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The sensitivity of the climate response to the magnitude and locationof freshwater forcing: last glacial maximum experiments

Bette L. Otto-Bliesner*, Esther C. BradyNational Center for Atmospheric Research, 1850 Table Mesa Drive, Boulder, Colorado 80305, USA

a r t i c l e i n f o

Article history:Received 31 January 2009Received in revised form3 June 2009Accepted 6 July 2009

a b s t r a c t

Proxy records indicate that the locations and magnitudes of freshwater forcing to the Atlantic Oceanbasin as iceberg discharges into the high-latitude North Atlantic, Laurentide meltwater input to the Gulfof Mexico, or meltwater diversion to the North Atlantic via the St. Lawrence River and other easternoutlets may have influenced the North Atlantic thermohaline circulation and global climate. We haveperformed Last Glacial Maximum (LGM) simulations with the NCAR Community Climate System Model(CCSM3) in which the magnitude of the freshwater forcing has been varied from 0.1 to 1 Sv and insertedeither into the subpolar North Atlantic Ocean or the Gulf of Mexico.

In these glacial freshening experiments, the less dense freshwater provides a lid on the ocean waterbelow, suppressing ocean convection and interaction with the atmosphere above and reducing theAtlantic Meridional Overturning Circulation (AMOC). This is the case whether the freshwater is addeddirectly to the area of convection south of Greenland or transported there by the subtropical andsubpolar gyres when added to the Gulf of Mexico. The AMOC reduction is less for the smaller freshwaterforcings, but is not linear with the size of the freshwater perturbation. The recovery of the AMOC froma ‘‘slow’’ state is w200 years for the 0.1 Sv experiment and w500 years for the 1 Sv experiment.

For glacial climates, with large Northern Hemisphere ice sheets and reduced greenhouse gases, thecold subpolar North Atlantic is primed to respond rapidly and dramatically to freshwater that is eitherdirectly dumped into this region or after being advected from the Gulf of Mexico. Greenland tempera-tures cool by 6–8 �C in all the experiments, with little sensitivity to the magnitude, location or duration ofthe freshwater forcing, but exhibiting large seasonality. Sea ice is important for explaining the responses.The Northern Hemisphere high latitudes are slow to recover. Antarctica and the Southern Ocean showa bipolar response, with warming and reduced sea ice. This warming continues after the cessation of thefreshwater forcing and shows a dependence on the duration of the freshwater forcing.

Equatorward of the expanded sea ice, the simulated temperature and salinity anomalies are sensitiveto the amount of colder and fresher waters that are advected out of the subpolar North Atlantic. In thetropical Atlantic, the recovery of the Intertropical Convergence Zone (ITCZ) from its more southerlyposition during the freshwater forcing is much more rapid than the recovery of the AMOC, and is morerelated to the recovery of low-latitude surface temperatures than Greenland temperature or sea ice.These results have implications for using proxy records as indirect measures of the AMOC.

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1. Introduction

Proxy records indicate large freshwater forcings to the AtlanticOcean basin during the last glacial interval. Ocean cores in thesubpolar North Atlantic between 40 and 55�N latitude displaydistinct layers of lithic grains in the sand fraction during seven timeintervals, so-called Heinrich events, during the time period

between 70 ka and 13 ka (Andrews, 1998; Broecker et al., 1992;Heinrich, 1988; Ruddiman, 1977). This ice-rafted debris (IRD) iscomposed of rock fragments with provenance likely from aroundthe Hudson Strait (but also with sources from other parts of theLaurentide ice sheet and from the Fennoscandian ice sheet), thatwas incorporated into icebergs or sea ice, transported across theNorth Atlantic, and deposited to the sea floor as the ice melted(Bond et al., 1992; Hemming, 2004). Two Heinrich events occurnear LGM: H2 with the magnetic susceptibility records suggestingdepositional phases centered at 23.5 ka and 25 ka occurred duringthe maximum extent of the ice sheets during the last glacial cycle,

* Corresponding author. Tel.: þ1 303 497 1723, fax: þ1 303 497 1348.E-mail address: [email protected] (B.L. Otto-Bliesner).

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Quaternary Science Reviews 29 (2010) 56–73


and H1 with depositional phases centered at 16 ka and 17.5 kaduring the early part of the deglaciation (Bard et al., 2000). Upliftedcoral reefs in Papua New Guinea suggest sea level excursions of 10–15 m corresponding to Heinrich events (Chappell, 2002; Yokoyamaet al., 2001). Hemming (2004) estimated from Heinrich layerrecords that the flux of meltwater ranged from 1.6 Sv over one yearto 0.3 Sv over 500 years, though with considerable uncertainty.Through modeling of oceanic d18O perturbations, Roche et al.(2004) obtain a duration of 250�150 yr and an ice release of2�1 m sea level equivalent for Heinrich event 4 occurring at about40 ka.

Abrupt meltwater pulses associated with ice sheet collapses, butwithout IRD layers, are also apparent in the geologic record of sealevel rise during the last deglaciation. Marine sediments ina shallow gulf along the Australian margin indicate a rapid rise ofsea level at 19 ka of 10–15 m in 100–500 years or 2 Sv to 0.25 Sv(Clark et al., 2004; Yokoyama et al., 2000). Meltwater pulse (MWP)1A occurring w14.5 ka has an estimated 20 m sea level rise overseveral hundred years (Bard et al., 1990; Clark et al., 2002a).Whether this meltwater came from widespread melting of theAntarctic or Northern Hemisphere ice sheets, or both, is still muchdebated (Clark et al., 2002a; Peltier, 2005). The cause of the YoungerDryas cold event, starting at w12.9 ka, is also highly debated, withthe preponderance of evidence suggesting the trigger to be fresh-water from Lake Agassiz flowing the North Atlantic, thoughconsensus on the pathway has not be reached (Broecker, 2003;Tarasov and Peltier, 2005; Broecker, 2006; Carlson et al., 2007).

In the Atlantic, the ocean conveyor circulation transports warm,saline subtropical waters northwards in the upper layers. At highlatitudes in the North Atlantic, the ocean loses heat to the atmo-sphere, becoming denser and sinking, particularly in theGreenland–Iceland–Norwegian (GIN) and Labrador Seas, with flowsouthward at deeper levels. This ocean circulation redistributesheat gained in the South Atlantic and tropics to the North Atlanticregion. It has been hypothesized that freshwater addition duringthe glacial period and deglaciation would have shutdown the NorthAtlantic part of the thermohaline circulation with influences on theglobal climate (Broecker et al., 1988; Broecker, 1994; Broecker,2003; Clark et al., 2002b).

A number of proxy records for the surface and deep oceansupport this interpretation. Measurements of 231Pa/230Th in coreGGC5 from the Bermuda Rise indicate a slowdown of the AtlanticMeridional Overturning Circulation (AMOC) starting at w19 ka,with a nearly complete shutdown from 17.5 ka until 15 ka(McManus et al., 2004). The onset of the slowdown correspondswith the meltwater pulse at w19 ka and the shutdown is concur-rent with the first of two depositional events identified with H1(Bard et al., 2000). A proxy record of magnetic susceptibility fromthe Eirik Drift, interpreted as a measure of North Atlantic DeepWater (NADW) flow intensity, confirms the near shutdown of theAMOC for 2000–3000 years during H1 (Stanford et al., 2006). Highresolution benthic foraminifera d13C records from 2000-m depthsin the North Atlantic also suggest perturbations in the AMOCduring Heinrich events, with major changes in the deep-watermasses, consistent with large reductions of deep-water formationand northward migration of southern ocean deep waters to 62�N(Elliott et al., 2002).

Sediments in cores from the Iberian margin and westernMediterranean indicate much colder sea surface temperature (SST)in phase with Heinrich events (Bard et al., 2000; Cacho et al., 1999).Marine core SU8188 from the Iberian margin also displays reducedsalinity during the past three Heinrich events (Bard et al., 2000),attributed to melting of icebergs advected to the margin duringHeinrich events. The salinity decrease during H1 is estimated to be1–2 psu based on the percentage of C37:4 among C37 alkenones and

2–3 psu based on d18O measurements in planktonic foraminiferafrom the same sediments. Surface salinity reductions of 2–3 psu forH1 have also been inferred from d18O for two cores farther north inthe Atlantic and within the IRD (Chapman and Maslin, 1999).Substantial reductions in salinity in the Gulf of Mexico during H1,2–4 psu, have been reconstructed for the Orca basin core (Floweret al., 2004).

Greenland ice cores and records from the nearby North Atlanticalso record millennial climate variability with rapid shifts duringDansgaard-Oeschger (DO) events from cold stadial to warm inter-stadial conditions. In Greenland, their amplitude reaches 8 to 16�Cin a few decades to centuries (Severinghaus and Brook, 1999;Landais et al., 2004; Huber et al., 2006), with shifts in deuteriumexcess from one year to the next during the last deglaciation (Stef-fensen et al., 2008). These DO events have occurred 25 times overthe last climatic cycle (North Greenland Ice Core Project Members,2004) with a bipolar seesaw response first identified for the largeevents (Blunier et al., 1998), then for the first large events of theglacial inception (Landais et al., 2006), and more generally over thepast 50 kyrs (EPICA Community Members, 2006) and for MIS5(Capron et al, 2010). Antarctic ice cores find that they are notrestricted to the last climatic cycle, with over 74 millennial changesin methane over the past 800,000 years (Loulergue et al., 2008).

Modeling results support the interpretation that increasedfreshwater additions during the glacial period and deglaciationwould have slowed or even shutdown the North Atlantic part of theoverturning circulation. Conceptual and simplified numericalmodels suggest that the AMOC may have two stable states, tran-sitioning from a strong AMOC to a weak or collapsed AMOCfollowing a freshwater addition to the North Atlantic (Rahmstorf,1995; Rahmstorf, 1996; Stocker and Wright, 1991). Fully coupledAOGCMs also indicate a slowing of the AMOC when freshwater isadded to the North Atlantic (Stouffer et al., 2006). When thefreshwater forcing is stopped, the AMOC recovers fully over 120–260 years in all but one of the AOGCMs, the GFDL_R30 model whichincluded flux adjustment to the ocean. These CMIP simulationswere done with present-day initial conditions.

Yet, the initial state and boundary conditions may have animportant effect on the response and recovery to freshwaterforcing. Ganopolski and Rahmstorf (2001) find that for the glacialclimate, the Atlantic thermohaline circulation declines graduallywhen freshwater inflow to the Atlantic is increased, in contrast toreaching a clear bifurcation point where the circulation goes to an‘‘off’’ mode for present climate in the CLIMBER model. The hyster-esis behavior for the glacial simulation is also much lesspronounced than in the simulation of present climate. Bitz et al.(2007) find that for glacial conditions the slow recovery from aninstantaneous freshwater pulse is related to the more extensive seaice and more stable density below the surface layer in the NorthAtlantic. In addition, with the Bering Strait closed at LGM, theexport of North Atlantic meltwater takes longer, having only anoutlet through the southern end of the North Atlantic, and thusbringing about a slower recovery of the AMOC (Hu et al., 2008).

There still exists though considerable uncertainty in the timing,location, magnitude and duration of freshwater perturbations intothe ocean during Heinrich and other meltwater events. Here, weuse simulations of the NCAR CCSM3 climate model with glacialboundary conditions to understand the sensitivity to freshwaterperturbations. We examine the regional and transient responses toexplore several questions: (1) What is the response and recovery ofthe AMOC to freshwater forcing? (2) How well do proxy records, inthe North Atlantic and more remotely, correspond to the AMOCresponse to and recovery from freshwater forcing? and (3) Howsensitive are the Atlantic and global responses to the rate, location,and duration of the freshwater perturbations?

B.L. Otto-Bliesner, E.C. Brady / Quaternary Science Reviews 29 (2010) 56–73 57


2. CCSM model and experiments

The National Center for Atmospheric Research (NCAR) CCSM3 isa global, coupled ocean–atmosphere–sea ice–land climate model(Collins et al., 2006a). The model used in this study is the samemodel used at higher resolution for the IPCC AR4 projections offuture climate. The atmospheric model is the NCAR CommunityAtmospheric Model version 3 (CAM3) and is a three-dimensionalprimitive equation model solved with the spectral method in thehorizontal and with 26 hybrid coordinate levels in the vertical(Collins et al., 2006b). The ocean model is the NCAR implementa-tion of the Parallel Ocean Program (POP), a three-dimensionalprimitive equation model in vertical z-coordinate (Gent et al.,2006). For this study, the atmospheric resolution is T42 (anequivalent grid spacing of approximately 2.8� in latitude andlongitude), and ocean grid is 320� 384 points with poles located inGreenland and Antarctica, and 40 levels extending to 5.5-km depth.The ocean horizontal resolution corresponds to a nominal gridspacing of approximately 1� in latitude and longitude with greaterresolution in the Tropics and high-latitude North Atlantic. The landmodel uses the same grid as the atmospheric model and includesa river routing scheme and specified but multiple land cover andplant functional types within a grid cell (Dickinson et al., 2006). Thesea ice model is a dynamic–thermodynamic formulation, whichincludes a subgrid-scale ice thickness distribution and elastic–viscous–plastic rheology (Briegleb et al., 2004). The sea ice modeluses the same horizontal grid and land mask as the ocean model.

The simulations have LGM forcings and boundary conditions, asdescribed by the Paleoclimate Modeling Intercomparison ProjectPhase 2 (PMIP2). The PMIP2 LGM forcings relative to modern arethe small change to insolation resulting from the slightly differentEarth’s orbit, which is set appropriate for 21 ka based on thecalculations of Berger (Berger, 1978), and the reduced concentra-tions of atmospheric carbon dioxide (CO2, 185 ppmv), methane(CH4, 350 ppbv), and nitrous oxide (N2O, 200 ppbv), as adoptedfrom the Greenland and Antarctic ice core records (Dallenbachet al., 2000; Fluckiger et al., 1999; Monnin et al., 2001). The PMIP2boundary conditions for the LGM simulations are the ICE-5G ice

sheet and topography (Peltier, 2004) with large continental icesheets over North America and northern Eurasia (Fig. 1). They alsoinclude the specification of additional land due to the lowering ofsea level by 120–130 m with the large amounts of water frozen inthe continental ice sheets. The lowering of sea level results in theclosing of the Bering Strait and more land in the Arctic. Vegetation,dust aerosols, and river routings are unchanged from the prein-dustrial conditions.

The LGM CCSM3 simulation has a global cooling of 4.5 �Ccompared to preindustrial conditions with amplification of thiscooling at high latitudes and over the continental ice sheets presentat LGM (Otto-Bliesner et al., 2006). Sea ice expands, reaching 45�Nin the western Atlantic in winter, but exhibits strong seasonalitywith the largely ice-free Nordic Seas during summer. CCSM3simulates reduced deep convection in the Nordic Seas andenhanced convection south of Greenland (Fig. 1a), in agreementwith proxies that suggest most GNAIW was likely formed south ofIceland (Duplessy et al., 1988; Pflaumann et al., 2003). The NorthAtlantic meridional overturning circulation weakens by 20% andshoals from a depth of 4 km at 40�N in the preindustrial CCSM3simulation to a depth of 2.4 km in the LGM simulation, withAntarctic Bottom Water (AABW) filling the deep ocean in the glacialNorth Atlantic (Fig. 2a), supporting the interpretation from paleo-nutrient tracers of a shallower North Atlantic Deep Water (NADW)(Lynch-Stieglitz et al., 2007). The deep Atlantic Ocean is muchcolder and saltier at LGM in the model (Otto-Bliesner et al., 2007),as also found in core fluid measurements from ODP cores (Adkinset al., 2002). The North Atlantic subtropical gyre shifts southwardand maintains a vigorous circulation under glacial forcings (Fig. 1b).The Antarctic Circumpolar Current (ACC) in CCSM3 increases byw60% at LGM as compared to preindustrial due to both an increasein the SH westerly wind stress over the Southern Ocean and anincrease in AABW formation and sea ice around Antarctica. Differ-ences in the strength of the SH westerlies is small among the LGMsimulations with the PMIP2 models (Menviel et al., 2008) but thesesimulations show substantial differences in LGM SH sea ice andAABW formation (Otto-Bliesner et al., 2007). Further details of theLGM simulation can be found in Otto-Bliesner et al. (2006).

Fig. 1. CCSM3 LGM control simulation. Annual (a) mixed layer depths (m) and (b) barotropic streamfunction (Sv), with positive indicating a clockwise circulation. Part (a) also showsthe locations of freshwater (blue) added to the ocean as given in Table 1 and the LGM ice sheet extents (purple).

B.L. Otto-Bliesner, E.C. Brady / Quaternary Science Reviews 29 (2010) 56–7358


In order to investigate the climate sensitivity to freshwaterdischarge into the oceans from melting ice sheets during glacialconditions, we have conducted several freshwater forcing experi-ments, so-called ‘‘hosing’’ experiments, using CCSM3 and byvarying amounts of freshwater added to two locations in theAtlantic Ocean. For our initial state, we started with the CCSM3simulation for LGM. As with all freshwater fluxes exchangedbetween the ocean and the other CCSM components, the imposedfreshwater was added to the CCSM3 ocean as a negative virtualsalinity flux to one of two regions: the subpolar North Atlantic andthe Gulf of Mexico (Fig. 1a). The subpolar North Atlantic region isdefined as 50–70�N, the same as the CMIP freshwater hosingsimulations for present-day forcings (Stouffer et al., 2006). FourNATL hosing simulations have been performed, with rates of 0.1–1.0 Sv (Table 1). The freshwater was added for 100 years starting atyear 400 of the LGM simulation. After 100 years, the hosing wasstopped and the climate system allowed to recover. The NATL_1.0

simulation was extended for an additional 500 years, while theother simulations were extended only 100 years past the hosingcessation. Similarly, two GOM hosing simulations, with the wateradded to the Gulf of Mexico, at rates of 0.28 and 0.5 Sv for 100 years,were completed. In addition, a simulation (NATL_0.1ex) in which0.1 Sv was added to the North Atlantic for 500 years was included totest the response to the equivalent volume of water as the NATL_0.5experiment but over a longer period and with a reduced rate.

3. The NATL_1.0 experiment

3.1. Mean annual response at end of freshwater forcing

By the end of the freshwater forcing in the NATL_1.0 experiment,the AMOC is greatly diminished; however it never shuts downcompletely and continues to transport some heat poleward in theNorth Atlantic (Fig. 2b). For the index of the intensity of the AMOCin this discussion, we chose the value of the positive overturningstreamfunction at 34�S, 814 m depth, which was the location of themaximum at 34�S in the LGM control simulation. This valuecorresponds to the transport of upper water that flows northwardinto the Atlantic basin and sinks in high latitudes to form themodel’s North Atlantic Deep Water. As the AMOC slows down,northward ocean heat transport is reduced at all latitudes in theAtlantic and reverses direction to a southward transport of heat atthe equator. The less dense freshwater acts as a lid on the surfaceocean suppressing deep ocean convection and mixed layer depthsare less than 450 m over all of the North Atlantic (not shown). Thisfreshwater lid also suppresses interaction of the deep ocean withthe atmosphere. The annual surface air temperatures cool in theregion of freshwater forcing (Fig. 3). Extensive sea ice covers theNorth Atlantic poleward of 45�N latitude in the annual mean. Thissea ice amplifies the cooling due to its high albedo and by furtherinsulating the atmosphere from the ocean. Greatest cooling ofannual surface air temperatures occurs over the North AtlanticOcean from south of Greenland and Iceland and extends into theGIN Seas, with cooling in excess of 10 �C (Fig. 3). Annual coolingover central Greenland is w6 �C.

A tongue of cooling of the atmosphere and ocean and fresheningof the ocean extends southwestward from Iberia into thesubtropical Atlantic and across the Isthmus of Panama (Figs. 3 and4c, d). This cooling is associated with the export of cold fresh waterfrom the subpolar region and advection equatorward by thesubtropical gyre. The reversal of the ocean heat transport tosouthward at the equator leads to a bipolar response in annualsurface air temperatures with warming in the Southern Hemi-sphere (SH), greater than 1 �C over a broad area of the subtropicalSouth Atlantic (Fig. 3). The Intertropical Convergence Zone (ITCZ)over the tropical Atlantic moves south into the warmer hemi-sphere. Precipitation increases over the tropical Atlantic and Brazilsouth of the equator and decreases over the tropical north Atlanticand northern Amazon (Fig. 4e, f). Precipitation also decreases overthe cold subpolar North Atlantic and over the tongue of cold,freshwater extending from Iberia into the Gulf of Mexico.

The southward ocean heat transport warms the SouthernOceans, reducing the sea ice fraction at 45–65�S which furtherenhances the warming in this region (Fig. 3). Warming over theAntarctic continent is modest and less than 1 �C. Warming alsooccurs over northern South America in association with thereduction in precipitation and cloudiness. Cooling of more than 2 �Cextends downstream from the North Atlantic into the mid-latitudesof Europe and Asia. Sea ice fraction increases over the North Pacificwith enhanced cooling associated with this sea ice. North Americashows no significant annual mean cooling, even a slight but notsignificant warming, during the hosing of the North Atlantic.

a

b

Fig. 2. Annual Atlantic meridional overturning circulation (Sv) for (a) LGM controlsimulation and (b) last 20 years of hosing in NATL_1.0 experiment. Positive indicatesa clockwise circulation.

Table 1Summary of experiments, ‘‘ESL’’ is the equivalent sea level change.

Experiment Region Amount (Sv) Years of hosing ESL equiv (m)

NATL_1.0 North Atlantic 1 100 9NATL_0.5 North Atlantic 0.5 100 4.5NATL_0.25 North Atlantic 0.25 100 2.25NATL_0.1 North Atlantic 0.1 100 0.9NATL_0.1ex North Atlantic 0.1 500 4.5GOM_0.5 Gulf of Mexico 0.5 100 4.5GOM_0.28 Gulf of Mexico 0.28a 100 2.8

a Because of an error in calculation, the experiment for the freshwater delivery tothe Gulf of Mexico with the smallest rate of hosing was 0.28 rather than 0.25 Sv.



The mid and high latitudes of the North Atlantic subsurface oceanwarm in response to the addition of freshwater to the surface of thesubpolar ocean (Fig. 4a, b). The fresh layer at the surface and the largeexpansion of insulating sea ice inhibit the transfer of heat from theocean to the atmosphere. Additionally, the AMOC diminishes ata lesser rate than the change in sea ice, still transporting heatnorthward into the North Atlantic during the hosing. This warmwater is subducted below the cold fresh surface water and sea ice.The subsurface warming, extending from 100 to 3000 m, acts to keepthe sea ice thickness in the GIN Seas only 1–2 m thick by enhancingbasal melting. In the Northern Hemisphere (NH) subtropics, the cold,freshwater at the surface is subducted into the southwestward flowof the subtropical gyre and appears as a cold, fresh anomalyextending to 600 m deep. In the SH tropics and subtropics, thesubsurface ocean warms due to reduced AMOC and reversed oceanheat transport (OHT) south of the equator and downward mixing ofheat from the thermocline (detailed analysis of ocean response willappear in Brady and Otto-Bliesner, in preparation).

3.2. Seasonal response to freshwater hosing

There is strong seasonality in the NH high-latitude temperatureand sea ice response in the NATL_1.0 experiment (Fig. 5). During NHwinter, persistent sea ice covers the sea surface from eastern US toIberia. December–January–February (DJF) surface air temperaturescool in excess of 15 �C in a broad region from the Labrador Sea to theBritish Isles. During NH summer, sea ice retreats dramatically, havingonly 50% coverage in the Labrador and Greenland–Iceland–Norwe-gian (GIN) Seas and along the northeast coast of Newfoundland. Thisexposes the freshwater cap to the summer insolation. Cooling is lessthan 4 �C over much of the northern North Atlantic, and June–July–August (JJA) cooling over Greenland is only 2 �C.

Significant seasonal differences of the surface temperatureresponses are also found over North and South America in theNATL_1.0 experiment. Over South America, the southward shift of

the ITCZ and reduced precipitation and cloud cover over northernSouth America enhances the warming in DJF. North America alsoshows a slight warming in DJF, extending from the southeast US toHudson Bay, though this warming is not significant in relation tothe interannual temperature variability of the LGM control at theselocations.

The low-latitude oceans do not show any significant seasonalityto their surface temperature responses. The cooling over the NHlow-latitude oceans is maintained by the southward advection ofcold fresh arctic water on the eastern side of the subtropical gyre,and over the SH low-latitude oceans by the reversal of the oceanheat transport at the equator, both of which have time scales longerthan the seasonal time scales.

There is significant seasonality in the SH high-latitude temper-ature response with greater warming during SH winter thansummer. JJA surface air temperatures warm in excess of 1 �C overthe Southern Ocean, associated with a reduction of sea ice, mostsignificantly at latitudes from 45 to 60�S (Fig. 5).

3.3. Transient behavior and recovery ofthe annual-averaged climate

The North Atlantic high-latitude climate responds swiftly to the1 Sv freshwater forcing in the NATL_1.0 experiment (Fig. 6). Sea iceforms rapidly and its areal extent increases by 50% reachinga maximum in the first decade of hosing. Greenland temperatures,as well as those over northwest Europe, also cool by 6 �C in the firstdecade. Greenland temperature shows a high degree of correlationwith the NH sea ice area, both in the mean response during thehosing and the shorter term variability, establishing the primaryrole of nearby sea ice in determining the response over Greenland,(Fig. 6b, c). The AMOC, by comparison, spins downs more slowly,with a decline to 80% of its pre-hosing strength after the firstdecade, 40% after 50 years, and then a more gradual but continueddecline over the next 50 years.

a c

b d

Fig. 3. Annual climate system response for LGM control and anomalies for last 20 years of hosing in NATL_1.0 experiment. (a, b) sea ice extent (fractional concentration) and (c, d)surface air temperature (�C).



In contrast to the rapid response during hosing, the high lati-tudes in the North Atlantic are slow to recover. Immediately afterthe freshwater forcing is stopped, the North Atlantic surface climatestarts to recover. Surface air temperatures over Greenland warm by5 �C in the first few years and sea ice retreats. This initial recovery isfueled by the resumption of convection in the ocean, withmaximum late winter-early spring mixed layer depths increasing to500–700 m in the Labrador and Irminger Seas. The warming atsubsurface depths is mixed up to the surface, warming sea surfacetemperature and reducing the sea ice extent. The warmer seasurface quickly exhausts its heat to the atmosphere as sensible andlatent heat. As well, precipitation increases, freshening the oceansurface. This recovery though is short-lived.

During the second decade after the freshwater forcing isstopped, NH sea ice area increases as the ocean cools and thesurface remains fresh supporting increased stability. The NorthAtlantic sea ice grows not only extensive in area as during thehosing, but also thicker with less interannual variability. Basalmelting of the sea ice, which continued throughout the hosing in

association with warmer subsurface temperatures, is reduced bya factor of two (not shown). GIN sea ice becomes much thicker (3–4 m). With a much reduced AMOC and northward ocean heattransport at the end of the hosing, the subsurface temperatures canonly very slowly warm. Fifty years into the recovery, the AMOCstrength reaches a minimum at about 20% of its pre-hosing LGMstrength and then recovers slowly, taking approximately 500 yearsto reach the pre-hosing glacial strength of w15 Sv. Greenlandtemperatures also recover slowly in concert with the recovery ofNH sea ice to its pre-hosing extent.

Fig. 8a shows the difficulty in using time series of Greenlandtemperature as a proxy for the AMOC temporal behavior. Duringthe first 5 years of the freshwater forcing, Greenland temperaturecools by w6 �C, most of its cooling in this experiment, while theAMOC decreases by less than 10%. Over the next 95 years offreshwater forcing, the AMOC continues to slow, while Greenlandtemperatures remain at about �40 �C, though with considerablevariability in concert with variability in the sea ice extent. The rapidbut aborted warming of Greenland temperatures immediately after

a

b

c

d

e

f

Fig. 4. Annual climate system response in Atlantic region for LGM control (top row) and anomalies for last 20 years of hosing in NATL_1.0 experiment (bottom row). (a, b) subsurfaceocean temperature (�C) averaged over depths of 300–1000 m; (c, d) surface salinity (psu); and (e, f) precipitation (mm/day).



the freshwater forcing is stopped is accompanied by only a smalltemporary increase in the AMOC. Over the next 50 years of therecovery, the AMOC returns to its weak state while Greenlandtemperatures remain between �40 and �42 �C. During theremaining 450 years of recovery, the AMOC and Greenlandtemperatures change both more synchronously and proportionally.

The climate in the tropics responds more gradually to thefreshwater input (Fig. 6). Cooling of Cariaco surface temperaturelags the AMOC response during the first decade of the hosing, thetime scale for the freshwater in the high latitudes to be advected tosubtropical and tropical latitudes by the gyre circulation. Temper-ature then cools by w3 �C during the next 40 years of the hosing,before stabilizing at a temperature of w19 �C for the remainder ofthe freshwater forcing period. With the shift in the Atlantic ITCZ,rainfall in northern South America decreases by w30% by the end ofthe hosing. The correspondence of the Amazon rainfall decreases

with the Cariaco cooling and South Atlantic warming (discussedlater) rather than the Greenland cooling suggests a low rather thanhigh latitude control on the shift in the ITCZ during the freshwaterperturbation. The shift in the ITCZ also affects the wind stresses inthe Caribbean, with upwelling in the Cariaco region increasing by30–50% during the freshwater forcing. This upwelling of coolsubsurface waters enhances the cooling of the Cariaco surfacetemperatures.

In contrast to Greenland temperature, Cariaco temperaturerecovers relatively quickly. Cariaco Basin, in the tropical NorthAtlantic, recovers two-thirds of its cooling in the first two decadesafter the end of the anomalous freshwater forcing in the subpolarNorth Atlantic. Similarly, northern Amazonia precipitation recoversrapidly, returning to almost normal rainfall in the first two decades.The rapid recovery is tied to the low-latitude temperatures ratherthan high-latitude temperature and sea ice conditions. Once the

a b

c d

Fig. 5. Seasonal anomalies for last 20 years of NATL_1.0 experiment. (a) DJF surface air temperature (�C), (b) JJA surface air temperature (�C), (c) DJF sea ice extent (fractionalconcentration), and (d) JJA sea ice extent (fractional concentration).



hosing is stopped, the tongue of cold, relatively fresh wateradvected by the subtropical gyre in the eastern basin disappears,and the subtropical surface ocean, with less cloudy skies andtherefore abundant solar absorption quickly warms. North Atlanticand Greenland temperatures, on the other hand, remain cold, and

sea ice extends to 40�N. Cariaco upwelling remains enhanced fora few decades after the cessation of the hosing before declininggradually to pre-hosing levels (Fig. 6).

Fig. 8b shows that the time evolution of Cariaco temperature(precipitation from northern South America shows similar

Fig. 6. Transient behavior of NATL_1.0 experiment before (negative years), during (years 0–100) and after (years 100–600) freshwater forcing for annual (top to bottom) AMOC,Greenland surface air temperature, Northern Hemisphere sea ice area, Cariaco surface air temperature, Cariaco upwelling at 50 m depth, and North Amazon precipitation. Regionsdefined in plot labels. Three-year running mean applied to annual averages.



behavior) is also not perfect as a measure of the AMOC temporalbehavior. Except for the short lag at the beginning of the freshwaterforcing, there is a linear relationship between the simulated Cariacotemperature and AMOC during hosing, but this relationship is notmaintained during recovery. In the first decade following the

cessation of the hosing, there is a rapid warming from w19 �C to20.4 �C about half of the recovery in temperature that is notaccompanied by a significant change in the strength of the AMOC(Fig. 6). For the next 50 years, the AMOC remains weak, while theCariaco temperature continues to increase by about 0.6 �C. For the

Fig. 7. Transient behavior of NATL_1.0 experiment before (negative years), during (years 0–100) and after (years 100–600) freshwater forcing for annual (top to bottom) AMOC,Hulu–Donge Caves surface air temperature, Hulu–Donge Caves precipitation, South Atlantic surface air temperature, Dronning Maud Land surface air temperature, and EPICA DomeC surface air temperature. Regions defined in plot labels. Three-year running mean applied to annual averages, except for the EPICA ice cores which have a six-year running meanapplied.



rest of the recovery, Cariaco temperature and the AMOC again arelinearly related, but the amount of Cariaco cooling for an AMOCdecrease during the freshwater forcing is three times greater thanthe amount of warming for an AMOC increase during this part ofthe recovery.

The South Atlantic warms by w1 �C during the hosing with thewarming starting only in the second decade of the freshwaterperturbation (Fig. 7). The lag of the response of the South Atlantic issuch that the South Atlantic temperatures do not start warminguntil the AMOC has already been reduced to 11 Sv, approximately70% of its pre-hosing strength (Fig. 8c). The temperatures warmrapidly, though only by about 1 �C as the AMOC continues todecrease. In contrast, the South Atlantic sea surface temperaturescool slowly with the same time scale as Greenland warms aftercessation of the freshwater perturbation.

Over the East Asian monsoon region, surface temperature ratherthan precipitation shows a response to the freshwater forcing(Fig. 7b, c). During the freshwater perturbation, surface airtemperatures cool gradually by w2.5 �C, in better correspondencewith the gradual slowing of the AMOC then the more rapid coolingin Greenland. After the cessation of the hosing, Hulu–Dongetemperatures warm gradually reaching the pre-hosing climatestate after w400 years. In contrast, annual precipitation in theHulu–Donge region shows considerable interannual variability but

no discernable signal in response to the freshwater forcing orcorrelation with the Atlantic changes.

Antarctica simulated surface air temperatures at EPICA DOME Cand Donning Maud Land warm throughout the freshwaterperturbation and for about 150–200 years after its cessation toabout 1 �C warmer than the pre-hosing temperatures but withsignificant interannual and decadal variability (Figs. 7 and 8d). Even500 years after the freshwater forcing has stopped, the EPICA DomeC temperature has not completely recovered to its initial state. Thismuch longer time scale for Antarctica agrees qualitatively with the‘‘thermal bipolar seesaw’’ conceptual model of Stocker and Johnsen(2003) which shows that although signals in the South Atlanticshould be in antiphase with the North Atlantic, the heat reservoir ofthe Southern ocean should lengthen its response.

4. Sensitivity to rate, location, and duration of freshwaterforcing

Fresh water added to the Labrador and GIN Seas has a direct andimmediate effect on the LGM convection south of Greenland, whilefreshwater added to the Gulf of Mexico has a more indirect effectwith the freshwater first needing to be transported by thesubtropical and subpolar gyres to the LGM convection sites. TheAMOC reduction is less for the smaller freshwater forcings, but is not

a b

c d

Fig. 8. Scatter plots for NATL_1.0 experiment of AMOC (Sv) versus surface air temperatures (�C) for (a) Greenland, (b) Cariaco, (c) South Atlantic, and (d) EPICA Dome C. Regionsdefined as in Fig. 6. Orange dots are for last 10 years of LGM control before freshwater forcing, blue dots for each year during the freshwater forcing, and green squares for the 500years of recovery after the cessation of the freshwater forcing.



linearly proportional with the size of the freshwater perturbation(Fig. 9a, Table 2). The NATL_1.0 and NATL_0.5 simulations both givelarge reductions in the AMOC. The AMOC is reduced less for a givensized forcing into the Gulf of Mexico compared to when the fresh-water is added directly into the subpolar North Atlantic. Not all thefresh water added to the GOM makes it into the high latitudes ofthe North Atlantic and thus active in the freshwater forcing of theAMOC. The percentage reduction in the AMOC for GOM_0.5 is

similar to the NATL_0.25 suggesting that GOM freshwater forcing isonly about 50% as effective as NATL freshwater forcing. Extendingthe freshwater hosing to 500 years in the NATL_0.1ex experimentresults in an additional reduction of the AMOC as compared to theNATL_0.1 experiment but a more vigorous AMOC than the NATLexperiments with greater rates of hosing.

All the freshwater forcing experiments show a reduction of thenorthward transport of heat in the Atlantic by the ocean as

a

b

Fig. 9. Comparison of experiments for varying rates and locations of freshwater forcing as defined in Table 1. (a) Transient behavior of AMOC before, during, and after freshwaterforcing, and (b) Atlantic northward ocean heat transport (PW) averaged over the last 20 years of hosing in each experiment.

Table 2Comparison of annual changes simulated by each of the experiments for several sites.

1 Sv, N.Atl 0.5 Sv N.Atl 0.25 Sv, N.Atl 0.1 Sv, N.Atl 0.1 Sv, N.Atl extended 0.5 Sv, GOM 0.28 Sv, GOM

Atlantic MOC, H1 wshutdowna �81% �70% �50% �29% �40% �51% �35%Greenland T, H1 �6 �Cb �5.6 �C �7.6 �C �7.6 �C �6.2 �C �5.8 �C �6.3 �C �5.5 �CIberian Margin – HE, SST �8 to �4 �C,

SSS �3 to �1 psuc�8.3 �C,�8.4 psu

�7.1 �C,�6.5 psu

�4.5 �C,�4.6 psu

�2.6 �C,�2.4 psu

�3.1 �C,�3.2 psu

�3.1 �C,�3.8 psu

�1.9 �C,�2.0 psu

Gulf of Mexico, SSS �4 to �2 psud �5.3 psu �1.8 psu �0.3 psu 0.3 psu 0.4 psu �11.0 psu �5.2 psuCariaco T, H1 �0.5 �Ce, YD �4 to �3 �C �2.7 �C �1.6 �C �0.6 �C �0.2 �C �0.2 �C �0.6 �C �0.2 �CN. Amazon Precipitation �32% �11% 4% 4% 1% �2% 2%Nordic Seas, HE Intermediate T þ2 to 4 �Cf 4.2 �C 3.9 �C 3.4 �C 2.0 �C 3.8 �C 0.9 �C 0.7 �CS. Atlantic, HE Mid-depth T þ1 to 3 �Cg 4.0 �C 2.8 �C 1.5 �C 0.5 �C 2.0 �C 2.1 �C 1.2 �CEPICA Dome C, AIM 2 to 8þ 0.5 to 3 �Ch 0.7 �C 0.3 �C 0.3 �C 0.3 �C 1.2 �C 0.6 �C 0 �C

The model values are annual anomalies for the last 20 years of the freshwater hosing.Model results are: % reduction of Atlantic MOC (overturning streamfunction, 34�S, 814 m), Greenland (surface air temperature, 70–80�N, 50–20�W), Iberian margin (SST andSSS, 30–40�N, 20–10�W), Gulf of Mexico (SSS, 20–30�N, 90–80�W), Cariaco (surface air temperature, 5–15�N, 70–50�W), Northern Amazon (% change of precipitation, 10�S–5�N, 70–50�W), Nordic Seas (ocean temperature, 60–70�N, 10–0�W, 730 m), South Atlantic (ocean temperature, 15–5�S, 5–15�E, 470 m), EPICA Dome C (surface airtemperature, 70–80�S, 110–140�E). Quantitative proxy interpretations that exist are tabulated in first column of table. For the proxy values, H1 denotes the Heinrich 1 event,HE denotes several Heinrich events during the glacial period, YD denotes the Younger Dryas, and AIM denotes Antarctic Isotope Maximum.Proxy data are from following.

a McManus et al., 2004.b Cuffey and Clow, 1997.c Bard et al., 2000.d Flower et al., 2004.e Lea et al., 2003.f Rasmussen and Thomsen, 2004.g Ruhlemann et al., 2004.h Stenni et al., 2003; EPICA Community Members, 2006.



compared to the LGM control (Fig. 9b). Poleward of 35�N latitude,all the freshwater forcing experiments show a much weakened(less than 0.25 PW) and very similar northward OHT. Equatorwardof 35�N, the reduction of Atlantic OHT is more sensitive to theamount of freshwater forcing. A reversal to southward transportnear the equator only occurs in the NATL_1.0 and NATL_0.5experiments. The ocean heat transport in the Atlantic is very similarfor the GOM_0.5/NATL_0.25 experiments and for the GOM_0.28/NATL_0.1 south of w20�N.

The subpolar glacial North Atlantic responds dramatically toadditional freshwater input of all magnitudes, locations, anddurations due to the large feedbacks among the ocean, atmosphereand sea ice. Owing to the very cold high latitude oceans with nearfreezing temperatures at LGM, extensive sea ice forms in the NorthAtlantic in all the experiments. Surface air temperatures are 10–15 �C colder from 50 to 70�N (Fig. 10). The cooling of Greenlandtemperatures is relatively insensitive to the amount or location ofthe hosing, ranging from 5.5 to 7.6 �C cooling, and is not propor-tional to the percentage reduction of the AMOC (Table 2). As theamount of freshwater added to the subpolar North Atlanticdecreases, the Greenland cooling and NH sea ice expansionbecomes less abrupt, though even in the NATL_0.25 simulationmost of the cooling and sea ice growth occurs within 15–20 years ofthe start of the freshwater perturbation (not shown). In theNATL_0.1 simulation, the Greenland cooling and NH sea iceexpansion occurs more gradually over the entire 100 years.Greenland temperatures show similar cooling for the GOM exper-iments but with a 20 year lag to the start of the hosing for thefreshwater to be advected into the North Atlantic convectionregions and sea ice and AMOC to respond (not shown).

The response of sea surface temperatures (SST) and salinities(SSS) in the Iberian margin are more sensitive to the rate andlocation of the freshwater forcing (Figs. 10 and 11; Table 2). Locatedjust south of the subpolar North Atlantic hosing area, the surfaceocean in the mid-latitudes of the northeast Atlantic is stronglyaffected by the tongue of cold, freshwater advected out of thesubpolar region, with increasing reductions in salinity andtemperature as the rate of freshwater added in the NATL simula-tions increases. The total volume of water added has less influenceon the SST and SSS responses at the Iberian margin, with theNATL_0.1 and NATL_0.1ex experiments having more similarreductions in temperature and salinity than the NATL_0.5 andNATL_0.1ex experiments. For the GOM simulations, the cooling andfreshening at the Iberian margin is less than the comparable NATLsimulations and lags the forcing by w25–30 years.

As the rate of freshwater addition to the subpolar North Atlanticdecreases, so does the decrease of temperature and salinity in thetongue of water advected southwestward from the Iberian margininto the Caribbean (Figs. 10 and 11, Table 2). The cooling of surfaceair temperature at Cariaco is less than 1 �C and the precipitation inthe northern Amazon is reduced by less than 10% for all casesexcept the cases with the larger freshwater forcings of 0.5 and 1 Svadded into the subpolar North Atlantic. The CCSM3 freshwaterforcing experiments suggest a threshold rate of freshwater forcingfor a detectable change above the natural variability at subtropicaland tropical sites. Salinity decreases in the Gulf of Mexico duringhosing are greatest for the GOM cases, as expected, but also showa significant salinity change of more than �5 psu in the NATL_1.0case (Fig. 11).

Surface warming found in the tropical South Atlantic is greaterfor larger freshwater forcings, but varies within a small range(Fig. 10). In the Nordic Seas, the subsurface ocean at mid-depthswarms in response to the addition of freshwater to the NorthAtlantic. Even when the freshwater is added to the Gulf of Mexico, ittakes only a decade to be transported to the subpolar North Atlantic

Fig. 10. Comparison of annual surface air temperature anomalies (�C) averaged overthe last 20 years of freshwater forcing for experiments with varying rates, locations,and duration of freshwater forcing.



and weaken deep convection. The subsurface warming is greatestfor the larger NATL forcings and weakest for the GOM forcings(Table 2). The total volume of water, rather than the rate of fresh-water forcing, is important for the subsurface responses, with theNATL_0.5 and NATL_0.1ex experiments showing comparablewarmings in the Nordic Seas and South Atlantic (Table 2). Alsointeresting and relevant to proxy records from the Gulf of Mexico isthat at mid-depths of w450 m, the NATL_1.0 experiment hascooling of 1–2 �C, while the NATL_0.25 and NATL_0.1 experimentshave warming of 1–2 �C (not shown).

Warming in the Southern Ocean and Antarctica shows thestrongest relationship with the duration of the forcing rather thaneither the rate or total volume of water added to the North Atlantic(Fig. 10). The NATL_0.1ex experiment with the duration of thefreshwater forcing increased from 100 to 500 years shows warmingof surface air temperatures over wide areas of the Southern Ocean.EPICA Dome C surface air temperatures have warmed by 1.2 �C bythe end of the hosing. The CCSM3 experiments only explore a smallsubset of possible freshwater forcing rates and durations thoughqualitatively agree with the EDML record which shows a linearrelationship between the amplitudes of Antarctic warmings andthe duration of the accompanying stadial in Greenland during MIS3 (EPICA Community Members, 2006).

Only the NATL_1.0 simulation has been run long enough to allowthe AMOC to recover to the pre-hosing state of w15 Sv. The recoverytime is w500 years. The other simulations only include the first 100years of recovery (Fig. 9a). The slope of the recovery starting twodecades after the end of the freshwater forcing is similar in all thesimulations. Extrapolating these curves, suggests that the AMOC inthe NATL_0.5 and GOM_0.5 simulations would recover in w350years, in the NATL_0.25 and GOM_0.28 simulations in w250 years,and the NATL_0.1 simulation in w200 years. The weaker cases seemto have shorter recovery times, because there is less to recover. Theyalso do not exhibit the aborted recovery of the NATL_1.0 experiment,but instead maintain a more continuous recovery.

5. Comparison to published proxy records

Although the freshwater simulations described in this paper arehighly idealized, it is instructive to compare the responses broadlyto published proxy reconstructions for Heinrich events and theYounger Dryas (Table 2). At the end of the freshwater forcing, allexperiments show reduction of the AMOC. The impact of thisslowdown of the AMOC on temperature and precipitation at proxylocations along a north–south transect in the Atlantic varies amonglocations and is not always linearly related to either the amount ofthe freshwater perturbation or the reduction of the AMOC.

Measurements of 231Pa/230Th in core GGC5 from the BermudaRise indicate a slowdown of the AMOC starting at w19 ka, witha nearly complete shutdown of the AMOC from 17.5 ka until 15 ka(McManus et al., 2004). Using an arbitrary cutoff of 50% reductionin the AMOC, all experiments except the NATL_0.1 and GOM_0.28give substantial slowdown of the AMOC by the end of 100 years offreshwater forcing. The weakest NATL experiment, NATL_0.1, whenextended for another 500 years of forcing, shows a further reduc-tion of the AMOC, �40% at year 500 as compared to �29% at year100. It is noteworthy that when sufficient freshwater is added to theGulf of Mexico, 0.5 Sv in our 100-year hosing experiments, theAMOC reduction is greater than 50%.

The simulated annual cooling of Greenland temperatures rangesfrom 5.5 to 7.6 �C. This cooling is comparable to the GISP2 d18O withborehole temperature calibration, which indicated that Greenlandannual mean temperatures cooled up to w4–6 �C during H1 and H2(Cuffey and Clow, 1997). The CCSM3 response also corroborates thestrong seasonality of the NH temperature response at high

Fig. 11. Comparison of annual surface salinity anomalies (psu) averaged over the last20 years of freshwater forcing for the experiments with varying rates, locations, andduration of freshwater forcing.



latitudes, as indicated by proxies (Denton et al., 2005). The Iberianmargin core indicates that Heinrich events 1–3 were associatedwith rapid surface coolings of 4–8 �C and reduced surface salinitiesof 1–3 psu (Bard et al., 2000). All the experiments give cooling atthe Iberian margin location, with the NATL experiments with ratesat least 0.25 Sv in better agreement with the proxy interpretation.The experiments generally overestimate the freshening of surfacesalinities as compared to the proxy estimates, except for theNATL_0.1 and GOM_0.28 experiments.

The simulated annual warming of Antarctic temperatures at theend of the freshwater forcing varies from no warming to þ1.2 �C.The experiments with warming greater than 0.5 �C compare wellwith the EPICA Dome C (EDC) and Dronning Maud Land (EDML) icecores, though on the low side of the warming indicated in thesecores (Stenni et al., 2003; EPICA Community Members, 2006). Thisis to be expected for two reasons. First, the simulated warmingcontinues after the cessation of the freshwater forcing. In theNATL_1.0 experiment, warming continues for an additional 150–200 years to about 1 �C warmer than the pre-hosing temperatures.Second, the freshwater forcing is applied for only 100–500 years,relatively short as compared to the duration of NGRIP stadials(EPICA Community Members, 2006). The NATL_0.1ex experiment,with the freshwater forcing applied for 500 years, gives the greatestwarming over Antarctica.

Surface salinities in the Gulf of Mexico show freshening in theexperiments where the water is added directly to this basin,although both GOM experiments suggest greater freshening thansuggested by the data. The strong salinity reductions in the GOMexperiments might result from the design of these experiments:the water is added only to the surface and for only 100 years. TheNATL_1.0 and NATL_0.5 also show substantial reduction of surfacesalinities in the Gulf of Mexico. These reductions are associatedwith the freshwater that is added between 50 and 70�N beingtransported south and west in the subtropical gyre. The NATL_0.5experiment shows the best correspondence with the salinityreductions suggested for H1 by Flower et al. (2004).

At low latitudes, in the tropical Atlantic, the record from theCariaco Basin on the northern Venezuelan shelf indicates decreasedriverine discharge from northern South America associated withthe colder North Atlantic during the H1 event and has been inter-preted as indicating a southward shift of the ITCZ and reducedprecipitation over northern South America (Peterson et al., 2000).Mg/Ca records suggest slightly cooler (0.5 �C) SSTs during the latterpart of H1 (Lea et al., 2003). Farther south, sediment cores in theBolivian Altiplano indicate wetter conditions and higher lake levelsduring H1 and H2, as well as the Younger Dryas (Baker et al., 2001),speleothem and travertine records for tropical northeastern Brazilfind wet periods that are synchronous with Heinrich events (Wanget al., 2004), and a sediment core from Brazil’s margin recordsincreased continental debris coincident with Heinrich events (Arzet al., 1998). The CCSM3 experiments show a qualitatively similarresponse compared to the proxy climate records, consistent witha shift southward of the ITCZ and a cooling at Cariaco. The responseis much weaker for the smaller forcing cases. The simulated coolingat Cariaco and decrease in precipitation in the northern Amazon arenot sensitive to the duration of the forcing period. None of theexperiments show cooling of the magnitude found for the YD event,but the NATL_0.25 agrees best to the H1 cooling found by proxies.

Eight cores from the North Atlantic and Nordic Seas taken atintermediate depths ranging from w850 to 1800 m indicatewarming of 2–4 �C during Heinrich events (Rasmussen andThomsen, 2004). The warming in the subsurface of the high-lati-tude North Atlantic Ocean is interpreted as the continuation of theNorth Atlantic drift during Heinrich events. The warm saline waterstill flows into the Nordic Seas but below the cold fresh layer

associated with the melting icebergs. All the NATL experimentsproduce subsurface warming in the Nordic Seas comparable to theproxy records. The total volume of water is important for producingthis warming with the NATL_0.5 experiment and NATL_0.1exexperiment showing similar warmings. The two GOM experimentsproduce much less warming in the Nordic Seas than suggested bythe proxy records.

Two sediment cores in the tropical Atlantic analyzed for H1 andthe Younger Dryas also indicate subsurface warming at mid-depthsof 1–3 �C (Ruhlemann et al., 2004). In the South Atlantic, the modelsimulations with 0.25 and 0.5 Sv of freshwater input, regardless oflocation of the forcing region, show subsurface warming in goodagreement to the proxy records (Table 2). The tropical subsurfacewarming is found to be sensitive to the duration of forcing with thewarming for NATL_0.1ex (duration of 500 years) greater than foundin NATL_0.1 (duration of 100 years).

6. Discussion and conclusions

In our glacial freshening experiments, the less dense fresh waterprovides a lid on the ocean water below, suppressing oceanconvection and interaction with the atmosphere above, resulting ina cooling of the surface ocean and atmosphere in the North Atlantic.This is the case whether the freshwater is added directly to the areaof convection south of Greenland or transported there by thesubtropical and subpolar gyres when added to the Gulf of Mexico,though in the latter case the freshwater becomes saltier as it istransported poleward and thus less active in the forcing of theAMOC. The cooling in the subpolar North Atlantic is also delayed bya decade after the start of the perturbation when the freshwater isput into the Gulf of Mexico. This delay time is the advective timescale for the fresh anomaly to be advected northward into thesubpolar oceans. This delay would most likely not be seen in theproxy records given their dating uncertainties and resolution. Theenhanced stability and near freezing temperatures of the subpolarsurface water allows sea ice to rapidly form in the North Atlantic,increasing to its maximum extent in the first decade in the NATLcases with large freshwater additions. The high albedo of the sea iceacts as a positive feedback, amplifying the cooling of the atmosphere.

Fig. 12. Difference of the annual implied northward heat transport by the global oceanand global atmosphere for the NATL_1.0 experiment as compared to the LGM control.



The AMOC reduction is less for the smaller freshwater forcings,but is not linear with the size of the freshwater perturbation,a result also found by Fluckiger et al. (2008) in the ECBILT-CLIOmodel. The NATL_1.0 and NATL_0.5 simulations both give largereductions in the AMOC, suggesting a possible effective thresholdin the maximum rate of freshwater delivery. The AMOC is stillreduced when freshwater is added into the Gulf of Mexico, thoughless for a given sized forcing into the Gulf of Mexico compared towhen the freshwater is added directly into the subpolar NorthAtlantic. The GOM results contrast with the hypothesized melt-water routing model of Clark et al. (2001) that suggested that whenmeltwater flowed to the Mississippi River the AMOC wouldincrease. Our CCSM3 results concur with other modeling results.Manabe and Stouffer (1997) found that the AMOC and climateevolved in a similar manner whether the freshwater was added tothe northern or subtropical North Atlantic, though the AMOCresponse was 4–5 times smaller in the subtropical experiment.Even when freshwater is added as a salinity anomaly in the deepwaters of the Gulf of Mexico to mimic sediment-laden meltwater,as has been done with the LOVECLIM model, the meltwater reachesthe convection sites and reduces the AMOC (Roche et al., 2007).

Greenland temperatures cool by 6–8 �C in all the experiments inthis study, with little sensitivity to the magnitude, location orduration of the freshwater forcing. The very similar coolings occurbecause extensive sea ice forms in the North Atlantic nearGreenland in all the experiments. Equatorward of the expanded seaice extent at the location of the Iberian Margin cores, the CCSM3experiments are sensitive to the magnitude and location of thefreshwater delivery to the North Atlantic. The simulated tempera-ture and salinity anomalies are affected by the colder and fresherwaters that are advected out of the subpolar North Atlantic. Ourresults agree with proxy data that indicate that the Iberian marginis strongly affected both by cooling and advection of low salinityArctic water masses during the last 3 Heinrich events (Bard et al.,2000).

Antarctic temperatures inferred from oxygen isotope ratiosmeasured in the high resolution ice core at Dronning Maud Landindicate an out-of-phase relationship between Greenland andAntarctic with broad warmings (Antarctic Isotope Maximums[AIMS]) associated with each of the Heinrich events over the last50 kyrs (EPICA Community Members, 2006). Simulated Antarcticand Southern Ocean temperatures confirm the bipolar responseindicated by ice core records and thermal bipolar seesaw model(Stocker and Johnsen, 2003). Comparison to interglacial hosingsimulations shows the circumpolar warming to be larger during ourglacial simulations, related to the larger simulated sea ice thicknessaround Antarctica at LGM (Morrill, personal communication). Oursimulations also suggest that the response is governed by theduration of the freshwater forcing, in qualitative agreement withthe EDML record which shows a linear relationship between theamplitudes of Antarctic warmings and the duration of the accom-panying stadial in Greenland during MIS 3 (EPICA CommunityMembers, 2006).

There is a large seasonality to the simulated surface temperatureresponse in the subpolar North Atlantic, with very large cooling inwinter and much less cooling in summer, also found by Fluckigeret al. (2008) with the ECBILT-CLIO model. The seasonal presence ofsea ice in the simulations is important for explaining the largeseasonal cycle of cooling. This concurs with the results of Dentonet al. (2005), who conclude that more marked changes of wintertemperatures and increased seasonality are implied by mismatchesbetween the ice core annual mean temperatures and moraine andsnowline changes in East Greenland, particularly during theYounger Dryas. They propose that large winter changes dominatedthe signal.

This interpretation of greater seasonality in the North Atlanticregion during Heinrich events is also confirmed by beetle remainswhich indicate a more continental climate in Britain than today (oreven the LGM), with coldest month temperatures w30–35 �Ccolder than warmest month temperatures (Atkinson et al., 1987).For glacial conditions, the summer–winter seasonal contrast insurface temperatures for the British Isles increases from w11 �C inthe LGM control simulations to w21 �C in the NATL_1.0 experiment.When the NATL_1.0 simulation is repeated for interglacial condi-tions (Morrill, personal communication), the summer–winterseasonal contrast in surface temperatures for the British Isles onlyincreases by a few degrees. Our CCSM3 simulations suggest that theextreme seasonality indicated by the proxy records can beexplained only when glacial boundary conditions and forcings arethe basis for the freshwater hosing simulations. This argues for thehigh-latitude North Atlantic playing a key role in abrupt climatechange during the last glacial cycle, with freshwater delivery to thesubpolar regions acting as the trigger.

The AMOC (34�S, 814 m) decreases at a rate slower than the seaice increase, continues to transport some heat poleward throughoutthe hosing, and never completely collapses. The freshwater forcing,rather than ‘‘jamming’’ the AMOC to the ‘‘off’’ position, graduallyweakens it with the AMOC reduced to less than 3 Sv in the NATL_1.0experiment. Globally, the northward transport of heat by the oceandecreases at all latitudes. The atmosphere partially compensates bytransporting more heat northward between 20�S and 35�N, butpoleward of 35�N the atmospheric transport of heat northward alsodecreases (Fig. 12). In the NATL_1.0 experiment, the Nordic Seaswarm at depths of w100–3000 m, with subduction of warm watersunder the cap of freshwater, similar to other modeling studies(Mignot et al., 2007) and in good comparison to proxy evidence. Thesubsurface warming melts the sea ice from below and keeps itrelatively thin at 1–2 m. The warming below the halocline issensitive to the magnitude, location and duration of the freshwaterforcing, being reduced as the freshwater forcing is reduced in theNATL experiments and smaller in the equivalent GOM experiments.For the NATL experiments, the total amount of freshwater ratherthan the rate of forcing determines this subsurface response.

It has been suggested that high latitude temperatures and icedrive the millennial variations of tropical and monsoon paleo-climates. Chiang and Bitz (2005) found that imposing anomalousland and sea ice in the Northern Hemisphere had similar effects onthe tropical Atlantic ITCZ, shifting it south over the oceans. In ourexperiments, the high northern latitude responses in surfacetemperature and sea ice extent are similar, regardless of theamount or location of the freshwater delivery. Yet we find muchdifferent responses in precipitation in the northern Amazon amongthe experiments, with decreases in precipitation more correlatedwith the changes in low-latitude temperatures. Recovery of theITCZ to its pre-hosing position also occurs more rapidly thanrecovery at high northern latitudes and is associated witha recovery of low-latitude temperature gradients. In the NATL_1.0experiment, following the end of the hosing, the ITCZ shiftsnorthward to its original position within decades and northernAmazonia precipitation increases, in sharp contrast to the slowrecovery of the northern high latitudes. Outside the Atlantic region,the Hulu Cave d18O speleothem record has been interpreted asa link between millennial-scale variations in the North Atlantic andthe Asian monsoon (Wang et al., 2001). Our simulations showsurface temperatures rather than precipitation in the Hulu-Dongeregion responding to the North Atlantic freshwater forcing.

The AMOC (34�S, 814 m) and Greenland surface air tempera-tures recover gradually, with recovery to pre-hosing states in200–500 years. In all cases, the AMOC recovers indicating thatCCSM3 may not have a stable ‘off’ state with no northern sinking or



deep water formation. The experiments here may be too short tohave forced a complete shutdown of the AMOC, though even inCCSM3 simulations with freshwater forcing for several thousandyears the AMOC does not completely shut down (Liu et al., 2009).Our sensitivity experiments suggest that the recovery time froma freshwater perturbation is sensitive to the magnitude of thefreshwater forcing; there is faster recovery with the lesser amountsbecause there is less water from which to recover. Knowing thehistory of the meltwater forcing is therefore critical to evaluatingthe ability of climate models to reproduce past events. The amount,duration and timing of freshwater forcing associated with Heinrichevents is only broadly known, and although IRD records confirmthe location as a belt extending from 40 to 55�N, the freshwaterforcing was likely not uniform in space or time. Indeed, our resultswith NATL_1.0 suggest that the freshwater forcing, e.g. the icebergsin Heinrich events, need not last for the entire period of reducedAMOC. In addition, hosing simulations simplify the process byadding freshwater rather than icebergs to the ocean. Models whichinclude iceberg trajectories, sedimentation rates and meltwaterinput are being developed and suggest a more complex meltwaterpattern than has been assumed in our experiments (De’Ath et al.,2006; Levine and Bigg, 2008).

Most numerical model investigations of the role of freshwateron the climate system have adopted modern conditions. The resultsfrom Renold et al., (2010) plus some unpublished results fromMorrill (personal communication) for interglacial boundaryconditions and forcings find that the AMOC has a two-stagerecovery with the second phase occurring in 60–80 years. Renoldet al. tie this rapid recovery to a density threshold in the GIN Seas,while Morrill finds a strong correlation with both GIN and Labradorseas winter sea ice extent. An idealized hosing experiment with thesame model and same resolution as Renold et al. but inserting thewater for glacial conditions at 19 ka shows a recovery of the AMOCas in our results and not a two-step recovery (Liu et al., 2009). Inaddition, the Bering Strait is closed at the LGM. The Bering Straitacts as an exhaust valve for North Atlantic freshwater anomalieswith flow through the strait reversed during the freshwaterperturbation (de Boer and Nof, 2004). In CCSM hosing simulationswhen the Bering Strait is closed, the recovery takes a century longerthan for an open Bering Strait (Hu et al., 2008).

Studies with simplified models suggest that the very rapid DOwarmings prominent in Stage 3 required background conditionsintermediate between full glacial and modern (Schulz, 2002; Wangand Mysak, 2006). Van Meerbeeck et al. (2009) argue that the LGMclimate should not be used to simulate the MIS 3 DO events. InLOVECLIM simulations, they find that although the overall strengthof the AMOC does not differ substantially between LGM and MIS 3,the MIS 3 convection sites shifted northward into the Labrador Seaand convection was enhanced in the Nordic Seas. The Labrador Seaconvection is expected to be more sensitive to meltwater from theLaurentide and Greenland ice sheets. Liu et al. (2009) find ina transient simulation with CCSM3 for 21–14.5 ka with transientdeglacial boundary conditions and forcings that the rapid BAwarming over Greenland is caused by not only the recovery of theAMOC and convective instability in the Nordic Seas, but also largeCO2 rise from H1 to the BA.

It is still then an open question if the models used to projectfuture climate change are capable of simulating the ‘‘size, speed,and extent’’ of past changes (National Research Council, 2002).Comprehensive climate models, while having more realistic feed-backs than conceptual models and models of intermediatecomplexity, still need to parameterize many processes that mayaffect the feedbacks between the atmosphere, ocean, and sea ice.Only EMICs and fully coupled models with flux corrections havebeen shown to exhibit multiple equilibrium (Stouffer et al., 2006;

Mikolajewicz et al., 2007). Past studies have also indicated that thestability of the THC is dependent on the choice of ocean mixingparameterization and parameter settings (Knutti et al., 2000) andthat the vertical subgrid-scale diffusion can affect whether the THCcollapse is stable (Manabe and Stouffer, 1999). The important role ofsea ice and its feedbacks on the ocean and atmosphere in this andother studies suggest that how well models capture this sensitivityis important to getting the right temporal response. Proxy datahave uncertainties in dating and calibration, and in particular,measure isotopes and other paleo tracers to infer past changes inclimate. New developments in climate models and computingpower will allow us to simulate the transient changes in proxies formore direct comparison with the data. These transient simulationswill require new histories of past meltwater forcings.

Acknowledgements

The National Center for Atmospheric Research is sponsored bythe National Science Foundation. We thank Bruce Briegleb, NanRosenbloom, and Bob Tomas for the running and processing of thesimulations, and C. Morrill for sharing unpublished results. We alsothank the two anonymous reviewers for their helpful suggestionswhich contributed to improving the manuscript.

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