Transient climate simulation for the last interglacial

U. Mikolajewicz, M. Gröger, E. Maier-Reimer, G. Schurgers, M. Vizca&íno & A. Winguth*
Max-Planck-Institute for Meteorology, Bundesstr. 53, 20146 Hamburg, Germany, mikolajewicz@dkrz.de
*) Center for Climatic Research, Madison, USA



1. Introduction

Model: MPI Earth System Model for paleo studies including atmosphere (ECHAM), ocean (LSG), ocean biogeochemistry (HAMOCC), and terrestrial biosphere (LPJ).
Experiments: The model was forced with time varying insolation corresponding to orbital parameters between 129 to 109 kyBP. Continental ice sheets were not coupled but prescribed according to potential pre-industrial conditions. Whereas changes in the total incoming solar radiation do not exceed 3 W/m2 in the course of the Last Interglacial seasonal, variations are quite large and comprise an evolution from maximum values around 127 kyr BP to minimum values at 116 kyr BP in the Northern Hemisphere. As a result, the yearly maximum insolation was increased by up to 10 % at 126 kyr BP and 60 °N compared to the present-day value in the northern hemisphere. The forcing is nearly inverse in respect to the two hemispheres. The major effect is a strengthening (attenuation) of seasonal contrasts in the northern (southern) hemisphere. The northern hemisphere summer insolation was enhanced while the winter received somewhat less solar radiation at 126 kyBP (fig. 1 left). At 115 kyBP nearly the same pattern is seen but with inverse premisses (Fig. 1 right) .

Insolation at the top of the atmosphere(high resolution):


Fig. 1: Annual cycle of prescribed insolation changes at 126 (left) and 115 kyBP compared to the control experment.

2. Changes in the Hydrological Cycle


Fig. 2: Changes in atmospheric moisture transports. Blue shading indicates freshwater gain at the Earth surface at 126 compared to 115 kyr BP. Color change (streamlines) from red to blue indicate air moisture increase at 126 kyr BP. (high resolution)

The strong heating leads to substantial changes in the hydrological cycle which is generally enhanced in the northern hemisphere. Large scale moisture transport changes result in increased freshwater input and increased salinities in the North Atlantic basin and decreased freshwater input/ increased salinities in the North Pacific during the warm optimum (Figure 2). This somewhat surprising result is caused by the combined effects of atmospheric interbasin freshwater transports and polward advection of tropical saline surface waters in the Atlantic. At 126 kyr BP the Atlantic/Arctic drainage basin north 30 oN receives about 28 mSv more freshwater than at 115 kyr BP. The surplus arises primarily from increased transports across the watersheds of North America (~68 mSv) which is partly compensated by a diminished export (~36 mSv) across the watersheds of Asia (Figure 2). However, in total the Atlantic Ocean north of 30 oS loses ~36 mSv freshwater mainly over South America and Panama where the export increases by ~64 mSv.

(high resolution):
Fig. 3:top: Evolution of sea surface salinity in the Atlantic north of 30 °N (left) and corresponding freshwater input (Sv;right) . Bottom: same for the Pacific. Black line indicates variability of the control run.

The import of freshwater over North America into the North Atlantic catchment area is related to the expense of the North Pacific, where the streamlines indicate moisture transfer from the subtropical high pressure system via the westerlies to the north where a part of the moistures escapes then across the Arctic/American the watershed (Figure 2). This export is clearly intensified during the warm optimum compared to 115 kyr BP (Figure 2). Strongly converging streamlines stimulate a large positive freshwater flux anomaly within a narrow band in the eastern tropical Pacific (Figure 2). The main source for this moisture is the subtropical South Pacific where increased uptake and transport via the trade winds leads to increased evaporation. Another important change is the enhanced export via South America and Panama. Whereas the export over Panama is at first order simply enhanced, farther south the transports are affected by substantial changes in the Walker Circulation that lead to strong east-west dipole pattern of freshwater fluxes in the tropical Atlantic and surrounding landmasses. During the northern hemisphere warm season strong heating over tropical Africa leads to stronger convection resulting in an anomalous moisture transport from the western and eastern tropical Atlantic to Africa as indicated by the anomalous streamlines in Figure 2. In the western tropical Atlantic strong divergent streamlines are seen over Amazonia, where moisture is exported for the one part to Africa and for the other part it escapes over South America to the tropical Pacific. These changes result in an extremely strong increase in evaporation in the tropical western Atlantic and a corresponding freshening of the eastern Atlantic which benefits by the increased runoff from the Congo and Niger catchment area.


2. Changes in Ocean Circulation
Due to the changes in solar radiation and the hydrological cycle the North Atlantic was warmer and more saline during the early Eemian (Fig. 4). In the Irminger Sea the temperature change effect on density is about twice as large com pared to the counteracting change in salinity. As a result, the overturning in the abyssal North Atlantic weakens. At 30 °N the weakening amounts up to more than 3 Sv (>10 %) between 2500 an 4000 m whereas the total overturning (which is modelled at around 1500 m) is only slightly and in significantly reduced. The altered circulation leads to changes in the nutrient distribution which are, on longer timescales, temporally and spatially negatively correlated to the salinity changes. In the North Pacific, both temperature and salinity changes

(high resolution):
Fig. 4: Left: Yearly SST anomaly (126 kyr BP - control). Right: Same for salinity in the the second model layer (50 -139 m)

(Fig. 5 left) have a decreasing effect on density of the upper layers. The stronger stratification weakens the large scale upwelling the North Pacific which is driven by vigorous AABW inflow from the south. The maximal convection depth in the central North Pacific is reduced by several hundred meters. As a result of these two changes the overturning circulation in the North Pacific is weakened in the upper limb. At 30 °N the weakening amounts to about 0.5 Sv (15 %) at 1500 m . However, the return flow near the continental margin of North America is strenghtened in the deeper layers so that the total overturning does not change significantly. This leads to a better ventilation and diminished nutrient concentrations in the North and East Pacific which strongly diminishes the productivity in areas with strong upwelling (Fig. 5 right).

(high resolution):
Fig. 5: Left: Zonally averaged density anomaly (kg/l) in the Pacific. Right: Anomaly of carbon export to abyssal layers (mol/m2)




2. Carbon Cycle

(high resolution):
Fig. 6: Changes in carbon storage (kg/m2) between 115 and 126 kyBP. Black line indicates the 95% confidence level for the ocean realm

The cooling in the course of the Eemian leads to a net release of 308 Gt of soil/vegetation carbon to the atmosphere. Almost all of this surplus is uptaken by the ocean (Fig. 6; Table 1) so that the atmospheric pCO2 change remains nearly below the noise level (Fig. 7). Most of carbon is stored as total dissolved CO2 which increases by 430 Gt. The carbonate sediment pool decreases by 140 Gt. Although the relative reduction of the carbonate sediment pool is almost identical in the Pacific (10.0 %) and Atlantic (8.3 %) the latter is, due to its larger pool (Fig. 7), the by far most important ocean. Between 126 and 115 kyr BP more than one half (76 Gt) of the 140 Gt reduction of the global carbonate pool is related to the Atlantic. It is followed by the Pacific which, because of its large extension accounts for at least one third (44 Gt). Only minor contributions come from the Indic with 8.9 % (12.5 Gt) and the Southern Ocean 5,5 %(7.9 Gt).

(high resolution):
Fig. 7: a) Inorganic carbon for selected ocean basins. b) top: modelled pCO2 and reconstruction from Vostok (indicated by star symbols). Below: Modelled carbon storage for the Ocean and terrestrial biosphere.Black lines indicate variability of the control run


UNDER CONSTRUCTION
Updated:
Matthias Gröger (groeger@dkrz.de)
Last modified: May 18 2006.