Ocean oxygen isotope constraints on mechanisms for millennial-scale climate variability

PALEOCEANOGRAPHY, VOL. 20,, doi:10.1029/2004pa001063, 2005 Ocean oxygen isotope constraints on mechanisms for millennial-scale climate variability Steffen Malskær Olsen Danish Meteorological Institute, Copenhagen, Denmark Gary Shaffer Danish Center for Earth System Science and Department of Geophysics, University of Copenhagen, Copenhagen, Denmark Department of Geophysics, University of Concepcion, Concepcion, Chile Christian J. Bjerrum Geological Institute, University of Copenhagen, Copenhagen, Denmark Received 23 June 2004; revised 27 October 2004; accepted 23 December 2004; published 26 February 2005. [1] Millennial-scale climate variability pervades the last several million years of Earth history with largest amplitudes during moderate glacial conditions. However, the processes behind such variability remain unclear. Here we present results from a simplified, coupled climate model that includes an explicit treatment of the atmosphere-ocean cycle of oxygen-18 isotope. The model exhibits a relatively strong North Atlantic overturning circulation for present day and last glacial maximum conditions. Self-sustained, millennial-scale oscillations with strong and weak overturning states are found for moderate glacial conditions. These oscillations reproduce observed features like sawtooth structure, Northern-Southern Hemisphere asynchrony, sensitivity to climate state, and oxygen isotope signals in the ocean. Amplitudes, structures, and phasing of observed millennial-scale variability in high-resolution, benthic oxygen isotope records from the North Atlantic Ocean are consistent with large heating/cooling cycles in the ocean interior, as found in the model. Glacial meltwater, changes in solar irradiance and high-frequency climate variability affect the timing, period, and persistence of the model climate cycles. However, the basic dynamics, structure, amplitude, and timescale of variability are due to internal, selfsustained oscillations in the model ocean-atmosphere system. Citation: Olsen, S. M., G. Shaffer, and C. J. Bjerrum (2005), Ocean oxygen isotope constraints on mechanisms for millennial-scale climate variability, Paleoceanography, 20,, doi:10.1029/2004pa001063. 1. Introduction [2] Ice core records from Greenland and Antarctica have revealed strong and pervasive, millennial-scale climate cycles during the last glacial period [Dansgaard et al., 1993; Blunier et al., 1998]. Ocean sediment records have documented significant millennial-scale variability during the last several million years with less, but persistent, variability during interglacial and maximum glacial conditions [McIntyre et al., 2001; McManus et al., 1999; Bond et al., 1997, 1999]. Greenland records of d 18 O w (d 18 O in water) show Dansgaard-Oeschger (DO) climate cycles, characterized by abrupt, decade-scale warming of 5 10 C to a warm, interstadial state, subsequent slow cooling and then abrupt, decade-scale cooling back to a cold, stadial state (Figure 1a). In contrast, Antarctic d 18 O w records indicate slow warming, before initial, rapid, northern warming, followed by slow cooling, simultaneous with slow, northern cooling (Figure 1b). This north-south structure and phasing is best defined during Heinrich (H) events (H1 H5 in Figure 1d) of increased ice rafted debris from massive iceberg discharges due to surging of the Laurentide ice sheet Copyright 2005 by the American Geophysical Union. 0883-8305/05/2004PA001063 [Bond et al., 1993]. However, the other DO cycles exhibit similar north-south structure and phasing. In Greenland, excursions to the warm state become shorter and less frequent toward the glacial maximum around 20 kyr before present (B.P.; Figure 1a) [Schulz et al., 1999]. Any explanation for millennial-scale climate variability should be able to account for these ice core observations [e.g., Ganopolski and Rahmstorf, 2002; Schulz et al., 2002]. However, since the ocean is a key player in such variability [Stocker and Johnsen, 2003], it is equally important that any such explanation can also account for high-resolution, ocean sediment observations. [3] Large decreases of 1% or more are found in planktonic d 18 O c (d 18 O in carbonate) in a high-resolution sediment core from north of Iceland (62 N) during H event stadials in Greenland (Figures 1a and 1c) [Rasmussen et al., 1996a]. A 1% decrease in d 18 O c could reflect a temperature increase of about 4 C, a salinity decrease of about 2% or some combination of the two. Smaller planktonic d 18 O c decreases less than 0.5% are found during other DO stadials. These d 18 O c features have a symmetric shape over time. Similar levels of millennial-scale, planktonic d 18 O c variability are found in high-resolution cores from north of 40 N in the North Atlantic with less variability in cores to 1of19

increases tied to rapid Greenland warming events. Features associated with H3 and H4 are exceptions to this pattern. Slow benthic d 18 O c increases of 0.5% or less at the 3000 m depth of the 37 N core off Portugal track slow cooling in Greenland and Antarctica (Figure 1e) [Shackleton et al., 2000]. In both benthic records, d 18 O c values decrease over about 1 kyr before rapid northern warming. These interrelationships are apparent for strong H events but also for weaker DO cycles, indicating similar dynamics for stronger and longer, millennial-scale climate cycles and weaker and shorter cycles [Bond et al., 1999]. Throughout the North Atlantic Ocean, millennial-scale benthic d 18 O c variability in high-resolution cores greatly increases from the deep ocean to intermediate depths (Figure 2). [5] Here we interpret both the ice and sediment core records with the help of simulations from a simplified, coupled climate model that includes the ocean-atmosphere d 18 O w cycle. Ocean d 18 O c is calculated explicitly from Figure 1. Millennial-scale variability in ice and deep sea sediment records. The d 18 O w records are from (a) a Greenland ice core [Grootes and Stuiver, 1997] and (b) an Antarctic ice core [Johnsen et al., 1972]. The d 18 O c records are from (c) high northern latitude planktonic foraminiferal calcite [Rasmussen et al., 1996a], (d) intermediate depth, high northern latitude benthic foraminiferal calcite [Rasmussen et al., 1996a], and (e) deep, middle northern latitude benthic foraminiferal calcite [Shackleton et al., 2000]. Figures 1a and 1b are on the GISP2 age scale [Blunier and Brook, 2001]. Figures 1c and 1d were fitted to this scale using magnetic susceptibility [Dokken and Jansen, 1999], and Figure 1e was fitted to this scale using planktonic d 18 O c [Shackleton et al., 2000]. Several Heinrich events are indicated (H2 H5) [Bond and Lotti, 1995]. Stippled lines aid visual correlation. Freshening and warming arrows are interpretations guided by our model d 18 O c calculations (see Figures 7 and 8). the south (Figure 2). An exception to this rule is strong variability in a planktonic d 18 O c record at 37 N off Portugal [Shackleton et al., 2000]. However, this may reflect strong zonal gradients in millennial-scale variability across the Atlantic Ocean in the 30 40 N latitude band [Vautravers et al., 2004]. [4] Large decreases of 1% or more are found in benthic d 18 O c at the 1000 m depth at 62 N during H event stadials. However, quite large benthic d 18 O c decreases of 0.5 1% coincide with other DO stadials (Figures 1a and 1d) [Rasmussen et al., 1996a]. Most of these d 18 O c features exhibit sawtooth shape and terminate with rapid Figure 2. Standard deviation (s) from high-resolution, filtered records (periods 12 kyr retained) of planktonic and benthic d 18 O c over the interval 50 20 kyr B.P. The one-sided error bars show estimated s prior to smoothing by moderate bioturbation, given core sedimentation rates [Anderson, 2001]. The records are from the North Atlantic sites, numbered as follows: 1, 6 N 44 W; 2, 5 N44 W [Curry et al., 1999; 3, 37 N 10 W[Shackleton et al., 2000]; 4, 4 N 43 W; 5, 43 N 30 W; 6, 55 N 14 W [Vidal et al., 1998]; 7, 61 N 23 W, [Curry et al., 1999]; 8, 67 N 04 E [Dokken and Jansen, 1999]; 9, 63 N 04 W [Rasmussen et al., 1996a]; and 10, 26 N 78 W [Curry et al., 1999]. Also shown are vertical s profiles for d 18 O c and d 18 O w from a model simulation shown in Figure 12. Model results were filtered as above before calculating s. 2of19

for ocean circulation and climate [Shaffer and Olsen, 2001; Toggweiler and Samuels, 1995]. Among the processes and features we consider are ocean vertical diffusion dependent on stratification, deep water formation off Antarctica by brine rejection and the seasonal cycle. We also include a novel explicit treatment of ocean-atmosphere d 18 O w cycle. We choose model parameter values guided by observations in the present day climate system and, for these parameter values, consider model solutions for present and past external forcing by atmospheric pco 2, icecap extent and insolation changes due to orbital variations [Berger and Loutre, 1991]. A detailed model description is presented in Appendix A. [7] A modern day on mode is found for preanthropogenic forcing, including relatively weak horizontal mixing to account for relatively weak modern winds (Figure 4a and Figure 3. Sketch of the geometry and components of the simple global climate model. The vertical striped bar at 40 S extends down to 2000 m depth and marks the model Drake Passage (see Appendix A for a detailed model description). Note that only the Atlantic Ocean circulation is modeled. modeled ocean temperature and d 18 O w distributions for comparison with high-resolution ocean sediment records. The model reproduces observations and reconstructions rather well for modern day and last glacial maximum forcing. Emphasis is placed on moderate glacial forcing, for which millennial-scale, self-sustained climate oscillations are found that compare well with d 18 O c from sediment cores. We also investigate the sensitivity of such oscillations to choices of model parameter values, climate noise and weak, periodic variations of solar irradiance. Model simulations are then carried out for the last glacial period, including simplified Heinrich events and slowly varying climate forcing. These simulations capture many important features from the ice and sediment core records. Finally, our results are discussed in the context of other recent work on millennial-scale climate variability. 2. Model Description and Validation 2.1. Modern Day Simulation [6] We use a simplified global climate model with atmosphere, ocean, land and sea ice components and four zone meridional resolution (Figure 3) [Olsen, 2002]. This model setup is similar to that of Gildor and Tziperman [2001]; however, our ocean submodel has fine (100 m) vertical resolution. This allows detailed treatment of ocean exchange and Southern Ocean processes, important Figure 4. Modeled modern day atmosphere and ocean conditions. (a) Meridional overturning circulation streamlines in Sv (1 Sv = 10 6 m 3 s 1 ) and (b) comparison of the seasonal cycle of zone mean, model air temperature (black lines) with observed air temperatures (gray lines). In Figure 4a the upper overturning cell carries North Atlantic Deep Water and a lower overturning cell carries Antarctic Bottom Water (positive contours: clockwise overturning). Yearly average sea ice extent is shown as white horizontal bars at the top. Note the brine rejection-driven circulation at 70 S and the surface layer, Ekman transport at 40 S, both injected into appropriate density surfaces in adjacent model zones. In Figure 4b the observed air temperatures at 1000 mbar are monthly means based on the last 40 years of the NCEP/NCAR CDAS/Reanalysis Project data. 3of19

Table 1. Climate Forcing Parameters for Model Time Slices Time Slice, kyr B.P. Atmospheric pco 2, ppmv Southernmost Ice Sheet Extent, f g, N Horizontal Mixing Scale, K h s,10 4 m 2 s 1 0 280 67.5 2.2 22 180 40 2.8 35 210 50 2.6 Table 1). This mode is characterized by deep convection in the northern North Atlantic Ocean and a rather strong upper overturning cell of the thermohaline circulation, carrying North Atlantic Deep Water (NADW). This cell fills the North Atlantic and recirculates in part by way of concentrated upwelling in the Southern Ocean and in part by more diffuse upwelling in the ocean interior. There is also a weaker, lower cell carrying Antarctic Bottom Water (AABW). This cell, forced by brine rejection in the Southern Ocean, extends northward across the equator but is mainly restricted to the Southern Ocean. These overturning cells bear considerable resemblance to, but are somewhat weaker than, observed overturning cells [e.g., Talley et al., 2003]. A comparison of the seasonal cycle of model air temperatures with zonal-averaged, observed air temperatures also shows considerable agreement (Figure 4b). The model reproduces the amplitude and phasing of wintertime cooling well in both hemispheres. However, the model underestimates summer warming at high northern latitudes and overestimates austral summer temperatures at low southern latitudes. Global mean air temperature for our modern day simulation is 16.3 C and this temperature increases by 1.9 C for a doubling of atmospheric pco 2. [8] A comparison of model ocean water mass properties with modern Atlantic Ocean observations reveals a number of common features (Figure 5). The upper ocean gradient of temperature (T) versus salinity (S) is slightly greater in the model than in data, reflecting somewhat too low T and/or too high S in the model main thermocline (Figure 5a). Likewise, the model underestimates T of NADW by 1 2 C and underestimates (overestimates) S of NADW (AABW) by about 0.2. These deficiencies may derive in part from the simple model configuration (i.e., no Pacific Ocean). On the other hand, there is good model-data agreement for the upper ocean gradient of d 18 O w versus S (Figure 5b). This lends confidence to the relatively simple, d 18 O w submodel used here (Appendix A). Formation of model shelf water from Southern Ocean surface water by brine rejection, associated with sea ice production around Antarctica, changes T and S but not d 18 O w (BR arrows in Figure 5). The admixture of glacial melt water freshens shelf water and depletes it in d 18 O w (GM arrows in Figure 5). Both processes combine to explain the appendix of low d 18 O w for S between 34.5 and 35 in model results and data. 2.2. Last Glacial Maximum Simulation [9] A glacial on mode was found for last glacial maximum (LGM) forcing at 22 kyr B.P., including larger horizontal mixing to account for stronger glacial winds [Shin et al., 2003] (Figure 6a and Table 1). This mode is also characterized by deep convection in the northern North Atlantic Ocean; however, now the upper overturning cell is about 30% weaker compared to the modern on mode and is significantly shallower, restricted to depths above 2500 3000 m throughout the North and South Atlantic oceans. Furthermore, this cell now only recirculates by way of concentrated upwelling in the Southern Ocean. The lower overturning cell of the glacial on mode transports more than twice as much AABW across the equator than for the modern on mode and AABW fills the entire deep North Atlantic below 3000 m. These features of the modeled glacial on mode compare well with LGM reconstructions [Sarnthein et al., 1994; Sarnthein et al., 2001; McManus et al., 2004]. Note that, for simplicity, neither sea ice production around Antarctica nor the Ekman transport at Drake Passage latitude were modified in the LGM simulation. Rather, the increased strength of the cross-equatorial part of the lower cell and the increased presence of AABW in the North Atlantic in the model appears to be coupled to weakening of the upper overturning cell. [10] Mean global air temperature in the LGM simulation is 13.6 C, 2.7 C colder than in the modern day simulation, whereby the Northern Hemisphere cools considerably more than the Southern Hemisphere (Figure 6b). This may be understood in terms of strong albedo cooling due to a large ice cover in the Northern Hemisphere and less ocean heat transport northward across the equator due to a weaker thermohaline circulation. However, our model may underestimate LGM cooling, particularly in the Southern Hemisphere, perhaps due in part to insufficient climate sensitivity to pco 2 changes. For example, LGM reconstructions show more extensive sea ice coverage in the Southern Ocean than we find in our model (Figure 6a) [Gersonde et al., 2003]. One important factor for the strength of the thermohaline circulation, and thereby climate, is the fresh water input to surface layer of the northern North Atlantic. We found model atmospheric water vapor transport across 40 N to be only slightly greater in the modern day simulation (0.635 sverdrup (Sv); 1 Sv = 10 5 m 3 s 1 ) than in the LGM simulation (0.614 Sv). The relative insensitivity of this transport to climate change in our model reflects near compensation of the effects on this transport by changes in meridional air temperature gradient and changes in water vapor content coupled to changing air temperatures (equation (A4) in Appendix A). [11] Modelled d 18 O w results for the LGM agree well with estimates from pore water data and with other model results [Adkins et al., 2002; Roche et al., 2004]. However, our results for LGM-modern day differences in d 18 O c for the deep (Atlantic) ocean fall short by about 0.5 0.7% of observed differences there, after assuming a mean LGM ocean value of 1.05% higher than at present [Duplessy et al., 2002]. This problem has also been found in other 4of19

modeling work [Roche et al., 2004]. The data may be interpreted as a cooling in the deep LGM Atlantic, 2 3 C greater still than the cooling of about 1 C found in our model there. Much of this discrepancy derives from the fact that our modern day NADW is 1 2 C too cold. Thus when NADW is replaced by AABW during the LGM in the model, cooling is underestimated. It is also possible that colder, windier conditions around Antarctica during the LGM may have led to more sea ice production, more brine rejection and thus more AABW formation there. These aspects are outside our present focus on moderate glacial conditions but present prospects for future study. 3. Model Climate Oscillations 3.1. Oscillation Structure and Dynamics [12] Millennial-scale climate oscillations are found in the model for external forcing at 35 kyr B.P., an intermediate value of horizontal mixing, but no external periodic forcing (Table 1). These self-sustained oscillations are coupled to large variability in surface layer salinity and interior temperature (Figures 7 and 8). Such deep-decoupling oscillations are also found in other simple models as well as 3-D Figure 5. Comparison of model ocean water mass properties for the modern day simulation with modern Atlantic Ocean observations. (a) Temperature (T) versus salinity (S). (b) The d 18 O w versus S. Heavy black lines are modeled, zone mean profiles from ocean surface to bottom. Arrows mark the transition of Southern Ocean surface layer water to shelf water (a precursor of Antarctic Bottom Water) by brine rejection during sea ice formation and export (BR) and by input of glacial meltwater (GM). Observational data are for the Atlantic Ocean between 60 S and 70 N and were taken from the Global Seawater Oxygen-18 Database (available at http://www.giss.nasa.gov/data/o18data). Note the good correspondence between the mean slope of model d 18 O w versus S and the data. Figure 6. Modeled last glacial maximum atmosphere and ocean conditions. (a) LGM meridional overturning circulation streamlines in Sv (positive contours: clockwise overturning). Yearly average sea ice extent is shown as white horizontal bars at the top. (b) Modeled, zone mean LGM temperatures minus modern day temperatures for the atmosphere (dashed lines) and the ocean surface layer (solid lines). 5of19

models [e.g., Winton and Sarachik, 1993; Winton, 1997; Sakai and Peltier, 1997; Oka et al., 2001; Schulz et al., 2002]. The behavior of our model demonstrates sensitivity to orbital forcing since pco 2, icecap extent and horizontal mixing at 35 kyr B.P. lie between modern (preindustrial) and LGM values while simulations for both of these periods yield steady states characterized by on mode conditions (Table 1). The results can be understood in terms of long Figure 7 6of19

Figure 8. Evolution of model variables for selected model zones and water depths over a self-sustained climate oscillation for 35 kyr B.P. external forcing. (a) Air temperature, (b) surface layer salinity, (c) surface layer d 18 O c, (d) surface layer d 18 O w, (e) intermediate and deep d 18 O c, and (f) intermediate and deep d 18 O w. Zone and depth identifications in Figures 8a, 8c, and 8e also apply to Figures 8b, 8d, and 8f, respectively. Ocean mean d 18 O w at 35 kyr B.P. is set to 0.75% but, for clarity, no glacial salinity offset is applied. Dots on vertical axes in Figures 8a and 8b indicate modern day simulation values (open dots are Southern Hemisphere values). Note that high latitude d 18 O c at 1000 m (Figure 8e) is in phase with surface layer d 18 O c (Figure 8c) due to temperature changes at 1000 m and salinity changes in the surface layer. term changes of low latitude (40 S 40 N) to high northern latitude (40 N 90 N) difference in yearly mean solar insolation (DQ; Figure 9). Whereas modern day and LGM orbital forcing are similar, DQ was 4 5% weaker at 35 kyr than at present or during the LGM. This weaker insolation difference leads to weaker thermal forcing of the thermohaline circulation at 35 kyr B.P., tending to destabilize the on mode. This new finding may help explain greater observed millennial-scale variability around this time (Figure 1) [see also Schulz et al., 1999]. [13] Model oscillations exhibit many observed features during DO cycles like rapid climate transitions, hemispheric asynchrony and, during stadials, warming of intermediate depth water in the tropical and North Atlantic and the invasion of the deep North Atlantic by Antarctic Bottom Water [Dansgaard et al., 1993; Blunier et al., 1998; Curry Figure 7. Model self-sustained climate oscillations for 35 kyr B.P. external forcing. (a) Time series versus depth of temperature and salinity (contoured) in the 40 90 N zone. Time slices of ocean distributions (b) after the onset of a cold phase, (c) during the off mode, (d) shortly before the onset of a warm phase, and (e) during a warm phase. Shown are overturning streamlines in Sv (top row), vertical diffusion coefficient (second row; ongoing convection is shaded), and anomalies of temperature (third row), salinity (fourth row), d 18 O w (fifth row), and d 18 O c (bottom row). Anomalies are calculated relative to local means over a complete oscillation (white contours mark zero anomaly). Sea ice extent is shown as white horizontal bars at the top of the second row. For clarity, only salinities above 34.0 are contoured and no glacial salinity offset is applied in Figure 7a. The period of the oscillation is 1.31 kyr. A sensitivity analysis (see Figure 10) shows that this period could be tuned toward the observed, 1.5 kyr period. Note how deep ocean temperature (Figure 7a) slowly increases during weak circulation conditions due to downward diffusion of heat into the low latitude ocean. 7of19

Figure 9. Time series of zone averaged, yearly mean solar insolation difference between the 40 S 40 N low zone and the 40 90 N zone from 50 kyr B.P. to the present. The vertical dashed lines mark the three different time slices considered here. The reduced thermal torque at 35 kyr B.P. permits self-sustained oscillations in the model for the prescribed atmospheric pco 2 and ice cap extent. et al., 1999; Oppo and Lehman, 1993; Rasmussen et al., 1996a]. When a halocline develops and convection ceases in the northern North Atlantic, the upper overturning cell weakens and shoals, giving way to the lower overturning cell that fills the deep Atlantic with cold Antarctic Bottom Water (Figures 7a and 7b). Without cooling by North Atlantic convection, the ocean interior above about 2500 m warms slowly by downward diffusion of heat at low latitudes (Figures 7a and 7c) [cf. Rühlemann et al., 2004]. Some of this heat is transported poleward in subsurface layers by the weak, shallow upper overturning cell and horizontal mixing. In the Southern Ocean, this heat is upwelled to the ocean surface, warming the atmosphere (Figures 7d and 8a). In the northern North Atlantic, this poleward heat transport creates a growing temperature inversion that weakens upper ocean stratification, promoting increased vertical mixing and, finally, the onset of convection there (Figures 7a, 7d, and 7e). The switch to convection and strong overturning is self-reinforcing as saltier, warmer water brought to the surface from below and from the south cools rapidly by air-sea exchange. Heat accumulated in the ocean interior is thereby released rapidly to the atmosphere and the sea ice edge moves from 60 N to 80 N within decades (Figure 7e). The heat output from the model ocean and the albedo decrease quickly and raise Northern Hemisphere air temperature to near present day levels despite low pco 2 and a large continental ice cover (Figure 8a). At the same time, convection cools the ocean interior above about 2500 m and the upper overturning cell fills the deep North Atlantic (Figure 7e). The cell gradually weakens because of cooling of the low latitude thermocline as warmer water there is transported poleward and replaced by cooler water from the south and from below. Less poleward salt transport in the weakened cell and continued fresh water input from the atmosphere promote gradual halocline buildup and, finally, convection shutdown in the northern North Atlantic. The oscillation period of order 1 kyr can be related to a diffusive timescale for the ocean thermocline, L 2 /k, where L is the ocean thermocline depth and k is the vertical diffusion coefficient in the thermocline. For example, L =110 3 m and k =2 10 5 m 2 s 1 [Ledwell et al., 1993] yields a timescale of about 1.5 kyr. 3.2. Comparisons With Observations [14] Air temperature amplitude, structure and north-south asynchrony over model cycles compare favorably with Greenland and Antarctic data (Figures 1a, 1b, and 8a). Model d 18 O c changes due to surface warming of 2 C in the 0 40 N model zone are consistent with low latitude planktonic d 18 O c data [Curry et al., 1999]. Warming at high southern latitudes during cool, off mode conditions in the north is often explained by weaker cross-equatorial heat transport in a weaker thermohaline circulation [Crowley, 1992]. Our results suggest that part of the observed Antarctic warming may also be explained by ocean interior warming at low latitudes, ocean transport to the Southern Ocean and upwelling into the surface layer there [cf. Rühlemann et al., 2004]. Large, abrupt increases in d 18 O c within the cycles at 1000 m depth in the 40 90 N zone reflect, in the model (Figure 8e) and likely in the data (Figure 1d), rapid cooling at the onset of convection (Figures 7a, 7d, and 7e). During the slow, interior oceanwarming phase of the cycles, model results show slow decreases in d 18 O c there. In general, data show a similar tendency although specific events deviate from this pattern, probably in part due to sedimentation rate assumptions. In the surface layer of the 40 90 N zone, d 18 O w changes, associated with fresh water balance (salinity) changes, dominate the d 18 O c signal (Figures 8b, 8c, and 8d). Thus, although model surface and intermediate depth d 18 O c in the northern North Atlantic are controlled in different ways, they vary in phase over the cycles (Figures 8c and 8e), as observed (Figures 1c and 1d) [Rasmussen et al., 1996a; Dokken and Jansen, 1999]. Slow increases and decreases in d 18 O c are found at 3000 m depth with minimum values at the onset of convection (Figure 8e). In the model, and possibly in the data (Figure 1e), this reflects interior heating and cooling cycles and, to a lesser extent, salinity and water mass changes as Antarctic Bottom Water invades the deep North Atlantic during the off phase of the cycles (Figures 7b 7d). Such circulation changes explain large observed, millennial-scale variability of benthic d 13 Cinthe deep North Atlantic during glacial time [Sarnthein et al., 2001; Curry et al., 1999]. 3.3. Sensitivity to Parameter Choices [15] We performed extensive studies of the sensitivity of our model results to choices of model parameter values for climate forcing of 35 kyr B.P. Results are quite sensitive to vertical diffusion and wind-driven circulation (estimated here by horizontal diffusion concentrated in the upper ocean), as in simpler and more complex, coupled models [Shaffer and Olsen, 2001; Oka et al., 2001]. For strong vertical and horizontal diffusion, the only model state is the 8of19

on mode (Figure 10a). Downward diffusion of heat at low latitudes enhances the thermal forcing of the upper overturning cell and the wind-driven, horizontal diffusion limits the meridional salinity gradient opposing this cell. For weak vertical and horizontal mixing, the only model state is the off mode. For sufficiently strong vertical but weak horizontal diffusion, both modes are possible steady state solutions. This multiple equilibrium behavior rests upon an advective salinity feedback on the upper cell, as shown by a similar result in a simpler coupled model [Shaffer and Olsen, 2001]. Model oscillations exist in a well-defined range of upper ocean vertical and horizontal diffusivity scales. For example, these oscillations are only found for small, upper ocean vertical diffusivity scales of less than about 3.5 10 5 m 2 s 1. Meridional overturning may be sensitive to vertical diffusion parameterizations [Nilsson et al., 2003]. However, in additional sensitivity studies (not shown) we found similar sensitivities as in Figure 10a for fixed low vertical diffusivity and for a stronger dependence of vertical diffusivity on stratification. [16] Oscillations are present in the model for a wide range of Southern Ocean forcing, demonstrating relative model insensitivity to changes in the values of Southern Ocean parameters (Figure 10b). Increased Ekman transport tends to favor the on mode while increased sea ice production and associated brine rejection tend to favor the off mode. Poleward heat transport in the atmosphere also increases when the atmospheric exchange coefficient for sensible heat is increased (Figure 10c). However, longwave radiation feedbacks in the overall atmospheric heat balance act to reduce the meridional air temperature gradient in this case. This leads to less poleward water vapor transport, favoring the on mode. The meridional air temperature gradient decreases slightly when the atmospheric exchange coefficient of water vapor/latent heat is increased; however, the net effect is an increase in poleward water vapor transport, favoring the off mode. This explains why model oscillations are rather insensitive to coupled changes of these exchange coefficients. Finally it should be noted that model oscillations are found for atmospheric water vapor trans- 9of19 Figure 10. Sensitivity of model climate regimes to parameter value changes for 35 kyr B.P. forcing. (a) Sensitivity to upper ocean, vertical, and horizontal diffusitivity scales, K v and K h, respectively, (b) sensitivity to Ekman transport, E, and sea ice formation/ice shelf melting, F i, and (c) sensitivity to exchange coefficients for sensible and latent heat in the atmosphere, K t and K q, respectively, expressed as % changes from standard values (Table A1). Five different model regimes are an on mode with strong overturning circulation in the North Atlantic, an off mode with weak overturning circulation in the North Atlantic, both on and off modes possible, 50 70 yr period oscillations in North Atlantic convection in a weak on mode situation (hatched area in Figure 10a) and millennial-scale oscillations (white; oscillation periods contoured in kyr). Other parameter values and forcing as for 35 kyr B.P. in Tables 1 and A1; dots at the figure centers mark the standard case of Figures 7 and 8.

Figure 11. Behavior of the model North Atlantic overturning circulation during a slow 30 kyr transition from warm to cold conditions forced by prescribed changes in atmospheric pco 2 and in f g, the southernmost latitude of ice cap extent. (a) Results for standard parameter values and 35 kyr orbital forcing and horizontal exchange ( wind-driven circulation ), (b) including zero-mean Gaussian noise on water vapor transport across 40 N with s = 0.0025 Sv, guided by observations over the Atlantic for a recent 40 year period [Walsh et al., 1994], otherwise as Figure 11a, (c) including a 1.5% modulation of the solar constant with a 1.5 kyr period, otherwise as Figure 11a, (d) including the weak forcings of Figure 11b and 11c, otherwise as Figure 11a. The gray vertical bars in Figure 11c and 11d mark weak solar forcing maxima. The range for which oscillations occur is broadened when noise (Figure 11b) or solar modulation (Figure 11c) are imposed. Note also phase-locking of the oscillation to the solar cycle at the perimeters of the oscillation range (Figures 11c and 11d). ports across 40 N in the range of 0.58 0.72 Sv, for our standard parameters values and 35 kyr B.P. forcing. 3.4. Influence of Weak, External Forcing [17] Millennial-scale, climate change may be influenced by pervasive, interannual- to decadal-scale variability in the real climate system. Such variability, when modeled as weak noise on the atmospheric water vapor transport across 40 N, broadens the range of climate forcing for which millennial-scale oscillations occur in the model (Figures 11a and 11b). Model oscillation periods are somewhat shorter in the presence of this noise, as on modes are triggered earlier in the cycles. Furthermore, other work has raised the possibility of weak solar forcing on millennial timescales [Björck et al., 2001; Bond et al., 2001]. The range of climate forcing for which oscillations occur is also broadened by forcing with a weak, prescribed, 1.5 kyr period modulation of the solar constant (Figure 11c). In addition, this weak forcing acts to phase-lock the oscillations into a 1.5 kyr period at the warm and cold ends of this range. Such phase-locking has also been found in other studies [Ganopolski and Rahmstorf, 2002; Timmermann et al., 2003]. Forcing by a combination of weak noise and periodic forcing leads to a still broader range of climate forcing for which oscillations occur. Again, phase-locking is found at the ends of this range; however, there is a broader interior segment without phase-locking (Figure 11d). It is important to note that the presence of weak forcing has little effect on oscillation amplitude and period, which are basically set by internal dynamics. System variables like Ekman transport and vertical and horizontal diffusion surely also vary over the real glacial-interglacial climate cycles and would also have influenced the climate range for which these oscillations are possible. The range-broadening effect of weak forcing may help explain the ubiquity of the millennium-scale climate variability in Quaternary climate records. 4. Last Glacial Period Simulations [18] A model simulation was carried out for the period 50 kyr to 20 kyr B.P. with standard model parameters, 10 of 19

Figure 12. Model simulation from 50 to 20 kyr B.P. of air temperature (T a ) and ocean d 18 O c.(a)t a in the 40 90 N zone. (b) T a in the 40 90 S zone. (c) Ocean surface d 18 O c in the 40 90 N zone. (d) The d 18 O c at 1000 m depth in the 40 90 N zone. (e) The d 18 O c at 3000 m depth in the 0 40 N zone. For this calculation the horizontal diffusivity scale, K s h, was increased linearly over this period from 2.5 to 2.8 10 4 m 2 s 1 to account for increasing wind speeds, atmospheric pco 2 was decreased linearly from 220 to 180 ppmv and the southernmost ice cap extent was advanced from 60 to 40 N such that ice cap area increased linearly with time over this period. Orbital forcing over this period was calculated from Berger and Loutre [1991]. The 40 90 N zone was forced by Heinrich meltwater cycles (d 18 O w = 30%) with 7 kyr buildup and 1 kyr release phases (both decreasing linearly to zero), guided by an ice cap surge model [MacAyeal, 1993]. Amplitudes were scaled to reflect an 8 m sea level change over the global ocean [Chappell, 2002]. The vertical stippled lines mark the transition from release to buildup phases of this forcing. For comparison with Figure 1, ocean mean d 18 O w was increased linearly from 0.5 to 1.0% over this period. Note how model fluctuations (Figures 12c, 12d, and 12e) compare with ocean sediment observations (Figures 1c 1e). slowly varying pco 2 and continental ice cover extent from observations, orbital forcing, slowly increasing horizontal mixing and imposed H events (Figure 12; see caption for details). The simulation facilitates a more direct comparison with the last glacial period observations, summarized in Figures 1 and 2. Model results mimic well many of the features in Figure 1 such as enhanced DO activity between 30 and 40 kyr B.P. and long on mode periods following H events. This enhanced DO activity in the simulation can be explained by a low difference in solar insolation between low latitudes and high northern latitudes during this period, as discussed above. Long on mode periods following H events in the model, and likely in the data, can be explained by the stabilization of the on mode via decreased fresh water flux to the northern North Atlantic surface layer, due to parameterized fresh water storage in growing ice caps. [19] Other observational features summarized in Figure 1 are also reproduced well in the simulation such as atmospheric warming at high southern latitudes during cold, off mode conditions in the northern North Atlantic and larger amplitudes in planktonic and benthic d 18 O c for H events than for DO events (Figures 1 and 12). As shown above, model atmospheric warming at high southern latitudes is largely associated with ocean interior warming during off mode conditions. Indeed, interior warming and cooling cycles explain similar, millennium-scale evolution of atmospheric temperature around Antarctica and deep, benthic d 18 O c in the North Atlantic Ocean in the model, and likely in the data (Figures 1b, 1e, 12b, and 12e). Larger planktonic d 18 O c amplitudes for H events than for DO events in the northern North Atlantic reflect iceberg meltwater of the H events in the model and in the data (Figures 1c and 12c). However, in the model and arguably in the data, greater benthic d 18 O c amplitudes for H events than for DO events in the North Atlantic are associated with longer, and thus greater, interior warming as surface layer freshening from iceberg melting during H events extends off mode conditions by suppressing open ocean convection (Figures 1d, 1e, 12d, and 12e). The model simulation also captures features in the data like in-phase variability for planktonic and benthic d 18 O c over H and DO events, symmetric event structure for planktonic d 18 O c and sawtooth event structure for benthic d 18 O c in the North Atlantic (Figures 1c 1e and 12c 12e). These features reflect in the model, and probably in the data, control of northern North Atlantic planktonic d 18 O c by surface layer salinity and control of North Atlantic benthic d 18 O c by interior temperature. [20] Model results show a strong increase in d 18 O c variability from the deep ocean up into the thermocline in the North Atlantic, as was found in the observations (Figure 2). A comparison of model d 18 O w and d 18 O c variability shows clearly that this strong increase is explained by a strong increase in model temperature variability. Model d 18 O c variability overestimates observed benthic d 18 O c variability within the North Atlantic thermocline and shows very large variability near 200 300 m, a depth range only in part captured by vertically migrating N. pachyderma (s) [Weinelt et al., 2001]. However, observed d 18 O c variability in the thermocline still greatly exceeds model d 18 O w there, suggesting that observed d 18 O c variability there reflects large, millennial-scale, temperature variability during the last glacial period. In contrast, model surface layer variability of d 18 O c in the northern North Atlantic is dominated by the meltwater signal from the H events and overestimates observed planktonic d 18 O c there. Greater variability in the 11 of 19

model than in the observations in the surface layer may reflect too large assumed H event amplitude, too low assumed d 18 O w in the iceberg meltwater, too weak near surface vertical exchange or some combination of these factors. 5. Discussion [21] One leading explanation for millennial-scale climate variability during glacial time involves switches among a modern on mode, a glacial on mode and a Heinrich off mode of the thermohaline circulation, forced by changes in the North Atlantic freshwater budget [Clark et al., 2002]. This explanation attributes DO cycles to switches between the modern and glacial on modes, associated with large horizontal displacements of deep convection sites. In this view, DO cycles are paced by millennial-scale, external forcing and perhaps amplified by stochastic resonance [Alley et al., 2001; Ganopolski and Rahmstorf, 2001]. However, for both modern and glacial on modes, deep convection would ventilate the North Atlantic with cold, salty water to well below 2000 m [Paul and Schäfer-Neth, 2003]. Therefore switches between modern on and glacial on modes would not lead to large d 18 O c variation above 2000m in the North Atlantic. Thus large benthic d 18 O c variability observed there over DO cycles need to be explained otherwise (Figures 1 and 2). [22] In other work, such observed benthic d 18 O c variability has instead been ascribed to off mode, deep ocean ventilation by d 18 O w -depleted water formed by brine rejection during sea ice formation in the northern North Atlantic [Vidal et al., 1998; Dokken and Jansen, 1999; van Kreveld et al., 2000]. However, the brine rejection hypothesis appears unlikely: In the formation of shelf water by brine rejection, surface layer salinity is raised by about 0.4 0.5 [Muench and Gordon, 1995]. In the Southern Ocean, upwelling of deep, salty water helps maintain relatively high, surface layer salinity (above 34 at present), facilitating deep water formation there by brine rejection. In the Arctic Ocean where deep upwelling does not occur at present, a more brackish surface layer forms (salinity well below 34). Then brine rejection creates lighter water that enters the shallow halocline, not the deep basin there [Bauch and Bauch, 2001]. Likewise, for an off mode with a brackish surface layer in the northern North Atlantic, brine rejection would most likely force shallow recirculation, not deepwater formation. [23] Another possible contribution to the observed benthic d 18 O c variability during H events might be glacial meltwater inputs associated with sea level changes of up to 10 m over these events [Chappell, 2002]. When spread over the global ocean, such events would only lead to a mean standard deviation of up to 0.03%. A vertical mean standard deviation in d 18 O w of up to 0.05% can be found for the North Atlantic Ocean in a simple two-box model with a conservative total exchange estimate of 10 Sv and the assumption that half of this meltwater enters the North Atlantic. In contrast, North Atlantic observations indicate a vertical mean variability of about 0.2% (Figure 2). Thus glacial meltwater effects fall short of explaining the large benthic d 18 O c variability above 3000 m in the North Atlantic during the last glacial period. [24] It appears then that large temperature changes in the interior of the North Atlantic Ocean are required to explain the large benthic d 18 O c variability observed there. To be consistent with observations, such interior temperature changes must be out of phase with atmospheric temperature changes in Greenland (Figure 1). Indeed faunal evidence has been interpreted as indicating significant warming at intermediate depths in the northern North Atlantic during the cold climate phase in Greenland [Rasmussen et al., 1996a, 1996b]. However, other interpretations in terms of food supply and oxygen content are possible, both probably related to circulation and there by temperature changes. In an off mode situation with a cold high latitude atmosphere, the ocean interior would warm as low latitude heating is no longer opposed by high latitude convective cooling [Rühlemann et al., 2004], as shown by our model results (Figure 7). This situation becomes more complicated because cold Antarctic Bottom Water invades the deep layers of the North Atlantic during an off mode. However, warming prevails above the influence of Antarctic Bottom Water. [25] There is general agreement that large meltwater pulses during H events forced off mode situations in the North Atlantic. From our model results and the discussion above, we interpret the large decreases in benthic d 18 O c at intermediate depths in the northern North Atlantic during these events to indicate warming of 4 5 C (Figure 1d). Furthermore, our model results and the above discussion also suggest that the large, but somewhat weaker, decreases in benthic d 18 O c during DO events are likely explained by somewhat weaker warming. Large warming-cooling cycles in the interior of the North Atlantic over DO events are consistent with transitions between states like on and off modes, but do not seem consistent with transitions between on modes with different convection sites [cf. Paul and Schäfer-Neth, 2003]. Thus we propose that millennial-scale climate variability during glacial time involves transitions among states like a glacial on mode, a glacial off mode and a Heinrich off mode of the thermohaline circulation. In this view, H events are essentially DO events that have been modified by changing fresh water forcing of the surface layer of the northern North Atlantic during iceberg discharge and subsequent ice cap buildup. Such an explanation of DO cycles in terms of deep-decoupling oscillations (free or forced in part by noise and weak solar forcing) is attractive since large temperature changes in the ocean interior, implied by benthic d 18 O c data and faunal composition, are a key dynamical component of the oscillations: Interior warming during the off mode eventually helps to destabilize near surface stratification in the northern North Atlantic, leading to increased vertical exchange and a rapid transition to open ocean convection and the on mode. [26] Our results also suggest that variable orbital forcing can significantly modulate the occurrence of millennialscale climate variability by modulating the thermal forcing of the thermohaline circulation (Figure 9). On the other 12 of 19

hand, our model may be too simplistic for a proper treatment of orbital forcing. For example, orbital forcing may also influence continental ice volume around the northern North Atlantic and, thereby, the strength of the thermohaline circulation by modulating fresh water inputs to the ocean surface layer there. Also, mean air temperatures in our model may be too sensitive to orbital forcing. Background air temperature in the 40 90 S zone is more strongly modulated than d 18 O w in the Byrd ice core by orbital forcing (Figures 1b and 12b). This may reflect model deficiencies like insufficient climate sensitivity to pco 2 changes or insufficient sea ice extent and export around Antarctica or may reflect orbital-modulation of source water for Antarctic ice. Our results and our suggestion that orbital forcing may significantly modulate thermal forcing of the thermohaline circulation should be further examined in the data from earlier periods during glacial time and by carrying out long runs with more complex, coupled models and orbital forcing. [27] Internal, millennial-scale variability has not been routinely found in coupled, 3-D ocean-atmosphere models. Our results suggest that this may reflect too strong vertical and/or horizontal diffusion in these models, often considerably higher than indicated by observations [Ledwell et al., 1993]. Overly diffusive models would smooth out the large temperature and salinity anomalies that lie at the heart of such oscillations. Sophisticated treatments of ocean diffusion and long coupled model runs without deep ocean acceleration of temperature and salinity are likely required for more realistic climate simulations. Likewise, high-resolution records of geochemical proxies sensitive only to temperature [e.g., Rosenthal et al., 1997; Skinner et al., 2003] are needed to confirm large temperature cycles in the ocean interior, proposed here as a hallmark of millennialscale climate change. 6. Conclusions [28] We conclude that millennial-scale climate variability during the last glacial period is best explained by deep decoupling oscillations, coupled to less frequent Heinrich events, based on constraints from modeling oxygen isotopes in water and calcite. In this view, most observed, millennialscale variability of benthic d 18 O c at intermediate depths in the North Atlantic Ocean during glacial time can be attributed to large temperature changes at these depths. Modulation of thermal forcing of the thermohaline circulation by orbital forcing may help explain why millennialscale climate variability was particularly active during the intermediate glacial conditions of Marine Isotope Stage 3 and why little such variability has been found during the Holocene or during the Last Glacial Maximum. The existence of noise in the climate system broadens the range of external forcing for which millennial-scale climate oscillations may occur. Any weak, millennial-scale, periodic changes in solar forcing would also broaden this range and also would tend to lock the oscillations into the period of such forcing, as has been found in earlier studies. However, the basic structure and amplitude of variability are set by the internal, self-sustained oscillations in the model ocean-atmosphere system and do not rely on external forcing. Appendix A [29] The simplified, global climate model contains atmosphere, ocean, land and sea ice components and is divided into four 360 wide zones bounded by 0, 40 N,S and the poles (Figure 3). The 4500 m deep, model ocean is continuously stratified with 100 m vertical resolution and consists of an Atlantic (60 wide, from 90 N to40 S) connected to a Southern Ocean (SO) (180 wide, from 40 S to70 S). Pacific Ocean heat transport is parameterized by incorporation into atmospheric transports. The model is designed for long integrations, possibly with a large number of tracers [see Olsen, 2002], and focuses on representing the vertical distributions of water masses. The division into high and low latitude zones only was motivated in part by the possibility to use simple, robust parameterizations of midlatitude atmospheric transports [e.g., Gildor and Tziperman, 2001]. A1. Atmosphere [30] We use a simple, zone mean, energy balance model for the near surface atmospheric temperature, T a ( C), forced by seasonally varying insolation, meridional transports and air-sea exchange. In combination with a sea ice parameterization, the model includes the ice albedo feedback, the insulating effect of sea ice and the seasonal cycle. The seasonal cycle permit variation in surface ocean density that leads to wintertime convection and deep water formation with the properties of the wintertime surface layer. [31] Prognostic equations for mean T a in the 0 40 and 40 90 zones, T l a and T h a, are obtained in each hemisphere by integrating the surface energy balance over the zones. Thus A l;h l;h l;h @Ta r 0 C p b @t ¼F merid a 2 Z 360 F toa F T 0 Z fm ;90 0;f m cosðfþdfdl; ða1þ where A l,h are surface areas and r 0 C p b l,h are the heat capacities for each atmospheric zone, expressed as water equivalent capacities, whereby C p is the specific heat capacity [4 10 3 J(kg C) 1 ], r 0 is the reference density of water (1 10 3 kg m 3 ), and b l,h are thicknesses (b l =5m, b h = 20 m), chosen to yield observed seasonal cycles of T l,h a. Furthermore, f m is the latitude of the zone division (40 N,S), F merid is the loss (low latitude) or gain (high latitude) of heat due to meridional transports across f m, a is the Earth s radius, and F toa and F T are the vertical fluxes of heat through the top of the atmosphere and the ocean surface. Cross-equatorial, atmospheric heat transport has been neglected in equation (A1) and a no flux boundary condition has been applied at the poles. [32] Latitudinal variations of T a in the model are represented by a second order Legendre polynomial in sine of latitude [Wang et al., 1999], T a ðfþ ¼ T 0 þ T 1 2 3 sin2 ðfþ 1 ða2þ 13 of 19