1. Introduction
The Last Glacial Maximum (LGM) was a cold period approximately 21 000 yr before present (21 ka) when massive ice sheets covered much of the Northern Hemisphere (NH) continents. Paleoclimate proxy records provide valuable information on the climate forcings and climate state at the Last Glacial Maximum, but they are limited in spatial and temporal resolution and can be difficult to interpret, especially if one wishes to deduce circulation features such as flow fields or energy fluxes. Climate models have proved to be increasingly useful as a tool for tackling some of the questions that cannot be answered using proxy data alone. Of interest in this study is the large-scale circulation of the atmosphere during glacial times and how it compares to the circulation observed in the present-day climate. In particular, we direct our attention toward the North Atlantic sector, a region of deep water formation that experienced large, abrupt climate events during the last glacial period.
In the 1990s, the Paleoclimate Model Intercomparison Project (PMIP) was undertaken to evaluate past climates using a collection of climate models. The first phase of the project (PMIP1) comprised, for the most part, uncoupled atmospheric general circulation models forced with LGM boundary conditions: a sea surface temperature (SST) and sea ice cover reconstruction from the Climate: Long-range Investigation, Mapping, and Prediction (CLIMAP) project (CLIMAP 1981) and the ICE-4G land ice reconstruction (Peltier 1994). The resulting simulations offered a glimpse into how atmospheric circulation may have been during LGM. Subsequent studies focused on describing and understanding specific features such as atmospheric heat transport, transient activity, and storm tracks (Hall et al. 1996; Kageyama et al. 1999; Kageyama and Valdes 2000). Their findings have been used to endorse the idea that the glacial world was a stormier world because of stronger equator-to-pole temperature gradients and enhanced production of baroclinic eddies. Revisiting the original studies, however, discloses some often-overlooked details.
Synthesizing results from all the European models in PMIP1, Kageyama et al. (1999) reported a northeastward extension of both Northern Hemisphere storm tracks across the board and an elongation of the Atlantic storm track in six of the seven models. They note “no systematic increase or decrease in the storminess from the present climate to the last glacial maximum,” although the figures in the paper do seem to indicate an increase in peak low-level transient eddy activity for most of the models (see Kageyama et al. 1999, their Figs. 1, 2, 6, and 7). In the specific case of the U.K. Universities’ Global Atmospheric Modelling Programme (UGAMP) prescribed SST simulation, the increase in Atlantic eddy activity is limited to low levels of the atmosphere and is associated with enhanced low-level baroclinicity, thus suggesting the presence of stronger but shallower synoptic waves (Hall et al. 1996). Furthermore, it was found that changes in the normal modes largely account for changes in the position and dominant wavenumber of the storms but not the amplitude of the actual storm tracks (Kageyama and Valdes 2000).
Included in the analysis of Kageyama et al. (1999) are simulations from select models that were run with a slab ocean, a configuration that allowed SSTs and sea ice to be computed based on their thermal response to the climate forcings. Results from these simulations are somewhat equivocal because the oceanic heat flux that must be specified for a slab ocean was set to present-day values, thereby determining, in large part, the SST and sea ice distributions as well as the location of the storm tracks. However, among slab ocean simulations, there is some support for increased storminess at low to middle levels of the atmosphere during LGM (Dong and Valdes 1998). Going one step further, a study involving an intermediate complexity model with a simple dynamical ocean model reports similar findings (Justino et al. 2005).
Mounting paleoclimate evidence has since made it clear that both the CLIMAP SST reconstruction and the ICE-4G land ice reconstruction used in PMIP1 are flawed. In view of such developments, the question of how jets and storminess changed during LGM is by no means settled. There are now simulations of the LGM by more complex, fully coupled climate models with updated climate forcings based on the paleoclimate record, some of which fall within the framework of the second phase of PMIP (PMIP2). The majority of the literature documenting these simulations is concerned with issues such as atmosphere–ocean processes in the tropics and subtropics, in which the CLIMAP reconstruction is known to have problems (Bush and Philander 1998; Broccoli 2000; Kitoh and Murakami 2002; Timmermann et al. 2004); the general features of LGM climate (Kitoh and Murakami 2001; Kim et al. 2003; Shin et al. 2003; Otto-Bliesner et al. 2006); and the role of ocean dynamics in the maintenance of this climate state (Dong and Valdes 1998; Hewitt et al. 2003). Initial PMIP2 intercomparison studies have found an improvement in many aspects of the new LGM simulations (Kageyama et al. 2006; Braconnot et al. 2007), however, considerable problems remain in reproducing the peak glacial–interglacial temperature change from the Greenland ice cores (Masson-Delmotte et al. 2006) and in simulating changes in the ocean thermohaline circulation (Otto-Bliesner et al. 2007).
In this study, we examine the atmospheric circulation of the LGM in a PMIP2 simulation from a state-of-the-art, fully coupled climate model, the Community Climate System Model, version 3 (CCSM3), that was developed at the National Center for Atmospheric Research (NCAR). The simulation was set up and performed by Otto-Bliesner et al. (2006) to contrast preindustrial, mid-Holocene, and LGM climates. One major objective of the present study is to characterize the mean state and variability of atmospheric circulation in the Atlantic sector during the LGM by examining the jet and transient eddies, with an emphasis on identifying differences between the LGM and present-day climates.
Interestingly, when the CCSM3 is forced with PMIP2 boundary conditions, the simulated LGM climate shows a strong, steady Atlantic jet and enhanced low-level baroclinicity, but diminished wintertime eddy activity at all levels of the atmosphere compared to the present day. These results appear to contradict the atmosphere-only simulations from PMIP1, and furthermore, they suggest the existence of an altered atmospheric circulation regime during the LGM compared to present day (PD). As an initial illustration of the differences, Fig. 1 shows contours of total horizontal wind speed at 200 mb for PD and LGM simulations from a selection of the higher-resolution PMIP1 models (details in Table 1) and from the CCSM3 coupled model. For each uncoupled model (Figs. 1c–h), the LGM simulation (forced with PMIP1 boundary conditions) produces an Atlantic jet with strength and orientation (southwest–northeast tilt) comparable to the present-day jet (both simulated and observed) but with a northeastward extension. This is in stark contrast to the CCSM3 LGM simulation (Fig. 1j), which features a strong Atlantic jet with a more zonal orientation. Two of the PMIP1 slab ocean simulations are also shown (Figs. 1a,b). Although there are difficulties in interpreting these simulations, they do serve to emphasize the qualitative difference between the CCSM3 LGM simulation and all the LGM simulations that use PMIP1 boundary conditions.
The paper is organized as follows. A brief description of the model and methods are contained in section 2. Section 3 presents the model results indicating that, during LGM, the atmosphere exhibited a strong, steady mean circulation with decreased eddy activity. Section 4 discusses some possible mechanisms for the suppression of eddy activity. Section 5 summarizes evidence, based on paleoclimate observations, that the CCSM3 simulation produces a more realistic representation of LGM climate than the PMIP1 simulations; the focus is on the roles of land ice forcing and the SST/sea ice distributions. Finally, the main results of this work are summarized in section 6.
2. Model description and methods
We investigate changes between PD and LGM climates as simulated by the CCSM3 (Collins et al. 2006a), a global coupled atmosphere–ocean–sea ice–land surface climate model developed at NCAR. The setup of and results from these model simulations are documented in detail in Collins et al. (2006a) and Otto-Bliesner et al. (2006). Briefly, the CCSM3 comprises the primitive equation Community Atmosphere Model, version 3 (CAM3) at T42 horizontal resolution with 26 hybrid coordinate vertical levels (Collins et al. 2006b); a land model with land cover and plant functional types, prognostic soil and snow temperature, and a river routing scheme (Dickinson et al. 2006); the NCAR implementation of the Parallel Ocean Program (POP) on a 320 × 384 dipole grid (nominal horizontal resolution of 1°) with 40 vertical levels (Smith and Gent 2002); and a dynamic–thermodynamic sea ice model on the same grid as the ocean model (Briegleb et al. 2004).
The forcings for the LGM simulation are in accordance with the protocols established by PMIP2 (http://pmip2.lsce.ipsl.fr): 21 ka insolation, atmospheric greenhouse gas concentrations based on ice core measurements (Flückiger et al. 1999; Dällenbach et al. 2000; Monnin et al. 2001), atmospheric aerosols at preindustrial values, and land ice and coastlines corresponding to the 120-m sea level depression from the ICE-5G reconstruction (Peltier 2004).
We use monthly-mean outputs from 50 yr of the simulations. Daily output for transient eddy analyses are taken from 25-yr branch runs. A bootstrap method was used to determine that 25-yr samples are adequate for stable eddy statistics. Using wintertime [December–March (DJFM)] Northern Hemisphere 850-mb 
3. Last Glacial Maximum climate in the coupled model
Of interest here is the large-scale atmospheric circulation in the model’s simulations of the LGM and PD climates. We will focus on the Northern Hemisphere Atlantic sector, in which differences in forcings between the LGM and PD, and consequently circulation features, are most dramatic. All differences between the LGM and PD climates discussed here are significant at the 95% confidence level unless otherwise noted.
a. Circulation and heat transport
Upper-level zonal wind and geopotential height provide a useful broad-brush picture of large-scale flow characteristics of the atmosphere. Figure 2 shows wintertime maps of these two fields for the LGM and PD simulations. Under the LGM forcings described in the previous section, we observe an enhanced stationary wave associated with the Laurentide ice sheet covering most of North America and a stronger, more zonal jet in the Atlantic sector downstream of the ice sheet. The stronger winds during LGM are consistent with the stronger equator-to-pole surface temperature gradient seen in Fig. 3. Changes in the Pacific are more subtle, showing a slight equatorward shift of the jet (Fig. 3) and the development of a split flow over Siberia, which is evident from the region of weak easterlies outlined by the thick zero contour in Fig. 2b.
The change in atmospheric circulation over the Atlantic sector during LGM is particularly striking. There is an inverse relationship between maximum jet strength and jet width in January for the PD simulation (red circles in Fig. 4), in which jet strength uATL is the maximum zonal wind in the sector and jet width is the latitude range over which uATL decreases to half its maximum value. The relationship is not evident in the LGM climate (blue triangles), which furthermore inhabits a sector of width–strength space separate from the PD climate. The LGM jet is 20% stronger on average, with half the variability in maximum speeds of the PD jet; it is also 30% narrower, with almost 6 times less variability in width. These results are robust to the choice of month(s) used to define winter and the exact longitude range used to define the Atlantic jet.
We find that the amount of heat transported toward the poles is remarkably similar in the PD and LGM simulations. Between the two model simulations, the LGM does, indeed, have slightly more transport, with the atmosphere helping to increase the peak NH value to 6.33 ± 0.01 PW, about 0.3 PW greater than in PD. The robustness of the total heat transport curve found across the CCSM3’s simulations of PD and LGM climates is corroborated by other coupled climate models (Hewitt et al. 2003; Shin et al. 2003) and from uncoupled general circulation models using a prescribed sea surface temperature forcing (Hall et al. 1996), with the magnitude of the LGM increase varying from 0.2 to 1.5 PW for peak NH values.
All else being equal, a back-of-the-envelope calculation tells us that this relatively modest 0.3 PW boost in heat transport at 35°N translates to a 3 W m−2 boost in heating rate north of this latitude circle. Assuming a midrange climate sensitivity of 0.5°C (W m−2)−1, this is equivalent to a 1.5°C warming of the mid- to high-latitude regions. Clearly, all else is not equal and the actual surface temperature difference between the two climates poleward of 35°N is closer to 10°C. Note that technically, the 0.5°C (W m−2)−1 value is a global climate sensitivity and should not be used to estimate the response of the polar cap to increased heat flux from the lower latitudes. However, the polar cap, as defined in our calculation (35°N to the North Pole), is a large region over which there is a net TOA radiation loss to space. Furthermore, the use of this global climate sensitivity to estimate the contribution of heat transport changes to the simulated temperature change is supported by an experiment in Seager et al. (2002). In this experiment, ocean heat transport was turned off, and this lead to a 1.3 PW reduction in energy that moved across 35°N, corresponding to a 13 W m−2 decrease in heating rate north of this latitude and a 6°C cooling of the mid- to high-latitude regions.
Upon closer inspection, there are interesting differences in the partitioning of these heat fluxes in the PD and LGM simulations. Overall, the net (atmosphere plus ocean) transport in the NH during LGM is very similar to that in the PD. The total AT in the LGM is also comparable to that in the PD, but there is a large difference in the balance of processes responsible for this AT. The pronounced ridging forced by the Laurentide ice sheet enhances the dry stationary wave heat transport contribution 
b. Transient eddy activity
Concentrating on boreal winter in the NH, the eddy fields reveal an LGM climate that is more quiescent than the PD (Fig. 6). In the Atlantic sector, the reduction in eddy activity from PD to LGM is observed at both low levels (15% decrease in sector-averaged υ′T ′) and upper levels [30% decrease in sector-averaged (u′2 + υ′2)/2]. Compared to PD, the LGM jets are strong and narrow (Fig. 3), and the eddy fluxes occupy a narrower latitudinal band hugging the axis of the jet core rather than a broad band perched on the poleward flank of the jet (e.g., compare Figs. 6b and 6c). The differences between the two climates can also be seen in Fig. 4, in which January poleward heat fluxes in the LGM simulation (blue triangles) span a narrow range of smaller values than in the PD simulation (red circles). Although we will not discuss the Pacific sector, it, too, exhibits changes in jet structure and eddy fluxes.
The measures of atmospheric flow in Table 2 provide another way to compare the Atlantic sector in today’s climate and glacial climates. The weak eddy activity in the LGM simulation coexists with a stronger, narrower Atlantic jet, as shown in Fig. 4. From this plot, we inferred a less variable upper-level flow field during LGM compared to PD. The winter season flow metrics provide additional and more direct ways to evaluate the steadiness of the LGM jet. From Table 2, we see a 35% decrease in monthly departures of kinetic energy from its climatological mean state (KEIATL in column three). Together, these results are consistent with the picture of global heat transport in Fig. 5, in which a slight overall increase in meridional energy flux during LGM is achieved by a greatly enhanced stationary wave, with the implication that the contribution from transient eddies must be weaker.
As a final remark on this topic, we note that the gross structure of the Atlantic jet and eddies described here for the PD simulation is consistent with observational data from the National Centers for Environmental Prediction (NCEP)–NCAR reanalysis (Kalnay et al. 1996) and the 40-yr European Centre for Medium-Range Weather Forecasts (ECMWF) Re-Analysis (ERA-40; Uppala et al. 2005). Both NCEP (not shown) and ERA-40 (Figs. 6a,d,g) show a broad wintertime jet with a southwest–northeast tilt and eddy activity peaking poleward of the jet, just as we have seen for the PD simulation (Figs. 6 b,e,h). One feature that does not compare well is the absolute jet strength, which the model overestimates in each winter month (as seen for January in Fig. 4) such that the winter season average of uATL is 20% too strong (Table 2). However, the variability in the large-scale flow field in the PD simulation is comparable to that in the two reanalyses products; furthermore, the PD variability is substantially different from that in the LGM simulation. For example, uATL spans a range of approximately 25 m s−1 in both the observations and in 50 yr of the PD simulation (Fig. 4), whereas the range in jet strengths in the LGM simulation is approximately half this value. Also, the variability as measured by the poleward heat flux υTATL and kinetic energy KEIATL indices (Table 2) appear to show characteristic values for the PD climate (3.42–3.56 K m s−1, 31–34 m2 s−2) that are distinct from those of the LGM climate (2.93 K m s−1, 22 m2 s−2).
c. Baroclinicity of the atmosphere
For each climate, regions of high Eady growth rate correspond in general to regions of enhanced eddy activity; across different climate states, however, Eady growth rates increase in regions where eddy amplitudes decrease. For example, there is a general increase in midlatitude growth rates during LGM relative to PD, especially over the Atlantic, which suggests that the glacial climate should be stormier than today’s climate. That the eddy diagnostics indicate decreased eddy activity during LGM despite more baroclinic conditions (larger σ) points to the complicated nature of the relationship between transient activity and the mean flow.
4. Suppression of eddy activity
The LGM climate as simulated by the CCSM3 coupled model exhibits stronger jets but weaker storms. This quiescent glacial climate is counter to the intuitive idea that a more vigorously circulating atmosphere should produce a more tempestuous world. It is also counter to more theoretically based indicators such as the stronger meridional temperature gradient and the increased low-level baroclinicity. However, a range of factors not captured by the Eady growth rate parameter also affect the life cycle of eddies. These factors range from diabatic heating (Hoskins and Valdes 1990) and seeding effects (Zurita-Gotor and Chang 2005) to “governors” that limit the ability of eddies to tap into the available baroclinicity (James 1987; Lee and Kim 2003; Harnik and Chang 2004). In this section, we discuss whether such factors may play a role in reducing the eddy activity in the coupled model LGM simulation. A prerequisite to this discussion is an assessment of the structure of the eddies.
a. Eddy structure
We examine regression structures at 200 and 850 mb in the CCSM3 simulations of PD and LGM climates (Fig. 8). The reference time series against which variables are linearly regressed is the 200-mb meridional wind field υ′200 at 43°N, 28°W for the PD climate and at 38°N, 28°W for the LGM climate. These points were selected based on the location of maximum variance in the DJFM υ′200 field. We normalize the b(i) coefficients such that the regression maps show amplitude, expressed in the units of the variable of interest, per standard deviation of the reference time series. Structures calculated using reference time series at different locations, reference time series based on different fields, and alternate definitions of the winter season are similar to these and are not shown here.
The eddy structure in both the PD and LGM climates resemble localized wave trains. The LGM eddy amplitudes peak just south of 40°N, the latitude marked by the solid gray line in Fig. 8, whereas the PD eddy amplitudes peak just north of this latitude. Although meridional wind variations υ′ are much stronger aloft (Figs. 8a,d) than at low levels (Figs. 8b,e) in both simulations (note the different contour intervals), there is a greater discrepancy in the LGM. Thus, eddies in glacial times appear to be trapped closer to the tropopause, with weaker low-level disturbance amplitudes. Furthermore, though the LGM has slightly weaker eddy amplitudes, the eddy structures are correlated over greater distances and are more persistent, as seen in lag regression analyses, in which the time series yt(i) is shifted by a certain number of days relative to xt (not shown). The larger spatial correlations are, in fact, consistent with the eastward extension of storm tracks seen in the PMIP1 simulations of Kageyama et al. (1999), however, the location and strength of the storm tracks themselves are not. These results suggest that differences in eddy structure may be an important part of why the LGM simulation exhibits such weak eddy activity compared to the PD simulation. There is a large body of theoretical work on why eddy activity might be suppressed in conditions that appear conducive to eddy growth. Next, we examine the relevance of these hypotheses for the reduction of eddy activity in the LGM.
b. Discussion
One established idea in the literature is the barotropic governor mechanism, in which horizontal shear is thought to limit baroclinic instability and hence storm growth (James 1987). The sharpening of the Atlantic jet (Fig. 3) makes this an attractive candidate, but some telltale signs of the barotropic governor are absent. If the barotropic governor were in operation, we would expect the eddies to bend and elongate into a boomerang or banana shape because they are deformed by the strongly sheared mean flow; this characteristic shape is not evident in the LGM simulation (Fig. 8).
Another explanation for the suppression of eddy activity is inspired by the existence of a modern analog to the strong jet–weak storms dichotomy, namely the Pacific midwinter suppression of storms (Nakamura 1992) that occurs every January when the jet is at its strongest (Fig. 9). According to Yin (2002), this phenomenon has to do with the southward migration of tropical convection in the western Pacific and the related strengthening of the subtropical jet during boreal winter. Yin (2002) diagnosed that these changes increase upper-level static stability and lower the tropopause height in a manner that discourages the growth of storms. In the modern climate, the climatological latitude of the Atlantic jet is 2°–3° north of the climatological latitude of the Pacific jet in winter, perhaps positioning it too far poleward to be affected by changes in tropical heating. In the LGM simulation, however, the more zonal orientation of the Atlantic jet pulls it toward the equator, in closer proximity to the tropical convection sites. Preliminary inspection reveals little difference in tropical heating or precipitation between the LGM and PD simulations in the Atlantic sector. Nonetheless, subtle changes in the tropics can have a large impact, so a more thorough investigation should be carried out before this mechanism for jet stability can be rejected.
In a related vein, Lee and Kim (2003) used idealized numerical experiments to demonstrate that flow regimes with strong (weak) subtropical jets exhibit weaker (stronger) eddy activity located just (far) poleward of the jet core. In a series of experiments, they varied the strength of the tropical convective heating and the high latitude cooling to strengthen or weaken the subtropical jet (Lee and Kim 2003; Son and Lee 2005). They showed that the strong subtropical jet regime supported more isotropic eddies and weaker barotropic conversion than the weak subtropical jet regime. The LGM simulation exhibits several features in common with their strong subtropical jet case, including relatively weak eddy fluxes located at roughly the same latitude as the tropospheric jet maximum (Fig. 6) and a slightly more isotropic eddy structure (Fig. 8). Rather than requiring changes in tropical heating, this idea allows for the possibility of a strong jet regime to be set up some other way, for example, by topographic effects from the Laurentide or additional high-latitude cooling because of extensive sea ice cover in the North Atlantic.
The Lorenz energy cycle (Lorenz 1955) provides further insight into why the eddy activity in the LGM simulation is suppressed compared to the PD simulation. The terms in the energy budget equation describe how energy is exchanged between the mean flow and the transient eddies as the eddy disturbances grow and decay. Different suppression mechanisms affect different parts of the eddy life cycle, and thus should have different signatures on these budget terms. Preliminary results (A. Donohoe 2007, personal communication) indicate that the LGM Atlantic mean state does indeed support faster eddy growth because of its sharper temperature gradients, despite competing influences from the increased static stability and the strong barotropic shear of the jet. However, the baroclinic conversion of potential energy into eddy kinetic energy is greatly reduced in the LGM compared to PD. The explanation for the reduced eddy activity appears to be related to the fact that there is substantially less seeding of the Atlantic baroclinic zone by upstream disturbances in the LGM, and that the disturbances that do enter the baroclinic zone are weak. The cause of the reduced seeding is under continuing investigation.
5. Evaluation of the CCSM3 simulation
The PMIP1 simulations of LGM climate share certain atmospheric flow characteristics that are very similar to those in the present-day climate, whereas the coupled model simulation of LGM exhibits fundamental differences when compared to the present-day climate. Specifically, in the PMIP1 models, the amount of storminess in the Atlantic sector during LGM is comparable to that in the PD climate, with slight increases in peak low-level transient eddy activity during the LGM (Kageyama et al. 1999). Furthermore, in the PMIP1 LGM simulations, the Atlantic jet shows the same northeastward extension as the Atlantic storm track with little change in strength (Fig. 1). Conversely, the CCSM3 LGM simulation exhibits a strong, steady jet with zonal orientation and weaker transient eddies at all levels of the atmosphere (Fig. 6).
We have reproduced the PMIP1 results with CAM3, the atmospheric component of the CCSM3, to illustrate more clearly the differences between the circulation simulated by the coupled model and the generic result obtained by the suite of uncoupled PMIP1 models (Fig. 10). The atmosphere model has now been forced with SST, sea ice and land ice prescribed exactly as specified in the PMIP1 experimental setup. The resulting simulation is the same that appears in Fig. 1h and hereafter will be referred to as PMIPa, where “a” means “atmosphere only.” Compared to the coupled model LGM simulation (Fig. 10c,d), PMIPa produces a stormier Atlantic sector and a weaker jet that exhibits a marked southwest-to-northeast tilt (Fig. 10e,f). Indeed, the eddy heat transport and jet orientation in PMIPa bear a closer resemblance to the coupled PD simulation (Fig. 10a,b) than to the coupled LGM simulation.
The source of the differences between the LGM climate simulated by the uncoupled models and the LGM climate simulated by the coupled model lies at the earth’s surface. Recall that the uncoupled PMIP1 models experience or “see” one set of prescribed land ice and SST/sea ice boundary conditions. The coupled model experiences a somewhat different set of land ice boundary conditions, those of the improved ICE-5G (Peltier 2004) reconstruction. In addition, the coupled model has interactive ocean and sea ice components that calculate their own SST/sea ice fields. The fact that CCSM3’s atmosphere model achieves a PMIP1-like LGM climate when run in the PMIP1 configuration is further support that these land and sea surface forcings, rather than a disparity in models or atmosphere model physics, are, indeed, key to creating the different LGM climates (Fig. 1h).
The question becomes one of evaluating which simulated LGM climate, the uncoupled PMIP1-like LGM or the coupled CCSM3-like LGM, is a more realistic representation of the actual climate of the Last Glacial Maximum. Detailed comparisons have been performed between PMIP1 model simulations and a selection of the fully coupled model simulations, including CCSM3, that participated in PMIP2. These studies conclude that the fully coupled PMIP2 simulations using ICE-5G land ice are in better agreement with the paleoclimate observations currently available, particularly in terms of surface air temperature, SST, and tropical climate (Braconnot et al. 2007), as well as regional temperatures over northern Eurasia and the North Atlantic (Kageyama et al. 2006). These intercomparisons include a number of partially coupled simulations from PMIP1 in which SSTs and sea ice were calculated by a slab ocean model. These will not be included in our discussion because their use of present-day oceanic heat transport in the slab ocean heavily influences the resulting SST patterns.
The following sections will point out certain aspects of the boundary conditions and model results where the CCSM3 simulation shows marked improvements, based on the paleoclimate record, over the PMIP1 simulations. (For a comprehensive treatment of this topic, the reader is referred to the intercomparison studies mentioned in this paper.) Also discussed are the potential consequences of these features for atmospheric dynamics in general and for the Atlantic jet and storm track in particular.
a. Land ice
The ICE-4G (Peltier 1994) and ICE-5G (Peltier 2004) reconstructions of deglaciation history are based in part on the theory of glacial isostatic adjustment, a process by which the external surface load of the continental ice sheets is compensated by changes in the surface of the earth. ICE-5G is widely considered to be an improvement over ICE-4G: its land ice reconstruction accounts for new refinements in the reconstruction of global sea level, and it also corrects many of the regional shortcomings of ICE-4G (see Peltier 2004, their references). As a result, at the time of the LGM, ICE-5G has more land-based ice than did ICE-4G (approximately 12 m sea level equivalent). In ICE-5G, there is much less land ice on Greenland and the Eurasian continent, whereas the Laurentide ice sheet complex over North America is significantly larger in volume compared to ICE-4G (15 m sea level equivalent, or 25% of the Laurentide ice volume), with the bulk of the extra ice forming the Keewatin Dome west of Hudson Bay (Fig. 11). The addition of this dome reconciles the original ICE-4G results with more recent observations, namely the large crustal uplift rates near Yellowknife, Canada, (Argus et al. 1999) and gravity measurements south and west of Hudson Bay (Lambert et al. 2001).
A higher Laurentide ice sheet forces a stronger stationary wave pattern in the Northern Hemisphere (Otto-Bliesner et al. 2006), and we expect changes in surface temperatures and in the interannual variability and seasonal cycle of surface temperature. In general, these features are better simulated in the PMIP2 runs using ICE-5G than in the PMIP1 runs using ICE-4G, with the exception of western European winter temperatures, which are still underestimated compared to pollen data (Kageyama et al. 2006). The altered stationary wave and the physical barrier of this larger ice sheet may also contribute to reduced seeding of the Atlantic jet (section 4b), and hence explain both the strengthening and stabilization observed in the LGM simulation.
b. SST and sea ice
The CCSM3 coupled model provides a self-consistent simulation of the Last Glacial Maximum climate. It produces a climatological SST distribution and sea ice coverage that differ from those determined in CLIMAP and used as boundary conditions in PMIP1 (CLIMAP 1981). At first glance, one might expect the observation-based CLIMAP dataset to provide a more trustworthy picture of the glacial ocean than a model. CLIMAP produced the first SST maps of the glacial ocean by using transfer functions to translate population distributions of fossil plankton species found in ice age marine sediments (CLIMAP 1981) into sea surface temperatures. The transfer functions (Imbrie and Kipp 1971) were derived from knowledge of how these same plankton species are distributed in today’s ocean. In the intervening decades, however, the proliferation of sediment core data and the advent of new statistical and geochemical methodologies for reconstructing past SSTs have challenged some of the assumptions and results of the CLIMAP method. Moreover, modeling studies have shown that the consequences of such SST errors on the atmospheric circulation are potentially dramatic, whether the errors themselves are in the tropics or in high latitudes (Yin and Battisti 2001; Toracinta et al. 2003; Hostetler et al. 2006).
Though the data still fall short of providing definitive SST and sea ice distributions in the glacial ocean, in regions where there is growing consensus, the CCSM3 simulation shows a clear improvement over the CLIMAP reconstruction. Among the more problematic areas of CLIMAP are the lack of cooling in tropical and subtropical regions, and the extensive sea ice in the North Atlantic (Fig. 12a). These features are very different in the CCSM3 simulation (Fig. 12b), and in fact in the coupled PMIP2 simulations in general (Kageyama et al. 2006; Braconnot et al. 2007, and references therein).
An intercomparison of eight coupled model simulations of the LGM shows a 1°–2°C annual mean cooling over the global tropics, which is in agreement with proxy data (Braconnot et al. 2007, their references). But more relevant than changes in the absolute temperature field in the tropics and subtropics are changes in the distribution of SSTs and SST gradients in the tropics. Evaluating whether CLIMAP or the CCSM3 simulation does a better job in this respect is difficult given the uncertainties in the SST reconstructions and the differences between them. However, several interesting features are robust enough in the simulation and the records to merit mention. At the Last Glacial Maximum, the Pacific warm pool is contracted meridionally in the CCSM3 simulation (Fig. 12), a feature that is consistent with the foraminifera census data (Trend-Staid and Prell 2002; Kucera et al. 2005); in contrast, CLIMAP has a warm pool that is expanded meridionally compared to the present-day climate. Several studies point to stronger zonal gradients across the tropical Pacific (Lea et al. 2000; Trend-Staid and Prell 2002; Kucera et al. 2005), but the reconstructed SST patterns themselves are quite dissimilar. Nevertheless, the CCSM3 simulation does exhibit stronger zonal and meridional tropical SST gradients compared to CLIMAP. In the northern subtropics, the SST simulated by the CCSM3 is more similar to the reconstruction of Trend-Staid and Prell (2002) than is the CLIMAP product, including more zonal isotherms in the regions of the Kuroshio and Gulf Stream current.
Another region in which there are significant differences between the CLIMAP reconstruction and the CCSM3-simulated SSTs is the North Atlantic (Fig. 13). Compared to CLIMAP, the simulated winter SSTs are 3°–5°C warmer along the west coast of Norway and 1°–3°C warmer in a broad swath of the Atlantic Ocean just south of the Greenland–Scotland ridge. These areas remain ice free in winter in CCSM3 and in other PMIP2 simulations (Kageyama et al. 2006), whereas in CLIMAP, perennial sea ice cover extends as far south as 50°N. The presence of open ocean and relatively high temperatures in portions of the eastern North Atlantic and, at least seasonally, in the Nordic Seas is supported by a large body of proxy data, much of which has been synthesized by the Multiproxy Approach for the Reconstruction of the Glacial Ocean (MARGO) project (additional information is available at http://margo.pangaea.de/). These data include estimates of SST and sea ice extent based on coccoliths (Hebbeln et al. 1994), biomarker pigments (Rosell-Melé and Koç 1997), alkenones (Rosell-Melé and Comes 1999), dinoflagellates (de Vernal and Hillaire-Marcel 2000), and foraminifera (Pflaumann et al. 2003; Sarnthein et al. 2003; Meland et al. 2005). The simulated North Atlantic conditions in CCSM3 are further corroborated by other coupled atmosphere–ocean models (Hewitt et al. 2003; Shin et al. 2003) and by inferred North Atlantic circulation patterns from foraminiferal assemblages (Lassen et al. 1999).
CLIMAP has very extensive sea ice in the Atlantic, with the winter ice edge reaching at least 45°N over the entire basin (Fig. 12). Consequently, as the storms track northeastward into the Nordic Seas in the PMIP1 simulations, they have access to the strong temperature gradients at the sea ice edge. In the CCSM3 coupled simulation of the LGM, sea ice in the western Atlantic spreads to similarly southern latitudes in winter, but in the eastern Atlantic the ice edge never ventures past 60°N (Fig. 12). With the sea ice so far north, the strongest midlatitude temperature gradient in the eastern Atlantic is now associated with the band of rapidly changing SSTs around 35°–40°N (see Kageyama et al. 2006; Fig. 12). The location and orientation of the jet and storm track in the coupled simulation are consistent with this SST gradient, which clearly points to its importance in determining the state of the Atlantic atmospheric circulation. However, the SSTs are most likely not the whole story. The size of the Laurentide ice sheet upstream seems to exert a large influence on the location of the North Atlantic storm track and jet and on the SST distribution in this region (Li 2007).
6. Concluding remarks
We have presented evidence for an altered atmospheric circulation regime in the Atlantic sector during the Last Glacial Maximum from a global coupled climate model. This LGM Atlantic circulation is characterized by a stronger, more zonally oriented Atlantic jet and yet reduced storminess. In other words, the coupled model produces a glacial climate that is quiescent compared to today’s climate. Analysis of the eddy structures suggests that the barotropic governor is unlikely to play a role in suppressing the storms. Reduced seeding of the Atlantic jet is currently under investigation as the most likely cause of the decrease in storminess during LGM.
The LGM climate simulated by the coupled model is remarkably different from the LGM climate simulated by atmosphere models forced with PMIP1 boundary conditions. Whereas the coupled simulation exhibits a strong, zonal jet with reduced eddies, the uncoupled simulations feature a North Atlantic circulation regime that is similar to the one in today’s climate. We have found that this discrepancy can be traced to differences in the ice sheet topography and sea surface conditions and not to differences in the atmosphere models and model physics. The LGM land ice was updated from the ICE-4G reconstruction in PMIP1 to the improved ICE-5G reconstruction in the coupled simulation. In addition, the SST and sea ice distributions seen by the atmosphere were prescribed to CLIMAP in PMIP1 but calculated by sophisticated ocean and sea ice models in CCSM3. The SST and sea ice distributions simulated by the coupled model are more consistent with the proxy data than are the prescribed SST and sea ice boundary conditions used in PMIP1. Thus, to the extent that changes in atmospheric circulation are linked to changes in orographic forcing and sea surface conditions, the CCSM3 coupled model simulation can be regarded as a better estimate of the LGM climate.
The existence of a different Atlantic atmospheric circulation regime during LGM has implications for our understanding of global heat transport and the stability of glacial climates. In particular, it can offer new outlooks on modes of climate variability that appear to be unique to glacial times, such as the abrupt Dansgaard–Oeschger warming events recorded in Greenland’s ice cores.
Acknowledgments
We wish to thank Joe Barsugli, Kerim Nisancioglu, Richard Seager, Mike Wallace, Justin Wettstein, and Jeff Yin for stimulating and constructive discussions, and Cecilia Bitz and Marc Michelson for assisting with the model simulations. We are also grateful to Michel Crucifix and two anonymous reviewers for their thoughtful comments. This research was supported by the Comer Fellowship Program. Data for this paper were accessed from http://pmip.lsce.ipsl.fr/, http://data.ecmwf.int/data/, and http://www.cdc.noaa.gov/cdc/data.ncep.reanalysis.html.
REFERENCES
Argus, D., W. Peltier, and M. Watkins, 1999: Glacial isostatic adjustment observed using very long baseline interferometry and satellite laser ranging geodesy. J. Geophys. Res., 104 , B12. 29077–29094.
Braconnot, P., and Coauthors, 2007: Results of PMIP2 coupled simulations of the Mid-Holocene and Last Glacial Maximum—Part 1: Experiments and large-scale features. Climate Past, 3 , 261–277.
Briegleb, B., C. Bitz, E. Hunke, W. Lipscomb, M. Holland, J. Schramm, and R. Moritz, 2004: Scientific description of the sea ice component in the Community Climate System Model Version 3. NCAR Tech. Note NCAR/TN-463+STR, 70 pp.
Broccoli, A. J., 2000: Tropical cooling at the Last Glacial Maximum: An atmosphere–mixed layer model simulation. Nature, 13 , 951–976.
Bush, A. B. G., and S. G. H. Philander, 1998: The role of ocean–atmosphere interactions in tropical cooling during the Last Glacial Maximum. Science, 279 , 1341–1344.
Chang, E. K. M., 1993: Downstream development of baroclinic waves as inferred from regression analysis. J. Atmos. Sci., 50 , 2038–2053.
CLIMAP, 1981: Seasonal Reconstructions of the Earth’s Surface at the Last Glacial Maximum. Geological Society of America Map and Chart Series, No. 36, 18 pp.
Collins, W. D., and Coauthors, 2006a: The Community Climate System Model version 3 (CCSM3). J. Climate, 19 , 2122–2143.
Collins, W. D., and Coauthors, 2006b: The formulation and atmospheric simulation of the Community Atmosphere Model version 3 (CAM3). J. Climate, 19 , 2144–2161.
Dällenbach, A., T. Blunier, J. Flückiger, B. Stauffer, J. Chappellaz, and D. Raynaud, 2000: Changes in the atmospheric CH4 gradient between Greenland and Antarctica during the Last Glacial and the transition to the Holocene. Geophys. Res. Lett., 27 , 1005–1008.
de Vernal, A., and C. Hillaire-Marcel, 2000: Sea-ice cover, sea-surface salinity and halo/thermocline structure of the northwest North Atlantic: Modern versus full glacial conditions. Quat. Sci. Rev., 19 , 65–85.
Dickinson, R. E., K. W. Oleson, G. B. Bonan, F. Hoffman, P. Thornton, M. Vertenstein, Z-L. Yang, and X. Zeng, 2006: The Community Land Model and its climate statistics as a component of the Community Climate System Model. J. Climate, 19 , 2302–2324.
Dong, B., and P. Valdes, 1998: Simulations of the Last Glacial Maximum climates using a general circulation model: Prescribed versus computed sea surface temperatures. Climate Dyn., 14 , 571–591.
Eady, E., 1949: Long waves and cyclone waves. Tellus, 1 , 33–52.
Flückiger, J., A. Dällenbach, B. Stauffer, T. Stocker, D. Raynaud, and J-M. Barnola, 1999: Variations of the atmospheric N2O concentration during abrupt climatic changes. Science, 285 , 227–230.
Hall, N. M. J., B. Dong, and P. J. Valdes, 1996: Atmospheric equilibrium, instability and energy transport at the last glacial maximum. Climate Dyn., 12 , 497–511.
Harnik, N., and E. K. Chang, 2004: The effects of variations in jet width on the growth of baroclinic waves: Implications for midwinter Pacific storm track variability. J. Atmos. Sci., 61 , 23–40.
Hebbeln, D., T. Dokken, E. Andersen, M. Held, and A. Elverhoi, 1994: Moisture supply for northern ice-sheet growth during the Last Glacial Maximum. Nature, 370 , 357–360.
Held, I. M., and B. J. Soden, 2006: Robust responses of the hydrological cycle to global warming. J. Climate, 19 , 5686–5699.
Hewitt, C., R. Stouffer, A. Broccoli, J. Mitchell, and P. J. Valdes, 2003: The effect of ocean dynamics in a coupled GCM simulation of the Last Glacial Maximum. Climate Dyn., 20 , 203–218. doi:10.1007/s00382-002-0272-6.
Hoskins, B. J., and P. J. Valdes, 1990: On the existence of storm-tracks. J. Atmos. Sci., 47 , 1854–1864.
Hostetler, S., N. Pisias, and A. Mix, 2006: Sensitivity of Last Glacial Maximum climate to uncertainties in tropical and subtropical ocean temperatures. Quat. Sci. Rev., 25 , 1168–1185.
Imbrie, J., and N. Kipp, 1971: A new micropaleontological method for quantitative paleoclimatology: Application to late Pleistocene Carribean core. The Late Cenozoic Glacial Ages, K. Turekian, Ed., Yale University Press, 71–181.
James, I., 1987: Suppression of baroclinic instability in horizontally sheared flows. J. Atmos. Sci., 44 , 3710–3720.
Justino, F., A. Timmermann, U. Merkel, and E. P. Souza, 2005: Synoptic reorganization of atmospheric flow during the Last Glacial Maximum. J. Climate, 18 , 2826–2846.
Kageyama, M., and P. Valdes, 2000: Synoptic-scale perturbations in AGCM simulations of the present and Last Glacial Maximum climates. Climate Dyn., 16 , 517–533.
Kageyama, M., P. Valdes, G. Ramstein, C. Hewitt, and U. Wyputta, 1999: Northern Hemisphere storm tracks in present day and Last Glacial Maximum climate simulations: A comparison of the European PMIP models. J. Climate, 12 , 742–760.
Kageyama, M., and Coauthors, 2006: Last Glacial Maximum temperatures over the North Atlantic, Europe and western Siberia: A comparison between PMIP models, MARGO sea-surface temperatures and pollen-based reconstructions. Quat. Sci. Rev., 25 , 2082–2102.
Kalnay, E., and Coauthors, 1996: The NCEP/NCAR 40-Year Reanalysis Project. Bull. Amer. Meteor. Soc., 77 , 437–471.
Kim, J-H., R. R. Schneider, S. Mulitza, and P. J. Müller, 2003: Reconstruction of SE trade-wind intensity based on sea-surface temperature gradients in the Southeast Atlantic over the last 25 kyr. Geophys. Res. Lett., 30 .2144, doi:10.1029/2003GL017557.
Kitoh, A., and S. Murakami, 2001: A simulation of the Last Glacial Maximum with a coupled atmosphere–ocean GCM. Geophys. Res. Lett., 28 , 2221–2224.
Kitoh, A., and S. Murakami, 2002: Tropical Pacific climate at the mid-Holocene and the Last Glacial Maximum simulated by a coupled ocean–atmosphere general circulation model. Paleoceanography, 17 .1047, doi:10.1029/2001PA000724.
Kucera, M., and Coauthors, 2005: Reconstruction of sea-surface temperatures from assemblages of planktonic foraminifera: Multi-technique approach based on geographically constrained calibration sets and its application to glacial Atlantic and Pacific Oceans. Quat. Sci. Rev., 24 , 951–998. doi:10.1016/j.quascirev.2004.07.014.
Lambert, A., N. Courtier, G. Sasegawa, F. Klopping, D. Winester, T. James, and J. Liard, 2001: New constraints on Laurentide postglacial rebound from absolute gravity measurements. Geophys. Res. Lett., 28 , 2109–2112.
Lassen, S., E. Jansen, K. Knudsen, A. K. M. Kristensen, and K. Christensen, 1999: Northeast Atlantic sea surface circulation during the past 30–1014 C kyr B.P. Paleoceanography, 14 , 616–625. doi:10.1029/1999PA900029.
Lea, D. W., D. K. Pak, and H. Spero, 2000: Climate impact of late Quaternary equatorial Pacific sea surface temperature variations. Science, 289 , 1719–1724. doi:10.1126/science.289.5485.1719.
Lee, S., and H-k Kim, 2003: The dynamical relationship between subtropical and eddy-driven jets. J. Atmos. Sci., 60 , 1490–1503.
Li, C., 2007: A general circulation modelling perspective on abrupt climate change during glacial times. Ph.D. thesis, University of Washington, 102 pp.
Lim, G. H., and J. M. Wallace, 1991: Structure and evolution of baroclinic waves as inferred from regression analysis. J. Atmos. Sci., 48 , 1718–1732.
Lindzen, R., and B. Farrell, 1980: A simple approximate result for the maximum growth rate of baroclinic instabilities. J. Atmos. Sci., 37 , 1648–1654.
Lorenz, E., 1955: Available potential energy and the maintenance of the general circulation. Tellus, 7 , 157–167.
Masson-Delmotte, V., and Coauthors, 2006: Past and future polar amplification of climate change: Climate model intercomparisons and ice-core constraints. Climate Dyn., 26 , 513–529.
Meland, M. Y., E. Jansen, and H. Elderfield, 2005: Constraints on SST estimates for the northern North Atlantic/Nordic Seas during the LGM. Quat. Sci. Rev., 24 , 835–852. doi:10.1016/j.quascirev.2004.05.011.
Monnin, E., A. Indermühle, A. Dällenbach, J. Flückiger, B. Stauffer, T. Stocker, D. Raynaud, and J-M. Barnola, 2001: Atmospheric CO2 concentrations over the last glacial termination. Science, 291 , 112–114.
Nakamura, H., 1992: Midwinter suppression of baroclinic wave activity in the Pacific. J. Atmos. Sci., 49 , 1629–1642.
Otto-Bliesner, B. L., E. Brady, G. Clauzet, R. Tomas, S. Levis, and Z. Kothavala, 2006: Last Glacial Maximum and Holocene climate in CCSM3. J. Climate, 19 , 2526–2544.
Otto-Bliesner, B. L., C. D. Hewitt, T. M. Marchitto, E. Brady, A. Abe-Ouchi, M. Crucifix, S. Murakami, and S. L. Weber, 2007: Last Glacial Maximum ocean thermohaline circulation: PMIP2 model intercomparisons and data constraints. Geophys. Res. Lett., 34 .L12706, doi:10.1029/2007GL029475.
Peltier, W. R., 1994: Ice age paleotopography. Science, 265 , 195–201.
Peltier, W. R., 2004: Global glacial isostasy and the surface of the ice-age Earth. Annu. Rev. Earth Planet. Sci., 32 , 111–149.
Pflaumann, U., M. Sarnthein, M. Chapman, J. D. H. Schulz, S. van Kreveld, E. Vogelsang, and M. Weinelt, 2003: Glacial North Atlantic: Sea-surface conditions reconstructed by GLAMAP 2000. Paleoceanography, 18 .1065, doi:10.1029/2002PA000774.
Raynor, N., D. Parker, E. Horton, C. K. Folland, L. V. Alexander, D. P. Rowell, E. Kent, and A. Kaplan, 2003: Global analyses of sea surface temperature, sea ice, and night marine air temperature since the late nineteenth century. J. Geophys. Res., 108 .4407, doi:10.1029/2002JD002670.
Rosell-Melé, A., and N. Koç, 1997: Paleoclimatic significance of the stratigraphic occurrence of photosynthetic biomarker pigments in the Nordic seas. Geology, 25 , 49–52.
Rosell-Melé, A., and P. Comes, 1999: Evidence for a warm Last Glacial Maximum in the Nordic seas or an example of shortcomings in UK 37’ and UK 37 to estimate low sea surface temperature? Paleoceanography, 14 , 770–776.
Sarnthein, M., U. Pflaumann, and M. Weinelt, 2003: Past extent of sea ice in the northern North Atlantic inferred from foraminiferal paleotemperature estimates. Paleoceanography, 18 .1047, doi:10.1029/2002PA000771.
Seager, R., D. S. Battisti, J. H. Y. N. Gordon, N. Naik, A. Clement, and M. Cane, 2002: Is the Gulf Stream responsible for Europe’s mild winters? Quart. J. Roy. Meteor. Soc., 128 , 2563–2586.
Shin, S-I., Z. Liu, B. Otto-Bliesner, E. Brady, J. Kutzbach, and S. Harrison, 2003: A simulation of the Last Glacial Maximum climate using the NCAR–CCSM. Climate Dyn., 20 , 127–151.
Smith, R., and P. Gent, 2002: Reference manual for the Parallel Ocean Program (POP) ocean component of the Community Climate System Model (CCSM2.0). Los Alamos National Laboratory Tech. Rep. LA-UR-02-2484, 73 pp.
Smith, T., and R. Reynolds, 2003: Extended reconstruction of global sea surface temperatures based on COADS data (1854–1997). J. Climate, 16 , 1495–1510.
Son, S-W., and S. Lee, 2005: The response of westerly jets to thermal driving in a primitive equation model. J. Atmos. Sci., 62 , 3741–3757.
Timmermann, A., F. Justino, F-F. Jin, and U. K. H. Goosse, 2004: Surface temperature control in the North and tropical Pacific during the last glacial maximum. Climate Dyn., 23 , 353–370. doi:10.1007/s00382-004-0434-9.
Toracinta, E. R., R. J. Oglesby, and D. Bromwich, 2003: Atmospheric response to modified CLIMAP ocean boundary conditions during the Last Glacial Maximum. J. Climate, 17 , 504–522.
Trenberth, K. E., 1991: Storm tracks in the Southern Hemisphere. J. Atmos. Sci., 48 , 2159–2178.
Trenberth, K. E., and J. M. Caron, 2001: Estimates of meridional atmosphere and ocean heat transports. J. Climate, 14 , 3433–3443.
Trend-Staid, M., and W. L. Prell, 2002: Sea surface temperature at the Last Glacial Maximum: A reconstruction using the modern analog technique. Paleoceanography, 17 .1065, doi:10.1029/2000PA000506.
Uppala, S. M., and Coauthors, 2005: The ERA-40 Re-Analysis. Quart. J. Roy. Meteor. Soc., 131 , 2961–3012.
Yin, J. H., 2002: The peculiar behavior of baroclinic waves during the midwinter suppression of the Pacific storm track. Ph.D. thesis, University of Washington, 121 pp.
Yin, J. H., and D. S. Battisti, 2001: The importance of tropical sea surface temperature patterns in simulations of Last Glacial Maximum climate. J. Climate, 14 , 565–581.
Zurita-Gotor, P., and E. K. M. Chang, 2005: The impact of zonal propagation and seeding on the eddy–mean flow equilibrium of a zonally varying two-layer model. J. Atmos. Sci., 62 , 2261–2273.
Wintertime jet position in simulations of the PD and LGM climate. Each panel shows the 70th percentile of 200-mb horizontal wind from one model with warm colors showing the present-day jet and cool colors showing the glacial jet. The contour intervals appear at the top right of the panel and were chosen to clearly delineate the jet position in each case. The color scales give the wind speed (m s−1). The results shown are (a)–(b) two PMIP1 models coupled to slab oceans, (c)–(g) five PMIP1 models forced with standard PMIP1 boundary conditions, and (h) the atmosphere component (CAM3) of CCSM3 forced with PMIP1 boundary conditions. Included for reference are (i) ERA-40 data for 1958–2001 and (j) the CCSM3 fully coupled simulation. Details of the models appear in Table 1.
Citation: Journal of Climate 21, 14; 10.1175/2007JCLI2166.1
Wintertime atmospheric circulation in CCSM3 simulations. Zonal wind at 250 mb from the (a) PD and (b) LGM simulations (10 m s−1 contours). The Atlantic jet is stronger and more zonal, whereas the Pacific jet is largely unchanged. Geopotential height at 500 mb from the (c) PD and (d) LGM simulations (120-m contours with an offset of −5400 m). The Laurentide ice sheet over North America forces a strong stationary wave that intensifies the flow downstream. The thick solid lines in all the maps denote the zero contour.
Citation: Journal of Climate 21, 14; 10.1175/2007JCLI2166.1
Wintertime zonal wind and temperature profiles by ocean sector. Zonal wind (contours) and temperature (colors) for DJFM averaged over (top) the Atlantic sector (90°W–0°) and (bottom) the Pacific sector (150°E–27°W). The contour interval is 10 m s−1, and the thick solid lines denote the zero contour. The right-most panels show the horizontal zonal wind shear in a 10° latitude band, which is marked on each panel by horizontal bars, equatorward of the jet core (red is PD, blue is LGM). The sharp midlatitude temperature front and strong jet in the LGM Atlantic are robust features no matter how the ocean sector is defined.
Citation: Journal of Climate 21, 14; 10.1175/2007JCLI2166.1
(left) Atlantic jet and eddy characteristics in CCSM3 coupled simulations and (right) from reanalysis data. These plots show the relationship between the width and strength of zonal mean zonal winds and northward eddy heat flux υTATL (K m s−1) at 850 mb over the Atlantic sector (90°W–0°). (top left) The monthly-mean jet width vs jet strength for 50 Januarys in the PD (red) and LGM (blue) simulations; (bottom left) the strength of the eddy heat flux vs jet strength for 25 Januarys. The horizontal line is the mean υTATL for each simulation, with the vertical dash marking the 95% confidence limits. The jet strength is the max zonal wind speed in the sector; the jet width is the lat range over which the jet speed decreases to half its max value. These results are robust to the choice of month(s) used to define winter and to whether the jet and eddy strengths are taken as the max in the zonal mean of different lon ranges straddling the jet/storm track core, or as the area-weighted mean of the largest 30–50 values at the jet/storm track core. (right) The corresponding plots for 44 Januarys of reanalysis data in the period 1958–2001.
Citation: Journal of Climate 21, 14; 10.1175/2007JCLI2166.1
Meridional heat transport from observations and simulations. (a) Recent estimates (Trenberth and Caron 2001) of RT (entire shaded area), AT (dark gray shading), and OT (light gray shading) heat transport using data from the Earth Radiation Budget Experiment (ERBE) and ERA-40. The sum of the dark and light gray areas represents the total heat transport by the climate system. (b) Comparison of heat transports from observations and in the PD and LGM simulations by CCSM3. The shaded curves are the same observational estimates shown in the top panel. Max uncertainties in the model simulations are <0.10 PW, which is less than one-third of the value of the largest differences between the simulations. (c) The dry stationary wave (SW) and total transient (TR) components of atmospheric heat transport in the CCSM3 simulations. The SW heat transport was estimated by taking the vertical integral of [
Citation: Journal of Climate 21, 14; 10.1175/2007JCLI2166.1
Wintertime jet position and eddy diagnostics from ERA-40 data (1958–2001) and in the CCSM3 simulations. Colors show (a)–(c) eddy temperature flux 
Citation: Journal of Climate 21, 14; 10.1175/2007JCLI2166.1
Wintertime baroclinicity in CCSM3 simulations. The Eady growth rate parameter, σ = 0.31f0N−1(∂uh/∂z), in the 850–700-mb layer for DJFM (0.2 day−1 contours with values greater than 0.8 day−1 shaded).
Citation: Journal of Climate 21, 14; 10.1175/2007JCLI2166.1
One-point regression analysis of wintertime eddy structure in CCSM3 simulations. DJFM regression maps for the (left) PD and (right) LGM climates: (a), (d) υ′200 (2 m s−1) meridional wind at 200 mb; (b), (e) υ′850 (1 m s−1) meridional wind at 850 mb; (c), (f) υ′850T ′850 (5 K m s−1) meridional heat flux at 850 mb. Black contours show positive correlations, gray contours show negative correlations, and the zero contour is omitted. The black crosses mark the locations of the reference time series, which is the υ′200 wind at 43°N, 28°W for the PD simulation and at 38°N, 28°W for the LGM simulation. The solid gray line marks 40°N. All eddy fields are calculated from daily data that have been high-pass filtered to retain variability at periods less than 8 days.
Citation: Journal of Climate 21, 14; 10.1175/2007JCLI2166.1
Midwinter suppression of the Pacific storm track. This is an update of the figure from Nakamura (1992) showing the midwinter suppression of the Pacific storm track from NCEP–NCAR reanalysis data 1948–2002. (top) The seasonal cycle of eddy temperature flux
Citation: Journal of Climate 21, 14; 10.1175/2007JCLI2166.1
Wintertime jet position and eddy diagnostics in CCSM3 simulations compared to the simulation PMIPa reproducing the PMIP1 results. Colors show (left) eddy temperature flux
Citation: Journal of Climate 21, 14; 10.1175/2007JCLI2166.1
Reconstructions of land ice topography over North America at LGM. (left) Profiles of the Laurentide ice sheet complex at 21 ka across latitude 55°N from the ICE-4G (Peltier 1994) and ICE-5G (Peltier 2004) deglaciation history datasets. (right) Topography and bathymetry in the North American sector from the ICE-5G dataset, with the white contour marking the extent of the ice sheet and the black line (A–A′) marking the latitude of the profiles in the left panel.
Citation: Journal of Climate 21, 14; 10.1175/2007JCLI2166.1
Sea surface conditions during LGM. (a) CLIMAP reconstruction compared to PD observations. (b) LGM as simulated by the coupled model forced by ICE-5G and 21 ka insolation and greenhouse gases compared to PD observations. (c) LGM compared to CLIMAP. Contours in top panels show annual mean SST (4°C), and colors show differences (°C). The white contours mark the 50% sea ice concentration line for August (dashed) and February (solid). Observed SSTs and sea ice are taken from the period 1943–2001 from the extended reconstructed sea surface temperature (ERSST; Smith and Reynolds 2003) and HadISST (Raynor et al. 2003) reconstructions.
Citation: Journal of Climate 21, 14; 10.1175/2007JCLI2166.1
Sea surface conditions in the North Atlantic during LGM from reconstructions and as simulated by CCSM3. The reconstructions are from the CLIMAP project (CLIMAP 1981), the Glacial Atlantic Ocean Mapping (GLAMAP) project (Sarnthein et al. 2003), and Meland et al. (2005). Colors show SST (°C), with the 50% sea ice concentration line marked by the thick white contour for (top) August and (bottom) February.
Citation: Journal of Climate 21, 14; 10.1175/2007JCLI2166.1
Atmosphere models shown in Fig. 1. The model abbreviations are Global Environmental and Ecological Simulation of Interactive Systems, version 2 (GENESIS2); the UGAMP; Canadian Centre for Climate Modeling and Analysis, version 2 (CCC GCM2); European Centre Hamburg Model, version 3.6 (ECHAM3.6); Laboratoire de Météorologie Dynamique, version 5.3 (LMD5.3); and CAM3. The second column indicates whether the model participated in PMIP1 and the remaining columns indicate the resolution of the model.
Jet and eddy characteristics in the Atlantic sector (15°–65°N, 90°W–0°) for DJFM winter. The term uATL is the max zonal wind at 200 mb in the sector, σu is the std dev of uATL, KEIATL is the sector-mean monthly departure from the climatological mean of the kinetic energy of the horizontal wind at 200 mb, and υTATL is the sector-mean northward eddy heat flux at 850 mb. The 95% confidence intervals of the seasonally averaged quantities were determined using a Student’s t test for the means and a chi-squared test for the std dev.