Info

Land - Ocean - Atmosphere Interactions and Monsoon Climate Change: A Paleo-Perspective

John E. Kutzbach

Center for Climatic Research, University of Wisconsin - Madison

Madison, Wisconsin

Sandy P. Harrison

Max Planck Institute for Bioyeoehemistry, lena, Germany

Michael T. Coe

Center for Sustitinitbihty and the Global F.nvironment, University of Wisconsin-Madison.

Madison, IV is con sin

Introduction 73

Response of the Monsoon to Orbital Forcing 75

Ocean Feedbacks on the Monsoon 76

Land-Surface Feedbacks on the Monsoon 77

Synergies between the Land, Ocean, and Atmosphere 81

The Role of Climate Variability 82

Final Remarks 83

References 83

1. Introduction

The climate system involves multiple interactions between the atmosphere, the land surface, and the oceans. Understanding both the physical and the biogeochemical linkages between these components is a fundamental challenge for earth system science. In addition to the complexity of the linkages and the existence of synergistic relationships between them (see, e.g., Berger, in press), there are very real difficulties in studying interactions which operate on timescales ranging from seconds to many millennia and in which, as a consequence, the relationship between cause and effect can be reversed (Schumm and Lichty, 1965). The seasonal cycle of atmospheric CO,, for example, is controlled by changes in the terrestrial biosphere as a function of plant phenology (Knorr and Heimann, 1995). On longer (i.e., multimillennial to glacial-interglacial) timesscales, changes in atmospheric [CO,] affcct the competitive balancc between C, and C4 plants, which dccreascs the productivity of the terrestrial biosphere and causcs massive redistributions of major vegetation types (Crowley and Baum, 1997; Levis ct al., 1999a). It is clear that a complete understanding of the physical and biogcochcmi-cal linkages in the earth system must includc how they have operated on longer timescales and the consequences of changes in their operation.

Furthermore, the recent geological past offers significant opportunities for using climate and earth system models to study geosphere-biosphere interactions. First, the observed changes in climate and paleoenvironmental conditions were large (COHMAP Members, 1988; Wright ct al, 1993). We can therefore expect to be able to resolve these changes even with the present somewhat limited earth system modeling capability. Second, the fundamental cause of these changes lies in changes in earth's orbital geometry (Hays et al, 1976; Imbrie, 1985; Berger, 1988), and the consequent changes in the seasonal and latitudinal distribution of incoming solar radiation (insolation) can be precisely specified (Berger, 1978; Berger and Loutre, 1991). Finally, at least for the most recent 10,000-30,000 years of the earth's history, the continental-scale to global-scale databases that have been assembled provide spatially- explicit reconstructions of climate and environmental parameters that can be used to benchmark model simulations (see, e.g., Kohfeld and Harrison, 2000). Thus, we can anticipate being able to study the changing role of physical and biogeochemical linkages between the atmosphere, the land, and the ocean within the climate system in an increasingly detailed fashion by combining pale o-observations and carefully designed model experiments.

Nevertheless, these studies (and more specifically the models required to make them) are still in their infancy. Until recently,

GLOBAL BIOGL.OCIIIi.MICAI. CYCt.LS IN litt. CLIMAi li SYS ILM

Copyright 'O 2001 by Ac.tdemic Press. Ali rights of reproduction in any form reserved.

a) Orbital Forcing a) Orbital Forcing max desert/steppe transition min '

max desert/steppe transition min '

b) Ocean Feedbacks desert/steppe transition min '

b) Ocean Feedbacks

LW BMRC

•ECHAM3 GEN2 GFDL GISS

d) Ocean + Land-Surface Feedbacks - + Orbital Forcing d) Ocean + Land-Surface Feedbacks - + Orbital Forcing

15 Latitude

R-C RVS-C RVSL-C RVSLW-C

(ocean-atmosphere vegetation) OA

(ocean-atmosphere) AV

(atmosphere vegetation)

(atmosphere only)

15 Latitude

e) 6000 yr B.P. Biome Distribution

Savannah

Xerophytic woods/scrub

Steppe

Desert

Modern Biome Distribution i i i ® i i i Savannah i-i-i—A-ik—i-i-1-1 Xerophytic woods/scrub i-1-1--i-1-1 Steppe

FIGURE 1 Zonally averaged simulated annual precipitation anomalies (6000 year B.P.—Control) versus latitude for northern Africa (land grid cells between 20°W and 30°E). Precipitation anomalies include the effects of: (a) radiative forcing (R) alone for the 18 climate models participating in the Pa-leoclimate Modeling Intercomparison Project (Joussaume et «/., 1999); (b) radiative forcing plus ocean feedbacks (ASST) for an asynchronous coupling of GENESIS2 and MOMl (Kutzbach and Liu, 1997); (c) radiative forcing plus land-surface feedbacks (soil, S; vegetation, V; lakes, L; and wetlands, W) simulated using CCM3 (Brostrom et al„ 1998); and (d) radiative forcing (A) plus ocean feedbacks (OA) from a fully coupled simulation with the 1PSL AOGCM, radiative forcing plus vegetation feedbacks (AV) from an AGCM simulation forced with 6000 yr B.P. vegetation derived by forcing BIOME 1 with the output from the OAGCM simulation, and radiative forcing plus ocean- and land-surface feedbacks from an asynchronous coupling of the IPSL AOGCM and BIOME1 (Braconnot et al„ 1999). The hatched lines in (a-d) represent upper and lower estimates of the additional precipitation (excess over modern) required to support the grassland vegetation observed in northern Africa at 6000 yr B.P. (see Joussaume et at, 1998). (e) Latitudinal distribution of biome types (desert, steppe, xerophytic, and savannah) for 6000 yr B.P. and 0 yr B.P. over the longitudes 20W-30E (Joussaume et al„ 1999).

past global changes have been primarily studied with atmospheric general circulation models (AGCMs). In simulations made with these models, other components of the earth system that were thought to have been important at a particular time (e.g., changes in the extent of land, ocean, and ice cover at the last glacial maximum, ca. 21,000 yr B.P.) are specified from observations. This approach is limited, both because of our inability to completely specify many paleoenvironmental boundary conditions (see, e.g., Broccoli and Marciniak, 1996) and, perhaps more importantly, because it ignores potential feedbacks between, e.g., land-surface, biospheric, or ocean changes on the atmosphere. The advent of fully coupled ocean-atmosphere models (e.g., Meehl, 1995; Murphy, 1995; Johns et al„ 1997; Braconnot ct al„ 1997; Gent ct al„ 1998) and atmosphere-vegetation models (e.g., Foley et al., 1998, in press; Levis et al., 1999a,b) allows these feedbacks to be included. However, full coupling between the ocean, atmosphere, and vegetation has only been achieved (to date) in very much simplified models (tire so-called EMICs or models of intermediate complexity: see, e.g., Gallee et al, 1992; Ganopolski et al., 1998). EMICs do not incorporate even all of those land-surface processes that are thought to impact on climate, and there are many otlrer aspects of the interaction between tire land, ocean, and atmosphere that have not been addressed in any coupled modeling scheme.

Although we are still a long way from having fully functional earth system models, some significant progress toward understanding the physical and biogeochemical linkages between the atmosphere, the land surface, and the oceans on geological timescales has been made during the past years. This chapter is not meant to provide a complete review of the state of knowledge, but rather (a) to illustrate the gains that have been made in understanding the linkages and feedbacks within the earth system by focusing particularly on the tropical monsoons and (b) to suggest some areas for future interdisciplinary work in this area.

2. Response of the Monsoon to Orbital Forcing

The expansion of the area influenced by the Afro-Asian summer monsoons during the early to mid-Holocene is one of the most striking features shown by palaeoenvironmental data (see, e.g., Street and Grove, 1976; Street-Perrott and Harrison, 1985; Street-Perrott et al, 1989; Roberts and Wright, 1993; Street-Perrott and Perrott, 1993; Winkler and Wang, 1993; Gasse and van Campo, 1994; Jolly et al, 1998a,b; Prentice and Webb, 1998; Yu et al, 1998, 2000; Kohfeld and Harrison, 2000; Prentice et al, 2000). The fundamental mechanism underlying these changes is well known (see, e.g., Kutzbach, 1981; Kutzbach and Otto-Bleisner, 1982; Kutzbach and Street-Perrott, 1985; Kutzbach and Guetter, 1986; Kutzbach and Gallimore, 1998; COHMAP Members, 1988; Kutzbach et al, 1993): the orbitally induced enhancement of

Northern Hemisphere summer insolation during the early to mid-Holocene (Berger, 1978) resulted in increased heating over the Northern Hemisphere continents and thus intensified the thermal contrast between the land and the ocean. The increased heating over the continents resulted in the northward displacement of the intertropical convergence zone (ITCZ) and hence of the monsoon front, while the enhanced land-sea contrast increased the flux of moisture from the ocean to the continent.

The response of the climate system to orbital forcing during the mid-Holocene (ca. 6000 Yr B.P.) has been investigated by a range of atmospheric general circulation models (AGCMs) within the Palaeoclimate Modelling Intercomparison Project (PMIP: Jous-saume and Taylor, 1995; 2000). In these simulations, the atmospheric [C02] was reduced (from 345 to 280 ppmv), but land-surface conditions and sea-surface temperatures (SSTs) were prescribed to be the same as today. The effect of the reduction in atmospheric [C02] is negligible given that the simulations were run with fixed (modern) SSTs; thus, the experiment can be viewed primarily as an examination of role of orbitally induced insolation changes on climate. The PMIP simulations confirm that orbital changes produce a significant enhancement of the Afro-Asian monsoons but show that the magnitude of the enhancement varies from model to model (Joussaume et al, 1999; see also individual simulations: Dong et al, 1996; Hewitt and Mitchell, 1996; Lorenz et al, 1996; Hall and Valdes, 1997; Masson and Joussaume, 1997; Vettoretti et al, 1998). The sensitivity of the monsoonal response to orbital forcing is a function of the climatological characteristics of the model: models whose African summer monsoon limit is farther north in the control simulation tend to demonstrate a larger northward extension of the monsoon limit in response to 6000 yr B.P. orbital forcing (Joussaume et al, 1999). One of the controls on the variation in the magnitude of the response between models appears to be the dynamical structures of regional subsidence and the subtropical anticyclone over northern Africa, which in turn are influenced by global-scale dynamics (de Noblet-Ducoudre et al, 2000) and are ultimately tied to the global-scale response to orographic and diabatic forcing. These differences in AGCM base-state dynamics apparently play a dominant role in determining the model response to changes in forcing (see, e.g., Masson et al, 1998) even when other components of the climate system, such as the ocean, are included in the simulation (Harrison et al, unpublished analyses).

Comparison of the simulated enhancement of the African monsoon with a variety of pale o-observations shows that the PMIP simulations (in common with earlier simulations of the response to orbital forcing) consistently underestimate both the northward shift in the monsoon belt shown by paleoenvironmental data and the magnitude of the precipitation required to produce the observed lake and vegetation changes in northern Africa. Comparisons of the spatial patterns in the simulated P — E fields with lake data from northern Africa (Yu and Harrison, 1996), for example, indicate that the PMIP simulations consistently underes timate the northward shift in the monsoon front. Similarly, when the changes in precipitation simulated in the PMIP experiments are used to drive an equilibrium vegetation model (BIOME3: Ilaxeltine and Prentice, 1996) in order to evaluate the likely response of vegetation to the simulated change in climate, the simulations consistently fail to reproduce the observed northward shift in the Sahara/Sahel boundary (Harrison et al., 1998). The precipitation required to generate the observed latitudinal distribution of grassland (steppe) vegetation in northern Africa at 6000 yr B.P. has been estimated using a combination of forward-modeling and inverse techniques. Joussaume et al. (1999) showed that the PMIP simulations underestimate the required precipitation at ca. 23°N by at least 100 mm (Fig. la), i.e., by ca. 50% of the minimum amount required to support grassland. When output from the PMIP experiments is used to simulate the extent of lakes across northern Africa using the HYDRA model (Coe, 1998; 2000), the observed area of Lake Chad (350,000 km2: Schneider, 1967; Pias, 1970) during the mid-Holocene is significantly underestimated by all of the models (Coe and Harrison, 2000). The failure of the PMIP simulations (in common with earlier AGCM simulations of the response to orbital forcing) to reproduce the observed changes in the African monsoon during the mid-Holocene provides strong support for the argument that the response to orbital forcing is mediated by feedbacks associated with changes in either the ocean or the land-surface.

3. Ocean Feedbacks on the Monsoon

Several studies indicate that ocean processes can produce feedbacks that enhance the monsoon response to 6000 yr B.P. orbital forcing (e.g., Kutzbach and Liu, 1997; Hewitt and Mitchell, 1998; Liu et al., 1999a,b; Otto-Bliesner, 1999; Texier et al., 2000; Braconnot et al., 2000). In the asynchronously coupled atmosphere-ocean experiments performed by Kutzbach and Liu (1997), for example, precipitation over northern Africa increases by 25% compared to simulations made with prescribed modern SSTs (Fig. lb). Monsoon enhancement is expressed by a northward shift of the monsoon front, increased precipitation and, at least in some cases, by an extension in the length of the monsoon season. A number of different processes appear to be involved. Radiative forcing alone, operating in a static column energy budget, cools the tropical Atlantic both north and south of the equator by as much as 0.5 °C in the spring (February through May) and raises the temperature by a comparable amount in autumn (August through November). The cooler ocean in the spring and early summer can enhance land-sea temperature contrast and thereby strengthen the African monsoon at onset (Hewitt and Mitchell, 1998). The fundamental changes in surface windflow associated with the orbitally forced enhancement of the southwesterly atmospheric inflow to West Africa can also act to decrease the normal north-easterly trades of the tropical North Atlantic, thereby reducing the total wind speed over the eastern Atlantic and consequently reducing evaporative cooling of the ocean surface. This change in the column energy budget preferentially increases SSTs north of the equator during the summer/autumn (Kutzbach and Liu, 1997). Most models show enhanced warming to the north of the equator, and some experiments even show a slight cooling south of the equator, during the summer/autumn months (Kutzbach and Liu, 1997; Hewitt and Mitchell, 1998; Otto-Bleisner, 1999; Braconnot et al., in press; Liu et al, 1999a). As shown by Hastenrath (1985) and others, a changed cross-equatorial SST gradient (warmer to the north) is of importance in producing a northward shift of the ITCZ in the North Atlantic and thereby increased advection of moisture into northern Africa from the west. Braconnot et al. (2000) have shown that increased south to north advection of heat within the upper ocean can also contribute to this dipole structure. Although most attention has been paid to the effects of orbitally forced ocean changes in the Atlantic on the African monsoon, it is possible that orbitally forced changes in the mean climate of the Pacific or Indian Oceans could, via teleconnection, influence the climate of northern Africa (Otto-Bleisner, 1999). In summary, models agree that there is a positive SST feedback effect on African monsoon precipitation in the mid-Holocene, although the relative importance of the various mechanisms that might contribute to this feedback requires further analysis. Furthermore, although the SST-driven enhancement is significant, the precipitation increase induced by the combined effect of orbital forcing and ocean feedbacks is not enough to support the observed grassland vegetation as far north as 23°N (Fig. lb).

All of the studies that have been conducted to date, whether the simulations are simple sensitivity tests made by prescribing stylized changes in ocean temperature in an AGCM (Texier et al, 2000) or with fully coupled ocean-atmosphere general circulation models (OAGCMs) (Hewitt and Mitchell, 1998; Otto-Bliesner, 1999; Braconnot et al, in press), show that ocean feedbacks enhance the northern African monsoon, although the relative importance of the mechanisms by which this enhancement is produced may vary from model to model. However, there is far less agreement about the role of ocean feedbacks on other monsoon systems. In coupled OAGCM simulations for both the early Holocene (11,000 yr B.P.: Liu et al, 1999a) and the mid-Holocene (6000 yr B.P.: Liu et al, 2000), the enhancement of the Indian monsoon is less than that produced by orbital changes alone (Fig. 2). In the 11,000 yr B.P. simulation, the reduction in precipitation due to ocean feedbacks is ca. 30% of the simulated increase due to the direct radiation effect. At 6000 yr B.P. the reduction is ca. 12% of the simulated increase due to the direct radiation effect. The negative ocean feedback on Indian monsoon rainfall appears to be caused by warming of the tropical Indian Ocean, which causes anomalous convergent flow over the Indian Ocean and hence increases precipitation over the ocean while decreasing precipitation over the Indian subcontinent. The simulated warming of the tropical Indian Ocean is partly a direct response to increased summer insolation, but is also partly due

Northern Africa 3

"O

FIGURE 2 Changes in the annual cycle of monthly mean precipitation (mm day-') over (a) northern Africa (land only, 5°W-35°E, 5-30°N) and (b) India (land only, 75-90°E, 10-25°N). The solid line (no circles) is the precipitation change forced by 11,000 year B.P. insolation (APnKl), obtained by holding SSTs at the modern (control) value. The solid line with circles is the precipitation change forced by the SST changes associated with 11,000 yr B.P. insolation (APSST). This SST feedback was isolated by differencing two 11,000 yr B.P. simulations: (1) an 11,000 yr B.P. simulation with SSTs specified from the results of an 11,000 yr B.P. coupled atmosphere-ocean simulation, and (2) an 11,000 yr B.P. simulation with SSTs specified from the coupled modern control simulation. All simulations for 11 ka used solar radiation values 11,000 yr B.P. based 11,000 yr B.P. orbital parameters, but other boundary conditions (atmospheric [CO,], ice sheets) were set at modern values. The dashed lines represent the standard deviation based on the internal variability of a long control simulation. The coupled atmosphere-ocean simulations used equilibrium asynchronous coupling (Liu et al., 1999a). The 11,000 year B.P. orbital forcing alone acts to enhance summer monsoon precipitation in both regions. The SST feedback is generally positive in the case of the African summer monsoon, with the main effects concentrated in spring and autumn, thereby increasing the length of the rainy season and the total precipitation. The SST feedback is generally negative in the case of the Indian summer monsoon.

to reduced evaporative cooling consequent on the weakening of the surface monsoon winds caused by the direct insolation response.

According to these AOGCM experiments, then, ocean feedbacks appear to somewhat damp the monsoon response to orbital forcing in India and increase the response in Africa. This difference in behavior may go some way to explaining why the observational evidence of monsoon changes in Africa is stronger than the response in India and over Asia more generally (though, admittedly, the amount of data from India is limited). A similar response (i.e., amplification of African monsoon precipitation and suppression of Indian monsoon precipitation) to ocean feedbacks is also shown in the experiments with prescribed SST changes by Texier et al. (2000). However, other OAGCM simulations (e.g., Hewitt and Mitchell, 1998; Braconnot et al, 2000) apparently do not demonstrate a comparable reduction in the strength of or-bitally induced enhancement of the Indian monsoon. Rather, in these simulations, ocean feedbacks further enhance the orbitally induced increase in the Indian monsoon in a fashion comparable to the enhancement of the African monsoon. The reasons for these differences between the response of the Indian monsoon to ocean feedbacks still need to be examined.

4. Land-Surface Feedbacks on the Monsoon

Changes in vegetation (and hence some soil characteristics, including organic matter content, and hence water-holding capacity and albedo) or the extent of surface water (lakes and wetlands) affect land-surface conditions through changing albedo (which determines the surface energy balance and hence surface heating), surface roughness (which affects both the water and energy fluxes between the land and the atmosphere), and moisture availability for recycling. The role of vegetation changes (and vegetation-induced soil changes) in enhancing orbitally induced changes in the monsoon circulation over northern Africa was originally studied by sensitivity experiments with stylized or quasi-realistic changes in vegetation and soil characteristics (e.g., Street-Perrott et al, 1990; Kutzbach et al., 1996; Brostrôm et al, 1998; Texier et al, 2000). These experiments produce a significant enhancement of the African monsoon compared to the effects of orbital-forcing alone. The vegetation-induced lowering of albedo increases surface heating (and hence amplifies the land-sea contrast, promoting increased advection of moisture into the continent). At the same time, the presence of vegetation and changes in the water-holding capacity of the soils leads to increased moisture recycling. In the Brostrôm et al (1998) experiments, the presence of vegetation (experiment RVS) leads to a substantial warming over northern Africa during the spring and early summer. As a result (Fig. lc), the onset of the monsoon occurs 2 months earlier than it does in response to orbital forcing alone (experiment R).

Northern Africa 3

"O

-

/Wad

"" N 1

Biogeochemical Cycle Grasslands
4 6

Month

Biogeochemical Cycle Grasslands

Boreal deciduous forest/woodland

Boreal evergreen mixed forest/woodland

Temperate/boreal mixed forest

Temperate conifererous forest

Temperate deciduous forest

Temperate broad-leaved evergreen forest

Tropical seasonal forest

Tropical rain forest

Tropical deciduous forest

Moist savannas

Dry savannas

Tall grassland

Short grassland

Xeric woodlands/scrub

Arid shrubland/steppe

Desert

Arctic/alpine tundra Polar desert

0 30E

Longitude

FIGURE 3 Simulated changes in precipitation and changes in hiome distributions at 6000 year B.P. relative to 0 year B.P., obtained with and without vegetation feedbacks: (a) annual cycle of monthly mean precipitation change (mm day"1) between simulations for 6000 yr B.P. and 0 year B.P. over northern Africa (land area, 11°W-34°E and 11-20°N) and (b) western portion of northern Africa (land area, ll°W-ll°E, ll-20°N). The solid line is the precipitation change forced by 6000 yr B.P. insolation with vegetation set at the modern control values. The dashed line is the precipitation change at 6000 yr B.P. forced by vegetation feedback. The vegetation feedback was isolated by differencing two 6000 yr B.P. simulations: (1) a 6000 yr B.P. simulation with vegetation specified from the results of a 6000 yr B.P. coupled atmosphere-vegetation simulation, and (2) a 6000 yr B.P. simulation with vegetation specified from the modern coupled control simulation (Doherty et al., 2000). The 6000 yr B.P. orbital forcing enhances the summer monsoon precipitation. Vegetation feedback enhances the precipitation in summer and autumn, with the largest effects in the western region. The overall effect is to lengthen the rainy season and increase the total precipitation, (c) Modern simulated biomes using BIOME 3 driven by present-day climatological values of the monthly mean annual cycle of temperature, precipitation, and solar radiation, and (d) 6000 yr B.P. simulated biomes using BIOME 3 driven by 6000 yr B.P. minus 0 yr B.P. differences in these climatic variables, taken from the coupled atmosphere-terrestrial vegetation model, and then combined with the present-day climatology (Doherty et al, 2000). (e) Observed distribution of vegetation during the mid-Holocene (from Jolly et al, 1998a,b).

The role of vegetation feedbacks has also been examined in a number of experiments with asynchronous coupling between an equilibrium vegetation model and an AGCM (e.g., Texier et al, 1997; Claussen and Gaylor, 1997; Pollard et al, 1998; de Noblet-Ducoudré et al, 2000) and, most recently, using the dynamically coupled GENESIS-IBIS atmosphere-vegetation model (Doherty et al, 2000; Fig. 3). A number of robust conclusions emerge from analyses of these simulations. Vegetation feedbacks increase precipitation during the peak of the monsoon season (July-August), by an amount comparable to the increase produced by orbital forcing alone but the absolute magnitude of this increase in precipitation is rather small. However, vegetation feedbacks have a significant impact on total precipitation by causing an extension of the monsoon season. Specifically, changcs in albedo caused by the prcscnce of vegetation lead to warming of the continent, thus enhancing land-sea contrast and increasing onshore advection, in spring and early summer and the onset of the monsoon therefore occurs 1-2 months (depending on the simulation) earlier than with orbital forcing alone (see, e.g., Brostrom et al, 1998; Texier et al, 2000; Doherty et al, 2000). Vegetation feedbacks also tend to prolong the monsoon into the autumn. Advection is relatively weak during the autumn, so that the extension of monsoon rains into the autumn appears to reflect enhanced moisture recycling in these simulations. In general, vegetation feedbacks appear to increase the importance of moisture recycling in maintaining the monsoon. The relative importance of the contributions of advection and recycling can be estimated from AP = A(P - E) + AE) (Braconnot et al, 1999), where A(P — E) represents the advection term and A£ the recycling term. In the Doherty et al (2000) experiments, the advection and recycling terms are of comparable magnitude in response to orbital changes (0.39 mm day-1 vs 0.42 mm day-1) but the increase in recycling (0.25 mm day-1) in response to vegetation feedbacks is larger than the increase in advection (0.18 mm day-1). The relative importance of advection and recycling in monsoon regions appears to vary between models, and the simulation of this partitioning has not been adequately evaluated. However, as models increasingly incorporate improved representations of vegetation and soil processes, there appears to be a significant increase in the importance of recycling relative to advection (see, e.g., Masson and Joussaume, 1997; Kleidon et al, 2000). Finally, vegetation feedbacks have a more significant role in enhancing monsoon precipitation in West Africa than in the eastern Sahara/Sahel. In the coupled GENESIS-IBIS simulations (Doherty et al, 2000), for example, vegetation feedbacks produce an increase in annual average precipitation of 0.63 mm day-1 comparable to the increase due to orbital forcing (0.66 mm day-1) over West Africa, whereas the comparable estimates over the eastern Sahara/Sahel region are 0.17 and 1.01 mm day-1. The regional differences in the impact of vegetation on the northern African monsoon are not so pronounced in e.g., the Texier et al (1997) simulations but are much larger in, e.g., the Claussen and Gaylor (1997) experiments.

Although most attention has been directed toward studying the possible feedbacks associated with vegetation, these were not the only landscape changes that would have affected land-surface characteristics. The mid-Holocene landscape of northern Africa was likely a mosaic of lakes, wetlands, and grasslands (Hoelzmann et al, 1998). Coe and Bonan (1997) used a model sensitivity experiment to illustrate that expanded lakes, specifically Palaeolake Chad and the lakes north of the Niger Bend, cause localized changes in circulation and some small enhancement of precipitation above and beyond that caused by orbital forcing alone. Prescribed additions of wetlands (with and without wetland vegetation) to a landscape of expanded grasslands also produced only small adjustments in large-scale precipitation (Carrington et al, in press). Brostrom et al (1998) analyses of the relative contribution of sequential changes in the extent of lakes (RVSL) and wetlands (RVSLW) compared to vegetation and vegetation-induced soil changes (RVS) across northern Africa confirm that the largest precipitation enhancement came from orbital forcing combined with changes in vegetation and soils; there was little or no additional enhancement of precipitation from wetlands or lakes (Fig. lc). It is perhaps not surprising that the relatively small areas covered by wetlands and lakes seem to produce only small-scale (local) climate perturbations. In some situations, P — E actually decreases over the expanded water surfaces because of increased evaporation, but the increase in precipitation over the surrounding catchment may be sufficient to maintain the lake and wetlands through local catchment-scale recycling. Nonetheless, these studies underscore the likelihood that only the relatively large-scale vegetation changes (and associated soil changes) interact to enhance regional precipitation, while the more localized areas of enhanced lakes and wetlands play only a minor role by comparison.

The amplification of the monsoon response to orbital forcing by land-surface feedbacks is apparently insufficient to explain the full observed expansion of the African monsoon. Thus, comparisons of the Brostrom et al (1998) simulations with benchmark data show that even when all possible land-surface feedbacks are taken into account (RVSLW), they are insufficient to produce the full observed northward expansion of grasslands into regions occupied today by desert (Fig. lc). There remains one further change in environmental conditions that might impact this question, however, namely the possibility that northward transport of excess runoff from the zone under the direct influence of the monsoon front could play a role. Simulations with the HYDRA model, driven by output from the GENESISIBIS coupled atmosphere-vegetation model (Figs. 4a, 4b) simulations for 6000 yr B.P., show that runoff is transported north from the zone of the monsoon front by a pale o-river network that is more extensive than today. Palaeodata (e.g., Hoelzmann et al, 1998) confirm the existence of a more extensive and active a)

FIGURE 4 Surface hydrology of northern Africa, simulated by HYDRA forced by runoff generated by the GF.NESIS2 AGCM coupled to the IBIS ecosystem model. HYDRA operates on a 5' X 5' (ca. 10 km) global grid to simulate the flow of water from land surfaces through a complex of rivers, lakes, and wetlands to the ocean or to inland drainage basins (such as closed lakes and interdunal wetlands), (a) Surface water area for 6000 year B.R simulated by HYDRA at the 5' X 5' horizontal (in black) showing pale o-lake Chad and other expanded pale o-lakes; and smoothened to 0.5° resolution (in pink) showing all regions with surface water area in excess of 10% of the 0.5° grid cell. The sum of the water areas at both resolutions is identical. (b) Change in annual mean discharge (in mm yr~') between simulations for 6000 and 0 yr B.P. over northern Africa. Only positive differences are shown. The colors represent those stream channels for which the discharge is increased in the 6000 yr B.P. experiment compared to modern. The results show the relatively large increase in runoff and stream flow in northern Africa (from 25-3000 mm yr~' increase). Greatest increases in discharge occur between about 15°-25°N and in Algeria. Paleostream channels occur throughout northern Africa where none exist today. Sheet-flow discharge across very flat terrain is also present in central Mali and in the northern basin of paleo-lake Chad. Simulated water areas of 6000 yr B.P. are shown in black.

FIGURE 4 Surface hydrology of northern Africa, simulated by HYDRA forced by runoff generated by the GF.NESIS2 AGCM coupled to the IBIS ecosystem model. HYDRA operates on a 5' X 5' (ca. 10 km) global grid to simulate the flow of water from land surfaces through a complex of rivers, lakes, and wetlands to the ocean or to inland drainage basins (such as closed lakes and interdunal wetlands), (a) Surface water area for 6000 year B.R simulated by HYDRA at the 5' X 5' horizontal (in black) showing pale o-lake Chad and other expanded pale o-lakes; and smoothened to 0.5° resolution (in pink) showing all regions with surface water area in excess of 10% of the 0.5° grid cell. The sum of the water areas at both resolutions is identical. (b) Change in annual mean discharge (in mm yr~') between simulations for 6000 and 0 yr B.P. over northern Africa. Only positive differences are shown. The colors represent those stream channels for which the discharge is increased in the 6000 yr B.P. experiment compared to modern. The results show the relatively large increase in runoff and stream flow in northern Africa (from 25-3000 mm yr~' increase). Greatest increases in discharge occur between about 15°-25°N and in Algeria. Paleostream channels occur throughout northern Africa where none exist today. Sheet-flow discharge across very flat terrain is also present in central Mali and in the northern basin of paleo-lake Chad. Simulated water areas of 6000 yr B.P. are shown in black.

river system in northern Africa during the mid-I Iolocene. It has been questioned whether the apparent expansion of steppe vegetation in northern Africa during the mid-I Iolocene reflected the presence of vegetation along water courses or in other well-watered locations, rather than the more general expansion implied by, e.g., the maps in Hoelzmann et al. (1998). If this were true, then the northward transport of excess runoff from the zone of the monsoon front, as shown in our simulations (Fig. 4b), would be a significant factor explaining the apparent mismatch between model simulations and observations. Although this suggestion cannot be entirely dismissed (given the limited number of sites documenting the mid-Holoccnc vegetation in northern Africa), it is unliklcy that the northward expansion of steppe vegetation was confined to water courses or other preferred locations. Vegetation records south of ca. 23°N, for example, do not contain pollen from any obligate desert species (I. C. Prentice and D. Jolly, unpublished analyses cited in Joussaume ct al., 1999). Some desert indicators would be expected to be present if the sites south of 23°N were representative of a landscape in which islands of steppe (in more well-watered sites) were set in a matrix of desert. On tire other hand, tire northward transport of excess runoff from further south could be responsible for the maintenance of wetlands and even lakes well beyond the limits of the monsoon front (Fig. 4a). Insofar as wetlands and lakes have an impact on local moisture recycling, as shown by Coe and Bonan (1997), Brostrom et al. (1998), and Carrington et al. (2000), any mechanism which increases their extent north of the monsoon front could potentially lead to feedbacks which might affect the spatial expression of monsoon enhancement. Additional experiments which take into account changes in the surface hydrological network, through either asynchronous or explicit coupling of a terrestrial hydrology model like HYDRA with an atmosphere-vegetation model (AVCGM), are required to test whether this mechanism might have an impact on the simulated monsoon climate.

5. Synergies between the Land, Ocean, and Atmosphere

Since neither land-surface nor ocean-surface feedbacks alone are sufficient to explain the observed expansion of the African monsoon during the mid-Holocene, synergistic feedbacks involving land-atmosphere-ocean interactions are likely to be involved (Ganopolski et al., 1998; Braconnot et al., 1999; Berger, in press). There have been only two attempts to examine this question. In simulations witlr an intermediate-complexity model, Ganopolski et al. (1998) showed tlrat vegetation feedbacks were more important than ocean feedbacks in the amplification of the African monsoon. This simulation may not be realistic, however, because the ocean model does not re solve the full dynamics of the equatorial ocean and the coarse spatial resolution of the atmospheric part of the model prevents the simulation of detailed regional monsoon changes. Braconnot et al. (1999) used asynchronous coupling between the IPSL OAGCM and an equilibrium biome model (BIOME1: Prentice et al., 1992) to examine the synergistic relationships between land and ocean feedbacks. This simulation makes it clear that incorporation of both kinds of feedbacks amplifies the orbitally induced enhancement of the African monsoon (Fig. Id). However, comparison of the amplification due either to the ocean alone (AO) or to vegetation alone (AV) in this model with the comparable feedback effects simulated by other models (e.g., respectively Fig. lb: Kutzbach and Liu, 1997; and Fig. lc: Brostrom et al., 1998) suggests that there is a strong model dependence in the magnitude of the simulated response. Thus, these experiments need to be revisited using a number of other coupled ocean-atmosphere-vegetation (OAVGCM) models to assess their robustness.

The omission of land-surface and ocean feedbacks and their possible interactions have been invoked to explain other mismatches between observations and climate simulations of the effects of orbital changes at 6000 yr B.P., including the degree of warming in the northern high latitudes (e.g., Foley et al, 1994; TEMPO, 1996; Texier et al, 1997) and the anomalous (i.e., opposite to the orbital forcing) winter warming in Europe (Prentice et al, 1998; Masson, 1998). Further improvements in the simulation of mid-Holocene climate changes will likely require the use of fully coupled OAVGCM models, which are now under development by several modeling groups.

6. The Role of Climate Variability

Changes in the mean climate state may be accompanied by changes in short-term (i.e., interannual to interdecadal) variability. The relationship between mean climate state and climate variability has not been extensively investigated, despite the fact that the impacts of climate change on earth systems may derive more from changes in variability than in the mean state. In large part this reflects the history of paleoclimate modeling. Simulations of Holocene paleoclimates using prescribed SSTs (as in PMIP) require only a short time (ca. 10-20 years) to reach equilibrium, and therefore attention naturally focused on describing only the mean climate since the simulation length was too short to permit studies of variability. As model simulations are necessarily extended with the advent of coupled OAGCMs and coupled atmosphere-vegetation general circulation models (AVGCMs), it has become both natural and important to focus on climate variability and the estimation of changes in climate variability.

Palaeoclimatic evidence suggests that short-term climate variability may have been substantially different from today during

Northern Africa China India

Northern Africa China India

4.6 5.0 5.4 5.8 6.2 6.6 6 6.5 7 7.5 8 8.5 9 1 2 3 4 5 6 7 8 9 10 11

Precipitation [mm day'1]

FIGURE 5 Frequency distributions of June-July-August (JJA) precipitation (mm day-1) at 0 year B.P. for (a) northern Africa (10°W-20°E and 10-20°N), (b) China (100- 120°E and 35-45°N), and (c) India (75-85°E and 25-30°N), for (d) northern Africa, (e) China and (f) India at 6000 year B.P. The simulated precipitation values were taken from the last 120 years of 150-year simulations with FOAM. The mean values (M) and standard deviation (SD) for each frequency distribution are shown. Especially in the case of the summer monsoon rains in northern Africa and China, the changes in the overall frequency distributions are very large. In northern Africa and China, the increases in mean JJA precipitation at 6000 year B.P. compared to 0 year B.P. pass two-tailed /-tests at the 95% level and the difference in the variances between 6000 yr B.P. and 0 yr B.P. pass an F- test at the 90% level. The largest variance (and the largest standard deviation) in northern Africa occurs at 6000 yr B.P., and the largest variance in China occurs at 0 yr B.P. The changes in mean and variance in the Indian region are not statistically significant.

4.6 5.0 5.4 5.8 6.2 6.6 6 6.5 7 7.5 8 8.5 9 1 2 3 4 5 6 7 8 9 10 11

Precipitation [mm day'1]

FIGURE 5 Frequency distributions of June-July-August (JJA) precipitation (mm day-1) at 0 year B.P. for (a) northern Africa (10°W-20°E and 10-20°N), (b) China (100- 120°E and 35-45°N), and (c) India (75-85°E and 25-30°N), for (d) northern Africa, (e) China and (f) India at 6000 year B.P. The simulated precipitation values were taken from the last 120 years of 150-year simulations with FOAM. The mean values (M) and standard deviation (SD) for each frequency distribution are shown. Especially in the case of the summer monsoon rains in northern Africa and China, the changes in the overall frequency distributions are very large. In northern Africa and China, the increases in mean JJA precipitation at 6000 year B.P. compared to 0 year B.P. pass two-tailed /-tests at the 95% level and the difference in the variances between 6000 yr B.P. and 0 yr B.P. pass an F- test at the 90% level. The largest variance (and the largest standard deviation) in northern Africa occurs at 6000 yr B.P., and the largest variance in China occurs at 0 yr B.P. The changes in mean and variance in the Indian region are not statistically significant.

the early to nrid-Holocene. Time-series of archaeological deposits in northern Peru (Sandweiss et al., 1996) and clastic deposits in an Andean lake in Ecuador (Rodbell et al., 1999) indicate less severe flooding events along the west coast of tropical South America during the early Holocene. The 5lsO records from the Sajama ice core in the tropical Andes also show less variability during the early Holocene than in the later Holocene (Thompson et al., 1998; Thompson, 2000). Records of fires in Australia (McGlone et al., 1992) and isotopic records from fossil corals in the western tropical Pacific (Gagan et al, 1998) have been interpreted as showing that monsoon rainfall was less variable during the first half of the Holocene than today.

Liu et al (1999b), in simulations examining the response of the Fast Ocean Atmosphere Model (FOAM: Jacob, 1998) to 11,000 yr B.P. orbital forcing, have shown that El Niño variability is reduced by ca. 20% and the spectral bandwidth of El Niño changes from the broad (3-10 year) peak characteristic of the modern simulations to a narrower peak (2-3 year) at 11,000 yr B.P. This reduction in the variability appears to be associated with the simulated increase in the Indian summer monsoon. There are, however, a number of mechanisms through which changes in the Indian summer monsoon circulation appear to impact the El Niño signal. First, the enhanced Indian monsoon strengthens the deep convection in the eastern Indian Ocean and western Pacific warm pool, increasing the strength of the easterly trades (by ca. 1 m s-1), and hence increasing the upwelling (and cooling) in the central and eastern Pacific. Ocean feedbacks further enhance this wind-driven cooling. In these simulations, the combination of forcing by changes in the Indian monsoon and positive ocean feedbacks leads to an SST cooling in the eastern Pacific of ca. 0.5°C in May-June-July. This cooling tends to suppress the growth of warm El Niño events during the Southern Hemisphere spring and therefore reduces their final amplitude later in the year. In the Liu et al (1999b) simulations, the annual mean trades are stronger (by ca. 2 s-1) than in the control simulation. This results in enhanced upwelling and therefore enhanced SST cooling throughout the year and provides a further mechanism for reducing El Niño variability. More recent simulations with the FOAM model (Liu et al, 2000) show that reduced El Niño variability is also produced in response to 6000 yr B.P. orbital forcing.

Otto-Bliesner (1999) found that the teleconnections relating the patterns of Pacific ENSO to Sahelian rainfall in the 6000 yr B.P. experiment are different from those in the control (modern) simulation. We have therefore reexamined the FOAM simulations specifically to determine whether there are changes in precipita-

tion variability associated with the mid-Holocene enhancement of the Afro-Asian monsoons. In northern Africa, the increased summer precipitation at 6000 yr B.P. is associated with a significant increase in interannual precipitation variability (Fig. 5). In India, the increase in mean precipitation during the monsoon season is also accompanied by increased interannual variability. However, the reverse is true in China. In our simulations, the enhancement of the Pacific monsoon leads to increased summer précipitation and reduced interannual variability (Fig. 5). The causes of these regional differences in the relationship between mean climate and climate variability required further analysis. However, it is clear that the increased/decreased variability at 6000 yr B.P. has the potential to significantly impact the regional paleoenvironmental response to the change in mean climate. For example, in relatively arid environments with comparable mean annual rainfall, sparse shrub or open woodland vegetation tends to be favored in regions with high interannual variability whereas steppe grasslands occur where the variability is less.

7. Final Remarks

The ability to correctly simulate past climates bears directly on whether we can confidently predict future climates (Joussaume, 1999). Comparisons of climate experiments with paleoenvironmental data have clearly demonstrated that the observed large changes in climate during the mid-Holocene (or at the LGM: see discussion in Kohfeld and Harrison, 2000) cannot be simulated without explicitly considering the feedbacks associated with the ocean, vegetation, and other components of the land-surface. There are a number of other feedbacks (e.g., radiative forcing by mineral aerosols at the LGM: Harrison et ai, in press; Claquin et ai, submitted) that are potentially important. COHMAP results (Kutzbach et al., 1998) indicating that the apparent mismatches between observed and simulated climate changes during the transition from glacial to interglacial conditions are greater than at either the LGM or the mid-Holocene suggest that the incorporation of these feedbacks may be even more important in attempts to simulate times of rapid climate change when there is a strong disequilibrium between insolation and other conditions. Thus, simulations of potential future climate changes need to be made using fully coupled ocean-atmosphere-biosphere models, taking into account the potential additional impact of changes in surface conditions on atmospheric aerosols.

References

Berger, A. ( 1978). Long-term variation of daily insolation and Quaternary climatic changes./. Anuos. Sci. 35, 2362-2367. Berger, A. (1988). Milankovitch theory and climate. Rev. Geophys. 26, 624-657.

Berger, A. (in press). The role of CO, and of the geosphere-biosphere interactions during the Milankovitch-forced glacial-interglacial cycles. In "Workshop on Geosphere-Biosphere Interactions and Climate." Pontifical Academy of Sciences, Vatican City.

Berger, A. and Loutre, M.-F. (1991). Insolation values for the climate of the last million years. Quaternary Sci. Rev. 10, 297-317.

Braconnot, P., Marti, O., and Joussaume, S. (1997). Adjustments and feedbacks in a global coupled ocean-atmosphere model. Climate Dynamics. 13, 507-519.

Braconnot, P., Joussaume, S., Marti, O., and de Noblet, N. (1999). Synergistic feedbacks from ocean and vegetation on the African monsoon response to mid-Holocene insolation. Geophys. Res. Lett. 16, 2481-2484.

Braconnot, P., Marti, O, Joussaume, S., and Leclainche, Y. (2000). Ocean feedback in response to 6 kyear BP insolation. /. Climate. 13, 1537-1553.

Broccoli, A. J. and Marciniak, E. P. (1996). Comparing simulated glacial climate and paleodata: a reexamination. Paleoceanography 11, 3-14.

Brostrom, A., Coe, M., Harrison, S. P., Gallimore, R., Kutzbach, J. E., Foley, J., Prentice, I. C., and Behling, P. (1998). Land surface feedbacks and pa-leomonsoons in northern Africa. Geophys. Res. Lett. 25, 3615- 3618.

Carrington, D., Gallimore, R. G., and Kutzbach, J. E. (2000). Climate sensitivity to wetlands and wetland vegetation in mid-Holocene north Africa. Climate Dynamics. 17, 151-157.

Claquin, T., Roelandt, C., Kohfeld, K. E., Harrison, S. P., Prentice, I. C., Balkanski, Y„ Bergametti, G„ Hansson, M„ Mahowald, N„ Rodhe, N„ and Schulz, M. (submitted). Radiative forcing of climate by ice-age dust. Nature.

Claussen, M. and Gaylor, V. (1997). The greening of the Sahara during the mid-Holocene: results of an interactive atmosphere-biome model. Global Ecol. Biogeography Lett. 6, 369-377.

Coe, M. T. (1998). A linked global model of terrestrial hydrologic processes: simulation of modern rivers, lakes, and wetlands. /. Geophys. Res. 103, 8885-8899.

Coe, M. T. (2000). Modeling terrestrial hydrological systems at the continental scale: testing the accuracy of an atmospheric GCM. /. Climate 13,686-704.

Coe, M. T. and Bonan, G. B. (1997). Feedbacks between climate and surface water in Northern Africa during the Middle Holocene. /. Geophys. Res. 102, 11087.

Coe, M. T. and Harrison, S. P. (2000). A comparison of the simulated surface water area in norhtern Africa for the 6000 year B.P. PMIP experiments. In "Paleoclimate Modeling Intercomparison Project, Proceedings of the Third Conference." (P. Braconnot, Ed.), WCRP.

COHMAP Members (1988). Climatic changes of the last 18,000 years: observations and model simulations. Science 241, 1043-1052.

Crowley, T. J. and Baum, S. K. (1997). Effect of vegetation on an ice-age climate model simulation./. Geophys. Res. 102, 16463-16480.

Doherty, R„ Kutzbach, J. E„ Foley, J.> and Pollard, D. (2000). Fully-coupled climate/dynamical vegetation model simulations over northern Africa during the mid-Holocene. Climate Dynamics. 16,561 -573.

Dong, B„ Valdes, P. J„ and Hall, N. M. J. (1996). The changes in mon-soonal climates due to Earth's orbital perturbations and ice age boundary conditions. Palaeoclimates: Data and Modelling 1,203-240.

Foley, J. A., Kutzbach, J. E., Coe, M. T„ and Levis, S. (1994). Climate and vegetation feedbacks during the mid-Holocene. Nature 371,52-54.

Foley, J. A., Levis, S., Prentice, I. C., Pollard, D., and Thompson, S ,L. (1998). Coupling dynamic models of climate and vegetation. Global Change Biol. 4, 561-579.

Foley, J. A., Levis, S„ Costa, M. H„ Doherty, R., Kutzbach, J. E„ and Pollard, D. (in press). Vegetation as an interactive part of global climate models. Ecol. Applications.

Cagan, M. K., Aylille, L., Hopley, D., Cali, I. A., Mortimer, G. E., Chappell, J., McCulloch, M. T., and Head, M J. (1998). Temperature and surface-ocean water balance of the Mid-Holocene tropical Western Pacific. Science 279, 1014-1018.

Gallee, J.-F., van Ypersele, J. P., Fichcfct, T., Marsiat, I., Tricot, C., and Berger, A. (1992). Simulation of the last glacial cycle by a coupled, sec-torially averaged climate-ice sheet model. Part 2. Response to insolation and CO, variations. /. Geophys. Res. 97, 15713-15740.

Ganopolski, A., Kubatzki, C., Claussen, M„ Brovkin, V., and Petoukhov, V. (1998). The influence of vegetation-atmosphere-ocean interaction on climate during the mid-Holocene. Science 280, 1916- 1919.

Gasse, F. and van Campo, E. (1994). Abrupt post-glacial climate events in West Asia and North Africa monsoon domains. Earth and Planetary Sci. Letl. 126,435-456.

Gent, P. R„ Bryan, F. O., Danabasoglu, G„ Doney, S.C., Holland, W. R„ Large, W. G., and McWilliams, J. C. (1998). The NCAR Climate System Model global ocean component. J. Climate 11,1287- 1306.

Hall, N. M. and Valdes, P. J. (1997). A GCM simulation of the climate 6000 years ago. J. Climate 10, 3-17.

Harrison, S. P., Jolly, D„ Laarif, F„ Abe-Ouchi, A., Dong, B„ Herterich, K„ Hewitt, C„ Joussaume, S., Kutzbach, J. E„ Mitchell, J., de Noblet, N„ and Valdes, P. (1998). Intercomparison of simulated global vegetation distribution in response to 6 kyr B.P. orbital forcing. /. Climate 11, 2721-2742.

Harrison, S. P., Kohfeld, K. E., Roelandt, C., and Claquin, T. (in press). The role of dust in climate today, at the last glacial maximum and in the future. Earth Sci. Rev.

Ilastenrath, S. (1985). "Climate and Circulation in the Tropics." Reidel, Norwell, MA., 455 p.

Haxeltine, A. and Prentice, I. C. (1996). BIOME3: An equilibrium terrestrial biosphere model based on ecophysiological constraints, resource availability, and competition among plant functional types. Global Bio-geochern. Cycles 10, 693-710.

Hays, J. D., Imbrie, J., and Shackleton, N. J. (1976). Variations in the earth's orbit: pacemakers of the ice ages. Science 194, 1121 -1132.

Hewitt, C. D. and Mitchell, J. F. B. (1996). GCM simulations of the climate of 6 k year BP: mean changes and interdecadal variability. J. Climate 9, 3505-3529.

Hewitt, C. D. and Mitchell, J. F. B. (1998). A fully coupled GCM simulation of the climate of the mid-Holocene. Geophys. Res. Lett. 25, 361-364.

Hoelzmann, P., Jolly, D„ Harrison, S. P., Laarif, F„ Bonnefille, R., and Pachur, H.-J. (1998). Mid-Holocene land-surface conditions in northern Africa and the Arabian peninsula: a data set for the analysis of bio-geophysical feedbacks in the climate system. Global Biogeochem. Cycles 12,35-51.

Imbrie, J. (1985). A theoretical framework for the Pleistocene Ice Ages. J. Geol. Soc. (London) 142, 417-432.

Jacob, R. L. (1998). Low frequency variability in a simulated atmosphere ocean system. PhD Thesis, University of Wisconsin-Madison, 155 p.

Johns, T. C, Carnell, R. E„ Crossley, J. E, Gregory, J. M„ Mitchell, J. F. B„ Senior, C. A., Tett, S. F. B„ and Wood, R. A. (1997). The second Iladley Centre coupled ocean-atmosphere GCM: model description, spinup and validation. Climate Dynamics 13, 103-134.

Jolly, D„ Harrison, S. P., Damnati, D. and Bonnefille, R. (1998a). Simulated climate and biomes of Africa during the Late Quaternary: comparison with pollen and lake status data. Quaternary Sci. Rev. 17, 629-657.

Jolly, D„ Prentice, I. C., Bonnefille, R„ Ballouche, A., Bengo, M„ Brénac, P.,

Buchet, G., Burney, D„ Cazet, J.-P., Cheddadi, R„ Edorh, T., Elenga, H„ Elmoutaki, S., Guiot, J., Laarif, F„ Lamb, H., Lezine, A. M„ Maley, J., Mbenza, M„ Peyearon, O., Reille, M„ Reynaud-Ferrara, I., Riollet, G. Ritchie, J. C., Roche, E„ Scott, L„ Ssemmanda, I., Staka, H., Umer, M„ Van Campo, E., Vilimumbala, S„ Vincens, A., and Waller, M. (1998b). Biome reconstruction from pollen and plant macrofossil data from Africa and the Arabian peninsula at 0 and 6 ka./. Biogeography 25, 1007-1027.

Joussaume, S. (1999). Modeling extreme climates of the past 20,000 years with general circulation models. In "Modeling the Earth's Climate and its Variability." (W. R. Holland, S. Joussaume, and F. David, Eds.), pp. 527-565, Elsevier, Amsterdam, Netherlands.

Joussaume, S. and Taylor, K. E. (1995). Status of the Paleoclimate Modeling Intercomparison Project (PMIP). In "Proceedings of the First International AMIP Scientific Conference, 15-19 May 1995." (W. L. Gates, Ed.), p. 532. Monterey, CA.

Joussaume, S. and Taylor, K. E. (2000). The Paleoclimate Modeling Intercomparison Project. In "Paleoclimate Modeling Intercomparison Project, Proceedings of the Third Conference." ( P. Braconnot, Ed.),WCRP.

Joussaume, S., Taylor, K. E., Braconnot, P., Mitchell, J. F. B., Kutzbach, J. E., Harrison, S. P., Prentice, I. C„ Broccoli, A. J., Abe-Ouchi, A., Bartlein, P. J., Bonfils, C„ Dong, B„ Guiot, J„ Herterich, K„ Hewitt, C. D„ Jolly, D„ Kim, J. W„ Kislov, A., Kitoh, A., Loutre, M. F„ Masson, V., McAvaney, B„ McFarlane, N„ deNoblet, N„ Peltier, W. R„ Peterschmitt, J. Y„ Pollard, D„ Rind, D., Royer, J. F., Schlesinger, M. E., Syktus, J., Thompson, S., Valdes, P., Vettoretti, G„ Webb, R. S„ and Wyputta, U. (1999). Monsoon changes for 6000 years ago: results of 18 simulations from the Paleoclimate Modeling Intercomparison Project (PMIP). Geophys. Res. Lett. 26, 859-862.

Kleidon, A., Fraedrich, K„ and Heimann, M. (2000). A green planet versus a desert world: estimating the maximum effect of vegetation on the land surface climate. Climatic Change 44, 471-493.

Knorr, W. and Heimann, M. (1995). Impact of drought stress and other factors on seasonal land biosphere C02 exchange studied through an atmospheric tracer transport model. Tellus Series B-Chem. Phys. Mete-orol. 47, 471-489.

Kohfeld, K. E. and Harrison, S. P. (2000). How well can we simulate past climates? Evaluating earth system models using global paleoenviron-mental datasets. Quaternary Sci. Rev. 19, 321-346.

Kutzbach, J. E. (1981). Monsoon climate of the eartly Holocene: climate experiment with the earth's orbital parameters for 9000 years ago. Sri-em: 214, 59-61.

Kutzbach, J. E. and Gallimore, R. G. (1998). Sensitivity of a coupled atmosphere/mixed layer ocean model to changes in orbital forcing at 9000 years B.P./. Geophys. Res. 93, 803-821.

Kutzbach, J. E„ and Guetter, P. J. (1986). The influence of changing orbital parameters and surface boundary conditions on climate simulations for the past 18000 years. /. Atmos. Sci. 4, 1726-1759.

Kutzbach, J. E. and Liu, Z. (1997). Response of the African monsoon to orbital forcing and ocean feedbacks in the Middle Holocene. Science 278, 440-443.

Kutzbach, J. E. and Otto-Bleisner, B. L. (1982). The sensitivity of the African-Asian monsoonal climate to orbital parameter changes for 9000 yr B.P. in a low-resolution general circulation model. /. Atmos. Sci. 39, 1177-1188.

Kutzbach, J. E. and Street-Perrott, F. A. (1985). Milankovitch forcing of fluctuations in the level of tropical lakes from 18 to 0 kyear B.P. Nature 317, 130-134.

Kutzbach, J. E„ Guetter, P. J., Behling, P. J., and Selin, R. (1993). Simulated climatic changes: Results of the COHMAP climate-model experi ments. In "Global Climates since the Last Glacial Maximum." (H. E. Wright Jr., J. E. Kutzbach, T. Webb III, W. F. Ruddiman, F. A. Street-Per-rott, and P. J. Bartlein, Eds.), pp. 24-93. University of Minnesota Press, Minneapolis.

Kutzbach, J. E., Bonan, G., Foley, J., and Harrison, S. P. (1996). Vegetation and soil feedbacks on the response of the African monsoon to forcing in the early to middle Holoccnc. Nature 384, 623-626.

Kutzbach, J. E., Gallimore, R., Harrison, S. P., Behling, P., Selin, R., and Laarif, F. (1998). Climate and biome simulations for the past 21,000 years. Quaternary Sci. Rev. 17, 473-506.

Levis, S„ Foley, J. A., and Pollard, D. ( 1999a). CO,, climate, and vegetation feedbacks at the Last Glacial Maximum. /. Geophys. Res. 104, 31191-31 198.

Levis, S„ Foley, J. A., and Pollard, D. (1999b). Potential high-latitude vegetation feedbacks on CO,-induced climate change. Geophys. Res. Lett. 26, 747-750.

Liu, Z„ Gallimore, R„ Kutzbach, J. E„ Xu, W„ Golubev, Y„ Behling, P., and Selin, R. (1999a). Modeling long-term climate changes with equilibrium asynchronous coupling. Climate Dynamics 15, 325-340.

Liu, Z„ Jacob, R., Kutzbach, J. E„ Harrison, S. P., and Anderson, J. (1999b). Monsoon impact on EI Niño in the early Holocene. PAGES Newslett. 7, 16-17.

Liu, Z„ Kutzbach, J. E„ and Wu, L. (2000). Mod

0 0

Post a comment