much larger and more vertically developed. When the growing cumulus resembles a head of cauliflower, it becomes a cumulus congestus, or towering cumulus (Tcu). Most often, it is a single large cloud, but, occasionally, several grow into each other, forming a line of towering clouds, as shown in Fig. 4.34. Precipitation that falls from a cumulus congestus is always showery with frequent changes in intensity.

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FIGURE 4.34 Cumulus congestus clouds are frequently called towering cumulus. These clouds are taller than cumulus clouds and are more likely to produce showers. Here a line of cumulus congestus clouds is building along Maryland’s eastern shore.

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FIGURE 4.35 A cumulonimbus cloud. Strong upper-level winds blowing from right to left produce a well-defined anvil shape. Sunlight scattered by falling ice crystals produces the white (bright) area beneath the anvil. Notice the heavy rainshower falling from the base of the cloud.

If a cumulus congestus continues to grow vertically, it develops into a giant cumulonimbus (Cb)—a thunderstorm cloud (see Fig. 4.35). While its dark base may be no more than 2000 ft above Earth’s surface, its top can extend upward to the tropopause, over 39,000 ft higher, and sometimes thousands of feet beyond that. A cumulonimbus can occur as an isolated cloud or as part of a line or “wall” of clouds. The tremendous amounts of energy released by the condensation of water vapor within a cumulonimbus results in the development of violent updrafts and downdrafts, which can exceed 70 knots. The lower (warmer) part of the cloud is usually composed only of water droplets. Higher up in the cloud, water droplets and ice crystals both abound, while, toward the cold top, there are only ice crystals. Swift winds at these higher altitudes can reshape the top of the cloud into a huge flattened anvil.* These great thunderheads may contain all forms of precipitation—large raindrops, snowflakes, snow pellets, and sometimes hailstones—all of which can fall to Earth in the form of heavy showers. Lightning, thunder, and even tornadoes are associated with cumulonimbus. (More information on the violent nature of thunderstorms and tornadoes is given in Chapter 10.) Cumulus congestus and cumulonimbus frequently look alike. However, you can usually distinguish them by looking at the top of the cloud. If the sprouting upper part of the cloud is sharply defined and not fibrous, it is usually a cumulus congestus; conversely, if the top of the cloud loses its sharpness and becomes fibrous in texture, it is usually a cumulonimbus. (Compare Fig. 4.34 with Fig. 4.35.) The weather associated with these clouds also differs: Lightning, thunder, and large hail typically occur with cumulonimbus. So far, we have discussed the ten primary cloud forms, summarized pictorially in Fig. 4.36. This figure, along *An anvil is a heavy block of iron or steel with a smooth, flat top on which metals are shaped by hammering.

with the cloud photographs and descriptions, should help you learn to identify the more common cloud forms. Don’t worry if you find it hard to estimate cloud heights. This is a difficult procedure, requiring much practice. You can use local objects (hills, mountains, tall buildings) of known height as references on which to base your height estimates. To better describe a cloud’s shape and form, a number of descriptive words can be used in conjunction with its name. We mentioned a few in the previous section; for example, a stratus cloud with a ragged appearance is a stratus fractus, and a cumulus cloud with marked vertical growth is a cumulus congestus. ▼Table 4.3 lists some of the more common terms that are used in cloud identification. SOME UNUSUAL CLOUDS Although the ten basic cloud forms are the most frequently seen, there are some unusual clouds that deserve mentioning. For example, moist air crossing a mountain barrier often forms into waves. The clouds that form in the wave crest usually have a lens shape and are, therefore, called lenticular clouds (see Fig. 4.37). Frequently, they form one above the other like a stack of pancakes, and at a distance they can resemble hovering spacecraft. Hence, it is no wonder that many UFO sightings take place when lenticular clouds are present. Similar to the lenticular cloud is the cap cloud, or pileus, which usually resembles a silken scarf capping the top of a sprouting cumulus cloud (see Fig. 4.38). Pileus clouds form when moist winds are deflected up and over

DID YOU KNOW? Cloud lovers have a newly identified type to watch for. Called asperitas, this dramatic cloud appears as a rolling, very turbulent, choppy wave cloud that looks very ominous, but doesn’t produce stormy weather. In 2016, the World Meteorological Organization (WMO) was evaluating whether to make asperitas the first new cloud formation to be officially recognized since 1951. HUMIDITY, CONDENSATION, AND CLOUDS

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FIGURE 4.36 A generalized illustration of basic cloud types based on height above the surface and vertical development.

▼ Table 4.3

Common Terms Used in Identifying Clouds

TERM

LATIN ROOT AND MEANING

DESCRIPTION

Lenticularis

(lens, lenticula, lentil)

Clouds having the shape of a lens; often elongated and usually with well-defined outlines, a term that applies mainly to cirrocumulus, altocumulus, and stratocumulus

Fractus

( (frangere, to break or fracture)

Clouds that have a ragged or torn appearance; applies only to stratus and cumulus

Humilis

(humilis, of small size)

Cumulus clouds with generally flattened bases and slight vertical growth

Congestus

(congerere, to bring together; to pile up)

Cumulus clouds of great vertical extent that, from a distance, may resemble a head of cauliflower

Undulatus

(unda, wave; having waves)

Clouds in patches, sheets, or layers showing undulations

Translucidus

(translucere, to shine through; transparent)

Clouds that cover a large part of the sky and are sufficiently translucent to reveal the position of the sun or moon

Mammatus

(mamma, mammary)

Baglike clouds that hang like a cow’s udder on the underside of a cloud; may occur with cirrus, altocumulus, altostratus, stratocumulus, and cumulonimbus

Pileus

( (pileus, cap)

A cloud in the form of a cap or hood above or attached to the upper part of a cumuliform cloud, particularly during its developing stage

Castellanus

(castellum, a castle)

Clouds that show vertical development and produce towerlike extensions, often in the shape of small castles

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© UCAR, Photo by Carlye Calvin

FIGURE 4.37 Lenticular clouds forming downwind of the Front Range of the Rocky Mountains, near Boulder, Colorado.

FIGURE 4.38 A pileus cloud forming above a developing

the top of a building cumulus congestus or cumulonimbus. If the air flowing over the top of the cloud condenses, a pileus often forms. Most clouds form in rising air, but the mammatus forms in sinking air. Mammatus clouds derive their name from their appearance—baglike sacs that hang beneath the cloud and resemble mammary glands (see Fig. 4.39). Although mammatus most frequently form on the underside of cumulonimbus, they may develop beneath cirrus, cirrocumulus, altostratus, altocumulus, and stratocumulus. Jet aircraft flying at high altitudes often produce a cirruslike trail of condensed vapor called a condensation trail or contrail (see Fig. 4.40). The condensation may come directly from the water vapor added to the air from engine exhaust. In this case, there must be sufficient mixing of the hot exhaust gases with the cold air to produce saturation. Contrails evaporate rapidly when the relative humidity of the surrounding air is low. If the relative

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cumulus cloud.

FIGURE 4.39 Mammatus clouds forming beneath a thunderstorm anvil.

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the summer at latitudes poleward of 50°, although they have been observed in recent years in Utah and Colorado. Studies reveal that these clouds are composed of tiny ice crystals. The water to make the ice may originate in meteoroids that disintegrate when entering the upper atmosphere or from the chemical breakdown of methane gas at high levels in the atmosphere. Our system of classifying clouds may never be completely finalized. One identified formation, asperitas, has only recently been recognized as a specific cloud type (see Did You Know on p. 103 and Fig. 4.43). FIGURE 4.40 A contrail forming behind a jet aircraft.

humidity is high, however, contrails can persist for many hours. Contrails can also form by a cooling process as the reduced pressure produced by air flowing over the wing causes the air to cool. Aside from the cumulonimbus cloud that sometimes penetrates into the stratosphere, all of the clouds described so far are observed in the lower atmosphere, the troposphere. Occasionally, however, clouds can be seen above the troposphere. For example, soft pearly looking clouds called nacreous clouds, or mother-of-pearl clouds, form in the stratosphere at altitudes above 30 km or 100,000 ft (see Fig. 4.41). They are best viewed in polar latitudes durdur ing the winter months when the sun, being just below the horizon, is able to illuminate them because of their high altitude. Their exact composition is not known, although they appear to be composed of water in either solid or liquid (supercooled) form. Wavy bluish-white clouds, so thin that stars shine brightly through them, can sometimes develop in the upper mesosphere, at altitudes above 75 km (46 mi). These clouds are at such a high altitude that they appear bright against a dark background. For this reason, they are called noctilucent clouds, meaning “luminous night clouds” (see Fig. 4.42). They are best seen at twilight, during

CLOUDS AND SATELLITE IMAGERY The weather satellite is a cloud-observing platform in Earth’s orbit. Satellites provide extremely valuable cloud images of areas where there are no ground-based observations. Because water covers over 70 percent of Earth’s surface, there are vast regions where few (if any) surface cloud observations are made. Before weather satellites were in use, tropical storms, such as hurricanes and typhoons, often went undetected until they moved dangerously near inhabited areas. Residents of the regions affected had little advance warning. Today, satellites spot these storms while they are still far out in the ocean and track them accurately. Two primary types of weather satellites are used for observing clouds. The first are called geostationary satellites (or geosynchronous satellites) because they orbit the equator at the same rate Earth spins and, hence, remain at nearly 36,000 km (22,300 mi) above a fixed spot on Earth’s surface (see Fig. 4.44). This positioning allows continuous monitoring of a specific region. Geostationary satellites are also important because they use a “real time” data system, meaning that the satellites transmit images to the receiving system on the ground as soon as the image is taken. Successive cloud images from these satellites can be put into a time-lapse movie sequence to show the cloud movement, dissipation, or development associated with weather fronts and storms. This information is a great help in forecasting the progress

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FIGURE 4.41 The clouds in this photograph are nacreous clouds. They form in the stratosphere and are most easily seen at high latitudes.

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FIGURE 4.42 The wavy clouds in this photograph are noctilucent clouds. They are usually observed at high latitudes, at altitudes above 75 km.

FIGURE 4.43 This asperitas, photographed over eastern Colorado, is a cloud that looks ominous but does not produce stormy weather.

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of large weather systems. Wind directions and speeds at various levels can also be approximated by monitoring cloud movement with the geostationary satellite. Complementing the geostationary satellites are polarorbiting satellites, which closely parallel Earth’s meridian (longitude) lines. These satellites pass over the north and south polar regions on each revolution. As Earth rotates to the east beneath the satellite, each pass monitors an area to the west of the previous pass (see Fig. 4.45). Eventually, the satellite covers the entire Earth. Polar-orbiting satellites have the advantage of scanning clouds directly beneath them. Thus, they provide sharp images in polar regions, where images from a geostationary satellite are distorted because of the low angle at which the satellite “sees” those regions. Polar orbiters also circle Earth at a much lower altitude (about 850 km, or 530 mi) than geostationary satellites. The lower altitude allows them to provide detailed images of phenomena such as violent storms and cloud systems. Continuously improved detection devices make weather observation by satellites more versatile than ever. Early satellites, such as TIROS I, launched in 1960, used television cameras to photograph clouds. Contemporary satellites use radiometers, which can observe clouds during both day and night by detecting radiation that emanates from the top of the clouds. When the radiometer measures only visible sunlight reflected from the cloud, the image is called a visible cloud image. Additionally, satellites have the capacity to obtain cloud images and, at the same time, provide vertical profiles of atmospheric temperature and moisture by detecting emitted radiation from atmospheric gases, such as water vapor. In modern satellites, a special type of advanced radiometer (called an imager) provides satellite images with much better resolution than did previous imagers. Moreover, another type of special radiometer (called a sounder) gives a more accurate profile of temperature and moisture at different levels in the atmosphere than did earlier instruments. The latest

FIGURE 4.44 The geostationary satellite moves through space at the same rate that Earth rotates, so it remains above a fixed spot on the equator and monitors one area constantly.

FIGURE 4.45 Polar-orbiting satellites scan from north to south, and on each successive orbit the satellite scans an area farther to the west. HUMIDITY, CONDENSATION, AND CLOUDS

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Figure 4.47 shows a visible satellite image (from a geostationary satellite) of a mid-latitude cyclonic storm in the eastern Pacific. Notice that all of the clouds in the image appear white. However, in the infrared image ( Fig. 4.48), taken on the same day (and at just about the same time), the clouds appear to have many shades of gray. In the visible image, the clouds covering part of Oregon and northern California appear relatively thin compared to the thicker, bright clouds to the west. Furthermore, these thin clouds must be high because they also appear bright in the infrared image. Along the elongated band of clouds associated with the occluded front and cold front, the clouds appear white and bright in both images, indicating a zone of thick, heavy clouds. Behind the front, the lumpy clouds are probably cumulus because they appear gray in the infrared image (Fig. 4.48), suggesting that their tops are low and relatively warm. When temperature differences are small, it is difficult to directly identify significant cloud and surface features on an infrared image. Some way must be found to increase the contrast between features and their backgrounds. This can be accomplished by a process called computer enhancement. Certain temperature ranges in the infrared image are assigned specific shades of gray, grading from black to

NOAA

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in the Geostationary Operational Environmental Satellite (GOES) series, GOES-R, includes many additional features, including higher spatial resolution and the ability to map lightning activity from space. Information on cloud thickness and height can be deduced from satellite images. Visible images show the sunlight reflected from a cloud’s upper surface. Because thick clouds have a higher reflectivity (albedo) than thin clouds, they appear brighter on a visible satellite image. However, middle and low clouds have just about the same reflectivity, so it is difficult to distinguish among them simply by viewing them in visible light. To make this distinction, infrared cloud images are used. Such pictures produce a better image of the actual radiating surface because they do not show the strong visible reflected light. Since warm objects radiate more energy than cold objects, high temperature regions can be artificially made to appear darker on an infrared image. Because the tops of low clouds are warmer than those of high clouds, cloud observations made in the infrared can distinguish between warm low clouds (dark) and cold high clouds (light)—see Fig. 4.46. Moreover, cloud temperatures can be converted by a computer into a threedimensional image of the cloud. These are the 3-D cloud photos presented on television by many weathercasters.

FIGURE 4.46 Generally, the lower the cloud, the warmer its top. Warm objects emit more infrared energy than do cold objects. Thus, an infrared satellite picture can distinguish warm, low (gray) clouds from cold, high (white) clouds.

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FIGURE 4.47 A visible image of the eastern Pacific Ocean taken at just about the same time on the same day as the image in Fig. 4.48. Notice that the clouds in the visible image appear white. Superimposed on the image are the cold, warm, and occluded fronts.

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NOAA

FIGURE 4.50 Infrared water-vapor image. The darker areas represent dry air aloft; the brighter the gray, the more moist the air in the middle or upper troposphere. Bright white areas represent dense cirrus clouds or the tops of thunderstorms. The area in color represents the coldest cloud tops. The swirl of moisture off the West Coast represents a well-developed mid-latitude cyclonic storm.

NOAA

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white. Often, clouds with cold tops, and those with tops near freezing, are assigned the darkest gray color. Figure 4.49 is an enhanced infrared image for the same day and area as shown in Figs. 4.47 and 4.48. To make cloud features more obvious, colors such as dark blue, red, or purple are assigned to clouds with the coldest (highest) tops. Hence, the dark red areas embedded along the occluded front and cold front in Fig. 4.49 represent the region where the coldest and, therefore, highest and thickest clouds are found. It is here where the stormiest weather is probably occurring. Also notice that, near the southern tip of the picture, the dark red blotches surrounded by areas of white are thunderstorms that have developed over warm tropical waters. They show up clearly as white, thick clouds in both the visible and infrared images. By examining the movement of these clouds on successive satellite images, forecasters can predict the arrival of clouds and storms, and the passage of weather fronts. In regions where there are no clouds, it is difficult to observe the movement of the air. To help with this situation, geostationary satellites are equipped with watervapor sensors that can profile the distribution of atmospheric water vapor in the middle and upper troposphere (see Fig. 4.50). In time-lapse films, the swirling patterns

FIGURE 4.48 Infrared satellite image of the eastern Pacific Ocean taken at just about the same time on the same day as the image in Fig. 4.47. Notice that the low clouds in this infrared image appear in various shades of gray.

FIGURE 4.49 An enhanced infrared image of the eastern Pacific Ocean taken on the same day as the images shown in Figs. 4.47 and 4.48.

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NASA Scientific Visualization Studio; data from NASA/JAXA GPM

of moisture clearly show wet regions and dry regions, as well as middle tropospheric swirling wind patterns and jet streams. Specialized satellites have gathered data on clouds and precipitation for more than 20 years. From 1997 to 2015, the long-lived Tropical Rainfall Measuring Mission (TRMM) satellite provided information on clouds and precipitation from about 35°S to 35°N. A joint venture of NASA and the Japan Aerospace Exploration Agency (JAXA), this satellite orbited Earth at an altitude of about 400 km (250 mi), sensing individual cloud features as small as about 1.5 miles in diameter. TRMM gathered three-dimensional images of clouds and storms, details on the intensity and distribution of precipitation, and data on Earth’s energy budget and lightning discharges within storms. TRMM has been succeeded by another NASA/JAXA project called the Global Precipitation Mission (GPM), whose core observatory was launched in 2014. GPM covers a much broader swath than TRMM—from about 65°S to 65°N—and it includes advanced sensors that can distinguish precipitation intensity and type as well as cloud characteristics (see Fig. 4.51). Another specialized satellite has also provided enhanced detail on clouds and precipitation. Launched in 2006, the NASA CloudSat satellite circles Earth in an orbit

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FIGURE 4.51 Precipitation in Hurricane Arthur as sensed by the Global Positioning Mission ((GPM) satellite on July 3, 2014, off the South Carolina coast. Colors on the surface ranging from light green to dark red to purple indicate areas ranging from low to high rainfall. The violet areas aloft indicate frozen precipitation.

FIGURE 4.52 (Top) Visible satellite image of super-typhoon Choi-Wan over the tropical eastern Pacific Ocean on September 15, 2009. (Bottom) CloudSat vertical radar profile through super-typhoon Choi-Wan. The location of the profile is shown by the red line in the top view.

about 700 km (435 mi) above the surface. Onboard CloudSat, a very sensitive radar (called Cloud Profiling Radar, or CPR) uses microwave radiation to peer into a cloud and unveil its very fine structures, including the altitude of the cloud’s top and base, its thickness, optical properties, the abundance of liquid and ice particles, along with the intensity of precipitation inside the cloud. CloudSat provides this information in a vertical view, as shown in Fig. 4.52. Such vertical profiling of a cloud’s makeup will potentially provide scientists with a better understanding of precipitation processes that go on inside the cloud and the role that clouds play in Earth’s global climate system.

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SUMMARY

KEY TERMS

In this chapter, we examined the hydrologic cycle and saw how water is circulated within our atmosphere. We then looked at some of the ways of describing humidity and found that relative humidity does not tell us how much water vapor is in the air but, rather, how close the air is to being saturated. A good indicator of the air’s actual water vapor content is the dew-point temperature. When the air temperature and dew point are close together, the relative humidity is high; when they are far apart, the relative humidity is low. When the air temperature drops below the dew point in a shallow layer of air near the surface, dew forms. If the dew freezes, it becomes frozen dew. Visible white frost forms when the air cools to a below-freezing dew-point temperature. As the air cools in a deeper layer near the surface, the relative humidity increases and water vapor begins to condense upon “water-seeking” hygroscopic condensation nuclei, forming haze. As the relative humidity approaches 100 percent, the air can become filled with tiny liquid droplets (or ice crystals) called fog. Upon examining fog, we found that it forms in two primary ways: (1) when air cools and evaporates and (2) when water evaporates mixes into the air. Condensation above Earth’s surface produces clouds. When clouds are classified according to their height and physical appearance, they are divided into four main groups: high, middle, and low clouds, and clouds with vertical development. Since each cloud type has physical characteristics that distinguish it from others, careful observation normally leads to correct identification. Satellites enable scientists to obtain a bird’s-eye view of clouds on a global scale. Polar-orbiting satellites obtain data covering Earth from pole to pole, while geostationary satellites located above the equator continuously monitor a desired portion of Earth. Both types of satellites use radiometers (imagers) that detect emitted radiation. As a consequence, clouds can be observed both day and night. Visible satellite images, which show sunlight reflected from a cloud’s upper surface, can distinguish thick clouds from thin clouds. Infrared images show an image of the cloud’s radiating top and can distinguish low clouds from high clouds. To increase the contrast between cloud features, infrared photographs are enhanced. Specialized instruments aboard satellites can be used to analyze cloud characteristics and precipitation.

The following terms are listed (with corresponding page numbers) in the order they appear in the text. Define each. Doing so will aid you in reviewing the material covered in this chapter. evaporation, 80 condensation, 80 precipitation, 80 hydrologic cycle, 80 saturated air, 81 condensation nuclei, 81 humidity, 82 actual vapor pressure, 83 saturation vapor pressure, 83 relative humidity, 84 supersaturated air, 84 dew-point temperature (dew point), 85 wet-bulb temperature, 88 heat stroke, 88 heat index (HI), 88 apparent temperature, 88 psychrometer, 90 hygrometer, 91 dew, 91 frost, 92 haze, 92 fog, 93 radiation fog, 93 advection fog, 94 upslope fog, 94 evaporation (mixing) fog, 95

cirrus clouds, 99 cirrocumulus clouds, 99 cirrostratus clouds, 99 altocumulus clouds, 99 altostratus clouds, 100 nimbostratus clouds, 100 stratocumulus clouds, 101 stratus clouds, 101 cumulus clouds, 102 cumulonimbus clouds, 103 lenticular clouds, 103 pileus clouds, 103 mammatus clouds, 105 contrail, 105 nacreous clouds, 106 noctilucent clouds, 106 geostationary satellites, 106 polar-orbiting satellites, 107

QUESTIONS FOR REVIEW . . . . .

Briefly explain the movement of water in the hydrologic cycle. How does condensation differ from precipitation? What are condensation nuclei and why are they important in our atmosphere? In a volume of air, how does the actual vapor pressure differ from the saturation vapor pressure? When are they the same? What does saturation vapor pressure primarily depend upon?

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. (a) What does the relative humidity represent? (b) When the relative humidity is given, why is it also important to know the air temperature? (c) Explain two ways the relative humidity can be changed. (d) During what part of the day is the relative humidity normally lowest? Normally highest? . Why do hot, humid summer days usually feel hotter than hot, dry summer days? . Why is cold polar air described as “dry” even when the relative humidity of that air is very high? . Why is the wet-bulb temperature a good measure of how cool human skin can become? . (a) What is the dew-point temperature? (b) How is the difference between dew point and air temperature related to the relative humidity? . How can you obtain both the dew point and the relative humidity using a sling psychrometer? . Explain how dew, frozen dew, and visible frost form. . List the two primary ways in which fog forms. . Describe the conditions that are necessary for the formation of: (a) radiation fog (b) advection fog . How does evaporation (mixing) fog form? . Clouds are most generally classified by height above Earth’s surface. List the major height categories and the cloud types associated with each. . How can you distinguish altostratus clouds from cirrostratus clouds? . Which clouds are normally associated with each of the following characteristics? (a) mackerel sky (b) lightning (c) halos (d) hailstones (e) mares’ tails (f) anvil top (g) light continuous rain or snow (h) heavy rain showers . Name the clouds that form above the troposphere. . How do geostationary satellites differ from polarorbiting satellites? . How can you distinguish a visible satellite image from an infrared satellite image? . Why are infrared satellite images enhanced?

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QUESTIONS FOR THOUGHT AND EXPLORATION . .

. .

.

. .

.

Use the concepts of condensation and saturation to explain why eyeglasses often fog up after you come indoors on a cold day. After completing a grueling semester of meteorological course work, you contact a travel agent to arrange a much-needed summer vacation. When the agent suggests a trip to the desert, you decline because of a concern that the dry air will make your skin feel uncomfortable. The travel agent assures you that almost daily “desert relative humidities are above 90 percent.” Could the agent be correct? Explain. Can the actual vapor pressure ever be greater than the saturation vapor pressure? Explain. Suppose while measuring the relative humidity using a sling psychrometer, you accidentally moisten both the dry-bulb and the wet-bulb thermometers. Will the relative humidity you determine be higher or lower than the air’s true relative humidity? A large family lives in northern Minnesota. This family gets together for a huge dinner three times a year: on Thanksgiving, on Christmas, and on the March equinox. The Thanksgiving and Christmas dinners consist of turkey, ham, mashed potatoes, and lots of boiled vegetables. The March dinner is pizza. The air temperature inside the home is about the same for all three meals (70°F), yet everyone remarks on how “warm, cozy, and comfortable” the air feels during the Thanksgiving and Christmas dinners, and how “cool” the inside air feels during the equinox meal. Explain to the family members why they might feel “warmer” inside the house during Thanksgiving and Christmas, and “cooler” during the March equinox. (The answer has nothing to do with the amount or type of food consumed.) Why is advection fog more common along the coast of southern California than along the coast of southern Virginia? With all other factors being equal, would you expect a lower minimum temperature on a night with cirrus clouds or on a night with stratocumulus clouds? Explain your answer. Explain why icebergs are frequently surrounded by fog.

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. While driving from cold air (well below freezing) into much warmer air (well above freezing), frost forms on the windshield of the car. Does the frost form on the inside or outside of the windshield? How can the frost form when the air is so warm? . Why do relative humidities seldom reach 100 percent in polluted air? . If all fog droplets gradually settle earthward, explain how fog can last (without disappearing) for many days at a time. . The air temperature during the night cools to the dew point in a deep layer, producing fog. Before the fog formed, the air temperature cooled each hour about 3°F. After the fog formed, the air temperature cooled by only 1°F each hour. Give two reasons why the air cooled more slowly after the fog formed.

. Why can you see your breath on a cold morning? Does the air temperature have to be below freezing for this to occur? . The sky is overcast and it is raining. Explain how you can tell if the cloud above you is a nimbostratus or a cumulonimbus. . You are sitting inside your house on a sunny afternoon. The shades are drawn and you look at the window and notice the sun disappears for about 10 seconds. The alternating light and dark periods last for nearly 30 minutes. Are the clouds passing in front of the sun cirrocumulus, altocumulus, stratocumulus, or cumulus? Give a reasonable explanation for your answer.

Go the the Basic Search field and search for news items using the keywords “cloud classification.” Review at least three news articles that describe techniques used to classify clouds based on data from weather satellites. What are the common elements among these techniques? How do they differ?

ONLINE RESOURCES Visit www.cengagebrain.com to view additional resources, including video exercises, practice quizzes, an interactive eBook, and more.

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CHAPTER

5

Cloud Development and Precipitation

Atmospheric Stability

T

Determining Stability

thought, it would be deep enough to cancel school, maybe for

Cloud Development and Stability

a day, possibly a week, or perhaps, forever. But clear skies and

Precipitation Processes

the voice from the back room that insisted, “Don’t “ even think

Contents

Precipitation Types

he young boy pushed his nose against the cold windowwindow pane, hoping to see snowflakes glistening in the light of

the streetlamp across the way. Perhaps erhaps if it snowed, he

a full moon gave little hope for snow on this evening. Nor did about snow. You know it won’t snow tonight; it’s too cold to snow.” With hopes dashed, the boy pondered: could it really be too cold to snow?

Measuring Precipitation

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Atmospheric Stability

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We know that most clouds form as air rises, expands, and cools. But why does the air rise on some occasions and not on others? And why do the size and shape of clouds vary so much when the air does rise? To answer these questions, let’s focus on the concept of atmospheric stability. When we speak of atmospheric stability, we are referring to a condition of equilibrium. For example, rock A resting in the depression in Fig. 5.1 is in stable equilibrium. If the rock is pushed up along either side of the hill and then let go of, it will quickly return to its original position. On the other hand, rock B, resting on the top of the hill, is in a state of unstable equilibrium, as a slight push will set it moving away from its original position. Applying these concepts to the atmosphere, we can see that air is in stable equilibrium when, after being lifted or lowered, it tends to return to its original position—it resists upward and downward air motions. Air that is in unstable equilibrium will, when given a little push, move farther away

FIGURE 5.1 When rock A is disturbed, it will return to its original position; rock B, however, will accelerate away from its original position.

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C

louds, spectacular features in the sky, add beauty and color to the natural landscape. Yet clouds are important for nonaesthetic reasons, too. As they form, vast quantities of heat are released into the atmosphere. Clouds help regulate Earth’s energy balance by reflecting and scattering solar radiation and by absorbing infrared energy from Earth’s surface. And, of course, without clouds there would be no precipitation. But clouds are also significant because they visually indicate the physical processes taking place in the atmosphere; to a trained observer, they are signposts in the sky. In the beginning of this chapter, we will look at the atmospheric processes these signposts point to, the first of which is atmospheric stability. Later, we will examine the different mechanisms responsible for the formation of most clouds. Toward the end of the chapter, we will peer into the tiny world of cloud droplets to see how rain, snow, and other types of precipitation form. And yes, we will answer the question raised in our opener, “Is it ever too cold to snow?”

FIGURE 5.2 The dry adiabatic rate. As long as the air parcel remains unsaturated, it expands and cools by 10°C per 1000 m; the sinking parcel compresses and warms by 10°C per 1000 m.

from its original position—it favors upward and downward motion. In order to explore the behavior of rising and sinking air, we must first review some concepts we learned in earlier chapters. Recall that a balloonlike blob of air is called an air parcel. (The concept of air parcels is illustrated in Fig. 4.4, p. 82, and in the Focus section on p. 32.) When an air parcel rises, it moves into a region where the air pressure surrounding it is lower. This situation allows the air molecules inside to push outward on the parcel walls, expanding it. As the air parcel expands, the air inside cools. If the same parcel is brought back to the surface, the increasing pressure around the parcel squeezes (compresses) it back to its original volume, and the air inside warms. Hence, a rising parcel of air expands and cools, while a sinking parcel is compressed and warms. If a parcel of air expands and cools, or compresses and warms, and there is no interchange of heat with its outside surroundings, this situation is called an adiabatic process. As long as the air in the parcel is unsaturated (the relative humidity is less than 100 percent), the rate of adiabatic cooling or warming remains constant and is about 10°C for every 1000 meters of change in altitude, or about 5.5°F for every 1000 feet. Since this rate of cooling or warming only applies to unsaturated air, it is called the dry adiabatic rate* (see Fig. 5.2). As the rising air cools, its relative humidity increases as the air temperature approaches the dew-point temperature. If the rising air cools to its dew-point temperature, the relative humidity becomes 100 percent. Further lifting results in condensation, a cloud forms, and latent heat is *For aviation purposes, the dry adiabatic rate is sometimes expressed as 3°C per 1000 ft.

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Determining Stability We determine the stability of the air by comparing the temperature of a rising parcel to that of its surroundings. If the rising air is colder than its environment, it will be more dense** (heavier) and tend to sink back to its original level. In this case, the air is stable because it resists upward movement. If the rising air is warmer and, therefore, less dense (lighter) than the surrounding air, it will continue to rise until it reaches the same temperature as its environment. This is an example of unstable air. To figure out the air’s stability, we need to measure the temperature both of the rising air and of its environment at various levels above Earth. A STABLE ATMOSPHERE Suppose we release a balloon-borne instrument called a radiosonde. (A photo of a radiosonde is found in Fig. 2 on p. 22.) As the balloon carries the radiosonde up into the atmosphere, it sends back temperature data, as shown in Fig. 5.3. Notice that the air temperature measured by the radiosonde decreases by 4°C for every 1000 meters rise in altitude. Remember *If condensed water or ice is removed from the rising saturated parcel, the cooling process is called an irreversible pseudoadiabatic process. **When, at the same level in the atmosphere, we compare parcels of air that are equal in size but vary in temperature, we find that cold air parcels are more dense than warm air parcels; that is, in the cold parcel, there are more molecules that are crowded closer together.

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released into the rising air. Because the heat added during condensation offsets some of the cooling due to expansion, the air no longer cools at the dry adiabatic rate but at a lesser rate called the moist adiabatic rate. (Because latent heat is added to the rising saturated air, the process is not really adiabatic.*) If a saturated parcel containing water droplets were to sink, it would compress and warm at the moist adiabatic rate because evaporation of the liquid droplets would offset the rate of compressional warming. Hence, the rate at which rising or sinking saturated air changes temperature—the moist adiabatic rate—is less than the dry adiabatic rate. Unlike the dry adiabatic rate, the moist adiabatic rate is not constant, but varies greatly with temperature and, hence, with moisture content, because warm saturated air produces more liquid water than cold saturated air. The added condensation in warm, saturated air liberates more latent heat. Consequently, the moist adiabatic rate is much less than the dry adiabatic rate when the rising air is quite warm; however, the two rates are nearly the same when the rising air is very cold. Although the moist adiabatic rate does vary, to make the numbers easy to deal with we will use an average of 6°C per 1000 m (3.3°F per 1000 ft) in most of our examples and calculations.

FIGURE 5.3 A stable atmosphere. Imagine that a helicopter could lift parcels of air, as shown above. An absolutely stable atmosphere exists when a rising air parcel is colder and heavier (i.e., more dense) than the air surrounding it. If given the chance (i.e., released), the air parcel in both situations would return to its original position, the surface. (In both situations, the helicopter shows that the air is being lifted. In the real world, this type of parcel lifting, of course, would be impossible.)

from Chapter 1 that the rate at which the air temperature changes with altitude is called the lapse rate. Because this rate is the one at which the air temperature surrounding us would be changing if we were to climb upward into the atmosphere, we refer to it as the environmental lapse rate. Notice in Fig. 5.3a that (with an environmental lapse rate of 4°C per 1000 m) a rising parcel of unsaturated, “dry” air is colder and heavier than the air surrounding it at all levels. Even if the parcel is initially saturated (Fig. 5.3b), as it rises it, too, will be colder than its environment at all levels. In both cases, the atmosphere is absolutely stable because the lifted parcel of air is colder and heavier than the air surrounding it. If released, the parcel will have a tendency to return to its original position. Since air in a stable atmosphere strongly resists upward vertical motion, it will, if forced to rise, tend to spread out horizontally. If clouds form in this rising air, they, too, will spread horizontally in relatively thin layers and usually have flat tops and bases. We might expect to see stratiform clouds—such as cirrostratus, altostratus, nimbostratus, or stratus—forming in a stable atmosphere. The atmosphere is stable when the environmental lapse rate is small, that is, when there is a relatively small difference in temperature between the surface air and the CLOUD DEVELOPMENT DEVELOPMENT AND PRECIPITATION

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The air aloft may warm as winds bring in warmer air or as the air slowly sinks over a large area. Recall that sinking (subsiding) air warms as it is compressed. The warming can produce an inversion, where the air aloft is actually warmer than the air at the surface. (Recall from Chapter 3 that an inversion represents an atmospheric condition where the air becomes warmer with height.) An inversion that forms by slow, sinking air is termed a subsidence inversion. Because inversions represent a very stable atmosphere, they act as a lid on vertical air motion. When an inversion exists near the ground, stratus, fog, haze, and pollutants are all kept close to the surface. In fact, as we will see in Chapter 14, most air pollution episodes occur with subsidence inversions. FIGURE 5.4 The initial environmental lapse rate in diagram (a) will become more stable (stabilize) as the air aloft warms and the surface air cools, as illustrated in diagram (b).

air aloft. Consequently, the atmosphere tends to become more stable—it stabilizes—as the air aloft warms or the surface air cools (see Fig. 5.4). The cooling of the surface air can be due to: . nighttime radiational cooling of the surface . an influx of cold surface air brought in by the wind . air moving over a cold surface

Robert Henson

So, on a typical day, the atmosphere is usually most stable in the early morning around sunrise, when the lowest surface air temperature is recorded. If the surface air becomes saturated in a stable atmosphere, a persistent layer of fog may form (see Fig. 5.5).

AN UNSTABLE ATMOSPHERE The atmosphere is unstable when the environmental air temperature decreases rapidly with height. For example, in Fig. 5.6, notice that the measured air temperature decreases by 11°C for every 1000-meter rise in altitude, which means that the environmental lapse rate is 11°C per 1000 meters. Also notice that a lifted parcel of unsaturated “dry” air in Fig. 5.6a, as well as a lifted parcel of saturated “moist” air in Fig. 5.6b, will, at each level above the surface, be warmer than the air surrounding them. Since, in both cases, the rising air parcels are warmer and less dense than the air around them, once the parcels start upward, they will continue to rise on their own, away from the surface. Thus, we have an absolutely unstable atmosphere. In an unstable environment, parcels of air are “buoyant” because the contrast between their warmer temperature and their cooler surroundings exerts an upward force on them. The warmer the air

FIGURE 5.5 On this morning near Boulder, Colorado, cold surface air has produced a stable atmosphere that inhibits vertical air motions and allows the fog to linger close to the ground.

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FIGURE 5.7 The initial environmental lapse rate in diagram (a) will become more unstable (that is, destabilize) as the air aloft cools and the surface air warms, as illustrated in diagram (b).

FIGURE 5.6 An unstable atmosphere. An absolutely unstable atmosphere exists when a rising air parcel is warmer and lighter (i.e., less dense) than the air surrounding it. If given the chance (i.e., released), the lifted parcel in both (a) and (b) would continue to move away (accelerate) from its original position. (As in Fig. 5.3, the lifting of air parcels by a helicopter would be impossible.)

parcels compared to their surroundings, the greater this buoyant force and the more rapidly they rise.* The atmosphere becomes more unstable as the environmental lapse rate steepens; that is, as the *A good example of buoyant force occurs when you submerge a balloon filled with air into a tub of water. As you push the lighter (less dense) balloon into the heavier (more dense) water, you can feel the upward-directed buoyant force acting on the balloon. Let go of the balloon and watch how rapidly it rises to the surface.

temperature of the air drops rapidly with increasing height. This circumstance can be brought on either by the air aloft becoming colder or by the surface air becoming warmer (see Fig. 5.7). The warming of the surface air may be due to: . daytime solar heating of the surface . an influx of warm surface air brought in by the wind . air moving over a warm surface The combination of cold air aloft and warm surface air can produce a steep lapse rate and an unstable atmosphere (see Fig. 5.8). Generally, then, as the surface air warms during the day, the atmosphere becomes more unstable—it destabilizes. The air aloft may cool as winds bring in colder air or as the air (or clouds) emit infrared radiation to space (radiational cooling). Just as sinking air produces

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FIGURE 5.8 The warmth from this forest fire in Idaho during August 2003 heats the air, causing instability near the surface. Warm, less-dense air (and smoke) bubbles upward, expanding and cooling as it rises. Eventually the rising air cools to its dew point, condensation begins, and a cumulus cloud forms. If the rising air parcels are large and strong enough, the resulting clouds (sometimes called “pyrocumulus”) may produce lightning and precipitation.

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warming and a more stable atmosphere, rising air, especially an entire layer where the top is dry and the bottom is humid, produces cooling and a more unstable atmosphere. The lifted layer becomes more unstable as it rises and stretches out vertically in the less-dense air aloft. This stretching effect steepens the environmental lapse rate as the top of the layer cools more than the bottom. Instability brought on by the lifting of air is often associated with the development of severe weather, such as thunderstorms and tornadoes, which are investigated more thoroughly in Chapter 10. It should be noted, however, that deep layers in the atmosphere are seldom, if ever, absolutely unstable. Absolute instability is usually limited to a very shallow layer near the ground on hot, sunny days. Here, the environmental lapse rate can exceed the dry adiabatic rate, and the lapse rate is called superadiabatic. A CONDITIONALLY UNSTABLE ATMOSPHERE Suppose an unsaturated (but humid) air parcel is somehow forced to rise from the surface, as shown in Fig. 5.9. (What causes the air parcel to rise will be covered in a later section.) As the parcel rises, it expands, and cools at the dry adiabatic rate until its air temperature cools to its dew point. At this level, the air is saturated, the relative humidity is 100 percent, and further lifting results in condensation and the formation of a cloud. The elevation above

the surface where the cloud first forms (in this example, 1000 meters) is called the condensation level. In Fig. 5.9, notice that above the condensation level, the rising saturated air cools at the moist adiabatic rate. Notice also that from the surface up to a level near 2000 meters, the rising, lifted air is colder than the air surrounding it. The atmosphere up to this level is stable. However, owing to the release of latent heat, the rising air near 2000 meters has actually become warmer than the air around it. Since the lifted air can rise on its own accord, the atmosphere is now unstable. The level in the atmosphere where the air parcel, after being lifted, becomes warmer than the air surrounding it, is called the level of free convection. The atmospheric layer from the surface up to 4000 meters in Fig. 5.9 has gone from stable to unstable because the rising air was humid enough to become saturated, form a cloud, and release latent heat, which warms the air. Had the cumulus cloud not formed, the rising air would have remained colder at each level than the air surrounding it. From the surface to 4000 meters, we have what is said to be a conditionally unstable atmosphere—the condition for instability being whether or not the rising air becomes saturated. Therefore, conditional instability means that if unsaturated stable air is somehow lifted to a level where it becomes saturated, instability may result. In Fig. 5.9, we can see that the environmental lapse rate is 9°C

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FIGURE 5.9 Conditionally unstable atmo atmosphere. The atmosphere is conditionally unstable when unsaturated, stable air is lifted to a level where it becomes saturated and warmer than the air surrounding it. If the atmosphere remains unstable, vertical developing cumulus clouds can build to great heights.

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BRIEF REVIEW

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Before going on to the next section, here is a brief review of some of the facts and concepts concerning atmospheric stability:

FIGURE 5.10 When the environmental lapse rate is greater than the dry adiabatic rate (blue region), the atmosphere is ab absolutely unstable. When the environmental lapse rate is less than the moist adiabatic rate (red region), the atmosphere is absolutely stable. And when the environmental lapse rate lies between the dry adiabatic rate and the moist adiabatic rate (green region), the atmosphere is conditionally unstable.

per 1000 meters. This value is between the dry adiabatic rate (10°C/1000 m) and the moist adiabatic rate (6°C/1000 m). Consequently, conditional instability exists whenever the environmental lapse rate is between the dry and moist adiabatic rates. Recall from Chapter 1 that the average lapse rate in the troposphere is about 6.5°C per 1000 m (3.6°F per 1000 ft). Since this value lies between the dry adiabatic rate and the average moist rate, the atmosphere is ordinarily in a state of conditional instability. Figure 5.10 summarizes how the three categories of stability (ab(ab solutely stable, conditionally unstable, and absolutely unstable) relate to the dry and moist adiabatic rates. At this point, it should be apparent that the stability of the atmosphere changes during the course of a day. In clear, calm weather around sunrise, surface air is normally colder than the air above it, a radiation inversion exists, and the atmosphere is quite stable, as indicated by smoke or haze lingering close to the ground. As the day progresses, sunlight warms the surface and the surface warms the air above. As the air temperature near the ground increases, the lower atmosphere gradually becomes more unstable, with maximum instability usually occurring during the hottest part of the day. On a humid summer afternoon this phenomenon can be witnessed by the development of cumulus clouds.

The air temperature in a rising parcel of unsaturated air decreases at the dry adiabatic rate, whereas the air temperature in a rising parcel of saturated air decreases at the moist adiabatic rate.

The dry adiabatic rate and moist adiabatic rate of cooling are different due to the fact that latent heat is released in a rising parcel of saturated air.

In a stable atmosphere, a lifted parcel of air will be colder (heavier) than the air surrounding it. Because of this fact, the lifted parcel will tend to sink back to its original position.

In an unstable atmosphere, a lifted parcel of air will be warmer (lighter) than the air surrounding it, and thus will continue to rise upward, away from its original position.

The atmosphere becomes more stable (stabilizes) as the surface air cools, the air aloft warms, or a layer of air sinks (subsides) over a vast area.

The atmosphere becomes more unstable (destabilizes) as the surface air warms, the air aloft cools, or a layer of air is lifted.

A conditionally unstable atmosphere exists when a parcel of air can be lifted to a level where it becomes saturated, a cloud forms, and the rising parcel becomes warmer than the air surrounding it.

The atmosphere is normally most stable in the early morning and most unstable in the afternoon.

Layered clouds tend to form in a stable atmosphere, whereas cumuliform clouds tend to form in a conditionally unstable atmosphere.

Cloud Development and Stability We know that most clouds form as air rises and cools and its water vapor condenses. Since air normally needs a “trigger” to start it moving upward, what is it that causes the air to rise so that clouds can form? The following mechanisms are primarily responsible for the development of the majority of clouds we observe: . surface heating and free convection . uplift along topography . widespread ascent due to the flowing together (convergence) of surface air . uplift along weather fronts (see Fig. 5.11). The first mechanism that can cause the air to rise is convection. Although we briefly looked at convection in Chapter 2 when we examined rising thermals and how they transfer heat upward into the atmosphere, we will now look at convection from a slightly different perspective: how rising thermals can produce cumulus clouds. CONVECTION AND CLOUDS Some areas of Earth’s surface are better absorbers of sunlight than others and,

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FIGURE 5.11 The primary ways clouds form: (a) surface heating and convection; (b) forced lifting along topographic barriers; (c) convergence of surface air; (d) forced lifting along weather fronts.

down Observe in Fig. 5.12 that the air motions are downward on the outside of the cumulus cloud. The downward motions are caused in part by evaporation around the outer edge of the cloud, which cools the air, making it heavy (more dense). Another reason for the downward motion is the completion of the convection current started by the thermal. Cool air slowly descends to replace the rising warm air. Therefore, we have rising air in the cloud and sinking air around it. Since subsiding air greatly inhibits the growth of thermals beneath it, small cumulus clouds usually have a great deal of blue sky between them (see Fig. 5.13).

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therefore, heat up more quickly. The air in contact with these “hot spots” becomes warmer than its surroundings. A hot “bubble” of air—a thermal—breaks away from the warm surface and rises, expanding and cooling as it ascends. As the thermal rises, it mixes with the cooler, drier air around it and gradually loses its identity. Its upward movement now slows. Before the thermal is completely diluted, subsequent rising thermals often penetrate it and help the air rise a little higher. If the rising air cools to its saturation point, the moisture will condense, and the thermal becomes visible to us as a cumulus cloud.

FIGURE 5.13 Cumulus clouds building on a warm summer afternoon. Each cloud represents a region where thermals are rising from the surface. The clear areas between the clouds are regions where the air is sinking. FIGURE 5.12 Cumulus clouds form as warm, invisible air bubbles detach themselves from the surface, then rise and cool to the condensation level. Below and within the cumulus clouds, the air is rising. Around the cloud, the air is sinking.

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FIGURE 5.14 Variations in the air’s stability, as indicated by the environmental lapse rate, greatly influence the growth of cumulus clouds.

TOPOGRAPHY AND CLOUDS Horizontally moving air obviously cannot go through a large obstacle, such as a mountain, so the air must go over it. Forced lifting along a topographic barrier is called orographic uplift. Often, large masses of air rise when they approach a long chain

of mountains such as the Sierra Nevada and Rockies. This lifting produces cooling, and if the air is humid, clouds form. Clouds produced in this manner are called orographic clouds. An example of orographic uplift and cloud development is given in Fig. 5.16. Notice that, after having risen over the mountain, the air at the surface on the leeward (downwind) side is considerably warmer than it was at the surface on the windward (upwind) side. The higher air temperature on the leeward side is the result of latent heat being converted into sensible heat during condensation on the windward side. In fact, the rising air at the top of the mountain is considerably warmer than it would have been had condensation not occurred. Notice also in Fig. 5.16 that the dew-point temperature of the air on the leeward side is lower than it was

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As the cumulus clouds grow, they shade the ground from the sun. This, of course, cuts off surface heating and upward convection. Without the continual supply of rising air, the cloud begins to erode as its droplets evaporate. Unlike the sharp outline of a growing cumulus, the cloud now has indistinct edges, with cloud fragments extending from its sides. As the cloud dissipates (or moves along with the wind), surface heating begins again and regenerates another thermal, which becomes a new cumulus. This is why you often see cumulus clouds form, gradually disappear, then reform in the same spot. The stability of the atmosphere plays an important part in determining the vertical growth of cumulus clouds. Notice in Fig. 5.14 that when a stable layer (such as an inversion) exists near the top of the cumulus cloud, the cloud would have a difficult time rising much higher, and it would remain as a “fair-weather” cumulus cloud, cumulus humilis. However, if a deep, conditionally unstable layer exists above the cloud, then the cloud can develop vertically into a towering cumulus congestus with a cauliflowerlike top. When the conditionally unstable air is several miles deep, the cumulus congestus can even develop into a cumulonimbus with a flat anvil-shaped top. Notice in Fig. 5.15 that the distant thunderstorm has an anvil-shaped shaped top. The cloud is shaped this way because it has reached the stable part of the atmosphere, and the rising air is unable to puncture very far into this stable layer, so the top of the cloud spreads laterally as high winds at this altitude (usually above 10 km or 33,000 ft) blow the cloud’s ice crystals horizontally. Atmospheric stability also plays a role in making many afternoons windier than mornings. This topic is discussed further in Focus section 5.1.

FIGURE 5.15 Cumulus clouds developing into thunderstorms in a conditionally unstable atmosphere over the Great Plains. Notice that, in the distance, the cumulonimbus with the flat anvil-shaped top has reached a stable layer of the atmosphere. CLOUD DEVELOPMENT DEVELOPMENT AND PRECIPITATION

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FOCUS

ON A SPE SPECIAL TOPIC 5.1

FIGURE 1 (a) During the early morning, there is little exchange between the surface winds and the winds aloft. (b) In the afternoon, when the atmosphere is usually most unstable, convection in the form of rising thermals links surface air with the air aloft, causing strong winds from aloft to reach the ground and produce strong, gusty surface winds.

circulation, it may pull some of the stronger winds aloft downward with it. If this sinking air should reach the surface, it produces a momentary gust of strong wind. This exchange of air also increases the average wind speed at the surface. Because this type of air exchange is greatest on a clear day in

the afternoon when the atmosphere is most unstable, we tend to experience the strongest, most gusty winds in the afternoon. At night, when the atmosphere stabilizes, the interchange between the surface air and the air aloft is at a minimum, and the winds at the surface tend to die down.

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On warm days when the weather is clear or partly cloudy, you may have noticed that the windiest time of the day is usually in the afternoon. Such windy afternoons occur because of several factors working together, including surface heating, convection, and atmospheric stability. We know that in the early morning the atmosphere is most stable, meaning that the air resists up-and-down motions. As an example, consider the flow of air in the early morning as illustrated in Fig. 1a. Notice that weak winds exist near the surface with much stronger winds aloft. Because the atmosphere is stable, there is little vertical mixing between the surface air and the air higher up. As the day progresses and the sun rises higher in the sky, the surface heats up and the lower atmosphere becomes more unstable. Over hot surfaces, the air begins to rise in the form of thermals that carry the slower-moving air with them (see Fig. 1b). At some level above the surface, the rising air links up with the faster-moving air aloft. If the air begins to sink as part of a convective

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Atmospheric Stability and Windy Afternoons—Hold On to Your Hat

FIGURE 5.16 Orographic uplift, cloud development, and the formation of a rain shadow. (Note: The reason for the decrease in the dew-point temperature of the rising, unsaturated air on the windward side, and for the increase in dew-point temperature of the unsaturated sinking air on the leeward side is given in the footnote at the bottom of p. 125.)

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. Air descending a mountain warms by compressional heating and, upon reaching the surface, can be much warmer than the air at the same level on the upwind side. . Air on the leeward side of a mountain is normally drier (has a lower dew point) than the air on the windward side. The lower dew point and higher air temperature on the leeward side produce a lower relative humidity, a greater potential for evaporation of water, and a rain shadow desert. Although clouds are more prevalent on the windward side of mountains, they may, under certain atmospheric conditions, form on the leeward side as well. For example, stable air flowing over a mountain often moves in a series of waves that may extend for several hundred miles on the leeward side. Such waves often resemble the waves that form in a river downstream from a large boulder. Recall *You may have noticed in Fig. 5.16 that the dew-point temperature of the rising unsaturated air on the windward side of the mountain beneath the cloud decreases by 2°C per 1000 m, and increases by 2°C per 1000 m in the descending unsaturated air on the leeward side. The decrease in dew-point temperature on the windward side is caused by the rapid decrease in air pressure of the rising air. Since the dew point is directly related to the actual vapor pressure, a decrease in total air pressure of the rising air causes a corresponding decrease in vapor pressure and, hence, a lowering of the dew-point temperature. Likewise, the rapid increase in air pressure of the sinking air on the leeward side causes an increase in the dew-point temperature.

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before the air was lifted over the mountain. The lower dew point and, hence, drier air on the leeward side is the result of water vapor condensing and then remaining as liquid cloud droplets and precipitation on the windward side.* This region on the leeward side of a mountain, where precipitation is noticeably low, and the air is often drier, is called a rain shadow. From Fig. 5.16 we have two important concepts to remember:

FIGURE 5.17 Lenticular clouds that form in the wave directly over the mountain are called mountain wave clouds, whereas those that form downwind of the mountain are called lee wave clouds. On the underside of the lee wave’s crest a turbulent rotor may form.

from Chapter 4 that wave clouds often have a characteristic lens shape and are called lenticular clouds. The formation of lenticular clouds is shown in Fig. 5.17. As moist air rises on the upwind side of the wave, it cools and condenses, producing a cloud. On the downwind side, the air sinks and warms, and the cloud evaporates. Viewed from the ground, the clouds appear motionless as the air rushes through them. When the air between the cloud-forming layers is too dry to produce clouds, lenticular clouds will form one above the other, sometimes extending into the stratosphere and appearing as a fleet of hovering spacecraft (see Fig. 5.18). Notice in Fig. 5.17 that beneath the lenticular cloud downwind of the mountain range, a large swirling eddy forms. The rising part of the swirling air may cool enough to produce a visible cloud called a rotor cloud. The air in the rotor is extremely turbulent and presents a major hazard

© Momatiuk-Eastcott/Corbis

FIGURE 5.18 Lenticular clouds tend to form over and downwind of mountains. They also tend to remain in one place as air rushes through them. Here, lenticular clouds are forming over mountainous terrain in Argentina’s Los Glaciares National Park.

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to aircraft in the vicinity. Dangerous flying conditions also exist near the lee side of the mountain, where strong downward air motions are present. Now, having examined the concept of stability and the formation of clouds, we are ready to see how minute cloud particles are transformed into rain and snow. The next section, therefore, takes a look at the processes that produce precipitation.

Precipitation Processes As we all know, cloudy weather does not necessarily mean that it will rain or snow. In fact, clouds may form, linger for many days, and never produce precipitation.* In Eureka, California, the August daytime sky is overcast more than 50 percent of the time, yet the average precipitation there for August is merely one-tenth of an inch. How, then, do cloud droplets grow large enough to produce rain? And why do some clouds produce rain, but not others? In Fig. 5.19, we can see that an ordinary cloud droplet is extremely small, having an average diameter of 0.02 millimeters (mm), which is less than one-thousandth of an inch. Also, notice in Fig. 5.19 that a typical cloud droplet is 100 times smaller in diameter than a typical raindrop. Clouds, then, are composed of many droplets too small to fall as rain. These minute droplets require only slight upward air currents to keep them suspended. Those

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*Recall from Chapter 4 that precipitation is any form of water (liquid or solid) that falls from a cloud and reaches the ground.

FIGURE 5.19 Relative sizes of raindrops, cloud droplets, and condensation nuclei, with diameters shown in millimeters (mm).

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droplets that do fall, descend slowly and evaporate in the drier air beneath the cloud. In Chapter 4, we learned that condensation begins on tiny particles called condensation nuclei. The growth of cloud droplets by condensation is slow and, even under ideal conditions, it would take several days for this process alone to create a raindrop. It is evident, then, that the condensation process by itself is entirely too slow to produce rain. Yet, observations show that clouds can develop and begin to produce rain in less than an hour. Since it takes about 1 million average-size cloud droplets to make an average-size raindrop, there must be some other process by which cloud droplets grow large and heavy enough to fall as precipitation. Even though all the intricacies of how rain is produced are not yet fully understood, two important processes stand out: (1) the collision-coalescence process and (2) the ice-crystal (or Bergeron) process. COLLISION AND COALESCENCE PROCESS In clouds with tops warmer than −15°C (5°F), the collisioncoalescence process can play a significant role in producing precipitation. To produce the many collisions necessary to form a raindrop, some cloud droplets must be larger than others. Larger drops can form on large condensation nuclei, such as salt particles, or through random collisions of droplets. Studies also suggest that turbulent mixing between the cloud and its drier environment can play a role in producing larger droplets. As cloud droplets fall, air slows them down. The amount of air resistance depends on the size of the drop and on its rate of fall: The greater its speed, the more air molecules the drop encounters each second. The speed of the falling drop increases until the air resistance equals the pull of gravity. At this point, the drop continues to fall, but at a constant speed, which is called its terminal velocity. Because larger drops have a smaller surface area–to-weight ratio, they must fall faster before reaching their terminal velocity. Thus, larger drops fall faster than smaller drops. Eventually, large droplets overtake and collide with smaller drops in their path. This merging of cloud droplets by collision is called coalescence. Within a cloud, there can be both updrafts and downdrafts. If an updraft is especially strong, then droplets of various sizes may all be pushed upward. Coalescence may now occur as smaller droplets are pushed upward more quickly, colliding with larger droplets in their path. Laboratory studies show that collision does not always guarantee coalescence; sometimes the droplets actually bounce apart during collision. For example, the forces that hold a tiny droplet together (surface tension) are so strong that if a droplet were to collide with another tiny droplet, chances are the two would not stick together (coalesce) (see Fig. 5.20). Coalescence

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appears to be enhanced if colliding droplets have opposite (hence, attractive) electrical charges.* An important factor influencing cloud droplet growth by the collision process is the amount of time the droplet spends in the cloud. Since rising air currents slow the rate at which droplets fall, a thick cloud with strong updrafts will maximize the time cloud droplets spend in a cloud and, hence, the size to which they grow. Clouds that have above-freezing temperatures at all levels are called warm clouds. In tropical regions, where warm cumulus clouds build to great heights, strong convective updrafts frequently occur. In Fig. 5.21, suppose a cloud droplet is caught in a strong updraft. As the droplet rises, smaller droplets rise more quickly and collide with it, allowing the droplet to reach a size of about 1 mm. At this point, the updraft in the cloud is just able to balance the pull of gravity on the drop, so the drop remains suspended until it grows just a little bigger. Once the fall velocity of the drop is greater than the updraft velocity in the cloud, the drop slowly descends. As the drop falls, some of the smaller droplets get caught in the airstream around it, and are swept aside. Larger cloud droplets are captured by the *It was once thought that atmospheric electricity played a significant role in the production of rain. Today, evidence suggests that the difference in electrical charge that exists between cloud droplets results from the bouncing collisions between them. It is felt that the weak separation of charge and the weak electrical fields in developing relatively warm clouds are not significant in initiating precipitation. However, studies show that coalescence is often enhanced in thunderstorms where strongly charged droplets exist in a strong electrical field.

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FIGURE 5.20 Collision and coalescence. (a) In a warm cloud composed only of small cloud droplets of uniform size, the drop droplets are less likely to collide as they all fall very slowly at about the same speed. Those droplets that do collide, frequently do not coalesce because of the strong surface tension that holds together each tiny droplet. (b) In a cloud composed of differentsize droplets, larger droplets fall faster than smaller droplets. Although some tiny droplets are swept aside, some collect on the larger droplet’s forward edge, while others (captured in the wake of the larger droplet) coalesce on the droplet’s backside.

FIGURE 5.21 A cloud droplet rising, then falling through a warm cumulus cloud can grow by collision and coalescence and emerge from the cloud as a large raindrop.

falling drop, which then grows larger. By the time this drop reaches the bottom of the cloud, it will be a large raindrop with a diameter of over 5 mm. Because raindrops of this size fall faster and reach the ground first, they typically occur at the beginning of a rainshower originating in these warm, convective cumulus clouds. So far, we have examined the way cloud droplets in warm clouds (that is, those clouds with temperatures above freezing) grow large enough by the collision-coalescence process to fall as raindrops. The most important factor in the production of raindrops is the cloud’s liquid water content. In a cloud with sufficient water, other significant factors are: . the range of droplet sizes . the cloud thickness . the updrafts of the cloud . the electric charge of the droplets and the electric field in the cloud Relatively thin stratus clouds with slow, upward air currents are, at best, only able to produce drizzle (the lightest form of rain), whereas the towering cumulus clouds associated with rapidly rising air can cause heavy showers. Now, let’s turn our attention to how clouds with temperatures below freezing are able to produce precipitation. ICE-CRYSTAL PROCESS The ice-crystal (or Bergeron) process* of rain formation proposes that both ice crystals and liquid cloud droplets are present in clouds at *The ice-crystal process is also known as the Bergeron process after the Swedish meteorologist Tor Bergeron, who proposed that essentially all raindrops begin as ice crystals. CLOUD DEVELOPMENT DEVELOPMENT AND PRECIPITATION

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FIGURE 5.22 The distribution of ice and water in a typical cumulonimbus cloud.

temperatures below freezing. This process of rain formation is extremely important in middle and high latitudes, where clouds are able to extend upwards into regions where air temperatures are below freezing. Such clouds are called cold clouds. Figure 5.22 illustrates a typical cumulonimbus cloud that has formed over the Great Plains of North America. In the warm region of the cloud (below the freezing level) where only water droplets exist, we might expect to observe cloud droplets growing larger by the collision and coalescence process described in the previous section. Surprisingly, in the cold air just above the freezing level, almost all of the cloud droplets are still composed of liquid water. Water droplets existing at temperatures below freezing are referred to as supercooled droplets. At higher levels, ice crystals become more numerous, but are still outnumbered by water droplets. Ice crystals exist overwhelmingly in the upper part of the cloud, where air temperatures drop to well below freezing. Why are there so few ice crystals in the middle of the cloud, even though temperatures there, too, are below freezing? Laboratory studies reveal that the smaller the amount of pure water, the lower the temperature at which water freezes. Since cloud droplets are extremely small, it takes very low temperatures to turn them into ice. Just as liquid cloud droplets form on condensation nuclei, ice crystals can form in subfreezing air if there are ice-forming particles present called ice nuclei. The number of ice-forming nuclei available in the atmosphere is small, especially at temperatures above −10°C (14°F). Although some uncertainty exists regarding the principal source of ice nuclei, it is known that certain clay minerals are excellent ice nuclei, as are some types of bacteria in decaying plant 128

leaf material, along with ice crystals themselves and other particles whose geometry resembles that of an ice crystal. We can now understand why there are so few ice crystals in the subfreezing region of some clouds. Liquid cloud droplets can freeze, but only at very low temperatures. Ice nuclei can initiate the growth of ice crystals, but they do not abound in nature. Therefore, we are left with a cold cloud that contains many more liquid droplets than ice particles, even at low temperatures. Neither the tiny liquid nor solid particles are large enough to fall as precipitation. How, then, does the ice-crystal process produce rain and snow? In the subfreezing air of a cloud, many supercooled liquid droplets will surround each ice crystal. Suppose that the ice crystal and liquid droplet in Fig. 5.23 are part of a cold (−15°C), supercooled, saturated cloud. Since the air is saturated, both the liquid droplet and the ice crystal are in equilibrium, meaning that the number of molecules leaving the surface of both the droplet and the ice crystal must equal the number of molecules returning. Observe, however, that there are more vapor molecules above the liquid, because molecules escape the surface of water much more easily than they escape the surface of ice. Consequently, more molecules escape the water surface at a given temperature, requiring more in the vapor phase to maintain saturation. Therefore, it takes more vapor molecules to saturate the air directly above the water droplet than it does to saturate the air directly above the ice crystal. Put another way, at the same subfreezing temperature, the saturation vapor pressure just above the water surface is greater than the saturation vapor pressure above the ice surface.* *This concept is illustrated in the insert in Fig. 4.5, p. 83.

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CLOUD SEEDING AND PRECIPITATION The primary goal in many experiments concerning cloud seeding is to inject (or seed) a cloud with small particles that will act as nuclei, so that the cloud particles will grow large enough to fall to the surface as precipitation. The first ingredient in any seeding project is, of course, the presence of clouds, as seeding does not generate clouds. In clouds where the air temperature is below freezing, at least a portion of the cloud (preferably the upper part) must be supercooled, because in this situation cloud seeding uses the ice-crystal process to cause the cloud particles to grow. The idea is to find clouds that have too low a ratio of ice crystals to droplets and then to add enough artificial ice nuclei so that the ratio of crystals to droplets is optimal (about 1:100,000) for producing precipitation. Some of the first experiments in cloud seeding were conducted by Vincent Schaefer and Irving Langmuir during the late 1940s. To seed a cloud, they dropped crushed pellets of dry ice (solid carbon dioxide) from a plane. Because dry ice has a temperature of −78°C (−108°F), it acts as a cooling agent. As the extremely cold pellets of dry ice fall through the cloud, they quickly cool the air around them. This cooling causes the air around the pellet to become supersaturated. In this supersaturated air, water

This difference in saturation vapor pressure causes water vapor molecules to move (diffuse) from the droplet toward the ice crystal. The removal of vapor molecules reduces the vapor pressure above the droplet. Since the droplet is now out of equilibrium with its surroundings, it evaporates to replenish the diminished supply of water vapor above it. This process provides a continuous source of moisture for the ice crystal, which absorbs the water vapor and grows rapidly (see Fig. 5.24). Hence, during the ice-crystal crystal (Bergeron) process, ice crystals grow larger at the expense of the surrounding water droplets. The ice crystals may now grow even larger. For example, in some clouds, ice crystals might collide with supercooled liquid droplets. Upon contact, the liquid droplets freeze into ice and stick together. This process of ice crystals growing larger as they collide with supercooled cloud droplets is called accretion. The icy matter that forms is called graupel (or snow pellets). As the graupel falls, it may fracture or splinter into tiny ice particles when it collides with cloud droplets. These splinters may then go on themselves to become new graupel, which, in turn, may produce more splinters. In colder clouds, the delicate ice crystals may collide with other crystals and fracture into smaller ice particles, or tiny seeds, which freeze hundreds of supercooled droplets on contact. In both cases a chain reaction can develop, producing many col ice crystals (see Fig. 5.25). As they fall, they may collide and stick to one another, forming an aggregate of ice crystals called a snowflake. If the snowflake melts before reaching the ground, it continues its fall as a raindrop. Much of the rain falling in middle and northern latitudes—even in summer—actually begins as snowflakes.

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FIGURE 5.23 In a saturated environment, the water droplet and the ice crystal are in equilibrium, as the number of water molecules leaving the surface of each droplet and ice crystal equals the number returning. There are more water vapor molecules above the droplet than above the ice, which produces a greater vapor pressure above the droplet. At saturation, then, the pressure exerted by the water molecules is greater over the water droplet than above the ice crystal.

FIGURE 5.24 The ice-crystal (Bergeron) process. (1) The greater number of water vapor molecules around the liquid droplet causes water molecules to diffuse from the liquid droplet toward the ice crystal. (2) The ice crystal absorbs the water vapor and grows larger, while (3) the water droplet grows smaller. CLOUD DEVELOPMENT DEVELOPMENT AND PRECIPITATION

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FIGURE 5.25 Ice particles in clouds.

© C. Donald Ahrens

vapor directly forms into many tiny cloud droplets. In the very cold air created by the falling pellets (below −40°C), the tiny droplets instantly freeze into tiny ice crystals. The newly formed ice crystals then grow larger by deposition as the water vapor molecules attach themselves to the ice crystals at the expense of the nearby liquid droplets and, upon reaching a sufficiently large size, fall as precipitation. In 1947, Bernard Vonnegut demonstrated that silver iodide (AgI) could be used as a cloud-seeding agent. Because silver iodide has a crystalline structure similar to an ice crystal, it acts as an effective ice nucleus at

FIGURE 5.26 Ice crystals falling from a dense cirriform cloud into a lower nimbostratus cloud. This photo was taken at an al altitude near 6 km (19,700 ft) above western Pennsylvania. At the surface, moderate rain was falling over the region.

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temperatures of −4°C (25°F) and lower. Silver iodide causes ice crystals to form in two primary ways: . Ice crystals form when silver iodide crystals come in contact with supercooled liquid droplets. . Ice crystals grow in size as water vapor is deposited onto the silver iodide crystal. Silver iodide is much easier to handle than dry ice, since it can be supplied to the cloud from burners located either on the ground or on the wing of a small aircraft. Although other substances, such as lead iodide and cupric sulfide, are also effective ice nuclei, silver iodide still remains the most commonly used substance in cloudseeding projects. (Additional information on the controversial topic of cloud seeding’s effectiveness is given in Focus section 5.2.) Under certain conditions, clouds can be seeded naturally. For example, when cirriform clouds lie directly above a lower cloud deck, ice crystals may descend from the higher cloud and seed the cloud below (see Fig. 5.26). As the ice crystals mix into the lower cloud, supercooled droplets are converted to ice crystals, and the precipitation process is enhanced. Sometimes the ice crystals in the lower cloud may settle out, leaving a clear area or “hole” in the cloud (see Fig. 2 in Focus section 5.2). When the cirrus clouds form waves downwind from a mountain chain, bands of precipitation often form—producing heavy precipitation in some areas and practically no precipitation in others (see Fig. 5.27).

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ON AN ENVIRONMENTAL ISSUE 5.2

Just how effective is artificial seeding with silver iodide or other substances in increasing precipitation? In any given year, dozens of cloud-seeding projects are taking place around the world. However, the likelihood that cloud seeding will reliably increase precipitation is a much-debated question among meteorologists. First of all, it is difficult to evaluate the results of any cloudseeding experiment. When a seeded cloud produces precipitation, the question always remains as to how much precipitation would have fallen had the cloud not been seeded. Other factors must be considered when evaluating cloud-seeding experiments: the type of cloud, its temperature and moisture content, the dropletsize distribution, and the updraft velocities in the cloud. Some experiments have suggested that cloud seeding in some areas, under the right conditions, may enhance precipitation by anywhere from 5 percent to 20 percent or more. However, the results vary from project to project, and it can be difficult to confirm or disprove claims of success. And so the controversy continues. Some cumulus clouds show an “explosive” growth after being seeded. The latent heat given off when the droplets freeze acts to warm the cloud, causing it to become more buoyant. It grows rapidly and becomes a longer-lasting cloud, which may produce more precipitation. The business of cloud seeding can be a bit tricky, though, since overseeding can produce too many ice crystals. When this happens, the cloud becomes glaciated (all liquid droplets become ice) and

© Alan Sealls/Weatherthings

Does Cloud Seeding Enhance Precipitation?

FIGURE 2 When an aircraft flies through a layer of altocumulus clouds composed of supercooled droplets, a hole in the cloud layer may form. The cirrus-type cloud in the center is probably the result of inadvertent cloud seeding by the aircraft.

the ice particles, being very small, do not fall as precipitation. Since few liquid droplets exist, the ice crystals cannot grow by the ice-crystal (Bergeron) process; rather, they disappear by changing from ice into water vapor (sublimating) and leaving a clear area in a thin, stratified cloud (see Fig. 2). Because dry ice can produce the most ice crystals in a supercooled cloud, it is the substance most suitable for deliberate overseeding. Hence, it is the substance most commonly used to dissipate cold fog at airports (see Chapter 4, p. 97). Warm clouds with temperatures above freezing have also been seeded in an attempt to produce rain. Tiny water drops and

PRECIPITATION IN CLOUDS In cold, strongly convective clouds, precipitation may begin only minutes after the cloud forms, and can be initiated by either the collisioncoalescence or the ice-crystal (Bergeron) process. Once either process begins, most precipitation growth is by accretion, as supercooled liquid droplets freeze on impact with snowflakes and ice crystals. Although precipitation is commonly absent in warm layered clouds, such as stratus, it is often associated with such cold-layered clouds as

particles of hygroscopic salt are injected into the base (or top) of the cloud. These particles (called seed drops ), when carried into the cloud by updrafts, create large cloud droplets, which grow even larger by the collision-coalescence process. Apparently, the seed-drop size plays a major role in determining the effectiveness of seeding with hygroscopic particles. To date, however, the results obtained using this method are inconclusive. In summary, cloud seeding in certain instances may lead to more precipitation; in others, to less precipitation; and, in still others, to no change in precipitation amounts. Many of the questions about cloud seeding have yet to be resolved.

nimbostratus and altostratus. This precipitation is thought to form principally by the ice-crystal (Bergeron) process because the liquid water content of these clouds is generally lower than that in convective clouds, thus making the collision-coalescence process much less effective. Nimbostratus clouds are normally thick enough to extend to levels where air temperatures are quite low, and they usually last long enough for the ice-crystal process to initiate precipitation. CLOUD DEVELOPMENT DEVELOPMENT AND PRECIPITATION

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and transformed into other forms of precipitation that can profoundly influence our environment.

FIGURE 5.27 Natural seeding by cirrus clouds may form bands of precipitation downwind of a mountain chain. Notice that heavy snow is falling only in the seeded areas.

BRIEF REVIEW In the last few sections we encountered a number of important concepts and ideas about how cloud droplets can grow large enough to fall as precipitation. Before examining the various types of precipitation, here is a summary of some of the important ideas presented so far: ●

Cloud droplets are very small, much too small to fall as rain.

Cloud droplets form on cloud condensation nuclei. Hygroscopic nuclei, such as salt, allow condensation to begin when the relative humidity is less than 100 percent.

Cloud droplets, in above-freezing air, can grow larger as fasterfalling, bigger droplets collide and coalesce with smaller droplets in their path.

In the ice-crystal (Bergeron) process of rain formation, both ice crystals and liquid cloud droplets must coexist at below-freezing temperatures. The difference in saturation vapor pressure between liquid droplets and ice crystals causes water vapor to diffuse from the liquid droplets (which shrink) toward the ice crystals (which grow).

Most of the rain that falls over middle latitudes results from melted snow that formed from the ice-crystal (Bergeron) process.

Cloud seeding with silver iodide can only be effective in coaxing precipitation from a cloud if the cloud is supercooled and the correct ratio of cloud droplets to ice crystals exists.

Precipitation Types Up to now, we have seen how cloud droplets can grow large enough to fall to the ground as rain or snow. While falling, raindrops and snowflakes can be altered by atmospheric conditions encountered beneath the cloud 132

RAIN Most people consider rain to be any falling drop of liquid water. To the meteorologist, however, that falling drop must have a diameter equal to, or greater than, 0.5 mm (0.02 in.) to be considered rain. Fine, uniform drops of water with diameters smaller than 0.5 mm (about one-half the width of the letter “o” in the print version of this page) are called drizzle. Most drizzle falls from stratus clouds; however, small raindrops may fall through air that is unsaturated, partially evaporate, and reach the ground as drizzle. Surfaces can be moistened by fog or mist, especially in windy conditions, even though the droplets in fog and mist are too tiny to fall to the ground as precipitation. Occasionally, the rain falling from a cloud never reaches the surface because the low humidity causes rapid evaporation. As the drops become smaller, their rate of fall decreases, and they appear to hang in the air as a rain streamer. These evaporating streaks of precipitation are called virga** (see Fig. 5.28). Raindrops may also fall from a cloud and not reach the ground if they encounter the rapidly rising air of an updraft. If the updraft weakens or changes direction and becomes a downdraft, the suspended drops will fall to the ground as a sudden rainshower. The showers falling from cumuliform clouds, such as cumulonimbus or cumulus congestus, are usually brief and sporadic as the cloud moves overhead and then drifts on by. If the shower is excessively heavy, it may be informally called a cloudburst. Beneath a cumulonimbus cloud, which normally contains strong, deep convection currents of rising and descending air, it is entirely possible for one side of a street to be dry (updraft side), while a heavy shower is occurring across the street (downdraft side) (see Fig. 5.29). Continuous rain, on the other hand, usually falls from a layered cloud that covers a large area and has weaker, more shallow vertical air currents. These are the conditions normally associated with nimbostratus clouds. Raindrops that reach Earth’s surface are seldom much larger than about 5 mm (0.2 in.), the reason being that the collisions (whether glancing or head-on) between raindrops tend to break them up into many smaller drops. Additionally, when raindrops grow too large they become unstable and break apart. What is the shape of the falling raindrop? Is it tear-shaped or is it round? You may be surprised at the answer, which is given in Focus section 5.3. After a rainstorm, visibility usually improves, primarily because precipitation removes (scavenges) many of the suspended particles. When rain combines with *Studies suggest that the “rain streamer” is actually caused by ice (which is more reflective) changing to water (which is less reflective). Apparently, most evaporation occurs below the virga line.

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© Ross DePaola

FIGURE 5.28 The streaks of falling precipitation that evapo evaporate before reaching the ground are called virga.

gaseous pollutants, such as oxides of sulfur and nitrogen, it becomes acidic. Acid rain, which has an adverse effect on plants and water resources, has become a major problem in many industrialized regions of the world over the past few decades. We will examine the acid rain problem more thoroughly in Chapter 14, which emphasizes air pollution.

spot the melting level when you look in the direction of the sun, if it is near the horizon. Because snow scatters incoming sunlight better than rain, the darker region beneath the cloud contains falling snow, while the lighter region is falling rain. The melting zone, then, is the transition between the light and dark areas (see Fig. 5.30). When the warmer air beneath the cloud is relatively dry, the snowflakes partially melt. As the liquid water evaporates, it chills the snowflake, which retards its rate of melting. Consequently, in air that is relatively dry, snowflakes can reach the ground even when the air temperature is considerably above freezing, even above 40°F. Is it ever “too cold to snow”? Although many believe this expression, the fact remains that it is never too cold to snow. For one thing, snow may actually fall from cold

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SNOW We know that much of the precipitation reaching the ground actually begins as snow. In summer, the freezing level is usually high and the snowflakes falling from a cloud melt before reaching the surface. In winter, however, the freezing level is much lower, and falling snowflakes have a better chance of survival. In fact, snowflakes can generally fall about 300 m (or 1000 ft) below the freezing level before completely melting. Occasionally, you can

FIGURE 5.29 Strong updrafts and downdrafts of a cumulonimbus cloud can cause rain to fall on one side of a street but not on the other.

FIGURE 5.30 Snow scatters sunlight more effectively than rain. Consequently, when you look toward the sun, the region of falling precipitation looks darker above the melting level than below it. CLOUD DEVELOPMENT DEVELOPMENT AND PRECIPITATION

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FOCUS

ON A SPE SPECIAL TOPIC 5.3

As rain falls, the drops take on a characteristic shape. Choose the shape in Fig. 3 that you think most accurately describes that of a falling raindrop. Did you pick number 1? The tear-shaped drop has been depicted by artists for many years. Unfortunately, raindrops are not tear-shaped. tear-shaped Actually, the shape depends on the drop size. Raindrops less than 2 mm (0.08 in.) in diameter are nearly spherical and look like raindrop number 2. The attraction among the molecules of the liquid (surface tension) tends to squeeze the drop into a shape that has the smallest surface area for its total volume—a sphere.

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Are Raindrops Tear-Shaped?

FIGURE 3 Which of the three drops shown here represents the real shape of a falling raindrop?

Large raindrops, with diameters exceeding 2 mm, take on a different shape as they fall. Believe it or not, they

air into extremely cold surface air. True, colder air cannot “hold” as much water vapor as warmer air, but no matter how cold the air becomes, it always contains some water vapor that could produce snow. In fact, tiny ice crystals have been observed falling at temperatures as low as −47°C (−53°F). We usually associate extremely cold air with “no snow” because the coldest winter weather occurs on clear, calm nights—conditions that normally prevail with strong high-pressure areas that have few if any clouds. When ice crystals and snowflakes fall from high cirrus clouds they are called fallstreaks. Fallstreaks behave in much the same way as virga: As the ice particles fall into drier air, they usually disappear as they change from ice

look like number 3, slightly elongated, flattened on the bottom, and rounded on top. As the larger drop falls, the air pressure against the drop is greatest on the bottom and least on the sides. The pressure of the air on the bottom flattens the drop, while the lower pressure on its sides allows it to expand a little. This mushroom shape has been described as resembling everything from a falling parachute to a loaf of bread, or even a hamburger bun. You may call it what you wish, but remember: It is not tear-shaped.

into vapor (called sublimation). Because the wind at higher levels moves the cloud and ice particles horizontally more quickly than do the slower winds at lower levels, fallstreaks often appear as dangling white streamers (see Fig. 5.31). Moreover, fallstreaks descending into lower, supercooled clouds may actually seed them. Snowflakes falling through moist air that is slightly above freezing slowly melt as they descend. A thin film of water forms on the edge of the flake, acting like glue when other snowflakes come in contact with it. In this way, several flakes can join to produce giant snowflakes that often measure an inch or more in diameter. These large, soggy snowflakes are associated with moist air

© C. Donald Ahrens

FIGURE 5.31 The dangling white streamers of ice crystals beneath these cirrus clouds are known as fallstreaks. The bending of the streaks is due to the changing wind speed with height.

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© Robert Robinson/iStockphoto.com

© Scott Cunazine/Photo Researchers, Inc.

FIGURE 5.32 Computer color-enhanced image of dendrite snowflakes.

and temperatures near freezing. However, when snowflakes fall through extremely cold air with a low moisture content, they do not readily stick together and small, powdery flakes of “dry” snow accumulate on the ground. If you catch falling snowflakes on a dark object and examine them closely, you will see that the most common snowflake form is a fernlike branching shape called dendrite (see Fig. 5.32). As ice crystals fall through a cloud, they are constantly exposed to changing temtem peratures and moisture conditions. Since many ice crystals can join together (aggregate) to form a much larger snowflake, ice crystals can assume many complex patterns. Snow falling from developing cumulus clouds is often in the form of flurries. These are usually light showers that fall intermittently for short durations and produce only light accumulations. A more intense snow shower is called a snow squall. These brief but heavy falls of snow are comparable to summer rainshowers and, like snow flurries, usually fall from cumuliform clouds. If the snow falls from intense cumuliform clouds producing thunder and lightning, the snow is often referred to as thundersnow. A more continuous snowfall (sometimes persisting for several ▼ Table 5.1

Snowfall Intensity

SNOWFALL DESCRIPTION

VISIBILITY

Light

Greater than ½ mile*

Moderate

Greater than ¼ mile, less than or equal to ½ mile

Heavy

Less than or equal to ¼ mile

*In the United States, the National Weather Service determines visibility (the greatest distance you can see) in miles.

FIGURE 5.33 High winds, blowing and falling snow, along with low temperatures, produced this blizzard over the Great Plains.

hours) accompanies nimbostratus and altostratus clouds. The intensity of snow is based on its reduction of horizontal visibility at the time of observation (see ▼Table 5.1). However, the intensity of snow as measured by visibility does not tell us how much water the snow is bringing to the surface. A moderate snow made up of dense, small flakes can deposit more frozen liquid than a heavy snow consisting of large, fluffy flakes. When a strong wind is blowing at the surface, snow can be picked up and deposited into huge drifts. Drifting snow is usually accompanied by blowing snow; that is, snow lifted from the surface by the wind and blown about in such quantities that horizontal visibility is greatly restricted. The combination of drifting and blowing snow, after falling snow has ended, is called a ground blizzard. A true blizzard is a weather condition characterized by low temperatures and strong winds (greater than 30 knots) bearing large amounts of fine, dry, powdery particles of snow that reduces visibility to less than one-quarter mile (and sometimes to as little as a few feet) for at least three hours (see Fig. 5.33). Figure 5.34 shows the annual average snowsnow fall across the United States and Canada. As you would expect, annual snowfall totals tend to be low in the southern United States and higher as you move north. Notice that in areas of the northeast United States, eastern

DID YOU KNOW? On a sunny day, when the atmosphere is conditionally unstable, thermals break away from the surface and rapidly rise into the atmosphere carrying surface air with them. If the thermals form over a field of alfalfa, a feed lot, or a garbage dump, the smell of the surface air can be carried upward for thousands of feet. CLOUD DEVELOPMENT DEVELOPMENT AND PRECIPITATION

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Brian Brettschneider

FIGURE 5.34 Average annual snowfall over the United States and Southern Canada.

Canada, and the mountainous west, annual snowfall totals exceed 183 cm (72 in.). In fact, Paradise Ranger Station on Mount Rainier, Washington, receives an annual average of 1758 cm (692 in.) of snow, making it one of the snowiest places in the world. Extremely heavy snowfalls also occur downwind from large bodies of water, such as the Great Lakes of North America. These so-called lake effect snows will be covered more completely in Chapter 8 (p. 215).

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SLEET AND FREEZING RAIN Consider the falling snowflake in Fig. 5.35. As it falls into warmer air, it begins to melt. When it falls through the deep subfreezing surface

FIGURE 5.35 Sleet forms when a partially melted snowflake or a cold raindrop freezes into a pellet of ice before reaching the ground.

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layer of air, the partially melted snowflake or cold raindrop turns back into ice, not as a snowflake, but as a tiny transparent (or translucent) ice pellet called sleet.* Generally, these ice pellets bounce when striking the ground and produce a tapping sound when they hit a window or piece of metal. The cold surface layer beneath a cloud may be too shallow to freeze raindrops as they fall. In this case, they reach the surface as supercooled liquid drops. Upon striking a cold object, the drops spread out and almost immediately freeze, forming a thin veneer of ice. This form of precipitation is called freezing rain, or glaze. If the drops are quite small, the precipitation is called freezing drizzle. When small, supercooled cloud or fog droplets strike an object whose temperature is below freezing, the tiny droplets freeze, forming an accumulation of white or milky granular ice called rime (see Fig. 5.36). Occasionally, light rain, drizzle, or supercooled fog droplets come into contact with such surfaces as bridges and overpasses that have cooled to a temperature below freezing. The tiny liquid droplets freeze on contact to road surfaces or pavements, producing a sheet of ice that often appears relatively dark. Such ice, usually called black ice, can produce extremely hazardous driving conditions. Freezing rain can create a beautiful winter wonderland by coating everything with silvery, glistening ice. At the same time, highways turn into skating rinks for automobiles, and the destructive weight of the ice—which *Occasionally, news media in the United States will use the term “sleet” to describe a mixture of rain and snow. While this is not the American definition, it is correct in some other nations, including the United Kingdom.

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FOCUS

ON AN O OBSERVATION 5.4

FIGURE 4 An aircraft undergoing de-icing during inclement winter weather.

(see Fig. 5.35). Even though rime ice redistributes the flow of air over the wing more than clear ice does, it is lighter in weight and is more easily removed with de-icers. Because the raindrops and cloud droplets in most clouds vary in size, a mixture of clear and rime ice usually forms on aircraft. Also, because concentrations of liquid water tend to be greatest in warm air, icing is usually heaviest and most severe when the air temperature is between 0°C and −10°C (32°F and 14°F). A major hazard to aviation, icing reduces aircraft efficiency by increasing weight. Icing has other adverse effects,

© UCAR

The formation of ice on an aircraft—called aircraft icing—can icing be extremely dangerous, sometimes leading to tragic acciacci dents. Icing is believed to be the probable cause for the crash-landing of a passenger plane as it approached the Detroit airport on January 9, 1997. All 29 people aboard the flight were killed. Fortunately, the number of icing-related aircraft fatalities in the United States has decreased subsub stantially in recent years with increased safety measures. Just how does aircraft icing develop? Consider an aircraft flying through an area of freezing rain or through a region of large supercooled droplets in a cumuliform cloud. As the large, supercooled drops strike the leading edge of the wing, they break apart and form a film of water, which quickly freezes into a solid sheet of ice. This smooth, transparent ice—called clear ice— is similar to the freezing rain or glaze that coats trees during ice storms. Clear ice can build up quickly; it is heavy and difficult to remove, even with modern de-icers. When an aircraft flies through a cloud composed of tiny, supercooled liquid dropdrop lets, rime ice may form. Rime ice forms when some of the cloud droplets strike the wing and freeze before they have time to spread, thus leaving a rough and brittle coating of ice on the wing. Because the small, frozen droplets trap air between them, rime ice usually appears white

© Annebique Bernard/Sygma/Corbis

Aircraft Icing

FIGURE 5.36 An accumulation of rime forms on tree branches as supercooled fog droplets freeze on contact in the belowfreezing air.

depending on where it forms. On a wing or fuselage, ice can disrupt the airflow and decrease the plane’s flying capability. When ice forms in the air intake of the engine, it robs the engine of air, causing a reduction in power. Icing may also affect the operation of brakes, landing gear, and instruments. Because of the hazards of ice on an aircraft, its wings are usually sprayed with a type of antifreeze before taking off during cold, inclement weather (see Fig. 4). In a snowstorm, the amount of liquid in the snow is the primary factor in determining how much de-icing fluid is needed for aircraft.

can be many tons on a single tree—breaks tree branches, power lines, and telephone cables. When there is a substantial accumulation of freezing rain, the result is an ice storm (see Fig. 5.37). A case in point is the catastrophic ice storm that struck an area from Oklahoma to West Virginia in January 2009. More than 2 million people lost power, and 65 people died. Many of the deaths were attributed to the use of emergency heaters without adequate ventilation, which led to carbon monoxide poisoning. The area most frequently hit by these storms extends over a broad region from Texas into Minnesota and eastward into the middle Atlantic states and New England. Such storms are extremely rare in most of California and Florida. (For information on freezing rain and its effect on aircraft, read Focus section 5.4.) CLOUD DEVELOPMENT DEVELOPMENT AND PRECIPITATION

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Syracuse Newspapers/Dick Blume/The Image Works

FIGURE 5.37 A heavy coating of freezing rain (glaze) covers Syracuse, New York, causing tree limbs to break and power lines to sag. This massive ice storm in January 1998 left millions without power and caused over $1 billion in damage in northern New England and southeast Canada.

HAIL Hailstones are pieces of ice, either transparent or partially opaque, ranging in size from that of small peas to that of golf balls or larger (see Fig. 5.40). Some are round; others take on irregular shapes. In the United States, the hailstone with the greatest measured circumference (18.7 inches) fell on Aurora, Nebraska, on June 22, 2003. This giant hailstone, almost as large as a soccer ball, had a measured diameter of 7 inches and probably weighed over 1.75 pounds. The hailstone in the United States with the

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In summary, Fig. 5.38 shows various winter tempertemper ature profiles and the type of precipitation associated with each. In profile (a), the air temperature is below freezing at all levels, and snowflakes reach the surface. In (b), a zone of above-freezing air causes snowflakes to partially melt; then, in the deep, subfreezing air at the surface, the liquid freezes into sleet. In the shallow sub-freezing surface air in (c), the melted snowflakes, now supercooled liquid drops, freeze on contact, producing freezing rain. In (d), the air temperature is above freezing in a sufficiently deep layer so that precipitation reaches the surface as rain.

SNOW GRAINS AND SNOW PELLETS Snow grains are small, opaque grains of ice, the solid equivalent of drizzle. They fall in small quantities from stratus clouds, and never in the form of a shower. Upon striking a hard surface, they neither bounce nor shatter. Snow pellets, on the other hand, are white, opaque grains of ice about the size of an average raindrop. They are sometimes confused with snow grains. The distinction is easily made, however, by remembering that, unlike snow grains, snow pellets are brittle, crunchy, and bounce (or break apart) upon hitting a hard surface. They usually fall as showers, especially from cumulus congestus clouds. Snow pellets form as ice crystals collide with supercooled water droplets that freeze into a spherical aggregate of icy matter (rime) containing many air spaces. When the ice particle accumulates a heavy coating of rime, it is called graupel. During the winter, when the freezing level is at a low elevation, the graupel reaches the surface as a snow pellet, a light, round clump of snowlike ice (see Fig. 5.39). On the surface, the accumulation of snow pellets sometimes gives the appearance of tapioca pudding; hence, it can be referred to as tapioca snow. In a thunderstorm, when the freezing level is well above the surface, graupel that reaches the ground is sometimes called soft hail. During summer, graupel may melt and reach the surface as large raindrops. In vigorously convective clouds, however, graupel may develop into full-fledged hailstones.

FIGURE 5.38 Vertical temperature profiles (solid red line) associated with different forms of precipitation.

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NOAA National Weather Service

FIGURE 5.40 Hailstones of varying sizes fell over west Texas during an intense thunderstorm.

FIGURE 5.41 This whopping hailstone fell on Vivian, South Dakota, on July 23, 2010. It had a record diameter of 8 inches, weighed a record 1.94 pounds, and had a circumference of 18.6 inches.

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largest diameter ever measured (8 inches) and the greatest weight (1.94 lb.) fell on Vivian, South Dakota, on July 23, 2010 (see Fig. 5.41). The world’s heaviest confirmed hailstone, at 2.25 pounds, fell in a Bangladesh thunderthunder storm on April 14, 1986. Needless to say, large hailstones are quite destructive. They can break windows, dent cars, batter roofs of homes. Even small hail can injure livestock and cause extensive damage to crops if it falls heavily, especially in a high wind. A single hailstorm can destroy a farmer’s crop in a matter of minutes, which is why farmers sometimes call it “the white plague.” Estimates are that, in the United States alone, hail damage amounts to hundreds of millions of dollars annually. Although hailstones are potentially lethal, only a few fatalities from falling hail have been documented in the United States since 1900. Deaths due to hail are somewhat more frequent in parts of Asia, where 92 people were killed in the Bangladesh thunderstorm of April 1986. Hail forms in a cumulonimbus cloud—usually an intense thunderstorm—when graupel, or large frozen raindrops, or just about any particles (even insects) act as embryos that grow by accretion, the accumulation of supercooled liquid droplets. For a hailstone to grow to golf ball size, it must remain in the cloud between five and golften minutes. Violent, upsurging air currents within the storm carry small ice particles high above the freezing level where the ice particles grow by colliding with supercooled liquid cloud droplets. Violent rotating updrafts in severe thunderstorms (especially the long-lived storms called supercells) are even capable of sweeping the growing ice particles laterally through the cloud. In fact, it appears that the best trajectory for hailstone growth is one that is nearly horizontal through the storm (see Fig. 5.42). As growing ice particles pass through regions of varyvary ing liquid water content, a coating of ice forms around them, causing them to grow larger and larger. In a strong updraft, the larger hailstones ascend very slowly, and may appear to “float” in the updraft, where they continue to grow rapidly by colliding with numerous supercooled liquid droplets. When winds aloft carry the large hailstones away from the updraft or when the hailstones reach appreciable size, they become too heavy to be supported by the rising air, and they begin to fall.

© 2008-Warren Faidley

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FIGURE 5.39 A snowflake becoming a rimed snowflake, then finally graupel (a snow pellet).

FIGURE 5.42 Hailstones begin as embryos (usually ice particles called graupel graupel) that remain suspended in the cloud by violent updrafts. The updrafts (or a single broad updraft as illustrated here) can sweep the ice particles horizontally through the cloud, producing the optimal trajectory for hailstone growth. Along their path, the ice particles collide with supercooled liquid droplets, which freeze on contact. The ice particles eventually grow large enough and heavy enough to fall toward the ground as hailstones. CLOUD DEVELOPMENT DEVELOPMENT AND PRECIPITATION

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FIGURE 5.43 A large hailstone cut and then photographed under regular light. The layered structure of the hailstone reveals that it traveled through a cloud of varying water content and temperature.

In the warmer air below the cloud, the hailstones begin to melt. Small hail often completely melts before reaching the ground, but in the violent thunderstorms of late spring and summer, hailstones often grow large enough to reach the surface before completely melting. Strangely, then, we find the largest form of frozen precipitation occurring during the warmest time of the year. Figure 5.43 shows a cut section of a very large hailhail stone. Notice that it has distinct concentric layers of milky white and clear ice. We know that a hailstone grows by accumulating supercooled water droplets. If the growing hailstone enters a region inside the storm where the liquid-water content is relatively low (called the dry growth regime), supercooled droplets will freeze immediately on the stone, producing a coating of white or opaque rime ice containing many air bubbles. Should the hailstone get swept into a region of the storm where the liquid-water content is higher (called the wet growth regime), supercooled water droplets will collect so rapidly on the stone that, due to the release of latent heat, the stone’s surface temperature remains at freezing, even though the surrounding air may be much colder. Now the supercooled droplets no longer freeze on impact; instead, they spread a coating of water around the hailstone, filling in the porous regions, which leaves a layer of clear ice around the stone. Therefore, as a hailstone passes through a thunderstorm of changing liquid-water content (the dry and wet growth regimes), alternating layers of opaque and clear ice form, as illustrated in Fig. 5.43. As the cumulonimbus cloud moves along, it may deposit its hail in a long, narrow band known as a hailstreak. If the cloud should remain almost stationary for a period of time, substantial accumulation of hail is possible. A rare British hailstorm in the town of Ottery St. Mary dropped 9 to 10 inches of hail in two hours in October 2008, and drifts of up to 6 feet of hail occurred 140

during an intense thunderstorm in Cheyenne, Wyoming, in August 1985. During November 2003, an unusual hailstorm dumped more than 12.5 cm (5 in.) of hail over sections of Los Angeles, California, causing gutters to clog and floods to occur. In addition to its destructive effect, accumulation of hail on a roadway is a hazard to traffic as when, for example, four people lost their lives near Soda Springs, California, in a 15-vehicle pileup on a hailcovered freeway in September 1989. Because hailstones are so damaging, various methods have been tried to prevent them from forming in thunderstorms. One method employs the seeding of clouds with large quantities of silver iodide. These nuclei freeze supercooled water droplets and convert them into ice crystals. The ice crystals grow larger as they come in contact with additional supercooled cloud droplets. In time, the ice crystals grow large enough to be called graupel, which then becomes a hailstone embryo. Large numbers of embryos are produced by seeding in hopes that competition for the remaining supercooled droplets will be so great that none of the embryos can grow into large and destructive hailstones. In the United States, the results of most hail-suppression experiments are still inconclusive, although such cloud seeding has been carried out in many areas for a number of years.

Measuring Precipitation INSTRUMENTS Any instrument that can collect and measure rainfall is called a rain gauge. A standard rain gauge consists of a funnel-shaped collector attached cross to a long measuring tube (see Fig. 5.44). The crosssectional area of the collector is 10 times that of the tube. Hence, rain falling into the collector is amplified tenfold in the tube, permitting measurements as precise as onehundredth (0.01) of an inch. An amount of rainfall less than one-hundredth of an inch is called a trace. Another instrument that measures rainfall is the tipping bucket rain gauge. In Fig. 5.45, notice that this gauge has a receiving funnel leading to two small metal collectors (buckets). The bucket beneath the funnel collects the rainwater. When it accumulates the equivalent of one-hundredth of an inch of rain, the weight of the water causes it to tip and empty itself. The second bucket immediately moves under the funnel to catch the water. When it fills, it also tips and empties itself, while the original bucket moves back beneath the funnel. Each time a bucket tips, an electric contact is made, causing a pen to register a mark on a remote recording chart. Adding up the total number of marks gives the rainfall for a certain time period. A problem with the tipping bucket rain gauge is that during each “tip” it loses some rainfall and, therefore, under-measures rainfall amounts, especially during

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DID YOU KNOW?

FIGURE 5.44 Components of the standard rain gauge.

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heavy downpours. The tipping bucket is the rain gauge used in the automated (ASOS) weather stations. Remote recording of precipitation can also be made with a weighingweighing type rain gauge. With this gauge, precipitation is caught in a cylinder and accumulates in a bucket. The bucket sits on a sensitive weighing platform. Special gears translate the accumulated weight of rain or snow into millimeters or inches of precipitation. The precipitation totals are recorded by a pen on chart paper, which covers a clock-driven drum. By using special electronic equipment, this information can be transmitted from various weighing types of rain gauges in remote areas to satellites or land-based stations, thus providing precipitation totals from previously inaccessible regions. Snow is challenging to measure, since accumulations can vary greatly from one spot to the next, especially when winds are strong. Traditionally, the depth of snow in a region is determined by physically measuring its depth at three or more representative areas. The amount of snowfall is defined as the average of these measurements. Snow accumulation rates can be measured at a single site by using a snow measuring board, a small wooden platform resting near the ground that is cleared off after each measurement (normally every six hours). Snow depth may also be measured by removing the collector and inner cylinder of a standard rain gauge and allowing snow to accumulate in the outer tube. Automated gauges are now used in many areas to collect snowfall and measure the amount of liquid water it contains. Typically, these gauges are surrounded by one or more octagonal fences that help to block wind and produce more accurate readings. Remote sensing techniques are becoming a popular way to measure snow depth, especially where harsh winter conditions make it difficult to reach observing stations. Laser beams and pulses of ultrasonic energy can be sent from a transmitter to a snowpack where the energy bounces off the snowpack back to the transmitter. This procedure allows the height of the snow accumulation to be measured in much the same way that radar measures the distance falling rain is from a transmitter. (The topic of radar and precipitation is discussed in the following section.) Snow height can also be obtained by measuring how long it takes

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At midday on January 28, 2014, snow began falling in the Atlanta metropolitan area. Although the surface air temperature was well below freezing, the snow melted upon striking the relatively warm roads, then quickly froze into sheets of ice just as schools and businesses were closing early. Thousands of people ended up stuck in gridlock for many hours, in some cases overnight. More than 2000 students were forced to spend the night at school, and at least three babies were born in cars that were stuck on area roads. All of this took place in a snowstorm that produced less than three inches.

FIGURE 5.45 The tipping bucket rain gauge. Each time the bucket fills with one-hundredth of an inch of rain, it tips, sending an electric signal to the remote recorder.

a GPS signal to travel from a satellite to a snow field and then be reflected upward to a receiver located in the snow field. On average, about 10 inches of snow will melt down to about 1 inch of water, giving a typical fresh snowpack a water equivalent* of 10:1. This ratio, however, will vary greatly, depending on the texture and packing of the snow. Knowing the water equivalent of snow can provide valuable information about spring runoff and the potential for flooding, especially in mountain areas. DOPPLER RADAR AND PRECIPITATION Radar (radio detection and ranging) is an essential tool of the atmospheric scientist, because it gathers information about storms and precipitation in otherwise inaccessible regions. Atmospheric scientists use radar to examine the inside of *Water Water equivalent is the depth of water that would result from the melting of a snow sample. CLOUD DEVELOPMENT DEVELOPMENT AND PRECIPITATION

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FIGURE 5.46 A microwave pulse is sent out from the radar transmitter. The pulse strikes raindrops and a fraction of its en energy is reflected back to the radar unit, where it is detected and displayed, as shown in Fig. 5.47.

*The Doppler shift (or effect) is the change in the frequency of waves that occurs when the observer or the target is moving toward or away from the other. As an example, suppose a high-speed train is approaching you. The higherpitched (higher frequency) whistle you hear as the train approaches will shift to a lower pitch (lower frequency) after the train passes.

NOAA/NWS

a cloud in much the same way that physicians use X rays to examine the inside of a human body. Essentially, the radar unit consists of a transmitter that sends out short microwave pulses. When this energy encounters a foreign object—called a target—a fraction of the energy is scattered back toward the transmitter and is detected by a receiver (see Fig. 5.46). The returning signal is amplified and displayed on a screen, producing an image or “echo” from the target. The elapsed time between transmission and reception indicates the target’s distance. The brightness of the echo is directly related to the amount (intensity) of rain, snow, or both falling through the cloud. So, the radar screen shows not only where precipitation is most likely occurring, but also how intense it is. Typically the radar image is displayed using various colors,

usually ranging from green or blue to dark red, to denote the intensity of precipitation within the range of the radar unit. During the 1990s, Doppler radar replaced the conventional radar units that were put into service shortly after World War II. Doppler radar is like conventional radar in that it can detect areas of precipitation and measure rainfall intensity (see Fig. 5.47a). Using special computer programs called algorithms, the rainfall intensity, over a given area for a given time, can be computed and displayed as an estimate of total rainfall over that particular area (see Fig. 5.47b). But the Doppler radar can do more than conventional radar. Because the Doppler radar uses the principle called Doppler shift,* it has the capacity to measure the speed at which falling rain is moving horizontally toward or away from the radar antenna. Falling rain moves with the wind. Consequently, Doppler radar allows scientists to peer into a tornado-generating thunderstorm and observe its wind. We will investigate these ideas further in Chapter 10, when we consider the formation of severe thunderstorms and tornadoes. In some instances, radar displays indicate precipitation where there is none reaching the surface. This situation happens because the radar beam travels in a nearly straight line while Earth’s surface curves away from it. Hence, the return echo is not necessarily that of precipitation reaching the ground, but is that of raindrops in the cloud. So, if Doppler radar indicates that it’s raining in your area, and outside you observe that it is not, remember that it is most likely raining, but the raindrops are probably evaporating before reaching the ground.

FIGURE 5.47 (a) Doppler radar display showing precipitation intensity over North Carolina for August 27, 2011, as Hurricane Irene moves onshore. (b) Doppler radar display showing one-hour rainfall estimates over North Carolina for August 27, 2011. Notice th that in some places Doppler radar estimated that more than 1.50 inches of rain had fallen in one hour.

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The most recent improvement for Doppler radar is polarimetric radar, also referred to as dual-polarization -polarization radar. This form of Doppler radar, which was installed around the United States in the early 2010s, transmits both a vertical and horizontal pulse that makes it easier to determine whether falling precipitation is in the form dual of rain or snow. Fig. 5.48 shows an example of a dualpolarization radar display. Another technology being explored is phased-array radar, which can gather a much greater amount of data using a grid of small transmitters instead of a single large one. As we discussed in Chapter 4 (p. 110), specialized satellites can also be used to observe precipitation from space. Up through 2015, the Tropical Rainfall Measuring Mission (TRMM) satellite gathered a massive amount of three-dimensional data on precipitation. The newer Global

SSAI/NASA/JAXA, Hal Pierce

NOAA/NWS

FIGURE 5.48 Doppler radar with dual-polarization technology shows precipitation and cloud particles inside a thunderstorm near Wichita, Kansas, on May 30, 2012. This technology was added to the NWS radars in order to more clearly identify different types of precipitation and to more precisely identify tornadic circulations.

FIGURE 5.49 A three-dimensional image of Hurricane Jimena over the Central Pacific Ocean, collected on September 1, 2015, by the Global Precipitation Mission satellite. Reds and oranges indicate the heaviest rainfall at the surface. The eye of Jimena is visible as the hollow area in the middle of intense showers and thunderstorms.

Precipitation Mission (GPM) satellite has taken the place of TRMM, gathering precipitation data over a much larger area. GPM is able to gather more detail on lighter precipitation than TRMM, and it includes a dual-polarization radar that can deduce the sizes of raindrops, hailstones, and other precipitation elements (see Fig. 5.49). Another satellite that gathers three-dimensional dimensional data on precipitaprecipita tion intensity, NASA’s CloudSat satellite, was launched in 2006. An image from CloudSat is presented in Fig. 4.52 on p. 110.

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SUMMARY In this chapter, we tied together the concepts of stability, cloud formation, and precipitation. We learned that because stable air tends to resist upward vertical motions, clouds forming in a stable atmosphere often spread horizontally and have a stratified appearance. A stable atmosphere can be caused either by the surface air being cooled or by the air aloft being warmed. An unstable atmosphere tends to favor vertical air currents and produce cumuliform clouds. Instability can be produced either by the surface air being warmed or by the air aloft being cooled. In a conditionally unstable atmosphere, rising unsaturated air can be lifted to a level where condensation begins, latent heat is released, and instability results. We learned that the development of most clouds results from either surface heating, uplift along topographic barriers (orographic uplift), convergence of surface air, or lifting along weather fronts. We looked at cloud droplets and found that, individually, they are too small and light to reach the ground as rain. They can grow in size as large cloud droplets falling through a cloud collide and merge with smaller droplets in their path. In clouds where the air temperature is below freezing, ice crystals can grow larger at the expense of the surrounding liquid cloud droplets. As an ice crystal begins to fall, it can grow larger by colliding with supercooled liquid droplets, which freeze on contact. In an attempt to coax more precipitation from them, some clouds are seeded with silver iodide. We examined the various forms of precipitation, from raindrops that freeze on impact (producing freezing rain) to raindrops that freeze into tiny ice pellets called sleet. We learned that strong updrafts in a cumulonimbus cloud can keep ice particles suspended above the freezing level, where they acquire an additional coating of ice and form destructive hailstones. We looked at instruments and found that although the rain gauge is still the most commonly used method of measuring precipitation, Doppler radar has become an important tool for determining precipitation intensity and estimating rainfall amount. Rainfall estimates can also be obtained from radar and microwave scanners onboard satellites.

KEY TERMS The following terms are listed (with corresponding page numbers) in the order they appear in the text. Define each. Doing so will aid you in reviewing the material covered in this chapter. adiabatic process, 116 dry adiabatic rate, 116 moist adiabatic rate, 117 environmental lapse rate, 117 144

absolutely stable atmosphere, 117 absolutely unstable atmosphere, 118

condensation level, 120 conditionally unstable atmosphere, 120 orographic uplift, 123 rain shadow, 125 precipitation, 126 collision-coalescence process, 126 coalescence, 126 ice-crystal (Bergeron) process, 127 supercooled droplet, 128 ice nuclei, 128 accretion, 129 cloud seeding, 129 rain, 132 drizzle, 132 virga, 132 shower (rain), 132

snow, 133 fallstreaks, 134 flurries (of snow), 135 snow squall, 135 thundersnow, 135 blizzard, 135 sleet, 136 freezing rain (glaze), 136 rime, 136 black ice, 136 ice storm, 137 snow grains, 138 snow pellets, 138 hailstones, 138 standard rain gauge, 140 trace (of precipitation), 140 water equivalent, 141 radar, 141 Doppler radar, 142

QUESTIONS FOR REVIEW . What is an adiabatic process? . How would one normally obtain the environmental lapse rate? . Why are the moist and dry adiabatic rates of cooling different? . How can the atmosphere be made more stable? More unstable? . If the atmosphere is conditionally unstable, what does this mean? What condition is necessary to bring on instability? . Explain why an inversion represents an extremely stable atmosphere. . What type of clouds would you most likely expect to see in a stable atmosphere? In a conditionally unstable atmosphere? . Why are cumulus clouds more frequently observed during the afternoon? . There are usually large spaces of blue sky between cumulus clouds. Explain why this is so. . Why do most thunderstorms have flat tops? . List four primary ways in which clouds form. . Explain why rain shadows form on the downwind (leeward) side of mountains. . On which side of a mountain (windward or leeward) would lenticular clouds most likely form? . What is the primary difference between a cloud droplet and a raindrop? . Why do typical cloud droplets seldom reach the ground as rain?

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. Describe how the process of collision and coalescence produces rain. . How does the ice-crystal (Bergeron) process produce precipitation? What is the main premise behind this process? . Explain the main principle behind cloud seeding. . Explain how clouds can be seeded naturally. . How does rain differ from drizzle? . Why do heavy showers usually fall from cumuliform clouds? Why does steady precipitation normally fall from stratiform clouds? . Why is it never too cold to snow? . How would you be able to distinguish between virga and fallstreaks? . What is the difference between freezing rain and sleet? . How do the atmospheric conditions that produce sleet differ from those that produce hail? . Describe how a standard rain gauge measures precipitation. . How can the depth of snow be measured? . (a) What is Doppler radar? (b) How does Doppler radar measure the intensity of precipitation? . How are satellites able to measure precipitation intensity inside clouds?

QUESTIONS FOR THOUGHT AND EXPLORATION .

Suppose a mountain climber is scaling the outside of a tall skyscraper. Two thermometers (shielded from the sun) hang from the climber’s belt. One thermometer hangs freely, while the other is enclosed in a partially inflated balloon. As the climber scales the building, describe the change in temperature measured by each thermometer. . Where would you expect the moist adiabatic rate to be greater: in the tropics or near the North Pole? Explain why.

. What changes in weather conditions near Earth’s surface are needed to transform an absolutely stable atmosphere into an absolutely unstable atmosphere? . In the middle latitudes, under what circumstances can a rain shadow be formed on the western side of a mountain range? . A major snowstorm occurred in northern New Jersey. Three volunteer weather observers measured the snowfall. Observer #1 measured the depth of newly fallen snow every hour. At the end of the storm, Observer #1 added up the measurements and came up with a total of 12 inches of new snow. Observer #2 measured the depth of new snow twice: once in the middle of the storm and once at the end, and came up with a total snowfall of 10 inches. Observer #3 measured the new snowfall only once, after the storm had stopped, and reported 8.4 inches. Which of the three observers do you feel has the correct snowfall total? List at least five possible reasons why the snowfall totals were different. . Why is a warm, tropical cumulus cloud more likely to produce precipitation than a cold, stratus cloud? . Suppose a thick nimbostratus cloud contains ice crystals and supercooled cloud droplets all about the same size. Which precipitation process will be most important in producing rain from this cloud? Why? . Clouds that form over water are usually more efficient in producing precipitation than clouds that form over land. Why do you think this is so? . Everyday in summer a blizzard occurs over the Great Plains. Explain where and why. . It is −12°C (10°F) in Albany, New York, and freezing rain is falling. Can you explain why? Draw a vertical profile of the air temperature (a sounding) that illustrates why freezing rain is occurring at the surface. . When falling snowflakes become mixed with sleet, why is this condition often followed by the snowflakes changing into rain? . Why are ice storms not associated with cumuliform clouds, such as cumulus congestus and cumulonimbus?

Go to the Academic Journals section of the Meteorology portal. Use the search box at the left to bring up journal articles related to “cloud seeding.” Find an article that describes a cloudseeding experiment. Do the authors claim that the seeding produced a measurable result? Is the result statistically significant? What other factors might have produced the result aside from the cloud seeding?

ONLINE RESOURCES Visit www.cengagebrain.com to view additional resources, including video exercises, practice quizzes, an interactive eBook, and more. CLOUD DEVELOPMENT DEVELOPMENT AND PRECIPITATION Copyright 2018 Cengage Learning. All Rights Reserved. May not be copied, scanned, or duplicated, in whole or in part. WCN 02-200-202

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CHAPTER

6

Air Pressure and Winds Contents Atmospheric Pressure

D

ecember 19, 1980, was a cool day in Lynn, Massachusetts, but not cool enough to dampen the spirits of more than

2000 people who gathered in Central Square—all quare—all hoping to

Surface and Upper-Air Charts

catch at least one of the 1500 dollar bills that would be

Why the Wind Blows

the aircraft circled the city and dumped the money onto the

Winds and Vertical Air Motions

people below. However, owever, to the dismay of the onlookers, a

Determining Wind Direction and Speed

dropped from a small airplane at noon. Right ight on schedule,

westerly wind caught the currency before it reached the ground and carried it out over the cold Atlantic Ocean. Had the pilot or the sponsoring leather manufacturer examined the weather charts beforehand, they might have been able to predict that the wind would ruin their advertising scheme.

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147 147

Atmospheric Pressure In Chapter 1, we learned several important concepts about atmospheric pressure. One is that air pressure is simply the mass of air above a given level. As we climb in altitude above Earth’s surface, there are fewer air molecules above us; hence, atmospheric pressure always decreases with increasing height. Another concept we learned is that our atmosphere is highly compressible. This means that most of our atmosphere is squeezed (compressed) close to Earth’s surface, a situation that causes air pressure to decrease with height, rapidly at first, then more slowly at higher altitudes. So one way to change air pressure is to simply move up or down in the atmosphere. But what causes the air pressure to change in the horizontal? And why does the air pressure change at the surface? HORIZONTAL PRESSURE VARIATIONS—A TALE OF TWO CITIES To answer these questions, we eliminate some of the complexities of the atmosphere by constructing models.. Figure 6.1 shows a simple atmospheric model, a column of air extending well up into the atmosphere. In the column, the dots represent air molecules. Our model assumes: (1) that the air molecules are not crowded close to the surface and, unlike in the real atmosphere, the air density remains constant from the surface up to the top of the column, (2) that the width of the column does not change with height, and (3) that the air cannot freely move into or out of the column. Suppose we somehow force more air into the column in Fig. 6.1. What would happen? If the air temperature in the column does not change, the added air would make 148

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T

he scenario on the previous page raises two questions: (1) Why does the wind blow? and (2) How can one tell its direction by looking at weather charts? Chapter 1 has already answered the first question: Air moves in response to horizontal differences in pressure. This is what happens when we open a vacuum-packed can: air rushes from the higher pressure region outside the can toward the region of lower pressure inside. In the atmosphere, the wind blows in an attempt to equalize imbalances in air pressure. Does this mean that the wind always blows directly from high to low pressure? Not really, because the movement of air is controlled not only by pressure differences but by other forces as well. In this chapter, we will first consider how and why atmospheric pressure varies. Then we will look at the forces that influence atmospheric motions aloft and at the surface. Through studying these forces, we will be able to tell how the wind should blow in a particular region by examining surface and upper-air charts.

FFIGURE 6.1 A model of the atmosphere where air density remains constant with height. The air pressure at the surface is related to the number of molecules above. When air of the same temperature is stuffed into the column, the surface air pressure rises. When air is removed from the column, the surface pressure falls. (In the actual atmosphere, unlike in this model, density decreases with height.)

the column more dense, and the added weight of the air in the column would increase the surface air pressure. Likewise, if a great deal of air were removed from the column, the surface air pressure would decrease. Consequently, to change the surface air pressure, we need to change the mass of air in the column above the surface. But how can this feat be accomplished? Look at the air columns in Fig 6.2a.* Suppose both columns are located at the same elevation, both have the same air temperature, and both have the same surface air pressure. There must therefore be the same number of molecules (same mass of air) in each column above both cities. Further suppose that the surface air pressure for both cities remains the same, while the air above city 1 cools and the air above city 2 warms (see Fig. 6.2b). As the air in column 1 cools, the molecules move more slowly and crowd closer together, so the air becomes more dense. In the warm air above city 2, the molecules move faster and spread farther apart, and the air becomes less dense. Since the width of the columns does not change (and if we assume an invisible barrier exists between the columns), the total number of molecules above each city remains the same, and the surface air pressure does not change. Therefore, in the more-dense cold air above city 1, the height of the column decreases, while in the less-dense warm air above city 2, the height of the column increases. We now have a cold, shorter, more-dense column of air above city 1 and a warm, taller, less-dense air column above city 2. From this, we can conclude that it takes a shorter column of cold, more-dense air to exert the same surface pressure as a taller column of warm, less-dense air. This concept has a great deal of meteorological significance. Atmospheric pressure decreases more rapidly with height in the cold column of air. In the cold air above city 1 (Fig. 6.2b), move up the column and observe how quickly you pass through the densely packed molecules. This activity indicates a rapid change in pressure. In the warmer, less-dense air, the pressure does not decrease as *We will keep our same assumption as in Fig. 6.1; that is, (1) the air molecules are not crowded close to the surface, (2) the width of each column does not change, and (3) air is unable to move into or out of each column.

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FFIGURE 6.2 Illustration of how variations in temperature can produce horizontal pressure forces. (Note that, for simplicity, this model assumes that the air density is constant with height, whereas in the actual atmosphere, density decreases with height.) ((a) Two air columns, each with identical mass, have the same surface air pressure. (b) Because it takes a shorter column of cold air to exert the same surface pressure as a taller column of warm air, as column 1 cools, it must shrink, and as column 2 warms, it must rise. (c) Because at the same level in the atmosphere there is more air above the H in the warm column than above the L in the cold column, warm air aloft is associated with high pressure and cold air aloft with low pressure. The pressure differences aloft create a force that causes the air to move from a region of higher pressure toward a region of lower pressure. The removal of air from column 2 causes its surface pressure to drop, whereas the addition of air into column 1 causes its surface pressure to rise. (The difference in height between the two columns is greatly exaggerated.)

circulation of air is established due to the heating and cooling of air columns. As we will see in Chapter 7, this type of thermal circulation is the basis for a wide range of wind systems throughout the world. In summary, we can see how heating and cooling columns of air can establish horizontal variations in air pressure both aloft and at the surface. It is these horizontal differences in air pressure that cause the wind to blow.

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rapidly with height, simply because you climb above fewer molecules in the same vertical distance. In Fig. 6.2c, move up the warm, red column until you come to the letter H. Now move up the cold, blue column the same distance until you reach the letter L. Notice that there are more molecules above the letter H in the warm column than above the letter L in the cold column. The fact that the number of molecules above any level is a measure of the atmospheric pressure leads to an important concept: Warm air aloft is normally associated with high atmospheric pressure, and cold air aloft is associated with low atmospheric pressure. In Fig. 6.2c, the horizontal difference in temperature creates a horizontal difference in pressure. The pressure difference establishes a force (called the pressure gradient force) that causes the air to move from higher pressure toward lower pressure. Consequently, if we remove the invisible barrier between the two columns near the top of Column 1 and allow the air aloft to move horizontally, the air will move from column 2 toward column 1. As the air aloft leaves column 2, the mass of the air in the column decreases, and so does the surface air pressure. Meanwhile, the accumulation of air in column 1 causes the surface air pressure to increase. Higher air pressure at the surface in column 1 and lower air pressure at the surface in column 2 causes the surface air to move from city 1 toward city 2 (see Fig. 6.3). As the surface air moves out away from city 1, the air aloft slowly sinks to replace this outwardly spreading surface air. As the surface air flows into city 2, it slowly rises to replace the depleted air aloft. In this manner, a complete

FFIGURE 6.3 The heating and cooling of air columns causes horizontal pressure variations aloft and at the surface. These pressure variations force the air to move from areas of higher pressure toward areas of lower pressure. In conjunction with these horizontal air motions, the air slowly sinks above the surface high and rises above the surface low. AIR PRESSURE AND AND WINDS

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ON A SSPECIAL TOPIC 6.1

The Atmosphere Obeys the Gas Law The relationship among the pressure, temperature, and density of air can be expressed by

This simple relationship, often referred to as the gas law (or equation of state), tells us that the pressure of a gas is equal to its temperature times its density times a constant. When we ignore the constant and look at the gas law in symbolic form, it becomes pT where, of course, p is pressure, T is temperature, and (the Greek letter rho, pronounced “row”) represents air density. The line is a symbol meaning “is proportional to.” A change in one variable causes a corresponding change in the other two variables. Thus, it will be easier to understand the behavior of a gas if we keep one variable from changing and observe the behavior of the other two. Suppose, for example, we hold the temperature constant. The relationship then becomes p (temperature constant). This expression says that the pressure of the gas is proportional to its density, as long as its temperature does not change. Consequently, if the temperature of a gas (such as air) is held constant: As the pressure increases the density increases, and

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Pressure temperature density constant.

FFIGURE 1 Air above a region of surface high pressure is more dense than air above a region of surface low pressure (at the same temperature). (The dots in each column represent air molecules.)

as the pressure decreases, the density decreases. In other words, at the same temperature, air at a higher pressure is more dense than air at a lower pressure. If we apply this concept to the atmosphere, then with nearly the same temperature and elevation, air above a region of surface high pressure is more dense than air above a region of surface low pressure (see Fig. 1). We can see, then, that for surface high-pressure areas (anticyclones) and surface low-pressure areas (mid-latitude cyclones) to form, the air density (mass of air) above these systems must change. We just considered how pressure and density are related when the temperature is not changing. What happens to the gas law when the pressure of a gas remains

Before we examine how air pressure is measured, you may wish to look at Focus section 6.1, which describes how air pressure, air density, and air temperature are interrelated. MEASURING AIR PRESSURE Up to this point, we have described air pressure as the mass of the atmosphere above any level. We can also define air pressure as the force exerted by the air molecules over a given area. Billions of air molecules constantly push on the human body. This force is exerted equally in all directions. We are not crushed by the force because billions of molecules inside the body push outward just as hard. Even though we do not actually feel the constant bombardment of air, we can detect quick changes in it. 150

constant? In shorthand notation, the relationship becomes (Constant pressure) constant T . This relationship tells us that when the pressure of a gas is held constant, the gas becomes less dense as the temperature goes up, and more dense as the temperature goes down. Therefore, at a given atmospheric pressure, air that is cold is more dense than air that is warm. Keep in mind that the idea that cold air is more dense than warm air applies only when we compare volumes of air at the same level, where pressure changes are relatively small in any horizontal direction.

For example, if we climb rapidly in elevation our ears may “pop.” This experience happens because air collisions outside the eardrum lessen as the air pressure decreases. The popping comes about as air collisions between the inside and outside of the ear equalize. Instruments that detect and measure pressure changes are called barometers, which literally means an instrument that measures bars. In meteorology, the bar is a unit of pressure that describes a force over a given area.* Because the bar is a relatively large unit, and because surface *By definition, a bar is a force of 100,000 newtons acting on a surface area of 1 square meter. A newton is the amount of force required to move an object with a mass of 1 kilogram so that it increases its speed at a rate of 1 meter per second each second.

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FFIGURE 6.4 Atmospheric pressure in inches of mercury and in millibars.

pressure changes are normally small, the unit of pressure most commonly found on surface weather maps is the millibar (mb), where one millibar is equal to onethousandth of a bar. The hectopascal (hPa)* is gradually replacing the millibar as the preferred unit of pressure on surface and upper-air maps. (No conversion is needed in moving between millibars and hectopascals.) A common pressure unit used in aviation and on television and radio weather broadcasts is inches of mercury (Hg). At sea level, standard atmospheric pressure** is 1013.25 mb = 1013.25 hPa = 29.92 in. Hg. Because we measure atmospheric pressure with an instrument called a barometer, atmospheric pressure is com also referred to as barometric pressure. Figure 6.4 compares pressure readings in millibars and in inches of mermer cury. Why do we refer to “inches of mercury”? Evangelista Torricelli, a student of Galileo’s, invented the mercury *The unit of pressure designed by the International System (SI) of measurement is the pascal (Pa), where 1 pascal is the force of 1 newton acting on a surface of 1 square meter. A more common unit is the hectopascal (hPa), as 1 hectopascal equals 1 millibar. (Additional pressure units and conversions are given in Appendix A.) **Standard atmospheric pressure at sea level is the pressure extended by a column of mercury 29.92 in. (760 mm) high, having a density of 1.36 104 kg/m3, and subject to an acceleration of gravity of 9.80 m/sec2.

FFIGURE 6.5 The mercury barometer. The height of the mercury column is a measure of atmospheric pressure.

barometer in 1643. His barometer, similar to those used today, consisted of a long glass tube open at one end and closed at the other (see Fig. 6.5). Removing air from the tube and covering the open end, Torricelli immersed the lower portion into a dish of mercury. He removed the cover, and the mercury rose up the tube to nearly 30 inches above the level in the dish. Torricelli correctly concluded that the column of mercury in the tube was balancing the weight of the air above the dish, and, hence, its height was a measure of atmospheric pressure. Mercury barometers are still used today in many settings, although in some areas, including Europe, they are no longer manufactured or sold because mercury use can pose a risk to human health. The most common type of home barometer—the aneroid barometer—contains no fluid. Inside this instrument is a small, flexible metal box called an aneroid cell. Before the cell is tightly sealed, air is partially removed, so

DID YOU KNOW? Although 1013.25 mb (29.92 in.) is the standard atmospheric pressure at sea level, it is not the average sea-level pressure. Earth’s average sea-level pressure is 1011.0 mb (29.85 in.). Because much of Earth’s surface is above sea level, Earth’s annual average surface pressure is estimated to be 984.43 mb (29.07 in.). AIR PRESSURE AND AND WINDS

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FFIGURE 6.6 The aneroid barometer.

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that small changes in external air pressure cause the cell to expand or contract. The size of the cell is calibrated to represent different pressures, and any change in its size is amplified by levers and transmitted to an indicating arm, which points to the current atmospheric pressure (see Fig. 6.6). Notice that the aneroid barometer often has descripdescrip tive weather-related words printed above specific pressure values. These descriptions indicate the most likely weather conditions when the needle is pointing to that particular pressure reading. Generally, the higher the reading, the more likely clear weather will occur, and the lower the reading, the better are the chances for inclement weather. This situation occurs because surface high-pressure areas are associated with sinking air and normally fair weather, whereas surface low-pressure areas are associated with rising air and usually cloudy, wet weather. A steady rise in atmospheric pressure (a rising barometer reading) usually indicates clearing weather or fair weather, whereas a steady drop in atmospheric pressure (a falling barometer reading) often signals the approach of a cyclonic storm with inclement weather. The altimeter and barograph are two types of aneroid barometers. Altimeters are aneroid barometers that measure pressure, but are calibrated to indicate altitude.

FFIGURE 6.7 A recording barograph.

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Barographs are recording aneroid barometers. Basically, the barograph consists of a pen attached to an indicating arm that marks a continuous record of pressure on chart paper. The chart paper is attached to a drum rotated slowly by an internal mechanical clock (see Fig. 6.7). Many aircraft use radar-based altimeters that take rapid meamea surements of the distance between the aircraft and ground. This is especially valuable during landing, when the aircraft’s altitude is changing rapidly. Digital barometers are becoming more common. This type of barometer uses a device called a transducer that detects the change in pressure exerted by the atmosphere on a precisely engineered surface. The change in pressure is then converted into electrical signals. Some digital barometers are small enough to be placed in smartphones, while others are designed for research settings. PRESSURE READINGS Obtaining the correct air pressure from a mercury barometer involves more than simply reading the height of the mercury column. Being a fluid, mercury is sensitive to changes in temperature; it will expand when heated and contract when cooled. Consequently, to obtain accurate pressure readings without the influence of temperature, all mercury barometers are corrected as if they were read at the same temperature. Also, because Earth is not a perfect sphere, the force of gravity is not a constant. Since small gravity differences influence the height of the mercury column, they must be taken into account when reading the barometer. Finally, each mercury barometer has its own “built-in” error, called instrument error, which is caused, in part, by the surface tension of the mercury against the glass tube. After being corrected for temperature, gravity, and instrument error, the barometer reading at a particular location and elevation is termed station pressure. Figure 6.8a gives the station pressure measured at four locations only a few hundred kilometers apart. The different station pressures of the four cities are due primarily to the cities being at different elevations above sea level. This fact becomes even clearer when we realize that atmospheric pressure changes much more quickly when we move upward than it does when we move sideways. A small vertical difference between two observation sites can yield a large difference in station pressure. Thus, to properly monitor horizontal changes in pressure, barometer readings must be corrected for altitude. Altitude corrections are made so that a barometer reading taken at one elevation can be compared with a barometer reading taken at another. Station pressure observations are normally adjusted to a level of mean sea level, the level representing the average surface height of the ocean. The adjusted reading is called sea-level pressure. The size of the correction depends primarily on how high the station is above sea level.

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FFIGURE 6.8 The top diagram (a) shows four cities (A, B, C, and D) at varying el elevations above sea level, all with different station pressures. The middle diagram (b) represents sea-level pressures of the four cities plotted on a sea-level chart. The bottom diagram (c) shows sea-level pressure readings of the four cities plus other sea-level pressure readings at locations not shown in (a) and (b). Isobars are drawn on the chart (gray lines) at intervals of 4 millibars.

Near Earth’s surface, atmospheric pressure decreases on the average by about 10 millibars (mb) for every 100 meters of increase in elevation (about 1 in. of mercury for each 1000-ft rise).* Notice in Fig. 6.8a that city A has a station pressure of 952 mb. Notice also that city A is 600 meters above sea level. Adding 10 mb per 100 m to its station pressure yields a sea-level pressure of 1012 mb (Fig. 6.8b). After all the station pressures are adjusted to sea level (Fig. 6.8b), we are able to see the horizontal variations in sea-level pressure—something we were not able to see from the station pressures alone in Fig. 6.8a. When more pressure data are added (Fig. 6.8c), the chart can be analyzed and the pressure pattern visualized. Isobars (lines connecting points of equal pressure) are drawn as solid dark lines at intervals of 4 mb, with 1000 mb being the base value. Note that the isobars do not pass through each point, but, rather, between many of them, with the exact values being interpolated from the data given on the chart. For example, follow the 1008-mb *This decrease in atmospheric pressure with height (10 mb/100 m) occurs when the air temperature decreases at the standard lapse rate of 6.5°C/1000 m. Because atmospheric pressure decreases more rapidly with height in cold (more dense) air than it does in warm (less dense) air, the vertical rate of pressure change is typically greater than 10 mb per 100 m in cold air and less than that in warm air.

line from the top of the chart southward and observe that there is no plotted pressure of 1008 mb. The 1008-mb isobar, however, comes closer to the station with a sealevel pressure of 1007 mb than it does to the station with a pressure of 1010 mb. With its isobars, the bottom chart (Fig. 6.8c) is now called a sea-level pressure chart, or, simply, a surface map. When weather data are plotted on the map, it becomes a surface weather map.

Surface and Upper-Air Charts Figure 6.9a is a simplified surface map that shows areas of high and low pressure and arrows that indicate wind

DID YOU KNOW? On October 12, 1962, one of the most powerful mid-latitude cyclonic storms ever to develop along the Pacific Northwest coast slammed into Oregon and Washington. With a central pressure of 960 mb (28.35 in.), the storm had such a strong pressure gradient that it produced winds on its eastern side in excess of 75 mi/hr along the Pacific Coast from California to British Columbia with a reported peak wind gust of 179 mi/hr at Cape Blanco, Oregon. AIR PRESSURE AND AND WINDS

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FFIGURE 6.9 (a) Surface map showing areas of high and low pressure. The solid lines are isobars drawn at 4-mb intervals. The arrows represent wind direction—the direction from which the wind is blowing. Notice that the wind blows across the isobars. (b) The upperlevel (500-mb) map for the same day as the surface map. Solid lines on the map are contour lines in meters above sea level. Dashed red lines are isotherms in °C. Arrows show wind direction. Notice that, on this upper-air map, the wind blows parallel to the contour lines.

direction, which is the direction from which the wind is blowing. The large blue H’s on the map indicate the centers of high pressure, which are also called anticyclones. The large red L’ L’s represent centers of low pressure, also known as mid-latitude cyclonic storms, or extratropical cyclones, because they form in the middle latitudes, outside of the tropics. The solid dark lines are isobars with units in millibars. Notice that the surface winds tend to blow across the isobars toward regions of lower pressure. In fact, as we briefly observed in Chapter 1, in the Northern Hemisphere the winds blow counterclockwise and inward toward the center of the lows and clockwise and outward from the center of the highs. Figure 6.9b shows an upper-air chart for the same day as the surface map in Fig. 6.9a. The upper-air map is a constant pressure chart because it is constructed to show height variations along a constant pressure (isobaric) surface, which is why these maps are also known as isobaric maps. This particular isobaric map shows height variations at a pressure level of 500 millibars and so is called a 500-millibar map. In Fig. 6.9b, the height of the 500-mb level varies from 5340 m to more than 5700 m. A typical height in mid-latitudes is about 5600 m, or 18,000 ft, above sea level. The solid dark lines on the map are contour lines, lines that connect points of equal altitude above sea level. Although the contour lines are height lines, they illustrate pressure much like isobars do. Consequently, contour lines of low height represent a region of lower pressure, and contour lines of high height represent a region of higher pressure. (Additional information on isobaric maps is given in Focus section 6.2.) Notice on the 500-mb map (Fig. 6.9b) that the contour lines typically decrease in value from south to north. This decrease in value is due to a decrease in temperature as illustrated by the dashed red lines, which 154

are isotherms—lines of equal temperature. Observe that colder air is generally to the north and warmer air to the south, and recall from our earlier discussion (see pp. 148-150) that cold air aloft is associated with low pressure, warm air aloft with high pressure. The contour lines are not straight, however. They bend and turn, indicating ridges (elongated highs) where the air is warmer and indicating depressions, or troughs (elongated lows) where the air is colder. The arrows on the 500-mb map show the wind direction. Notice that, unlike the surface winds that cross the isobars in Fig. 6.9a, the winds on the 500-mb chart tend to flow parallel to the contour lines in a wavy west-to-east direction. Surface and upper-air charts are valuable tools for the meteorologist. Surface maps describe where the centers of high and low pressure are found, as well as the winds and weather associated with these systems, whereas upper-air charts are extremely important in forecasting the weather. The upper-level winds not only determine the movement of surface pressure systems but, as we will see in Chapter 8, they determine whether these surface systems will intensify or weaken. At this point, however, our interest lies mainly in the movement of air. Now that we have looked at surface and upper-air maps, we will use them to study why the wind blows the way it does, at both the surface and aloft.

Why the Wind Blows Our understanding of why the wind blows stretches back through several centuries, with many scientists contributing to our knowledge. When we think of the movement of air, however, one great scholar stands out: Isaac Newton (1642– 1727), who formulated several fundamental laws of motion.

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FOCUS

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Isobaric Maps

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FFIGURE 2 The area shaded gray in the diagram represents a surface of constant pressure. Because of the changes in air density, a surface of constant pressure rises in warm, lessdense air and lowers in cold, more-dense air. These changes in height of a constant pressure (500-mb) surface show up as contour lines on a constant pressure (isobaric) 500-mb map.

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Figure 2 shows a column of air where warm, less-dense air lies to the south and cold, more-dense air lies to the north. The area shaded gray at the top of the column represents a constant pressure (isobaric) surface, where the atmospheric pressure at all points along the surface is 500 millibars. Notice that the height of the pressure surface varies. In the warmer air, a pressure reading of 500 mb is found at a higher level, while in the colder air, 500 mb is observed at a much lower level. The variations in height of the 500-mb constant pressure surface are shown as contour lines on the constant pressure (500-mb) map, situated at the bottom of the column. Each contour line tells us the altitude above sea level at which we would obtain a pressure reading of 500 mb. As we would expect from our earlier discussion of pressure in Figs. 6.2 and 6.3, p. 149, the contour lines on the chart are higher in the warm air and lower in the cold air. Although contour lines are height lines, keep in mind that they illustrate pressure in the same manner as do isobars, as contour lines of high height (warm air aloft) represent regions of higher pressure, and contour lines of low height (cold air aloft) represent regions of low pressure. In many instances, the contour lines on an isobaric map are not straight, but rather appear as wavy lines. Figure 3 illustrates how these wavy contours on the map relate to the change in altitude of the isobaric surface.

FFIGURE 3 The wavelike patterns of an isobaric surface reflect the changes in air temperature. An elongated region of warm air aloft shows up on an isobaric map as higher heights and a ridge; the colder air shows as lower heights and a trough.

NEWTON’S LAWS OF MOTION Newton’s first law of motion states that an object at rest will remain at rest and an object in motion will remain in motion (and travel at a constant velocity along a straight line) as long as no force is exerted on the object. For example, a baseball in a pitcher’s hand will remain there until a force (a push) acts upon the ball. Once the ball is pushed (thrown), it would continue to move in that direction forever if it were not for the force of air friction (which slows it down), the force of gravity (which pulls it toward the ground), and the catcher’s mitt

(which exerts an equal but opposite force to bring it to a halt). Similarly, to start air moving, to speed it up, to slow it down, or even to change its direction requires the action of an external force. This brings us to Newton’s second law. Newton’s second law states that the force exerted on an object equals its mass times the acceleration produced.* In symbolic form, this law is written as F = ma. *Newton’s second law can also be stated in this way: The acceleration of an object (times its mass) is caused by all of the forces acting on it. AIR PRESSURE AND AND WINDS

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We will first study the forces that influence the flow of air aloft. Then we will see which forces modify winds near the ground.

FFIGURE 6.10 The higher water level creates higher fluid pressure at the bottom of tank A and a net force directed toward the lower fluid pressure at the bottom of tank B. This net force causes water to move from higher pressure toward lower pressure.

From this relationship we can see that, when the mass of an object is constant, the force acting on the object is directly related to the acceleration that is produced. A force in its simplest form is a push or a pull. Acceleration is the speeding up, the slowing down, and/or the changing of direction of an object. (More precisely, acceleration is the change of velocity* over a period of time.) Because more than one force may act upon an object, Newton’s second law always refers to the net, or total, force that results. An object will always accelerate in the direction of the total force acting on it. Therefore, to determine in which direction the wind will blow, we must identify and examine forces that affect the horizontal movement of air. These forces include: . pressure gradient force . Coriolis force . friction

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*Velocity specifies both the speed of an object and its direction of motion.

FFIGURE 6.11 The pressure gradient between point 1 and point 2 is 4 mb per 100 km. The net force directed from higher toward lower pressure is the pressure gradient force.

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FORCES THAT INFLUENCE THE WIND We have already learned that horizontal differences in atmospheric pressure cause air to move and, hence, the wind to blow. Since air is an invisible gas, it may be easier to see how pressure differences cause motion if we examine a visible fluid, such as water. In Fig. 6.10, the two large tanks are connected by a pipe. Tank A is two-thirds full and tank B is only one-half full. Since the water pressure at the bottom of each tank is proportional to the weight of water above, the pressure at the bottom of tank A is greater than the pressure at the bottom of tank B. Moreover, since fluid pressure is exerted equally in all directions, there is a greater pressure in the pipe directed from tank A toward tank B than from B toward A. Since pressure is force per unit area, there must also be a net force directed from tank A toward tank B. This force causes the water to flow from left to right, from higher pressure toward lower pressure. The greater the pressure difference, the stronger the force, and the faster the water moves. In a similar way, horizontal differences in atmospheric pressure cause air to move. Figure 6.11 shows a region Pressure Gradient Force of higher pressure on the map’s left side, with lower prespres sure on the right. The isobars show how the horizontal pressure is changing. If we compute the amount of pressure change that occurs over a given distance, we have the pressure gradient; thus Pressure gradient 5

difference in pressure . distance

In Fig. 6.11, the pressure gradient between points 1 and 2 is 4 millibars per 100 kilometers. Suppose the pressure in Fig. 6.11 were to change, and the isobars become closer together. This condition would produce a rapid change in pressure over a relatively short distance, or what is called a steep (or strong) strong pressure gradient. However, if the pressure were to change such that the isobars spread farther apart, then the difference in pressure would be small over a relatively large distance. This condition is called a gentle (or weak) pressure gradient. Notice in Fig. 6.11 that when differences in horizontal air pressure exist there is a net force acting on the air. This force, called the pressure gradient force (PGF), is directed from higher toward lower pressure at right angles to the isobars. The magnitude of the force is directly related to the pressure gradient. Steep pressure gradients correspond to strong pressure gradient forces and vice versa. Figure 6.12 shows the relationship between pressure gradient and pressure gradient force.

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The pressure gradient force is the force that causes the wind to blow. Closely spaced isobars on a weather chart indicate steep pressure gradients, strong forces, and high winds. In contrast, widely spaced isobars indicate gentle pressure gradients, weak forces, and light winds. An example of a steep pressure gradient and strong winds is illustrated on the surface weather map in Fig. 6.13. Notice that the tightly packed isobars around Hurricane Sandy are producing a steep pressure gradient and strong surface winds, gusting to more than 80 knots over portions of Long Island. If the pressure gradient force were the only force acting upon air, we would always find winds blowing directly from higher toward lower pressure. However, the moment air starts to move, it is deflected in its path by the Coriolis force.

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Coriolis Force The Coriolis force describes an apparent force that is due to the rotation of Earth. To understand

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FFIGURE 6.12 The closer the spacing of the isobars, the greater the pressure gradient. The greater the pressure gradient, the stronger the pressure gradient force (PGF). The stronger the PGF, PGF the greater the wind speed. The red arrows represent the relative magnitude of the force, which is always directed from higher toward lower pressure.

FFIGURE 6.14 On nonrotating platform A, the thrown ball moves in a straight line. On platform B, which rotates counter counterclockwise, the ball continues to move in a straight line. However, platform B is rotating while the ball is in flight; thus, to the person throwing the ball on platform B, the ball appears to deflect to the right of its intended path.

how it works, consider two people playing catch as they sit opposite one another on the rim of a merry-go-round (see Fig. 6.14, platform A). If the merry-go-round is not moving, each time the ball is thrown, it moves in a straight line to the other person. Suppose the merry-go-round starts turning counterclockwise, the same direction Earth spins as viewed from above the North Pole. If we watch the game of catch from above, we see that the ball moves in a straight-line path just as before. However, to the people playing catch on the merry-go-round, the ball seems to veer to its right each time it is thrown, always landing to the right of the point intended by the thrower (see Fig. 6.14, platform B). FFIGURE 6.13 Surface map for 4 p.m. (EST) Monday, October 29, 2012, as Hurricane (Superstorm) Sandy approaches the New Jersey shore from the east. Isobars are dark gray lines with units in millibars. The interval between isobars is 4 mb. Sandy’s central pressure is 943 mb or 27.85 in. The tightly packed isobars are associated with a strong pressure gradient, a strong pressure gradient force, and high winds, gusting to over 80 knots over portions of Long Island, New York. Wind directions at each station are shown by lines that parallel the wind. Wind speeds are indicated by barbs and flags, where a wind indicated by the symbol would be a wind from the northwest at between 23 and 27 knots. (See blue insert.) Wind speeds over the ocean, where no observations are available, are based on computer model calculations. AIR PRESSURE AND AND WINDS

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FFIGURE 6.15 Except at the equator, a free-moving object heading either east or west (or any other direction) will appear from Earth to deviate from its path as Earth rotates beneath it. The deviation (Coriolis force) is greatest at the poles and decreases to zero at the equator. Notice that the aircraft’s deviation from its intended destination is greatest at high latitudes and nonexistent at the equator.

This perspective is accounted for by the fact that, while the ball moves in a straight-line path, the merrygo-round rotates beneath it, and by the time the ball reaches the opposite side, the catcher has moved. To anyone on the merry-go-round, it seems as if there is some force causing the ball to deflect to the right of its intended path. This apparent force is called the Coriolis force after Gaspard Coriolis, a nineteenth-century French scientist who worked it out mathematically. (Because it is an apparent force due to the rotation of Earth, it is also called the Coriolis effect.) This effect occurs not only on merrygo-rounds but on rotating Earth, too. All free-moving objects, such as ocean currents, aircraft, artillery projectiles, and air molecules seem to deflect from a straight-line path because Earth rotates under them. The Coriolis force causes the wind to deflect to the right of its intended path in the Northern Hemisphere and to the left of its intended path in the Southern Hemisphere. To illustrate, consider a satellite in polar circular orbit. If Earth were not rotating, the path of the satellite would be observed to move directly from north to south, parallel to Earth’s meridian lines. However, Earth does rotate, carrying us and meridians eastward with it. Because of this rotation, in the Northern Hemisphere we see the satellite moving southwest instead of due south; it seems to veer off its path and move toward its right. In the Southern Hemisphere, Earth’s direction of rotation is clockwise as viewed from above the South Pole. Consequently, a satellite moving northward from the South Pole would appear to move northwest and, hence, would veer to the left of its path. As the wind speed increases, the Coriolis force increases; hence, the stronger the wind, the greater the deflection. Additionally, the Coriolis force increases for all wind 158

speeds from a value of zero at the equator to a maximum at the poles. This phenomenon is illustrated in Fig. 6.15 where three aircraft, each at a different latitude, are flying along a straight-line path, with no external forces acting on them. The destination of each aircraft is due east and is marked on the illustration in Fig. 6.15a. Each plane travels in a straight path relative to an observer positioned at a fixed spot in space. Earth rotates beneath the moving planes, causing the destination points at latitudes 30° and 60° to change direction slightly when seen by an observer in space (see Fig. 6.15b). To an observer standing on Earth, however, it is the plane that appears to deviate. The amount of deviation is greatest toward the pole and nonexistent at the equator. Therefore, the Coriolis force has a far greater effect on the plane at high latitudes (large deviation) than on the plane at low latitudes (small deviation). On the equator, it has no effect at all. The same is true of its effect on winds. In summary, to an observer on Earth, objects moving in any direction (north, south, east, or west) are deflected to the right of their intended path in the Northern Hemisphere and to the left of their intended path in the Southern Hemisphere. The amount of deflection depends upon: . the rotation of Earth . the latitude . the object’s speed In addition, the Coriolis force acts at right angles to the wind, only influencing wind direction and never wind speed. The Coriolis force is present in all motions relative to Earth’s surface. However, because its strength is directly related to the speed of motion, the Coriolis force tends to be weak over relatively small areas. In most of our everyday

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BRIEF REVIEW In summary, we know that: ●

Atmospheric (air) pressure is the pressure exerted by the mass of air above a region.

A change in surface air pressure can be brought about by changing the mass (amount of air) above the surface.

Heating and cooling columns of air can establish horizontal variations in atmospheric pressure aloft and at the surface.

A difference in horizontal air pressure produces a horizontal pressure gradient force.

The pressure gradient force is always directed from higher pressure toward lower pressure and it is the pressure gradient force that causes the air to move and the wind to blow.

Steep pressure gradients (tightly packed isobars on a weather map) indicate strong pressure gradient forces and high winds; gentle pressure gradients (widely spaced isobars) indicate weak pressure gradient forces and light winds.

Once the wind starts to blow, the Coriolis force causes it to bend to the right of its intended path in the Northern Hemisphere and to the left of its intended path in the Southern Hemisphere.

STRAIGHT-LINE FLOW ALOFT Earlier in this chapter, we saw that the winds aloft on an upper-level chart blow more or less parallel to the isobars or contour lines. We can see why this phenomenon happens by carefully looking at Fig. 6.16, which shows a map in the Northern Hemisphere, above Earth’s frictional influence,* at an altitude of about 1 kilometer above Earth’s surface. Horizontal pressure changes are shown by isobars. The evenly spaced isobars indicate a constant pressure gradient force (PGF) directed from south toward north as indicated by the red arrow at the left. Why, then, does the map show a wind blowing from the west? We can answer this question *The friction layer (the layer where the wind is influenced by frictional interaction with objects on Earth’s surface) usually extends from the surface up to about 1000 m (3300 ft) above the ground.

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experiences, the force is so weak (compared to other forces involved in those experiences) that it is negligible. Contrary to popular belief, the Coriolis force does not cause water to turn clockwise or counterclockwise when draining from a sink. The Coriolis force is also minimal on small-scale winds, such as those that blow inland along coasts in summer. Here, the Coriolis force itself might be strong because of high winds, but the force cannot produce much deflection over the relatively short distances. Only where winds blow over long distances is the effect significant. With this information in mind, we will first examine how the pressure gradient force and the Coriolis force produce straightline winds aloft, above the frictional influence of Earth’s surface. We will then look at what other forces come into play as winds blow along a curved path.

FFIGURE 6.16 Above the level of friction, air initially at rest will accelerate until it flows parallel to the isobars at a steady speed with the horizontal pressure gradient force (PGF) balanced by the Coriolis force (CF). Wind blowing under these conditions is called geostrophic.

by placing a parcel of air at position 1 in the diagram and watching its behavior. At position 1, the PGF acts immediately upon the air parcel, accelerating it northward toward lower pressure. However, the instant the air begins to move, the Coriolis force deflects the air toward its right, curving its path. As the parcel of air increases in speed (positions 2, 3, and 4), the magnitude of the Coriolis force increases (as shown by the blue arrows), bending the wind more and more to its right. Eventually, the wind speed increases to a point where the Coriolis force just balances the PGF. At this point (position 5), the wind no longer accelerates because the net force is zero. Here the wind flows in a straight path, parallel to the isobars at a constant speed.* This flow of air is called a geostrophic (geo: (geo: “earth;” strophic: “turning”) wind. Notice that the geostrophic wind blows in the Northern Hemisphere with lower pressure to its left and higher pressure to its right. When the flow of air is purely geostrophic, the isobars (or contour lines) are straight and evenly spaced, and the wind speed is constant. In the real world, isobars are rarely straight or evenly spaced, and the wind normally changes speed as it flows along. So, the geostrophic wind is usually only an approximation of the actual wind. However, the *At first, it may seem odd that the wind blows at a constant speed with no net force acting on it. But when we remember that the net force is necessary only to accelerate (F = ma) the wind, it makes more sense. For example, it takes a considerable net force to push a car and get it rolling from rest. But once the car is moving, it only takes a force large enough to counterbalance friction to keep it going. There is no net force acting on the car, yet it rolls along at a constant speed. AIR PRESSURE AND AND WINDS

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FFIGURE 6.17 The isobars and contours on an upper-level chart are like the banks along a flowing stream. When they are widely spaced, the flow is weak; when they are narrowly spaced, the flow is stronger. The increase in winds on the chart results in a stronger Coriolis force (CF), which balances a larger pressure gradient force (PGF).

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approximation is generally close enough to help us more clearly understand the behavior of the winds aloft. As we would expect from our previous discussion of winds, the speed of the geostrophic wind is directly related to the pressure gradient. In Fig. 6.17, we can see that a geostrophic wind flowing parallel to the isobars is similar to water in a stream flowing parallel to its banks. At position 1, the wind is blowing at a low speed; at position 2, the pressure gradient increases and the wind speed picks up. Notice also that at position 2, where the wind speed is greater, the Coriolis force is greater and balances the stronger pressure gradient force. Because the geostrophic wind blows parallel to isobars (or contour lines), we can estimate the geostrophic wind directly by observing the orientation of the isobars on an upper-level chart. Notice in Fig. 6.17 that the west-to-east orientation of the isobars produces a westerly wind. In a similar way, it is possible to estimate the wind flow and pressure patterns aloft by watching the movement of clouds. This concept is explained further in Focus section 6.3 (p. 161). We know that the winds aloft do not always blow in a straight line; frequently, they curve and bend into meandering loops. In the Northern Hemisphere, winds blow counterclockwise around areas of low pressure and clockwise around areas of high pressure. The next section explains why.

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CURVED WINDS AROUND LOWS AND HIGHS ALOFT Because lows are also known as cyclones, the flow of air around them (which is counterclockwise in the Northern Hemisphere) is often called cyclonic flow. Likewise, the flow of air around a high (which is clockwise in the Northern Hemisphere) is called anticyclonic flow. Look at the wind flow around the upper-level low (Northern Hemisphere) in Fig. 6.18a. At first, it appears as though the wind is defying the Coriolis force by bending to the left as it moves counterclockwise around the system. Let’s see why the wind blows this way. Suppose we consider a parcel of air initially at rest at position 1 in Fig. 6.18a. The pressure gradient force accelerates the air inward toward the center of the low and the Coriolis force deflects the moving air to its right, until the air is moving parallel to the isobars at position 2. If the isobars were straight lines ahead of this point, the wind would move northward at a constant speed, bringing it to position 3. In reality, the isobars are curved, and the wind follows that curvature (position 4). A wind that blows at a constant speed parallel to curved isobars above the level of frictional influence is termed a gradient wind. Why does the wind follow this curved path? Look closely at position 2 (Fig. 6.18a), where a parcel is moving northward, and

FFIGURE 6.18 Winds and related forces around an area of low pressure above the friction level in the Northern Hemisphere. Notice that the pressure gradient force (PGF) is in red, while the Coriolis force (CF) is in blue

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FOCUS

ON AN O OBSERVATION 6.3

Both the wind direction and the orientation of the isobars aloft can be estimated by obob serving middle- and high-level clouds from Earth’s surface. Suppose, for example, we are in the Northern Hemisphere watching clouds directly above us move from southsouth west to northeast at an elevation of about 3000 m or 10,000 ft (see Fig. 4a). This indicates that the geostrophic wind at this level is southwesterly. Looking downwind, the geostrophic wind blows parallel to the isobars with lower pressure on the left and higher pressure on the right. Thus, if we stand with our backs to the direction from which the clouds are moving, lower pressure aloft will always be to our left and higher prespres sure to our right. right From this observation, we can draw a rough upper-level chart (Fig. 4b), which shows isobars and wind direction for an elevation of approximately 10,000 ft. The isobars aloft will not continue in a southwest-northeast direction indefinitely;

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Estimating Wind Direction and Pressure Patterns Aloft by Watching Clouds

FFIGURE 4 The movement of clouds can help determine the geostrophic wind direction and orientation of the isobars. Upper-level clouds moving from the southwest (a) indicate isobars and winds aloft (b). When extended horizontally, the upper-level chart appears as in (c), where lower pressure is to the west and higher pressure is to the east.

rather, they will often bend into wavy patterns. We may carry our observation one step further, then, by assuming a bending of the lines (Fig. 4c). Thus, with a southwesterly wind aloft, a trough of low pressure will be found to our west and a ridge

observe that if the wind were in geostrophic balance, the inward-directed pressure gradient force (PGF) PGF) would be in PGF balance with the outward-directed Coriolis force (CF) CF) and CF the wind should continue blowing in a straight line toward position 3. Suppose the parcel reaches position 3. Notice in Fig. 6.18b that here the PGF would now be angled toward the southwest, because the PGF is always directed toward lower pressure at right angles to the isobars. This means that part of the PGF will be directed southward, against the northward motion of the parcel. This situation slows

of high pressure to our east. What would be the pressure pattern if the winds aloft were blowing from the northwest? Answer: A trough would be to the east and a ridge to the west.

the wind down just a bit. Recall that the Coriolis force is directly related to wind speed, so the slower wind weakens the Coriolis force. As a result of the weaker Coriolis force, the PGF causes the wind to bend the parcel toward the left and thus move in a circular, counterclockwise path, parallel to curved isobars around the low, as illustrated in Fig. 6.18c. Look at Fig. 6.19a and notice that above the level of frictional influence, the winds blow clockwise around an area of high pressure. The same spacing of the isobars tells us that the magnitude of the PGF is the same as in

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FFIGURE IGURE 6.19 Winds and related forces around an area of high prespres sure above the friction level in the Northern Hemisphere. Notice that the pressure gradient force (PGF) is in red, while the Coriolis force (CF) is in blue.

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Fig. 6.18a. Suppose at position 2 in Fig. 6.19a the wind is in geostrophic balance with the PGF directed outward and the Coriolis force directed inward. If the wind were in geostrophic balance, the air would move in a straight line parallel to straight-line isobars. But the isobars are curved. Let’s further suppose that the air parcel moves from position 2 directly southward to position 3. We can see in Fig. 6.19b that at position 3 the PGF crosses the isobars toward the southeast, producing a southerly component of the PGF, which increases the wind speed just a bit. The increase in wind speed increases the magnitude of the Coriolis force, thereby bending the parcel to the right, which causes the wind to blow clockwise, parallel to the isobars in a circular path around the high.* The greater Coriolis force around the high as compared to around the low results in an interesting relationship. Because the Coriolis force (at any given latitude) can increase only when the wind speed increases, we can see that for the same pressure gradient (the same spacing

WINDS ON UPPER-LEVEL CHARTS On the upper-level 500-mb map ( Fig. 6.20), notice that, as we would expect, the winds tend to parallel the contour lines in a wavy west-to-east direction. Notice also that the contour lines tend to decrease in height from south to north. This situation occurs because the air at this level is warmer to the south and colder to the north. On the map, where horizontal temperature contrasts are large, there is also a large height gradient—the contour lines are close together and the winds are strong. Where the horizontal temperature contrasts are small, there is a small height gradient—the contour lines are spaced farther apart and the winds are weaker. In general, on maps such as this we find stronger north-to-south temperature contrasts in winter than in summer, which is why the winds aloft are usually stronger in winter. In Fig. 6.20, the wind is geostrophic where it blows in a straight path, parallel to evenly spaced contour lines; it

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*Earlier in this chapter we learned that an object accelerates when there is a change in its speed or direction (or both). Therefore, the gradient wind blowing around the low-pressure center is constantly accelerating because it is constantly changing direction. This acceleration, called the centripetal acceleration, is directed at right angles to the wind, inward toward the low center. Remember from Newton’s second law that, if an object is accelerating, there must be a net force acting on it. In this case, the net force acting on the wind must be directed toward the center of the low, so that the air will keep moving in a circular path. This inward-directed force is called the centripetal force (centri: “center;” petal: “to push toward”). In some cases, it is more convenient to express the centripetal force (and the centripetal acceleration) as the centrifugal force, an apparent force that is equal in magnitude to the centripetal force, but directed outward from the center of rotation.

of the isobars), the winds around a high-pressure area (or a ridge) must be greater than the winds around a low-pressure area (or a trough). Typically, however, this difference in speed is overshadowed by the stronger winds that occur around low-pressure areas (cyclonic storms) because of the closer spacing of the isobars and stronger pressure gradients associated with them. In the Southern Hemisphere, the pressure gradient force starts the air moving and the Coriolis force deflects the moving air to the left, thereby causing the wind to blow clockwise around lows and counterclockwise around highs. So far we have seen how winds blow in theory, but how do they appear on an actual map?

FFIGURE 6.20 An upper-level 500-mb map showing wind direction, as indicated by lines that parallel the wind. Wind speeds are indicated by barbs and flags. (See the blue insert.) Solid gray lines are contours in meters above sea level. Dashed red lines are iisotherms in °C.

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FFIGURE 6.21 (a) The effect of surface friction is to slow down the wind so that, near the ground, the wind crosses the isobars and blows toward lower pressure. (b) This phenomenon at the surface produces an inflow of air around a low and an outflow of air around a h high. Aloft, the winds blow parallel to the lines, usually in a wavy west-to-east pattern. Both diagram (a) and (b) are in the Northern Hemisphere.

is gradient where it blows parallel to curved contour lines. Where the wind flows in large, looping meanders, following a more or less north-south trajectory (as it does in Fig. 6.20 off the west coast of North America), the wind-flow pattern is called meridional. Where the winds are blowing in a west-to-east direction (as seen in Fig. 6.20 over the eastern third of the United States), the flow is termed zonal. Because the winds aloft in middle and high latitudes generally blow from west to east, planes flying in this direction have a beneficial tail wind, which explains why a flight from San Francisco to New York City takes, on average, about 30 to 45 minutes less than the return flight. If the flow aloft is zonal, clouds, storms, and surface anticyclones tend to move more rapidly from west to east. However, where the flow aloft is meridional, as we will see in Chapter 8, surface storms tend to move more slowly, often intensifying into major storm systems. We know that because of the pressure gradient force and the Coriolis force (which, as we have seen, bends moving air to the right in the Northern Hemisphere), the winds aloft in the middle latitudes of the Northern Hemisphere tend to blow in a west-to-east pattern. Because the Coriolis force bends moving air to the left in the Southern Hemisphere, does this situation mean that the winds aloft in the Southern Hemisphere blow from east to west? The answer to this question is given in Focus section 6.4. Take a minute and look back at Fig. 6.13 on p. 157 and observe that the winds on the surface map tend to cross the isobars, blowing from higher pressure toward lower pressure. Observe also that the tightly packed isobars are indicating strong winds. Yet this same distribution of pressure (same pressure gradient) and same temperature on an upper-level chart would produce even stronger winds. Why, then, do surface winds normally cross the isobars and why do they blow more slowly than the winds aloft? The answer to both of these questions is friction.

SURFACE WINDS Because of surface friction, winds on a surface weather map do not blow exactly parallel to the isobars; instead, they cross the isobars, moving from higher to lower pressure. The angle at which the wind crosses the isobars varies, but averages about 30°. The frictional drag of the ground slows the wind down. Because the effect of friction decreases as we move away from Earth’s surface, wind speeds tend to increase with height above the ground. The atmospheric layer that is influenced by friction, called the friction layer (or planetary boundary layer), usually extends upward to an altitude near 1000 m or 3000 ft above the surface, but this altitude can vary somewhat since both strong winds and rough terrain can extend the region of frictional influence. In Fig. 6.21a, the wind aloft is blowing at a level above the frictional influence of the ground. At this level, the wind is approximately geostrophic and blows parallel to the isobars with the horizontal pressure gradient force (PGF) on its left balanced by the Coriolis force (CF) on its right. Notice, however, that at the surface the wind speed is slower. Apparently, the same pressure gradient force aloft will not produce the same wind speed at the surface, and the wind at the surface will not blow in the same direction as it does aloft. Near the surface, friction reduces the wind speed, which in turn reduces the Coriolis force. Consequently, the weaker Coriolis force no longer balances the pressure gradient force, and the wind blows across the isobars toward lower

DID YOU KNOW? The difference in air pressure from the base to the top of a New York skyscraper about 0.5 km (1600 ft) tall is typically near 55 millibars, a value much greater than the typical horizontal difference in air pressure between New York and Miami, Florida—a distance of more than 1800 km (1100 mi). AIR PRESSURE AND AND WINDS

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FOCUS

ON AN O OBSERVATION 6.4

Winds Aloft in the Southern Hemisphere FFIGURE 5 Upperlevel chart that extends over the Northern and Southern Hemispheres. Solid gray lines on the chart are isobars.

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In the Southern Hemisphere, just as in the Northern Hemisphere, the winds aloft blow because of horizontal differences in pressure. The pressure differences, in turn, are due to variations in temperature. Recall from an earlier discussion of pressure that warm air aloft is associated with high pressure and cold air aloft with low pressure. Look at Fig. 5. It shows an upper-level chart that extends from the Northern Hemisphere into the Southern Hemisphere. Over the equator where the air is warmer, the pressure aloft is high. North and south of the equator, where the air is colder, the pressure aloft is lower. Let’s assume, to begin with, that there is no wind on the chart. In the Northern HemiHemi sphere, the pressure gradient force directed northward starts the air moving toward lower pressure. Once the air is set in motion, the CoCo riolis force bends it to the right until it is a west wind wind, blowing parallel to the isobars.

In the Southern Hemisphere, the pressure gradient force directed southward starts the air moving south. But notice that the Coriolis force in the Southern Hemisphere bends the moving

pressure. The pressure gradient force is now balanced by the sum of the frictional force and the Coriolis force. Therefore, in the Northern Hemisphere, we find surface winds blowing counterclockwise and into a low; they flow clockwise and out of a high (see Fig. 6.21b). In the Southern Hemisphere, winds blow clockwise and inward around surface lows, counterclockwise and outward around surface highs. See Fig. 6.22, which shows a surface weather map and the general wind-flow pattern on a day in December in South America.

Winds and Vertical Air Motions Up to this point, we have seen that surface winds blow in toward the center of low pressure and outward away from the center of high pressure. As air moves inward toward the center of a low-pressure area (see Fig. 6.23), it must go somewhere. This converging air cannot go into the ground, so it slowly rises. Above the surface low (at about 6 km or so), the air begins to spread apart (diverge) to compensate for the converging surface air. As long as the upper-level diverging air balances the converging surface air, the central pressure in the low does not change. However, the surface pressure will change if upper-level divergence and surface convergence 164

air to its left, until the wind is blowing parallel to the isobars from the west. Hence, in the middle and high latitudes of both hemispheres, we generally find westerly winds aloft.

are not in balance. For example, as we saw earlier in this chapter, when we examined the air pressure above two cities, the surface air pressure will change if the mass of air above the surface changes. Consequently, if upper-level divergence exceeds surface convergence (that is, more air is removed at the top than is taken in at the surface), the air pressure at the center of the low will decrease, and isobars around the low will become more tightly packed. This situation increases the pressure gradient (and, hence, the pressure gradient force), which, in turn, increases the surface winds. Surface winds move outward (diverge), away from the center of a high-pressure area. Observe in Fig. 6.23 that to replace this laterally spreading surface air the air aloft converges and slowly descends. Again, as long as upperlevel converging air balances surface diverging air, the air pressure in the center of the high will not change. Should surface diverging air exceed upper-level converging air (more air removed at the surface than is brought in at the top), the air pressure at the center of the high will decrease, the pressure gradient force will weaken, and the surface winds will blow more slowly. (Convergence and divergence of air are so important to the development or weakening of surface pressure systems that we will examine this topic again in Chapter 8, when we look more closely at the vertical structure of pressure systems.)

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FFIGURE IGURE 6.22 (a) Surface weather map showing isobars and winds on a day in December in South America. (b) This boxed area shows the idealized flow around surface pressure systems in the Southern Hemisphere.

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The rate at which air rises above a low or descends above a high is small compared to the horizontal winds that spiral about these systems. Generally, the vertical motions are usually only about several centimeters per second, or about 1.5 km (or 1 mi) per day. Earlier in this chapter we learned that air moves in response to pressure differences. Because air pressure decreases rapidly with increasing height above the surface, there is always a strong pressure gradient force directed upward, much stronger than in the horizontal. Why, then, doesn’t the air rush off into space? Air does not rush off into space because the upwarddirected pressure gradient force is nearly always exactly balanced by the downward force of gravity. When these two forces are in exact balance, the air is said to be in hydrostatic equilibrium. When air is in hydrostatic equilibrium, there is no net vertical force acting on it,

FFIGURE 6.23 Winds and air motions associated with surface highs and lows in the Northern Hemisphere.

and so there is no net vertical acceleration. (Remember the example of the moving car in the footnote on p. 159.) Most of the time, the atmosphere approximates hydrostatic balance, even when air slowly rises or descends at a constant speed. However, this balance does not exist in violent thunderstorms and tornadoes, where the air shows appreciable vertical acceleration. Such storms occur over relatively small vertical distances, considering the total vertical extent of the atmosphere.

Determining Wind Direction and Speed Wind is characterized by its direction, speed, and gustiness. Because air is invisible, we cannot really see it. Rather, we see things being moved by it. Thus, we can determine wind direction by watching the movement of objects as air passes them. For example, the rustling of small leaves, smoke drifting near the ground, and flags waving on a pole all indicate wind direction. In a light breeze, a tried and true method for an observer to determine wind direction is to raise a wet finger into the air. The dampness quickly evaporates on the windward side, cooling the skin. Traffic sounds carried from nearby railroads or airports can also be used to help figure out the direction of the wind. Even your nose can alert you to the wind direction as the smell of fried chicken or broiled hamburgers drifts with the wind from a local restaurant. We already know that wind direction is given as the direction from which it is blowing—a north wind blows AIR PRESSURE AND AND WINDS

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FIGURE 6.24 An onshore wind blows from water to land, whereas an offshore wind blows from land to water.

from the north toward the south. However, near large bodies of water and in hilly regions, wind direction is expressed differently. For example, wind blowing from the water onto the land is referred to as an onshore wind (or onshore breeze), whereas wind blowing from land to water is called an offshore wind (or offshore breeze). ). (See Fig. 6.24.) Air moving uphill is an upslope wind; air moving downhill is a downslope wind. The wind direction can also be given as degrees about a 360° circle. These directions are expressed by the numbers shown in Fig. 6.25. For example: A wind direction of 360° is a north wind; an east wind is 90°; a south wind is 180°; and calm is expressed as zero. It is also common practice to express the wind direction in terms of compass points, such as N, NW, NE, and so on.

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THE INFLUENCE OF PREVAILING WINDS At many locations, the wind blows more frequently from one direction than from any other. The prevailing wind is the name given to the wind direction most often observed

FIGURE 6.25 Wind direction can be expressed in degrees about a circle or as compass points.

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during a given time period. Prevailing winds can greatly affect the climate of a region. For example, where the prevailing winds are upslope, the rising, cooling air makes clouds, fog, and precipitation more likely than where the winds are downslope. Prevailing onshore winds in summer carry moisture, cool air, and fog into coastal regions, whereas prevailing offshore breezes carry warmer and drier air into the same locations. In city planning, the prevailing wind is typically considered when planning where industrial centers, factories, and city dumps should be built. All of these, of course, must be located so that the wind will not carry pollutants into populated areas. Sewage disposal plants must be situated downwind from large housing developments, and major runways at airports must be aligned with the prevailing wind to assist aircraft in taking off or landing. In the high country, strong prevailing winds can bend and twist tree branches toward the downwind side, producing sculpted “flag” trees (see Fig. 6.26). The prevailing wind can even be a significant factor in building an individual home. In the northeastern half of the United States, the prevailing wind in winter is northwest and in summer it is southwest. Thus, to maximize comfort and energy efficiency, houses built in the northeastern United States should have windows facing southwest to provide summertime ventilation and few, if any, windows facing the cold winter winds from the northwest. The northwest side of the house should be thoroughly insulated and, perhaps, even protected by a windbreak. The prevailing wind can be represented by a wind rose (see Fig. 6.27), which indicates the percentage of time the wind blows from different directions. Extensions from the center of a circle point to the wind direction, and the length of each extension indicates the percentage of time the wind blew from that direction. Wind speed does not affect the length of each extension. However, the strength of the wind blowing from different directions is very helpful to know, so wind roses may include this information, as shown by the wind rose illustrated in Fig. 6.27. In this case, the prevailing wind is from the south, but winds from the west and

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© C. Donald Ahrens

FIGURE 6.26 In the high country of Colorado, these trees standing unprotected from the wind are often sculpted into “flag” trees.

north are almost as common. While this wind rose covers all four seasons and all times of day, a wind rose can be made for any particular time of the day, and it can represent the wind direction for any month or season of the year.

(see Fig. 6.30). These anemometers measure changes in ultrasonic signals that are produced when the wind blows across three sets of transmitters and receivers. The changes in signal speed are then converted into wind speeds in each of the three directions. The wind-measuring instruments described thus far are “ground-based” and only give wind speed or direction

© Jan Null

WIND INSTRUMENTS A very old, yet reliable, weather instrument for determining wind direction is the wind vane. Most wind vanes consist of a long arrow with a tail, which is allowed to move freely about a vertical post (see Fig. 6.28). The arrow always points into the wind and, hence, always gives the wind direction. Wind vanes can be made of almost any material. At airports, a cone-shaped bag, opened at both ends so that it extends horizontally as the wind blows through it, is situated near the runway. This form of wind vane, called a wind sock, enables pilots to tell the surface wind direction when landing. The instrument that measures wind speed is the anemometer. Cup anemometers consist of three (or more) hemispheric cups mounted on a vertical shaft as shown in Fig. 6.28. The difference in wind pressure from one side of a cup to the other causes the cups to spin about the shaft. The rate at which they rotate is directly proportional to the speed of the wind. The spinning of the cups is usually translated through a system of gears into wind speed, which can be read from a dial or transmitted to a recorder. The aerovane (skyvane) is an instrument that indicates both wind speed and direction. It consists of a bladed propeller that rotates at a rate proportional to the wind speed. Its streamlined shape and a vertical fin keep the blades facing into the wind (see Fig. 6.29). When attached to a recorder, a continuous record of both wind speed and direction can be obtained. Over the last few years, the national ASOS network of automated weather stations has been replacing the original cup anemometers with sonic anemometers

FIGURE 6.27 This wind rose represents the percentage of time the wind blew from various directions at the International Airport in Louisville, Kentucky, between 1977 and 2006. The most common directions are south (S) and north (N); the least common is northnorthwest (NNW). Each bar includes color categories showing how often the wind blew at a particular speed from a given direction. The wind was variable or calm about 9 percent of the time (not shown in the wind rose). (Midwest Regional Climate Center)

FIGURE 6.28 A wind vane and a cup anemometer. AIR PRESSURE AND AND WINDS

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© C. Donald Ahrens

FIGURE 6.29 The aerovane (skyvane).

© Jan Null

at a particular fixed location. But the wind is influenced by local conditions, such as buildings, trees, and so on. Also, wind speed normally increases rapidly with height above the ground. Thus, wind instruments should be exposed to freely flowing air well above the roofs of buildings. In practice, unfortunately, anemometers are placed at various levels often resulting in erratic wind observations. Thus, placement of anemometers is critical in order to avoid erratic and unrepresentative wind readings. Wind information can also be obtained by using instruments that either ascend or descend through the atmosphere. In the first example, a balloon rises from the surface carrying a radiosonde (an instrument package designed to measure the vertical profile of temperature, pressure, and humidity—see Chapter 1, p. 22). Equipment located on the ground constantly tracks the balloon, measuring its vertical and horizontal angles as well as its height above the ground. From this information, a computer determines and prints

FIGURE 6.30 The sonic anemometer shown here measures wind speed as part of the ASOS system.

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the vertical profile of wind from the surface up to where the balloon normally pops, typically in the stratosphere near 30 km or 100,000 ft. The observation of winds using a radiosonde balloon is called a rawinsonde observation. Rawindsonde data, taken from balloons launched around the world every 12 hours has been an important tool in tracking upper-level winds for more than 70 years. On special occasions, such as during field studies or when hurricanes are being monitored, an airplane is used to deploy an instrument package called a dropsonde, which falls to the ground via parachute. The resulting data are dropwindsonde observations. With the aid of an upward-pointing Doppler radar, a vertical profile of wind speed and direction up to an altitude of 16 km or so above the ground can be obtained. Such a profile is called a wind sounding, and the radar, a wind profiler (or simply a profiler). Doppler radar, like conventional radar, emits pulses of microwave radiation that are returned (backscattered) from various targets, in this case the irregularities in moisture and temperature created by turbulent, twisting eddies that move with the wind. Doppler radar works on the principle that, as these eddies move toward or away from the receiving antenna, the returning radar pulse will change in frequency. The Doppler radar wind profilers are so sensitive that they can translate the backscattered energy from these eddies into a vertical picture of wind speed and direction in a column of air 16 km (10 mi) thick.Wind profilers have been used at NASA’s Kennedy Space Center in Florida and across the Great Plains. Observations of upper-level winds are also made by satellites. As we saw in Chapter 4, geostationary satellites positioned above a particular location can show the movement of clouds, which is translated into wind direction and speed. Satellites now measure surface winds above the ocean by observing the roughness of the sea. An instrument called a scatterometer sends out a microwave pulse of energy that travels through the clouds, down to the sea surface. A portion of this energy is bounced back to the satellite. The amount of energy returning depends on the roughness of the sea: rougher seas return more energy. Since the sea’s roughness depends upon the strength of the wind blowing over it, the intensity of the returning energy can be translated into the surface wind speed and op direction, as illustrated in Fig. 6.31. Scatterometers operate aboard two European satellites called MetOp-A and MetOp-B, and a NASA scatterometer called RapidScat was placed aboard the International Space Station in 2014. The wind is an essential weather element that affects our environment in many ways. It can shape the landscape, transport material from one area to another, and generate ocean waves. It can also turn the blades of a windmill and blow down a row of trees. The reason the wind is capable of such feats is that, as the wind blows against an object, it exerts a force upon it. The amount of force exerted by

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FOCUS

ON A SSPECIAL TOPIC 6.5

For centuries, thousands of small windmills—their arms spinning in a stiff breeze—pumped water, sawed wood, and even supplemented the electrical needs of homes and farms around the world. It was not until the energy crisis of the early 1970s, however, that we seriously considered wind-driven turbines, called wind tur turbines, to run generators that produce elecbines tricity. Recently, the amount of wind energy in the United States has doubled every few years. As of 2015, only China had more capacity in its wind energy installations. Wind seems an attractive way of producing energy. It is nonpolluting and, unlike the sun, is not restricted to daytime use. A single commercial wind turbine, which may cost several million dollars to build and install, can generate power for hundreds of homes or businesses at a cost that is comparable in some areas to electricity generated by fossil fuels. However, the sight of large wind turbines is not to everyone’s liking. (Probably, though, it is no more of an eyesore than other types of energy-related infrastructure, such as oil pumps and electrical transmission towers.) Unfortunately, each year the blades of spinning turbines kill countless birds. To help remedy this problem, many wind turbine companies hire avian specialists to study bird behavior,

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Wind Energy

FIGURE 6 A wind farm near Palm Springs, California, generates electricity that is sold to Southern California.

and some turbines are actually shut down during nesting time. And the blades of modern high-capacity turbines turn more slowly, thereby helping birds to avoid them. If the wind turbine is to produce electricity, there must be wind—and not just any wind, but a flow of air neither too weak nor too strong. A slight breeze will not turn the blades, and a powerful wind gust could severely damage the machine. Thus, regions with the greatest potential for wind-generated power have moderate, steady winds. So much of the Great Plains of the United States is well-suited to wind power, so much so that the region has been dubbed “the Saudi Arabia of wind energy.”

the wind over an area increases as the square of the wind velocity. So when the wind velocity doubles, the force it exerts on an object goes up by a factor of four. To harness some of the wind’s energy and turn it into electricity, large

As of 2015, more than 48,000 wind turbines had the capacity of generating more than 74,000 megawatts of electricity in the United States, which is enough energy to supply the annual needs of more than 15 million homes. In California alone, there are thousands of wind turbines, many of which are on wind farms, clusters of 50 or more wind turbines (see Fig. 6). Wind provided more than 4 percent of the elecelec tricity needs in the United States in 2015, roughly double the percentage of 2010. The United States Department of Energy has estimated that wind energy using current technology could provide 20 percent of the nation’s electricity by 2030.

wind turbines and wind farms are becoming increasingly common around the world. More information on this topic is given in Focus section 6.5.

EUMETSAT

FFIGURE 6.31 A satellite image of wind speed and cloud-top temperature associated with Cyclone Ita on April 10, 2014. This in intense tropical cyclone eventually struck far northeast Australia (left side of image). The colored background shows the temperature of the cyclone’s cloud tops, which is correlated with the intensity of convection. Each wind barb represents 10 m/s (19 knots), with the satellite-observed winds reaching 50 m/s (95 knots) near the center of Ita. Wind speeds were gathered by the European Advanced Scatterometer ((ASCAT ASCAT)) aboard the MetOp-A satellite. ASCAT

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SUMMARY This chapter gives us a broad view of how and why the wind blows. Aloft where horizontal variations in temperature exist, there is a corresponding horizontal change in pressure. The difference in pressure establishes a force, the pressure gradient force (PGF), which starts the air moving from higher toward lower pressure. Once the air is set in motion, the Coriolis force bends the moving air to the right of its intended path in the Northern Hemisphere and to the left in the Southern Hemisphere. Above the level of surface friction, the wind is bent enough so that it blows nearly parallel to the isobars, or contours. Where the wind blows in a straight-line path, and a balance exists between the pressure gradient force and the Coriolis force, the wind is termed geostrophic. Where the wind blows parallel to curved isobars (or contours), the wind is called gradient wind. When the windflow pattern aloft is more west-to-east, the flow is called zonal; when the wind-flow pattern aloft is more northto-south or south-to-north, the flow is called meridional. The interaction of the forces causes the wind in the Northern Hemisphere to blow clockwise around regions of high pressure and counterclockwise around areas of low pressure. In the Southern Hemisphere, the wind blows counterclockwise around highs and clockwise around lows. The effect of surface friction is to slow down the wind. This causes the surface air to blow across the isobars from higher pressure toward lower pressure. Consequently, in both hemispheres, surface winds blow outward, away from the center of a high, and inward, toward the center of a low. We also looked at various methods and instruments used to determine and measure wind speed and direction.

KEY TERMS The following terms are listed (with corresponding page numbers) in the order they appear in the text. Define each. Doing so will aid you in reviewing the material covered in this chapter. air pressure, 148 millibar, 151 standard atmospheric pressure, 151 barometer, 151 mercury barometer, 151 aneroid barometer, 151 station pressure, 152 sea-level pressure, 152 170

isobar, 153 surface map, 153 anticyclone, 154 mid-latitude cyclonic storm, 154 isobaric map, 154 contour line, 154 ridge, 154 trough, 154

pressure gradient, 156 pressure gradient force, 156 Coriolis force, 157 geostrophic wind, 159 gradient wind, 161 meridional (wind flow), 163 zonal (wind flow), 163 friction layer, 163 hydrostatic equilibrium, 165

onshore wind, 166 offshore wind, 166 prevailing wind, 166 wind rose, 166 wind vane, 167 anemometer, 167 aerovane, 167 wind profiler, 168

QUESTIONS FOR REVIEW . Explain why atmospheric pressure always decreases with increasing altitude. . What might cause the air pressure to change at the bottom of an air column? . Why is the decrease of air pressure with increasing altitude more rapid when the air is cold? . What is considered standard sea-level atmospheric pressure in millibars? In inches of mercury? In hectopascals? . With the aid of a diagram, describe how a mercury barometer works. . How does an aneroid barometer measure atmospheric pressure? . How does sea-level pressure differ from station pressure? Can the two ever be the same? Explain. . Why will Denver, Colorado, always have a lower station pressure than Chicago, Illinois? . What are isobars? In what increment are they usually drawn on a surface weather map? . On an upper-level map, is cold air aloft generally associated with low or high pressure? What about warm air aloft? . What do Newton’s first and second laws of motion tell us? . What does a steep (or strong) pressure gradient mean? How would it appear on a surface map? . What is the name of the force that initially sets the air in motion and, hence, causes the wind to blow? . Explain why, on a map, closely spaced isobars (or contours) indicate strong winds, and widely spaced isobars (or contours) indicate weak winds. . What does the Coriolis force do to moving air (a) in the Northern Hemisphere? (b) in the Southern Hemisphere? . Explain how each of the following influences the Coriolis force: (a) wind speed (b) latitude.

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. Why do upper-level winds in the middle latitudes of both hemispheres generally blow from west to east? . What is a geostrophic wind? On an upper-level chart, how does it blow? . What are the forces that affect the horizontal movement of air? . Describe how the wind blows around highpressure areas and low-pressure areas aloft and near the surface (a) in the Northern Hemisphere; and (b) in the Southern Hemisphere. . How does zonal flow differ from meridional flow? . If the clouds overhead are moving from north to south, would the upper-level center of low pressure be to the east or west of you? . On a surface map, why do surface winds tend to cross the isobars and flow from higher pressure toward lower pressure? . Since there is always an upward-directed pressure gradient force, why doesn’t air rush off into space? . List as many ways as you can of determining wind direction and wind speed. . Below is a list of instruments. Describe how each one measures wind speed, wind direction, or both. (a) wind vane (b) cup anemometer (c) aerovane (skyvane) (d) radiosonde (e) satellite (f) wind profiler . An upper wind direction is reported as 225°. From what compass direction is the wind blowing?

QUESTIONS FOR THOUGHT AND EXPLORATI RA ON RATI .

The gas law states that pressure is proportional to temperature times density. Use the gas law to

. .

.

. . . . .

.

explain why a basketball seems to deflate when placed in a refrigerator. Can the station pressure ever exceed the sea-level pressure? Explain. The pressure gradient force causes air to move from higher pressures toward lower pressures (perpendicular to the isobars), yet actual winds rarely blow in this fashion. Explain why they don’t. The Coriolis force causes winds to deflect to the right of their intended path in the Northern Hemisphere, yet around a surface low-pressure area, winds blow counterclockwise, appearing to bend to their left. Explain why. Explain why, on a sunny day, an aneroid barometer would indicate “stormy” weather when carried to the top of a hill or mountain. Pilots often use the expression “high to low, look out below.” In terms of upper-level temperature and pressure, explain what this can mean. If Earth were not rotating, how would the wind blow with respect to centers of high and low pressure? Why are surface winds that blow over the ocean closer to being geostrophic than those that blow over the land? In the Northern Hemisphere, you observe surface winds shift from N to NE to E, then to SE. From this observation, you determine that a west-to-east moving high-pressure area (anticyclone) has passed north of your location. Describe with the aid of a diagram how you were able to come to this conclusion. As a cruise ship crosses the equator, the entertainment director exclaims that water in a tub will drain in the opposite direction now that the ship is in the Southern Hemisphere. Give two reasons to the entertainment director why this assertion is not so.

Go to the Wind Energy portal. Under the Primary Sources section, open the most recent “Year-End Market Report” from the American Wind Energy Association. Examine the map of wind power capacity by state. Which states show the most and least capacity? Do you think the differences are mainly due to the availability of strong, reliable winds? What other factors might be playing an important role?

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CHAPTER

7

Atmospheric Circulations Contents Scales of Atmospheric Motion Eddies—Big and Small

O

n December 28, 1997, a United nited Airlines’ Boeing 747 carrying 374 passengers was over the Pacific acific Ocean en route to

Hawaii awaii from Japan. About two hours into the flight, the aircraft was at a cruising altitude of 31,000 feet when suddenly, east of Tokyo, okyo, this routine, uneventful flight turned harrowing.

Local Wind Systems

Seat-belt eat-belt signs were turned on because of reports of severe air

Global Winds

turbulence nearby. The he plane hurtled upward, then quickly

Atmosphere-Ocean Interactions

Screaming, creaming, terrified passengers not fastened to their seats

dropped by about 30 meters (100 feet) before stabilizing. were flung against the walls of the aircraft, then dropped. Bags, serving trays, and luggage that slipped out from under the seats were flying about inside the plane. Within seconds, the entire ordeal was over. A total of 160 people were injured. Tragically, ragically, there was one fatality: A 32-year-old woman who had been hurled against the ceiling of the plane died of severe head injuries. What sort of atmospheric phenomenon could cause such turbulence?

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T

he aircraft in our opener on the previous page encountered a turbulent eddy—an “air pocket”—in perfectly clear weather. Such eddies are not uncommon, especially in the vicinity of jet streams. In this chapter, we will examine a variety of eddy circulations. First, we will look at the formation of small-scale winds. Then we will examine slightly larger circulations—local wind—such as the sea breeze and the chinook, describing how they form and the type of weather they generally bring. Finally, we will look at the general wind-flow pattern around the world.

Scales of Atmospheric Motion The air in motion—what we commonly call wind—is a powerful phenomenon. It is invisible, yet we see evidence of it nearly everywhere we look. It sculpts rocks, moves leaves, blows smoke, and lifts water vapor upward to where it can condense into clouds. The wind is with us wherever we go. On a hot day, it can cool us off; on a cold day, it can make us shiver. A breeze can sharpen our appetite when it blows the aroma from the local bakery in our direction. The wind is a powerful element. The workhorse of weather, it moves storms and large fair-weather systems around the globe. It transports heat, moisture, dust, insects, bacteria, and pollens from one area to another. Circulations of all sizes exist within the atmosphere. Little whirls form inside bigger whirls, which encompass even larger whirls—one huge mass of turbulent, twisting eddies.* For clarity, meteorologists arrange circulations according to their size. This hierarchy of motion from tiny gusts to giant storms is called the scales of motion. Consider smoke rising into otherwise clean air from a chimney in an industrial section of a large city

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*Eddies are spinning globs of air that have a life history of their own.

(see Fig. 7.1a). Within the smoke, small chaotic motions—tiny eddies—cause it to tumble and turn. These eddies constitute the smallest scale of atmospheric motion—the microscale. At the microscale level, eddies with diameters of a few meters or less not only disperse smoke, but they also sway branches and swirl dust and papers into the air. They form by convection or by the wind blowing past obstructions and are usually short-lived, lasting only a few minutes at best. In Fig. 7.1b, observe that, as the smoke rises, it drifts many kilometers downwind. This circulation of city air constitutes the next larger scale—the mesoscale ( “middle scale”). Typical mesoscale circulations range from a few kilometers to about a hundred kilometers in diameter. Generally, they last longer than microscale motions, often many minutes, hours, or in some cases as long as a day. Mesoscale circulations include local winds (which form along shorelines and mountains), as well as thunderstorms, tornadoes, and small tropical cyclones. When we look for the smokestack on a surface weather map (Fig. 7.1c), neither the smokestack nor the circulation of city air shows up. All that we see are the circulations around high- and low-pressure areas— the cyclones and anticyclones of the middle latitudes—as well as the large tropical cyclones of lower latitudes. We are now looking at the synoptic scale, or weather-map scale. Circulations of this magnitude dominate regions of hundreds to even thousands of square kilometers and, although the life spans of these features vary, they typically last for days and sometimes weeks. When we look at wind patterns over the entire Earth, we are looking at the global scale, or planetary scale. Together, the synoptic and global scales are referred to as the macroscale—the largest scale of atmospheric motion. Figure.7.2 summarizes the various scales of motion and their average life span.

FIGURE 7.1 Scales of atmospheric motion. The tiny microscale motions constitute a part of the larger mesoscale motions, FI which, in turn, are part of the much larger macroscale. Notice that as the scale becomes larger, motions observed at the smaller scale are no longer visible.

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© Cengage Learning®.

FIGURE 7.2 The scales of atmospheric motion with the phenomenon’s average size and life span. (Because the actual size of certain features can vary, some of the features fall into more than one category.)

crest. These so-called rotors have violent vertical motions that produce extreme turbulence and hazardous flying conditions. Strong winds blowing over a mountain in stable air can produce a mountain wave eddy on the downwind side, with a reverse flow near the ground. On a much smaller scale, the howling of wind on a blustery night is believed to be caused by eddies that are constantly being shed around obstructions, such as chimneys and roof corners. Turbulent eddies form aloft as well as near the surface. Turbulence aloft can occur suddenly and unexpectedly, especially where the wind changes its speed or direction (or both) abruptly. Such a change is called wind shear.

When the wind encounters a solid object, a whirl of air, or eddy, forms on the object’s downwind side.* The size and shape of the eddy often depend upon the size and shape of the obstacle and on the speed of the wind. Light winds produce small stationary eddies. Wind moving past trees, shrubs, and even your body produces small eddies. (You may have had the experience of dropping a piece of paper on a windy day only to have it carried away by a swirling eddy as you bend down to pick it up.) Air flowing over a building produces larger eddies that will, at best, be about the size of the building. Strong winds blowing past an open sports stadium can produce eddies that may rotate in such a way as to create surface winds on the playing field that move in a direction opposite to the wind flow above the stadium. Wind blowing over a fairly smooth surface produces few eddies, but when the surface is rough, many eddies form. The eddies that form downwind from obstacles can produce a variety of interesting effects. For instance, wind moving over a mountain range in stable air with a speed greater than 40 knots usually produces waves and eddies, such as those shown in Fig. 7.3. We can see that eddies form both close to the mountain and beneath each wave *The irregular, disturbed flow of gusty winds and eddies is called turbulence.

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Eddies—Big and Small

FIGURE 7.3 Under stable conditions, air flowing past a mountain range can create eddies many kilometers downwind of the mountain itself. ATMOSPHERIC CIRCULATIONS

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At the surface, the air pressure remains unchanged until the air aloft begins to move. As the air aloft moves from south to north, air leaves the southern area and “piles up” above the northern area. This redistribution of air reduces the surface air pressure to the south and raises it to the north. Consequently, a pressure gradient force is established at Earth’s surface from north to south and, hence, surface winds begin to blow from north to south. We now have a distribution of pressure and temperature and a circulation of air, as shown in Fig. 7.4c. As the cool surface air flows southward, it warms and becomes less dense. In the region of surface low pressure, the warm air slowly rises, expands, cools, and flows out the top at an elevation of about 1 km (3300 ft) above the surface. At this level, the air flows horizontally northward toward lower pressure, where it completes the circulation by slowly sinking and flowing out the bottom of the surface high. Circulations brought on by changes in air temperature, in which warmer air rises and colder air sinks, are termed thermal circulations. The regions of surface high and low atmospheric pressure created as the atmosphere either cools or warms are called thermal (cold core) highs and thermal (warm core) lows. In general, they are shallow systems, usually extending no more than a few kilometers above the ground.

DID YOU KNOW? On a blustery night, the howling of the wind can be caused by eddies. As the wind blows past chimneys and roof corners, small eddies form. These tiny swirls act like pulses of compressed air that ultimately reach your eardrum and produce the sound of howling winds.

The shearing creates forces that produce eddies along a mixing zone. If the eddies form in clear air, this form of turbulence is called clear air turbulence, or CAT. When an airplane is flying through such turbulence, the bumpiness can range from small vibrations to violent up-anddown motions that force passengers against their seats and toss objects throughout the cabin. (Additional information on this topic is given in Focus section 7.1.)

Local Wind Systems

THERMAL CIRCULATIONS Consider the vertical distribution of pressure shown in Fig. 7.4a. The isobars* all lie parallel to Earth’s surface; thus, there is no horizontal variavaria tion in pressure (or temperature), and there is no pressure gradient and no wind. Suppose in Fig. 7.4b the atmosphere is cooled to the north and warmed to the south. In the cold, dense air above the surface, the isobars bunch closer together, while in the warm, less-dense air, they spread farther apart. This dipping of the isobars produces a horizontal pressure gradient force (PGF) aloft that causes the air to move from higher pressure toward lower pressure. *The isobars depicted here actually represent a surface of constant pressure (an isobaric surface) rather than a line, or isobar. Information on isobaric surfaces is given in Focus section 6.2 on p. 155.

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Every summer, millions of people flock to the New Jersey shore, hoping to escape the oppressive heat and humidity of the inland region. On hot, humid afternoons, these travelers often encounter thunderstorms about twenty miles or so from the ocean, thunderstorms that invariably last for only a few minutes. In fact, by the time the vacationers arrive at the beach, skies are generally clear and air temperatures are much lower, as cool ocean breezes greet them. If the travelers return home in the afternoon a few days later, these “mysterious” showers often occur at just about the same location as before. The showers are not really mysterious, of course. They are caused by a local wind system, the sea breeze. As cooler ocean air pours inland, it forces the warmer (less dense), unstable humid air to rise and condense, producing majestic clouds and rainshowers along a line that separates the contrasting temperatures. The sea breeze forms as part of a thermally driven circulation, so we will begin our study of local winds by examining the formation of thermal circulations.

FIGURE 7.4 A thermal circulation produced by the heating and cooling of the atmosphere near the ground. The H’s and L’s refer to atmospheric pressure. The lines represent surfaces of constant pressure (isobaric surfaces), where 1000 is 1000 millibars.

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FOCUS

ON AN O OBSERVATION 7.1

Eddies and “Air Pockets” extend upward into the stratosphere. As we learned earlier, when these huge eddies develop in clear air, this form of turbulence is referred to as clear air turbulence, or CAT. CAT The eddies that form in clear air can have diameters ranging from a couple of meters to several hundred meters. An unsuspecting aircraft entering such a region can be in for more than just a bumpy ride. If the aircraft flies into a zone of descending air, it can drop suddenly, producing the sensation that there is no air to support the wings. Consequently, these regions have come to be known as air pockets. Commercial aircraft entering air pockets have dropped hundreds of meters, injuring passengers and flight attendants not strapped into their seats. Five passengers and crew members had to be hospitalized in February 2014 after severe turbulence struck a Boeing 737 jetliner as it approached Billings, Montana. In April 1981, a DC-10 jetliner flying at 11,300 m (37,000 ft)

over central Illinois encountered a region of severe clear air turbulence and reportedly plunged about 600 m (2000 ft) toward Earth before stabilizing. Twenty-one of the 154 people aboard were injured; one person sustained a fractured hip and another person, after hitting the ceiling, jabbed himself in the nose with a fork, then landed in the seat in front of him.* Clear air turbulence has occasionally caused structural damage to aircraft by breaking off vertical stabilizers and tail structures. Fortunately, the effects are usually not this dramatic. The potential adverse effects of clear air turbulence is one important reason why passengers are frequently told to “fasten your seat belts” while flying, even when there are no thunderstorms or obvious hazards in sight.

*Another example of an aircraft that experienced severe turbulence as it flew into an air pocket is given in the opening vignette on p. 173.

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To better understand how eddies form along a zone of wind shear, imagine that, high in the atmosphere, there is a stable layer of air having vertical wind speed shear (changing wind speed with height) as depicted in Fig. 1a. The top half of the layer slowly slides over the bottom half, and the relative speed of both halves is low. As long as the wind shear between the top and bottom of the layer is small, few if any eddies form. However, if the shear and the corresponding relative speed of these layers increases (Figs. 1b and 1c), wavelike unduundu lations may form. When the shearing exceeds a certain value, the waves break into large swirls, with significant vertical movement (Fig. 1d). Eddies such as these often form in the upper troposphere near jet streams, where large wind speed shears exist. If wavelike clouds form in the region of wind shear, they are often called billow clouds (see Fig. 2). Turbulent eddies also occur in conjuncconjunc tion with mountain waves, which may

FIGURE 1 The formation of clear air turbulence (CAT) along a boundary of increasing wind speed shear. The wind in the top layer FI increases in speed from (a) through (d) as it flows over the bottom layer.

© C. Donald Ahrens

FIGURE 2 Billow clouds forming in a region of rapidly changing wind speed, called wind shear.

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SEA AND LAND BREEZES The sea breeze is a type of thermal circulation. The uneven heating rates of land and water (described in Chapter 3) cause these mesoscale coastal winds. During the day, the land heats more quickly than the adjacent water, and the intensive heating of the air above produces a shallow thermal low. The air over the water remains cooler than the air over the land; hence, a shallow thermal high exists above the water. The overall effect of this pressure distribution is a sea breeze that blows at the surface from the sea toward the land as illustrated in Fig. 7.5. Since the strongest gradients of temperature and pressure occur near the land-water boundary, the strongest winds typically occur right near the beach and diminish inland. Further, since the greatest contrast in temperature between land and water usually occurs in the afternoon, sea breezes are strongest at this time. (The same type of breeze that develops along the shore of a large lake is called a lake breeze.) At night, the land cools more quickly than the water, and the air above the land becomes cooler than the air over the water, producing a distribution of pressure such as the one shown in Fig. 7.5b. With higher surface pressure now

over the land, the wind reverses itself and becomes a land breeze, a breeze at the surface that flows from the land toward the water. Temperature contrasts between land and water are generally much smaller at night; hence, land breezes are usually weaker than their daytime counterpart, the sea breeze. In regions where greater nighttime temperature contrasts exist, stronger land breezes occur over the water, off the coast. They are not usually noticed much onshore, but are frequently observed by ships in coastal waters. Look at Fig. 7.5 again and observe that the rising air is over the land during the day and over the water during the night. Therefore, along the humid east coast of the United States, daytime clouds tend to form over land and nighttime clouds over water. This explains why, at night, distant lightning flashes are sometimes seen over the ocean. The leading edge of the sea breeze is called the sea breeze front. As the front moves inland, a rapid drop in temperature usually occurs just behind it. In some locations, this temperature change can be 5°C (9°F) or more during the first hours—a refreshing experience on a hot, sultry day. In regions where the water temperature is warm, the cooling effect of the sea breeze is hardly evident. Since cities near

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FIGURE 7.5 Development of a sea FI breeze and a land breeze. (a) At the surface, a sea breeze blows from the water onto the land, whereas (b) the land breeze blows from the land out over the water. Notice that the pressure at the surface changes more rapidly with the sea breeze. This situation indicates a stronger pressure gradient force and higher winds with a sea breeze.

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FIGURE 7.6 Typically, during the summer over Florida, convergFI ing sea breezes in the afternoon produce uplift that enhances thunderstorm development and rainfall. However, when westerly surface winds dominate and a ridge of high pressure forms over the area, thunderstorm activity diminishes, and dry conditions prevail.

the ocean usually experience the sea breeze by noon, their highest temperature usually occurs much earlier than in inland cities. Along the east coast of North America, the passage of the sea breeze front is marked by a wind shift, usually from west to east. In the cool ocean air, the relative humidity rises as the temperature drops. If the relative humidity increases beyond about 70 percent, water vapor begins to condense upon particles of sea salt or industrial smoke, producing haze. When the ocean air is highly concentrated with pollutants, the sea breeze front may meet relatively clear air and thus appear as a smoke front, or a smog front. If the ocean air becomes saturated, a mass of low clouds and fog will mark the leading edge of the marine air. When there is a sharp contrast in air temperature across the frontal boundary, the warmer, lighter air will converge and rise. In many regions, this makes for good sea breeze glider soaring. If this rising air is sufficiently moist, a line of cumulus clouds will form along the sea

breeze front, and, if the air is also conditionally unstable, thunderstorms may form. As previously mentioned, on a hot, humid day one can drive toward the shore, encounter heavy showers several miles from the ocean, and arrive at the beach to find it sunny with a steady onshore breeze. When cool, dense, stable marine air encounters an obstacle, such as a row of hills, the heavy air tends to flow around them rather than over them. When the opposing breezes meet on the opposite side of the obstruction, they form what is called a sea breeze convergence zone. Such conditions are common along the mountainous Pacific coast of North America. Sea breezes in Florida help produce that state’s abundant summertime rainfall. On the Atlantic side of the state, the sea breeze blows in from the east; on the Gulf con shore, it moves in from the west (see Fig. 7.6). The convergence of these two moist wind systems, coupled with daytime convection, produces cloudy conditions and showery weather over the land (see Fig. 7.7). Over the water (where cooler, more stable air lies close to the sursur face), skies often remain cloud-free. On many days during June and July 1998, however, Florida’s converging wind system did not materialize. The lack of converging surface air and its accompanying showers left much of the state parched. Huge fires broke out over northern and central Florida, which left hundreds of people homeless and burned many thousands of acres of grass and woodlands. A weakened sea breeze and dry conditions have produced wildfires on numerous other occasions, including during the spring of 2006. Convergence of coastal breezes is not restricted to ocean areas. Both Lake Michigan and Lake Superior are capable of producing well-defined lake breezes. In upper Michigan, where these large bodies of water are separated by a narrow strip of land, the two breezes push inland and converge near the center of the peninsula, creating afternoon clouds and showers, while the lakeshore areas remains sunny, pleasantly cool, and dry.

© T. Ansel Toney

FIGURE 7.7 Surface heating and FI lifting of air along a converging sea breeze combine to form thunderstorms almost daily during the summer in southern Florida.

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FIGURE 7.8 Valley FI breezes blow uphill dur during the day; mountain breezes blow downhill at night. (The L’s and H’s represent pressure, whereas the purple lines represent surfaces of constant pressure.)

their maximum strength in the early afternoon, cloudiness, showers, and even thunderstorms are common over mountains during the warmest part of the day—a fact well known to seasoned hikers and climbers. KATABATIC WINDS Although any downslope wind is technically a katabatic wind, the name is usually reserved for downslope winds that are much stronger than mountain breezes. Katabatic (or fall) winds can rush down elevated slopes at hurricane speeds, but most are not that intense and many are on the order of 10 knots or less. The ideal setting for a katabatic wind is an elevated plateau surrounded by mountains, with an opening that slopes rapidly downhill (see Fig. 7.10). When winter snows accumulate on the plateau, the overlying air grows extremely cold. Along the edge of the plateau the cold, dense air begins to descend through gaps and saddles in the hills, usually as a gentle or moderate cold breeze. If

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MOUNTAIN AND VALLEY BREEZES Mountain and valley breezes develop along mountain slopes. Observe in Fig. 7.8 that, during the day, sunlight warms the valley walls, which in turn warm the air in contact with them. The heated air, being less dense than the air of the same altitude above the valley, rises as a gentle upslope wind known as a valley breeze. At night, the flow reverses. The mountain slopes cool quickly, chilling the air in contact with them. The cooler, more-dense air glides downslope into the valley, providing a mountain breeze. (Because gravity is the force that directs these winds downhill, they are also referred to as gravity winds, or nocturnal drainage winds.) This daily cycle of wind flow is best developed in clear, summer weather when prevailing winds are light. When the upslope valley winds are well developed and have sufficient moisture, they can reveal themselves through cumulus clouds that build above mountain summits (see Fig. 7.9). Since valley breezes usually reach

FIGURE 7.9 As mountain slopes warm during the day, air rises and often condenses into cumuliform clouds, such as the ones shown FI here.

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*Information on geographic features and their location in North America is provided at the back of the book.

FIGURE 7.10 Strong katabatic winds can form where cold FI winds rush downhill from an elevated plateau covered with snow.

air downslope. As the air descends, it is compressed and warms. So the main source of warmth for a chinook is compressional heating, as potentially warmer (and drier) air is brought down from aloft. Clouds and precipitation on the mountain’s windward side can enhance the chinook. For example, as the cloud forms on the upwind side of the mountain in Fig. 7.11, the release of latent heat inside the cloud supplesupple ments the compressional heating on the downwind side. This phenomenon makes the descending air at the base of the mountain on the downwind side warmer than it was before it started its upward journey on the windward side. The air is also drier, since much of its moisture was removed as precipitation on the windward side. (More information on temperature changes associated with chinooks is given in Focus section 7.2.) Along the Front Range of the Rockies, a bank of clouds forming on the windward side of the mountains is a telltale sign of an impending chinook. This cloud feature, called a foehn wall, usually remains stationary as air rises, condenses, and then rapidly descends the leeward slopes, often causing strong winds in foothill communities. These

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CHINOOK (FOEHN) WINDS The chinook wind is a warm, dry, downslope wind that descends the eastern slope of the Rocky Mountains. The region of the chinook is rather narrow and extends from northeastern New Mexico northward into Canada. Similar winds occur along the leeward slopes of mountains in other regions of the world. The general term for chinook wind is the foehn (a name that originated in the European Alps), but there are many local names, such as the zonda in Argentina. When foehn winds move through an area, the temperature rises sharply, sometimes 20°C (36°F) or more in less than an hour, and a corresponding sharp drop in the relative humidity occurs, occasionally to less than 5 percent. In North America, chinooks occur when strong westerly winds aloft flow over a north-south–trending mountain range, such as the Rockies and Cascades. Such conditions can produce a trough of low pressure on the mountain’s eastern side, a trough that tends to force the

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the breeze, however, is confined to a narrow canyon or channel, the flow of air can increase, often destructively, as cold air rushes downslope like water flowing over a fall. Katabatic winds are observed in various regions of the world. For example, along the northern Adriatic coast in the former Yugoslavia, a polar invasion of cold air from Russia descends the slopes from a high plateau and reaches the lowlands as the bora, a cold, gusty, northeasterly wind with speeds sometimes in excess of 100 knots. A similar, but often less violent, cold wind known as the mistral descends the western mountains into the Rhone Valley of France, and then out over the Mediterranean Sea. It frequently causes frost damage to exposed vineyards and makes people bundle up in the otherwise mild climate along the Riviera. Strong, cold katabatic winds also blow downslope off the ice sheets in Greenland and Antarctica, occasionally with speeds greater than 100 knots. In North America, when cold air accumulates over the Columbia Plateau of Idaho, Oregon, and Washington,* it may flow westward through the Columbia River Gorge as a strong, gusty, and sometimes violent wind. Even though the sinking air warms by compression, it is so cold to begin with that it reaches the ocean side of the Cascade Mountains much colder than the marine air it replaces. The Columbia Gorge wind (called the coho) is often the harbinger of a prolonged cold spell. Strong downslope katabatic-type winds funneled through a mountain canyon can do extensive damage. In January 1984, a ferocious downslope wind blew through Yosemite National Park in California at speeds estimated at 100 knots. The wind toppled trees and, unfortunately, caused a fatality when a tree fell on a park employee sleeping in a tent.

FIGURE 7.11 A chinook wind can be enhanced when clouds FI form on the mountain’s windward side. Heat added and moisture lost on the upwind side produce warmer and drier air on the downwind side. ATMOSPHERIC CIRCULATIONS

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FOCUS

ON A SSPECIAL TOPIC 7.2

Chinooks are thirsty winds. As they move over a heavy snow cover, they can melt and evaporate a foot of snow in less than a day. This situation has led to some tall tales about these so-called “snow eaters.” Canadian folklore has it that a sled-driving traveler once tried to outrun a chinook. During the entire ordeal, as the story has it, his front runners were supposedly in snow while his back runners were on bare soil. Actually, the chinook is important economically. It not only brings relief from the winter cold, but it also uncovers prairie grass, so that livestock can graze on the open range. Also, these warm winds can help keep railroad tracks clear of snow, so that trains can keep running. On the other hand, the drying effect of a chinook can create an extreme fire hazard. And when a chinook follows spring planting, the seeds can die in the parched soil. Along with the dry air comes a buildup of static electricity, making a simple handshake a shocking experience. These warm, dry winds have sometimes adversely affected human behavior. During periods of chinook winds some people feel irritable and depressed and others become ill. The exact reason for this phenomenon is not clearly understood. Chinook winds have been associated with very rapid temperature changes. On

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Snow Eaters and Rapid Temperature Changes

FIGURE 3 Cities near the warm air–cold air boundary can experience sharp FI temperature changes if cold air should slosh back and forth like water in a bowl.

January 11, 1980, the air temperature in Great Falls, Montana, rose from –32°F to 17°F (a 49°F rise in temperature) in just seven minutes. The nation’s largest temperature swing within a 24-hour period occurred in Loma, Montana, on January 15, 1972, when the temperature soared from –54°F to 48°F. How such rapid changes in temperature can occur is illustrated in Fig. 3. Notice that a shallow layer of extremely cold air has moved out of Canada and is now resting against the Rocky Mountains. The cold air behaves just as any fluid, and, in some cases, atmospheric conditions can cause the air to slosh up and down much like water does when a bowl is rocked back and forth. This rocking motion

strong winds are especially notorious in winter in Boulder, Colorado, where windstorms have caused millions of dollars in damage. Figure 7.12 shows how a foehn wall appears as one looks west toward the Rockies from the Colorado plains. The photograph was taken on a winter afternoon with the air temperature about –7°C (20°F). That evening, the chinook moved downslope at high speeds through foothill valleys, picking up sand and pebbles (which dented cars and cracked windshields). The chinook spread out over the plains like a warm blanket, raising the air temperature the following day to a mild 15°C (59°F). The chinook and its wall of clouds remained for several days, bringing with it a welcomed break from the cold grasp of winter. SANTA ANA WINDS A warm, dry wind that blows downhill from the east or northeast into southern California is the Santa Ana wind. As the air descends from the elevated desert plateau, it funnels through 182

can cause extreme temperature variations for cities located at the base of the hills along the periphery of the cold air–warm air boundary, as they are alternately in and then out of the cold air. Such a situation is probably responsible for the extremely rapid two-minute temperature change of 49°F recorded at Spearfish, South Dakota, during the morning of January 22, 1943. On the same morning, in nearby Rapid City, the temperature fluctuated from –4°F at 5:30 A .M. to 54°F at 9:40 A .M., then down to 11°F at 10:30 A .M. and up to 55°F just 15 minutes later. At nearby cities, the undulating cold air produced similar temperature variations that recurred over several hours.

mountain canyons in the San Gabriel and San Bernardino Mountains, finally spreading over the Los Angeles Basin and San Fernando Valley and out over the Pacific Ocean (see Fig. 7.13). The wind often blows with exceptional speed—occasionally over 90 knots— in the Santa Ana Canyon (the canyon from which it derives its name). These warm, dry winds develop as a region of high pressure builds over the Great Basin. The clockwise circulation around the anticyclone forces air downslope from the high plateau. Thus, compressional heating provides the primary source of warming. The air is dry, since it originated in the desert, and it dries out even more as it is heated. Figure 7.14 shows a typical wintertime Santa Ana situation. As the wind rushes through canyon passes, it lifts dust and sand and dries out vegetation, which sets the stage for serious brush fires, especially in autumn, when

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© C. Donald Ahrens

FIGURE 7.12 A foehn wall forming over the Colorado Rockies (viewed from the plains). FI

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FIGURE 7.13 Warm, dry Santa FI Ana winds sweep downhill through mountain canyons into Southern California. The large H represents higher air pressure over the elevated desert.

chaparral-covered hills are already parched from the dry summer.* One such fire in November 1961—the infamous Bel Air fire—burned for three days, destroying 484 homes and causing over $25 million in damage in 1961 dollars (close to $200 million in today’s dollars). During October 2003, massive wildfires driven by strong Santa Ana winds swept through Southern California. The fires charred more than 750,000 acres, destroyed over 2800 homes, took 20 lives, and caused over $2 billion in property damage. Only four years later (after one of the driest years on record) in October 2007, wildfires broke out again in

FIGURE 7.14 Surface weather map showing Santa Ana condiFI tions in January. Maximum temperatures for this particular day are given in °F. Observe that the downslope winds blowing into Southern California raised temperatures into the upper 80s, while elsewhere temperature readings were much lower.

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*Chaparral denotes a shrubby environment, in which many of the plant species contain highly flammable oils.

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NASA/MODIS Rapid Response

rains wash away topsoil and, in some areas, create serious mudslides. The adverse effects of a wind-driven Santa Ana fire may be felt long after the fire itself has been put out.

FIGURE 7.15 Satellite view showing strong northeasterly Santa FI Ana winds on October 23, 2007, blowing smoke from massive wild fires (red dots) across Southern California and out over the Pacific Ocean.

Southern California. Pushed on by hellacious Santa Ana winds that gusted to over 80 knots, the fires raced through dry vegetation, scorching virtually everything in their paths. The fires, which extended from north of Los Angeles to the Mexican border (see Fig. 7.15), burned over 500,000 acres, destroyed more than 1800 homes, and took 9 lives. The total costs of the fires exceeded $1.5 billion. Four hundred miles to the north in Oakland, California, a ferocious Santa Ana–type wind was responsible for the disastrous Oakland hills fire during October 1991, which damaged or destroyed over 3000 dwellings, caused over $1.5 billion in damage, and took 25 lives protec (see Fig. 7.16). When a Santa Ana fire removes the protective cover of vegetation, the land is ripe for erosion, as winter

DESERT WINDS Winds of all sizes develop over the desert. Huge dust storms form in dry regions, where strong winds are able to lift and fill the air with particles of fine dust. In February 2001, an exceptionally large dust storm (about the size of Spain) formed over the African Sahara and then swept westward off the African coast, then northeastward for thousands of miles. During the drought years of the 1930s, large dust storms formed over the Great Plains of the United States. Some individual storms lasted for three days and spread dust for hundreds of miles to the east, to the Atlantic coast and beyond. In desert areas where loose sand is more prevalent, sandstorms develop, as high winds enhanced by surface heating rapidly carry sand particles close to the ground. A spectacular example of a storm composed of dust or sand is the haboob (from Arabic habb: “to blow”). The haboob forms as cold downdrafts along the leading edge of a thunderstorm lift dust or sand into a huge, tumbling dark cloud that can extend horizontally for more than 100 kilometers and rise vertically to the base of the thunderstorm. Haboobs are most common in the African Sudan (where about twenty-four occur each year) and in the Desert Southwest of the United States, especially in southern Arizona. A particularly strong haboob swept into Phoenix, Arizona, in July 2011 (see Fig. 7.17), moving vast amounts of dust from a landscape parched by drought. On a smaller scale, in dry areas, the wind can also produce rising, spinning columns of air that pick up dust or sand from the ground. Called dust devils or whirlwinds,* these rotating vortices generally form on clear, hot days over a dry surface where most of the sunlight goes into *In Australia, the Aboriginal word willy-willy refers to a dust devil.

© Jim Pire

FIGURE 7.16 From atop his roof, a FI resident of Oakland’s Rockridge district looks on in disbelief as his neighbors’ homes are consumed in a raging firestorm on October 20, 1991.

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AP Photo/Amanda Lee Myers

FIGURE 7.17 A large haboob FI (dust storm) moves through Phoenix, Arizona, on July 5, 2011.

Having diameters of only a few meters and heights usually of less than 100 m (330 ft), most dust devils last only a short time (see Fig. 7.19). However, some dust devils reach sizable dimension, extending upward from the surface for several hundred meters. Such whirlwinds are capable of considerable damage; winds exceeding 75 knots can overturn mobile homes and tear the roofs off buildings. Fortunately, the majority of dust devils are small. Also keep in mind that dust devils are not tornadoes. The circulations of most tornadoes (as we will see in Chapter 10) develop near the base of a thunderstorm,

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heating the surface, rather than into evaporating water from vegetation. The atmosphere directly above the hot surface becomes unstable, convection sets in, and the heated air rises, often lifting dust, sand, and dirt high into the air. Wind, often deflected by small topographic barriers, flows into this region, rotating the rising air as depicted in Fig. 7.18. Depending on the nature of the topographic feature, the spin of a dust devil around its central core may be cyclonic or anticyclonic, and both directions occur with about equal frequency. (Dust devils are too small and fleeting to be substantially influenced by the Coriolis force.)

FIGURE 7.18 The formation of a dust devil. On a hot, dry day, the atmosphere next to the ground becomes unstable. As the heated FI air rises, wind blowing past an obstruction twists the rising air, forming a rotating column, or dust devil. Air from the sides rushes into the rising column, lifting sand, dust, leaves, or any other loose material from the surface. ATMOSPHERIC CIRCULATIONS Copyright 2018 Cengage Learning. All Rights Reserved. May not be copied, scanned, or duplicated, in whole or in part. WCN 02-200-202

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© Nilton Renno

FIGURE 7.19 A well-developed dust devil moves over a hot FI desert landscape on a clear summer day.

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whereas the circulation of a dust devil begins at the surface, normally in sunny weather (although some dust devils form beneath convective-type clouds). Desert winds are not confined to planet Earth; they form on the planet Mars as well. Most of the Martian dust storms are small, and only cover a relatively small portion of that planet. However, some dust storms can actually grow large enough to encircle Mars with a dusty haze, as happened in 2001. Dust devils also form on Mars when high winds sweep over uneven terrain.

SEASONALLY CHANGING WINDS—THE MONSOON On our planet there are thermal circulations that are much larger than those of the more local sea and land breezes described earlier. A good example of such a circulation is the monsoon, which derives from the Arabic mausim (“seasons”).A monsoon wind system is one that changes direction seasonally, blowing from one direction in summer and from the opposite direction in winter. This seasonal reversal of winds is especially well developed in eastern and southern Asia. In some ways, the monsoon is similar to a large-scale sea breeze. During the winter, the air over northern Asia becomes much colder than the air over the ocean. A large, shallow high-pressure area develops over continental Siberia, producing a clockwise circulation of air that flows out over the Indian Ocean and South China Sea (see Fig. 7.20a). Subsiding air of the anticyclone and the downslope movement of northeasterly winds from the inland plateau provide eastern and southern Asia with generally fair weather and the dry season. Hence, the winter monsoon means clear skies and generally dry weather, with surface winds that blow from land to sea. In summer, the wind-flow pattern reverses itself as air over the continents becomes warmer than air above the water. A shallow thermal low develops over the continental interior. The heated air within the low rises, and the surrounding air responds by flowing counterclockwise into the low center. This condition brings moisture-bearing winds that sweep into the continent from the ocean. The humid air converges with drier continental air, causing the humid air to rise; further lifting is provided by hills and mountains. Lifting cools the air to its saturation point, resulting in heavy showers and thunderstorms. Thus, the summer monsoon of southeastern Asia, which lasts from about June through September, means wet, rainy weather (the wet season) with surface winds that blow from sea to land (see Fig. 7.20b). Although the majority of rain falls during the wet season, it does not rain all the time. In fact, rainy periods of between 15 to 40 days are often followed

FIGURE 7.20 Changing annual surface wind-flow patterns associated with the winter and summer Asian monsoons. FI

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FIGURE 7.21 Average annual precipitation for Cherrapunji, FI India. Note the abundant rainfall during the summer monsoon (April through October) and the lack of rainfall during the winter monsoon (November through March).

NOAA

by 3 to 15 days of hot, sunny weather known as monsoon breaks. The strength of the Indian monsoon appears to be related to the reversal of surface air pressure that occurs at irregular intervals about every two to seven years at opposite ends of the tropical South Pacific Ocean. As we will see later in this chapter, this reversal of pressure (which is known as the Southern Oscillation) is linked to the ocean warming phenomenon known as El Niño. During an El Niño event, surface water near the equator becomes warmer over the central and eastern Pacific. Over the region of warm water we find rising air, huge convective clouds, and heavy rain. Meanwhile, to the west of the warm water (over the region influenced by the summer monsoon), sinking air inhibits cloud formation and convection. Hence, during El Niño years, monsoon rainfall is more likely to be deficient. Summer monsoon rains over southern Asia can be truly extreme. The town of Cherrapunji (also known as Sohra), located about 300 km inland on the southern slopes of the Khasi Hills in northeastern India, receives an average of 1176 cm (463 in.) of rainfall each year, most of it between April and October (see Fig. 7.21). The town also holds world records for the heaviest rainfall measured anywhere in a 12-month period—2,647 cm (1042 in.) from August 1860 to July 1861—and the heaviest 48-hour rainfall, 249 cm (98 in.) on June 15–16, 1995. The summer monsoon rains are essential to the agriculture of southern and eastern Asia. More than two billion people rely on the summer rains so that crops will grow. The people also depend on the rains for drinking water. Unfortunately, the monsoon can be unreliable in both duration and intensity, and these are difficult to predict. Since the monsoon is vital to the survival of so many people, it is no wonder that meteorologists have investigated it extensively. They have tried to develop methods of accurately forecasting the intensity and duration of the monsoon. With the aid of current research projects and the latest climate models (which tie in the interaction of ocean and atmosphere), there is hope that monsoon forecasts will improve in accuracy. Monsoon wind systems exist in other regions of the world, such as in Australia, Africa, and North and South America, where large contrasts in temperature develop between oceans and continents. Usually, however, these systems are not as pronounced as in southeast Asia. For example, the North American monsoon affects northwest Mexico and the southwestern United States, especially Arizona, New Mexico, Nevada, and the southern part of California. In this region, spring and early summer are normally dry, as warm westerly winds predominate. By mid-July, however, southerly or southeasterly winds are more common, and so are afternoon showers and thunderstorms (see Fig. 7.22 and Fig. 7.23).

FIGURE 7.22 Enhanced infrared satellite image with heavy FI arrows showing strong monsoonal circulation. Moist, southerly winds from the Gulf of California and southeasterly winds (far right arrow) from the Gulf of Mexico are causing showers and thunderstorms (yellow and red areas) to form over the southwestern section of the United States during July 2001. ATMOSPHERIC CIRCULATIONS

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© C. Donald Ahrens

FIGURE 7.23 Clouds and thunFI derstorms forming over Arizona, as humid monsoonal air flows northward over the region during July 2007.

BRIEF REVIEW Before moving on to the next section here is a brief review of some of the main points so far: ●

The size of atmospheric circulations range from the smallest microscale to the larger mesoscale, to the largest macroscale.

Thermal pressure systems are shallow pressure systems that are driven by the unequal heating and cooling of Earth’s surface.

The sea breeze and the land breeze are types of thermal circulations that are due to uneven heating and cooling rates of land and water.

At the surface, a sea breeze blows from water to land; whereas a land breeze blows from land to water.

A valley breeze blows uphill during the day and a mountain breeze blows downhill at night.

Chinook (foehn) winds are warm, dry winds that blow downhill along the eastern side of the Rocky Mountains.

The main source of warmth for the chinook is compressional heating.

Santa Ana winds are warm, dry downslope winds that warm by compressional heating and blow from the east or northeast into Southern California.

Dust devils tend to form over dry terrain on clear, hot days. They are not tornadoes, although the winds of a large dust devil may cause minor damage to structures.

Monsoon winds are winds that change direction seasonally. In southern Asia, the winter monsoon, which blows from land to water, is dry; the summer monsoon, which blows from water to land, is wet.

Global Winds Up to now, we have seen that local winds vary considerably from day to day and from season to season. As you may suspect, these winds are part of a much larger circulation, the little whirls within larger whirls that we spoke of earlier in this chapter. Indeed, if the rotating high- and low-pressure areas in our atmosphere are like spinning eddies in a huge river, then the flow of air around the globe 188

is like the meandering river itself. When winds throughout the world are averaged over a long period of time, the local wind patterns vanish, and what we see is a picture of the winds on a global scale—what is commonly called the general circulation of the atmosphere. GENERAL CIRCULATION OF THE ATMOSPHERE Before we study the general circulation, we must remember that it only represents the average air flow around the world. Actual winds at any one place and at any given time may vary considerably from this average. Nevertheless, the average can answer why and how the winds blow around the world the way they do—why, for example, prevailing surface winds are northeasterly in Honolulu, Hawaii, and westerly in New York City. The average can also give a picture of the mechanism driving these winds, as well as a model of how heat is transported from equatorial regions poleward, keeping the climate in middle latitudes tolerable. The underlying cause of the general circulation is the unequal heating of Earth’s surface. We learned in Chapter 2 that, averaged over the entire earth, incoming solar radiation is roughly equal to outgoing earth radiation. However, we also know that this energy balance is not maintained for each latitude, since the tropics experience a net gain in energy, while polar regions suffer a net loss. To balance these inequities, the atmosphere transports warm air poleward and cool air equatorward. Although seemingly simple, the actual flow of air is complex; certainly not everything is known about it. In order to better understand it, we will first look at some models (that is, artificially constructed analogies) that eliminate some of the complexities of the general circulation. SINGLE-CELL MODEL The first model is the single-cell model, in which we assume that: . Earth’s surface is uniformly covered with water (so that differential heating between land and water does not come into play).

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FIGURE 7.24 Diagram (a) shows the general circulation of air on the side of Earth facing the sun on a nonrotating Earth FI uniformly covered with water and with the sun directly above the equator. (Vertical air motions are highly exaggerated in the vertical.) Diagram (b) shows the names that apply to the different regions of the world and their approximate latitudes.

. The sun is always directly over the equator (so that the winds will not shift seasonally). . Earth does not rotate (so that the only force we need deal with is the pressure gradient force). With these assumptions, the general circulation of the atmosphere on the side of Earth facing the sun would look much like the representation in Fig. 7.24a— a huge thermally driven convection cell in each hemisphere. (For reference, the names of the different regions of the world and their approximate latitudes are given in Figure 7.24b.) The circulation of air described in Fig. 7.24a is the Hadley cell (named after the eighteenth-century English meteorologist George Hadley, who first proposed the idea). It is driven by energy from the sun. Excessive heating of the equatorial area produces a broad region of surface low pressure, while at the poles excessive cooling creates a region of surface high pressure. In response to the horizontal pressure gradient, cold surface polar air flows equatorward, while at higher levels air flows toward the poles. The entire circulation consists of a closed loop with rising air near the equator, sinking air over the poles, an equatorward flow of air near the surface, and a return flow aloft. In this manner, some of the excess energy of the tropics is transported as sensible and latent heat to the regions of energy deficit at the poles. Such a simple cellular circulation does not actually exist on Earth. For one thing, Earth rotates, so the Coriolis force would deflect the southwardmoving surface air in the Northern Hemisphere to the right, producing easterly surface winds at practically all latitudes north of the equator. We know that this does not happen and that

prevailing winds in middle latitudes actually blow from the west. Therefore, observations alone tell us that a closed circulation of air between the equator and the poles is not an accurate model for a rotating earth. But this model does show us how a non-rotating planet would balance an excess of energy at the equator and a deficit at the poles. How, then, does the wind blow on a rotating planet? To answer, we will keep our model simple by retaining our first two assumptions—that is, that Earth is covered with water and that the sun is always directly above the equator. THREE-CELL MODEL If we allow Earth to spin, the simple convection system breaks into a series of cells as shown in Fig. 7.25. Although this model is considerably more complex than the single-cell model, there are some similarities. The tropical regions still receive an excess of heat and the poles a deficit. In each hemisphere, three cells instead of one have the task of energy redistribution. A surface high-pressure area is located at the poles, and a broad trough of surface low pressure still exists at the equator. From the equator to latitude 30°, the circulation is the Hadley cell. Let’s look at this model more closely by examining what happens to the air above the equator. (Refer to Fig. 7.25 as you read the following section.) Over equatorial waters, the air is warm, horizontal pressure gradients are weak, and winds are light. This region is referred to as the doldrums. (The monotony of the weather in this area has given rise to the expression “down in the doldrums.”) Here, warm, humid air rises, often condensing into huge cumulus clouds and thunderstorms ATMOSPHERIC CIRCULATIONS

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FIGURE 7.25 The idealized wind FI and surface-pressure distribution over a uniformly water-covered rotating earth.

that liberate an enormous amount of latent heat. This heat makes the air more buoyant and provides energy to drive the Hadley cell. The rising air reaches the tropopause, which acts like a barrier, causing the air to move laterally toward the poles. The Coriolis force deflects this poleward flow toward the right in the Northern Hemisphere and to the left in the Southern Hemisphere, providing westerly winds aloft in both hemispheres. (We will see later that these westerly winds reach maximum velocity and produce jet streams near 30° latitude and 60° latitude.) Air aloft moving poleward from the tropics constantly cools by giving up infrared radiation, and at the same time it also begins to converge, especially as it approaches the middle latitudes.* This convergence (piling up) of air aloft increases the mass of air above the surface, which in turn causes the air pressure at the surface to increase. Hence, at latitudes near 30°, the convergence of air aloft produces belts of high pressure called subtropical highs (or anticyclones). As the converging, relatively dry air above the highs slowly descends, it warms by compression. This subsiding air produces generally clear skies and warm surface temperatures; hence, on earth it is here that we find the major deserts of the world, such as the Sahara of Africa and the Sonoran of North America (see Fig. 7.26). Over the ocean, the weak pressure gradients in the center of the high produce only weak winds. According *You can see why the air converges if you have a globe of the world. Put your fingers on meridian lines at the equator and then follow the meridians poleward. Notice how the lines and your fingers bunch together in the middle latitudes.

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to legend, sailing ships traveling to the New World were frequently becalmed in this region; and, as food and supplies dwindled, horses were either thrown overboard or eaten. Because of this, the region is sometimes called the horse latitudes. From the horse latitudes near latitude 30°, some of the surface air moves back toward the equator. It does not flow straight back, however, because the Coriolis force deflects the air, causing it to blow from the northeast in the Northern Hemisphere and from the southeast in the Southern Hemisphere. These steady winds provided sailing ships with an ocean route to the New World; hence, these winds are called the trade winds. Near the equator, the northeast trades converge with the southeast trades along a boundary called the intertropical convergence zone (ITCZ). In this region of surface convergence, air rises and continues its cellular journey. Along the ITCZ, it is usually very wet as the rising air develops into huge thunderstorms that drop copious amounts of rain in the form of heavy showers (see Fig. 7.27). Meanwhile, at latitude 30°, not all of the surface air moves equatorward. Some air moves toward the poles

DID YOU KNOW? Christopher Columbus was a lucky man. The year he set sail for the New World, the trade winds had edged unusually far north, and a steady northeast wind eased his ships along. Only for about ten days did he encounter the light and variable wind more typical of this notorious region (30°N)— the horse latitudes.

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Image copyright Harris Shiffman, 2010. Used under license from Shutterstock.com

FIGURE 7.26 Subtropical deserts, FI such as the Sonoran Desert shown here, are mainly the result of sinking air associated with subtropical high-pressure areas.

and deflects toward the east, resulting in a more or less westerly air flow—called the prevailing westerlies or, simply, westerlies—in both hemispheres. Consequently, from Texas northward into Canada, it is much more common to experience winds blowing out of the west than from the east. The westerly flow in the real world is not constant as migrating areas of high and low pressure break up the surface flow pattern from time to time. In the middle latitudes of the Southern Hemisphere, where the surface is mostly water, winds blow more steadily from the west. As this mild surface air travels poleward from latitude 30°, it encounters cold air moving down from the poles. These two air masses of contrasting temperature do not readily mix. They are separated by a boundary called the polar front, a zone of low pressure—the subpolar

low—where surface air converges and rises, and storms and clouds develop. In our model in Fig. 7.25, some of the rising air returns at high levels to the horse latitudes, where it sinks back to the surface in the vicinity of the subtropical high. This middle cell (called the Ferrel cell, after the American meteorologist William Ferrel) is completed when surface air from the horse latitudes flows poleward toward the polar front. Notice in Fig. 7.25, p. 190, that in the Northern Hemisphere, behind the polar front the cold air from the poles is deflected by the Coriolis force, so that the general flow of air is from the northeast. Hence, this is the region of the polar easterlies. In winter, the polar front, with its cold air, can move into middle and subtropical latitudes, producing a cold polar outbreak. Along the front, a portion of the

NASA

FIGURE 7.27 The solid red line in FI this visible satellite image marks the position of the ITCZ in the eastern Pacific. The bright white clouds are huge thunderstorms forming along the ITCZ.

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rising air moves poleward, and the Coriolis force deflects the air into a westerly wind at high levels. Air aloft eventually reaches the poles, slowly sinks to the surface, and flows back toward the polar front, completing the weak polar cell. We can summarize all of this by referring back to Fig. 7.25 on p. 190 and noting that, at the surface, there are two major areas of high pressure and two major areas of low pressure. Areas of high pressure exist near latitude 30° and the poles; areas of low pressure exist over the equator and near 60° latitude in the vicinity of the polar front. Knowing the way the surface winds blow around these pressure systems on the three-cell model gives us a generalized picture of how surface winds blow throughout the world. The trade winds extend from the subtropical high to the equator, the westerlies from the subtropical high to the polar front, and the polar easterlies from the poles to the polar front. How does this three-cell model compare with actual observations of winds and pressure in the real world? We know, for example, that upper-level winds at middle latitudes generally blow from the west. The middle cell in our model, however, suggests an east wind aloft as air flows equatorward. Discrepancies do exist between this model and atmospheric observations, but the model does agree closely with the winds and pressure distribution at the surface, so we will examine this next. AVERAGE SURFACE WINDS AND PRESSURE: THE REAL WORLD When we examine the real world with its continents and oceans, mountains and ice fields, we obtain an average distribution of sea-level pressure and winds for January and July, as shown in Figs. 7.28a and 7.28b. Look closely at both maps and observe that there are regions where pressure systems appear to persist throughout the year. These systems are referred to as semipermanent highs and lows because they move only slightly during the course of a year. In Fig. 7.28a, we can see that there are four semipermanent pressure systems in the Northern Hemisphere during January. In the eastern Atlantic, between latitudes 25° and 35°N is the Bermuda-Azores high, often called the Bermuda high, and, in the Pacific Ocean, its counterpart, the Pacific high. These are the subtropical anticyclones that develop in response to the convergence of air aloft. Since surface winds blow clockwise around these systems, we find the trade winds to the south and the prevailing westerlies to the north. In the Southern Hemisphere, where there is relatively less land area, there is less contrast between land and water, and the subtropical highs show up as well-developed systems with a clearly defined circulation. Where we would expect to observe the polar front (between latitudes 40° and 65°), there are two semipermanent subpolar lows. In the North Atlantic, there is the Greenland-Icelandic low or simply Icelandic low, 192

which covers Iceland and southern Greenland, while the Aleutian low sits over the Gulf of Alaska and the Bering Sea near the Aleutian Islands in the North Pacific. These zones of cyclonic activity actually represent regions where numerous storms, having traveled eastward, tend to converge, especially in winter. In the Southern Hemisphere, where there is very little land to disrupt the flow, the subpolar low forms a continuous trough that completely encircles the globe. The January map (Fig. 7.28a) shows other pressure systems that are not semipermanent in nature but are still observed often. Over Asia, for example, there is a huge (but shallow) thermal anticyclone called the Siberian high, which forms because of the intense cooling of the land. South of this system, the winter monsoon shows up clearly, as air flows away from the high across Asia and out over the ocean. A similar (but less intense) anticyclone (called the Canadian high) is evident over North America. As summer approaches, the land warms and the cold, shallow highs disappear. In some regions, areas of surface low pressure replace areas of high pressure. The lows that form over the warm land are shallow thermal lows. On the July map (Fig. 7.28b), warm thermal lows are found over the Desert Southwest of the United States, over the plateau of Iran, and north of India. As the thermal low over India intensifies, warm, moist air from the ocean is drawn into it, producing the wet summer monsoon so characteristic of India and Southeast Asia. When we compare the January and July maps, we can see several changes in the semipermanent pressure systems. The strong subpolar lows so well developed in January over the Northern Hemisphere are hardly discernible on the July map. The subtropical highs, however, remain dominant in both seasons. Because the sun is overhead in the Northern Hemisphere in July and overhead in the Southern Hemisphere in January, the zone of maximum surface heating shifts seasonally. In response to this shift, the major pressure systems, wind belts, and ITCZ (heavy red line in Fig. 7.28) shift toward the north in July and toward the south in January.* THE GENERAL CIRCULATION AND PRECIPITATION PATTERNS The position of the major features of the general circulation and their latitudinal displacement (which annually averages about 10° to 15°) strongly influence the climate of many areas. For example, on the global scale, we would expect abundant rainfall where the air rises and very little where the air sinks. Thus, areas of high rainfall exist in the tropics, where humid air rises in conjunction with the ITCZ, and between about 40° and 55° latitude, where middle-latitude cyclonic storms and the polar front *An easy way to remember the seasonal shift of surface pressure systems is to think of birds. In the Northern Hemisphere, they migrate south in the winter and north in the summer.

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FIGURE 7.28 Average sea-level pressure distribution and surface wind-flow patterns for January (a) and for July (b). The solid red FI line represents the position of the ITCZ. ATMOSPHERIC CIRCULATIONS Copyright 2018 Cengage Learning. All Rights Reserved. May not be copied, scanned, or duplicated, in whole or in part. WCN 02-200-202

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FIGURE 7.31 Average annual precipitation for Los Angeles, FI California (Los Angeles International Airport), and Atlanta, Georgia (Hartsfield-Jackson International Airport), for the climatological period 1981–2010.

FIGURE 7.29 Rising and sinking air associated with the major FI pressure systems of Earth’s general circulation. Where the air rises, precipitation tends to be abundant (blue shade); where the air sinks, drier regions prevail (tan shade). Note that the sinking air of the subtropical highs produces the major desert regions of the world.

force air upward. Areas of low precipitation are found near 30° latitude in the vicinity of the subtropical highs and in polar regions where the air is cold and dry (see Fig. 7.29). During the summer, the Pacific high drifts northnorth ward to a position off the California coast (see Fig. 7.30). Sinking air on its eastern side produces a strong upperlevel subsidence inversion, which tends to keep summer weather along the West Coast relatively dry. The rainy season typically occurs in winter when the high moves south and storms can penetrate the region. Observe in Fig. 7.30 that along the East Coast, the clockwise circulation of winds around the Bermuda high brings warm, tropical air northward into the United States and southern Canada from the Gulf of Mexico and the Atlantic Ocean.

Because sinking air is not as well developed on this side of the high, the humid air can rise and condense into towering cumulus clouds and thunderstorms. In part, then, it is the air motions associated with the subtropical highs that keep summer weather dry in California and moist in Georgia. (Compare the patterns for Los Angeles, California, and Atlanta, Georgia, in Fig. 7.31.) WESTERLY WINDS AND THE JET STREAM In Chapter 6, we learned that the winds above the middle latitudes in both hemispheres blow in a wavy west-to-east direction. The reason for these westerly winds is that, aloft, we generally find higher pressure over equatorial regions and lower pressures over polar regions. Where these upperlevel winds tend to concentrate into narrow bands, we find rivers of fast-flowing air—what we call jet streams. Characteristics of Jet Streams Atmospheric jet streams are swiftly flowing air currents hundreds of miles long, normally less than several hundred miles wide, and typically less than a mile thick. Wind speeds in the

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FIGURE 7.30 During the summer, the FI Pacific high moves northward. Sinking air along its eastern margin (over California) produces a strong subsidence inversion, which causes relatively dry weather to prevail. Along the western margin of the Bermuda high, southerly winds bring in humid air, which rises, condenses, and produces abundant rainfall.

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FIGURE 7.32 Average position of the FI polar jet stream and the subtropical jet stream, with respect to a model of the general circulation in winter. Both jet streams are flowing from west to east.

*The subtropical jet stream is normally found between 20° and 30° latitude.

flowing air, or jet core, is represented by the heavy dark arrows. The map shows a strong polar jet stream sweeping south over the Great Plains with an equally strong subtropical jet over the Gulf states. Notice that the polar jet has a number of loops, with one off the west coast of North America and another over eastern Canada. Observe in the satellite image (Fig. 7.34b) that the polar jet stream (blue arrow) is directing cold, polar air into the Plains states, while the subtropical jet stream (orange arrow) is sweeping subtropical moisture, in the form of a dense cloud cover, over the southeastern states. The looping pattern of the polar jet stream has an important function. In the Northern Hemisphere, where

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central core of a jet stream often exceed 100 knots and occasionally 200 knots. Jet streams are usually found at the tropopause at elevations between 10 and 14 km (33,000 and 46,000 ft) although they can occur at both higher and lower altitudes. Jet streams were first encountered by high-flying military aircraft during World War II, but their existence was suspected before that time. Ground-based observations of fast-moving cirrus clouds had revealed that westerly winds aloft must be moving rapidly indeed. Figure 7.32 illustrates the average position of two jet streams, the tropopause, and the general circulation of air for the Northern Hemisphere in winter. Both jet streams are located at tropopause gaps, where mixing between tropospheric and stratospheric air takes place. The jet stream situated near 30° latitude at about 13 km (43,000 ft) above the subtropical high is the subtropical jet stream.* To the north, the jet stream situated at a lower altitude of about 10 km (33,000 ft) near the polar front is known as the polar front jet stream or, simply, the polar jet stream. In Fig. 7.32, the wind in the center of the jet stream would be flowing as a westerly wind away from the viewer. This direction, of course, is only an average, as jet streams of often flow in a wavy west-to-east pattern. When the polar jet stream flows in broad loops that sweep north and south, it may even merge with the subtropical jet. Occasionally, the polar jet splits into two jet streams. The jet stream to the north is often called the northern branch of the polar jet, whereas the one to the south is called the southern branch. Figure 7.33 illustrates how the polar jet stream and the subtropical jet stream might appear as they sweep around Earth in winter. We can better see the looping pattern of the jet by po studying Fig. 7.34a, which shows the position of the polar jet stream and the subtropical jet stream at the 300-mb level (near 9 km or 30,000 ft) on March 9, 2005. The fastest

FIGURE 7.33 Jet streams are swiftly flowing currents of air FI that move in a wavy west-to-east direction. The figure shows the position of the polar jet stream and subtropical jet stream in winter. Although jet streams are shown as one continuous river of air, in reality they are discontinuous, with their position varying from one day to the next. ATMOSPHERIC CIRCULATIONS

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FIGURE 7.34 (a) Position of the polar jet stream and the subtropical jet stream at the 300-mb level (about 9 km or 30,000 ft above FI sea level) on March 9, 2005. Solid lines are lines of equal wind speed (isotachs) in knots. (b) Satellite image showing clouds and positions of the jet streams for the same day.

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the air flows southward, swiftly moving air directs cold air equatorward; where the air flows northward, warm air is carried toward the poles. Jet streams, therefore, play a major role in the global transfer of heat. Moreover, since jet

FIGURE 7.35 Diagram (a) is a model that shows a vertical 3-D FI view of the polar front in association with a sharply dipping 500-mb pressure surface, an isotherm (dashed line), and the position of the polar front jet stream in winter. The diagram is highly exaggerated in the vertical. Diagram (b) represents a 500-mb chart that cuts through the polar front as illustrated by the dipping 500-mb surface in (a). Sharp temperature contrasts along the front produce tightly packed contour lines and strong winds (contour lines are in meters above sea level).

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streams tend to meander around the world, we can easily understand how pollutants or volcanic ash injected into the atmosphere in one part of the globe could eventually settle to the ground many thousands of kilometers downwind. And, as we will see in Chapter 8, the looping nature of the polar jet stream has an important role in the development of mid-latitude cyclonic storms. The Formation of Jet Streams Since jet streams are bands of strong winds, they form in the same manner as all winds do—from horizontal differences in air pressure. In Fig. 7.35a, notice that the polar jet stream forms along the polar front where sharp contrasts in temperature propro duce rapid horizontal pressure changes and strong winds. Notice also in Fig. 7.35a that as the 20°C isotherm crosses the frontal boundary, it dips sharply. This rapid change in temperature causes the constant pressure (isobaric) 500-mb surface to bend sharply as it passes through the front. The bending of the 500-mb surface in Fig. 7.35a shows up as tightly packed contour lines and strong winds along the front on the 500-mb chart (Fig. 7.35b). Because northto-south temperature contrasts along the front are greater in winter than they are in summer, the polar jet stream shows seasonal variations. In winter, the polar jet stream winds are stronger and the jet moves farther south, sometimes as far south as Florida and Mexico. In summer, the polar jet stream is weaker and forms over higher latitudes. Look back at Fig. 7.32 on p. 195 and see that the subtropical jet stream forms on the poleward (north) side of the Hadley cell, at a higher altitude than the polar jet stream. Here, warm air aloft carried poleward by the Hadley cell produces sharp temperature differences, strong pressure gradients, and high winds. Although the polar and subtropical jets are the two most frequently in the news, there are other jet streams that deserve mentioning. For example, there is a low-level jet stream that forms just above the Central Plains of the

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in the tropics, from the west in the middle latitudes, and from the east in polar regions.

DID YOU KNOW? The jet stream is in part responsible for the only American casualties by enemy attack on the continental United States in World War II. During the war, when the existence of the jet stream was first confirmed, the Japanese attempted to drop bombs on the United States mainland by launching balloons that carried explosives and incendiary devices. The hydrogenfilled balloons drifted east from Japan for thousands of miles across the Pacific Ocean at an altitude above 30,000 feet. Unfortunately, a group of six picnickers in Oregon found a balloon bomb in the woods and attempted to move it, which caused it to explode, killing all six people. Estimates are that as many as 300 balloon bombs may still be scattered throughout regions of the western United States.

United States. During the summer, this jet (which can attain wind speeds of more than 60 knots) often contributes to the formation of nighttime thunderstorms by transporting moisture and warm air northward. Higher up in the atmosphere, over the subtropics, a summertime easterly jet called the tropical easterly jet forms at the base of the tropopause. And during the dark polar winter, a stratospheric polar jet forms near the top of the stratosphere.

BRIEF REVIEW Before going on to the next section, which describes the many interactions between the atmosphere and the ocean, here is a review of some of the important concepts presented so far: ●

The two major semipermanent subtropical highs that influence the weather of North America are the Pacific high situated off the west coast and the Bermuda high situated off the southeast coast.

The polar front is a zone of low pressure where cyclonic storms often form. It separates the mild westerlies of the middle latitudes from the cold, polar easterlies of the high latitudes.

In equatorial regions, the intertropical convergence zone (ITCZ) is a boundary where air rises in response to the convergence of the northeast trades and the southeast trades. Along the ITCZ huge thunderstorms produce heavy rain showers.

In the Northern Hemisphere, the major global pressure systems and wind belts shift northward in summer and southward in winter.

The northward movement of the Pacific high in summer tends to keep summer weather along the west coast of North America relatively dry.

Jet streams exist where strong winds become concentrated in narrow bands. The polar-front jet stream is associated with the polar front. The polar jet meanders in a wavy west-to-east pattern, becoming strongest in winter when the contrast in temperature along the front is greatest.

The subtropical jet stream is found on the poleward side of the Hadley cell, between 20° and 30° latitude. It is normally observed at a higher altitude than the polar jet stream.

The general flow of air around the globe in both the Northern and Southern Hemispheres finds surface winds blowing from the east

At the surface, in both the Northern and Southern Hemispheres, areas of high pressure tend to persist at the poles and along a belt near 30° latitude. Areas of surface low pressure are generally found near the equator and between about 40° and 55° latitude.

Atmosphere-Ocean Interactions The atmosphere and oceans are both dynamic fluid systems that interact with one another in many complex ways. For example, evaporation of ocean water provides the atmosphere with surplus water that falls as precipitation. The latent heat that is taken up by the water vapor during evaporation goes into the atmosphere during condensation to fuel storms. The storms, in turn, produce winds that blow over the ocean, which causes waves and currents. The currents, in turn, can modify the weather and climate of a region by bringing in vast quantities of warm or cold water. The complexity of the interaction between the atmosphere and ocean makes our scientific understanding of how one influences the other on a global scale far from complete. What we will focus on in the remainder of this chapter is what we do know, beginning with ocean currents. Later, we will concentrate on some of the most important weather and climate oscillations that result from atmosphere-ocean interactions. GLOBAL WIND PATTERNS AND SURFACE OCEAN CURRENTS As the wind blows over the oceans, it causes the surface water to drift along with it. The moving water gradually piles up, creating pressure differences within the water itself. This leads to further motion several hundreds of meters down into the water. In this manner, the general wind flow around the globe starts the major surface ocean currents moving. The relationship between the general wind flow and ocean currents can be seen by comparing Fig. 7.36 with Fig. 7.37. Because of the larger frictional drag in water, ocean curcur rents move more slowly than the prevailing winds above. Typically, they range in speed from several kilometers per day to several kilometers per hour. However, comparing Fig. 7.36 with Fig. 7.37, we can see that ocean currents do not follow the wind pattern exactly; rather, they spiral in semiclosed circular whirls. On the eastern edge of continents there is usually a warm current that flows from the equator to the pole. For example, in the North Atlantic, flowing northward along the east coast of the United States, is a tremendous warm water current called the Gulf Stream, which carries vast quantities of tropical water into higher latitudes. Off the coast of North Carolina, the Gulf Stream provides warmth and moisture for developing mid-latitude cyclonic storms. ATMOSPHERIC CIRCULATIONS

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FIGURE 7.36 Annual average global wind patterns and surface FI high-pressure areas over the oceans.

WINDS AND UPWELLING Earlier, we saw that the cool California Current flows roughly parallel to the west coast of North America. From this, we might conclude that summer surface water temperatures will be cool along the coast of Washington and gradually warmer as

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Notice in Fig. 7.37 that as the Gulf Stream moves northward, the prevailing westerlies steer it away from the coast of North America and eastward toward Europe. Generally, it widens and slows as it merges into the broader North Atlantic Drift. As this current approaches Europe, part of it flows northward along the coasts of Great Britain and Norway, bringing with it warm water (which helps keep winter temperatures much warmer than one would expect this far north). The other part flows southward as the Canary Current, which transports cool, northern water equatorward. In the Pacific Ocean, the counterpart

to the Canary Current is the California Current that carries cool water southward along the coastline of the western United States. Hence, on the western edge of the major continents, there is usually a cool current that flows from the pole toward the equator. Up to now, we have seen that atmospheric circulations and ocean circulations are closely linked; wind blowing over the oceans produces surface ocean currents. The currents, along with the wind, transfer heat from tropical areas, where there is a surplus of energy, to polar regions, where there is a deficit. This helps to equalize the latitudinal energy imbalance with about 40 percent of the total heat transport in the Northern Hemisphere coming from surface ocean currents. The environmental implications of this heat transfer are tremendous. If the energy imbalance were to go unchecked, yearly temperature differences between low and high latitudes would increase greatly, and the climate would gradually change.

FIGURE 7.37 Average position and extent of the major surface ocean currents. Cold currents are shown in blue; warm currents FI are shown in red.

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FIGURE 7.38 Average sea surface temperatures (°F) along the FI west coast of North America during August.

we move south. A quick glance at the water temperatures along the west coast of the United States during August (see Fig. 7.38) quickly alters that notion. The coldest water is observed along the northern California coast near Cape Mendocino. The reason for the cold, coastal water is upwelling—the rising of cold water from below. For upwelling to occur, the wind must flow more or sum less parallel to the coastline. Notice in Fig. 7.39 that summer winds tend to parallel the coastline of California. As the wind blows over the ocean, the surface water beneath it is set in motion. As the surface water moves, it bends slightly to its right due to the Coriolis effect. (Remember, it would bend to the left in the Southern Hemisphere.) The water beneath the surface also moves, and it too bends slightly to its right. The net effect of this phenomenon is that a rather shallow layer of surface water moves at right angles to the surface wind and heads seaward. As the surface water drifts away from the coast, cold, nutrient-rich

water from below rises (upwells) to replace it. Upwelling is strongest and surface water is coolest where the wind parallels the coast, such as it does in summer along the coast of northern California. Because of the cold coastal water, summertime weather along the West Coast often consists of low clouds and fog, as the air over the water is chilled to its saturation point. Upwelling produces good fishing conditions, however, as higher concentrations of nutrients are brought to the surface. But swimming is only for the hardiest of souls, as the average surface water temperature along the coast of northern California in summer is nearly 10°C (18°F) colder than the average coastal water temperature found at the same latitude along the Atlantic coast. Between the ocean surface and the atmosphere, there is an exchange of heat and moisture that depends, in part, on temperature differences between water and air. In winter, when air-water temperature contrasts are greatest, there is a substantial transfer of sensible and latent heat from the ocean surface into the atmosphere. This energy helps to maintain the global air flow. Consequently, even a relatively small change in surface ocean temperatures can modify atmospheric circulations and have far-reaching effects on global weather and climate patterns. The next section describes how weather events can be linked to surface ocean temperature changes in the tropical Pacific. EL NIÑO, LA NIÑA, AND THE SOUTHERN OSCILLATION Along the west coast of South America, where the cool Peru Current sweeps northward, southerly winds promote upwelling of cold, nutrient-rich water that gives rise to large fish populations, especially anchovies. The abundance of fish supports a large population of sea birds whose droppings (called guano) produce huge phosphaterich deposits, a valuable source of fertilizer. Every two to five years or so, a warm current of nutrient-poor tropical water moves southward, replacing the cold, nutrient-rich surface water. Because this condition frequently occurs

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FIGURE 7.39 As winds blow FI parallel to the west coast of North America, surface water is transported to the right (out to sea). Cold water moves up from below (upwells) to replace the surface water. The large H represents the position of the Pacific high in summer. Blue arrows show the movement of water.

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around Christmas, local fishermen (more than a century ago) called this warm current Corriente del Niño, which translated means “current of the Christ Child;” hence, the warm current’s name—El Niño. It was once thought that El Niño was a local event that occurred only along the west coast of Peru and Ecuador. It is now known that the ocean-warming can cover an area of the tropical Pacific much larger than the continental United States. Although in recent decades the term El Niño has gained global prominence, the large, prolonged warming that develops at irregular intervals every three to seven years is often referred to as an El Niño event. During such an event, the surface water temperature across much of the tropical Pacific rises by 0.5°C (0.9°F) or more for periods of a few months to a year or more. During a El Niño event, because of the unusually warm water, large numbers of fish and marine plants may die and other marine species may travel far from the tropics. A very strong El Niño event in 2015–2016 resulted in whale and hammerhead sharks moving up the California coast, and California squid were reported in southeast Alaska. The El Niño of 1972–1973 was one of the first to be studied in depth. It reduced the annual Peruvian anchovy catch by more than 50 percent in 1972. Since much of the harvest of this fish is converted into fishmeal and exported for use in feeding livestock and poultry, the world’s fish meal production in 1972 was greatly reduced. Countries such as the United States that rely on fish meal for animal feed had to use soybeans as an alternative. This raised poultry prices in the United States by more than 40 percent. Peru’s fishing industry is now more carefully managed, so that the decline of anchovy during El Niño is not causing such severe repercussions. Why does the ocean become so warm over such a large area during an El Niño event? Normally in the tropical Pacific Ocean, trade winds are persistently blowing westward from a region of higher pressure over the eastern Pacific toward a region of lower pressure over the western Pacific, centered near Indonesia (see Fig. 7.40a). The trades create upwelling that brings cool water to the sursur face over the eastern tropical Pacific and along the west coast of South America. As this cool water moves westward, it is heated by sunlight and the atmosphere. Consequently, surface water along the equator usually is cooler in the eastern Pacific and warmer in the western Pacific. In addition, the dragging of surface water by the trades raises sea level by a few inches in the western Pacific and lowers it in the eastern Pacific, which produces a thick layer of warm water over the tropical western Pacific Ocean. The higher level of water in the western Pacific causes warm water to flow slowly eastward toward South America as a weak, narrow ocean current called the countercurrent. Every few years, the surface atmospheric pressure patterns break down. Air pressure rises over the western 200

Pacific and falls over the eastern Pacific (see Fig. 7.40b). This change in pressure weakens the trades, and over a period of a few months, easterly winds may be replaced by westerly winds that strengthen the countercurrent. Warm water gradually extends farther east across the tropical Pacific. The warm surface water provides energy for the development of huge convective clouds. Hence, the location of showers and thunderstorms may shift east as well. If the ocean warming reaches a certain threshold and remains at that level long enough, an El Niño event is declared to be in place. Toward the end of the warming period, which typically lasts about a year but may recur for another year or more, atmospheric pressure over the eastern Pacific reverses and begins to rise, whereas, over the western Pacific, it falls. This seesaw pattern of reversing surface air pressure at opposite ends of the Pacific Ocean is called the Southern Oscillation. Because the reversals in air pressure and the ocean warming are more or less simultaneous, scientists call this phenomenon the El Niño–Southern Oscillation or ENSO for short. Although most El Niño episodes follow a similar evolution, each event has its own personality, differing in both strength and behavior. During especially strong El Niño events (such as in 1982–1983,1997–1998, and 2015–2016) the easterly trades may actually become westerly winds, as illustrated in Fig. 7.40b. As these winds push eastward, they drag surface water with them. This dragging raises sea level in the eastern Pacific and lowers sea level in the western Pacific. The eastward-moving water gradually warms under the tropical sun, becoming as much as 6°C (11°F) warmer than normal in the eastern equatorial Pacific. Gradually, a thick layer of warm water pushes into coastal areas of Ecuador and Peru, choking off the upwelling that supplies cold, nutrient-rich water to South America’s coastal region. The unusually warm water can extend from South America’s coastal region for many thousands of kilometers westward along the equator (see Fig. 7.41a). Following an El Niño event, the trade winds usually return to normal. However, if the trades are exceptionally strong, then unusually cold surface water develops over the central and eastern Pacific, and warm water and rainy weather are confined mainly to the western tropical Pacific (see Figure 7.41b). This cold-water episode is termed La Niña (“the girl child”). In some ways, La Niña can be thought of as the opposite of El Niño. Water temperatures in the eastern tropical Pacific are colder than average during La Niña, while they are warmer than average during El Niño. Keep in mind, however, that while an El Niño event features a reversal of the typical east-to-west wind flow and ocean currents across the tropical Pacific, La Niña features a strengthening of the same east-to-west wind flow and ocean currents. La Niña events sometimes occur immediately after El Niño events, but they can also develop independently.

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FIGURE 7.40 In diagram (a), under FI non-El Niño conditions higher pressure over the southeastern Pacific and lower pressure near Indonesia produce easterly trade winds along the equator. These winds promote upwelling and cooler ocean water in the eastern Pacific, while warmer water prevails in the western Pacific. The trades are part of a circulation that typically finds rising air and heavy rain over the western Pacific and sinking air and generally dry weather over the eastern Pacific. When the trades are exceptionally strong, water along the equator in the eastern Pacific becomes quite cool. This cool event is called La Niña. During El Niño conditions—diagram (b)—atmospheric pressure decreases over the eastern Pacific and rises over the western Pacific. This change in pressure causes the trades to weaken or reverse direction, which enhances the countercurrent that carries warm water from the west over a vast region of the eastern tropical Pacific.

The large areas of abnormally warm water associated with El Niño and abnormally cold water associated with La Niña can have effects far beyond South America. During El Niño, the usual region of warm water in the western tropical Pacific is moved thousands of kilometers to the east. There, the water fuels the atmosphere with additional warmth and moisture, which feeds into storminess and rainfall. Along with the added warmth from the oceans, latent heat is released during condensation. All of these factors help change the regional atmospheric circulation, effects that can propagate for thousands of miles, rearranging the locations where rising and sinking air tend to predominate. As a result, some regions of the world experience too much rainfall during El Niño, whereas others have too little. Over the warm tropical central Pacific, the

frequency of typhoons usually increases. However, over the tropical Atlantic, between Africa and Central America, the winds aloft tend to disrupt the organization of thunderstorms that is necessary for hurricane development; hence, there are usually fewer hurricanes in this region during strong El Niño events. And, as we saw earlier in this chapter, during a strong El Niño, summer monsoon conditions tend to weaken over India. Although the actual mechanism by which changes in surface ocean temperatures influence global wind patterns is not fully understood, the by-products are plain to see. For example, during exceptionally warm El Niños, drought is normally felt in Indonesia, southern Africa, and Australia, while heavy rains and flooding often occur in Ecuador and Peru. In the Northern Hemisphere, a strong ATMOSPHERIC CIRCULATIONS

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NOAA/PHEL/TAO

FIGURE 7.41 (a) Average sea surface temperature departures from normal as measured by satellite. During El Niño conditions, FI upwelling is greatly diminished and warmer than normal water (deep red color) extends from the coast of South America westward, across the Pacific. (b) During La Niña conditions, strong trade winds promote upwelling, and cooler than normal water (dark blue color) extends over the eastern and central Pacific.

several tenths of a degree Fahrenheit for a few months. The very strong El Niño of 2015–2016 helped produce recordwarm global surface temperatures in 2015 and 2016. Figure 7.42 shows warm events (El Niño years) in red and cold events (La Niña years) in blue. Notice in Fig. 7.42 that it is possible to have two or more El Niño or

FIGURE 7.42 The Oceanic Niño Index (ONI). The numbers FI on the left side of the diagram represent a running 3-month mean for sea surface temperature variations (from normal, in degrees Celsius) over the tropical Pacific Ocean from latitude 5°N to 5°S and from longitude 120°W to 170°W (area outlined in adjacent map). Warm EI Niño episodes are in red; cold La Niña episodes are in blue. Warm and cold events occur when the deviation from the normal is 0.5 or greater (faint red and blue shading in graph). An index value between 0.5 and 0.9 is considered weak; an index value between 1.0 and 1.4 is considered moderate; an index value between 1.5 and 2.0 is considered strong; and an index value greater than 2.0 is considered very strong.

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NOAA and Jan Null

NOAA and Jan Null

subtropical westerly jet stream often directs mid-latitude cyclonic storms into California and heavy rain into the Gulf Coast states. The total global impact of an El Niño event due to flooding, winds, and drought may exceed many billions of dollars. A major El Niño event can also pump so much heat into the atmosphere that global temperatures rise by

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FIGURE 7.43 Typical winter weather patterns across North America during an El Niño warm event (a) and during a La Niña cold FI event (b). During El Niño conditions, a persistent trough of low pressure forms over the north Pacific and, to the south of the low, the jet stream (from off the Pacific) steers wet weather and cyclonic storms into California and the southern part of the United States. During La Niña conditions, a persistent high-pressure area forms south of Alaska forcing the polar jet stream and accompanying cold air over much of western North America. The southern branch of the polar jet stream directs moist air from the ocean into the Pacific Northwest, producing a wet winter for that region.

La Niña periods in a row, separated by neutral conditions. El Niño events typically develop in the autumn (Northern Hemisphere), peak in the winter, and weaken in spring and summer. Although El Niño events rarely last more than a year, La Niña conditions are more likely to recur or persist for two or three years in a row. Over a long time frame (several decades), the tropical Pacific is divided roughly equally among El Niño, La Niña, and neutral conditions. As we have seen, El Niño and the Southern Oscillation are part of a large-scale ocean-atmosphere interaction that can take several years to run its course. During this time, certain regions in the world can expect significant climatic responses to an ENSO event. Figure 7.43 shows how typical winter weather patterns over North America will change between El Niño conditions and La Niña conditions. Such ocean-atmosphere interactions, where a warmer or colder ocean surface can influence weather patterns in distant parts of the world, are called teleconnections. As an example of teleconnections, note in Fig. 7.43 that in the Pacific Northwest, La Niña winters tend to be wetter than usual, while drier-than-average winters are more likely during El Niño. Notice also in Fig. 7.43 that in some areas, the teleconnections related to La Niña are opposite to those of El Niño. However, this is not the case in every location, because the mechanisms behind El Niño and La Niña are not exact opposites. Also, it is important to keep in mind that these preferred patterns are not guaranteed to occur with every El Niño or La Niña event, since each one is dif different. For example, Los Angeles, California, had a drierthan-average winter during the strong El Niño event of 2015–2016, even though strong El Niño winters tend to be wetter than average in Los Angeles. Over the long term, though, the effects shown in Fig. 7.43 are the most likely ones to expect. So while we cannot say with certainty that an El Niño winter will bring more rainfall than average to

California, we can say that the odds are higher during a strong El Niño than they might otherwise be. For several decades, atmospheric scientists have developed and improved techniques for predicting the future state of the El Niño–Southern Oscillation and the onset of El Niño and La Niña events. This is an especially difficult challenge because ENSO events are not like typical day-to-day weather features. Instead, they unfold in a much more gradual way through a complex sequence of events involving both the ocean and atmosphere. Using long-range climate models, the National Oceanic and Atmospheric Administration and several other agencies around the world now issue outlooks giving the probability that El Niño or La Niña will develop over the next few months. When it becomes clear in the Northern Hemisphere summer or autumn that an El Niño or La Niña event is actually developing, then forecasters can provide several months’ notice of the type of conditions that are most likely to occur in the winter, the time of year when the events usually reach their full strength. Up to this point, we have looked at El Niño and the Southern Oscillation, and at how the reversal of surface ocean temperatures and atmospheric pressure combine to influence regional and global weather and climate patterns. There are other atmosphere-ocean interactions that can have an effect on large-scale weather patterns. Some of these are described in the following section. OTHER ATMOSPHERE-OCEAN INTERACTIONS Over the North Pacific Ocean, periodic changes in surface ocean temperatures can influence weather along the west coast of North America for much longer periods of time than El Niño and La Niña. The Pacific Decadal Oscillation (PDO) is like ENSO in that it has a warm phase and a cool phase, and the temperature patterns produced by the PDO are similar in some locations to those produced ATMOSPHERIC CIRCULATIONS

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FIGURE 7.44 Typical winter sea FI surface temperature departure from normal in °C during the Pacific Decadal Oscillation’s warm phase (a) and cool phase (b).

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impact of an El Niño event. Likewise, La Niña events are generally reinforced when the PDO is in its cool phase. The PDO was generally positive from 1922 to 1947, negative from 1947 to 1977, positive from 1977 to 1998, and negative from 1998 to 2013. In the mid-2010s, the PDO appeared to be entering a new long-term positive phase. Scientists are researching the factors that cause the PDO to shift from one long-term phase to the other. Over the Atlantic, a periodic reversal of pressure called the North Atlantic Oscillation (NAO) has a substantial effect on the weather in Europe and along the east coast of North America, especially during the winter. The NAO is measured by the difference in atmospheric pressure between the vicinity of the Icelandic low and the region of the Bermuda-Azores high. (Note that the NAO is defined by atmospheric change, whereas ENSO and the PDO are defined by oceanic change.) When the air pressure in the North Atlantic is unusually high near the Azores and unusually low near Iceland, the increased pressure gradient leads to stronger westerlies. These winds, in turn, drive frequent, powerful cyclonic storms into northern Europe, which tends to make winters wet and mild. During this positive phase of the NAO, winters in the eastern United States also tend to be wet and relatively mild, while northern Canada and eastern Europe are typically cold and dry (see Fig. 7.45a). The negative phase of the NAO occurs when the atmospheric pressure in the vicinity of the Icelandic low rises, while the pressure drops in the region of the Bermuda high (see Fig. 7.45b). This results in a reduced pressure gradient and weaker westerlies, which steer fewer and weaker winter storms across the Atlantic. The weaker jet stream also allows for storms to move or develop unusually far south, bringing wet weather to southern Europe and other areas

(Source: JISAO, University of Washington, http://jisao.washington.edu/pdo.)

by ENSO. However, the PDO has more of an influence in the mid-latitudes of the North Pacific than in the tropical Pacific, and it operates on a much longer time scale than ENSO. Each PDO phase tends to predominate for 20 to 30 years before switching to the other phase. During the warm (or positive) phase of the PDO, unusually warm surface water exists along the west coast of North America, while over the central North Pacific, cooler than normal surface water prevails (see Fig. 7.44a). At the same time, the Aleutian low in the Gulf of Alaska strengthens, which causes more Pacific storms to move into Alaska and California. This situation causes winters, as a whole, to be warmer and drier over northwestern North America. Elsewhere, winters tend to be drier over the Great Lakes, and cooler and wetter in the southern United States. Cool (or negative) PDO phases have cooler-thanaverage surface water along the west coast of North America and an area of warmer-than-normal surface water extending from Japan into the central North Pacific (see Fig. 7.44b). Winters in the cool phase tend to be cooler and wetter than average over northwestern North America, wetter over the Great Lakes, and warmer and drier in the southern United States. These climate patterns only represent average conditions, as individual years within either PDO phase may vary considerably, sometimes reverting to the opposite phase for short periods. These variations, which can last several months to a year or more, make it more difficult to decipher exactly when the PDO has changed from one long-term phase to the other. Knowing the long-term phase of the PDO can help improve seasonal climate prediction, because the warm phase of PDO tends to reinforce the climate patterns of El Niño, which helps strengthen the

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FIGURE 7.45 Change in surface atmospheric pressure and typical winter weather patterns associated with the (a) positive phase FI and (b) negative phase of the North Atlantic Oscillation.

near the Mediterranean Sea. The weaker, more variable jet stream during a negative NAO phase also allows cold air masses to move southward more readily into northern Europe and the eastern United States, which tends to make these regions generally colder and drier than usual (although the presence of cold air can sometimes lead to intense winter storms in the eastern United States). Far eastern Canada and eastern Europe often experience milder-than-average winters during a negative NAO. Although the NAO often varies from month to month, it can also tend to favor one phase or the other for several years. Closely related to the North Atlantic Oscillation, but analyzed at a more northerly latitude, is the Arctic Oscillation (AO).The AO is defined by variations in atmospheric pressure between the Arctic and the North Pacific and Atlantic. During the positive (warm) phase of the AO, higher pressures to the south and lower pressures across the Arctic produce strong westerly winds aloft. These winds wrap around the semipermanent zone of upper-level low pressure over the North Pole that is sometimes referred to as the polar vortex. When a positive AO is in place, the polar vortex is typically stronger than usual, and cold arctic air generally remains bottled up in and near the polar regions. Thus, winters in the United States tend to be warmer than normal, while winters over Newfoundland and Greenland tend to be very cold. Meanwhile, strong winds over the

Atlantic direct storms into northern Europe, bringing with them wet, mild weather. During the negative (cold) phase of the AO, pressure differences are smaller between the Arctic and regions to the south, leading to weaker and more variable westerly winds aloft. The weaker westerlies mean that the polar vortex can more easily shift to lower latitudes, and cold arctic air is now able to penetrate farther south, often producing colder-than-normal winters over much of the United States. Cold air also invades northern Europe and Asia, while Newfoundland and Greenland normally experience warmer than normal winters. Much like the NAO, the AO switches from one phase to another on an irregular basis, and one phase may predominate for several years in a row, bringing with it a succession of either cold or mild winters. Unlike El Niño and La Niña, the NAO and AO can switch modes over just a few weeks’ time, and these changes cannot be predicted more than a couple of weeks in advance. This influence makes it more challenging to anticipate winter conditions over eastern North America and Europe than over western North America, where the more slowly varying ENSO and the PDO play a larger role. As our knowledge of the interactions between the ocean and atmosphere improves, we can expect scientists to gain skill at predicting shifts in all of these phenomena, as well as the resulting influences they have on regional weather and climate.

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SUMMARY In this chapter, we examined a variety of atmospheric circulations. We looked at small-scale winds and found that eddies can form in a region of strong wind shear, especially in the vicinity of a jet stream. On a slightly larger scale, land and sea breezes blow in response to local pressure differences created by the uneven heating and cooling rates of land and water. Monsoon winds change direction seasonally, while mountain and valley winds change direction daily. A warm, dry wind that descends the eastern side of the Rocky Mountains is the chinook. The same type of wind in the Alps is the foehn. A warm, dry downslope wind that blows into southern California is the Santa Ana wind. Local intense heating of the surface can produce small rotating winds, such as the dust devil, while downdrafts in a thunderstorm are responsible for the desert haboob. The largest pattern of winds that persists around the globe is called the general circulation. At the surface in both hemispheres, winds tend to blow from the east in the tropics, from the west in the middle latitudes, and from the east in polar regions. Where upper-level westerly winds tend to concentrate into narrow bands, we find jet streams. The annual shifting of the major pressure systems and wind belts—northward in July and southward in January—strongly influences the annual precipitation of many regions. Toward the end of the chapter we examined the interaction between the atmosphere and oceans. Here we found the interaction to be an ongoing process where everything, in one way or another, seems to influence everything else. On a large scale, winds blowing over the surface of the water drive the major ocean currents; the oceans, in turn, release energy to the atmosphere, which helps to maintain the general circulation of winds. When atmospheric circulation patterns change over the tropical Pacific, and the trade winds weaken or reverse direction, warm tropical water is able to flow eastward toward South America where it chokes off upwelling. When the warm water extends over a vast area of the tropical Pacific, and persists for several months to a year or more, the warming is called an El Niño event, and the associated reversal of pressure over the Pacific Ocean is called the Southern Oscillation. The large-scale interaction between the atmosphere and the ocean during El Niño and the Southern Oscillation (ENSO) affects global atmospheric circulation patterns, resulting in too much rain in some areas and not enough in others. Over the northern central Pacific and along the west coast of North America the surface water temperature reverses every 20 to 30 years, a phenomenon called the Pacific Decadal Oscillation (PDO). Over the Atlantic Ocean there is a periodic reversal of air pressure called the North Atlantic Oscillation that influences weather in various regions of the world. Atmospheric pressure changes over the Arctic produce a related phenomenon, the Arctic Oscillation which causes 206

winter weather patterns to change across the United States, Greenland, and Europe. Researchers are studying how the interchange between atmosphere and ocean can produce such events.

KEY TERMS The following terms are listed (with corresponding page numbers) in the order they appear in the text. Define each. Doing so will aid you in reviewing the material covered in this chapter. scales of motion, 174 microscale, 174 mesoscale, 174 synoptic scale, 174 global scale, 174 macroscale, 174 rotor, 175 wind shear, 175 clear air turbulence (CAT), 176 thermal circulation, 176 sea breeze, 178 land breeze, 178 valley breeze, 180 mountain breeze, 180 katabatic wind, 180 chinook wind, 181 Santa Ana wind, 182 haboob, 184 dust devils (whirlwinds), 184 monsoon, 186 monsoon wind system, 184 general circulation of the atmosphere, 188 Hadley cell, 189 doldrums, 189 subtropical highs, 190

trade winds, 190 intertropical convergence zone (ITCZ), 190 westerlies, 191 polar front, 191 subpolar low, 191 polar easterlies, 191 Bermuda high, 192 Pacific high, 192 Icelandic low, 192 Aleutian low, 192 Siberian high, 192 jet stream, 195 subtropical jet stream, 195 polar front jet stream, 195 Gulf Stream, 197 upwelling, 199 El Niño, 200 Southern Oscillation, 200 ENSO, 200 La Niña, 200 teleconnections, 203 Pacific Decadal Oscillation (PDO), 203 North Atlantic Oscillation (NAO), 204 Arctic Oscillation (AO), 205 polar vortex, 205

QUESTIONS FOR REVIEW . Describe the various scales of motion and give an example of each. . What is wind shear and how does it relate to clear air turbulence? . Using a diagram, explain how a thermal circulation develops. . Why does a sea breeze at the surface blow from sea to land and a land breeze from land to sea? . Which wind will produce clouds: a valley breeze or a mountain breeze? Why? . What are katabatic winds? How do they form? . Explain why chinook winds are warm and dry. . Describe how dust devils usually form.

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. (a) Briefly explain how the monsoon wind system develops over eastern and southern Asia. (b) Why in India is the summer monsoon wet and the winter monsoon dry? . Draw a large circle. Now, place the major surface semipermanent pressure systems and the wind belts of the world at their appropriate latitudes. . According to Fig. 7.25 (p. 190), most of the United States is located in what wind belt? . Explain how and why the average surface pressure features shift from summer to winter. . Explain why summers along the West Coast of the United States tend to be dry, whereas along the East Coast summers tend to be wet. . How does the polar front influence the development of the polar front jet stream? . Why is the polar jet stream typically stronger in winter? . Explain the relationship between the general circulation of air and the circulation of surface ocean currents. . Describe how the winds along the west coast of North America produce upwelling. . What is an El Niño event? What happens to the surface pressure at opposite ends of the Pacific Ocean during the Southern Oscillation? . Describe how an ENSO event may influence the weather in different parts of the world. . What are the conditions over the tropical eastern and central Pacific Ocean during La Niña? . Describe the ocean surface temperatures associated with the Pacific Decadal Oscillation. . How does the positive (warm) phase of the Northern Atlantic Oscillation differ from the negative (cold) phase? . During the negative (cold) phase of the Arctic Oscillation when Greenland is experiencing mild winters, what type of winters (cold or mild) is Northern Europe usually experiencing?

QUESTIONS FOR THOUGHT AND EXPLORATION A ATION . Suppose you are fishing in a mountain stream during . . .

.

. . . . .

the early morning. Is the wind more likely to be blowing upstream or downstream? Explain why. Why, in Antarctica, are winds on the high plateaus usually lighter than winds in steep, coastal valleys? What atmospheric conditions must change so that the westerly flowing polar front jet stream reverses direction and becomes an easterly flowing jet stream? After a snowstorm, Cheyenne, Wyoming, reports a total snow accumulation of 48 cm (19 in.), while the maximum depth in the surrounding countryside is only 28 cm (11 in.). If the storm’s intensity and duration were practically the same for a radius within 50 km of Cheyenne, explain why Cheyenne received so much more snow. The prevailing winds in southern Florida are northeasterly. Knowing this, would you expect the strongest sea breezes to be along the east or west coast of southern Florida? What about the strongest land breezes? Explain why icebergs tend to move at right angles to the direction of the wind. Give two reasons why pilots would prefer to fly in the core of a jet stream rather than just above or below it. Why do the major ocean currents in the North Indian Ocean reverse direction between summer and winter? Why is the surface water along the northern California coast warmer in winter than in summer? The Coriolis force deflects moving water to the right of its intended path in the Northern Hemisphere and to the left of its intended path in the Southern Hemisphere. Why, then, does upwelling tend to occur along the western margin of continents in both hemispheres?

Go to the Reference section of the Global Environment Watch: Meteorology portal. Search under the term “El Niño” and consult the document “El Niño and La Niña,” from Environmental Science: In Context (or other reference works in this portal) to answer these questions: What is the earliest recorded use of the term El Niño? Have scientists found evidence for El Niño and La Niña in prehistoric times? Are El Niño and La Niña expected to change as Earth’s climate warms?

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CHAPTER

8

Air Masses, Fronts, and Middle-Latitude Cyclones

Air Masses

A

Fronts

70 miles an hour, accompanied by a deep bellowing sound,

Mid-Latitude Cyclonic Storms

with its icy blast, swept over the land, and everything was

Contents

bout two o’clock in the afternoon it began to grow dark from a heavy, black cloud which was seen in the northnorth

west. Almost instantly the strong wind, traveling at the rate of

frozen hard. The he water in the little ponds in the roads froze in waves, sharp-edged and pointed, as the gale had blown it. The chickens, pigs, and other small animals were frozen in their tracks. Wagon wheels ceased to roll, frozen to the ground. Men, going from their barns or fields a short distance from their homes, in slush and water, returned a few minutes later walking on the ice. Those caught out on horseback were frozen to their saddles and had to be lifted off and carried to the fire to be thawed apart. TTwo wo young men were frozen to death near Rushville. One ne of them was found with his back against a tree, with his horse’s bridle over his arm and his horse frozen in front of him. The he other was partly in a kneeling position, with a tinder box in one hand and a flint in the other, with both eyes wide open, as if intent on trying to strike a light. Many other casualties were reported. As to the exact temperature, however, no instrument has left any record; but the ice was frozen in the stream, as variously reported, from six inches to a foot in thickness in a few hours. John Moses, Illinois: Historical and Statistical

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T

he opening quotation of this chapter details the passage of a spectacular cold front as it moved through Illinois on December 21, 1836. While some of the incidents described are likely exaggerated, at least one observing station, in Augusta, Illinois, reported air temperatures dropping from 40°F at dawn to 0°F by 2 p.m., and many people and animals are believed to have perished in the sudden cold. Fortunately, temperature changes of this magnitude with cold fronts are uncommon. In this chapter, we will examine the more typical weather associated with cold fronts and warm fronts. We will address questions such as: Why are cold fronts usually associated with showery weather? How can warm fronts during the winter cause freezing rain and sleet to form over a vast area? And how can one read the story of an approaching warm front by observing its clouds? We will also see how weather fronts are an integral part of a midlatitude cyclonic storm. But, first, so that we may better understand fronts and storms, we will examine air masses. We will look at where and how they form and the type of weather usually associated with them.

Air Masses An air mass is an extremely large body of air whose properties of temperature and humidity are fairly similar in any horizontal direction at any given altitude. A single air mass may cover more than a million square kilometers. In Fig. 8.1, a large winter air mass, associated with a

high-pressure area, covers over half of the United States. Note that although the surface air temperature and dew point vary somewhat, everywhere the air is cold and dry, with the exception of the zone of snow showers on the eastern shores of the Great Lakes. This cold, shallow anticyclone will drift eastward, carrying with it the temperature and moisture characteristic of the region where the air mass formed; hence, in a day or two, cold air will be located over the central Atlantic Ocean. Part of weather forecasting is, then, a matter of determining air mass characteristics, predicting how and why they change, and in what direction the systems will move. SOURCE REGIONS Regions where air masses originate are known as source regions. In order for a huge mass of air to develop uniform characteristics, its source region should be generally flat and of uniform composition, with light surface winds. The longer the air remains stagnant over its source region, or the longer the path over which the air moves, the more likely it will acquire properties of the surface below. Ideal source regions are usually those areas dominated by surface high pressure, which include the ice- and snow-covered arctic plains in winter and subtropical oceans in summer. The middle latitudes, where surface temperatures and moisture characteristics vary considerably, are not good source regions. Instead, the middle latitudes are a transition zone where air masses with different physical properties move in, clash, and produce an exciting array of weather activity.

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FIGURE 8.1 Here, a large, extremely cold winter air mass is dominating the weather over much of the United States. At almost all cities, the air is cold and dry. Upper number is air temperature (°F); bottom number is dew point (°F).

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▼ Table 8.1

Air Mass Classification and Characteristics

SOURCE REGION

ARCTIC REGION (A)

POLAR (P)

TROPICAL (T)

Land

cA

cP

cT

Continental (c)

extremely cold, dry, stable; ice- and snow-covered surface

cold, dry, stable

hot, dry, stable air aloft; unstable surface air

Water

mP

mT

Maritime (m)

cool, moist, unstable

warm, moist; usually unstable

CLASSIFICATION Air masses are usually classified according to their temperature and humidity, both of which usually remain fairly uniform in any horizontal direction. There are cold and warm air masses, humid and dry air masses. Air masses are grouped into five general categories according to their source region (see ▼ Table 8.1). Air masses that originate in polar latitudes are designated by the capital letter “P” (for polar); those that form in warm tropical regions are designated by the capital letter “T” (for tropical). If the source region is land, the air mass will be dry and the lowercase letter “c” (for continental) precedes the P or T. If the air mass originates over water, it will be moist—at least in the lower layers—and the lowercase letter “m” (for maritime) precedes the P or T. We can now see that polar air originating over land will be classified cP on a surface weather map, whereas tropical air originating over water will be marked as mT. In winter, an extremely cold air mass that forms over the Arctic is designated as cA, continental arctic. Sometimes, however, it is difficult to distinguish between arctic and polar air masses, especially when the arctic air mass has traveled over warmer terrain. Table 8.1 lists the five basic air masses.

After the air mass spends some time over its source region, it may begin to move in response to strengthening winds aloft. As it moves away from its source region, the air mass encounters surfaces that may be warmer or colder than itself. When the air mass is colder than the underlying surface, it is warmed from below, which produces instability at low levels. In this case, increased convection and turbulent mixing near the surface usually result in good visibility, cumuliform clouds, and showers of rain or snow. On the other hand, when the air mass is warmer than the surface below, the lower layers are chilled by contact with the cold Earth. Warm air above cooler air produces stable air with little vertical mixing. This situation causes the accumulation of dust, smoke, and pollutants, which restricts surface visibilities. In moist air, stratiform clouds accompanied by drizzle or fog may form. AIR MASSES OF NORTH AMERICA The principal air masses (with their source regions) that enter the United States are shown in Fig. 8.2. We are now in a position to study the formation and modification of each of these air masses and the variety of weather that accompanies them.

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FIGURE 8.2 Air mass source regions and their paths.

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FIGURE 8.3 A shallow but large dome of extremely cold air— a continental arctic air mass—moves slowly southeastward across the upper plains. The leading edge of the air mass is marked by a cold front. (Numbers represent air temperature, °F.)

Continental Polar (cP) and Continental Arctic (cA) Air Masses The bitterly cold weather that invades southern Canada and the United States in winter is associated with continental polar and continental arctic air masses. These air masses originate over the ice- and snow-covered regions of the Arctic, northern Canada, and Alaska where long, clear nights allow for strong radiational cooling of the surface. Air in contact with the surface becomes quite cold and stable. Since little moisture is added to the air, it is also quite dry. Eventually a portion of this cold air breaks away and, under the influence of the airflow aloft, moves southward as an enormous shallow high-pressure area, as illustrated in Fig. 8.3. (As we will see later in this chapter, air masses can be more than 100 times wider than they are tall.) As the cold air moves into the interior plains, there are no topographic barriers to restrain it, so it continues southward, bringing with it frigid temperatures. As the air mass moves over warmer land to the south, the air temperature moderates slightly. Even during the afternoon, when the surface air is most unstable, cumulus clouds are rare because of the extreme dryness of the air. At night, when the winds die down, rapid radiational surface cooling and clear skies combine to produce low minimum temperatures. If the cold air moves as far south as central or southern Florida, or south Texas, the winter vegetable crop may be severely damaged. When the cold, dry air mass moves over a relatively warm body of water, such as the Great Lakes, heavy snow showers—called lake-effect snows—often form on the downwind shores. (More information on lake-effect snows is provided in Focus section 8.1.) In winter, the generally fair weather accompanying polar continental and arctic air masses is due to the stable nature of the atmosphere aloft. Sinking air develops above the large dome of high pressure. The subsiding air warms by compression and creates warmer air, which lies above colder surface air, often causing a strong upper-level temperature inversion to form. Should the anticyclone 212

stagnate over a region for several days, visibility gradually drops as pollutants become trapped in the cold air near the ground. Usually, however, winds aloft move the cold air mass either eastward or southeastward. The Rockies, Sierra Nevada, and Cascades normally protect the Pacific Northwest from the onslaught of arctic air, but, occasionally, very cold air masses do invade these regions. When the upper-level winds over Washington and Oregon blow from the north or northeast on a trajectory beginning over northern Canada or Alaska, cold air can slip over the mountains and extend its icy fingers all the way to the Pacific Ocean. As the air moves off the high plateau, over the mountains, and on into the lower valleys, compressional heating of the sinking air causes its temperature to rise, so that by the time it reaches the lowlands, it is considerably warmer than it was originally. However, in no way would this air be considered warm. In some cases, the subfreezing temperatures slip over the Cascades and extend southward into the coastal areas of southern California. A similar but less dramatic warming of continental polar and arctic air masses occurs along the east coast of the United States. Air rides up and over the lower Appalachian Mountains. Turbulent mixing and compressional heating increase the air temperatures on the downwind side, with the result that cities located to the east of the Appalachian Mountains usually do not experience temperatures as low as those on the west side. In Fig. 8.1 on p. 210, notice that for the same time of day—in this case 7 a.m. (EST)— Philadelphia, on the eastern side of the mountains, with an air temperature of 14°F, is 16°F warmer than Pittsburgh, at –2°F, on the western side of the mountain. Figure 8.4 shows two upper-air patterns that led to extremely cold outbreaks of arctic air during December 1989 and 1990. Upper-level winds typically blow from west to east, but, in both of these cases, the flow, as indicated by the heavy, dark arrows, had a strong north-south (meridional) trajectory. The H represents the positions of the cold surface anticyclones. Numbers on the map represent minimum temperatures (°F) recorded during the cold spells. East of the Rocky Mountains, over 350 record low temperatures were set between December 21 and 24, 1989, with the arctic outbreak causing an estimated $480 million in damage to the fruit and vegetable crops in Texas and Florida. Along the West Coast, the frigid air during December 1990 caused over $300 million in damage to the vegetable and citrus crops, as temperatures over parts of California plummeted to their lowest readings in more than fifty years. Notice in both cases how the upperlevel wind directs the paths of the air masses. Continental polar air that moves into the United States in summer has properties much different from its winter counterpart. The source region remains the same but the air is now accompanied by long summer days that melt

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FOCUS

ON A SPECIAL TOPIC 8.1 O

buoyant and less stable. Rapidly, the air sweeps up moisture, soon becoming saturated. Out over the water, the vapor condenses into steam fog. As the air continues to warm, it rises and forms billowing cumuliform clouds, which continue to grow as the air becomes more unstable. Eventually, these clouds produce heavy showers of snow, which make the lake seem like a snow factory. Once the air and clouds reach the downwind side of the lake, additional lifting is provided by low hills and the convergence of air as it slows down over the rougher terrain. In late winter, the frequency and intensity of lake-effect snows often taper off as the temperature contrast between water and air diminishes and larger portions of the lakes freeze. Generally, the longer the stretch of water over which the air mass travels (the longer the fetch), the greater the amount of warmth and moisture derived from the lake, and the greater the potential for heavy snow showers. Consequently, forecasting lake-effect snowfalls depends to a large degree on determining the trajectory of the air as it flows over the lake. Regions that experience heavy lake-effect snowfalls are shown in Fig. 2.* As the cold air moves farther east, the heavy snow showers usually taper off; *Buffalo, New York, is a city that experiences heavy lakeeffect snows. Visit the National Weather Service website in Buffalo at http://www.weather.gov/buf/lakeeffect and read about lake-effect snowstorms measured in feet, and other interesting weather stories.

FIGURE 2 Areas shaded white show regions that experience heavy lake-effect snows.

however, the western slope of the Appalachian Mountains produces further lifting, enhancing the possibility of more and heavier showers. The heat given off during condensation warms the air and, as the air descends the eastern slope, compressional heating warms it even more. Snowfall ceases, and by the time the air arrives in Philadelphia, New York, or Boston, the only remaining trace of the snow showers occurring on the other side of the mountains is the puffy cumulus clouds drifting overhead. Lake-effect (or enhanced) snows are not confined to the Great Lakes. In fact, any large unfrozen lake (such as the Great Salt Lake) can enhance snowfall when cold, relatively dry air sweeps over it. Moreover, a type of lake-effect snow occurs when cold air moves over a relatively warm ocean, then lifts slightly as it moves over a landmass. Such oceaneffect snows are common over Cape Cod, Massachusetts, in winter.

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During the winter, when the weather in the Midwest is dominated by clear and cold polar or arctic air, people living on the eastern shores of the Great Lakes brace themselves for heavy snow showers. Snowstorms that form on the downwind side of one of these lakes are known as lake-effect snows. (Since the lakes are responsible for enhancing the amount of snow that falls, these snowstorms are also called lake-enhanced snows, especially when the snow is associated with a cold front or mid-latitude cyclone.) Such storms are highly localized, extending from just a few kilometers to more than 100 km inland. The snow usually falls as a heavy shower or squall in a concentrated zone. So centralized is the region of snowfall that one part of a city may accumulate many inches of snow, while, in another part, the ground is bare. The amount of snow that falls can be enormous: for example, 65 inches fell near Buffalo, New York (on the eastern side of Lake Ontario), in less than 48 hours during November 2014. Lake-effect snows are most numerous from November to January. During these months, cold air moves over the lakes when they are relatively warm and often not frozen. The contrast in temperature between water and air can be as much as 25°C (45°F). Studies show that the greater the contrast in temperature, the greater the potential for snow showers. In Fig. 1, we can see that as the cold air moves over the warmer water, the air mass is quickly warmed from below, making it more

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Lake-Effect (Enhanced) Snows

FIGURE 1 The formation of lake-effect snows. Cold, dry air crossing the lake gains moisture and warmth from the water. The more buoyant air now rises, forming clouds that deposit large quantities of snow on the lake’s leeward (downwind) shores.

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FIGURE 8.4 Average upper-level wind flow (heavy arrows) and surface posi position of anticyclones (H) associated with two extremely cold outbreaks of arctic air during late December 1989 and 1990. Numbers on the map represent some minimum temperatures (°F) measured during each cold snap.

snow and warm the land. The air is only moderately cool, and surface evaporation adds water vapor to the air. A summertime continental polar air mass usually brings relief from the oppressive heat often occurring in the central and eastern states, as cooler air lowers the air temperature to more comfortable levels. Daytime heating warms the lower layers, producing surface instability. With its added moisture, the water vapor in the rising air may condense and create a sky dotted with fair-weather cumulus clouds. When an air mass moves over a large body of water, its original properties can change considerably. For instance, cold, dry continental polar air moving over the Gulf of Mexico warms rapidly and gains moisture. The air quickly assumes the qualities of a maritime air mass. Notice in Fig. 8.5 that rows of cumulus clouds are forming over the Gulf of Mexico parallel to northerly surface winds as continental polar air is being warmed by the water beneath

it, causing the air mass to destabilize. As the air continues its journey southward into Mexico and Central America, strong, moist northerly winds build into heavy clouds (bright white area) and showers along the northern coast. In this way, a once cold, dry, and stable air mass can be modified to such an extent that its original characteristics are no longer discernible. When this happens, the air mass is often given a new designation. In summary, polar and arctic air masses are responsible for the bitterly cold winter weather that can cover wide sections of North America. When the air mass originates over the Canadian Northwest Territories, frigid air can bring record-breaking low temperatures. Such was the case on Christmas Eve 1983, when arctic air covered most of North America. (A detailed look at this air mass and its accompanying record-setting low temperatures is given in Focus section 8.2.)

NOAA

FIGURE 8.5 Visible satellite image showing the modification of cold continental polar air as it moves over the warmer Gulf of Mexico and the Atlantic Ocean.

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The winter of 1983–1984 was one of the coldest on record across North America. Unusually frigid weather arrived in December, which ended up as the coldest December for the contiguous 48 states since records began in 1895 (a title it still holds). During the first part of the month, continental polar air covered most of the northern and central plains. As the cold air moderated slightly, far to the north a huge mass of bitterly cold arctic air was forming over the frozen reaches of the Canadian Northwest Territories. By mid-month, the frigid air, associated with a massive high-pressure area, covered all of northwest Canada. Meanwhile, aloft, strong northerly winds directed the leading edge of the frigid air southward over the prairie provinces of Canada and southward into the United States. The extraordinarily cold air was accompanied in some regions by winds gusting to 45 knots. At least one news reporter labeled the onslaught, “The Siberian Express.” (The term has since become common in news coverage of cold waves in the United States, although it does not necessarily mean the air mass originates in Siberia.) In many locations across the United States, the Siberian Express dropped temperatures to the lowest readings ever recorded during the month of December. On December 22, Elk Park, Montana, recorded an unofficial low of –64°F, only 6°F higher than the all-time low of –70°F for the United States (excluding Alaska), which was recorded at Rogers Pass, Montana, on January 20, 1954. The center of the massive anticyclone gradually pushed southward out of Canada. By December 24, its center was over eastern Montana (see Fig. 3), where the sea-level pressure at Miles City reached an incredible 1064 mb (31.42 in.), a record for the 48 contiguous states. An enormous ridge of high pressure stretched from the Canadian arctic coast to the Gulf of Mexico. On the east side of the ridge, cold westerly winds brought lake-effect snows to the eastern shores of the Great Lakes. To the south of the high-pressure center, cold easterly winds, rising along the elevated plains, brought light amounts of upslope

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The Return of the Siberian Express

FIGURE 3 Surface weather map for 7 a.m., EST, December 24, 1983. Solid lines are isobars. Areas shaded white represent snow. An extremely cold arctic air mass covers nearly 90 percent of the United States. (Weather symbols for the surface map are given in Appendix C.)

snow* to sections of the Rocky Mountain states. Notice in Fig. 3 that, on Christmas Eve, arctic air covered almost 90 percent of the United States. As the cold air swept eastward and southward, a hard freeze caused hundreds of millions of dollars in damage to the fruit and vegetable crops in Texas, Louisiana, and Florida. On Christmas Day, 125 record low temperature readings were set in twenty-four states. That afternoon, at 1:00 p.m., it was actually colder in Atlanta, Georgia, at 9°F, than it was in Fair Fairbanks, Alaska (10°F). One of the worst cold waves to occur in December during the twentieth century continued through the week, as many new record lows were established in the Deep South from Texas to Louisiana. By January 1, the extreme cold had moderated, as the upper-level winds became more westerly. These winds brought milder Pacific air eastward into the Great Plains. The warmer pattern continued until about January 10, when the Siberian Express decided to make a return visit. Driven by strong upper-level northerly winds, impulse after impulse of arctic air from Canada swept across the United *Upslope snow forms as cold air moving from east to west gradually rises (and cools even more) as it approaches the Rocky Mountains.

States. On January 18, a low of –65°F was recorded at Middle Sinks, Utah. On January 19, temperatures plummeted to a new low of –7°F for the airports in Philadelphia and Baltimore, tying the coldest readings observed on any date at these locations. Toward the end of the month, the upper-level winds once again became more westerly. Over much of the nation, the cold air moderated. But the Express was to return at least one more time. The beginning of February saw relatively warm air covering much of the United States from California to the Atlantic coast. However, on February 4, an arctic outbreak spread southward and eastward across the United States. Although freezing air extended southward into central Florida, the Express ran out of steam, and a February heat wave soon engulfed most of the United States east of the Rocky Mountains as warm, humid air from the Gulf of Mexico spread northward. Even though February 1984 was a warmer-than-normal month over much of the United States, the winter of 1983–1984 (December, January, and February) will go down in the record books as one of the coldest winters for the United States as a whole since reliable record keeping began.

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NOAA

FIGURE 8.6 Clouds and airflow aloft (large blue arrow) associated with maritime polar air moving into California. The large L shows the position of an upper-level low. Regions experiencing precipitation are also shown. The small, white clouds over the open ocean are cumulus clouds forming in the conditionally unstable air mass. (Precipitation symbols are given in Appendix C at the back of the book.)

Maritime Polar (mP) Air Masses During the winter, polar and arctic air originating over Asia and frozen polar regions is carried eastward and southward over the Pacific Ocean by the circulation around the prevailing Aleutian low. The ocean water modifies these cold air masses by adding warmth and moisture to them. Since this air travels over water many hundreds or even thousands of kilometers, it gradually changes into a maritime polar air mass. By the time this air mass reaches the Pacific coast, it is cool, moist, and conditionally unstable. The ocean’s effect is to keep air near the surface warmer than the air aloft. Temperature readings in the 40s and 50s (°F) are common near the surface, while air at an altitude of about a kilometer or so may be at the freezing point. Within this colder air, characteristics of the original cold, dry air mass may still prevail. As the air moves inland, coastal mountains force it to rise,

and much of its water vapor condenses into rain-producing clouds. In the colder air aloft, the rain changes to snow, with heavy amounts accumulating in mountain regions. Over the relatively warm open ocean, the cool moist air mass produces cumulus clouds that show up as tiny white splotches on a visible satellite image (see Fig. 8.6). When maritime polar air from the Pacific moves into North America, it loses much of its moisture as it crosses a series of mountain ranges. Beyond these mountains, it travels over a cold, elevated plateau that chills the surface air and slowly transforms the lower level into dry, stable continental polar air. East of the Rockies this air mass is referred to as Pacific air (see Fig. 8.7). Here, it often brings fair weather and temperatures that are cool but not nearly as cold as the continental polar and arctic air that invades this region from northern Canada. In fact, when Pacific air from the west replaces retreating cold air from the north, chinook winds often develop (see p. 181). Furthermore, when the modified maritime polar air replaces moist subtropical air, thunderstorms can form along the boundary separating the two air masses. Along the east coast of North America, maritime polar air originates in the North Atlantic when modified continental polar air moves over a large body of water. Because the water of the North Atlantic is very cold and the air mass travels only a short distance over water, wintertime Atlantic maritime polar air masses are usually much colder than their Pacific counterparts. Because the prevailing winds aloft are westerly, Atlantic maritime polar air masses are also much less likely to move into the United States, although they often affect Europe. Figure 8.8 illustrates a typical late winter or early spring surface weather pattern that carries maritime polar air from the Atlantic into the New England and middle Atlantic states. A slow-moving, cold anticyclone drifting to the east (north of New England) causes a northeasterly onshore flow of cold, moist air to the south. The boundary separating this invading colder air from warmer air even farther south is marked by a stationary front. North of this front, northeasterly winds provide generally undesirable weather, consisting of damp air and low, thick clouds from which light precipitation falls in the form of rain, drizzle, or snow. As we will see

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FIGURE 8.7 After crossing several mountain ranges, cool, moist, mari maritime polar air from off the Pacific Ocean descends the eastern side of the Rockies as modified, relatively dry Pacific air.

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lows (see Fig. 8.28a, p. 232), sometimes swing northeastward along the coast, where they become northeasters (or nor’easters) bringing with them strong northeasterly winds, heavy rain or snow, and coastal flooding. (We will examine northeasters later in this chapter when we examine mid-latitude cyclonic storms.)

FIGURE 8.8 Winter and early spring surface weather pattern that usually prevails during the invasion of cold, moist maritime polar (mP) air into the mid-Atlantic and New England states. (Green-shaded area represents light rain and drizzle; pink-shaded region represents freezing rain and sleet; white-shaded area is experiencing snow.)

NOAA

later in this chapter, when upper atmospheric conditions are right, mid-latitude cyclonic storms may develop along the stationary front, move eastward, and intensify near the shores of Cape Hatteras. Such storms, called Hatteras

Maritime Tropical (mT) Air Masses The wintertime source region for Pacific maritime tropical air masses is the subtropical central and eastern Pacific Ocean. (Look back at Fig. 8.2, p. 211.) Air from this region must travel over many kilometers of water before it reaches the California coast. Consequently, these air masses are of often very warm and moist by the time they arrive along the West Coast. In winter, the warm air produces heavy precipitation usually in the form of rain, even at high elevations. Melting snow and rain quickly fill rivers, which overflow into the low-lying valleys. The rapid snowmelt leaves local ski slopes barren, and the heavy rain can cause disastrous mud slides in the steep canyons. Figure 8.9 shows maritime tropical air (usually referred to as subtropical air) streaming into northern California on January 1, 1997. The flow of humid, subtropical air, which originated near the Hawaiian Islands, was termed by at least one forecaster the “Pineapple Express,” a term that has since become common in media coverage. After battering the Pacific Northwest with heavy rain, the Pineapple Express roared into northern and central California, causing catastrophic floods that sent over 100,000 people fleeing from their homes, landslides that closed roads, property damage (including crop losses)

FIGURE 8.9 An infrared satellite image that shows maritime tropical air (heavy yellow arrow) moving into northern California on January 1, 1997. The warm, humid airflow (sometimes called the “Pineapple Express”) produced heavy rain and extensive flooding in northern and central California. AIR MASSES, FRONTS, AND MIDDLE-LATITUDE CYCLONES Copyright 2018 Cengage Learning. All Rights Reserved. May not be copied, scanned, or duplicated, in whole or in part. WCN 02-200-202

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NOAA/NWS

FIGURE 8.10 An atmospheric river streaming north from the Gulf of Mexico brought torrential rain to parts of Louisiana and nearby states on March 10, 2016. Rainfall totals for the month of March (not shown) exceeded 20 inches in parts of Louisiana.

that amounted to more than $1.5 billion, and eight fatalities. Yosemite National Park, which sustained over $170 million in damages due mainly to flooding, was forced to close for more than two months. The Pineapple Express is one example of the feature known as an atmospheric river. Like a river of moisture far above ground, an atmospheric river transports huge amounts of water vapor from one place to another, typically from the warm, humid tropics and subtropics into the mid-latitudes. A typical atmospheric river is 400 to 600 km wide and a few kilometers deep. The strongest atmospheric rivers can transport up to 15 times more moisture in the form of water vapor than the amount of water flowing into the Gulf of Mexico at the mouth of the Mississippi River. Along the west coast of the United States, a few atmospheric river events each year can provide as much as half of all annual precipitation. A powerful atmospheric river in early December 2012 brought more than 20 in. of rain to parts of northern California, with wind gusts as high as 150 mi/hr. Atmospheric rivers also flow north from lower latitudes to strike the central and eastern United States, such as the one that brought 10 to 20 in. of rain and devastating floods to parts of Texas and Louisiana in March 2016 (see Fig. 8.10). The humid subtropical air that influences much of the weather east of the Rockies originates over the Gulf of Mexico and Caribbean Sea. In winter, cold polar and arctic air tends to dominate the continental weather scene, so maritime tropical air is usually confined to the Gulf and extreme southern states. Occasionally, a slow-moving cyclonic storm system over the Central Plains draws warm, humid air northward. Gentle south or southwesterly 218

winds carry this air into the central and eastern parts of the United States in advance of the system. Since the land is still extremely cold, air near the surface is chilled to its dew point. Fog and low clouds form in the early morning, dissipate by midday, and re-form in the evening. This mild winter weather in the Mississippi and Ohio valleys lasts, at best, only a few days. Soon cold air will move down from the north behind the eastward-moving storm system. Along the boundary between the two air masses, the warm, humid air is lifted above the more-dense cold, polar air, which often leads to heavy and widespread precipitation and storminess. When a large, mid-latitude cyclonic storm system stalls over the Central Plains, a constant supply of warm, humid air from the Gulf of Mexico can bring recordbreaking maximum temperatures to the eastern half of the country. Sometimes the air temperatures are higher in the mid-Atlantic states than they are in the Deep South, as compressional heating warms the air even more as it moves downslope after crossing the Appalachian Mountains. Figure 8.11 shows a surface weather map and the associated upper-level airflow (heavy arrow) that brought unseasonably warm maritime tropical air into the central and eastern states during March 2012. A large surface high-pressure area centered off the east coast coupled with a strong southwesterly flow aloft carried warm, moist air into the Midwest and East, causing a record-breaking March heat wave. The flow aloft prevented the surface low and the cooler air behind it from making much eastward progress, so that the warm spell lasted a week or more in some locations. Michigan saw its first ever 90°F reading in March, and Burlington, Vermont, reached 80°F. More

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FIGURE 8.11 Weather conditions during an unseasonably warm spell in the eastern portion of the United States that occurred in the second half of March 2012. The surface low-pressure area, fronts, and maximum temperatures are shown for March 21. The heavy arrow is the average upper-level flow during the warm period.

than 7000 daily record highs were set or tied in the United States during March. Note that, on the west side of the surface low, the winds aloft funneled cool Pacific air into the western states, bringing chilly conditions from California to the Rockies. Hence, while people in the Northwest were huddled around heaters, others several thousand miles away in the Midwest and Northeast were soaking up sunshine and warmth. We can see that it is the upper-level flow, directing cooler air southward and warm subtropical air northward, that makes these contrasts in temperature possible. In summer, the circulation of air around the Bermuda High (which sits off the southeast coast of North America—see Fig. 7.30, p. 194) pumps warm, humid maritime tropical air northward from off the Gulf of Mexico and from off the Atlantic Ocean into the eastern half of the United States. As this humid air moves inland, it warms even more, rises, and frequently condenses into cumuliform clouds, which produce afternoon showers and thunderstorms. You can count on thunderstorms developing along the Gulf Coast almost every summer afternoon. As evening approaches, thunderstorm activity typically dies off. Nighttime cooling lowers the temperature of this hot, muggy air only slightly. Should the air become saturated, fog or low clouds usually form, and these normally dissipate by late morning as surface heating warms the air again. A weak, but often persistent, flow around an upperlevel anticyclone in summer will spread warm, humid tropical air from the Gulf of Mexico or from the Gulf of California into the southern and central Rockies, where it causes afternoon thunderstorms. Occasionally, this easterly flow can work its way even farther west, producing shower activity in the otherwise dry southwestern desert.

Humid subtropical air that originates over the southeastern Pacific and Gulf of California remains south of the United States during most of the year. However, during the summer monsoon period, a weak upper-level southerly flow can spread this humid air northward into the southwestern United States, most often in Arizona, Nevada, and southern California. In many places, the moist, conditionally unstable air aloft only shows up as middle and high cloudiness. However, where the moist flow meets a mountain barrier, it tends to rise and condense into towering shower-producing clouds. If the air is sufficiently moist, it can cause heavy showers over a broad area. (For an illustration of exceptionally strong flow of subtropical air into this region, see Fig. 7.22 on p. 187.) Continental Tropical (cT) Air Masses The only real source region for hot, dry continental tropical air masses in North America is found during the summer in northern Mexico and the adjacent arid southwestern United States (see Fig. 8.2, p. 211). Here, the air mass is hot, dry, and conditionally unstable at low levels, with frequent dust devils forming during the day. Because of the low relative humidity (typically less than 10 percent during

DID YOU KNOW? A continental tropical air mass, stretching from southern California to the heart of Texas, brought record warmth to the Desert Southwest during the last week of June 1990. The temperature, which on June 26 soared to a sweltering all-time record high of 122°F in Phoenix, Arizona, caused officials to suspend aircraft takeoffs at Sky Harbor Airport because of uncertainty over how planes would perform in such conditions. AIR MASSES, FRONTS, AND MIDDLE-LATITUDE CYCLONES

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unstable, cool maritime air accompanied by widely scattered showers dominates the weather for several days or more. The real weather action, however, usually occurs not within air masses but at their margins, where air masses with sharply contrasting properties meet—in the zone marked by weather fronts.*

BRIEF REVIEW

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Before we examine fronts, here is a review of some of the important facts about air masses:

FIGURE 8.12 From July 14 through July 22, 2005, continental tropical air covered a large area of the southwestern United States. Numbers on the map represent maximum temperatures (°F) during this period. The large H with the isobar shows the upper-level position of the subtropical high. Sinking air associated with the high contributed to the hot, dry weather. Winds aloft were weak, with the main flow over central Canada.

the afternoon), air must rise more than 3000 m (10,000 ft) before condensation begins. Furthermore, an upper-level ridge usually produces sinking air over the region, tending to make the air aloft rather stable and the surface air even warmer. Consequently, skies are generally clear, the weather is hot, and rainfall is practically nonexistent where continental tropical air masses prevail. If this air mass moves outside its source region and into the Great Plains and stagnates over that region for any length of time, a severe drought may result. Fig. 8.12 shows an instance where continental tropical air covered a large portion of the southwestern United States and produced hot, dry weather during July 2005. So far, we have examined the various air masses that enter North America annually. The characteristics of each depend upon the air mass source region and the type of surface over which the air mass moves. The winds aloft determine the trajectories of these air masses. Occasionally, an air mass will control the weather in a region for some time. These persistent weather conditions are sometimes referred to as airmass weather. Airmass weather is especially common in the southeastern United States during summer as, day after day, humid subtropical air from the Gulf brings sultry conditions and afternoon thunderstorms. It is also common in the Pacific Northwest in winter when conditionally 220

An air mass is a large body of air whose properties of temperature and humidity are fairly similar in any horizontal direction.

Source regions for air masses tend to be generally flat, of uniform composition, and in an area of light winds dominated by surface high pressure.

Continental air masses form over land. Maritime air masses form over water. Polar air masses originate in cold, polar regions, and extremely cold arctic air masses form over arctic regions. Tropical air masses originate in warm, tropical regions.

Continental polar (cP) air masses are cold and dry; continental arctic (cA) air masses are extremely cold and dry. Continental arctic air masses produce the extreme cold of winter as they move across North America.

Continental tropical (cT) air masses are hot and dry, and are responsible for the heat waves of summer in the western half of the United States.

Maritime polar (mP) air masses are cold and moist; these air masses are responsible for the cold, damp, and often wet weather along the northeast coast of North America, as well as for the cold, rainy winter weather along the west coast of North America.

Maritime tropical (mT) air masses are warm and humid, and are responsible for the hot, muggy weather that frequently plagues the eastern half of the United States in summer.

Fronts Although we briefly looked at fronts in Chapter 1, we are now in a position to study them in depth, which will aid us in forecasting the weather. We will now learn about the general nature of fronts—how they move and what weather patterns are associated with them. A front is the transition zone between two air masses of different densities. Since density differences are most often caused by temperature differences, fronts usually separate air masses with contrasting temperatures. Often, they separate air masses with different humidities as well. Remember that air masses have both horizontal and vertical extent; consequently, the upward extension of a front is referred to as a frontal surface, or a frontal zone. *The word front is used to denote the clashing or meeting of two air masses, probably because it resembles the fighting in Western Europe during World War I, when the term originated.

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FIGURE 8.13 The polar front represents a cold frontal boundary that separates colder air from warmer air at the surface and aloft. The more shallow arctic front separates cold air from extremely cold air.

STATIONARY FRONTS A stationary front has essentially no movement.* On a colored weather map, it is drawn as an alternating red and blue line. Semicircles face toward colder air on the red line and triangles point toward warmer air on the blue line. The stationary front between points A and B in Fig. 8.14 marks the boundary where cold, dense continental polar (cP) air from Canada butts up against the north-south trending Rocky Mountains. Unable to cross the barrier, the cold air shows little *They are usually called quasi-stationary fronts because they can show some movement.

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Figure 8.13 illustrates the vertical extent of two frontal zones—the polar front and the arctic front. The polar front boundary, which extends upward more than 5 km (3 mi), separates warm, humid air to the south from cold polar air to the north. The arctic front, which separates cold air from extremely cold arctic air, is much shallower than the polar front and only extends upward to an

altitude of about one or two kilometers. In the next several sections, as we examine individual fronts on a flat surface weather map, keep in mind that all fronts have horizontal and vertical extent. Figure 8.14 shows a surface weather map illustratillustrat ing four different fronts. Notice that the fronts are assoasso ciated with lower pressure and that the fronts separate differing air masses. As we move from west to east across the map, the fronts appear in the following order: a stationary front between points A and B; a cold front between points B and C; a warm front between points C and D; and an occluded front between points C and L. Let’s examine the properties of each of these fronts.

FIGURE 8.14 A weather map showing surface pressure systems, air masses, fronts, and isobars (in millibars) as solid gray lines. Large arrows in color show airflow. (Green-shaded area represents rain; pink-shaded area represents freezing rain and sleet; white-shaded area represents snow.) AIR MASSES, FRONTS, AND MIDDLE-LATITUDE CYCLONES Copyright 2018 Cengage Learning. All Rights Reserved. May not be copied, scanned, or duplicated, in whole or in part. WCN 02-200-202

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or no westward movement. The stationary front is drawn along a line separating the continental polar air from the milder, more humid maritime polar air to the west. Notice that the surface winds tend to blow parallel to the front, but in opposite directions on either side of it. Moreover, upper-level winds often blow parallel to a stationary front. The weather along the front is clear to partly cloudy, with much colder air lying on its eastern side. Because both air masses are relatively dry, there is no precipitation. This is not, however, always the case. When warm, moist air rides up and over the cold air, widespread cloudiness with light precipitation can cover a vast area. These are the conditions that prevail north of the east-west running stationary front depicted in Fig. 8.8, p. 217. In some cases when a stationary front butts up against a mountain range, as shown in Fig. 8.14, winds blowing upslope can generate light rain or snow (called upslope precipitation) if there is enough moisture in the air. If the colder air to the east begins to retreat and is replaced by the warmer air to the west, the front in Fig. 8.14 will no longer remain stationary; it will become a warm front. If, on the other hand, the colder air slides up over the mountain and replaces the warmer air on the other side, the front will become a cold front. If either a cold front or a warm front stops moving, it becomes a stationary front.

. sharp temperature changes over a relatively short distance *Locating any front on a weather map is not always a clear-cut process. Even meteorologists can disagree on an exact position.

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In Fig. 8.15, we can see a large contrast in air temperature and dew point on either side of the front. There is also a wind shift from southwesterly ahead of the front, to northwesterly behind it. Notice that each isobar kinks as it crosses the front, forming an elongated area of low pressure—a trough—which accounts for the wind shift. Since surface winds normally blow across the isobars toward lower pressure, we find winds with a southerly component ahead of the front and winds with a northerly component behind it. Since the cold front is a trough of low pressure, sharp changes in pressure can be significant in locating the front’s position. One important fact to remember is that the lowest pressure usually occurs just as the front passes a station. Notice that, as you move toward the front, the pressure drops, and, as you move away from it, the pressure rises. The precipitation pattern along the cold front in Fig. 8.15 might appear similar to the Doppler radar image shown in Fig. 8.16. The region in color extending from northeast to southwest represents precipitation along a cold front. Notice that light-to-moderate rain (colored in green) occurs over a wide area along the front, while the heavier precipitation (yellow) tends to occur in a narrow

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COLD FRONTS The cold front between points B and C on the surface weather map in Fig. 8.14 represents a zone where cold, dry, stable polar air is replacing warm, moist, conditionally unstable subtropical air. The front is drawn as a solid blue line with the triangles along the front showing its direction of movement. How did the meteorologist know to draw the front at that location? A closer look at the situation will give us the answer. The weather in the immediate vicinity of this cold front in the southern United States is shown in Fig. 8.15. The data plotted on the map represent the current weather at selected cities. The station model used to represent the data at each reporting station is a simplified one that shows temperature, dew point, present weather, cloud cover, sealevel pressure, and wind direction and speed. The little line in the lower right-hand corner of each station shows the pressure change—the pressure tendency, whether rising (/) or falling (\)—during the last three hours. With all of this information, the front can be properly located.* We can use the following criteria to locate a front on a surface weather map:

. changes in the air’s moisture content (as shown by marked changes in the dew point) . shifts in wind direction . pressure and pressure changes . clouds and precipitation patterns

FIGURE 8.15 A closer look at the surface weather associated with the cold front situated in the southern United States in Fig. 8.14. (Gray lines are isobars. Green-shaded area represents rain; white-shaded area represents snow.)

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DID YOU KNOW?

NOAA

One of the strongest cold fronts ever observed in the United States swept across the Great Plains and Midwest on November 10–11, 1911. On November 10, the air temperature in Rapid City, South Dakota, dropped an incredible 75°F in just two hours—from 62°F at 6 p.m. to –13°F at 8 p.m. On November 11, Oklahoma City set a record high of 83°F before the front passed, followed by a record low of 17°F just before midnight. More than a century later, both records were still standing.

FIGURE 8.16 A Doppler radar image showing precipitation patterns along a cold front similar to the cold front in Fig. 8.15. Green represents light-to-moderate precipitation; yellow represents heavier precipitation; and red the most likely areas for thunderstorms. (The cold front is superimposed on the radar image.)

band along the front itself. Strong thunderstorms (red) are found only in certain areas along the front. The cloud and precipitation patterns in Fig. 8.15 along the line X–X’ are shown in a side view of the front in Fig. 8.17. We can see in Fig. 8.17 that, at the front, the cold, dense air wedges under the warm air, forcing the warm air upward, much like a snow shovel forces snow upward as it glides through the snow. As the moist, conditionally unstable air rises, it condenses into a series of cumuliform clouds. Strong, upper-level westerly winds blow the delicate ice crystals (which form near the top of the cumulonimbus) into cirrostratus (Cs) and cirrus (Ci). These clouds usually appear far in advance of the approaching front. At the front itself, a relatively narrow band of thunderstorms (Cb) produces heavy showers with gusty winds. Behind the front, the air cools quickly. (Notice how the

freezing level dips as it crosses the front.) The winds shift from southwesterly to northwesterly, pressure rises, and precipitation ends. As the air dries out, the skies clear, except for a few lingering cumulus clouds. Observe that the leading edge of the front is steep. The steepness is due to friction, which slows the airflow near the ground. The air aloft pushes forward, blunting the frontal surface. Even so, the leading edge of the front does not extend very high into the atmosphere. If we could walk from where the front touches the surface back into the cold air, a distance of 50 km, the front would be about 1 km above us. Thus, the slope of the front—the ratio of vertical rise to horizontal distance—is 1:50. This is typical for a fast-moving cold front, a front that moves at about 25 knots or more. In a slower-moving cold front—one that moves at about 15 knots—the slope is much more gentle. With slow-moving cold fronts, clouds and precipitation usually cover a broad area behind the front. When the ascending warm air is stable, stratiform clouds, such as nimbostratus, become the predominant cloud type and fog may even develop in the rainy area. Occasionally, along a fast-moving front, a line of active showers and thunderstorms, called a squall line, develops parallel to and often ahead of the advancing front. Scattered thunderstorms may also occur well ahead of a cold front. So far, we have considered the general weather patterns of “typical” cold fronts. There are, of course, many exceptions. In fact, no two fronts are exactly alike. In some,

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FIGURE 8.17 A model representing vertical view of a model representing weather across the cold front in Fig. 8.15 along the line X–X´.

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FIGURE 8.18 A “back door” cold front moving into New England during the spring. Notice that, behind the front, the weather is cold and damp with drizzle, while to the south, ahead of the front, the weather is partly cloudy and warm.

the cold air is shallow; in others, it is much deeper. If the rising warm air is dry and stable, scattered clouds are all that form, and there is no precipitation. In extremely dry weather, a marked change in the dew point, accompanied by a slight wind shift, may be the only clue to a passing cold front. During the winter, a series of cold polar or arctic outbreaks may travel across the United States so quickly that warm air is unable to develop ahead of the front. In this case, frigid arctic air associated with the arctic front usually replaces cold polar air, and a drop in temperature is the only indication that a cold front has moved through your area. Along the West Coast, the Pacific Ocean modifies the air so much that typical cold fronts, such as those ▼ Table 8.2

described in the previous section, are rarely if ever seen. In fact, as a cold front moves inland from the Pacific Ocean, the surface temperature contrast across the front may be quite small. Topographic features usually distort the wind pattern so much that locating the position of the front and the time of its passage are exceedingly difficult. In this case, the pressure tendency is the most reliable indication of a frontal passage. Most cold fronts move toward the south, southeast, or east. But sometimes they will move southwestward out of Canada into the northeastern United States. Cold fronts that move in from the east, or northeast, are called “back door” cold fronts. Typically, as the front passes, westerly surface winds shift to easterly or northeasterly, and temperatures drop (see Fig. 8.18). Even though cold-front weather patterns have many exceptions, learning these patterns can be to your advantage if you live in an area that experiences well-defined cold fronts. Knowing them improves your own ability to make short-range weather forecasts. For your reference, ▼ Table 8.2 summarizes idealized cold-front weather in winter in the Northern Hemisphere. WARM FRONTS In Fig. 8.14, p. 221, a warm front is drawn along the solid red line running from points C to D. Here, the leading edge of advancing warm, moist subtropical air from the Gulf of Mexico replaces the retreating cold maritime polar air from the North Atlantic. The direction of frontal movement is given by the half circles, which point into the cold air; this front is heading toward the northeast. As the cold air recedes, the warm front slowly advances. The average speed of a warm front is about 10 knots, or about half that of an average cold front. During the day, as mixing occurs on both sides of the front, its movement may be much faster. Warm fronts often move in a series of rapid jumps, which show up on

Typical Weather Conditions Associated with a Cold Front in Winter in the Northern Hemisphere

WEATHER ELEMENT

BEFORE PASSING

WHILE PASSING

AFTER PASSING

Winds

South or southwest

Gusty, shifting

West or northwest

Temperature

Warm

Sudden drop

Steadily dropping

Pressure

Falling steadily

Minimum, then sharp rise

Rising steadily

Clouds

Increasing Ci, Cs, then either Tcu* or Cb*

Tcu or Cb

Often Cu, Sc* when ground is warm

Precipitation

Short period of showers

Heavy showers of rain or snow, sometimes with hail, thunder, and lightning

Decreasing intensity of showers, then clearing

Visibility

Fair to poor in haze

Poor, followed by improving

Good, except in showers

Dew point

High; remains steady

Sharp drop

Lowering

*Tcu stands for towering cumulus, such as cumulus congestus; whereas Cb stands for cumulonimbus. Sc stands for stratocumulus.

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FIGURE 8.19 Surface weather associated with a typical warm front in winter. A vertical view along the dashed line P–P’ is shown in Fig. 8.20. (Green-shaded area represents rain; pinkshaded area represents freezing rain and sleet; white-shaded area represents snow.)

successive weather maps. At night, however, radiational cooling creates cool, dense surface air behind the front. This inhibits both lifting and the front’s forward progress. When the forward surface edge of the warm front passes a station, the wind shifts, the temperature rises, and the overall weather conditions improve. To see why, we will examine the weather commonly associated with the warm front both at the surface and aloft. Figure 8.19 is a surface weather map showing the position of a warm front in winter and its associated weather. Figure 8.20 is a vertical view of a model of the warm front in Fig. 8.19. Look at these two figures and observe that, in the model, the warmer, less-dense air rides

up and over the colder, more-dense surface air. This rising of warm air over cold (called overrunning in this model) produces clouds and precipitation well in advance of the front’s surface boundary. The warm front that separates the two air masses has an average slope of about 1:300—a much more gentle or inclined shape than that of a typical cold front.* Suppose we are standing at the position marked P´ in Fig. 8.19 and Fig. 8.20. Note that we are over 1200 km (750 mi) ahead of where the warm front is touching the surface. Here, the surface winds are light and variable, the air is cold, and about the only indication of an approaching warm front is the high cirrus clouds overhead. We know the front is moving slowly toward us and that within a day or so it will pass our area. Suppose that, instead of waiting for the front to pass us, we drive toward it, observing the weather as we go. Heading toward the warm front, we notice that the cirrus (Ci) clouds gradually thicken into a thin, white veil of cirrostratus (Cs) whose ice crystals cast a halo around the sun.** Almost imperceptibly, the clouds thicken and lower, becoming altocumulus (Ac) and altostratus (As) through which the sun shows only as a faint spot against an overcast gray sky. Snowflakes begin to fall, and we are still over 600 km (370 mi) from the surface front. The snow increases, and the clouds thicken into a sheetlike covering of nimbostratus (Ns). The winds become brisk out of the southeast, while the atmospheric pressure slowly falls. Within 400 km (250 mi) of the front, the cold surface air mass is now quite shallow. The surface air temperature moderates and, as we approach the front, the light snow changes first into sleet. It then becomes freezing rain and finally rain and drizzle as the air temperature climbs above freezing. Overall, the precipitation remains light or moderate but covers a broad area. Moving still closer to the front, warm, moist air mixes with cold, moist air producing ragged windblown stratus (St) and fog. (As you might deduce, flying in the vicinity of a warm front can be quite hazardous.) *This slope of 1:300 is a more gentle slope than that of most warm fronts. Typically, the slope of a warm front is on the order of 1:150 to 1:200. **If the warm air is relatively unstable, ripples or waves of cirrocumulus clouds will appear as a “mackerel sky.”

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FIGURE 8.20 A model illustrating a vertical view of clouds, precipitation, and winds across the warm front in Fig. 8.19 along the line P–P.

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▼ Table 8.3

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Finally, after a trip of over 1200 km, we reach the warm front’s surface boundary. As we cross the front, the weather changes are noticeable, but much less pronounced than those experienced with the cold front; they show up more as a gradual transition rather than a sharp change. On the warm side of the front, the air temperature and dew point rise, the wind shifts from southeast to south or southwest, and the atmospheric pressure stops falling. The light rain ends and, except for a few stratocumulus, the fog and low clouds vanish. This scenario of an approaching warm front represents average, if not idealized, warm-front weather in winter. In some instances, the weather can differ from this example dramatically. For example, if the overrunning warm air is relatively dry and stable, only high and middle clouds will form, and no precipitation will occur. On the other hand, if the warm air is relatively moist and conditionally unstable (as is often the case during the summer), heavy showers can develop as thunderstorms become embedded in the cloud mass. Some of these thunderstorms may have bases at a relatively high level above the surface; these are called elevated storms. Along the West Coast, the Pacific Ocean significantly modifies the surface air so that warm fronts are difficult to locate on a surface weather map. Also, not all warm fronts move northward or northeastward. On rare occasions, a warm front will move into the eastern seaboard from the Atlantic Ocean as it spins all the way around a deep cyclonic storm positioned off the coast. Cold northeasterly winds ahead of the front usually become warm northeasterly winds behind it. Even with these exceptions, knowing the normal sequence of warm-front weather can be useful, especially if you live where warm fronts become well developed. You can look for certain cloud and weather patterns and make reasonably accurate short-range forecasts of your own. ▼ Table 8.3 summarizes typical winter warm-front weather.

FIGURE 8.21 A dryline represents a narrow boundary where there is a steep horizontal change in moisture as indicated by a rapid change in dew-point temperature. Here, a dryline moving across Texas and Oklahoma separates warm, moist air from warm, dry air during an afternoon in May.

DRYLINES In the southern Great Plains, warm, humid air may be separated from warm, dry air along a boundary called a dryline. Because dew-point temperatures may drop along this boundary by as much as 10°C (18°F) per kilometer, drylines have been referred to as dew-point fronts.* Although drylines can occur in the United States as far north as the Dakotas, and as far east as the TexasLouisiana border, they are most frequently observed in the western half of Texas, Oklahoma, and Kansas, especially during spring and early summer. In these locations, drylines tend to move eastward during the day, then westward dry toward evening. Figure 8.21 shows a well-developed dryline moving across Texas and Oklahoma during May 2001. Cumulus clouds and thunderstorms often form along or to the east of the dryline. This cloud development is *Recall from Chapter 4 that the dew-point temperature is a measure of the amount of water vapor in the air.

Typical Weather Conditions Associated with a Warm Front in the Northern Hemisphere

WEATHER ELEMENT

BEFORE PASSING

WHILE PASSING

AFTER PASSING

Winds

South or southeast

Variable

South or southwest

Temperature

Cool to cold, slow warming

Steady rise

Warmer, then steady

Pressure

Usually falling

Leveling off

Slight rise, followed by fall

Clouds

In this order: Ci, Cs, As, Ns, St, and fog; occasionally Cb in summer

Stratus-type

Clearing with scattered Sc, especially in summer; occasionally Cb in summer

Precipitation

Light-to-moderate rain, snow, sleet, or drizzle; showers in summer

Drizzle or none

Usually none; sometimes light rain or showers

Visibility

Poor

Poor, but improving

Fair in haze

Dew point

Steady rise

Steady

Rise, then steady

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OCCLUDED FRONTS If a cold front catches up to and overtakes a warm front, the frontal boundary created between the two air masses is called an occluded front, or, simply, an occlusion (meaning “closed off ”). There are two main types of occlusions, cold occlusions and warm occlusions. On a surface weather map, an occluded front is represented as a purple line with alternating cold-front triangles and warm-front half circles; both symbols point in the direction toward which the front is moving. Look back at Fig. 8.14, p. 221, and notice that the air behind the occluded front is colder than the air ahead of it. This is known as a cold-type occluded front front, or cold occlusion. Let’s see how this front develops. The classic model of a developing cold-type occluded front is shown in Fig. 8.22. Along line A–A’, the cold front is rapidly approaching the slower-moving warm front. Along B–B’, the cold front overtakes the warm front, and as we can see in the vertical view across C–C’, it underrides and lifts both the warm front and the warm air mass off the ground. This situation allows colder air to replace cool air at the surface. As a cold occluded front approaches, the weather sequence is similar to that of a warm front, with high clouds lowering and thickening into middle and low clouds and precipitation forming well in advance of the surface front. Since the front represents a trough of low pressure, southeasterly winds and falling atmospheric pressure occur ahead of it. The frontal passage, however, can bring weather similar to that of a cold front: heavy, showery precipitation, with winds shifting to west or northwest. After a period of wet weather, the sky begins to clear, atmospheric pressure rises, and the air turns colder. The most active weather usually occurs where the cold front is just overtaking the warm front, at the point of occlusion, where the greatest contrast in temperature occurs. The classic model of a developing warm-type occontinen cluded front is shown in Fig. 8.23. Notice that continental polar air over eastern Washington and Oregon is much colder than milder maritime polar air moving inland from the Pacific Ocean. Observe also that the surface air ahead of the warm front is colder than the air behind the cold

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caused in part by daytime convection and a sloping terrain. The Central Plains area of North America is higher to the west and lower to the east. Convection over the elevated western plains carries dry air high above the surface. Westerly winds sweep this dry air eastward over the lower plains where it overrides the slightly cooler but more humid air at the surface. This situation sets up a potentially unstable atmosphere that finds warm, dry air above warm, moist air. In regions where the air rises, cumulus clouds and organized bands of thunderstorms can form. We will look more closely at drylines and their effect on developing thunderstorms in Chapter 10.

FIGURE 8.22 A model that shows the formation of a coldoccluded front. The faster-moving cold front in (a) catches up to the slower-moving warm front in (b) and forces it to rise off the ground in (c). (Green-shaded area on the map represents precipitation.)

front. Consequently, when the cold front catches up to and overtakes the warm front, the milder, lighter air behind the cold front is unable to lift the colder, heavier air off the ground. As a result, the cold front rides “piggyback” along the sloping warm front, producing a warm-type occluded front. The surface weather associated with the warm AIR MASSES, FRONTS, AND MIDDLE-LATITUDE CYCLONES

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differences ahead of the advancing warm front and behind the cold front has been challenged. Studies show that the stability of the air (rather than the surface air temperature) is the most important factor in determining the type of occluded front that forms. The air within the warm frontal zone is typically more stable than the air within the cold frontal zone, so the air behind the cold front is more likely to ride up and over the warm front than to flow beneath it. For this reason, warm-type occluded fronts are far more common than cold-type occluded fronts. In the world of weather fronts, occluded fronts are the mavericks. In fact, they may show up on a surface chart as a trough of low pressure separating two cold air masses. Because of this, locating and defining occluded fronts at the surface is often difficult for the meteorologist.* Similarly, you too may find it hard to recognize an occlusion. In spite of this, we will assume that the weather associated with occluded fronts in North America behaves in a similar way to that shown in ▼ Table 8.4. The frontal systems described so far are actually part of a much larger storm system: the middle-latitude cyclone. Figure 8.24 shows the cold front, warm front,

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*In some countries, such as Canada, the surface occlusion is seldom analyzed on a surface weather map. Instead, the location of the occluded front aloft— where the cold air lifts the warm air above the surface—is marked by a trowal (which stands for “trough of warm air aloft”). So the position of the trowal on the weather map (as indicated by a “hook”) marks the location where the cold and warm fronts intersect aloft.

occlusion is similar to that of a warm front. In addition, the relatively mild winter air that moves into Europe from the North Atlantic causes many of the occlusions that move into this region in winter to be of the warm variety. Contrast Fig. 8.22 and Fig. 8.23. Note that the primary difference between the cold- and warm-type occluded front is the location of the upper-level front. In a cold occlusion, there is an upper-level warm front that follows the surface occluded front, whereas in a warm occlusion, there is an upper-level cold front that precedes the surface occluded front. So far, the ideas we have presented on the formation of occluded fronts are based on the surface air temperature on either side of the front. In recent years, the classic model of cold and warm occluded fronts forming due to temperature 228

NOAA

FIGURE 8.23 A model illustrating the formation of a warmtype occluded front. The faster-moving cold front in (a) over overtakes the slower-moving warm front in (b). The lighter air behind the cold front rises up and over the denser air ahead of the warm front. The bottom illustration represents a surface map of a warm occlusion.

FIGURE 8.24 A visible satellite image showing a mid-latitude cyclonic storm with its weather fronts over the Atlantic Ocean during March 2005. Superimposed on the image is the position of the surface cold front, warm front, and occluded front. Precipitation symbols indicate where precipitation is reaching the surface.

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▼ Table 8.4

Typical Winter Weather Most Often Associated with Occluded Fronts in North America

WEATHER ELEMENT

BEFORE PASSING

WHILE PASSING

AFTER PASSING

Winds

East, southeast, or south

Variable

West or northwest

(a) Cold-type occluded

Cold or cool

Dropping

Colder

(b) Warm-type occluded

Cold

Rising

Milder

Pressure

Usually falling

Low point

Usually rising

Clouds

In this order: Ci, Cs, As, Ns

Ns, sometimes Tcu and Cb

Ns, As, or scattered Cu

Precipitation

Light, moderate, or heavy precipitation

Light, moderate, or heavy continuous precipitation or showers

Light-to-moderate precipitation followed by general clearing

Visibility

Poor in precipitation

Poor in precipitation

Improving

Dew point

Steady

Usually slight drop, especially if cold-occluded

Slight drop, although may rise a bit if warm-occluded

Temperature

and occluded front with such a storm. Notice, as we would expect, clouds and precipitation form in a rather narrow band along the cold front and in a much wider band along the warm front and the occluded front. The next section explains where, why, and how midlatitude cyclones form.

Mid-Latitude Cyclonic Storms Early weather forecasters were aware that precipitation generally accompanied falling barometers and areas of low pressure. However, it was not until the early part of the twentieth century that scientists began to piece together the information that yielded the ideas of modern meteorology and cyclonic storm development. Working largely from surface observations, a group of scientists in Bergen, Norway, developed a model explaining the life cycle of an extratropical, or middle-latitude cyclonic storm; that is, a storm that forms at middle and high latitudes outside of the tropics. This extraordinary group of meteorologists included Vilhelm Bjerknes, his son Jakob, Halvor Solberg, and Tor Bergeron. They published their Norwegian Cyclone Model shortly after World War I. It was widely acclaimed and became known as the “polar front theory of a developing wave cyclone,” or, simply, the polar front theory. What these meteorologists gave to the world was a working model of how a mid-latitude cyclone progresses through the stages of birth, growth, and decay. An important part of the model involved the development of weather along the polar front. As new information became available, the original work was modified, so that, today, it serves as a convenient way to describe the structure

and weather associated with a migratory middle-latitude cyclonic storm system, such as the one shown in Fig. 8.24. POLAR FRONT THEORY The development of a midlatitude cyclone, according to the Norwegian model, begins along the polar front. Remember (from our discussion of the general circulation in Chapter 7) that the polar front is a semicontinuous global boundary separating cold polar air from warm subtropical air. Because the midlatitude cyclone forms and moves along the polar front in a wavelike manner, the developing storm is referred to as a wave cyclone. The stages of a developing wave cyclone from a surface perspective are illustrated in the sequence of surface weather maps shown in Fig. 8.25. Figure 8.25a shows a segment of the polar front as a stationary front. It represents a trough of lower pressure with higher pressure on both sides. Cold air to the north and warm air to the south flow parallel to the front, but in opposite directions. This type of flow sets up a cyclonic wind shear. You can conceptualize the shear more clearly if you place a pen between the palms of your hands and move your left hand toward your body; the pen turns counterclockwise, cyclonically. Under the right conditions, a wavelike kink forms on the front, as shown in Fig. 8.25b. The wave that forms is known as a frontal wave. Watching the formation of a frontal wave on a weather map is like watching a water wave from its side as it approaches a beach: It first builds, then breaks, and finally dissipates. This pattern is the reason a mid-latitude cyclonic storm system is known as a wave cyclone. Figure 8.25b shows the newly formed wave with a cold front pushing southward and a warm front moving AIR MASSES, FRONTS, AND MIDDLE-LATITUDE CYCLONES

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FIGURE 8.25 The idealized life cycle of a mid-latitude cyclonic storm (a through f) in the Northern Hemisphere based on the polar front theory. As the life cycle progresses, the system moves northeastward in a dynamic fashion. The small arrow next to each L shows the direction of storm movement.

northward. The region of lowest pressure is at the junction of the two fronts. As the cold air displaces the warm air upward along the cold front, and as warm air rises ahead of the warm front, a narrow band of precipitation forms (shaded green area). Steered by the winds aloft, the system typically moves east or northeastward and gradually becomes a fully developed open wave in 12 to 24 hours (Fig. 8.25c). The central pressure of the wave cyclone is now much lower, and several isobars encircle the wave’s apex. These more tightly packed isobars create a stronger cyclonic flow, as the winds swirl counterclockwise and inward toward the low’s center. Precipitation forms in a wide band ahead of the warm front and along a narrow band of the cold front. The region of warm air between the cold and warm fronts is known as the warm sector. Here, the weather tends to be partly cloudy, although scattered showers and thunderstorms may develop if the air is conditionally unstable. Energy for the storm is derived from several sources. As the air masses try to attain equilibrium, warm air rises and cold air sinks, transforming potential energy into kinetic energy (that is, energy of motion). Condensation supplies energy to the system in the form of latent heat. And, as the surface air converges toward the low center, wind speeds may increase, producing an increase in kinetic energy. As the open wave moves eastward, central pressures continue to decrease, and the winds blow more vigorously. The faster-moving cold front constantly inches closer to the warm front, squeezing the warm sector into a smaller area (as shown in Fig. 8.25d), and the wave quickly develops into a mature cyclone. In this model, the cold front eventually overtakes the warm front and the system becomes occluded. At this point, the storm is usually most intense, with clouds and precipitation covering a large area. The area of most intense weather is normally found to the northwest of the storm’s center. Here, strong winds 230

and blowing and drifting snow can create blizzard conditions in winter. The intense storm system shown in Fig. 8.25e gradually dissipates, because cold air now lies on both sides of the occluded front. Without the supply of energy provided by the rising warm, moist air, the old storm system dies out and gradually disappears (Fig. 8.25f). Occasionally, however, a new wave will form on the westward end of the trailing cold front. We can think of the sequence of a developing wave cyclone as a whirling eddy in a stream of water that forms behind an obstacle, moves with the flow, and gradually vanishes downstream. The entire life cycle of a wave cyclone can last from a few days to more than a week. Take a second and look back at the mid-latitude cyclonic storm depicted in the satellite image in Fig. 8.24. According to what you have just read, what is the stage of development of this storm? (Answer given in footnote below.*) Figure. 8.26 shows a series of wave cyclones at various stages of development along the polar front in winwin ter. Such a succession of storms is known as a “family” of cyclones. Observe that to the north of the front are cold anticyclones; to the south over the Atlantic Ocean is the warm, semipermanent Bermuda high. The polar front itself has developed into a series of loops, and at the apex of each loop is a cyclone. The cyclone over the northern plains (Low 1) is just forming; the one along the east coast (Low 2) is an open wave; and the system near Iceland (Low 3) is dying out. If the average rate of movement of a wave cyclone from birth to decay is 25 knots, then it is entirely possible for a storm to develop over the central part of the United States, intensify into a large storm over New England, become occluded over the ocean, and reach the coast of England in its dissipating stage less than a week after it forms. *The storm shown in Fig. 8.24 is in its occluded stage, and would be classified as a mature cyclone.

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FIGURE 8.26 A series of wave cyclones (a “family” of cyclones) forming along the polar front.

Up to now, we have considered the polar front model of a developing wave cyclone, which represents a rather simplified version of the stages that a mid-latitude cyclonic storm system must go through. In fact, even though few (if any) storms adhere to the model exactly, it still can serve as a good foundation for understanding the structure of cyclonic storms. So keep the model in mind as you read the following sections. WHERE DO MID-LATITUDE CYCLONES TEND TO FORM? Any development or strengthening of a midlatitude cyclone is called cyclogenesis. Certain regions of North America show a propensity for cyclogenesis, including the eastern slopes of the Rockies. As westerly winds blow over a mountain range, the air expands vertically on the downwind (lee) side, which can help intensify any pre-existing areas of low pressure. Troughs and developing cyclonic storms that form in this manner are called ex lee-side lows (see Fig. 8.27). Additional areas that exhibit cyclogenesis are the Great Basin, the Gulf of Mexico, and the Atlantic Ocean east of the Carolinas. Near Cape

DID YOU KNOW? In mid-January 1888, a ferocious mid-latitude cyclonic storm swept across the Great Plains from Texas to the Dakotas and into Wisconsin. Strong winds, extremely low temperatures, and heavy snow on the storm’s western side wiped out the Plains’ free-range livestock and took 237 lives. This infamous storm has come to be known as the “Children’s Blizzard” because of the many dozens of schoolchildren frozen to death on their way to school.

Hatteras, North Carolina, for example, warm Gulf Stream water can supply moisture and warmth to the region south of a stationary front, thus increasing the contrast between air masses to a point where storms can suddenly spring up along the front. As noted earlier, these cyclones normally move northeastward along the Atlantic coast, bringing high winds and heavy snow or rain to coastal areas. Before the age of modern satellite imagery and weather prediction such coastal storms would often go undetected during their formative stages; and sometimes an evening weather forecast of “fair and colder” along the eastern seaboard would have to be changed to “heavy snowfall” by morning. Fortunately, with today’s weather information gathering and forecasting techniques, these storms rarely strike by surprise. (Storms that form along the eastern seaboard of the United States and then move northeastward are called nor’easters or northeasters. Additional information on nor’easters is given in Focus section 8.3) Figure 8.28 shows the typical paths taken in winter by mid-latitude cyclones and anticyclones (high-pressure areas). Notice in Fig. 8.28a that some of the lows are named after the region where they form, such as the Hatteras Low, which develops off the coast near Cape Hatteras, North Carolina. The Alberta Clipper forms (or redevelops) on the eastern side of the Canadian Rockies in Alberta, then rapidly skirts across the northern tier states. Similarly, the Colorado Low forms (or redevelops) on the eastern side of the United States Central Rockies. Notice that the lows generally move eastward or northeastward, whereas the highs (Fig. 8.28b) typically move southeastward, then eastward. AIR MASSES, FRONTS, AND MIDDLE-LATITUDE CYCLONES

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are many surface conditions that influence the formation of a mid-latitude cyclonic storm, including mountain ranges and land-ocean temperature contrasts. However, the real key to the development of a wave cyclone is found in the upper-wind flow, in the region of the high-level westerlies. Therefore, before we can arrive at a reasonable answer to our question, we need to see how the winds aloft influence surface pressure systems. DEVELOPING MID-LATITUDE CYCLONES AND ANTICYCLONES In Chapter 7, we learned that thermal pressure systems are shallow systems that are typically weaker with increasing height above the surface. On the other hand, developing mid-latitude cyclonic storms are dynamic lows that are usually stronger with height. This means that a surface low-pressure area will appear on an upper-level chart as either a closed low or a trough. Suppose the upper-level low is directly above the surface low, as illustrated in Fig. 8.29. Notice that only at the surface (because of friction) do the winds blow inward toward the low’s center. As these winds converge (flow together), the air “piles up.” This piling up of air, called convergence, causes air density to increase directly above the surface low. This increase in mass causes surface pressures to rise; gradually, the low fills and the surface low dissipates. The same reasoning can be applied to surface anticyclones. Winds blow outward away from the center of a surface high. If a closed high or ridge lies directly over the surface anticyclone, divergence (the spreading out of air) at the surface will remove air from the column directly above the high. The decrease in mass causes the surface pressure to fall and the surface high-pressure area to weaken. Consequently, it appears that, if upper-level pressure systems were always located directly above those at the surface (such as shown in Fig. 8.29), cyclones and anticyclones would die out soon after they form (if they could form at all). What, then, is it that allows these systems to develop and intensify? (Before reading on, you may wish to review the additional information on convergence and divergence given in Focus section 8.4.).

FIGURE 8.27 As westerly winds blow over a mountain range, the air can expand vertically on the downwind (lee) side, enhancing the development of a trough or cyclonic storm, called a lee-side low.

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Some frontal waves form suddenly, grow in size, and develop into huge cyclonic storms, then slowly dissipate, with the entire process taking several days to a week to complete. Other frontal waves remain small and never grow into giant weather-producers. Why is it that some frontal waves develop into huge cyclonic storms, whereas others simply dissipate in a day or so? This question poses one of the real challenges in weather forecasting. The answer is complex. Indeed, there

FIGURE 8.28 (a) Typical paths of winter mid-latitude cyclones. The lows are named after the region where they form. (b) Typical paths of winter anticyclones.

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FIGURE 8.29 If lows and highs aloft were always directly above lows and highs at the surface, the surface systems would quickly dissipate.

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FOCUS

ON A SPECIAL TOPIC 8.3 O

Northeasters, commonly called nor’easters, are east coast mid-latitude cyclonic storms that develop or intensify off the eastern seaboard of North America during the fall, winter, and spring. They usually move northeastward along the coast, often bringing strong northeasterly winds to coastal areas, hence the name, nor’easter. In addition to strong winds, these storms can bring heavy rain, snow, and sleet. Most often they deepen and become most intense off the coast of New England. Nor’easters are fueled by the large temperature gradient between the warm ocean and the cold continental landmass. They gain additional energy from the moisture over the ocean, especially the warm Gulf Stream that flows northeastward parallel to the east coast of the United States. The ferocious northeaster of January 22–24, 2016 (shown in Fig. 4), produced the heaviest snow totals on record at several locations, including New York’s Kennedy International Airport (30.6 in.) and Baltimore-Washington International Airport (29.2 in.). Much of the snow fell in less than 24 hours. High winds and 20-foot waves pounded the coastlines of Delaware and New Jersey, and high storm tides put many coastal areas and highways under water. Studies suggest that some of the nor’easters that batter the coastline in winter may actually possess some of the characteristics of a tropical hurricane. For example, a strong nor’easter that dumped between 10 and 40 inches of snow over parts of the Northeast and New England in February 2013 developed an eye-like feature. One of the most dramatic combinations of tropical and extratropical cyclone processes ever observed was Hurricane

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Nor’easters

FIGURE 4 The surface weather map for 7 a. m. (EST), January 23, 2016, shows an intense low-pressure area (central pressure 987 mb, or 29.15 in.), which is generating strong northeasterly winds and heavy precipitation (areas shaded green for rain, white for snow, and salmon for sleet or freezing rain) from the mid-Atlantic states into New England. More than 30 million people were placed under blizzard warnings as a result of this storm, which dropped more than two feet of snow from Washington, D.C., to New York. Damage estimates were as high as $3 billion.

Sandy in late October 2012. Sandy briefly attained Category 3 strength before striking Cuba, weakened, then restrengthened as it approached the northeast United States. The storm grew to an immense size, with some circulation features comparable to a strong nor’easter,

The Role of Converging and Diverging Air For midlatitude cyclones and anticyclones to maintain themselves or intensify, the winds aloft must blow in such a way that zones of converging and diverging air form. For example, notice in Fig. 8.30 that the surface winds are converging about the center of the low; while aloft, directly above the low, the winds are diverging. For the surface low to develop into a major storm system, upper-level divergence of air must be greater than surface convergence of air; that is,

while it also maintained hurricane characteristics until just several hours before it struck New Jersey, when it was reclassified as a post-tropical cyclone. (We will examine hurricanes and their characteristics, including Sandy, in more detail in Chapter 11.)

more air must be removed above the storm than is brought in at the surface. When this phenomenon happens, surface air pressure decreases, and we say that the storm system is intensifying, or deepening. If the reverse should occur (more air flowing in at the surface than is removed at the top), surface pressure will rise, and the storm system will weaken and gradually dissipate in a process called filling. Notice also in Fig. 8.30 that surface winds are diverging about the center of the high, while aloft, directly above AIR MASSES, FRONTS, AND MIDDLE-LATITUDE CYCLONES

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FOCUS

ON A SPECIAL TOPIC 8.4 O

FIGURE 5 The formation of convergence (CON) and divergence (DIV) of air with a constant wind speed (indicated by flags) in the upper troposphere. Circles represent air parcels that are moving parallel to the contour lines on a constant pressure chart. Below the area of convergence the air is sinking, and we find the surface high (H). Below the area of divergence the air is rising, and we find the surface low (L).

marching in a band. When the marchers in front slow down, the rest of the band members squeeze together, causing convergence; when the marchers in front start to run, the band members spread apart, or diverge. In summary, speed convergence takes place when the wind speed decreases downwind, and speed divergence takes

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We know that convergence is the piling up of air above a region, while divergence is the spreading out of air above some region. Convergence and divergence of air may result from changes in wind direction and wind speed. For example, convergence occurs when moving air is funneled into an area, much in the way cars converge when they enter a crowded freeway. Divergence occurs when moving air spreads apart, much as cars spread out when a congested two-lane freeway becomes three lanes. On an upper-level chart, this type of convergence (also called confluence) occurs when contour lines move closer together, as a steady wind flows parallel to them (see the upper-level chart in Fig. 5). On the same chart, this type of divergence (also called difdif fluence) occurs when the contour lines move apart as a steady wind flows parallel to them. Notice that below the area of divergence lies the surface middle-latitude cyclonic storm. Convergence and divergence may also result from changes in wind speed. Speed convergence occurs when the wind slows down as it moves along, whereas speed divergence occurs when the wind speeds up. We can grasp these relationships more clearly if we imagine air molecules to be

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A Closer Look at Convergence and Divergence

FIGURE 8.30 Convergence, divergence, and vertical motions associated with surface pressure systems. Notice that for the surface storm to intensify, the upper trough of low pressure must be located to the left (or west) of the surface low. Gray lines on the upper-air chart are contour lines.

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place when the wind speed increases downwind. As you continue to read about midlatitude cyclonic storms, keep in mind from Fig. 5 that converging air aloft is most likely found to the left (west) of the upper trough, whereas diverging air is most likely found to the right (east) of the trough.

the anticyclone, they are converging. In order for the surface high to strengthen, upper-level convergence of air must exceed low-level divergence of air (more air must be brought in above the anticyclone than is removed at the surface). When this occurs, surface air pressure increases, and we say that the high-pressure area is building. In Fig. 8.29, the convergence of air aloft causes an accumulation of air above the surface high, which allows the air to sink slowly and replace the diverging surface air. Above the surface low, divergence allows the converging surface air to rise and flow out the top of the column. We can see from Fig. 8.30 that when an upper-level trough is as sufficiently deep as is illustrated here, a region of converging air usually forms on the west side of the trough and a region of diverging air forms on the east side. (For reference, compare Fig. 8.30 with Fig. 5 above.) Aloft, the area of diverging air is directly above the surface low, and the area of convergence is directly

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FIGURE 8.31 (a) Upper-air chart showing a longwave with three shortwaves (heavy dashed lines) embedded in the flow. (b) Twenty-four hours later the shortwaves have moved rapidly around the longwave. Notice that the shortwaves labeled 1 and 3 tend to deepen th the longwave troughs, while shortwave 2 has weakened as it moves into a ridge. Dashed red lines are isotherms (lines of constant temperature) in °C. Solid gray lines are contours. Blue arrows indicate cold advection and red arrows, warm advection.

above the surface high. This configuration means that, for a surface mid-latitude cyclone to intensify, the upper-level trough of low pressure must be located behind (or to the west of) the surface low. When the upper-level trough is in this position, the atmosphere is able to redistribute its mass, as regions of low-level convergence are compensated for by regions of upper-level divergence, and vice versa. Winds aloft steer the movement of the surface pressure systems. Since the winds above the surface low in Fig. 8.30 are blowing from the southwest, the surface low should move northeastward. The northwesterly winds above the surface high should direct it toward the southeast. These paths are typical of the average movement of surface pressure systems in the eastern two-thirds of the United States, as shown in Fig. 8.28 on p. 232. Waves in the Westerlies Waves that form in the flow aloft can play a major role in the development of midlatitude cyclones. Recall from Chapter 6 that the flow above the middle latitudes usually consists of a series of waves in the form of troughs and ridges. The distance from trough to trough (or ridge to ridge) is known as the wavelength. When the wavelength is on the order of many thousands of kilometers, the wave is called a longwave.. Observe in Fig. 8.31a that the length of the longwave is greater than the width of North America. Typically, at any given time, there are between three and six longwaves looping around Earth. These longwaves are also known as Rossby waves, after C. G. Rossby, a famous meteorologist who carefully studied their motion. In Fig. 8.31a, we can see that embedded in longwaves are shortwaves, which are small disturbances or ripples that move with the wind flow.

By comparing Fig. 8.31a with Fig. 8.31b, we can see that while the longwaves move eastward very slowly, the shortwaves move fairly quickly around the longwaves. Generally, shortwaves deepen (that is, increase in size) when they approach a longwave trough and weaken (become smaller) when they approach a ridge. Also notice in Fig. 8.31b that when a shortwave moves into a longwave trough, the trough tends to deepen. The upper flow is now capable of providing the necessary ingredients for the development or intensification of surface low- and high-pressure areas as illustrated in Fig. 8.30. Notice in Fig. 8.31b that there are regions where the winds (blue and red arrows) cross the isotherms (dashed red lines). In these regions, there is an important process taking place called temperature advection.* Where the wind crosses the isotherms in such a way that colder air is replacing warmer air, the transport of colder air into a region is called cold advection. In Fig. 8.31b, cold advection is represented by blue arrows. Where the wind crosses the isotherms in such a way that warmer air is replacing colder air, the transport of warmer air into a region is called warm advection. On the map, warm advection is represented by red arrows. Because temperature advection plays a major part in the development of a mid-latitude cyclonic storm, we will examine its important role in the following section. Upper-Air Support for the Developing Storm To better understand how a wave cyclone can develop and intensify into a huge mid-latitude cyclonic storm, we need to examine atmospheric conditions at the surface and aloft. Suppose that a portion of a longwave *Recall from Chapter 2 that the term advection refers to the horizontal transfer of any atmospheric property by the wind. AIR MASSES, FRONTS, AND MIDDLE-LATITUDE CYCLONES

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FIGURE 8.32 An idealized 3-D view of the formation of a mid-latitude cyclone. (a) A longwave trough at 500 mb lies parallel to and directly above the surface stationary front. (b) A shortwave (not shown) disturbs the flow aloft, initiating temperature advecti advection (blue arrow, cold advection; red arrow, warm advection). The upper trough intensifies and provides the necessary vertical motions (as shown by vertical arrows) for the development of the surface wave cyclone. (c) As the surface storm moves northeastward, it occludes, and without upper-level diverging air to compensate for surface converging air, the cyclonic storm system dissipates.

trough at the 500-mb level lies directly above a surface stationary front, as illustrated in Fig. 8.32a. On the 500-mb chart, contour lines (solid lines) and isotherms (dashed lines) parallel each other and are crowded close together. Colder air is located in the northern half of the map, while warmer air is located to the south. Winds at this level are blowing at fairly high velocities. Suppose a shortwave moves through this region, disturbing the flow as shown in Fig. 8.32b. As the flow aloft becomes disturbed, it begins to lend support for the intensification of surface pressure systems, as a region of converging air forms above position 1 in Fig. 8.32b and a region of diverging air forms above position 2. The converging air aloft causes the surface air pressure to rise in the region marked H in Fig. 8.32b. Surface winds begin to blow out away from the region of higher pressure, and the air aloft gradually sinks to replace it. Meanwhile, diverging air aloft causes the surface air pressure to decrease beneath position 2, in the region marked L on the surface map. This initiates rising air, as the surface winds blow in

DID YOU KNOW? The great mid-latitude cyclonic storm of March 1993 that moved northeastward along the east coast of the United States set record low barometric pressure readings in a dozen states. The huge storm produced wind gusts exceeding 90 knots (104 mi/hr) from New England to Florida and deposited 50 billion tons of snow over the east coast of the United States. Some cities that set all-time 24-hour record snowfall totals in the storm include Syracuse, New York (35.6 in.); Beckley, West Virginia (28.2 in.); and Birmingham, Alabama (13 in.). In just those 24 hours, Birmingham received more snow than it had ever recorded across any entire winter.

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toward the region of lower pressure. As the converging surface air develops cyclonic spin, cold air flows southward and warm air northward. We can see in Fig. 8.32b that the western half of the stationary front is now a cold front and the eastern half a warm front. Cold air moves in behind the cold front, while warm air slides up along the warm front. These regions of cold and warm advection occur all the way up to the 500-mb level, about 18,000 feet above sea level. On the 500-mb chart in Fig. 8.32b, cold advection is occurring at position 1 (blue arrow) as the wind crosses the isotherms, bringing cold air into the trough. The cold advection makes the air more dense, which has the effect of deepening the trough. The deepening of the upper trough causes the contour lines to crowd closer together and the winds aloft to increase. Meanwhile, at position 2 warm advection is taking place (red arrows), which has the effect of strengthening the ridge. Therefore, the overall effect of differential temperature advection is to amplify the upper-level wave. As the trough aloft deepens, its curvature increases, which in turn increases the region of divergence above the developing surface storm. At this point, the surface mid-latitude cyclone rapidly develops as surface pressures fall. In regions where there is cold advection, some of the cold, heavy air sinks; where there is warm advection, some of the warm, light air rises. The sinking of cold air and the rising of warm air provide energy for a developing cyclone, as potential energy is transformed into kinetic energy. Further, if clouds form, condensation in the ascending air releases latent heat, which warms the air. The warmer air lowers the surface pressure, which strengthens the surface low even more. So, we now have a full-fledged

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FIGURE 8.33 (a) As the polar jet stream and its area of maximum winds (the jet streak, or core) swings over a developing mid-latitude cyclone, an area of divergence (D) draws warm surface air upward, and an area of convergence (C) allows cold air to sink. The jet stream removes air above the surface storm, which causes surface pressures to drop and the storm to intensify. (b) When the surface storm moves northeastward and occludes, it no longer has the upper-level support of diverging air, and the surface storm gradually dies out.

middle-latitude cyclonic storm with all of the necessary ingredients for its development. Eventually, the warm air curls around the north side of the low, and the storm system occludes (see Fig. 8.32c). Some storms may continue to deepen, but most do not as they move out from under the region of upper-level divergence. Additionally, at the surface the storm weakens as the supply of warm air is cut off and cold, dry air behind the cold front (called a dry slot) is drawn in toward the surface low. The Role of the Jet Stream Streams can play an additional part in the formation of surface mid-latitude cyclones. When the polar jet stream flows in a wavy westto-east pattern, deep troughs and ridges exist in the flow aloft. Notice in Fig. 8.33a that, in the trough, the area shaded orange represents a strong core of winds called the jet stream core, or jet streak. The curving of the jet stream coupled with the changing wind speeds around the jet streak produce regions of strong convergence and divergence along the flanks of the jet. The region of diverging air above the surface low (marked D in Fig. 8.33a) draws warm surface air upward to the jet stream, which quickly sweeps the air downstream. Since the air above the mid-latitude cyclone is being removed more quickly than converging surface winds can supply air to the storm’s center, the storm rapidly intensifies and surface winds increase. Above the high-pressure area, a region of converging air (marked C in Fig. 8.33a) feeds cold air downward into the anticyclone to replace the diverging surface air. Hence, we find the polar jet stream

removing air above the surface cyclonic storm and supplying air to the surface anticyclone. As the jet stream steers the cyclonic storm along (toward the northeast, in this case), the surface storm occludes, and cold air surrounds the surface low (see Fig. 8.33b). Since the surface low has moved out from under the pocket of diverging air aloft, the occluded storm gradually fills as the surface air flows into the system. Since the Northern Hemisphere’s polar jet stream is strongest and moves farther south in winter, we can see why mid-latitude cyclonic storms are better developed and move more quickly during the coldest months. During the summer when the polar jet stream shifts northward, developing mid-latitude cyclonic storm activity shifts northward as well, occurring principally in Canada over the province of Alberta and the Northwest Territories. In general, we now have a fairly good picture as to why some surface lows intensify into huge mid-latitude cyclones while others do not. For a surface cyclonic storm to intensify, there must be an upper-level counterpart— a trough of low pressure—that lies to the west of the surface low. As shortwaves disturb the flow aloft, they cause regions of differential temperature advection to appear, leading to an intensification of the upper-level trough. At the same time, the polar jet stream forms into waves and swings slightly south of the developing storm. When these conditions exist, zones of converging and diverging air, along with rising and sinking air, provide energy conversions for the storm’s growth. With this

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atmospheric situation, storms can form even where there are no pre-existing fronts. In regions where the upper-level flow is not disturbed by shortwaves or where no upper trough or jet stream exists, the necessary vertical and horizontal motions are insufficient to enhance cyclonic storm

development and we say that the surface storm does not have the necessary upper-air support. The horizontal and vertical motions, cloud patterns, and weather that typically occur with a developing open-wave cyclone are summarized in Fig. 8.34.

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FIGURE 8.34 Summary of clouds, weather, verti vertical motions, and upper-air support associated with a developing mid-latitude cyclone. Dark green area represents precipitation.

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SUMMARY In this chapter, we considered the different types of air masses and the various weather each brings to a particular region. Continental arctic air masses are responsible for the extremely cold (arctic) outbreaks of winter, whereas continental polar air masses are responsible for cold, dry weather in winter and cool, pleasant weather in summer. Maritime polar air, having traveled over an ocean for a considerable distance, brings to a region cool, moist weather. The hot, dry weather of summer is associated with continental tropical air masses, whereas warm, humid conditions are due to maritime tropical air masses. Where air masses with sharply contrasting properties meet, we find weather fronts. Along the leading edge of a cold front, where colder air replaces warmer air, showers are prevalent, especially if the warmer air is moist and conditionally unstable. Along a warm front, warmer air rides up and over colder surface air, producing widespread cloudiness and light-to-moderate precipitation that can cover thousands of square kilometers. When the rising air is conditionally unstable (such as it often is in summer), showers and thunderstorms may form ahead of the advancing warm front. Occluded fronts, which are often difficult to locate and define on a surface weather map, may have characteristics of both cold and warm fronts. We learned that fronts are actually part of the midlatitude cyclone. We examined where, why, and how these storms form and found that a mid-latitude cyclone goes through a series of stages from birth, to maturity, to death as an occluded storm. An important influence on the development of a mid-latitude cyclonic storm is the upperair flow, including the jet stream. We learned that when an upper-level low lies to the west of the surface low, and the polar jet stream bends and then dips south of the surface storm, an area of divergence above the surface low provides the necessary ingredients for the surface midlatitude cyclone to develop into a deep low-pressure area.

KEY TERMS The following terms are listed (with corresponding page numbers) in the order they appear in the text. Define each. Doing so will aid you in reviewing the material covered in this chapter. air mass, 210 source regions (for air masses), 210 continental polar (air mass), 212

continental arctic (air mass), 212 continental tropical, 218 lake-effect snows, 212 maritime polar (air mass), 000 Pacific air, 216 maritime tropical (air mass), 217 continental tropical (air mass), 219 front, 220 stationary front, 221 cold front, 222 “back door” cold fronts, 224 warm front, 224 dryline, 226 overrunning, 227 occluded front, 227 occlusion, 227 polar front theory, 229 wave cyclone, 229 frontal wave, 229 open wave, 230 cyclogenesis, 231 lee-side low, 231 nor’easter, 231 convergence, 232 divergence, 232 longwave (in westerly flow aloft), 235 shortwave (in westerly flow aloft), 235 cold advection, 235 warm advection, 235 jet streak, 237

QUESTIONS FOR REVIEW . (a) What is an air mass? (b) If an area is described as a “good air-mass source region,” what information can you give about it? . How does a continental arctic air mass differ from a continental polar air mass? . Why is continental polar air not welcome to the Central Plains in winter yet very welcome in summer? . What are lake-effect snows and how do they form? On which side of a lake do they typically occur? . Explain why the central United States is not a good air-mass source region. . List the temperature and moisture characteristics of each of the major air mass types. AIR MASSES, FRONTS, AND MIDDLE-LATITUDE CYCLONES

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. Which air mass only forms in summer over the southwestern United States? . Why are maritime polar air masses along the east coast of the United States usually colder than those along the nation’s west coast? Why are they also less prevalent? . Explain how the airflow aloft regulates the movement of air masses. . The boundaries between neighboring air masses tend to be more distinct during the winter than during the summer. Explain why. . What type of air mass would be responsible for the weather conditions listed below? (a) hot, muggy summer weather in the Midwest and the East; (b) refreshing, cool, dry breezes after a long summer hot spell on the Central Plains; (c) persistent cold, damp weather with drizzle along the East Coast; (d) drought with high temperatures over the Great Plains; (e) record-breaking low temperatures over a large portion of North America; (f) cool weather with showers over the Pacific Northwest; (g) daily afternoon thunderstorms along the Gulf Coast . Describe the typical characteristics of: (a) a warm front (b) a cold front (c) an occluded front . Sketch side views of a model showing a typical cold front, warm front, and cold-occluded front. Include in each diagram cloud types and patterns, areas of precipitation, surface winds, and relative temperature on each side of the front. . Describe the stages of a developing mid-latitude cyclonic storm using the polar front theory. . Why do mid-latitude cyclones tend to develop along the polar front? . List four regions in North America where midlatitude cyclones tend to develop. . Why is it important that for a surface low to develop or intensify, its upper-level counterpart must be to the left (or west) of the surface storm? . If upper-level diverging air above a surface area of low pressure exceeds converging air around the surface low, will the surface low weaken or intensify? Explain. . Describe some of the necessary ingredients (upper-air support) for a wave cyclone to develop into a huge mid-latitude cyclonic storm system. . Explain the role that upper-level diverging air plays in the development of a mid-latitude cyclone. 240

. How does the polar jet stream influence the formation of a mid-latitude cyclone? . Explain why, in the eastern half of the United States, a mid-latitude cyclonic storm often moves eastward or northeastward.

QUESTIONS FOR THOUGHT AND EXPLORATION A ATION . If Lake Erie freezes over in January, is it still possible to have lake-effect snow on its eastern shores in February? Explain your answer. . Explain how an autumn anticyclone can bring record low temperatures and continental polar air to the southeastern United States and, only a day or so later, bring record high temperatures and maritime tropical air to the same region. . During the winter, cold-front weather is typically more violent than warm-front weather. Why is this so? Explain why this is not necessarily true during the summer. . You are in upstate New York and observe the wind shifting from the east to the south. This wind shift is accompanied by a sudden rise in both air temperature and dew-point temperature. What type of front is passing? . Why does the same cold front typically produce more rain over Kentucky than over western Kansas? . Explain why the boundaries between neighboring air masses tend to be more distinct during the winter than during the summer? . Sketch a Southern Hemisphere mid-latitude cyclonic storm, complete with isobars and at least two types of fronts. Compare and contrast this Southern Hemisphere cyclone with its Northern Hemisphere counterpart. . Why are mid-latitude cyclones described as waves? . Explain how this can happen: At the same time a mid-latitude cyclonic storm over the eastern United States is moving northeastward, a large surface high-pressure area over the northern plains is moving southeastward. . Would a wave cyclone intensify or dissipate if the upper trough were located to the east of the surface low-pressure area? Explain your answer with the aid of a diagram.

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Under the Websites and Blogs section of Global Environment Watch: Meteorology, go to the site “Arctic Climatology and Meteorology PRIMER for Newcomers to the North” (National Snow and Ice Data Center). Within the Basics section of this site, consult the “Arctic Climate” page. What are the factors that can influence local weather within a particular arctic air mass? Consult the “Optical and Acoustic Phenomena” page. Why is it that conversations can sometimes be heard more than a mile away within a continental arctic air mass?

ONLINE RESOURCES Visit www.cengagebrain.com to view additional resources, including video exercises, practice quizzes, an interactive eBook, and more.

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CHAPTER

9

Weather Forecasting Contents Weather Observations

S

ometimes there is no job security in weather forecasting. In fact, at least one weather forecaster actually lost his job for

not altering his prediction. On April pril 15, 2001, a function honorhonor

Acquisition of Weather Information

ing a well-known radio talk show host was scheduled for outout

Weather Forecasting Tools

that a local forecaster at the radio station that sponsored the

Weather Forecasting Methods

such a forecast might discourage people from attending the

Time Range of Forecasts

forecast and predict a greater possibility of sunshine. The fore-

Accuracy and Skill in Weather Forecasting

doors at the Madera, California, alifornia, fairgrounds. The story goes event had called for a “chance of rain” on April 15. Upset that function, the station manager told the forecaster to alter his caster refused and was promptly fired. Apparently, retribution reigned supreme—it poured on the event.

Weather Forecasting Using Surface Charts Using Forecasting Tools to Predict the Weather

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W

eather forecasts are issued to save lives, to save property and crops, and to tell us what to expect in our atmospheric environment. In addition, knowing what the weather will be like in the future is vital to many human activities. For example, a summer forecast of extended heavy rain and cool weather would have construction supervisors planning work under protective cover, department stores advertising umbrellas instead of bathing suits, and ice cream vendors vacationing as their business declines. The forecast would alert farmers to harvest their crops before their fields became too soggy to support the heavy machinery needed for the job. And the commuter knows that prolonged rain could mean clogged gutters, flooded highways, stalled traffic, blocked railway lines, and late dinners. Put yourself in the shoes of a weather forecaster: It is your responsibility to predict the weather accurately so that thousands (possibly millions) of people in your area will know whether to carry an umbrella, wear an overcoat, or prepare for a winter storm. Since weather forecasting is not an exact science, your predictions will occasionally be incorrect. If your erroneous forecast misleads many people, you may become the target of jokes, insults, and even anger. There are even people who expect you to be able to predict the unpredictable. For example, on Monday you may be asked whether two Mondays from now will be a nice day for a picnic. And, of course, what about next winter? Will it be bitterly cold? Unfortunately, accurate answers to such questions are beyond meteorology’s present technical capabilities. Will forecasters ever be able to answer such questions confidently? If so, what steps are being taken to improve the forecasting art? How are forecasts made, and why do they sometimes go awry? These are just a few of the questions we will address in this chapter.

Weather Observations Weather forecasting basically entails predicting how the present state of the atmosphere will change. Consequently, if we wish to make a weather forecast, we must know the present weather conditions over a large area. A network of observing stations located across the world provides the forecaster with this information. Forecasters have access to many maps and charts showing the present conditions at various atmospheric heights, as well as vertical profiles (called soundings) of temperature, dew point, and winds. Also available are visible and infrared satellite images, as well as Doppler radar information that can detect and monitor the severity of precipitation and thunderstorms. All of these sources are used by forecasters to monitor current weather and anticipate future conditions. Many of these observations are also brought into computer-based 244

atmospheric models that project the weather forward, as we will see later in this chapter. More than 10,000 land-based stations and hundreds of ships and buoys provide surface weather information at least four times a day. Most airports observe conditions hourly, and hundreds of automated stations send reports even more often. To sample the atmosphere above ground level, radiosondes are launched at more than 800 locations around the world, including almost 100 U.S sites. Special launches may occur in association with research projects or to help provide more data when a major weather threat is looming, such as a high-impact winter storm or an outbreak of tornadic thunderstorms. Data on upper-air conditions may also be gathered and provided by some aircraft as they travel their usual routes. In addition, many types of observations collected by satellites are available to forecasters, providing a clearer representation of the atmosphere (see Fig. 9.1). A network of more than 100 Doppler radar units covcov ers nearly all of the 48 contiguous United States. These radars provide round-the-clock information on the evolution of rain, snow, sleet, and hail. Because Doppler radar can track winds as well as precipitation, the network is also a valuable tool in providing warnings of destructive windstorms and tornadoes, as we will see in the following chapters. Some of the most advanced computer forecast models are now being designed to incorporate information from radars, which could help improve forecasts significantly.

Acquisition of Weather Information Collecting weather data is only the beginning of the process that leads to a forecast. Meteorologists at government weather services and private firms across the world rely on an accurate supply of weather data in order to make predictions for their areas. A United Nations agency— the World Meteorological Organization (WMO), which includes more than 175 nations—is responsible for the international exchange of weather data. The WMO certifies that the observation procedures do not vary among nations, an extremely important task since the observations must be comparable. Weather information from all over the world is transmitted electronically to government meteorological centers worldwide. This includes the National Centers for Environmental Prediction (NCEP), a branch of the National Weather Service (NWS) located near the University of Maryland in College Park, just outside Washington, D.C. Here, the massive job of analyzing the data, running models, preparing weather maps and charts, and predicting the weather on a global and national basis begins.* From NCEP, *By international agreement, data are plotted using symbols illustrated in Appendix C.

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to make their own forecasts assisted by data and products from NCEP, or to modify an NWS forecast. Other stations hire meteorologically untrained announcers who paraphrase the NWS forecast, or read them word for word. On the web and on smartphones, the public can access local forecasts from the NWS and from private firms, often presented with eye-catching graphics. (Before going on, you may wish to read Focus section 9.1, which describes how TV weather forecasters present weather visuals.)

Weather Forecasting Tools

NASA/MSFC Earth Science Office

In the course of a single workday, a typical forecaster may examine and compare dozens or even hundreds of individual weather maps. To help forecasters handle all the available charts and maps, the NWS employs the highspeed Advanced Weather Interactive Processing System (AWIPS). A second-generation version, called AWIPS II, was adopted by the NWS starting in 2013 (see Fig. 9.2). The AWIPS II system has data communications, storage, processing, and display capabilities (including graphical overlays) to better help the individual forecaster extract and assimilate information from the mass of available data. In addition, AWIPS is able to process information received from satellites and surface stations as well as from the Doppler radar system (the WSR-88D), which now includes dual-polarization technology. Much of the information from ASOS* and Doppler radar is processed by software according to predetermined formulas, or algorithms, before it goes to the forecaster. Certain criteria or combinations of measurements can alert the forecaster to an impending weather situation. A software component of AWIPS II called the Graphical Forecast Editor allows forecasters to look at the daily prediction of weather elements, such as temperature and

FIGURE 9.1 These three satellite images were each collected FIG by the GOES-East satellite at 11:15 a.m. on June 16, 2014. (a) Visible imagery, which detects solar radiation being reflected, is useful for detecting areas of snow cover as well as low-level cloud features, such as the cumulus developing in Nebraska that became thunderstorms later in the day. (b) Infrared imagery analyzes the temperature at the top of clouds, which is related to the amount of infrared energy being emitted; the bright colors over eastern South Dakota and Iowa indicate the high, cold tops of severe thunderstorms. (c) Water-vapor imagery detects the amount of energy absorbed at a particular wavelength that reveals moisture present in the middle and upper troposphere. Lighter colors indicate more moisture, while darker colors (such as in western Kansas and eastern Colorado) show where the upper troposphere is drier.

NOAA

observations and computer model output are transmitted to U.S. private forecasting firms and public agencies. Many of NCEP’s products are also posted on the web. Across the nation, dozens of NWS Weather Forecast Offices (WFOs) use the information to issue local and regional weather forecasts. Standard forecasts are prepared every 12 hours and updated as needed in between these intervals. The public gets weather forecasts through a variety of channels, including radio, television, computers, and smartphones. Many broadcasting stations hire private meteorological companies or professional meteorologists

*Additional information on ASOS is found in Chapter 3, on p. 73.

FIGURE 9.2 Two NWS forecasters testing the AWIPS II system, FIG adopted by the NWS in 2013. WEATHER FORECASTING

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FOCUS

ON AN OB OBSERVATION 9.1

TV WEATHERCASTERS—HOW DO THEY DO IT? As you watch the TV weathercaster, you typically see a person describing and pointpoint ing to specific weather information, such as satellite and radar images, and weather maps, as shown in Fig. 1. What you may not know is that in many instances the weathercaster is actually pointing to a blank board (usually green or blue) on which there is nothing ( Fig. 2).* This

should not wear green clothes because such clothing would not be picked up by the camera; what you would see on your home screen would be a head and hands moving about the weather graphics! How, then, does a TV weathercaster know where to point on the blank wall? Positioned on each side of the green wall are TV monitors (look carefully at Fig. 2) that weathercasters watch so that they know where to point.

KSHB-TV Kansas City

KSHB-TV Kansas City

*On some stations, forecasters point to weather information that appears on a very large TV screen.

process of electronically superimposing weather information in the TV camera against a blank wall is called color-separation overlay, or chroma key key. The chroma key process works because the studio camera is constructed to pick up all colors except (in this case) green. The various maps, charts, satellite photos, and other graphics are electronically inserted from a computer into this green area of the color spectrum. The person in the TV studio

FIGURE 1 On your home television, this weathercaster appears FIG to be pointing to weather information directly behind him.

FIGURE 2 In the studio, however, the weathercaster is actually FIG standing in front of a blank green board.

dew point, in a gridded format with spacing as small as 2.5 km (1.6 mi). Presenting the data in this format allows the forecaster to predict the weather more precisely over a relatively small area. With so much information at the forecaster’s disposal, it is essential that the data be easily accessible and in a format that allows several weather variables to be viewed at one time. The meteogram is a chart that shows how one or more weather variables has changed at a station over a given period of time. As an example, the chart may represent how air temperature, dew point, and sea-level pressure have changed over the past five days, or it may illustrate how these same variables are projected to change over the next five days (see Fig. 9.3). Another aid in weather forecasting is the use of soundings. A sounding is a two-dimensional vertical profile of temperature, dew point, and winds espe (see Fig. 9.4).* The analysis of a sounding can be especially helpful when making a short-range forecast that

covers a relatively small area, such as the mesoscale. The forecaster examines the sounding from the immediate area (or closest proximity), as well as the soundings from sites upwind, to see how the atmosphere might be changing. Computer programs automatically calculate from the sounding a number of meteorological indexes that can aid the forecaster in determining the likelihood of smallerscale weather phenomena, such as thunderstorms, tornadoes, and hail. Soundings also provide information that can aid in the prediction of fog, air pollution alerts, and the downwind mixing of strong winds. Satellite information is also a valuable tool for the forecaster. Visible, enhanced infrared, and water vapor images provide a wealth of information, some of which comes from inaccessible regions, that can be examined to analyze fast-changing conditions.* Special instruments aboard satellites can detect a wide variety of important phenomena, including lightning, sea surface temperature, and smoke from forest fires. Many kinds of satellite observations are incorporated into forecast models.

*A sounding is obtained from a radiosonde or from satellite data. For additional information on the radiosonde, see the Focus section in Chapter 1 on p. 22.

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*Information provided by satellites is discussed in various sections of this book. For example, see Chapter 4, p. 106.

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FIGURE 9.3 Meteogram FIG illustrating predicted weather at the surface and aloft at St. Louis, Missouri, from 6 a.m., November 19, 2007, to noon on November 21, 2007. The forecast is derived from the Global Forecast System (GFS) model. (NOAA)

Weather Forecasting Methods Up to this point, we have examined some of the weather data and tools a forecaster might use in making a weather prediction. With all of this information available to the forecaster, including countless charts and maps, just how does a meteorologist make a weather forecast?

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THE COMPUTER AND WEATHER FORECASTING: NUMERICAL WEATHER PREDICTION Each day the many thousands of observations transmitted to NCEP are depicted in surface and upper-air charts. Meteorologists interpret the weather patterns and then correct any errors that may be present. The final chart is referred to as an analysis. FIGURE 9.4 A sounding of air temperature, dew point, and FIG winds near Oklahoma City, Oklahoma, during the evening of May 20, 2013, on the same day a violent tornado (pictured in Fig. 9.12) tore through the town of Moore, Oklahoma, which lies just south of Oklahoma City. WEATHER FORECASTING Copyright 2018 Cengage Learning. All Rights Reserved. May not be copied, scanned, or duplicated, in whole or in part. WCN 02-200-202

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Computers not only plot and analyze data, but they also predict the weather—a much more challenging task. Today’s supercomputers can analyze large quantities of data extremely quickly, carrying out trillions of calculations per second. Because the atmosphere is so complex, some of the most powerful supercomputers on Earth are devoted to weather and climate prediction. The routine daily forecasting of weather by the computer using mathematical equations is known as numerical weather prediction. The many weather variables are constantly changing, so meteorologists have devised atmospheric models that describe the present state of the atmosphere. These are not physical models that paint a picture of a developing storm; they are, rather, mathematical models consisting of many equations that describe how atmospheric temperature, pressure, winds, and moisture will change with time. The models do not fully represent the actual atmosphere, since the processes occurring around the world at every instant are too complex to represent completely. Instead, the models are very useful approximations, formulated to retain the most important aspects of the atmosphere’s behavior. How do these models actually work? The equations are translated into complex software, and surface and upper-air observations of temperature, pressure, moisture, winds, and air density are fed into the equations at regular intervals. The process of integrating these data into numerical models is called data assimilation. As more and more types of data are assimilated into the models, the quality of the model forecasts often improves. To determine how each of these key meteorological variables will change, each equation is solved for a small increment of future time—say, five minutes—for a large number of locations called grid points, each situated a given distance from the next.* In addition, each equation is solved for as many as 50 levels in the atmosphere. The results of these computations are then fed back into the original equations. The computer again solves the equations with the new “data,” thus predicting weather over the following five minutes. This procedure is done repeatedly until it reaches some desired time in the future. For example, one standard NWS model produces weather depictions every hour out to 18 hours. Another covers a longer period, but with larger steps in between; it produces snapshots of weather conditions every 3 hours out to 84 hours (3.5 days). And one model even forecasts the state of the atmosphere 384 hours (16 days) into the future. Once the calculations of future weather are completed, the computer then analyzes the data and draws the projected positions of pressure systems with their isobars or contour lines. The final forecast chart representing the atmosphere *Some models have a grid spacing smaller than 0.5 km, whereas the spacing in others exceeds 100 km. There are models that actually describe the atmosphere using a set of mathematical equations with wavelike characteristics rather than a set of discrete numbers associated with grid points.

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at a specified future time is called a prognostic chart, or, simply, a prog. Supercomputers can solve the equations of atmospheric motion far more quickly and efficiently than could possibly be done by hand. For example, just to produce a 24-hour forecast chart for the Northern Hemisphere requires many hundreds of millions of mathematical calculations. It would take a group of meteorologists working full time with hand calculators years to produce a single chart; by the time the forecast was available, the weather for that day would already be ancient history. Today, computer models are taken for granted in weather forecasting. They have become so important that our current level of skill in weather forecasting would be impossible without them. In fact, in some cases a computer model may do just as well as a human being in predicting high and low temperatures up to several days in advance during tranquil weather. However, forecasters must be careful not to take the prediction of any model as the gospel truth. If forecasters rely too much on models, without bringing their own knowledge and experience into the forecasting process, they may become victims of what has been termed “meteorological cancer.” During the most unusual and threatening weather situations, the best forecasts occur when the output of computer models is carefully adjusted as needed by an experienced, knowledgeable meteorologist. An ever-increasing variety of models (and, hence, progs) is now available to forecasters, each producing a slightly different interpretation of the weather for the same projected time and atmospheric level. Figure 9.5 shows four progs for different levels in the atmosphere 24 hours into the future. How the forecaster might use each prog in making a prediction is given in ▼ Table 9.1. The difference between progs can result from the way the models use the equations, or from the distance between grid points, called resolution. Some models predict certain features better than others: One model may work best in predicting the position of troughs on upper-level charts, whereas another may forecast the position of surface lows quite well. Although a model with a higher resolution can provide more detail, such enhanced detail alone does not necessarily mean the model is more accurate. A good forecaster knows the idiosyncrasies of each model and carefully scrutinizes all the progs. The forecaster then makes a prediction based on the guidance from the computer, his or her personalized practical interpretation of the weather situation, and any local geographic features that influence the weather within the specific forecast area. Currently, forecast models predict the weather reasonably well 4 to 6 days into the future, with decreasing accuracy for longer intervals. These models tend to do a better job of predicting temperature and jet-stream patterns than predicting precipitation. However, even with all

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of the modern advances in weather forecasting provided by ever more powerful computers, forecasts by the National Weather Service and by private firms are sometimes wrong. (An example of making a 24-hour forecast using different progs is given toward the end of this chapter on p. 264.)

NOAA/National Weather Service

WHY COMPUTER-BASED FORECASTS CAN GO AWRY AND STEPS TO IMPROVE THEM Why do forecasts sometimes go wrong? There are a number of reasons for this unfortunate situation. For one, computer models have inherent flaws that limit the accuracy of weather forecasts. For example, computer forecast models idealize the real atmosphere, meaning that each model makes certain assumptions about the atmosphere. These assumptions may be on target for some weather situations and be way off for others. Consequently, the computer may produce a prog that comes quite close to describing the actual state of the atmosphere on one day and not so close on another. A forecaster who bases a prediction on an “off day” computer prog may find a forecast of “rain and windy” turning out to be a day of “clear and colder.” Another forecasting problem arises because the majority of models are not global in their coverage, and errors can creep in along the model’s boundaries. For example, a model that predicts the weather for North America may not accurately treat weather systems that move in along its boundary from the western Pacific. Obviously, a global model would always be preferred. But a global model of similar sophistication with a high resolution requires an incredible number of computations. Even though many thousands of weather observations are taken worldwide each day, there are still regions where observations are sparse, particularly over the oceans and at higher latitudes. To help alleviate this problem, newer satellites are providing a more accurate profile of temperature and humidity for the computer models. Wind information now comes from a variety of sources, such as Doppler radar, commercial aircraft, buoys, and satellites that translate ocean surface roughness into surface wind speed. (See Chapter 6, p. 167.)

FIGURE 9.5 Computer-drawn 24-hour progs using the GFS FIG (Global Forecast Systems) model for 850 mb, 700 mb, 500 mb, and 300 mb. The progs were drawn on March 5, 2013, and became valid on March 6 at 7 a.m. Solid lines represent height contours. Orange shade shows the wind speed in knots. Green shade shows where the predicted relative humidity is 70 percent or greater.

Grid Spacing Earlier, we saw that the computer solves the equations that represent the atmosphere at many locations called grid points, each spaced from 100 km to as little as 0.5 km apart. On computer models with large spacing between grid points (say 60 km), larger weather systems, such as extensive mid-latitude cyclones and anticyclones, show up on computer progs, whereas much smaller systems, such as thunderstorms, do not. Because the model is too coarse to generate showers and thunderstorms directly, these features are instead parameterized, meaning that they are approximated for broad areas instead of being predicted for specific points. The computer models that forecast for a large area such as North America are, WEATHER FORECASTING

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▼ Table 9.1

The Use of Various V Charts as a Forecasting Tool

FORECAST CHART

APPROXIMATE ALTITUDE ABOVE SEA LEVEL

Surface map

ELEMENTS THAT MAY BE SPOTTED AND TRACKED

• • • •

Location and motion of frontal systems, centers of high and low pressure Areas of cloudiness, precipitation, high wind, and fog Cross-isobar winds that indicate strengthening or weakening low pressure Warm, moist air that can foster shower and thunderstorm development if conditional instability is present

850 mb

1500 m (4900 ft)

• • • •

High moisture values that can contribute to heavy precipitation Convergent winds associated with strengthening low pressure areas A low-level jet stream that can help intensify thunderstorm development Temperatures that determine whether precipitation will fall as snow, rain, sleet, or a mixture

700 mb

3000 m (9800 ft)

• • • •

Moisture to feed precipitation and mid-latitude storm systems Temperature advection that could strengthen or weaken fronts A dry, warm layer that can inhibit thunderstorm development Temperatures that help determine ice crystal and snowfall type

500 mb

5600 m (18,400 ft)

• General steering flow for mid-latitude storm systems, hurricanes, and tropical cyclones • Location and motion of ridges, troughs, and short waves that generate and strengthen surface features • Areas of cold advection that can help increase conditional instability and support thunderstorm development • Large areas of high or low heights that correspond to unusually warm or cold conditions at the surface, depending on region and time of year

300 mb

9180 m (30,100 ft)

• Location of core of jet stream • Jet streaks and areas of divergence within jet stream that may correspond to intensifying low pressure at the surface • Areas of high pressure and light, divergent wind in the tropics and subtropics that can support hurricane development

therefore, better at predicting the widespread precipitation associated with a large cyclonic storm than predicting where localized showers and thunderstorms will occur. In summer, when much of the precipitation falls as local showers, a computer prog may have indicated fair weather while outside it is pouring rain. To capture the smaller-scale weather features as well as the terrain of the region, the distance between grid points on some models is being reduced. For example, the forecast model known as High-Resolution Rapid Refresh (HRRR) has a grid spacing as low as 3 km. Instead of parameterizing showers and thunderstorms, a model like HRRR with its small grid spacing (high resolution) can actually incorporate radar information and simulate how showers and thunderstorms might evolve. The problem with high resolution models is that, as the horizontal spacing between grid points decreases, the number of computations increases. When the distance is halved, there are 8 times as many computations to perform, and the time (and computational expense) required to run the model goes up by a factor of 16. 250

Another forecasting problem is that many computer models cannot adequately interpret many of the factors that influence surface weather, such as the interactions of water, ice, surface friction, and local terrain on weather systems. Many large-scale models now take mountain regions and oceans into account. Some models (such as HRRR) take even smaller factors into account, features that large-scale computers miss due to their larger grid spacing. Given the effect of local terrain, as well as the impact of some of the other problems previously mentioned, computer models that forecast the weather over a vast area do an inadequate job of predicting local weather

DID YOU KNOW? When a weather forecast calls for “fair weather,” does the “fair” mean that the weather is better than “poor” but not up to being “good”? According to the National Weather Service, the subjective term “fair” implies a rather pleasant weather situation where there is no precipitation, temperatures are seasonable, winds are light, and visibility is good, with less than 40 percent of the sky covered by opaque clouds, such as stratus.

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conditions, such as surface temperatures, winds, and precipitation. Even with better observing techniques and highquality computer models, countless small, unpredictable atmospheric fluctuations, referred to as chaos, limit model accuracy. For example, tiny eddies, much smaller than the grid spacing on the computer model, go unaccounted for in the model. These small disturbances, as well as small errors (uncertainties) in the data, generally amplify with time as the computer tries to project the weather further and further into the future. After a number of days, these initial imperfections tend to dominate, and the forecast shows little or no accuracy in predicting the behavior of the real atmosphere. In essence, what happens is that the small uncertainty in the initial atmospheric conditions eventually leads to a huge uncertainty in the model’s forecast. There is, therefore, a limit to how far into the future we will ever be able to accurately forecast the weather at a specific place and time. However, it is still possible to make climatological projections that give the likelihood of particular types of weather far into the future.

Ensemble Forecasts Because of the atmosphere’s chaotic nature, meteorologists have turned to a technique called ensemble forecasting to improve short- and medium-range forecasts. The ensemble approach is based on running several forecast models—or different versions (simulations) of a single model—each beginning with slightly different weather information to reflect the errors inherent in the measurements. Suppose, for example, a forecast model predicts the state of the atmosphere 24 hours into the future. For the ensemble forecast, the entire model simulation is repeated, but only after the initial conditions are “tweaked” just a little. The “tweaking,” of course, represents the degree of uncertainty in the observations. Repeating this process several times creates an ensemble of forecasts for a range of small initial changes. Figure 9.6 shows an ensemble 500-mb forecast for March 1, 2013 (96 hours or 4 days into the future). The chart is constructed by running the model 17 different times, each time with a slightly different initial condition. Notice that the ensemble numbers are in strong agreement in some locations, such as the eastern Pacific, while there is large uncertainty in the other locations, such as

NOAA

FIGURE 9.6 Ensemble 500-mb FIG forecast chart for March 1, 2013, issued on the morning of February 25. The blue lines represent the 5460-meter contour; the red lines, the 5640-meter contour; and the green lines, a 16-year average, represents the average locations of the two contours. The yellow lines show where the main or “operational” member of the ensemble placed the contours.

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for total snowfall in Boston, Massachusetts. Notice that the forecasts vary tremendously, from only 3.5 inches to an eye-opening 51 inches. The solid black line indicates the ensemble mean, which is 24 inches. Since the actual snow total in Boston in this storm was 24.9 inches, a forecast that relied on the ensemble mean would have been an excellent one. The ensemble mean is not a perfect predictor in every case, but on average it tends to perform as well or better than any single model. In summary, imperfect numerical weather predictions may result from flaws in the computer models, from errors that creep in along the models’ boundaries, from the sparseness of data, and/or from inadequate representation of many pertinent processes, interactions, and inherently chaotic behavior that occurs within the atmosphere. However, by watching carefully for potential model errors and by using ensembles made up of a number of different model runs, a forecaster can increase the likelihood of making an accurate prediction. © Cengage Learning®.

OTHER FORECASTING TECHNIQUES Because the weather affects every aspect of our daily lives, attempts to predict it accurately have been made for centuries. One of the earliest attempts was undertaken by Theophrastus, a pupil of Aristotle, who in 300 bc compiled all sorts of weather indicators in his Book of Signs. A dominant influence in the field of weather forecasting for 2000 years, this work consists of ways to foretell the weather by examining natural signs, such as the color and shape of clouds, and the intensity at which a fly bites. Some of these signs have validity and are a part of our own weather folklore— ”A halo around the moon portends rain” is one of these. Today, we realize that the halo is caused by the bending of light as it passes through ice crystals and that ice crystaltype clouds (cirrostratus) are often the forerunners of an approaching cyclonic storm (see Fig. 9.8). If you keep

the midwestern United States. As the forecast goes further and further into the future, the lines usually look more and more like scrambled spaghetti, which is why an ensemble forecast chart is often referred to as a spaghetti plot. If, at the end of a specific time, the progs, or model runs, match each other fairly well, the forecast is considered robust. This situation allows the forecaster to issue a prediction with a high degree of confidence. If the progs disagree, the forecaster has less faith in the computer model prediction and will issue a forecast with more limited confidence. In essence, the less agreement among the progs, or model runs, the less predictable the weather. Consequently, it would not be wise to make outdoor plans for Saturday when on Monday the weekend forecast calls for “sunny and warm” with a low degree of confidence. Most everyday weather forecasts do not indicate the degree of forecaster confidence, but it is sometimes mentioned in TV weather reports or special NWS advisories, especially when threatening weather is a possibility. The example in Figure 9.7 shows how widely the members of an ensemble can diverge for a critical weather situation, as with the blizzard that struck New England in early February 2013. The ensemble includes three different models, each run six times with slightly different initial conditions, plus two other runs. The result is 20 predictions 252

© C. Donald Ahrens

FIGURE 9.7 An ensemble of forecasts of total snowfall in FIG Boston, Massachusetts, for February 8–9, 2013, issued early on February 6. Each trace represents the forecast from a different model run, all carried out at the same time but with slightly different initial conditions.

FIGURE 9.8 A halo around the sun (or moon) means that FIG rain is on the way, a weather forecast made by simply observing the sky.

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your eyes open and your senses keenly tuned to your environment, you should, with a little practice, be able to make fairly good short-range local weather forecasts by interpreting the messages written in the weather elements. Official weather forecasting activities were launched by the governments of many nations in the late 1800s and expanded as observational techniques improved in the 1900s. During the years before the advent of computer-based weather models, many forecasting methods were based largely on the experience of the forecaster. Many of these techniques were of value, but typically they gave a more general overview of what the weather should be like rather than a specific forecast. As late as the mid-1950s, all weather maps and charts, including those depicting current as well as future conditions, were plotted by hand and analyzed by individuals. Meteorologists predicted the weather using certain rules that related to the particular weather system in question. For short-range forecasts of six hours or less, surface weather systems were moved along at a steady rate. Upper-air charts were used to predict where surface storms would develop and where pressure systems aloft would intensify or weaken. The predicted positions of these systems were extrapolated into the future using linear graphical techniques and current maps. Experience played a major role in making the forecast. In many cases, these forecasts turned out to be amazingly accurate. However, with the advent of modern supercomputers, along with our present observing techniques, today’s forecasts are considerably better. Probably the easiest weather forecast to make is a persistence forecast, which is simply a prediction that future weather will be the same as present weather. If it is snowing today, a persistence forecast would call for snow through tomorrow. Such forecasts are most accurate for time periods of several hours and become less and less accurate after that. Persistence forecasts are also more useful at those times and places where the weather tends to change less dramatically. Another method of forecasting is the steady-state, or trend, forecast. The principle involved here is that surface weather systems tend to move in the same direction and at approximately the same speed as they have been moving, providing no evidence exists to indicate otherwise. Suppose, for example, that a cold front is moving eastward at an average speed of 30 mi/hr and it is 90 mi west of your home. Using the steady-state method, we might extrapolate and predict that the front should pass through your area in three hours. The analog method is yet another form of weather forecasting. Basically, this method relies on the fact that existing features on a weather chart (or a series of charts) may strongly resemble features that produced certain weather conditions sometime in the past. To the forecaster, the weather map “looks familiar,” and for this reason the analogue method is often referred to as pattern

recognition. A forecaster might look at a prog and say “I’ve seen this weather situation before, and this happened.” Previous weather events can then be utilized as a guide to the future. The problem here is that, even though weather situations may appear similar, they are never exactly the same. There are always sufficient differences in the variables to make applying this method a challenge. Even so, the analog method can be used to predict a number of weather elements, such as maximum temperature. Suppose that in New York City the average maximum temperature on a particular date for the past 30 years is 10°C (50°F). By statistically relating the maximum temperatures on this date to other weather elements—such as the wind, cloud cover, and humidity—a relationship between these variables and maximum temperature can be drawn. By comparing these relationships with current weather information, the forecaster can predict the maximum temperature for the day. Statistical forecasts of weather elements are based on the past performance of computer models. Known as Model Output Statistics, or MOS, these predictions, in effect, are statistically weighted analogue forecast corrections incorporated into the computer model output. For example, a forecast of tomorrow’s maximum temperature for a city might be derived from a statistical equation that uses a numerical model’s forecast of relative humidity, cloud cover, wind direction, and air temperature. When the Weather Service issues a forecast calling for rain, it is usually followed by a probability. For example: “The chance of rain is 60 percent.” Does this mean (a) that it will rain on 60 percent of the forecast area or (b) that there is a 60 percent chance that it will rain within the forecast area? Neither one! The expression means that there is a 60 percent chance that any random place in the forecast area, such as your home, will receive measurable rainfall. Looking at the forecast in another way, if the forecast for 10 days calls for a 60 percent chance of rain, it should rain where you live on 6 of those days. The verification of the forecast (whether it actually rained or not) is usually made at the Weather Service office, but remember that the computer models forecast for a given region, not for an individual location. When the National Weather Service issues a forecast calling for a “slight chance of rain,” what is the probability (percentage) that it will rain? ▼ Table 9.2 provides this information. An example of a probability forecast using climatological data is given in Fig. 9.9. The map shows the probability of a “White Christmas”—1 inch or more of snow on the ground—across the United States. The map is based on the average of 30 years of data and gives the likelihood of snow in terms of a probability. For instance, the chances are greater than 90 percent (9 Christmases out of 10) that portions of northern Minnesota, Michigan, and Maine will experience a White Christmas. In Chicago, it is close to 50 percent; and in Washington, D.C., about WEATHER FORECASTING

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FIGURE 9.9 Probability of a FIG “White Christmas”—one inch or more of snow on the ground— based on a 30-year average, from 1981 to 2010, inclusive. The probabilities do not include all of the mountainous areas in the western United States. (NOAA)

20 percent. Many places in the far west and south have probabilities less than 5 percent, but nowhere is the probability exactly 0, for there is always some chance (no matter how small) that a mantle of white will cover the ground on Christmas Day. For example, on December 24–25, 2004 (see Fig. 9.10), Corpus Christi, Texas, reported 4.4 inches of snowfall, and Brownsville, Texas, at the very southern part of the state, had 1.5 inches of snow, making it the first snowfall in Brownsville since 1899. For both cities, it was the heaviest snowfall on record for any date. Predicting the weather by weather types employs the analog method. In general, weather patterns are categorized

into similar groups or “types,” using such criteria as the position of the subtropical highs, the upper-level flow, and the prevailing storm track. As an example, when the Pacific high is weak or depressed southward and the flow aloft is zonal (west-to-east), surface storms tend to travel rapidly eastward across the Pacific Ocean and into the United States without developing into deep systems. But when the Pacific high is to the north of its normal position and the upper

Forecast wording used by the National Weather Service to describe the percentage probability of measurable precipitation (0.01 inch or greater) for steady precipitation and for convective, showery precipitation. FORECAST WORDING FOR STEADY PRECIPITATION

FORECAST WORDING FOR SHOWERY PRECIPITATION

10 to 20 percent

Slight chance of precipitation

Isolated showers

30 to 50 percent

Chance of precipitation

Scattered showers

60 to 70 percent

Precipitation likely

Numerous showers

≥ 80 percent

Precipitation,* rain, snow

Showers**

PERCENT PROBABILITY OF PRECIPITATION

*A forecast that calls for an 80 percent chance of rain in the afternoon might read like this: “. . . cloudy today with rain this afternoon. . . .” For an 80 percent chance of rain showers, the forecast might read “. . . cloudy today with rain showers this afternoon. . . .” **The 60 percent chance of rain does not apply to a situation that involves rain showers. In the case of showers, the percentage refers to the expected area over which the showers will fall.

254

NASA/GSFC

▼ Table 9.2

FIGURE 9.10 Satellite view of South Texas along the Gulf Coast FIG on Christmas Day, 2004. The white area covering Corpus Christi and Brownsville is snow. The probability of measurable snow on the ground in either of these two cities on Christmas Day is less than one percent (see Fig. 9.9). Yet, Corpus Christi received over 4 inches of snow and Brownsville about 1.5 inches. Just days later, the temperature climbed into the 80s (°F).

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Time Range of Forecasts Weather forecasts are normally grouped according to how far into the future the forecast extends. For example, a weather forecast for up to a few hours (usually not more than 6 hours) is called a nowcast (a very short range forecast). The techniques used in making such a forecast normally involve subjective interpretations of surface observations, satellite imagery, and Doppler radar information. Often the forecaster moves weather systems along by the steady state or trend method of forecasting, with human experience and pattern recognition coming into play. When severe or hazardous weather is likely or is occurring, the National Weather Service issues shortrange alerts in the form of weather watches, warnings, and advisories. A watch indicates that atmospheric conditions favor hazardous weather occurring over a particular region during a specified time period. These hazards may or may not actually develop, and their timing and location are uncertain, so a watch simply means to be on “watch” for that threat and to be prepared to act if necessary. When hazardous weather has developed, or is about to develop, two types of alerts may be issued. A warning indicates that the hazard now occurring or imminent is considered to be a threat to life and/or property (such as tornadoes, flash floods, severe thunderstorms, and winter storms). An advisory is similar to a warning, except that *The climate of a region represents the total accumulation of daily and seasonal weather events for a specific interval of time, most often 30 years.

© Cengage Learning®.

airflow is meridional (north-south), looping waves form in the flow with surface lows usually developing into huge storms. Since upper-level longwaves move slowly, usually remaining almost stationary for perhaps a few days to a week or more, the particular surface weather at different positions around the wave is likely to persist for some time. Figure 9.11 presents an example of weather conditions most likely to prevail with a winter meridional weather type. A forecast based on the climate* of a particular region is known as a climatological forecast. Anyone who has lived in Los Angeles for a while knows that July and August are practically rain-free. In fact, rainfall data for the summer months taken over many years reveal that rainfall amounts of more than a trace occur in Los Angeles about 1 day in every 90, or only about 1 percent of the time. Therefore, if we predict that it will not rain on a particular date next year during July or August in Los Angeles, our chances are nearly 99 percent that the forecast will be correct based on past records. Since it is unlikely that this pattern will significantly change in the near future, we can confidently make the same forecast for the year 2025.

FIGURE 9.11 Winter weather type showing upper airflow FIG (heavy arrow), surface position of Pacific high, and general weather conditions that should prevail.

it is used to indicate hazards that are usually less severe (such as light snow, light freezing rain, or dense fog). Note that sometimes even an advisory-level hazard can produce major problems, such as when light snow falls, melts, and quickly freezes on road surfaces. Weather forecasts that range from about 12 hours to a few days (generally up to 3 days or 72 hours) are called short-range forecasts. The forecaster may incorporate a variety of techniques in making a short-range forecast, such as satellite imagery, Doppler radar, surface weather maps, upper-air wind data, and pattern recognition. As the forecast period extends beyond about 12 hours, the forecaster tends to weight the forecast heavily on computerdrawn progs and statistical information, such as Model Output Statistics (MOS). A medium-range forecast is one that extends from about 3 to 8 days (192 hours) into the future. Mediumrange forecasts are almost entirely based on computerderived products, such as forecast progs and statistical forecasts (MOS). A forecast that extends beyond 3 days is often called an extended forecast.

DID YOU KNOW? Weather presentations have come a long way since the early days of television. From the 1950s through the 1970s, before computer graphics were available, fads and gimmicks were common. Both men and women delivered forecasts in attire that could range from bathing suits to clown costumes. Puppets and cartoon characters were often used to deliver the forecast or to react to it. At a number of stations, the weatherperson actually wrote temperatures and drew maps while standing behind a clear Plexiglas screen. Because of the workings of the TV camera, these weathercasters had to write backwards. WEATHER FORECASTING

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National Weather Service/NOAA

FIGURE 9.12 The 90-day outlook issued by NOAA in mid-May 2014 for (a) precipitation and (b) temperature for the period June FIG through August 2014. For precipitation (a), the darker the green color, the greater the probability of precipitation being above normal, whereas the deeper the brown color, the greater the probability of precipitation being below normal. For temperature (b), the darker the orange/red colors, the greater the probability of temperatures being above normal, whereas the darker the blue color, the greater the probability of temperatures being below normal. On both maps, the letter A stands for “above normal” and the letter B for “below normal.”

A forecast that extends beyond about 8 days (192 hours) is called a long-range forecast. Although computer progs are available for up to 16 days into the future, they are not accurate in predicting local temperature and precipitation, and at best only show the broad-scale weather features. The NOAA Climate Prediction Center summarizes these general trends in products called outlooks that cover 6- to 10-day and 8- to 14-day periods. These are not forecasts in the strict sense, but rather give an overview of how the expected precipitation and temperature patterns may compare with average conditions. Figure 9.12 gives a typical 90-day outlook. NOAA also issues seasonal outlooks every month. These cover three-month periods that overlap and extend out to roughly a year. Again, rather than depicting specific weather features, these outlooks show the odds that a given area might experience temperatures or precipitation that are above or below average. Initially, these outlooks were based mainly on the relationship between the projected average upper-air flow and the surface weather conditions that the type of flow would create. Today, long-range forecasts call on models that link the atmosphere with sea-surface temperature, such as the Climate Forecast System version 2 (CFSv2). Many of the outlooks also take into account persistence statistics that carry over the general weather pattern from immediately preceding months, seasons, and years. In Chapter 7, we saw how a vast warming (El Niño) or cooling (La Niña) of the equatorial tropical Pacific can 256

affect the weather in different regions of the world. These interactions, where a warmer or cooler tropical Pacific can influence rainfall in California, are called teleconnections.* These types of interactions between widely separated regions are identified through statistical correlations. For example, over regions of North America, where temperature and precipitation patterns tend to depart from normal during El Niño and La Niña events, the Climate Prediction Center can issue—months in advance—a seasonal outlook of an impending wetter or drier winter. Seasonal outlooks using teleconnections have become increasingly useful. Up to now we have looked at how weather forecasts are made and how forecasts can influence our daily lives. For a look at how weather forecasts can influence the marketplace, read Focus section 9.2. In most locations throughout North America, the weather is fair more often than rainy. Consequently, there is a forecasting bias toward fair weather, which means that if you make a forecast of “no rain” where you live for each day of the year, your forecast will be correct more than 50 percent of the time. But did you show any skill in making your correct forecast? What constitutes skill anyway? And how accurate are the forecasts issued by the National Weather Service? *Teleconnections include not only El Niño and La Niña but other indices, such as the Pacific Decadal Oscillation, the North Atlantic Oscillation, and the Arctic Oscillation. For more information on these indices, see Chapter 7, pp. 203-205.

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FOCUS

ON A SSPECIAL TOPIC 9.2

A good forecast can not only make or break plans for a picnic, but it can also spell the difference between profit and loss for an entire business. Weather predictions are a critical tool for many parts of the economy. Short-term forecasts can help an orange grower deal with the threat of a hard freeze or tip off a construction company to the risk of work delays. On a broader scale, the prices of stocks and commodities* can swing up or down based on the approach of a major storm, the forecasts of its behavior, and the damage left behind. For example, the price of frozen concentrated orange juice rose more than 40 percent in the month after Hurricane Charley struck many of Florida’s citrus groves in August 2004. For many companies, seasonal outlooks are even more important than dayto-day forecasts. A corporation that makes bread or pasta might pay close attention to long-term outlooks for temperature and precipitation across the wheat-growing areas of North America in order to anticipate potential drops in supply. For energy companies, even a small seasonal shift can play a huge role in the demand for summer cooling or winter heating. An unusually mild winter might provide a boost to airlines and trucking companies, which would suffer fewer delays from snow and ice, but it could also cut into the sales of coldweather clothing. Long-term outlooks for El Niño and La Niña can provide months of

valuable lead time on where winter temperatures in the United States are likely to run warmer or cooler than average. The most direct protection against the risk of weather-related financial downturns comes from insurance for hail, flooding, drought, and the like. Weather insurance typically covers only the most dire meteorological threats, much like a catastrophic health-care plan that covers heart attacks but not chronic illness. Several other tools can help a company use weather predictions to smooth out the potential ups and downs in profit linked to the atmosphere. Many commodities can be traded through contracts called futures (a type of derivative*). Futures contracts are agreements to buy or sell a commodity at a fixed price at some later date. For example, a bread-baking company might buy wheat futures based on a projected precipitation outlook. This forecast would help the company plan with more confidence, knowing that the cost it will pay for wheat won’t change even if a drought should strike and the price of wheat goes up dramatically. It is also possible to trade futures contracts based on indices of the weather itself. Rather than specifying the future cost of a commodity, a weather derivative contract puts a price tag on a particular weather outcome, such as a record-hot summer that boosts demand for air conditioning. Many

*Commodities represent a vast array of goods bought and sold in large quantities, from oranges to oil.

*Derivatives are contracts that derive their value from some other quantity, such as the price of a commodity.

Accuracy and Skill in Weather Forecasting In spite of the complexity and ever-changing nature of the atmosphere, forecasts made by the National Weather Service out to between 12 and 24 hours are usually quite accurate. Those made for between 2 and 5 days are fairly good. Beyond about 7 days, computer prog forecast accuracy falls off rapidly because of the chaotic nature of the atmosphere. Although weather predictions made for up to 3 days are by no means perfect, they are far better than simply flipping a coin. But just how accurate are they?

Pichugin Dmitry/Shutterstock

WEATHER PREDICTION AND THE MARKETPLACE

FIGURE 3 A baking company might FIG arrange to buy this wheat in advance at a guaranteed price if long-range weather forecasts point toward a poor crop.

such contracts are based on heating- or cooling-degree days, described in Chapter 3. Many investors and speculators try to make a profit on the twists and turns of the atmosphere, often by looking at weather predictions and by buying and selling futures or weather derivatives. Traders keep a close eye on both seasonal weather projections and short-term forecasts, such as the track of a hurricane that could knock out oil and gas production. For example, as tropical storm Rita gathered strength on September 20, 2005, and forecasts called for Rita to approach the Gulf Coast as a major hurricane, the price of oil rose by the largest single-day amount on record—$4.39, or about 7 percent.

One problem with determining forecast accuracy is deciding what constitutes a right or wrong forecast. Suppose tomorrow’s forecast calls for a minimum temperature of 35°F. If the official minimum turns out to be 37°F, is the forecast incorrect? Is it as incorrect as one 10 degrees off? By the same token, what about a forecast for snow over a large city when the snow line cuts the city in half, with the southern portion receiving heavy amounts and the northern portion none? Is the forecast right or wrong? Or what if the snow started just 3 hours earlier than predicted, falling during the morning rush hour instead of at mid-morning? At present, there is no clear-cut answer to the question of determining forecast accuracy, so WEATHER FORECASTING

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VINCENT DELIGNY/AFP/Getty Images

FIGURE 9.13 This violent torFIG nado ripped through the town of Moore, Oklahoma, on March 20, 2013. Even with a lead time of more than 15 minutes, the tornado took 24 lives, including 10 children, and injured 377 others. Even though predicting the precise location where a tornado will form is impossible at this time, it is possible to predict up to several days into the future where and when violent tornado outbreaks are likely.

meteorologists use a variety of mathematical techniques to measure the quality of their predictions. These might take into account how much the weather naturally varies at a given location and time of year, or how much data is actually available to determine whether a forecast was correct. How does forecast accuracy compare with forecast skill? Suppose you are forecasting the daily summertime weather in Los Angeles. It is not raining today and your forecast for tomorrow calls for “no rain.” Suppose that tomorrow it doesn’t rain. You made an accurate forecast, but did you show any skill in so doing? Earlier, we saw that the chance of measurable rain in Los Angeles on any summer day is very small indeed; chances are good that day after day it will not rain. For a forecast to show skill, it should be better than one based solely on the current weather ((persistence) or on the “normal” weather (climatology) for a given region. Therefore, during the summer in Los Angeles, a forecaster will have many accurate forecasts calling for “no measurable rain,” but will need skill to predict correctly on which summer days it will rain. If on a sunny July day in Los Angeles you happen to forecast rain for tomorrow and it rains, you have not only made an accurate forecast, but you have also shown skill in making your forecast because your forecast was better than both persistence and climatology. A meteorological forecast, then, shows skill when it is more accurate than a forecast utilizing only persistence or climatology. Persistence forecasts are usually difficult to improve upon for a period of time of several hours or less. Weather forecasts ranging from 12 hours to a few days generally show much more skill than persistence forecasts. However, as the range of the forecast period increases, the skill level drops quickly because of the effects of chaos discussed earlier in this chapter. The 6- to 14-day mean outlooks both show some skill (which has been increasing over the last several decades) in predicting temperature 258

and precipitation, although the accuracy of precipitation forecasts is less than that for temperature. Today, 7-day forecasts of major weather features are roughly as skillful as 3- to 4-day forecasts were in the 1990s. Beyond 15 days, specific forecasts are only slightly better than climatology. However, the level of skill in making forecasts of average monthly temperature and precipitation approximately doubled from 1995 to 2006. Forecasting large-scale weather events several days in advance (such as the disruptive Groundhog Day blizzard of 2011 across the Midwest) is far more accurate than forecasting the precise evolution and movement of small-scale, short-lived weather systems, such as tornadoes and severe thunderstorms. In fact, 3-day forecasts of the development and movement of a major low-pressure system show more skill today than 36-hour forecasts did in the 1990s. Even though determining the precise location where a tornado will form is presently beyond modern forecasting techniques, the regions where tornadic storms are likely to form can often be predicted several days in advance. With improved observing systems, such as Doppler radar and advanced satellite imagery, the lead time of watches and warnings for severe storms has increased. In fact, the lead time* for tornado warnings has more than doubled since the 1980s, with the average lead time now being close to 15 minutes, and the lead time for the most deadly and destructive tornadoes often more than 30 minutes (see Fig. 9.13). Although scientists may never be able to skillfully predict the weather beyond about 15 days using available observations, the prediction of climatic trends is more promising. Whereas individual weather systems vary *Lead time is the interval of time between the issue of the warning and actual observance of the event, in this case, the tornado.

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greatly and are difficult to forecast very far in advance, global-scale patterns of winds and pressure frequently show a high degree of persistence and predictable change over periods of a few weeks to a month or more. With the latest generation of high-speed supercomputers, general circulation models (GCMs) are doing a far better job of predicting large-scale atmospheric behavior than did the earlier models. (In Chapter 13, we will examine in more detail the climatic predictions based on numerical models.)

BRIEF REVIEW Up to this point, we have looked at the various methods of weather forecasting. Before going on, here is a review of some of the important ideas presented so far: ●

Available to the forecaster are a number of tools that can be used when making a forecast, including surface and upper-air maps, computer progs, meteograms, soundings, Doppler radar, and satellite information.

The forecasting of weather by high-speed computers is known as numerical weather prediction. Mathematical models that describe how atmospheric temperature, pressure, winds, and moisture will change with time are programmed into the computer. The computer then draws surface and upper-air charts, and produces a variety of forecast charts called progs.

Imperfections of the computer models—atmospheric chaos and small errors in the data—greatly limit the accuracy of weather forecasts for periods beyond a few days.

Ensemble forecasting is a technique based on running several forecast models (or different versions of a single model), each beginning with slightly different weather information to approximate errors in the measurements.

A persistence forecast is a prediction that future weather will be the same as the present weather, whereas a climatological forecast is based on the climatology of a particular region.

For a forecast to show skill, it must be better than a persistence forecast or a climatological forecast.

Weather forecasts that range from about 12 hours to about 3 days are called short-range forecasts. Those that extend from about 3 days to 8 days are called medium-range forecasts, and forecasts that extend beyond about 8 days are called long-range forecasts.

Seasonal outlooks provide an overview of how temperature and precipitation patterns may compare with normal conditions.

Weather Forecasting Using Surface Charts The best forecasts incorporate information about multiple layers of the atmosphere using numerical modeling, as we saw earlier in this chapter. However, even when computer models are skilled at advancing large-scale weather features forward in time, a capable forecaster also needs

a strong sense of how surface features typically evolve, which can help him or her when the progs disagree. Suppose that we wish to make a short-range weather forecast and the only information available is a surface weather map. Can we make a forecast from such a chart? Most definitely. And our chances of that forecast being correct improve markedly if we have maps available from several days back. We can use these past maps to locate the previous position of surface features and predict their movement. A simplified surface weather map is shown in Fig. 9.14. The map portrays early winter weather condicondi tions on Tuesday morning at 6 a.m. (CST). A single isobar is drawn around the pressure centers to show their positions without cluttering the map. Note that an open wave cyclone is developing over the Central Plains with showers forming along a cold front and light rain, snow, and sleet ahead of a warm front. The dashed lines on the map represent the position of the weather systems 6 hours ago. Our first question is: How will these systems move? DETERMINING THE MOVEMENT OF WEATHER SYSTEMS There are several rules of thumb we can use in forecasting the movement of surface pressure systems and fronts: . For short time intervals, mid-latitude cyclonic storms and fronts tend to move in the same direction and at approximately the same speed as they did during the previous 6 hours (providing, of course, there is no evidence to indicate otherwise). . Low-pressure areas tend to move in a direction that parallels the isobars in the warm air (the warm sector) ahead of the cold front. . Lows tend to move toward the region of the greatest drop in surface pressure, whereas highs tend to move toward the region of the greatest rise in surface pressure. . Surface pressure systems tend to move in the same direction as the wind at 5500 m (18,000 ft)—the 500-mb level. The speed at which surface systems move is about half the speed of the winds at this level. When the surface map (Fig. 9.14) is examined carefully and when rules of thumb 1 and 2 are applied, it appears that—based on present trends—the low-pressure area over the Central Plains should move northeast. When we observe the 500-mb upper-air chart ( Fig. 9.15), it too suggests that the surface low should move northeast at a speed of about 25 knots. A FORECAST FOR SIX CITIES We are now in a position to make a weather forecast for six cities. To do so, we will project the surface pressure systems, fronts, and current weather into the future by assuming steady-state WEATHER FORECASTING

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FIGURE 9.14 Surface weather map for 6 a.m. (CST). Tuesday. Dashed lines indicate positions of weather features six hours ago. Areas FIG shaded green are receiving rain, while areas shaded white are receiving snow, and those shaded pink, freezing rain or sleet.

FIGURE 9.15 A 500-mb chart for 6 a.m. (CST) Tuesday, showing wind flow. The light orange L represents the position of the surFIG face low. The winds aloft tend to steer surface pressure systems along and, therefore, indicate that the surface low should move northeastward at about half the speed of the winds at this level, or 25 knots. Solid lines are contours in meters above sea level.

conditions. Fig. 9.16 gives the 12- and 24-hour projected positions of these features. A word of caution before we make our forecasts. We are assuming that the pressure systems and fronts are moving at a constant rate, which may or may not occur. 260

Low-pressure areas, for example, tend to accelerate until they occlude, after which their rate of movement slows. Furthermore, the direction of moving systems can change due to “blocking” highs and lows that exist in their path or because of shifting upper-level wind patterns. We

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FIGURE 9.16 Projected 12- and 24-hour movement of fronts, pressure systems, and precipitation from 6 a.m. (CST) FIG Tuesday until 6 a.m. (CST) Wednesday. (The dashed lines represent frontal positions 6 hours ago.)

will assume a constant rate of movement and forecast accordingly, always keeping in mind that the longer our forecasts extend into the future, the more susceptible they are to error. If we move the low- and high-pressure areas eastward, as illustrated in Fig. 9.16, we can make a basic weather forecast for various cities. For example, the cold front moving into north Texas on Tuesday morning is projected to pass Dallas by that evening, so a forecast for the Dallas area would be “warm with showers, then turning colder.” But we can do much better than this. Knowing the weather conditions that accompany advancing pressure areas and fronts, we can make more detailed weather forecasts that will take into account changes in temperature, pressure, humidity, cloud cover, precipitation, and winds. Our forecast will include the 24-hour period from Tuesday morning to Wednesday morning for the cities of Augusta, Georgia; Washington, D.C.; Chicago, Illinois; Memphis, Tennessee; Dallas, Texas; and Denver, Colorado. We will begin with Augusta. Weather Forecast for Augusta, Georgia On Tuesday morning, cold, dry polar air associated with a highpressure area brought freezing temperatures and fair weather to the Augusta area (see Fig. 9.14). Clear skies, light winds, and low humidities allowed rapid nighttime cooling so that, by morning, temperatures were in the low 30s (°F). Now look closely at Fig. 9.16 and observe that the high-pressure area is moving slowly eastward, away from Augusta. Southerly winds on the western side of this system will bring warmer and more humid air to the

region. Therefore, afternoon temperatures will be warmer than those of the day before. As the warm front approaches from the west, clouds will increase, appearing first as cirrus, then thickening and lowering into the normal sequence of warm-front clouds. Barometric pressure should fall. Clouds and high humidity should keep minimum temperatures well above freezing on Tuesday night. Note in Fig. 9.16 that the projected area of precipitation (greenshaded region) does not quite reach Augusta. With all of this in mind, our forecast might read something like this: Clear and cold this morning with moderating temperatures by afternoon. Increasing high clouds with skies becoming overcast by evening. Cloudy and not nearly as cold tonight and tomorrow morning. Winds will be light and out of the south or southeast. Barometric pressure will fall slowly.

Wednesday morning we discover that the weather in Augusta is foggy with temperatures in the upper 40s (°F). But fog was not in the forecast. What went wrong? We

DID YOU KNOW? Groundhog Day (February 2) is the day that is supposed to represent the midpoint of winter—halfway between the winter solstice and the vernal equinox. Years ago, in an attempt to forecast what the remaining half of winter would be like, people placed the burden of weather prognostication on various animals, such as the groundhog, which is actually a woodchuck. Folklore says that if the groundhog emerges from his burrow and sees (or casts) his shadow on the ground and then returns to his burrow, there will be six more weeks of winter weather. One can only wonder whether it is really the groundhog’s shadow that drives him back into his burrow or the people standing around gawking at him. WEATHER FORECASTING

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forgot to consider that the ground was still cold from a recent cold snap. The warm, moist air moving over the cold surface was chilled below its dew point, resulting in fog. Above the fog were the low clouds we predicted. The minimum temperatures remained higher than anticipated because of the release of latent heat during fog formation and the absorption of infrared energy by the fog droplets. Not bad for a start. Now we will forecast the weather for Washington, D.C. Rain or Snow for Washington, D.C.? Look at Fig. 9.16 and observe that the low-pressure area over the Central Plains is slowly approaching Washington, D.C., from the west. Hence, the clear weather, light southwesterly winds, and low temperatures on Tuesday morning (Fig. 9.14) will gradually give way to increasing cloudiness, winds becoming southeasterly, and slightly higher temperatures. By Wednesday morning, the projected band of precipitation will be over the city. Will it be in the form of rain or snow? Without a sounding (a vertical profile of temperature), this question is difficult to answer. We can see in Fig. 9.16, however, that on Tuesday morning cities south of Washington, D.C.’s latitude are receiving snow. So a reasonable forecast would call for snow, possibly changing to rain as warm air moves in aloft in advance of the approaching fronts. A 24-hour forecast for Washington, D.C., might sound like this: Increasing clouds today and continued cold. Snow beginning by early Wednesday morning, possibly changing to rain. Winds will be out of the southeast. Atmospheric pressure will fall.

Wednesday morning a friend in Washington, D.C., calls to tell us that the sleet began to fall but has since changed to rain. Sleet? Another fractured forecast! Well, almost. What we forgot to account for this time was the intensification of the storm. As the low-pressure area moved eastward, it deepened; central pressure lowered, pressure gradients tightened, and southeasterly winds blew stronger than anticipated. As air moved inland off the warmer Atlantic, it rode up and over the colder surface air. Snow falling into this warm layer at least partially melted; it then refroze as it entered the colder air near ground level. The influx of warmer air from the ocean slowly raised the surface temperatures, and the sleet soon became rain. Although we did not see this possibility when we made our forecast, a forecaster more familiar with local surroundings would have. Let’s move on to Chicago. Big Snowstorm for Chicago From Figs. 9.14 and 9.16, it appears that Chicago is in for a major snowstorm. Overrunning of warm air has produced a wide area of snow which, from all indications, is heading directly for the Chicago area. Since cold air north of the low’s center will be over Chicago, precipitation reaching the ground should be frozen. On Tuesday morning (Fig. 9.16) the 262

leading edge of precipitation is less than 6 hours away from Chicago. Based on the projected path of the low-pressure area (Fig. 9.16) light snow should begin to fall around noon on Tuesday. By evening, as the storm intensifies, snowfall should become heavy. It should taper off and finally end around midnight as the center of the low moves on east. If it snows for a total of 12 hours—6 hours as light snow (around 1 inch every 3 hours) and 6 hours as heavy snow (around 1 inch per hour)—then the total expected accumulation will be between 6 and 10 inches. As the low moves eastward, passing south of Chicago, winds on Tuesday will gradually shift from southeasterly to easterly, then northeasterly by evening. Since the storm system is intensifying, it should produce strong winds that will swirl the snow into huge drifts, which may bring traffic to a crawl. The winds will continue to shift as they become northerly and finally northwesterly by Wednesday morning. By then the storm center will probably be far enough east so that skies should begin to clear. Cold air moving in from the northwest behind the storm will cause temperatures to drop further. Barometer readings during the storm will fall as the low’s center approaches and reach a low value sometime Tuesday night, after which they will begin to rise. A weather forecast for Chicago might be: Cloudy and cold with light snow beginning by noon, becoming heavy by evening and ending by Wednesday morning. Total accumulations will range between 6 and 10 inches. Winds will be strong and gusty out of the east or northeast today, becoming northerly tonight and northwesterly by Wednesday morning. Barometric pressure will fall sharply today and rise tomorrow.

A text message Wednesday morning from a friend in Chicago reveals that our forecast was correct except that the total snow accumulation so far is 13 inches. We were off in our forecast because the storm system slowed as it became occluded. We did not consider this because we moved the system by the steady-state forecast method. At this time of year (early winter), Lake Michigan is not quite frozen over, and the added moisture picked up from the lake by the strong easterly and northeasterly winds also helped to enhance the snowfall. Again, a knowledge of the local surroundings would have helped make a more accurate forecast. The weather about 500 miles south of Chicago should be much different from this. Mixed Bag of Weather for Memphis Observe in Fig. 9.16 that, within 24 hours, both a warm and a cold front should move past Memphis, Tennessee. The light rain that began Tuesday morning should saturate the cool air, creating a blanket of low clouds and fog by midday. The warm front, as it moves through sometime Tuesday afternoon, should cause temperatures to rise slightly as winds shift to the south or southwest. At night, clear to partly

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cloudy skies should allow the ground and air above to cool, offsetting any tendency for a rapid rise in temperature. Falling pressures should level off in the warm air, then fall once again as the cold front approaches. According to the projection in Fig. 9.16, the cold front should arrive sometime before midnight on Tuesday, bringing with it gusty northwesterly winds, showers, the possibility of thunderstorms, rising pressures, and colder air. Taking all of this into account, our weather forecast for Memphis will be:

slowed and became stationary along a line stretching from the Gulf of Mexico westward through southern Texas and northern Mexico. (From the surface map alone, we had no way of knowing this would happen.) Along the stationary front a wave of low pressure formed. This wave caused warm, moist Gulf air to slide northward up and over the cold surface air. Clouds formed, minimum temperatures did not go as low as expected, and we are left with a fractured forecast. Let’s try Denver.

Cloudy and cool with light rain, low clouds, and fog early today, becoming partly cloudy and warmer by late this afternoon. Clouds increasing with possible showers and thunderstorms later tonight and turning colder. Winds southeasterly this morning, becoming southerly or southwesterly this evening and shifting to northwesterly tonight. Pressures falling this morning, leveling off this afternoon, then falling again, but rising after midnight.

Clear but Cold for Denver In Fig. 9.14, we can see that, based on our projections, the cold high-pressure area will be centered slightly to the south of Denver by Wednesday morning. Sinking air aloft associated with this high-pressure area should keep the sky relatively free of clouds. Weak pressure gradients will produce only weak winds and this, coupled with dry air, will allow for intense radiational cooling. Minimum temperatures will probably drop to well below 0°F. Our forecast should therefore read:

A friend who lives near Memphis e-mailed us on Wednesday to inform us that our forecast was correct except that the thunderstorms did not materialize and that Tuesday night dense fog formed in low-lying valleys, but by Wednesday morning it had dissipated. Apparently, in the warm air, winds were not strong enough to mix the cold, moist air that had settled in the valleys with the warm air above. It’s on to Dallas. Cold Wave for Dallas From Fig. 9.16, it appears that our weather forecast for Dallas should be straightforward, since a cold front is expected to pass the area around noon on Tuesday. Weather along the front (Fig. 9.14) is showery with a few thunderstorms developing; behind the front the air is clear but cold. By Wednesday morning it looks as if the cold front will be far to the east and south of Dallas and an area of high pressure will be centered over southern Colorado. North or northwesterly winds on the east side of the high will bring cold arctic air into Texas, dropping temperatures as much as 40°F within a 24-hour period. With minimum temperatures well below freezing, Dallas will be in the grip of a cold wave. Our weather forecast should therefore read something like this: Increasing cloudiness and mild this morning with the possibility of showers and thunderstorms this afternoon. Clearing and turning much colder tonight and tomorrow. Winds will be southwesterly today, becoming gusty north or northwesterly this afternoon and tonight. Pressures falling this morning, then rising later today.

How did our forecast turn out? A quick check of Dallas weather on your smartphone on Wednesday morning reveals that the weather there is cold but not as cold as expected, and the sky is overcast. Cloudy weather? How can this be? The cold front moved through on schedule Tuesday afternoon, bringing showers, gusty winds, and cold weather with it. Moving southward, the front gradually

Clear and cold through tomorrow. Northerly winds today becoming light and variable by tonight. Temperatures tomorrow morning will be below zero. Barometric pressure will continue to rise.

Almost reluctantly Wednesday morning, we look up the weather conditions at Denver and find that it is clear and very cold. A successful forecast at last! We find out, however, that the minimum temperature did not go below zero; in fact, 13°F was as cold as it got. A downslope wind coming off the mountains to the west of Denver kept the air mixed and the minimum temperature higher than expected. Again, a forecaster familiar with the local topography of the Denver area would have foreseen the conditions that lead to such downslope winds and would have taken this into account when making the forecast. A complete picture of the surface weather systems for 6 a.m. Wednesday morning is given in Fig. 9.17. By comparing this chart to Fig. 9.16, we can summarize why our forecasts did not turn out exactly as we had predicted. For one thing, the center of the low-pressure area over the Central Plains moved slower than expected. This slow movement allowed a southeasterly flow of mild Atlantic air to overrun cooler surface air ahead of the storm while, behind the low, cities remained in the snow area for a longer time. The weak wave that developed along the trailing cold front over South Texas brought cloudiness and precipitation to Texas and prevented the really cold air from penetrating deep into the south. Farther west, the high-pressure area originally over Montana moved more southerly than southeasterly, which set up a pressure gradient that brought westerly downslope winds to eastern Colorado. The subjective forecasting techniques demonstrated for these six cities are those one might use when making WEATHER FORECASTING

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FIGURE 9.17 Surface weather map for 6 a.m. CDT Wednesday. FIG

a short-range weather prediction with limited resources. The following section describes how a weather forecaster might predict the weather in a region where, to the west, surface weather features are greatly modified by a vast body of water, and only scanty surface and upper-air data are available. Here the forecaster must rely heavily on experience as well as more sophisticated tools, which include satellite data, upper-air charts, Doppler radar, and computer progs.

Using Forecasting Tools to Predict the Weather It is late afternoon, and outside the weather forecast office near San Francisco the meteorologist mulls over what is going on in the sky. Overhead is a thin covering of cirrostratus; to the west, draped over the foothills, is the everpresent stratus and fog. The air is cool and the winds are westerly. It is Sunday, March 25, and the forecaster’s task is to make a 24-hour weather forecast for the coastal area of central California. What will tomorrow’s weather be like? Will it be similar to today’s or will it change markedly? A slowly falling barometer of 1016 mb (30 in.) and the high clouds moving in from the west point to an approaching storm system. A persistence forecast might be good for the next several hours, but what about tomorrow morning or tomorrow 264

afternoon? One of the biggest challenges for modern forecasters is to zero in on the elements that are most important on any given day. The forecast funnel (illustrated in Fig. 9.18) outlines the steps used by forecasters to steer their attention from large scales to smaller scales and from short time frames to longer periods. The forecast funnel allows forecasters to produce the best possible predictions in a limited amount of time, by starting out with an examination of large-scale features (top of the funnel) and ending with local-scale forecasts (bottom of the funnel). Before looking at the large scale, however, let’s take a quick look at current surface conditions. The surface map for 4 p.m. (PST) Sunday, Fig. 9.19, shows there are no weather fronts approaching the West Coast. In fact, the nearest front is a stationary one that has stalled over the Rockies. There is, however, a region of low pressure centered about 1100 km (700 mi) west of San Francisco, which (according to previous maps) has been there for several days. With a central pressure of only about 1012 mb (29.88 in.), the system is fairly weak. Could this weak storm system be causing the increase in high cloudiness and the falling barometer? And will this pattern lead to rain tomorrow? A look at the 500-mb chart may help with these questions. Figure 9.20 shows HELP FROM THE 500-MB CHART the 500-mb analysis for 4 p.m. Sunday afternoon. While examining the chart, the meteorologist recognizes certain

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FIGURE 9.19 Surface weather map for 4 p.m. (PST) Sunday, FIG March 25.

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clues that will aid in making the forecast. For one thing, the 5640-m height contour is over northern California. The forecaster knows that when this contour line is situated here or farther south, the statistical probability of receiving measurable rainfall over central California increases greatly. West of San Francisco the flow is meridional with a warm, upper high situated just south of Alaska. To the south both east and west of the high are troughs. Because the shape of this flow around the high resembles the Greek letter omega (Ω), the high and its accompanying ridge is known as an omega high. The forecaster recognizes the omega high as a blocking high, one that tends to persist in the same geographic location for many days. This blocking pattern also tends to keep the troughs in their respective positions, which has been the case for several days now. But the chart indicates that the cold upper trough located west of San Francisco may be changing somewhat. Observe the spacing of the contour lines around this trough. Even with a limited number of actual wind observations, the close spacing of the contours to the west and northwest of the trough, and the more widely spaced contours to the east of the trough, hint that stronger winds

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FIGURE 9.20 The 500-mb chart for 4 p.m. (PST) Sunday, March 25. FIG Arrows indicate wind flow. Red arrows indicate warm advection and blue arrows, cold advection. Solid lines are height contours where 564 equals 5640 meters above sea level. Dashed lines are isotherms (lines of equal temperature) in °C. The heavy dashed purple line shows position of a shortwave trough.

FIGURE 9.18 The forecast funnel, applied to a region of the FIG northeast coast of the United States.

exist to the west of the trough. The forecaster knows from past experience that this usually means the trough will deepen. Also note that on the west side of the trough, cold air is flowing southward (blue arrows), indicating that cold advection is occurring here. The heavy dashed purple line on the west side of the trough represents the position of a shortwave trough, which is moving rapidly southward. The injection of cold air and the shortwave into the main trough should cause it to intensify. To the east of the main trough, warm moist air is moving northeastward (red arrows), indicating that warm advection is occurring. It is the lifting and condensing of this moist air that is producing the high clouds over San Francisco. All of these conditions—high wind speeds, cold air moving southward, and warm air WEATHER FORECASTING

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moving northward, and a shortwave moving into a longwave trough—manifest themselves as a deepening of the longwave trough. As the upper trough deepens, it should be capable of providing the necessary conditions favorable for the development of the surface low into a major midlatitude cyclonic storm. The forecaster will likely consult maps for other levels of the atmosphere as well, such as 700 and 850 mb (see p. 250). These maps can reveal, for example, where moisture and warm air is flowing into the developing low-pressure center at lower levels of the atmosphere. Such knowledge gives the forecaster information on how quickly and strongly the surface storm will develop. As we saw in Chapter 8, one of the main ingredients necessary for the development and intensification of a surface low is divergence of the airflow aloft. The forecaster knows that divergence aloft is associated with a decrease in surface pressure. This decrease, in turn, causes surface air to converge and rise, and its moisture to potentially condense into widespread cloudiness. But where will regions of divergence, convergence, and rising air be found on tomorrow’s map? And how will tomorrow’s map be different from today’s? This is where the computer and the forecaster work together to come up with a prediction.

*Explaining the differences among the three models is beyond the scope of this book. Each model treats the atmosphere in a slightly different way. Some models have closer grid points. Some models have better resolution in the lower part of the atmosphere, whereas others have better resolution in the higher regions of the troposphere. The idea in this forecasting example is not to illustrate the different models in use, but rather to show how a forecaster might use any numerical computer model as a forecasting tool.

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THE MODELS PROVIDE ASSISTANCE The computer progs predict the future positions of weather systems. Some of the progs also predict where shortwave troughs will be located. It is important to know where the shortwaves will be found, because to the east of them there is usually upper-level divergence, lower-level convergence, rising air, clouds, and precipitation. Hence, predicting

the position of a shortwave means predicting regions of inclement weather. Three 24-hour forecast models that predict the positions of the shortwaves, upper-level pressure systems, and the flow aloft at the 500-mb level for Monday, March pre 26, at 4 p.m. (PST) are shown in Fig. 9.21.* (Each prediction is made on Sunday afternoon.) Observe that there is good agreement among the models in that each model moves the upper trough slowly eastward and positions it off the coast. However, the actual positioning of the trough and the shortwaves (heavy dashed lines) differ only slightly for each model. For example, model A and model C move the upper trough eastward more quickly than does model B. After examining each prog carefully, the forecaster must decide which model most accurately describes the future state of the atmosphere. Over the years, the forecaster knows that model A has performed well in predicting the positions of upper troughs that develop off the coast. Likewise, model C—because it uses more closely spaced grid points and a greater number of data points, and thus has better resolution—has done an admirable job of forecasting the positions of both upper troughs and shortwaves. On the other hand, although model B has its own strengths, it tends to move

FIGURE 9.21 Three computer-drawn progs (model A, model B, and model C) that show the 24-hour projected 500-mb chart. Solid FIG lines are contours, where 564 represents a height of 564 decameters (5640 meters) above sea level. Dashed lines represent proje projected positions of shortwaves. (Predictions were made on Sunday, March 25, at 4 p.m., PST.)

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FIGURE 9.22 Surface weather map for 4 a.m. (PST) Monday, FIG March 26.

the shortwaves along too slowly. Consequently, the forecaster puts more confidence into model A and model C for this particular situation. Using experience and the progs, the meteorologist sets out to predict the weather for the next 24 hours. The 24-hour progs for both model A and model C show a shortwave (labeled 1) approaching the California coast. As the shortwave approaches the coastline, clouds will increase and thicken and the likelihood of rain will increase. Therefore, a forecast for the next 24 hours might read like this: Increasing cloudiness Sunday night with rain beginning Monday morning and continuing through Monday afternoon.

A VALID FORECAST By early Monday morning, the maps begin to show the changes that the computer progs predicted. The surface map for 4 a.m. (PST) Monday morning ( Fig. 9.22) shows that the surface low in the Pacific has moved eastward and developed into a broad trough west of California. (Compare its position with Fig. 9.19.) The surface low has deepened considerably, as indicated by its central pressure of 1004 mb (29.65 in.). The approach of the storm is evidenced in San Francisco by thick middle clouds, southerly winds, and a falling barometer, nearly 4 mb lower than 12 hours ago. All these signs suggest that rain is on the way. On the 500-mb chart for 4 a.m. Monday morning ( Fig. 9.23), we can see that the movement of cold air southward around the upper trough, along with the swift movement of the shortwave, has caused the upper trough to deepen. Note that the height contours are now displaced farther south and that the contour in the middle of the trough is lower than on the previous 500-mb map (Fig. 9.20). Compare Fig. 9.20 with the 24-hour progs in

FIGURE 9.23 The 500-mb analysis for 4 a.m. (PST) Monday, FIG March 26. Heavy dashed lines show position of shortwaves. Solid lines are height contours where 564 equals 5640 meters above sea level. (Compare with Fig. 9.21, the 24-hour progs for model A and model C.)

Fig. 9.21 and notice that both model A and model C are projecting the movement of the shortwave (number 1) quite well. The forecaster made a wise choice in showing confidence in these two computer models as they did a good job predicting the position of the upper-level low and shortwave. Since the shortwave is moving with the flow toward San Francisco, it should rain today. But at what time will the rain begin? The forecaster must now go deeper into the forecast funnel and examine local conditions in greater detail. Here is where satellite and radar information come in. SATELLITE AND UPPER-AIR ASSISTANCE The infrared satellite image taken at 6:45 a.m. Monday (see Fig. 9.24) shows that the middle clouds presently over California will soon give way to an organized band of cumuliform clouds in the shape of a comma. Such comma clouds tell the forecaster that the surface low-pressure area off the coast is developing into a mature mid-latitude cyclone. Observe that this comma-shaped cloud band is associated with the surface low shown in Fig. 9.22. A quick glance at the 300-mb chart for 4 a.m. Monday morning ( Fig. 9.25) shows strong jet stream winds over Northern California and off the coast, suggesting that southwesterly winds aloft will carry the large commashaped cloud and its weather directly into California. And an area of strong divergence aloft (pink and red color on the map) associated with a jet streak will aid in deepening the cyclonic storm. By examining the movement of the cloud mass on successive satellite images, the forecaster can predict its arrival time and, hence, when rainfall will begin. According to satellite images, the leading edge of the comma WEATHER FORECASTING

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NOAA NOAA

FIGURE 9.24 Infrared satellite image taken at 6:45 a.m. (PST) FIG Monday, March 26. The cloud in the shape of a comma indicates that the mid-latitude cyclonic storm is deepening. (The heavy dashed line shows the tail of the comma cloud.)

NOAA

cloud should be just offshore by Monday afternoon. Also, Doppler radar indicates that, just off the coast, light rain is now falling from the middle cloud layer. The position of the upper shortwave trough and the jet stream above the area of surface low pressure should provide enough

FIGURE 9.25 The 300-mb chart for 4 a.m. (PST) Monday, FIG March 26. Solid lines are height contours, where 900 equals 9000 meters above sea level. The darkest color on the map indicates the jet stream core, or jet streak.

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FIGURE 9.26 Doppler radar image showing precipitation FIG falling across northern and central California at 5 p.m. (PST) on Monday, March 26. Areas in blue and green indicate light to moderate precipitation; yellow indicates heavier precipitation; orange and red, heaviest precipitation.

support for the low to continue to intensify. Pressure gradients around the low will likely increase, creating strong, gusty winds from the south as the storm approaches. An amended forecast for San Francisco might read: Rain beginning this morning, becoming heavy by this afternoon. Strong and gusty southerly winds.

A DAY OF RAIN AND WIND The first raindrops falling from altostratus clouds dampen city streets nea