Download Chapter 4: Climate change

Chapter 4
CHAPTER 4
Climate Change
CLIMATE CHANGE....................................................................................................... 1
4.1
INTRODUCTION ................................................................................................................................ 2
4.2
CLIMATE HISTORY OF THE LAST 250,000 YEARS ............................................................................. 2
4.3
MECHANISMS OF CLIMATE CHANGE ................................................................................................ 4
4.3.1
Plate tectonics......................................................................................................................... 5
4.3.2
Orbital changes ...................................................................................................................... 8
4.3.3
Greenhouse effect ................................................................................................................... 9
4.3.4
Freshwater runoff and thermohaline circulation.................................................................. 16
4.3.5
Solar variability.................................................................................................................... 17
4.3.6
Aerosols ................................................................................................................................ 18
4.4
ANTHROPOGENIC CLIMATE CHANGE ............................................................................................. 20
4.5
CLIMATE FEEDBACKS .................................................................................................................... 24
4.6
TABLES ......................................................................................................................................... 28
4.7
FIGURE LEGENDS .......................................................................................................................... 32
Ecological Climatology
4.1 Introduction
Climate has changed over the course of Earth’s history. Just 18 000 years ago, Earth was in the
grips of a prolonged cold period in which much of North America, northern Europe and northern Russia
were covered with ice. Over the past two million years there have been numerous such ice ages lasting tens
of thousands of years separated by shorter warm interglacial periods. Our current climate is that of a warm
interglacial and history suggests that over the next several thousand years prolonged cooling culminating in
another ice age is possible in the absence of human influences. The geologic record also reveals numerous
rapid climate changes over periods as short as decades or centuries. Climate change is the result of changes
in the external forcing of the climate system by the Sun and internal physical, chemical, and biological
feedbacks among the atmospheric, oceanic, and terrestrial components of the climate system.
4.2 Climate history of the last 250 000 years
Over the past several hundred thousand years, glaciers have cyclically advanced and retreated in
the Northern Hemisphere with a period of about 100 000 years. Glaciation typically takes 90 000 years to
complete while deglaciation occurs over a period of 10 000 years. Ice cores extracted deep below the
surface in Greenland and Antarctica record Earth’s climate history and reveal these ice ages. Atmospheric
gases are trapped in ice as the glaciers grow, providing a record of the temperature and chemical
composition of the atmosphere at the time the ice formed. A 2755-m deep ice core extracted from
Antarctica illustrates climate change over the past 250 000 years (Figure 4.1). The record shows two
distinct cold periods from about 190 000 to 140 000 years ago and from about 115 000 to 18 000 years ago.
During these glacial periods, temperatures were several degrees colder than now and ice advanced
southwards in the Northern Hemisphere. The record also shows two warm periods following these ice ages
from about 140 000 to 125 000 years ago and presently beginning about 15 000 years ago. During these
interglacial periods, climate warmed by several degrees and glaciers retreated northwards.
The last 150 000 years of this record illustrate climate change over a full glacial cycle (Crowley
and North 1991). In the ice age preceding this period, glaciers reached their maximum extent about 140 000
to 150 000 years ago. About 140 000 years ago, climate warmed rapidly, increasing in temperature by 8 °C
over a period of 10 000 years. The rapid deglaciation marked the end of a prolonged glacial period and the
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Chapter 4 – Climate Change
beginning of a warm interglacial period. The onset of the last glaciation began about 115 000 years ago,
when temperatures rapidly dropped by several degrees. A cold period about 75 000 years ago ushered in
the main glacial phase. The subsequent buildup of glaciers culminated in a glacial maximum at about 18
000 to 21 000 years before present (Figure 4.2). At the height of the last ice age, northern North America
was covered by the Cordilleran ice sheet in the west and the Laurentide ice sheet in the east. Ice also
covered Greenland and Iceland. In Europe, glaciers covered Scotland, Fennoscandinavia, the Barents Sea
north of Finland and European Russia, the Kara Sea north of western Siberia, and east Siberia. These ice
sheets were at least one kilometer thick and in many regions were two or more kilometers thick. Then
climate warmed rapidly and glaciers melted over a period of about 6000 years (Kutzbach and Guetter 1986;
COHMAP 1988; Huntley and Webb 1988; Wright et al. 1993). Plants colonized the newly exposed soil,
and vegetation that had been restricted to southern locations migrated northwards in response to the
warming.
The warming arose from changes in solar radiation reaching Earth. At 18 000 years ago, the
amount of solar radiation in the Northern Hemisphere was similar to present conditions. Over the next
several thousand years summertime solar radiation in the Northern Hemisphere increased while wintertime
radiation decreased (Figure 4.3). The increased summer radiation warmed the continents and melted the
glaciers. The greatest increase in summer radiation occurred between 12 000 and 6 000 years ago, when the
seasonality of Northern Hemisphere radiation was greatly enhanced compared with present. During this
time, summer solar radiation in the Northern Hemisphere was about 8% (30 W m-2) greater than present
and winter radiation decreased by a similar amount. Regions of the Northern Hemisphere were warmer
than present. The climate of northern Africa was wetter than present. Since then, summer radiation
decreased and winter radiation increased to present conditions.
Ice cores and other climate records show large, rapid climate changes since the last glacial
maximum (Alley and Clark 1999; Alley 2000; Overpeck and Webb 2000). Changes in temperature of
several degrees throughout large regions of the Northern Hemisphere occurred numerous times in periods
as short as years to decades. One such climate change was the Younger Dryas cold period 11 000 to 10 000
years ago. Around 11 000 years ago, following prolonged warming that brought an end to the last ice age,
temperatures abruptly cooled in North America and Europe. Newly emerged forests reverted back to glacial
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Ecological Climatology
tundra. Glaciers again advanced southwards and down mountains. This cold period lasted for about 1000
years before warming began again. The climate changes were not limited to northern latitudes. Much of
tropical and subtropical Africa experienced increased aridity.
Climate also changes over shorter timescales (Bradley and Jones 1992a; Jones et al. 1998; Mann
et al. 1998). Figure 4.4 shows temperature for the Northern Hemisphere since 1000 A.D. Climate was
relatively mild from about 1000 A.D. to 1200 A.D. It was during this time, known as the ‘Medieval warm
period’, that the Vikings colonized Iceland, Greenland, and North America. Climate change was not limited
to the middle and high latitudes. During the Medieval warm period, the climate of equatorial east Africa
was drier than today (Verschuren et al. 2000). Then climate began to fluctuate, with periods of cold winters
interspersed with periods of warm winters. During this time of colder climate, the Norse colonies in North
America were abandoned. Beginning about 1550 A.D., climate entered a prolonged period of cold
temperatures known as the ‘Little Ice Age’ that lasted until 1850, with a main phase from 1550 to 1700
(Lamb 1995). Winters were long and cold; summers were short. Alpine glaciers advanced to lower
elevations (Lowell 2000). A wetter climate than today prevailed in equatorial east Africa (Verschuren et al.
2000). At about 1700, temperatures began to warm in an erratic recovery from the Little Ice Age.
Temperatures warmed by over 0.5 °C between the mid-1800s and 1940s. Temperatures then cooled over
the next 25 years and have been warming since.
4.3 Mechanisms of climate change
Climate changes naturally and from human influences (Burroughs 1997; Graedel and Crutzen
1997; Houghton 1997; Karl et al. 1997; Schneider 1990, 1997; Karl and Trenberth 1999). Natural
processes include:
•
plate tectonics, which changes the distribution of continents and oceans and creates mountain
ranges;
•
the geometry of Earth’s orbit around the Sun, which changes the amount of solar radiation
received on Earth;
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Chapter 4 – Climate Change
•
the chemical composition of the atmosphere, whereby increasing concentrations of water
vapor, CO2, methane, nitrous oxide, and other radiatively important gases in the atmosphere
warm climate through the greenhouse effect;
•
changes in the cycling of water over land, which alter the runoff of freshwater to oceans and
the thermohaline circulation;
•
solar variability, in which the amount of radiation emitted by the Sun varies in relation to
sunspot activity; and
•
aerosols in the atmosphere, which alter Earth’s radiative balance.
Changes in the biogeographical distribution and functioning of terrestrial ecosystems alter biogeochemical
cycles and surface energy exchange and are other sources of climate change. Humans alter climate through
emission of greenhouse gases and aerosols and through changes in land use and vegetation cover.
4.3.1 Plate tectonics
Plate tectonics, or continental drift, refers to the slow movement of continents at rates of a few
centimeters per year (Kump et al. 1999). About 540 to 500 million years ago, the continents were widely
dispersed along the equator. They drifted together and collided over time so that by 300 million years ago
the continents were gathered into a large supercontinent called Pangaea that was centered on the equator.
Pangaea began to break apart 200 million years ago, and since then the continents have slowly drifted to
their current locations. Our current continental geography is part of a larger cycle in which continents
assemble into a supercontinent that slowly breaks apart only to reassemble again. This cycle of
supercontinent formation and destruction takes about 500 million years to complete. Given that Pangaea
formed 300 million years ago, the next supercontinent is expected to form in 200 million years as the
Pacific Ocean closes.
The period about 80 million years ago, known as the mid-Cretaceous, illustrates the climate
altering influences of plate tectonics (Crowley and North 1991). At this time three continental blocks
formed large contiguous land areas: North America-Greenland-Eurasia; South America-Antarctica-IndiaAustralia; and Africa (Figure 4.5). The oceans consisted of a single large Pacific basin. North America and
Europe were closer together as were South America and Africa. As a result, the Atlantic Ocean did not yet
exist. India and Australia were attached to Antarctica, and the Indian Ocean had not yet formed. Sea level
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Ecological Climatology
was about 100 m to 200 m higher than present. This flooded portions of western Europe, northern Africa,
and interior North America with shallow seas. The southerly location of Africa, India, and Australia
precluded development of a circumpolar Antarctic current. Instead, open passageways between North and
South America and Eurasia and Africa led to the development of a shallow equatorial seaway known as the
Tethys Sea. This seaway provided a circumglobal ocean circulation and is one of the most distinctive
geographic features of the period. The climate during this time was warmer than present, with global mean
annual temperature several degrees warmer. The warming was especially prominent at high latitudes so that
the equator-to-pole temperature gradient was reduced. Atmospheric CO2 levels were two to nine times
higher than present, which contributed to the warming through the greenhouse effect. Changes in oceanic
heat transport are also thought to have contributed to the warmer climate and reduced equator-to-pole
temperature gradient (Barron and Peterson 1989; Brady et al. 1998).
Since this time, the continents have slowly moved to their modern geographic locations. Resultant
changes in ocean surface currents and the thermohaline circulation altered heat transport. The westward
drift of North America and South America led to the formation of the Atlantic Ocean. Africa moved
northwards to converge with Europe, and the Indian subcontinent moved northwards to converge with
Asia. This, combined with the northward movement of Australia and the emergence of the Central
American isthmus, closed the Tethys equatorial seaway. The formation of the North Atlantic, South
Atlantic, and Indian Oceans allowed the development of subtropical gyres at these latitudes. The opening of
the Drake Passage between South America and Antarctica and the northward movement of India and
Australia led to a circumpolar ocean circulation around Antarctica.
Closing of seaways may have been a precursor to the Pleistocene glaciation some 2.75 million
yeas ago. In particular, the Isthmus of Panama gradually closed 13 to 2 million years ago. Beginning some
4 to 5 million years ago, the isthmus was closed sufficiently to affect deep ocean circulation and Northern
Hemisphere climate (Driscoll and Haug 1998; Haug and Tiedemann 1998). Closure redirected Atlantic
Ocean surface currents, enhanced the Gulf Stream, and transported warm surface water to high northern
latitudes. The warmer high latitude climate may have initiated a subsequent glacial period. The warm
climate enhanced precipitation over Eurasia, which increased freshwater runoff to the Arctic Ocean via
Siberian rivers. By decreasing salinity, the input of freshwater facilitated the formation of sea ice, which
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Chapter 4 – Climate Change
cooled climate. This, in turn, may have reduced the efficiency of the thermohaline circulation, allowing for
further cooling. Closure of the Indonesian seaway 3 to 4 million years ago may also have greatly altered
climate (Cane and Molnar 2001). The northward displacement of New Guinea switched flow through
Indonesia from warm South Pacific water to colder North Pacific water. Cooler temperatures in the Indian
Ocean may have lead to reduced rainfall over eastern Africa. Moreover, reduced atmospheric heat transport
from the tropics to high latitudes may have cooled climate and contributed to glaciation.
Mountain building also changes climate. Locally, temperatures cool due to higher elevation of the
land. Regions downwind from mountains typically are in a rain shadow and receive only sparse
precipitation, while upwind slopes receive heavy precipitation. In addition, regional and hemispheric
climate change occurs as atmospheric circulation is altered by the presence of mountains. High mountains
and plateaus block the west-to-east flow of the jet streams in middle latitudes. Meanders in the jet streams
are enhanced because the eastward flow of air is diverted around the high terrain. Seasonal monsoon
circulations are also altered due to changes in the summer heating and winter cooling of air over the
uplifted terrain. In summer, the Sun heats the high plateaus, warming the air above. The warm air becomes
less dense and rises, creating a surface low pressure. Air is drawn in from adjacent regions, flowing along
the surface in towards the low pressure cell. In winter the opposite circulation occurs. Cold, dense air sinks
over the high plateaus, creating a high surface pressure and outward surface flow.
Geologic evidence shows that major uplifting in several regions throughout the world over the
past 40 to 10 million years altered regional temperature and precipitation (Ruddiman et al. 1989, 1997;
Kutzbach et al. 1989; Ruddiman and Kutzbach 1989, 1991; Prell and Kutzbach 1992; Zhisheng et al.
2001). The Tibetan Plateau in Southeast Asia covers more than 2 million km2 with an average elevation of
about 4.5 km. The Himalayan Mountains form a narrow mountain range along the southern edge of the
plateau. The high topography of this region is a result of the collision of India with Asia some 50 to 40
million years ago and subsequent uplifting of the land. In western United States, major uplifting over the
past 15 million years formed the Sierra Nevada and Rocky Mountains and the high Great Basin and
Colorado Plateau between them. Changes in atmospheric circulation as a result of this uplifting created
greater regional contrasts in climate. Whereas the pre-uplift climate was relatively warm and wet yearround, regions became seasonally colder or warmer and drier or wetter in relation to the geography of
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mountains. In particular, the formation of the Himalayas and the Tibetan Plateau had a pronounced effect
on the climate of Southeast Asia by affecting the strength of the Indian monsoon. Before India collided
with Asia, the small size of the Asian continent and its low elevation prevented a strong land-sea
temperature contrast, which dampened the monsoon. Increased elevation following uplifting strengthened
the monsoon and brought more precipitation onto the continent. In addition, there was greater precipitation
contrast west and east of Tibet. The climate to the west became drier while the east became wetter.
4.3.2 Orbital changes
One reason for the recurring waxing and waning of glaciers is that the amount of solar radiation
received on Earth varies as the eccentricity of Earth’s orbit around the Sun, its angle of tilt, and the time of
year when closest to the Sun cyclically change over time in a process known as the Milankovitch cycles
(Hays et al. 1976; Berger 1988; Berger et al. 1993; Shackleton 2000). The eccentricity of Earth’s orbit
around the Sun varies from elliptical to nearly circular and back to elliptical with a period of about 100 000
years (Figure 4.6, top). The current orbit results in about a ±3% variation in solar radiation between when
Earth is closest and farthest from the Sun. A greater eccentricity produces greater variation and changes the
lengths of seasons. A second change is that the angle of tilt (also known as obliquity), which gives us
seasons, varies from 22° to 24.5° with a period of about 41 000 years (Figure 4.6, bottom). A smaller tilt
results in less seasonal variation between winter and summer in middle and high latitudes; winters are
milder and summers cooler. A larger tilt amplifies the seasons at high latitudes. The third effect, known as
the precession of the equinoxes, changes the distance between Earth and Sun during any given season
(Figure 4.7). The location of the equinoxes and solstices slowly shift celestial location in Earth’s orbit
around the Sun with a period of about 22 000 years. Currently, Earth is closest to the Sun in January and
farthest in July. In about 11 000 years Earth will be closer in July. More radiation than present will be
received in July and less in January. Compared with the present climate, the Southern Hemisphere winter
will be warmer and summer cooler while the Northern Hemisphere winter will be colder and summer
warmer.
Figure 4.8 shows the changes in these three orbital parameters over the past 500 000 years and for
100 000 years in the future. Eccentricity changes over the longest time scale – a 100 000-year period. There
is also evidence of a 400 000-year period with relatively low eccentricity 400 000 years ago and today and
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Chapter 4 – Climate Change
higher eccentricity in between. Earth’s orbit is currently only modestly elliptical and is becoming even less
elliptical over time. Obliquity or angle of tilt changes much faster over time. The current tilt of 23.5° is
slowly decreasing to an angle of about 22.6° in about 10 000 years. Precession changes most rapidly over
time, and over the next several thousand years Earth will be closer to the Sun in July than in January.
Figure 4.9 shows how these three orbital parameters determine the amount of solar radiation on
Earth. For most of the past 150 000 years, less solar radiation than present was received in January and
more solar radiation than present was received in July. Large positive anomalies in summer solar radiation
at about 125 000 and 10 000 years ago correspond to interglacial periods. The last ice age from about 115
000 to 18 000 years ago had less solar radiation in January than present. Indeed, the Vostok ice core reveals
a close correspondence between temperature and solar radiation in which several well-marked temperature
minima correspond to minima in annual solar radiation (Figure 4.1).
4.3.3 Greenhouse effect
The atmosphere is comprised primarily of nitrogen and oxygen, which together account for 99%
of the volume of the atmosphere (Table 4.1). Many other gases occur in trace amounts that when combined
comprise less than 1% of the volume of the atmosphere. Although they occur in minor quantities, some of
these gases play an important role in Earth’s radiation balance through the greenhouse effect. Chief among
these gases are water vapor, CO2, methane, and nitrous oxide. Ozone and halocarbons (chlorofluorocarbons
(CFCs), hydrofluorocarbons (HFCs), and other carbon compounds containing fluorine, chlorine, bromine,
or iodine) are other important greenhouse gases. These gases are poor absorbers of solar radiation, but are
strong absorbers of longwave radiation. As a result, the Sun’s radiation passes through the atmosphere and
heats the surface, but the longwave radiation emitted by the surface is absorbed by greenhouse gases in the
atmosphere (Figure 2.1). Some of this radiation is lost to space, but the vast majority (324 W m-2) is
emitted back to the surface. This re-emission of longwave radiation back to the surface is the greenhouse
effect that warms the surface. Twice as much energy (324 W m-2) is absorbed at Earth’s surface from
atmospheric longwave radiation as from solar radiation (168 W m-2). Without this, Earth would have an
effective temperature of –19 °C instead of the observed 15 °C.
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Ecological Climatology
The role of greenhouse gases in warming Earth can be quantified by a simple planetary energy
balance model in which the energy emitted by Earth annually is equal to the annual energy absorbed by
Earth. Otherwise, Earth would gain (lose) energy and climate would warm (cool). Indeed, Earth’s energy
budget shows that solar radiation absorbed by the atmosphere is balanced by the longwave radiation
emitted at the top of the atmosphere to space (Figure 2.1). Mathematically, the energy balance at the top of
the atmosphere is
S
c (1 − r ) = εσ (T + 273.15) 4
4
The left-hand side of this equation represents the solar radiation absorbed by the atmosphere and surface.
The right-hand side of this equation is the outgoing longwave radiation at the top of the atmosphere. In this
equation, Sc = 1368 W m-2, known as the solar constant, is the amount of radiation emitted by the Sun. The
2
2
division by four arises because an area of π r intercepts solar radiation but a surface area of 4π r emits
e
e
longwave radiation, where re is the radius of Earth. Hence S / 4 = 342 W m-2 is the incoming solar
c
radiation at the top of the atmosphere averaged over Earth’s surface area. The term r is the planetary albedo
– the fraction of solar radiation reflected to space. Currently, 107 W m-2 of the incoming 342 W m-2 is
reflected so that r = 0.31. The emission of longwave radiation to space is given in terms of an effective
surface temperature T (in units of degrees Celsius, which is converted to the Kelvin scale by adding
273.15), where σ = 5.67 × 10-8 W m-2 K-4. Emissivity (ε) accounts for the extent to which Earth is not a
perfect emitter of longwave radiation. Atmospheric gases absorb some of the longwave radiation emitted at
the surface. If Earth were a perfect emitter, ε = 1 and the surface temperature would be –19 °C. However,
the amount of radiation lost to space (235 W m-2) as a portion of surface emission (390 W m-2) is only 60%;
the remaining 40% is absorbed by greenhouse gases. Hence, the actual emissivity is only 0.60 and Earth’s
effective temperature is 15 °C.
Water vapor is the dominant greenhouse gas both in concentration and radiative warming. The
amount of water vapor in the atmosphere varies geographically and seasonally and can be as high as 4% of
the atmosphere, with more water vapor in the warm tropics than the colder polar regions and more water
vapor in warm seasons than cold seasons. Atmospheric water is part of a continual cycling of water among
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Chapter 4 – Climate Change
atmosphere, ocean, and land. Rates of precipitation and evaporation depend on temperature and other
climatic factors so that as climate changes the amount of water vapor in the atmosphere also changes.
Carbon dioxide is the most widely discussed greenhouse gas. Its current concentration in the
atmosphere is about 365 parts per million by volume (ppm), which is the result of geological processes,
biological processes, and human activities. Of the some 1023 g of carbon on Earth, all but a small portion is
buried in sedimentary rocks. Only about 0.04% of this carbon (40 000 × 1015 g) is in biologically active
pools near Earth’s surface, and of this the vast majority (38 000 × 1015 g) is contained in oceans. Plants and
soil contain 2060 × 1015 g C. The atmosphere has the least carbon: 750 × 1015 g C.
Over a timescale of millions of years, CO2 cycles among atmosphere, oceans, and continents in
response to the chemical weathering of rocks (Figure 4.10). Carbon dioxide is taken up by plants during
photosynthesis and released in soil as plant debris falls to the ground and decomposes. The CO2 in the soil
reacts with water to form carbonic acid, which slowly weathers and disintegrates rocks over time through a
series of chemical reactions. In carbonate rocks, carbonic acid dissolves rocks to yield one calcium and two
bicarbonate ions. Silicate rocks also dissolve to yield one calcium and two bicarbonate ions. The dissolved
products of weathering are carried by rivers to oceans. Here, marine organisms use the calcium and
bicarbonate ions to build skeletons or shells of calcium carbonate. This releases CO2 back to the
atmosphere. When these marine organisms die, calcium carbonate and other organic material are deposited
on the sea floor. Over time, they become buried deep underground. At high temperatures deep in the earth,
the calcium carbonate reacts with silicon dioxide to form calcium silicate and CO2. This CO2 is vented into
the atmosphere through volcanic eruptions and soda springs. Rates of weathering increase with warmer
temperatures and more precipitation so that the geochemical carbon cycle varies as climate changes.
Carbon fluxes involved in the geochemical carbon cycle are small (typically less than 0.5 × 1015 g
C per year), but weathering of rocks can constitute a significant uptake of atmospheric CO2 over geologic
timescales. For carbonate rocks, all the atmospheric CO2 taken up during weathering is returned to the
atmosphere. One molecule of CO2 dissolves carbonate minerals to yield two bicarbonate ions. One
bicarbonate ion is transformed by marine life into calcium carbonate and buried on the sea floor to become
rock. The other is transformed to CO2 and released back to the atmosphere. The weathering of silicate
rocks, however, requires two molecules of CO2, but only one molecule of CO2 is returned to the
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Ecological Climatology
atmosphere by marine organisms. Metamorphism of rocks deep underground and venting to the atmosphere
through volcanoes is required before the second molecule of CO2 is returned to the atmosphere. Thus,
silicate weathering results in a net loss of atmospheric CO2 in the absence of volcanoes.
Over shorter timescales of days, seasons, years, and decades, CO2 is also cycling among the
atmosphere and marine and terrestrial organisms through biological processes (Figure 4.11). On land,
carbon is stored in plant material such as leaves, twigs, branches, and roots, where about 45% of dry weight
is carbon, and in the soil as decomposing plant debris. Most of the biologically active carbon is stored in
soil. It is estimated that plants contain 560 × 1015 g C worldwide while soils hold almost three times as
much carbon. The carbon in living plant tissue is obtained from the atmosphere during plant
photosynthesis. In an average year, 120 × 1015 g C are taken up by terrestrial plants during photosynthesis.
Some of this carbon is used during metabolic processes that maintain tissues and grow new plant material.
This process, known as respiration, returns 60 × 1015 g C to the atmosphere annually on average. The
difference, 60 × 1015 g C, is the annual net productivity of terrestrial plants. Leaves, twigs, and other plant
material fall to the ground and are slowly decomposed over time by microorganisms living in the soil.
Decomposition of plant debris returns another 60 × 1015 g C to the atmosphere annually. In this
representation of the carbon cycle, terrestrial ecosystems, in the absence of human influences, are not
thought to uptake or release carbon annually on average; the amount absorbed during photosynthesis is
balanced by the amount released during respiration and decomposition.
Oceans, on the other hand, are thought to have an average net annual carbon uptake of about 2 ×
1015 g (Figure 4.11). Carbon dioxide is continually being exchanged between air and sea (Najjar 1992;
Falkowski et al. 1998; Field et al. 1998). Physical and chemical processes regulate some of this exchange.
Carbon dioxide dissolves in water as a function of its concentration in the air. As atmospheric levels
increase, the oceans respond by dissolving more CO2. Carbon dioxide is also more soluble in cold water
than warm water. As a result, for water in equilibrium with the atmosphere the concentration of dissolved
CO2 increases as temperature decreases. Increasing wind speeds increase the rate of air-sea exchange. In
addition, biological activity by phytoplankton and zooplankton – small plants and animals that live in
surface waters – removes CO2 from the atmosphere. The carbon is buried in sediments as the organisms die
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Chapter 4 – Climate Change
and settle on the ocean floor. The biological pump is an important regulator of atmospheric CO2. Without
it, atmospheric CO2 concentrations would increase to 900-1000 ppm (Sarmiento and Wofsy 1999, p. 16).
The annual carbon fluxes in the geological carbon cycle are small compared with those in the
biological carbon cycle – about one-tenth to one-ten-thousandth in magnitude. For example, rivers carry
about 0.80 × 1015 g C per year as dissolved and particulate matter to oceans. Some 0.15 × 1015 g C are
buried in ocean sediments annually. Volcanoes inject 0.02 to 0.05 × 1015 g C into the atmosphere annually.
The biological cycle, then, superimposes large rapid changes on the small, slow changes of the geological
cycle. Over the timescale of decades to centuries, changes in ecosystem productivity and ocean biology are
important natural causes of changes in atmospheric CO2, but over timescales of millions of years, the
geological carbon cycle is most important. It is the geological carbon cycle that has maintained the low
concentration of CO2 in the atmosphere through Earth’s history. In the absence of volcanic activity, rock
weathering is estimated to deplete atmospheric CO2 over a period of one million years (Schlesinger 1997,
p. 368).
Human activities have a significant impact on the global carbon cycle. The burning of oil, coal,
and wood to generate heat and electricity, the combustion of gasoline for transportation, and other
industrial processes release CO2 to the atmosphere. Over the period 1980 to 1995, the annual emission from
the combustion of fossil fuels was 5.7 × 1015 g C yr-1. In addition, clearing of forested lands for agriculture
released an additional 1.9 × 1015 g C each year. The total emission of CO2 from fossil fuel combustion and
land use practices is small compared with natural fluxes. However, although the exchanges of carbon
between atmosphere-land and atmosphere-ocean are large, the gains are nearly offset by the losses so that
the net fluxes are quite small. In the carbon cycle depicted in Figure 4.11, terrestrial ecosystems neither
store nor release carbon annually, and marine ecosystems take up about 2 × 1015 g C each year. Human
emissions are large compared with these net fluxes.
Our understanding of the global carbon cycle is incomplete. With total anthropogenic emissions of
7.6 × 1015 g C per year, with no net annual uptake or release of carbon by terrestrial ecosystems, and a net
annual uptake of 2.1 × 1015 g C yr-1 by oceans, atmospheric CO2 should increase by 5.5 × 1015 g C per year
(Figure 4.11). This number is not supported by measurements of atmospheric CO2 made at numerous
locations throughout the world, which show the actual increase is only 3.2 × 1015 g C per year. The
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Ecological Climatology
remainder, 2.3 × 1015 g C yr-1, must be removed from the atmosphere during biological activity and stored
in the oceans and terrestrial ecosystems. Although the exact fate of this ‘missing’ carbon is uncertain,
analyses of atmospheric CO2 and oxygen data suggest that a large portion (up to 2 × 1015 g C per year) is
sequestered in temperate and boreal ecosystems in the Northern Hemisphere (Tans et al. 1990; Ciais et al.
1995, 2000; Keeling et al. 1996; Rayner et al. 1999). Ecological mechanisms for this uptake include
climate variability and change, recovery of forests from disturbance, and stimulation of plant growth by
increasing CO2 in the atmosphere and increasing deposition of nitrogen on land (Houghton et al. 1998;
Lloyd 1999; Prentice et al. 2000a).
Uncertainty in the carbon budget arises because the net carbon flux in terrestrial ecosystems is the
small difference between large uptake during photosynthesis and large releases during respiration and
decomposition. None of these fluxes are known precisely, and small errors in the estimated fluxes can
result in large differences in net ecosystem uptake or release. For example, a ±5% error in our estimate of
photosynthesis, which in fact is much smaller than the detection limit given current measurement
techniques, means that terrestrial ecosystems could have a net annual uptake of 6 × 1015 g C or could have a
net annual release of 6 × 1015 g C. A ±10% error in estimated decomposition, which is harder to measure
than photosynthesis, could also result in a similar ±6 × 1015 g C uncertainty in the annual uptake or release
of carbon by terrestrial ecosystems. In fact, much of our current understanding of the terrestrial carbon
cycle comes not from direct measurements of photosynthesis, respiration, or decomposition, but rather
from measurements of CO2 in the atmosphere and the inferred terrestrial fluxes needed to match the
observations.
Methane is another important greenhouse gas that cycles among atmosphere, ocean, and terrestrial
pools. Its atmospheric concentration, like CO2, has increased since the beginning of the Industrial
Revolution (Table 4.1). Although the concentration of methane in the atmosphere is two orders of
magnitude less than that of CO2 (1.725 ppm versus 365 ppm), it is a more important greenhouse gas
because one molecule of methane causes 21 times the greenhouse warming of one molecule of CO2 (Shine
et al. 1995). The largest natural source of methane is from wetlands, where anaerobic decomposition in
waterlogged soils produces methane (Table 4.2). A little more than one-half of this emission is thought to
come from tropical wetlands, which are spatially extensive and have high emission rates due to warm
14
Chapter 4 – Climate Change
temperatures. Termites and other insects also release methane to the atmosphere as a result of anaerobic
decomposition of organic matter in their bodies. However, the greater source of methane is from human
activities, which release more than twice as much methane annually as do natural sources. Once in the
atmosphere, most methane (485 × 1012 g per year) is transformation through a series of chemical reactions
with hydroxyl radicals (OH). A much smaller amount diffuses naturally from the atmosphere into soils.
Similar to CO2, the methane budget has an imbalance when compared with the observed atmospheric
increase in methane. Our current understanding of methane production and consumption shows that
atmospheric methane should increase at a rate of 20 × 1012 g per year, but observations show it is increasing
at a faster rate (30 × 1012 g per year).
Despite its low concentration in the atmosphere, nitrous oxide is another important greenhouse
gas. Its concentration in the atmosphere has increased 16% since the Industrial Revolution (Table 4.1), and
the temperature warming caused by one molecule of nitrous oxide is 206 times that of one molecule of CO2
(Shine et al. 1995). Currently, human activities contribute about one-third of the annual emission of nitrous
oxide to the atmosphere (Table 4.3). Soils naturally emit 6 × 1012 g N annually from nitrification and
denitrification during the global nitrogen cycle. One-half of this occurs in wet tropical forests. Oceans also
emit nitrous oxide to the atmosphere as part of the marine nitrogen cycle. Human activities emit 5.7 × 1012
g N annually. Much of this emission is related to fertilizer application during agriculture. Photolysis of
nitrous oxide in the stratosphere destroys 12.3 × 1012 g N per year, producing nitrogen (N2) and nitric oxide
(NO). As with the other gases, our understanding of nitrous oxide consumption and production is
incomplete. Atmospheric measurements show nitrous oxide is increasing at a rate of 3.9 × 1012 g N per year
rather than 3.4 × 1012 g implied by the estimated budget.
The concentration of CO2 and methane in the atmosphere has varied over the past 150 000 years in
relation to temperature changes (Figure 4.1). Carbon dioxide levels were lower during the ice age (200-240
ppm) than during the current or preceding interglacials (260-280 ppm). Atmospheric methane ranged from
a low of about 0.4 ppm in the last glacial, increasing to 0.7 ppm as climate warmed and the glaciers melted.
The concentration of nitrous oxide in the atmosphere has also varied over time in parallel with temperature
changes (Flückiger et al. 1999). Low concentrations occurred during cold periods and high concentrations
occurred during warm periods. The co-occurrence in time of low atmospheric concentrations of CO2,
15
Ecological Climatology
methane, and nitrous oxide and cold temperatures is not coincidental. Low concentrations of these gases
cooled climate by reducing the greenhouse effect while high concentrations warmed climate. In addition,
the rates of biological uptake and release of these gases depend on climate. For example, the increased
amounts of methane in the atmosphere with deglaciation are likely to have resulted from increased
emissions from tropical or northern wetlands (Chappellaz et al. 1990, 1993; Brook et al. 2000).
4.3.4 Freshwater runoff and thermohaline circulation
Oceans transport vast quantities of heat poleward from the tropics. One such circulation is the
North Atlantic thermohaline circulation (Broecker et al. 1985; Broecker 1997; Manabe and Stouffer 1999),
though the Pacific Ocean is also important (Cane 1998; Pierrehumbert 2000). This density-driven ocean
circulation combines with wind-driven surface currents to carry warm shallow water from the North Pacific
Ocean across the Indian Ocean, around Africa, and into the North Atlantic (Figure 2.13). Heat carried from
tropical waters northwards into the North Atlantic Ocean is released to eastward-moving air masses, which
keeps northern European winter temperatures milder than expected from their high latitude. A weak
circulation reduces this heat transport and causes colder winters.The circulation is maintained by the
formation of ‘deep water’ in the North Atlantic, where cold, salty water sinks to the ocean bottom and
flows southward to the southern tip of Africa. Here, it joins a deep river of water that flows around
Antarctica, eventually flowing into the Indian and Pacific Oceans, where it begins its long journey back to
the Atlantic. Increases in freshwater input to the North Atlantic due to long-term changes in continental
runoff or glacial melt affect the poleward transport of heat by making the surface water less salty, so that
the water is less dense and does not sink to the ocean bottom.
Climate records reconstructed throughout the North Atlantic region show frequent large and rapid
climate changes over the past several thousand years that are related to rapid reorganization of the North
Atlantic thermohaline circulation. The Younger Dryas cold period 11 000 to 10 000 years before present
appears to have been caused by a shutdown of the thermohaline circulation as a result of diversion of runoff
from the Mississippi River drainage basin to the St. Lawrence River (Broecker et al. 1989). Prior to 11 000
years ago, runoff from melting glaciers in Canada drained into the Gulf of Mexico through the Mississippi
River (Figure 4.12). As the glaciers retreated northward, melt water was channeled into the Great Lakes
and flowed through the St. Lawrence River into the North Atlantic. The northward transport of heat in the
16
Chapter 4 – Climate Change
North Atlantic Ocean diminished and climate cooled. By about 10 000 years ago, drainage into the
Mississippi River and Gulf of Mexico re-established when the southward-advancing ice once again
dammed eastern outflow channels. Another cold period occurred some 8 000 years ago, when temperatures
cooled 4-8 °C in Greenland and 1-3 °C in northern Europe. Glaciers again advanced southward in Norway
and to lower elevation in the Austrian Alps. This cold period may have been caused by massive outflow of
freshwater from Hudson Bay into the Labrador Sea (von Grafenstein et al. 1998; Barber et al. 1999;
Renssen et al. 2001). Prior to this cold episode, remnant ice occupied Hudson Bay, creating a large ice dam
for glacial lakes draining northwards (Figure 4.12). Rapid melting of this ice cleared Hudson Strait and
allowed the glacial lakes to the south to drain northward through Hudson Strait into the Labrador Sea and
North Atlantic. The draining of these large remnant glacial lakes is estimated to have released more than
100 000 km3 of water. The Little Ice Age may also have been a result of a breakdown in the thermohaline
circulation (Broecker 2000; Keigwin and Boyle 2000). The climate effects of a reduced thermohaline
circulation are not limited to the North Atlantic regions but extend to the northern tropics, where prolonged
droughts and greater ocean productivity coincide with these large freshwater intrusions into the North
Atlantic (Street-Perrott and Perrott 1990; Lamb et al. 1995; Hughen et al. 1996).
4.3.5 Solar variability
The Sun is often described as producing a constant amount of radiation. Indeed, this is embodied
in the concept of the ‘solar constant’. This number, currently about 1368 W m-2, is the amount of radiation
received at the mean Earth-Sun distance by one square meter of surface area at the top of the atmosphere
oriented perpendicular to the Sun. In fact, however, the amount of radiation has varied from 1364 to 1368
W m-2 since the early 1600s (Figure 4.13). Variation in the intensity of solar radiation is related to the
appearance of dark ‘spots’ on the Sun’s surface. During times of maximum sunspot activity, the Sun emits
more radiation. Although sunspots are regions with colder than average temperatures, and hence emit less
radiation, they are accompanied by bright regions called faculae that are hotter than average and which emit
more radiation. The net effect is greater solar radiation emission during times of high sunspot number and
less solar radiation emission during times of few sunspots. Because they are easily observed, a long record
of sunspot occurrence exists. This record shows an irregular cycle in sunspot abundance with a period of
about 11 years. Sunspots were virtually absent during the period 1645 to 1715, during which time less solar
17
Ecological Climatology
radiation was received on Earth. This decrease in solar radiation, known as the Maunder Minimum,
corresponds to the Little Ice Age and may explain why temperatures were abnormally cold during this time
(Eddy 1976). Since the Maunder Minimum, solar radiation has generally increased, especially since the late
1800s (Figure 4.13).
4.3.6 Aerosols
In addition to gases such as nitrogen, oxygen, water vapor, CO2, methane, and nitrous oxide, the
atmosphere contains microscopic airborne particles known as aerosols (Table 4.4). Primary aerosols enter
the atmosphere directly as dust from land, salts from oceans, debris from fires, and volcanic ash. Wind
erosion from arid and semi-arid environments carries the largest single amount of aerosols into the
atmosphere (1500 × 1012 g per year). Sea salt formed from ocean spray is a similarly large source of
aerosols. Secondary aerosols are formed by chemical reactions in the atmosphere that convert emitted gases
to particles. Sulfate and organic particles are emitted naturally through biological processes on land and in
oceans. Sulfate aerosols form when dimethylsulfide ([CH3]2S) is produced in the oceans during
decomposition of phytoplankton. This is the largest natural emission of sulfur gas to atmosphere, where it
is readily converted to sulfate aerosols. Over land, organic condensates form during chemical reactions
involving the emission of non-methane hydrocarbons from terrestrial vegetation. A wide variety of
particles are produced through human activities, though the total emission (386 × 1012 g per year) is only
about one-tenth of the natural emissions. Sulfate aerosols produced through combustion of sulfurcontaining fossil fuels are particularly important. Airborne particles such as dust and sulfate aerosols
directly affect climate by absorbing or scattering solar radiation. In addition, sulfate aerosols alter climate
indirectly by influencing the number and size of cloud droplets. In this way, clouds are brightened,
reducing the sunlight reaching the surface.
The amount of aerosols in the atmosphere has changed considerably over time. For example, the
Vostok ice core reveals large changes in dust deposition over Antarctica during the past 185 000 years
(Figure 4.1). The lowest dust fluxes occurred during the current interglacial and the preceding interglacial
at about 130 000 years ago. Large dust spikes occurred about 20 000, 60 000, and between 140 000 and 180
000 years ago during glacial periods. These high dust fluxes are likely related to the expansion of deserts
18
Chapter 4 – Climate Change
(Petit et al. 1990; Jouzel et al. 1993). In particular, increased aridity and low atmospheric CO2 likely
reduced plant cover and growth, expanding the source region for dust emissions (Mahowald et al. 1999).
Volcanoes are a particularly important source of short-term climate change. Volcanoes inject dust,
debris, and gases high into the atmosphere at altitudes of 15 km to 25 km where they can remain for many
months. Strong winds at these altitudes rapidly blow the volcanic material around the world. This material
alters the radiation balance and by reflecting more solar radiation to space can cool surface temperatures
worldwide. In particular, sulfur dioxide, the primary gas emitted by volcanoes, combines with oxygen and
water to form sulfuric acid. This gas then condenses into fine sulfate aerosols that scatter some of the
incoming solar radiation back to space.
The period from 1600 to 1900 saw repeated volcanic eruptions (Figure 4.14). The 1600s were an
especially active time for volcanism. Following a period of relatively little volcanism from 1920 to 1960,
the last few decades have again seen increased volcanic activity. Two of the largest eruptions in the 1900s
were the April 1982 eruption of El Chichón in southern Mexico and the June 1991 eruption of Mount
Pinatubo in the Philippines. Major volcanic eruptions such as these cool global mean annual temperature by
about 0.2 °C for one to two years following eruption (Robock and Mao 1995; Hansen et al. 1996; Jones and
Kelly 1996; Briffa et al. 1998). Repeated, significant volcanic activity in the 1600s and 1700s, in
conjunction with decreased solar radiation as a result of few sunspots, may explain the cooling of the Little
Ice Age (Free and Robock 1999). Smaller volcanic eruptions have minimal impact on global temperatures
but large impact on regional temperatures. For example, the volcanic plume injected into the atmosphere
during the May 1980 eruption of Mount St. Helens in Washington state cooled daytime temperatures in
northwestern United States immediately after the eruption by as much as 8 °C (Robock 1981; Robock and
Mass 1982; Mass and Robock 1982). Low-level volcanic dust increased nighttime temperatures by up to 8
°C. This effect quickly diminished the next day as the debris dissipated and was transported to the east.
One of the more widespread and devastating eruptions occurred in April 1815 when Mount
Tambora in Indonesia exploded in the largest known volcanic eruption. The ash from this eruption,
estimated to be 100 km3 of debris, was carried into the upper atmosphere, where it blocked the Sun’s rays
from reaching Earth’s surface for the next 18 months. Temperatures in the following year were below
normal over much of the Northern Hemisphere (Figure 4.15, color plate). Cooler temperatures caused a
19
Ecological Climatology
particularly miserable summer in 1816 – often called the year without a summer (Hughes 1976; Stommel
and Stommel 1979; Watson 1990; Burroughs 1997). The summer of 1816 was especially cold in Europe
and New England, where it snowed as far south as Massachusetts in June. In both regions, there was
widespread crop failure, famine, and misery. The dismal weather in the European Alps is thought to have
inspired Mary Shelley to write Frankenstein during her summer travels (Lamb 1995; Watson 1990).
The June 1991 eruption of Mount Pinatubo provided an opportunity to monitor the geographic
spread of volcanic haze and its effect on climate (Hansen et al. 1992; Fiocco et al. 1996; Stenchikov et al.
1998; Kirchner et al. 1999). Prior to the eruption, there was little aerosol in the stratosphere (Figure 4.16,
color plate). By July, a large aerosol plume encircled the world in the tropics between latitudes 10° S and
30° N. In this region, the aerosol optical depth (a measure of aerosol amount) was ten times that prior to the
eruption. Most of this debris was found between 20 km and 25 km in the atmosphere. These aerosols
spread further north and south over time so that by September much of the planet was under a thick
volcanic haze. Aerosol concentrations in the stratosphere remained high for the next two years. A similar
rapid (over a three week period) circumglobal transport of volcanic haze was seen following the eruption of
El Chichón (Robock and Matson 1983).
4.4 Anthropogenic climate change
One of the more scientifically and politically debated climate changes is the warming of Earth’s
surface since the mid-1800s (Figure 4.17). Observations show Earth’s annual global mean surface
temperature warmed by about 0.6 °C between 1861 and 1997 (Hansen et al. 1999; Jones et al. 1999). Ten
of the 12 warmest years on record occurred between 1987 and 1998. This warming was not continuous
from year to year, but rather occurred in two distinct periods (1920s to the 1940s, 1970s to present)
separated by a period of little temperature change (1940s to 1970s). The instrumental record dates back to
the mid to late 1800s. Prior to that, tree rings, coral, ice cores, and other temperature-sensitive proxy data
have been used to reconstruct global temperatures (Figure 4.4). Such temperature reconstructions show that
the 1900s stand out as an exceptionally warm century with an unprecedented rate of warming (Overpeck et
al. 1997; Mann et al. 1998).
20
Chapter 4 – Climate Change
Air temperature is not the only indicator that Earth’s climate is changing (Easterling et al. 2000b).
Oceans have warmed by 0.3 °C in the top 300 m since the mid-1950s (Levitus et al. 2000). Analyses of
ground temperatures throughout the world show Earth’s surface temperature warmed by about 0.5 °C in the
1900s and that the 1900s were the warmest of the past five centuries (Pollack et al. 1998; Huang et al.
2000; Harris and Chapman 2001). In response to this warming, spring snow cover is decreasing (Groisman
et al. 1994; Brown and Goodison 1996; Frei et al. 1999; Brown 2000); Northern Hemisphere lakes and
rivers are freezing later in autumn and thawing earlier in spring (Magnuson et al. 2000); alpine glaciers are
melting (Oerlemans and Fortuin 1992; Oerlemans 1994; Warrick et al. 1996; Lowell 2000); the Greenland
ice sheet is melting (Krabill et al. 1999, 2000); permafrost is melting (Overpeck et al. 1997; Serreze et al.
2000); and the Arctic ice pack is shrinking geographically and decreasing in thickness (Johannessen et al.
1999; Rothrock et al. 1999; Vinnikov et al. 1999; Serreze et al. 2000). Sea level is rising, in part because
ice melt adds water to the oceans (Warrick et al. 1996). Diurnal temperature range (the difference between
daily maximum and minimum temperatures) has decreased worldwide because daily minimum
temperatures have warmed at a greater rate than daily maximum temperatures (Easterling et al. 1997,
2000b; Jones et al. 1999). This could be a result of increased cloud cover, which reduces daytime heating
by the Sun (Dai et al. 1999b). Between 1900 and 1988, annual precipitation increased at a rate of about 2.5
mm per decade over North America and much of the world except the tropics (Dai et al. 1997). In
response, riverflow increased across the United State over the past several decades (Lettenmaier et al.
1994; Lins and Slack 1999). Climate change is likely to be most noticed in the frequency of extreme events
(Mearns et al. 1984; Karl et al. 1999; Meehl et al. 2000). Indeed, temporal trends in precipitation show not
only an increase in the mean but also in the occurrence of extreme events (Karl et al. 1995; Dai et al. 1997;
Karl and Knight 1998; Kunkel et al. 1999a; Easterling et al. 2000a,b). Changes in droughts, floods, cold
freezes, heatwaves, and intense rainfall are the basis for integrated indices of climate change (Karl et al.
1996; Hansen et al. 1998a).
Whether these changes are a result of natural climate variability or global warming due to
increased greenhouse gases is the subject of considerable scientific research. Following recovery from the
last ice age, atmospheric CO2 concentrations remained at about 270 to 280 ppm for several thousand years
(Friedli et al. 1986; Barnola et al. 1987). Since the mid-1800s, the concentration of CO2, methane, nitrous
21
Ecological Climatology
oxide and other greenhouse gases in the atmosphere has increased (Figure 4.18). Although these gases
naturally cycle among atmosphere, land, and oceans, they have emissions that are directly attributable to
human activities such as fossil fuel combustion, land use, and agriculture.
Prior to 1850, the emission of CO2 in fossil fuel combustion was relatively minor. Since then, as
society industrialized and population increased, the annual emission rate has increased (Figure 4.19). Now,
human activities emit about 6.5 × 1015 g C each year to the atmosphere. Since 1751, over 250 × 1015 g C
were released to the atmosphere during the combustion of fossil fuels. Half of these emissions have
occurred since the mid-1970s. These emissions are largely restricted to industrialized regions of the world
(Figure 4.20, color plate). Combustion of fossil fuels accounts for most of the increase in atmospheric CO2.
Land use practices (e.g., deforestation, reforestation) cause a net release of an additional amount equal to
15% of the fossil fuel emission. However, only about one-half of the total anthropogenic emission of CO2
remains in the atmosphere (Figure 4.11).
The concentrations of methane and nitrous oxide in the atmosphere have also increased. Methane
is emitted during production and transportation of coal, natural gas, and oil. Methane is also emitted during
agricultural activities and during the decomposition of wastes in landfills. One-half of the annual
anthropogenic methane emissions (190 × 1012 g) is related to the production and use of fossil fuels, the
disposal of wastes in landfills, and sewage treatment (Table 4.2). Digestion by grazing animals releases
some 85 × 1012 g CH4 annually. Rice paddies release methane similar to natural wetlands. Forest fires
produce a small amount of methane annually due to incomplete combustion of organic material. Nitrous
oxide is emitted during agricultural and industrial activities and during combustion of fossil fuels.
Conversion of forest to cultivated lands and the application of fertilizer or manure increase nitrous oxide
emissions (Table 4.3). Combustion of fossil fuels releases relatively small amounts of nitrous oxide, but
industrial production of chemicals (primarily nylon) releases significant nitrous oxide.
The balance of evidence suggests that global mean surface temperature warmed by 0.6 °C during
the 1900s and that although natural climate variability contributed to this warming, global mean surface air
temperature will warm 1-6°C by 2100 in response to increased atmospheric concentrations of greenhouse
gases (IPCC 2001; Houghton et al. 2001). The controversy and uncertainty in estimating anthropogenic
greenhouse climate change stems from the complexity of the climate system and the numerous physical,
22
Chapter 4 – Climate Change
chemical, and biological feedbacks, both positive and negative, among atmospheric, oceanic, and terrestrial
components of the system that enhance or mitigate the expected warming (Schneider 1990, 1997;
Burroughs 1997; Houghton 1997; Karl et al. 1997; Karl and Trenberth 1999). Moreover, detecting the
expected warming signal in the instrumental temperature record is difficult because the record is short,
because of natural variability in the climate system, and because of confounding human effects that cool
climate.
Whereas the CO2 released in the burning of fossil fuels is expected to warm climate, sulfate
aerosols also released in fossil fuel combustion are thought to cool climate, both directly by reflecting more
solar radiation to space and indirectly by increasing the brightness of clouds (Charlson et al. 1991, 1992;
Kiehl and Briegleb 1993; Kiehl et al. 2000). The combustion of fossil fuels emits sulfur dioxide to the
atmosphere. Emissions are strongest in highly populated and industrialized regions of eastern North
America, Europe, and Japan (Figure 4.21, color plate). China also has large annual emissions of sulfur from
burning coal. Through a variety of chemical reactions, sulfur dioxide is converted to sulfate aerosol in the
atmosphere. Sulfate aerosols have greatest concentrations in the Northern Hemisphere in summer (Barth et
al. 2000; Rasch et al. 2000). Relatively little sulfate aerosol is found in the Southern Hemisphere or in the
Northern Hemisphere in winter. Sulfate aerosols remain close to emission sources, but plumes extend
downwind of the emission region. The seasonal cycle in the Northern Hemisphere is only weakly
influenced by seasonal variation in emissions. Rather, it is controlled by seasonal variation in atmospheric
transport and chemistry.
The climate of the past 150 years must be understood in terms of the combined effects of all
radiative forcings, which constitute the changes in energy available to the climate system (Figure 4.22).
Increases in atmospheric concentrations of CO2, methane, nitrous oxide, and halocarbons are a positive
radiative forcing that has warmed climate. Ozone is another important greenhouse gas, influencing both
solar and terrestrial radiation. Increasing amounts of ozone in the troposphere is another positive radiative
forcing that has warmed climate while loss of ozone in the stratosphere is a negative forcing that cooled
climate. The combined direct effect of aerosols (sulfates, fossil fuel soot, and biomass burning) is a
negative radiative forcing, primarily due to sulfate aerosols. The indirect effect of aerosols as a result of
23
Ecological Climatology
altered clouds is not as well known but is thought to cool climate. Increased solar irradiance over the past
150 years is a positive radiative forcing that has warmed climate.
Figure 4.23 shows Northern Hemisphere temperature and associated forcings since the early 1600s
(Mann et al. 1998). Time-dependent correlation of temperature with CO2, solar irradiance, and volcanic
aerosols show that each of these factors contributed to temperature variability over the past 400 years. Solar
irradiance was a significant forcing from the mid-1600s to early 1700s, which corresponds to the Little Ice
Age. The steady increase in solar irradiance from the early 1800s through mid-1900s corresponds to the
overall warming trend. Greenhouse gases emerged as the dominant forcing during the 1900s.
4.5 Climate feedbacks
Numerous physical, chemical, and biological processes within the atmosphere, oceans, and land
regulate Earth’s climate, and many of these processes feed back to accentuate or mitigate warming or
cooling. Two types of feedbacks exist. A positive feedback accentuates a perturbation within the system
while a negative feedback mitigates the perturbation. Water vapor is an example of a positive feedback. As
temperature increases, the amount of water vapor that air can hold increases. Evaporation increases in
response and more water is pumped into the atmosphere. Since water vapor is a powerful greenhouse gas,
any increase in water vapor in the atmosphere as a result of warming will feed back to accentuate the
warming. Clouds are another important feedback within the climate system because they increase the
planetary albedo (cooling climate) and reduce longwave radiation lost to space (warming climate). The
effect depends on the type of cloud. A low cloud reflects more solar radiation back to space with little
change in outgoing longwave radiation. An increase in low clouds may, therefore, act as a negative
feedback that mitigates warming. A thin cloud high in the atmosphere may warm climate by reducing
longwave emission to space with little effect on incoming solar radiation.
One large positive feedback is related to the amount of snow and ice covering the surface. Sea ice
reflects more solar radiation than open water. Glaciers and snow reflect more solar radiation than soil or
vegetation. By reflecting more solar radiation, sea ice, glaciers, and snow cool the surface climate because
less radiation is available to heat the surface. As climate warms, sea ice and glaciers begin to melt. This
reinforces the warming because more solar radiation is absorbed at the surface. Such a feedback can be
24
Chapter 4 – Climate Change
seen over the past 20 years, during which spring snow cover in the Northern Hemisphere retreated, albedo
increased, and temperatures warmed (Groisman et al. 1994). Conversely, expanding sea ice and glaciers as
a result of cooling feed back to accentuate the cooling. This is thought to be an important feedback
necessary for ice ages (Gallimore and Kutzbach 1995).
Feedbacks within the climate system can take a variety of forms. One feedback is related to the
growth of phytoplankton in oceans (Charlson et al. 1987). The production of dimethylsulfide, an important
precursor of sulfate aerosols in the atmosphere, is related to the growth of phytoplankton. Warmer
temperatures lead to greater marine productivity, which leads to greater emission of dimethylsulfide and
higher concentrations of sulfate aerosols in the atmosphere. These aerosols may result in increased and
brighter clouds so that more solar radiation is reflected to space and climate cools. Hence, phytoplankton
may act as a negative feedback, mitigating climate warming.
The geologic record suggests large climate feedbacks in Earth’s history. The world was frozen for
tens of million years from the poles to the equator at least twice and possibly as many as four times
between 750 million and 580 million years ago (Hoffman et al. 1998; Hoffman and Schrag 2000; Hyde et
al. 2000; Crowley et al. 2001). During these ice ages, CO2 levels in the atmosphere decreased, and Earth
cooled to –50 °C. Oceans froze to a depth of one kilometer. Ice that covered land and oceans reinforced the
cooling by reflecting more solar radiation to space. Recovery from glaciation was provided by the carbon
cycle. Volcanoes vented CO2 into the atmosphere, but this CO2 was not removed during chemical
weathering of rocks. Rock weathering was prevented because there was little liquid water in the bitter cold.
As a result, CO2 slowly accumulated in the atmosphere. Over millions of years, atmospheric CO2 increased
to 350 times that of today. A vigorous greenhouse warming then melted the ice over a period as short as a
few hundred years. Surface temperatures warmed to about 50 °C. The melting ice exposed the underlying
oceans and land, which reinforced the warming as more solar radiation was absorbed by these surfaces.
Higher rates of evaporation with the warmer temperatures pumped more moisture into the atmosphere,
which further fueled the warming. The carbon cycle again provided negative feedback, now from a
runaway greenhouse effect. More moisture in the atmosphere drove an enhanced hydrologic cycle, which
weathered rock exposed when the glaciers melted. Atmospheric CO2 concentrations decreased. How Earth
began to freeze is unknown. It is possibly related to the unusual clustering of land along the equator. The
25
Ecological Climatology
warm climate accelerated erosion, reducing atmospheric CO2 concentrations and initiating cooling through
a reduced greenhouse effect.
Many of the feedbacks in the climate system are related to physical and biological processes that
occur on land. Greater discharge of freshwater into the North Atlantic by rivers shuts down the
thermohaline circulation, triggering cooling. Melting of glaciers provides much of this freshwater, but also
influences climate through the ice albedo feedback (Clark et al. 1999). The weathering of rocks provides a
negative feedback against excessive rise in atmospheric CO2. High levels of CO2 in the atmosphere warm
temperatures, resulting in faster rates of weathering that scrub CO2 from the atmosphere. High dust
emissions during glacial times may deposit more iron in oceans, which might have stimulated
phytoplankton growth, reduced atmospheric CO2, and cooled temperatures further (Watson et al. 2000;
Maher and Dennis 2001). Terrestrial ecosystems, by altering biogeochemical cycles and energy flows, are
an important regulator of climate. Solar and longwave radiation are absorbed by leaves, stems, soil, and
other elements of the land surface. This heat is dissipated when solar radiation is reflected back to space,
objects emit longwave radiation, heat is carried away by winds, and through evaporation. The properties
that influence these energy exchanges vary among vegetation types. Grasslands reflect more solar radiation
than forests. Forests are taller and dissipate heat by winds more efficiently than short grasslands. Moist
landscapes evaporate more water and are cooler than dry landscapes. In addition, terrestrial ecosystems are
an important regulator of biogeochemical cycles related to CO2, methane, nitrous oxide, and dust
concentrations in the atmosphere.
These biogeophysical and biogeochemical processes are themselves influenced by climate,
directly in terms of their rates and indirectly because the geographic distribution of vegetation is controlled
by climate. Changes in the rate of carbon uptake during plant productivity or carbon loss during soil
decomposition as a result of climate change can feed back to accentuate or mitigate the climate change
depending on whether carbon is removed from or released to the atmosphere. Over centuries or more,
changes in the geographic distribution of vegetation alter the biogeochemical cycles, the hydrologic cycle,
and surface energy fluxes. The Arctic is thought to be particularly sensitive to climate change due to
numerous terrestrial feedbacks. Melting of sea ice, glaciers, and permafrost and release of carbon from
26
Chapter 4 – Climate Change
terrestrial ecosystems are all thought to reinforce current warming trends (Overpeck et al. 1997; Serreze et
al. 2000).
27
Ecological Climatology
4.6 Tables
Table 4.1. Composition of the atmosphere
Percent (by volume)
Gas
Chemical symbol
Current (1998)
Nitrogen
N2
78.08%
Oxygen
O2
20.95%
Argon
Ar
0.93%
Trace gases
Pre-industrial (1760)
< 1%
Water vapor
H2O
Variable (0 to 4%)
Carbon dioxide
CO2
365 ppm
278 ppm
Methane
CH4
1.725 ppm
0.700 ppm
Nitrous oxide
N2O
0.314 ppm
0.270 ppm
Note: Of the many gases that occur in trace amounts, only the four major greenhouse gases are shown. 1
ppm = 0.0001%.
28
Chapter 4 – Climate Change
Table 4.2. Annual production and consumption of methane (CH4)
Amount
(1012 g CH4 per year)
Production
Natural sources
Wetlands
115
Termites
20
Oceans and freshwater
15
Geological processes
10
Total emissions
160
Anthropogenic sources
Fossil fuels (coal mines, gas and oil industries)
100
Waste management (landfills, sewage)
90
Grazing animals (cows)
85
Rice paddies
60
Biomass burning
40
Total emissions
375
Consumption
Chemical reactions in atmosphere
485
Uptake by soils
30
Total loss
515
Net atmospheric increase
20
Observed atmospheric increase
30
Source: Data from Schlesinger (1997, p. 373).
29
Ecological Climatology
Table 4.3. Annual production and consumption of nitrous oxide (N2O)
Amount
(1012 g N per year)
Production
Natural sources
Soils
6
Oceans
4
Total emissions
10
Anthropogenic sources
Cultivated soils
3.5
Industrial production of chemicals
1.3
Biomass burning
0.5
Cattle and feed lots
0.4
Total emissions
5.7
Consumption
Chemical reactions in atmosphere
12.3
Total loss
12.3
Net atmospheric increase
3.4
Observed atmospheric increase
3.9
Source: Data from Schlesinger (1997, p. 395).
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Chapter 4 – Climate Change
Table 4.4. Annual sources of aerosols
Amount
(1012 g per year)
Sources
Natural sources
Primary aerosols
Soil dust
1500
Sea salt
1300
Organic debris
50
Volcanic dust
33
Secondary aerosols
Sulfates from natural precursors
102
Organic condensates
55
Nitrates from NOx
22
Total emissions
3062
Anthropogenic sources
Primary aerosols
Industrial particles
100
Forest fires
80
Soot
20
Secondary aerosols
Sulfates from SO2
140
Nitrates from NOx
36
Organic condensates
10
Total emissions
386
Total atmospheric sources
3448
Source: Data from Schlesinger (1997, p. 57). See also Andreae and Crutzen (1997).
31
Ecological Climatology
4.7 Figure Legends
Figure 4.1. Climate history reconstructed from the Vostok ice core over the past 250 000 years: (a)
Temperature deviation from present (Jouzel et al. 1996). (b) Annual solar radiation at latitude 60° N
(Berger 1978; Berger and Loutre 1991). This latitude is used because solar radiation at high latitudes in the
Northern Hemisphere is critical to glacier dynamics. Arrows show periods of low solar radiation. (c)
Atmospheric CO2 concentration in parts per million by volume (Barnola et al. 1987). (d) Atmospheric
methane concentration in parts per billion (Chappellaz et al. 1990). (e) Dust flux (Petit et al. 1990; Jouzel et
al. 1993). Data provided by the National Geophysical Data Center (National Oceanic and Atmospheric
Administration, Boulder, Colorado).
Figure 4.2. Geographic distribution of glaciers. Top: Present. Bottom: 18 000 years before present. Data
from Peltier (1994) and provided by the National Geophysical Data Center (National Oceanic and
Atmospheric Administration, Boulder, Colorado).
Figure 4.3. Solar radiation, atmospheric CO2, and glacier volume over the past 18 000 years. Land ice is as
a percent of the ice volume 18 000 years ago. Northern Hemisphere solar radiation (dashed lines) is shown
for summer (June through August) and winter (December through February) as a percent difference from
present values. Adapted from Kutzbach and Guetter (1986), COHMAP (1988), and Kutzbach and Webb
(1993).
Figure 4.4. Temperature of the Northern Hemisphere over the past 1000 years as a deviation from the 1961
to 1990 mean. Temperatures were reconstructed from tree rings, ice cores, corals, and historical documents.
Data from Jones et al. (1998) and provided by the National Geophysical Data Center (National Oceanic and
Atmospheric Administration, Boulder, Colorado).
Figure 4.5. Location of land and water 80 million years ago. Continental outlines show modern geography.
Data from Hay et al. (1999) and provided by the National Center for Atmospheric Research (Boulder,
Colorado).
32
Chapter 4 – Climate Change
Figure 4.6. Changes in eccentricity and obliquity over time. Top: Changes in the eccentricity of Earth’s
orbit over the course of 100 000 years. Bottom: Changes in angle of tilt over the course of 41 000 years.
Figure 4.7. Precession of the equinoxes. Diagrams show changes in the celestial location of the solstices
and equinoxes over the course of 22 000 years. The eccentricity of Earth’s orbit is highly exaggerated.
Figure 4.8. Periodic changes in the geometry of Earth’s orbit around the Sun over the past 500 000 years
and for 100 000 years into the future. Top: Eccentricity. Middle: Obliquity. Bottom: Precession. Data from
Berger (1978) and Berger and Loutre (1991) and provided by the National Geophysical Data Center
(National Oceanic and Atmospheric Administration, Boulder, Colorado).
Figure 4.9. Solar radiation over the past 150 000 years. Solar radiation is the difference from present as a
function of latitude (vertical axis) and time (horizontal axis). Latitudes are positive in the Northern
Hemisphere and negative in the Southern Hemisphere. Negative deviations (lower solar radiation than
present) are shaded. Top: January. Bottom: July. Data from Berger (1978) and Berger and Loutre (1991)
and provided by the National Geophysical Data Center (National Oceanic and Atmospheric Administration,
Boulder, Colorado).
Figure 4.10. Geologic carbon cycle arising from the weathering of rocks, sedimentation, and volcanoes.
Adapted from Berner and Lasaga (1989) and Schlesinger (1997, p. 9). Annual fluxes are from Schlesinger
(1997, p. 368).
Figure 4.11. Global carbon cycle for the period 1980 to 1995. Geologic fluxes for rivers, volcanoes, and
ocean burial are included for comparison. Adapted from Schlesinger (1997, p. 359) with updated land use
(Houghton 1999, 2000a). Annual fluxes for ocean burial and volcanoes are from Schlesinger (1997, p.
368). See also Schimel (1995), Houghton (2000a), and Wigley and Schimel (2000).
Figure 4.12. Diversion of freshwater runoff during deglaciation. Top: 11 000 to 10 000 years before
present. The shaded region shows the approximate location of the Laurentide ice sheet. Bottom: 8 400 to 8
000 years before present. The combined shaded region shows the approximate location of ice 8 900 years
ago. The two darkly shaded regions show the ice extent 8 200 years ago.
33
Ecological Climatology
Figure 4.13. Solar irradiance from 1610 to 1994 reconstructed from observations of the number of
sunspots. Data from Lean et al. (1995) and provided by the National Geophysical Data Center (National
Oceanic and Atmospheric Administration, Boulder, Colorado).
Figure 4.14. Volcanic activity from 1610 to 1995. Data from Mann et al. (1998). See also Bradley and
Jones (1992b). Robock and Free (1995) review indices of volcanic activity.
Figure 4.15. Effect of the Tambora volcanic eruption on temperatures in 1816. Temperatures were
reconstructed from tree-ring, coral, ice-melt, and instrumental records and are the anomaly from the 1902
to 1980 reference period. Data from Mann et al. (1998) and provided by the National Geophysical Data
Center (National Oceanic and Atmospheric Administration, Boulder, Colorado).
Figure 4.16. Injection of aerosols in the stratosphere by the eruption of Mount Pinatubo in June 1991. The
three right panels show stratospheric optical depth in May 1991 prior to eruption and July and September
1991 following eruption. High optical depths indicate high amounts of aerosols. The left panels show the
same three months in 1990, one year prior to eruption. July 1991 and September 1991 optical depths are ten
times that of other months. Images provided by the NASA Goddard Institute for Space Studies (New York
City, New York). See also McCormick and Veiga (1992).
Figure 4.17. Annual global mean surface temperature for 1856 to 1998 based on station, ship, and buoy
measurements. Temperatures have been adjusted for changes over time in instrumentation, measurement
techniques, exposure of instruments, station location, time of observation, and urbanization. The left axis
shows the anomaly from the 1961 to 1990 mean. The right axis shows the actual temperature. Data from
Jones et al. (1999) and provided by the Climatic Research Unit (University of East Anglia, Norwich).
Hansen et al. (1999) has another dataset that shows similar temporal trends
Figure 4.18. Concentration of greenhouse gases in the atmosphere from 1850 to 1998. Top: Carbon
dioxide. Middle: Methane (CH4). Bottom: Nitrous oxide (N2O). Data from Hansen et al. (1998b) and
provided by the NASA Goddard Institute for Space Studies (New York City, New York).
34
Chapter 4 – Climate Change
Figure 4.19. Annual CO2 emission from fossil fuel combustion from 1751 to 1996. Units are 1015 g C per
year. The left axis shows annual emission. The right axis is the cumulative emission. Data from Andres et
al. (1996, 1999) and provided by the Carbon Dioxide Information Analysis Center (Oak Ridge National
Laboratory, Oak Ridge, Tennessee).
Figure 4.20. Geographic distribution of CO2 emission for 1990 in units of 1012 g C per year. Data from
Andres et al. (1996, 1999) and provided by the Carbon Dioxide Information Analysis Center (Oak Ridge
National Laboratory, Oak Ridge, Tennessee).
Figure 4.21. Geographic distribution of anthropogenic sulfur emission for 1985 in units of 106 g per year.
Data from Benkovitz et al. (1996) and provided by the National Center for Atmospheric Research
(Boulder, Colorado).
Figure 4.22. Global mean radiative forcing in 2000 relative to 1750. Vertical bars indicate estimates for
various forcings. Vertical lines indicate a range of possible estimates for each forcing. Adapted from IPCC
(2001, p. 8).
Figure 4.23. Northern Hemisphere temperature and associated climate forcings from 1610 to 1995.
Northern Hemisphere temperature was reconstructed from tree-ring, coral, ice-melt, and instrumental
records as an anomaly from the 1902 to 1980 reference period. Data from Mann et al. (1998). See also
Jones et al. (1999), Hansen et al. (1998b, 1999), and Tett et al. (1999).
35