This article appeared in a journal published by Elsevier. The attached copy is furnished to the author for internal non-commercial research and

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2 Quaternary Science Reviews 29 (2010) Contents lists available at ScienceDirect Quaternary Science Reviews journal homepage: The sensitivity of the climate response to the magnitude and location of freshwater forcing: last glacial maximum experiments Bette L. Otto-Bliesner *, Esther C. Brady National Center for Atmospheric Research, 1850 Table Mesa Drive, Boulder, Colorado 80305, USA article info abstract Article history: Received 31 January 2009 Received in revised form 3 June 2009 Accepted 6 July 2009 Proxy records indicate that the locations and magnitudes of freshwater forcing to the Atlantic Ocean basin as iceberg discharges into the high-latitude North Atlantic, Laurentide meltwater input to the Gulf of Mexico, or meltwater diversion to the North Atlantic via the St. Lawrence River and other eastern outlets may have influenced the North Atlantic thermohaline circulation and global climate. We have performed 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 inserted either 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 water below, suppressing ocean convection and interaction with the atmosphere above and reducing the Atlantic Meridional Overturning Circulation (AMOC). This is the case whether the freshwater is added directly to the area of convection south of Greenland or transported there by the subtropical and subpolar gyres when added to the Gulf of Mexico. The AMOC reduction is less for the smaller freshwater forcings, but is not linear with the size of the freshwater perturbation. The recovery of the AMOC from a 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, the cold subpolar North Atlantic is primed to respond rapidly and dramatically to freshwater that is either directly dumped into this region or after being advected from the Gulf of Mexico. Greenland temperatures cool by 6 8 C in all the experiments, with little sensitivity to the magnitude, location or duration of the 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 show a bipolar response, with warming and reduced sea ice. This warming continues after the cessation of the freshwater 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 sensitive to the amount of colder and fresher waters that are advected out of the subpolar North Atlantic. In the tropical Atlantic, the recovery of the Intertropical Convergence Zone (ITCZ) from its more southerly position during the freshwater forcing is much more rapid than the recovery of the AMOC, and is more related 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. Ó 2009 Elsevier Ltd. All rights reserved. 1. Introduction Proxy records indicate large freshwater forcings to the Atlantic Ocean basin during the last glacial interval. Ocean cores in the subpolar North Atlantic between 40 and 55 N latitude display distinct layers of lithic grains in the sand fraction during seven time intervals, so-called Heinrich events, during the time period * Corresponding author. Tel.: þ , fax: þ address: ottobli@ucar.edu (B.L. Otto-Bliesner). between 70 ka and 13 ka (Andrews, 1998; Broecker et al., 1992; Heinrich, 1988; Ruddiman, 1977). This ice-rafted debris (IRD) is composed of rock fragments with provenance likely from around the Hudson Strait (but also with sources from other parts of the Laurentide ice sheet and from the Fennoscandian ice sheet), that was incorporated into icebergs or sea ice, transported across the North Atlantic, and deposited to the sea floor as the ice melted (Bond et al., 1992; Hemming, 2004). Two Heinrich events occur near LGM: H2 with the magnetic susceptibility records suggesting depositional phases centered at 23.5 ka and 25 ka occurred during the maximum extent of the ice sheets during the last glacial cycle, /$ see front matter Ó 2009 Elsevier Ltd. All rights reserved. doi: /j.quascirev

3 B.L. Otto-Bliesner, E.C. Brady / Quaternary Science Reviews 29 (2010) and H1 with depositional phases centered at 16 ka and 17.5 ka during the early part of the deglaciation (Bard et al., 2000). Uplifted coral reefs in Papua New Guinea suggest sea level excursions of m corresponding to Heinrich events (Chappell, 2002; Yokoyama et al., 2001). Hemming (2004) estimated from Heinrich layer records that the flux of meltwater ranged from 1.6 Sv over one year to 0.3 Sv over 500 years, though with considerable uncertainty. Through modeling of oceanic d 18 O perturbations, Roche et al. (2004) obtain a duration of yr and an ice release of 2 1 m sea level equivalent for Heinrich event 4 occurring at about 40 ka. Abrupt meltwater pulses associated with ice sheet collapses, but without IRD layers, are also apparent in the geologic record of sea level rise during the last deglaciation. Marine sediments in a shallow gulf along the Australian margin indicate a rapid rise of sea level at 19 ka of m in 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 over several hundred years (Bard et al., 1990; Clark et al., 2002a). Whether this meltwater came from widespread melting of the Antarctic or Northern Hemisphere ice sheets, or both, is still much debated (Clark et al., 2002a; Peltier, 2005). The cause of the Younger Dryas cold event, starting at w12.9 ka, is also highly debated, with the preponderance of evidence suggesting the trigger to be freshwater from Lake Agassiz flowing the North Atlantic, though consensus 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 high latitudes in the North Atlantic, the ocean loses heat to the atmosphere, becoming denser and sinking, particularly in the Greenland Iceland Norwegian (GIN) and Labrador Seas, with flow southward at deeper levels. This ocean circulation redistributes heat gained in the South Atlantic and tropics to the North Atlantic region. It has been hypothesized that freshwater addition during the glacial period and deglaciation would have shutdown the North Atlantic part of the thermohaline circulation with influences on the global climate (Broecker et al., 1988; Broecker, 1994; Broecker, 2003; Clark et al., 2002b). A number of proxy records for the surface and deep ocean support this interpretation. Measurements of 231 Pa/ 230 Th in core GGC5 from the Bermuda Rise indicate a slowdown of the Atlantic Meridional 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 corresponds with the meltwater pulse at w19 ka and the shutdown is concurrent with the first of two depositional events identified with H1 (Bard et al., 2000). A proxy record of magnetic susceptibility from the Eirik Drift, interpreted as a measure of North Atlantic Deep Water (NADW) flow intensity, confirms the near shutdown of the AMOC for years during H1 (Stanford et al., 2006). High resolution benthic foraminifera d 13 C records from 2000-m depths in the North Atlantic also suggest perturbations in the AMOC during Heinrich events, with major changes in the deep-water masses, consistent with large reductions of deep-water formation and northward migration of southern ocean deep waters to 62 N (Elliott et al., 2002). Sediments in cores from the Iberian margin and western Mediterranean 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 reduced salinity during the past three Heinrich events (Bard et al., 2000), attributed to melting of icebergs advected to the margin during Heinrich events. The salinity decrease during H1 is estimated to be 1 2 psu based on the percentage of C 37:4 among C 37 alkenones and 2 3 psu based on d 18 O measurements in planktonic foraminifera from the same sediments. Surface salinity reductions of 2 3 psu for H1 have also been inferred from d 18 O for two cores farther north in the 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 (Flower et al., 2004). Greenland ice cores and records from the nearby North Atlantic also record millennial climate variability with rapid shifts during Dansgaard-Oeschger (DO) events from cold stadial to warm interstadial conditions. In Greenland, their amplitude reaches 8 to 16 C in a few decades to centuries (Severinghaus and Brook, 1999; Landais et al., 2004; Huber et al., 2006), with shifts in deuterium excess from one year to the next during the last deglaciation (Steffensen et al., 2008). These DO events have occurred 25 times over the last climatic cycle (North Greenland Ice Core Project Members, 2004) with a bipolar seesaw response first identified for the large events (Blunier et al., 1998), then for the first large events of the glacial inception (Landais et al., 2006), and more generally over the past 50 kyrs (EPICA Community Members, 2006) and for MIS5 (Capron et al, 2010). Antarctic ice cores find that they are not restricted to the last climatic cycle, with over 74 millennial changes in methane over the past 800,000 years (Loulergue et al., 2008). Modeling results support the interpretation that increased freshwater additions during the glacial period and deglaciation would have slowed or even shutdown the North Atlantic part of the overturning circulation. Conceptual and simplified numerical models suggest that the AMOC may have two stable states, transitioning from a strong AMOC to a weak or collapsed AMOC following a freshwater addition to the North Atlantic (Rahmstorf, 1995; Rahmstorf, 1996; Stocker and Wright, 1991). Fully coupled AOGCMs also indicate a slowing of the AMOC when freshwater is added to the North Atlantic (Stouffer et al., 2006). When the freshwater forcing is stopped, the AMOC recovers fully over years in all but one of the AOGCMs, the GFDL_R30 model which included flux adjustment to the ocean. These CMIP simulations were done with present-day initial conditions. Yet, the initial state and boundary conditions may have an important effect on the response and recovery to freshwater forcing. Ganopolski and Rahmstorf (2001) find that for the glacial climate, the Atlantic thermohaline circulation declines gradually when freshwater inflow to the Atlantic is increased, in contrast to reaching a clear bifurcation point where the circulation goes to an off mode for present climate in the CLIMBER model. The hysteresis behavior for the glacial simulation is also much less pronounced than in the simulation of present climate. Bitz et al. (2007) find that for glacial conditions the slow recovery from an instantaneous freshwater pulse is related to the more extensive sea ice and more stable density below the surface layer in the North Atlantic. In addition, with the Bering Strait closed at LGM, the export of North Atlantic meltwater takes longer, having only an outlet through the southern end of the North Atlantic, and thus bringing 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 into the ocean during Heinrich and other meltwater events. Here, we use simulations of the NCAR CCSM3 climate model with glacial boundary conditions to understand the sensitivity to freshwater perturbations. We examine the regional and transient responses to explore several questions: (1) What is the response and recovery of the AMOC to freshwater forcing? (2) How well do proxy records, in the North Atlantic and more remotely, correspond to the AMOC response to and recovery from freshwater forcing? and (3) How sensitive are the Atlantic and global responses to the rate, location, and duration of the freshwater perturbations?

4 58 B.L. Otto-Bliesner, E.C. Brady / Quaternary Science Reviews 29 (2010) CCSM model and experiments The National Center for Atmospheric Research (NCAR) CCSM3 is a global, coupled ocean atmosphere sea ice land climate model (Collins et al., 2006a). The model used in this study is the same model used at higher resolution for the IPCC AR4 projections of future climate. The atmospheric model is the NCAR Community Atmospheric Model version 3 (CAM3) and is a three-dimensional primitive equation model solved with the spectral method in the horizontal and with 26 hybrid coordinate levels in the vertical (Collins et al., 2006b). The ocean model is the NCAR implementation of the Parallel Ocean Program (POP), a three-dimensional primitive equation model in vertical z-coordinate (Gent et al., 2006). For this study, the atmospheric resolution is T42 (an equivalent grid spacing of approximately 2.8 in latitude and longitude), and ocean grid is points with poles located in Greenland and Antarctica, and 40 levels extending to 5.5-km depth. The ocean horizontal resolution corresponds to a nominal grid spacing of approximately 1 in latitude and longitude with greater resolution in the Tropics and high-latitude North Atlantic. The land model uses the same grid as the atmospheric model and includes a river routing scheme and specified but multiple land cover and plant functional types within a grid cell (Dickinson et al., 2006). The sea ice model is a dynamic thermodynamic formulation, which includes a subgrid-scale ice thickness distribution and elastic viscous plastic rheology (Briegleb et al., 2004). The sea ice model uses the same horizontal grid and land mask as the ocean model. The simulations have LGM forcings and boundary conditions, as described by the Paleoclimate Modeling Intercomparison Project Phase 2 (PMIP2). The PMIP2 LGM forcings relative to modern are the small change to insolation resulting from the slightly different Earth s orbit, which is set appropriate for 21 ka based on the calculations of Berger (Berger, 1978), and the reduced concentrations of atmospheric carbon dioxide (CO 2, 185 ppmv), methane (CH 4, 350 ppbv), and nitrous oxide (N 2 O, 200 ppbv), as adopted from the Greenland and Antarctic ice core records (Dallenbach et al., 2000; Fluckiger et al., 1999; Monnin et al., 2001). The PMIP2 boundary conditions for the LGM simulations are the ICE-5G ice sheet and topography (Peltier, 2004) with large continental ice sheets over North America and northern Eurasia (Fig. 1). They also include the specification of additional land due to the lowering of sea level by m with the large amounts of water frozen in the continental ice sheets. The lowering of sea level results in the closing of the Bering Strait and more land in the Arctic. Vegetation, dust aerosols, and river routings are unchanged from the preindustrial conditions. The LGM CCSM3 simulation has a global cooling of 4.5 C compared to preindustrial conditions with amplification of this cooling at high latitudes and over the continental ice sheets present at LGM (Otto-Bliesner et al., 2006). Sea ice expands, reaching 45 N in the western Atlantic in winter, but exhibits strong seasonality with the largely ice-free Nordic Seas during summer. CCSM3 simulates reduced deep convection in the Nordic Seas and enhanced convection south of Greenland (Fig. 1a), in agreement with proxies that suggest most GNAIW was likely formed south of Iceland (Duplessy et al., 1988; Pflaumann et al., 2003). The North Atlantic meridional overturning circulation weakens by 20% and shoals from a depth of 4 km at 40 N in the preindustrial CCSM3 simulation to a depth of 2.4 km in the LGM simulation, with Antarctic Bottom Water (AABW) filling the deep ocean in the glacial North Atlantic (Fig. 2a), supporting the interpretation from paleonutrient tracers of a shallower North Atlantic Deep Water (NADW) (Lynch-Stieglitz et al., 2007). The deep Atlantic Ocean is much colder and saltier at LGM in the model (Otto-Bliesner et al., 2007), as also found in core fluid measurements from ODP cores (Adkins et al., 2002). The North Atlantic subtropical gyre shifts southward and maintains a vigorous circulation under glacial forcings (Fig. 1b). The Antarctic Circumpolar Current (ACC) in CCSM3 increases by w60% at LGM as compared to preindustrial due to both an increase in the SH westerly wind stress over the Southern Ocean and an increase in AABW formation and sea ice around Antarctica. Differences in the strength of the SH westerlies is small among the LGM simulations with the PMIP2 models (Menviel et al., 2008) but these simulations show substantial differences in LGM SH sea ice and AABW formation (Otto-Bliesner et al., 2007). Further details of the LGM 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 shows the locations of freshwater (blue) added to the ocean as given in Table 1 and the LGM ice sheet extents (purple).

5 B.L. Otto-Bliesner, E.C. Brady / Quaternary Science Reviews 29 (2010) simulation was extended for an additional 500 years, while the other simulations were extended only 100 years past the hosing cessation. Similarly, two GOM hosing simulations, with the water added 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 which 0.1 Sv was added to the North Atlantic for 500 years was included to test the response to the equivalent volume of water as the NATL_0.5 experiment but over a longer period and with a reduced rate. Fig. 2. Annual Atlantic meridional overturning circulation (Sv) for (a) LGM control simulation and (b) last 20 years of hosing in NATL_1.0 experiment. Positive indicates a clockwise circulation. In order to investigate the climate sensitivity to freshwater discharge into the oceans from melting ice sheets during glacial conditions, we have conducted several freshwater forcing experiments, so-called hosing experiments, using CCSM3 and by varying amounts of freshwater added to two locations in the Atlantic Ocean. For our initial state, we started with the CCSM3 simulation for LGM. As with all freshwater fluxes exchanged between the ocean and the other CCSM components, the imposed freshwater was added to the CCSM3 ocean as a negative virtual salinity flux to one of two regions: the subpolar North Atlantic and the Gulf of Mexico (Fig. 1a). The subpolar North Atlantic region is defined as N, the same as the CMIP freshwater hosing simulations for present-day forcings (Stouffer et al., 2006). Four NATL hosing simulations have been performed, with rates of Sv (Table 1). The freshwater was added for 100 years starting at year 400 of the LGM simulation. After 100 years, the hosing was stopped and the climate system allowed to recover. The NATL_1.0 Table 1 Summary of experiments, ESL is the equivalent sea level change. Experiment Region Amount (Sv) Years of hosing ESL equiv (m) NATL_1.0 North Atlantic NATL_0.5 North Atlantic NATL_0.25 North Atlantic NATL_0.1 North Atlantic NATL_0.1ex North Atlantic GOM_0.5 Gulf of Mexico GOM_0.28 Gulf of Mexico 0.28 a a Because of an error in calculation, the experiment for the freshwater delivery to the Gulf of Mexico with the smallest rate of hosing was 0.28 rather than 0.25 Sv. a b 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 down completely and continues to transport some heat poleward in the North Atlantic (Fig. 2b). For the index of the intensity of the AMOC in this discussion, we chose the value of the positive overturning streamfunction at 34 S, 814 m depth, which was the location of the maximum at 34 S in the LGM control simulation. This value corresponds to the transport of upper water that flows northward into the Atlantic basin and sinks in high latitudes to form the model s North Atlantic Deep Water. As the AMOC slows down, northward ocean heat transport is reduced at all latitudes in the Atlantic and reverses direction to a southward transport of heat at the equator. The less dense freshwater acts as a lid on the surface ocean suppressing deep ocean convection and mixed layer depths are less than 450 m over all of the North Atlantic (not shown). This freshwater lid also suppresses interaction of the deep ocean with the atmosphere. The annual surface air temperatures cool in the region of freshwater forcing (Fig. 3). Extensive sea ice covers the North Atlantic poleward of 45 N latitude in the annual mean. This sea ice amplifies the cooling due to its high albedo and by further insulating the atmosphere from the ocean. Greatest cooling of annual surface air temperatures occurs over the North Atlantic Ocean from south of Greenland and Iceland and extends into the GIN Seas, with cooling in excess of 10 C(Fig. 3). Annual cooling over central Greenland is w6 C. A tongue of cooling of the atmosphere and ocean and freshening of the ocean extends southwestward from Iberia into the subtropical Atlantic and across the Isthmus of Panama (Figs. 3 and 4c, d). This cooling is associated with the export of cold fresh water from the subpolar region and advection equatorward by the subtropical gyre. The reversal of the ocean heat transport to southward at the equator leads to a bipolar response in annual surface air temperatures with warming in the Southern Hemisphere (SH), greater than 1 C over a broad area of the subtropical South Atlantic (Fig. 3). The Intertropical Convergence Zone (ITCZ) over the tropical Atlantic moves south into the warmer hemisphere. Precipitation increases over the tropical Atlantic and Brazil south of the equator and decreases over the tropical north Atlantic and northern Amazon (Fig. 4e, f). Precipitation also decreases over the 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 Southern Oceans, reducing the sea ice fraction at S which further enhances the warming in this region (Fig. 3). Warming over the Antarctic continent is modest and less than 1 C. Warming also occurs over northern South America in association with the reduction in precipitation and cloudiness. Cooling of more than 2 C extends downstream from the North Atlantic into the mid-latitudes of Europe and Asia. Sea ice fraction increases over the North Pacific with enhanced cooling associated with this sea ice. North America shows no significant annual mean cooling, even a slight but not significant warming, during the hosing of the North Atlantic.

6 60 B.L. Otto-Bliesner, E.C. Brady / Quaternary Science Reviews 29 (2010) 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). The mid and high latitudes of the North Atlantic subsurface ocean warm in response to the addition of freshwater to the surface of the subpolar ocean (Fig. 4a, b). The fresh layer at the surface and the large expansion of insulating sea ice inhibit the transfer of heat from the ocean to the atmosphere. Additionally, the AMOC diminishes at a lesser rate than the change in sea ice, still transporting heat northward into the North Atlantic during the hosing. This warm water is subducted below the cold fresh surface water and sea ice. The subsurface warming, extending from 100 to 3000 m, acts to keep the sea ice thickness in the GIN Seas only 1 2 m thick by enhancing basal melting. In the Northern Hemisphere (NH) subtropics, the cold, freshwater at the surface is subducted into the southwestward flow of the subtropical gyre and appears as a cold, fresh anomaly extending to 600 m deep. In the SH tropics and subtropics, the subsurface ocean warms due to reduced AMOC and reversed ocean heat transport (OHT) south of the equator and downward mixing of heat from the thermocline (detailed analysis of ocean response will appear in Brady and Otto-Bliesner, in preparation). the ITCZ and reduced precipitation and cloud cover over northern South America enhances the warming in DJF. North America also shows a slight warming in DJF, extending from the southeast US to Hudson Bay, though this warming is not significant in relation to the interannual temperature variability of the LGM control at these locations. The low-latitude oceans do not show any significant seasonality to their surface temperature responses. The cooling over the NH low-latitude oceans is maintained by the southward advection of cold fresh arctic water on the eastern side of the subtropical gyre, and over the SH low-latitude oceans by the reversal of the ocean heat transport at the equator, both of which have time scales longer than the seasonal time scales. There is significant seasonality in the SH high-latitude temperature response with greater warming during SH winter than summer. JJA surface air temperatures warm in excess of 1 Cover the Southern Ocean, associated with a reduction of sea ice, most significantly at latitudes from 45 to 60 S(Fig. 5) Seasonal response to freshwater hosing There is strong seasonality in the NH high-latitude temperature and sea ice response in the NATL_1.0 experiment (Fig. 5). During NH winter, persistent sea ice covers the sea surface from eastern US to Iberia. December January February (DJF) surface air temperatures cool in excess of 15 C in a broad region from the Labrador Sea to the British Isles. During NH summer, sea ice retreats dramatically, having only 50% coverage in the Labrador and Greenland Iceland Norwegian (GIN) Seas and along the northeast coast of Newfoundland. This exposes the freshwater cap to the summer insolation. Cooling is less than 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 temperature responses are also found over North and South America in the NATL_1.0 experiment. Over South America, the southward shift of 3.3. Transient behavior and recovery of the annual-averaged climate The North Atlantic high-latitude climate responds swiftly to the 1 Sv freshwater forcing in the NATL_1.0 experiment (Fig. 6). Sea ice forms rapidly and its areal extent increases by 50% reaching a maximum in the first decade of hosing. Greenland temperatures, as well as those over northwest Europe, also cool by 6 C in the first decade. Greenland temperature shows a high degree of correlation with the NH sea ice area, both in the mean response during the hosing and the shorter term variability, establishing the primary role 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 first decade, 40% after 50 years, and then a more gradual but continued decline over the next 50 years.

7 B.L. Otto-Bliesner, E.C. Brady / Quaternary Science Reviews 29 (2010) a c e b d 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) subsurface ocean temperature ( C) averaged over depths of m; (c, d) surface salinity (psu); and (e, f) precipitation (mm/day). In contrast to the rapid response during hosing, the high latitudes in the North Atlantic are slow to recover. Immediately after the freshwater forcing is stopped, the North Atlantic surface climate starts to recover. Surface air temperatures over Greenland warm by 5 C in the first few years and sea ice retreats. This initial recovery is fueled by the resumption of convection in the ocean, with maximum late winter-early spring mixed layer depths increasing to m in the Labrador and Irminger Seas. The warming at subsurface depths is mixed up to the surface, warming sea surface temperature and reducing the sea ice extent. The warmer sea surface quickly exhausts its heat to the atmosphere as sensible and latent heat. As well, precipitation increases, freshening the ocean surface. This recovery though is short-lived. During the second decade after the freshwater forcing is stopped, NH sea ice area increases as the ocean cools and the surface remains fresh supporting increased stability. The North Atlantic sea ice grows not only extensive in area as during the hosing, but also thicker with less interannual variability. Basal melting of the sea ice, which continued throughout the hosing in association with warmer subsurface temperatures, is reduced by a factor of two (not shown). GIN sea ice becomes much thicker (3 4 m). With a much reduced AMOC and northward ocean heat transport at the end of the hosing, the subsurface temperatures can only very slowly warm. Fifty years into the recovery, the AMOC strength reaches a minimum at about 20% of its pre-hosing LGM strength and then recovers slowly, taking approximately 500 years to reach the pre-hosing glacial strength of w15 Sv. Greenland temperatures also recover slowly in concert with the recovery of NH sea ice to its pre-hosing extent. Fig. 8a shows the difficulty in using time series of Greenland temperature as a proxy for the AMOC temporal behavior. During the first 5 years of the freshwater forcing, Greenland temperature cools by w6 C, most of its cooling in this experiment, while the AMOC decreases by less than 10%. Over the next 95 years of freshwater forcing, the AMOC continues to slow, while Greenland temperatures remain at about 40 C, though with considerable variability in concert with variability in the sea ice extent. The rapid but aborted warming of Greenland temperatures immediately after

8 62 B.L. Otto-Bliesner, E.C. Brady / Quaternary Science Reviews 29 (2010) 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 (fractional concentration), and (d) JJA sea ice extent (fractional concentration). the freshwater forcing is stopped is accompanied by only a small temporary increase in the AMOC. Over the next 50 years of the recovery, the AMOC returns to its weak state while Greenland temperatures remain between 40 and 42 C. During the remaining 450 years of recovery, the AMOC and Greenland temperatures change both more synchronously and proportionally. The climate in the tropics responds more gradually to the freshwater input (Fig. 6). Cooling of Cariaco surface temperature lags the AMOC response during the first decade of the hosing, the time scale for the freshwater in the high latitudes to be advected to subtropical and tropical latitudes by the gyre circulation. Temperature then cools by w3 C during the next 40 years of the hosing, before stabilizing at a temperature of w19 C for the remainder of the freshwater forcing period. With the shift in the Atlantic ITCZ, rainfall in northern South America decreases by w30% by the end of the hosing. The correspondence of the Amazon rainfall decreases with the Cariaco cooling and South Atlantic warming (discussed later) rather than the Greenland cooling suggests a low rather than high latitude control on the shift in the ITCZ during the freshwater perturbation. The shift in the ITCZ also affects the wind stresses in the Caribbean, with upwelling in the Cariaco region increasing by 30 50% during the freshwater forcing. This upwelling of cool subsurface waters enhances the cooling of the Cariaco surface temperatures. In contrast to Greenland temperature, Cariaco temperature recovers relatively quickly. Cariaco Basin, in the tropical North Atlantic, recovers two-thirds of its cooling in the first two decades after the end of the anomalous freshwater forcing in the subpolar North Atlantic. Similarly, northern Amazonia precipitation recovers rapidly, returning to almost normal rainfall in the first two decades. The rapid recovery is tied to the low-latitude temperatures rather than high-latitude temperature and sea ice conditions. Once the

9 B.L. Otto-Bliesner, E.C. Brady / Quaternary Science Reviews 29 (2010) Fig. 6. Transient behavior of NATL_1.0 experiment before (negative years), during (years 0 100) and after (years ) 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. Regions defined in plot labels. Three-year running mean applied to annual averages. hosing is stopped, the tongue of cold, relatively fresh water advected by the subtropical gyre in the eastern basin disappears, and the subtropical surface ocean, with less cloudy skies and therefore abundant solar absorption quickly warms. North Atlantic and Greenland temperatures, on the other hand, remain cold, and sea ice extends to 40 N. Cariaco upwelling remains enhanced for a few decades after the cessation of the hosing before declining gradually to pre-hosing levels (Fig. 6). Fig. 8b shows that the time evolution of Cariaco temperature (precipitation from northern South America shows similar

10 64 B.L. Otto-Bliesner, E.C. Brady / Quaternary Science Reviews 29 (2010) Fig. 7. Transient behavior of NATL_1.0 experiment before (negative years), during (years 0 100) and after (years ) 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 Dome C 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 mean applied. behavior) is also not perfect as a measure of the AMOC temporal behavior. Except for the short lag at the beginning of the freshwater forcing, there is a linear relationship between the simulated Cariaco temperature and AMOC during hosing, but this relationship is not maintained during recovery. In the first decade following the cessation of the hosing, there is a rapid warming from w19 Cto 20.4 C about half of the recovery in temperature that is not accompanied by a significant change in the strength of the AMOC (Fig. 6). For the next 50 years, the AMOC remains weak, while the Cariaco temperature continues to increase by about 0.6 C. For the

11 B.L. Otto-Bliesner, E.C. Brady / Quaternary Science Reviews 29 (2010) 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. Regions defined 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 500 years of recovery after the cessation of the freshwater forcing. rest of the recovery, Cariaco temperature and the AMOC again are linearly related, but the amount of Cariaco cooling for an AMOC decrease during the freshwater forcing is three times greater than the amount of warming for an AMOC increase during this part of the recovery. The South Atlantic warms by w1 C during the hosing with the warming starting only in the second decade of the freshwater perturbation (Fig. 7). The lag of the response of the South Atlantic is such that the South Atlantic temperatures do not start warming until the AMOC has already been reduced to 11 Sv, approximately 70% of its pre-hosing strength (Fig. 8c). The temperatures warm rapidly, though only by about 1 C as the AMOC continues to decrease. In contrast, the South Atlantic sea surface temperatures cool slowly with the same time scale as Greenland warms after cessation of the freshwater perturbation. Over the East Asian monsoon region, surface temperature rather than precipitation shows a response to the freshwater forcing (Fig. 7b, c). During the freshwater perturbation, surface air temperatures cool gradually by w2.5 C, in better correspondence with the gradual slowing of the AMOC then the more rapid cooling in Greenland. After the cessation of the hosing, Hulu Donge temperatures warm gradually reaching the pre-hosing climate state after w400 years. In contrast, annual precipitation in the Hulu Donge region shows considerable interannual variability but no discernable signal in response to the freshwater forcing or correlation with the Atlantic changes. Antarctica simulated surface air temperatures at EPICA DOME C and Donning Maud Land warm throughout the freshwater perturbation and for about years after its cessation to about 1 C warmer than the pre-hosing temperatures but with significant interannual and decadal variability (Figs. 7 and 8d). Even 500 years after the freshwater forcing has stopped, the EPICA Dome C temperature has not completely recovered to its initial state. This much 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 Atlantic should be in antiphase with the North Atlantic, the heat reservoir of the Southern ocean should lengthen its response. 4. Sensitivity to rate, location, and duration of freshwater forcing Fresh water added to the Labrador and GIN Seas has a direct and immediate effect on the LGM convection south of Greenland, while freshwater added to the Gulf of Mexico has a more indirect effect with the freshwater first needing to be transported by the subtropical and subpolar gyres to the LGM convection sites. The AMOC reduction is less for the smaller freshwater forcings, but is not

12 66 B.L. Otto-Bliesner, E.C. Brady / Quaternary Science Reviews 29 (2010) 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 freshwater forcing, and (b) Atlantic northward ocean heat transport (PW) averaged over the last 20 years of hosing in each experiment. linearly proportional with the size of the freshwater perturbation (Fig. 9a, Table 2). The NATL_1.0 and NATL_0.5 simulations both give large reductions in the AMOC. The AMOC is reduced less for a given sized forcing into the Gulf of Mexico compared to when the freshwater is added directly into the subpolar North Atlantic. Not all the fresh water added to the GOM makes it into the high latitudes of the North Atlantic and thus active in the freshwater forcing of the AMOC. The percentage reduction in the AMOC for GOM_0.5 is similar to the NATL_0.25 suggesting that GOM freshwater forcing is only about 50% as effective as NATL freshwater forcing. Extending the freshwater hosing to 500 years in the NATL_0.1ex experiment results in an additional reduction of the AMOC as compared to the NATL_0.1 experiment but a more vigorous AMOC than the NATL experiments with greater rates of hosing. All the freshwater forcing experiments show a reduction of the northward transport of heat in the Atlantic by the ocean as Table 2 Comparison 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 wshutdown a 81% 70% 50% 29% 40% 51% 35% Greenland T, H1 6 C b 5.6 C 7.6 C 7.6 C 6.2 C 5.8 C 6.3 C 5.5 C Iberian Margin HE, SST 8 to 4 C, SSS 3 to 1 psu c 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 psu d 5.3 psu 1.8 psu 0.3 psu 0.3 psu 0.4 psu 11.0 psu 5.2 psu Cariaco T, H1 0.5 C e,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 C N. Amazon Precipitation 32% 11% 4% 4% 1% 2% 2% Nordic Seas, HE Intermediate T þ2 to4 C f 4.2 C 3.9 C 3.4 C 2.0 C 3.8 C 0.9 C 0.7 C S. Atlantic, HE Mid-depth T þ1 to3 C g 4.0 C 2.8 C 1.5 C 0.5 C 2.0 C 2.1 C 1.2 C EPICA Dome C, AIM 2 to 8 þ 0.5 to 3 C h 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, N, W), Iberian margin (SST and SSS, N, W), Gulf of Mexico (SSS, N, W), Cariaco (surface air temperature, 5 15 N, W), Northern Amazon (% change of precipitation, 10 S 5 N, W), Nordic Seas (ocean temperature, N, 10 0 W, 730 m), South Atlantic (ocean temperature, 15 5 S, 5 15 E, 470 m), EPICA Dome C (surface air temperature, S, 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., b Cuffey and Clow, c Bard et al., d Flower et al., e Lea et al., f Rasmussen and Thomsen, g Ruhlemann et al., h Stenni et al., 2003; EPICA Community Members, 2006.

13 B.L. Otto-Bliesner, E.C. Brady / Quaternary Science Reviews 29 (2010) 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. Equatorward of 35 N, the reduction of Atlantic OHT is more sensitive to the amount of freshwater forcing. A reversal to southward transport near the equator only occurs in the NATL_1.0 and NATL_0.5 experiments. The ocean heat transport in the Atlantic is very similar for 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 to additional freshwater input of all magnitudes, locations, and durations due to the large feedbacks among the ocean, atmosphere and sea ice. Owing to the very cold high latitude oceans with near freezing temperatures at LGM, extensive sea ice forms in the North Atlantic in all the experiments. Surface air temperatures are C colder from 50 to 70 N(Fig. 10). The cooling of Greenland temperatures is relatively insensitive to the amount or location of the hosing, ranging from 5.5 to 7.6 C cooling, and is not proportional to the percentage reduction of the AMOC (Table 2). As the amount of freshwater added to the subpolar North Atlantic decreases, the Greenland cooling and NH sea ice expansion becomes less abrupt, though even in the NATL_0.25 simulation most of the cooling and sea ice growth occurs within years of the start of the freshwater perturbation (not shown). In the NATL_0.1 simulation, the Greenland cooling and NH sea ice expansion occurs more gradually over the entire 100 years. Greenland temperatures show similar cooling for the GOM experiments but with a 20 year lag to the start of the hosing for the freshwater to be advected into the North Atlantic convection regions 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 and location of the freshwater forcing (Figs. 10 and 11; Table 2). Located just south of the subpolar North Atlantic hosing area, the surface ocean in the mid-latitudes of the northeast Atlantic is strongly affected by the tongue of cold, freshwater advected out of the subpolar region, with increasing reductions in salinity and temperature as the rate of freshwater added in the NATL simulations increases. The total volume of water added has less influence on the SST and SSS responses at the Iberian margin, with the NATL_0.1 and NATL_0.1ex experiments having more similar reductions in temperature and salinity than the NATL_0.5 and NATL_0.1ex experiments. For the GOM simulations, the cooling and freshening at the Iberian margin is less than the comparable NATL simulations and lags the forcing by w25 30 years. As the rate of freshwater addition to the subpolar North Atlantic decreases, so does the decrease of temperature and salinity in the tongue of water advected southwestward from the Iberian margin into the Caribbean (Figs. 10 and 11, Table 2). The cooling of surface air temperature at Cariaco is less than 1 C and the precipitation in the northern Amazon is reduced by less than 10% for all cases except the cases with the larger freshwater forcings of 0.5 and 1 Sv added into the subpolar North Atlantic. The CCSM3 freshwater forcing experiments suggest a threshold rate of freshwater forcing for a detectable change above the natural variability at subtropical and tropical sites. Salinity decreases in the Gulf of Mexico during hosing are greatest for the GOM cases, as expected, but also show a significant salinity change of more than 5 psu in the NATL_1.0 case (Fig. 11). Surface warming found in the tropical South Atlantic is greater for larger freshwater forcings, but varies within a small range (Fig. 10). In the Nordic Seas, the subsurface ocean at mid-depths warms in response to the addition of freshwater to the North Atlantic. Even when the freshwater is added to the Gulf of Mexico, it takes only a decade to be transported to the subpolar North Atlantic Fig. 10. Comparison of annual surface air temperature anomalies ( C) averaged over the last 20 years of freshwater forcing for experiments with varying rates, locations, and duration of freshwater forcing.

14 68 B.L. Otto-Bliesner, E.C. Brady / Quaternary Science Reviews 29 (2010) and weaken deep convection. The subsurface warming is greatest for the larger NATL forcings and weakest for the GOM forcings (Table 2). The total volume of water, rather than the rate of freshwater forcing, is important for the subsurface responses, with the NATL_0.5 and NATL_0.1ex experiments showing comparable warmings in the Nordic Seas and South Atlantic (Table 2). Also interesting and relevant to proxy records from the Gulf of Mexico is that at mid-depths of w450 m, the NATL_1.0 experiment has cooling of 1 2 C, while the NATL_0.25 and NATL_0.1 experiments have warming of 1 2 C (not shown). Warming in the Southern Ocean and Antarctica shows the strongest relationship with the duration of the forcing rather than either the rate or total volume of water added to the North Atlantic (Fig. 10). The NATL_0.1ex experiment with the duration of the freshwater forcing increased from 100 to 500 years shows warming of surface air temperatures over wide areas of the Southern Ocean. EPICA Dome C surface air temperatures have warmed by 1.2 Cby the end of the hosing. The CCSM3 experiments only explore a small subset of possible freshwater forcing rates and durations though qualitatively agree with the EDML record which shows a linear relationship between the amplitudes of Antarctic warmings and the duration of the accompanying stadial in Greenland during MIS 3(EPICA Community Members, 2006). Only the NATL_1.0 simulation has been run long enough to allow the AMOC to recover to the pre-hosing state of w15 Sv. The recovery time is w500 years. The other simulations only include the first 100 years of recovery (Fig. 9a). The slope of the recovery starting two decades after the end of the freshwater forcing is similar in all the simulations. Extrapolating these curves, suggests that the AMOC in the NATL_0.5 and GOM_0.5 simulations would recover in w350 years, 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 seem to have shorter recovery times, because there is less to recover. They also 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 Fig. 11. Comparison of annual surface salinity anomalies (psu) averaged over the last 20 years of freshwater forcing for the experiments with varying rates, locations, and duration of freshwater forcing. Although the freshwater simulations described in this paper are highly idealized, it is instructive to compare the responses broadly to published proxy reconstructions for Heinrich events and the Younger Dryas (Table 2). At the end of the freshwater forcing, all experiments show reduction of the AMOC. The impact of this slowdown of the AMOC on temperature and precipitation at proxy locations along a north south transect in the Atlantic varies among locations and is not always linearly related to either the amount of the freshwater perturbation or the reduction of the AMOC. Measurements of 231 Pa/ 230 Th in core GGC5 from the Bermuda Rise indicate a slowdown of the AMOC starting at w19 ka, with a nearly complete shutdown of the AMOC from 17.5 ka until 15 ka (McManus et al., 2004). Using an arbitrary cutoff of 50% reduction in the AMOC, all experiments except the NATL_0.1 and GOM_0.28 give substantial slowdown of the AMOC by the end of 100 years of freshwater forcing. The weakest NATL experiment, NATL_0.1, when extended for another 500 years of forcing, shows a further reduction of the AMOC, 40% at year 500 as compared to 29% at year 100. It is noteworthy that when sufficient freshwater is added to the Gulf of Mexico, 0.5 Sv in our 100-year hosing experiments, the AMOC reduction is greater than 50%. The simulated annual cooling of Greenland temperatures ranges from 5.5 to 7.6 C. This cooling is comparable to the GISP2 d 18 O with borehole temperature calibration, which indicated that Greenland annual mean temperatures cooled up to w4 6 C during H1 and H2 (Cuffey and Clow, 1997). The CCSM3 response also corroborates the strong seasonality of the NH temperature response at high