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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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Page 48 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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Page 49 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

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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.)

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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.)

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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

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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

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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.

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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,

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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).

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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.

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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

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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).

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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

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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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Page 103 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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Page 104 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

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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.

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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

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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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Page 108 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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Page 109 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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Page 110 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.