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5
Climate-System Components
The climate attributes that influence society, as noted earlier,
are themselves influenced by a broad range of physical and
biogeochemical processes, or components (including forcings) of our
climate system. Therefore, to improve our understanding of how
changes in these attributes manifest themselves over
decade-to-century time scales, we must address the issues involving
those components that will most efficiently advance this
understanding.
While the existence of climate patterns offers hope that some
fraction of the variability in the climate attributes may be
related to the state of these patterns, ultimately we must
understand the physics that control both the evolution of the
climate system and the patterns themselves. A relationship between
climate patterns and climate attributes may afford us some
statistical forecasting capabilities, but only of configurations or
types of changes already documented. Forecasting future variations
demands that we understand the physical and biogeochemical
interactions controlling climate response and feedbacks, and
identify the slow components of the system in which predictability
resides.
This chapter briefly describes our current understanding of how
physics and biogeochemistry influence climate, particularly the six
climate attributes outlined in Chapter 2, and presents the primary
issues that must be resolved to advance most expeditiously and
cost-effectively our understanding of climate change and
variability on dec-cen time scales. The six sections of this
chapter present the components and forcings of the climate system
in discipline-based discussions. This division is somewhat
arbitrary, since dec-cen-scale change and variability in the
atmosphere involve considerable couplings with and feedbacks from
the oceans, land, and cryosphere. Consequently, the study of
dec-cen change and variability entails multi- and interdisciplinary
issues, and highly coupled systems. Past study of climate and its
components has generally proceeded along disciplinary boundaries,
however, and the funding sources for such study have been similarly
partitioned. Much as we would have liked to have organized this
chapter into the new cross-disciplinary structures that will
ultimately be needed for future advances in dec-cen climate
research, it proved quite difficult to determine an ideal, or even
acceptable, cross-disciplinary structure that would conveniently
present the multitude of issues, both disciplinary and
cross-disciplinary, in a logical progression. We have chosen
instead to indicate by cross-referencing the relationships that may
guide future cross-disciplinary organizational structures.
This chapter begins with an overview of the atmospheric
composition and radiative forcing, which is fundamental to
externally forced (natural and anthropogenic) variability and
change. External forcing of the climate system, while not properly
a component of climate, is included here. Because this document
articulates a plan for addressing the science of dec-cen climate
change and variability, external forcing must be included for
completeness, and to provide the necessary foundation for
subsequent discussion in the report. Given the thoroughness of the
topic's coverage in the IPCC assessment process, and the
accessibility of the IPCC reports, we do not attempt to replicate
that review. Rather, we draw from it and build on it in order to
provide an overview of the atmospheric composition and radiative
forcing most relevant to dec-cen climate issues.
The remaining sections of this chapter focus on five distinct
components of the climate system. The first two, which are closely
related, involve two aspects of the atmosphere: atmospheric
circulation and the hydrologic cycle. (Of course, the latter
section's scope involves more than just the atmosphere, since it
discusses the storage of water and its movement through the
atmosphere and boundaries.) These two sections are followed by the
three atmospheric boundary components from which most internal
dec-cen variability originates: the oceans, the cryosphere, and
land and vegetation. Interdisciplinary aspects of the components'
interactions are presented throughout the sections when
appropriate, and several of the broader crosscutting issues that
defy traditional disciplinary categorization are presented in
Chapter 6.
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Representative terms from entire chapter:
atmospheric circulation
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In each of the six sections of this chapter, the discussion is
partitioned into subsections dealing with the influence of the
particular climate-system component on climate attributes, the
evidence of variability and change of that component on dec-cen
time scales, and the mechanisms through which that component
operates within the climate system. At the end of each section
there is a discussion of the principal outstanding issues
associated with that climate-system component, as well as an
overview of some of the key observational and modeling priorities
that will help resolve the outstanding issues. The discussion of
the requisite observational and modeling strategies is not intended
to be comprehensive; rather, it provides a broad perspective on the
types of research initiatives that are most likely to be
productive.
Finally, we wish to emphasize that this chapter deals with all
the components of the climate system that influence dec-cen
variability, whether that variability be natural, anthropogenically
induced, or anthropogenically modified natural.
Atmospheric Composition and Radiative
Forcing
Changes in solar output—either in terms of total radiative
flux (the solar constant), or in terms of the spectral distribution
of this radiation—will directly influence the radiative
environment and energy budget at the Earth's surface, the response
of the climate system, and the response of many life forms.
Moreover, changes in the atmospheric concentration of a number of
trace constituents directly influence the transfer of radiative
energy throughout the atmospheric column, and therefore the energy
balance in the atmosphere, including the temperature at the Earth's
surface. Such direct climate influences are modified by myriad
feedbacks that indirectly affect surface temperatures and radiative
fluxes, the hydrologic cycle, storm frequency and intensity, sea
level, and ecosystem structure and functioning. Increasing the
skill with which such feedbacks can be quantified is a principal
challenge for earth system science over the next decade.
The primary reason for the current widespread concern about
global climate change is that human activities are increasing the
greenhouse effect of the atmosphere and the tropospheric aerosol
burden, and weakening the stratospheric ozone shield against
ultraviolet radiation. Greenhouse gases (e.g., H2O, CO2,
CH4, N2O, chlorofluorocarbons, and O3 in the troposphere) warm the Earth' s
surface by trapping a portion of the outgoing longwave-radiation
flux. Atmospheric aerosols tend to cause surface cooling by
scattering solar radiation back into space (although they can
produce the opposite effect if they consist of very dark material
or if they are over a bright surface such as snow or ice), and they
exert indirect effects by providing nucleation sites for the
formation of cloud droplets. The net influence of the myriad
feedbacks responding to changes in atmospheric gas and aerosol
content has yet to be determined. Better understanding of these
climatic influences will be fundamental to our ability to predict
the nature and magnitude of the climate' s response to
anthropogenic change in any of the forcing factors.
Radiative forcing is affected not only by anthropogenic changes,
but also by natural variations in the sun's output and by the input
and distribution of volcanic aerosols. Largely unpredictable, these
elements exert measurable influence over the Earth's radiative
budget and atmospheric chemical interactions, and account for some
of the natural dec-cen variability in the Earth's climate. Solar
output, volcanic aerosol contributions, and atmospheric gases and
aerosols thus represent the main forcings, natural and
anthropogenic, of the climate system. In this respect, they are
distinct from the components of the climate system discussed in the
other sections of this chapter, and changes in them will drive
responses in those other components. Ultimately we need to be able
to differentiate climate variations driven by changes in the
forcings (internal or external) from variations that are the
expression of internal or coupled modes of variability, which will
occur even when forcing is steady. Our efforts to understand the
behavior of climate variations may be furthered by the fact that
the forcings and responses may vary with latitude or regional
characteristics, possibly relating specific forcings to specific
responses or climatic fingerprints. For example, the stratospheric
warming by volcanic aerosols in the Northern Hemisphere winter is
greater in low latitudes than in high latitudes (Labitzke and
Naujokat, 1983; Labitzke and McCormick, 1992). The differential
heating produces a larger pole-to-equator temperature gradient,
which in turn increases the zonal winds and enhances the
stratospheric polar vortex. The stronger polar vortex may affect
the vertically propagating tropospheric planetary waves, and so
modify the tropospheric circulation and alter surface air
temperature (Mao and Robock, 1998). Thus, radiative influences
associated with aerosols may differ from those driven by other
types of radiative forcing in the high latitudes.
Influence on Attributes
The solar radiation striking the Earth, however it may be
modified by the atmosphere's components, fundamentally mediates the
Earth's energy budget and climate through a complex array of
feedbacks. In the process, it influences all of the climate
attributes discussed in Chapter 2. These feedbacks include changing
the atmospheric concentration of water vapor, itself the major
greenhouse gas; changing cloudiness; changing the surface albedo
due to changes in snow, ice, and vegetative cover; changing source
and sink rates for carbon dioxide, methane, and nitrous oxide;
changing the formation rates for tropospheric ozone and aerosols;
and changing the transport and storage of heat in the oceans. Each
of these feedbacks further influences the surface temperature and
radiative fields, which in turn alter the evaporation of water
from, and precipitation onto, land and water
Page 50
surfaces, as well as the water balance of glaciers, ice caps,
and snow fields. Soil moisture and runoff are affected, influencing
the water quality and quantity of surface waters and the salinity
of surface layers of the ocean. Sea level responds to the heat
content of the oceans and the distribution of heat in the oceans,
as well as reflecting the proportion of the Earth's total water
mass that resides in the oceans. Changes in the radiation budget
also affect ocean transport and storage of heat and carbon, further
modifying surface temperatures and the hydrologic cycle. Changes in
energy and water fluxes may also alter the pattern or strength of
pressure systems in the atmosphere, thereby modifying the tracks,
intensity, and frequency of storms.
Ecosystems are influenced by changes in radiation through a
variety of related processes and reactions. For example, ozone is
important to ecosystems and society, in part, because it filters
UV-B radiation, as mentioned in Chapter 2. Ozone depletion
increases the surface flux of UV-B, which increases health and
ecosystem risks. Increased UV-B has been explicitly linked to
damage to marine phytoplankton (Smith, 1995), which form the base
of the marine trophic system and organic carbon cycle. Industrial
and natural aerosols in the lower troposphere reduce the quality of
the air we breathe, increasing risks to human health and to
ecosystems. Changes in atmospheric carbon dioxide directly
influence vegetation through both fertilization and changes in
response to water stress. A number of chemical feedbacks associated
with changes in aerosol and ozone levels can affect ecosystems. For
example, tropospheric ozone controls the oxidizing capacity of the
troposphere and its ability to remove other pollutants. It has also
been implicated in reduced crop growth (see, e.g., Reich and
Amundson, 1985).
Evidence of Decade-to-Century-Scale
Variability and Change
External forcings of the climate system, in a number of cases,
vary on dec-cen time scales. Some of these forcings are the result
of human activities (e.g., emissions of chlorofluorocarbons), some
are natural in origin (e.g., solar variability), and some are both
anthropogenic and natural (e.g., aerosols). The primary types of
radiative forcing exhibiting dec-cen variability are outlined
below; each of these three classes is represented.
Greenhouse Gases
Carbon dioxide is the most important of the greenhouse gases
emitted as a result of our activities. Not only is it responsible
for a little over half of the current direct anthropogenic
greenhouse forcing (IPCC, 1995) but its long atmospheric residence
time assures that any enhancement of atmospheric concentration will
persist for many centuries. Methane, which is the second greatest
contributor to direct anthropogenic greenhouse forcing, is
characterized by shorter residence times, but more rapid growth in
atmospheric concentration than CO2
(IPCC, 1995). The increases in atmospheric carbon dioxide and
methane over the last thousand years, as measured from ice cores
and directly in the atmosphere, are depicted in Figure 5-1. The
relative constancy of both gases until the turn of the twentieth
century indicates that their natural variability in the atmosphere
has been relatively small over the last millennium. During the last
glacial maximum (about 18,000 BP) CO2 and CH4
were respectively about 70 percent and 45 percent of the more
recent pre-industrial levels (Barnola et al., 1987; Jouzel et al.,
1993; Nakazawa et al., 1993; Chappellaz et al., 1993a). Extensive
analyses of sources and sinks for both of these gases (e.g., Wigley
and Schimel, 1994; IPCC, 1995), leave no doubt that their steep
rises during the latter part of this century, coinciding with the
human population explosion, is the result of human activities. The
rate of CO2 emissions from
fossil-fuel burning has increased approximately 250 percent in the
past 30 years (Figure 5-2, upper curve). Although the net global
CO2 uptake rate exhibits substantial
interannual variability in response to climatic variations (Figure
5-2, lower curve), it has generally increased as the concentration
of atmospheric CO2 has risen (Figure
5-1, solid curve). Atmospheric methane's rate of growth varies
substantially from year to year, but that rate has generally been
decreasing over the last two decades (Figure 5-3), for reasons that
are not entirely clear.
Evidence from ice cores indicates a strong coupling between
global surface temperature and the concentration of atmospheric
methane since at least 40,000 BP (Chappellaz et al., 1993a;
Severinghaus et al., 1998). It is believed that
Figure 5-1
Atmospheric carbon dioxide and methane during the last 1,000 years.
CO2 (solid curve) refers to the vertical scale on the left; CH4 (dashed
curve) refers to the scale on the right. The CO2 curve is based on long
-term CO2 data from Etheridge et al. (1996) and modem CO2 data from
Conway et al. (1994). The CH4 curve is based on long-term CH4 data
from Blunier et al. (1993) and Nakazawa et al. (1993), and more recent
CH4 data from Dlugokencky et al. (1994) and Etheridge et al. (1992). (Figure
courtesy of P. Tans, NOAA/CMDL.)
Page 51
Figure 5-2
The upper curve represents the rate of CO2 emissions from fossil-fuel burning
(Marland et al., 1994). The lower curve represents the net global uptake rate of
CO2 by the oceans and the terrestrial biosphere. This uptake rate was derived
using the assumption that the Mauna Loa CO2 record is representative of the
atmosphere as a whole. The difference between the lower and upper curves is
the rate of atmospheric CO2 increase (corrected for the seasonal cycle). (Figure
courtesy of P. Tans, NOAA/CMDL.)
Figure 5-3
The rate of increase of atmospheric methane over the last 150 years.
Based on data from Etheridge et al. (1996) and Dlugokencky et
al. (1994). (Upper panel courtesy of P. Tans, NOAA/CMDL; lower courtesy
of E. Dlugokensky, NOAA/ CMDL.)
the changes in methane may have been responding, at least in
part, to changes in soil moisture and wetland extent (which
partially control methane emissions), driven by re-organizations of
the climate system. Although the precise nature of the mechanisms
that have caused temperature and methane to co-vary in the past are
somewhat uncertain, these paleorecords indicate the possibility
that temperature and methane may also co-vary in response to future
climate changes.
Changes in tropospheric ozone, a third greenhouse gas, are not
well documented. We have a limited number of discontinuous surface
records that indicate tropospheric ozone may have doubled since the
1950s or at least since the nineteenth century (Figure 5-4). The
data on free tropospheric ozone that are available from selected
sites since 1970 show no consistent trends, however.
Stratospheric Ozone
One of the best-known changes in atmospheric composition
observed over the last several decades is the dramatic
Figure 5-4
Measurements of surface ozone from different locations in Europe showing
increasing concentrations from before the end of the 1950s (circles) to 1990-1991
(triangles) during August and September, as a function of altitude. (From
Staehelin et al., 1994; reprinted with permission of Elsevier Science.)
Page 52
decrease in stratospheric ozone over Antarctica. The comparison
of the annual cycle in column ozone between Arctic (Resolute) and
Antarctic (Halley Bay) locations is shown in Figure 5-5.
Measurements in Antarctica between 1956 and 1965 showed a
difference of 200 Dobson units (one DU equals 10-2 mm·atmosphere of column ozone)
between Arctic and Antarctic springtime values (Dobson, 1966). This
difference is due to the differing meteorologies of the two
regions, particularly the isolation of the Antarctic vortex much
later in the spring. More recent measurements, made at Halley Bay
by the British Antarctic Survey, demonstrate an additional 200 DU
deficit, commonly called the Antarctic ''ozone hole.'' This
dramatic decrease in Antarctic stratospheric ozone has been a
regular feature since 1989; it represents a major decadal change in
our planet. In the past few boreal springs, significant decreases
in Arctic ozone have been noted as well (NOAA, 1995, 1996, 1997).
Although these Arctic levels have generally not been much lower
than typical tropical values (~250 DU), they do constitute a
significant anomaly for that region.
Evidence of stratospheric ozone depletion over dec-cen time
scales is also indicated in other records. Figure 5-6
Figure 5-5
Annual cycle of column ozone from the Arctic (Resolute) and Antarctic (Halley Bay)
for 1956-1965 and 1994. The top and middle curves are smoothed representations of
the Arctic and Antarctic data, respectively. The points towards the bottom of the figure
illustrate the magnitude of the ozone "hole" at Halley Bay in 1994. Units are Dobson
units (DU). The Southern Hemisphere time scale (bottom axis) has been shifted by 6
months to line up with that of the Northern Hemisphere (top axis). (Figure courtesy of
R. Stolarski. Halley Bay data from J.D. Shanklin of the British Antarctic Survey.
After Dobson, 1966; reprinted with permission of the Royal Meteorological Society.)
Figure 5-6
Anomalies from 1926-1996 of total ozone measured over Arosa, Switzerland,
relative the 1926-1969 mean of 339 DU. The dotted line shows the 5-year moving
average and the solid line shows the annual mean. The downward trend since
1978 averaged 1.12% per decade. (From Staehelin et al., 1998; reprinted with permission
of the American Geophysical Union.)
shows annually-averaged deviations in ozone over Arosa,
Switzerland, since the 1930s; a decline over the last two decades
is apparent. Losses of total ozone (i.e., the mass of ozone
vertically integrated through the entire atmosphere) have been
greatest in the higher latitudes, with very little change in the
tropics (WMO, 1995).
The eruption of Mt. Pinatubo in June of 1991 provided a nearly
hundred-fold increase in the surface area available for
heterogeneous chemical processing in the stratosphere. Observations
following the eruption indicated significant reductions in NO2 (Johnston et al., 1992; Koike et al.,
1994) along with increased concentrations of HNO3 (Rinsland et al., 1994). These changes
suggest that reactive nitrogen species (e.g., NO2) were repartitioned into less reactive
forms, which in turn helped to temporarily enhance the levels of
active, ozone-depleting chlorine radicals (e.g., ClO) relative to
those of the more inert chlorine reservoirs (e.g., HCl). Although
the predicted massive ozone loss in the volcanic cloud did not
occur (Prather, 1992), observations at that time showed evidence of
greater ozone depletion than that expected in response to the
continued growth in stratospheric chlorine abundance (Hofmann et
al., 1994; Komhyr et al., 1993). A 6-8 percent loss of ozone in the
tropics immediately after the eruption is more likely to have been
associated with the vertical lofting that accompanied the strong
stratospheric heating by the aerosols (Kinne et al., 1992).
Overall, observations by the Total Ozone Mapping Spectrometer
(TOMS) showed an additional global ozone deficit of about 2-3
percent by mid- 1992 that might be attributed to Mt. Pinatubo
(Gleason et al., 1993). The atmosphere had mostly returned to
normal a couple of years after the volcanic perturbation, and it is
difficult to determine how much of
Page 53
the ozone depletion in these years should be attributed to
chlorine increases and how much to volcanic aerosols.
The decrease in polar stratospheric-ozone concentrations that
has been documented over the past 30 years is strongly related, at
least in part, to the increase in atmospheric chlorine (Figures 5-7
and 5-8). While levels of chlorinated compounds in the atmosphere
are still high, their growth rates tend to be decreasing, and in
some cases are negative.
Aerosols
Volcanic aerosols can have a significant influence on the
radiative balance (defined as the difference between absorbed solar
radiation and outgoing longwave radiation) of both the stratosphere
and the Earth's climate system. For instance, Labitzke et al.
(1983) showed that the aerosols produced by the eruption of El
Chichón, which achieved a peak concentration at 24 km
between about 10ºS and 30ºN (for
Figure 5-7
Atmospheric trends of chlorinated compounds controlled under the Montreal
Protocol from 1977 to 1995. The mixing ratios from surface measurements are
reported as monthly means in parts per trillion (ppt) in dry air. CFC-11 and CFC-
12 data are updated from Elkins et al. (1993); methyl chloroform (CH3CCl3),
carbon tetrachloride (CCl4), and CFC-113 (CCl3F-CClF3) data are updated from Montzka
et al. (1996). (Figure courtesy of NOAA/CMDL.)
Figure 5-8
Stratospheric trend of HCl from 1991 to 1995. HALOE is the Halogen Occultation
Experiment. (From Russell et al., 1996; reprinted with permission of Macmillan Magazines, Ltd.)
the first six months), warmed this region of the atmosphere by a
few degrees. Following the eruption of Mt. Pinatubo, substantial
changes to the planetary albedo were observed (Minnis et al.,
1993). In addition, substantial heating in the tropical
stratosphere was observed immediately after Pinatubo's eruption.
This heating was sufficient to cause tropical stratospheric
temperatures at 30 hPa to increase as much as three standard
deviations above the 26-year mean (Labitzke and McCormick, 1992).
On the other hand, global surface temperature was also observed to
decrease in the months following the Pinatubo eruption as a result
of the increased planetary albedo (see, e.g., Dutton and Christy,
1992), and the temperature remained suppressed through 1993, as
predicted (Hansen et al., 1996).
In addition to these radiative effects of volcanic aerosol,
recent work by Solomon et al. (1996) demonstrates that the observed
aerosol variability can influence the modeled ozone trends. Periods
of peak aerosol loading appear to correlate better with additional
ozone depletion than with a trend fitted to the dominant driving
force in ozone depletion, stratospheric chlorine levels. It is
difficult to interpret this trend in ozone over the 15 years of
TOMS data without including the concurrent variations in
stratospheric aerosols.
Since the late 1970s, near-global monitoring of stratospheric
aerosol distribution has been carried out by in situ (Wilson et
al., 1992), ground-based (e.g., Osborn et al., 1995), and
satellite-based instrumentation (SAM II and SAGE measurements; see,
e.g., Thomason et al., 1997b). Over this period, the primary source
of stratospheric aerosol variability has been periodic injections
of aerosol, or of gaseous aerosol precursors such as SO2, by volcanic eruptions. In general,
stratospheric aerosols are produced in situ by processes that
include the photochemical transformation of gaseous SO2 into H2SO4
aerosol. For example, the composite SAM II/SAGE/SAGE II record of
stratospheric-aerosol optical depth shows large effects from the
eruptions of El
Page 54
Chichón in 1982 and Mount Pinatubo in 1991, as well as
effects of several smaller eruptions such as Mount St. Helens in
1980, Nevada del Ruiz in 1985, and Kelut in 1990 (McCormick et al.,
1993). The eruption of Mt. Pinatubo may have caused the largest
perturbation to stratospheric aerosol loading of any eruption since
Krakatau in 1883.
The meridional distribution and the residence time of volcanic
aerosols are strongly dictated by the latitude of the eruption, the
altitude reached by the eruption plume, time of year, and phase of
the quasi-biennial oscillation at the time of the initial aerosol
injection (Trepte et al., 1993). As a result, any reconstructions
of stratospheric aerosol loading resulting from volcanic eruptions
are subject to significant uncertainties if they are extended
backwards before the start of global measurements in 1978 (see,
e.g., Sato et al., 1993), since many of the aforementioned
parameters are poorly known. Additional indications of aerosol
concentrations over longer time periods can be obtained from the
longer records of Sato et al. (1993) (Figure 5-9) and from ice-core
analyses (e.g., those of Zielinski et al., 1994). It has been
suggested that the non-volcanic background stratospheric aerosol
mass has increased by 5 percent annually over the period from 1978
to 1989 (Hofmann, 1990), though SAGE-based evaluations tend to
argue against this increase (Thomason et al., 1997a).
Solar Radiation
The sun is the driving force of climate; even small variations
in the amount of energy that the Earth receives can apparently have
significant impact (for instance, see the last section in this
chapter for a discussion of the role of solar
Figure 5-9
Stratospheric aerosols as a function of time. For the period 1883-1990, aerosol
optical depths are estimated from optical extinction data, whose quality increases
with time over that period. For the period 1850-1882, aerosol optical depths are
more crudely estimated from volcanological evidence for the volume of ejecta from
major known volcanoes. (From Sato et al., 1993; reprinted with permission of the
American Geophysical Union.)
Figure 5-10
Total solar irradiance from 1975-1995 measured by the Active Cavity Radiometer Irradiance
Monitor/Solar Maximum Mission and Upper Atmosphere Research Satellite (ACRIM/SMM
and ACRIM/UARS). The dotted line is a model of the total irradiance variability obtained
from a parameterization of the influence of sunspot darkening and facular brightening, which
are recognized as the two primary mechanisms of irradiance variability during the 11-year
solar cycle. (After Lean et al., 1995; reprinted with permission of the American Geophysical Union.)
variations in the waxing and waning of the great ice ages). By
comparison, a doubling of CO2 in the
atmosphere would generate a radiative forcing equivalent to a 1.8
percent increase in solar irradiance. Best estimates derived from
solar proxies suggest dec-cen changes in solar irradiance on the
order of 0.25 percent over the past 400 years (Nesme-Ribes et al.,
1993; Lean et al., 1995; Hoyt and Schatten, 1993). However, the
only direct record of solar-irradiance variability we have covers
only the last one and one-half solar cycles; as Figure 5-10
illustrates, the recent range of variations is about 0.1 percent.
The total-irradiance record shown in Figure 5-10 is based on
satellite observations, and involves a modeled reconstruction over
this period. (It has long been known from indirect measures of
solar radiation that the variability of the sun's UV radiation has
an 11-year period.)
Although UV radiation constitutes only a small portion of the
total solar irradiance, it is more variable by at least an order of
magnitude than the visible-radiation portion, and therefore
contributes significantly to total solar variability. This UV
variability has special relevance to chemical interactions in the
upper atmosphere, where the temperature structure depends partly on
the absorption of UV radiation by O3, O2,
N2, and other gases. This
relationship was highlighted by Hood and McCormack (1992), who
showed a strong correlation between O3 and UV radiation on the 11-year solar
cycle.
Additional records of the sun's activity are derived from
observations, beginning early in the seventeenth century, of the
occurrence of dark spots on the face of the sun; they are not a
direct measurement of solar irradiance, but over the period for
which we have direct irradiance measures, high sunspot activity
correlates strongly with increased irradiance.
Page 55
Sunspots are associated with bright faculae that surround the
dark spot. Although the spots themselves are areas of decreased
irradiance, the faculae are longer-lived and more areally
extensive, leading to an overall increase in total irradiance at
times of sunspot maxima. Sunspot observations indicate that solar
activity has varied on an 11-year cycle for the past 300 years
(since about AD 1690). But longer-term variation has been inferred
from observations of sunspots made over the last several centuries.
For instance, during the Maunder Minimum (1650-1690) no sunspots
were observed (Lean, 1991). Longer- and shorter-period variance
also occurs. The sunspot record exhibits an 80-100 year period
known as the Gleissberg cycle, and the apparent alternation of
stronger and weaker 11-year cycles produces a concentration of
variance with a 22-year period. Over shorter periods, the sun
exhibits variations associated with its rotation (which has a
27-day period), and monthly and yearly variations are seen within
the envelope of the 11-year activity cycle.
Indirect indicators of solar activity, such as sunspots and the
abundance of cosmogenic nuclides (e.g.,14C and10Be), have considerably longer records
than direct observations. Figure 5-11 shows two different indices
that are commonly used to infer some measure of solar activity
(e.g., solar wind), and are known to correlate with irradiance over
the last solar cycle (see, e.g., Wilson and Hudson, 1991). These
longer, proxy records show distinct long-term shifts in solar
activity over the past several centuries; for example, such shifts
can be seen in the record of10Be
found in ice cores. Production of10Be by galactic cosmic-ray particles in
the Earth's atmosphere is modulated by the solar wind; this
long-lived radionuclide is removed from the atmosphere by
precipitation and preserved in ice cores. Ice-care10Be abundances were significantly
higher in the fifteenth and late seventeenth centu-
Figure 5-11
Time series of sunspot number and
10Be in ice cores, which are both known to
correlate with irradiance over the past few solar cycles. (After Beer et al., 1988; reprinted
with permission of Macmillan Magazines, Ltd.)
ries, implying that the solar wind was much weaker then than it
is today. The relationship between solar wind and solar irradiance
has been calibrated for the last two solar cycles; the
extrapolation for conditions outside the range of direct
observations of total solar irradiance—if
applicable—implies a dec-cen solar irradiance variation with
periods in which irradiance may be lower by as much as 0.25
percent.
Mechanisms
The sun's radiation, volcanic eruptions, and human emissions of
greenhouse gases and aerosols are sources of variability and change
that are external to the climate system. Except for the sunspot
cycle, they are not predictable at this time. A number of internal
and coupled modes of variability within the climate system,
however, influence concentrations of trace gases and aerosols in
the Earth's atmosphere. Understanding the mechanisms and forcings
behind these modes of variability will enhance our ability to
predict climate variations.
Greenhouse Gases and the Carbon
Cycle
The major externally forced causes of the observed CO2 increase are the burning of fossil fuels
(Marland et al., 1994) and forest destruction (Houghton et al.,
1987). Internally, carbon is transferred relatively rapidly among
three major "mobile" reservoirs—the oceans, the atmosphere,
and the biosphere. About one-seventh of the atmospheric CO2 enters the oceans each year, and half as
much is fixed into organic material by photosynthesis on land.
These fluxes are almost balanced by the amounts leaving the oceans
or returned to the atmosphere through microbial decay,
respectively. We know that more CO2
is entering these reservoirs than leaving, however, because the
rate of atmospheric increase is only about half as large as the
global production of CO2 through the
combustion of fossil fuels. From year to year imbalances manifest
themselves in interannual variations of the rate at which
atmospheric CO2 increases; the
swings of net global CO2 uptake
shown in Figure 5-2 are related to known climate variations such as
El Niño. Several studies have found correlations on
different time scales between the rate of atmospheric CO2 increase and global average temperature,
as well as ENSO indicators (Elliot et al., 1991; Dai and Fung,
1993; Keeling et a1., 1995).
Although carbon dioxide cycles quickly among the mobile
reservoirs, it leaves the ocean-atmosphere system only very slowly,
through the burial of organic matter and deposition of carbonate
rocks. Dissolution of calcite, also a very slow process, adds
carbon to the mobile reservoirs, but increases the carbon-holding
capacity of the oceans even more by changing their alkalinity.
Therefore, the rates at which future anthropogenic CO2 is removed from the atmosphere will
depend mostly on how the additional carbon from fossil-fuel burning
is partitioned between the mobile reservoirs,
Page 56
which can essentially be considered to constitute a closed
system. Thus, the important factors controlling increases in
atmospheric CO2 concentrations are:
first, the rate of fossil-fuel consumption; second, the circulation
of the ocean and, to a lesser extent, how the circulation affects
marine biological productivity; third, management of the land; and
fourth, carbon storage by ecosystems, possibly stimulated by
increased CO2 and anthropogenic
deposition of nitrogen. Climate change will affect all of these
natural processes.
The magnitude of the ocean's role in the partitioning of CO2 is dependent on the ocean's chemical
capacity to take up CO2. This
capacity is determined by the amounts of carbonate and borate ions,
which can be "titrated" by newly dissolved CO2 into bicarbonate and boric acid,
respectively. It takes many centuries for most of this capacity to
be accessible to the atmosphere, however, because the ocean turns
over very slowly. The ocean's pH could be lowered by a full point
if "all" (defined as 400×1015 mol) fossil-fuel carbon were burned
(Tans, 1998). (Since the pre-industrial era, we have consumed about
5 percent of that amount.) These estimates are based on the
assumption that the ocean's "biological pump'' keeps operating as
it does today. This pump represents the photosynthesis in the
sunlit surface layer of the ocean and the sinking of organic
particles that keeps the carbon content and CO2 partial pressure lower in surface waters
than in the deep oceans. Without any ocean biology, the partial
pressure of CO2 in the atmosphere
would be two to three times higher than it is today (Najjar, 1992).
At low and temperate latitudes almost all the available phosphate
and nitrate are consumed, but this process is only partially
effective at high latitudes. Changes in the effectiveness of the
biological pump at high latitudes, which result from changes in the
balance between the rates of thermohaline overturning and the rates
of biological production, have been invoked in attempts to explain
the atmospheric CO2 concentration
differences between glacial and interglacial periods (see, e.g.,
Knox and McElroy, 1984, and Sarmiento and Toggweiler, 1984).
The "solubility pump" is another process by which the ocean
maintains a vertical gradient of carbon. Because deep-water
formation sites are cold and the solubility of CO2 is inversely related to temperature,
water with a high inorganic carbon content is "pumped" to the deep
oceans at deep-water formation sites. A vertical CO2 gradient is thereby produced between the
deep water and the warmer, overlying waters of the remainder of the
ocean' s surface. The strength of the solubility pump is affected
by changes in alkalinity, air-sea gas contrast, and ocean
temperature. Without the solubility and biological pumps, the
concentration of atmospheric CO2
would be three to four times higher than it is today (Najjar,
1992).
The carbon delivered to the deep ocean by these pumps is
exchanged with that in the atmosphere on time scales of centuries
and longer. The strengths of both pumps may change with changes in
mixed-layer characteristics and upwelling. The latter will alter
both the nutrient supply and the time available for surface
phytoplankton to utilize these nutrients, as well as affecting the
surface temperature, mixed-layer thickness, and air-sea gas
contrast.
Deforestation—which before 1940 or so occurred principally
in temperate latitudes, but more recently has been taking place
mostly in the tropics—has long been considered a large source
of atmospheric CO2 (Houghton et al.,
1987). In the global balance, tropical deforestation is compensated
for through increased net uptake by terrestrial ecosystems,
principally at temperate latitudes (Tans et al., 1990; Wofsy et
al., 1993; Ciais et al., 1995; Battle et al., 1996). Even in the
tropics there could be large areas of net CO2 uptake (Grace et al., 1995). Possible
explanations for the observed uptake are fertilization of plants by
higher atmospheric CO2 levels (see,
e.g., Mooney et al., 1991) and fertilization by atmospheric
nitrogen deposition (see, e.g., Schindler and Bailey, 1993, and
Townsend et al., 1996). An additional complication in the internal
and coupled mechanisms that produce variations in the carbon cycle
is the fact that the balance between source and sink may shift as
climate changes (Dai and Fung, 1993), which may account for some
increase in terrestrial CO2 uptake.
For example, there is some evidence that the Arctic tundra, once a
net sink for atmospheric CO2, may
have turned into a source during the last decades as a result of
Arctic warming (Oechel et al., 1993). These results still need to
be confirmed by data from additional sampling sites.
Unlike long-lived CO2, methane
has an atmospheric lifetime of about 10 years. The increasing
atmospheric methane burden reflects the growth of CH4 sources in recent decades, with about 60
to 80 percent of this increase attributable to human activities
(IPCC, 1996a). The increasing atmospheric concentration of methane
directly affects the radiative balance and the chemistry of the
troposphere; it accounts for approximately 20 percent of the
increase in radiative forcing since the pre-industrial era (IPCC,
1995). In addition, it has an indirect effect on the stratosphere,
because, once oxidized, it is an important source of stratospheric
water vapor. The most important sources of atmospheric methane are
wetlands, rice agriculture, cattle and sheep, biomass burning,
fossil fuels, landfills and waste, and termites (see, e.g., Fung et
al., 1991).
The atmospheric fate of methane is largely determined by the
level of ultraviolet radiation in the troposphere and the
concentrations of other key trace gases (e.g., ozone, water vapor,
nitrogen oxides, carbon monoxide) responsible for the production
and recycling of OH radicals, which in turn initiate the oxidation
of methane. Perhaps somewhat surprisingly, no significant decadal
trend in OH concentration itself has been detected, in that no
change has been seen in the inferred atmospheric lifetime of the
synthetic industrial compound chemical methylchloroform, which is
also attacked by OH (Prinn et al., 1995). Accurate predictions of
future CH4 levels need to take into
account the effects of the relationships between CH4, CO, and OH (Prather, 1994).
Page 57
Changes in the concentrations of atmospheric chlorofluorocarbons
and similar fully halogenated industrial compounds with no natural
sources are controlled primarily by emissions. The lifetimes of the
individual gases are determined by transport to and photolysis in
the stratosphere. The loss of the equally long-lived nitrous oxide
(N2O) is likewise limited by
stratospheric chemistry, but its emissions come from a mix of both
natural biogenic sources and anthropogenic perturbations to the
nitrogen cycle. Both sets of gases build up in and decay from the
atmosphere on a time scale of centuries. Ozone, on the other hand,
is also an important greenhouse gas, but its atmospheric
concentration is determined primarily by a balance between in situ
production (mainly in the stratosphere) and photochemical losses
that occur on a time scale of a year or less. The balance is tipped
in favor of losses when chlorofluorocarbon (CFC) and N2O concentrations increase, because they
catalyze ozone destruction in the stratosphere.
The source and sinks of anthropogenic radiatively active gases
remain a primary concern in determining the extent of the effect of
greenhouse gases on global climate. However, an equally important
(yet poorly understood) influence on climate is the distribution of
increased atmospheric water vapor associated with a speeding up of
the hydrologic cycle (discussed in detail later). The atmospheric
portion of the hydrologic cycle is complex, and operates on both
short and long time scales. The fast processes associated with this
mechanism, such as cloud formation and the related intra-and
inter-cloud radiative impacts, influence cloud nucleation, longwave
radiation, albedo feedbacks, and ultimately the surface energy
balance. On dec-cen time scales, the impact of increased water
vapor is realized through alterations in large-scale cloud
distribution (shown earlier in Figure 2-12), which reflect both the
water-vapor distribution and the hydrologic cycle's response. Also,
the latent heating of clouds, and its radiative effects, influence
the large-scale atmospheric circulation and hydrologic
cycle—additional complexities that need to be better
understood.
Stratospheric Ozone
Our understanding of the cause of the Antarctic ozone hole has
grown considerably thanks to ground-based and aircraft campaigns,
and, more recently, satellite missions. The evidence is clear:
Observations of high levels of reactive chlorine species coincide
with the observations of rapid ozone depletion. The only identified
cause of this ozone depletion since the 1970s is the rise in
stratospheric chlorine levels, which is driven by the increasing
abundance of chlorofluorocarbons and other halocarbons in the lower
atmosphere. Laboratory studies have identified and quantified the
reactions of chlorine radicals that catalytically destroy ozone,
and numerical models of the stratospheric circulation and chemistry
predict similar losses. This loss in stratospheric ozone was likely
responsible for the recent cooling of the lower stratosphere, and
model results indicate that the ozone loss could be expected to
have a general cooling effect on the climate (IPCC, 1995). An
enhancement of stratospheric ozone destruction would reduce the
stratospheric source of tropospheric ozone and increase the UV
radiation that drives tropospheric photochemistry. Present models
using the best laboratory physics and chemistry can simulate such
ozone loss.
CFCs and related halocarbons have no known natural sources, and
their atmospheric concentrations in 1950 were negligible. The much
higher concentrations currently measured reflect a history in which
halocarbons are emitted at the Earth's surface, propagate
vertically through the troposphere, and, over the course of five
years, ultimately reach the upper levels of the stratosphere. The
rise in tropospheric chlorine loading from CFCs and related
halocarbons is documented in Figure 5-7. (Note that the
concentrations of all of these gases except CFC-12 have begun to
fall since 1993, as a result of declining halocarbon emissions.)
Upon reaching the stratosphere, these compounds dissociate and
release chlorine, which in the upper stratosphere is predominantly
in the form of HCl. Figure 5-8 shows the increase in stratospheric
HCl observed by satellite during the early 1990s. The total content
and magnitude match those of the tropospheric halocarbon sources
(allowing for the 5-year lag). The recent decline in tropospheric
chlorine should be visible in the HCl record over the next several
years.
In addition to their effect on stratospheric ozone, CFCs are
potent greenhouse gases; they have contributed to approximately
one-quarter of the increase in greenhouse-gas radiative forcing
over the past decade (11 percent of the total increase since
pre-industrial times). The decline in radiative forcing that CFCs
induce through stratospheric ozone depletion is likely somewhat
less than their direct radiative contribution (IPCC, 1995).
Aerosols
Since 1978, a series of low-latitude, high-altitude injections
of volcanic aerosols has maintained a maximum in aerosol loading in
the tropics, centered at altitudes between 20 and 27 km. Although
the primary controlling mechanism is external, there are internal
mechanisms that serve to limit the spatial distribution and
temporal longevity of these injected aerosols. For example, the
latitudinal wind gradient in the subtropics impedes transport
between the tropics and mid-latitudes. Thus, the maximum in
aerosols following a large volcanic eruption in the tropics (e.g.,
Mt. Pinatubo or El Chichón) remains for a few years in the
tropical stratosphere as a long-lived source of aerosol for the
middle and high latitudes (Trepte and Hitchman, 1992; Thomason et
al., 1997b). Non-volcanic sources of stratospheric aerosols, such
as natural organic carbonyl sulfide (OCS) and industrially derived
SO2, also tend to support the
presence of a tropical aerosol.
Page 58
While explosive volcanic eruptions are the most significant
source of stratospheric aerosols, a non-volcanic background level
of stratospheric aerosols appears to be present. It has been
suggested that this may result from the diffusion of tropospheric
OCS into the stratosphere (Crutzen, 1976). However, recent research
(Chin and Davis, 1995) suggests that OCS has likely produced only
negligible amounts of the stratospheric aerosols observed since
1978. Hofmann (1990) has proposed that the possible 5 percent
annual increase in the non-volcanic background stratospheric
aerosol mass from 1978 to 1989 could be related to the increase in
sulfur emissions from commercial aircraft or other anthropogenic
sources.
Tropospheric aerosols also influence the overall surface
radiative balance through their light-scattering and absorption
properties. Tropospheric aerosols, being relatively short-lived in
the atmosphere, show regional variability related to the
distribution of their sources. Consequently, the non-uniform
distribution of these aerosols yields spatially heterogeneous
radiative forcing, even given a uniform greenhouse-gas (or natural)
forcing. Examples of tropospheric aerosol sources and typical
aerosol types are: deserts, which produce mineral dust; vegetation,
which produces particulate organic carbon and sulfate aerosols;
biomass burning, which produces soot; oceans, which produce sea
salt; and industrial centers, which produce dust, soot, and
sulfates. The characteristics of the Earth's surface play a role
not only in determining the type of aerosols that are emitted, but
in determining the dispersal of aerosols. For example, the wind
friction caused by surface roughness, which is greatly influenced
by the stature and density of the vegetation, affects turbulent
mixing of air in the planetary boundary layer.
Depending on the absorption characteristics of a given aerosol,
its effect on radiative forcing can be positive or negative. For
instance, low-albedo soot aerosols produced from biomass and
fossil-fuel burning tend to produce surface warming. However,
higher-albedo sulfate and dust aerosols tend to produce surface
cooling, and are believed to exceed the radiative influence of
darker aerosols on a globally averaged basis by 0 to 1.5 W m-2 (IPCC, 1996a).
Predictability
Greenhouse Gases
The primary uncertainty regarding predictions of future warming
associated with increased concentrations of greenhouse gases comes
from uncertainties in the emission scenarios, as well as tremendous
gaps in our understanding of, and ability to represent in models,
the myriad feedback processes that may act to enhance or diminish
any direct warming. One of the most important feedback processes is
the interaction between atmospheric water vapor, clouds, and the
surface radiation balance. The details of this complex interaction
are still poorly understood. Until this understanding is improved,
and better parameterizations constructed, a fundamental question
regarding the impact of greenhouse gases cannot be adequately
addressed.
A great variety of global carbon-cycle models have been used to
estimate future burdens of atmospheric CO2 on the basis of projected rates of
fossil-fuel combustion and associated greenhouse warming. Important
model parameters are calibrated such that the past behavior of
atmospheric CO2, radiocarbon (14C), and the13C/12C
isotopic ratio are reasonably well reproduced. Representation of
the processes responsible for the partitioning of carbon between
the mobile reservoirs is often very crude, sometimes speculative,
and involves many assumptions. These models are best viewed as
extrapolative tools, and their predictive power beyond the next few
decades is tenuous.
Prediction of future atmospheric methane concentrations depends
primarily on prediction of the CH4
sources, and sec-ondarily on how the oxidative capability of the
atmosphere evolves. The main sources of CH4 have been identified, but the range of
uncertainty in emissions rate on regional and global scales is
typically a factor of two or three for each source process.
Stratospheric Ozone
The year-to-year variations of stratospheric ozone at a given
location are not yet predictable, and modulation of the depletion
on 2- to-3-year periods by major volcanic eruptions (as observed)
cannot be predicted. However, the decadal trend in ozone depletion
can be predicted from the evolving load of stratospheric chlorine
(e.g., Figure 5-7b) and bromine: The tropospheric abundance of
their source gases is now declining for the first time, and we can
expect stratospheric levels to follow in a few years.
While the behavior of the Antarctic ozone hole on a decadal time
scale is fairly straightforward, it is predictable in the long term
only if Antarctic meteorology remains more or less the same as
today's. If it does, the ozone hole will persist until chlorine
levels drop somewhere below about 2 ppb, which is expected to occur
around 2050 at best if the phaseout of CFCs and related compounds
is successful. This prediction is fairly certain, although the slow
decay of CFCs and the degree to which the Montreal Protocol is
followed makes the exact date at which we drop below the ozone-hole
threshold uncertain within about 20 years. Fortunately, CFC and
chlorine increases are very predictable, given the future releases
of CFCs and related halocarbons.
Aerosols
Prediction of the long-term influence of tropospheric aerosols
may be possible, but a better understanding will be needed of how
the spatially heterogeneous, but semi-stationary, distribution of
tropospheric aerosols influences the larger-scale climate response.
The stratospheric effects of volcanic aerosols are necessarily
unpredictable, however, because the occurrence and magnitude of
eruptions are not
Page 100
scale, an ice-core record from the Guliya Ice Cap (also on the
Qinghai-Tibetan Plateau) provides evidence of regional climatic
conditions over the last glacial cycle.36Cl data suggest that the deepest 20 m
of this 308.6 m core may be more than 500,000 years old. The d18O change
for the most recent deglaciation is ~5.4 per mil, similar to
changes shown in cores from Huascarán (Peru) and the poles
(Thompson et al., 1997). The oxygen isotopes vary in a pattern
similar to that of the CH4 records
from the polar ice cores indicating that the global CH4 levels and the tropical hydrologic cycle
are linked.
In 1990 a joint U.S.-U.S.S.R. team visited the Gregoriev Ice Cap
(42ºN, 78ºE, 4660 masl) in the Tian Shan region of
Kirghizia. They obtained two 20 m cores containing records of
climatic variation extending back to 1940 (Thompson et al., 1993b).
The d18O profiles in both cores indicate a
warming trend since the mid- 1970s. The team also measured borehole
temperatures. They found a temperature of -2.0ºC in a borehole
20 m deep at 4660 masl, whereas a 1962 Soviet expedition measured a
temperature of -4.2ºC at 20 m even though their borehole
location was at only 4400 masl. This indicates a warming of several
degrees in the near-surface mean annual air temperatures since the
early 1960s. (Because some refreezing of meltwater occurs on
Gregoriev, this 2.2ºC difference is considered an upper
limit.)
There appears to be evidence of significant variability in the
distribution of the snow fields in Canada between 1915 and 1992
(Brown and Goodison, 1996). Remotely sensed data from the Advanced
Very High Resolution Radiometer (AVHRR) also show decadal-scale
changes in the areal extent of snow cover over both Eurasia and
North America (Robinson et al., 1993; Walland and Simmonds, 1997).
The AVHRR-based data, which begin in the early 1970s, indicate more
extensive snow cover in the 1970s to mid-1980s relative to the
later part of the time series; the decline occurs over a span of
approximately five years, from 1985-1986 to 1989-1990. Walland and
Simmonds (1997) found significant co-variability between Eurasia
and North America, with the Eurasian signal lagging behind the
North American signal by over a year. The approximate 10 percent
reduction in Northern Hemisphere snow cover that occurred in the
latter part of the 1980s has likely increased the radiative balance
and surface temperature (up to 1.5ºC) over the northern
extratropical land area, particularly in the spring (Groisman et
al., 1994a,b). The Eurasian winter snow fields also appear to
co-vary with sea-ice distribution north of Siberia (Maslanik et
al., 1996).
Tropical Evidence
There is mounting evidence for a recent, strong warming in the
tropics, which is signaled by the rapid retreat and even
disappearance of ice caps and glaciers at high elevations. These
ice masses are particularly sensitive to small changes in ambient
temperatures, since they already exist very close to the melting
point. One of the best-studied tropical ice caps is Quelccaya
(13ºS, 70ºW, 5670 masl) in southern Peru. In 1983 two ice
cores that went down to bedrock were recovered from there,
providing the first conclusive tropical evidence of the Little Ice
Age (Thompson et al., 1986). Since 1976 Quelccaya has been visited
repeatedly for extensive monitoring. In 1991 and again in 1995
shallow cores were drilled at the summit, near the sites of the
1983 deep drilling and a 15 m coring in 1976. Comparison of the
d18O
records from these four cores (1976, 1983, 1991, and 1995) reveals
that the seasonally resolved paleoclimatic record, formerly
preserved as d18O variations, is no longer being
retained within the currently accumulating snowfall on the ice cap.
The percolation of meltwater throughout the accumulating snowpack
is vertically homogenizing the d18O.
The extent and volume of Quelccaya' s largest outlet glacier,
Qori Kalis, was measured six times between 1963 and 1995. These
observations documented a drastic retreat that has accelerated over
time. Brecher and Thompson (1993) reported that the rate of retreat
from 1983 to 1991 was three times that from 1963 to 1983, and in
the most recent period (1993 to 1995) the retreat was five times
faster. Associated with this retreat was a sevenfold increase in
the rate of volume loss, determined by comparing the 1963-to-1978
volume-loss rate to that of 1993 to 1995. Observations made in 1995
confirmed Qori Kalis' accelerating retreat (see the upper portion
of Color plate 4), as well as further retreat of the margins of the
Quelccaya ice cap, and the development of three adjacent lakes
since 1983.
In 1993 two cores were drilled from the col of Huascarán
(9ºS, 77ºW, 6048 masl), a mountain in the north-central
Andes of Peru (Thompson et al., 1995). The
d18O data from these cores
(see the lower portion of Color plate 4) indicate that the
nineteenth and twentieth centuries were the warmest in the last
5,000 years. Their d18O record and meteorological
observations made in the region reveal an accelerated rate of
warming since 1970, concurrent with the rapid retreat of ice masses
throughout the Cordillera Blanca and of the Qori Kalis glacier.
Additional evidence exists for recent warming in the tropics.
Hastenrath and Kruss (1992) reported that the total ice cover on
Mount Kenya has decreased by 40 percent between 1963 and 1987.
Kaser and Noggler (1991) reported that the Speke Glacier in the
Ruwenzori Range of Uganda has retreated substantially since it was
first observed in 1958. The shrinking of these ice masses in the
high mountains of Africa is consistent with similar observations at
high elevations along the South American Andes, and indeed
throughout most of the world (see Figure 5-29). This general
retreat of tropical glaciers is concurrent with an increase in the
water-vapor content of the tropical middle troposphere, which may
have led to warming in the tropical troposphere (Flohn and Kapala,
1989; Diaz and Graham, 1996).
Page 101
Figure 5-29
Changes in global ice cover during the twentieth century. (Figure courtesy of Lonnie G. Thompson, Byrd Polar Research Center.)
Mechanisms
The cryosphere is thought to be one of the most sensitive
components of the climate system. Its variability is driven by
external forcing, as well as by internal and coupled modes. Its
response to external forcing is most dramatically demonstrated by
the waxing and waning of the glacial ice sheets, which is paced by
changes in solar irradiation. These irradiance changes, which are
caused by changes in the Earth's orbital parameters (see, e.g.,
Hays et al., 1976), are small in magnitude, but significant in
their seasonal distribution, specifically in the high Northern
Hemisphere latitudes. These changes in total solar irradiance
appear to be responsible for the largest variations in climate
experienced over the last several million years: the ice-age
cycles. While the details by which this small change in forcing is
amplified are still unknown, Imbrie et al. (1992) suggest that the
Arctic sea-ice fields respond directly to the radiative changes.
The resulting change in freshwater exported to the thermohaline
source regions of the North Atlantic influences the effectiveness
of the thermohaline circulation system, thereby communicating the
regional changes elsewhere in the globe. (For instance, large
changes may be induced in the Antarctic sea-ice fields). While this
particular scenario is controversial, the fact that the extent of
the continental ice sheets shows strong correlation with the level
of external radiative forcing (see, e.g., Imbrie et al., 1984)
suggests a high sensitivity to that external forcing.
Volcanic activity can also have a strong influence on the
cryosphere (Overpeck et al., 1997). Volcanic-ash deposition can
change ice and snow cover from being one of the most highly
reflective media in the climate system (and thus greatly resistant
to direct solar radiative melting) to one of the most absorptive
(and thus highly susceptible to radiative melting at the surface).
This change in albedo can have a large influence on climate when
ash is deposited on the relatively thin sea-ice cover; when the ice
is melted, the low-albedo ocean surface is exposed, and further
surface warming will result. Sea-ice cover greatly reduces the
ocean-atmosphere exchange of heat, moisture, and radiatively active
gases, so its removal would alter all of those properties. The
reduction in the heat flux associated
Page 102
with sea-ice removal is typically between a factor of 20 (thin
Antarctic ice) and 100 (thick Arctic ice), levels that can
significantly influence regional warming and gas exchange.
Finally, the cryosphere's sensitivity to anthropogenic forcing
is not yet known. GCM simulations often show a high sensitivity to
changes in sea-ice fields in global-warming simulations. For
example, recent numerical simulations of global climate, under
conditions of doubled atmospheric CO2, show that an impressive 38 percent of
the annual average global warming could be attributed to the
response of sea ice (Rind et al., 1995). Note, however, that while
sea-ice changes seem to constitute a considerable positive feedback
in the warming, it is highly uncertain what the response of the sea
ice to global warming will actually be. This uncertainty is most
clearly revealed by similar model experiments regarding the
response of the Southern Ocean sea-ice fields to a doubling of
atmospheric CO2: smaller sea-ice
changes occur in simulations using the coupled atmosphere-ocean GCM
of Manabe and Stouffer (1994) than in the coupled model of
Washington and Meehl (1995). Currently, the details of the
anthropogenic forcing mechanism and its net influence are still
very much in question.
Clearly the most important mechanisms influencing cryospheric
variability are its couplings to the atmosphere, ocean, and land
surface. They lead to a set of geographically unique polar
feedbacks such as the ice/snow-albedo feedback, ice-cloud feedback,
ice-ocean feedback (the effects of which apply to a variety of
scales, from those influencing the sea-ice distribution to those
influencing the vigor of the global thermohaline circulation), and
ice-sheet-ocean feedback, including associated instabilities. Each
of these feedbacks is discussed below.
The ice-albedo feedback encompasses an entire suite of processes
that influence the concentration, spatio-temporal distribution, and
surface characteristics of ice. All these characteristics influence
surface albedo, which in turn alters the surface radiation balance,
which feeds back onto the process itself. For example, sea-ice
concentration—which reflects a balance between the advective
divergence/convergence of ice and the thermodynamic decay/growth of
ice (both processes open/close leads, areas of open
water)—influences air-sea heat flux and albedo. In turn, both
the air-sea heat flux and albedo influence the thermodynamic
growth/decay rate, thereby further altering the ice
concentration.
The surface characteristics of ice and snow, which can change
albedos from a high of near 0.6 to a low of 0.3, are influenced by
moisture content, seawater flooding, ice ridging, age, thickness,
and crystalline properties. These are a function of dynamic
conditions forced by the winds and ocean currents, ice thickness
controlled by the atmosphere and ocean heat fluxes, atmospheric
temperatures, snow loads, atmospheric surface history, and more.
Many of these processes and characteristics are themselves a
function of the albedo.
A comparable set of feedbacks involving land, snow, and
atmosphere can influence the distribution of vast snowfields on
high-latitude land masses. Because of snowfields' high albedo and
great areal extent, even small changes in their distribution can
have a considerable influence on regional and hemispheric
conditions. The snow insulates the ground, controlling its
temperature profile and heat content. These in turn influence the
productivity of the local ecosystems, as well as the snow thickness
and albedo, by moderating the snow's basal heat flux. The
interaction between permafrost, vegetation, and the storage and
release of methane also can lead to climatic feedbacks on longer
time scales.
Snow-cover variations on the high Tibetan Plateau may be
important because of the effect of surface albedo on the strength
of the Asian monsoon (Sirocko et al., 1993). The intensities of E1
Niño events may also be influenced by plateau snow cover. On
longer time scales, model simulations indicate that increases in
snow and ice cover during the last glacial maximum on the Tibetan
Plateau, along with the resulting increases in albedo, may have
caused weakening of the monsoonal circulation (Kutzbach et al.,
1998).
In addition to the surface albedo of ice and snow, the surface
and basal boundary conditions and internal thermodynamics of ice
and snow control the surface temperature, which in turn controls
the surface longwave back-radiation, and thus atmospheric
conditions. The details of this coupled mechanism and the
sensitivity of climate to it are not yet known, though model
experiments suggest that subtle errors in surface temperature can
introduce a systematic bias leading to unrestrained temperature
growth in the models.
The ice-cloud feedback includes a variety of processes that
determine the local cloud formation and distribution, primarily as
a function of the atmospheric column structure, moisture content,
and radiative balance. These processes and properties are
intimately tied to the surface conditions. For example, increased
ice concentration reduces the area of exposed ocean surface,
decreasing the surface heat and moisture sources. Surface heat
plays a fundamental role in determining planetary boundary-layer
characteristics, and surface moisture determines the local
availability of water for cloud formation. Visual images of the
Antarctic obtained from satellites show that large areas of open
ocean (associated with polynyas) in the frigid ice fields lead to
considerable convection and local cumulus generation, which are
more typical of tropical regions. Even the presence of surface-melt
ponds throughout the Arctic pack-ice fields in summer leads to
extensive ground fogs, which have considerable influence on the
surface heat balance. Limited studies in the polar regions suggest
that many of the surface-cloud feedbacks that are observed in
lower-latitude regions behave differently in ice-covered regions
(Curry et al., 1996).
The ice-ocean feedback influences the formation and ventilation
of global deep and bottom waters, while significantly constraining
the ice's thickness and spatial/temporal distribution. This
feedback reflects the important role that
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salinity plays in determining the density of water in cold
regions, although it is moderated to some extent by the associated
latent-heat effects. Sea ice is the predominant source of mobile
freshwater, while the growth of sea ice drives surface salinity
fluxes. As sea ice grows, it rejects seawater salt, which works its
way back into the surface ocean as a brine solution flowing through
drain channels (whose density and effectiveness may be a function
of temperature). Ice growth thus serves as the chief source of
surface salt, which is the dominant determinant of the surface
buoyancy flux in winter. Salt rejection is one of the mechanisms
that effect the density changes that lead to ocean convection, both
shallow and deep. Convection ventilates the ocean and controls the
formation of the intermediate, deep, and bottom waters that drive
the global thermohaline circulation. Even shallow convection can
mix a considerable amount of heat into the surface of the ocean,
and heat content strongly influences ice thickness and
concentration and moderates the extent and seasonality of lead
formation.
The huge margins of the ice sheets, which float on the sea,
contribute considerable freshwater to the ocean, through both basal
melting and ''calving" of icebergs. When this meltwater is injected
at depths significantly below sea level (i.e., from the bottom of a
marginal ice shelf) in the Antarctic, the high compressibility of
the cold water leads to the formation of dense plumes of deep
water. When the meltwater is injected at the surface, from
icebergs, the freshwater can stabilize the surface water and
inhibit deep-water formation. The rates at which all these
processes occur will change with ocean temperature, which can
reflect local, regional, or global influences.
Both sea ice and its snow cover, which represent the predominant
source of freshwater in the polar regions, stabilize the locations
of polar ocean convection. Ice growth in one location—along a
continental shelf, say—will tend to enhance convection there
through salinity rejection. As that ice drifts offshore due to
local winds and later melts, the salinity decrease will tend to
suppress convection. In either case, convection and melt zones are
stabilized by this freshwater input. If the freshwater budget is
altered through variations in ice flux, then the vertical stability
of the ocean may change, leading to a reduction or enhancement of
deep-water formation and ventilation. The significance of this
surface freshwater flux is determined by the underlying ocean
structure, which is governed by the regional wind-stress forcing
and local stratification.
The ice-sheet-ocean feedback is a function of the sensitivity of
the huge ice sheets (formidable reservoirs of freshwater), of sea
level, and of ocean temperatures. The rate at which an ice sheet
flows into the sea is controlled in part by internal pressure
gradients reflecting the thickness and height of ice near its
center of accumulation relative to that at the margins, and in part
by the effectiveness of the frictional coupling at its lower
boundary. As glacial ice flows out onto the ocean, the friction is
removed, and the ice floats and thins as it spreads away from the
continental "grounding line." If sea level is raised, this
grounding line may be shifted inland, causing the ice to decouple
from the bedrock and move more rapidly out into the ocean. The rate
of this acceleration will depend on the nature of the coupling with
the bedrock: If the ice is frozen to the Earth, the friction is
considerable, and its elimination will have a large relative
effect; if the coupling was moderate to begin with due to an
unconsolidated Earth base, or due to basal melting from the
pressure, the elimination of the friction leads to a smaller
relative effect. The slope of the continental landmass inland will
also determine the effect of a sea-level rise. A shallow slope can
lead to a dramatic shift in grounding line, but the shallow slope
will lead to a weaker driving pressure gradient. A steep topography
may result in a minimal shift of grounding line, but the
considerable pressure gradient may result in great destabilization
and rapid drainage of the ice sheet. In either case, the change in
sea level can influence the rate of ice drainage and thus the rate
of additional sea-level change.
Internal mechanisms of cryospheric variability are usually
significant only for the large ice sheets. Internal ice dynamics
may be responsible for altering the basal friction coupling and
internal flow dynamics of large ice sheets; however, the details of
these processes are not well known. Dynamic instabilities can
influence the rate at which ice sheets drain into the ocean, and
thus the rates of both sea-level rise and freshwater supply to the
surface oceans. Ice-sheet surges induced by these instabilities may
account for the vast armadas of icebergs that roamed the North
Atlantic during the last glacial period, leading to the episodic,
brief Heinrich events in that region (MacAyeal, 1994). Some suggest
that the influence of Heinrich events reaches well beyond the North
Atlantic (see, e.g., Bond and Lotti, 1995). It has been speculated
that the West Antarctic ice sheet has collapsed in the past,
possibly because of the aforementioned instabilities; if such a
collapse were to happen again, its impact on global sea level over
centennial time scales would be tremendous.
Yet another set of polar feedbacks is associated with sea-ice
rheology. Internal ice deformation and flow influence lead
formation (affecting ice concentration), ridging events (affecting
ice-surface conditions, seawater flooding, and ice thickness and
concentration), and ice-flow directions (affecting freshwater
distribution and ice concentration). Each of these will influence
the albedo, ice-cloud feedback, and ice-ocean feedback. The
relative importance of these internal feedbacks for cryospheric
variability is not fully known at this time.
Remaining Issues and Questions
• How have the sea-ice, snow, and permafrost fields
changed on dec-cen time scales, and what is the relationship of
these changes to dec-cen patterns of atmosphere, ocean,
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and land-surface variability? As discussed earlier, the
NAO and PNA extend into the Arctic, and their indices often reflect
changes originating in the polar low-pressure cells. It is
therefore important to determine the degree of co-variability
between changes in ice and snow fields and patterns such as these.
Sea-ice changes have been implicated in changes in the thermohaline
circulation on a variety of time scales; how do they co-vary? What
spatial and temporal patterns of contemporary polar climate change
are manifested in changes in permafrost temperature profiles?
• Through what mechanisms do sea-ice fields, atmosphere,
and ocean interact on dec-cen time scales? Do regional or even
local changes in ice divergence alter the albedo, surface heat and
moisture fluxes, upper-ocean conditions, and cloud formation enough
to influence the Icelandic or Aleutian lows, and thus the NAO and
PNA? How do similar near-surface changes influence atmospheric
circulation in the Southern Hemisphere? Do changes in such
patterns, in large-scale planetary waves, or in ocean circulation
alter polar conditions enough to drive other polar changes that may
result in changes to other parts of the climate system? For
example, would a change in NAO influence the volume of freshwater
exported from the Arctic in the form of sea ice enough to
significantly alter the thermohaline circulation? Observational
evidence says it might (Dickson et al., 1997), as do model
experiments (Tremblay, 1997). Also, the thermohaline circulation is
sensitive to surface buoyancy fluxes in source regions. These
sources lie predominantly in the polar regions, where the growth,
decay, and spatial redistribution of ice play dominant roles in the
buoyancy flux, and thus may exert a strong influence on the process
of mid- and deep-water formation. The ice in turn is highly
dependent on the stability of the underlying water column, setting
the stage for considerable feedbacks and interactions among the
system components.
• What are the mechanisms of interaction among the snow
fields, permafrost, atmosphere, and land systems on dec-cen time
scales? Do snow-related changes in surface albedo, surface heat
and moisture fluxes, soil moisture, vegetation cover, and cloud
formation significantly influence atmospheric patterns or
large-scale planetary waves, and thus drive long-term feedbacks in
the climate system? For example, do changes in the seasonal or
spatial distribution of extensive winter snowfields alter the
surface vegetation or soil moisture enough to drive longer-term
influences elsewhere in the climate system? What are the physical
relationships between permafrost surface temperature, surface air
temperature, and other climatic parameters, and what are the
mechanisms controlling these relationships?
• What are the mechanisms by which changes in the
cryosphere of the polar regions are linked or teleconnected to
mid-latitude and tropical regions? Model results suggest that
changes in the sea-ice fields alter the nature of the Hadley cell
through their influence on the equator-to-pole meridional
temperature gradient. Observations suggest that the Antarctic
Circumpolar Wave co-varies with ENSO and Indian Ocean monsoons
through mechanisms not yet understood. Changes in the thermohaline
circulation may be related to changes in the surface freshwater
balance associated with the growth and transport of sea ice. The
ocean's interaction with ice shelves can alter the surface volume
(and thus the gyre characteristics) in the subtropical regions,
which alter SST without any change in surface forcing.
• What are the historical and current global budgets of
glacial ice and snow, and what are the primary mechanisms
controlling those budgets? Because glacial ice and snow budgets
directly affect sea level, we need to better quantify the mass
balance of the continental ice sheets, alpine glaciers, and
permanent snow fields. In particular, the ice mass balance at the
base of the floating ice shelves is in considerable question, and
whether the Greenland and Antarctic ice sheets are gaining or
losing mass is still uncertain. Establishing how this ice and snow
budget has varied through time will give some indication of the
range, rate, and rapidity of change experienced through natural
variability. The IPCC (1996a) lists four major gaps that need to be
filled to obtain better estimates of glacier contribution to
sea-level rise:
1) development of models that link meteorology to glacier mass
balance and dynamic response;
2) extension of models to those glaciers expected to have the
largest influence on sea level (the valley and piedmont glaciers of
Alaska, Patagonian ice caps, and monsoon-fed Asian glaciers);
3) quantification of the refreezing of meltwater inside
glaciers; and
4) better understanding of iceberg calving and its interaction
with glacier flow dynamics.
Other areas of uncertainty concerning ice budgets are the
controls on the melt/growth rate at the base of floating ice
sheets, including the rate of ice-sheet drainage as a function of
sea level (which alters friction), as well as the precipitation
response to cold-region climatic changes.
Processes and Parameterizotions
The internal dynamics and thermodynamics of sea ice and ice
sheets are generally fairly well understood, and can be readily
parameterized despite their complexity. The largest uncertainties
in predicting the extent and thickness of sea ice and glacial ice
lie in the treatment of the boundaries. For example, the surface
albedos of ice and snow under a variety of conditions, and how
those conditions arise, are still poorly resolved and understood.
Heat fluxes across the boundaries can be reasonably parameterized,
but for sea ice, the partitioning of lateral versus vertical heat
fluxes at the edge and base of sea ice with the ocean is still not
understood. This
Page 105
particular distribution dictates the partitioning between
lateral growth/decay of the ice (controlling its geographic
distribution) and vertical growth/decay (controlling its
thickness). This partitioning in turn controls the lead area and
thus the extent of the ice cover, which then affects the albedo,
ice-cloud feedback, and the effectiveness of the insulating ice
cover and air-sea heat flux. Even if this partitioning were
understood, ice modeling would remain problematic because the
prediction of lateral wall area for a given concentration of ice,
or surface forcing, is conceptually difficult.
Considerable uncertainties are associated with the prognostic
treatment of polar clouds. Prognosis is further complicated by the
possibility that polar clouds respond differently from clouds in
other regions to changes in surface and forcing conditions.
Better-polar cloud modeling is needed to more realistically depict
the important ice-cloud feedback discussed above, and an extensive,
fundamental observational data base is required as well. The
observations will improve our theoretical understanding of polar
processes as well as our ability to predict polar-cloud
behavior.
Other aspects of the surface energy balance need to be better
understood to improve long-term cryosphere and climate predictions.
These include the details controlling the spatial heterogeneity of
ice surface conditions and their net influence on surface fluxes,
and the extent to which seawater flooding of thin, seasonal ice
cover affects their melting and albedo, particularly in the
Antarctic polar oceans. Seawater flooding not only may alter the
thermodynamics of the system, but may be corrupting interpretations
of satellite images of ice concentration, confounding our ability
to monitor the ice and evaluate the models.
Models have treated the basal boundary condition of ice sheets
as a function of underlying surface composition and temperature, as
well as of the ocean-ice-sheet interaction along ice shelves. The
observations needed to test these formulations are lacking,
however. Correct model treatment of ice streams, which are
small-scale features with high flow rates, will also require
further study. These streams represent a major path through which
ice sheets are drained, and their distribution may greatly alter
estimates of average ice drainage and ice-sheet stability.
Some of the larger-scale polar feedbacks—for instance, the
export of ice from the Arctic, its role in the formation of North
Atlantic Deep Water, and long-time-scale feedbacks into the polar
regions from any such process—are still a long way from being
fully understood. Likewise, neither the long-term feedbacks between
the climate system and the polar regions, nor the various local and
regional feedbacks already discussed, are well understood—in
particular, how the small-scale feedbacks affect the larger-scale
climate processes. Finally, the prognostic treatment of snow, like
that of precipitation in general, is still a difficult prospect.
Some progress has recently been made, however, although
considerable effort will be required to obtain even a statistically
correct representation of the spatial and temporal distribution of
the snow fields and their dependencies.
Observations
Several types of observations are critical to the issues
articulated above. Long-term monitoring of sea-surface salinity
along with SST is important, since salinity represents the dominant
control over the water density of high-latitude regions. The
sea-ice distribution, motion fields, and thickness need to be known
in order to determine the associated freshwater transports and
buoyancy fluxes. Permafrost temperature profiles provide unique
indications of integrated dec-cen climate change over vast
geographic regions; more such profiles should be collected in order
to better define the spatial and temporal distribution of change.
Consistent monitoring of iceberg calving and an observational
system for determining the basal melt or growth of sea ice (e.g.,
an array of moored buoys measuring temperature and salinity across
the floating ice shelves) must be established before the sea-ice
budget can be closed. Finally, both field and satellite studies are
needed to refine the mass budgets of the Greenland and Antarctic
ice sheets. On-site studies focused on changes in ice flow,
melting, and calving should be continued and extended. Observations
of water-vapor net flux (divergence) will help to pin down the
source of the ice sheets' mass. A laser altimeter on a
polar-orbiting satellite is needed to augment the existing radar
altimetry. These instruments will provide accurate estimates of
ice-sheet volume and give early warning of possible ice-sheet
collapse.
Model parameterizations must be improved to better represent the
ice-albedo feedback, snow-climate feedbacks, ice-cloud feedback,
ice-ocean feedback, ice-sheet-ocean feedback, and ice-sheet
instabilities. Also, simulation of sea-ice and snow distribution
and related impacts must be improved. Randall et al. (1998) have
described some of the observational requirements necessary to
improve our ability to model these processes on large scales, and
some of the existing research programs that have been designed to
fulfill these requirements.
The feedbacks among the hydrologic cycle (including river runoff
into the Arctic), the atmospheric circulation, and the thermohaline
circulation must be better understood on a variety of scales,
because such larger-scale feedbacks may play a fundamental role in
polar climate. The potential for extracting high-resolution records
of past climate change from polar sediments along the Antarctic
continental shelves and slopes and in polar fjords and Arctic lakes
and estuaries (the latter being the primary focus of the
Paleoclimates of Arctic Lakes and Estuaries (PALE) program of the
Arctic System Science initiative) should be evaluated, and pursued
if proven feasible.
Page 106
Land and Vegetation
Influence on Attributes
The state of the land and its vegetation affect the climate in a
number of ways. The fraction of solar irradiance absorbed by
different landscapes depends on vegetation. For example, deserts
reflect a greater portion of the incoming solar radiation than
vegetated regions do. Of the vegetated regions, grasslands reflect
more radiation than surfaces covered by forests. The influence of
vegetation on albedo is amplified when the angle of incident solar
radiation is low, as it is during high-latitude winters, and when
snow covers the ground; forests present a light-absorbing layer
above the snow, whereas bare ground and grasses do not. To assess
the sensitivity of high latitude regions to the albedo effect of
boreal vegetation, Bonan et al. (1992) computed the climate
response of a GCM in which the forests north of 45ºN were
replaced by bare ground. The zonally-averaged temperatures in this
deforestation simulation were generally between 3 and 10ºC
colder in the mid- to high-latitudes, relative to a simulation
conducted with the same GCM in which forests north of 45ºN
were present. Whitlock and Bartlein (1997) suggest that changes in
vegetation may have played a significant role in climate changes in
northwest America over the last 125,000 years. In a set of GCM
simulations of Cretaceous-era climate, Otto-Bliesner and Upchurch
(1997) found that the globally averaged temperature was over
2ºC warmer in a model run that included a best-guess estimate
of global vegetation cover than in one in which bare soil covered
the land surface. The vegetation decreased the surface albedo,
causing high-latitude areas to warm and delaying sea-ice formation,
which in turn further decreased albedo and increased
temperatures.
Processes in the soil and plants both absorb and produce
long-lived greenhouse gases (CO2,
CH4, N2O), thereby influencing the atmosphere's
infrared-radiation budget. Vegetation emits chemically reactive
organic gases (terpenes, isoprene, methanol, etc.) that are
involved in atmospheric reactions that lead to production of ozone
in the troposphere. Photochemical processes in the lower atmosphere
also cause small particles to be created from hydrocarbons emitted
by plants. These particles scatter light, causing a visible bluish
haze that decreases the transmission of solar radiation to the
ground. Soil and mineral dust, whose atmospheric entrainment is
influenced by vegetation cover, also affects the scattering of
light. Both surface temperature and turbulent mixing of air in the
planetary boundary layer are functions of wind friction at the
Earth's surface. Surface roughness is greatly influenced by both
the stature and density of vegetation, in addition to the effects
of topography.
Plants also partly control the hydrologic cycle through
evapotranspiration, as noted earlier in this chapter. Leaves open
their stomata during photosynthesis, causing them to lose water
vapor into the atmosphere while taking up CO2. Roughly two-thirds of precipitation
over land is recycled water vapor from plants. Model simulations of
the Amazon confirm that the vegetation plays an active role in
maintaining the regional hydrologic regime; simulated deforestation
resulted in dramatically decreased precipitation and increased
temperature and evaporation (Shukla et al., 1990). Soils, which are
a very slowly created mixture of rock-weathering products and
organic material derived from plants, function as a reservoir of
water. They thus influence the timing of evaporation from the land
surface. Therefore, plants indirectly influence surface temperature
through their effect on soil moisture, which has a large heat
capacity, and thus influences latent heating. Evapotranspiration
also changes the balance between the fluxes of sensible and latent
heat at the surface, causing local surface cooling. When plants are
water-stressed, their stomata may close to reduce transpiration and
conserve water, thereby warming the surrounding air. The expected
physiological response of plants to a high-CO2 world would be to close their stomata
somewhat, reducing their evaporative loss, but furthering warming
over the continents (Sellers et al., 1996).
The urban landscape has a marked influence on climate; a
recognized problem in studies of long-term temperature change is
that many meteorological measurement sites have gradually become
absorbed into expanding metropolitan areas, known as ''urban heat
islands." The artificial heat output of the greater New York
metropolitan area is about one-eighth of the solar energy absorbed
there on the ground. Furthermore, wind speed has diminished,
particle loadings have increased, anthropogenic emissions of many
trace gases have increased, and precipitation and other weather
features have changed markedly for tens of miles downwind from many
urban areas (Barry and Chorley, 1992).
Evidence of Decade-to-Century-Scale
Variability and Change
The major ecosystem zones (biomes) of the Earth, such as tundra,
temperate grassland, and wet tropical forest, are determined in
part by the range and variability of a region's temperature and
precipitation. The type of vegetation prevalent in the past at a
given location is sometimes recorded in pollen buried in ancient
soils and sediments. Such data show that large vegetation changes
have occurred in many areas in response to climate change. For
example, pollen data tell us that large parts of the Sahara,
although currently completely barren, supported vegetation (savanna
woodland and desert grassland) from about 9500 to 4500 BP (see,
e.g., Ritchie et al., 1985). Evidence has been found of increased
lake levels in the area during the same time period; both
conditions have been linked to a strengthened monsoon circulation
in that period (Kutzbach and Street-Perrott, 1985). In western
Europe, many tree species such as pine, elm, and oak migrated
Page 107
northward and westward with surprising rapidity (typically 150
to 500 m per year) after the close of the last ice age, replacing a
shrub-dominated vegetation (Huntley, 1988). Species adapted to
Arctic and alpine tundra suffered a crisis in western Europe during
the warm period in the mid-Holocene, around 6000 BP, when their
habitat was at a minimum. More recently, the succession in which
the dominant tree species changed from beech to oak to pine during
the Little Ice Age has been recorded in pollen in southern Ontario
(Campbell and McAndrews, 1993).
Between one-third and one-half of the Earth' s surface has been
transformed by human actions (Vitousek et al., 1997). The evidence
of ecosystem variations is especially pronounced during the last
150 years. Vast tracts of temperate forest were cut down during the
nineteenth and early twentieth centuries. The location of greatest
deforestation has shifted to the tropics in the most recent
decades. One-fifth of the tropical forest area was lost between
1960 and 1990, and it is estimated that the remaining area is being
lost at a rate of 7 percent per decade (WRI, 1996). Today, almost
40 percent of the Earth's land area (excluding Antarctica) is
devoted to cropland and permanent pasture (WRI, 1996). Most of this
agricultural expansion has occurred at the expense of forests and
grasslands; only a few small patches of original prairie remain on
the North American continent. The majority of wetlands in the
United States have been drained (Kusler et al., 1994) during the
last half-century. Aerial photography shows clearly how dominant
society's influence over the land is—human settlements and
structures, roads, a checkerboard of croplands, artificial lakes,
coastal modifications, and so on. Currently about 8 percent of the
land in Western Europe is (sub)urbanized or covered by roads, and
45 percent is devoted to cropland and pasture; in the United States
the corresponding figures are 4 percent and 45 percent,
respectively (WRI, 1996). It is possible, of course, that the
relatively large portion of the land surface that is managed in
some way by humans may permit us to exert a modest amount of
deliberate climate control, since we can control the reflective and
absorptive properties of man-made structures.
Evidence for changes in the amount of carbon stored in
vegetation and soils derives principally from our knowledge of
changes in land use. Deforestation results in the loss of carbon in
standing wood, and causes oxidation of part of the organic carbon
stored in forest soils. Agricultural practices and reforestation
also affect the carbon balance. Combining the recorded global
history of land use with time-dependent models of carbon dynamics,
Houghton et al. (1987) estimated the loss of carbon to the
atmosphere that can be attributed to direct human intervention to
be 1.0-2.6 Gt of carbon per year in 1980; this flux has varied
through time since at least 1850 (Houghton, 1993; IPCC, 1996a).
While land-use changes are important, recent climate variability
has probably also led to substantial changes in vegetation-related
carbon fluxes. Using historical temperature and precipitation data
in conjunction with a carbon-cycle model, Dai and Fung (1993) found
that climate may have caused significant interdecadal variations in
regional and global terrestrial carbon storage since 1940.
Observations of increasing amplitude of intra-seasonal atmospheric
CO2 variations (Keeling et al.,
1996a) and remotely-sensed, large-scale increases in terrestrial
photosynthetic activity (Myneni et al., 1997) suggest that plant
growth has increased in recent years.
Measurements of CO2
concentrations in ice cores provide a very clear record of changes
in carbon storage between glacial and interglacial times, but these
measurements do not directly distinguish the respective roles
played by the oceans and the terrestrial systems in causing the
atmospheric concentration changes. Obviously changes in carbon
storage can be expected when the geography of vegetation is
significantly altered (Prentice and Sykes, 1995; Friedlingstein et
al., 1995). It has also been inferred from ice-core records that
the emissions of CH4 varied between
glacial and interglacial periods (Chappellaz et al., 1993b;
Thompson et al., 1993a), and that they have strongly increased in
recent years. Changes in ecosystem types and land use (wetlands,
rice paddies, cattle grazing, etc.) clearly have had a major impact
on these emissions. Global N2O
emissions have also increased during recent decades, but there is
still considerable uncertainty as to the cause.
Mechanisms
Past vegetation changes have been driven by natural climate
variations. Regional and global models of vegetation dynamics are
based on the sensitivity of species and ecosystems to variables
such as the mean coldest-month temperature, the annual accumulated
temperature over 5ºC, precipitation, and soil moisture
capacity (see, e.g., Prentice et al., 1992; VEMAP, 1995). These
variables reflect vegetation characteristics, such as: most woody
tropical plants are killed when the temperature drops below
0ºC, and for a species to sustain growth, the air temperature
must exceed a species-specific minimum value for a species-specific
minimum length of time (expressed as growing degree days). The
climate warming that has been projected for the coming centuries
could induce changes to natural vegetation as great as those at the
end of the last ice age; species distributions in North America
could be shifted by as much as 500 or 1,000 km (Overpeck et al.,
1991). The variability and types of disturbance are another
significant factor in determining ecosystem composition and
distribution. For instance, the frequency and severity of wind
storms and fires affect the migration and establishment of species
and ecosystems, and need to be taken into account in predicting the
geography of future ecosystems (Overpeck et al., 1990).
At present, the dominant reason for changes in vegetation is
direct human intervention, both purposeful and inadvertent. More
than half of the ice-free surface of the continents has been
altered substantially by human uses (Kates et al.,
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1990). Our need for food and resources drives land-use patterns
that result, either rapidly or gradually, in land-cover changes.
Vitousek et al. (1986) estimate that 31 percent of all net primary
production on land directly serves humans as fiber, food, or fuel;
2.3 percent is actually consumed by us or by animals that we use
for food. As human populations grow and place even more demands on
our natural environment, this already pervasive influence of humans
on the Earth's biota is likely to concomitantly increase.
Anthropogenic change in ecosystem functioning can also result from
the removal of predators or the introduction of invasive
species.
Both direct and indirect effects of CO2 have been recognized as mechanisms of
change in the interaction of, and competition between, species
composing the vegetation on undisturbed land. Fertilization of
plant growth by higher atmospheric CO2, and by moderate amounts of wet and dry
deposition of nitric acid, has a demonstrable influence on
vegetation. To test the response of intact ecosystems to CO2 changes more realistically than in
laboratory settings, so-called free-air CO2 enrichment (FACE) experiments are being
carded out, in which CO2 is pumped
over ecosystems in their natural environment. One such experiment,
carded out in Chesapeake Bay wetlands, is reported by Drake (1992).
Long-term CO2-enrichment studies
with manipulated microclimate have also been carried out in
enclosures, an example of which is the assessment by Tissue and
Oechel (1987) of the effects of temperature and CO2 change on Arctic tundra. The responses
to elevated CO2 in the Arctic and
Chesapeake Bay cases were quite different, which suggests that
nutrient (especially nitrogen) availability may play an important
role in regulating response to increased CO2 and temperature (Rastetter et al.,
1992). Another effect of enhanced levels of CO2 is that plant root-to-shoot ratio tends
to increase (Rogers et al., 1994). Schindler and Bailey (1993)
estimated that the amount of carbon storage stimulated by
anthropogenic nitrogen deposition may be between 1.0 and 2.3 Gt of
carbon per year. Others, such as Asner et al. (1997), consider the
potential of this effect to be lower.
Not only does enhanced CO2
fertilization tend to increase the amount of carbon stored in live
vegetation, but it can alter the species balance of ecosystems. For
instance, under greater ambient CO2
concentrations, C3 plants (the
majority of crops) tend to be favored over C4 plants (some essential warm-weather
crops, including corn and sugar cane) (Poorter, 1993). Shifts in
species composition may also arise from changes in the availability
of nutrients (Wedin and Tilman, 1996). By affecting climate,
elevated CO2 levels may also
indirectly influence the number and balance of species in
ecosystems (Davis and Zabinski, 1992; Barry et al., 1995).
There is increasing awareness that future ecosystems may not
represent a simple, climatically driven redistribution of
ecosystems as they are currently composed. Rather, mechanisms such
as those mentioned above must be taken into account in predictions
about future ecosystems. Furthermore, factors such as changes in
land use and management, which are quite difficult to predict, are
likely to be at least as important as climate-related changes.
Direct human intervention and climate change do not act as
independent agents of vegetation change. Both the U.S. Dust Bowl of
the 1930s and the desertification in the Sahel are examples of how
unfavorable climatic conditions and societal demands may
synergistically lead to environmental degradation.
Variations in fire frequency and intensity can often be related
to variations in climatic conditions. For instance, the widespread
fires in southeast Asia in 1997 have been attributed to the extreme
E1 Niño-related drought at that time. In addition to having
a direct economic impact when fires affect forestry and personal
property, such changes can also influence ecosystems in a number of
ways. Active suppression of forest fires, while benefiting humans
in many ways, can be detrimental to certain species that depend on
fire for various reasons (e.g., facilitating germination).
Moreover, fire suppression can lead to age homogenization, in which
forests tend to become dominated by single-age stands. Although
uniform forests are often high in timber productivity, the decrease
in diversity leaves them generally more vulnerable to fire,
windstorms, disease, and other naturally occurring events (Noss and
Cooperrider, 1994). Increased fire frequency can also cause local
extinctions of species, even in mature forest stands (Gill,
1994).
Acid deposition has led to widespread dieback of trees,
especially at higher elevations. Elevated surface ozone can reduce
photosynthesis, increase respiration, and lead to leaf senescence
earlier in the season (Chameides et al., 1994), all of which reduce
productivity. Increased UV-B radiation has been shown to reduce
photosynthesis and growth in many species in greenhouses, although
the effects are less marked under field conditions where light
levels are high (Allen and Amthor, 1995).
By causing changes in the vegetation and the soils, the
above-described processes will have an impact on the biogeochemical
cycles. Because climate depends in part on the chemistry of the
atmosphere, large-scale atmospheric chemistry-vegetation-climate
feedbacks may exist. For instance, higher atmospheric CO2 concentrations may directly (via the
fertilization effect) and indirectly (via climate-induced changes)
increase carbon sequestration in vegetation (see, e.g., Woodwell
and Mackenzie, 1995), yielding a negative feedback to the level of
atmospheric CO2. Several other
chemistry-vegetation-climate feedbacks have been proposed, many of
which are discussed in Woodwell and MacKenzie.
Predictability
Future changes in the composition and distribution of
ecosystems, and the accompanying biogeochemical cycles of carbon
and nitrogen, are hard to predict. Not only are a large number of
factors simultaneously undergoing change, but we cannot be certain
of future human actions. Pollution, fer-
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tilization, climate change, land use, succession, the use of
pesticides, species extinction and the introduction of new species,
and the fragmentation of once-widespread ecosystems are all
occurring at once, often making it difficult to decipher cause and
effect. For example, atmospheric measurements have established the
existence of a carbon sink of appreciable magnitude at temperate
latitudes on the continents of the Northern Hemisphere (Tans and
White, 1998), but it has proven difficult to choose among the
competing explanatory hypotheses (Houghton et al., 1998).
Candidates are CO2 fertilization,
nitrogen fertilization, afforestation, and a climate-driven
increase in carbon storage (see IPCC, 1995, for an overview).
Changes in land use do not follow a predictable progression. The
progression will depend on local factors, most of them economic,
social, and technological. For example, population growth has in
many cases contributed to the conversion of forested land to
farmland, but in the eastern United States and western Europe the
process has been reversed during the last 50 years (McKibben,
1995). Accurate modeling of climate and atmospheric chemistry
require the accurate specification of land-surface parameters that
are intimately tied to the fluxes of heat, water vapor, and trace
gases. Although attempts have been made to predict land-use changes
(Zuidema et al., 1994), our skill in this regard is still low,
largely because we lack sufficient insight into what has been
called the human dimension of global change. The International
Geosphere-Biosphere Programme's Human Dimensions Project has
outlined a science/research plan designed to increase our skill in
predicting the progression from human needs to land-cover change
(Turner et al., 1993). The plan proposes to classify the world's
land area by similar social and environmental circumstances into a
manageable number of categories, and probe the causal connections
in each category in more detail according to a common protocol or
framework.
Remaining Issues and Questions
• What are the effects of human activity and climate
change on ecosystem structure and function? From paleoclimatic
records, we know that both vegetation and animal species respond to
climate variations according to their individual tolerances.
Competitive and trophic interactions among species are thereby
altered, redefining where organisms can survive and reproduce and
changing ecosystem compositions. The ability of organisms to
respond to future climate variations or change will be greatly
influenced by human land-use patterns and other anthropogenic
influences. Associated with structural changes in ecosystems are
changes in the biogeochemical cycling of carbon and nutrients, in
ways that remain difficult to anticipate. Finally, the distribution
of disease-carrying organisms will change with ecosystem
restructuring and redistribution (IPCC, 1996b).
• What are the relative contributions of the different
processes by which vegetation and soils store or lose carbon?
Vegetation and soils store three times as much carbon as the
atmosphere or upper ocean, yet large uncertainties remain regarding
the quantitative contributions of various processes. The carbon
sink in the Northern Hemisphere has increased over recent decades;
forest regrowth resulting from changing land-use patterns, or
perhaps increased fertilization by CO2 and nitrogen, or simply climate change
may have been factors in this increase.
• At what rates will vegetation and soils emit
CH4, N2O, and volatile organic carbon (VOC)
compounds in the future? CH4
production in soils depends strongly on moisture conditions
(including the extent of the permafrost, which is slowly melting).
N2O production is a result of
denitrification processes that occur in soils. The rates at which
VOC compounds (ozone precursors) are emitted depend heavily on the
species involved. Changes in these emissions will depend on a
combination of factors involving both ecosystems and climate.
• How do dec-cen-scale changes in land use and land
cover affect the energy balance of the land surface on dec-cen time
scales? The nature of land cover, which determines its
reflectivity, is expected to change with changing climate and human
activities. For example, a warmer high-latitude climate will favor
the expansion of boreal forest into tundra-dominated regions, with
a concomitant lowering of the albedo. Desertification, which may
result from human or natural activity or both, increases surface
albedo. The thermal structure, moisture content, and dynamics of
the atmosphere are influenced by the proportions of sensible and
latent heat transferred from the surface, which is a function of
the type and extent of land cover.
• How does vegetation influence the transfer of
freshwater through the land surface on dec-cen time scales? The
extent of stomatal opening influences the rate of
evapotranspiration from the land surface. Higher atmospheric
CO2 concentrations will cause
CO2 to more readily enter plants;
plants will then be able to keep their stomata somewhat more
closed, which will decrease their transpiration losses and increase
their water-use efficiency. An increase in vegetation density tends
to decrease runoff and increase evaporative fluxes, resulting in
greater atmospheric water-vapor content and precipitation over
land.
• How does changing vegetation cover influence the
loading and composition of atmospheric aerosols on dec-cen time
scales? Vegetation naturally emits aerosol precursors (e.g.,
non-methane hydrocarbons), and the nature and amount of these
compounds depends on the species. The distribution of aerosol
precursors will therefore change as ecosystems and species respond
to climate variations and human perturbations. Biomass burning
generates aerosols (particularly soot) that influence the regional
radiation balance. Desertification produces mineral dust that is
transported into the troposphere and exerts a regional
radiative
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forcing. The distribution of all these aerosols can be expected
to vary on dec-cen time scales in response to climatic and human
influences.
Processes, Parameterizations, and
Observations
Changes in land-surface characteristics—including surface
vegetation, topsoil extent, and soil moisture—must be
monitored on a long-term basis. Not only do these changes alter the
distribution of surface reservoirs of radiatively active gases and
the surface-atmosphere exchange of those gases, they also influence
albedo and, through stress effects on plant evapotranspiration
efficiency, the hydrologic cycle.
Long-term monitoring of near-surface aerosol distributions will
be required to assess whether perturbations of stable gradients of
these aerosols could induce stationary changes in the surface
radiation balance, which could lead to large-scale alteration of
circulation.
In order to improve models' abilities to predict dec-cen-scale
variability, we need to more realistically parameterize many
land-surface processes, such as: interactions between soil and
vegetation under various conditions (including frozen soils);
surface-atmosphere gas exchange and net uptake (including
biogeochemical and physical feedbacks); and the effect of
land-surface processes on atmospheric conditions, (including
evaporation and precipitation). Clearly our understanding of most
of these processes must be improved first.
Land-surface characteristics and radiatively active atmospheric
constituents are vital sets of climate-model parameters, and are
generally not prognostic variables that can be used interactively
by models. At present, because changes in these factors cannot yet
be adequately predicted, they are considered to be an external
forcing in most models, and their characteristics must be specified
in advance. Even in the absence of any significant skill in
predicting land-cover change, however, we can usefully run
different vegetation scenarios in physical global-change models.
This approach would at least yield some insight into likely
climatic and environmental consequences of those scenarios, and
provide some guidance for setting environmental-policy goals
pertaining to land cover. In addition, as with greenhouse gases,
the transient evolution of land cover (including wetlands) under a
slowly changing climate and rapidly exploding population must be
monitored to provide the boundary conditions needed for model
simulations and assessment of plausible future trends.