climate and climate change [john p. rafferty]

7/17/2019 Climate and Climate Change [John P. Rafferty]

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Published in 2011 by Britannica Educational Publishing(a trademark of Encyclopædia Britannica, Inc.)in association with Rosen Educational Services, LLC29 East 21st Street, New York, NY 10010.

Copyright © 2011 Encyclopædia Britannica, Inc. Britannica, Encyclopædia Britannica,and the Thistle logo are registered trademarks of Encyclopædia Britannica, Inc. Allrights reserved.

Rosen Educational Services materials copyright © 2011 Rosen Educational Services, LLC.All rights reserved.

Distributed exclusively by Rosen Educational Services.For a listing of additional Britannica Educational Publishing titles, call toll free (800) 237-9932.

First Edition

Britannica Educational PublishingMichael I. Levy: Executive Editor

J.E. Luebering: Senior Manager

Marilyn L. Barton: Senior Coordinator, Production ControlSteven Bosco: Director, Editorial TechnologiesLisa S. Braucher: Senior Producer and Data EditorYvette Charboneau: Senior Copy EditorKathy Nakamura: Manager, Media Acquisition

John P. Rafferty: Associate Editor, Earth and Life Sciences

Rosen Educational ServicesHope Lourie Killcoyne: Senior Editor and Project ManagerNelson Sá: Art DirectorCindy Reiman: Photography ManagerMatthew Cauli: Designer, Cover Design

Introduction by Nancy Finton

Library of Congress Cataloging-in-Publication Data

Climate and climate change / edited by John P. Rafferty. p. cm. -- (The living earth)“In association with Britannica Educational Publishing, Rosen Educational Services.”Includes bibliographical references and index.ISBN 978-1-61530-388-5 (eBook )1. Climatology. 2. Climatic changes. I. Rafferty, John P.QC981.C623 2011551.5—dc22

2010015842

On the cover: The United States Geological Survey reports that as a result of global warming conditions, melting ice in the Arctic Sea could result in a loss of two-thirds of the world's polar bear population by the middle of the 21st century. Ice floes, such as the onepictured here, are vital for polar bears as a place to rest and breed, and as a launching pointfor hunting. Decreasing summer sea ice forces bears to swim greater distances than normal,leaving them exhausted, underweight, and vulnerable to drowning. Shutterstock.com

On pages v, 1, 102, 135, 168, 227, 330, 332, 339: The distinctive funnel-shaped cloud of atornado. Charles Doswell III/Stone/Getty Images

On page x: Flooded train tracks in New Orleans, La., the result of Hurricane Gustav, a 2008tropical cyclone that cut a swath across Hispaniola, Jamaica, Cuba, and the U.S. Gulf Coast.Stephen Morton/Getty Images

On pages xiv and xv: The major climatic groups are based on patterns of averageprecipitation, average temperature, and the natural vegetation found on Earth. This mapdepicts the world distribution of climate types based on the classification originallyinvented by Wladimir Köppen in 1900. Copyright Encyclopaedia Britannica; rendering for thisedition by Rosen Educational Services



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Introduction x

Chapter 1: Climate 1Solar Radiation and Temperature 2 The Distribution of Radiant Energy

from the Sun 2 The Effects of the Atmosphere 3

Average Radiation Budgets 6 Surface-Energy Budgets 7 Climatology 9Temperature 9 The Global Variation of Mean

Temperature 10 Diurnal, Seasonal, and Extreme

Temperatures 11

Temperature Variation with Height 13

Circulation, Currents, and Ocean-Atmosphere Interaction 14

Short-Term Temperature Changes 17Atmospheric Humidity and Precipitation 18 Atmospheric Humidity 18 Precipitation 30Atmospheric Pressure and Wind 56 Atmospheric Pressure 57 Wind 59 Maritime Continent 82 Monsoons 82 Upper-Level Winds 86

Chapter 2: Climatic

Classification 102

Approaches to Climatic Classification 105

CONTENTS



Genetic Classifications 106 Empirical Classifications 108The World Distribution of Major

Climatic Types 114 Type A Climates 114 Type B Climates 118 Type C and D Climates 123 Type E Climates 130 Type H Climates 133

Chapter 3: Climate and Life 135

The Gaia Hypothesis 136The Evolution of Life and the Atmosphere 138The Role of the Biosphere in the Earth-Atmosphere System 140 The Biosphere and Earth’s Energy

Budget 140

The Cycling of Biogenic Atmospheric Gases 143 Bioclimatology 149 Biosphere Controls on the Structure

of the Atmosphere 150 Biosphere Controls on the Planetary

Boundary Layer 151 Biosphere Controls on Maximum

Temperatures by Evaporation and Transpiration 153

Biosphere Controls on Minimum Temperatures 154

Climate and Changes in the Albedo of the Surface 157

The Effect of Vegetation Patchiness on Mesoscale Climates 158

Biosphere Controls on SurfaceFriction and Localized Winds 159

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Biosphere Impacts on Precipitation Processes 160

Climate, Humans, and Human

Affairs 164

Chapter 4: Climate Change 168

The Earth System 169Evidence for Climate Change 174Causes of Climate Change 176 Solar Variability 176 Volcanic Activity 177 Tectonic Activity 179 Orbital (Milankovitch)

Variations 180 Greenhouse Gases 182 Feedback Within the Earth

System 182 Human Activities 184

Climate Change Within a Human Life Span 186 Seasonal Variation 187 Interannual Variation 189 Decadal Variation 192 El Niño 193Climate Change Since the Emergence of Civilization 197

Centennial-Scale Variation 198 Millennial and Multimillennial

Variation 200Climate Change Since the Emergence of Humans 204 Recent Glacial and Interglacial

Periods 205 Paleoclimatology 207 Glacial and Interglacial Cycles of the

Pleistocene 209

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The Last Great Cooling 212Climate Change Through Geologic Time 213

Cenozoic Climates 214 Phanerozoic Climates 217 The Climates of Early Earth 219Abrupt Climate Changes in Earth History 223

Chapter 5: Global Warming 227

Causes of Global Warming 230The Greenhouse Effect 230 Radiative Forcing 233 The Influences of Human Activity

on Climate 234 Carbon Sequestrian 249 Natural Influences on Climate 252 Feedback Mechanisms and Climate

Sensitivity 260Climate Research 265 Modern Observations 266 Prehistorical Climate Records 268 Theoretical Climate Models 269Potential Effects of Global Warming 274 Simulations of Future Climate

Change 276 Environmental Consequences of

Global Warming 282 Socioeconomic Consequences of

Global Warming 286Global Warming and Public Policy 290 Kyoto Protocol 291 The IPCC and the Scientific

Consensus 294

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The UN Framework Convention and the Kyoto Protocol 296

Future Climate-Change Policy 299 An Inconvenient Truth 302Public Awareness and Action 305 Green Architecture 305 Electric and Hybrid Vehicles 316 Biofuels 323Conclusion 328

Glossary 330

Bibliography 332

Index 339

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INTRODUCTIONINTRODUCTION



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7 Introduction 7

Since at least the 1970s, Earth’s changing climate hasbeen a popular news topic. Even though scientists have

been studying varying aspects of climate for centuries,much about climate remains murky to the general public,and understandably so. The forces that drive the world’s cli-mates are complex and dependent on factors as diverse as

geography, topography, fluctuations in solar radiation, and volcanism. Feedback loops are numerous: a region’s climatedetermines the types of vegetation found there; however,

vegetative cover in turn affects the temperature and rainfall

in an area. In addition, climatic conditions can fluctuate inperiods as short as a day and as long as millions of years.

Before tackling climate change, readers of this book will first receive a thorough grounding in the dynamicsof Earth’s climate, and how climate interacts with liv-ing things and other parts of the Earth system. Readers

will then have an opportunity to explore climate change

throughout Earth’s history before investigating the causes,effects, and current scientific and public policy responsesto the phenomenon of global warming.

Simply put, climate refers to the “average” weatheroccurring within a given geographic location over a longperiod of time. A location’s climate is made up of the stan-dard weather conditions occurring in different seasons,

the variability of weather at that location across differentperiods of time, and the frequency of special atmosphericphenomena, such as tropical cyclones (hurricanes) andtornadoes. More specifically, climate refers to all the ele-ments that cause changes in the atmosphere, such as solarradiation, humidity, cloud cover, and wind and ocean cur-rents. Such a discussion would not be complete without an

exploration into how Earth’s hydrosphere (the wateryregions of the planet), lithosphere (the crust and upper-most portion of the mantle), and biosphere (the regionpopulated by living things) interact with the atmosphere



xii

and with one another. Humans and their activities alsoaffect Earth’s climate.

The ultimate driver of climate is the Sun, Earth’s localstar that emits all types of radiation. Some of this radia-tion reaches Earth’s atmosphere and surface. The climateof a particular location on Earth’s surface is primarilydetermined by the amount of solar radiation it receives.Heating is maximized when incoming solar radiationstreams at a 90° angle toward Earth’s surface. The radia-tion angle, and thus the amount of radiation, varies with a

location’s latitude, the seasons, and the time of day.In general, temperature rises when the atmosphere

absorbs solar radiation emitted by Earth’s surface. Manyfactors affect how the energy from incoming solar radia-tion is distributed throughout a region. Cloud cover canscatter radiation before it is absorbed. The amount of

water vapour and other gases in the atmosphere influence

how much radiation is absorbed. Surface conditions alsoplay a role. For example, snow cover and ice cover reflect alarge portion of solar radiation, while surface waterabsorbs much more. As a result, energy stored in bodies of

water can subsequently heat the atmosphere above.At any given time of day, half of the planet is bathed

in sunlight, so the Earth system absorbs solar radiation

constantly. However, the temperature of the atmosphereremains relatively stable, because Earth also emits heat, orthermal radiation, back into space. The difference betweenabsorbed and emitted radiation at any given point onEarth’s surface is the location’s radiation budget. The radia-tion budget is the main factor in a location’s total energy budget. Other components of a location’s total energy bud-

get include the quantity of energy stored and the quantitytransferred in and out by wind and ocean currents.For more than 2,000 years, since the classical Greek

period, people have been looking for ways to classify Earth’s

7 Climate and Climate Change 7



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climates. Since then, more than 100 climate classificationshave been developed. All can be sorted into either geneticor empirical classification systems. Genetic systems classifyclimates by their contributing factors, such as energy bud-

gets, wind patterns, or geographical factors. Scientists find genetic schemes appealing because they account for causalfactors (such as the movements of air masses, solar radia-tion, the influence of topography, etc.). They are, however,more challenging to use, because they rely on a complexinterpretation of climate observations and often yield

results that conflict with the more commonly used empiri-cal classifications. Empirical schemes rely exclusively onobserved weather and environmental data and can be basedon one or more factors, such as temperature, precipitation,and humidity. Many popular empirical systems have classi-fied regions according to their dominant vegetation type.

One of the most well-known climate classifications is an

empirical system first published in 1900 by German mete-orologist and climatologist Wladimir Köppen. Köppendevised formulas that could describe newly mapped zonesof vegetation, using temperature as the primary definingcriterion. The Köppen system has been criticized for fail-ing to acknowledge several factors, such as sunshine, wind,periodic droughts or other extreme events, and environ-

mental change. Still, it remains a widely used means ofclassifying Earth’s climates, partly because newer systemsare more complex and difficult to work with.

As long as there has been life on Earth, living thingsand climate have influenced one another. Plants, ani-mals, and other life-forms exchange energy and matter

with the atmosphere. Before the evolution of life, the

atmosphere was believed to have been mainly made upof carbon dioxide and water vapour. Both of these gasesare known as greenhouse gases, because they are rela-tively transparent to incoming shortwave radiation, and

7 Introduction 7



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7 Introduction 7




xvi

more opaque to long-wave radiation (most of which isexpressed as heat) emitted from Earth. Four billion yearsago, at the start of the Archean Eon (the earlier of thetwo formal divisions of Precambrian time), the Sun pro-duced only 25 percent as much energy as it does today.However, scientists maintain that large concentrationsof greenhouse gases in the atmosphere slowed the emis-sion of long-wave radiation enough to make Earth’stemperature similar to that of today. By metabolizingcarbon dioxide, early forms of bacteria and other, more

primitive organisms gradually changed the composi-tion of Earth’s atmosphere. Over time, oxygen levelsincreased, and carbon dioxide levels fell. Today, carbondioxide makes up about 0.04 percent of all atmospheric

gases—a huge reduction. In addition to affecting Earth’sradiation budget, the biosphere influences wind, tran-spiration and evaporation, nighttime temperatures, and

other factors that affect climate.Compared to trillions of bacteria found in Earth’s bio-

sphere, humans have little influence on the atmospherethrough their own metabolic processes. However, peoplehave a greater influence on Earth’s atmosphere throughtheir economic and cultural activities. Further, peoplerelease carbon into the atmosphere when they burn fossil

fuels, such as oil, coal, natural gas, or biomass (trees or grasses). Increasing the carbon concentration in the atmo-sphere increases the opacity of Earth’s atmosphere toheat, and thus smaller amounts are released into space.

In addition, people can affect Earth’s climate by chang-ing Earth’s surface. The degree to which the Sun’s energyheats the atmosphere depends, in part, on the reflectiv-

ity—known as albedo—of Earth’s surface. Clearing aforest or grassland for agriculture or building construc-tion, for example—increases the location’s albedo andmore energy is reflected back into space than would



xvii

otherwise be retained. Even slight increases in the albedosignificantly decrease the amount of rainfall in an area: a0.1 percent increase in albedo has been linked to a 20 per-cent decline in thunderstorm-driven rainfall.

To fully understand climate, scientists must sort and weigh the vast number of interactions that occur withinthe Earth system. The atmosphere interacts with all ofEarth’s surface features (e.g., oceans, vegetation cover,ice masses, topography, etc.) as well as others that occurbeneath the surface (e.g., volcanism). At present, many

scientists working in disciplines as varied as climatology,ecology, glaciology, geology, and the social sciences coop-erate within a comprehensive discipline known as Earthsystem science. Earth system scientists also include paleo-ecologists, paleoclimatologists, and other scholars whostudy past periods of geologic time.

The study of climatic change involves tracking local and

regional climates over long time periods, sometimes overmillions of years. Different types of evidence are used toreconstruct climate histories. Widespread, reliable weatherdata exists for the past two centuries or so in the form ofmeteorological measurements and historical documents. Inrare cases, documents such as court records, ships’ logs, anddiaries provide weather reports that are several centuries old.

Paleoclimatologists have many other sources of evidence,however. Climate affects the growth and success of partic-ular plant and animal species in a region, the weathering ofminerals, the accumulation of ice, and much more. Plant andanimal fossils, mineral deposits, and ice cores are commonsources of information on past climatic conditions. This evi-dence is generally circumstantial, and Earth system scientists

search for multiple forms of evidence to verify their conclu-sions and work to develop better tools to analyze their data.Earth has experienced many episodes of climate

change throughout its long history. Some of these changes

7 Introduction 7



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have been gradual, whereas others have been relativelyrapid, warming or cooling large portions of the planet in

years or decades. Climate can be influenced by many fac-tors; however, those factors that change the amount ofsolar radiation striking Earth’s surface or change how thisincoming radiation is distributed across Earth’s surfaceare the most important. The amount of solar radiationthat reaches Earth can vary across short timescales (suchas with the frequency of solar storms) or longer timescales(such as with the steadily growing brightness). Large

amounts of gases and particulates catapulted into Earth’sstratosphere from violent volcanic eruptions can blockincoming solar radiation, producing a significant coolingeffect on the planet. The year 1816 was known as the “year

without summer,” following the eruption of Indonesia’sMount Tambora the year before. Over periods of thou-sands of years, changes to Earth’s axis and the shape of its

orbit around the Sun change how incoming solar radiationis distributed across Earth’s surface. In addition, the move-ments of Earth’s tectonic plates over millions of yearschange the orientation of continents, mountains, andoceans thus altering patterns of air and water circulation.

Another important factor influencing climate involveshow well Earth’s atmosphere retains the energy it receives

from the Sun. The “greenhouse effect” is the name of theprocess that slows the release of solar energy into space.More specifically, after Earth’s surface absorbs incomingsolar radiation, much of it is emitted, or re-radiated, backinto space; however, the emission process is hinderedby the presence of greenhouse gases. The atmosphere’sheat-trapping ability increases as the concentration of

greenhouse gases rises. Earth’s climate has fluctuatedthroughout time as levels of greenhouse gases—primar-ily carbon dioxide, methane, and water vapour—haveslowly risen or fallen. Vegetation, soil, and the oceans can




xix

store or release these gases, and volcanism can increase greenhouse gas levels by adding carbon dioxide to theatmosphere. Since the dawn of the Industrial Revolution,the widespread burning of fossil fuels by humans andtheir activities, along with deforestation, has increased

greenhouse gas levels. Climatologists agree that such ris-ing greenhouse gas concentrations have affected Earth’sradiation budget and that some of the effects of climaticchange are occurring as a result of these budgetary adjust-ments. Nevertheless, how Earth’s climate will continue to

react remains a topic of research, study, and debate.Global warming refers to any increase in the tem-

perature of near-surface air during the past two centuriesthat can be traced to human activities. According to theIntergovernmental Panel on Climate Change (IPCC),Earth’s average surface temperature rose about 0.6° C (1.1°F) in the 20th century. The IPCC contends that much of

this rise is attributed to human activities. The amountthat Earth’s temperature will actually rise in coming years,the severity of the consequences, and the best courses ofaction are being hotly debated around the world.

The future of human-influenced global warming, andthe human response to it, is not clear. International pactsintended to address the issue, such as the 1997 Kyoto

Protocol and the 2009 Copenhagen Accord, have beenthe subject of much contention. Actions to reduce global warming are often seen to conflict with a country’s eco-nomic interests. Various countries differ in their responses,

with the European Union leading the way in setting clearreduction goals and setting up the world’s first multilateralsystem for trading carbon dioxide emissions. Despite the

debates, a few ideas are reaching wide acceptance, chiefamong them that the amounts and types of energy peopleuse—now and in the future—will alter the composition ofthe atmosphere, and Earth’s climate.

7 Introduction 7



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Weather is what we pay attention to on a day-to-daybasis: temperature, humidity, precipitation (type,

frequency, and amount), atmospheric pressure, and wind(speed and direction). Climate, on the other hand, is thecondition of the atmosphere at a particular location overan extended period of time (from one month to many mil-lions of years, but generally 30 years). Climate therefore is

the long-term summation of the atmospheric elements(and their variations) that, over short time periods, consti-tute weather.

From the ancient Greek origins of the word ( klíma , “aninclination or slope”—e.g., of the Sun’s rays; a latitudezone of the Earth; a clime) and from its earliest usage inEnglish, climate has been understood to mean the atmo-

spheric conditions that prevail in a given region or zone.In the older form, clime , it was sometimes taken to includeall aspects of the environment, including the natural veg-etation. The best modern definitions of climate regard itas constituting the total experience of weather and atmo-spheric behaviour over a number of years in a given region.Climate is not just the “average weather” (an obsolete, and

always inadequate, definition). It should include not onlythe average values of the climatic elements that prevail atdifferent times but also their extreme ranges and variabil-ity and the frequency of various occurrences. Just as one

year differs from another, decades and centuries are foundto differ from one another by a smaller, but sometimes sig-nificant, amount. Climate is therefore time-dependent,

and climatic values or indexes should not be quoted with-out specifying what years they refer to.This book treats the factors that produce weather and

climate and the complex processes that cause variations in

CLIMATE




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both. Other major points of coverage include global cli-matic types and microclimates. The book also considersthe impact of climate on human life—as well as the effectsof human activities on climate.

SOLAR RADIATIONAND TEMPERATURE

Air temperatures have their origin in the absorption ofradiant energy from the Sun. They are subject to many

influences, including those of the atmosphere, ocean, andland, and are modified by them. As variation of solar radia-tion is the single most important factor affecting climate,it is considered here first.

The Distribution of Radiant Energyfrom the Sun

Nuclear fusion deep within the Sun releases a tremendousamount of energy that is slowly transferred to the solarsurface, from which it is radiated into space. The planetsintercept minute fractions of this energy, the amountdepending on their size and distance from the Sun. A1-square-metre (11-square-foot) area perpendicular (90°)

to the rays of the Sun at the top of Earth’s atmosphere, forexample, receives about 1,365 watts of solar power. (Thisamount is comparable to the power consumption of a typ-ical electric heater.) Because of the slight ellipticity ofEarth’s orbit around the Sun, the amount of solar energyintercepted by Earth steadily rises and falls by ±3.4 percentthroughout the year, peaking on January 3, when Earth is

closest to the Sun. Although about 31 percent of thisenergy is not used as it is scattered back to space, theremaining amount is sufficient to power the movement of



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7 Climate 7

atmospheric winds and oceanic currents and to sustainnearly all biospheric activity.

Most surfaces are not perpendicular to the Sun, andthe energy they receive depends on their solar elevationangle. (The maximum solar elevation is 90° for the over-head Sun.) This angle changes systematically with latitude,the time of year, and the time of day. The noontime eleva-tion angle reaches a maximum at all latitudes north of theTropic of Cancer (23.5° N) around June 22 and a minimumaround December 22. South of the Tropic of Capricorn

(23.5° S), the opposite holds true, and between the twotropics, the maximum elevation angle (90°) occurs twice a

year. When the Sun has a lower elevation angle, the solarenergy is less intense because it is spread out over a largerarea. Variation of solar elevation is thus one of the mainfactors that accounts for the dependence of climaticregime on latitude. The other main factor is the length of

daylight. For latitudes poleward of 66.5° N and S, thelength of day ranges from zero (winter solstice) to 24 hours(summer solstice), whereas the Equator has a constant12-hour day throughout the year. The seasonal range oftemperature consequently decreases from high latitudesto the tropics, where it becomes less than the diurnalrange of temperature.

The Effects of the Atmosphere

Of the radiant energy reaching the top of the atmosphere,46 percent is absorbed by Earth’s surface on average, butthis value varies significantly from place to place, depend-ing on cloudiness, surface type, and elevation. If there is

persistent cloud cover, as exists in some equatorial regions,much of the incident solar radiation is scattered back tospace, and very little is absorbed by Earth’s surface. Water



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Average exchange of energy between the surface, the atmosphere, and space, as percentages of incident solar radiation (1 unit = 3.4 watts per square metre).Encyclopædia Britannica, Inc.



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surfaces have low reflectivity (4–10 percent), except in lowsolar elevations, and are the most efficient absorbers. Snowsurfaces, on the other hand, have high reflectivity (40–80percent) and so are the poorest absorbers. High-altitudedesert regions consistently absorb higher-than-averageamounts of solar radiation because of the reduced effectof the atmosphere above them.

An additional 23 percent or so of the incident solarradiation is absorbed on average in the atmosphere, espe-cially by water vapour and clouds at lower altitudes and by

ozone (O3 ) in the stratosphere. Absorption of solar radia-tion by ozone shields the terrestrial surface from harmfulultraviolet light and warms the stratosphere, producingmaximum temperatures of −15 to 10 °C (5 to 50 °F) at analtitude of 50 km (30 miles). Most atmospheric absorptiontakes place at ultraviolet and infrared wavelengths, somore than 90 percent of the visible portion of the solar

spectrum, with wavelengths between 0.4 and 0.7 µm(0.00002 to 0.00003 inch), reaches the surface on a cloud-free day. Visible light, however, is scattered in varyingdegrees by cloud droplets, air molecules, and dust parti-cles. Blue skies and red sunsets are in effect attributable tothe preferential scattering of short (blue) wavelengths byair molecules and small dust particles. Cloud droplets

scatter visible wavelengths impartially (hence, clouds usu-ally appear white) but very efficiently, so the reflectivity ofclouds to solar radiation is typically about 50 percent andmay be as high as 80 percent for thick clouds.

The constant gain of solar energy by Earth’s surface issystematically returned to space in the form of thermallyemitted radiation in the infrared portion of the spectrum.

The emitted wavelengths are mainly between 5 and 100µm (0.0002 and 0.004 inch), and they interact differ-ently with the atmosphere compared with the shorter

7 Climate 7




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wavelengths of solar radiation. Very little of the radia-tion emitted by Earth’s surface passes directly throughthe atmosphere. Most of it is absorbed by clouds, car-bon dioxide, and water vapour and is then reemitted inall directions. The atmosphere thus acts as a radiativeblanket over Earth’s surface, hindering the loss of heatto space. The blanketing effect is greatest in the pres-ence of low clouds and weakest for clear cold skies thatcontain little water vapour. Without this effect, the meansurface temperature of 15 °C (59 °F) would be some 30 °C

(86 °F) colder. Conversely, as atmospheric concentrationsof carbon dioxide, methane, chlorofluorocarbons, andother absorbing gases continue to increase, in large partowing to human activities, surface temperatures shouldrise because of the capacity of such gases to trap infraredradiation. The exact amount of this temperature increase,however, remains uncertain because of unpredictable

changes in other atmospheric components, especiallycloud cover. An extreme example of such an effect (com-monly dubbed the greenhouse effect) is that produced bythe dense atmosphere of the planet Venus, which resultsin surface temperatures of about 475 °C (887 °F). This con-dition exists in spite of the fact that the high reflectivity ofthe Venusian clouds causes the planet to absorb less solar

radiation than Earth.

Average Radiation Budgets

The difference between the solar radiation absorbedand the thermal radiation emitted to space determinesEarth’s radiation budget. Since there is no appreciable

long-term trend in planetary temperature, it may beconcluded that this budget is essentially zero on a globallong-term average. Latitudinally, it has been found that



7

much more solar radiation is absorbed at low latitudesthan at high latitudes. On the other hand, thermal emis-sion does not show nearly as strong a dependence onlatitude, so the planetary radiation budget decreases sys-tematically from the Equator to the poles. It changes frombeing positive to negative at latitudes of about 40° N and40° S. The atmosphere and oceans, through their generalcirculation, act as vast heat engines, compensating for thisimbalance by providing nonradiative mechanisms for thetransfer of heat from the Equator to the poles.

While Earth’s surface absorbs a significant amount ofthermal radiation because of the blanketing effect of theatmosphere, it loses even more through its own emissionand thus experiences a net loss of long-wave radiation.This loss is only about 14 percent of the amount emittedby the surface and is less than the average gain of totalabsorbed solar energy. Consequently, the surface has on

average a positive radiation budget.By contrast, the atmosphere emits thermal radiation

both to space and to the surface, yet it receives long-waveradiation back from only the latter. This net loss of ther-mal energy cannot be compensated for by the modest gainof absorbed solar energy within the atmosphere. Theatmosphere thus has a negative radiation budget, equal in

magnitude to the positive radiation budget of the surfacebut opposite in sign. Nonradiative heat transfer againcompensates for the imbalance, this time largely by verti-cal atmospheric motions involving the evaporation andcondensation of water.

Surface-Energy Budgets

The rate of temperature change in any region is directlyproportional to the region’s energy budget and inversely

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proportional to its heat capacity. While the radiation budget may dominate the average energy budget of manysurfaces, nonradiative energy transfer and storage also are

generally important when local changes are considered.Foremost among the cooling effects is the energy

required to evaporate surface moisture, which producesatmospheric water vapour. Most of the latent heat con-tained in water vapour is subsequently released to theatmosphere during the formation of precipitating clouds,although a minor amount may be returned directly to

the surface during dew or frost deposition. Evaporationincreases with rising surface temperature, decreasingrelative humidity, and increasing surface wind speed.Transpiration by plants also increases evaporation rates,

which explains why the temperature in an irrigated fieldis usually lower than that over a nearby dry road surface.

Another important nonradiative mechanism is the

exchange of heat that occurs when the temperature of theair is different from that of the surface. Depending on

whether the surface is warmer or cooler than the air nextto it, heat is transferred to or from the atmosphere by tur-bulent air motion (more loosely, by convection). Thiseffect also increases with increasing temperature differ-ence and with increasing surface wind speed. Direct heat

transfer to the air may be an important cooling mecha-nism that limits the maximum temperature of hot drysurfaces. Alternatively, it may be an important warmingmechanism that limits the minimum temperature of coldsurfaces. Such warming is sensitive to wind speed, so calmconditions promote lower minimum temperatures.

In a similar category, whenever a temperature differ-

ence occurs between the surface and the medium beneaththe surface, there is a transfer of heat to or from themedium. In the case of land surfaces, heat is transferred by



9

conduction, a process where energy is conveyed through amaterial from one atom or molecule to another. In the

case of water surfaces, the transfer is by convection andmay consequently be affected by the horizontal transportof heat within large bodies of water.

TEMPERATURE

The temperature of the air is one of the most important

determinants of Earth’s climate. Although the global aver-age air temperature does not change much from day today or year to year, the mean temperature of any particular

Climatology is a branch of the atmospheric sciences concerned withboth the description of climate and the analysis of the causes of cli-matic differences and changes and their practical consequences.Climatology treats the same atmospheric processes as meteorology,but it seeks as well to identify the slower-acting influences and longer-term changes of import, including the circulation of the oceans andthe small yet measurable variations in the intensity of solar radiation.

From its origins in 6th-century-BCE Greek science, climatologyhas developed along two main lines: regional climatology and physicalclimatology. The first is the study of discrete and characteristic

weather phenomena of a particular continental or subcontinentalregion. The second involves a statistical analysis of the various weatherelements, principally temperature, moisture, atmospheric pressure,and wind speed, and a detailed examination of the basic relationshipsbetween such elements. Since the 1960s a third main branch, dynamicmeteorology, has emerged. It deals primarily with the numerical simu-lation of climate and climatic change, employing models ofatmospheric processes based on the fundamental equations of

dynamic meteorology. Other significant subdisciplines of climatologyinclude bioclimatology and paleoclimatology.

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location may differ substantially from other points onEarth’s surface. A location’s mean temperature dependson geographic factors and how air and water distributeheat to that location. The temperature of a place at anysingle point in time varies according to time of day, season,the altitude of the location, and other factors.

The Global Variation of Mean Temperature

Global variations of average surface-air temperatures are

largely due to latitude, continentality, ocean currents, andprevailing winds.

The effect of latitude is evident in the large north-south gradients in average temperature that occur atmiddle and high latitudes in each winter hemisphere.These gradients are due mainly to the rapid decrease ofavailable solar radiation but also in part to the higher sur-

face reflectivity at high latitudes associated with snow andice and low solar elevations. A broad area of the tropicalocean, by contrast, shows little temperature variation.

Continentality is a measure of the difference betweencontinental and marine climates and is mainly the result ofthe increased range of temperatures that occurs over landcompared with water. This difference is a consequence of

the much lower effective heat capacities of land surfacesas well as of their generally reduced evaporation rates.Heating or cooling of a land surface takes place in a thinlayer, the depth of which is determined by the ability ofthe ground to conduct heat. The greatest temperaturechanges occur for dry, sandy soils, because they are poorconductors with very small effective heat capacities and

contain no moisture for evaporation. By far the greatesteffective heat capacities are those of water surfaces, owingto both the mixing of water near the surface and the



11

penetration of solar radiation that distributes heating todepths of several metres. In addition, about 90 percent ofthe radiation budget of the ocean is used for evaporation.Ocean temperatures are thus slow to change.

The effect of continentality may be moderated byproximity to the ocean, depending on the direction andstrength of the prevailing winds. Contrast with oceantemperatures at the edges of each continent may be fur-ther modified by the presence of a north- or south-flowingocean current. For most latitudes, however, continental-

ity explains much of the variation in average temperatureat a fixed latitude as well as variations in the differencebetween January and July temperatures.

Diurnal, Seasonal, and Extreme Temperatures

The diurnal range of temperature generally increases with

distance from the sea and toward those places where solarradiation is strongest—in dry tropical climates and on highmountain plateaus (owing to the reduced thickness of theatmosphere to be traversed by the Sun’s rays). The averagedifference between the day’s highest and lowest tempera-tures is 3 °C (5 °F) in January and 5 °C (9 °F) in July in thoseparts of the British Isles nearest the Atlantic. The differ-

ence is 4.5 °C (8 °F) in January and 6.5 °C (12 °F) in July onthe small island of Malta. At Tashkent, Uzbekistan, it is 9°C (16 °F) in January and 15.5 °C (28 °F) in July, and atKhartoum, Sudan, the corresponding figures are 17 °C (31°F) and 13.5 °C (24 °F). At Kandahar, Afghanistan, which liesmore than 1,000 metres (about 3,300 feet) above sea level,it is 14 °C (25 °F) in January and 20 °C (36 °F) in July. There,

the average difference between the day’s highest and low-est temperatures exceeds 23 °C (41 °F) in September andOctober, when there is less cloudiness than in July. Near

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WORLD TEMPERATURE EXTREMEShighest recorded air temperature

temperaturecontinent or

region

place (with

elevation*)

degrees

C

degrees

F

Africa Al-`Aziziyah, Libya (112 mor 367 ft)

57.7 136

Antarctica Lake Vanda 77 degrees 32minutes S, 161 degrees 40

minutes E (99 m or 325 ft)

15 59

Asia Tirat Zevi, Israel (–300 mor –984 ft)

53.9 129

Australia Cloncurry, Queensland(193 m or 633 ft)

53.1 127.5

Europe Sevilla, Spain (39 m or 128 ft) 50 122

North America Death Valley (Greenland

Ranch), California, U.S.(–54 m or –177 ft)

56 134.5

South America Rivadavia, Argentina (205m or 672 ft)

48.9 120

Tropical Pacificislands

Echague, Luzon, Republicof the Philippines (78 m or257 ft)

40.5 105

lowest recorded air temperature

temperature

continent or

region

place (with

elevation*)

degrees

C

degrees

F

Africa Ifrane, Morocco (1,635 mor 5,363 ft)

–23.9 –11

Antarctica Vostok 78 degrees 27minutes S, 106 degrees 52minutes E (3,420 m or11,218 ft)

–89.2 –128.6



13

Asia Oymyakon, Russia (806 mor 2,644 ft)

–67.7 –89.9

Australia Charlotte Pass, New SouthWales (1,780 m or 5,840 ft)

–22.2 –8

Europe Ust-Shchuger, Russia(85 mor 279 ft)

–55 –67

North America Snag, Yukon Terr., Canada(646 m or 2,119 ft)

–62.8 –81

South America Colonia, Sarmiento,Argentina (268 m or 879 ft)

–33 –27

Tropical Pacificislands

Haleakala, Hawaii, U.S.(2,972 m or 9,748 ft)

–7.8 18

*Above or below sea level.

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the ocean at Colombo, Sri L., the figures are 8 °C (14 °F) in January and 4.5 °C (8 °F) in July.

The seasonal variation of temperature and the magni-tudes of the differences between the same month indifferent years and different epochs generally increasetoward high latitudes and with distance from the ocean.

Temperature Variation with Height

There are two main levels where the atmosphere isheated—namely, at Earth’s surface and at the top of theozone layer (about 50 km, or 30 miles, up) in the strato-sphere. Radiation balance shows a net gain at these levelsin most cases. Prevailing temperatures tend to decrease

with distance from these heating surfaces (apart from theionosphere and the outer atmospheric layers, where other

processes are at work). The world’s average lapse rate oftemperature (change with altitude) in the lower atmo-sphere is 0.6 to 0.7 °C per 100 metres (about 1.1 to 1.3 °Fper 300 feet). Lower temperatures prevail with increasing




14

height above sea level for two reasons: (1) because there isa less favourable radiation balance in the free air, and (2)because rising air—whether lifted by convection currentsabove a relatively warm surface or forced up over moun-tains—undergoes a reduction of temperature associated

with its expansion as the pressure of the overlying atmo-sphere declines. This is the adiabatic lapse rate oftemperature, which equals about 1 °C per 100 metres(about 2 °F per 300 feet) for dry air and 0.5 °C per 100metres (about 1 °F per 300 feet) for saturated air, in which

condensation (with liberation of latent heat) is producedby adiabatic cooling. The difference between these ratesof change of temperature (and therefore density) of risingair currents and the state of the surrounding air deter-mines whether the upward currents are accelerated orretarded—i.e., whether the air is unstable, so vertical con-

vection with its characteristically attendant tall cumulus

cloud and shower development is encouraged or whetherit is stable and convection is damped down.

For these reasons, the air temperatures observed onhills and mountains are generally lower than on low

ground, except in the case of extensive plateaus, whichpresent a raised heating surface (and on still, sunny days,

when even a mountain peak is able to warm appreciably

the air that remains in contact with it).

Circulation, Currents, andOcean-Atmosphere Interaction

The circulation of the ocean is a key factor in air tempera-ture distribution. Ocean currents that have a northward

or southward component, such as the warm Gulf Streamin the North Atlantic or the cold Peru (Humboldt) Currentoff South America, effectively exchange heat between low



15

and high latitudes. In tropical latitudes the ocean accountsfor a third or more of the poleward heat transport; at lati-tude 50° N, the ocean’s share is about one-seventh. In theparticular sectors where the currents are located, theirimportance is of course much greater than these figures,

which represent hemispheric averages.A good example of the effect of a warm current is that

of the Gulf Stream in January, which causes a strongeast-west gradient in temperatures across the eastern edgeof the North American continent. The relative warmth of

the Gulf Stream affects air temperatures all the way acrossthe Atlantic, and prevailing westerlies extend the warmingeffect deep into northern Europe. As a result, Januarytemperatures of Tromsø, Nor. (69°40’ N), for example,average 24 °C (43 °F) above the mean for that latitude. TheGulf Stream maintains a warming influence in July, but it isnot as noticeable because of the effects of continentality.

The ocean, particularly in areas where the surface is warm, also supplies moisture to the atmosphere. This inturn contributes to the heat budget of those areas in whichthe water vapour is condensed into clouds, liberatinglatent heat in the process. This set of events occurs fre-quently in high latitudes and in locations remote from theocean where the moisture was initially taken up.

The great ocean currents are themselves wind-driven—set in motion by the drag of the winds over vastareas of the sea surface, especially where the tops of wavesincrease the friction with the air above. At the limits of the

warm currents, particularly where they abut directly upona cold current—as at the left flank of the Gulf Stream inthe neighbourhood of the Grand Banks off Newfoundland

and at the subtropical and Antarctic convergences in theoceans of the Southern Hemisphere—the strong thermal gradients in the sea surface result in marked differences

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in the heating of the atmosphere on either side of theboundary. These temperature gradients tend to positionand guide the strongest flow of the jet stream in the atmo-sphere above and thereby influence the development andsteering of weather systems.

Interactions between the ocean and the atmosphereproceed in both directions. They also operate at differentrates. Some interesting lag effects, which are of value inlong-range weather forecasting, arise through the consid-erably slower circulation of the ocean. Thus, enhanced

strength of the easterly trade winds over low latitudes ofthe Atlantic north and south of the Equator impels more

water toward the Caribbean and the Gulf of Mexico, pro-ducing a stronger flow and greater warmth in the GulfStream approximately six months later. Anomalies in theposition of the Gulf Stream–Labrador Current boundary,

which produce a greater or lesser extent of warm water

Major surface currents of the world’s oceans. Subsurface currents also movevast amounts of water, but they are not known in such detail. EncyclopaediaBritannica, Inc.



17

near the Grand Banks, so affect the energy supply to theatmosphere and the development and steering of weathersystems from that region that they are associated withrather persistent anomalies of weather pattern over theBritish Isles and northern Europe. Anomalies in the equa-torial Pacific and in the northern limit of the KuroshioCurrent (also called the Japan Current) seem to haveeffects on a similar scale. Indeed, through their influenceon the latitude of the jet stream and the wavelength (thatis, the spacing of cold trough and warm ridge regions) in

the upper westerlies, these ocean anomalies exercise aninfluence over the atmospheric circulation that spreads toall parts of the hemisphere.

Sea-surface temperature anomalies that recur in theequatorial Pacific at variable intervals of two to seven

years can sometimes produce major climatic perturba-tions. One such anomaly is known as El Niño (Spanish for

“The Child”; it was so named by Peruvian fishermen whonoticed its onset during the Christmas season).

During an El Niño event, warm surface water flowseastward from the equatorial Pacific, in at least partialresponse to weakening of the equatorial easterly winds,and replaces the normally cold upwelling surface water offthe coast of Peru and Ecuador that is associated with the

northward propagation of the cold Peru Current. Thechange in sea-surface temperature transforms the coastalclimate from arid to wet. The event also affects atmo-spheric circulation in both hemispheres and is associated

with changes in precipitation in regions of North America,Africa, and the western Pacific.

Short-Term Temperature Changes

Many interesting short-term temperature fluctuationsalso occur, usually in connection with local weather

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disturbances. The rapid passage of a mid-latitude coldfront, for example, can drop temperatures by 10 °C (18 °F)in a few minutes and, if followed by the sustained move-ment of a cold air mass, by as much as 50 °C in 24 hours,

with life-threatening implications for the unwary.Temperature increases of up to 40 °C in a few hours alsoare possible downwind of major mountain ranges when airthat has been warmed by the release of latent heat on the

windward side of a range is forced to descend rapidly onthe other side (such a wind is variously called chinook,

foehn, or Santa Ana). Changes of this kind, however,involve a wider range of meteorological processes thandiscussed in this chapter.

ATMOSPHERIC HUMIDITYAND PRECIPITATION

Atmospheric humidity, which is the amount of water vapour or moisture in the air, is another leading climaticelement, as is precipitation. All forms of precipitation,including drizzle, rain, snow, ice crystals, and hail, are pro-duced as a result of the condensation of atmosphericmoisture that forms clouds in which some of the particles,by growth and aggregation, attain sufficient size to fall

from the clouds and reach the ground.

Atmospheric Humidity

At 30 °C (86 °F), 4 percent of the volume of the air may beoccupied by water molecules, but, where the air is colderthan –40 °C (–40 °F), less than one-fifth of 1 percent of the

air molecules can be water. Although the water vapourcontent may vary from one air parcel to another, theselimits can be set because vapour capacity is determined by



19

temperature. Temperature has profound effects uponsome of the indexes of humidity, regardless of the pres-ence or absence of vapour.

The connection between an effect of humidity and anindex of humidity requires simultaneous introduction ofeffects and indexes. Vapour in the air is a determinantof weather, because it first absorbs the thermal radiationthat leaves and cools Earth’s surface and then emits ther-mal radiation that warms the planet. Calculation ofabsorption and emission requires an index of the mass

of water in a volume of air. Vapour also affects the weatherbecause, as indicated above, it condenses into clouds andfalls as rain or other forms of precipitation. Tracing themoisture-bearing air masses requires a humidity indexthat changes only when water is removed or added.

Humidity Indexes

There are three indexes that express humidity. Of these,relative humidity, which is the amount of water vapour inthe air relative to the amount the air can hold, is the mostcommonly used.

Absolute Humidity

Absolute humidity is the vapour concentration or density

in the air. If mv is the mass of vapour in a volume of air,then absolute humidity dv is simply dv = mv/ V , in which V is the volume and dv is expressed in grams per cubic metre.This index indicates how much vapour a beam of radia-tion must pass through. The ultimate standard in humiditymeasurement is made by weighing the amount of water

gained by an absorber when a known volume of air passes

through it; this measures absolute humidity, which may vary from 0 gram per cubic metre in dry air to 30 gramsper cubic metre (0.03 ounce per cubic foot) when the

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vapour is saturated at 30 °C (86 °F). The dv of a parcel of airchanges, however, with temperature or pressure eventhough no water is added or removed, because, as the gasequation states, the volume V increases with the absolute,or Kelvin, temperature and decreases with the pressure.

Specific Humidity

The meteorologist requires an index of humidity that doesnot change with pressure or temperature. A property ofthis sort will identify an air mass when it is cooled or when

it rises to lower pressures aloft without losing or gaining water vapour. Because all the gases will expand equally, theratios of the weight of water to the weight of dry air, or thedry air plus vapour, will be conserved during such changesand will continue identifying the air mass.

The mixing ratio r is the dimensionless ratio r = mv/ ma, where ma is the mass of dry air, and the specific humidity

q is another dimensionless ratio q = mv/ ( ma + mv ). Because mv is less than 3 percent of ma at normal pressure and tem-peratures cooler than 30 °C, r and q are practically equal.These indexes are usually expressed in grams per kilogrambecause they are so small; the values range from 0 grams perkilogram in dry air to 28 grams per kilogram in saturatedair at 30 °C. Absolute and specific humidity indexes have

specialized uses, so they are not familiar to most people. Relative Humidity

Relative humidity ( U ) is so commonly used that a state-ment of humidity, without a qualifying adjective, can beassumed to be relative humidity. U can be defined, then, interms of the mixing ratio r that was introduced above. U =

100 r / rw, which is a dimensionless percentage. The divisor rw is the saturation mixing ratio, or the vapour capacity.Relative humidity is therefore the water vapour content



21

of the air relative to its content at saturation. Because thesaturation mixing ratio is a function of pressure, and espe-cially of temperature, the relative humidity is a combinedindex of the environment that reflects more than watercontent. In many climates the relative humidity rises toabout 100 percent at dawn and falls to 50 percent by noon.A relative humidity of 50 percent may reflect many differ-ent quantities of vapour per volume of air or gram of air,and it will not likely be proportional to evaporation.

An understanding of relative humidity thus requires

a knowledge of saturated vapour, which will be dis-cussed in the following section, on the relation betweentemperature and humidity. At this point, however, therelation between U and the absorption and retentionof water from the air must be considered. Small poresretain water more strongly than large pores; thus, when aporous material is set out in the air, all pores larger than

a certain size (which can be calculated from the relativehumidity of the air) are dried out.

The water content of a porous material at air tem-perature is fairly well indicated by the relative humidity.The complexity of actual pore sizes and the viscosityof the water passing through them makes the relationbetween U and moisture in the porous material imper-

fect and slowly achieved. The great suction also strainsthe walls of the capillaries, and the consequent shrinkageis used to measure relative humidity.

The absorption of water by salt solutions is also relatedto relative humidity without much effect of temperature.The air above water saturated with sodium chloride ismaintained at 75 to 76 percent relative humidity at a tem-

perature between 0 and 40 °C (32 and 104 °F).In effect, relative humidity is a widely used environ-mental indicator, but U does respond drastically to changes

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22

in temperatures as well as moisture, a response caused bythe effect of temperature upon the divisor r W in U .

The Relationship Between Temperature and Humidity

Tables that show the effect of temperature upon the satu-ration mixing ratio r w are readily available. Humidity ofthe air at saturation is expressed more commonly, how-ever, as vapour pressure. Thus, it is necessary to understand

vapour pressure and in particular the gaseous nature of

water vapour.The pressure of the water vapour, which contributes

to the pressure of the atmosphere, can be calculated fromthe absolute humidity d v by the gas equation:

in which R is the gas constant, T the absolute tempera-ture, M w the molecular weight of water, and e the water vapour pressure in millibars (mb).

Relative humidity can be defined as the ratio of the vapour pressure of a sample of air to the saturation pres-sure at the existing temperature. Further, the capacity for

vapour and the effect of temperature can now be pre-sented in the usual terms of saturation vapour pressure.Within a pool of liquid water, some molecules are con-

tinually escaping from the liquid into the space above, while more and more vapour molecules return to the liq-uid as the concentration of vapour rises. Finally, equalnumbers are escaping and returning; the vapour is then

saturated, and its pressure is known as the saturation vapour pressure, e w . If the liquid and vapour are warmed,relatively more molecules escape than return, and e w rises.



23

There is also a saturation pressure with respect to ice. The vapour pressure curve of water has the same form asthe curves for many other substances. Its location is fixed,however, by the boiling point of 100 °C (212 °F), where thesaturation vapour pressure of water vapour is 1,013 mb (1standard atmosphere), the standard pressure of the atmo-sphere at sea level. The decrease of the boiling point withaltitude can be calculated. For example, the saturation

vapour pressure at 40 °C (104 °F) is 74 mb (0.07 standardatmosphere), and the standard atmospheric pressure near

18,000 metres (59,000 feet) above sea level is also 74 mb;thus, it is where water boils at 40° C.

The everyday response of relative humidity to temper-ature can be easily explained. On a summer morning, thetemperature might be 15 °C (59 °F) and the relative humid-ity 100 percent. The vapour pressure would be 17 mb (0.02standard atmosphere) and the mixing ratio about 11 parts

per thousand (11 grams of water per kilogram of air by weight). During the day the air could warm to 25 °C (77 °F), while evaporation could add little water. At 25 °C the satu-ration pressure is fully 32 mb (0.03 standard atmosphere).If, however, little water has been added to the air, its

vapour pressure will still be about 17 mb. Thus, with nochange in vapour content, the relative humidity of the air

has fallen from 100 to only 53 percent, illustrating whyrelative humidity does not identify air masses.The meaning of dew-point temperature can be illus-

trated by a sample of air with a vapour pressure of 17 mb.If an object at 15 °C is brought into the air, dew will formon the object. Hence, 15 °C is the dew-point temperatureof the air—i.e., the temperature at which the vapour pres-

ent in a sample of air would just cause saturation or thetemperature whose saturation vapour pressure equalsthe present vapour pressure in a sample of air. Below

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freezing, this index is called the frost point. There is aone-to-one correspondence between vapour pressureand dew point. The dew point has the virtue of being eas-ily interpreted because it is the temperature at which ablade of grass or a pane of glass will become wet with dewfrom the air. Ideally, it is also the temperature of fog orcloud formation.

The clear meaning of dew point suggests a meansof measuring humidity. A dew-point hygrometer wasinvented in 1751. For this instrument, cold water was

added to water in a vessel until dew formed on the vessel,and the temperature of the vessel, the dew point, pro-

vided a direct index of humidity. The greatest use of thecondensation hygrometer has been to measure humid-ity in the upper atmosphere, where a vapour pressureof less than a thousandth millibar makes other meansimpractical.

Another index of humidity, the saturation deficit,can also be understood by considering air with a vapourpressure of 17 mb. At 25 °C the air has (31 − 17), or 14, mbless vapour pressure than saturated vapour at the sametemperature; that is, the saturation deficit is 14 mb (0.01standard atmosphere).

The saturation deficit has the particular utility of

being proportional to the evaporation capability of the air.The saturation deficit can be expressed as

and, because the saturation vapour pressure e w rises with

rising temperature, the same relative humidity will corre-spond to a greater saturation deficit and evaporation at warm temperatures.



25

Humidity and Climate

The small amount of water in atmospheric vapour, relativeto water on Earth, belies its importance. Compared withone unit of water in the air, the seas contain at least100,000 units, the great glaciers 1,500, the porous earthnearly 200, and the rivers and lakes 4 or 5. The effective-ness of the vapour in the air is magnified, however, by itsrole in transferring water from sea to land by the media ofclouds and precipitation and that in absorbing radiation.

The vapour in the air is the invisible conductor thatcarries water from sea to land, making terrestrial life pos-sible. Fresh water is distilled from the salt seas and carriedover land by the wind. Water evaporates from vegetation,and rain falls on the sea too, but the sea is the bigger source,and rain that falls on land is most important to humans.The invisible vapour becomes visible near the surface as

fog when the air cools to the dew point. The usual noctur-nal cooling will produce fog patches in cool valleys. Or the

vapour may move as a tropical air mass over cold land orsea, causing widespread and persistent fog, such as occursover the Grand Banks off Newfoundland. The delivery of

water by means of fog or dew is slight, however.When air is lifted, it is carried to a region of lower

pressure, where it will expand and cool as described by the gas equation. It may rise up a mountain slope or overthe front of a cooler, denser air mass. If condensationnuclei are absent, the dew point may be exceeded by thecooling air, and the water vapour becomes supersaturated.If nuclei are present or if the temperature is very low, how-ever, cloud droplets or ice crystals form, and the vapour is

no longer in the invisible guise of atmospheric humidity.The invisible vapour has another climatic role—namely, absorbing and emitting radiation. The

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26

temperature of Earth and its daily variation are deter-mined by the balance between incoming and outgoingradiation. The wavelength of the incoming radiationfrom the Sun is mostly shorter than 3 µm (0.0001 inch).It is scarcely absorbed by water vapour, and its receiptdepends largely upon cloud cover. The radiationexchanged between the atmosphere and Earth’s sur-face and the eventual loss to space is in the form of long

waves. These long waves are strongly absorbed in the 3- to8.5-µm band and in the greater than 11-µm range, where

vapour is either partly or wholly opaque. As noted above,much of the radiation that is absorbed in the atmosphereis emitted back to Earth, and the surface receipt of long

waves, primarily from water vapour and carbon dioxidein the atmosphere, is slightly more than twice the directreceipt of solar radiation at the surface. Thus, the invis-ible vapour in the atmosphere combines with clouds and

the advection (horizontal movement) of air from differ-ent regions to control the surface temperature.

The world distribution of humidity can be portrayedfor different uses by different indexes. To appraise thequantity of water carried by the entire atmosphere,the moisture in an air column above a given point onEarth is expressed as a depth of liquid water. It varies

from 0.5 mm (0.02 inch) over the Himalayas and 2 mm(0.08 inch) over the poles in winter to 8 mm (0.3 inch)over the Sahara, 54 mm (2 inches) in the Amazon region,and 64 mm (2.5 inches) over India during the wet season.During summer the air over the United States trans-ports 16 mm (0.6 inch) of water vapour over the GreatBasin and 45 mm (1.8 inches) over Florida.

The humidity of the surface air may be mapped as vapour pressure, but a map of this variable looks much likethat of temperature. Warm places are moist, and cool ones



27

are dry; even in deserts the vapour pressure is normally 13mb (0.01 standard atmosphere), whereas over the north-ern seas it is only about 4 mb (0.004 standard atmosphere).Certainly the moisture in materials in two such areas willbe just the opposite, so relative humidity is a more widelyuseful index.

Average Relative Humidity

The average relative humidity for July reveals the humid-ity provinces of the Northern Hemisphere when aridity is

at a maximum. At other times the relative humidity gener-ally will be higher. The humidities over the SouthernHemisphere in July indicate the humidities that compa-rable regions in the Northern Hemisphere will attain in

January, just as July in the Northern Hemisphere suggeststhe humidities in the Southern Hemisphere during

January. A contrast is provided by comparing a humid cool

coast to a desert. The midday humidity on the Oregoncoast, for example, falls only to 80 percent, whereas in theNevada desert it falls to 20 percent. At night the contrastis less, with averages being over 90 and about 50 percent,respectively.

Although the dramatic regular decrease of relativehumidity from dawn to midday has been attributed largely

to warming rather than declining vapour content, the con-tent does vary regularly. In humid environments, daytimeevaporation increases the water vapour content of the air,and the mixing ratio, which may be about 12 grams perkilogram, rises by 1 or 2 grams per kilogram in temperateplaces and may attain 16 grams per kilogram in a tropicalrainforest. In arid environments, however, little evapora-

tion moistens the air, and daytime turbulence tends tobring down dry air; this decreases the mixing ratio by asmuch as 2 grams per kilogram.

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Humidity also varies regularly with altitude. On theaverage, fully half the water in the atmosphere lies below0.25 km (about 0.2 mile), and satellite observations overthe United States in April revealed 1 mm (0.04 inch) orless of water in all the air above 6 km (4 miles). A crosssection of the atmosphere along 75° W longitude showsa decrease in humidity with height and toward the poles.The mixing ratio is 16 grams per kilogram just northof the Equator, but it decreases to 1 gram per kilogram at50° N latitude or 8 km (5 miles) above the Equator. The

transparent air surrounding mountains in fair weather is very dry indeed.

Closer to the ground, the water vapour content alsochanges with height in a regular pattern. When water

vapour is condensing on Earth’s surface at night, the con-tent is greater aloft than at the ground; during the day thecontent is, in most cases, less aloft than at the ground

because of evaporation.

Evaporation and Humidity

Evaporation, mostly from the sea and from vegetation,replenishes the humidity of the air. It is the change of liq-uid water into a gaseous state, but it may be analyzed asdiffusion. The rate of diffusion, or evaporation, will be

proportional to the difference between the pressure ofthe water vapour in the free air and the vapour that is nextto, and saturated by, the evaporating liquid. If the liquidand air have the same temperature, evaporation is propor-tional to the saturation deficit. It is also proportional tothe conductivity of the medium between the evaporatorand the free air. If the evaporator is open water, the con-

ductivity will increase with ventilation. But if theevaporator is a leaf, the diffusing water must pass throughthe still air within the minute pores between the water



29

Global distribution of mean annual evaporation (in centimetres).Encyclopaedia Britannica, Inc.

inside and the dry air outside. In this case the porositymay modify the conductivity more than ventilation.

The temperature of the evaporator is rarely the sameas the air temperature, however, because each gram ofevaporation consumes about 600 calories (2,500 joules)and thus cools the evaporator. The availability of energy toheat the evaporator is therefore as important as the satu-ration deficit and conductivity of the air. Outdoors, someof this heat may be transferred from the surrounding airby convection, but much of it must be furnished by radia-

tion. Evaporation is faster on sunny days than on cloudyones not only because the sunny day may have drier air butalso because the Sun warms the evaporator and thus raisesthe vapour pressure at the evaporator. In fact, accordingto the well-known Penman calculation of evaporation (an

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equation that considers potential evaporation as a func-tion of humidity, wind speed, radiation, and temperature),this loss of water is essentially determined by the net radi-ation balance during the day.

Precipitation

Precipitation is one of the three main processes (evapora-tion, condensation, and precipitation) that constitute thehydrologic cycle, the continual exchange of water between

the atmosphere and Earth’s surface. Water evaporatesfrom ocean, land, and freshwater surfaces, is carried aloftas vapour by the air currents, condenses to form clouds,and ultimately is returned to Earth’s surface as precipita-tion. The average global stock of water vapour in theatmosphere is equivalent to a layer of water 2.5 cm (1 inch)deep covering the whole Earth. Because Earth’s average

annual rainfall is about 100 cm (39 inches), the averagetime that the water spends in the atmosphere, between itsevaporation from the surface and its return as precipita-tion, is about ¼0 of a year, or about nine days. Of the water

vapour that is carried at all heights across a given region bythe winds, only a small percentage is converted into pre-cipitation and reaches the ground in that area. In deep and

extensive cloud systems, the conversion is more efficient,but even in thunderclouds the quantities of rain and hailreleased amount to only some 10 percent of the total mois-ture entering the storm.

In the measurement of precipitation, it is necessary todistinguish between the amount—defined as the depth ofprecipitation (calculated as though it were all rain) that has

fallen at a given point during a specified interval of time—and the rate or intensity, which specifies the depth of water that has fallen at a point during a particular interval



31

of time. Persistent moderate rain, for example, might fallat an average rate of 5 mm per hour (0.2 inch per hour)and thus produce 120 mm (4.7 inches) of rain in 24 hours.A thunderstorm might produce this total quantity of rainin 20 minutes, but at its peak intensity the rate of rainfallmight become much greater—perhaps 120 mm per hour(4.7 inches per hour), or 2mm (0.08 inch) per minute—fora minute or two.

The amount of precipitation falling during a fixedperiod is measured regularly at many thousands of

places on Earth’s surface by rather simple rain gauges.Measurement of precipitation intensity requires a record-ing rain gauge, in which water falling into a collector ofknown surface area is continuously recorded on a movingchart or a magnetic tape. Investigations are being carriedout on the feasibility of obtaining continuous measure-ments of rainfall over large catchment areas by means

of radar.Apart from the trifling contributions made by dew,

frost, and rime, as well as desalination plants, the solesource of fresh water for sustaining rivers, lakes, and alllife on Earth is provided by precipitation from clouds.Precipitation is therefore indispensable and overwhelm-ingly beneficial to humankind, but extremely heavy rainfall

can cause great harm: soil erosion, landslides, and flood-ing. Hailstorm damage to crops, buildings, and livestockcan prove very costly.

The Origin of Precipitation in Clouds

Clouds are formed by the lifting of damp air, which coolsby expansion as it encounters the lower pressures existing

at higher levels in the atmosphere. The relative humidityincreases until the air has become saturated with water vapour, and then condensation occurs on any of the

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aerosol particles suspended in the air. A wide variety ofthese exist in concentrations ranging from only a few percubic centimetre in clean maritime air to perhaps 1 mil-lion per cubic cm (16 million per cubic inch) in the highlypolluted air of an industrial city. For continuous condensa-tion leading to the formation of cloud droplets, the airmust be slightly supersaturated. Among the highly effi-cient condensation nuclei are sea-salt particles and theparticles produced by combustion (e.g., natural forest firesand man-made fires). Many of the larger condensation

nuclei over land consist of ammonium sulfate. These areproduced by cloud and fog droplets absorbing sulfur diox-ide and ammonia from the air. Condensation onto thenuclei continues as rapidly as water vapour is made avail-able through cooling; droplets about 10 µm (0.0004 inch)in diameter are produced in this manner. These dropletsconstitute a nonprecipitating cloud.

Cloud Types

The meteorologist classifies clouds mainly by their appear-ance, according to an international system similar to oneproposed in 1803. But because the dimensions, shape,structure, and texture of clouds are influenced by the kindof air movements that result in their formation and growth

and by the properties of the cloud particles, much of what was originally a purely visual classification can now be jus-tified on physical grounds.

The first International Cloud Atlas was published in 1896.Developments in aviation during World War I stimulatedinterest in cloud formations and in their importance as anaid in short-range weather forecasting. This led to the pub-

lication of a more extensive atlas, the International Atlasof Clouds and States of Sky, in 1932 and to a revised editionin 1939. After World War II, the World Meteorological



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Different types of clouds form at different heights. EncyclopædiaBritannica, Inc.

Organization published a new International Cloud Atlas (1956) in two volumes. It contains 224 plates, describing

10 main cloud genera (families) subdivided into 14 speciesbased on cloud shape and structure. Nine general variet-ies, based on transparency and geometric arrangement,also are described. The genera, listed according to theirheight, are as follows:

1. High: mean heights from 5 to 13 km, or 3 to 8 miles

a. Cirrusb. Cirrocumulusc. Cirrostratus

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2. Middle: mean heights 2 to 7 km, or 1 to 4 milesa. Altocumulusb. Altostratusc. Nimbostratus

3. Low: mean heights 0 to 2 km, or 0 to 1.2 miles

a. Stratocumulusb. Stratusc. Cumulusd. Cumulonimbus

Heights given are approximate averages for temperate

latitudes. Clouds of each genus are generally lower in thepolar regions and higher in the tropics.

Four principal classes are recognized when clouds areclassified according to the kind of air motions that producethem: (1) layer clouds formed by the widespread regular

ascent of air, (2) layer clouds formed by widespread irregu-lar stirring or turbulence, (3) cumuliform clouds formed bypenetrative convection, and (4) orographic clouds formedby the ascent of air over hills and mountains.

The widespread layer clouds associated with cyclonicdepressions, near fronts and other inclement-weathersystems, are frequently composed of several layers that

may extend up to 9 km (5.6 miles) or more, separated byclear zones that become filled in as rain or snow develops.These clouds are formed by the slow, prolonged ascentof a deep layer of air, in which a rise of only a few cen-timetres per second is maintained for several hours. Inthe neighbourhood of fronts, vertical velocities becomemore pronounced and may reach about 10 cm (4 inches)

per second.Most of the high cirrus clouds visible from the groundlie on the fringes of cyclonic cloud systems, and, though



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due primarily to regular ascent, their pattern is oftendetermined by local wave disturbances that finally triggertheir formation after the air has been brought near its sat-uration point by the large-scale lifting.

On a cloudless night, the ground cools by radiatingheat into space without heating the air adjacent to the

ground. If the air were quite still, only a very thin layer would be chilled by contact with the ground. More usually,however, the lower layers of the air are stirred by motionover the rough ground, so the cooling is distributed

through a much greater depth. Consequently, when the airis damp or the cooling is great, a fog a few hundred metresdeep may form, rather than a dew produced by condensa-tion on the ground.

In moderate or strong winds, the irregular stirringnear the surface distributes the cooling upward, and thefog may lift from the surface to become a stratus cloud,

which is not often more than 600 metres (about 2,000feet) thick.

Radiational cooling from the upper surfaces of fogsand stratus clouds promotes an irregular convection

within the cloud layer and causes the surfaces to have a waved or humped appearance. When the cloud layer isshallow, billows and clear spaces may develop; it is then

described as stratocumulus instead of stratus.Usually, cumuliform clouds appearing over land areformed by the rise of discrete masses of air from near thesunlight-warmed surface. These rising lumps of air, or ther-mals, may vary in diameter from a few tens to hundreds ofmetres as they ascend and mix with the cooler, drier airabove them. Above the level of the cloud base, the release

of latent heat of condensation tends to increase the buoy-ancy of the rising masses, which tower upward and emergeat the top of the cloud with rounded upper surfaces.

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Middle clouds

(Top) Altocumulus undulatus , a layer of shaded, regularly arranged rolls.Richard Jepperson/Photo Researcher (Bottom) Altocumulus perlu-cidus , a white and gray layer in which there are spaces between the elements.H. von Meiss-Teuffen/Photo Researcher



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Middle clouds

(Top) Altostratus translucidus , showing the Sun as if seen through ground glass. Lawrence Smith/Photo Researcher (Bottom) Altocumulus radia-tus , a layer with laminae arranged in parallel bands. Russ Kinne/PhotoResearcher




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At any moment a large cloud may contain a number ofactive thermals and the residues of earlier ones. A newthermal rising into a residual cloud will be partially pro-tected from having to mix with the cool, dry environmentand therefore may rise farther than its predecessor. Oncea thermal has emerged as a cloud turret at the summit orthe flanks of the cloud, rapid evaporation of the dropletschills the cloud borders, destroys the buoyancy, and pro-duces sinking. A cumulus thus has a characteristicpyramidal shape and, viewed from a distance, appears to

have an unfolding motion, with fresh cloud masses con-tinually emerging from the interior to form the summitand then sinking aside and evaporating.

In settled weather, cumulus clouds are well scatteredand small; horizontal and vertical dimensions are only akilometre or two. In disturbed weather, they cover a largepart of the sky, and individual clouds may tower as high as

10 km (6 miles) or more, often ceasing their growth onlyupon reaching the stable stratosphere. These clouds pro-duce heavy showers, hail, and thunderstorms .

At the level of the cloud base, the speed of the risingair masses is usually about 1 metre (3.3 feet) per second butmay reach 5 metres (16 feet) per second, and similar valuesare measured inside smaller clouds. The upcurrents in

thunderclouds, however, often exceed 5 metres per secondand may reach 30 metres (98 feet) per second or more.The rather special orographic clouds are produced by

the ascent of air over hills and mountains. The air streamis set into oscillation when it is forced over the hill, andthe clouds form in the crests of the (almost) stationary

waves. There may thus be a succession of such clouds

stretching downwind of the mountain, which remain sta-tionary relative to the ground in spite of strong winds thatmay be blowing through the clouds. The clouds have very



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smooth outlines and are called lenticular (lens-shaped) or“wave” clouds. Thin wave clouds may form at great heights(up to 10 km, even over hills only a few hundred metreshigh) and occasionally are observed in the stratosphere (at20 to 30 km [12 to 19 miles]) over the mountains of Norway,Scotland, Iceland, and Alaska. These atmospheric waveclouds are known as nacreous or “mother-of-pearl” cloudsbecause of their brilliant iridescent colours.

The Mechanisms of Precipitation Release

Growing clouds are sustained by upward air currents, which may vary in strength from a few centimetres persecond to several metres per second. Considerable growthof the cloud droplets (with falling speeds of only about 1cm, or 0.4 inch, per second) is therefore necessary if theyare to fall through the cloud, survive evaporation in theunsaturated air below, and reach the ground as drizzle or

rain. The production of a few large particles from a largepopulation of much smaller ones may be achieved in oneof two ways. The first of these depends on the fact thatcloud droplets are seldom of uniform size; droplets formon nuclei of various sizes and grow under slightly differentconditions and for different lengths of time in differ-ent parts of the cloud. A droplet appreciably larger than

average will fall faster than the smaller ones and so willcollide and fuse (coalesce) with some of those that it over-takes. Calculations show that, in a deep cloud containingstrong upward air currents and high concentrations ofliquid water, such a droplet will have a sufficiently longjourney among its smaller neighbours to grow to rain-drop size. This coalescence mechanism is responsible for

the showers that fall in tropical and subtropical regionsfrom clouds whose tops do not reach altitudes where airtemperatures are below 0 °C (32 °F) and therefore cannot

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contain ice crystals. Radar evidence also suggests thatshowers in temperate latitudes may sometimes be ini-tiated by the coalescence of waterdrops, although theclouds may later reach heights at which ice crystals mayform in their upper parts.

The second method of releasing precipitation canoperate only if the cloud top reaches elevations at whichair temperatures are below 0 °C and the droplets in theupper cloud regions become supercooled. At temperaturesbelow –40 °C (–40 °F), the droplets freeze automatically

or spontaneously. At higher temperatures, they can freezeonly if they are infected with special minute particles calledice nuclei. The origin and nature of these nuclei are notknown with certainty, but the most likely source is clay-silicate particles carried up from the ground by the wind.As the temperature falls below 0 °C, more and more icenuclei become active, and ice crystals appear in increasing

numbers among the supercooled droplets. Such a mix-ture of supercooled droplets and ice crystals is unstable,however. The cloudy air is usually only slightly supersatu-rated with water vapour with respect to the droplets andis strongly oversaturated with respect to ice crystals; thelatter thus grow more rapidly than the droplets. After sev-eral minutes, the growing crystals acquire falling speeds

of tens of centimetres per second, and several of themmay become joined to form a snowflake. In falling into the warmer regions of the cloud, this flake may melt and hit ground as a raindrop.

The deep, extensive, multilayer cloud systems, from which precipitation of a widespread persistent characterfalls, are generally formed in cyclonic depressions (lows)

and near fronts. Cloud systems of this type are associ-ated with feeble upcurrents of only a few centimetresper second that last for at least several hours. Although



41

the structure of these great rain-cloud systems is beingexplored by aircraft and radar, it is not yet well under-stood. That such systems rarely produce rain, as distinctfrom drizzle, unless their tops are colder than about −12°C (10 °F) suggests that ice crystals are mainly respon-sible. This view is supported by the fact that the radarsignals from these clouds usually take a characteristicform that has been clearly identified with the melting ofsnowflakes.

Showers, Thunderstorms, and Hail Precipitation from shower clouds and thunderstorms,

whether in the form of raindrops, pellets of soft hail, ortrue hailstones, is generally of great intensity and shorterduration than that from layer clouds and is usually com-posed of larger particles. The clouds are characterized bytheir large vertical depth, strong vertical air currents, and

high concentrations of liquid water, all factors favouringthe rapid growth of precipitation elements by the accre-tion of cloud droplets.

In a cloud composed wholly of liquid water, raindropsmay grow by coalescence. For example, a droplet beingcarried up from the cloud base grows as it ascends bysweeping up smaller droplets. When it becomes too heavy

to be supported by the upcurrents, the droplet falls, con-tinuing to grow by the same process on its downwardjourney. Finally, if the cloud is sufficiently deep, the drop-let will emerge from its base as a raindrop.

In a dense, vigorous cloud several kilometres deep, thedrop may attain its limiting stable diameter (about 6 mm[0.2 inch]) before reaching the cloud base and thus will

break up into several large fragments. Each of these maycontinue to grow and attain breakup size. The number ofraindrops may increase so rapidly in this manner that after

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a few minutes the accumulated mass of water can no longer be supported by the upcurrents and falls as a heavyshower. These conditions occur more readily in tropicalregions. In temperate regions where the freezing level (0

°C) is much lower in elevation, conditions are more favour-able for the ice-crystal mechanism.The hailstones that fall from deep, vigorous clouds

in warm weather consist of a core surrounded by sev-eral alternate layers of clear and opaque ice. When the

growing particle traverses a region of relatively high airtemperature or high concentration of liquid water, or

both, the transfer of heat from the hailstone to the aircannot occur rapidly enough to allow all of the deposited water to freeze immediately. This results in the formation

The structure of a thunderstorm. When the atmosphere becomes unstableenough to form large, powerful updrafts and downdrafts (as indicated bythe red and blue arrows), a towering thundercloud is built up. At times theupdrafts are strong enough to extend the top of the cloud into the tropopause,the boundary between the troposphere (or lowest layer of the atmosphere)

and the stratosphere. Encyclopædia Britannica, Inc.



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of a wet coating of slushy ice, which may later freeze toform a layer of compact, relatively transparent ice. If thehailstone then enters a region of lower temperature andlower water content, the impacting droplets may freezeindividually to produce ice of relatively low density withair spaces between the droplets. The alternate layers areformed as the stone passes through regions in which thecombination of air temperature, liquid-water content, andupdraft speed allows alternately wet and dry growth.

It is held by some authorities that lightning is closely

associated with the appearance of precipitation, especiallyin the form of soft hail, and that the charge and the strongelectric fields are produced by ice crystals or cloud drop-lets striking and bouncing off the undersurfaces of the hailpellets.

Types of Precipitation

Several forms of precipitation exist. Precipitation com-monly reaches Earth’s surface as a liquid; however, it mayalso occur as ice crystals, snowflakes, or ice pellets.

Drizzle

Liquid precipitation in the form of very small drops, withdiameters between 0.2 and 0.5 mm (0.008 and 0.02 inch)

and terminal velocities between 70 and 200 cm per second(28 and 79 inches per second), is defined as drizzle. It formsby the coalescence of even smaller droplets in low-layerclouds containing weak updrafts of only a few centimetresper second. High relative humidity below the cloud base isrequired to prevent the drops from evaporating beforereaching the ground; drizzle is classified as slight, moder-

ate, or thick. Slight drizzle produces negligible runofffrom the roofs of buildings, and thick drizzle accumulatesat a rate in excess of 1 mm per hour (0.04 inch per hour).

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Rain and Freezing Rain

Liquid waterdrops with diameters greater than those ofdrizzle constitute rain. Raindrops rarely exceed 6 mm (0.2inch) in diameter because they become unstable whenlarger than this and break up during their fall. The ter-

minal velocities of raindrops at ground level range from 2metres per second (7 feet per second) for the smallest toabout 10 metres per second (30 feet per second) for thelargest. The smaller raindrops are kept nearly sphericalby surface-tension forces, but, as the diameter surpassesabout 2 mm (0.08 inch), they become increasingly flat-tened by aerodynamic forces. When the diameter reaches

6 mm, the undersurface of the drop becomes concavebecause of the airstream, and the surface of the dropis sheared off to form a rapidly expanding “bubble” or

A rain shaft piercing a tropical sunset as seen from Man-o’-War Bay, Tobago,

Caribbean Sea. NOAA



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“bag” attached to an annular ring containing the bulk ofthe water. Eventually the bag bursts into a spray of finedroplets, and the ring breaks up into a circlet of millime-tre-sized drops.

Rain of a given intensity is composed of a spectrum ofdrop sizes, the average and median drop diameters beinglarger in rains of greater intensity. The largest drops, whichhave a diameter greater than 5 mm (0.2 inch), appear onlyin the heavy rains of intense storms.

When raindrops fall through a cold layer of air (colder

than 0 °C, or 32 °F) and become supercooled, freezing rainoccurs. The drops may freeze on impact with the groundto form a very slippery and dangerous “glazed” ice that isdifficult to see because it is almost transparent.

Snow and Sleet

Snow in the atmosphere can be subdivided into ice crys-

tals and snowflakes. Ice crystals generally form on icenuclei at temperatures appreciably below the freezingpoint. Below –40 °C (–40 °F) water vapour can solidify

without the presence of a nucleus. Snowflakes are aggre- gates of ice crystals that appear in an infinite variety ofshapes, mainly at temperatures near the freezing pointof water.

In British terminology, sleet is the term used todescribe precipitation of snow and rain together or ofsnow melting as it falls. In the United States, it is used todenote partly frozen ice pellets.

Snow crystals generally have a hexagonal pattern, often with beautifully intricate shapes. Three- and 12-branchedforms occur occasionally. The hexagonal form of the

atmospheric ice crystals, their varying size and shape notwithstanding, is an outward manifestation of an internalarrangement in which the oxygen atoms form an open

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Classification of frozen precipitation. Vincent J. Schaefer



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lattice (network) with hexagonally symmetrical structure.According to a recent internationally accepted classifica-tion, there are seven types of snow crystals: plates, stellars,columns, needles, spatial dendrites, capped columns, andirregular crystals. The size and shape of the snow crystalsdepend mainly on the temperature of their formation andon the amount of water vapour that is available for depo-sition. The two principal influences are not independent;the possible water vapour admixture of the air decreasesstrongly with decreasing temperature. The vapour pres-

sure in equilibrium, or state of balance, with a level surfaceof pure ice is 50 times greater at –2 °C (28 °F) than at –42°C (–44 °F), the likely limits of snow formation in the air.

At temperatures above about –40 °C (–40 °F), the crys-tals form on nuclei of very small size that float in the air(heterogeneous nucleation). The nuclei consist predomi-nantly of silicate minerals of terrestrial origin, mainly clay

minerals and micas. At still lower temperatures, ice mayform directly from water vapour (homogeneous nucle-ation). The influence of the atmospheric water vapourdepends mainly on its degree of supersaturation withrespect to ice.

If the air contains a large excess of water vapour, thesnow particles will grow fast, and there may be a tendency

for dendritic (branching) growth. With low temperature,the excess water vapour tends to be small, and the crystalsremain small. In relatively dry layers, the snow particles

generally have simple forms. Complicated forms of crys-tals will cling together with others to form snowflakes thatconsist occasionally of up to 100 crystals; the diameter ofsuch flakes may be as large as 2.5 cm (1 inch). This process

will be furthered if the crystals are near the freezing pointand wet, possibly by collision with undercooled waterdroplets. If a crystal falls into a cloud with great numbers

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of such drops, it will sweep some of them up. Coming intocontact with ice, they freeze and form an ice cover aroundthe crystal. Such particles are called soft hail or graupel.

Snow particles constitute the clouds of cirrus type—namely cirrus, cirrostratus, and cirrocumulus—andmany clouds of alto type. Ice and snow clouds originatenormally only at temperatures some degrees below thefreezing point; they predominate at –20 °C (–4 °F). Intemperate and low latitudes these clouds occur in thehigher layers of the troposphere. In tropical regions

they hardly ever occur below 4,570 metres (15,000 feet).On high mountains and particularly in polar regions,they can occur near the surface and may appear as icefogs. If cold air near the ground is overlain by warmer air(a very common occurrence in polar regions, especiallyin winter), mixture at the border leads to supersatu-ration in the cold air. Small ice columns and needles,

“diamond dust,” will be formed and will float down, glit-tering, even from a cloudless sky. In the coldest partsof Antarctica, where temperatures near the surface arebelow −50 °C (–58 °F) on the average and rarely above −30°C (–22 °F), the formation of diamond dust is a commonoccurrence. The floating and falling ice crystals pro-duce in the light of the Sun and the Moon the manifold

phenomena of atmospheric optics, halos, arcs, circles,mock suns, some coronas, and iridescent clouds. Mostof the different optical appearances can be explained bythe shapes of the crystals and their position with respectto the light source.

Most of the moderate to heavy rain in temperate lati-tudes depends on the presence of ice and snow particles

in clouds. In the free atmosphere, droplets of fluid watercan be undercooled considerably; typical ice clouds origi-nate mainly at a temperature near –20 °C. At an identical



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temperature below the freezing point, the water mol-ecules are kept more firmly in the solid than in the fluidstate. The equilibrium pressure of the gaseous phase issmaller in contact with ice than with water. At –20 °C,

which is the temperature of the formation of typical iceclouds (cirrus), the equilibrium pressure with respectto undercooled water (relative humidity 100 percent) is22 percent greater than the equilibrium pressure of the

water vapour in contact with ice. Hence, with an excessof water vapour beyond the equilibrium state, the ice par-

ticles tend to incorporate more water vapour and to growmore rapidly than the water droplets.

Being larger and so less retarded by friction, the iceparticles fall more rapidly. In their fall they sweep up some

water droplets, which on contact become frozen. Thus, acloud layer originally consisting mainly of undercooled

water with few ice crystals is transformed into an ice cloud.

The development of the anvil shape at the top of a tower-ing cumulonimbus cloud shows this transformation veryclearly. The larger ice particles overcome more readily therising tendency of the air in the cloud. Falling into lowerlevels they grow, aggregating with other crystals and pos-sibly with waterdrops, melt, and form raindrops whennear-surface temperatures permit.

Hail

Solid precipitation in the form of hard pellets of ice thatfall from cumulonimbus clouds is called hail. It is conve-nient to distinguish between three types of hail particles.

The first is soft hail, or snow pellets, which are whiteopaque rounded or conical pellets as large as 6 mm (0.2

inch) in diameter. They are composed of small cloud drop-lets frozen together, have a low density, and are readilycrushed.

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The second is small hail (ice grains or pellets), whichare transparent or translucent pellets of ice that arespherical, spheroidal, conical, or irregular in shape, withdiameters of a few millimetres. They may consist of frozenraindrops, of largely melted and refrozen snowflakes, or ofsnow pellets encased in a thin layer of solid ice.

True hailstones, the third type, are hard pellets ofice, larger than 5 mm (0.2 inch) in diameter, that maybe spherical, spheroidal, conical, discoidal, or irregularin shape and often have a structure of concentric layers

of alternately clear and opaque ice. A moderately severestorm may produce stones a few centimetres in diameter,

whereas a very severe storm may release stones with amaximum diameter of 10 cm (4 inches) or more. Largedamaging hail falls most frequently in the continentalareas of middle latitudes (e.g., in the Nebraska-Wyoming-Colorado area of the United States, in South Africa,

and in northern India) but is rare in equatorial regions.Terminal velocities of hailstones range from about 5metres (16 feet) per second for the smallest stones to per-haps 40 metres (130 feet) per second for stones 5 cm (2inches) in diameter.

World Distribution of Precipitation

Precipitation does not fall uniformly across Earth’s sur-face. Some areas receive tremendous amounts, whereasothers receive much less. The amount and regularity ofprecipitation at a given location depends greatly on thelatitude, local topographic features, and prevailing windpatterns of that location.

Regional and Latitudinal Distribution

The yearly precipitation averaged over the whole Earth isabout 100 cm (39 inches), but this is distributed very



51

unevenly. The regions of highest rainfall are found in theequatorial zone and the monsoon area of Southeast Asia.Middle latitudes receive moderate amounts of precipita-tion, but little falls in the desert regions of the subtropicsand around the poles.

If Earth’s surface were perfectly uniform, the long-term average rainfall would be distributed in distinctlatitudinal bands, but the situation is complicated bythe pattern of the global winds, the distribution of landand sea, and the presence of mountains. Because rainfallresults from the ascent and cooling of moist air, the areasof heavy rain indicate regions of rising air, whereas the des-

erts occur in regions in which the air is warmed and driedduring descent. In the subtropics, the trade winds bringplentiful rain to the east coasts of the continents, but the

Global distribution of mean annual rainfall (in centimetres). EncyclopaediaBritannica, Inc.

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west coasts tend to be dry. On the other hand, in high lati-tudes the west coasts are generally wetter than the eastcoasts. Rain tends to be abundant on the windward slopesof mountain ranges but sparse on the lee sides.

In the equatorial belt, the trade winds from both hemi-spheres converge and give rise to a general upward motionof air, which becomes intensified locally in tropical stormsthat produce very heavy rains in the Caribbean, the Indianand southwest Pacific oceans, and the China Sea and inthunderstorms that are especially frequent and active over

the land areas. During the annual cycle, the doldrumsmove toward the summer hemisphere, so outside a centralregion near the Equator, which has abundant rain at allseasons, there is a zone that receives much rain in summerbut a good deal less in winter.

The dry areas of the subtropics—such as the desertregions of North Africa, the Arabian Peninsula, South

Africa, Australia, and central South America—are due tothe presence of semipermanent subtropical anticyclonesin which the air subsides and becomes warm and dry.These high-pressure belts tend to migrate with the sea-sons and cause summer dryness on the poleward side and

winter dryness on the equatorward side of their meanpositions. The easterly trade winds, having made a long

passage over the warm oceans, bring plentiful rains to theeast coasts of the subtropical landmasses, but the westcoasts and the interiors of the continents, which are oftensheltered by mountain ranges, are very dry.

In middle latitudes, weather and rainfall are domi-nated by traveling depressions and fronts that yield a gooddeal of rain in all seasons and in most places except the far

interiors of the Asian and North American continents.Generally, rainfall is more abundant in summer, except onthe western coasts of North America, Europe, and NorthAfrica, where it is higher during the winter.



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At high latitudes and especially in the polar regions,the low precipitation is caused partly by subsidence of airin the high-pressure belts and partly by the low tempera-tures. Snow or rain occur at times, but evaporation fromthe cold sea and land surfaces is slow, and the cold air haslittle capacity for moisture.

The influence of oceans and continents on rainfall isparticularly striking in the case of the Indian monsoon.During the Northern Hemisphere winter, cool dry airfrom the interior of the continent flows southward and

produces little rain over the land areas. After the air hastraveled some distance over the warm tropical ocean,however, it releases heavy shower rains over the EastIndies. During the northern summer, when the monsoonblows from the southwest, rainfall is heavy over India andSoutheast Asia. These rains are intensified where the air isforced to ascend over the windward slopes of the Western

Ghats and the Himalayas.The combined effects of land, sea, mountains, and

prevailing winds show up in South America. There thedesert in southern Argentina is sheltered by the Andesfrom the westerly winds blowing from the Pacific Ocean,and the west-coast desert not only is situated under theSouth Pacific subtropical anticyclone but is also pro-

tected by the Andes against rain-bearing winds from theAtlantic.

Amounts and Variability

The long-term average amounts of precipitation for a sea-son or a year give little information on the regularity with

which rain may be expected, particularly for regions where

the average amounts are small. For example, at Iquique,a city in northern Chile, four years once passed withoutrain, whereas the fifth year gave 15 mm (0.6 inch); the five-

year average was therefore 3 mm (0.1 inch). Clearly, such

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averages are of little practical value, and the frequencydistribution or the variability of precipitation also mustbe known.

The variability of the annual rainfall is closely relatedto the average amounts. For example, over the BritishIsles, which have a very dependable rainfall, the annualamount varies by less than 10 percent above the long-termaverage value. A variability of less than 15 percent is typicalof the mid-latitude cyclonic belts of the Pacific andAtlantic oceans and of much of the wet equatorial regions.

In the interiors of the desert areas of Africa, Arabia, andCentral Asia, however, the rainfall in a particular year maydeviate from the normal long-term average by more than40 percent. The variability for individual seasons ormonths may differ considerably from that for the year as a

whole, but again the variability tends to be higher wherethe average amounts are low.

The heaviest annual rainfall in the world was recordedat the village of Cherrapunji, India, where 26,470 mm(1,042 inches) fell between August 1860 and July 1861. Theheaviest rainfall in a period of 24 hours was 1,870 mm (74inches), recorded at the village of Cilaos, Réunion, in theIndian Ocean on March 15–16, 1952. The lowest recordedrainfall in the world occurred at Arica, a port city in north-

ern Chile. An annual average, taken over a 43-year period, was only 0.5 mm (0.02 inch).Although past records give some guide, it is not possi-

ble to estimate very precisely the maximum possibleprecipitation that may fall in a given locality during a spec-ified interval of time. Much will depend on a favourablecombination of several factors, including the properties of

the storm and the effects of local topography. Thus, it ispossible only to make estimates that are based on analysesof past storms or on theoretical calculations that attempt



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to maximize statistically the various factors or the mosteffective combination of factors that are known to con-trol the duration and intensity of the precipitation. Formany important planning and design problems, however,estimates of the greatest precipitation to be expected at a

given location within a specified number of years arerequired.

In the designing of a dam, the highest 24-hour rainfallto be expected once in 30 years over the whole catchmentarea might be relevant. For dealing with such problems, a

great deal of work has been devoted to determining frompast records the frequency with which rainfalls of givenintensity and total amount may be expected to reoccur atparticular locations and also to determining the statisticsof rainfall for a specific area from measurements made atonly a few points.

The Effects of PrecipitationRain, snow, and other forms of precipitation may alter theland upon which it falls. The sheer force of falling rain maycompact or erode soils. In addition, precipitation thatneither evaporates nor becomes absorbed by the soil runsoff into surface waters.

Raindrop Impact and Soil ErosionLarge raindrops, up to 6 mm (0.2 inch) in diameter, haveterminal velocities of about 10 metres (30 feet) per secondand so may cause considerable compaction and erosion ofthe soil by their force of impact. The formation of a com-pacted crust makes it more difficult for air and water toreach the roots of plants and encourages the water to run

off the surface and carry away the topsoil with it. In hillyand mountainous areas, heavy rain may turn the soil intomud and slurry, which may produce enormous erosion by

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mudflow generation. Rainwater running off hard impervi-ous surfaces or waterlogged soil may cause local flooding.

Surface Runoff

The rainwater that is not evaporated or stored in the soileventually runs off the surface and finds its way into rivers,streams, and lakes or percolates through the rocks andbecomes stored in natural underground reservoirs. A givencatchment area must achieve an overall balance such thatprecipitation ( P ) less evaporation of moisture from the

surface ( E ) will equal storage in the ground ( S ) and runoff( R ). This may be expressed: P − E = S + R. The runoff maybe determined by measuring the flow of water in the rivers

with stream gauges, and the precipitation may be mea-sured by a network of rain gauges, but storage andevaporation are more difficult to estimate.

Of all the water that falls on Earth’s surface, the rel-

ative amounts that run off, evaporate, or seep into the ground vary so much for different areas that no firmfigures can be given for Earth as a whole. It has beenestimated, however, that in the United States 10 to 50percent of the rainfall runs off at once, 10 to 30 percentevaporates, and 40 to 60 percent is absorbed by thesoil. Of the entire rainfall, 15 to 30 percent is estimated

to be used by plants, either to form plant tissue or intranspiration.

ATMOSPHERIC PRESSURE AND WIND

Atmospheric pressure and wind are both significant con-trolling factors of Earth’s weather and climate. Although

these two physical variables may at first glance appear tobe quite different, they are in fact closely related. Windexists because of horizontal and vertical differences



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(gradients) in pressure, yielding a correspondence thatoften makes it possible to use the pressure distributionas an alternative representation of atmospheric motions.Pressure is the force exerted on a unit area, and atmo-spheric pressure is equivalent to the weight of air abovea given area on Earth’s surface or within its atmosphere.This pressure is usually expressed in millibars (mb; 1 mbequals 1,000 dynes per square cm) or in kilopascals (kPa; 1kPa equals 10,000 dynes per square cm). Distributions ofpressure on a map are depicted by a series of curved lines

called isobars, each of which connects points of equalpressure.

Atmospheric Pressure

At sea level the mean pressure is about 1,000 mb (100kPa), varying by less than 5 percent from this value at

any given location or time. Since charts of atmosphericpressure often represent average values over several days,pressure features that are relatively consistent day afterday emerge, while more transient, short-lived featuresare removed. Those that remain are known as semi-permanent pressure centres and are the source regionsfor major, relatively uniform bodies of air known as air

masses. Warm, moist maritime tropical (mT) air formsover tropical and subtropical ocean waters in association with the high-pressure regions prominent there. Cool,moist maritime polar (mP) air, on the other hand, formsover the colder subpolar ocean waters just south and eastof the large, winter oceanic low-pressure regions. Overthe continents, cold dry continental polar (cP) air and

extremely cold dry continental arctic (cA) air forms in thehigh-pressure regions that are especially pronounced in winter, while hot dry continental tropical (cT) air forms

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over hot desertlike continental domains in summer inassociation with low-pressure areas, which are sometimescalled heat lows.

Sea-level pressure is dominated by closed high- andlow-pressure centres, which are largely caused by differ-ential surface heating between low and high latitudes andbetween continental and oceanic regions. High pressuretends to be amplified over the colder surface features.And because of seasonal changes in surface heating, thepressure centres exhibit seasonal changes in their charac-

teristics. For example, the Siberian High, Aleutian Low,and Icelandic Low that are so prominent in the winter

virtually disappear in summer as the continental regions warm relative to surrounding bodies of water. At the sametime, the Pacific and Atlantic highs amplify and migratenorthward.

At altitudes well above Earth’s surface, the monthly

average pressure distributions show much less tendencyto form in closed centres but rather appear as quasi-con-centric circles around the poles. This more symmetricalappearance reflects the dominant role of meridional(north-south) differences in radiative heating and cooling.Excess heating in tropical latitudes, in contrast to polarareas, produces higher pressure at upper levels in the trop-

ics as thunderstorms transfer air to higher levels. Inaddition, the greater heating/cooling contrast in winter yields stronger pressure differences during this season.Perfect symmetry between the tropics and the poles isinterrupted by wavelike atmospheric disturbances associ-ated with migratory and semipermanent high- andlow-pressure surface weather systems. These weather sys-

tems are most pronounced over the Northern Hemisphere, with its more prominent land-ocean contrasts and orographic (high-elevation) features.



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Wind

Winds occur as the result of changes in atmospheric pres-sure. Other factors, such as the Coriolis force and frictionat Earth’s surface, can alter both the speed and directionof winds. Winds may appear as cyclones and anticyclones,and local and zonal phenomena.

The Relationship of Wind to Pressure and

Governing Forces

The changing wind patterns are governed by IsaacNewton’s second law of motion, which states that thesum of the forces acting on a body equals the productof the mass of that body and the acceleration caused bythose forces. The basic relationship between atmosphericpressure and horizontal wind is revealed by disregardingfriction and any changes in wind direction and speed to

yield the mathematical relationship

where u is the zonal wind speed (+ eastward), v the merid-ional wind speed (+ northward), f = 2ω sin ϕ (Coriolis

parameter), ω the angular velocity of Earth’s rotation,ϕ the latitude, ρ the air density (mass per unit volume), p the pressure, and x and y the distances toward theeast and north, respectively. This simple non-acceler-ating flow is known as geostrophic balance and yields amotion field known as the geostrophic wind. Equation(1) expresses, for both the x and y directions, a balance

between the force created by horizontal differences inpressure (the horizontal pressure-gradient force) and anapparent force that results from Earth’s rotation (the

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Coriolis force). The pressure-gradient force expresses thetendency of pressure differences to effectuate air move-ment from higher to lower pressure. The Coriolis forcearises because the air motions are observed on a rotatingnearly spherical body. The total motion of a parcel of airhas two parts: (1) the motion relative to Earth as if theplanet were fixed, and (2) the motion given to the par-cel of air by the planet’s rotation. When the atmosphereis viewed from a fixed point in space, Earth’s rotation isapparent. An observer in space would witness the total

motion of the atmosphere. Conversely, an observer onthe ground sees and measures only the relative motionof the atmosphere, because he is also rotating and can-not see directly the rotational motion applied by Earth.Instead, the observer on the ground sees the effect of therotation as a deviation applied to the relative motion.The quantity that describes this deviation is the Coriolis

force. Because the Coriolis force results from a ground-level frame of reference on a rotating planet, it is not atrue force.

More specifically, the observer on the ground experi-ences the Coriolis force as a deflection of the relativemotion to the right in the Northern Hemisphere and tothe left in the Southern Hemisphere. Of particular signifi-

cance in this simple model of wind-pressure relationshipsis the fact that the geostrophic wind blows in a directionparallel to the isobars, with the low pressure on the observ-er’s left as he looks downwind in the Northern Hemisphereand on his right in the Southern Hemisphere.

Wind speed increases as the distance between isobarsdecreases (or pressure gradient increases). Curvature (i.e.,

changes in wind direction) can be added to this model withrelative ease in a flow representation known as the gradi-ent wind. The basic wind-pressure relationships, however,remain qualitatively the same. Of greatest importance is



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the fact that large-scale, observed winds tend to behavemuch as the geostrophic- or gradient-flow models predictin most of the atmosphere. The most notable exceptionsoccur in low latitudes, where the Coriolis parameterbecomes very small—equation (1) cannot be used to pro-

vide a reliable wind estimate—and in the lowest kilometreof the atmosphere, where friction becomes important.The friction induced by airflow over the underlying sur-face reduces the wind speed and alters the simple balanceof forces such that the wind blows with a component

toward lower pressure.

Cyclones and Anticyclones

Cyclones and anticyclones are regions of relatively lowand high pressure, respectively. They occur over most of

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A top view and vertical cross section of a tropical cyclone. EncyclopædiaBritannica, Inc.



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Earth’s surface in a variety of sizes ranging from the verylarge semipermanent examples described above to smaller,highly mobile systems. The latter are the focus of discus-sion in this section.

Common to both cyclones and anticyclones are thecharacteristic circulation patterns. The geostrophic-

wind and gradient-wind models dictate that, in theNorthern Hemisphere, flow around a cyclone—cycloniccirculation—is counterclockwise, and flow around ananticyclone—anticyclonic circulation—is clockwise.

Circulation directions are reversed in the SouthernHemisphere. In the presence of friction, the superimposedcomponent of motion toward lower pressure produces a“spiraling” effect toward the low-pressure centre and awayfrom the high-pressure centre.

The cyclones that form outside the equatorial belt,known as extratropical cyclones, may be regarded as large

eddies in the broad air currents that flow in the generaldirection from west to east around the middle and higherlatitudes of both hemispheres. They are an essential partof the mechanism by which the excess heat received fromthe Sun in Earth’s equatorial belt is conveyed towardhigher latitudes. These higher latitudes radiate more heatto space than they receive from the Sun, and heat must

reach them by winds from the lower latitudes if their tem-perature is to be continually cool rather than cold. If there were no cyclones and anticyclones, the north-south move-ments of the air would be much more limited, and there

would be little opportunity for heat to be carried polewardby winds of subtropical origin. Under such circumstancesthe temperature of the lower latitudes would increase,

and the polar regions would cool; the temperature gradi-ent between them would intensify.Strong horizontal gradients of temperature are par-

ticularly favourable for the formation and development of



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cyclones. The temperature difference between polarregions and the Equator builds up until it becomes suffi-ciently intense to generate new cyclones. As theirassociated cold fronts sweep equatorward and their warmfronts move poleward, the new cyclones reduce the tem-perature difference. Thus, the wind circulation on Earthrepresents a balance between the heating effects of solarradiation occurring in the polar regions and at the Equator.Wind circulation, through the effect of cyclones, anticy-clones, and other wind systems, also periodically destroys

this temperature contrast.Cyclones of a somewhat different character occur

closer to the Equator, generally forming in latitudesbetween 10° to 30° N and S over the oceans. They gener-ally are known as tropical cyclones when their winds equalor exceed 74 miles (119 km) per hour. They are also knownas hurricanes if they occur in the Atlantic Ocean and the

Caribbean Sea, as typhoons in the western Pacific Oceanand the China Sea, and as cyclones off the coasts ofAustralia. These storms are of smaller diameter than theextratropical cyclones, ranging from 100 to 500 km (60 to300 miles) in diameter, and are accompanied by winds thatsometimes reach extreme violence. These storms are morefully described the last chapter of this book, in the section

on tropical cyclones. Extratropical Cyclones

Of the two types of large-scale cyclones, extratropicalcyclones are the most abundant and exert influence on thebroadest scale; they affect the largest percentage of Earth’ssurface. Furthermore, this class of cyclones is the princi-

pal cause of day-to-day weather changes experienced inmiddle and high latitudes and thus is the focal point ofmuch of modern weather forecasting. The seeds for manycurrent ideas concerning extratropical cyclones were

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sown between 1912 and 1930 by a group of Scandinavianmeteorologists working in Bergen, Nor. This so-calledBergen school, founded by Norwegian meteorologist andphysicist Vilhelm Bjerknes, formulated a model for acyclone that forms as a disturbance along a zone of strongtemperature contrast known as a front, which in turn con-stitutes a boundary between two contrasting air masses.In this model the masses of polar and mid-latitude airaround the globe are separated by the polar front (thetransition region separating warmer tropical air from

colder polar air). This region possesses a strong tempera-ture gradient, and thus it is a reservoir of potential energythat can be readily tapped and converted into the kineticenergy associated with extratropical cyclones.

For this reservoir to be tapped, a cyclone (called a wave, or frontal, cyclone) must develop much in the way shown in the diagram on page 65. The feature that

is of primary importance prior to cyclone development(cyclogenesis) is a front, represented in the initial stage(A) as a heavy black line with alternating triangles orsemicircles attached to it. This stationary or very slow-moving front forms a boundary between cold and warmair and thus is a zone of strong horizontal temperature

gradient (sometimes referred to as a baroclinic zone).

Cyclone development is initiated as a disturbance alongthe front, which distorts the front into the wavelike con-figuration (B; wave appearance). As the pressure withinthe disturbance continues to decrease, the disturbanceassumes the appearance of a cyclone and forces pole-

ward and equatorward movements of warm and coldair, respectively, which are represented by mobile fron-

tal boundaries. As depicted in the cyclonic circulationstage (C), the front that signals the advancing cold air(cold front) is indicated by the triangles, while the front



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Evolution of a wave (frontal) cyclone. Encyclopædia Britannica, Inc.

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Cross section of clouds and precipitation. The direction of frontal movement isindicated by the arrows.

corresponding to the advancing warm air (warm front) isindicated by the semicircles. As the cyclone continues tointensify, the cold dense air streams rapidly equatorward,

yielding a cold front with a typical slope of 1 to 50 and apropagation speed that is often 8 to 15 metres per second(about 18 to 34 miles per hour) or more. At the same time,the warm less-dense air moving in a northerly directionflows up over the cold air east of the cyclone to produce a

warm front with a typical slope of 1 to 200 and a typicallymuch slower propagation speed of about 2.5 to 8 metres

per second (6 to 18 miles per hour). This difference inpropagation speeds between the two fronts allows thecold front to overtake the warm front and produce yetanother, more complicated frontal structure, known asan occluded front. An occluded front (D) is representedby a line with alternating triangles and semicircles on thesame side. This occlusion process may be followed by fur-

ther storm intensification. The separation of the cyclonefrom the warm air toward the Equator, however, even-tually leads to the storm’s decay and dissipation (E) in aprocess called cyclolysis.



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The life cycle of such an event is typically several days,during which the cyclone may travel from several hundredto a few thousand kilometres. In its path and wake occurdramatic weather changes. Warm frontal weather is mostfrequently characterized by stratiform clouds, whichascend as the front approaches and potentially yield rainor snow. The passing of a warm front brings a rise in airtemperature and clearing skies. The warmer air, however,may also harbour the ingredients for rain shower or thun-derstorm formation, a condition that is enhanced as the

cold front approaches.The passage of the cold front is marked by the influx

of colder air, the formation of stratocumulus clouds withsome lingering rain or snow showers, and then even-tual clearing. While this is an oft-repeated scenario,

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Satellite image of a large dust storm in the Takla Makan Desert, northwesternChina. MODIS Rapid Response Team/NASA/GFSC



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it is important to recognize that many other weathersequences can also occur. For example, the stratiformclouds of a warm front may have imbedded cumulus for-mations and thunderstorms; the warm sector might bequite dry and yield few or no clouds; the pre-cold-front

weather may closely resemble that found ahead of the warm front; or the post-cold-front air may be completelycloud-free. Cloud patterns oriented along fronts and spi-raling around the cyclone vortex are consistently revealedin satellite pictures of Earth.

The actual formation of any area of low pressurerequires that mass in the column of air lying above Earth’ssurface be reduced. This loss of mass then reduces the sur-face pressure. In the late 1930s and early ’40s, threemembers of the Bergen school—Norwegian Americanmeteorologists Jacob Bjerknes and Jørgen Holmboe andSwedish American meteorologist Carl-Gustaf Rossby—

recognized that transient surface disturbances wereaccompanied by complementary wave features in the flowin the middle and higher atmospheric layers associated

with the jet stream. These wave features are accompaniedby regions of mass divergence and convergence that sup-port the growth of surface-pressure fields and direct theirmovement.

While extratropical cyclones form and intensify inassociation with fronts, there are small-scale cyclones thatappear in the middle of a single air mass. A notable exam-ple is a class of cyclones, generally smaller than the frontal

variety, that form in polar air streams in the wake of a fron-tal cyclone. These so-called polar lows are most prominentin subpolar marine environments and are thought to be

caused by the transfer of heat and moisture from the warmer water surface into the overlying polar air and bysupporting middle-tropospheric circulation features.



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Other small-scale cyclones form on the lee side of moun-tain barriers as the general westerly flow is disturbed bythe mountain. These “lee cyclones” may produce major

windstorms and dust storms downstream of a mountainbarrier.

Anticyclones

While cyclones are typically regions of inclement weather,anticyclones are usually meteorologically quiet regions.Generally larger than cyclones, anticyclones exhibit per-

sistent downward motions and yield dry stable air thatmay extend horizontally many hundreds of kilometres.

In most cases, an actively developing anticycloneforms over a ground location in the region of cold airbehind a cyclone as it moves away. This anticyclone formsbefore the next cyclone advances into the area. Such ananticyclone is known as a cold anticyclone. A result of

the downward air motion in an anticyclone, however, iscompression of the descending air. As a consequence ofthis compression, the air is warmed. Thus, after a fewdays, the air composing the anticyclone at levels 2 to 5 km(1 to 3 miles) above the ground tends to increase in tem-perature, and the anticyclone is transformed into a warmanticyclone.

Warm anticyclones move slowly, and cyclones arediverted around their periphery. During their transforma-tion from cold to warm status, anticyclones usually moveout of the main belt followed by cyclones in middle latitudesand often amalgamate with the quasi-permanent bands ofrelatively high pressure found in both hemispheres aroundlatitude 20° to 30°—the so-called subtropical anticy-

clones. On some occasions the warm anticyclones remainin the belt normally occupied by the mid-latitude westerly winds. The normal cyclone tracks are then considerably

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modified; atmospheric depressions (areas of low pressure)are either blocked in their eastward progress or divertedto the north or south of the anticyclone. Anticyclonesthat interrupt the normal circulation of the westerly windin this way are called blocking anticyclones, or blockinghighs. They frequently persist for a week or more, and theoccurrence of a few such blocking anticyclones may domi-nate the character of a season. Blocking anticyclones areparticularly common over Europe, the eastern Atlantic,and the Alaskan area.

The descent and warming of the air in an anticyclonemight be expected to lead to the dissolution of cloudsand the absence of rain. Near the centre of the anticy-clone, the winds are light and the air can become stagnant.Air pollution can build up as a result. The city of LosAngeles, for example, often has poor air quality because itis frequently under a stationary anticyclone. In winter the

ground cools, and the lower layers of the atmosphere alsobecome cold. Fog may be formed as the air is cooled to itsdew point in the stagnant air. Under other circumstances,the air trapped in the first kilometre above Earth’s surfacemay pick up moisture from the sea or other moist sur-faces, and layers of cloud may form in areas near the

ground up to a height of about 1 km (0.6 mile). Such layers

of cloud can be persistent in anticyclones (except over thecontinents in summer), but they rarely grow thick enoughto produce rain. If precipitation occurs, it is usually drizzleor light snow.

Anticyclones are often regions of clear skies and sunny weather in summer; at other times of the year, cloudy andfoggy weather—especially over wet ground, snow cover,

and the ocean—may be more typical. Winter anticyclonesproduce colder than average temperatures at the surface,particularly if the skies remain clear. Anticyclones are



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responsible for periods of little or no rain, and such peri-ods may be prolonged in association with blocking highs.

Cyclone and Anticyclone Climatology

Migrating cyclones and anticyclones tend to be distrib-uted around certain preferred regions, known as tracks,that emanate from preferred cyclogenetic and anticyclo-

genetic regions. Favoured cyclogenetic regions in theNorthern Hemisphere are found on the lee side of moun-tains and off the east coasts of continents. Cyclones then

track east or southeast before eventually turning towardthe northeast and decaying. The tracks are displaced far-ther northward in July, reflecting the more northwardposition of the polar front in summer. Continentalcyclones usually intensify at a rate of 0.5 mb (0.05 kPa) perhour or less, although more dramatic examples can befound. Marine cyclones, on the other hand, often experi-

ence explosive development in excess of 1 mb (0.1 kPa) perhour, particularly in winter.

Anticyclones tend to migrate equatorward out of thecold air mass regions and then eastward before decay-ing or merging with a warm anticyclone. Like cyclones,

warm anticyclones also slowly migrate poleward withthe warm season.

In the Southern Hemisphere, where most of Earth’ssurface is covered by oceans, the cyclones are distributedfairly uniformly through the various longitudes. Typically,cyclones form initially in latitudes 30° to 40° S and movein a generally southeastward direction, reaching matu-rity in latitudes near 60° S. Thus, the Antarctic continentis usually ringed by a number of mature or decaying

cyclones. The belt of ocean from 40° to 60° S is a regionof persistent, strong westerly winds that form part of thecirculation to the north of the main cyclone centres;

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these are the “roaring forties,” where the westerly windsare interrupted only at intervals by the passage south-eastward of developing cyclones.

Local Winds

Organized wind systems occur in spatial dimensions rang-ing from tens of metres to thousands of kilometres andpossess residence times that vary from seconds to weeks.The concept of scale considers the typical size and life-time of a phenomenon. Since the atmosphere exhibits

such a large variety of both spatial and temporal scales,efforts have been made to group various phenomena intoscale classes. The class describing the largest and longest-lived of these phenomena is known as the planetary scale.Such phenomena are typically a few thousand kilometresin size and have lifetimes ranging from several days to sev-eral weeks. Examples of planetary-scale phenomena

include the semipermanent pressure centres discussedearlier and certain globe-encircling upper-air waves.

A second class is known as the synoptic scale. Spanningsmaller distances, a few hundred to a few thousand kilo-metres, and possessing shorter lifetimes, a few to severaldays, this class contains the migrating cyclones andanticyclones that control day-to-day weather changes.

Sometimes the planetary and synoptic scales are com-bined into a single classification termed the large-scale,or macroscale. Large-scale wind systems are distinguishedby the predominance of horizontal motions over verticalmotions and by the preeminent importance of the Coriolisforce in influencing wind characteristics. Examples oflarge-scale wind systems include the trade winds and the

westerlies.There is a third class of phenomena of even smallersize and shorter lifetime. In this class, vertical motions



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Typical sea-breeze (afternoon) and land-breeze (night) circulations with asso-ciated cloud formations.

may be as significant as horizontal movement, and theCoriolis force often plays a less important role. Known asthe mesoscale, this class is characterized by spatial dimen-sions of ten to a few hundred kilometres and lifetimes of aday or less. Because of the shorter time scale and because

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the other forces may be much larger, the effect of theCoriolis force in mesoscale phenomena is sometimesneglected.

Two of the best-known examples of mesoscalephenomena are the thunderstorm and its devastating by-product, the tornado. The present discussion focuses onless intense, though nevertheless commonly observed,

wind systems that are found in rather specific geographiclocations and thus are often referred to as local windsystems.

The so-called sea and land breeze circulation is a local wind system typically encountered along coastlines adja-cent to large bodies of water and is induced by differencesthat occur between the heating or cooling of the watersurface and the adjacent land surface. Water has a higherheat capacity (i.e., more units of heat are required to pro-duce a given temperature change in a volume of water)

than do the materials in the land surface. Daytime solarradiation penetrates to several metres into the water, the

water vertically mixes, and the volume is slowly heated. Incontrast, daytime solar radiation heats the land surfacemore quickly because it does not penetrate more than afew centimetres below the land surface. The land surface,now at a higher temperature relative to the air adjacent to

it, transfers more heat to its overlying air mass and createsan area of low pressure. It should be noted that the surfaceflow is from the water toward the land and thus is called asea breeze.

Since the landmass possesses a lower heat capacitythan water, the land cools more rapidly at night than doesthe water. Consequently, at night the cooler landmass

yields a cooler overlying air mass and creates a zone of rel-atively higher pressure. This produces a circulation cell with air motions opposite to those found during the day.



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When the valley floor warms during the day, warm air rises up the slopes of

surrounding mountains and hills to create a valley breeze. At night, denser cool air slides down the slopes to settle in the valley, producing a mountain breeze.Encyclopædia Britannica, Inc.

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This flow from land to water is known as a land breeze.The land breeze is typically shallower than the sea breezesince the cooling of the atmosphere over land is confinedto a shallower layer at night than the heating of the air dur-ing the day.

Sea and land breezes occur along the coastal regions ofoceans or large lakes in the absence of a strong large-scale

wind system during periods of strong daytime heating ornighttime cooling. Those who live within 10 to 20 km (6 to12 miles) of the coastline often experience the cooler 19-

to 37-km-per-hour (12- to 23-mile-per-hour) winds of thesea breeze on a sunny afternoon only to find it turn into asultry land breeze late at night. One of the features of thesea and land breeze is a region of low-level air convergencein the termination region of the surface flow. Such conver-

gence often induces local upward motions and cloudformations. Thus, in sea and land breeze regions, it is not

uncommon to see clouds lying off the coast at night; theseclouds are then dissipated by the daytime sea breeze,

which forms new clouds, perhaps with showers occurringover land in the afternoon.

Another group of local winds is induced by the pres-ence of mountain and valley features on Earth’s surface.One subset of such winds, known as mountain winds or

breezes, is induced by differential heating or coolingalong mountain slopes. During the day, solar heating ofthe sunlit slopes causes the overlying air to move upslope.These winds are also called anabatic flow. At night, as theslopes cool, the direction of airflow is reversed, and cooldownslope drainage motion occurs. Such winds may berelatively gentle or may occur in strong gusts, depend-

ing on the topographic configuration. These winds areone type of katabatic flow. In an enclosed valley, the coolair that drains into the valley may give rise to a thick fog



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condition. Fog persists until daytime heating reverses thecirculation and creates clouds associated with the upslopemotion at the mountain top.

Another subset of katabatic flow, called foehn winds(also known as chinook winds east of the Rocky Mountainsand as Santa Ana winds in southern California), is inducedby adiabatic temperature changes occurring as air flowsover a mountain. Adiabatic temperature changes are thosethat occur without the addition or subtraction of heat;they occur in the atmosphere when bundles of air are

moved vertically. When air is lifted, it enters a region oflower pressure and expands. This expansion is accompa-nied by a reduction of temperature (adiabatic cooling).When air subsides, it contracts and experiences adiabatic

warming. As air ascends on the windward side of themountain, its cooling rate may be moderated by heat thatis released during the formation of precipitation. However,

having lost much of its moisture, the descending air on theleeward side of the mountain adiabatically warms fasterthan it was cooled on the windward ascent. Thus, theeffect of this wind, if it reaches the surface, is to produce

warm, dry conditions. Usually, such winds are gentle andproduce a slow warming. On occasion, however, foehn

winds may exceed 185 km (115 miles) per hour and produce

air-temperature increases of tens of degrees (sometimesmore than 20 °C [36 °F]) within only a few hours.Other types of katabatic wind can occur when the

underlying geography is characterized by a cold plateauadjacent to a relatively warm region of lower elevation.Such conditions are satisfied in areas in which major icesheets or cold elevated land surfaces border warmer large

bodies of water. Air over the cold plateau cools and formsa large dome of cold dense air. Unless held back by background wind conditions, this cold air will spill over into

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General patterns of atmospheric circulation over an idealized Earth with auniform surface (top) and the actual Earth (bottom). Both horizontal andvertical patterns of atmospheric circulation are depicted in the diagram of the

actual Earth. Encyclopædia Britannica, Inc.



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the lower elevations with speeds that vary from gentle (afew kilometres per hour) to intense (93 to 185 km [58 to 115miles] per hour), depending on the incline of the slope ofthe terrain and the distribution of the background pres-sure field. Two special varieties of katabatic wind are wellknown in Europe. One is the bora, which blows from thehighlands of Croatia, Bosnia and Herzegovina, andMontenegro to the Adriatic Sea; the other is the mistral,

which blows out of central and southern France to theMediterranean Sea. Creating blizzard conditions, intense

katabatic winds often blow northward off the AntarcticIce Sheet.

Zonal Surface Winds

Measurements of seasonal mean sea-level pressure revealthat, on the average, certain geographic locations canexpect to experience winds that emanate from one pre-

vailing direction largely dictated by the presence of majorsemipermanent pressure systems. Such prevailing windshave long been known in marine environments because oftheir influence on the great sailing ships.

Tropical and subtropical regions are characterized by a general band of low pressure lying near the Equator. Thisband is bounded by centres of high pressure that may

extend poleward into the middle latitudes. Between theselow- and high-pressure regions is the region of the tropical winds. Of these the most extensive are the trade winds. Sonamed because of their favourable influence on tradeships traveling across the subtropical North Atlantic,trade winds flow westward and somewhat in the directionof the Equator on the equatorward side of the subtropical

high-pressure centres. The “root of the trades,” occurringon the eastern side of a subtropical high-pressure centre,is characterized by subsiding air. This produces the very

warm, dry conditions above a shallow layer of oceanic

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stratus clouds found in the eastern extremes of the sub-tropical Atlantic and Pacific ocean basins. As the trade

winds progress westward, however, subsidence abates, theair mass becomes more humid, and scattered showersappear. These showers occur particularly on islands withelevated terrain features that interrupt the flow of the

warm moist air. The equatorward flow of the trade windsof the Northern and Southern hemispheres often resultsin a convergence of the two air streams in a region knownas the intertropical convergence zone (ITCZ). Deep con-

vective clouds, showers, and thunderstorms occur alongthe ITCZ.

When the air reaches the western extreme of the high-pressure centre, it turns poleward and then eventuallyreturns eastward in the middle latitudes. The poleward-moving air is now warm and laden with moist maritimetropical air (mT); it gives rise to the warm, humid, show-

ery climate characteristic of the Caribbean region, easternSouth America, and the western Pacific island chains. The

westerlies are associated with the changeable weathercommon to the middle latitudes. Migrating extratropi-cal cyclones and anticyclones associated with contrasting

warm moist air moving poleward from the tropics and colddry air moving equatorward from polar latitudes yield peri-

ods of rain (sometimes with violent thunderstorms), snow,sleet, or freezing rain interrupted by periods of dry, sunny,and sometimes bitterly cold conditions. Furthermore,these patterns are seasonally dependent, with moreintense cyclones and colder air prevailing in winter but

with a higher incidence of thunderstorms common inspring and summer. In addition, these migrations and the

associated climate are complicated by the presence oflandmasses and major mountain features, particularly inthe Northern Hemisphere.



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Monsoon storm, Apache Junction, Ariz. © Corbis

The westerlies lie on the equatorward side of the semi-permanent subpolar centres of low pressure. Poleward ofthese centres, the surface winds turn westward again oversignificant portions of the subpolar latitudes. As in themiddle latitudes, the presence of major landmasses, nota-bly in the Northern Hemisphere, results in significant

variations in these polar easterlies. In addition, the windsystems and the associated climate are seasonally depen-dent. During the short summer season, the wind systemsof the polar latitudes are greatly weakened. During thelong winter months, these systems strengthen, and peri-ods of snow alternate with long intervals of dry cold aircharacteristic of continental polar or continental arctic

air masses.These major regions of surface circulation and theirassociated pressure fields are related to mean meridional

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(north-south) circulation patterns as well. Although theirpresence is discernible in long-term mean statistics accu-mulated over a hemisphere, such cells are often difficult todetect on a daily basis at any given longitude.

Monsoons

Particularly strong seasonal pressure variations occur overcontinents. Such seasonal fluctuations, commonly calledmonsoons, are more pronounced over land surfacesbecause these surfaces are subject to more significant sea-sonal temperature variations than are water bodies. Since

land surfaces both warm and cool faster than water bodies,they often quickly modify the temperature and densitycharacteristics of air parcels passing over them.

In meteorology the maritime continent is the region made up of partsof Southeast Asia and the islands of Indonesia and the Philippines. It

is not a true continent but an area made up of thousands of islands of various sizes and numerous shallow bodies of water. It is named forthe widespread interaction between land and water occurring there.The relief on many of the islands and peninsulas is significant, and thesurrounding seas possess some of the highest sea surface tempera-tures on Earth. These characteristics help to spawn numerousthunderstorms generated by sea-breeze convergence and convection.

At the global scale, the maritime continent is a key driver of

atmospheric circulation because of its enormous ability to transferheat to the air. It is strongly associated with the El Niño/SouthernOscillation (ENSO) during the dry season, and it is the heat source forthe Australian monsoon. In addition, the maritime continent’s topographic complexities often cause atmospheric models tounderestimate the region’s true temperature and rainfall patterns.



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Monsoons blow for approximately six months fromthe northeast and six months from the southwest, princi-pally in South Asia and parts of Africa; however, similarconditions also occur in Central America and the areabetween Southeast Asia and Australia . Summer monsoonshave a dominant westerly component and a strong ten-dency to converge, rise, and produce rain. Wintermonsoons have a dominant easterly component and astrong tendency to diverge, subside, and cause drought.Both are the result of differences in annual temperature

trends over land and sea.

The Diurnal Variability of Monsoons

Landmasses in regions affected by monsoons warm up very rapidly in the afternoon hours, especially on days with cloud-free conditions; surface air temperaturesbetween 35 and 40 °C (95 and 104 °F) are not uncom-

mon. Under such conditions, warm air is slowly andcontinually steeped in the moist and cloudy environ-ment of the monsoon. Consequently, over the course ofa 24-hour period, energy from this pronounced diurnal,or daily, change in terrestrial heating is transferred to thecloud, rain, and diurnal circulation systems. The scaleof this diurnal change extends from that of coastal sea

breezes to that of continent-sized processes. Satelliteobservations have confirmed that the effects of rapiddiurnal temperature change occur at continental scales.For example, air from surrounding areas is drawn intothe lower troposphere over warmer land areas of SouthAsia during summer afternoon hours. This buildup ofafternoon heating is accompanied by the production

of clouds and rain. In contrast, a reverse circulation,characterized by suppressed clouds and rain, is noted inthe early morning hours.

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(Bottom) Graph of monsoon rainfall in India, 1871–1981. Annual rainfall amounts are depicted as percentages departing from the 110-year average. The dashed line superimposed on the graph suggests a recurring trend over this time period. Encyclopædia Britannica, Inc.

(Top) Graph depicting the influence of El Niño/Southern Oscillation (ENSO)on rainfall produced by the Indian summer monsoon. During years when

ENSO is active, monsoon-driven precipitation over India often declines.Encyclopædia Britannica, Inc.

The Intra-Annual Variability of Monsoons

Monsoon rainfall and dry spells alternate on several times-

cales. One such well-known timescale is found aroundperiods of 40–50 or 30–60 days. This is called the Madden- Julian oscillation (MJO), named for American atmosphericscientists Roland Madden and Paul Julian in 1971. Thisphenomenon comes in the form of alternating cyclonicand anticyclonic regions that enhance and suppress rain-fall, respectively, and flow eastward along the Equator inthe Indian and Pacific oceans. The MJO has the ability toinfluence monsoonal circulation and rainfall by addingmoisture during its cyclonic (wet) phase and reducing con-

vection during its anticyclonic (dry) phase. At the surfacein monsoon regions, both dry and wet spells result. Theseperiods may alternate locally on the order of two or more

weeks per phase.

The Interannual Variability of MonsoonsThe variability of monsoon-driven rainfall in the IndianOcean and Australia appears to parallel El Niño episodes.During El Niño events, which occur about every twoto seven years, ocean temperatures rise over the centralequatorial Pacific Ocean by about 3 °C (5.4 °F). Atypicalconditions characterized by increased rising air motion,

convection, and rain are created in the western equato-rial Pacific. At the same time, a compensating lobe of

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descending air, producing below-normal rainfall, appearsin the vicinity of eastern Australia, Malaysia, and India.

Many other factors, aside from equatorial PacificOcean surface temperatures, contribute to the interan-nual variability of monsoon rainfall. Excessive spring snowand ice cover on the Plateau of Tibet is related to the defi-cient monsoon rainfall that occurs during the followingsummer season in India. Furthermore, strong evidenceexists that relates excessive snow and ice cover in westernSiberia to deficient Asian summer rainfall. Warmer than

normal sea surface temperatures over the Indian Oceanmay also contribute somewhat to above-normal rainfall inSouth Asia. The interplay among these many factorsmakes forecasting monsoon strength a challenging prob-lem for researchers.

A rather clear signature on the decadal variability ofIndian rainfall has been documented by the Indian

Weather Service. Periods of heavier-than-normal rainfallare followed by decades of somewhat less rainfall.

Upper-Level Winds

The flow of air around the globe is greatest in the higheraltitudes, or upper levels. Upper-level airflow occurs in

wavelike currents that may exist for several days beforedissipating. Upper-level wind speeds generally occur onthe order of tens of metres per second and vary withheight. The characteristics of upper-level wind systems

vary according to season and latitude and to some extenthemisphere and year. Wind speeds are strongest in themidlatitudes near the tropopause and in the mesosphere.

The Characteristics of Upper-Level Wind Systems

Upper-level wind systems, like all wind systems, may bethought of as having parts consisting of uniform flow,



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Rossby wave patterns over the North Pole depicting the formation of anoutbreak of cold air over Asia.

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A weather balloon being released at a weather station at the South Pole. Balloon-borne instrument packages designed to track upper-level winds andcapable of being tracked by radar are called rawinsondes. NOAA

rotational flow (with cyclonic or anticyclonic curvature),convergent or divergent flow (in which the horizontal areaof masses of air shrinks or expands), and deformation (by

which the horizontal area of air masses remains constant while experiencing a change in shape). Upper-level windsystems in the midlatitudes tend to have a strong compo-nent of uniform flow from west to east (“westerly” flow),though this flow may change during the summer. A seriesof cyclonic and anticyclonic vortices superimposed on theuniform west-to-east flow make up a wavetrain (a succes-

sion of waves occurring at periodic intervals). The wavesare called Rossby waves after Swedish American meteorol-ogist C.G. Rossby, who first explained fundamental aspectsof their behaviour in the 1930s. Waves whose wavelengths



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are around 6,000 km (3,700 miles) or less are called “short waves,” while those with longer wavelengths are called“long waves.” In addition, short waves progress in thesame direction as the mean airflow, which is from west toeast in the midlatitudes; long waves retrogress (that is,move in the opposite direction of the mean flow). Althoughthe undulating current of air is composed of a number of

waves of varying wavelength, the dominant wavelengthis usually around several thousand kilometres. Near andunderneath the tropopause, regions of divergence are

found over regions of gently rising air at the surface, whileregions of convergence aloft are found over regions ofsinking air below. These regions are usually much more dif-ficult to detect than the regions of rotational and uniformflow. While the horizontal wind speed is typically in therange of 10–50 metres per second (about 20–110 miles perhour), the vertical wind speed associated with the waves is

only on the order of centimetres per second.The characteristics of upper-level wind systems are

known mainly from an operational worldwide network ofrawinsonde observations. (A rawinsonde is a type of radio-sonde designed to track upper-level winds and whoseposition can be tracked by radar.) Winds measured fromDoppler-radar wind profilers, aircraft navigational sys-

tems, and sequences of satellite-observed cloud imageryhave also been used to augment data from the rawinsondenetwork; the latter two have been especially useful fordefining the wind field over data-sparse regions, such asover the oceans.

The winds at upper levels, where surface friction doesnot occur, tend to be approximately geostrophic. In other

words, there is a near balance between the pressure gradi-ent force, which directs air from areas of relatively highpressure to areas of relatively low pressure, and the

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Coriolis force, which deflects air from its straight-linepath to the right in the Northern Hemisphere and to theleft in the Southern Hemisphere. An important conse-quence of this geostrophic balance is that the winds blowparallel to isobars (cartographic lines indicating areas ofequal pressure), and, according to Buys Ballott’s law, lowerpressures will be found to the left of the direction of the

wind in the Northern Hemisphere and to the right ofthe wind in the Southern Hemisphere. Furthermore, windspeed increases as the spacing between isobars decreases.

In a wavetrain of westerly flow, the regions of cyclonicflow are associated with troughs of low pressure, whereasanticyclonic flow are characterized by ridges of high pres-sure. Rising motions tend to be found downstream fromthe troughs and upstream from the ridges, while sinkingmotions tend to be found downstream from the ridgesand upstream from the troughs. The areas of rising motion

tend to be associated with clouds and precipitation(inclement weather), whereas the areas of sinking motiontend to be associated with clear skies (fair weather).

The vertical variation of the structure of the wavesdepends upon the temperature pattern. In general, owingto the net difference in incoming shorter-wavelength solarradiation and outgoing longer-wavelength infrared radia-

tion between the polar and the equatorial regions, there isa horizontal temperature gradient in the troposphere. Atboth the surface and upper levels, the troposphere is

warmest at low latitudes and coldest at high latitudes. Theatmosphere is mainly in hydrostatic balance, or equilib-rium, between the upward-directed pressure gradientforce and the downward-directed force of gravity. This cir-

cumstance is expressed in the following relationship:∂ p/ ∂ z = – ρ g (1) where ∂ p/ ∂ z is the partial derivative of p withrespect to z, p is the pressure, z is the height, ρ is the



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density of the air, and g is the acceleration of gravity. Aconsequence of this hydrostatic relationship is that thepressure at any level is equal to the weight of the columnof air above. According to the ideal gas law, p = ρ RT (2) where R is the gas constant and T is the temperature.At any given pressure, the density varies inversely withtemperature. Therefore, relatively cold air is heavier thanrelatively warm air at the same pressure. It follows from (1)and (2) that pressure decreases more rapidly with height athigh latitudes in the colder air than it does at lower lati-

tudes in the warmer air. If there is a westerly geostrophic wind at midlevels in the troposphere, then pressuredecreases with increasing latitude. Consequently, the hor-izontal spacing between isobars decreases with height.Thus, the geostrophic wind speed, which approximatesthe actual wind speed, increases with height. Above thetropopause the pole-to-Equator temperature gradient is

reversed as air temperature increases with height, so thatthe westerlies decrease in intensity in the stratosphere.Thus, the strongest westerly current of winds is locatednear the tropopause.

The aforementioned relationship can be analyzedquantitatively by considering the vertical variation in the

geostrophic wind, which is found from the hydrostatic

equation (1), the ideal gas law (2), and the geostrophic windformula, approximately as follows. ∂ug / ∂ z = – g / fT ∂T / ∂ y and ∂vg / ∂ z = – g / fT ∂T / ∂ z, (3) where ∂ug / ∂ z is the partialderivative of ug with respect to z, ug and vg are the compo-nents of the geostrophic wind in the zonal (straight from

west to east) and meridional (north to south) directions,respectively, and f is the Coriolis parameter. The equations

given in (3) are known as the thermal-wind relations. Thedifference between the geostrophic wind at some higherlevel and the geostrophic wind below is called the thermal

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wind. It follows that the thermal wind vector is orientedso that in the colder air it lies to the left in the NorthernHemisphere and to the right in the Southern Hemisphere.

In addition to the general pole-to-Equator tempera-ture gradient found in the troposphere, there are zonallyoriented temperature variations that are wavelike. In fact,to a first approximation, the isotherms (cartographic linesindicating areas of equal temperature) are nearly paral-lel to the isobars in the upper levels of the troposphere.Most frequently, relatively cold air lies just upstream from

upper-level troughs and just downstream from upper-levelridges, while relatively warm air lies just upstream fromupper-level ridges and just downstream from upper-leveltroughs. The thermal-wind relation (3) indicates thatthe wavetrain of troughs and ridges tilts with height to the

west. In the midlatitudes during the summer, and in somelocations within the midlatitudes during the winter, the

meridional temperature gradient weakens so much thatthe westerlies become weak or nonexistent. As a result,the wavelike wind field disappears and the flow pattern isthat of cyclones and anticyclones “cut off” from the flow.When cold air is colocated with the upper-level cyclonesand warm air is colocated with the upper-level anticy-clones, according to (3), both circulation patterns increase

in intensity with height and are called cold-core and warm-core systems, respectively. Tropical cyclones, on the otherhand, are warm-core systems that are most intense at thesurface and that decrease in intensity with height.

The vertical structure of upper-level waves has animportant effect on smaller-scale features that may beembedded within them. The susceptibility of the atmo-

sphere to vertical overturning (a mixing of lower-level warmer air with higher-level colder air) through deepcumulus convection (e.g., thunderstorms) depends on the



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rate at which temperature decreases with height. Whenregions of relatively cold air aloft associated with upper-level troughs or cyclones become superimposed during the

winter over relatively warm ocean surfaces or duringthe summer over hot and humid landmasses, then convec-tive storms can form. The type of mesoscale convectivesystem (MCS) that can form depends in large part on the

vertical wind shear. When the vertical shear is very strong,supercells and tornadoes may be spawned, especially dur-ing the warmer months. During the winter, bands of

precipitation sometimes line up along the vertical shear vector through a process known as slantwise convection.

The Propagation and Development of Waves

Upper-level waves in the westerlies in midlatitudes usuallymove from west to east, in part as a result of advection (aprocess in which the airflow transports a property of the

atmosphere [warmth, cold, etc.] downstream) and in partas a result of propagation, which acts in the oppositedirection, toward the west. Rossby showed that to a goodapproximation, c = U – β / (2π/ L )2, (4) where c is the phasespeed of the waves, U is the speed from west to east of thecomponent of upper-level wind due to uniform flow, β isthe meridional, or north-south, gradient of the Coriolis

parameter ( f ), and L is the zonal wavelength (the distancebetween successive troughs or ridges). According to (4),since the magnitude of f increases toward the poles, β ispositive, and hence waves whose wavelengths are shorthave a relatively small component due to propagation. Inthis situation, advection overwhelms the effect of propa-

gation and the waves move on downstream. On the other

hand, if in midlatitudes the wavelength is very long, thenthe effects of propagation may exactly cancel the effectsof advection, and the waves may become stationary; or if

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the wavelength becomes even longer, then the waves maybecome retrograde. From (4) it can be seen that Rossby

waves owe their propagation characteristics to the north-south variation of f . In nature, temperature effects andheating and cooling over warm and cold surfaces can mod-ify (4) somewhat.

The physical basis for (4) and for the development ofupper-level systems and how they relate to surface sys-tems is described by an elegant theory developed in thelate 1940s called quasigeostrophic theory. A measure of

the tendency for a fluid to rotate is known as vorticityand is given by the following equation: ζ = ∂v/ ∂ x – ∂u/ ∂ y (5) where ζ is the relative vorticity with respect to Earth’ssurface. The variables x and y are the coordinate axes forspace and correspond to the measurements to the east andnorth, respectively. The variables u and v are zonal andmeridional components (the components of motion in

the easterly and northerly directions), respectively, of the wind. On the rotating Earth, the vorticity is the sum ofthe relative vorticity with respect to Earth’s surface, givenby the aforementioned expression, and Earth’s vorticity,

given by f , the Coriolis parameter. Troughs are associated with cyclonic vorticity, and ridges are associated withanticyclonic vorticity. In a wavetrain, the pressure falls

downstream from troughs, where the wind is directed fromthe region of maximum vorticity along the trough to theregion of minimum vorticity, which is along the ridge, andthe pressure rises downstream from ridges. On the otherhand, pressure can rise east of troughs (and west of ridges)

where there is component of motion from the Equator tothe pole. For example, pressure rises from regions of low

magnitude of f to higher magnitude of f (from low valuesof Earth’s vorticity to higher values of Earth’s vorticity).Likewise, pressure falls west of troughs (and east of ridges)



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where there is a component of motion from one of thepoles to the Equator—from relatively high magnitude of

f to lower magnitude of f . The effect of pressure increasesand decreases is greatest when the wavelength is relatively

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Vertical cross sections through a wave system depicting typical divergence andconvergence distributions for non-tilting and tilting systems. EncyclopædiaBritannica, Inc.

short, such as when the effects of the advection of Earth’s vorticity are overwhelmed by the effects of advection of

relative vorticity.The development and amplification of Rossby wavesis typically a result of the advection of warmer or colder



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air at low levels. When warm air is advected underneatha layer of air not experiencing much, if any, advection,the pressure at the top of the layer rises. Conversely,pressure falls when cold air advects under a similar layerof air. If the wavetrain tilts to the west with height sothat cold air lies to the west of troughs and thus eastof ridges, the pressure aloft in the troughs decreases.Similarly, when warm air lies to the east of troughs andthus west of ridges, the pressure aloft in the ridgesincreases. As a result, the amplitude of the waves in the

wavetrain increases, thereby enhancing the temperatureadvection process, so that there is a positive feedbackmechanism that makes the waves continue to amplify.In this process, called “baroclinic instability,” potentialenergy is converted into kinetic energy—which occurs as

wind—as warm, light air rises and cold, heavy air sinks.Since baroclinic instability is associated with horizontal

temperature gradients, according to the thermal windrelation (3), there must be vertical wind shear.

It is also possible for Rossby waves to amplify througha process called barotropic instability. Barotropic instabil-ity, however, requires horizontal shear, not vertical shear;kinetic energy for the waves comes from the mean kineticenergy associated with the westerly wind current. The

waves grow in amplitude at the expense of the mean flow.Barotropic instability can occur when the horizontal shear varies with latitude such that the sum of Earth’s vorticityand the relative vorticity associated with the horizontalshear is small with respect to latitude.

The Relationships Between Upper-Level Winds

and Surface Features

Rossby waves propagating through the upper and middletroposphere cause disturbances to form at the surface.



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According to quasigeostrophic theory, when there is a wavetrain embedded within a zone of pole-to-Equatortemperature gradient, air rises east of upper-level troughs(and west of upper-level ridges) and sinks west of upper-level troughs (and east of upper-level ridges). These verticalair motions are required to maintain the approximate geo-

strophic and hydrostatic balance, which are necessary forquasigeostrophic equilibrium. Air converges at the surfaceunderneath the rising current of air to compensate for theupward loss of mass and diverges at the surface underneatha sinking current of air to compensate for the downward

gain of mass. As a consequence of the lateral deviationof the air by the Coriolis force, Earth’s vorticity is con-

verted into cyclonic relative vorticity where air convergesand anticyclonic relative vorticity where air diverges.According to the geostrophic wind relation, cyclonic gyres

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Positions of jet streams in the atmosphere. Arrows indicate directions of mean motions in a meridional plane. Encyclopædia Britannica, Inc.



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are associated with low-pressure centres, whereas anticy-clonic gyres are connected with areas of high pressure.Thus, low-pressure areas form at the surface downstreamfrom upper-level troughs and upstream from upper-levelridges, whereas the reverse is true for high-pressure areas.These surface low- and high-pressure areas thereby createa westward tilt with height of the waves in pressure. Sincethere tends to be a pole-to-Equator-directed geostrophic

wind west of surface lows and east of surface highs, andan Equator-to-pole-directed geostrophic wind east of

surface lows and west of surface highs, there is cold advec-tion underneath upper-level troughs and warm advectionunderneath upper-level ridges; the baroclinic instabilityprocess is thus facilitated.

Jet Streams

The upper-level wind flow described above is frequently

concentrated into relatively narrow bands called jetstreams, or jets. The jets, whose wind speeds are usually inexcess of 30 metres per second (about 70 miles per hour)but can be as high as 107 metres per second (about 240miles per hour), act to steer upper-level waves. Jet streamsare of great importance to air travel because they affectthe ground speed, the velocity relative to the ground, of

aircraft. Since strong upper-level flow is usually associated with strong vertical wind shear, jet streams in midlatitudesare accompanied by strong horizontal temperature gradi-ents, as required by the thermal wind relation (3). Someregions of high vertical wind shear are marked by clear-airturbulence (CAT). Jet streams whose extents are relativelyisolated are called jet streaks. Well-defined circulation

patterns of rising and sinking air are usually found justupstream and downstream, respectively, from jet streaks(that are not too curved). Rising motion is found to the



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Meridional cross section of the atmosphere to a height of 60 kilometres(37.2 miles) in summer and winter hemispheres, showing seasonal changes.

Numerical values for wind are in units of metres per second and typical of

the Northern Hemisphere, but the structure is much the same in the Southern Hemisphere. Positive and negative signs indicate winds of opposite directional sense. Encyclopædia Britannica, Inc.

left and right just downstream and upstream, respectively,and sinking motion is found to the right and left justdownstream and upstream, respectively. Jets tend to bestrongest near the tropopause where the horizontal tem-perature gradient reverses.

The polar front jet moves in a generally westerly direc-tion in midlatitudes, and its vertical wind shear whichextends below its core is associated with horizontaltemperature gradients that extend to the surface. As a con-sequence, this jet manifests itself as a front that marks the

division between colder air over a deep layer and warmerair over a deep layer. The polar front jet can be baroclini-cally unstable and break up into waves. The subtropical jetis found at lower latitudes and at slightly higher elevation,

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owing to the increase in height of the tropopause at lowerlatitudes. The associated horizontal temperature gradi-ents of the subtropical jet do not extend to the surface,so that a surface front is not evident. In the tropics aneasterly jet is sometimes found at upper levels, especially

when a landmass is located poleward of an ocean, so thetemperature increases with latitude. The polar front jetand the subtropical jet play a role in maintaining Earth’s

general circulation. They are slightly different in eachhemisphere because of differences in the distribution of

landmasses and oceans.

Winds in the Stratosphere and Mesosphere

The winds in the stratosphere and mesosphere are usuallyestimated from temperature data collected by satellites.The winds at these high levels are assumed to be geo-strophic. Overall, in the midlatitudes, they have a westerly

component in the winter and an easterly component inthe summer. The highest zonal winds are around 60–70metres per second (135–155 miles per hour) at 65–70 km (37– 43 miles) above Earth’s surface. The west-wind componentis stronger during the winter in the Southern Hemisphere.The axes of the strongest easterly and westerly wind com-ponents in the Southern Hemisphere tilt toward the south

with increased altitude during the Northern Hemisphere winter and the Southern Hemisphere summer. The zonalcomponent of the thermal wind shear is in accord with thezonal distribution of temperature.

During the winter there is, in the mean, an intensecyclonic vortex about the poles in the lower stratosphere.Over the North Pole this vortex has an embedded mean

trough over northeastern North America and over north-eastern Asia, whereas over the Pacific there is a weakanticyclonic vortex. The winter cyclonic vortex over the



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South Pole is much more symmetrical than the one overthe North Pole. During the summer there is an anticycloneabove each pole that is much weaker than the wintertimecyclone.

In the stratosphere, deviations from the mean behav-iour of the winds occur during events called sudden

warmings, when the meridional temperature gradientreverses on timescales as short as several days. This alsohas the effect of reversing the zonal wind direction.Sudden warmings tend to occur during the early and

middle parts of the winter and the transition periodfrom winter to spring. The latter marks the change-over from the cold winter polar cyclone to the warm sum-mer polar anticyclone. It is noteworthy that long wavesfrom the troposphere can propagate into the stratosphereduring the winter when westerlies and sudden warmingsoccur, but this is not the case during the summer when

easterly winds prevail.The zonal component of the winds in the stratosphere

above equatorial and tropical regions is, in the mean, rela-tively weak. This is not necessarily the case at any giventime, because they reverse direction on the average every13–14 months. This phenomenon, which is known as thequasi-biennial oscillation (QBO), is caused by the interac-

tion of vertically propagating waves with the mean flow.Its effect is greatest around 27 km (17 miles) above Earth’ssurface in the equatorial region. The strongest easterliesare stronger than the strongest westerlies.

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CLIMATIC CLASSIFICATION

The climate of an area is the synthesis of the environ-mental conditions (soils, vegetation, weather, etc.)

that have prevailed there over a long period of time. Thissynthesis involves both averages of the climatic ele-ments and measurements of variability (such as extreme

values and probabilities). Climate is a complex, abstractconcept involving data on all aspects of Earth’s environ-

ment. As such, no two localities on Earth may be said tohave exactly the same climate. Nevertheless, it is readilyapparent that, over restricted areas of the planet, cli-mates vary within a limited range and that climaticregions are discernible within which some uniformity isapparent in the patterns of climatic elements. Moreover,

widely separated areas of the world possess similar cli-

mates when the set of geographic relationships occurringin one area parallels that of another. This symmetry andorganization of the climatic environment suggests anunderlying worldwide regularity and order in the phe-nomena causing climate (e.g., patterns of incoming solarradiation, vegetation, soils, winds, temperature, and airmasses), which were discussed in chapter 1.

Climate classification is an attempt to formalize thisprocess of recognizing climatic similarity, of organizing,simplifying, and clarifying the vast amount of environ-mental data, and of systematizing the long-term effects ofinteracting climatic processes to enhance scientific under-standing of climates. Users of climate classificationsshould be aware of the limitations of the procedure,

however.First, climate is a multidimensional concept, andit is not an obvious decision as to which of the many



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observed environmental variables should be selected asthe basis of the classification. This choice must be madeon a number of grounds, both practical and theoretical.For example, using too many different elements opensup the possibilities that the classification will have toomany categories to be readily interpreted and that manyof the categories will not correspond to real climates.Moreover, measurements of many of the elements of cli-mate are not available for large areas of the world or havebeen collected for only a short time. The major excep-

tions are soil, vegetation, temperature, and precipitationdata, which are more extensively available and have beenrecorded for extended periods of time.

The choice of variables also is determined by the pur-pose of the classification (e.g., to account for distributionof natural vegetation, to explain soil formation processes,or to classify climates in terms of human comfort). The

variables relevant in the classification will be determinedby this purpose, as will the threshold values of the vari-ables chosen to differentiate climatic zones.

A second difficulty results from the generally gradualnature of changes in the climatic elements over Earth’ssurface. Except in unusual situations due to mountainranges or coastlines, temperature, precipitation, and

other climatic variables tend to change only slowlyover distance. As a result, climate types tend to changeimperceptibly as one moves from one locale on Earth’ssurface to another. Choosing a set of criteria to distin-

guish one climatic type from another is thus equivalentto drawing a line on a map to distinguish the climaticregion possessing one type from that having the other.

While this is in no way different from many other clas-sification decisions that one makes routinely in dailylife, it must always be remembered that boundaries




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between adjacent climatic regions are placed somewhat arbitrarily through regions of continuous, gradualchange and that the areas defined within these boundar-ies are far from homogeneous in terms of their climaticcharacteristics.

Most classification schemes are intended for global-or continental-scale application and define regionsthat are major subdivisions of continents hundreds tothousands of kilometres across. These may be termedmacroclimates. Not only will there be slow changes

(from wet to dry, hot to cold, etc.) across such a regionas a result of the geographic gradients of climatic ele-ments over the continent of which the region is a part,but there will exist mesoclimates within these regionsassociated with climatic processes occurring at a scaleof tens to hundreds of kilometres that are created byelevation differences, slope aspect, bodies of water, dif-

ferences in vegetation cover, urban areas, and the like.Mesoclimates, in turn, may be resolved into numerousmicroclimates, which occur at scales of less than 0.1 km(0.06 mile), as in the climatic differences between for-ests, crops, and bare soil, at various depths in a plantcanopy, at different depths in the soil, on different sidesof a building, and so on.

These limitations notwithstanding, climate classifi-cation plays a key role as a means of generalizing the geographic distribution and interactions among cli-matic elements, of identifying mixes of climaticinfluences important to various climatically dependentphenomena, of stimulating the search to identify thecontrolling processes of climate, and, as an educational

tool, to show some of the ways in which distant areas ofthe world are both different from and similar to one’sown home region.



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APPROACHES TO CLIMATICCLASSIFICATION

The earliest known climatic classifications were those ofclassical Greek times. Such schemes generally divided theEarth into latitudinal zones based on the significant paral-lels of 0°, 23.5°, and 66.5° of latitude and on the length ofday. Modern climate classification has its origins in themid-19th century, with the first published maps of tem-perature and precipitation over Earth’s surface, which

permitted the development of methods of climate group-ing that used both variables simultaneously.Many different schemes of classifying climate have

been devised (more than 100), but all of them may bebroadly differentiated as either empiric or genetic meth-ods. This distinction is based on the nature of the dataused for classification. Empirical methods make use

of observed environmental data, such as temperature,humidity, and precipitation, or simple quantities derivedfrom them (e.g., evaporation). In contrast, a genetic

Map of climatic zones as envisioned by the Ancient Greeks. EncyclopaediaBritannica, Inc..





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method classifies climate on the basis of its causal ele-ments, the activity and characteristics of all factors (airmasses, circulation systems, fronts, jet streams, solar radi-ation, topographic effects, and so forth) that give rise tothe spatial and temporal patterns of climatic data. Hence,

while empirical classifications are largely descriptive ofclimate, genetic methods are (or should be) explanatory.Unfortunately, genetic schemes, while scientifically moredesirable, are inherently more difficult to implementbecause they do not use simple observations. As a result,

such schemes are both less common and less successfuloverall. Moreover, the regions defined by the two types ofclassification schemes do not necessarily correspond; inparticular, it is not uncommon for similar climatic formsresulting from different climatic processes to be groupedtogether by many common empirical schemes.

Genetic Classifications

Genetic classifications group climates by their causes.Among such methods, three types may be distinguished:(1) those based on the geographic determinants of climate,(2) those based on the surface energy budget, and (3) thosederived from air-mass analysis.

In the first class are a number of schemes (largely the work of German climatologists) that categorize climatesaccording to such factors as latitudinal control of temper-ature, continentality versus ocean-influenced factors,location with respect to pressure and wind belts, andeffects of mountains. These classifications all share a com-mon shortcoming: they are qualitative, so that climatic

regions are designated in a subjective manner rather thanas a result of the application of some rigorous differentiat-ing formula.



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An interesting example of a method based on theenergy balance of Earth’s surface is the 1970 classifica-tion of Werner H. Terjung, an American geographer.His method utilizes data for more than 1,000 locations

worldwide on the net radiation received at the surface,the available energy for evaporating water, and the avail-able energy for heating the air and subsurface. The annualpatterns are classified according to the maximum energyinput, the annual range in input, the shape of the annualcurve, and the number of months with negative magni-

tudes (energy deficits). The combination of characteristicsfor a location is represented by a label consisting of severalletters with defined meanings, and regions having similarnet radiation climates are mapped.

Probably the most extensively used genetic systems,however, are those that employ air-mass concepts. Airmasses are large bodies of air that, in principle, possess rel-

atively homogeneous properties of temperature, humidity,etc., in the horizontal. Weather on individual days may beinterpreted in terms of these features and their contrastsat fronts.

Two American geographer-climatologists have beenmost influential in classifications based on air mass. In1951, Arthur N. Strahler described a qualitative classifi-

cation based on the combination of air masses presentat a given location throughout the year. Some years later(1968 and 1970), John E. Oliver placed this type of clas-sification on a firmer footing by providing a quantitativeframework that designated particular air masses andair mass combinations as “dominant,” “subdominant,”or “seasonal” at particular locations. He also provided

a means of identifying air masses from diagrams ofmean monthly temperature and precipitation plottedon a “thermohyet diagram,” a procedure that obviates





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the need for less common upper-air data to make theclassification.

Empirical Classifications

Most empirical classifications are those that seek to groupclimates based on one or more aspects of the climate sys-tem. While many such phenomena have been used in this

way, natural vegetation stands out as one of prime impor-tance. The view held by many climatologists is that natural

vegetation functions as a long-term integrator of the cli-mate in a region; the vegetation, in effect, is an instrumentfor measuring climate in the same way that a thermometermeasures temperature. Its preeminence is apparent in thefact that many textbooks and other sources refer to cli-mates using the names of vegetation, as, for example,rainforest, taiga, or tundra.

Wladimir Köppen, a German botanist-climatologist,developed the most popular (but not the first) of these

vegetation-based classifications (see pages xiv and xv). Hisaim was to devise formulas that would define climaticboundaries in such a way as to correspond to those of the

vegetation zones that were being mapped for the firsttime during his lifetime. Köppen published his first

scheme in 1900 and a revised version in 1918. He contin-ued to revise his system of classification until his death in1940. Other climatologists have modified portions ofKöppen’s procedure on the basis of their experience in

various parts of the world.Köppen’s classification is based on a subdivision of

terrestrial climates into five major types, which are rep-

resented by the capital letters A, B, C, D, and E. Each ofthese climate types except for B is defined by tempera-ture criteria. Type B designates climates in which the



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controlling factor on vegetation is dryness (rather thancoldness). Aridity is not a matter of precipitation alonebut is defined by the relationship between the precipi-tation input to the soil in which the plants grow andthe evaporative losses. Since evaporation is difficult toevaluate and is not a conventional measurement at mete-orological stations, Köppen was forced to substitute aformula that identifies aridity in terms of a temperature-precipitation index (i.e., evaporation is assumed to becontrolled by temperature). Dry climates are divided into

arid (BW) and semiarid (BS) subtypes, and each may bedifferentiated further by adding a third code, h for warmand k for cold.

As noted above, temperature defines the other fourmajor climate types. These are subdivided, with additionalletters again used to designate the various subtypes. Type Aclimates (the warmest) are differentiated on the basis of the

seasonality of precipitation: Af (no dry season), Am (shortdry season), or Aw (winter dry season). Type E climates (thecoldest) are conventionally separated into tundra (ET) andsnow/ice climates (EF). The mid-latitude C and D climatesare given a second letter, f (no dry season), w (winter dry),or s (summer dry), and a third symbol (a, b, c, or d [the lastsubclass exists only for D climates]), indicating the warmth

of the summer or the coldness of the winter.The Köppen classification has been criticized on many grounds. It has been argued that extreme events, such as aperiodic drought or an unusual cold spell, are just as signif-icant in controlling vegetation distributions as the meanconditions upon which Köppen’s scheme is based. It alsohas been pointed out that factors other than those used in

the classification, such as sunshine and wind, are importantto vegetation. Moreover, it has been contended that natu-ral vegetation can respond only slowly to environmental





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change, so that the vegetation zones observable today arein part adjusted to past climates. Many critics have drawnattention to the rather poor correspondence between theKöppen zones and the observed vegetation distributionin many areas of the world. In spite of these and otherlimitations, the Köppen system remains the most popularclimatic classification in use today.

A major contribution to climate grouping was madeby the American geographer-climatologist C. WarrenThornthwaite in 1931 and 1948. He first used a vegetation-

based approach that made use of the derived concepts oftemperature efficiency and precipitation effectivenessas a means of specifying atmospheric effects on vegeta-tion. His second classification retained these conceptsin the form of a moisture index and a thermal efficiencyindex but radically changed the classification criteria andrejected the idea of using vegetation as the climatic inte-

grator, attempting instead to classify “rationally” on thebasis of the numerical values of these indices. His 1948scheme is encountered in many climatology texts, but ithas not gained as large a following among a wide audienceas the Köppen classification system, perhaps because ofits complexity and the large number of climatic regionsit defines.

While vegetation-based climate classifications could beregarded as having relevance to human activity through what they may indicate about agricultural potential and nat-ural environment, they cannot give any sense of how humanbeings would feel within the various climate types. Terjung’s1966 scheme was an attempt to group climates on the basisof their effects on human comfort. The classification makes

use of four physiologically relevant parameters: tempera-ture, relative humidity, wind speed, and solar radiation. Thefirst two are combined in a comfort index to express



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CLASSIFICATION OF MAJORCLIMATIC TYPES ACCORDING TO THE

KÖPPEN-GEIGER-POHL SCHEMEletter symbol

1st 2nd 3rd criterion

A Temperature of coolest month18 degrees Celsius or higher

f Precipitation in driest monthat least 60 mm

m Precipitation in driest monthless than 60 mm but equal toor greater than 100 – ( r /25)1

w Precipitation in driest monthless than 60 mm and less than100 – ( r /25)1


atmospheric conditions in terms perceived as extremelyhot, hot, oppressive, warm, comfortable, cool, keen, cold,

very cold, extremely cold, and ultra cold. Temperature, windspeed, and solar radiation are combined in a wind effectindex expressing the net effect of wind chill (the coolingpower of wind on exposed surfaces) and addition of heat tothe human body by solar radiation. These indices are com-bined for different seasons in different ways to express howhumans feel in various geographic areas on a yearly basis.Terjung visualized that his classification would find applica-

bility in medical geography, climatological education,tourism, housing, clothing, and as a general analytical tool.

Many other specialized empirical classifications havebeen devised. For example, there are those that differenti-ate between types of desert and coastal climates, thosethat account for different rates of rock weathering or soilformation, and those based on the identification of simi-

lar agricultural climates.




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B2 70% or more of annual pre-cipitation falls in the summerhalf of the year and r less than

20t + 280, or 70% or more ofannual precipitation falls inthe winter half of the year and

r less than 20t, or neither halfof the year has 70% or more ofannual precipitation and r lessthan 20t + 1403

W r is less than one-half of theupper limit for classification asa B type (see above)

S r is less than the upper limitfor classification as a B typebut is more than one-half ofthat amount

h t equal to or greater than 18degrees Celsius

k t less than 18 degrees Celsius

C Temperature of warmestmonth greater than or equal to10 degrees Celsius, and tem-perature of coldest month lessthan 18 degrees Celsius but

greater than –3 degrees Celsiuss Precipitation in driest month

of summer half of the year isless than 30 mm and less thanone-third of the wettestmonth of the winter half

w Precipitation in driest month

of the winter half of the yearless than one-tenth of theamount in the wettest monthof the summer half



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f Precipitation more evenlydistributed throughout year;

criteria for neither s nor wsatisfied

a Temperature of warmest month22 degrees Celsius or above

b Temperature of each of four warmest months 10 degreesCelsius or above but warmest

month less than 22 degreesCelsius

c Temperature of one to threemonths 10 degrees Celsius orabove but warmest month lessthan 22 degrees Celsius

D Temperature of warmest

month greater than or equal to10 degrees Celsius, andtemperature of coldest month

–3 degrees Celsius or lower

s Same as for type C

w Same as for type C

f Same as for type C

a Same as for type Cb Same as for type C

c Same as for type C

d Temperature of coldest monthless than –38 degrees Celsius(d designation then usedinstead of a, b, or c)

E Temperature of warmestmonth less than 10 degreesCelsius





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THE WORLD DISTRIBUTION OFMAJOR CLIMATIC TYPES

The following discussion of the climates of the world isbased on groupings of Köppen’s climatic types (see pages

xiv and xv). Highland climates are also depicted.

Type A Climates

Köppen’s A climates are found in a nearly unbroken beltaround the Earth at low latitudes, mostly within 15° N and

T Temperature of warmestmonth greater than 0 degreesCelsius but less than 10degrees Celsius

F Temperature of warmest month0 degrees Celsius or below

H4 Temperature and precipitationcharacteristics highly depen-dent on traits of adjacentzones and overall elevation—

highland climates may occurat any latitude

1In the formulas above, r is average annual precipitation total(mm) and t is average annual temperature (degrees Celsius). Allother temperatures are monthly means (degrees Celsius), andall other precipitation amounts are mean monthly totals (mm).2Any climate that satisfies the criteria for designation as a B

type is classified as such, irrespective of its other characteristics.3The summer half of the year is defined as the monthsApril–September for the Northern Hemisphere andOctober–March for the Southern Hemisphere.4Most modern climate schemes consider the role of altitude.The highland zone has been taken from Trewartha (1968).



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S. Their location within a region in which available netsolar radiation is large and relatively constant from monthto month ensures both high temperatures (generally inexcess of 18 °C [64 °F]) and a virtual absence of thermalseasons. Typically, the temperature difference betweenday and night is greater than that between the warmestand the coolest month, the opposite of the situation inmid-latitudes. The terms winter and summer have littlemeaning, rather, in many locations annual rhythm is pro-

vided by the occurrence of wet and dry seasons. Type A

climates are controlled mainly by the seasonal fluctua-tions of the trade winds, the intertropical convergencezone (ITCZ), and the Asian monsoon.


Lowland rain forest along the northern coast of Ecuador. Tropical lowland rain forests are vegetation types found in the ever-wet tropics that are dominated by broad-leaved evergreen trees. They grow primarily in South andCentral America, West and Central Africa, Indonesia, parts of Southeast Asia,

and northern Australia. © Victor Englebert



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Wet Equatorial Climate (Af)

Within about 12° latitude of the Equator lies a region ofconsistently high temperatures (around 30 °C [86 °F]),

with plentiful precipitation (150–1,000 cm [59–394inches]), heavy cloud cover, and high humidity, with verylittle annual temperature variation. Such regions lie withinthe influence of the ITCZ in all months; the converging,ascending air spawns convectional thunderstorm activity

with much of the rainfall occurring in late afternoon or

early evening when the atmosphere is most susceptibleto thunderstorms. While precipitation is profuse in allmonths, variations do occur in response to the precise loca-tion of the ITCZ—drier months result when the ITCZmoves away from the region in question. Other zones ofAf climate are found beyond the usual margins of ITCZactivity, in coastal Madagascar, southeast coastal Brazil,

and much of Central America and western Colombia, where trade winds blow onshore all year to encountercoastlines backed by mountain barriers that stimulate theformation of precipitation as warm, moist tropical air isforced to ascend and cool. Some of these regions also mayreceive precipitation from tropical disturbances, includ-ing hurricanes.

Tropical Monsoon and

Trade-Wind Littoral Climates (Am)

These climates resemble the Af in most characteristics, with small annual temperature ranges, high tempera-tures, and plentiful precipitation (often more than Afclimates in annual total). They differ from the latter,

however, in that they exhibit a short dry season, usuallyin the low-sun (“winter”) season. The highest tempera-tures generally occur at the end of this clear spell. Two



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distinct processes can give rise to Am climate types. Thelargest areas, mostly in southern and southeastern Asia,result from the Asian monsoon circulation that bringsconvective and orographic precipitation in the summer

when warm, moist, maritime tropical air moves overland to converge into the low-pressure zone north of theHimalayas. In winter, by contrast, cool, dry air divergesout of the Siberian anticyclone to the north, bring-ing a cooler, drier, and clearer period of variable length.In the Americas and in Africa, Am climates are of the

trade-wind variety. These areas receive precipitation onnarrow coastal strips through orographic effects as themoist air of the trade winds ascends mountain chains.Seasonal migrations and changes in the intensity of these

winds give rise to short, moderately dry seasons. Summer


The sun sets on a savanna in the African country of Kenya. © Digital Vision/ Getty Images



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precipitation may be enhanced by tropical disturbancestraveling in the trade winds.

Tropical Wet-Dry Climate (Aw)

This climate has distinct wet and dry seasons, with mostof the precipitation occurring in the high-sun (“summer”)season. Total amounts of rainfall are less than in the previ-ous two climate types (50–175 cm [20–69 inches]), most of

which occur in convectional thunderstorms. The dry sea-son is longer than in the Am climates and becomes

progressively longer as one moves poleward through theregion. Temperatures are high throughout the year butshow a greater range than Af and Am climates (19–20 °C[66–68 °F] in winter and 24–27 °C [75–81 °F] in summer).Throughout most of the region, the cause of the seasonalcycle is the shift in the tropical circulation throughout the

year. During the high-sun season, the ITCZ moves pole-

ward and brings convergent and ascending air to theselocations, which stimulates convective rainfall. Duringthe low-sun season, the convergence zone moves offto the winter hemisphere and is replaced by the peripheryor core of the subtropical anticyclone, with its subsiding,stable air resulting in a period of dry, clear weather,the intensity and length of which depend on latitude. The

subtropical anticyclone occurs in the descending portionof the Hadley cell.

Type B Climates

Arid and semiarid climates cover about a quarter of Earth’sland surface, mostly between 50° N and 50° S, but they are

mainly found in the 15–30° latitude belt in both hemi-spheres. They exhibit low precipitation, great variabilityin precipitation from year to year, low relative humidity,



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high evaporation rates (when water is available), clearskies, and intense solar radiation.

Tropical and Subtropical Desert Climate (BWh,Part of BWk)

Most of Earth’s tropical, true desert (BW) climates occurbetween 15° and 30° latitude, at the poleward end of theHadley cell circulation. These regions are dominated inall months by the subtropical anticyclone, with itsdescending air, elevated inversions, and clear skies. This is

an atmospheric environment that inhibits precipitation.The most extreme arid areas also are far removed fromsources of moisture-bearing winds in the interiors of con-tinents and are best developed on the western sides ofcontinents, where the subtropical anticyclone shows itsmost intense development. An exception to the generaltendency for aridity to be associated with subsidence is in

the so-called Horn of Africa region, where the dryness ofSomalia is caused more by the orientation of the land-mass in relation to the atmospheric circulation. Both thehigh- and low-sun monsoonal winds blow parallel tothe coast, so that moisture-laden maritime air can pene-trate over land only infrequently. In most low-latitudedeserts, cloud cover is uncommon (fewer than 30 days per

year have clouds in some areas). Precipitation amountsare mostly in the range 0–25 cm (0–10 inches), althoughthe unreliability of precipitation is more significantthan the small totals. Average figures have little meaning;a location with a 10-year mean of 5 cm (2 inches), forexample, might have received 50 cm (about 20 inches) inone year as a result of an unusual intrusion of moist air,

followed by nine years with no measurable precipitation.Temperatures are high, with monthly means in therange 21–32 °C (70–90 °F). Daily temperature variations




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are extreme. Ranges of 35 °C (63 °F) are not unknown whendaytime maxima in excess of 40 °C (104 °F) are followed bya rapid nocturnal temperature drop brought about by thelimited capacity of the dry, cloudless desert air to emitinfrared radiation to the ground to offset radiation loss

from the surface at night. The highest air temperaturesrecorded on Earth have been in the BWh regions; forexample, in shaded, well-ventilated locations, DeathValley in the western United States has reached 57 °C (135°F), while al-‘Azızıyah in Libya has had a recorded high of58 °C (136 °F). Actual surface temperatures may reach 82 °C(180 °F) on dry sand under intense sunshine.

An interesting variant of tropical and subtropical des-erts are the so-called West Coast Desert areas found onthe western coastal margins of the regions discussed above

Kerzaz oasis on Wadi Saoura, western Sahara, Alg. Victor Englebert



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(e.g., in the Sonoran Desert of North America, the Peruand Atacama deserts of South America, and the Sahara[Moroccan part] and Namib deserts of Africa). Theseareas are much cooler than their latitude would suggest(monthly mean temperatures of only 15–21 °C [59–70 °F]),and parts are classified as BWk in Köppen’s scheme. Thecooling results from airflow off adjacent coastal waters

where upwelling of the ocean gives rise to cold currents.Deserts of this sort are subject to frequent fog and low-level clouds; yet they are extremely arid. Some parts of the

Atacama Desert, for example, have recorded no precipita-tion for 20 years.

Tropical and Subtropical Steppe Climate (BSh)

The low-latitude semiarid (or steppe) climate occurs pri-marily on the periphery of the true deserts treated above.It is transitional to the Aw climate on the equatorward

side (showing a summer rainfall maximum associated withthe ITCZ and a small annual temperature range) andto the Mediterranean climate on its poleward margin(with a cooler, wetter winter resulting from the higher lati-tude and mid-latitude frontal cyclone activity). Annualprecipitation totals are greater than in BW climates (38– 63 cm [15–25 inches]). Yearly variations in amount are not

as extreme as in the true deserts but are nevertheless large.Mid-Latitude Steppe and

Desert Climate (BSk, Part of BWk)

Although these climates are contiguous with the tropi-cal dry climates of North and South America and ofcentral Asia, they have different origins. Cool true des-

erts extend to 50° latitude and cool steppes reach nearly60° N in the Canadian Prairies, well beyond the limitsof the subtropical anticyclone. These climates owe their




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Mid-latitude grasslands and steppes, like this one in Kansas, U.S., are typically found in areas far from sources of moist, maritime air. © MedioImages/ Getty Images

origins to locations deep within continental interiors,far from the windward coasts and sources of moist,maritime air. Remoteness from sources of water vapouris enhanced in some regions (e.g., the Great Plains

of the United States) by mountain barriers upwind.Temperature conditions are extremely variable, withannual means decreasing and annual ranges increasingpoleward. In the higher latitudes, winters are severelycold, with meager precipitation (much of it in the formof snow) associated with polar and arctic air masses.Summer precipitation is more often convective, arriving

in the form of scattered thunderstorm activity broughtabout by irregular incursions of moist air. Both BWkand BSk climates in the mid-latitudes owe their origins



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to these mechanisms, but the steppe type tends to belocated peripheral to the true desert, either adjacent tothe moister C and D climates or at the poleward extentof the range, where reduced evaporation under coolerconditions makes more of the scarce precipitation avail-able as soil moisture for plant growth.

Type C and D Climates

Through a major portion of the middle and high latitudes

(mostly from 25° to 70° N and S) lies a group of climates clas-sified within the Köppen scheme as C and D types. Mostof these regions lie beneath the upper-level, mid-latitude

westerlies throughout the year, and it is in the seasonal variations in location and intensity of these winds and theirassociated features that the explanation of their climaticcharacter must be sought. During summer, the polar-front

and its jet stream move poleward, and air masses of tropicalorigin are able to extend to high latitudes. During winter,as the circulation moves equatorward, tropical air retreatsand cold polar outbreaks influence weather, even withinthe subtropical zone. The relative frequency of these airmasses of different origins varies gradually from low tohigh latitude and is largely responsible for the observed

temperature change across the belt (which is most markedin winter). The air masses interact in the frontal systemscommonly found embedded within the traveling cyclonesthat lie beneath the polar-front jet stream. Ascent inducedby convergence into these low-pressure cells and by upliftat fronts induces precipitation, the main location of whichshifts with the seasonal circulation cycle. Other important

sources of precipitation are convection, mainly in tropi-cal air, and forced uplift at mountain barriers. Monsooneffects modify this general pattern, while the subtropical





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anticyclone plays a role in the explanation of climate onthe western sides of the continents in the subtropics.

Humid Subtropical Climate (Cfa, Cwa)

These climates are found on the eastern sides of the con-tinents between 20° and 35° N and S latitude. Most showa relatively uniform distribution of precipitation through-out the year (the Cfa types), with totals in the range 75–150cm (30–59 inches). In summer, these regions are largelyunder the influence of moist, maritime airflow from the

western side of the subtropical anticyclonic cells over low-latitude ocean waters. Temperatures are high; the warmestmonths generally average about 27 °C (81 °F), with meandaily maxima from 30 °C to 38 °C (86 °F to 100 °F) and

warm, oppressive nights. Summers are usually somewhat wetter than winters, with much of the rainfall coming fromconvectional thunderstorm activity; tropical cyclones

also enhance warm-season rainfall in some regions. Thecoldest month is usually quite mild (5–12 °C [41–54 °F]),although frosts are not uncommon, and winter precipita-tion is derived primarily from frontal cyclones along thepolar front. In North America, the spring and early sum-mer seasons, when the front begins its northward return,are notorious for the outbreak of tornadoes associated

with frontal thunderstorms along the zone of interactionbetween tropical and polar air. In eastern and southernAsia, the monsoon influence results in a modified humidsubtropical climate (Cwa) that has a clearly defined dry

winter when air diverges from the Siberian anticyclone,and the polar front and cyclone paths are deflected aroundthe region. These areas generally lie on the poleward side

of Am and Aw climates and exhibit a somewhat largerannual temperature range than Cfa types. Winters aresunny and rather cool. Annual precipitation totals average



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about 100 cm (39 inches) but vary from 75 to 200 cm (30to 79 inches).

Mediterranean Climate (Csa, Csb)

Between about 30° and 45° latitude on the western sidesof the continents is found a series of climates that showthe unusual combination of hot, dry summers and cool,

wet winters. Poleward extension and expansion of thesubtropical anticyclonic cells over the oceans bring sub-siding air to the region in summer, with clear skies and

high temperatures. When the anticyclone moves equatorward in winter, it is replaced by traveling, frontal cyclones with their attendant precipitation. Annual temperatureranges are generally smaller than those found in the Cfaclimates, since locations on the western sides of conti-nents are not well positioned to receive the coldest polarair, which develops over land rather than over the ocean.

Mediterranean climates also tend to be drier than humidsubtropical ones, with precipitation totals ranging from 35to 90 cm (14 to 35 inches); the lowest amounts occur ininterior regions adjacent to the semiarid steppe climates.Some coastal locations (e.g., southern California in the

western United States) exhibit relatively cool summerconditions and frequent fogs where cold offshore currents

prevail. Only in Europe, where the latitude for this climatetype fortuitously corresponds to an ocean basin (that ofthe Mediterranean, from which this climate derives itsname), does this climate type extend eastward away fromthe coast for any significant distance.

Marine West Coast Climate (Cfb, Cfc)

Poleward of the Mediterranean climate region on the western sides of the continents, between 35° and 60° Nand S latitude are regions that exhibit ample precipitation





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in all months. Unlike their equatorward neighbours, theseareas are located beyond the farthest poleward extent ofthe subtropical anticyclone, and they experience the mid-latitude westerlies and traveling frontal cyclones all year.Precipitation totals vary somewhat throughout the year inresponse to the changing location and intensity of thesestorm systems, but annual accumulations generally rangefrom 50 to 250 cm (20 to 98 inches), with local totals exceed-ing 500 cm (197 inches) where onshore winds encountermountain ranges. Not only is precipitation plentiful but

it is also reliable and frequent. Many areas have rainfallmore than 150 days per year, although the precipitationis often of low intensity. Fog is common in autumn and

winter, but thunderstorms are infrequent. Strong gales with high winds may be encountered in winter. Theseare equable climates with few extremes of temperature.Annual ranges are rather small (10–15 °C or [50–59 °F]),

about half those encountered farther to the east in thecontinental interior at the same latitude. Mean annual tem-peratures are usually 7–13 °C (45–55 °F) in lowland areas, the

winters are mild, and the summers are relatively moderate,rarely having monthly temperatures above 20 °C (68 °F).

In North and South America, Australia, and NewZealand, north–south mountain ranges backing the west

coasts of the landmasses at these latitudes confine themarine west coast climate to relatively narrow coastalstrips (but enhance precipitation). By contrast, in Europethe major mountain chains (the Alps and Pyrenees) runeast–west, permitting Cfb and Cfc climates to extendinland some 2,000 km (about 1,250 miles) into easternGermany and Poland.

Humid Continental Climate (Dfa, Dfb, Dwa, Dwb)

The D climates are primarily Northern Hemispheric phe-nomena, since landmasses are absent at the significant



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Like much of western Russia and parts of eastern Europe, the Midwest resideswithin the humid continental climate zone. One example of a landscape foundwithin this climate type is shown by this photograph of a dairy farm in south-central Wisconsin. ©Robert Frerck/Odyssey Productions

latitudes in the Southern Hemisphere. The humid conti-nental subgroup occupies a region between 30° and 60° Nin central and eastern North America and Asia in themajor zone of conflict between polar and tropical airmasses. These regions exhibit large seasonal temperaturecontrasts with hot summers and cold winters. Precipitationtends to be ample throughout the year in the Df section,being derived both from frontal cyclones and, in summermonths, from convectional showers when maritime tropi-cal air pushes northward behind the retreating polar front.

Many areas show a distinct summer precipitation maxi-mum because of this convective activity, although moreuniform patterns are not uncommon. Severe thunder-storms and tornadoes are an early summer occurrence





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when the polar front is in the southern margin of the Dfaregion. Winter precipitation often occurs in the form ofsnow, and a continuous snow cover is established for fromone to four months in many parts of the region, especiallyin the north. This snow often arrives in conjunction withhigh winds from an intense frontal cyclone, giving rise to ablizzard.

Winters tend to be cold but are subject to occasionalfrigid or mild spells brought about by periodic incursionsof arctic or tropical air. Indeed the changeable nature of

weather in all seasons is a characteristic feature of the cli-mate, especially in such areas as the eastern United Statesand Canada where there are few topographic barriers tolimit the exchange of air masses between high and lowlatitudes. Mean temperatures are typically below freezingfrom one to several months, and the frost-free season var-ies from fewer than 150 to 200 days per year. Annual

precipitation totals range from 50 to 125 cm (about 20 to50 inches), with higher amounts in the south of the regionand in the uplands.

In eastern Asia (Manchuria and Korea), a monsoonal variant of the humid continental climate (Dwa, Dwb)occurs. This climate type has a pronounced summer pre-cipitation maximum and a cold, dry winter dominated by

continental polar air diverging out of the nearby Siberiananticyclone.

Continental Subarctic Climate

(Dfc, Dfd, Dwc, Dwd)

North of the humid continental climate, from about50° to 70° N, in a broad swath extending from Alaska

to Newfoundland in North America and from north-ern Scandinavia to Siberia in Eurasia, lie the continentalsubarctic climates. These are regions dominated by the



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Boreal coniferous forest dominated by spruce trees ( Picea ). Boreal conif-erous forests are evergreen coniferous forests that often grow just south ofthe tundra in the Northern Hemisphere where winters are long and cold

and days are short. In North America the boreal forest stretches from Alaska across Canada to Newfoundland; it stops just south of the northernCanadian border. The vast taiga of Asia extends across Russia into north-

eastern China and Mongolia. In Europe it covers most of Finland, Sweden, Norway, and regions in the Scottish Highlands. Erwin & Peggy Bauer/ Bruce Coleman Ltd.

winter season, a long, bitterly cold period with short, cleardays, relatively little precipitation (mostly in the form ofsnow), and low humidity. In Asia the Siberian anticyclone,the source of continental polar air, dominates the interior,and mean temperatures 40–50 °C (40–58 °F) below freez-ing are not unusual. The North American representativeof this climate is not as severe but is still profoundly cold.Mean monthly temperatures are below freezing for six toeight months, with an average frost-free period of only





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50–90 days per year, and snow remains on the ground formany months. Summers are short and mild, with long daysand a prevalence of frontal precipitation associated withmaritime tropical air within traveling cyclones. Meantemperatures in summer only rarely exceed 16 °C (61 °F),except in interior regions where values near 25 °C (77 °F)are possible. As a result of these temperature extremes,annual temperature ranges are larger in continental sub-arctic climates than in any other climate type on Earth,up to 30 °C (54 °F) through much of the area and more

than 60 °C (108 °F) in central Siberia, although coastalareas are more moderate. Annual precipitation totals aremostly less than 50 cm (about 20 inches), with a concen-tration in the summer. Higher totals, however, occur inmarine areas near warm ocean currents. Such areas alsoare generally somewhat more equable and may be des-ignated marine subarctic climates. Areas with a distinct

dry season in winter, which results in the Köppen climatetypes Dwc and Dwd, occur in eastern Siberia, both in theregion where the wintertime anticyclone is establishedand in the peripheral areas subject to dry, divergent air-flow from it.

Type E Climates

Köppen’s type E climates are controlled by the polar andarctic air masses of high latitudes (60° N and S and higher).These climates are characterized by low temperaturesand precipitation and by a surprisingly great diversity ofsubtypes.

Tundra Climate (ET)

Tundra climates occur between 60° and 75° of latitude,mostly along the Arctic coast of North America and



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Thawed surface of the permafrost on the tundra in summer, Taymyr Peninsula,

Siberia. © John Hartley/NHPA

Eurasia and on the coastal margins of Greenland. Meanannual temperatures are below freezing and annualranges are large (but not as large as in the adjacent conti-nental subarctic climate). Summers are generally mild,

with daily maxima from 15 to 18 °C (59 °F to 64 °F),although the mean temperature of the warmest monthmust be less than 10 °C (50 °F). Days are long (a result ofthe high latitude), but they are often cloudy. The snowcover of winter melts in the warmer season (though inplaces with mean annual temperatures of –9 °C [16 °F] orless the ground at depth remains permanently frozen as

permafrost); however, frosts and snow are possible inany month. Winters are long and cold (temperaturesmay be below 0 °C [32 °F] for 6 to 10 months), especially





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in the region north of the Arctic Circle where, for atleast one day in the year, the Sun does not rise. Winterprecipitation generally consists of dry snow, with sea-sonal totals less than in the summer when cyclonicstorms that develop along the boundary between theopen sea and sea ice yield rainfall. Typical annual totalsare less than 35 cm (about 14 inches), but a range from 25to 100 cm (10 to 39 inches) is possible, with higher totalsin upland areas.

Snow and Ice Climate (EF)This climate occurs poleward of 65° N and S latitudeover the ice caps of Greenland and Antarctica and overthe permanently frozen portion of the Arctic Ocean.Temperatures are below freezing throughout the year,and annual temperature ranges are large but again notas large as in the continental subarctic climates. Winters

are frigid, with mean monthly temperatures from –20°C to –65 °C (–4 °F to –85 °F); the lowest temperaturesoccur at the end of the long polar night. The EF climateholds the distinction for the lowest recorded tempera-tures at Earth’s surface: the Vostok II research stationin Antarctica holds the record for the lowest extremetemperature (–89.2 °C [–129 °F]), while the Plateau

Station has the lowest annual mean (–56 °C [–69 °F]).Daily temperature variations are very small, becausethe presence of snow and ice at the surface refrigeratesthe air. Precipitation is meager in the cold, stable air (inmost cases, 5 to 50 cm [2 to 20 inches]), with the largestamounts occurring on the coastal margins. Most of thisprecipitation results from the periodic penetration of a

cyclone into the region, which brings snow and ice pelletsand, with strong winds, blizzards. High winds also occurin the outer portions of the Greenland and Antarctic EF



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climates, where cold, dense air drains off the higher, cen-tral sections of the ice caps as katabatic winds.

Type H Climates

The major highland regions of the world (the Cascades,Sierra Nevadas, and Rockies of North America, theAndes of South America, the Himalayas and adjacentranges and the Tibetan Highlands [or Plateau] of Asia,the eastern highlands of Africa, and the central portions

of Borneo and New Guinea) cannot be classified real-istically at this scale of consideration, since the effectsof altitude and relief give rise to myriad mesoclimatesand microclimates. This diversity over short horizontaldistances is unmappable at the continental scale. Verylittle of a universal nature can be written about suchmountain areas except to note that, as a rough approxi-

mation, they tend to resemble cooler, wetter versions ofthe climates of nearby lowlands in terms of their annualtemperature ranges and seasonality of precipitation.Otherwise, only the most general characteristics maybe noted.

With increasing height, temperature, pressure, atmo-spheric humidity, and dust content decrease. The reduced

amount of air overhead results in high atmospherictransparency and enhanced receipt of solar radiation(especially of ultraviolet wavelength) at elevation. Altitudealso tends to increase precipitation, at least for the first4,000 metres (about 13,100 feet). The orientation ofmountain slopes has a major impact on solar radiationreceipt and temperature and also governs exposure to

wind. Mountains can have other effects on the wind cli-mate; valleys can increase wind speeds by “funneling”regional flows and may generate mesoscale mountain- and





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valley-wind circulations as well. Cold air also may drainfrom higher elevations to create “frost pockets” in low-lying valleys. Furthermore, mountains can act as barriersto the movement of air masses, can cause differences inprecipitation amounts between windward and leewardslopes (the reduced precipitation on and downwind fromlee slopes is called a rain shadow), and, if high enough, cancollect permanent snow and ice on their peaks and ridges;the snow line varies in elevation from sea level in the sub-arctic to about 5,500 metres (about 18,000 feet) at 15–25°

N and S latitude.



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The connection between climate and life arises froma two-way exchange of mass and energy between the

atmosphere and the biosphere. In Earth’s early history,before life evolved, only geochemical and geophysicalprocesses determined the composition, structure, anddynamics of the atmosphere. Since life evolved on Earth,biochemical and biophysical processes have played a role

in the determination of the composition, structure, anddynamics of the atmosphere. Humans, Homo sapiens , areincreasingly shouldering this role by mediating interac-tions between the biosphere and the atmosphere.

The living organisms of the biosphere use gases from,and return “waste” gases to, the atmosphere, and thecomposition of the atmosphere is a product of this gas

exchange. It is very likely that, prior to the evolution oflife on Earth, 95 percent of the atmosphere was made upof carbon dioxide, and water vapour was the second mostabundant gas. Other gases were present in trace amounts.This atmosphere was the product of geochemical and geo-physical processes in the interior of Earth and was mediatedby volcanic outgassing. It is estimated that the great mass

of carbon dioxide in this early atmosphere gave rise to anatmospheric pressure 60 times that of modern times. Todayonly about 0.035 percent of Earth’s atmosphere is carbondioxide. Much of the carbon dioxide present in Earth’s firstatmosphere has been removed by photosynthesis, chemo-synthesis, and weathering. Currently, most of the carbondioxide now resides in Earth’s limestone sedimentary

rocks, in coral reefs, in fossil fuels, and in the living compo-nents of the present-day biosphere. In this transformation,the atmosphere and the biosphere coevolved through con-tinuous exchanges of mass and energy.

CLIMATE AND LIFE



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The planet Earth. NASA

Biogenic gases are gases critical for, and produced by,living organisms. In the contemporary atmosphere, theyinclude oxygen, nitrogen, water vapour, carbon dioxide,carbon monoxide, methane, ozone, nitrogen dioxide, nitricacid, ammonia and ammonium ions, nitrous oxide, sulfurdioxide, hydrogen sulfide, carbonyl sulfide, dimethyl sul-fide, and a complex array of non-methane hydrocarbons.

Of these gases, only nitrogen and oxygen are not “green-house gases.” Added to this roster of biogenic gases is amuch longer list of human-generated gases from industrial,commercial, and cultural activities that reflect the diver-sity of the human enterprise on Earth.

THE GAIA HYPOTHESIS

The notion that the biosphere exerts important controlson the atmosphere and other parts of the Earth system



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7 Climate and Life 7

has increasingly gained acceptance among earth and eco-system scientists. While this concept has its origins in the

work of American oceanographer Alfred C. Redfield inthe mid-1950s, it was English scientist and inventor JamesLovelock that gave it its modern currency in the late1970s. Lovelock initially proposed that the biospherictransformations of the atmosphere support the biospherein an adaptive way through a sort of “genetic group selec-tion.” This idea generated extensive criticism and spawneda steady stream of new research that has enriched the

debate and advanced both ecology and environmental sci-ence. Lovelock called his idea the “Gaia hypothesis” anddefined Gaia as

a complex entity involving Earth’s biosphere, atmosphere,

oceans, and soil; the totality constituting a feedback of cyber-

netic systems which seeks an optimal physical and chemical

environment for life on this planet.

The Greek word Gaia, or Gaea, meaning “MotherEarth,” is Lovelock’s name for Earth, which is envi-sioned as a “superorganism” engaged in planetarybiogeophysiology. The goal of this superorganism is toproduce a homeostatic, or balanced, Earth system. The

scientific process of research and debate will eventuallyresolve the issue of the reality of the “Gaian homeo-static superorganism,” and Lovelock has since revisedhis hypothesis to exclude goal-driven genetic groupselection. Nevertheless, it is now an operative norm incontemporary science that the biosphere and the atmo-sphere interact in such a way that an understanding of

one requires an understanding of the other. Furthermore,the reality of two-way interactions between climate andlife is well recognized.




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THE EVOLUTION OF LIFE ANDTHE ATMOSPHERE

Life on Earth began at least as early as 3.5 billion years agoduring the middle of the Archean Eon (about 4 billion to2.5 billion years ago). It was during this interval that lifefirst began to exercise certain controls on the atmosphere.The atmosphere’s prebiological state is often character-ized as being rich in water vapour and carbon dioxide.Though some nitrogen was also present, little if any oxy-

gen was available. Chemical reactions with hydrogensulfide, hydrogen, and reduced compounds of nitrogenand sulfur precluded any but the shortest lifetime for freeoxygen in the atmosphere. As a result, life evolved in anatmosphere that was reducing (high hydrogen content)rather than oxidizing (high oxygen content). In additionto their chemically reducing character, the predominant

gases of this prebiotic atmosphere, with the exception ofnitrogen, were largely transparent to incoming sunlightbut opaque to outgoing terrestrial infrared radiation. As aresult, these gases are called, perhaps improperly, green-house gases because they are able to slow the release ofoutgoing radiation back into space.

In the Archean Eon, the Sun produced as much as

25 percent less light than it does today; however, Earth’stemperature was much like that of today. This is possiblebecause the greenhouse gas–rich Archean atmosphere

was effective in retarding the loss of terrestrial radiationto space. The resulting long residence time of energy

within the Earth-atmosphere system resulted in a warmeratmosphere than would have been possible otherwise.

The average temperature of Earth’s surface in the earlyArchean Eon was warmer than the modern global aver-age. It was, according to some sources, probably similar



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to temperatures found in today’s tropics. Depending onthe amount of nitrogen present during the Archean Eon,it has been suggested that the atmosphere may haveheld more than 1,000 times as much carbon dioxide as itdoes today.

Archean organisms included photosynthetic andchemosynthetic bacteria, methane-producing bacteria,and a more primitive group of organisms now called the“Archaea” (a group of prokaryotes more related to eukary-otes than to bacteria and found in extreme environments).

Through their metabolic processes, organisms of theArchean Eon slowly changed the atmosphere. Hydrogenrose from trace amounts to about 1 part per million (ppm)of dry air. Methane concentrations increased from nearzero to about 100 ppm. Oxygen increased from near zeroto 1 ppm, whereas nitrogen concentrations rose to encom-pass 99 percent of all atmospheric molecules excluding

water vapour. Carbon dioxide concentrations decreasedto only 0.3 percent of the total; however, this was nearly10 times the current concentration. The composition ofthe atmosphere, its radiation budget, its thermodynam-ics, and its fluid dynamics were transformed by life fromthe Archean Eon.

American geochemist Robert Garrels calculated that,

in the absence of life and given the burial rate of carbonin rocks, oxygen would be unavailable to form water, andfree hydrogen would be lost to space. Without the pres-ence of life and compounded by this loss of hydrogen,there would be no oceans, and Earth would have becomemerely a dusty planet by the middle of the Archean Eon.By the end of the Archean Eon 2.5 billion years ago, both

the pigment chlorophyll and photosynthetic organismshad evolved such that the production of oxygen increasedrapidly. The atmosphere became transformed from a





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reducing atmosphere with carbon dioxide, limited oxygen, and anaerobic organisms (that is, life-forms that donot require oxygen for respiration) in control to one withan oxidizing atmosphere that was rich in oxygen, poor incarbon dioxide, and dominated by aerobic organisms (thatis, life-forms requiring oxygen for respiration).

With the decline in carbon dioxide and a rise in oxygen,the greenhouse warming capacity of Earth’s atmosphere

was sharply reduced; however, this happened over a periodof time when the energy produced by the Sun increased

systematically. These compensating changes resulted in arelatively constant planetary temperature over much ofEarth’s history.

THE ROLE OF THE BIOSPHERE INTHE EARTH-ATMOSPHERE SYSTEM

Although Earth’s biosphere is primarily known as theenvironmental sphere containing life, much of it overlaps

with Earth’s atmosphere. In addition, the biosphere alsoplays significant roles in Earth’s energy budget and incycling biogenic atmospheric gases.

The Biosphere and Earth’s Energy Budget

Biogenic gases in the atmosphere play a role in thedynamics of Earth’s planetary radiation budget, thethermodynamics of the planet’s moist atmosphere, and,indirectly, the mechanics of the fluid flows that are Earth’splanetary wind systems. In addition, human cultural andeconomic activities add a new dimension to the relation-

ship between the biosphere and the atmosphere. Whilehumans are biologically trivial compared with bacteriain the exchange of gases with the atmosphere, chemical



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compounds produced from human industrial activitiesand other economic enterprises are changing the gaseouscomposition of the atmosphere in climatically significant

ways. The largest changes involve the harvesting of ancientcarbon stores. This organic material has been transformedinto fossil fuels (coal, petroleum, natural gas, and others)by geologic processes acting upon the remains of plantsand animals over many millions of years. Different formsof carbon may be burned and thus used as energy sources.In so doing, organic carbon is converted into carbon diox-

ide. Additionally, humans are also burning trees, grasses,and other biomass for cooking purposes and clearing theland for agriculture and other activities. The combinationof burning both fossil fuels and biomass is enriching theatmosphere with carbon dioxide and adding to the essen-tial reservoir of greenhouse gases.

Earth’s atmosphere is largely transparent to sunlight.

Of the sunlight absorbed by the entire Earth-atmospheresystem, about one-third is absorbed by the atmosphereand two-thirds by Earth’s surface. Sunlight is absorbed bythe molecules of the atmosphere, by cloud droplets, andby dust and debris. Though oxygen and nitrogen makeup nearly 99 percent of the atmosphere, these diatomicmolecules do not vibrate in a way that permits them to

absorb terrestrial radiation. They are largely transparentto outgoing terrestrial radiation as well as to incomingsolar radiation.

Over the continents, the surface cover of vegetationis the principal absorbing medium of Earth’s surface,although other surfaces such as bare rock, sand, and wateralso absorb solar radiation. At night, absorption at the

surface (that is, below 1.2 metres [4 feet]) is reradiated,in the form of long-wave infrared radiation, away fromEarth’s surface back toward space. Most of this infrared





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radiation is absorbed by the principal biogenic trace gasesof the atmosphere—the so-called greenhouse gases: water

vapour, carbon dioxide, and methane. Without these biogenic greenhouse gases, Earth would be 33 °C (59 °F) colderon average than it is. A moderate-emission scenario fromthe 2007 Intergovernmental Panel on Climate Change(IPCC) report predicts that the continued addition of

greenhouse gases from fossil fuels will increase the aver-age global temperature by between 2.3 and 4.3 °C (4.1 and7.7 °F) over the next century. Other scenarios, predicting

greater greenhouse gas emissions, forecast even greater global warming.

AVERAGE COMPOSITION OF THEATMOSPHERE

gas composition

by volume(ppm)*

composition

by weight(ppm)*

total mass

(1020 g)

Nitrogen 780,900 755,100 38.648

Oxygen 209,500 231,500 11.841

Argon 9,300 12,800 0.655

Carbon dioxide 386 591 0.0299

Neon 18 12.5 0.000636

Helium 5.2 0.72 0.000037

Methane 1.5 0.94 0.000043

Krypton 1.0 2.9 0.000146

Nitrous oxide 0.5 0.8 0.000040

Hydrogen 0.5 0.035 0.000002

Ozone** 0.4 0.7 0.000035

Xenon 0.08 0.36 0.000018*ppm = parts per million. **Variable, increases with height.



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THE CYCLING OF BIOGENICATMOSPHERIC GASES

The cycling of oxygen, nitrogen, water vapour, and carbondioxide, as well as the trace gases—methane, ammonia,

various oxides of nitrogen and sulfur, and non-methanehydrocarbons—between the atmosphere and the bio-sphere results in relatively constant proportions of thesecompounds in the atmosphere over time. Without thecontinuous generation of these gases by the biosphere,

they would quickly disappear from the atmosphere.The carbon cycle, as it relates to the biosphere, is sim-

ple in its essence. Inorganic carbon (carbon dioxide) isconverted to organic carbon (the molecules of life). Tocomplete the cycle, organic carbon is then converted backto inorganic carbon. Ultimately, the carbon cycle is pow-ered by sunlight as green plants and cyanobacteria

(blue-green algae) use sunlight to split water into oxygenand hydrogen and to fix carbon dioxide into organic car-bon. Carbon dioxide is removed from the atmosphere,and oxygen is added. Animals engage in aerobic respira-tion, in which oxygen is consumed and organic carbon isoxidized to manufacture inorganic carbon dioxide. Itshould be noted that chemosynthetic bacteria, which are

found in deep-ocean and cave ecosystems, also fix carbondioxide and produce organic carbon. Instead of using sun-light as an energy source, these bacteria rely on theoxidation of either ammonia or sulfur.

The carbon cycle is fully coupled to the oxygen cycle.Each year, photosynthesis fixes carbon dioxide andreleases 100,000 megatons of oxygen to the atmosphere.

Respiration by animals and living organisms consumesabout the same amount of oxygen and produces carbondioxide in return. Oxygen and carbon dioxide are thus



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The carbon cycle is the complex path that carbon follows through the atmosphere, oceans, soil, and plants and animals. Encyclopædia Britannica, Inc.



145

Green plants such as trees use carbon dioxide, sunlight, and water to cre-

ate sugars. Sugars provide the energy that makes plants grow. The processcreates oxygen, which people and other animals breathe. EncyclopædiaBritannica, Inc.

coupled in two linked cycles. On a seasonal basis, anenrichment of atmospheric carbon dioxide occurs in the

winter half of the year, whereas a drawdown of atmo-spheric carbon dioxide takes place during the summer.

The nitrogen cycle begins with the fixing of inorganicatmospheric nitrogen (N2 ) into organic compounds. Thesenitrogen-containing compounds are used by organismsand, through the process of denitrification, are convertedback to inorganic atmospheric nitrogen. Ammonia andammonium ions are the products of nitrogen fixation





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and may be incorporated into some organisms as organicnitrogen-containing molecules. In addition, ammoniumions may be oxidized to form nitrites, which can be fur-ther oxidized by nitrifying bacteria into nitrates. Thoughboth nitrites and nitrates may be used to make organicnitrogen-containing molecules, nitrates are especially use-ful for plant growth and are key compounds that supportboth terrestrial and aquatic food chains. Nitrates may alsobe denitrified by bacteria to produce nitrogen gas. Thisprocess completes the nitrogen cycle.

The nitrogen cycle is coupled to both the carbon andoxygen cycles. Seventy-eight percent of the gases of theatmosphere by volume is diatomic nitrogen. Diatomicnitrogen is the most stable of the nitrogen-containing

gases of the atmosphere. Only 300 megatons of nitrogenmust be produced each year by denitrifying bacteria toaccount for losses. These losses mostly occur during light-

ning discharges and during nitrogen-fixation activities byblue-green algae and nitrogen-fixing bacteria. (The latterare found in the root nodules of certain plants calledlegumes.) Nitrogen-fixing bacteria use atmospheric nitro-

gen to produce oxides of nitrogen. Denitrifying bacteriaconvert nitrates in soils and wetlands to nitrogen gas,

which is then returned to the atmosphere.

Nitrous oxide occurs in trace amounts (0.3 ppm) in theatmosphere. Between 100 and 300 megatons of nitrousoxide are produced by soil and marine bacteria each yearto maintain this concentration. In the atmosphere, nitrousoxide is short-lived because it is quickly broken down byultraviolet light. Nitric oxide (NO), a minor contributorin the breakdown of stratospheric ozone, is the by-prod-

uct of this reaction.Like nitrous oxide, ammonia is also produced insoils and marine waters by bacteria and escapes into



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The nitrogen cycle. Encyclopædia Britannica, Inc.





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the atmosphere. Nearly 1,000 megatons are added to theatmosphere each year. Ammonia decreases the acidity ofprecipitation and serves as a nutrient for plants whenreturned to the land via precipitation.

There are six major sulfur-containing biogenic atmo-spheric gases that are part of the sulfur cycle. They includehydrogen sulfide, carbon disulfide (CS2 ), carbonyl sulfide(COS), dimethyl sulfide (DMS; C2H6S), dimethyl disulfide([CH3S]2 ), and methyl mercaptan (CH3SH). Sulfur dioxide(SO2 ) is an oxidation product of these sulfur gases, and

it is also added to the atmosphere by volcanoes, burningbiomass, and anthropogenic sources (i.e., smelting metalsand coal ignition). SO2 is removed from the atmosphereand returned to the biosphere in rainfall. The increasedacidity of rain and snow from anthropogenic additionsof SO2 and oxides of nitrogen is often referred to as “acidprecipitation.” The acidity of this precipitation and other

phenomena, such as “acid fog,” is partly cancelled by therelease of ammonia in the atmosphere.

The concentration of methane at any one time in theatmosphere is only about 1.7 ppm. Though only a trace

gas, it is highly reactive and plays a key role in the chemicalreactions that control the composition of the atmosphere.Methanogenic bacteria in wetland sediments decompose

organic matter and release 1,000 megatons of gaseousmethane to the atmosphere per year. In the lower atmo-sphere, methane reacts with oxygen to produce waterand carbon dioxide. Each year 2,000 megatons of oxygenare removed from the atmosphere by this mechanism.This loss of oxygen must be replaced by photosynthesis.Some methane reaches the upper stratosphere, where its

interaction with oxygen is a major source of upper strato-spheric moisture. Within wetlands, bacteria producemethyl halide compounds (methyl chloride [CH3Cl] and



149

methyl iodide [CH3I] gases), whereas these same methylhalides are produced in forests by fungi. These gases, uponreaching the stratosphere, regulate the production ofstratospheric ozone by contributing to its natural break-down. Without the continual production of methane by

methanogenic bacteria, the oxygen concentration of theatmosphere would increase by 1 percent in only 12,000 years. Dangerously high levels of oxygen in the atmo-sphere would greatly increase the incidence of wildfires. Ifthe oxygen concentration of Earth’s atmosphere rose fromits current concentration of 21 percent to 25 percent, evendamp twigs and grass would easily ignite. Non-methane

hydrocarbons of terrestrial origin are generally well mixedin the free atmosphere above the planetary boundarylayer (PBL). These organic particles weaken incoming


The branch of climatology that deals with the effects of the physicalenvironment on living organisms over an extended period of time iscalled bioclimatology. Although Hippocrates touched on these mat-ters 2,000 years ago in his treatise on Air, Waters, and Places, thescience of bioclimatology is relatively new. It developed into a signifi-cant field of study during the 1960s owing largely to a growing concernover the deteriorating environment.

Because almost every aspect of climate and weather has someeffect on living organisms, the scope of bioclimatology is almost limit-less. Certain areas are emphasized more than others, however, among

them studies of the influence of weather and climate on small plantorganisms and insects responsible for the development of plant, ani-mal, and human diseases; the influence of weather and climate onphysiological processes in normal healthy humans and their diseases;the influence of microclimate in dwellings and urban centres onhuman health; and the influence of past climatic conditions on thedevelopment and distribution of plants, animals, and humans.



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solar radiation as it passes through the atmosphere, andreductions of 1 percent have been recorded.

Biosphere Controls on the Structure ofthe Atmosphere

Because the biosphere plays a key role in the flux ofenergy from the surface to the atmosphere, it also con-tributes to the structure of the atmosphere. Three majorfluxes are important: the direct transfer of heat from the

surface to the atmosphere by conduction and convection(sensible heating), the energy flux to the atmosphere car-ried by water vapour via evaporation and transpirationfrom the surface (latent heat energy), and the flux of radi-ant energy from the surface to the atmosphere (infraredterrestrial radiation). These fluxes differ in the altitude at

which the heating of the air takes place and thus contrib-

ute to the thermal structuring of the atmosphere. Sensibleheating primarily warms the PBL of the atmosphere. Inmarine areas, the PBL occurs in the lowest 1 km (3,300feet); in heavily vegetated areas, the PBL occurs in thelowest 1 to 2 km (3,300 to 6,600 feet); and in arid regions,it occurs in the lowest 4 or 5 km (13,100 to 16,400 feet) ofthe atmosphere. In contrast, the latent heat of the atmo-

sphere is released when the water vapour is convertedinto cloud droplets by condensation. Heating by latentenergy release generally occurs above the PBL.

On the other hand, heating of the atmosphere by radi-ation from the surface depends on the density of theatmosphere and its water vapour content. Radiative heat-ing from the surface declines with increasing altitude. The

availability of water to evaporate from the surface limitsthe sensible heating of the air near the surface and solimits the maximum daytime surface air temperature.



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Biosphere Controls on the PlanetaryBoundary Layer

The top of the PBL can be visually marked by the eleva-tion of the base of the clouds. In addition, the PBL canalso be denoted by a thin layer of haze often seen by pas-sengers aboard airplanes during takeoff from airports.During the day, the air within the PBL is thoroughly mixedby convection induced by the heating of Earth’s surface.The thickness of the PBL depends on the intensity of this

surface heating and the amount of water evaporated intothe air from the biosphere. In general, the greater theheating of the surface, the deeper the PBL. Over deserts,the PBL may extend up to 4 or 5 km (13,100 or 16,400 feet)in altitude. In contrast, the PBL is less than 1 km (0.6 mile)thick over ocean areas, since little surface heating takesplace there because of the vertical mixing of water. The

wetter the air advected into the region and the greater theadditional water added by evaporation and transpiration,the lower the height of the top of the PBL. For every 1 °C(1.8 °F) increase in daily maximum surface temperature fora well-mixed PBL, the top of the PBL is elevated 100metres (about 325 feet). In New England forests during thedays following the spring leafing, it has been shown that

the top of the PBL is lowered to between 200 and 400metres (650 and 1,300 feet). By contrast, during themonths before the leafing out, the PBL thickens fromsolar heating as the sun rises higher in the sky and daylength increases.

If convective mixing of the air in the PBL is vigorous,convection currents may penetrate through the tempera-

ture inversion at the top of the PBL. The cooling of thelifting air initiates the condensation of water vapour andthe development of miniscule particles of liquid water





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called cloud droplets. The small clouds just above the PBLare known as planetary boundary layer clouds. Theseclouds scatter direct sunlight. As the ratio of diffuse sun-light to direct beam sunlight increases, greater levels ofphotosynthetic productivity are favoured in the biospherebelow. The result is a dynamic synergy between the atmo-sphere and biosphere.

The landscapes of most human-dominated ecosystemsare decidedly “patchy” in their geography. Cities, suburbs,fields, forests, lakes, and shopping centres both heat and

evaporate water into the air of the PBL according to thenature of the surfaces involved. Convection and the pros-pect of breaking through the top of the PBL vary markedlyacross such heterogeneous landscapes. These upward anddownward currents or vertical eddies within the PBLtransfer mass and energy upward from the surface. Thefrequency, timing, and strength of convective weather ele-

ments, including thunderstorms, vary according to thepatchiness of the land use and land cover pattern of thearea. In general, the greater the patchiness of the land-scape and the earlier the hour in the day, the more frequentand more intense these rain-producing systems become.

In the absence of an organized storm in the region, theair above the PBL sinks gently and the air below lifts. At

the top of the PBL, a small inversion, where temperaturesincrease with height, develops. This inversion essentiallybecomes a stable layer in the atmosphere. Emissions fromthe biosphere below are thus contained within the PBL andmay build up below this layer over time. Consequently, thePBL may become quite turbid, hazy, or filled with smog.

When the sinking from above is vigorous, the PBL

inversion grows in thickness. This situation has theeffect of hindering the development of thunderstorms, which depend on rapidly rising air. This often occurs over



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southern California, and thus the chance of thunder-storms forming there is small. Emissions from both thebiosphere and from anthropogenic activities accumulatein this part of the atmosphere, and pollution may build upto such an extent that health warnings may be required. Inlocations free of temperature inversions, convection pro-cesses are strong enough, particularly during the summermonths, that emissions are scavenged and quickly liftedby thunderstorms to regions high above the PBL. Often,acidic compounds from these emissions are returned to

the surface in the precipitation that falls.

Biosphere Controls on MaximumTemperatures by Evaporation andTranspiration

Solar radiation is converted to sensible and latent heat at

Earth’s surface. A change in sensible heat results in achange in the temperature of a medium, whereas energystored as latent heat is used to drive a process, such as aphase change in a substance from its liquid to its gaseousstate, and does not produce a change in temperature.Thus, the daily maximum surface temperature at a givenlocation is dependent on the amount of radiant energy

converted to sensible heat. Water available for evapora-tion increases latent heating by adding water vapour tothe atmosphere. As a result, relatively little energy remainsto heat the air, and thus the sensible heating of the air nearthe ground is minimized. In addition, daily maximumtemperatures are not as high in locations with stronglatent heating.

As day length increases from winter to summer, sensi-ble heating and maximum surface temperatures rise. Inthe U.S. Midwest, prior to the leafing out of vegetation





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in the springtime and the resulting rise in evaporation andtranspiration, sensible heating causes an average increasein maximum surface temperatures of only about 0.3 °C(0.5 °F) per day. The process of leaf production creates asurge in evaporation and transpiration and results inincreased latent heating and reduced sensible heating.After leafing, since most of the available thermal energy isused to convert liquid water to water vapour rather than toheat the air, the average day-to-day rise in daily maximumtemperatures is reduced to about 0.1 °C (0.2 °F) per day.

This effect extends upward through the atmosphere.Prior to leafing out, the one-kilometre-thick layer occur-ring between the 850-to-750-millibar pressure level (whichtypically occurs between 1,650 and 2,750 metres [5,400and 9,000 feet]) in the Midwest warmed at the rate of 0.1°C (0.2 °F) per day. Following leafing out, the warming ratefell to 0.02 °C (0.04 °F) per day. Scientists have used com-

puter models of the atmosphere to study the effect oftranspiration from vegetation on maximum surface airtemperatures. In these models, the variable controllingtranspiration by vegetation was “turned off,” and the char-acter of the resulting modeled climate was studied. Bysubtracting the effect of transpiration, temperatures incentral North America and on the other continents were

predicted to equilibrate at a very hot 45 °C (113 °F). Such warming is nearly realized in desert areas where moistureis unavailable for transpiration.

Biosphere Controls on MinimumTemperatures

During the late 1860s, British experimental physicist JohnTyndall, based on his studies of the infrared radiationabsorption by atmospheric gases, concluded that night-time minimum temperatures were dependent on the



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concentration of trace gases in the atmosphere. Of these gases, water vapour had the greatest impact. To emphasizethe significance of water vapour on decreases in air tem-perature during the night, he wrote that if all the water

vapour in the air over England was removed even for asingle night, it would be “attended by the destruction ofevery plant which a freezing temperature could kill.” As aresult, it follows that the greater the water content of theatmosphere, the lower the radiative loss of energy to thesky and the less the surface atmosphere is cooled. Thus,

locations with substantial amounts of water vapour expe-rience reduced nocturnal cooling.

Water vapour in the atmosphere also limits the extentto which temperatures fall at night. This limiting temper-ature is known as the dew point, which is defined as thetemperature at which condensation begins. Over NorthAmerica east of the 100th meridian (a line of longitude

traditionally dividing the moist eastern part of NorthAmerica from drier western areas), average nighttimeminimum temperatures are within a degree or two of thedew point temperature. Upon nocturnal cooling, the dewpoint is reached, condensation begins, and latent energyis converted to heat. Additional temperature falls areretarded by this release of heat to the atmosphere. A sig-

nificant fraction of the water in the atmosphere over thecontinents comes from the evaporation of water from soilsand the transpiration from vegetation. Transpired waterdirectly moderates temperature by increasing humid-ity and thus raising the dew point. As a consequence, theamount of outgoing terrestrial radiation released to spaceis reduced. This results in the elevation of the minimum

temperature of the air above what it would otherwise be.The effect of spring leafing on the buildup of humidityin the lower atmosphere has received the attention ofresearchers in recent years. In the late 1980s, American




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Graph of atmospheric vapour pressure and expected minimum temperature for 56 days prior to and following the average day of the leafing out of lilac plants, based on data compiled by M.D. Schwartz and T.R. Karl, 1990.Encyclopædia Britannica, Inc.

climatologists M.D. Schwartz and T.R. Karl used the

superimposed epoch method to study the climate beforeand after the leafing out of lilac plants in the spring in theU.S. Midwest. (This method uses time series data frommultiple locations, which can be compared to one anotherby adjusting each data set around the respective onsetdate of lilac blooming.) Prior to the average date of leafing,the atmospheric humidity (vapour pressure) is relatively

constant and minimum temperatures hover near freezing.At leafing, there is an abrupt increase in atmospherichumidity. Following leafing, daily minimum temperaturesalso increase abruptly. Although frosts are possible until



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June 10th in many parts of the Midwest, the chances offrost decline as the atmosphere is humidified.

Climate and Changes in the Albedo ofthe Surface

The amount of solar energy available at the surface forsensible and latent heating of the atmosphere dependson the albedo, or the reflectivity, of the surface. Surfacealbedos vary by location, season, and land cover type. The

albedo of unvegetated ground devoid of snow ranges from0.1 to 0.6 (10 to 60 percent), while the albedo of fully for-ested lands ranges from 0.08 to 0.15. An increase of 0.1in regional albedo has been associated with a 20 percentdecline in rainfall events connected with thunderstorms.Equivalent reductions in both evaporation and transpira-tion have also been reported in areas with sudden increases

in albedo.The greatest changes in albedo occur in regions under-

going desertification and deforestation. Depending onthe albedo of the underlying soil, reductions in vegetativeland cover may give rise to albedo increases of as much as0.2. Model studies of the vegetative zone known as theSahel in Africa reveal that albedo increased from 0.14 to

0.35 due to desertification occurring during the 20th cen-tury. This coincided with a 40 percent decrease in rainfall.In addition, it is likely that the clearing of forests and prai-ries for agricultural crops over the past several hundred

years has altered the albedo of extensive regions of themiddle latitudes.

Contemporary agricultural practices give rise to large

variations in albedo from season to season as the landpasses through the cycle of tilling, planting, crop growth,and harvest. At larger scales, an agricultural mosaic oftenemerges as each different plot of ground is covered by





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plantings of a single species. Viewed from the air, land-scapes in the middle latitudes appear as a heterogeneousmix of forests, grasslands, meadows, water bodies, farm-lands, wetlands, and urban types. The resultant patchinessin the landscape produces a patchiness in surface albedo.The mosaic of land use types creates a mix in the fluxes ofsensible and latent heat to the atmosphere. Such changesto the heat flux have been shown to cause changes in thetiming, intensity, and frequency of summer thunderstorms.

The Effect of Vegetation Patchiness onMesoscale Climates

The establishment of vegetation bands or patches 50to 100 km (30 to 60 miles) in width in semiarid regionscould increase atmospheric convection and precipitationbeyond that expected over areas of uniform vegetation.

This convection creates spatial differences in the upwardand downward wind velocities and contributes to thedevelopment of mesoscale (20 to 200 km [12 to 120 miles])circulation in the atmosphere. For example, when cre-ating models for forecasting atmospheric conditionson the Great Plains and along the Front Range of theRocky Mountains, the mix of land cover and vegetation

types must be specified to properly relate the fluxes ofmomentum and sensible and latent heat to the larger-scalecirculation of the atmosphere. Proper calculations are alsonecessary to estimate rainfall. In addition, the specificlocation and hour of the day that thunderstorms occurdepend on the heterogeneity of the vegetation cover ofthis region. Field observations have shown that the het-

erogeneity of surface roughness (small-scale irregularitiesin topography), soil moisture, forest coverage, and tran-spiration affect the location and pace of the formation of



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convective clouds and rainfall. Both convection and thun-derstorm development tend to occur earlier in the day inheterogeneous landscapes.

Biosphere Controls on Surface Friction and Localized Winds

Averaged annually over Earth’s entire surface, the Sunprovides about 345 watts per square metre of energy.About 30 percent of this energy is reflected away to space

and is never used in the Earth-atmosphere system. Of that which remains, a little less than 1 percent (3.1 watts persquare metre) accelerates the air by generating winds. Anequal amount of energy must eventually be lost, or else

wind speeds would perpetually increase.Earth as a thermodynamic system is dissipative—the

mechanical energy of the winds is eventually converted to

heat through friction. Over the continents, it is the com-bination of terrain and the veneer of vegetation that offersthe frictional roughness to dissipate the surface winds andconvert this kinetic energy into heat. Marine windsapproaching the British Isles average about 12 metres persecond (27 miles per hour), but they are decelerated to 6metres per second (13 miles per hour) because of the fric-

tion of the landscape’s surface shortly after the windsmake landfall. Without vegetation cover, the continents would offer much less friction to the wind, and windspeeds in unvegetated landscapes would be nearly twice asfast as those in vegetated landscapes.

The correct specification of Earth’s surface roughnessdue to vegetation, for use in computer models of the atmo-

sphere, is critical to proper model performance. If theheight of the terrain and vegetation is not specified cor-rectly, the patterns of Earth’s winds, global geography, and





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rainfall will be poorly modeled. When modeling newlydesertified areas, such as the Sahel, it is important tounderstand that desertification creates vegetation oflower stature and thus lower surface roughness values. Asa result, both wind velocities and wind direction couldchange from previous patterns over landscapes with taller

vegetation.The extent of this impact of the biosphere on the

atmosphere is revealed in climate model studies. Onesuch study modeled the influence of reduced vegetation

on surface roughness over the Indian subcontinent andprovided evidence for a weaker monsoon and reducedrainfall. Given that much of the northwest third of Indiaunderwent a severe desertification and cultural collapsenear the beginning of historical times, the role culturesplay in vegetation reduction and climate change shouldnot be ignored.

The vegetation cover of the continents is not passivein response to the winds. Greenhouse-grown trees sub-jected to mechanical forces designed to mimic the windslay down new woody tissue called “reaction wood,” whichresults in a stiffer tree over time. This material helps treesbecome more resilient and offer more frictional resistanceto wind. This negative feedback, where increased winds

result in stiffer vegetation and thereby subsequentlyreduced wind speeds, might well apply at the global scaleby balancing the energy used to heat and accelerate theair (3.1 watts per square metre) with the surface frictionneeded to dissipate it.

Biosphere Impacts on Precipitation

Processes

The formation and subsequent freezing of cloud dropletsdepend on the presence of cloud condensation nuclei and



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ice nuclei, respectively. Significantly, the biosphere is amajor source of both of these kinds of nuclei. Over thecontinents, condensation nuclei are readily available andare of biogenic as well as anthropogenic origin. Examplesof condensation nuclei include sea salt, small soil particles,and dust. In addition, through transpiration in plants, thebiosphere contributes to the recycling of rainfall.

Cloud Condensation Nuclei

As atmospheric convection increases with the heating of

the day, cloud condensation nuclei are mixed into andabove the planetary boundary layer and into the tropo-sphere. In the bottom 0.5 km (the lowest 1,600 feet or so)of the atmosphere, nuclei typically number 2.2 × 1010 percubic metre. In the next 0.5 km (between 1,600 and 3,300feet) above, half as many nuclei are found. The number ofcondensation nuclei continues to decline with increased

altitude. Furthermore, in general, the number of nuclei inthe air over land is 10 times higher than over the oceans.

Cloud condensation nuclei are generally abundant.They do not limit cloud formation over the continents;however, low numbers of condensation nuclei over theoceans may limit cloud formation there. In addition tonatural sources, particulates from fuel combustion and

sulfur dioxide gas resulting from high sulfur fuels alsocontribute to the load of condensation nuclei over thecontinents. Both the number and kind of condensationnuclei present in the atmosphere affect the cloudinessand the brightness of clouds in a given region. In this way,condensation nuclei play a significant role in determiningboth regional and global albedo.

There is a type of condensation nuclei that forms inthe marine air over the margins of continents. Thoughthese nuclei are often few in number, they play a largerole in cloud formation near the coasts of continents and





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may contribute significantly to both planetary albedo and global average temperature. Typically, sources of conden-sation nuclei in marine air are sulfate aerosols formed fromthe biogenic production of dimethyl sulfide (DMS) bymarine algae. Given that DMS production increases withsea surface temperatures, a negative feedback may result.The central idea in this feedback hypothesis is that warmer

waters result in the increased production of condensationnuclei by phytoplankton and thus produce more clouds.Increased cloudiness shades the ocean surface and results

in lower temperatures that limit condensation nucleiproduction. It is estimated that a 30 percent increase inmarine condensation nuclei would increase planetaryalbedo by 0.005 (0.5 percent) or produce a 0.7 percentreduction in solar radiation and a planetary average tem-perature decrease of 1.3 °C (2.3 °F). The sensitivity of thisnegative feedback on planetary temperatures remains in

active debate.

Biogenic Ice Nuclei

As water vapour condenses onto condensation nuclei, thedroplets grow in size. Growth proceeds at relative humid-ity as low as 70 percent, but the rate of growth is veryslow. Growth by condensation is most rapid where the

air is slightly supersaturated with water vapour. At thispoint, cloud droplets typical of the size of fog dropletsarise. Should temperatures fall to the level where freezingbegins, the temperature difference between the drop-let and the surrounding air (the vapour pressure deficit)strongly favours rapid condensation into the crystallinelattice of an ice particle. Ice particles that grow rapidly

soon reach sizes where they begin to fall. As they fall,they collide and merge with smaller droplets and thereby grow larger.



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The formation of ice is of critical importance. Adroplet of pure water, such as distilled water, will auto-matically freeze in the atmosphere at a temperature of

–40 °C (–40 °F). Freezing at warmer temperatures requiresa substance upon which ice crystallization can take place.The common clay mineral kaolinite, a contaminant of thedroplet, raises this freezing point to around –25 °C (–13 °F).Furthermore, silver iodide, often used in cloud seeding toencourage rainfall, and sea salts also cause ice to form at

–25 °C. Freezing at still warmer temperatures is most com-

mon with biogenic ice nuclei. Upon ice formation, heatenergy on the order of 80 calories per gram of water fro-zen are released. This energy increases the sensible heatof the air and causes the air to become more buoyant. Theprocess of ice formation encourages convection, cloudi-ness, and precipitation from clouds.

The decomposition of organic matter is a major

source of biogenic ice nuclei. Ice crystal formation hasbeen shown to occur at temperatures as warm as –2 to –3°C (28.5 to 26.6 °F) when biogenic ice nuclei are involved.The common freezing temperature for biogenic nuclei

varies systematically according to biome and latitude.The coldest freezing-temperature nuclei occur abovethe tropics, whereas the warmest occur above the Arctic.

Freezing produces greater buoyancy of the particles andhelps them to reach higher vertical velocities within theclouds. The vertical motions and the larger droplet sizethat occur with biogenic materials favour the chargeseparation needed to produce lightning. Subsequently,oceanic areas with few biogenic ice nuclei are also areasof low lightning frequency. The production of biogenic

nuclei from organic matter decomposition is greatestduring the warm months when bacterial decompositionis greatest.





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Recycled Rainfall

The water that is transpired into the atmosphere fromthe biosphere is eventually returned to the surface as pre-cipitation. This vegetation-transpiration component ofthe hydrologic cycle is referred to as “recycled rainfall.”While the oceans are the major source of atmospheric

water vapour and rainfall, water from plant transpira-tion is also significant. For example, in the 1970s and ’80s,analyses performed by American meteorologist Michael

Garstang on the city of Manaus, Braz., in the Amazonbasin revealed that around 20 percent of the precipitationcame from water transpired by vegetation; the remain-ing 80 percent of this precipitation (an estimate madeby German American meteorologist Heinz Lettau in the1970s) was generated by the Atlantic Ocean. Isotopicstudies of rainwater collected at various points in the

Amazon basin indicated that nearly half of the total raincame from water originating in the ocean and half trans-pired through the vegetation. Evidence of the proportionof transpired water in rainfall reaching as high as 88 per-cent has been reported for the Amazon foothills of theAndes. General climate circulation models indicate that,

without transpired water from plants, rainfall in the cen-

tral regions of the continents would be greatly reduced. Asa general rule, the farther the distance from oceanic watersources, the higher the fraction of rainwater originatingfrom transpiration.

CLIMATE, HUMANS,AND HUMAN AFFAIRS

In the late 1960s and early ’70s, climatologists envisionedthe start of a new ice age because it was becoming clear



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that a cycle of planetary cooling was detected in the post-1940s record. The Central Intelligence Agency of theUnited States commissioned studies of the global politicalstresses that would ensue with a 1 °C (1.8 °F) temperaturedecline. The central question revolved around the loca-tions of the political hot spots in a cooling world. Theresults of these studies were published under the titleThe Weather Conspiracy. In the late 1970s, HarvardUniversity’s Center for International Affairs addressedthese issues in a book by English diplomat and environ-

mentalist Crispin Tickell titled Climatic Change and World Affairs. Tickell sounded a warning:

A shift of 2 °C in mean temperatures leads either to ice ages or

to melting of the polar ice caps, either of which would destroy

much of present civilization.

In the late 1970s the global warming concerns aris-ing from the burning of fossil fuels were still a decadeaway, but the Harvard report provided the impetus toresearch the possible links between the burning of fossilfuels and global warming. In academia during this time,climatologists and historians began working together toreexamine the past connections between climate and

history.Since prehistoric times, humans have altered the landcover of the continents to suit their economic and culturalenterprises. In so doing, they have recast the two-way bal-ance of mass and energy exchange between the atmosphereand the biosphere. Since the characteristics of Earth’s cli-mate in the absence of cities, agriculture, and other human

land uses are unknown, comparing the dynamics of thepresent Earth-atmosphere system with that of preurbanand preagricultural times is very difficult.





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In any case, humans are subject to the same climatic variations and changes that affect other life-forms. Foodand fibre resources are climatically sensitive and vary inprice accordingly. The cost of hauling coal by railroaddepends on the temperature of the rails and on the weight

gain that occurs when coal is wetted by rainfall. The con-sumption of fossil fuels for heating and cooling purposes isalso climate-dependent. Damage from hurricanes, floods,droughts, snowstorms, tornadoes, and other weather phe-nomena have real costs. Climate conditions also affect

decisions concerning travel and leisure.As an exceptionally adaptable species, humans are

found in all corners of the planet save the coldest polarregions, highest mountain peaks, and lowest oceantroughs. As a result, there remain few places on Earththat are not in some respect aptly classified as human-dominated ecosystems. No region is untouched by human

influence. The release of waste products from domesticand economic enterprises (burning fossil fuels, syntheticchemical use, trash production, etc.) alters the composi-tion of the atmosphere, and gases and particulates relatedto these activities travel to all parts of the globe. In con-trast, land clearing and development often permanentlyalter the surface of the planet and modify patterns of both

surface heating and local weather. Economically, theseenterprises are purposeful actions; however, they areinadvertent when it comes to the realized changes in theenvironment. Air pollution, ozone depletion, acid precipi-tation, global warming, desertification, smog production,and deforestation are but a few of the human impacts onthe climate system that arise from the alteration of the

mass and energy exchange with the atmosphere.The Gaia hypothesis, introduced at the beginningof this chapter, remains controversial in the scientific



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community. Nevertheless, the question of whether Homo sapiens possesses the capacity and the will to maintain theEarth-atmosphere system in a sort of relative homeostaticbalance is intriguing. The Montreal Protocol on SubstancesThat Deplete the Ozone Layer, which has resulted in thephasing out of chlorofluorocarbons (CFCs)—a groupof industrial compounds that react with and disassoci-ate ozone molecules—is a collective adaptive responseby humans to a perceived and predicted threat to lifefrom stratospheric ozone depletion. The Environmental

Protection Acts ratified by the United Kingdom andAustralia and the Kyoto Protocol to the United NationsFramework Convention on Climate Change are someexamples of attempts to combat deleterious environmen-tal change associated with the release of additional carbondioxide into the air. If humans are to maintain the Earth-atmosphere system, it is through the social institutions of

research and education that they will recognize environ-mental threats, come to wise decisions, and attempt toexercise responsible stewardship of the planet.




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CLIMATE CHANGE

C limate change is periodic modification of Earth’s cli-mate brought about as a result of changes in the

atmosphere as well as interactions between the atmo-sphere and various other geologic, chemical, biological,and geographic factors within the Earth system.

The atmosphere is a dynamic fluid that is continuallyin motion. Both its physical properties and its rate and

direction of motion are influenced by a variety of factors,including solar radiation, the geographic position of con-tinents, ocean currents, the location and orientation ofmountain ranges, atmospheric chemistry, and vegetation

growing on the land surface. All these factors changethrough time. Some factors, such as the distribution ofheat within the oceans, atmospheric chemistry, and sur-

face vegetation, change at very short timescales. Others,such as the position of continents and the location andheight of mountain ranges, change over very long time-scales. Therefore, climate, which results from the physicalproperties and motion of the atmosphere, varies at everyconceivable timescale.

Climate is often defined loosely as the average weather

at a particular place, incorporating such features as tem-perature, precipitation, humidity, and windiness. A morespecific definition would state that climate is the meanstate and variability of these features over some extendedtime period. Both definitions acknowledge that the

weather is always changing, owing to instabilities inthe atmosphere. And as weather varies from day to day, so

too does climate vary, from daily day-and-night cycles upto periods of geologic time hundreds of millions of yearslong. In a very real sense, climate variation is a redundantexpression—climate is always varying. No two years are



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7 Climate Change 7

exactly alike, nor are any two decades, any two centuries,

or any two millennia.This chapter addresses the concept of climatic varia-

tion and change within the set of integrated naturalfeatures and processes known as the Earth system. Thenature of the evidence for climate change is explained, asare the principal mechanisms that have caused climatechange throughout the history of Earth. Finally, a detailed

description is given of climate change over many differenttimescales, ranging from a typical human life span to all of geologic time.

THE EARTH SYSTEM

The atmosphere is influenced by and linked to other fea-

tures of Earth, including oceans, ice masses (glaciers andsea ice), land surfaces, and vegetation. Together, they makeup an integrated Earth system, in which all componentsinteract with and influence one another in often complex

A series of photographs of the Grinnell Glacier taken from the summit of

Mount Gould in Glacier National Park, Montana, in 1938, 1981, 1998, and 2006 (from left to right). In 1938 the Grinnell Glacier filled the entire area atthe bottom of the image. By 2006 it had largely disappeared from this view.1938-T.J. Hileman/Glacier National Park Archives, 1981 - Carl Key/ USGS, 1998 - Dan Fagre/USGS, 2006 - Karen Holzer/USGS



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ways. For instance, climate influences the distribution of vegetation on Earth’s surface (e.g., deserts exist in aridregions, forests in humid regions), but vegetation in turninfluences climate by reflecting radiant energy back intothe atmosphere, transferring water (and latent heat) from

soil to the atmosphere, and influencing the horizontalmovement of air across the land surface.Earth scientists and atmospheric scientists are still

seeking a full understanding of the complex feedbacksand interactions among the various components ofthe Earth system. This effort is being facilitated by thedevelopment of an interdisciplinary science called Earth

system science. Earth system science is composed of a wide range of disciplines, including climatology (the studyof the atmosphere), geology (the study of Earth’s surface

Drought-resistant plants in Repetek Preserve in the southeastern Karakum Desert, Turkmenistan. © Rodger Jackman/Oxford Scientific Films Ltd.



171

and underground processes), ecology (the study of howEarth’s organisms relate to one another and their envi-ronment), oceanography (the study of Earth’s oceans),

glaciology (the study of Earth’s ice masses), and the socialsciences (the study of human behaviour in its social andcultural aspects).

A full understanding of the Earth system requiresknowledge of how the system and its components havechanged through time. The pursuit of this understand-ing has led to development of Earth system history, an

interdisciplinary science that includes not only the contri-butions of Earth system scientists but also paleontologists(who study the life of past geologic periods), paleoclima-tologists (who study past climates), paleoecologists (whostudy past environments and ecosystems), paleoceanog-raphers (who study the history of the oceans), and otherscientists concerned with Earth history. Because different

components of the Earth system change at different ratesand are relevant at different timescales, Earth system his-tory is a diverse and complex science. Students of Earthsystem history are not just concerned with document-ing what has happened; they also view the past as a seriesof experiments in which solar radiation, ocean currents,continental configurations, atmospheric chemistry, and

other important features have varied. These experimentsprovide opportunities to learn the relative influencesof and interactions between various components of theEarth system. Studies of Earth system history also specifythe full array of states the system has experienced in thepast and those the system is capable of experiencing inthe future.

Undoubtedly, people have always been aware ofclimatic variation at the relatively short timescales of sea-sons, years, and decades. Biblical scripture and other early

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documents refer to droughts, floods, periods of severecold, and other climatic events. Nevertheless, a full appre-ciation of the nature and magnitude of climatic changedid not come about until the late 18th and early 19th cen-turies, a time when the widespread recognition of thedeep antiquity of Earth occurred. Naturalists of this time,including Scottish geologist Charles Lyell, Swiss-bornnaturalist and geologist Louis Agassiz, English naturalistCharles Darwin, American botanist Asa Gray, and Welshnaturalist Alfred Russel Wallace, came to recognize geo-

logic and biogeographic evidence that made sense only inthe light of past climates radically different from thoseprevailing today.

Geologists and paleontologists in the 19th and early20th centuries uncovered evidence of massive climaticchanges taking place before the Pleistocene—that is,before some 2.6 million years ago. For example, red beds

indicated aridity in regions that are now humid (e.g.,England and New England), whereas fossils of coal-swampplants and reef corals indicated that tropical climates onceoccurred at present-day high latitudes in both Europe andNorth America. Since the late 20th century the develop-ment of advanced technologies for dating rocks, together

with geochemical techniques and other analytical tools,

have revolutionized the understanding of early Earth sys-tem history.The occurrence of multiple epochs in recent Earth

history during which continental glaciers, developed athigh latitudes, penetrated into northern Europe and east-ern North America was recognized by scientists by thelate 19th century. Scottish geologist James Croll proposed

that recurring variations in orbital eccentricity (the devia-tion of Earth’s orbit from a perfectly circular path) wereresponsible for alternating glacial and interglacial periods.



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The precession of Earth’s axis. Encyclopædia Britannica, Inc.

Croll’s controversial idea was taken up by Serbian math-

ematician and astronomer Milutin Milankovitch in theearly 20th century. Milankovitch proposed that the mech-anism that brought about periods of glaciation was drivenby cyclic changes in eccentricity as well as two other orbitalparameters: precession (a change in the directional focusof Earth’s axis of rotation) and axial tilt (a change in theinclination of Earth’s axis with respect to the plane of its

orbit around the Sun). Orbital variation is now recognizedas an important driver of climatic variation throughoutEarth’s history.

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EVIDENCE FOR CLIMATE CHANGE

All historical sciences share a problem: As they probe far-ther back in time, they become more reliant on fragmentaryand indirect evidence. Earth system history is no excep-tion. High-quality instrumental records spanning the pastcentury exist for most parts of the world, but the recordsbecome sparse in the 19th century, and few records pre-date the late 18th century. Other historical documents,including ship’s logs, diaries, court and church records,

and tax rolls, can sometimes be used. Within strict geographic contexts, these sources can provide informationon frosts, droughts, floods, sea ice, the dates of monsoons,and other climatic features—in some cases up to severalhundred years ago.

Fortunately, climatic change also leaves a varietyof signatures in the natural world. Climate influences

the growth of trees and corals, the abundance and geographic distribution of plant and animal species, thechemistry of oceans and lakes, the accumulation of ice incold regions, and the erosion and deposition of materialson Earth’s surface. Paleoclimatologists study the tracesof these effects, devising clever and subtle ways to obtaininformation about past climates. Most of the evidence of

past climatic change is circumstantial, so paleoclimatol-ogy involves a great deal of investigative work. Whereverpossible, paleoclimatologists try to use multiple linesof evidence to cross-check their conclusions. They arefrequently confronted with conflicting evidence, butthis, as in other sciences, usually leads to an enhancedunderstanding of the Earth system and its complex

history. New sources of data, analytical tools, and instru-ments are becoming available, and the field is movingquickly. Revolutionary changes in the understanding of



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Earth’s climate history have occurred since the 1990s,and coming decades will bring many new insights andinterpretations.

Ongoing climatic changes are being monitored bynetworks of sensors in space, on the land surface, andboth on and below the surface of the world’s oceans.Climatic changes of the past 200–300 years, especiallysince the early 1900s, are documented by instrumentalrecords and other archives. These written documentsand records provide information about climate change

in some locations for the past few hundred years. Some very rare records date back over 1,000 years. Researchersstudying climatic changes predating the instrumentalrecord rely increasingly on natural archives, which arebiological or geologic processes that record some aspectof past climate. These natural archives, often referredto as proxy evidence, are extraordinarily diverse; they

include, but are not limited to, fossil records of pastplant and animal distributions, sedimentary and geo-chemical indicators of former conditions of oceans andcontinents, and land surface features characteristic ofpast climates. Paleoclimatologists study these natu-ral archives by collecting cores, or cylindrical samples,of sediments from lakes, bogs, and oceans; by studying

surface features and geological strata; by examining treering patterns from cores or sections of living and deadtrees; by drilling into marine corals and cave stalag-mites; by drilling into the ice sheets of Antarctica andGreenland and the high-elevation glaciers of the Plateauof Tibet, the Andes, and other montane regions; and bya wide variety of other means. Techniques for extracting

paleoclimatic information are continually being devel-oped and refined, and new kinds of natural archives arebeing recognized and exploited.

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CAUSES OF CLIMATE CHANGE

It is much easier to document the evidence of climate variability and past climate change than it is to determinetheir underlying mechanisms. Climate is influenced by amultitude of factors that operate at timescales rangingfrom hours to hundreds of millions of years. Many of thecauses of climate change are external to the Earth system.Others are part of the Earth system but external to theatmosphere. Still others involve interactions between

the atmosphere and other components of the Earth sys-tem and are collectively described as feedbacks withinthe Earth system. Feedbacks are among the most recentlydiscovered and challenging causal factors to study.Nevertheless, these factors are increasingly recognized asplaying fundamental roles in climate variation. The mostimportant mechanisms are described in this chapter.

Solar Variability

The luminosity, or brightness, of the Sun has been increas-ing steadily since its formation. This phenomenon isimportant to Earth’s climate, because the Sun providesthe energy to drive atmospheric circulation and consti-

tutes the input for Earth’s heat budget. Low solarluminosity during Precambrian time underlies the faint young Sun paradox, described in the section near the endof this chapter, the Climates of Early Earth.

Radiative energy from the Sun is variable at very smalltimescales, owing to solar storms and other disturbances,but variations in solar activity, particularly the frequency

of sunspots, are also documented at decadal to millennialtimescales and probably occur at longer timescales as well.The “Maunder minimum,” a period of drastically reduced



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sunspot activity between 1645 and 1715 CE, has been sug-

gested as a contributing factor to the Little Ice Age.

Volcanic Activity

Volcanic activity can influence climate in a number of ways at different timescales. Individual volcanic eruptionscan release large quantities of sulfur dioxide and other

aerosols into the stratosphere, reducing atmospherictransparency and thus the amount of solar radiation reach-ing Earth’s surface and troposphere. A recent example is

The Sun as imaged in extreme ultraviolet light by the Earth-orbiting Solar and Heliospheric Observatory (SOHO) satellite. A massive loop-shapederuptive prominence is visible at the lower left. Nearly white areas are the

hottest; deeper reds indicate cooler temperatures. NASA

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the 1991 eruption in the Philippines of Mount Pinatubo, which had measurable influences on atmospheric circula-tion and heat budgets. The 1815 eruption of MountTambora on the island of Sumbawa had more dramaticconsequences, as the spring and summer of the following

year (1816, known as “the year without a summer”) wereunusually cold over much of the world. New England andEurope experienced snowfalls and frosts throughout thesummer of 1816.

Volcanoes and related phenomena, such as oceanrifting and subduction, release carbon dioxide intoboth the oceans and the atmosphere. Emissions are

low; even a massive volcanic eruption such as MountPinatubo releases only a fraction of the carbon dioxideemitted by fossil-fuel combustion in a year. At geologic time-scales, however, release of this greenhouse gas can have

Iceland's Fimmvorduhals volcano is shown here March 27, 2010, days after itbegan spewing lava for the first time in two centuries. Home to approximately

200 volcanoes of various types, Iceland has seen increasing volcanic activity

since the 1970s. AFP/Getty Images



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important effects. Variations in carbon dioxide release by volcanoes and ocean rifts over millions of years can alterthe chemistry of the atmosphere. Such changeability incarbon dioxide concentrations probably accounts formuch of the climatic variation that has taken place duringthe Phanerozoic Eon.

Tectonic Activity

Tectonic movements of Earth’s crust have had pro-

found effects on climate at timescales of millions to tensof millions of years. These movements have changedthe shape, size, position, and elevation of the conti-nental masses as well as the bathymetry of the oceans.Topographic and bathymetric changes in turn have hadstrong effects on the circulation of both the atmosphereand the oceans. For example, the uplift of the Tibetan

Plateau during the Cenozoic Era affected atmosphericcirculation patterns, creating the South Asian monsoonand influencing climate over much of the rest of Asiaand neighbouring regions.

Tectonic activity also influences atmospheric chemis-try, particularly carbon dioxide concentrations. Carbondioxide is emitted from volcanoes and vents in rift zones

and subduction zones. Variations in the rate of spreadingin rift zones and the degree of volcanic activity near platemargins have influenced atmospheric carbon dioxide con-centrations throughout Earth’s history. Even the chemical

weathering of rock constitutes an important sink for car-bon dioxide. (A carbon sink is any process that removescarbon dioxide from the atmosphere by the chemical con-

version of CO2 to organic or inorganic carbon compounds.)Carbonic acid, formed from carbon dioxide and water, is areactant in dissolution of silicates and other minerals.Weathering rates are related to the mass, elevation, and

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exposure of bedrock. Tectonic uplift can increase all thesefactors and thus lead to increased weathering and carbondioxide absorption. For example, the chemical weatheringof the rising Tibetan Plateau may have played an impor-tant role in depleting the atmosphere of carbon dioxideduring a global cooling period in the late Cenozoic Era.

Orbital (Milankovitch) Variations

The orbital geometry of Earth is affected in predictable

ways by the gravitational influences of other planets in thesolar system. Three primary features of Earth’s orbit areaffected, each in a cyclic, or regularly recurring, manner.First, the shape of Earth’s orbit around the Sun varies fromnearly circular to elliptical (eccentric), with periodicitiesof 100,000 and 413,000 years. Second, the tilt of Earth’saxis with respect to the Sun, which is primarily respon-

sible for Earth’s seasonal climates, varies between 22.1°and 24.5° from the plane of Earth’s rotation around theSun. This variation occurs on a cycle of 41,000 years. In

general, the greater the tilt, the greater the solar radiationreceived by hemispheres in summer and the less receivedin winter. The third cyclic change to Earth’s orbital geom-etry results from two combined phenomena: (1) Earth’s

axis of rotation wobbles, changing the direction of theaxis with respect to the Sun, and (2) the orientation ofEarth’s orbital ellipse rotates slowly. These two processescreate a 26,000-year cycle, called precession of the equi-noxes, in which the position of Earth at the equinoxes andsolstices changes. Today Earth is closest to the Sun (peri-helion) near the December solstice, whereas 9,000 years

ago perihelion occurred near the June solstice.These orbital variations cause changes in the latitudinaland seasonal distribution of solar radiation, which in turndrive a number of climate variations. Orbital variations play



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major roles in pacing glacial-interglacial and monsoonalpatterns. Their influences have been identified in climaticchanges over much of the Phanerozoic. For example,cyclothems—which are interbedded marine, fluvial, andcoal beds characteristic of the Pennsylvanian Subperiod

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In this illustration of the greenhouse effect on Earth, Some incoming sunlight is reflected by the Earth’s atmosphere and surface, but most is absorbedby the surface, which is warmed. Infrared (IR) radiation is then emitted

from the surface. Some IR radiation escapes to space, but some is absorbed bythe atmosphere’s greenhouse gases (especially water vapour, carbon dioxide,

and methane) and reradiated in all directions, some to space and some backtoward the surface, where it further warms the surface and the lower atmo-

sphere. Encyclopædia Britannica, Inc.



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(318.1 million to 299 million years ago)—appear to repre-sent Milankovitch-driven changes in mean sea level.

Greenhouse Gases

Greenhouse gases are gas molecules that have the propertyof absorbing infrared radiation (net heat energy) emittedfrom Earth’s surface and reradiating it back to Earth’s sur-face, thus contributing to the phenomenon known as the

greenhouse effect. Carbon dioxide, methane, and water

vapour are the most important greenhouse gases, and theyhave a profound effect on the energy budget of the Earthsystem despite making up only a fraction of all atmospheric

gases. Concentrations of greenhouse gases have variedsubstantially during Earth’s history, and these variationshave driven substantial climate changes at a wide range oftimescales. In general, greenhouse gas concentrations have

been particularly high during warm periods and low duringcold phases. A number of processes influence greenhouse

gas concentrations. Some, such as tectonic activities, oper-ate at timescales of millions of years, whereas others, suchas vegetation, soil, wetland, and ocean sources and sinks,operate at timescales of hundreds to thousands of years.Human activities—especially fossil-fuel combustion since

the Industrial Revolution—are responsible for steadyincreases in atmospheric concentrations of various green-house gases, especially carbon dioxide, methane, ozone,and chlorofluorocarbons (CFCs).

Feedback Within the Earth System

Perhaps the most intensively discussed and researchedtopic in climate variability is the role of interactions andfeedbacks among the various components of the Earthsystem. The feedbacks involve different components that



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Surface reflectance (albedo) of solar energy under different patterns of landuse. (Top) In a preagricultural landscape, large forest-covered areas of low

surface albedo alternate with large open areas of high albedo. (Bottom) In an agricultural landscape, a patchwork of smaller forested and open areas exists,each with its characteristic albedo. Encyclopædia Britannica, Inc.



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operate at different rates and timescales. Ice sheets, seaice, terrestrial vegetation, ocean temperatures, weatheringrates, ocean circulation, and greenhouse gas concentra-tions are all influenced either directly or indirectly bythe atmosphere; however, they also all feed back into theatmosphere, thereby influencing it in important ways.For example, different forms and densities of vegetationon the land surface influence the albedo, or reflectivity,of Earth’s surface, thus affecting the overall radiationbudget at local to regional scales. At the same time, the

transfer of water molecules from soil to the atmosphere ismediated by vegetation, both directly (from transpirationthrough plant stomata) and indirectly (from shading andtemperature influences on direct evaporation from soil).This regulation of latent heat flux by vegetation can influ-ence climate at local to global scales. As a result, changesin vegetation, which are partially controlled by climate,

can in turn influence the climate system. Vegetation alsoinfluences greenhouse gas concentrations; living plantsconstitute an important sink for atmospheric carbon diox-ide, whereas they act as sources of carbon dioxide whenthey are burned by wildfires or undergo decomposition.These and other feedbacks among the various componentsof the Earth system are critical for both understanding

past climate changes and predicting future ones.

Human Activities

Recognition of global climate change as an environmen-tal issue has drawn attention to the climatic impact ofhuman activities. Most of this attention has focused

on carbon dioxide emission via fossil-fuel combustionand deforestation. Human activities also yield releasesof other greenhouse gases, such as methane (from rice



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cultivation, livestock, landfills, and other sources) andchlorofluorocarbons (from industrial sources). There islittle doubt among climatologists that these greenhouse

gases will affect the radiation budget of Earth; the natureand magnitude of the climatic response are a subject ofintense research activity. Paleoclimate records from treering, coral, and ice core records indicate a clear warmingtrend spanning the 20th century. In fact, the 20th century

was the warmest of the past 10 centuries, and the 1990s were the warmest decade of the entire period. Many cli-

matologists have pointed to this warming pattern as clearevidence of human-induced climate change resulting fromthe production of greenhouse gases.

A second type of human impact, the conversion of veg-etation by deforestation, afforestation, and agriculture, isreceiving mounting attention as a further source of cli-mate change. It is becoming increasingly clear that human

impacts on vegetation cover can have local, regional,and even global effects on climate, owing to changes inthe sensible and latent heat flux to the atmosphere and

The global average surface temperature range for each year from 1861 to 2000

is shown by solid gray bars, with the confidence range in the data for each year shown by thin whisker bars. The average change over time is shown by the solid curve. Encyclopædia Britannica, Inc.

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Abandoned farmstead showing the effects of wind erosion in the Dust Bowl,Texas county, Okla., 1937. USDA Photo

the distribution of energy within the climate system. Theextent to which these factors contribute to recent andongoing climate change is an important, emerging areaof study.

CLIMATE CHANGE WITHIN AHUMAN LIFE SPAN

Regardless of their locations on the planet, all humansexperience climate variability and change within their life-

times. The most familiar and predictable phenomena arethe seasonal cycles, to which people adjust their clothing,outdoor activities, thermostats, and agricultural practices.However, no two summers or winters are exactly alike inthe same place; some are warmer, wetter, or stormier thanothers. This interannual variation in climate is partly



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responsible for year-to-year variations in fuel prices, crop yields, road maintenance budgets, and wildfire hazards.Single-year, precipitation-driven floods can cause severeeconomic damage, such as those of the upper MississippiRiver drainage basin during the summer of 1993, and lossof life, such as those that devastated much of Bangladeshin the summer of 1998. Similar damage and loss of life canalso occur as the result of wildfires, severe storms, hurri-canes, heat waves, and other climate-related events.

Climate variation and change may also occur over lon-

ger periods, such as decades. Some locations experiencemultiple years of drought, floods, or other harsh condi-tions. Such decadal variation of climate poses challengesto human activities and planning. For example, multiyeardroughts can disrupt water supplies, induce crop failures,and cause economic and social dislocation, as in the caseof the Dust Bowl droughts in the midcontinent of North

America during the 1930s. Multiyear droughts may evencause widespread starvation, as in the Sahel drought thatoccurred in northern Africa during the 1970s and ’80s.

Seasonal Variation

Every place on Earth experiences seasonal variation in cli-

mate (though the shift can be slight in some tropicalregions). This cyclic variation is driven by seasonal changesin the supply of solar radiation to Earth’s atmosphere andsurface. Earth’s orbit around the Sun is elliptical; it iscloser to the Sun ( 147 million km [about 91 million miles])near the winter solstice and farther from the Sun (152 mil-lion km [about 94 million miles]) near the summer solstice

in the Northern Hemisphere. Furthermore, Earth’s axis ofrotation occurs at an oblique angle (23.5°) with respect toits orbit. Thus, each hemisphere is tilted away from theSun during its winter period and toward the Sun in its

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A diagram shows the position of Earth at the beginning of each season in the Northern Hemisphere. Encyclopædia Britannica, Inc.

summer period. When a hemisphere is tilted away fromthe Sun, it receives less solar radiation than the oppositehemisphere, which at that time is pointed toward the Sun.Thus, despite the closer proximity of the Sun at the wintersolstice, the Northern Hemisphere receives less solarradiation during the winter than it does during the sum-mer. Also as a consequence of the tilt, when the Northern

Hemisphere experiences winter, the Southern Hemisphereexperiences summer.Earth’s climate system is driven by solar radiation;

seasonal differences in climate ultimately result fromthe seasonal changes in Earth’s orbit. The circulation of airin the atmosphere and water in the oceans responds to sea-sonal variations of available energy from the Sun. Specific

seasonal changes in climate occurring at any given locationon Earth’s surface largely result from the transfer of energyfrom atmospheric and oceanic circulation. Differences insurface heating taking place between summer and winter



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cause storm tracks and pressure centres to shift positionand strength. These heating differences also drive seasonalchanges in cloudiness, precipitation, and wind.

Seasonal responses of the biosphere (especially vegeta-tion) and cryosphere (glaciers, sea ice, snowfields) alsofeed into atmospheric circulation and climate. Leaf fall bydeciduous trees as they go into winter dormancy increasesthe albedo (reflectivity) of Earth’s surface and may lead to

greater local and regional cooling. Similarly, snow accumu-lation also increases the albedo of land surfaces and often

amplifies winter’s effects.

Interannual Variation

Interannual climate variations, including droughts,floods, and other events, are caused by a complex array offactors and Earth system interactions. One important

feature that plays a role in these variations is the periodicchange of atmospheric and oceanic circulation patternsin the tropical Pacific region, collectively known as ElNiño–Southern Oscillation (ENSO) variation. Althoughits primary climatic effects are concentrated in the tropi-cal Pacific, ENSO has cascading effects that often extendto the Atlantic Ocean region, the interior of Europe and

Asia, and the polar regions. These effects, called telecon-nections, occur because alterations in low-latitudeatmospheric circulation patterns in the Pacific regioninfluence atmospheric circulation in adjacent and down-stream systems. As a result, storm tracks are diverted andatmospheric pressure ridges (areas of high pressure)and troughs (areas of low pressure) are displaced from

their usual patterns.As an example, El Niño events occur when the easterlytrade winds in the tropical Pacific weaken or reverse direc-tion. This shuts down the upwelling of deep, cold waters

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During years when the North Atlantic Oscillation (NAO) is in its positive phase, the eastern United States, southeastern Canada, and northwestern Europe experience warmer winter temperatures, whereas colder temperatures are found in these locations during its negative phase. When the El Niño/ Southern Oscillation (ENSO) and NAO are both in their positive phase,

European winters tend to be wetter and less severe; however, beyond this general tendency, the influence of the ENSO upon the NAO is not well understood. Encyclopædia Britannica, Inc.

off the west coast of South America, warms the easternPacific, and reverses the atmospheric pressure gradient inthe western Pacific. As a result, air at the surface moveseastward from Australia and Indonesia toward the centralPacific and the Americas. These changes produce highrainfall and flash floods along the normally arid coast ofPeru and severe drought in the normally wet regionsof northern Australia and Indonesia. Particularly severe El



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Niño events lead to monsoon failure in the Indian Oceanregion, resulting in intense drought in India and EastAfrica. At the same time, the westerlies and storm tracksare displaced toward the Equator, providing Californiaand the desert Southwest of the United States with wet,stormy winter weather and causing winter conditions inthe Pacific Northwest, which are typically wet, to become

warmer and drier. Displacement of the westerlies alsoresults in drought in northern China and from northeast-ern Brazil through sections of Venezuela. Long-term

records of ENSO variation from historical documents,tree rings, and reef corals indicate that El Niño eventsoccur, on average, every two to seven years. However, thefrequency and intensity of these events vary through time.

The North Atlantic Oscillation (NAO) is anotherexample of an interannual oscillation that producesimportant climatic effects within the Earth system and

can influence climate throughout the NorthernHemisphere. This phenomenon results from variation inthe pressure gradient, or the difference in atmosphericpressure between the subtropical high, usually situatedbetween the Azores and Gibraltar, and the Icelandic low,centred between Iceland and Greenland. When the pres-sure gradient is steep due to a strong subtropical high and

a deep Icelandic low (positive phase), northern Europeand northern Asia experience warm, wet winters with fre-quent strong winter storms. At the same time, southernEurope is dry. The eastern United States also experiences

warmer, less snowy winters during positive NAO phases,although the effect is not as great as in Europe. The pres-sure gradient is dampened when NAO is in a negative

mode—that is, when a weaker pressure gradient existsfrom the presence of a weak subtropical high andIcelandic low. When this happens, the Mediterraneanregion receives abundant winter rainfall, while northern

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Europe is cold and dry. The eastern United States is typi-cally colder and snowier during a negative NAO phase.

The ENSO and NAO cycles are driven by feedbacksand interactions between the oceans and atmosphere.Interannual climate variation is driven by these and othercycles, interactions among cycles, and perturbations inthe Earth system, such as those resulting from large injec-tions of aerosols from volcanic eruptions. One example ofa perturbation due to volcanism is the 1991 eruptionof Mount Pinatubo in the Philippines, which led to a

decrease in the average global temperature of approxi-mately 0.5 °C (0.9 °F) the following summer.

Decadal Variation

Climate varies on decadal timescales, with multiyear clus-ters of wet, dry, cool, or warm conditions. These multiyear

clusters can have dramatic effects on human activities and welfare. For instance, a severe three-year drought in thelate 16th century probably contributed to the destructionof Sir Walter Raleigh’s “Lost Colony” at Roanoke Island in

what is now North Carolina, and a subsequent seven-yeardrought (1606–12) led to high mortality at the JamestownColony in Virginia. Also, some scholars have implicated

persistent and severe droughts as the main reason for thecollapse of the Maya civilization in Mesoamerica between750 and 950 CE; however, discoveries in the early 21st cen-tury suggest that war-related trade disruptions played arole, possibly interacting with famines and other drought-related stresses.

Although decadal-scale climate variation is well docu-

mented, the causes are not entirely clear. Much decadal variation in climate is related to interannual variations.For example, the frequency and magnitude of ENSOchange through time. The early 1990s were characterized



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This phenomenon, which translates to “The Christ Child,” is the

anomalous appearance, every few years, of unusually warm ocean con-ditions along the tropical west coast of South America. This event isassociated with adverse effects on fishing, agriculture, and local weather from Ecuador to Chile and with far-field climatic anomaliesin the equatorial Pacific and occasionally in Asia and North Americaas well. The Oceanic Niño Index (ONI), a measure of the departurefrom normal sea surface temperature in the east-central PacificOcean, is the standard means by which each El Niño episode is deter-

mined, gauged, and forecast. El Niño episodes are indicated by seasurface temperature increases of more than 0.5 °C (0.9 °F) for at leastfive successive overlapping three-month seasons.

The name El Niño was originally used during the 19th century bythe fishermen of northern Peru in reference to the annual flow of warm equatorial waters southward around Christmas time. Peruvianscientists later noted that more intense changes occurred at intervalsof several years and were associated with catastrophic seasonal flood-ing along the normally arid coast, while the thermal anomalies lastedfor a year or more. The more unusual episodes gained world attentionduring the 20th century, and the original annual connotation of thename was replaced by that of the anomalous occurrence.

The timing and intensity of El Niño events vary widely. The firstrecorded occurrence of unusual desert rainfall was in 1525, when theSpanish conquistador Francisco Pizarro landed in northern Peru.Historians suggest that the desert rains and vegetation encounteredby the Spaniards may have facilitated their conquest of the Inca

empire. The intensity of El Niño episodes varies from weak thermalanomalies (2–3 °C [about 4–5 °F]) with only moderate local effects to very strong anomalies (8–10 °C [14–18 °F]) associated with worldwideclimatic perturbations. El Niño events occur irregularly at two- toseven-year intervals, and the strong events are less common. Theintermittency varies widely, however, and the phenomenon is neitherperiodic nor predictable in the sense that ocean tides are.

Beginning with the work of Sir Gilbert Walker in the 1930s, cli-

matologists recognized a similar interannual change in the tropicalatmosphere, which Walker termed the Southern Oscillation (SO).El Niño and the Southern Oscillation appear to be the oceanic and

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The upwelling process in the ocean along the coast of Peru. A thermocline and a nutricline separate the warm, nutrient-deficient upper layer from thecool, enriched layer below. Under normal conditions (top) , these interfaces are

shallow enough that coastal winds can induce upwelling of the lower-layer nutrients to the surface, where they support an abundant ecosystem. During

an El Niño event (bottom) , the upper layer thickens so that the upwelledwater contains fewer nutrients, thus contributing to a collapse of marine pro-

ductivity. Encyclopædia Britannica, Inc.



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atmospheric components of a single large-scale, coupled interac-tion—the El Niño/Southern Oscillation (ENSO). During the warm

phase of ENSO, the South Pacific trade-wind system undergoes achange of state, or “seesaw,” in which the westward-blowing trades weaken along the Equator as the normally high pressure in the east-ern South Pacific decreases and the low pressure over northernAustralia and Indonesia rises. The pressure change and diminishedtrade winds cause warm surface water to move eastward along theEquator from the western Pacific, while the warm surface layer inthe east becomes thicker. Under normal conditions, the northward-

blowing winds off South America cause nutrient-rich waters toupwell from below the shallow, warm surface layer. The nutrients(mainly phosphates and nitrates) provide a plentiful supply of foodfor photosynthesizing plankton, on which the fish feed. During ElNiño, however, the thicker surface layer acts as a barrier to effec-tive upwelling by the coastal winds. The unenriched surface watersare poor in nutrients and cannot support the normally productivecoastal ecosystem. Fish populations are decimated as great numbers

migrate to less-affected areas in search of food, resulting in tempo-rarily reduced yields for the countries in the region. In 1972–73 thisled not only to local economic setbacks but to repercussions in the world commodity markets as well.

The warm ocean conditions in the equatorial Pacific induce large-scale anomalies in the atmosphere. Rainfall increases manyfold inEcuador and northern Peru, causing coastal flooding and erosion andconsequent hardships in transportation and agriculture. Additionally,strong El Niño events are associated with droughts in Indonesia,

Australia, and northeastern South America and with altered patternsof tropical storms in the tropical belt. During the stronger El Niñoepisodes, the atmospheric “teleconnections” are extensive enough tocause unusually severe winter weather at the higher latitudes of Northand South America.

The El Niño episodes of 1982–83 and 1997–98 were the mostintense of the 20th century. The 1982–83 episode lasted from mid-1982 to mid-1983. Sea surface temperatures in the eastern tropical

Pacific and much of the equatorial zone farther west were 5–10 °C(9–18 °F) above normal. Australia was hit by severe drought; typhoonsoccurred as far east as Tahiti; and central Chile suffered from record

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by repeated El Niño events, and several such clusters have

been identified as having taken place during the 20th cen-tury. The steepness of the NAO gradient also changes atdecadal timescales; it has been particularly steep sincethe 1970s.

Recent research has revealed that decadal-scale variations in climate result from interactions betweenthe ocean and the atmosphere. One such variation is the

Pacific Decadal Oscillation (PDO), also referred to asthe Pacific Decadal Variability (PDV), which involveschanging sea surface temperatures (SSTs) in the NorthPacific Ocean. The SSTs influence the strength and posi-tion of the Aleutian Low, which in turn strongly affectsprecipitation patterns along the Pacific Coast of NorthAmerica. PDO variation consists of an alternation

between “cool-phase” periods, when coastal Alaska is rela-tively dry and the Pacific Northwest relatively wet (e.g.,1947–76), and “warm-phase” periods, characterized by rel-atively high precipitation in coastal Alaska and low

rainfall and flooding. Also, the west coast of North America wasunusually stormy during the winter of 1982–83, and fish catches were

dramatically altered from Mexico to Alaska.The El Niño episode of 1997–98 is regarded by some scientists asthe strongest such event of the 20th century and has the distinctionof being the first episode monitored from beginning to end by scien-tific instrumentation. Although sea surface temperatures and weatherpatterns paralleled the 1982–83 event, the ONI for the 1997–98 epi-sode was the highest on record. The 1997–98 event produced droughtconditions in Brazil, Indonesia, Malaysia, and the Philippines andbrought heavy rains to the dry seacoast of Peru. In the United Statesthe southeastern states and California experienced significantincreases in winter rainfall, and record-breaking warm temperaturesin the upper Midwest caused some journalists to label the period “the year without a winter.”



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precipitation in the Pacific Northwest (e.g., 1925–46,1977–98). Tree ring and coral records, which span at leastthe last four centuries, document PDO variation.

A similar oscillation, the Atlantic MultidecadalOscillation (AMO), occurs in the North Atlantic andstrongly influences precipitation patterns in easternand central North America. A warm-phase AMO (rela-tively warm North Atlantic SSTs) is associated withrelatively high rainfall in Florida and low rainfall in muchof the Ohio Valley. However, the AMO interacts with the

PDO, and both interact with interannual variations, suchas ENSO and NAO, in complex ways. Such interactionsmay lead to the amplification of droughts, floods, or otherclimatic anomalies. For example, severe droughts overmuch of the conterminous United States in the first few

years of the 21st century were associated with warm-phaseAMO combined with cool-phase PDO. The mechanisms

underlying decadal variations, such as PDO and AMO,are poorly understood, but they are probably related toocean-atmosphere interactions with larger time constantsthan interannual variations. Decadal climatic variationsare the subject of intense study by climatologists andpaleoclimatologists.

CLIMATE CHANGE SINCE THEEMERGENCE OF CIVILIZATION

Human societies have experienced climate change sincethe development of agriculture some 10,000 years ago.These climate changes have often had profound effects onhuman cultures and societies. They include annual and

decadal climate fluctuations such as those described ear-lier, as well as large-magnitude changes that occur overcentennial to multimillennial timescales. Such changes arebelieved to have influenced and even stimulated the initial

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cultivation and domestication of crop plants, as well as thedomestication and pastoralization of animals. Humansocieties have changed adaptively in response to climate

variations, although evidence abounds that certain societ-ies and civilizations have collapsed in the face of rapid andsevere climatic changes.

Centennial-Scale Variation

Historical records as well as proxy records (particularly

tree rings, corals, and ice cores) indicate that climate haschanged during the past 1,000 years at centennial times-cales; that is, no two centuries have been exactly alike.During the past 150 years, the Earth system has emergedfrom a period called the Little Ice Age, which was charac-terized in the North Atlantic region and elsewhere byrelatively cool temperatures. The 20th century in particu-

lar saw a substantial pattern of warming in many regions.Some of this warming may be attributable to the transi-tion from the Little Ice Age or other natural causes.However, many climate scientists believe that much of the20th-century warming, especially in the later decades,resulted from atmospheric accumulation of greenhouse

gases (especially carbon dioxide, CO2 ).

The Little Ice Age is best known in Europe and theNorth Atlantic region, which experienced relatively coolconditions between the early 14th and mid-19th centuries.This was not a period of uniformly cool climate, sinceinterannual and decadal variability brought many warm

years. Furthermore, the coldest periods did not alwayscoincide among regions; some regions experienced rela-

tively warm conditions at the same time others weresubjected to severely cold conditions. Alpine glaciersadvanced far below their previous (and present) limits,obliterating farms, churches, and villages in Switzerland,



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France, and elsewhere. Frequent cold winters and cool, wet summers ruined wine harvests and led to crop failuresand famines over much of northern and central Europe.The North Atlantic cod fisheries declined as ocean tem-peratures fell in the 17th century. The Norse colonies onthe coast of Greenland were cut off from the rest of Norsecivilization during the early 15th century as pack ice andstorminess increased in the North Atlantic. The westerncolony of Greenland collapsed through starvation, andthe eastern colony was abandoned. In addition, Iceland

became increasingly isolated from Scandinavia.The Little Ice Age was preceded by a period of rela-

tively mild conditions in northern and central Europe.This interval, known as the Medieval Warm Period,occurred from approximately 1000 CE to the first half ofthe 13th century. Mild summers and winters led to goodharvests in much of Europe. Wheat cultivation and vine-

yards flourished at far higher latitudes and elevations thantoday. Norse colonies in Iceland and Greenland prospered,and Norse parties fished, hunted, and explored the coastof Labrador and Newfoundland. The Medieval WarmPeriod is well documented in much of the North Atlanticregion, including ice cores from Greenland. Like the LittleIce Age, this time was neither a climatically uniform

period nor a period of uniformly warm temperatureseverywhere in the world. Other regions of the globe lackevidence for high temperatures during this period.

Much scientific attention continues to be devoted to aseries of severe droughts that occurred between the 11thand 14th centuries. These droughts, each spanning severaldecades, are well documented in tree-ring records across

western North America and in the peatland records of theGreat Lakes region. The records appear to be related toocean temperature anomalies in the Pacific and Atlanticbasins, but they are still inadequately understood. The

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information suggests that much of the United States issusceptible to persistent droughts that would be devastat-ing for water resources and agriculture.

Millennial and Multimillennial Variation

The climatic changes of the past thousand years are super-imposed upon variations and trends at both millennialtimescales and greater. Numerous indicators from easternNorth America and Europe show trends of increased cool-

ing and increased effective moisture during the past 3,000 years. For example, in the Great Lakes–St. Lawrenceregions along the U.S.-Canadian border, water levels of thelakes rose, peatlands developed and expanded, moisture-loving trees such as beech and hemlock expanded theirranges westward, and populations of boreal trees, such asspruce and tamarack, increased and expanded southward.

These patterns all indicate a trend of increased effectivemoisture, which may indicate increased precipitation,decreased evapotranspiration (water loss) due to cooling,or both. The patterns do not necessarily indicate a mono-lithic cooling event; more complex climatic changesprobably occurred. For example, beech expanded north-

ward and spruce southward during the past 3,000 years in

both eastern North America and western Europe. Thebeech expansions may indicate milder winters or longer growing seasons, whereas the spruce expansions appearrelated to cooler, moister summers. Paleoclimatologistsare applying a variety of approaches and proxies to helpidentify such changes in seasonal temperature and mois-ture during the Holocene Epoch.

Just as the Little Ice Age was not associated with coolconditions everywhere, so the cooling and moisteningtrend of the past 3,000 years was not universal. Someregions became warmer and drier during the same time



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period. For example, northern Mexico and the Yucatanexperienced decreasing moisture in the past 3,000 years.Heterogeneity of this type is characteristic of climaticchange, which involves changing patterns of atmosphericcirculation. As circulation patterns change, the transportof heat and moisture in the atmosphere also changes. Thisfact explains the apparent paradox of opposing tempera-ture and moisture trends in different regions.

The trends of the past 3,000 years are just the latestin a series of climatic changes that occurred over the past

11,700 years or so—the interglacial period referred to asthe Holocene Epoch. At the start of the Holocene, rem-nants of continental glaciers from the last glaciation stillcovered much of eastern and central Canada and partsof Scandinavia. These ice sheets largely disappeared by6,000 years ago. Their absence—along with increas-ing sea surface temperatures, rising sea levels (as glacial

meltwater flowed into the world’s oceans), and especiallychanges in the radiation budget of Earth’s surface owingto Milankovitch variations (changes in the seasons result-ing from periodic adjustments of Earth’s orbit aroundthe Sun)—affected atmospheric circulation. The diversechanges of the past 10,000 years across the globe are dif-ficult to summarize in brief, but some general highlights

and large-scale patterns are worthy of note. These includethe presence of early to mid-Holocene thermal maximain various locations, variation in ENSO patterns, and anearly to mid-Holocene amplification of the Indian Oceanmonsoon.

Thermal Maxima

Many parts of the globe experienced higher temperaturesthan today some time during the early to mid-Holocene. Insome cases the increased temperatures were accompaniedby decreased moisture availability. Although the thermal

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maximum has been referred to in North America and elsewhere as a single widespread event (variously referred toas the “Altithermal,” “Xerothermic Interval,” “ClimaticOptimum,” or “Thermal Optimum”), it is now recognizedthat the periods of maximum temperatures varied amongregions. For example, northwestern Canada experiencedits highest temperatures several thousand years earlierthan central or eastern North America. Similar heteroge-neity is seen in moisture records. For instance, the recordof the prairie-forest boundary in the Midwestern region of

the United States shows eastward expansion of prairie inIowa and Illinois 6,000 years ago (indicating increasinglydry conditions), whereas Minnesota’s forests expanded

westward into prairie regions at the same time (indicat-ing increasing moisture). The Atacama Desert, locatedprimarily in present-day Chile and Bolivia, on the westernside of South America, is one of the driest places on Earth

today, but it was much wetter during the early Holocene when many other regions were at their driest.

The primary driver of changes in temperature andmoisture during the Holocene was orbital variation, whichslowly changed the latitudinal and seasonal distribution ofsolar radiation on Earth’s surface and atmosphere.However, the heterogeneity of these changes was caused

by changing patterns of atmospheric circulation and oceancurrents.

ENSO Variation in the Holocene

Because of the global importance of ENSO variationtoday, Holocene variation in ENSO patterns and inten-sity is under serious study by paleoclimatologists. The

record is still fragmentary, but evidence from fossil corals,tree rings, lake records, climate modeling, and otherapproaches is accumulating that suggests that (1) ENSO



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variation was relatively weak in the early Holocene, (2)ENSO has undergone centennial to millennial variationsin strength during the past 11,700 years, and (3) ENSOpatterns and strength similar to those currently in placedeveloped within the past 5,000 years. This evidence isparticularly clear when comparing ENSO variation overthe past 3,000 years to today’s patterns. The causes oflong-term ENSO variation are still being explored, butchanges in solar radiation owing to Milankovitch varia-tions are strongly implicated by modeling studies.

The Amplification of the Indian Ocean Monsoon

Much of Africa, the Middle East, and the Indian subcon-tinent are under the strong influence of an annual climaticcycle known as the Indian Ocean monsoon. The climateof this region is highly seasonal, alternating between clearskies with dry air (winter) and cloudy skies with abundant

rainfall (summer). Monsoon intensity, like other aspectsof climate, is subject to interannual, decadal, and centen-nial variations, at least some of which are related to ENSOand other cycles. Abundant evidence exists for large varia-tions in monsoon intensity during the Holocene Epoch.Paleontological and paleoecological studies show thatlarge portions of the region experienced much greater

precipitation during the early Holocene (11,700–6,000 years ago) than today. Lake and wetland sediments datingto this period have been found under the sands of parts ofthe Sahara Desert. These sediments contain fossils of ele-phants, crocodiles, hippopotamuses, and giraffes, together

with pollen evidence of forest and woodland vegetation. Inarid and semiarid parts of Africa, Arabia, and India, large

and deep freshwater lakes occurred in basins that are nowdry or are occupied by shallow, saline lakes. Civilizationsbased on plant cultivation and grazing animals, such as the

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Harappan civilization of northwestern India and adjacentPakistan, flourished in these regions, which have sincebecome arid.

These and similar lines of evidence, together withpaleontological and geochemical data from marine sedi-ments and climate-modeling studies, indicate that theIndian Ocean monsoon was greatly amplified duringthe early Holocene, supplying abundant moisture farinland into the African and Asian continents. This ampli-fication was driven by high solar radiation in summer,

which was approximately 7 percent higher 11,700 yearsago than today and resulted from orbital forcing (changesin Earth’s eccentricity, precession, and axial tilt). Highsummer insolation resulted in warmer summer air tem-peratures and lower surface pressure over continentalregions and, hence, increased inflow of moisture-ladenair from the Indian Ocean to the continental interiors.

Modeling studies indicate that the monsoonal flow wasfurther amplified by feedbacks involving the atmosphere,

vegetation, and soils. Increased moisture led to wettersoils and lusher vegetation, which in turn led to increasedprecipitation and greater penetration of moist air intocontinental interiors. Decreasing summer insolation dur-ing the past 4,000–6,000 years led to the weakening of

the Indian Ocean monsoon.

CLIMATE CHANGE SINCE THEEMERGENCE OF HUMANS

The history of humanity—from the initial appearanceof genus Homo over 2,000,000 years ago to the advent

and expansion of the modern human species ( Homo sapi-ens ) beginning some 150,000 years ago—is integrallylinked to climate variation and change. Homo sapiens has



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experienced nearly two full glacial-interglacial cycles, butits global geographical expansion, massive populationincrease, cultural diversification, and worldwide ecologicaldomination began only during the last glacial period andaccelerated during the last glacial-interglacial transition.The first bipedal apes appeared in a time of climatic tran-sition and variation, and Homo erectus, an extinct speciespossibly ancestral to modern humans, originated duringthe colder Pleistocene Epoch and survived both the tran-sition period and multiple glacial-interglacial cycles. Thus,

it can be said that climate variation has been the midwifeof humanity and its various cultures and civilizations.

Recent Glacial and Interglacial Periods

With glacial ice restricted to high latitudes and altitudes,Earth 125,000 years ago was in an interglacial period simi-

lar to the one occurring today. During the past 125,000 years, however, the Earth system went through an entire glacial-interglacial cycle, only the most recent of manytaking place over the last million years. The most recentperiod of cooling and glaciation began approximately120,000 years ago. Significant ice sheets developed andpersisted over much of Canada and northern Eurasia.

After the initial development of glacial conditions,the Earth system alternated between two modes, one ofcold temperatures and growing glaciers and the otherof relatively warm temperatures (although much coolerthan today) and retreating glaciers. These Dansgaard-Oeschger (DO) cycles, recorded in both ice cores andmarine sediments, occurred approximately every 1,500

years. A lower-frequency cycle, called the Bond cycle, issuperimposed on the pattern of DO cycles; Bond cyclesoccurred every 3,000–8,000 years. Each Bond cycle is

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characterized by unusually cold conditions that takeplace during the cold phase of a DO cycle, the subsequentHeinrich event (which is a brief dry and cold phase), andthe rapid warming phase that follows each Heinrich event.During each Heinrich event, massive fleets of icebergs

were released into the North Atlantic, carrying rockspicked up by the glaciers far out to sea. Heinrich eventsare marked in marine sediments by conspicuous layers oficeberg-transported rock fragments.

Many of the transitions in the DO and Bond cycles

were rapid and abrupt, and they are being studied intenselyby paleoclimatologists and Earth system scientists tounderstand the driving mechanisms of such dramatic cli-matic variations. These cycles now appear to result frominteractions between the atmosphere, oceans, ice sheets,and continental rivers that influence thermohaline circu-lation (the pattern of ocean currents driven by differences

in water density, salinity, and temperature, rather than wind). Thermohaline circulation, in turn, controls oceanheat transport, such as the Gulf Stream.

During the past 25,000 years, the Earth system hasundergone a series of dramatic transitions. The mostrecent glacial period peaked 21,500 years ago duringthe Last Glacial Maximum, or LGM. At that time, the

northern third of North America was covered by theLaurentide Ice Sheet, which extended as far south as DesMoines, Iowa; Cincinnati, Ohio; and New York City. TheCordilleran Ice Sheet covered much of western Canada as

well as northern Washington, Idaho, and Montana in theUnited States. In Europe the Scandinavian Ice Sheet satatop the British Isles, Scandinavia, northeastern Europe,

and north-central Siberia. Montane glaciers were exten-sive in other regions, even at low latitudes in Africa andSouth America. Global sea level was 125 metres ( 410 feet)



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Paleoclimatology is the scientific study of the climatic conditions of

past geologic ages. Paleoclimatologists seek to explain climate varia-tions for all parts of the Earth during any given geologic period,beginning with the time of the Earth’s formation. Many related fieldscontribute to the field of paleoclimatology, but the basic research dataare drawn mainly from geology and paleobotany; speculative attemptsat explanation have come largely from astronomy, atmospheric phys-ics, meteorology, and geophysics.

Two major factors in the study of both ancient and present-dayclimatic conditions of the Earth are the changes in the relationshipbetween the Earth and the Sun (e.g., the slight alteration in the con-figuration of the Earth’s orbit) and the changes in the surface of theplanet itself (such phenomena as volcanic eruptions, mountain-build-ing events, the transformations of plant communities, and thedispersal of the continents after the breakup of the supercontinentPangea). Some of the questions that were studied in the past havebeen largely explained. Paleoclimatologists found, for example, thatthe warmth of the northern hemispheric landmasses during at least

90 percent of the last 570 million years is mainly due to the drift of thecontinents across the latitudes; until about 150 million years ago, bothNorth America and Europe were much closer to the Equator thanthey are today. Other questions, such as the reasons behind the irregu-lar advances and retreats of the ice sheets (i.e., glacial and interglacialepisodes), are much more difficult to explain, and no completely satis-factory theory has been presented.

below modern levels, because of the long-term net transferof water from the oceans to the ice sheets. Temperaturesnear Earth’s surface in unglaciated regions were about 5°C (9 °F) cooler than today. Many Northern Hemisphereplant and animal species inhabited areas far south of theirpresent ranges. For example, jack pine and white spruce

trees grew in northwestern Georgia, 1,000 km (600 miles)south of their modern range limits in the Great Lakesregion of North America.

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The Last Deglaciation

The continental ice sheets began to melt back about20,000 years ago. Drilling and dating of submerged fossilcoral reefs provide a clear record of increasing sea levels asthe ice melted. The most rapid melting began 15,000 yearsago. For example, the southern boundary of the LaurentideIce Sheet in North America was north of the Great Lakesand St. Lawrence regions by 10,000 years ago, and it hadcompletely disappeared by 6,000 years ago.

The warming trend was punctuated by transientcooling events, most notably the Younger Dryas climateinterval of 12,800–11,600 years ago. The climatic regimesthat developed during the deglaciation period in manyareas, including much of North America, have no mod-ern analog (i.e., no regions exist with comparable seasonalregimes of temperature and moisture). For example, in the

interior of North America, climates were much more con-tinental (that is, characterized by warm summers and cold

winters) than they are today. Also, paleontological studiesindicate assemblages of plant, insect, and vertebrate spe-cies that do not occur anywhere today. Spruce trees grew

with temperate hardwoods (ash, hornbeam, oak, and elm)in the upper Mississippi River and Ohio River regions. In

Alaska, birch and poplar grew in woodlands, and there were very few of the spruce trees that dominate the present-day Alaskan landscape. Boreal and temperate mammals,

whose geographic ranges are widely separated today, coex-isted in central North America and Russia during thisperiod of deglaciation. These unparalleled climatic condi-tions probably resulted from the combination of a unique

orbital pattern that increased summer insolation andreduced winter insolation in the Northern Hemisphereand the continued presence of Northern Hemisphere ice



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sheets, which themselves altered atmospheric circulationpatterns.

Climate Change and the Emergence ofAgriculture

The first known examples of animal domesticationoccurred in western Asia between 11,000 and 9,500 yearsago when goats and sheep were first herded, whereas exam-ples of plant domestication date to 9,000 years ago when

wheat, lentils, rye, and barley were first cultivated. This

phase of technological increase occurred during a time ofclimatic transition that followed the last glacial period.A number of scientists have suggested that, although cli-mate change imposed stresses on hunter-gatherer-foragersocieties by causing rapid shifts in resources, it also pro-

vided opportunities as new plant and animal resourcesappeared.

Glacial and Interglacial Cycles ofthe Pleistocene

The glacial period that peaked 21,500 years ago was onlythe most recent of five glacial periods in the last 450,000

years. In fact, the Earth system has alternated between

glacial and interglacial regimes for more than two mil-lion years, a period of time known as the Pleistocene.The duration and severity of the glacial periods increasedduring this period, with a particularly sharp change occur-ring between 900,000 and 600,000 years ago. Earth iscurrently within the most recent interglacial period, theHolocene Epoch.

The continental glaciations of the Pleistocene leftsignatures on the landscape in the form of glacial depos-its and landforms; however, the best knowledge of the

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The blue areas are those that were covered by ice sheets in the past. The Kansan and Nebraskan sheets overlapped almost the same areas, and the Wisconsin and Illinoisan sheets covered approximately the same territory. In the high altitudes of the West are the Cordilleran ice sheets. An area at the junction ofWisconsin, Minnesota, Iowa, and Illinois was never entirely covered with ice.Encyclopaedia Britannica, Inc.

magnitude and timing of the various glacial and intergla-cial periods comes from oxygen isotope records in oceansediments. These records provide both a direct measureof sea level and an indirect measure of global ice volume.Water molecules composed of a lighter isotope of oxygen,16O, are evaporated more readily than molecules bearinga heavier isotope, 18O. Glacial periods are characterizedby high 18O concentrations and represent a net transferof water, especially with 16O, from the oceans to the icesheets. Oxygen isotope records indicate that interglacial

periods have typically lasted 10,000–15,000 years, andmaximum glacial periods were of similar length. Mostof the past 500,000 years—approximately 80 percent—have been spent within various intermediate glacial states



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Europe, like North America, had four periods of glaciation. Successive ice caps reached limits that differed only slightly. The area covered by ice at any timeis shown in white. Encyclopaedia Britannica, Inc.

that were warmer than glacial maxima but cooler thaninterglacials. During these intermediate times, substantial

glaciers occurred over much of Canada and probably cov-ered Scandinavia as well. These intermediate states werenot constant; they were characterized by continual, mil-lennial-scale climate variation. There has been no averageor typical state for global climate during Pleistocene andHolocene times; the Earth system has been in continualflux between interglacial and glacial patterns.

The cycling of the Earth system between glacial and

interglacial modes has been ultimately driven by orbital variations. However, orbital forcing is by itself insuf-ficient to explain all of this variation, and Earth systemscientists are focusing their attention on the interactionsand feedbacks between the myriad components of theEarth system. For example, the initial development of a

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continental ice sheet increases albedo over a portion ofEarth, reducing surface absorption of sunlight and leadingto further cooling. Similarly, changes in terrestrial vegeta-tion, such as the replacement of forests by tundra, feedback into the atmosphere via changes in both albedo andlatent heat flux from evapotranspiration. Forests—par-ticularly those of tropical and temperate areas, with theirlarge leaf area—release great amounts of water vapour andlatent heat through transpiration. Tundra plants, whichare much smaller, possess tiny leaves designed to slow

water loss; they release only a small fraction of the water vapour that forests do.

The discovery in ice core records that atmosphericconcentrations of two potent greenhouse gases, carbondioxide and methane, have decreased during past glacialperiods and peaked during interglacials indicates impor-tant feedback processes in the Earth system. Reduction of

greenhouse gas concentrations during the transition to a glacial phase would reinforce and amplify cooling alreadyunder way. The reverse is true for transition to intergla-cial periods. The glacial carbon sink remains a topic ofconsiderable research activity. A full understanding of gla-cial-interglacial carbon dynamics requires knowledge ofthe complex interplay among ocean chemistry and circula-

tion, ecology of marine and terrestrial organisms, ice sheetdynamics, and atmospheric chemistry and circulation.

The Last Great Cooling

The Earth system has undergone a general cooling trendfor the past 50 million years, culminating in the develop-

ment of permanent ice sheets in the Northern Hemisphereabout 2.75 million years ago. These ice sheets expandedand contracted in a regular rhythm, with each glacial



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maximum separated from adjacent ones by 41,000 years(based on the cycle of axial tilt). As the ice sheets waxedand waned, global climate drifted steadily toward coolerconditions characterized by increasingly severe glacia-tions and increasingly cool interglacial phases. Beginningaround 900,000 years ago, the glacial-interglacial cyclesshifted frequency. Ever since, the glacial peaks have been100,000 years apart, and the Earth system has spent moretime in cool phases than before. The 41,000-year period-icity has continued, with smaller fluctuations superimposed

on the 100,000-year cycle. In addition, a smaller, 23,000- year cycle has occurred through both the 41,000-year and100,000-year cycles.

The 23,000-year and 41,000-year cycles are drivenultimately by two components of Earth’s orbital geome-try: the equinoctial precession cycle (23,000 years) and theaxial-tilt cycle (41,000 years). Although the third param-

eter of Earth’s orbit, eccentricity, varies on a 100,000-yearcycle, its magnitude is insufficient to explain the 100,000-

year cycles of glacial and interglacial periods of the past900,000 years. The origin of the periodicity present inEarth’s eccentricity is an important question in currentpaleoclimate research.

CLIMATE CHANGE THROUGHGEOLOGIC TIME

The Earth system has undergone dramatic changesthroughout its 4.5-billion-year history. These haveincluded climatic changes diverse in mechanisms, mag-nitudes, rates, and consequences. Many of these past

changes are obscure and controversial, and some havebeen discovered only recently. Nevertheless, the historyof life has been strongly influenced by these changes,

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some of which radically altered the course of evolution.Life itself is implicated as a causative agent of some ofthese changes, as the processes of photosynthesis andrespiration have largely shaped the chemistry of Earth’satmosphere, oceans, and sediments.

Cenozoic Climates

The Cenozoic Era—encompassing the past 65.5 million years, the time that has elapsed since the mass extinction

event marking the end of the Cretaceous Period—has abroad range of climatic variation characterized by alter-nating intervals of global warming and cooling. Earth hasexperienced both extreme warmth and extreme cold dur-ing this period. These changes have been driven by tectonicforces, which have altered the positions and elevations ofthe continents as well as ocean passages and bathymetry.

Feedbacks between different components of the Earthsystem (atmosphere, biosphere, lithosphere, cryosphere,and oceans in the hydrosphere) are being increasingly rec-ognized as influences of global and regional climate. Inparticular, atmospheric concentrations of carbon dioxidehave varied substantially during the Cenozoic for reasonsthat are poorly understood, though its fluctuation must

have involved feedbacks between Earth’s spheres.Orbital forcing is also evident in the Cenozoic,although, when compared on such a vast era-level times-cale, orbital variations can be seen as oscillations againsta slowly changing backdrop of lower-frequency climatictrends. Descriptions of the orbital variations have evolvedaccording to the growing understanding of tectonic and

biogeochemical changes. A pattern emerging from recentpaleoclimatologic studies suggests that the climaticeffects of eccentricity, precession, and axial tilt have been



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amplified during cool phases of the Cenozoic, whereasthey have been dampened during warm phases.

The meteor impact that occurred at or very close tothe end of the Cretaceous came at a time of global warm-ing, which continued into the early Cenozoic. Tropicaland subtropical flora and fauna occurred at high latitudesuntil at least 40 million years ago, and geochemical recordsof marine sediments have indicated the presence of warmoceans. The interval of maximum temperature occurredduring the late Paleocene and early Eocene epochs (58.7 to

40.4 million years ago). The highest global temperaturesof the Cenozoic occurred during the Paleocene-EoceneThermal Maximum (PETM), a short interval lastingapproximately 100,000 years. Although the underlyingcauses are unclear, the onset of the PETM about 56 mil-lion years ago was rapid, occurring within a few thousand

years, and ecological consequences were large, with wide-

spread extinctions in both marine and terrestrialecosystems. Sea surface and continental air temperaturesincreased by more than 5 °C (9 °F) during the transitioninto the PETM. Sea surface temperatures in the high-lati-tude Arctic may have been as warm as 23 °C (73 °F),comparable to modern subtropical and warm-temperateseas. Following the PETM, global temperatures declined

to pre-PETM levels, but they gradually increased to near-PETM levels over the next few million years during aperiod known as the Eocene Optimum. This temperaturemaximum was followed by a steady decline in global tem-peratures toward the Eocene-Oligocene boundary, whichoccurred about 33.9 million years ago. These changes are

well-represented in marine sediments and in paleontologi-

cal records from the continents, where vegetation zonesmoved Equator-ward. Mechanisms underlying the coolingtrend are under study, but it is most likely that tectonic

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movements played an important role. This period saw the gradual opening of the sea passage between Tasmania andAntarctica, followed by the opening of the Drake Passagebetween South America and Antarctica. The latter, whichisolated Antarctica within a cold polar sea, produced

global effects on atmospheric and oceanic circulation.Recent evidence suggests that decreasing atmosphericconcentrations of carbon dioxide during this period mayhave initiated a steady and irreversible cooling trend overthe next few million years.

A continental ice sheet developed in Antarctica duringthe Oligocene Epoch, persisting until a rapid warmingevent took place 27 million years ago. The late Oligoceneand early to mid-Miocene epochs (28.4 to 13.8 million

years ago) were relatively warm, though not nearly as warmas the Eocene. Cooling resumed 15 million years ago, andthe Antarctic Ice Sheet expanded again to cover much

of the continent. The cooling trend continued throughthe late Miocene and accelerated into the early PlioceneEpoch, 5.3 million years ago. During this period theNorthern Hemisphere remained ice-free, and paleobo-tanical studies show cool-temperate Pliocene floras athigh latitudes on Greenland and the Arctic Archipelago.The Northern Hemisphere glaciation, which began 3.2

million years ago, was driven by tectonic events, such asthe closing of the Panama seaway and the uplift of theAndes, the Tibetan Plateau, and western parts of NorthAmerica. These tectonic events led to changes in the cir-culation of the oceans and the atmosphere, which in turnfostered the development of persistent ice at high north-ern latitudes. Small-magnitude variations in carbon

dioxide concentrations, which had been relatively lowsince at least the mid-Oligocene (28.4 million years ago),are also thought to have contributed to this glaciation.



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Phanerozoic Climates

The Phanerozoic Eon (542 million years ago to thepresent), which includes the entire span of complex, multi-cellular life on Earth, has witnessed an extraordinary arrayof climatic states and transitions. The sheer antiquity ofmany of these regimes and events renders them difficultto understand in detail. However, a number of periodsand transitions are well known, owing to good geologi-cal records and intense study by scientists. Furthermore,

a coherent pattern of low-frequency climatic variation isemerging, in which the Earth system alternates between

warm (“greenhouse”) phases and cool (“icehouse”) phases.The warm phases are characterized by high temperatures,high sea levels, and an absence of continental glaciers.Cool phases in turn are marked by low temperatures, lowsea levels, and the presence of continental ice sheets, at

least at high latitudes. Superimposed on these alterna-tions are higher-frequency variations, where cool periodsare embedded within greenhouse phases and warm peri-ods are embedded within icehouse phases. For example,

glaciers developed for a brief period (between 1 millionand 10 million years) during the late Ordovician and earlySilurian ( about 430 million years ago), in the middle of the

early Paleozoic greenhouse phase (560 million to 350 mil-lion years ago). Similarly, warm periods with glacial retreatoccurred within the late Cenozoic cool period during thelate Oligocene and early Miocene epochs.

The Earth system has been in an icehouse phase forthe past 30 million to 35 million years, ever since the devel-opment of ice sheets on Antarctica. The previous major

icehouse phase occurred between about 350 million and 250million years ago, during the Carboniferous and Permianperiods of the late Paleozoic Era. Glacial sediments dating

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to this period have been identified in much of Africa as well as in the Arabian Peninsula, South America, Australia,India, and Antarctica. At the time, all these regions werepart of Gondwana, a high-latitude supercontinent inthe Southern Hemisphere. The glaciers atop Gondwanaextended to at least 45° S latitude, similar to the latitudereached by Northern Hemisphere ice sheets during thePleistocene. Some late Paleozoic glaciers extended evenfurther Equator-ward—to 35° S. One of the most strik-ing features of this time period are cyclothems, repeating

sedimentary beds of alternating sandstone, shale, coal,and limestone. The great coal deposits of North America’sAppalachian region, the American Midwest, and northernEurope are interbedded in these cyclothems, which mayrepresent repeated transgressions (producing limestone)and retreats (producing shales and coals) of ocean shore-lines in response to orbital variations.

The two most prominent warm phases in Earth his-tory occurred during the Mesozoic and early Cenozoiceras (approximately 250 million to 35 million years ago)and the early and mid-Paleozoic ( approximately 500 mil-lion to 350 million years ago). Climates of each of these

greenhouse periods were distinct; continental positionsand ocean bathymetry were very different, and terrestrial

vegetation was absent from the continents until relativelylate in the Paleozoic warm period. Both of these periodsexperienced substantial long-term climate variation andchange; increasing evidence indicates brief glacial epi-sodes during the mid-Mesozoic.

Understanding the mechanisms underlying icehouse- greenhouse dynamics is an important area of research,

involving an interchange between geologic recordsand the modeling of the Earth system and its compo-nents. Two processes have been implicated as drivers ofPhanerozoic climate change. First, tectonic forces caused



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changes in the positions and elevations of continents andthe bathymetry of oceans and seas. Second, variations in

greenhouse gases were also important drivers of climate,though at these long timescales they were largely con-trolled by tectonic processes, in which sinks and sourcesof greenhouse gases varied.

The Climates of Early Earth

The pre-Phanerozoic interval, also known as Precambrian

time, comprises some 88 percent of the time elapsedsince the origin of Earth. The pre-Phanerozoic is a poorlyunderstood phase of Earth system history. Much of thesedimentary record of the atmosphere, oceans, biota, andcrust of the early Earth has been obliterated by erosion,metamorphosis, and subduction. However, a number ofpre-Phanerozoic records have been found in various parts

of the world, mainly from the later portions of the period.Pre-Phanerozoic Earth system history is an extremelyactive area of research, in part because of its importancein understanding the origin and early evolution of lifeon Earth. Furthermore, the chemical composition ofEarth’s atmosphere and oceans largely developed duringthis period, with living organisms playing an active role.

Geologists, paleontologists, microbiologists, planetary geologists, atmospheric scientists, and geochemists arefocusing intense efforts on understanding this period.Three areas of particular interest and debate are the “faint

young Sun paradox,” the role of organisms in shapingEarth’s atmosphere, and the possibility that Earth wentthrough one or more “snowball” phases of global glaciation.

The Faint Young Sun Paradox

Astrophysical studies indicate that the luminosity of theSun was much lower during Earth’s early history than it

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has been in the Phanerozoic. In fact, radiative output was low enough to suggest that all surface water on Earthshould have been frozen solid during its early history,but evidence shows that it was not. The solution to this“faint young Sun paradox” appears to lie in the presenceof unusually high concentrations of greenhouse gasesat the time, particularly methane and carbon dioxide.As solar luminosity gradually increased through time,concentrations of greenhouse gases would have to havebeen much higher than today. This circumstance would

have caused Earth to heat up beyond life-sustaininglevels. Therefore, greenhouse gas concentrations musthave decreased proportionally with increasing solarradiation, implying a feedback mechanism to regulate

greenhouse gases. One of these mechanisms might havebeen rock weathering, which is temperature-dependentand serves as an important sink for, rather than source

of, carbon dioxide by removing sizable amounts of this gas from the atmosphere. Scientists are also lookingto biological processes (many of which also serve ascarbon dioxide sinks) as complementary or alterna-tive regulating mechanisms of greenhouse gases on the

young Earth.

Photosynthesis and Atmospheric Chemistry

The evolution by photosynthetic bacteria of a new photo-synthetic pathway, substituting water (H2O) for hydrogensulfide (H2S) as a reducing agent for carbon dioxide, haddramatic consequences for Earth system geochemistry.Molecular oxygen (O2 ) is given off as a by-product ofphotosynthesis using the H2O pathway, which is energeti-

cally more efficient than the more primitive H2S pathway.Using H2O as a reducing agent in this process led to thelarge-scale deposition of banded-iron formations, or



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BIFs, a source of 90 percent of present-day iron ores.Oxygen present in ancient oceans oxidized dissolvediron, which precipitated out of solution onto the oceanfloors. This deposition process, in which oxygen was usedup as fast as it was produced, continued for millions of

years until most of the iron dissolved in the oceans wasprecipitated. By approximately 2 billion years ago, oxygen

was able to accumulate in dissolved form in seawater andto outgas to the atmosphere. Although oxygen does nothave greenhouse gas properties, it plays important indi-

rect roles in Earth’s climate, particularly in phases of thecarbon cycle. Scientists are studying the role of oxygenand other contributions of early life to the developmentof the Earth system.

The “Snowball Earth” and

“Slushball Earth” Hypotheses

Geochemical and sedimentary evidence indicates thatEarth experienced as many as four extreme cooling eventsbetween 750 million and 580 million years ago. Geologistshave proposed that Earth’s oceans and land surfaces werecovered by ice from the poles to the Equator during theseevents. This “Snowball Earth” hypothesis is a subject ofintense study and discussion. Two important questions

arise from this hypothesis. First, how, once frozen, couldEarth thaw? Second, how could life survive periods of global freezing? A proposed solution to the first questioninvolves the outgassing of massive amounts of carbondioxide by volcanoes, which could have warmed the plan-etary surface rapidly, especially given that major carbondioxide sinks (rock weathering and photosynthesis) would

have been dampened by a frozen Earth. A possible answerto the second question may lie in the existence of present-day life-forms within hot springs and deep-sea vents,

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The Younger Dryas event was characterized by a substantial and relatively sudden drop in temperature between 12,800 and 11,600 years ago. In additionto cold regions, the evidence of this temperature change has been discovered intropical and subtropical regions. Encyclopaedia Britannica, Inc.

which would have persisted long ago despite the frozenstate of Earth’s surface.

A counter-premise known as the “Slushball Earth”hypothesis contends that Earth was not completely fro-zen over. Rather, in addition to massive ice sheets coveringthe continents, parts of the planet (especially ocean areasnear the Equator) could have been draped only by a thin,

watery layer of ice amid areas of open sea. Under this sce-nario, photosynthetic organisms in low-ice or ice-freeregions could continue to capture sunlight efficiently and

survive these periods of extreme cold.



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ABRUPT CLIMATE CHANGES INEARTH HISTORY

An important new area of research, abrupt climate change,has developed since the 1980s. This research has beeninspired by the discovery, in the ice core records ofGreenland and Antarctica, of evidence for abrupt shifts inregional and global climates of the past. These events,

which have also been documented in ocean and continen-tal records, involve sudden shifts of Earth’s climate system

from one equilibrium state to another. Such shifts are ofconsiderable scientific concern because they can revealsomething about the controls and sensitivity of the cli-mate system. In particular, they point out nonlinearities,the so-called “tipping points,” where small, gradualchanges in one component of the system can lead to alarge change in the entire system. Such nonlinearities arise

from the complex feedbacks between components of theEarth system. For example, during the Younger Dryasevent a gradual increase in the release of fresh water to theNorth Atlantic Ocean led to an abrupt shutdown of thethermohaline circulation in the Atlantic basin. Abrupt cli-mate shifts are of great societal concern, for any suchshifts in the future might be so rapid and radical as to out-

strip the capacity of agricultural, ecological, industrial,and economic systems to respond and adapt. Climate sci-entists are working with social scientists, ecologists, andeconomists to assess society’s vulnerability to such “cli-mate surprises.”

The Younger Dryas event (12,800 to 11,600 years ago)is the most intensely studied and best-understood exam-

ple of abrupt climate change. The event took place duringthe last deglaciation, a period of global warming when theEarth system was in transition from a glacial mode to aninterglacial one. The Younger Dryas was marked by a

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sharp drop in temperatures in the North Atlantic region;cooling in northern Europe and eastern North America isestimated at 4 to 8 °C (7.2 to 14.4 °F). Terrestrial and marinerecords indicate that the Younger Dryas had detectableeffects of lesser magnitude over most other regions ofEarth. The termination of the Younger Dryas was veryrapid, occurring within a decade. The Younger Dryasresulted from an abrupt shutdown of the thermohalinecirculation in the North Atlantic, which is critical for thetransport of heat from equatorial regions northward

(today the Gulf Stream is a part of that circulation). Thecause of the shutdown of the thermohaline circulation isunder study; an influx of large volumes of freshwater frommelting glaciers into the North Atlantic has been impli-cated, although other factors probably played a role.

Paleoclimatologists are devoting increasing attentionto identifying and studying other abrupt changes. The

Dansgaard-Oeschger cycles of the last glacial period arenow recognized as representing alternation between twoclimate states, with rapid transitions from one state tothe other. A 200-year-long cooling event in the NorthernHemisphere approximately 8,200 years ago resulted fromthe rapid draining of glacial Lake Agassiz into the NorthAtlantic via the Great Lakes and St. Lawrence drainage.

This event, characterized as a miniature version of theYounger Dryas, had ecological impacts in Europe andNorth America that included a rapid decline of hemlockpopulations in New England forests. In addition, evi-dence of another such transition, marked by a rapid dropin the water levels of lakes and bogs in eastern NorthAmerica, occurred 5,200 years ago. It is recorded in ice

cores from glaciers at high altitudes in tropical regions as well as tree-ring, lake-level, and peatland samples fromtemperate regions.



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Abrupt climatic changes occurring before thePleistocene have also been documented. A transientthermal maximum has been documented near thePaleocene-Eocene boundary (55.8 million years ago), andevidence of rapid cooling events are observed near theboundaries between both the Eocene and Oligoceneepochs (33.9 million years ago) and the Oligocene andMiocene epochs (23 million years ago). All three of theseevents had global ecological, climatic, and biogeochemicalconsequences. Geochemical evidence indicates that the

warm event occurring at the Paleocene-Eocene boundary was associated with a rapid increase in atmospheric car-bon dioxide concentrations, possibly resulting from themassive outgassing and oxidation of methane hydrates(a compound whose chemical structure traps methane

within a lattice of ice) from the ocean floor. The two cool-ing events appear to have resulted from a transient series

of positive feedbacks among the atmosphere, oceans, icesheets, and biosphere, similar to those observed in thePleistocene. Other abrupt changes, such as the Paleocene-Eocene Thermal Maximum, are recorded at various pointsin the Phanerozoic.

Abrupt climate changes can evidently be caused by a variety of processes. Rapid changes in an external factor

can push the climate system into a new mode. Outgassingof methane hydrates and the sudden influx of glacialmeltwater into the ocean are examples of such externalforcing. Alternatively, gradual changes in external factorscan lead to the crossing of a threshold; the climate systemis unable to return to the former equilibrium and passesrapidly to a new one. Such nonlinear system behaviour is

a potential concern as human activities, such as fossil-fuelcombustion and land-use change, alter important compo-nents of Earth’s climate system.

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Humans and other species have survived countless cli-matic changes in the past, and humans are a notablyadaptable species. Adjustment to climatic changes,

whether it is biological (as in the case of other species) orcultural (for humans), is easiest and least catastrophic

when the changes are gradual and can be anticipated tolarge extent. Rapid changes are more difficult to adaptto and incur more disruption and risk. Abrupt changes,especially unanticipated climate surprises, put human cul-tures and societies, as well as both the populations of

other species and the ecosystems they inhabit, at consid-erable risk of severe disruption. Such changes may well be

within humanity’s capacity to adapt, but not withoutpaying severe penalties in the form of economic, ecologi-cal, agricultural, human health, and other disruptions.Knowledge of past climate variability provides guidelineson the natural variability and sensitivity of the Earth

system. This knowledge also helps identify the risksassociated with altering the Earth system with greenhouse

gas emissions and regional to global-scale changes inland cover.



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GLOBAL W ARMING

G lobal warming is the phenomenon of increasing aver-age air temperatures near the surface of Earth over

the past one to two centuries. Since the mid-20th century,climate scientists have gathered detailed observations of

various weather phenomena (such as temperature, precip-itation, and storms) and of related influences on climate(such as ocean currents and the atmosphere’s chemical

composition). These data indicate that Earth’s climate haschanged over almost every conceivable timescale sincethe beginning of geologic time and that, since at least thebeginning of the Industrial Revolution, the influence ofhuman activities has been deeply woven into the very fab-ric of climate change.

Giving voice to a growing conviction of most of the

scientific community, the Intergovernmental Panel onClimate Change (IPCC) reported that the 20th centurysaw an increase in global average surface temperature ofapproximately 0.6 °C (1.1 °F). The IPCC went on to statethat most of the warming observed over the second half ofthe 20th century could be attributed to human activities,and it predicted that by the end of the 21st century the

average surface temperature would increase by another 1.8to 4.0 °C (3.2 to 7.2 °F), depending on a range of possiblescenarios. Many climate scientists agree that significanteconomic and ecological damage would result if globalaverage temperatures rose by more than 2 °C [3.6 °F] insuch a short time. Such damage might include increasedextinction of many plant and animal species, shifts in

patterns of agriculture, and rising sea levels. The IPCCreported that the global average sea level rose by some 17cm (6.7 inches) during the 20th century, that sea levels rose



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During the second half of the 20th century and early part of the 21st century, global average surface temperature increased and sea level rose. Over the same period, the amount of snow cover in the Northern Hemisphere decreased.



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faster in the second half of that century than in the firsthalf, and that—again depending on a wide range of sce-narios—the global average sea level could rise by another18 to 59 cm (7 to 23 inches) by the end of the 21st century.Furthermore, the IPCC reported that average snow coverin the Northern Hemisphere declined by 4 percent, or 1.5million square km (580,000 square miles), between 1920and 2005.

The scenarios referred to above depend mainlyon future concentrations of certain trace gases, called

greenhouse gases, that have been injected into the loweratmosphere in increasing amounts through the burningof fossil fuels for industry, transportation, and residentialuses. Modern global warming is the result of an increasein magnitude of the so-called greenhouse effect, a warm-ing of Earth’s surface and lower atmosphere caused by thepresence of water vapour, carbon dioxide, methane, and

other greenhouse gases. Of all these gases, carbon dioxideis the most important, both for its role in the greenhouseeffect and for its role in the human economy. It has beenestimated that, at the beginning of the industrial age inthe mid-18th century, carbon dioxide concentrationsin the atmosphere were roughly 280 parts per million(ppm). By the end of the 20th century, carbon dioxide

concentrations had reached 369 ppm (possibly the high-est concentrations in at least 650,000 years), and, if fossilfuels continue to be burned at current rates, they areprojected to reach 560 ppm by the mid-21st century—essentially, a doubling of carbon dioxide concentrationsin 300 years. It has been calculated that an increase ofthis magnitude alone (that is, not accounting for possible

effects of other greenhouse gases) would be responsiblefor adding 2 to 5 °C (3.6 to 9 °F) to the global average sur-face temperatures that existed at the beginning of theindustrial age.




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A vigorous debate is in progress over the extent andseriousness of rising surface temperatures, the effects ofpast and future warming on human life, and the need foraction to reduce future warming and deal with its conse-quences. Although climate scientists state that evidenceof global warming is unequivocal, sustained by decades ofresearch involving numerous independent investigations,they face continued criticism by skeptics, who contendthat the mechanisms of climate are not fully understood.Skeptics often state that climate variations can be

explained by natural cycles alone, noting that there is littleevidence linking climate changes to human activities.They often contest the temperature and rainfall projec-tions made by climate scientists.

CAUSES OF GLOBAL WARMING

Global warming is largely the product of rising greenhouse gas concentrations in the atmosphere; however, manyother factors also contribute. The concentration of volca-nic aerosols in the atmosphere, the Sun’s ever-increasingsolar energy output, and cyclical changes in Earth’s orbitare natural influencers on Earth’s climate. In addition, sev-eral factors can interact with one another to produce

feedback loops, which can accelerate or dampen Earth’snear-surface air temperatures.

The Greenhouse Effect

The average surface temperature of Earth is maintainedby a balance of various forms of solar and terrestrial radia-

tion. Solar radiation is often called “shortwave” radiationbecause the frequencies of the radiation are relatively highand the wavelengths relatively short—close to the vis-ible portion of the electromagnetic spectrum. Terrestrial



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radiation, on the other hand, is often called “longwave”radiation because the frequencies are relatively lowand the wavelengths relatively long—somewhere in theinfrared part of the spectrum. Downward-moving solarenergy is typically measured in watts per square metre.The energy of the total incoming solar radiation at thetop of Earth’s atmosphere (the so-called “solar constant”)amounts roughly to 1,366 watts per square metre annually.Adjusting for the fact that only one-half of the planet’s sur-face receives solar radiation at any given time, the average

surface insolation is 342 watts per square metre annually.The amount of solar radiation absorbed by Earth’s sur-

face is only a small fraction of the total solar radiationentering the atmosphere. For every 100 units of incomingsolar radiation, roughly 30 units are reflected back to spaceby either clouds, the atmosphere, or reflective regionsof Earth’s surface. This reflective capacity is referred to as

Earth’s planetary albedo, and it need not remain fixed overtime, since the spatial extent and distribution of reflectiveformations, such as clouds and ice cover, can change. The70 units of solar radiation that are not reflected may beabsorbed by the atmosphere, clouds, or the surface. In theabsence of further complications, in order to maintainthermodynamic equilibrium, Earth’s surface and atmo-

sphere must radiate these same 70 units back to space.Earth’s surface temperature (and that of the lower layer ofthe atmosphere essentially in contact with the surface) istied to the magnitude of this emission of outgoing radia-tion according to the Stefan-Boltzmann law.

Earth’s energy budget is further complicated by the greenhouse effect. Trace gases with certain chemical

properties—the so-called greenhouse gases, mainly car-bon dioxide (CO2 ), methane (CH4 ), and nitrous oxide(N2O)—absorb some of the infrared radiation pro-duced by Earth’s surface. Because of this absorption,

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some fraction of the original 70 units does not directlyescape to space. Because greenhouse gases emit the sameamount of radiation they absorb and because this radia-tion is emitted equally in all directions (that is, as muchdownward as upward), the net effect of absorption by

greenhouse gases is to increase the total amount of radia-tion emitted downward toward Earth’s surface and loweratmosphere. To maintain equilibrium, Earth’s surface andlower atmosphere must emit more radiation than theoriginal 70 units. Consequently, the surface temperature

must be higher. This process is not quite the same as that which governs a true greenhouse, but the end effect issimilar. The presence of greenhouse gases in the atmo-sphere leads to a warming of the surface and lower partof the atmosphere (and a cooling higher up in the atmo-sphere) relative to what would be expected in the absenceof greenhouse gases.

It is essential to distinguish the “natural,” or background, greenhouse effect from the “enhanced” greenhouseeffect associated with human activity. The natural green-house effect is associated with surface warming propertiesof natural constituents of Earth’s atmosphere, especially

water vapour, carbon dioxide, and methane. The existenceof this effect is accepted by all scientists. Indeed, in its

absence, Earth’s average temperature would be approxi-mately 33 °C (59 °F) colder than today, and Earth would be afrozen and likely uninhabitable planet. What has been sub-ject to controversy is the so-called enhanced greenhouseeffect, which is associated with increased concentrationsof greenhouse gases caused by human activity. In particu-lar, the burning of fossil fuels raises the concentrations of

the major greenhouse gases in the atmosphere, and thesehigher concentrations have the potential to warm theatmosphere by several degrees.



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Since 1750 the concentration of carbon dioxide and other greenhouse gases

has increased in Earth’s atmosphere. As a result of these and other factors, Earth’s atmosphere retains more heat than in the past. EncyclopaediaBritannica, Inc.

Radiative Forcing

In light of this discussion of the greenhouse effect, it is

apparent that the temperature of Earth’s surface and loweratmosphere may be modified in three ways: (1) through anet increase in the solar radiation entering at the top ofEarth’s atmosphere, (2) through a change in the fractionof the radiation reaching the surface, and (3) through achange in the concentration of greenhouse gases in theatmosphere. In each case the changes can be thought of

in terms of “radiative forcing.” As defined by the IPCC,radiative forcing is a measure of the influence a given cli-matic factor has on the amount of downward-directed

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radiant energy impinging upon Earth’s surface. Climaticfactors are divided between those caused primarily byhuman activity (such as greenhouse gas emissions andaerosol emissions) and those caused by natural forces(such as solar irradiance); then, for each factor, so-calledforcing values are calculated for the time period between1750 and the present day. “Positive forcing” is exerted byclimatic factors that contribute to the warming of Earth’ssurface, whereas “negative forcing” is exerted by factorsthat cool Earth’s surface.

On average about 342 watts of solar radiation strikeeach square metre of Earth’s surface per year, and thisquantity can in turn be related to a rise or fall in Earth’ssurface temperature. Temperatures at the surface may alsorise or fall through a change in the distribution of terrestrialradiation (that is, radiation emitted by Earth) within theatmosphere. In some cases, radiative forcing has a natural

origin, such as during explosive eruptions from volcanoes where vented gases and ash block some portion of solarradiation from the surface. In other cases, radiative forc-ing has an anthropogenic, or exclusively human, origin.For example, anthropogenic increases in carbon dioxide,methane, and nitrous oxide are estimated to account for2.3 watts per square metre of positive radiative forcing.

When all values of positive and negative radiative forcingare taken together and all interactions between climaticfactors are accounted for, the total net increase in surfaceradiation due to human activities since the beginning ofthe Industrial Revolution is 1.6 watts per square metre.

The Influences of Human Activity on Climate

Human activity has influenced global surface tempera-tures by changing the radiative balance governing theEarth on various timescales and at varying spatial scales.



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Petroleum refinery at Ras Tanura, Saudi Arabia. Herbert Lanks—Shostal

The most profound and well-known anthropogenic influ-ence is the elevation of concentrations of greenhouse

gases in the atmosphere. Humans also influence climateby changing the concentrations of aerosols and ozone andby modifying the land cover of Earth’s surface.

Greenhouse GasesAs discussed above, greenhouse gases warm Earth’s surfaceby increasing the net downward longwave radiation reach-ing the surface. The relationship between atmosphericconcentration of greenhouse gases and the associatedpositive radiative forcing of the surface is different foreach gas. A complicated relationship exists between the

chemical properties of each greenhouse gas and the rela-tive amount of longwave radiation that each can absorb.What follows is a discussion of the radiative behaviour ofeach major greenhouse gas.

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The present-day surface hydrologic cycle, in which water is transferred fromthe oceans through the atmosphere to the continents and back to the oceans over

and beneath the land surface. The values in parentheses following the various

forms of water (e.g., ice) refer to volumes in millions of cubic kilometres; those following the processes (e.g., precipitation) refer to their fluxes in millions ofcubic kilometres of water per year. Encyclopædia Britannica, Inc.

Water Vapour

Water vapour is the most potent of the greenhouse gasesin Earth’s atmosphere, but its behaviour is fundamentallydifferent from that of the other greenhouse gases. Theprimary role of water vapour is not as a direct agent ofradiative forcing but rather as a climate feedback—thatis, as a response within the climate system that influ-ences the system’s continued activity. This distinctionarises from the fact that the amount of water vapour in

the atmosphere cannot, in general, be directly modifiedby human behaviour but is instead set by air temperatures.The warmer the surface, the greater the evaporation rate



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In this illustration of the carbon cycle, Carbon is transported in various formsthrough the atmosphere, the hydrosphere, and geologic formations. One ofthe primary pathways for the exchange of carbon dioxide (CO2 ) takes placebetween the atmosphere and the oceans; there a fraction of the CO2 combineswith water, forming carbonic acid (H 2CO3 ) that subsequently loses hydrogen

ions (H + ) to form bicarbonate (HCO3− ) and carbonate (CO32− ) ions. Mollusk shells or mineral precipitates that form by the reaction of calcium or other metal ions with carbonate may become buried in geologic strata and eventually release CO2 through volcanic outgassing. Carbon dioxide also exchangesthrough photosynthesis in plants and through respiration in animals. Dead

and decaying organic matter may ferment and release CO2 or methane (CH 4 )or may be incorporated into sedimentary rock, where it is converted to fos-

sil fuels. Burning of hydrocarbon fuels returns CO2 and water (H 2O) to the atmosphere. The biological and anthropogenic pathways are much faster

than the geochemical pathways and, consequently, have a greater impacton the composition and temperature of the atmosphere. EncyclopædiaBritannica, Inc.

of water from the surface. As a result, increased evapora-tion leads to a greater concentration of water vapour inthe lower atmosphere capable of absorbing longwave radi-ation and emitting it downward.

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Carbon Dioxide

Of the greenhouse gases, carbon dioxide (CO2

) is most sig-nificant. Natural sources of atmospheric CO2 includeoutgassing from volcanoes, the combustion and naturaldecay of organic matter, and respiration by aerobic (oxy-

gen-using) organisms. These sources are balanced, onaverage, by a set of physical, chemical, or biological pro-cesses, called “sinks,” that tend to remove CO2 from theatmosphere. Significant natural sinks include terrestrial

vegetation, which takes up CO2 during the process ofphotosynthesis.

A number of oceanic processes also act as carbon sinks.One such process, called the “solubility pump,” involvesthe descent of surface sea water containing dissolved CO2.Another process, the “biological pump,” involves theuptake of dissolved CO2 by marine vegetation and phyto-

plankton (small, free-floating, photosynthetic organisms)living in the upper ocean or by other marine organismsthat use CO2 to build skeletons and other structures madeof calcium carbonate (CaCO3 ). As these organisms expireand fall to the ocean floor, the carbon they contain istransported downward and eventually buried at depth. Along-term balance between these natural sources and

sinks leads to the background, or natural, level of CO2 inthe atmosphere.In contrast, human activities increase atmospheric

CO2 levels primarily through the burning of fossil fuels(principally oil and coal, and secondarily natural gas, foruse in transportation, heating, and the generation ofelectrical power) and through the production of cement.

Other anthropogenic sources include the burning offorests and the clearing of land. Anthropogenic emis-sions currently account for the annual release of about 7



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gigatons (7 billion tons) of carbon into the atmosphere.Anthropogenic emissions are equal to approximately 3percent of the total emissions of CO

2

by natural sources,and this amplified carbon load from human activitiesfar exceeds the offsetting capacity of natural sinks (byperhaps as much as 2–3 gigatons per year). CO2 has conse-quently accumulated in the atmosphere at an average rateof 1.4 parts per million (ppm) by volume per year between1959 and 2006, and this rate of accumulation has been lin-ear (that is, uniform over time). However, certain current

sinks, such as the oceans, could become sources in thefuture. This may lead to a situation in which the concen-tration of atmospheric CO2 builds at an exponential rate.

The natural background level of carbon dioxide varieson timescales of millions of years due to slow changes inoutgassing through volcanic activity. For example, roughly100 million years ago, during the Cretaceous Period, CO2

concentrations appear to have been several times higherthan today (perhaps close to 2,000 ppm). Over the past700,000 years, CO2 concentrations have varied over a farsmaller range (between roughly 180 and 300 ppm) in asso-ciation with the same Earth orbital effects linked to thecoming and going of the Pleistocene ice ages. By the early21st century, CO2 levels reached 384 ppm, which is approx-

imately 37 percent above the natural background level ofroughly 280 ppm that existed at the beginning of theIndustrial Revolution. According to ice core measure-ments, this level (384 ppm) is believed to be the highest inat least 650,000 years.

Radiative forcing caused by carbon dioxide varies in anapproximately logarithmic fashion with the concentration

of that gas in the atmosphere. The logarithmic relation-ship occurs as the result of a saturation effect whereinit becomes increasingly difficult, as CO2 concentrations

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increase, for additional CO2 molecules to further influ-ence the “infrared window” (a certain narrow band of

wavelengths in the infrared region that is not absorbed byatmospheric gases). The logarithmic relationship predictsthat the surface warming potential will rise by roughly thesame amount for each doubling of CO2 concentration. Atcurrent rates of fossil-fuel use, a doubling of CO2 concen-trations over preindustrial levels is expected to take placeby the middle of the 21st century (when CO2 concentra-tions are projected to reach 560 ppm). A doubling of CO2

concentrations would represent an increase of roughly 4 watts per square metre of radiative forcing. Given typi-cal estimates of “climate sensitivity” in the absence ofany offsetting factors, this energy increase would leadto a warming of 2 to 5 °C (3.6 to 9 °F) over preindustrialtimes. The total radiative forcing by anthropogenic CO2 emissions since the beginning of the industrial age is

approximately 1.66 watts per square metre.

Methane

Methane (CH4 ) is the second most important greenhouse gas. CH4 is more potent than CO2 because the radiativeforcing produced per molecule is greater. In addition, theinfrared window is less saturated in the range of wave-

lengths of radiation absorbed by CH4, so more moleculesmay fill in the region. However, CH4 exists in far lowerconcentrations than CO2 in the atmosphere, and its con-centrations by volume in the atmosphere are generallymeasured in parts per billion (ppb) rather than ppm. CH4 also has a considerably shorter residence time in the atmo-sphere than CO2 (the residence time for CH4 is roughly 10

years, compared with hundreds of years for CO2 ).Natural sources of methane include tropical andnorthern wetlands, methane-oxidizing bacteria that feedon organic material consumed by termites, volcanoes,



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seepage vents of the seafloor in regions rich with organicsediment, and methane hydrates trapped along the conti-nental shelves of the oceans and in polar permafrost. Theprimary natural sink for methane is the atmosphere itself,as methane reacts readily with the hydroxyl radical (OH- ) within the troposphere to form CO2 and water vapour(H2O). When CH4 reaches the stratosphere, it is destroyed.Another natural sink is soil, where methane is oxidized bybacteria.

As with CO2, human activity is increasing the CH4

concentration faster than it can be offset by natural sinks.Anthropogenic sources currently account for approxi-mately 70 percent of total annual emissions, leading tosubstantial increases in concentration over time. Themajor anthropogenic sources of atmospheric CH4 are ricecultivation, livestock farming, the burning of coal and nat-ural gas, the combustion of biomass, and the decomposition

of organic matter in landfills. Future trends are particu-larly difficult to anticipate. This is in part due to anincomplete understanding of the climate feedbacks asso-ciated with CH4 emissions. In addition, as humanpopulations grow, it is difficult to predict how possiblechanges in livestock raising, rice cultivation, and energyutilization will influence CH4 emissions.

It is believed that a sudden increase in the concentra-tion of methane in the atmosphere was responsible for a warming event that raised average global temperaturesby 4–8 °C (7.2–14.4 °F) over a few thousand years duringthe so-called Paleocene-Eocene Thermal Maximum, orPETM. This episode took place roughly 55 million yearsago, and the rise in CH4 appears to have been related to

a massive volcanic eruption that interacted with meth-ane-containing flood deposits. As a result, large amountsof gaseous CH4 were injected into the atmosphere. It isdifficult to know precisely how high these concentrations

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were or how long they persisted. At very high concen-trations, residence times of CH4 in the atmosphere canbecome much greater than the nominal 10-year residencetime that applies today. Nevertheless, it is likely that theseconcentrations reached several ppm during the PETM.

Methane concentrations have also varied over asmaller range (between roughly 350 and 800 ppb) in asso-ciation with the Pleistocene ice age cycles. Preindustriallevels of CH4 in the atmosphere were approximately700 ppb, whereas early 21st-century levels exceeded 1,770

ppb. (These concentrations are well above the naturallevels observed for at least the past 650,000 years.) Thenet radiative forcing by anthropogenic CH4 emissions isapproximately 0.5 watt per square metre—or roughly one-third the radiative forcing of CO2.

Surface-Level Ozone and Other Compounds

The next most significant greenhouse gas is surface, orlow-level, ozone (O3 ). Surface O3 is a result of air pollution;it must be distinguished from naturally occurring strato-spheric O3, which has a very different role in the planetaryradiation balance. The primary natural source of surfaceO3 is the subsidence of stratospheric O3 from the upperatmosphere. In contrast, the primary anthropogenic

source of surface O3 is photochemical reactions involvingthe atmospheric pollutant carbon monoxide (CO). Thebest estimates of the concentration of surface O3 are 50ppb, and the net radiative forcing due to anthropogenicemissions of surface O3 is approximately 0.35 watt persquare metre.

Nitrous Oxides and Fluorinated Gases

Additional trace gases produced by industrial activity thathave greenhouse properties include nitrous oxide (N2O)and fluorinated gases (halocarbons), the latter including



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sulfur hexafluoride, hydrofluorocarbons (HFCs), and per-fluorocarbons (PFCs). Nitrous oxide is responsible for0.16 watt per square metre radiative forcing, while fluori-nated gases are collectively responsible for 0.34 watt persquare metre. Nitrous oxides have small background con-centrations due to natural biological reactions in soil and

water, whereas the fluorinated gases owe their existencealmost entirely to industrial sources.

Aerosols

The production of aerosols represents an importantanthropogenic radiative forcing of climate. Collectively,aerosols block—that is, reflect and absorb—a portion ofincoming solar radiation, and this creates a negative radia-tive forcing. Aerosols are second only to greenhouse gasesin relative importance in their impact on near-surface airtemperatures. Unlike the decade-long residence times of

the “well-mixed” greenhouse gases, such as CO2 and CH4,aerosols are readily flushed out of the atmosphere withindays, either by rain or snow (wet deposition) or by settlingout of the air (dry deposition). They must therefore becontinually generated in order to produce a steady effecton radiative forcing. Aerosols have the ability to influenceclimate directly by absorbing or reflecting incoming solar

radiation, but they can also produce indirect effects on cli-mate by modifying cloud formation or cloud properties.Most aerosols serve as condensation nuclei (surfaces upon

which water vapour can condense to form clouds); how-ever, darker-coloured aerosols may hinder cloud formationby absorbing sunlight and heating up the surrounding air.Aerosols can be transported thousands of kilometres from

their sources of origin by winds and upper-level circula-tion in the atmosphere.Perhaps the most important type of anthropogenic

aerosol in radiative forcing is sulfate aerosol. It is produced

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from sulfur dioxide (SO2 ) emissions associated with theburning of coal and oil. Since the late 1980s, global emis-sions of SO

2

have decreased from about 73 million tons toabout 54 million tons of sulfur per year.

Nitrate aerosol is not as important as sulfate aerosol,but it has the potential to become a significant source ofnegative forcing. One major source of nitrate aerosol issmog (the combination of ozone with oxides of nitrogenin the lower atmosphere) released from the incompleteburning of fuel in internal-combustion engines. Another

source is ammonia (NH3 ), which is often used in fertilizersor released by the burning of plants and other organicmaterials. If greater amounts of atmospheric nitrogen areconverted to ammonia and agricultural ammonia emis-sions continue to increase as projected, the influence ofnitrate aerosols on radiative forcing is expected to grow.

Both sulfate and nitrate aerosols act primarily by

reflecting incoming solar radiation, thereby reducing theamount of sunlight reaching the surface. Most aerosols,unlike greenhouse gases, impart a cooling rather than

warming influence on Earth’s surface. One prominentexception is carbonaceous aerosols such as carbon blackor soot, which are produced by the burning of fossil fuelsand biomass. Carbon black tends to absorb rather than

reflect incident solar radiation, and so it has a warmingimpact on the lower atmosphere, where it resides. Becauseof its absorptive properties, carbon black is also capable ofhaving an additional indirect effect on climate. Throughits deposition in snowfall, it can decrease the albedo ofsnow cover. This reduction in the amount of solar radia-tion reflected back to space by snow surfaces creates a

minor positive radiative forcing.Natural forms of aerosol include windblown mineraldust generated in arid and semiarid regions and sea saltproduced by the action of waves breaking in the ocean.



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Changes to wind patterns as a result of climate modi-fication could alter the emissions of these aerosols. Theinfluence of climate change on regional patterns of arid-ity could shift both the sources and the destinations ofdust clouds. In addition, since the concentration of seasalt aerosol, or sea aerosol, increases with the strength ofthe winds near the ocean surface, changes in wind speeddue to global warming and climate change could influencethe concentration of sea salt aerosol. For example, somestudies suggest that climate change might lead to stronger

winds over parts of the North Atlantic Ocean. Areas withstronger winds may experience an increase in the concen-tration of sea salt aerosol.

Other natural sources of aerosols include volcaniceruptions, which produce sulfate aerosol, and biogenicsources (e.g., phytoplankton), which produce dimethylsulfide (DMS). Other important biogenic aerosols, such

as terpenes, are produced naturally by certain kindsof trees or other plants. For example, the dense forests ofthe Blue Ridge Mountains of Virginia in the United Statesemit terpenes during the summer months, which in turninteract with the high humidity and warm temperaturesto produce a natural photochemical smog. Anthropogenicpollutants such as nitrate and ozone, both of which serve

as precursor molecules for the generation of biogenicaerosol, appear to have increased the rate of productionof these aerosols severalfold. This process appears to beresponsible for some of the increased aerosol pollution inregions undergoing rapid urbanization.

Human activity has greatly increased the amount ofaerosol in the atmosphere compared with the background

levels of preindustrial times. In contrast to the globaleffects of greenhouse gases, the impact of anthropogenicaerosols is confined primarily to the Northern Hemisphere,

where most of the world’s industrial activity occurs. The

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pattern of increases in anthropogenic aerosol over time isalso somewhat different from that of greenhouse gases.During the middle of the 20th century, there was a sub-stantial increase in aerosol emissions. This appears to havebeen at least partially responsible for a cessation of surface

warming that took place in the Northern Hemispherefrom the 1940s through the 1970s. Since that time, aerosolemissions have leveled off due to antipollution measuresundertaken in the industrialized countries since the 1960s.Aerosol emissions may rise in the future, however, as a

result of the rapid emergence of coal-fired electric power generation in China and India.

The total radiative forcing of all anthropogenic aero-sols is approximately –1.2 watts per square metre. Of thistotal, –0.5 watt per square metre comes from direct effects(such as the reflection of solar energy back into space), and

–0.7 watt per square metre comes from indirect effects

(such as the influence of aerosols on cloud formation).This negative radiative forcing represents an offset ofroughly 40 percent from the positive radiative forcingcaused by human activity. However, the relative uncer-tainty in aerosol radiative forcing (approximately 90percent) is much greater than that of greenhouse gases. Inaddition, future emissions of aerosols from human activi-

ties, and the influence of these emissions on future climatechange, are not known with any certainty. Nevertheless, itcan be said that, if concentrations of anthropogenic aero-sols continue to decrease as they have since the 1970s, asignificant offset to the effects of greenhouse gases will bereduced, opening future climate to further warming.

Land-Use Change

There are a number of ways in which changes in land usecan influence climate. The most direct influence is throughthe alteration of Earth’s albedo, or surface reflectance. For



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example, the replacement of forest by cropland and pas-ture in the middle latitudes over the past several centurieshas led to an increase in albedo, which in turn has led to

greater reflection of incoming solar radiation in thoseregions. This replacement of forest by agriculture hasbeen associated with a change in global average radiativeforcing of approximately –0.2 watt per square metre since1750. In Europe and other major agricultural regions, suchland-use conversion began more than 1,000 years ago andhas proceeded nearly to completion. For Europe, the neg-

ative radiative forcing due to land-use change has probablybeen substantial, perhaps approaching –5 watts per squaremetre. The influence of early land use on radiative forcingmay help to explain a long period of cooling in Europe thatfollowed a period of relatively mild conditions roughly1,000 years ago. It is generally believed that the mild tem-peratures of this “medieval warm period,” which was

followed by a long period of cooling, rivaled those of 20th-century Europe.

Land-use changes can also influence climate throughtheir influence on the exchange of heat between Earth’ssurface and the atmosphere. For example, vegetationhelps to facilitate the evaporation of water into the atmo-sphere through evapotranspiration. In this process, plants

take up liquid water from the soil through their rootsystems. Eventually this water is released through transpi-ration into the atmosphere, as water vapour through thestomata in leaves. While deforestation generally leads tosurface cooling due to the albedo factor discussed above,the land surface may also be warmed as a result of therelease of latent heat by the evapotranspiration process.

The relative importance of these two factors, one exert-ing a cooling effect and the other a warming effect, variesby both season and region. While the albedo effect islikely to dominate in middle latitudes, especially during

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the period from autumn through spring, the evapotrans-piration effect may dominate during the summer in themidlatitudes and year-round in the tropics. The latter caseis particularly important in assessing the potential impactsof continued tropical deforestation.

The rate at which tropical regions are deforested isalso relevant to the process of carbon sequestration, thelong-term storage of carbon in underground cavities andbiomass rather than in the atmosphere. By removing car-bon from the atmosphere, carbon sequestration acts

to mitigate global warming. Deforestation contributes to global warming as fewer plants are available to take up car-bon dioxide from the atmosphere. In addition, as fallentrees, shrubs, and other plants are burned or allowed toslowly decompose, they release as carbon dioxide the car-bon they stored during their lifetimes. Furthermore, anyland-use change that influences the amount, distribution,

or type of vegetation in a region can affect the concentra-tions of biogenic aerosols; however, the impact of suchchanges on climate is indirect and relatively minor.

Stratospheric Ozone Depletion

Since the 1970s the loss of ozone (O3 ) from the strato-sphere has led to a small amount of negative radiative

forcing of the surface. This negative forcing represents acompetition between two distinct effects caused by thefact that ozone absorbs solar radiation. In the first case, asozone levels in the stratosphere are depleted, more solarradiation reaches Earth’s surface. In the absence of anyother influence, this rise in insolation would represent apositive radiative forcing of the surface. However, there is

a second effect of ozone depletion that is related to its greenhouse properties. As the amount of ozone in thestratosphere is decreased, there is also less ozone toabsorb longwave radiation emitted by Earth’s surface.



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The long-term storage of carbon in plants, soils, geologic formations,and the ocean is called carbon sequestration. This process occursboth naturally and as a result of anthropogenic activities and typicallyrefers to the storage of carbon that has the immediate potential tobecome carbon dioxide gas. In response to growing concerns aboutclimate change resulting from increased carbon dioxide concentra-tions in the atmosphere, considerable interest has been drawn to thepossibility of increasing the rate of carbon sequestration throughchanges in land use and forestry and also through geoengineeringtechniques such as carbon capture and storage.

Carbon sources and carbon sinks

Anthropogenic activities such as the burning of fossil fuels have releasedcarbon from its long-term geologic storage as coal, petroleum, and nat-ural gas and have delivered it to the atmosphere as carbon dioxide gas.

Carbon dioxide is also released naturally, through the combustion anddecomposition of plants and animals. The amount of carbon dioxide inthe atmosphere has increased since the beginning of the industrial age,

Forests, such as this one found in the Adirondack Mountains near Keene Valley, New York, are vast storehouses of carbon. Jerome Wyckoff

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and this increase is caused mainly by the burning of fossil fuels. Carbondioxide is a very effective greenhouse gas—that is, a gas that absorbs

infrared radiation emitted from Earth’s surface. As carbon dioxide con-centrations rise in the atmosphere, more infrared radiation is retained,and the average temperature of Earth’s lower atmosphere rises. Thisprocess is referred to as global warming.

Reservoirs that retain carbon and keep it from entering Earth’satmosphere are known as carbon sinks. For example, deforestation isa source of carbon emission into the atmosphere, but forest regrowthis a form of carbon sequestration, the forests themselves serving ascarbon sinks. Carbon is transferred naturally from the atmosphere to

terrestrial carbon sinks through photosynthesis; it may be stored inaboveground biomass as well as in soils. Beyond the natural growth ofplants, other terrestrial processes that sequester carbon include growth of replacement vegetation on cleared land, land-managementpractices that absorb carbon, and increased growth due to elevatedatmospheric carbon dioxide levels and enhanced nitrogen deposition.It is important to note that carbon sequestered in soils and above ground vegetation could be released again to the atmosphere through

land-use or climatic changes. For example, combustion (which iscaused by fires) or decomposition (which results from microbe infes-tation) can cause the release of carbon stored in forests to the

The generalized carbon cycle. Encyclopædia Britannica, Inc.



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atmosphere. Both processes join oxygen in the air with carbon storedin plant tissues to produce carbon dioxide gas.

If the terrestrial sink becomes a significant carbon source throughincreased combustion and decomposition, it has the potential to addlarge amounts of carbon to the atmosphere and oceans. Globally, thetotal amount of carbon in vegetation, soil, and detritus is roughly2,200 gigatons (1 gigaton = 1 billion tons), and it is estimated that theamount of carbon sequestered annually by terrestrial ecosystems isapproximately 2.6 gigatons. The oceans themselves also accumulatecarbon, and the amount found just under the surface is roughly 920

gigatons. The amount of carbon stored in the oceanic sink exceeds theamount in the atmosphere (about 760 gigatons). Of the carbon emit-ted to the atmosphere by human activities, only 45 percent remains inthe atmosphere; about 30 percent is taken up by the oceans, and theremainder is incorporated into terrestrial ecosystems.

Carbon sequestration and climate change mitigation

The Kyoto Protocol under the United Nations Framework Conventionon Climate Change allows countries to receive credits for their car-bon-sequestration activities in the area of land use, land-use change,and forestry as part of their obligations under the protocol. Suchactivities could include afforestation (conversion of nonforested landto forest), reforestation (conversion of previously forested land to for-est), improved forestry or agricultural practices, and revegetation.According to the Intergovernmental Panel on Climate Change(IPCC), improved agricultural practices and forest-related mitigationactivities can make a significant contribution to the removal of carbondioxide from the atmosphere at relatively low cost. These activitiescould include improved crop and grazing land management—forinstance, more efficient fertilizer use to prevent the leaching of unusednitrates, tillage practices that minimize soil erosion, the restoration oforganic soils, and the restoration of degraded lands.

Carbon capture and storage

Some policy makers, engineers, and scientists seeking to mitigate global warming have proposed new technologies of carbon

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With less absorption of radiation by ozone, there is a cor-responding decrease in the downward re-emission ofradiation. This second effect overwhelms the first and

results in a modest negative radiative forcing of Earth’ssurface and a modest cooling of the lower stratosphere byapproximately 0.5 °C (0.9 °F) per decade since the 1970s.

Natural Influences on Climate

There are a number of natural factors that influence Earth’s

climate. These factors include external influences suchas explosive volcanic eruptions, natural variations in theoutput of the Sun, and slow changes in the configuration

sequestration. These technologies include a geoengineering proposalcalled carbon capture and storage (CCS). In CCS processes, carbon

dioxide is first separated from other gases contained in industrialemissions. It is then compressed and transported to a location that isisolated from the atmosphere for long-term storage. Suitable storagelocations might include geologic formations such as deep saline for-mations (sedimentary rocks whose pore spaces are saturated with water containing high concentrations of dissolved salts), depleted oiland gas reservoirs, or the deep ocean. Although CCS typically refersto the capture of carbon dioxide directly at the source of emission

before it can be released into the atmosphere, it may also includetechniques such as the use of scrubbing towers and “artificial trees” toremove carbon dioxide from the surrounding air.

There are many economic and technical challenges to implement-ing carbon capture and storage on a large scale. The IPCC has estimatedthat carbon capture and storage would increase the cost of electricity generation by about one to five cents per kilowatt-hour, depending onthe fuel, technology, and location. Leakage of carbon from reservoirsis also a concern, but it is estimated that properly managed geological

storage is very likely (that is, 66–90 percent probability) to retain 99percent of its sequestered carbon dioxide for over 1,000 years.



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of Earth’s orbit relative to the Sun. In addition, there arenatural oscillations in Earth’s climate that alter globalpatterns of wind circulation, precipitation, and surfacetemperatures. One such phenomenon is the El Niño/ Southern Oscillation (ENSO), a coupled atmospheric andoceanic event that occurs in the Pacific Ocean every threeto seven years. In addition, the Atlantic MultidecadalOscillation (AMO) is a similar phenomenon that occursover decades in the North Atlantic Ocean. Other typesof oscillatory behaviour that produce dramatic shifts

in climate may occur across timescales of centuries andmillennia.

Volcanic Aerosols

Explosive volcanic eruptions have the potential to injectsubstantial amounts of sulfate aerosols into the lowerstratosphere. In contrast to aerosol emissions in the

lower troposphere, aerosols that enter the stratospheremay remain for several years before settling out, becauseof the relative absence of turbulent motions there.Consequently, aerosols from explosive volcanic eruptionshave the potential to affect Earth’s climate. Less explosiveeruptions, or eruptions that are less vertical in orienta-tion, have a lower potential for substantial climate impact.

Furthermore, because of large-scale circulation patterns within the stratosphere, aerosols injected within tropicalregions tend to spread out over the globe, whereas aero-sols injected within midlatitude and polar regions tendto remain confined to the middle and high latitudes ofthat hemisphere. Tropical eruptions, therefore, tend tohave a greater climatic impact than eruptions occurring

toward the poles. In 1991 the moderate eruption of MountPinatubo in the Philippines provided a peak forcing ofapproximately –4 watts per square metre and cooled the

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climate by about 0.5 °C (0.9 °F) over the following few years. By comparison, the 1815 Mount Tambora eruptionin present-day Indonesia, typically implicated for the 1816“year without a summer” in Europe and North America, isbelieved to have been associated with a radiative forcingof approximately –6 watts per square metre.

While in the stratosphere, volcanic sulfate aerosolactually absorbs longwave radiation emitted by Earth’ssurface, and absorption in the stratosphere tends to resultin a cooling of the troposphere below. This vertical pat-

tern of temperature change in the atmosphere influencesthe behaviour of winds in the lower atmosphere, primarilyin winter. Thus, while there is essentially a global coolingeffect for the first few years following an explosive vol-canic eruption, changes in the winter patterns of surface

winds may actually lead to warmer winters in some areas,such as Europe. Some modern examples of major erup-

tions include Krakatoa (Indonesia) in 1883, El Chichón(Mexico) in 1982, and Mount Pinatubo in 1991. There isalso evidence that volcanic eruptions may influence otherclimate phenomena such as ENSO.

The Variations in Solar Output

Direct measurements of solar irradiance, or solar out-

put, have been available from satellites only since thelate 1970s. These measurements show a very small peak-to-peak variation in solar irradiance (roughly 0.1 percentof the 1,366 watts per square metre received at the top ofthe atmosphere, for approximately 0.12 watt per squaremetre). However, indirect measures of solar activity areavailable from historical sunspot measurements dating

back through the early 17th century. Attempts have beenmade to reconstruct graphs of solar irradiance variationsfrom historical sunspot data by calibrating them against



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the measurements from modern satellites; however,since the modern measurements span only a few of themost recent 11-year solar cycles, estimates of solar output

variability on 100-year and longer timescales are poorly

correlated. Different assumptions regarding the relation-ship between the amplitudes of 11-year solar cycles andlong-period solar output changes can lead to considerabledifferences in the resulting solar reconstructions. Thesedifferences in turn lead to fairly large uncertainty in esti-mating positive forcing by changes in solar irradiance since1750. (Estimates range from 0.06 to 0.3 watt per square

metre.) Even more challenging, given the lack of any mod-ern analog, is the estimation of solar irradiance during theso-called Maunder Minimum, a period lasting from the

Twelve solar X-ray images obtained by Yohkoh between 1991 and 1995. The solar coronal brightness decreases by a factor of about 100 during a solar cycle

as the Sun goes from an “active” state (left) to a less active state (right). G.L.Slater and G.A. Linford; S.L. Freeland; the Yohkoh Project

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mid-17th century to the early 18th century when very fewsunspots were observed. While it is likely that solar irra-diance was reduced at this time, it is difficult to calculateby how much. However, additional proxies of solar outputexist that match reasonably well with the sunspot-derivedrecords following the Maunder Minimum; these may beused as crude estimates of the solar irradiance variations.

In theory it is possible to estimate solar irradianceeven farther back in time, over at least the past millen-nium, by measuring levels of cosmogenic isotopes such ascarbon-14 and beryllium-10. Cosmogenic isotopes are

The trend shown in the longer reconstruction was inferred by Lean (2000) from modeling the changes in the brightness of stars similar to the Sun. Thetrend depicted in the shorter reconstruction by Y. Wang et al. (2005) was basedon a magnetic flux model that simulated the long-term evolution of faculae(bright granular structures on the Sun’s surface). Both models track a slightincrease in solar irradiance since 1900.



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isotopes that are formed by interactions of cosmic rays with atomic nuclei in the atmosphere and that subse-

quently fall to Earth, where they can be measured in theannual layers found in ice cores. Since their productionrate in the upper atmosphere is modulated by changes insolar activity, cosmogenic isotopes may be used as indirectindicators of solar irradiance. However, as with the sun-spot data, there is still considerable uncertainty in theamplitude of past solar variability implied by these data.

Solar forcing also affects the photochemical reactionsthat manufacture ozone in the stratosphere. Throughthis modulation of stratospheric ozone concentrations,changes in solar irradiance (particularly in the ultraviolet

Monthly satellite measurements of total solar irradiance since 1980 compar-ing NASA’s ACRIMSAT data by Willson and Mordvinov (2003) with the

Physikalisch-Meteorologisches Observatorium Davos (PMOD) composite developed by Fröhlich and Lean (2004). The PMOD composite combines ACRIM data collected by the Solar Maximum Mission (SMM) and Upper Atmosphere Research Satellite (UARS) with those provided by the Solar and Heliospheric Observatory (SOHO) and Nimbus 7 satellites.

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The layers of Earth’s atmosphere. The yellow line shows the response of airtemperature to increasing height. Encyclopædia Britannica, Inc.

portion of the electromagnetic spectrum) can modify howboth shortwave and longwave radiation in the lower strato-sphere are absorbed. As a result, the vertical temperatureprofile of the atmospheric can change, and this change inturn can influence phenomena such as the strength of the

winter jet streams.



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Earth’s axis of rotation itself rotates, or precesses, completing one circle every 26,000 years. Consequently, Earth’s North Pole points toward different stars(and sometimes toward empty space) as it travels in this circle. This precession

is so slow that it is not noticeable in a person’s lifetime, though astronomers must consider its effect when studying ancient sites such as Stonehenge.Encyclopædia Britannica, Inc.

The Variations in Earth’s Orbit

On timescales of tens of millennia, the dominant radiativeforcing of Earth’s climate is associated with slow varia-tions in the geometry of Earth’s orbit about the Sun. These

variations include the precession of the equinoxes (that is,

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changes in the timing of summer and winter), occurringon a roughly 26,000-year timescale; changes in the tiltangle of Earth’s rotational axis relative to the plane ofEarth’s orbit around the Sun, occurring on a roughly41,000-year timescale; and changes in the eccentricity(the departure from a perfect circle) of Earth’s orbitaround the Sun, occurring on a roughly 100,000-year tim-escale. Changes in eccentricity slightly influence the meanannual solar radiation at the top of Earth’s atmosphere,but the primary influence of all the orbital variations listed

above is on the seasonal and latitudinal distribution ofincoming solar radiation over Earth’s surface. The majorice ages of the Pleistocene Epoch were closely related tothe influence of these variations on summer insolation athigh northern latitudes. Orbital variations thus exerted aprimary control on the extent of continental ice sheets.However, Earth’s orbital changes are generally believed to

have had little impact on climate over the past few millen-nia, and so they are not considered to be significant factorsin present-day climate variability.

Feedback Mechanisms and Climate Sensitivity

There are a number of feedback processes important to

Earth’s climate system and, in particular, its responseto external radiative forcing. The most fundamental ofthese feedback mechanisms involves the loss of longwaveradiation to space from the surface. Since this radiativeloss increases with increasing surface temperaturesaccording to the Stefan-Boltzmann law, it represents a sta-bilizing factor (that is, a negative feedback) with respect

to near-surface air temperature.Climate sensitivity can be defined as the amountof surface warming resulting from each additional wattper square metre of radiative forcing. Alternatively, it is



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sometimes defined as the warming that would result froma doubling of CO2 concentrations and the associated addi-tion of 4 watts per square metre of radiative forcing. In theabsence of any additional feedbacks, climate sensitivity

would be approximately 0.25 °C (0.45 °F) for each addi-tional watt per square metre of radiative forcing. Statedalternatively, if the CO2 concentration of the atmospherepresent at the start of the industrial age (280 ppm) weredoubled (to 560 ppm), the resulting additional 4 watts persquare metre of radiative forcing would translate into a

1 °C (1.8 °F) increase in air temperature. However, thereare additional feedbacks that exert a destabilizing, ratherthan stabilizing, influence, and these feedbacks tend toincrease the sensitivity of climate to somewhere between0.5 and 1.0 °C (0.9 and 1.8 °F) for each additional watt persquare metre of radiative forcing.

Water Vapour Feedback Unlike concentrations of other greenhouse gases, the con-centration of water vapour in the atmosphere cannotfreely vary. Instead, it is determined by the temperature ofthe lower atmosphere and surface through a physical rela-tionship known as the Clausius-Clapeyron equation,named for 19th-century German physicist Rudolf Clausius

and 19th-century French engineer Émile Clapeyron.Under the assumption that there is a liquid water surfacein equilibrium with the atmosphere, this relationshipindicates that an increase in the capacity of air to hold

water vapour is a function of increasing temperature ofthat volume of air. This assumption is relatively good overthe oceans, where water is plentiful, but not over the con-

tinents. For this reason the relative humidity (the percentof water vapour the air contains relative to its capacity) isapproximately 100 percent over ocean regions and muchlower over continental regions (approaching 0 percent in

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arid regions). Not surprisingly, the average relative humid-ity of Earth’s lower atmosphere is similar to the fraction ofEarth’s surface covered by the oceans (that is, roughly 70percent). This quantity is expected to remain approxi-mately constant as Earth warms or cools. Slight changes to

global relative humidity may result from human land-usemodification, such as tropical deforestation and irriga-tion, which can affect the relative humidity over land areasup to regional scales.

The amount of water vapour in the atmosphere will

rise as the temperature of the atmosphere rises. Since water vapour is a very potent greenhouse gas, even morepotent than CO2, the net greenhouse effect actuallybecomes stronger as the surface warms, which leads toeven greater warming. This positive feedback is known asthe “water vapour feedback.” It is the primary reason thatclimate sensitivity is substantially greater than the previ-

ously stated theoretical value of 0.25 °C (0.45 °F) for eachincrease of 1 watt per square metre of radiative forcing.

Cloud Feedbacks

It is generally believed that as Earth’s surface warms andthe atmosphere’s water vapour content increases, globalcloud cover increases; however, the effects on near-sur-

face air temperatures are complicated. In the case of lowclouds, such as marine stratus clouds, the dominant radi-ative feature of the cloud is its albedo. Here any increasein low cloud cover acts in much the same way as anincrease in surface ice cover: more incoming solar radia-tion is reflected and Earth’s surface cools. On the otherhand, high clouds, such as the towering cumulus clouds

that extend up to the boundary between the troposphereand stratosphere, have a quite different impact on thesurface radiation balance. The tops of cumulus clouds areconsiderably higher in the atmosphere and colder than



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their undersides. Cumulus cloud tops emit less longwaveradiation out to space than the warmer cloud bottomsemit downward toward the surface. The end result of theformation of high cumulus clouds is greater warming atthe surface.

The net feedback of clouds on rising surface tempera-tures is therefore somewhat uncertain. It represents acompetition between the impacts of high and low clouds,and the balance is difficult to determine. Nonetheless,most estimates indicate that clouds on the whole repre-

sent a positive feedback and thus additional warming.

Ice Albedo Feedback

The so-called ice albedo feedback arises from the simplefact that ice is more reflective (that is, has a higher albedo)than land or water surfaces. Therefore, as global ice coverdecreases, the reflectivity of Earth’s surface decreases,

more incoming solar radiation is absorbed by the surface,and the surface warms. This feedback is considerablymore important when there is relatively extensive globalice cover, such as during the height of the last ice age,roughly 25,000 years ago. On a global scale the importanceof ice albedo feedback decreases as Earth’s surface warmsand there is relatively less ice available to be melted.

Carbon Cycle Feedbacks

Another important set of climate feedbacks involves the global carbon cycle. In particular, the two main reservoirsof carbon in the climate system are the oceans and the ter-restrial biosphere. These reservoirs have historically takenup large amounts of anthropogenic CO2 emissions.

Roughly 50–70 percent is removed by the oceans, whereasthe remainder is taken up by the terrestrial biosphere.Global warming, however, could decrease the capacity ofthese reservoirs to sequester atmospheric CO2. Reductions

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in the rate of carbon uptake by these reservoirs wouldincrease the pace of CO2 buildup in the atmosphere andrepresent yet another possible positive feedback toincreased greenhouse gas concentrations.

In the world’s oceans, this feedback effect might takeseveral paths. First, as surface waters warmed, they wouldhold less dissolved CO2. Second, if more CO2 was addedto the atmosphere and taken up by the oceans, bicarbon-ate ions (HCO3

– ) would multiply and ocean acidity wouldincrease. Since calcium carbonate (CaCO3 ) is broken

down by acidic solutions, rising acidity would threatenocean-dwelling fauna that incorporate CaCO3 into theirskeletons or shells. As it became increasingly difficult forthese organisms to absorb oceanic carbon, there wouldbe a corresponding decrease in the efficiency of the bio-logical pump that helps to maintain the oceans as a carbonsink (as described in the section Carbon Dioxide). Third,

rising surface temperatures might lead to a slowdown inthe so-called thermohaline circulation, a global patternof oceanic flow that partly drives the sinking of surface

waters near the poles and is responsible for much of theburial of carbon in the deep ocean. A slowdown in thisflow due to an influx of melting fresh water into what arenormally saltwater conditions might also cause the solu-

bility pump, which transfers CO2 from shallow to deeper waters, to become less efficient. Indeed, it is predictedthat if global warming continued to a certain point, theoceans would cease to be a net sink of CO2 and wouldbecome a net source.

As large sections of tropical forest are lost because ofthe warming and drying of regions such as Amazonia, the

overall capacity of plants to sequester atmospheric CO2 would be reduced. As a result, the terrestrial biosphere,though currently a carbon sink, would become a car-bon source. Ambient temperature is a significant factor



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affecting the pace of photosynthesis in plants, and manyplant species that are well-adapted to their local climaticconditions have maximized their photosynthetic rates. Astemperatures increase and conditions begin to exceed theoptimal temperature range for both photosynthesis andsoil respiration, the rate of photosynthesis would decline.As dead plants decompose, microbial metabolic activity (aCO2 source) would increase and would eventually outpacephotosynthesis.

Under sufficient global warming conditions, methane

sinks in the oceans and terrestrial biosphere also mightbecome methane sources. Annual emissions of methaneby wetlands might either increase or decrease, dependingon temperatures and input of nutrients, and it is possiblethat wetlands could switch from source to sink. There isalso the potential for increased methane release as a resultof the warming of Arctic permafrost (on land) and further

methane release at the continental margins of the oceans(a few hundred metres below sea level). The current aver-age atmospheric methane concentration of 1,750 ppb isequivalent to 3.5 gigatons (3.5 billion tons) of carbon. Thereare at least 400 gigatons of carbon equivalent stored inArctic permafrost and as much as 10,000 gigatons (10 tril-lion tons) of carbon equivalent trapped on the continental

margins of the oceans in a hydrated crystalline formknown as clathrate. It is believed that some fraction ofthis trapped methane could become unstable with addi-tional warming, although the amount and rate of potentialemission remain highly uncertain.

CLIMATE RESEARCH

Modern research into climatic variation and change isbased on a variety of empirical and theoretical lines ofinquiry. One line of inquiry is the analysis of data that

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record changes in atmosphere, oceans, and climate fromroughly 1850 to the present. In a second line of inquiry,information describing paleoclimatic changes is gath-ered from “proxy,” or indirect, sources such as ocean andlake sediments, pollen grains, corals, ice cores, and treerings. Finally, a variety of theoretical models can be usedto investigate the behaviour of Earth’s climate under dif-ferent conditions. These three lines of investigation aredescribed in this section.

Modern Observations

Although a limited regional subset of land-based recordsis available from the 17th and 18th centuries, instrumentalmeasurements of key climate variables have been collectedsystematically and at global scales since the mid-19th toearly 20th centuries. These data include measurements of

surface temperature on land and at sea, atmospheric pres-sure at sea level, precipitation over continents and oceans,sea-ice extents, surface winds, humidity, and tides. Suchrecords are the most reliable of all available climate data,since they are precisely dated and are based on well-under-stood instruments and physical principles. Correctionsmust be made for uncertainties in the data (for instance,

gaps in the observational record, particularly during ear-lier years) and for systematic errors (such as an “urban heatisland” bias in temperature measurements made on land).

Since the mid-20th century a variety of upper-airobservations have become available (for example, oftemperature, humidity, and winds), allowing climaticconditions to be characterized from the ground upward

through the upper troposphere and lower stratosphere.Since the 1970s these data have been supplementedby polar-orbiting and geostationary satellites and by



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The Keeling curve, named after American climate scientist Charles David Keeling, tracks changes in the concentration of carbon dioxide (CO2 ) in Earth’s

atmosphere at a research station on Mauna Loa in Hawaii. Although theseconcentrations experience small seasonal fluctuations, the overall trend showsthat CO2 is increasing in the atmosphere. Encyclopædia Britannica, Inc.

platforms in the oceans that gauge temperature, salin-ity, and other properties of seawater. Attempts have beenmade to fill the gaps in early measurements by using various

statistical techniques and “backward prediction” modelsand by assimilating available observations into numeri-cal weather prediction models. These techniques seekto estimate meteorological observations or atmospheric

variables (such as relative humidity) that have been poorlymeasured in the past.

Modern measurements of greenhouse gas concentra-

tions began with an investigation of atmospheric carbondioxide (CO2 ) concentrations by American climate sci-entist Charles Keeling at the summit of Mauna Loa in

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Hawaii in 1958. Keeling’s findings indicated that CO2 concentrations were steadily rising in association withthe combustion of fossil fuels, and they also yielded thefamous “Keeling curve,” a graph in which the longer-termrising trend is superimposed on small oscillations relatedto seasonal variations in the uptake and release of CO2 from photosynthesis and respiration in the terrestrialbiosphere. Keeling’s measurements at Mauna Loa applyprimarily to the Northern Hemisphere.

Taking into account the uncertainties, the instru-

mental climate record indicates substantial trends overthe past century consistent with a warming Earth. Thesetrends include a rise in global surface temperature of 0.6°C (1.1 °F), an associated elevation of global sea level of17 cm (6.7 inches), and a decrease in snow cover in theNorthern Hemisphere of approximately 1.5 million squarekm (580,000 square miles). Increases in global sea level are

attributed to a combination of seawater expansion due toocean heating and freshwater runoff caused by the melt-ing of terrestrial ice. Reductions in snow cover are theresult of warmer temperatures favouring a steadily shrink-ing winter season.

Prehistorical Climate Records

In order to reconstruct climate changes that occurredprior to about the mid-19th century, it is necessary to use“proxy” measurements—that is, records of other naturalphenomena that indirectly measure various climate con-ditions. Some proxies, such as most sediment cores andpollen records, glacial moraine evidence, and geothermal

borehole temperature profiles, are coarsely resolved ordated and thus are only useful for describing climatechanges on long timescales. Other proxies, such as growthrings from trees or oxygen isotopes from corals and ice



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cores, can provide a record of yearly or even seasonal cli-mate changes.

The data from these proxies should be calibrated toknown physical principles or related statistically to therecords collected by modern instruments, such as satel-lites. Networks of proxy data can then be used to inferpatterns of change in climate variables, such as the behav-iour of surface temperature over time and geography.Yearly reconstructions of climate variables are possibleover the past 1,000 to 2,000 years using annually dated

proxy records, but reconstructions farther back in timeare generally based on more coarsely resolved evidencesuch as ocean sediments and pollen records. For these,records of conditions can be reconstructed only on times-cales of hundreds or thousands of years. In addition, sincerelatively few long-term proxy records are available for theSouthern Hemisphere, most reconstructions focus on

the Northern Hemisphere.The various proxy-based reconstructions of the aver-

age surface temperature of the Northern Hemispherediffer in their details. These differences are the result ofuncertainties implicit in the proxy data themselves andalso of differences in the statistical methods used to relatethe proxy data to surface temperature. Nevertheless, all

studies as reviewed in the Fourth Assessment Report ofthe Intergovernmental Panel on Climate Change (IPCC)indicate that the average surface temperature sinceabout 1950 is higher than at any time during the previous1,000 years.

Theoretical Climate Models

To understand and explain the complex behaviour ofEarth’s climate, modern climate models incorporateseveral variables that stand in for materials passing

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To understand and explain the complex behaviour of Earth’s climate, modernclimate models incorporate several variables that stand in for materials pass-ing through Earth’s atmosphere and oceans and the forces that affect them.Encyclopaedia Britannica, Inc.

through Earth’s atmosphere and oceans and the forcesthat affect them.

Theoretical models of Earth’s climate system can beused to investigate the response of climate to externalradiative forcing as well as its own internal variability. Twoor more models that focus on different physical processes



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may be coupled or linked together through a common fea-ture, such as geographic location. Climate models varyconsiderably in their degree of complexity. The simplestmodels of energy balance describe Earth’s surface as a

globally uniform layer whose temperature is determinedby a balance of incoming and outgoing shortwave andlongwave radiation. These simple models may also con-sider the effects of greenhouse gases. At the other end ofthe spectrum are fully coupled, three-dimensional, globalclimate models. These are complex models that solve for

radiative balance; for laws of motion governing the atmo-sphere, ocean, and ice; and for exchanges of energy andmomentum within and between the different compo-nents of the climate. In some cases, theoretical climatemodels also include an interactive representation ofEarth’s biosphere and carbon cycle.

Even the most-detailed climate models cannot resolve

all the processes that are important in the atmosphereand ocean. Most climate models are designed to gauge thebehaviour of a number of physical variables over space andtime, and they often artificially divide Earth’s surface intoa grid of many equal-sized “cells.” Each cell may neatlycorrespond to some physical process (such as summernear-surface air temperature) or other variable (such as

land-use type), and it may be assigned a relatively straight-forward value. So-called “sub-grid-scale” processes, suchas those of clouds, are too small to be captured by the rela-tively coarse spacing of the individual grid cells. Instead,such processes must be represented through a statisticalprocess that relates the properties of the atmosphere andocean. For example, the average fraction of cloud cover

over a hypothetical “grid box” (that is, a representative volume of air or water in the model) can be estimated fromthe average relative humidity and the vertical tempera-ture profile of the grid cell. Variations in the behaviour

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of different coupled climate models arise in large partfrom differences in the ways sub-grid-scale processes aremathematically expressed.

Despite these required simplifications, many theoreti-cal climate models perform remarkably well whenreproducing basic features of the atmosphere, such as thebehaviour of midlatitude jet streams or Hadley cell circu-lation. The models also adequately reproduce importantfeatures of the oceans, such as the Gulf Stream. In addi-tion, models are becoming better able to reproduce the

main patterns of internal climate variability, such as thoseof El Niño/Southern Oscillation (ENSO). Consequently,periodically recurring events—such as ENSO and otherinteractions between the atmosphere and ocean cur-rents—are being modeled with growing confidence.

Climate models have been tested in their ability toreproduce observed changes in response to radiative

forcing. In 1987 a team at NASA’s Goddard Institutefor Space Studies in New York City used a fairly primi-tive climate model to predict warming patterns thatmight occur in response to three different scenarios ofanthropogenic radiative forcing. Warming patterns wereforecasted for subsequent decades. Of the three scenar-ios, one very closely followed the actual near-surface air

temperature pattern that occurred through the 1990sand into the following decade, and it predicted the tem-perature rise of roughly 0.5 °C (0.9 °F) that actually tookplace during that interval. The NASA team also used aclimate model to successfully predict that global meansurface temperatures would cool by about 0.5 °C for oneto two years after the 1991 eruption of Mount Pinatubo

in the Philippines.More recently, so-called “detection and attribution”studies have been performed. These studies compare



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Graph of the predicted increase in the concentration of carbon dioxide (CO2)in Earth’s atmosphere according to a series of climate change scenarios that

assume different levels of economic development, population growth, and fos-

sil-fuel use. Encyclopaedia Britannica, Inc.

predicted changes in near-surface air temperature andother climate variables with patterns of change thathave been observed for the past one to two centuries.The simulations have shown that the observed pat-terns of warming of Earth’s surface and upper oceans,

as well as changes in other climate phenomena such asprevailing winds and precipitation patterns, are con-sistent with the effects of an anthropogenic influencepredicted by the climate models. In addition, climatemodel simulations have shown success in reproduc-ing the magnitude and the spatial pattern of cooling inthe Northern Hemisphere between roughly 1400 and

1850—during the Little Ice Age, which appears to haveresulted from a combination of lowered solar outputand heightened explosive volcanic activity.

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PROJECTED RANGE OF SEA-LEVEL RISE BYCLIMATE CHANGE SCENARIO*

scenario temperaturechange (°C) in

2090–99 relativeto 1980–99

sea-level rise (m) in2090–99 relative to

1980–99

B1 1.1–2.9 0.18–0.38

A1T 1.4–3.8 0.20–0.45

B2 1.4–3.8 0.20–0.43

A1B 1.7–4.4 0.21–0.48A2 2.0–5.4 0.23–0.51

A1Fl 2.4–6.4 0.26–0.59

*Ranges of sea-level rise are based on various models of climate changethat exclude the possibility of future rapid changes in ice flow, such asthe melting of the Greenland and Antarctic ice caps. Source:Intergovernmental Panel on Climate Change Fourth Assessment Report

POTENTIAL EFFECTS OF GLOBALWARMING

The path of future climate change will depend on whatcourses of action are taken by society—in particular the

emission of greenhouse gases from the burning of fos-sil fuels. A range of alternative emissions scenarios havebeen proposed by the IPCC to predict future climatechanges. These scenarios depend on various assumptionsconcerning future rates of human population growth,economic development, energy demand, technologi-cal advancement, and other factors. Scenarios with the

smallest increases in greenhouse gases are associated withlow-to-moderate economic growth and increasing useof more environmentally friendly and resource-efficienttechnologies or with rapid economic growth and use of



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alternative energy sources. In each of these scenarios,CO2 concentrations at the year 2100 are held near orbelow twice the levels of the preindustrial age—that is,at or below 600 ppm. At the other extreme are scenariosthat assume either more intensive use of fossil fuels orunabated growth in world population. In these scenarios

CO2 concentrations approach or exceed 900 ppm by 2100.The intermediate scenarios reflect so-called “business-as-usual” (BAU) activities, where civilization continues toburn fossil fuels at early 21st-century rates. In the BAU

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Graph of the predicted increase in Earth’s average surface temperature according to a series of climate change scenarios that assume different levelsof economic development, population growth, and fossil-fuel use. The assump-tions made by each scenario are given at the bottom of the graph.



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scenarios CO2 concentrations approach levels of 700–750ppm by 2100.

Simulations of Future Climate Change

The differences between the various simulations arisefrom disparities between the various climate models usedand from assumptions made by each emission scenario.For example, best estimates of the predicted increases in

global surface temperature between the years 2000 and

2100 range from about 1.8 °C (3.2 °F) to about 4 °C (7.2 °F),depending on which emission scenario is assumed and

which climate model is used. These projections are con-servative in that they do not take into account potentialpositive carbon cycle feedbacks discussed above. Onlythe lower-end emissions scenarios are likely to hold addi-tional global surface warming by 2100 to less than 2 °C

(3.6 °F)—a level considered by many scientists to be thethreshold above which pervasive and extreme climaticeffects will occur.

Patterns of Warming

The greatest increase in near-surface air temperature isprojected to occur over the polar region of the Northern

Hemisphere because of the melting of sea ice and theassociated reduction in surface albedo. Greater warmingis predicted over land areas than over the ocean. Largelydue to the delayed warming of the oceans and their greaterspecific heat, the Northern Hemisphere—with less than40 percent of its surface area covered by water—isexpected to warm faster than the Southern Hemisphere.

Some of the regional variation in predicted warming isexpected to arise from changes to wind patterns and oceancurrents in response to surface warming. For example, the

warming of the region of the North Atlantic Ocean just



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south of Greenland is expected to be slight. This anomalyis projected to arise from a weakening of warm northwardocean currents combined with a shift in the jet streamthat will bring colder polar air masses to the region.

Precipitation Patterns

The climate changes associated with global warming arealso projected to lead to changes in precipitation patternsacross the globe. Increased precipitation is predicted inthe polar and subpolar regions, whereas decreased precip-

itation is projected for the middle latitudes of bothhemispheres as a result of the expected poleward shift inthe jet streams. Whereas precipitation near the Equator ispredicted to increase, it is thought that rainfall in the sub-tropics will decrease. Both phenomena are associated

with a forecasted strengthening of the tropical Hadley cellpattern of atmospheric circulation.

Changes in precipitation patterns are expected toincrease the chances of both drought and flood conditionsin many areas. Decreased summer precipitation in NorthAmerica, Europe, and Africa, combined with greater ratesof evaporation due to warming surface temperatures, isprojected to lead to decreased soil moisture and droughtin many regions. Furthermore, since anthropogenic cli-

mate change will likely lead to a more vigorous hydrologiccycle with greater rates of both evaporation and precipita-tion, there will be a greater probability for intenseprecipitation and flooding in many regions.

Regional Predictions

Regional predictions of future climate change remain lim-

ited by uncertainties in how the precise patterns ofatmospheric winds and ocean currents will vary withincreased surface warming. For example, some uncer-tainty remains in how the frequency and magnitude of El

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Niño/Southern Oscillation (ENSO) events will adjust toclimate change. Since ENSO is one of the most promi-nent sources of interannual variations in regional patternsof precipitation and temperature, any uncertainty in howit will change implies a corresponding uncertainty in cer-tain regional patterns of climate change. For example,increased El Niño activity would likely lead to more win-ter precipitation in some regions, such as the desertsouthwest of the United States. This might offset thedrought predicted for those regions, but at the same time

it might lead to less precipitation in other regions. Rising winter precipitation in the desert southwest of the UnitedStates might exacerbate drought conditions in locationsas far away as South Africa.

Ice Melt and Sea-Level Rise

A warming climate holds important implications for other

aspects of the global environment. Because of the slowprocess of heat diffusion in water, the world’s oceans arelikely to continue to warm for several centuries in responseto increases in greenhouse concentrations that havetaken place so far. The combination of seawater’s thermalexpansion associated with this warming and the meltingof mountain glaciers is predicted to lead to an increase in

global sea level of 0.21–0.48 metre (0.7–1.6 feet) by 2100under the BAU emissions scenario. However, the actualrise in sea level could be considerably greater than this. Itis probable that the continued warming of Greenland willcause its ice sheet to melt at accelerated rates. In addition,this level of surface warming may also melt the ice sheetof West Antarctica. Paleoclimatic evidence suggests that

an additional 2 °C (3.6 °F) of warming could lead to theultimate destruction of the Greenland Ice Sheet, an eventthat would add another 5 to 6 metres (16 to 20 feet) topredicted sea-level rise. Such an increase would submerge



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a substantial number of islands and lowland regions.Selected lowland regions include substantial parts of theU.S. Gulf Coast and Eastern Seaboard (including roughlythe lower third of Florida), much of the Netherlandsand Belgium (two of the European Low Countries), andheavily populated tropical areas such as Bangladesh. In

addition, many of the world’s major cities—such as Tokyo,New York, Mumbai (Bombay), Shanghai, and Dhaka—arelocated in lowland regions vulnerable to rising sea levels.With the loss of the West Antarctic ice sheet, additional

NASA image showing locations on Antarctica where temperatures hadincreased between 1959 and 2009. Red represents areas where temperatures

had increased the most over the period, particularly in West Antarctica, while dark blue represents areas with a lesser degree of warming. Temperaturechanges are measured in degrees Celsius. GSFC Scientific VisualizationStudio/NASA

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Thermohaline circulation transports and mixes the water of the oceans. Inthe process it transports heat, which influences regional climate patterns. The

density of seawater is determined by the temperature and salinity of a volumeof seawater at a particular location. The difference in density between one location and another drives the thermohaline circulation. EncyclopaediaBritannica, Inc.

sea-level rise would approach 10.5 metres (34 feet). Whilethe current generation of models predicts that such globalsea-level changes might take several centuries to occur,

it is possible that the rate could accelerate as a result ofprocesses that tend to hasten the collapse of ice sheets.One such process is the development of moulins, or large,

vertical shafts in the ice that allow surface meltwater topenetrate to the base of the ice sheet. A second processinvolves the vast ice shelves off Antarctica that buttressthe grounded continental ice sheet of Antarctica’s inte-

rior. If these ice shelves collapse, the continental ice sheetcould become unstable, slide rapidly toward the ocean,and melt, thereby further increasing mean sea level. Thus



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far, neither process has been incorporated into the theo-retical models used to predict sea-level rise.

Ocean Circulation Changes

Another possible consequence of global warming is adecrease in the global ocean circulation system known asthe “thermohaline circulation” or “great ocean conveyorbelt.” This system involves the sinking of cold saline watersin the subpolar regions of the oceans, an action that helpsto drive warmer surface waters poleward from the sub-

tropics. As a result of this process, a warming influence iscarried to Iceland and the coastal regions of Europe thatmoderates the climate in those regions. Some scientistsbelieve that global warming could shut down this oceancurrent system by creating an influx of fresh water frommelting ice sheets and glaciers into the subpolar NorthAtlantic Ocean. Since fresh water is less dense than saline

water, a significant intrusion of fresh water would lowerthe density of the surface waters and thus inhibit the sink-ing motion that drives the large-scale thermohalinecirculation. It has also been speculated that, as a conse-quence of large-scale surface warming, such changes couldeven trigger colder conditions in regions surrounding theNorth Atlantic. Experiments with modern climate mod-

els suggest that such an event would be unlikely. Instead, amoderate weakening of the thermohaline circulationmight occur that would lead to a dampening of surface

warming—rather than actual cooling—in the higher lati-tudes of the North Atlantic Ocean.

Tropical Cyclones

One of the more controversial topics in the scienceof climate change involves the impact of global warm-ing on tropical cyclone activity. It appears likely that

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rising tropical ocean temperatures associated with global warming will lead to an increase in the intensity (and theassociated destructive potential) of tropical cyclones.In the Atlantic a close relationship has been observedbetween rising ocean temperatures and a rise in thestrength of hurricanes. Trends in the intensities of tropi-cal cyclones in other regions, such as in the tropical Pacificand Indian oceans, are more uncertain due to a paucity ofreliable long-term measurements.

While the warming of oceans favours increased tropi-

cal cyclone intensities, it is unclear to what extent risingtemperatures affect the number of tropical cyclones thatoccur each year. Other factors, such as wind shear, couldplay a role. If climate change increases the amount of windshear—a factor that discourages the formation of tropicalcyclones—in regions where such storms tend to form, itmight partially mitigate the impact of warmer tempera-

tures. On the other hand, changes in atmospheric windsare themselves uncertain—because of, for example, uncer-tainties in how climate change will affect ENSO.

Environmental Consequences ofGlobal Warming

Global warming and climate change have the potential toalter biological systems. More specifically, changes to near-surface air temperatures will likely influence ecosystemfunctioning and thus the biodiversity of plants, animals,and other forms of life. The current geographic ranges ofplant and animal species have been established by adapta-tion to long-term seasonal climate patterns. As global

warming alters these patterns on timescales considerablyshorter than those that arose in the past from natural cli-mate variability, relatively sudden climatic changes maychallenge the natural adaptive capacity of many species.



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It has been estimated that one-fifth to one-third of allplant and animal species are likely to be at an increasedrisk of extinction if global average surface temperaturesrise another 1.5 to 2.5 °C (2.7 to 4.5 °F) by the year 2100.This temperature range falls within the scope of the loweremissions scenarios projected by the IPCC. Species-lossestimates climb to as much as 40 percent for a warming inexcess of 4.5 °C (8.1 °F)—a level that could be reached inthe IPCC’s higher emissions scenarios. A 40 percentextinction rate would likely lead to major changes in the

food webs within ecosystems and have a destructiveimpact on ecosystem function.

Surface warming in temperate regions is likely to leadchanges in various seasonal processes—for instance, ear-lier leaf production by trees, earlier greening of vegetation,altered timing of egg-laying and hatching, and shifts in theseasonal migration patterns of birds, fishes, and other

migratory animals. In high-latitude ecosystems, changesin the seasonal patterns of sea ice threaten predators suchas polar bears and walruses. (Both species rely on brokensea ice for their hunting activities.) Also in the high lati-tudes, a combination of warming waters, decreased seaice, and changes in ocean salinity and circulation is likelyto lead to reductions or redistributions in populations of

algae and plankton. As a result, fish and other organismsthat forage upon algae and plankton may be threatened.On land, rising temperatures and changes in precipitationpatterns and drought frequencies are likely to alter pat-terns of disturbance by fires and pests.

Other likely impacts on the environment include thedestruction of many coastal wetlands, salt marshes, and

mangrove swamps as a result of rising sea levels and theloss of certain rare and fragile habitats that are often hometo specialist species that are unable to thrive in other envi-ronments. For example, certain amphibians limited to

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Cross section of a generalized coral polyp. Encyclopædia Britannica, Inc.

isolated tropical cloud forests either have become extinctalready or are under serious threat of extinction. Cloudforests—tropical forests that depend on persistent conden-sation of moisture in the air—are disappearing as optimal



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condensation levels move to higher elevations in responseto warming temperatures in the lower atmosphere.

In many cases a combination of stresses caused byclimate change as well as human activity represents aconsiderably greater threat than either climatic stressesor nonclimatic stresses alone. A particularly importantexample is coral reefs, which contain much of the ocean’sbiodiversity. Rising ocean temperatures increase thetendency for coral bleaching (a condition where zooxan-thellae, or yellow-green algae, living in symbiosis with

coral either lose their pigments or abandon the coral pol- yps altogether), and they also raise the likelihood of greaterphysical damage by progressively more destructive tropi-cal cyclones. In many areas coral is also under stress fromincreased ocean acidification, marine pollution, runofffrom agricultural fertilizer, and physical damage by boatanchors and dredging.

Another example of how climate and nonclimaticstresses combine is illustrated by the threat to migratoryanimals. As these animals attempt to relocate to regions

with more favourable climate conditions, they are likely toencounter impediments such as highways, walls, artificial

waterways, and other man-made structures.Warmer temperatures are also likely to affect the

spread of infectious diseases, since the geographicranges of carriers, such as insects and rodents, are oftenlimited by climatic conditions. Warmer winter condi-tions in New York in 1999, for example, appear to havefacilitated an outbreak of West Nile virus, whereas thelack of killing frosts in New Orleans during the early1990s led to an explosion of disease-carrying mosqui-

toes and cockroaches. Warmer winters in the Koreanpeninsula and southern Europe have allowed the spreadof the Anopheles mosquito, which carries the malariaparasite, whereas warmer conditions in Scandinavia in

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recent years have allowed for the northward advance ofencephalitis.

In the southwestern United States, alternationsbetween drought and flooding related in part to the ENSOphenomenon have created conditions favourable for thespread of hantaviruses by rodents. The spread of mos-quito-borne Rift Valley fever in equatorial East Africa hasalso been related to wet conditions in the region associ-ated with ENSO. Severe weather conditions conducive torodents or insects have been implicated in infectious dis-

ease outbreaks—for instance, the outbreaks of choleraand leptospirosis that occurred after Hurricane Mitchstruck Central America in 1998. Global warming couldtherefore affect the spread of infectious disease throughits influence on ENSO or on severe weather conditions.

Socioeconomic Consequences of

Global Warming

Socioeconomic impacts of global warming could be sub-stantial depending on the actual temperature increasesover the next century. Models predict that a net global

warming of 1 to 3 °C (1.8 to 5.4 °F) beyond the late-20th-century global average would produce economic losses in

some regions (particularly the tropics and high latitudes)and economic benefits in others. For warming beyondthese levels, benefits would tend to decline and costsincrease. For warming in excess of 4 °C (7.2 °F), modelspredict that costs will exceed benefits on average, with

global mean economic losses estimated between 1 and 5percent of gross domestic product. Substantial disrup-

tions could be expected under these conditions, specificallyin the areas of agriculture, food and forest products, waterand energy supply, and human health.



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Agricultural productivity might increase modestly intemperate regions for some crops in response to a local

warming of 1–3 °C (1.8–5.4 °F), but productivity will gener-ally decrease with further warming. For tropical andsubtropical regions, models predict decreases in crop pro-ductivity for even small increases in local warming. Insome cases, adaptations such as altered planting practicesare projected to ameliorate losses in productivity for mod-est amounts of warming. An increased incidence ofdrought and flood events would likely lead to further

decreases in agricultural productivity and to decreases inlivestock production, particularly among subsistencefarmers in tropical regions. In regions such as the AfricanSahel, decreases in agricultural productivity have alreadybeen observed as a result of shortened growing seasons,

which in turn have occurred as a result of warmer and drierclimatic conditions. In other regions, changes in agricul-

tural practice, such as planting crops earlier in the growingseason, have been undertaken. The warming of oceans ispredicted to have an adverse impact on commercial fisher-ies by changing the distribution and productivity of

various fish species, whereas commercial timber produc-tivity may increase globally with modest warming.

Water resources are likely to be affected substantially

by global warming. At current rates of warming, a 10–40percent increase in average surface runoff and water avail-ability has been projected in higher latitudes and in certain

wet regions in the tropics by the middle of the 21st cen-tury, while decreases of similar magnitude are expected inother parts of the tropics and in the dry regions in the sub-tropics. This would be particularly severe during the

summer season. In many cases water availability is alreadydecreasing or expected to decrease in regions that havebeen stressed for water resources since the turn of the 21st

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century. Such regions as the African Sahel, western NorthAmerica, southern Africa, the Middle East, and westernAustralia continue to be particularly vulnerable. In theseregions drought is projected to increase in both magni-tude and extent, which would bring about adverse effectson agriculture and livestock raising. Earlier and increasedspring runoff is already being observed in western NorthAmerica and other temperate regions served by glacial orsnow-fed streams and rivers. Fresh water currently storedby mountain glaciers and snow in both the tropics and

extratropics is also projected to decline and thus reducethe availability of fresh water for more than 15 percentof the world’s population. It is also likely that warmingtemperatures, through their impact on biological activityin lakes and rivers, may have an adverse impact on waterquality, further diminishing access to safe water sourcesfor drinking or farming. For example, warmer waters

favour an increased frequency of nuisance algal blooms, which can pose health risks to humans. Risk-managementprocedures have already been taken by some countries inresponse to expected changes in water availability.

Energy availability and use could be affected in at leasttwo distinct ways by rising surface temperatures. In gen-eral, warmer conditions would favour an increased demand

for air-conditioning; however, this would be at least par-tially offset by decreased demand for winter heating intemperate regions. Energy generation that requires watereither directly, as in hydroelectric power, or indirectly, asin steam turbines used in coal-fired power plants or incooling towers used in nuclear power plants, may becomemore difficult in regions with reduced water supplies.

As discussed above, it is expected that human health will be further stressed under global warming condi-tions by potential increases in the spread of infectiousdiseases. Declines in overall human health might occur



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with increases in the levels of malnutrition due to disrup-tions in food production and by increases in the incidenceof afflictions. Such afflictions could include diarrhea,cardiorespiratory illness, and allergic reactions in the mid-latitudes of the Northern Hemisphere as a result of risinglevels of pollen. Rising heat-related mortality, such as thatobserved in response to the 2003 European heat wave,might occur in many regions, especially in impoverishedareas where air-conditioning is not generally available.

The economic infrastructure of most countries is

predicted to be severely strained by global warming andclimate change. Poor countries and communities withlimited adaptive capacities are likely to be dispropor-tionately affected. Projected increases in the incidenceof severe weather, heavy flooding, and wildfires associ-ated with reduced summer ground moisture in manyregions will threaten homes, dams, transportation net-

works and other facets of human infrastructure. Inhigh-latitude and mountain regions, melting permafrostis likely to lead to ground instability or rock avalanches,further threatening structures in those regions. Risingsea levels and the increased potential for severe tropicalcyclones represent a heightened threat to coastal com-munities throughout the world. It has been estimated

that an additional warming of 1–3 °C (1.8–5.4 °F) beyondthe late-20th-century global average would threatenmillions more people with the risk of annual flooding.People in the densely populated, poor, low-lying regionsof Africa, Asia, and tropical islands would be the most

vulnerable, given their limited adaptive capacity. In addi-tion, certain regions in developed countries, such as the

Low Countries of Europe and the Eastern Seaboard andGulf Coast of the United States, would also be vulner-able to the effects of rising sea levels. Adaptive stepsare already being taken by some governments to reduce

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the threat of increased coastal vulnerability through theconstruction of dams and drainage works.

GLOBAL WARMING ANDPUBLIC POLICY

Since the 19th century, many researchers working acrossa wide range of academic disciplines have contributed toan enhanced understanding of the atmosphere and the

global climate system. Concern among prominent climate

scientists about global warming and human-induced (or“anthropogenic”) climate change arose in the mid-20thcentury, but most scientific and political debate over theissue did not begin until the 1980s. Today, leading climatescientists agree that many of the ongoing changes to the

global climate system are largely caused by the releaseinto the atmosphere of greenhouse gases, or gases that

enhance Earth’s natural greenhouse effect. Most green-house gases are released by the burning of fossil fuels forheating, cooking, electrical generation, transportation,and manufacturing, but they are also released as a result of



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The Kyoto Protocol to the United Nations Framework Conventionon Climate Change is an international treaty, named for the Japanese

city in which it was adopted in December 1997, that aimed to reducethe emission of gases that contribute to global warming. In force since2005, the protocol called for reducing the emission of six greenhouse gases in 36 countries to 5.2 percent below 1990 levels in the “commit-ment period” 2008–12. It was widely hailed as the most significantenvironmental treaty ever negotiated, though some critics questionedits effectiveness.

Background and provisionsThe Kyoto Protocol was adopted as the first addition to the UnitedNations Framework Convention on Climate Change (UNFCCC), aninternational treaty that committed its signatories to develop nationalprograms to reduce their emissions of greenhouse gases. Such gases,including carbon dioxide and methane, affect the energy balance of the global atmosphere in ways expected to lead to an overall increase in global average temperature, known as global warming. According to theIntergovernmental Panel on Climate Change, established by the UnitedNations Environment Programme and the World MeteorologicalOrganization in 1988, the long-term effects of global warming wouldinclude a general rise in sea level around the world, resulting in theinundation of low-lying coastal areas and the possible disappearanceof some island states; the melting of glaciers, sea ice, and Arctic per-mafrost; an increase in the number of extreme climate-related events,such as floods and droughts, and changes in their distribution; and an

increased risk of extinction for 20 to 30 percent of all plant and animalspecies. The Kyoto Protocol committed most of the Annex I signato-ries to the UNFCCC to mandatory emission-reduction targets, which varied depending on the unique circumstances of each country. Othersignatories to the UNFCCC and the protocol, consisting mostly ofdeveloping countries, were not required to restrict their emissions.The protocol entered into force in February 2005, 90 days after beingratified by at least 55 Annex I signatories that together accounted for at

least 55 percent of total carbon-dioxide emissions in 1990.The protocol provided several means for countries to reach theirtargets. One approach was to make use of natural processes, called“sinks,” that remove greenhouse gases from the atmosphere. The




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planting of trees, which take up carbon dioxide from the air, would bean example. Another approach was the international program calledthe Clean Development Mechanism (CDM), which encouraged devel-

oped countries to invest in technology and infrastructure inless-developed countries, where there were often significant opportu-nities to reduce emissions. Under the CDM, the investing countrycould claim the effective reduction in emissions as a credit towardmeeting its obligations under the protocol. An example would be aninvestment in a clean-burning natural gas power plant to replace a pro-posed coal-fired plant. A third approach was emissions trading, whichallowed participating countries to buy and sell emissions rights and

thereby placed an economic value on greenhouse-gas emissions.European countries initiated an emissions-trading market as a mecha-nism to work toward meeting their commitments under the KyotoProtocol. Countries that failed to meet their emissions targets wouldbe required to make up the difference between their targeted andactual emissions, plus a penalty amount of 30 percent, in the subse-quent commitment period, beginning in 2012; they would also beprevented from engaging in emissions trading until they were judgedto be in compliance with the protocol. The emission targets for com-mitment periods after 2012 were to be established in future protocols.

Challenges

Although the Kyoto Protocol represented a landmark diplomaticaccomplishment, its success was far from assured. Indeed, reportsissued in the first two years after the treaty took effect indicated thatmost participants would fail to meet their emission targets. Even if

the targets were met, however, the ultimate benefit to the environ-ment would not be significant, according to some critics, since theUnited States, the leading emitter of greenhouse gases (approximately25 percent of the total), was not party to the protocol, and China, thesecond leading emitter, was not required to restrict its emissionsbecause of its status as a developing country. Other critics claimedthat the emission reductions called for in the protocol were too mod-est to make a detectable difference in global temperatures in the

subsequent several decades, even if fully achieved with U.S. participa-tion. Meanwhile, some developing countries argued that improvingadaptation to climate variability and change was just as important asreducing greenhouse-gas emissions.



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the natural decomposition of organic materials, wildfires,deforestation, and land-clearing activities. Opponents ofthis view have often stressed the role of natural factors inpast climatic variation and have accentuated the scientificuncertainties associated with data on global warming andclimate change. Nevertheless, a growing body of scientistshas called upon governments, industries, and citizens toreduce their emissions of greenhouse gases.

All countries emit greenhouse gases, but highly indus-trialized countries and more populous countries emit

significantly greater quantities than others. Countries inNorth America and Europe that were the first to undergothe process of industrialization have been responsible forreleasing most greenhouse gases in absolute cumulativeterms since the beginning of the Industrial Revolutionin the mid-18th century. Today these countries are beingjoined by large developing countries such as China and

India, where rapid industrialization is being accom-panied by a growing release of greenhouse gases. TheUnited States, possessing approximately 5 percent of the

global population, emitted almost 21 percent of global greenhouse gases in 2000. The same year, the then 25member states of the European Union (EU)—possessinga combined population of 450 million people—emitted 14

percent of all anthropogenic greenhouse gases. This fig-ure was roughly the same as the fraction released by the1.2 billion people of China. In 2000 the carbon footprint(that is, the amount of carbon dioxide emissions associ-ated with all the activities of a person or other entity) ofan average American was 24.5 tons of greenhouse gases,the carbon footprint of the average person living in the

EU was 10.5 tons, and the carbon footprint of the aver-age person living in China was only 3.9 tons. AlthoughChina’s per capita greenhouse gas emissions remained

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significantly lower than those of the EU and the UnitedStates, it was the largest greenhouse gas emitter in 2006in absolute terms.

The IPCC and the Scientific Consensus

An important first step in formulating public policy on global warming and climate change is the gathering ofrelevant scientific and socioeconomic data. In 1998 theIntergovernmental Panel on Climate Change (IPCC) was

established by the World Meteorological Organizationand the United Nations Environment Programme. TheIPCC is mandated to assess and summarize the latestscientific, technical, and socioeconomic data on climatechange and to publish its findings in reports presentedto international organizations and national governmentsall over the world. Many thousands of the world’s leading

scientists and experts in the areas of global warming andclimate change have worked under the IPCC to producemajor sets of assessments in 1990, 1995, 2001, and 2007.These reports have evaluated the scientific basis of global

warming and climate change, the major issues relating tothe reduction of greenhouse gas emissions, and the pro-cess of adjusting to a changing climate.

The first IPCC report, published in 1990, stated that a good deal of data showed that human activity affected the variability of the climate system; nevertheless, the authorsof the report could not reach a consensus on the causesand effects of global warming and climate change at thattime. The 1995 IPCC report stated that the balance of evi-dence suggested “a discernible human influence on the

climate.” The 2001 IPCC report confirmed earlier find-ings and presented stronger evidence that most of the warming over the previous 50 years was attributable tohuman activities. The 2001 report also noted that observed



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changes in regional climates were beginning to affectmany physical and biological systems and that there wereindications that social and economic systems were alsobeing affected.

The IPCC’s fourth assessment, issued in 2007, reaf-firmed the main conclusions of earlier reports, but theauthors also stated—in what was regarded as a conserva-tive judgment—that they were at least 90 percent certainthat most of the warming observed over the previous halfcentury had been caused by the release of greenhouse

gases through a multitude of human activities. Both the2001 and 2007 reports stated that during the 20th cen-tury there had been an increase in global average surfacetemperature of 0.6 °C (1.1 °F), within a margin of error of±0.2 °C (0.4 °F). Whereas the 2001 report forecasted anadditional rise in average temperature by 1.4 to 5.8 °C (2.5to 10.4 °F) by 2100, the 2007 report refined this forecast

to an increase of 1.8–4.0 °C (3.2–7.2 °F) by the end of the21st century. These forecasts were based on examinationsof a range of scenarios that characterized future trends in

greenhouse gas emissions.Each IPCC report has helped to build a scientific con-

sensus that elevated concentrations of greenhouse gases inthe atmosphere are the major drivers of rising near-surface

air temperatures and their associated ongoing climaticchanges. In this respect, the current episode of climaticchange, which began about the middle of the 20th century,is seen to be fundamentally different from earlier periodsin that critical adjustments have been caused by activitiesresulting from human behaviour rather than nonanthro-pogenic factors. The IPCC’s 2007 assessment projected

that future climatic changes could be expected to includecontinued warming, modifications to precipitation pat-terns and amounts, elevated sea levels, and “changes in thefrequency and intensity of some extreme events.” Such

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changes would have significant effects on many societiesand on ecological systems around the world.

The UN Framework Convention and theKyoto Protocol

The reports of the IPCC and the scientific consensus theyreflect have provided one of the most prominent bases forthe formulation of climate-change policy. On a globalscale, climate-change policy is guided by two major trea-

ties: the United Nations Framework Convention onClimate Change (UNFCCC) of 1992 and the associated1997 Kyoto Protocol to the UNFCCC (named after thecity in Japan where it was concluded).

The UNFCCC was negotiated between 1991 and 1992.It was adopted at the United Nations Conference onEnvironment and Development in Rio de Janeiro in June

1992 and became legally binding in March 1994. In Article2 the UNFCCC sets the long-term objective of “stabiliza-tion of greenhouse gas concentrations in the atmosphereat a level that would prevent dangerous anthropogenicinterference with the climate system.” Article 3 establishesthat the world’s countries have “common but differenti-ated responsibilities,” meaning that all countries share an

obligation to act—though industrialized countries have aparticular responsibility to take the lead in reducing emis-sions because of their relative contribution to the problemin the past. To this end, the UNFCCC Annex I lists 40specific industrialized countries and countries with econ-omies in transition plus the European Community (EC;formally succeeded by the EU in 2009), and Article 4 states

that these countries should work to reduce their anthro-pogenic emissions to 1990 levels. However, no deadlineis set for this target. Moreover, the UNFCCC does not



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assign any specific reduction commitments to non-AnnexI countries (that is, developing countries).

The follow-up agreement to the UNFCCC, the KyotoProtocol, was negotiated between 1995 and 1997 and wasadopted in December 1997. The Kyoto Protocol regulatessix greenhouse gases released through human activities:carbon dioxide (CO2 ), methane (CH4 ), nitrous oxide(N2O), perfluorocarbons (PFCs), hydrofluorocarbons(HFCs), and sulfur hexafluoride (SF6 ). Under the KyotoProtocol, Annex I countries are required to reduce their

aggregate emissions of greenhouse gases to 5.2 percentbelow their 1990 levels by no later than 2012. Toward this

goal, the protocol sets individual reduction targets foreach Annex I country. These targets require the reduc-tion of greenhouse gases in most countries, but they alsoallow increased emissions from others. For example, theprotocol requires the then 15 member states of the EU

and 11 other European countries to reduce their emis-sions to 8 percent below their 1990 emission levels,

whereas Iceland, a country that produces relatively smallamounts of greenhouse gases, may increase its emissionsas much as 10 percent above its 1990 level. In addition,the Kyoto Protocol requires three countries—NewZealand, Ukraine, and Russia—to freeze their emissions

at 1990 levels.The Kyoto Protocol outlines five requisites by whichAnnex I parties can choose to meet their 2012 emissiontargets. First, it requires the development of national poli-cies and measures that lower domestic greenhouse gasemissions. Second, countries may calculate the benefitsfrom domestic carbon sinks that soak up more carbon than

they emit. Third, countries can participate in schemes thattrade emissions with other Annex I countries. Fourth,signatory countries may create joint implementation

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programs with other Annex I parties and receive credit forsuch projects that lower emissions. Fifth, countries mayreceive credit for lowering the emissions in non-Annex Icountries through a “clean development” mechanism, suchas investing in the building of a new wind power project.

In order to go into effect, the Kyoto Protocol had to beratified by at least 55 countries, including enough Annex Icountries to account for at least 55 percent of that group’stotal greenhouse gas emissions. More than 55 countriesquickly ratified the protocol, including all the Annex I

countries except for Russia, the United States, andAustralia. It was not until Russia, under heavy pressurefrom the EU, ratified the protocol that it became legallybinding in February 2005.

The most-developed regional climate-change policy todate has been formulated by the EU in part to meet itscommitments under the Kyoto Protocol. By 2005 the 15

EU countries that have a collective commitment underthe protocol reduced their greenhouse gas emissions to 2percent below their 1990 levels, though it is not certainthat they will meet their 8 percent reduction target by2012. In 2007 the EU set a collective goal for all 27 mem-ber states to reduce their greenhouse gas emissions by 20percent below 1990 levels by the year 2020. As part of its

effort to achieve this goal, the EU in 2005 established the world’s first multilateral trading scheme for carbon diox-ide emissions, covering more than 11,500 large installationsacross its member states.

In the United States, by contrast, Pres. George W.Bush and a majority of senators rejected the KyotoProtocol, citing the lack of compulsory emission reduc-

tions for developing countries as a particular grievance.At the same time, U.S. federal policy does not set anymandatory restrictions on greenhouse gas emissions, andU.S. emissions increased over 16 percent between 1990



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and 2005. Partly to make up for a lack of direction at thefederal level, many individual U.S. states have formulatedtheir own action plans to address global warming and cli-mate change and have taken a host of legal and politicalinitiatives to curb emissions. These initiatives include:capping emissions from power plants, establishing renew-able portfolio standards requiring electricity providersto obtain a minimum percentage of their power fromrenewable sources, developing vehicle emissions and fuelstandards, and adopting “green building” standards.

Future Climate-Change Policy

Countries differ in opinion on how to proceed with inter-national policy after the commitment period of the KyotoProtocol ends in 2012. The EU supports a continuation ofa legally based collective approach in the form of another

treaty, but other countries, including the United States,support more voluntary measures, as in the Asia-PacificPartnership on Clean Development and Climate that wasannounced in 2005. Long-term goals formulated in Europeand the United States seek to reduce greenhouse gas emis-sions by up to 80 percent by the middle of the 21st century.Related to these efforts, the EU set a goal of limiting tem-

perature rises to a maximum of 2 °C (3.6 °F) abovepreindustrial levels. (Many climate scientists and otherexperts agree that significant economic and ecologicaldamage will result should the global average of near-sur-face air temperatures rise more than 2 °C [3.6 °F] abovepreindustrial temperatures in the next century.)

Despite differences in approach, countries began

negotiations on a new treaty in 2009 based on an agree-ment made at the United Nations Climate ChangeConference in 2007 in Bali, Indon., designed to replacethe Kyoto Protocol after it expires. The 15th Conference

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of the Parties was convened in Copenhagen with the goal of developing a binding global agreement on green-house gas emissions. After two weeks of frenetic andheated negotiation, 193 attending countries recognized anonbinding agreement to keep the rise in global averagetemperatures under 2 °C (3.6 °F). Attendees also agreed toprovide $30 billion in short-term aid to less-developedcountries until 2012.

A growing number of the world’s cities are initiating amultitude of local and subregional efforts to reduce their

emissions of greenhouse gases. Many of these municipali-ties are taking action as members of the InternationalCouncil for Local Environmental Initiatives and its Citiesfor Climate Protection program, which outlines principlesand steps for taking local-level action. In 2005 the U.S.Conference of Mayors adopted the Climate ProtectionAgreement, in which cities committed to reduce emis-

sions to 7 percent below 1990 levels by 2012. In addition,many private firms are developing corporate policies toreduce greenhouse gas emissions. One notable example ofan effort led by the private sector is the creation of theChicago Climate Exchange as a means for reducing emis-sions through a trading process.

As public policies relative to global warming and cli-

mate change continue to develop globally, regionally,nationally, and locally, they fall into two major types. Thefirst type, mitigation policy, focuses on different ways toreduce emissions of greenhouse gases. As most emissionscome from the burning of fossil fuels for energy and trans-portation, much of the mitigation policy focuses onswitching to less carbon-intensive energy sources (such as

wind, solar, and hydropower), improving energy efficiencyfor vehicles, and supporting the development of newtechnology. In contrast, the second type, adaptationpolicy, seeks to improve the ability of various societies to



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face the challenges of a changing climate. For example,some adaptation policies are devised to encourage groupsto change agricultural practices in response to seasonalchanges, whereas other policies are designed to preparecities located in coastal areas for elevated sea levels.

In either case, long-term reductions in greenhouse gasdischarges will require the participation of both industrialcountries and major developing countries. In particular,the release of greenhouse gases from Chinese and Indiansources is rising quickly in parallel with the rapid industri-

alization of those countries. In 2006 China overtook theUnited States as the world’s leading emitter of greenhouse

gases in absolute terms (though not in per capita terms),largely because of China’s increased use of coal and otherfossil fuels. Indeed, all the world’s countries are faced withthe challenge of finding ways to reduce their greenhouse

gas emissions while promoting environmentally and

socially desirable economic development (known as “sus-tainable development” or “smart growth”). Whereas someopponents of those calling for corrective action continueto argue that short-term mitigation costs will be too high,a growing number of economists and policy makers arguethat it will be less costly, and possibly more profitable, forsocieties to take early preventive action than to address

severe climatic changes in the future. Many of the mostharmful effects of a warming climate are likely to takeplace in developing countries. Combating the harmfuleffects of global warming in developing countries will beespecially difficult, as many of these countries are alreadystruggling and possess a limited capacity to meet chal-lenges from a changing climate.

It is expected that each country will be affected differ-ently by the expanding effort to reduce global greenhouse gas emissions. Countries that are relatively large emitters will face greater reduction demands than will smaller

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An Inconvenient Truth is an American documentary film, released in2006, featuring the multimedia presentation of former U.S. vice presi-dent Al Gore that formed the basis for his traveling lecture tour onthe emerging human challenge of global warming and climate change.

From the stage of a small theatre in Los Angeles, Gore juxtaposesthe science behind global warming with elements from his own per-sonal and political life, mixing seriousness with humour to convey hismessage. His presentation uses a collection of graphs, photographs,

and other imagery to describe the greenhouse effect, changes inatmospheric carbon dioxide concentrations throughout history,human energy use and population growth, and how all of these forcescontribute to global warming. In the second half of the film, Goredescribes some of the projected effects of global warming, citing someof the early signs of changing conditions in the Arctic and Antarctic.The film concludes with Gore addressing the common misconcep-tions surrounding global warming and challenging viewers to bring

about needed changes to reduce greenhouse gas emissions. The cred-its provide viewers with suggestions on how they can combat global warming in their own communities.

An Inconvenient Truth

Al Gore in An Inconvenient Truth (2006). © Eric Lee/ParamountClassics, a division of Paramount Pictures; all rights reserved



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Movie poster from An Inconvenient Truth (2006), directed by Davis

Guggenheim and starring Al Gore. Paramount Classics and Participant Productions. © Paramount Classics, a division of Paramount Pictures;all rights reserved

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emitters. Similarly, countries experiencing rapid economic growth are expected to face growing demands to controltheir greenhouse gas emissions as they consume increas-ing amounts of energy. Differences will also occur across

industrial sectors and even between individual companies.For example, producers of oil, coal, and natural gas—

which in some cases represent significant portions ofnational export revenues—may see reduced demand orfalling prices for their goods as their clients decrease theiruse of fossil fuels. In contrast, many producers of new,more climate-friendly technologies and products (such as

generators of renewable energy) are likely to see increasesin demand.To address global warming and climate change, societ-

ies must find ways to fundamentally change their patternsof energy use in favour of less carbon-intensive energy

generation, transportation, and forest and land use man-agement. A growing number of countries have taken on

this challenge, and there are many things individuals toocan do. For instance, consumers have more options topurchase electricity generated from renewable sources.

Gore’s descriptions of climatic processes and warming extrapola-tions in the film have been criticized, particularly by skeptics of global

warming. Many climate scientists agree that some of the statementsmade in the film exaggerate projections or gloss over the nuancesassociated with the science of climate change; however, they maintainthat the science depicted in the film is largely accurate. An Inconvenient

Truth won Academy Awards in 2006 for best feature-length documen-tary and best song. Since its release, the film has been made part of thecurricula of many schools around the world, but some local school dis-tricts have voiced their disapproval of its use without the presentationof opposing views.



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Additional measures that would reduce personal emis-sions of greenhouse gases and also conserve energy includethe operation of more energy-efficient vehicles, the use ofpublic transportation when available, and the transitionto more energy-efficient household products. Individualsmight also improve their household insulation, learnto heat and cool their residences more effectively, andpurchase and recycle more environmentally sustainableproducts.

PUBLIC AWARENESS AND ACTION

In response to ecological and socioeconomic threats asso-ciated with global warming, a number of initiatives inconstruction and transportation have been developed toreduce the production of greenhouse gases. In additionto slowing the pace of global warming, these initiatives are

designed to reduce overall energy use, which carries eco-nomic, as well as environmental, benefits.

Green Architecture

The philosophy of architecture that advocates sustain-able energy sources, the conservation of energy, the

reuse and safety of building materials, and the siting of abuilding with consideration of its impact on the environ-ment is known as green architecture. Its emergence wasa reflection of rising ecological consciousness during the20th century in response to resource-intensive buildingpractices.

In the early 21st century the building of shelter (in

all its forms) consumed more than half of the world’sresources—translating into 16 percent of the Earth’s freshwater resources, 30–40 percent of all energy supplies, and

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50 percent by weight of all the raw materials withdrawnfrom Earth’s surface. Architecture was also responsiblefor 40–50 percent of waste deposits in landfills and 20–30percent of greenhouse gas emissions.

Many architects after the post-World War II buildingboom were content to erect emblematic civic and corpo-rate icons that celebrated profligate consumption andomnivorous globalization. At the turn of the 21st century,however, a building’s environmental integrity—as seen inthe way it was designed and how it operated—became an

important factor in how it was evaluated.

The Rise of Eco-Awareness

In the United States, environmental advocacy, as an orga-nized social force, gained its first serious momentum aspart of the youth movement of the 1960s. In rebellionagainst the perceived evils of high-rise congestion and

suburban sprawl, some of the earliest and most dedicatedeco-activists moved to rural communes, where they livedin tentlike structures and geodesic domes. In a certainsense, this initial wave of green architecture was based onadmiration of the early Native American lifestyle and itsminimal impact on the land. At the same time, by isolatingthemselves from the greater community, these youthful

environmentalists were ignoring one of ecology’s mostimportant principles: that interdependent elements workin harmony for the benefit of the whole.

Influential pioneers who supported a more integra-tive mission during the 1960s and early ’70s includedthe American architectural critic and social philoso-pher Lewis Mumford, the Scottish-born American

landscape architect Ian McHarg, and the British scien-tist James Lovelock. They led the way in defining greendesign, contributing significantly to the popularization of



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environmental principles. For example, in 1973 Mumfordproposed a straightforward environmental philosophy:

The solution of the energy crisis would seem simple: transform

solar energy via plants and produce enough food power and

manpower in forms that would eliminate the wastes and per-

versions of power demanded by our high-energy technology. In

short, plant, eat, and work!

McHarg, who founded the department of landscape

architecture at the University of Pennsylvania, laid the ground rules for green architecture in his seminal book Design with Nature (1969). Envisioning the role of humanbeings as stewards of the environment, he advocated anorganizational strategy, called “cluster development,” that

would concentrate living centres and leave as much natu-ral environment as possible to flourish on its own terms. In

this regard McHarg was a visionary who perceived Earthas a self-contained and dangerously threatened entity.

This “whole Earth” concept also became the basisof Lovelock’s Gaia hypothesis (a concept described inchapter 3). Named after the Greek Earth goddess, hishypothesis defined the entire planet as a single unifiedorganism, continuously maintaining itself for survival.

Lovelock described this organism as

a complex entity involving the Earth’s biosphere, atmosphere,

oceans, and soil; the totality constituting a feedback or cyber-

netic system which seeks an optimal physical and chemical

environment for life on this planet.

During the 1970s the Norwegian environmental phi-losopher Arne Naess proposed a theory of “deep ecology”(or “ecosophy”), asserting that every living creature in

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nature is equally important to Earth’s precisely balancedsystem. Working in exact opposition to this philosophy,the politics and economics of that decade accelerated thedevelopment of green awareness. The lack of business reg-ulation in the United States meant unlimited consumptionof fossil fuels. Meanwhile, the 1973 OPEC oil crisis broughtthe cost of energy into sharp focus and was a painfulreminder of worldwide dependence on a very small num-ber of petroleum-producing countries. This crisis, in turn,brought into relief the need for diversified sources of

energy and spurred corporate and government investmentin solar, wind, water, and geothermal sources of power.

Green Design Takes Root

By the mid-1980s and continuing through the ’90s, thenumber of environmental advocacy societies radicallyexpanded; groups such as Greenpeace, Environmental

Action, the Sierra Club, Friends of the Earth, and theNature Conservancy all experienced burgeoning member-ships. For architects and builders a significant milestone

was the formulation in 1994 of Leadership in Energy andEnvironmental Design (LEED) standards, establishedand administered by the U.S. Green Building Council.These standards provided measurable criteria for the

design and construction of environmentally responsiblebuildings. The basic qualifications are as follows:

1. Sustainable site development involves, when-ever possible, the reuse of existing buildingsand the preservation of the surrounding envi-ronment. The incorporation of earth shelters,

roof gardens, and extensive planting through-out and around buildings is encouraged.2. Water is conserved by a variety of means includ-

ing the cleaning and recycling of gray (previously



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used) water and the installation of building-by-building catchments for rainwater. Water usageand supplies are monitored.

3. Energy efficiency can be increased in a varietyof ways, for example, by orienting buildings totake full advantage of seasonal changes in thesun’s position and by the use of diversified andregionally appropriate energy sources, whichmay—depending on geographic location—include solar, wind, geothermal, biomass, water,

or natural gas.4. The most desirable materials are those that are

recycled or renewable and those that requirethe least energy to manufacture. They ideallyare locally sourced and free from harmfulchemicals. They are made of nonpolluting rawingredients and are durable and recyclable.

5. Indoor environmental quality addresses theissues that influence how the individual feels ina space and involves such features as the senseof control over personal space, ventilation,temperature control, and the use of materialsthat do not emit toxic gases.

The 1980s and early ’90s brought a new surge ofinterest in the environmental movement and the riseto prominence of a group of more socially responsiveand philosophically oriented green architects. TheAmerican architect Malcolm Wells opposed the legacyof architectural ostentation and aggressive assaultson the land in favour of the gentle impact of under-

ground and earth-sheltered buildings—exemplified byhis Brewster, Mass., house of 1980. The low impact,in both energy use and visual effect, of a structurethat is surrounded by earth creates an almost invisible

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architecture and a green ideal. As Wells explained, thiskind of underground building is “sunny, dry, and pleas-ant” and “offers huge fuel savings and a silent, greenalternative to the asphalt society.”

The American physicist Amory Lovins and his wife,Hunter Lovins, founded the Rocky Mountain Institutein 1982 as a research centre for the study and promotionof the “whole system” approach favoured by McHarg andLovelock. Years before the LEED standards were pub-lished, the institute, which was housed in a building that

was both energy-efficient and aesthetically appealing,formulated the fundamental principle of authentic greenarchitecture: to use the largest possible proportion ofregional resources and materials. In contrast to the con-

ventional, inefficient practice of drawing materials andenergy from distant, centralized sources, the Lovins teamfollowed the “soft energy path” for architecture—i.e., they

drew from alternative energy sources.The Center for Maximum Potential Building Systems

(Max Pot; founded in 1975 in Austin, Texas, by theAmerican architect Pliny Fisk III) in the late 1980s joined

with others to support an experimental agricultural com-munity called Blueprint Farm, in Laredo, Texas. Its broadermission—with applications to any geographic location—

was to study the correlations between living conditions,botanical life, the growing of food, and the economic-eco-logical imperatives of construction. This facility was builtas an integrative prototype, recognizing that naturethrives on diversity. Fisk concluded that single-enterpriseand one-crop territories are environmentally dysfunc-tional—meaning, for example, that all of a crop’s predators

converge, natural defenses are overwhelmed, and chemi-cal spraying to eliminate insects and weeds becomesmandatory. In every respect, Blueprint Farm stood fordiversified and unpredictable community development.



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The crops were varied, and the buildings were constructedof steel gathered from abandoned oil rigs and combined

with such enhancements as earth berms, sod roofs, andstraw bales. Photovoltaic panels, evaporative cooling,and wind power were incorporated in this utopian demon-stration of the symbiotic relationships between farmingand green community standards.

The American architect William McDonough rose to green design fame in 1985 with his Environmental DefenseFund Building in New York City. That structure was one

of the first civic icons for energy conservation resultingfrom the architect’s close scrutiny of all of its interiorproducts, construction technology, and air-handling sys-tems. Since then, McDonough’s firm established valuableplanning strategies and built numerous other green build-ings—most significantly, the Herman Miller factory andoffices (Holland, Mich., 1995), the corporate offices of

Gap, Inc. (San Bruno, Calif., 1997), and Oberlin College’sAdam Joseph Lewis Center for Environmental Studies(Oberlin, Ohio, 2001).

McDonough’s main contribution to the evolution ofsustainable design was his commitment to what he hascalled “ecologically intelligent design,” a process thatinvolves the cooperation of the architect, corporate lead-

ers, and scientists. This design principle takes into accountthe “biography” of every aspect of manufacture, use, anddisposal: the choice of raw ingredients, transport of mate-rials to the factory, fabrication process, durability of goodsproduced, usability of products, and recycling potential.McDonough’s latest version of the principle—referred toas “cradle-to-cradle” design—is modeled after nature’s

own waste-free economy and makes a strong case for the goal of reprocessing, in which every element that is usedin or that results from the manufacturing process has itsown built-in recycling value.

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The Principles of Building Green

The advances in research and in building techniquesachieved by the above-mentioned green design lumi-naries have been compiled into a reliable database ofenvironmental construction methods and sustainablematerials—some of which have been in use for thousandsof years yet remain the basis for contemporary advancesin environmental technology. For private residences of the21st century, the essential green design principles are as

follows:

• Alter native energy sources. Whenever feasible,build homes and communities that supply theirown power; such buildings may operate entirelyoff the regional power grid, or they may be ableto feed excess energy back onto the grid. Wind

and solar power are the usual alternatives. Thequality of solar collectors and photovoltaicpanels continues to improve with the advanceof technology; practical considerations forchoosing one supplier over another includeprice, durability, availability, delivery method,technology, and warranty support.

• Energy conservation. Weatherize buildings formaximum protection against the loss of warmor cool air. Major chemical companies havedeveloped responsibly manufactured, depend-able, moisture-resistant insulating materialsthat do not cause indoor humidity problems.

Laminated glass was also radically improvedat the end of the 20th century; some windowsprovide the same insulation value as traditional



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stone, masonry, and wood construction. Inregions that experience extreme heat, straw-bale or mud-brick construction—used sinceancient times—is a good way to save moneyand energy.

• Reuse of materials. Use recycled building materi-als. Although such products were scarce in theearly 1990s, since the early 21st century theyhave been readily available from a burgeoning

number of companies that specialize in salvag-ing materials from demolition sites.

• Careful siting . Consider using underground orearth-sheltered architecture, which can beideal for domestic living. Starting at a depth ofabout 1.5 metres (5 feet) below the surface, the

temperature is a constant 52 °F (11 °C)—whichmakes the earth itself a dependable source ofclimate control.

Individual, corporate, and governmental efforts tocomply with or enforce LEED standards include recyclingat household and community levels, constructing smaller

and more efficient buildings, and encouraging off-the-gridenergy supplies. Such efforts alone cannot preserve the global ecosystem, however. On the most basic level, theultimate success of any globally sanctioned environmentalmovement depends as much on its social, psychological,and aesthetic appeal as on its use of advanced technologies.

The environmental movement in the 21st century can

succeed only to the extent that its proponents achieve abroad-based philosophical accord and provide the samekind of persuasive catalyst for change that the Industrial

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Revolution offered in the 19th century. This means shap-ing a truly global (as well as optimistic and persuasive)philosophy of the environment. Much depends on thebuilding arts and integrative thinking. Architects will haveto abandon 20th-century specialization and reliance ontechnology and, with builders and clients, help support

grassroots, community-oriented, and globally unifyingobjectives. In the words of Earth Day founder GaylordNelson,

The ultimate test of man’s conscience may be his willingness to sacrifice something today for future generations whose words

of thanks will not be heard.

The Challenges to Architecture

If architecture is to become truly green, then a revolutionof form and content—including radical changes in the

entire look of architecture—is essential. This can onlyhappen if those involved in the building arts create a fun-damentally new language that is more contextuallyintegrative, socially responsive, functionally ethical, and

visually germane.The potentialities of environmental science and tech-

nology must be creatively examined. Already there exists a

rich reservoir of ideas from science and nature—cyber-netics, virtual reality, biochemistry, hydrology, geology,and cosmology, to mention a few. Furthermore, just as theIndustrial Revolution once generated change in manyfields in the 19th century, so too the information revolu-tion, with its model of integrated systems, serves as aconceptual model in the 21st century for a new approach

to architecture and design in the broader environment.As community governments begin to legislate state-of-the-art green standards, they must encourage appropriateartistic responses to such regional attributes as surrounding



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topography, indigenous vegetation, cultural history, andterritorial idiosyncrasy. For instance, communities mightencourage innovative fusions of architecture with land-scape—where trees and plants become as much a part ofarchitectural design as construction materials—so thatbuildings and their adjacent landscapes essentially merge.In such thinking, buildings are not interpreted as isolatedobjects, and the traditional barriers between inside andoutside and between structure and site are challenged.

Likewise, green architecture in the 21st century has

similar obligations to the psychological and physical needsof its inhabitants. Buildings are most successful when theyrespond to multiple senses—meaning that truly greendesign engages touch, smell, and hearing as well as sight inthe design of buildings and public spaces.

Continuing advances in environmental technologyhave significantly strengthened the goals of sustainable

architecture and city planning over the last decade. Yetmany people consider the environmental crisis beyondtheir comprehension and control. Though technologicalsolutions are necessary, they represent only one facet ofthe whole. Indeed, the transfer of responsibility to engi-neers and scientists threatens the social and psychologicalcommitment needed for philosophical unity.

Increasing numbers of people seek new symbioticrelationships between their shelter and the broader ecol-ogy. This growing motivation is one of the most promisingsigns in the development of a consensus philosophy of theenvironment. As the environmental movement gainsmomentum, it underlines the anthropologist MargaretMead’s observation:

Never doubt that a small group of thoughtful, committed citi-

zens can change the world. Indeed, it is the only thing that

ever has.

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Component systems of a typical electric automobile and hybrid gasoline-elec-tric automobile. Encyclopædia Britannica, Inc.

Electric and Hybrid Vehicles

Since the 1990s interest has risen in the development of vehicles that can be powered by sources other than gaso-line. Such alternatives include fully electric vehicles,

gasoline-electric hybrids, and vehicles powered by diesel,ethanol, and other fuels.

Modern electric cars and trucks have been manufac-tured in small numbers in Europe, Japan, and the UnitedStates since the 1980s. However, electric propulsion is

only possible for relatively short-range vehicles, usingpower from batteries or fuel cells. In a typical system, a

group of lead-acid batteries connected in a series powerselectric alternating-current (AC) induction motors to pro-pel the vehicle. When nickel–metal hydride batteries aresubstituted, the driving range is doubled. A solid-state rec-tifier, or power inverter, changes the direct current (DC)

supplied by the battery pack to an AC output that is con-trolled by the driver using an accelerator pedal to vary theoutput voltage. Because of the torque characteristics of



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electric motors, conventional gear-type transmissions arenot needed in most designs. Weight and drag reduction, as

well as regenerative systems to recover energy that wouldotherwise be lost, are important considerations in extend-ing battery life. Batteries may be recharged in six hoursfrom a domestic electrical outlet.

Conventional storage-battery systems do not havehigh power-to-weight ratios for acceleration or energy-to-

weight ratios for driving range to match gasoline-powered general-purpose vehicles. Special-purpose applications,

however, may be practical because of the excellent low-emission characteristics of the system. Such systems havebeen used to power vehicles on the Moon and in special-ized small vehicles driven within factories.

Several hybrid vehicles are now being produced. Theycombine an efficient gasoline engine with a lightweight,high-output electric motor that produces extra power

when needed. During normal driving, the motor becomesa generator to recharge the battery pack. This eliminatesthe need to plug the car into an electrical outlet forrecharging. The primary advantage of hybrids is that thesystem permits downsizing the engine and always operat-ing in its optimum efficiency range through the use ofadvanced electronic engine and transmission controls.

Experimental Systems

The U.S. automotive industry spends $18 billion or moreon research and development of future products in atypical year—the most spent by any industry in theUnited States. Increasing pressure from various gov-ernments is requiring manufacturers to develop very

low-emission vehicles. Authorities in California esti-mate that motor vehicles produce 30 percent of the greenhouse gases that they consider to be responsiblefor ozone layer degradation in the upper atmosphere.

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To meet this challenge, manufacturers are working on atimetable to produce more efficient vehicle designs.

Expansion of the total potential automotive marketin the future and concern for the environment may beexpected to change cars of the future. Special-purpose

vehicles designed for specific urban or rural functions, with appropriate power systems for each type of use, maybe needed. Possibilities include solar, steam, gas turbine,new hybrid combinations, and other power sources.

Steam power plants have been reexamined in the light

of modern technology and new materials. The continu-ous-combustion process used to heat the steam generatoroffers potentially improved emission characteristics.

Gas turbines have been tested extensively and have good torque characteristics, operate on a wide variety offuels, have high power-to-weight ratios, meet emissionstandards, and offer quiet operation. Some studies have

shown that the advantages of the system are best realizedin heavy-duty vehicles operating on long, nearly constantspeed runs. Efficiencies and operating characteristics canbe improved by increasing operating temperatures. Thismay become commercially feasible utilizing ceramicmaterials that are cost-effective. Successful designs requireregenerative systems to recover energy from hot exhaust

gas and transfer it to incoming air. This improves fueleconomy, reduces exhaust temperatures to safer levels,and eliminates the need for a muffler in some designs.

A number of other designs have been studied involving variations of engine combustion cycles such as turbo-charged gasoline and diesel (two- and four-stroke) designs.Rotary engines have been produced in Germany and

Japan, but they have been discontinued because of exhaustemission control complexity. Variable valve timing canoptimize performance and economy and provide a moreconstant engine torque output at different engine speeds.



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By delaying the opening of the engine exhaust valve,exhaust gas is effectively recirculated to reduce tailpipeemissions. Electro-hydraulic valves that totally replacethe complexity of camshaft designs, or idlers that may bemoved to change the geometry of the camshaft timingchain and retard valve opening, may be used for thispurpose.

Solar-powered electric demonstration vehicles havebeen built by universities and manufacturers. Solar collec-tor areas have proved to be too large for conventional cars,

however. Development continues on solar cell design.Microprocessors have become increasingly impor-

tant in improving fuel economy and reducing undesirableexhaust emissions for all vehicle types. For example, analogcomputers have been employed for many years in carbure-tors that change fuel rate in response to directly measuredmanifold air pressure. In addition, research to develop

so-called intelligent vehicles that can assist the driver andeven operate without driver intervention, at least on spe-cial roads, has made some progress. These developmentshave been made possible by highly reliable solid-statedigital computers and similarly reliable electronic sensors.The automobile industry has worked with governmen-tal bodies to link vehicles to their environments using

advanced telecommunication signals, electronic systems,and digital computers, both within the vehicle and aboardsatellites and in other remote locations. Applications maybe divided into functions for basic vehicle system assis-tance, safety and security applications, and informationand entertainment systems.

The automobile industry is responsible for about

two-thirds of the rubber, one-third of the platinum, one-third of the aluminum, one-seventh of the steel, andone-tenth of the copper consumed in the United Stateseach year. About three-quarters of the material in a car is

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recyclable, and in the United States 19 out of 20 scrappedcars are recycled. Because the automobile is likely toremain an important part of the transportation system, itrequires continuing improvement in safety and emissioncontrol as well as performance and cost.

Alternative-Fuel Vehicles

In addition to electricity, automotive manufacturers canreplace standard gasoline engines with engines that run onother fuels or engines and power systems that can switch

back and forth between gasoline and electricity.

Diesel Vehicles

After World War II the diesel engine, particularly for lighttrucks and taxis, became popular in Europe because of itssuperior fuel economy and various tax incentives. Duringthe 1970s General Motors converted some gasoline

passenger-car engines to the more economical compres-sion-ignition diesel operation, and Mercedes-Benz,Volkswagen, and Peugeot marketed diesel lines in Americathat derived from their European models. The ebbing offuel shortages and the easing of gasoline prices, combined

with various drawbacks to diesel engines (noise, poorcold-weather starting, limited fuel and service in some

communities), reduced American demand by the early1980s. Europe, which had not embraced diesels for privatepassenger cars, reversed course with the development ofenvironmentally friendly common rail direct-injectiondiesel engines in the late 1990s. By 2005 diesel cars rep-resented roughly half of all European passenger car sales.

Electric Vehicles

The first of the fuel crises, in 1973–74, rekindled interestin electric vehicles in America. Numerous experimentersand entrepreneurs began work on battery electric cars, the



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most successful being the CitiCar built by a Florida com-pany, Sebring Vanguard, Inc. The CitiCar had a plastic,

wedge-shaped, two-seater body over a welded aluminumchassis. Lead-acid batteries supplied power to a 3.5-horse-power General Electric motor. With about 2,600 builtbetween 1974 and 1976 (and another 2,000 of its successor,the ComutaCar, built between 1978 and 1981), the CitiCar

was the most prolific of the late-20th-century electrics.Ultimately, the falling price of oil put an end to electriccar sales.

Subsequent alternative propulsion programs weredriven by environmental concerns. In 1990 the CaliforniaAir Resources Board mandated that within eight yearsall auto manufacturers were to ensure that 2 percent oftheir sales in the state be “zero emission” vehicles. For allpractical purposes this meant battery electrics. GeneralMotors took this edict most seriously, beginning work on

an aluminum backbone frame, composite plastic body,and low-rolling-resistance tires. Introduced in 1996 asthe General Motors EV1, it was offered on lease throughSaturn dealers in Arizona and California. Only 800 werecontracted for, and production halted in 2000, with 100remaining in service through 2005. In 1998–99 GeneralMotors and Ford also offered battery electric pickup

trucks, most of which were placed with governmentfleets. The shortcoming of all these battery electrics wastheir limited range—less than 100 miles with lead-acidbatteries. More capable nickel–metal hydride cells wereinordinately expensive. The faltering efforts resulted inrelaxation of the California mandate.

Electric-Gasoline Hybrids

In 1997 Toyota introduced its four-passenger Prius hybridto the Japanese market. Combining a small gasoline engineand an electric motor through a sophisticated control

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2006 Toyota Prius. In 1997 Toyota introduced the Prius, an electric-gasoline hybrid vehicle. Toyota Motor Sales, USA

system, the Prius uses gasoline power only when neces-sary to supplement electric propulsion or to recharge itsbatteries. (That same year in Europe, the hybrid AudiDuo was introduced, but its poor sales led European man-ufacturers to focus on diesel designs.) Honda was the firstmanufacturer to offer a hybrid in the American market,

the two-passenger Insight in December 1999. In order toestablish hybrid technology in the American marketplace,Toyota initially offered substantial discounts on the Prius

when it introduced it to the United States in 2000; the U.S.and some state governments also offered tax incentivesand other perquisites (such as unlimited use of commuterlanes and exemption from paying parking meters) to

encourage production and sales of alternative-fuel vehi-cles. Although the Prius offered only a relatively modestincrease in fuel economy, the removal of any need to plug



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the batteries in for recharging overcame the chief draw-back to pure electric vehicles. The Prius was an immediatehit with trend-conscious Californians—with many celeb-rities choosing to drive hybrids instead of luxury cars—andprospective buyers often had to wait months for delivery.In 2004 the Ford Escape Hybrid (SUV) became the firstAmerican hybrid, beating two General Motors trucks, theChevrolet Silverado and the GMC Sierra, to market byone year. The first luxury hybrid, the Toyota Lexus SUV,

was introduced in the United States in 2006.

Ethanol and Fuel Cells

In 1999 Brazil mandated that by 2003 all new cars sold inthe country had to be FlexFuel vehicles (FFVs)—vehiclescertified to run on gasoline containing up to 85 percentethanol (ethyl alcohol), marketed as E85. This initiativeled numerous American, European, and Japanese manu-

facturers to certify some of their models as E85-compliant, which is indicated by the eighth character in the vehicleidentification number, or VIN.

General Motors and DaimlerChrysler primarily haveconcentrated on fuel cell development, assisted by U.S.

government grants. However, usable technology for the general public is still years away.

Biofuels

A biofuel is any fuel that is derived from biomass—that is,plant material or animal waste. Since such feedstock mate-rial can be replenished readily, biofuel is considered to bea source of renewable energy, unlike fossil fuels such as

petroleum, coal, and natural gas. Biofuel is perceived byits advocates as a cost-effective and environmentallybenign alternative to petroleum and other fossil fuels,

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particularly within the context of rising petroleum pricesand increased concern over the contributions made byfossil fuels to global warming. Many critics express con-cerns about the scope of the expansion of certain biofuelsbecause of the economic and environmental costs associ-ated with the refining process and the removal of vastareas of arable land from food production.

Types of Biofuels

Some long-exploited biofuels, such as wood, can be used

directly as a raw material that is burned to produce heat.The heat, in turn, can be used to run generators in a powerplant to produce electricity. A number of existing powerfacilities burn grass, wood, or other kinds of biomass.

Liquid biofuels are of particular interest becauseof the vast infrastructure already in place to use them,especially for transportation. The liquid biofuel in

greatest production is ethanol (ethyl alcohol), whichis made by fermenting starch or sugar. Brazil and theUnited States are among the leading producers of eth-anol. In the United States, ethanol biofuel is madeprimarily from corn (maize) grain, and it is typicallyblended with gasoline to produce “gasohol,” a fuel thatis 10 percent ethanol. In Brazil, ethanol biofuel is made

primarily from sugarcane, and it is commonly used as a 100-percent-ethanol fuel or in gasoline blends containing85 percent ethanol.

The second most common liquid biofuel is biodiesel, which is made primarily from oily plants (such as the soy-bean or oil palm) and to a lesser extent from other oilysources (such as waste cooking fat from restaurant deep-

frying). Biodiesel, which has found greatest acceptance inEurope, is used in diesel engines and usually blended withpetroleum diesel fuel in various percentages.



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Other biofuels include methane gas—which can bederived from the decomposition of biomass in the absenceof oxygen—and methanol, butanol, and dimethyl ether—

which are in development.At present, much focus is on the development of meth-

ods to produce ethanol from biomass that possesses a highcellulose content. This cellulosic ethanol could be pro-duced from abundant low-value material, including woodchips, grasses, crop residues, and municipal waste. Themix of commercially used biofuels will undoubtedly shift

as these fuels are developed, but the range of possibilitiespresently known could furnish power for transportation,heating, cooling, and electricity.

Economic and Environmental Considerations

In evaluating the economic benefits of biofuels, the energyrequired to produce them has to be taken into account. For

example, the process of growing corn to produce ethanolconsumes fossil fuels in farming equipment, in fertilizermanufacturing, in corn transportation, and in ethanoldistillation. In this respect ethanol made from corn rep-resents a relatively small energy gain; the energy gain fromsugarcane is greater and that from cellulosic ethanol couldbe even greater.

Biofuels also supply environmental benefits but,depending on how they are manufactured, can alsohave serious environmental drawbacks. As a renewableenergy source, plant-based biofuels in principle makelittle net contribution to global warming and climatechange; the carbon dioxide (a major greenhouse gas) thatenters the air during combustion will have been removed

from the air earlier as growing plants engage in photosyn-thesis. Such a material is said to be “carbon neutral.” Inpractice, however, the industrial production of agricultural

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biofuels can result in additional emissions of greenhouse gases that may offset the benefits of using a renewable fuel.These emissions include carbon dioxide from the burningof fossil fuels during the production process and nitrousoxide from soil that has been treated with nitrogen fertil-izer. In this regard, cellulosic biomass is considered to bemore beneficial.

Land use is also a major factor in evaluating the bene-fits of biofuels. Corn and soybeans are important foods,and their use in producing fuel can therefore affect the

economics of food price and availability. By 2007 aboutone-fifth of the corn output in the United States was allo-cated to the production of biofuel, and one study showedthat even if all U.S. corn land was used to produce ethanol,it could replace just 12 percent of gasoline consumption.In addition, crops grown for biofuel can compete for the

world’s natural habitats. For example, emphasis on etha-

nol derived from corn is shifting grasslands and brushlandsto corn monocultures, and emphasis on biodiesel is bring-ing down ancient tropical forests to make way for palmplantations. Loss of natural habitat can change the hydrol-ogy, increase erosion, and generally reduce biodiversity of

wildlife areas. The clearing of land can also result in thesudden release of a large amount of carbon dioxide as the

plant matter that it contains is burned or allowed to decay.Some of the disadvantages of biofuels apply mainly tolow-diversity biofuel sources—corn, soybeans, sugarcane,oil palms—which are traditional agricultural crops. Onealternative involves the use of highly diverse mixtures ofspecies, with the North American tallgrass prairie as a spe-cific example. Converting degraded agricultural land that

is out of production to such high-diversity biofuel sourcescould increase wildlife area, reduce erosion, cleanse waterborne pollutants, store carbon dioxide from the air



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as carbon compounds in the soil, and ultimately restorefertility to degraded lands. Such biofuels could be burneddirectly to generate electricity or converted to liquid fuelsas technologies develop.

The proper way to grow biofuels to serve all needssimultaneously will continue to be a matter of muchexperimentation and debate, but the fast growth in biofuelproduction will likely continue. In the European Union,for example, biofuels are planned to account for 5.75 per-cent of transport fuels by 2010, and 10 percent of European

vehicles are expected to run exclusively on biofuels by2020. In the United States the Energy Independence andSecurity Act of 2007 mandated the use of 136 billion litres(36 billion gallons) of biofuels annually by 2020, more thana sixfold increase over 2006 production levels. The legisla-tion also requires, with certain stipulations, that 79 billionlitres (21 billion gallons) of the total amount be biofuels

other than corn-derived ethanol, and it continued cer-tain government subsidies and tax incentives for biofuelproduction. In addition, the technology for producingcellulosic ethanol is being developed at a number of pilotplants in the United States.

One distinctive promise of biofuels is that, in combi-nation with an emerging technology called carbon capture

and storage, the process of producing and using biofuelsmay be capable of perpetually removing carbon dioxidefrom the atmosphere. Under this vision, biofuel crops

would remove carbon dioxide from the air as they grow,and energy facilities would capture the carbon dioxide

given off as biofuels are burned to generate power.Captured carbon dioxide could be sequestered (stored) in

long-term repositories such as geologic formationsbeneath the land, in sediments of the deep ocean, or con-ceivably as solids such as carbonates.

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CONCLUSION

As far as we know, Earth is the only planet that supportslife. One of the most important features of the Earth sys-tem that allows life to exist is the planet’s hospitableclimate. Neither too hot nor too cold for living things,animals in particular can be found from Earth’s polarregions to its Equator. The reliability of Earth’s climate isessential to the ecological well-being of the planet’s life-forms. All plants, animals, and other organisms alive today

have successfully adapted to the ecological conditions in which they live, and these conditions are heavily influ-enced by temperature and rainfall patterns across longperiods of time. When these patterns are upset, ecosys-tems change, and the living things that inhabit ecosystemsin flux must adapt to the new conditions or perish.

Climate change, like ecological change, is constant. For

extended periods of time, the climate characteristics of aparticular ecosystem may remain within a certain range.Over time, however, the natural drivers of climate mayalter the temperature and precipitation patterns enoughto transform how energy and nutrients move through theecosystem. Usually, these changes occur slowly. Whilesome species die out or migrate out of the transformed

ecosystem, many can acclimate to the new conditions.Some intervals of geologic time, however, have beencharacterized by tremendous climate upheavals, whichresulted in the demise of large numbers of species.

Most scientists agree that until about 1750 CE natu-ral forces were the only factors that affected climate. Bythe middle of the 18th century the Industrial Revolution

had commenced, and human beings began to burn largeamounts of coal to smelt metals and power machines. Asthe countries of Europe and North America switched



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from agrarian to industrial societies over the following250 years, the use of coal and other fossil fuels, such aspetroleum and natural gas, expanded, which resulted inthe release of tremendous amounts of carbon dioxide andother gases.

Many scientists assert that the release of these greenhouse gases—which strengthen the atmosphere’sheat-trapping ability—is partially responsible for anincrease in global average temperatures in recent decades.The rising concentration of greenhouse gases in the atmo-

sphere slows the emission of longwave radiation fromEarth’s surface back to space, and thus near-surface airtemperatures rise. As this additional heat is distributedaround the globe, climate patterns change, increasing therisk of environmental, as well as economic, disruption.

The world is responding to the threat of global warm-ing. As efforts to develop international agreements that

limit the emission of carbon dioxide and other greenhouse gases continue, many cities around the world have alreadyadopted policies to curb emissions. In addition, technolo-

gies and practices are emerging that promise to reduceemissions of greenhouse gases. Some of these methods

greatly increase the efficiency of machines that use fossilfuels, whereas others produce no net emissions or effec-

tively remove greenhouse gases from the atmosphere.

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ameliorate To make better; to improve.anomalous Deviating from what is normal or expected.anthropogenic Produced by humans.bathymetric Related to the measurement of depths of

water in oceans, seas, and lakes.biogenic Produced by living organisms.cellulose A complex carbohydrate, or polysaccharide,

consisting of 3,000 or more glucose units.chlorofluorocarbon Any of several organic compounds

containing carbon, fluorine, and chlorine.contiguous Adjacent.denitrification The removal or loss of nitrogen from a

compound.desalination Removal of dissolved salts from seawater

and from the salty waters of inland seas, highly miner-alized groundwaters, and municipal wastewaters.

diurnal Recurring on a daily basis.hydrostatics Branch of physics that deals with the char-

acteristics of fluids at rest, particularly with thepressure in a fluid or exerted by a fluid (gas or liquid)on an immersed body.

isobar Line on a weather map of constant barometricpressure drawn on a given reference surface.isotope One of two or more atoms with the same atomic

number (number of protons in the nucleus) but withdifferent atomic weights (numbers of neutrons in thenucleus).

katabatic wind Wind that blows down a slope because

of gravity. It occurs at night, when the highlands radi-ate heat and are cooled. Also called downslope windor gravity wind.



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7 Glossary 7

latent heat Characteristic amount of energy absorbed orreleased by a substance during a change in physicalstate that occurs without a change in temperature.

luminosity The quality or state of giving off light, espe-cially self-generated light.

orographic That which pertains to or is caused bymountains, such as orographic clouds or orographicprecipitation.

perihelion The point in the path of a celestial body(such as a planet) that is nearest to the Sun. (The aph-

elion is the point farthest from the Sun.)protocol Code of correct conduct.proxy An authorized substitute.rime White, opaque, granular deposit of ice crystals

formed on objects that are at a temperature below thefreezing point.

stalagmite An icicle-shaped rock made of dissolved cal-

cium minerals formed on the floor of a cave by thedrip of calcium-rich water. (A stalactite, similarly con-stituted, hangs from the ceiling or sides of a cave.)

stoma (plural stomata) From the Greek for mouth, asimple, minute opening in the surface layer of aplant’s leaves and stems.

stratosphere Layer of the atmosphere above the

troposphere.subduction A geologic action or process in which oneedge of one crustal plate descends below the edge ofanother.

taiga Beginning where the tundra ends, a subarcticbiome characterized by conifers and a lichen-coveredforest floor.

torque A force that tends to turn the body to which it isapplied.troposphere Lowest layer of the atmosphere.



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CLIMATE

Introductory works include Edward Aguado and James E.Burt, Understanding Weather and Climate , 4th ed. (2007); P.Kabat et al. (eds.), Vegetation, Water, Humans, and theClimate: A New Perspective on an Interactive System (2004);Intergovernmental Panel on Climate Change: Working

Group I, Climate Change 2007: The Physical Science Basis(2007); and National Research Council (U.S.), Panel onClimate Change Feedbacks, Understanding Climate Change

Feedbacks (2003). Two excellent comprehensive reference works are John E. Oliver (ed.), The Encyclopedia of WorldClimatology (2005); and Stephen H. Schneider (ed.),

Encyclopedia of Climate and Weather , 2 vol. (1996).

Definitions of meteorological terms are provided in ToddS. Glickman (ed.), Glossary of Meteorology , 2nd ed. (2000);and Secretariat of the World Meteorological Organization,

International Meteorological Vocabulary , 2nd ed. (1992),including nomenclature in English, French, Russian, andSpanish.

Current research is reported in the following journals:

Bulletin of the American Meteorological Society (monthly);Climatic Change (6/yr.); International Journal of Biometeorology (6/yr.); Journal of Applied Meteorology (monthly);

International Journal of Climatology (15/yr.); Journal of Meteorological Research (bimonthly); International Journal of Meteorology (10/yr.); Monthly Weather Review ; Quarterly Journal of the Royal Meteorological Society ; Russian Meteorology

and Hydrology (monthly); Global Climate Review (monthly);Weather (monthly); Weatherwise (bimonthly); and WMO Bulletin (quarterly).



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CLIMATIC CLASSIFICATION

Useful introductory discussions on the classificationof climatic zones can be found in Alan H. Strahler andArthur N. Strahler, Introducing Physical Geography, 4th ed.(2006); and Tom L. McKnight and Darrel Hess, PhysicalGeography: A Landscape Appreciation, 9th ed. (2007).Advanced treatments are provided in P.J. Robinson andA. Henderson-Sellers, Contemporary Climatology, 2nded. (1999); and Roger G. Barry and Richard J. Chorley,

Atmosphere, Weather, and Climate, 8th ed. (2003). The globaldistribution of major climate types is the subject of GlennR. McGregor and S. Nieuwolt, Tropical Climatology: An

Introduction to the Climates of the Low Latitudes, 2nd ed.(1998). The foundational work on modern climate clas-sification is contained in Glenn T. Trewartha and Lyle H.Horn, An Introduction to Climate, 5th ed. (1980). Particular

regions are examined in H.E. Landsberg (ed.), WorldSurvey of Climatology, 16 vol. in 18 (1969–2001); and theoft-celebrated account contained in Glenn T. Trewartha,The Earth’s Problem Climates, 2nd ed. (1981).

CLIMATE AND LIFE

A brief history of the coevolution of life and the atmo-sphere combined with the environmental consequencesof changing the composition of the modern atmosphereis presented in Karl K. Turekian, Global EnvironmentalChange: Past, Present, and Future (1996). Additionally, the

work by William R. Cotton and Roger A. Pielke, Sr., Human Impacts on Weather and Climate, 2nd ed. (2007), chronicles

the effects of human-induced landscape modificationupon the dynamics of regional weather and climate.Interactions between organisms, the ecological sys-

tems they inhabit, and the atmosphere are detailed in




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William P. Lowry and Porter P. Lowry, Fundamentals of Biometeorology: Interactions of Organisms and the Atmosphere,2 vol. (2001); and Stephen H. Schneider, Encyclopedia ofClimate and Weather , 2 vol. (1996). The Gaia hypothesis isoutlined in James E. Lovelock, Gaia: A New Look at Life on

Earth (2000); and the role played by the biosphere in con-trolling the atmosphere and the oceans throughout

geologic time is explained in James E. Lovelock, The Agesof Gaia: A Biography of Our Living Earth, rev. ed. (1995).Causes of climatic change are explained in detail in Bert

Bolin et al. (eds.), The Greenhouse Effect, Climatic Change, and Ecosystems (1989). The first systematic treatment of thechemical and physical significance of atmospheric trace

gases produced by the biosphere, including most of theearly work on the greenhouse effect, is performed in JohnTyndall, Heat Considered As a Mode of Motion (1871).

CLIMATE CHANGE

Overviews of Earth’s climate system combined with a gen-eral treatment of climate variation since the PleistoceneEpoch are provided in William F. Ruddiman, Earth’sClimate: Past and Future, 2nd ed. (2008); Tjeerd H. vanAndel, New Views on an Old Planet: A History of Global

Change, 2nd ed. (1994); and Richard B. Alley, The Two-MileTime Machine: Ice Cores, Abrupt Climate Change, and Our

Future (2000). The impacts of recent climate variation andchange upon society are considered in César N. Caviedes,

El Niño in History: Storming Through the Ages (2001); BrianFagan, Floods, Famines, and Emperors: El Niño and the Fateof Civilizations (1999); and Mike Davis, Late Victorian

Holocausts: El Niño Famines and the Making of the Third World (2001). A discussion of the cultural and historical effects ofthe Medieval Warm Period and Little Ice Age is provided



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in Brian Fagan, The Little Ice Age: How Climate Made History,1300–1850 (2000).

More advanced treatments of recent climate vari-ation include Vera Markgraf (ed.), InterhemisphericClimate Linkages (2001); Neil Roberts, The Holocene: An

Environmental History, 2nd ed. (1998); H.E. Wright, Jr., etal. (eds.), Global Climates Since the Last Glacial Maximum (1993); Jean M. Grove, Little Ice Ages: Ancient and Modern,2nd ed. (2004); Anson Mackay et al. (eds.), Global Changein the Holocene (2005); and The National Research Council,Surface Temperature Reconstructions for the Last 2,000 Years (2006). Past climate variation within the context of futureclimate change is presented in the Intergovernmental Panelon Climate Change: Working Group I, Climate Change

2007: The Physical Science Basis: Summary for Policymakers: Fourth Assessment Report (2007). A general review of abruptclimate change is provided in John D. Cox, Climate Crash:

Abrupt Climate Change and What It Means for Our Future (2007); and a more technical treatment can be found inThe National Research Council, Abrupt Climate Change:

Inevitable Surprises (2002).The development of early life and its effects on

Earth’s oceans and atmosphere are discussed in AndrewH. Knoll, Life on a Young Planet: The First Three Billion Years

of Evolution on Earth (2005). Pre-Pleistocene climates aresummarized in Thomas J. Crowley and Gerald R. North, Paleoclimatology (1991); and Lawrence A. Frakes, Jane E.Francis and Jozef I. Syktus, Climate Modes of the Phanerozoic:The History of Earth’s Climate over the Past 600 Million Years (1992). Methods of pre-Pleistocene climate investiga-tion are thoroughly discussed in Judith Totman Parrish,

Interpreting Pre-Quaternary Climate from the Geologic Record (1998). A highly readable narrative that considers the 19th-century and early to mid-20th-century development of




336

ideas surrounding the orbital forcing of climate is provided in John Imbrie and Katherine Palmer Imbrie, Ice Ages: Solving the Mystery (1979, reissued 1986).

GLOBAL WARMING SCIENCE

An excellent general overview of the factors governingEarth’s climate over all timescales is presented in WilliamRuddiman, Earth’s Climate: Past and Future (2000). In addi-tion, Richard C.J. Somerville, The Forgiving Air:

Understanding Environmental Change (1996, reissued 1998),is a readable introduction to the science of climate and

global environmental change. John Houghton, GlobalWarming: The Complete Briefing (1997), also offers an acces-sible treatment of the science of climate change as well asa discussion of the policy and ethical overtones of climatechange as an issue confronting society. Spencer Weart,

Discovery of Global Warming (2003), provides a reasonedaccount of the history of climate change science.

A somewhat more technical introduction to the sci-ence of climate change is provided in David Archer, GlobalWarming: Understanding the Forecast (2006). More advancedtreatments of the science of global warming and cli-mate change are included in Intergovernmental Panel on

Climate Change: Working Group I, Climate Change 2007:The Physical Science Basis: Summary for Policymakers: Fourth

Assessment Report (2007); and Intergovernmental Panelon Climate Change: Working Group II, Climate Change

2007: Climate Change Impacts, Adaptations, and Vulnerability: Fourth Assessment Report (2007). Possible solutions to thechallenges of global warming and climate change are

detailed in Intergovernmental Panel on Climate Change:Working Group III, Climate Change 2007: Mitigation ofClimate Change: Fourth Assessment Report (2007).



337

7 Bibliography 7

A number of books present thoughtful discussions of global warming as an environmental and societal issue. Stillprescient is an early account provided in Bill McKibben,The End of Nature (1989). Other good treatments includeStephen Schneider, Laboratory Earth (2001); AlbertGore, An Inconvenient Truth (2006); Elizabeth Kolbert,

Field Notes from a Catastrophe (2006); Eugene Linden,The Winds of Change (2006); Tim Flannery, The Weather

Makers (2006); and Mike Hulme, Why We Disagree AboutClimate Change: Understanding Controversy, Inaction and

Opportunity (2009). An excellent exposition for youngerreaders is found in Andrew Revkin, The North Pole Was

Here (2007).

GLOBAL WARMING POLICY

Stephen H. Schneider, Armin Rosencranz, and John

O. Niles (eds.), Climate Change Policy: A Survey (2002),is a primer on various aspects of the policy debate thatexplains alternatives for dealing with climate change. Abroad analysis of the climate change debate is impartedin Andrew E. Dessler and Edward A. Parson, The Science

and Politics of Global Climate Change: A Guide to the Debate (2006). A summary of the quantitative aspects of green-

house gas emissions designed to assist stakeholders andpolicy makers is provided in Kevin A. Baumert, TimothyHerzog, and Jonathan Pershing, Navigating the Numbers:Greenhouse Gas Data and International Climate Policy (2005).

John T. Houghton, Global Warming: The Complete Briefing ,3rd ed. (2004), offers a perspective on climate changefrom one of the leading participants in the IPCC pro-

cess. Daniel Sarewitz and Roger Pielke, Jr., “Breakingthe Global-Warming Gridlock,” The Atlantic Monthly,286(1):55–64 (2000), presents an alternative view on how




338

to make progress on climate policy by focusing on reduc-ing vulnerability to climate impacts.

Thoughtful discussions of the politics underlying theissue of climate change are provided in Ross Gelbspan,

Boiling Point (2004); Mark Lynas, High Tide (2004); andRoss Gelbspan, The Heat Is On (1998). The social justiceimplications involved in adapting the human popula-tion to changing climatic conditions are presented in W.Neil Adger et al. (eds.), Fairness in Adaptation to ClimateChange (2006).



339339

A

adiabatic temperature changes, 77aerosol emissions, and global

warming, 234, 243 – 246Agassiz, Louis, 172alternative-fuel vehicles, 320 – 323anticyclones, 61– 62, 69 – 71, 72,

124, 125, 129, 130Archean Eon, Earth during,138 – 140

atmospheric humidity, 18 – 30atmospheric pressure, 56 – 58, 59automobiles, electric and hybrid,

316 – 323

BBergen school, 64, 68bioclimatology, 149biofuels, 323 – 327biogenic gases, 136, 142

cycling of, 140, 143 – 164biosphere

controls on maximum

temperatures, 153 – 154controls on minimum

temperatures, 154– 157controls on planetary bound-

ary layer, 151– 153controls on structure of

atmosphere, 150controls on surface friction

and localized winds, 159 – 160and Earth’s energy budget,140 – 142

impacts on precipitationprocesses, 160 – 164

role of in the Earth-atmosphere system, 140 – 142

Bjerknes, Jacob, 68Bjerknes, Vilhelm, 64

Ccarbon cycle, 143 – 145, 146

feedbacks, 263 – 265, 276carbon dioxide, as greenhouse

gas, 238 – 240, 249 – 250, 291carbon sequestration, 248,

249 – 252carbon sinks, 250 – 251, 292Cenozoic Era, 179, 180, 217, 218

climates of, 214– 216climate

and changes in albedo of thesurface, 157 – 158

defined, 1, 102, 168of early Earth, 219 – 222and humans, 136, 140 – 141,

164– 167, 184– 186, 197 – 198,

225– 226, 227, 230, 232, 234– 252and life, 135– 167natural influences on, 252 – 260

climate change, 168 – 226abrupt climate changes in

Earth history, 223 – 226causes of, 176 – 186centennial-scale variation,

198 – 200decadal variations, 192 – 197defined, 168




340

and Earth system, 169 – 173and emergence of

agriculture, 209

evidence for, 174– 175interannual variation, 189 – 192millennial and multimillennial

variation, 200 – 204seasonal variation, 187 – 189simulations of future, 276 – 282since the emergence of

civilization, 197 – 204

since the emergence ofhumans, 204– 213through geologic time,

213 – 222 within a human life span,

186 – 197climate research, 265– 273

modern observations, 266 – 268prehistorical climate records,

268 – 269theoretical climate models,

269 – 273climatic classification, 102 – 134

approaches to, 105– 111empirical classifications, 106,

108 – 111 genetic classifications, 106 – 108

limitations of, 102 – 104climatic types, world distribu-

tion of, 114– 134type A, 114– 118type B, 118 – 123type C and D, 123 – 130type E, 130 – 133type H, 133 – 134

climatology, about, 9, 149cloud condensation nuclei,160 – 162

cloudsfeedbacks, 262 – 263origin of precipitation in,

31– 32types of, 32 – 39continentality, effect on tem-

peratures, 10 – 11Coriolis force, 59 – 60, 61, 72, 73,

74, 90, 97Cretaceous Period, 214, 215Croll, James, 172 – 173

cycling of biogenic atmospheric gases, 140, 143 – 164cyclones, 61 – 69, 72, 124, 125,

128, 130climatology of, 71– 72extratropical, 62, 63 – 69 global warming and, 281– 282

DDarwin, Charles, 172deforestation, 157, 166, 184, 185,

247 – 248, 250dew point, 23 – 24droughts, 83, 109, 166, 172, 174,

187, 189, 192, 196, 197,199 – 200, 286, 287, 288

EEarth

abrupt climate changes inhistory of, 223 – 226

albedo of surface and changesin climate, 157 – 158, 246 – 248

climates of early, 219 – 222 variations in orbit and global

warming, 259 – 260Earth system, 168, 169 – 173,

176, 182



341

feedback within, and climatechange, 182 – 184

Earth system history, 171, 172, 174

Earth system science, 170 – 171electric vehicles, 316 – 317,

320 – 321El Niño, 17, 82, 85, 189 – 191,

193 – 196El Niño/Southern Oscillation

(ENSO), 82, 189 – 192,193 – 194, 196, 197, 202 – 203,

253, 254, 272, 277 – 278, 282, 286Environmental ProtectionActs, 167

Eocene Epoch, 215, 216, 225evolution of life and the atmo-

sphere, 134, 138 – 140

F

faint young Sun paradox,219 – 220feedbacks, 176, 182 – 184, 211, 236

and climate sensitivity,260 – 265

Fisk, Pliny, III, 310flooding, 31, 56, 166, 172, 174, 187,

189, 193, 196, 197, 286, 287fossil fuels, use of, 141, 142, 165,

166, 182, 184, 225, 229, 232,238, 240, 249, 268, 274, 275,290, 300, 323 – 324

frost point, 24

GGaia hypothesis, 136 – 137,

166 – 167, 307Garrels, Robert, 139Garstang, Michael, 164

glacial-interglacial cycles, 172,181, 201, 205, 212 – 213

of the Pleistocene, 209 – 212,

242, 260recent, 205– 209

global warming, 142, 165, 227–327

causes of, 230 – 265and climate research, 265– 273debate about, 230, 232, 304defined, 227

effects of, 227 – 229, 274– 290, 291environmental consequencesof, 282 – 286

in the future, 227, 246, 274–282

human activities and, 227, 230,232, 234– 252, 290

public awareness and action,305– 327

and public policy, 290 – 305since Industrial Revolution,

227, 229, 234, 240, 293,328 – 329

socioeconomic consequencesof, 286 – 290

in 20th century, 227Gore, Al, 302 – 304

Gray, Asa, 172 green architecture, 305– 315

beginnings of, 308 – 311challenges to, 314– 315principles of, 312 – 314and rise of eco-awareness,

306 – 308 greenhouse effect, 230 – 232,

233, 290 greenhouse gases, 136, 138, 140,141, 142, 220, 221

7 Index 7




342

and climate change, 178 – 179,182, 184– 185, 229, 230,231– 232, 233, 234, 235– 243,

245, 274, 295public policy on reducing, 290,

291, 293 – 294, 296, 297 – 299,300, 301, 304– 305

Gulf Stream, 15, 16, 272

H

hail, 30, 31, 41, 42 – 43, 49 – 50Holmboe, Jørgen, 68Holocene Epoch, 200, 201– 203,

204, 209, 211humans and climate, 136, 140 –

141, 164– 167, 182, 184– 186,225– 226, 227, 230, 232, 234

climate change since emergence of civilization,197 – 204

climate change since emergence of humans, 204– 213

climate change within ahuman lifespan, 186 – 197

influence of human activity onclimate, 234– 252

humidity, 18 – 30

absolute, 19 – 20average relative, 27 – 28climate and, 25– 27evaporation and, 28 – 30relationship between tempera-

ture and, 22 – 24relative, 19, 20 – 22, 27specific, 20

humidity indexes, 19 – 22hybrid vehicles, 317, 321– 323hydrologic cycle, 30

Iice ages/continental glaciation,

172, 181, 201, 205, 217, 218ideas about, 172 – 173

ice albedo feedback, 263ice nuclei, biogenic, 161, 162 – 163

Inconvenient Truth, An, 302 – 304Indian Ocean monsoon,

203 – 204Intergovernmental Panel on

Climate Change (IPCC),

227 – 229, 233 – 234, 251, 252,269, 274, 291

and scientific consensus on global warming, 294– 296

intertropical convergence zone(ITCZ), 80, 115, 116, 118, 121

J

jet streams, 98 – 100, 123 Julian, Paul, 85

KKarl, T.R., 156katabatic flow, 76 – 79Keeling, Charles, 267 – 268Köppen, Wladimir, 108 – 110, 114,

121, 123, 130Kyoto Protocol, 167, 251, 291–

292, 296, 297 – 299

Llandslides, 31land-use change, and global

warming, 246 – 248latitude, effect ontemperatures, 10



343

Lettau, Heinz, 164life

climate and, 135– 167

evolution of, and the atmo-sphere, 138 – 140

Little Ice Age, 177, 198 – 199,200, 273

Lovelock, James, 137, 306 – 307, 310

Lovins, Amory and Hunter, 310Lyell, Charles, 172

M

macroclimates, 104Madden, Roland, 85Madden-Julian oscillation

(MJO), 85maritime continent, 82Maunder Minimum, 176,

255– 256McDonough, William, 311McHarg, Ian, 306 – 307, 310Medieval Warm Period, 199mesoclimates, 104Mesozoic Era, 218methane, as greenhouse gas,

240 – 242, 291

Milankovitch, Milutin, 173Milankovitch variations, 182,

201, 203Miocene Epoch, 216, 217monsoons, 82 – 86, 115, 123, 174,

191, 203 – 204diurnal variability of, 83interannual variability of,

85– 86intra-annual variability of, 85Mumford, Lewis, 306 – 307

NNaess, Arne, 307 – 308nitrogen cycle, 145– 146North Atlantic Oscillation

(NAO), 191– 192, 196, 197

OOligocene Epoch, 215, 216, 217, 225Oliver, John E., 107orbital (Milankovitch) variations,

and climate change, 172 – 173,180 – 182, 201, 203, 211

oxygen cycle, 143 – 145, 146ozone

as greenhouse gas, 242stratospheric depletion of,

248, 252

PPaleocene-Eocene Thermal

Maximum (PETM), 215,241– 242

Paleocene Epoch, 215, 225paleoclimatology, about, 207Paleozoic Era, 218Penman calculation of

evaporation, 29 – 30

Phanerozoic Eon, 179, 181, 220, 225climates of, 217 – 219

photosynthesis and atmosphericchemistry on early Earth,220 – 221

planetary boundary layer (PBL),151– 153

Pleistocene Epoch, 172, 205, 225

glacial and interglacial cyclesof, 209 – 212, 242, 260

Pliocene Epoch, 216

7 Index 7




344

Precambrian time, 219precipitation, 18, 30 – 56

effects of, 55– 56

mechanisms of release, 39 – 41origin of in clouds, 31– 32types of, 43 – 50 world distribution of, 50 – 55

R radiative forcing, 233 – 234, 236,

239, 240, 243, 246, 247, 259

negative, 243, 246, 247, 248, 252rain, 30, 41– 43, 44– 45, 51– 52, 54

impact of, 31, 55– 56recycling of rainfall, 161, 164

recycled rainfall, 161, 164Redfield, Alfred C., 137Rossby, Carl-Gustaf, 68, 88Rossby waves, 88 – 89, 93 – 96

run off, surface, 55, 56

Ssaturation deficit, 24Schwartz, M.D., 156Slushball Earth hypothesis, 222snow and sleet, 45– 49, 55Snowball Earth hypothesis,

221– 222soil erosion, 31, 55– 56solar output, variations in and

global warming, 254– 258solar radiation and

temperature, 2 – 9average radiation budgets, 6 – 7distribution of radiant energy

from the sun, 2 – 3effects of the atmosphere, 3 – 6surface energy budgets, 7 – 9

solar variability, and climatechange, 176 – 177

Strahler, Arthur N., 107

Ttectonic activity, and climate

change, 179 – 180, 182temperature, 9 – 18

circulation, currents, andocean-atmospheric interac-

tion, 14– 17diurnal, season, and extremetemperatures, 11– 13

global variation of meantemperature, 10 – 11

relationship between humidityand, 22 – 24

short-term temperaturechanges, 17 – 18

solar radiation and, 2 – 9temperature variation with

height, 13 – 14Terjung, Werner H., 107thermal maxima, 201– 202Thornthwaite, C. Warren, 110Tickell, Crispin, 165tornadoes, 74, 166trade winds, 72, 79 – 80Tyndall, John, 154– 155

U

United Nations FrameworkConvention on ClimateChange, 251, 291, 296 – 299

upper-level winds, 86 – 101characteristics of, 86 – 93jet streams, 98 – 100



propagation and developmentof waves, 93 – 96

relationship with surface

features and, 96 – 98in the stratosphere and

mesosphere, 100 – 101

V vegetation patchiness, effect of

on mesoscale climates,

158 – 159 volcanic activity, and climatechange, 177 – 179, 192, 252

volcanic aerosols in atmosphere,and global warming, 230,253 – 254

W

Walker, Gilbert, 193Wallace, Alfred Russel, 172

water vapourfeedback, 261– 262as greenhouse gas, 236 – 237, 262

Wells, Malcolm, 309 – 310 westerlies, 72, 81, 123, 191 wind, 56 – 57, 59 – 82

local, 72 – 79, 159 – 160relationship to pressure and

governing forces, 59 – 61upper-level, 86 – 101zonal surface, 79 – 82

wind systems, 72 – 74, 140macroscale/large-scale, 72mesoscale, 72 – 74sea and land breeze circula-

tion, 74– 76upper-level, 86 – 101

Y

Younger Dryas event, 208,223 – 224

7 Index 7

climate and climate change [john p. rafferty]

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