Transient simulation of the last glacial inception. Part II: sensitivity and feedback analysis
- First Online:
- Cite this article as:
- Calov, R., Ganopolski, A., Petoukhov, V. et al. Climate Dynamics (2005) 24: 563. doi:10.1007/s00382-005-0008-5
- 127 Views
The sensitivity of the last glacial-inception (around 115 kyr BP, 115,000 years before present) to different feedback mechanisms has been analysed by using the Earth system model of intermediate complexity CLIMBER-2. CLIMBER-2 includes dynamic modules of the atmosphere, ocean, terrestrial biosphere and inland ice, the last of which was added recently by utilising the three-dimensonal polythermal ice-sheet model SICOPOLIS. We performed a set of transient experiments starting at the middle of the Eemiam interglacial and ran the model for 26,000 years with time-dependent orbital forcing and observed changes in atmospheric CO2 concentration (CO2 forcing). The role of vegetation and ocean feedback, CO2 forcing, mineral dust, thermohaline circulation and orbital insolation were closely investigated. In our model, glacial inception, as a bifurcation in the climate system, appears in nearly all sensitivity runs including a run with constant atmospheric CO2 concentration of 280 ppmv, a typical interglacial value, and simulations with prescribed present-day sea-surface temperatures or vegetation cover—although the rate of the growth of ice-sheets growth is smaller than in the case of the fully interactive model. Only if we run the fully interactive model with constant present-day insolation and apply present-day CO2 forcing does no glacial inception appear at all. This implies that, within our model, the orbital forcing alone is sufficient to trigger the interglacial–glacial transition, while vegetation, ocean and atmospheric CO2 concentration only provide additional, although important, positive feedbacks. In addition, we found that possible reorganisations of the thermohaline circulation influence the distribution of inland ice.
Nowadays, it is generally accepted that the reduction of boreal summer insolation drives the build-up of the ice masses. In the first part of this paper (Calov et al. 2005), the capability of our model to simulate the last glacial inception with prescribed orbital and CO2 forcing was demonstrated. Further, it was shown that the snow albedo feedback is the primary amplifier of the orbital forcing. Here, we analyse the role of other potentially important feedbacks and mechanisms related to terrestrial vegetation, atmospheric CO2 concentration, mineral dust and changes in the thermohaline circulation.
Among several other sensitivity studies in their paper, Galleé et al. (1992) demonstrated the temperature-lowering effect of expanded tundra (at the expense of boreal forest) for LGM (last glacial maximum, 21 kyr BP) climate conditions with the zonally averaged LLN climate model. Later, Gallimore and Kutzbach (1996) compared time-slices at 115,000 kyr BP from the NCAR atmospheric GCM (general circulation model) CCM1 coupled to a mixed layer ocean. They showed that increased albedo representing expanded tundra area considerably lowers the Northern Hemisphere’s temperature leading to a more persistent snow cover. DeNoblet et al. (1996) coupled interactively a biome (Prentice et al. 1992) global model with the LMD 5.3 GCM (fixed ocean), pinpointing with subsequent iterations of the biome model and the LMD-5.3 model the feedback mechanism through lowered temperature influencing the vegetation types and vice versa. Pollard and Thompson (1997) coupled the GENESIS2 atmospheric GCM, which includes mixed layer ocean and land surface, asynchronously to a vertically integrated ice-sheet model to simulate glacial inception with fixed 116 kyr orbital insolation and constant atmospheric CO2 concentration. Despite these very favourable inception conditions, their North American inland-ice cover does not extend far enough to the south. A feedback analysis including vegetation using the Earth system model of intermediate complexity MoBidiC was carried out by Crucifix and Loutre (2002). But their model excludes the feedback through dynamic ice sheets.
In Khodri et al. (2001, 2003) it was shown, using the IPSL-CM2 atmospheric-ocean GCM, that an increased meridional atmospheric moisture transport and changes in the ocean circulation might facilitate the growth of ice sheets. We add to their results our investigation of a hypothesised southward shift of the convection sites in the North Atlantic and the implications of this for glacial inception, focusing especially on Scandinavia. The role of the thermohaline circulation of the North Atlantic has also been investigated by Wang and Mysak (2002) with another Earth system model of intermediate complexity.
The common picture of CO2 reconstructions is of a more or less rapid drop of about 40 ppmv between 115 kyrs BP and 105 kyrs BP with a duration in the range from 5,000 years to 9,000 years (Barnola et al. 1987; Petit et al. 1999; Fischer et al. 1999). Futhermore, the atmospheric CO2 concentration in the palaeo-records starts to drop after the boreal summer insolation decreases. More than 10 years ago, Verbitsky and Oglesby (1992) analysed the effect of a reduction in the atmospheric CO2 concentration on the initiation of the ice sheets. They derived snowfall fields with the NCAR GCM1 by assuming an initial snow cover of 10 m for the Northern Hemisphere. This snowfall was used as input for a vertically integrated ice-sheet model, which then calculated the distribution of the ice sheets in the Northern Hemisphere. Verbitsky and Oglesby (1992) found in a simulation with a very low atmospheric CO2 concentration of 100 ppmv a very low ice volume compared to the sea-level proxies. Although neither the change in orbital insolation nor the feedback of the ice sheets into the climate system are accounted for in their computations, they draw the conclusion that the atmospheric CO2 concentration alone is not the “major player” in the initiation of glaciation during the Pleistocene. Vettoretti and Peltier (2004) investigated the importance of different orbital parameters and the CO2 forcing with the CCCma AGCM2, which includes a mixed-layer ocean model. They concluded after a detailed analysis that for glacial inception the eccentricity-precession forcing has approximately the same magnitude as the CO2 forcing. On the other hand, Vettoretti and Peltier concede that their model did not match the observed rate of sea-level drop during the last glacial inception, because mechanisms such as vegetation feedback are still missing in their model.
The potential importance of mineral dust for the ablation of ice sheets during the last glacial cycle and, in particular, for the glacial termination has been discussed by Peltier and Marshall (1995). In their simulations with an ice-sheet model coupled to an energy-balance model they mimicked the increase of ablation through dust by increasing ablation artificially by a factor of three. In Calov et al. (2005) it was shown with a physically based model for the dependence of snow albedo on dust (Warren and Wiscombe 1980) that during the latter stage of a glacial cycle, when the volume of the ice sheet is large enough, an increase of atmospheric dust could provide a negative feedback preventing further ice-sheet growth. Whether this is applicable to the initial stage of glaciation or only to the latter stages is not clear, however. Nonetheless, evidence has been found for an increased dustiness during glacial inception in the north-western Pacific (Porter 2001).
We start with a short summary of the model description (see Calov et al. 2005 for further details) and specification of the experimental set-up (Sect. 2). The role of mineral dust for glacial inception is discussed in Sect. 3. Simulated climatic and glaciological fields in the course of glacial inception are discussed in Sect. 4. A detailed inspection and analysis of the role of terrestrial vegetation and the ocean follow in Sect. 5. In addition, the role of atmospheric CO2 concentration and orbital insolation will be studied in Sect. 6. The implications of possible changes of the thermohaline circulation in the Atlantic during last glacial inception are examined closely in Sect. 7. We conclude with a discussion of our findings (Sect. 8) and a summary of our results together with concluding remarks (Sect. 9).
2 The model and the experimental set-up
2.1 The model
The Earth system model of intermediate complexity CLIMBER-2 (Petoukhov et al. 2000) encompasses the atmosphere, ocean, terrestrial biosphere and ice sheets. The latter have been recently integrated into the model by inclusion of the polythermal ice-sheet model SICOPOLIS (Greve 1997a, b) and a newly developed bi-directional coupling module named SEMI (Surface Energy and Mass balance Interface). A detailed description of the models and the coupling is provided in Calov et al. (2005).
CLIMBER-2 is based on a statistical-dynamical atmosphere model with a low spatial resolution (7 longitudinal sectors and 18 latitudinal belts), and the zonally averaged ocean module accounts for three separate ocean basins, the Atlantic, Indian and Pacific. Atmosphere and ocean interact through the surface fluxes of heat, fresh water and momentum.
SICOPOLIS simulates the time-dependent extent, thickness, velocity, temperature, water content and age for grounded ice sheets. Furthermore, possible basal layers of temperate ice are detected by fulfilling the Stefan-type conditions at the transition surface between cold and temperate ice. The bedrock responds to the load of the ice through the buoyancy forces of the asthenosphere.
The coupling module (SEMI) calculates the energy and mass balance on the fine grid of SICOPOLIS—the Northern Hemisphere is resolved by a grid spacing of 1.5° for the longitude and 0.75° for the latitude—using interpolated atmospheric characteristics (air temperature and humidity, long-wave and short-wave radiation, precipitation) from the CLIMBER-2 model and accounting for the orography on the fine grid. In turn, SICOPOLIS provides CLIMBER-2 with the temporal change of orography and fractions of land and glaciers.
2.2 The experimental set-up
Notation of the experiments
Atmosphere, ocean, vegetation and ice sheets
Atmosphere, ocean, vegetation, ice sheets and dust
Atmosphere, ocean, ice sheets and dust. Vegetation fixed
Atmosphere, vegetation, ice sheets and dust. Ocean fixed
Atmosphere, ice sheets and dust. Ocean and vegetation fixed
Atmosphere, ocean, vegetation, ice sheets and dust. CO2 fixed
Atmosphere, ocean, vegetation, ice sheets and dust. Thermohaline circulation changed by a fresh-water perturbation in the Nordic Seas
Atmosphere, ocean, vegetation, ice sheets and dust. Isolation fixed
Steady-state run with constant pre-industrial conditions
where t is the model time. DPD and DLGM are the present-day dust-deposition rate and the LGM dust-deposition rate, respectively. The weights in Eq. 1 are determined through w(t)=V(t)/VLGM. We compute the evolution of the inland-ice volume of the Northern Hemisphere V with the ice-sheet model SICOPOLIS. VLGM=48 million km3 is an empirical estimate of the LGM Northern Hemisphere’s inland-ice volume corresponding roughly to 110 m sea level drop (approximately the LGM volume of the Northern Hemisphere ice sheets excluding Greenland). The dust-deposition rates DPD and DLGM, interpolated to the grid of the ice-sheet model, are taken from Mahowald et al. (1999). Mahowald et al. (1999) computed the dust-deposition rates with an entrainment and transport model employing wind and precipitation fields of a GCM and applying dust sources which are modelled with the BIOME3 terrestrial biosphere model. Equation 1 is a heuristic approach to parameterise the approximate time dependence of the dust-deposition rate during last glacial inception. We argue as follows. With increased ice volume the sea level drops and the continental shelf regions are exposed to the air. These regions can serve as additional dust sources. Therefore, it appears to be reasonable to use the ice volume as a parameter. It is natural to assume a linear dependence as a first approach.
The experiments where the ice volume is dependent on dust-deposition rate are denoted by inclusion of the letter “D”, e.g. experiment AOVID. Further experiments were with fixed present-day vegetation cover (AOID) or with “fixed ocean” (seasonal variations of sea-surface temperature and sea-ice cover were held on their present-day values, AVID) or with fixed atmospheric CO2 concentration at present-day level (experiment AOVID_280). Additionally, we introduced an artificial fresh-water perturbation to the Atlantic thermohaline circulation, which leads to a southward shift of the North Atlantic convection site (experiment AOVID_THC).
3 The role of mineral dust
In this section, we assess the potential role of the dust impact on snow albedo for the Northern Hemisphere’s glaciation by comparison of experiments AOVID and AOVI. The physical basis of the dependence of the albedo of snow on mineral dust has been pointed out in Calov et al. (2005).
A large part of the differences between results obtained in experiments AOVID and AOVI can be explained by the direct impact of the dust distribution on the albedo of snow; with a higher dust-deposition rate the albedo of snow is reduced, and the snowmelt is amplified—recall that the dust-deposition rate in experiment AOVID is the weighted mean (Eq. 1) between the present-day and the LGM values. Regions like north-eastern Eurasia and northern Alaska, where the dust-deposition rate during LGM is high, show little or even no glaciation in experiment AOVID. Regions with a moderate LGM dust-deposition rate mostly show a comparable inland-ice cover in experiments AOVID and AOVI (see the Laurentide, the Cordilleran and the Fennoscandian ice sheets in Fig. 3). The glaciation in north-eastern Eurasia is highly sensitive to dust in our simulation, because snowfall is low there. The low snowfall enhances the sensitivity of the ice sheets to mineral dust, because the dust concentration in snow for a given dust-deposition rate is inversely proportional to the snowfall. In addition, the snow albedo is a function of the dust concentration in the snow. Therefore, changes in the dust-deposition rate have the strongest impact on the surface albedo in areas with low snowfall. In experiment AOVI, where the dust-deposition rate is held fixed to its modern value, the snow-albedo feedback (Calov et al. 2005) causes, under reduced boreal summer insolation, a rapid expansion of inland-ice cover in north-eastern Eurasia, similar to that in northern North America (see Fig. 3e). The big ice sheet in north-eastern Eurasia is the reason why the sea level curve of experiment AOVI has little tendency to level off (Fig. 2). In experiment AOVID, there is only minor glaciation in north-eastern Eurasia, because the glacial dust-deposition rate there is much higher than the modern one, which leads to stronger melting in the ablation zone of the ice sheet as compared to experiment AOVI. This hampers the snow-albedo feedback. As experiment AOVID produces more realistic results than experiment AOVI, we decided to apply time-dependent dust for all further simulations in this paper.
4 Glacial inception and climate changes
Our model simulates the onset of glaciation at the same time as the Earth system model of intermediate complexity of Wang and Mysak (2002). The total ice volume in one of their simulations, named run 1 in their paper, is about a factor of 2/3 smaller than in our experiment AOVID. This (moderate) difference could partly be due to the missing of vegetation feedback in their model. However, there are larger differences in the distribution of continental ice cover between the model of Wang and Mysak (2002) and our model. While in the model of Wang and Mysak (2002) the ice volume at 110 kyr BP is distributed half-and-half in North America and Eurasia, our model simulates much more inland ice in North America than in Eurasia then; at 110 kyr BP, experiment AOVID simulates about four times more ice volume in North American than in Eurasia. This is partly explainable with the inclusion of dust in our simulation.
The summer precipitation at 118 kyr BP (Fig. 6b) decreases in most parts of the Earth. A reduction of the Asian summer monsoon by 1.0 mm day−1 can clearly be seen. Interestingly enough, An (2000) reports that such a weakening is qualitatively seen—at least for the present-day to LGM contrast—in the proxy data. A southward shift of the Intertropical Convergence Zone can be detected by the two bands of positive and negative precipitation anomalies.
At 115 kyr BP, the anomalies of boreal summer temperature (Fig. 6c) have increased by a combination of processes: the boreal summer insolation decreases substantially at 115 kyr BP and the atmospheric CO2 concentration has become lower by about 40 ppmv. Over northern North America, the summer temperature drops by 12°C, a decrease that can be partly attributed to local cooling due to the surface elevation and albedo of the Laurentide ice sheet. Over north-eastern Asia, the temperature decreases by 6°C. The temperature drop of 2°C on the southern Hemisphere cannot be explained by either the limited glaciation there, or by the negligible decrease of summer (JJA) insolation (about 0.3 Wm−2 at 65°S). In our model, the drop of temperature over the Southern Hemisphere is caused by lowered atmospheric CO2 concentration and by the intensification of interhemispheric exchange of heat via the ocean circulation.
At 115 kyr BP, the precipitation anomalies (Fig. 6d) over Canada double in comparison with that at 118 kyr BP (0.4 mm day−1). The Asian summer monsoon further weakens.
5 The role of vegetation and ocean feedbacks
6 The role of atmospheric CO2 concentration and orbital insolation
As shown by Loutre and Berger (2000), we have confirmed that the changes of atmospheric CO2 concentration during the last glacial inception alone are not sufficient to take the system into glacial state. In experiment AOVID_PDI (PDI stands for present-day insolation), the orbital insolation is held at its present-day value throughout the whole simulation. The magnitude of the present-day boreal summer insolation lies approximately between that at 125 kyr BP (Eemian optimum) and at 115 kyr BP. In experiment AOVID_PDI, the inland-ice area in North America and Eurasia remains negligibly small; the drop in the prescribed atmospheric CO2 concentration of nearly 40 ppmv does not push the system into the glacial state (Fig. 7b). Nonetheless, the change of atmospheric CO2 concentration has to be accounted for, which follows from comparison of experiments AOVID and AOVID_280 (Fig. 7b). In the former experiment, the atmospheric CO2 concentration after Barnola et al. (1987) is prescribed, while in the latter one, it is held constant at its pre-industrial value of 280 ppmv through the whole simulation. Of course, in experiments AOVID and AOVID_280 (and all others except AOVID_PDI) the change of orbital insolation according to Berger (1978) is used. In experiment AOVID_280, the climate system runs into glacial state, but the inland-ice cover is smaller than in experiment AOVID. At 110 kyr BP, experiment AOVID simulates about a factor of 1.5 more ice area than experiment AOVID_280.
Experiments AOID and AOVID_280 both simulate lower ice volume and area (Fig. 5) than experiment AOVID does, which illustrates the impacts of variable atmospheric CO2 concentration and interactive vegetation (see Fig. 8). By comparing experiment AOID with experiment AOVID_280 it can be seen that the inland-ice cover reacts in North America in a different way than in northern Europe. Between about 118 kyr BP and 105 kyr BP, experiment AOID simulates less ice volume than experiment AOVID_280 in North America, while the ice volume in northern Europe is larger or (approximately) equal in the same runs during the same time span. This is because vast areas of North America are soon covered with ice, which replaces the vegetation there, while in northern Europe (and Eurasia) the area where vegetation in the course of expanding inland-ice cover may evolve, remains much larger. Between 118 kyr and 113 kyr the large drop of the atmospheric CO2 concentration (experiment AOID) and the retreat of boreal forest (experiment AOVID_280) lead to the same inland-ice cover in northern Europe. In North America there is an additional feedback through the strong increase of inland-ice cover. After about 113 kyr BP, the forest fraction in experiment AOVID_280 increases, thereby leading to warming in northern Europe and less ice volume there than in experiment AOID. Figure 5b shows that the glacial inception in experiment AOVID_280 appears slightly later than that in experiment AOVID, which is attributed to the drop of the atmospheric CO2 concentration and subsequent greenhouse cooling. Experiment AOID has a slightly later glacial inception than either AOVID or AOVID_280, because the vegetation feedback is neglected here.
7 The possible influence of the thermohaline ocean circulation
It was proposed in several studies (Tarasov and Peltier 1997; Khodri et al. 2001, 2003) that reorganisation of Atlantic thermohaline circulation via a reduction of northward heat transport might serve as an additional positive feedback facilitating glacial inception. In particular, Khodri et al. (2001), using a coupled climate model, have found that the orbital configuration corresponding to 115 kyr BP leads to a reduction of the thermohaline circulation and less intensive convection in the Nordic Seas as compared to the present climate state. However, because the model used in the study of Khodri et al. (2001, 2003) did not include ice-sheet dynamics, it was not possible to quantify whether these changes in the ocean circulation could substantially contribute to the build up of the ice sheets in the Northern Hemisphere. To address this issue, we performed the sensitivity experiment AOVID_THC in which we imposed a considerably larger change in the ocean circulation than that simulated in Khodri et al. (2001).
As it was shown in previous studies (e.g. Ganopolski and Rahmstorf 2001), CLIMBER-2 reveals two stable modes of operation under full glacial conditions (similar to LGM): a “cold” (stadial) mode and a “warm” (interstadial) mode. These two modes of operation are characterised by different locations of North Atlantic Deep Water (NADW) formation—in the cold mode, NADW is formed about 15° farther south than in the warm mode. In our experiments, both volume and area of ice sheets remain considerably smaller than at LGM. Therefore, a transition between warm and cold modes did not occur naturally during transient simulations. To mimic such transitions, we imposed the change of one of the model parameter—the so-called “freshwater bypass parameter” (see Ganopolski and Rahmstorf 2001) at 115.5 kyr BP. A change in this parameter leads to an increased fresh-water flux to the Nordic Seas, which causes a rapid transition from the warm to the cold mode of the thermohaline circulation, a weakening of NADW formation and considerable cooling over the North Atlantic.
In conclusion, the effect of a change in the Atlantic thermohaline circulation on the Northern Hemisphere’s inland-ice cover is quite complex. While the temperature drop through a southward shift of convection site influences both Europe and North America, the precipitation drop over Europe is stronger than over northern North America (Fig. 10). This happens because the western Pacific and the mid-latitude Atlantic are the moisture sources for North America and the northern North Atlantic is the main moisture source for Scandinavia (see e.g. Peixoto and Oort 1992). Therefore, there is more inland ice (Fig. 5) in North America for a cold-mode ocean than for a warm-mode ocean. The existence of a large ice sheet in North America additionally cools (and dries) the whole climate system, in particular the Scandinavian region—an effect which additionally contributes to the thinning and southwestward shift of the Fennoscandian ice sheet. Interestingly, our experiment AOVID_THC with two ice domes at 110 kyr BP agrees better with a reconstruction after Kleman et al. (1997) than experiment AOVID.
In contrast to what happens over Scandinavia, ice volume and ice area in North America both show a greater increase in experiment AOVID_THC than in experiment AOVID (Fig. 5a, b), which is consistent with the differences in temperature and precipitation between AOVID_THC and AOVID (Fig. 10a, b); in northern North America the precipitation drop is smaller than in Scandinavia, while the temperature decreases in both regions are comparable. Thus, our results suggest that changes in the Atlantic thermohaline circulation might affect the rate of growth of the ice sheets, but it is unlikely to be a major positive feedback.
We have presented transient simulations of the last glacial inception. Experiment AOVI, which uses the present-day dust-deposition rate, displays the main features of the last glacial inception, i.e., a rapid expansion of inland ice on the Northern Hemisphere at about 115 kyr BP. However, experiment AOVI yields a somewhat larger inland-ice cover in northeastern Eurasia and Alaska than what is indicated by geological evidence. Such an overestimation of inland-ice cover in eastern Eurasia had already been noticed by others (e.g. Marsiat 1994; Calov and Marsiat 1998). In experiment AOVID, where the dust deposition rate was allowed to vary between modern and LGM values, both sea-level change and geographical extent of ice sheets are in much better agreement with palaeodata. It might be possible that our parameterisation (Eq. 1) overestimates the dust-deposition rate during the last glacial inception. However, the data sources for dust do not give a homogenous picture—the dust time series rather depend on the location where the data were taken. In particular, Porter (2001) reports an increased mineral dust concentration in the Northwestern Pacific during glacial inception, which supports our approach. Our model does not account for the radiative effect of dust (Claquin et al. 2003). The impact of the radiative effect of mineral dust on the interglacial–glacial climate change will be investigated in a forthcoming study. Here, we demonstrate the importance of one aspect of mineral dust for ice sheets in the climate system, namely its relevance for ablation of snow and ice.
It is noteworthy that our simulated ice volume agrees well with SPECMAP data for MIS 5e to 5c (Fig. 2). Coral data indicate less ice volume at MIS 5c than the SPECMAP reconstruction (Fig. 2). Since it is known that δ18O data represent a mixed signal of sea-level and deep-ocean temperature changes, it is probable that the sea-level rise during MIS 5c was considerably larger in reality than displayed by the δ18 O signal. If this is the case, then our model underestimates the decrease of ice volume after glacial inception. In this respect, it is interesting to note that in experiments with the LLN climate model, the ice volume returns to the interglacial level during MIS 5c. The reasons for such a differing sensitivity of climate-ice sheet models to the orbital forcing are still unknown.
In our model, glacial inception is triggered by a decrease in boreal summer insolation. Once a critical threshold is crossed, the snow-albedo feedback pushes the system from a interglacial to a glacial state (Calov et al. 2005). Other mechanisms in the climate system such as feedbacks through atmospheric CO2 concentration, vegetation and ocean amplify the strength of the interglacial–glacial transition. Our results suggest the existence of strong synergy in the climate system regarding vegetation and ocean feedbacks during glacial inception.
Vegetation has been recognised as one of the important components of the climate system (by, e.g., Claussen 2004). Our vegetation model (VECODE as a part of CLIMBER-2) displays a broad picture of the dynamics of vegetation. At the Eemian interglacial, CLIMBER-2 shows an increase of forest cover in the northern high latitudes while during the last glacial inception, a retreat of boreal forest is simulated. Our model shows that the dynamics of vegetation together with the ocean feedback amplifies glacial inception. This corroborates earlier results by deNoblet et al. (1996). In their simulations, which were done without inland-ice dynamics, the vegetation feedback helped to create favourable conditions for the growth of the mid-latitude inland ice. Crucifix and Loutre (2002) showed that in their model, MoBiDiC, interactive vegetation cover is crucial to yield a perennial snow cover north of 60°N after approximately 122 kyr BP. Indeed, MoBiDiC has about two times higher sensitivity to changes of forest cover in the latitudinal zone between 60° and 70°N than CLIMBER-2 (Brovkin et al. 2003). Recently, Kageyama et al. (2004) presented a study of last glacial inception with CLIMBER-2 coupled to the ice-sheet model GREMLINS. The sensitivity to interactive vegetation in their coupled model appears to be higher than that in our model. This is because Kageyama et al. (2004) kept the vegetation fixed at the rather warm condition at 126 kyr BP in their atmosphere–ocean-ice-sheets simulation and not at present-day condition as in our corresponding experiment AOVI. Meissner et al. (2003) presented results from simulations with the UVic Earth System model using a constant orbital forcing (representative for 116 kyr BP) and a constant atmospheric CO2 concentration of 240 ppmv. Meissner et al. (2003) confirm the significance of vegetation feedbacks. In their model, synergies and feedbacks related to vegetation dynamics appear to double atmospheric cooling during the last glacial inception. Interestingly enough, Meissner et al. (2003) find that without feedbacks related to vegetation dynamics, monthly averaged September snow, which the authors refer to as perennial snow, still occurs in their model over Baffin Island. This suggests the possibility of glacial inception without any vegetation feedback as seen in the study presented here.
Whether or not changes in the Atlantic circulation took place at the last glacial inception, as simulated in experiment AOVID_THC, cannot yet be answered. Wang and Mysak (2002) found an intensified thermohaline circulation during the last glacial inception in their Earth system model of intermediate complexity. Such an intensification of the thermohaline circulation is similarly reproduced in our experiment AOVID too; but this is not shown explicitly in our paper. In experiment AOVID_THC, we investigate the impact of an induced decrease of the thermohaline circulation, which is even stronger than that reported by Khodri (2001). Kageyama et al. (1999) simulated an increase in precipitation over the North Atlantic at 115 kyr BP in most of their experiments. Such an increase of precipitation could contribute to a weakening of the thermohaline circulation in the North Atlantic. Despite the large decrease of the thermohaline circulation between experiment AOVID and experiment AOVID_THC (about 11 Sv difference in maximum Atlantic overturning at 115 kyr BP), the response of our model in terms of ice volume is rather small. Hence, feedbacks related to changes in the thermohaline circulation, as those related to the vegetation, are also not likely to have caused the last glacial inception. However, these feedbacks appear to be a strong amplifier once the inception has started.
9 Summary and conclusions
An Earth system model of intermediate complexity, CLIMBER-2, has been used to explore the processes, feedbacks and synergisms that led to the last glacial inception some 120 to 115 kyr BP. The fully coupled atmosphere-ocean- vegetation-inland-ice model successfully simulated the reconstructed ice-volume change both in amplitude and evolution.
The last glacial inception appeared in the model as a bifurcation in the physical climate system once a threshold in maximum boreal summer insolation was crossed (Calov et al. 2005). The ice-albedo feedback is the major mechanism, which leads to a rapid expansion of the area covered by inland ice. Feedbacks and synergisms related to vegetation dynamics, ocean dynamics and changes in atmospheric CO2 concentration serve as important amplifiers, but they are not essential to trigger the last glacial inception in our model. These positive feedbacks considerably contribute to the increase of inland-ice cover.
The transition from interglacial to glacial climate appears in our model when pre-industrial atmospheric CO2 concentration is prescribed instead of changes in atmospheric CO2 concentration during the last glacial inception according to reconstructions by Barnola et al. (1987). In turn, there was no sign of a glacial inception in the simulations when the orbital insolation was kept at its present-day value and the atmospheric CO2 concentration from the palaeo-records (Barnola et al. 1987) was prescribed.
The role of mineral dust has been investigated by comparing experiments with fixed present-day dust deposition and with varying dust deposition, respectively. The experiments in which the dust deposition rate was allowed to vary between modern and LGM values—using the modelled ice volume as a weighting factor—were in better agreement regarding geological evidence of inland ice extent than the experiment in which dust deposition was kept constant at present-day values. This indicates that an increase of dust deposition on inland ice may act as regional negative feedback. In this study, only the impact of dust on the albedo of snow, which affects the ablation of the ice sheets, is considered. Radiative effects of airborne dust were ignored in this study.
In our simulations, most of the inland ice is formed in North America, while the European ice sheet is relatively small and restricted to Scandinavia. The growth of ice sheets causes a large amount of cooling in the Northern Hemisphere and considerable changes in all components of the climate system. In particular, precipitation is considerably reduced over North America, the Amazon region, North Africa and large parts of Eurasia.
The inland-ice cover over Scandinavia has a subtle dependence on changes in the thermohaline circulation of the Atlantic Ocean. Depending on whether the convection site in the North Atlantic is located at a more southern or a more northern position, the climate over Scandinavia is rather moist and warm or, alternatively, dry and cold. In the former case, the ice over Scandinavia builds up in the model to a rather thick ice sheet which is located more in the northern part of this region while in the latter case, the simulated ice sheet is rather thin and located more to the south.
The authors wish to thank Ralf Greve for providing us with his polythermal ice-sheet model SICOPOLIS as well as for his assistance. We thank Natalie Mahowald, who sent us her reconstruction of present-day and LGM distributions of dust deposition, and Claire Waelbroeck, who gave us her sea-level record. We further thank two anonymous referees whose thorough reviews improved our earlier manuscript. Alison Schlums helped in editing our manuscript. This work was supported by the Deutsche Forschungsgemeinschaft (research grant CL 178/2-1 and CL 178/2-2) and partly supported by the German Climate Programme DEKLIM (subcontract to BMBF project 01 LD 0041).