K. J. Meissner Æ R. Gerdes Coupled climate modelling of ocean circulation...

K. J. Meissner Æ R. Gerdes

Coupled climate modelling of ocean circulationchanges during ice age inception

Received: 21 August 2000 /Accepted: 3 July 2001 / Published online: 7 December 2001! Springer-Verlag 2001

Abstract Freshening of high latitude surface waters canchange the large-scale oceanic transport of heat andsalt. Consequently, atmospheric and sea ice perturba-tions over the deep water production sites excite alarge-scale response establishing an oceanic ‘‘tele-connection’’ with time scales of years to centuries. Tostudy these feedbacks, a coupled atmosphere-ocean-seaice model consisting of a two dimensional atmosphericenergy and moisture balance model (EMBM) coupledto a thermodynamic sea ice model and an ocean gen-eral circulation model is utilised. The coupled modelreproduces many aspects of the present oceanic circu-lation. We also investigate the climate impact ofchanges in fresh water balance during an ice age in-itiation. In this experiment part of the precipitationover continents is stored within continental ice sheets.During the buildup of ice sheets the oceanic stratifica-tion in the North Atlantic is weakened by a reducedcontinental run-o! leading to an enhanced thermoha-line circulation. Under these conditions salinity is re-distributed such that deep water is more saline thanunder present conditions. Once the ice sheets built up,we simulate an ice age climate without net fresh waterstorage on the continents. In this case the coupledmodel reproduces the shallow and weak overturningcell, an ice edge advance insulating the upper ocean,and many other aspects of the glacial circulation.

1 Introduction

Changes in solar radiation and its geographical andseasonal distribution are not su"cient to explain theobserved glaciation cycles (Crowley and North 1991).Positive feedbacks in the climate system are essentialamplifiers of the external forcing (Broecker and Denton1990). One important mechanism involves the inland icesheets (Imbrie et al. 1992). As summer insolationdecreases to values low enough to prevent snow meltover large areas during summer, the ice-albedo feedbackaids the glaciation process (Budyko 1969). During iceage inception, atmospheric moisture fluxes must carrythe necessary moisture for ice sheet growth in theNorthern Hemisphere. This moisture originates mainlyfrom the subtropical and subpolar Atlantic Ocean(Ruddiman and McIntyre 1981). Evaporation dependson the di!erence between sea surface temperature (SST)and surface atmospheric temperature (SAT). It thusdepends on the oceanic heat transport that is governedby the thermohaline circulation (THC). The heattransport not only increases SSTs but also keeps largeparts of the northern North Atlantic ice free, allowinge!ective moisture flux into the atmosphere. Ruddimanand McIntyre (1981) thus postulated that surface watersover large parts of the subpolar and the northern sub-tropical Atlantic remained warm and saline for severalthousand years into the major ice-growth phase of thelast glaciation.

Deep water formation in the northern North Atlanticis a major part of the large-scale THC (Broecker 1991).In regions of deep water formation, relatively smallperturbations in the sea surface salinity (SSS) can en-hance or weaken convection, having considerable con-sequences for the general circulation (Rahmstorf 1999).Large-scale changes, such as ice sheet growth or melting,seriously a!ect the fresh water fluxes in the most vul-nerable regions for the THC (Liccardi et al. 1999). Theseinteractions are an example of the feedback complexity

Climate Dynamics (2002) 18: 455–473DOI 10.1007/s00382-001-0192-x

K. J. Meissner (&)Alfred Wegener Institute for Polar and Marine Research,Germany, now at the School of Earth and Ocean Sciences,University of Victoria, P O Box 3055,Stn CSC, Victoria, BC, V8W 3P6, CanadaE-mail: [email protected]

R. GerdesAlfred Wegener Institute for Polarand Marine Research, Bussestr.24,D-27570 Bremerhaven, Germany

between the atmosphere, sea ice, and ocean duringmajor changes in glacier volume at the inception ortermination of glaciations.

Duplessy and Labeyrie (1992) interpret the dis-appearance of subpolar planctonic foraminifera in theNorwegian Sea during the second half of the last inter-glacial (marine isotope stage 5e or the Eemian) as anindication for the cooling of Norwegian Sea surfacewaters that preceded the growth of continental ice sheets(see also Cortijo et al. 1994). Marked variations innorthward oceanic heat flux existed in high northernlatitudes in the absence of Northern Hemisphere icesheet forcing (Fronval and Jansen 1996). Apparently,the high latitudes of the Northern Hemisphere re-sponded early to orbital climate forcing, when inter-glacial conditions still prevailed in Europe and theNorth Atlantic. On the other hand, summer SST andSSS were stable or increased during the same time periodat subtropical and subpolar North Atlantic sites (Cortijoet al. 1999). This implies a high SST contrast betweenthe subpolar North Atlantic and the Nordic Seas thatcould have lead to enhanced cyclogenesis and strongernorthward moisture transport in the atmosphere, sup-porting the subsequent growth of inland ice sheets(Khodri et al. 2001).

While cooling in the high northern latitudes of theNordic Seas preceded the growth of major ice sheets, thesituation was di!erent in the Atlantic. Here, ice sheetgrowth during the transition from marine isotope stage5e to 5d preceded surface cooling as well as the shut downof the THC in the Atlantic that is thought to have ter-minated the last interglacial (Broecker 1998). At 54"N inthe easternNorth Atlantic (deep sea core V29-191), warmsurface ocean conditions prevailed long after the increasein dO18 in benthic foraminifera that announced thegrowth of inland ice sheets (Kukla et al. 1997). Referringto pollen records on land, Kukla et al. (1997) concludethat the climatic optimum in western Europe was con-temporaneous with the first half of marine isotope stage5d. Later, ice rafted debris appeared at that site togetherwith cold water foraminifera. This implies that sub-stantial amounts of continental ice were already presentas the eastern North Atlantic started to cool. Broecker(1998) speculates that release of fresh water from a newlyformed ice cap in eastern Canada was responsible for thefinal slowdown of the THC in the Atlantic.

In this paper, we pursue two aims: first we simulatethe e!ect of the inception of glaciation on the oceaniccirculation; second, we model the glacial circulationwithout relying on insu"ciently known boundary values(Fieg and Gerdes 2001). SSS estimates for the last glacialmaximum (LGM) contain large uncertainties, especiallywhere melting and freezing occurs and in marginal ba-sins (Rohling and Bigg 1998; Winguth et al. 1999). Dueto the sensitivity of ocean models to changes in SSS inhigh latitudes (Maier-Reimer et al. 1993; Gerdes andKoberle 1995) small errors in those salinity estimates canlead to large errors in large-scale ocean circulation. Theapproach used here is to begin integration from the

current climate state. We then change the boundaryconditions (e.g. continental run-o!) such that the systemmakes a transition to a glacial state. Unlike modelsforced with mixed boundary conditions (e.g. Bryan1986) we instead utilise a fully coupled atmosphere-ocean-sea ice model.

In Sect. 2 we begin by discussing the coupled climatemodel employed, and by describing the coupling pro-cedure of each component. In Sect. 3 the results of acontrol run for the present day climate are comparedwith observations. Sections 4 and 5 contain descriptionsof our glacial inception experiments. Finally, Sects. 6and 7 consist of a discussion and conclusions.

2 Model description and coupling

The atmospheric component of our model is a global two dimen-sional energy and moisture balance model (see Meissner 1999 andAppendix A for a detailed description). Its thermodynamics arebased on the models of Budyko (1969) and Sellers(1969); the ver-tically integrated balance of atmospheric moisture and energy isconsidered, with advection represented by a simple di!usive ap-proximation. The model’s prognostic variables are temperature andhumidity, allowing the prediction of latent, sensible and radiativeheat fluxes, as well as moisture fluxes between the ocean or the seaice and the atmosphere.

A thermodynamic sea ice model regulates heat and salt fluxes tothe ocean when the energy balance at the atmosphere-ocean in-terface produces SSTs at the freezing point ()1.9 "C) or below(Semtner 1976). The model computes ice thicknesses and sea icesurface temperatures and is detailed in Appendix B.

The third component of the coupled system is an oceangeneral circulation model (OGCM), MOM2 (Modular OceanModel, version 2, 1995), developed at the Geophysical FluidDynamics Laboratory (GFDL) (Pacanowski 1995). The hor-izontal resolution is 3" in latitude and 4" in longitude, whilethere are 20 levels in the vertical with the thickness increasingwith depth (see Table 1). The topography is taken from Dana-basoglu and McWilliams (1995). Smoothing and a minimumdepth requirement in their model removes the Arctic shelf seasand leaves the Greenland-Iceland-Scotland ridge area too deep.Because this results in unrealistic large-scale circulation featuresin the present coupled model, we reduce the ocean depth to 400m in the Barents Sea and 1100 m for the Greenland-IcelandRidge (Meissner 1999). The values of the horizontal and verticalmixing parameters for momentum as well as the time steps fortracers, barotropic momentum and baroclinic momentumare summarised in Table 2. Di!erent time steps are utilised fortracers (temperature and salinity) and momentum. With thisasynchronous integration technique, the propagation speeds ofthe fastest waves (i.e. the internal gravity waves and the externalRossby waves) are reduced (Bryan 1984).

Tracer advection is implemented via the FCT (Flux CorrectedTransport, Zalesak 1979) scheme, which eliminates undesirablecomputational ‘‘noise’’ within the models tracer fields while redu-cing di!usive e!ects on the tracer fields as far as possible (seeGerdes et al. 1991; Weaver and Eby 1997).

The three modules, atmospheric, oceanic and sea ice model, arecoupled with di!ering time steps. For this procedure di!ering timesegments are defined, during which each model computes its fluxeswith fixed boundary conditions given by the two other components.The time segment for the atmosphere and the ice model covers 12 h,while the ocean model integrates 24 h between transfers. During anintegration segment the submodels average the values in time thatwill then be interchanged.

Figure 1 indicates which variables are exchanged within eachtransfer. Since the atmospheric model does not predict the pressure

456 Meissner and Gerdes: Coupled climate modelling of ocean circulation changes

field, the wind stress at the ocean surface is given by climatologicalvalues from Hellerman and Rosenstein (1983). The only other ex-ternal forcing field in the coupled model is the daily computedshortwave incoming radiation (Berger 1978). The ocean model re-ceives heat fluxes from the atmosphere and the sea ice (Foa and Qw,respectively), as well as fresh water fluxes from precipitation minusevaporation (P–E), continental run-o!, and melting or freezing ofsea ice (Fs). Over continents, precipitation is assumed to im-mediately run-o! according to the distribution of the 75 catchmentbasins (see Fig. 3). To calculate heat fluxes and evaporation theatmosphere needs the SST and the sea ice surface temperature (Ts

and Tis) as well as the fractional sea ice cover in a grid box. The seaice model receives the heat flux at the ice-atmosphere interface (Fia)and the snow fallen on ice-covered surfaces from the atmosphericmodel. The ocean model delivers the SST (Ts) for the calculation ofthe heat flux between the ocean and the sea ice. The ocean and thesea ice model are configured on a 4" · 3"grid, while the atmospherehas a resolution of 10" · 10". Consequently, within each exchangebetween model components the variables are interpolated linearlyon the respective grid, conserving energy and moisture (fresh wa-ter). Underneath an atmospheric grid box there can be oceanboxes, continental boxes as well as boxes with fractional sea icecover. Thus, the fractional land and sea ice cover must be con-sidered during exchanges. The atmospheric model calculates twoindependent heat fluxes according to 100% ocean or sea ice cover,which are transferred proportionally to the sea ice model and the

ocean model. Also the fresh water fluxes are proportioned, ac-cording to the current sea ice cover, between the sea ice model andthe ocean model (see Fig. 2 for an example).

According to Aagaard and Carmack (1989), 2790 km3 of freshwater per year is transported through Fram Strait into the NorthAtlantic by advection of sea ice. Furthermore, 920 km3 per year istransported via the Canadian Archipelago into the Labrador Sea.In the current model, the Canadian Archipelago is closed due to thesimplified topography of the ocean model; additionally, the sea icemodel is purely thermodynamic and cannot simulate an advectivesea ice transport. In order to incorporate this fresh water flux,precipitation fallen over sea ice is immediately transferred to theEast Greenland Current and the Labrador Sea. This procedure alsoensures that sea ice will not acquire unrealistic thickness simply dueto accumulation of snow on multi-year ice. The parametrisation issimilar to the treatment of sea ice export recently introduced byGanopolski and Rahmstorf (2001) in their zonally averaged ocean

Table 2 Time steps andmixing parameters of theocean model

Description Parameter Value Units

Horizontal viscosity Am 106 m2s–1

Vertical viscosity jm 10–3 m2s–1

Model time step (tracer) sTS 43,200 sModel time step (momentum, baroclinic) suv 1,728 sModel time step (momentum, barotropic) sY 864 s

Table 1 Vertical resolution inmetres. Dz = layer thickness,z = depth of the scalar andvector points

Layer Dz z Layer Dz z

1 51.23 25.00 11 265.64 1,354.382 56.13 77.46 12 296.55 1,635.673 65.79 137.52 13 326.30 1,947.474 80.00 209.05 14 354.18 2,288.275 98.39 297.25 15 379.49 2,655.826 120.51 405.82 16 401.61 3,047.257 145.82 538.27 17 420.00 3,459.058 173.70 697.47 18 434.21 3,887.259 203.45 885.67 19 443.87 4,327.4610 234.36 1,104.38 20 448.77 4,775.00

Fig. 1 Schematic representation of exchanges between componentsof the coupled model

Fig. 2 Example of overlaying grids. The atmospheric grid isrepresented by the solid line and the oceanic grid by the dottedline. Underneath an atmospheric grid box there can be severalocean grid boxes covered by land, sea ice, or water

Meissner and Gerdes: Coupled climate modelling of ocean circulation changes 457

model. In the polar areas of the Southern Hemisphere such atreatment is not necessary as practically all sea ice melts duringsummer. Di!erential surface fresh water fluxes due to sea icetransport probably contribute to dense water formation overAntarctic shelves. This e!ect is not incorporated in the currentmodel. Consistent with the simplicity of the sea ice model, snowfallen over sea ice is immediately transferred to the underlyingocean grid boxes without imposing a lateral transport of freshwater.

Due to the model resolution, the circulation of the Arctic Oceanis very sensitive to fresh water inputs. Hence, some of the catch-ment areas of East Siberian and Canadian rivers are assigned toestuaries in the Pacific Ocean. This procedure was devised byWeaver et al. (1998) to overcome the deficiencies of coarse re-solution models to reproduce the land geometry and to sustainNorth Atlantic Deep Water (NADW) production. The entirecontinental discharge into the Arctic Ocean is 6,735 km3 per yearwith this rearrangement and is still higher than the estimated realvalue of 3,300 km3 per year (Aagaard and Carmack 1989).

3 Control run

Since the sensitivity of the model climate ultimately de-pends on the model’s climate state itself, the importanceof properly simulating the present day climate is para-mount. We therefore begin the discussion with themodel’s equilibrium state obtained after 3,700 years ofintegration. The large-scale temperature patterns at adepth of 210 m (Fig. 4) are similar to the observed cli-matological distributions (e.g. Levitus et al. 1994).Highest temperatures prevail in the centres of the sub-tropical gyres. At their eastern edge, we find lowertemperatures that are due to the equatorward advectionof subpolar mode waters and the upwelling of coldwater, forced by the prescribed winds. The northeast

extension of the subtropical warm pool in the NorthAtlantic can be regarded as an imprint of the model’sTHC. Advection of warm water with the Gulf Stream/North Atlantic Current system yields relatively hightemperatures in the northeast North Atlantic and theNordic Seas. As in other coarse resolution modelsimulations, the path of the Gulf Stream follows theNorth American coast too closely and too far north. Onthe other hand, the path of the North Atlantic Currentalong approximately 50"N to the eastern North Atlanticand the Nordic Seas is well represented (Kase andKrauss 1996).

In the North Pacific, the subtropical warm pool oc-cupies too large an area and curtails the subpolar do-main. The Kuroshio flows along the Asian coast almostup to 50"N. This causes the formation of North PacificIntermediate Water to be underrepresented in themodel. The salinities reflect the characteristic asymmetrybetween the Atlantic and the Pacific Ocean (see Fig. 5).Strong north-eastward transports in the North Atlanticlead to a subtropical salt signal that penetrates far to thenorth. However, the asymmetry is less distinct than inclimatological data; the salinities are too high in thesubtropical Pacific and too low in the entire Atlantic.The salt tongue in the western part of the North Atlanticis well represented; however, the subtropical sourcewaters are already too fresh such that salinities in thenortheastern Atlantic turn out lower than observed. Thiscauses the formation of relatively fresh and low densityNADW. The tropical region and the subtropics of theSouthern Hemisphere closely resemble the observedclimatology, although the Pacific equatorial cold tongueis only weakly pronounced. Both the Atlantic and the

Fig. 3 Run-o! scheme of thecoupled model. Grid boxesa"liated to the same catchmentarea are filled with the samegrey shade. Arrows indicate gridboxes in which the run-o! isallocated to the ocean. Catch-ment areas without arrows areoverlapping with underlyingocean grid boxes


Pacific feature an Equatorial Undercurrent, somewhatshallower in the Atlantic than in the Pacific.

The Antarctic Circumpolar Current (ACC), with atransport of 160 Sv at Drake Passage, dominates theSouthern Ocean circulation. The modelled transport isthus slightly higher than estimates based on observations(Whitworth III et al. 1982; Orsi et al. 1995). The cyclonicgyre in the Weddell Sea is reproduced, while the north-ward deflection of the Antarctic Circumpolar Currenton the eastern side of the Drake Passage is also similarto the behaviour of the observed circulation. The modelreproduces the Southern Ocean salinity structure verywell, including the salinity minimum band between thesubtropics and the immediate vicinity of the Antarcticcontinent. South of 40"S, salinities are generally too lowand temperatures too high compared to climatology.Precipitation minus evaporation over the SouthernOcean computed by the atmospheric model does notexceed 300 mm per year and lies near the magnitudesgiven by Peixoto and Oort (1992) (evaporation rate of200 mm per year and precipitation rate of 500 mm peryear in between 50"S and 70"S). A more likely cause forthe low SSS is the high run-o! from the Antarctic con-tinent. The zonal mean of the total freshwater flux dueto precipitation, evaporation, ice melting and freezing,and run-o! is shown in Fig. 6. The atmospheric model istuned with precipitation minus evaporation rates givenby the COADS data set (Da Silva et al. 1994, dottedline). The net precipitation rates between 45"S and 75"Sare unrealistically large, leading to a maximum in themodel’s precipitation at 70"S (solid line). This causes themodel’s high continental discharge from Antarctica.Figure 6 also provides the corresponding data fromOberhuber(1988) (dashed line). Other than in the polarregions, the model results closely fit the climatologicaldata of Oberhuber (1988).

To demonstrate the water mass distribution in themodel we show in Fig. 7 global zonal averages of po-tential temperature and salinity in comparison with cli-matology. The model’s Southern Ocean contains a largepool of fresh water that leads into a very pronouncedvolume of Antarctic Intermediate Water (AAIW). Thiswater is too fresh in the model and covers a large depthinterval. The model AAIW spreads as far north as 40"N;it is not impeded by Mediterranean Water that in natureoccupies a similar depth range in the North Atlantic. Inthe extreme south, a stable halocline prohibits the for-mation of Antarctic Bottom Water (AABW). Accord-ingly, the simulated abyssal ocean lacks very lowtemperatures. Except for the excessive continental run-o!, other processes may be important in preventingdense water formation in the extreme south. Namely, thesea ice model employed is purely thermodynamic; hence,once sea ice of enough thickness forms, it insulates theocean surface from the atmosphere, preventing heat lossfrom the ocean and formation of dense water. In reality,surface divergence in this area due to the wind forcingwould allow for the formation of leads and hence heatloss from the ocean, additional sea ice formation, brinerejection, and sustained dense water formation.

In the northern polar latitudes subsurface waters aretoo warm and fresh in the model. Here, the coarse re-solution and the simplified topography of the oceanmodel prevent the proper water mass formation andspreading (Wiebe and Weaver 1999).

The warm water sphere occupies a large area in themodel due to an overemphasis on downward Ekmanpumping of warm and saline waters compared to hor-izontal transport processes. Experience with the FCT-advection scheme used here (e.g. Gerdes et al. 1991)shows that implicit di!usivities in the ocean’s interiorare very small such that the tracer balance there is

Fig. 4 Annual mean potentialtemperatures and velocities at adepth of 210 m at the end of thecontrol run. Temperatures aregiven in "C and velocities incm s–1


entirely advective. The vertical advection is mostlyprescribed by the wind stress and must be balanced byhorizontal advection. This only occurs when the hor-izontal gradients are su"ciently large, resulting in toodeep a thermocline in the model.

Amongst the most important features are the prop-erties and the pathways of simulated NADW. The hightemperatures in the formation areas of NADW result inhigh temperatures over a large volume of the ocean.These temperatures are also responsible for relativelylight NADW that spreads at relatively shallow depths inthe model. To better examine the spreading of NADW,we introduced an artificial surface tracer. The con-centration of the tracer is held at unity in the top layerover a period of 1,000 years, enabling an investigation ofthe extent to which water masses in deeper layers are

replaced by surface waters during the last thousandyears of model integration. The concentration and ve-locities of the tracer at a depth of 1,950 m are re-presented in Fig. 8. The highest concentrations can befound in the western part of the North Atlantic, close tothe convection sites that are situated in the Irminger,Labrador and Greenland Seas. The deep westernboundary current carries NADW southwards along theAmerican continental slope. In the southern Atlantic,NADW reaches the northern flank of the ACC where itis transformed to Circumpolar Deep Water (CDW) andspreads eastward. Several northward branches of CDWare visible in the Indian and Pacific Oceans. This deepwater circulation corresponds to present concepts of theredistribution of NADW (Schmitz 1996).

The zonally integrated meridional heat transports ofthe coupled system are shown in Fig. 9. The atmosphericheat transport amounts to 3.7 Petawatts (PW) at 45"Nand to –4.2 PW at 35"S, which is about 25% higher thanestimated by Peixoto and Oort (1992). Thus the atmo-spheric model counterbalances the relatively weakoceanic heat transport that reaches 1.16 PW at 20"N and0.52 PW (southward) at 10"S. Weak meridional oceanicheat transports are a common problem for ocean modelswith coarse resolution (e.g. Fanning and Weaver 1996).A large part of the heat transport in the ocean is due tothe meridional overturning motion. For the NorthAtlantic we show the stream function for the zonallyintegrated mass transport in Fig. 13a. The southwardbranch of the meridional overturning carries NADWand only extends to about 2,000 m depth. The maximumvalue of 18 Sv, however, corresponds to estimatesbased on measurements (e.g. Dickson and Brown 1994;Schmitz 1995).

The equilibrium present day atmospheric tempera-ture (SAT) and precipitation minus evaporation areshown in Figs. 10 and 11. The SAT is in reasonablygood agreement with NCEP reanalysis data (Kalnay

Fig. 5 Annual mean salinities(units in psu) at a depth of210 m at the end of the controlrun

Fig. 6 Zonally averaged fresh water flux (units in mm per year) atthe end of the control run (solid line). Climatological data fromOberhuber (1988) (dashed line) and Da Silva et al. (1994)(dottedline) are shown for comparison


et al. 1996) on large scales. In some local regions overland, the model is not able to simulate cold tempera-tures due to high elevations (e.g. the Rocky Mountains,Andes, and Himalayas). Temperatures over inland icesheets (Greenland and Antarctica) are too warm for thesame reason. The simulated precipitation minus eva-poration pattern shows the Intertropical ConvergenceZone as well as the subtropical regions with high eva-poration rates. The two dimensional pattern as well asthe values are in good agreement with climatologicaldata (Kalnay et al. 1996) (see Figs. 11 and 6, respec-tively).

4 Glacial transition

The coupled model is driven only by the imposed solarradiation (Berger 1978) and wind stress at the sea surface(Hellerman and Rosenstein 1983). As such our coupledclimate model is suited to investigate scenarios which

deviate from present conditions. At the beginning of thelast ice age 115 kyr BP, solar radiation at mid-latitudeswas weaker than today (Milankovitch 1930). Moreover,Kageyama et al. (1999) suggest that during the inceptionof the last glaciation a more north-eastward position ofthe tail of the storm-track favoured precipitation overnorthwestern Europe and in particular over the site ofgrowth of the Fennoscandian ice sheet. Thus, the airtemperatures decreased and the proportion of snowfalling over continents during winter (which did not meltduring summer), increased during the transition frominterglacial to glacial conditions. Continental ice sheetsbegan to expand, which again caused a local tempera-ture decrease due to a higher surface albedo (Budd andSmith 1981). The ocean’s reaction to these modifiedboundary conditions is not well known and remains akey issue in climate research. Marine data show a sub-stantial weakening of the THC in the northern Atlanticduring the LGM (e.g. Sarnthein et al. 1994). The causesof this weakening and the behaviour of the oceanic cir-

Fig. 7a–d Annual mean globalzonal averages: a potentialtemperature from the model(in "C), b potential temperaturefrom the climatological data(in "C) (Levitus et al. 1994),c salinity from the model (inpsu), d salinity from theclimatological data (in psu)(Levitus et al. 1994)


culation during the growth of continental ice sheets isuncertain. An important factor a!ecting the oceaniccirculation during ice ages is the global increase andspatial redistribution of salinity, caused by the storage offresh water on the continents within continental icesheets. Oxygen isotope analysis of deep sea cores in-dicates that continental glaciations occurred generally inless than 10,000 years (McIntyre et al. 1978). To examinethe influence of this modified fresh water balance on theocean, an experiment in which half of the precipitationover North America and Scandinavia was stored ingrowing continental ice sheets (over a period of 4,742years) was conducted.

In order to investigate the transition to a glacial statewe begin from the control run (which has reached a sta-tistical equilibrium after 3,672 years, period T1). We thenconsider a buildup of continental ice in North Americaand Scandinavia by storing half of the precipitation inthese regions (year 3,673 to 8,414, period T2). We regardthis as a sensitivity study for the ocean circulation. Thedynamics of continental ice sheets and the dynamicresponse of the atmosphere are not considered here. Theincoming shortwave radiation remains unchanged, whilethe surface albedo over the continental ice is set to 0.7. Thesurface albedo at the end of this phase (Fig. 12) indicatesthe extent of the continental ice sheets. With this proce-dure, we assume that the direct e!ect of changes in thesolar radiation on the oceanic circulation is secondary tothe indirect e!ect of the changes in the hydrological cycle(Weaver et al. 1998). In the following period T3 (a further3,713 years of integration), all atmospheric fresh waterfluxes are discharged into the ocean again; however, thepreviously accumulated continental ice sheets and thesurface albedo over continental ice remain constant. Thisphase simulates a stationary ice sheet during full glacialconditionswhere accumulation and ablation balance eachother, and where the cover and thickness of the con-tinental ice sheets are assumed constant. For simplicity,we employ the same run-o! scheme during all phases,although the changing land surface elevation due to icesheets would redirect certain fresh water pathways intothe ocean.

An additional 2.847 Æ 1016 m3 of fresh water is storedon the continents during experiment T3. Flint (1971)quantifies the maximum amount of fresh water accu-mulated in continental ice sheets as 7.056 Æ 1016 m3. Thedi!erence between this figure and the observed amounttoday (2.406 Æ 1016 m3) is 4.65 Æ 1016 m3. In period T3 themodel thus contains approximately 60% of thecontinental ice sheets during the LGM. This is

Fig. 9 Annual mean zonally integrated meridional heat transport,in Petawatts, from the coupled system. The total oceanic andatmospheric heat transports are represented by the solid and dottedlines, respectively. Transports of sensible (dashed line) and latent(dashed-dotted line) heat by atmospheric transient eddies are shownas well

Fig. 8 Concentration of anartificial tracer (see text) andvelocities (in cm s–1) at a depthof 1,950 m at the end of thecontrol run


approximately the ice volume acquired during isotopestage 5d (Chappell and Shackelton 1986).

5 Results

During the buildup of continental ice sheets, the freshwater fluxes at the ocean surface are considerably re-duced. The surface salinity increases during T2 whichcauses a destabilization of the water column and en-hanced deep water formation in the northern NorthAtlantic. The Atlantic overturning cell (which is a proxyfor deep water production) during T1, T2 and T3 isrepresented in Fig. 13. At the end of T2 the maximum ofthe NADW cell exceeds 24 Sv (Fig. 13b), 6 Sv more thanin the control run (Fig. 13a). The southward export

ranges between 14 Sv and 16 Sv, strengthened by about4 Sv compared to T1. The overturning cell extends to adepth of 4,000 m during T2, virtually filling the abyssalAtlantic (Fig. 13b).

SST (Fig. 14c) and SSS (Fig. 14d) distributions dur-ing T2 are similar to those during period T1. The sea icemargin(visible as the –2 "C isotherm in Fig. 14c) in theNordic Seas occupies a similar position as at the end ofT1. However, the northwestern North Atlantic featureshigh temperatures and salinities. Thus, the pronouncedzonal gradient of salinity in the reference run can nolonger be detected (compare Fig. 14b). The verticaldensity distribution (not shown) indicates that convec-tion sites in the North Atlantic have shifted westwardinto the Labrador Sea and the western Greenland Sea.This shift is reflected in the SST and SSS patterns.

Fig. 10 Annual mean atmo-spheric surface temperature, in"C, at the end of the control run

Fig. 11 Annual mean precipi-tation minus evaporation, inmm per year, at the end of thecontrol run


The salinity at the end of T2, zonally averaged overthe Atlantic basin, is represented in Fig. 15b. The for-mation of NADW is clearly visible between 60"N and70"N and extends to great depths, filling the Atlanticbasin as far south as 60"S. The intensified formation ofNADW leads to rising temperatures between 1,000 mdepth and the bottom. In the polar regions the salinityincreases and the halocline retreats to high northernlatitudes. This feature can also be seen in Fig. 14d,where the strong salinity gradient between polar regionsand mid-latitudes is displaced northwards to 70"N.

In experiment T3 the amount of inland ice remainsconstant and all of the precipitation over the continentsis discharged into the ocean as in experiment T1. Thesurface layer of the northern Atlantic becomes fresherand the water column is thus stabilised, resulting in areduction in deep water formation. Figure 13crepresents the meridional overturning streamfunctionduring T3. The maximum of the cell does not exceed 12Sv, being reduced by about 14 Sv in comparison withT2, and is even 8 Sv weaker than the maximum of thereference run T1. The southward transport does notreach 8 Sv, 8 Sv and 4 Sv weaker than during T2 and T1,respectively. A similar reduction was found by Weaveret al. (1998). Additionally, the overturning cell is shal-low, reaching only 2,000 m depth.

The regions of deep water formation shift south-eastwards during T3, moving west of Ireland. In theeastern part of the northern Atlantic, the salt tonguespreads up to 70"N (just as far as during T2). Contraryto the results from T2, SST (Fig. 14e) and SSS (Fig. 14f)are characterised by strong zonal gradients in the NorthAtlantic. The western part of the northern North

Atlantic is covered by sea ice north of 45"N. The polarhalocline is located as far south as 42"N in this region. Aweak overturning cell transports less warm water fromthe subtropics into the North Atlantic. SST decreasesand allows sea ice formation in the Labrador Sea and o!the north-eastern coast of North America. In sea icecovered regions a lack of evaporation, fresh water inputsin the form of continental run-o!, and precipitationcause decreasing salinity in the upper ocean layer.

Fig. 12 Surface albedo for full glacial conditions that are reachedat the end of integration phase T2 (see text)

Fig. 13a–c Annual mean of the meridional overturning motion inthe Atlantic in Sv: a year 3,707 (experiment T1), b year 8,139(experiment T2), c year 12,127 (experiment T3)


Another consequence is an oceanic heat loss reduction inice covered regions and hence a further weakening of theTHC. However, even during full glacial conditions(experiment T3) we find open water in the eastern partof the Nordic Seas. While this contradicts earlierreconstructions (CLIMAP Project Members 1981), it isconsistent with the recent estimates of Sarnthein et al.(1994) and Vernal et al. (2000).

The zonally averaged salinity over the Atlantic basinat the end of T3 is shown in Fig. 15c. A saline watermass, already formed during T2, occupies the entiredeep ocean basin below 3,000 m. The high density of thiswater inhibits vertical exchange; therefore, the deep partof the Atlantic basin is no longer ventilated, forming anisolated deep water mass below 3,000 m depth. NADW

formed during T3 is fresher than that of T2; its char-acteristic salinity is 34.7 psu and it fills the basin between1,000 m and 3,000 m depth. In this region the potentialtemperature is higher than in T2; during T3 the con-vection sites have shifted to the south and the NADWbecomes warmer.

6 Discussion

Initiation of a glaciation through the continental storageof fresh water over the North American and Scandina-vian land masses permits the reproduction of somecharacteristics of the glacial oceanic circulation. The

Fig. 14a–f Annual mean seasurface temperatures in "C:a year 3,707 (experiment T1),c year 8,139 (experiment T2),and e year 12,127 (experimentT3); and annual mean seasurface salinities in psu: b year3,707 (experiment T1), d year8,139 (experiment T2), andf year 12,127 (experiment T3)


results of integration period T3 are in good agree-ment with numerous paleoclimatological investigations.Lynch-Stieglitz et al. (1999) utilise oxygen-isotope ratiosof benthic foraminifera, which lived along the oceanmargins along the Florida Current during the LGM, toshow that the overturning cell was indeed weaker duringglacial times. On the other hand, Yu et al. (1996) deductonly a relatively small decrease in the oceanic large-scaleoverturning motion based on 231Pa/230Th ratios in sedi-ment cores. Rutberg et al. (2000) point out that enhanced

biological activity along the North Atlantic continentaland ice margins could reduce the signal in the radio-chemical data, hiding a glacial-interglacial reorganisa-tion of the THC (see also Boyle and Rosenthal 1996).Previous investigations with coupled ocean-atmosphere-sea ice models (Weaver et al. 1998), as well as modelstudies of Fichefet et al. (1994) and Duplessy et al. (1996)driven by LGM data (CLIMAP Project Members 1981),all find a weakening of the THC. Fichefet et al. (1994)find a maximum of the overturning cell of 10 Sv, reducedfrom 16 Sv in their control run. The overturning cell isshallower during the simulation of the LGM and reachesonly a depth of 2,000 m. As in our results, the convectionsites shift to the south in this model simulation, with noformation of NADW north of 55"N, unlike the presentclimate. Ganopolski et al. (1998) use a global coupledocean-atmosphere model of intermediate complexity tosimulate the equilibrium climate of the LGM. Althoughtheir overall rate of NADW formation is only slightlyreduced in comparison to their control run, other char-acteristics, such as a southward shift of convection sites,are similar to our results.

The lack of AABW in the control run and the presentmodel’s di"culty in simulating dense water formationover Antarctic shelves lead to an underestimate of gla-cial AABW. It is conceivable that the minor response ofAABW to the glacial changes is an artifact of the model.It should be kept in mind, however, that the only glacialforcing employed is the growth and subsequent main-tenance of Northern Hemisphere continental ice masses.

An important result of our investigations is the re-inforcement of the THC during the glacial transition(experiment T2). The transition period is essential forthe glacial circulation state in that dense water is se-questered in the deep ocean basins during this phase. Inthe following glacial state such dense water is no longerformed at the surface and the deep water masses re-main unventilated. The glacial state is not stationarybecause di!usion will slowly erode the deep stratifica-tion and eventually allow newly formed water to againreach the deepest parts of the ocean. This kind of be-haviour would include violent flushes of the deep oceanassociated with sudden increases in the strength of theoverturning (Weaver 1999). In this context, the di!u-sivities in the deep ocean become important. Recentresults from tracer release experiments (Ledwell et al.2000) indicate that di!usivities are small in the interiorof the ocean, but can reach large values over roughtopography in the deep basins. Since tides arepresumably the source of the mechanical energy (Munkand Wunsch 1998), this kind of mixing does notdepend on the mean large scale circulation itself. Onthe other hand, Webb and Suginohara (2001) point outthat upwelling of NADW in the Southern Oceanreduces the need for relatively large diapycnal mixingin the interior of the ocean that was long thoughtessential to close the THC.

Ruddiman et al. (1980), utilising oxygen isotope data,argue that the North Atlantic Ocean (from at least 40"N

Fig. 15a–c Annual mean salinity zonally averaged over theAtlantic basin in psu: a year 3,707 (experiment T1), b year 8,139(experiment T2), and c year 12,127 (experiment T3)


to 60"N) maintained warm SSTs during much of therapid inland ice growth as the climate system moved intoa fully developed glaciation. Summer SST and SSS atdi!erent subtropical and subpolar North Atlantic siteseven increased during the initial phase of continental icebuildup (Cortijo et al. 1999). Ice sheet growth during thetransition from marine isotope stage 5e to 5d precededsurface cooling and the shut down of the THC in theAtlantic (Broecker 1998). At 54"N in the eastern NorthAtlantic (deep sea core V29-191), warm surface oceanconditions prevailed much beyond the increase in dO18

in benthic foraminifera which is a proxy for the volumeof inland ice sheets (Kukla et al. 1997). In the north-western Atlantic, southeast of Newfoundland (deep seacore CH69-K09), dC13 in benthic foraminifera increasedduring the period of rising benthic dO18 values at thetransition from stage 5e to stage 5d (Labeyrie et al.1999). This indicates increasing strength of the mer-idional overturning circulation in the North Atlanticduring the growth of the ice sheets. These findings areconsistent with the scenario developed here (inland icegrowth phase T2). The THC increases during the initialgrowth of inland ice sheets; it is reduced only later whenfresh water run-o! again reaches the ocean at the equi-librium rate.

There is little direct evidence of an increase in thestrength of the THC during ice age inceptions. Increasingsurface temperatures in the northern North Atlanticcould also be due to a redirection of ocean currents;however, other model results show the potential for en-hanced THC during interglacial – glacial transitions andunder glacial conditions. Wang and Mysak (2000) si-mulate an initiation of glaciation with a coupled atmo-sphere-ocean-sea ice-land surface model by slowlyreducing the solar radiation and by increasing theplanetary emissivity in the northern high latitudes.Although their approach is quite di!erent from the ex-perimental design presented here, the results are similar.When land ice is growing, the THC of their zonallyaveraged ocean model intensifies, resulting in a warmsubpolar North Atlantic Ocean. The additional north-ward heat transport associated with the stronger THC isinsu"cient to reverse the process of northern glaciation.With LGM atmospheric forcing, Bigg et al. (1998) findan oceanic ‘‘northern sinking state’’ (NSS) whereNADW formation was much stronger than today. Theyargue that this state is consistent with paleoclimate data.Bigg et al. (1998) speculate that the NSS is representativeof the glacial ocean between Heinrich events. A state withmuch weaker overturning circulation in the Atlantic isrepresentative of the ocean immediately after fresh waterevents. On the other hand Khodri et al. (2001) find areduction in North Atlantic meridional overturning in acoupled ocean-atmosphere simulation where the insola-tion is changed to correspond with the Earth’s orbitalparameters at 115 kyr BP. This is due to enhanced mer-idional moisture transport in the atmosphere, a weakerIcelandic low, and presumably higher winter tempera-tures in the northern North Atlantic. As only a short

integration (100 years) is performed and no changes ininland ice mass are considered, the e!ect discussed here isneglected in their study.

There is an apparent discrepancy between ourmodel results and observations in that the Nordic Seascooled and freshened well before the major ice sheetsdeveloped (Duplessy and Labeyrie 1992; Cortijo et al.1994; Fronval and Jansen 1996). Cooler conditions inhigh northern latitudes are not necessarily inconsistentwith a stronger THC; a shift in the major deep waterformation sites to south of the Greenland-Scotland-Ridge could easily account for a cooling at highnorthern latitudes while the subpolar North Atlanticwould stay warm and saline. On the other hand, it ispossible that not all relevant geographical details arereproduced in a coarse resolution model such as ours.The still relatively low Greenland-Scotland-Ridge inthe model could lead to an unrealistically large inflowof Atlantic water into the Nordic Seas when the THCis strong. Another factor is the simple assumptionregarding the changes in continental run-o! related tochanges in inland ice volume. Fronval and Jansen(1997) find that cooling of the Nordic Seas and initialice buildup on the surrounding continents commencedwithin the Eemian, before the main increase in globalice volume. Thus, the run-o! history into the NordicSeas di!ers from that of the subpolar Atlantic. Run-o! into the Nordic Seas may have resumed fullstrength after the buildup of local ice sheets before theend of the Eemian, generating a stable halocline thatsuppressed sinking in the Nordic Seas. This is notcaptured by our assumptions about the run-o! historyand would require a much more sophisticated inlandice model.

7 Conclusion

In this study a coupled atmosphere-ocean-sea ice modelconsisting of a two dimensional EMBM, which calcu-lates both temperature and moisture prognostically,coupled to a thermodynamic sea ice model and anocean general circulation model (the GFDL MOM2,Pacanowski 1995) is utilised to investigate the climaticimpact of the changes in fresh water balance during anice age initiation. While other processes (e.g. themodification of the wind field by the continental icesheets (Mayewski et al. 1997; Andersen et al. 1998),and changes in the carbon cycle (Petit et al. 1999) arenot taken into account, we perturb the fresh waterbalance in order to examine the e!ect of the modifiedfresh water flux as an isolated mechanism. During thebuildup of ice sheets (our experiment T2) the coupledmodel shows an enhancement of the thermohaline cir-culation (THC) leading to a warm and saline NorthAtlantic, in agreement with observations (Ruddiman etal. 1980; Kukla et al. 1997; Labeyrie et al. 1999). In asecond experiment, a simulation of an ice age climate isinduced by allowing all precipitation over continents to


discharge into the ocean again. The ice sheet (i.e. thesurface albedo and the stored fresh water) remainsconstant during this period (our experiment T3). Al-though integrating under similar boundary conditionsas in the control run (experiment T1), the model re-produces main characteristics of the glacial circulation.The redistribution of salt within the ocean during T2leads to a weakening of the THC and a shift of theconvection sites south-eastwards.

Acknowledgements The authors would like to thank Holger F. G.Brix and Augustus F. Fanning for many useful comments on anearlier version of this manuscript. We thank Gokhan Danabasoglufor providing the bottom topography data from his model.Furthermore, we would like to thank the three anonymousreviewers of this manuscript for their constructive comments.Daithı A. Stone is gratefully acknowledged for editing Englishgrammar.

Appendix A

The atmospheric energy and moisture balance model

A1 Energy and moisture balance

The deduction of the following equations of mass conservation andenergy balance can be found in Lohmann and Gerdes (1998) whileconstants and parameters of the atmospheric model can be foundin Table 3.

The mass conservation of water vapour qv is expressed by

Hqqair@qv@t

!Z

dpgrp " #uhqv$ % E & P #1$

where qair is the air density in the moisture containing layer, g thegravity acceleration, p the pressure, #p the gradient operator forconstant pressure p, uh the horizontal velocity field, E the eva-poration, and P the precipitation. Hq denotes a constant scaleheight for the specific humidity (see Table 3).

The vertically integrated balance for the atmospheric tempera-ture can be written as in Lohmann and Gerdes (1998)Z

dpgCp

@Ta@t

!Z

dpgCprp " #uhTa$ % Qt

R &QbR !QS ! LvP #2$

where Cp is the specific heat at constant pressure, Ta the atmo-spheric temperature, Qt

R and QbR the net radiation at the top of the

atmosphere (incoming shortwave radiation minus outgoing

shortwave and longwave radiation) and at the interface betweenocean or sea ice and atmosphere, respectively, QS the sensibleheat flux and Lv the latent heat of condensation. The heat andfresh water fluxes between ocean, sea ice and atmosphere arecomputed using the atmospheric surface temperature Ta(p0) whilethe prognostic equations are vertically integrated. An estimationproposed by Lohmann and Gerdes (1998) of Ta(p0) as a functionof the vertically integrated temperature is utilised, in which thevariation of bottom temperature dTa(p0) is related to the changeof the vertically averaged temperature (Rennick 1977; Chen et al.1995) (d represents a small variation):

dZ

dpgTa#p$ % b1dTa#p0$ #3$

b1 %5708 kgm&2 ; Ta#p0$ ' 275K

3966 kgm&2 Ta#p0$ < 275K

(

The e!ective change of the humidity content of the atmosphere inrelation to the surface temperature Ta(p0) is parametrised by:

dZ

dpgqv#Ta; p$ % b2dTa#p0$ #4$

The model neglects the energy exchanges in the energy balancedue to phase conversion in the global energy balance. This ap-proximation can be justified by applying a scale analysis (Lvb2 .10–13 Jm–2K–1, Cpb1 . 103 Jm–2K–1). Therefore, the atmosphericenergy balance is given by

Cpb1@Ta#p0$

@t!Z

dpgCprp " #uhTa$

!Z

dpgLvrp " #uhqv$ % Qt

R ! Foa ! Fia #5$

where p0 is the reference pressure and Foa and Fia are the heatfluxes at the interface between ocean or sea ice and atmosphere,respectively.

The atmospheric model is thus governed by two prognosticequations determining moisture Eq. (1) and surface temperatureEq. (5).

A2 Transport parametrisation

The poleward transport of heat at low latitudes is primarily ac-complished by the Hadley cell. The prevailing process in middleand high latitudes is baroclinic instability (Peixoto and Oort1992). The di!erence between these modes of transport can bedemonstrated by splitting the velocity, temperature, or humidityvariables into large scale, long term quantities (uh; Ta; qv) and their

Table 3 Parameters of theatmospheric model Parameter Description Value Source

aa Albedo of a cloudless atmosphere 0.08 3es Emissivity of sea ice 0.97l Bulk coe"cient 16 Wm–2K–1 2qair Air density 1.225 kgm–3

r Stefan-Boltzmann constant 5.67 Æ10–8 Wm–2 K–4

s Time step 864 sAa Absorptivity of the atmosphere 0.18 3cE Bulk coe"cient 1.5 Æ10–3 1cH Bulk coe"cient 0.83 Æ10–3 1Cp Specific heat at constant pressure 1,004 Jkg–1K–1

g Gravity constant 9.81 ms–2

Hq Scale height 1,800 m 4Lv Latent heat of condensation 2.5 Æ106 Jkg–1

p0 Reference pressure 1,000 hPa

Sources: 1, Gill (1982); 2, Grosfeld (personal communication); 3, London (1957); 4, Peixoto and Oort(1992)


deviations (u0h; T0a; q

0v). Here, the large-scale transport due to

transient eddies is parametrised based on the theory of baroclinicinstability (Green 1970; Stone 1972), whereas the mean circulationis held fixed to climatological data (Kalnay et al. 1996; Da Silvaet al. 1994). The sensible eddy heat transport is calculated interms of the horizontal gradient of the surface temperature Ta(p0):Z

dpgCpu0hT 0

a % &CpKs#/$jrpTa#p0$jrpTa#p0$ #6$

where / denotes latitude. The transport of moisture by transienteddies is parametrised in analogous manner to the sensible eddyheat transport (Vallis 1982):Z

dpgu0hq0v % &KL#/$jrpTa#p0$jrpqv#p0$ #7$

The time constant components Ks#/$ and KL#/$ ofthe di!usion coe"cients Ks#/$jrpTa#p0$j and KL#/$ jrpTa#p0$j arefitted to reproduce the present-day climate, with air moisture andtemperature values taken from the NCEP/NCAR reanalysis(Kalnay et al. 1996) and the zonally averaged data for energytransport by transient eddies from Peixoto and Oort (1992).Taking into account that the variance of the zonal wind com-ponent in transient eddies has the same magnitude as the varianceof the meridional component (Oort and Rasmusson 1971), weassume that the atmosphere is isotropic and therefore take thesame coe"cients for meridional and zonal transports. The valuesof these coe"cients are displayed in Fig. 16.

A3 Sinks and sources

The sinks and sources for the atmospheric model are representedon the right sides of Eqs. (1) and (5). They occur at the top of theatmosphere and at the ocean or sea ice interface.

A3.1 Radiation balance

On top of the atmosphere the shortwave incoming radiationQtop;down

SW #/$ is computed daily (Berger 1978). The reflected short-wave radiation at the top of the atmosphere Qtop;up

SW #k;/; t$ is afunction of longitude (k), latitude (/) and time (t). It is calculatedusing the albedo of the whole earth-atmosphere system (planetaryalbedo ap(k, /, t)), the latter consisting of the albedo of clouds ac(k,/), the albedo of a cloudless atmosphere aa (due to air molecules,dust and water vapour scattering), the absorptivity of the atmo-sphere Aa, and the surface albedo as(k, /, t) (Haney 1971)

ap % aa ! ac ! as!

1& #aa ! ac ! Aa$"

#8$

The albedo of clouds, the albedo of the cloudless atmosphere, andthe absorptivity of the atmosphere are temporally constant (aa =0.08 and Aa = 0.18, London 1957). The surface albedo is computedas a function of sea ice or snow cover with the isotherm Ta(p0) = –10 "C as the limit of snow cover over continents (Budyko 1969).This parametrisation is only appropriate for annual mean tem-peratures; if the seasonal cycle is taken into account, sea ice andsnow cover correspond to higher values of the surface temperature,typically –2 to –4 "C (DeBlonde et al. 1992). Figure 17 shows thezonal averages of the surface albedo and the sum of atmosphericalbedo and absorptivity in comparison with the values of Haney(1971).

The net longwave radiation QtopLW is modelled by considering the

planet as a grey body with an emissivity ep(k, /). The infraredemission is then given as in Fanning and Weaver (1996) andStocker et al. (1992) by

QtopLW % ep#k;/$rTa#p0$4 #9$

where r = 5.67 Æ 10–8 Wm–2K–4 is the Stefan-Boltzmann constant.The zonal mean of the tuned emissivity is represented in Fig. 18. Itis determined by fitting it to climatological data of the NCEP/NCAR 40-Year Reanalysis Project (Kalnay et al. 1996).

Fig. 16 Di!usion coe"cients used to calculate the heat transportsby transient eddies. Grey: jl(/) = KL(/) Æ Lv/a

2 Æ 10–7, black:js(/) = Ks(/) Æ cp/a

2 Æ 10–4

Fig. 17 Zonally averaged albedo. Values according to Haney(1971) are represented by the dotted line; the solid line contours thealbedo used by the atmospheric model

Fig. 18 Zonally averaged planetary emissivity ep. The values givenby Stocker et al. (1992) and Fanning and Weaver (1996) arerepresented by the dash-dotted and the dotted lines, respectively.The solid line represents the emissivity used in this study


The radiation balance at the top of the atmosphere is then givenby

QtR % #1& ap$Qtop;down

SW & QtopLW : #10$

A3.2 Fluxes at the ocean-sea ice-atmosphere interface

The heat flux between atmosphere and ocean (Foa) or sea ice (Fia) iscomposed of incoming shortwave radiation Qi, net longwave ra-diation QB, sensible heat flux QS, and the heat loss due to evapora-tion LvE over ice free regions:

Foa;ia % &Qi ! QB ! LvE ! QS :

Qi is given by Haney (1971) as:

Qi % #1& as$Qtop;downSW

!

1& #aa ! ac ! Aa$"

#11$

For the longwave radiation di!erent formulations have to be takeninto account depending on the nature of the surface. We chose theexpression from Gill (1982) for ice free regions:

QB;water % 0:9858 " rT 4s #0:39& 0:05

#####

eap

$#1& 0:6n2c$ #12$

where Ts represents the SST in "C, ea the vapour pressure of watervapour, and nc the cloud cover (nc=ac/0.6, Haney 1971). Over sea iceQB is calculatedwith the parametrisationof Idso and Jackson (1969):

QB;ice % es#rT 4is & QLW ;down$ #13$

QLW ;down % rT 4a #p0$

$

1& 0:261 exp

%

&7:77 " 10&4

( #273& Ta#p0$$2&'

#1! 0:275nc$ #14$

Tis indicates the ice surface temperature and es the emissivity of seaice (es=0.97, Fischer 1995). Evaporation and sensible heat flux overocean are computed with bulk formulas (Gill 1982):

QS;water % qairCpcHu!

Ts & Ta#p0$"

#15$

E % qaircEu#qs & qv$ #16$

where u denotes the wind speed at 10 m height, qs the saturatedhumidity forTa(p0)=Ts, and qv the water vapour. cH and cE are bulkcoe"cients (cH = 0.83 Æ10–3 and cE = 1.5 Æ10–3, Gill 1982). Over seaice the sensible heat flux is computed as

QS;ice % l!

Tis & Ta#p0$"

#17$

where l is the atmospheric constant for heat flux (l=16 Wm–2 K–1,Grosfeld personal communication). Precipitation is parametrised as inFanning and Weaver (1996) and Weber (1998):

P % max

$

qairHq

s

%

qv & rhmax " qs#Ta$&

; 0

'

#18$

where rhmax = 0.85 represents the maximum relative humidity and sthe length of the time step.

A4 Numerics

The meridional and zonal grid spacing is 10"and the integration isperformed with the Euler scheme with centring spatial di!erences.In spite of the small time step needed by the explicit integrationscheme, the atmospheric model uses only 6% of the integrationtime of the entire system.

Appendix B

The sea ice model

The parameters of the sea ice model are summarised in Table 4.The energy balance is split into two terms (Semtner 1976). Re-garding the atmospheric interface, we can write:

qiceLf

%

@h@t

&

a% Fia & Qc #19$

qice density of sea ice, Lf the latent heat of fusion, h the ice thickness,Fia the heat flux between atmosphere and sea ice, and Qc the heatflux through the ice. The index a denotes the atmospheric interface.There is no freezing on the sea ice surface (Fia–Qc £ 0). If the icesurface temperature Tis is below the surface freezing point (0 "C),there is no melting (Fia–Qc = 0). The energy balance for theoceanic interface can be expressed by:

qiceLf

%

@h@t

&

w% Qc & Qw #20$

where Qw is the heat flux from the ocean into the sea ice and theindex w indicates the oceanic interface. The dynamics of sea ice arenot considered in this model, and so the energy balance for the seaice can be expressed by:

qiceLf@h@t

% qiceLf

%

@h@t

&

a! qiceLf

%

@h@t

&

w% Fia & Qw #21$

Computing the surface temperature of sea ice, we distinguish twocases. Melting on the surface:

Tis % 0oC and Qc ) Fia #22$

otherwise:

Tis < 0oC and Qc % Fia #23$

The heat fluxes Qc and Fia are functions of Tis; the surface tem-perature of sea ice can thus be calculated from these equations usingthe Newton-Raphson iteration method (Press et al. 1992). The heatflux between ocean and sea ice is parametrised with a bulk formula.The coe"cient c (c = ChÆDu) is the product of a coe"cient Ch

(Ch=3Æ10–4, Omstedt and Wettlaufer 1992) and the di!erencebetween sea ice and ocean velocities Du. This di!erence is assumedto be constant in this study (Du = 10 cms–1) Grosfeld et al. 1997.The bottom temperature of the sea ice is constant (Tib = –1.9 "C).qw andCpw are the density and the specific heat capacity of sea water,

Table 4 Parameters of the seaice model Parameter Description Value Source

as Surface albedo for sea ice 0.7c Coe"cient for turbulent mixing 3 Æ10–5 ms–1 1qice Sea ice density (5 psu, –10 "C) 873 kgm–3

qw Sea water density 1,028 kgm–3

s Time step 43,200 sCpw Specific heat for sea water 3,950 Jkg–1K–1

Ki Coe"cient for di!usion 2.04 Wm–1K–1 2Lf Heat of fusion 345 kJkg–1

Tib Bottom temperature of the sea ice )1.9"C

Sources: 1, Ohmstedt and Wettlaufer (1992); 2, Unterseiner (1964)


respectively. Furthermore, we assume the coe"cient of di!usion forthe heat flux through the sea ice to be constant Ki = 2.04 Wm–1K–1

(Untersteiner 1964).

Fia % &Qi ! QB;ice ! QS;ice #24$

Qc % Ki#Tib & Tis$

h#25$

Qw % qwCpwc#Tib & Ts$ #26$

If the sea ice thickness exceeds a threshold of Hmax=4 m, the ice isassumed to be insulating (Qc = 0).

The ocean model uses a rigid lid as an upper boundary condi-tion and hence requires the conversion of freshwater fluxes toequivalent salt fluxes. Supposing that the sea ice consists of freshwater, the salt flux Fs can be expressed by

Fs % Soqice@h@t

#27$

with So being the salinity of sea water in psu. The fractional icecover in a grid box is defined as a linear function of the icethickness. An ice thickness of 1 cm corresponds to an ice cover ofone percent. The ice cover is 100% when the ice thickness reaches1 m.

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K. J. Meissner Æ R. Gerdes Coupled climate modelling of ocean circulation...

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