Post-Eemian Glacial Inception. Part II: Elements of a

912 VOLUME 16J O U R N A L O F C L I M A T E

q 2003 American Meteorological Society

Post-Eemian Glacial Inception. Part II: Elements of a Cryospheric Moisture Pump

G. VETTORETTI AND W. R. PELTIER

Department of Physics, University of Toronto, Toronto, Ontario, Canada

(Manuscript received 8 January 2002, in final form 20 June 2002)

ABSTRACT

This paper extends the analyses of the glacial inception process described in a previous paper (‘‘Part I: TheImpact of Summer Seasonal Temperature Bias’’). The analyses described therein were based upon the use ofthe Canadian Centre for Climate Modelling and Analysis (CCCma) GCMII. Three simulations of the modernclimate system were described that were, respectively, warm biased, unbiased, and cold biased with respect tothe set of Atmospheric Model Intercomparison Project 2 SSTs and land surface temperatures in summer. Thesethree control models were perturbed by the modification of the orbital insolation regime appropriate to the time116 000 years before present (116 kyr BP) during which the most recent period of continental glaciation began.Two of the three simulations do deliver perennial snow cover in polar latitudes. Analyses of the land surfaceenergy balance, hydrological cycle, and energetics of the atmosphere in the Northern Hemisphere polar regionat 116 kyr BP discussed in greater detail herein reveal a set of positive feedback mechanisms favoring glaciation.It is proposed that these feedbacks are coupled to the main Milankovitch ice–albedo feedback that has heretoforebeen assumed to be the key to understanding the initiation of widespread continental glaciation. In particular,it is demonstrated that the polar surface energy balance plays an important role in summer snowmelt and in theannual maintenance of perennial snow cover. Furthermore, increased water vapor transport into the NorthernHemisphere summer polar regions at 116 kyr BP increases the net annual snow accumulation in these post-Eemian climate simulations through the action of an atmospheric–cryospheric feedback mechanism. An expla-nation for the absence of perennial snow cover in Alaska during the post-Eemian period is proposed. It issuggested that the transport of latent and sensible heat into this region is increased under 116 kyr BP orbitalforcing, which therefore acts to maintain sufficient summer snowmelt that is vitally important in preventingglacial initiation.

1. Introduction

On the basis of oxygen isotopic records derived fromdeep sea sedimentary cores (e.g., Hays et al. 1976; Im-brie et al. 1984) that are interpreted as providing a proxyfor global land ice volume, it is clear that land ice vol-ume and thus sea level are directly influenced by var-iations in the orbital obliquity (at 41 000 yr, or 41 kyr)and eccentricity-precession (at 23, 22, and 19 kyr) in-solation cycles that arise from gravitational n-body ef-fects in the solar system on the geometry of the earth’sorbit around the sun (e.g., Laskar et al. 1993). The 100kyr eccentricity cycle, whose frequency matches that ofthe dominant mode of ice volume variability followingthe onset of the mid-Pleistocene climate transition at900 kyr before present (BP) does not, however, induceany significant variation of solar forcing at the sameperiod. It is therefore clear that the origins of the 100kyr glacial–interglacial cycle of continental ice volumevariability must be associated with nonlinear processesthat arise internally to the climate system itself. A num-

Corresponding author address: Dr. W. R. Peltier, University ofToronto, Department of Physics, Toronto, ON M5S 1A7, Canada.E-mail: [email protected]

ber of such nonlinear feedback processes are expectedto be active during post-Eemian glacial inception thatoccurred at approximately 116 000 calendar years be-fore present. It was following this inception event thatthe vast continental ice sheets developed that subse-quently grew to cover all of Canada and much of north-western Europe.

On the basis of a variety of more direct measures ofrelative sea level including those based on corals (e.g.,Fairbanks 1989; Bard et al. 1990), the connection be-tween land ice volume and orbital insolation variationshas been further reinforced. Land ice is inferred to ex-pand most rapidly when radiative forcing is reduced insummer. The most significant reductions in summer in-solation arise during periods of high eccentricity andlow obliquity that are characterized by a precession an-gle that aligns the summer solstice with aphelion. Thisunique configuration results in a seasonal cycle that ischaracterized by a strong reduction in Northern Hemi-sphere summer high-latitude insolation as comparedwith modern. Milankovitch (1941) proposed that it isprecisely the resulting reduction in summer tempera-tures that would allow for the previous winter snowfallto persist through the summer season at high latitudes

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through a reduction in summer snowmelt and/or in-creased snowfall. This would lead to a situation in whichhigh-latitude snow fields would be continuously aug-mented over periods of thousands of years. This snowmass would then further increase in response to positivefeedbacks in the earth’s climate system. The main snow–ice albedo feedback would lead to a further decrease insummer temperatures and thus to a further increase incontinental ice sheet area and volume. Over many mil-lennia the snowpacks would combine into continentalice sheets that eventually covered the high-latitude con-tinents.

A number of reasons have been suggested to explainwhy the majority of atmospheric GCMs (AGCMs) failto produce glacial inception when the model insolationparameters are adjusted to those of the post-Eemian ep-och [see the companion paper (Vettoretti and Peltier2003, hereafter Part I) for a review]. There are alsoquestions as to why some of the models that do initiateglaciation have nucleation centers that are discordantwith the geologic record. The consensus view is ap-parently that these models either have inaccurate mod-ern climatologies or lack critical features of the realclimate system that are crucial in simulating the physicalinteractions that contribute to producing perennial snowfields at high northern latitudes. Others argue that modeldeficiencies simply derive from the use of overly lowAGCM horizontal and vertical resolutions of the dy-namics, physics, and boundary conditions. Some of thecomponents that AGCMs either account for only crude-ly or fail to represent altogether include land surfaceand/or biosphere processes, ocean dynamics, or cryos-pheric processes.

Many AGCM sensitivity studies of greenhouse gas–induced warming have revealed a large sensitivity tochanges in the doubling of atmospheric CO2 concen-trations. The range of global average temperature chang-es simulated in AGCMs of the generation used in thisstudy span the temperature range from approximately1.58 to 4.58C (Houghton et al. 1996). The large sensi-tivity in temperature is a result of the type of parame-terizations that are used in the models and the feedbacksthat occur within the model simulation, many of whichare not well understood. The model used in this studyhas a CO2 doubling sensitivity in global average tem-perature of 3.08C. Under changes in paleoclimate forc-ing and orbital perturbations a similar range of tem-perature change is observed within a set of models thatwere also used in CO2 doubling experiments (e.g., Pinotet al. 1999; Vettoretti et al. 2000a,b). This suggests thatit may be easier to initiate glaciation in models withlarger sensitivities to changes in boundary conditions.It is more likely, however, that models that have a cold-biased control climate would be subject to the onset ofinitiation more so than other comparable models [e.g.,see Joussaume et al. (1999) and Part I].

Long-term climate system variability as inferred onthe basis of proxy climate records (e.g., Imbrie et al.

1984; Barnola et al. 1987) suggests that atmosphericbehavior is significantly influenced by surface biomassvariability, ocean dynamics, and ice sheet growth/decay.In the context of one recent study of glacial inception(Dong and Valdes 1995), it has been suggested that thesurface thermal response of the ocean together with therepresentation of high-elevation continental orographymay be the most important requirements for a successfulsimulation of perennial snowfall. Others have suggestedthat migration of the distribution of vegetation may haveled to a positive (tundra–taiga) feedback in the climatesystem, which was responsible for initiating the peren-nial snow fields required for ice sheet growth (Gallimoreand Kutzbach 1996; de Noblet et al. 1996). Each ofthese prior analyses provide useful insights into the is-sues involved in understanding glacial inception.

In this second part of our study of this process wewill describe the results obtained from three AGCMsensitivity studies of the post-Eemian climate state at116 kyr BP. In the results section of Part I, we dem-onstrated that the reduction in summer seasonal inso-lation that occurred at 116 kyr BP was able to inducea state in which there is significant perennial snow coverat high northern latitudes in two of the three 116 kyrBP experiments. Here we will illustrate the action ofseveral positive feedback mechanisms that may be in-strumental in maintaining perennial snow cover con-ditions in this model. The focus of this study will beon the changes in surface energy balance and changesin the dynamics of the atmosphere at 116 kyr BP thatconspire to aid and abet this process. In the concludingsection we will review the results presented and discussthem in the context of ongoing work to produce a com-plete model of the 100 kyr ice age cycle in which thegeneral circulation model is explicitly coupled to a mod-ern three-dimensional thermo-mechanical model of con-tinental ice sheet evolution.

2. ResultsAccording to the premise in Part I, considerable un-

derstanding of the mechanisms involved in glacial in-ception can be achieved by investigating changes inclimate north of the Arctic Circle. Several of the mech-anisms that influence the accumulation of snow coverat 116 kyr BP will be investigated in what follows inan attempt to address the complete climate system re-sponse to the Northern Hemisphere decrease in high-latitude summer orbital insolation forcing [see Fig. 1 ofPart I for an illustration of the top-of-the-atmosphere(TOA) 116 kyr BP insolation regime]. The mechanismsto be discussed involve processes that impact the surfaceenergy balance and those that are dynamics related.These will be discussed in the following sections.

a. Feedbacks associated with the surface energybalance

The surface energy balance in AGCMs is known tobe significantly altered by the representation of dynamic



FIG. 1. (a) Topographic anomalies over Canada (in meters) between T32 spectral and 18 3 18gridpoint representation of the land–ocean surface in the model. (b) A cross section at 788Ndemonstrating the elevation difference and spatial misrepresentation between spectral and grid-point inferences of major continental features in the northern polar hemisphere.

and topographic boundary conditions within the model[e.g., see Holzer (1996) for a discussion]. In models ofthe kind employed in paleoclimate analyses, the bound-ary conditions are represented using spherical harmonicexpansions that are truncated at degree and order in therange from 15 to 42. There are therefore inadequaciesin properly representing the elevation in mountainousareas where high-elevation regions tend to exhibit agreat deal of orographic variance. This is well known,of course, and the implications of this effect have beendiscussed to some degree in previous studies (e.g., Pol-lard and Thompson 1997).

Two important problems concerning the semispectral

model used in this study are associated with spectralunderresolution of the topography. The first concernsthe spectral elevation anomalies at high latitudes andthe effect that these have on surface temperature whenthe atmospheric lapse rate is taken into account basedupon the difference between the actual elevation andthe model elevation (Fig. 1a). The second involves theinfluence of spectral topographic underresolution uponthe land–ocean mask used in the AGCM. In particular,there are large displacements of high northern latitudehigh-elevation land features that are in close proximityto other landmasses with high elevation such as is thecase with the highlands on Baffin Island and Ellesmere



FIG. 2. Arctic time series of (a) shortwave absorbed radiation, (b)surface net longwave radiation, and (c) latent and (d) sensible heatflux anomalies over land in the CB experiment. The contour intervalfor the shortwave radiation and heat fluxes is 10 W m22 while thatfor the longwave radiation is 5 W m22. Negative values are dashed.Ocean points are masked in gray.

Island near the Greenland ice sheet (Fig. 1b). WhenGreenland is spectrally decomposed to wavenumber 32in the model we employ, the steep slopes and high pla-teau are not adequately represented. Coastal elevationsare too high while the central regions of the ice sheettend to be too low (Fig. 1b). Since the spectral decom-position gives more weight to the 3-km high Greenlandice sheet, the 500-m average elevation of Baffin Islandthat lies directly to the west of Greenland is subjectedto large spectral ripples, which at times contradict themean sea level inferred from the model land–oceanmask. The first Gibbs oscillation has the effect of plac-ing the central portion of Baffin Island almost 600 mbelow its true elevation, which is below sea level (Fig.1a). When considering an average atmospheric lapserate of about 78C it becomes apparent that regions suchas Baffin Island are on average 48C warmer than wouldbe expected if the true elevation were included withinthe model. This is certain to affect mass balance in thisregion of the model. Baffin Island, which has mountainpeaks that reach 2000 m in elevation, is expected tohave been a nucleation center for the North Americanice sheet during post-Eemian glacial inception (Clark etal. 1993, and references therein) and is therefore notlikely to have its role in glacial inception properly cap-tured in a model of this nature. The western cordilleraby and large is resolved more accurately than the ele-vated regions to the east (Fig. 1b) but still suffers from500-m positive and negative height anomalies (Fig. 1a).This is likely to impact upon the mass balance that wasdiscussed in Part I of this study. In particular, it is in-teresting to note that the temperature over the St. EliasMountains would be approximately 48C colder if lapserate affects were taken into consideration. This mightbe enough to reduce summer snowmelt to the extentthat a local patch of perennial snow cover would developat 116 kyr BP in this region. It is also interesting tonote that the 500-m positive anomalies to the northwestand southeast of the St. Elias Mountains would act toinhibit perennial snow cover, in particular over Alaska,at 116 kyr BP after the positive 48C lapse rate correctionwas considered. In later studies, especially in coupledAGCM–ice sheet model (ISM) simulations, this effectshould be taken into consideration [e.g., see Marshalland Clark (1999) for detailed discussions].

The energy balance in the land surface scheme in thismodel is governed by the balance of absorbed solarradiation at the surface, the subsequent emission of ther-mal radiation, and the exchange of latent and sensibleheat fluxes from the surface to the atmosphere [see Vet-toretti et al. (2000a) for a complete description]. Thechange in soil moisture between liquid and solid phasesalong with the melting of snow cover also influence thesurface energy balance. The changes from 116 kyr BPto modern over land for the solar absorbed and net long-wave radiation along with latent and sensible heat fluxesare displayed in Figs. 2a–d for the cold-biased (CB)experiment. The patterns for the warm-biased (WB) and

unbiased (UB) experiments are similar in nature to theCB experiment but are of weaker magnitude than thechanges seen in the CB experiment. The decrease inabsorbed shortwave at the surface (Fig. 2a) is substan-tially greater than that characteristic of the 116 kyr BP(TOA) insolation anomaly. The July anomalies displayminima at 758N that reach 280 W m22 in the CB ex-periment. Changes in albedo, which result from increas-es in snow cover, appear to explain much of the decreaseobserved in the absorbed shortwave radiation at 116 kyrBP and will be discussed later. The changes in totalcloud cover in July over land range between 24% and26% in the three experiments (not shown), which im-plies that this component of the radiative balance allowsfor more shortwave absorption in summer at 116 kyrBP. This again indicates that albedo plays a dominantrole in the shortwave radiation absorbed at these highlatitudes.

The decreases in the net terrestrial longwave radiationemitted from the land surface over the course of theArctic annual cycle are illustrated in Fig. 2b for the CBexperiment. The positive anomaly (maximum) in thisfigure indicates that there is a decrease in the net upwardemitted longwave radiation at 116 kyr BP. The maxi-mum in the anomaly, however, occurs 2 months afterthe minimum in absorbed shortwave radiation, which



FIG. 3. Arctic time series of (a) surface temperature, (b) surfacealbedo, (c) precipitation, and (d) evaporation anomalies over land inthe CB experiment. The contour interval for the temperature is 28Cwhile that for the albedo is 0.1. The contour interval for precipitationand evaporation is 0.2 mm day21. Negative values are dashed. Oceanpoints are masked in gray.

occurs in July. The 2-month delay in the decrease innet longwave from the surface as compared with theshortwave radiation may simply result from the fact thatthere is more snow cover in August–October in the gla-cial inception experiments and thus less terrestrial emis-sion from the cooler surface at 116 kyr BP. The increasedamounts of soil moisture of 20–30 kg m22 that are pre-sent at 116 kyr BP in August and September (not shown)are likely to influence the amount of emitted longwaveradiation. Increased soil moisture would alter the energybalance at this time of the year by increasing the thermalheat capacity of the land at 116 kyr BP. The changes inevaporation and precipitation that will be discussed inwhat follows may also play a role in the longwave ra-diative phase lag.

The changes in latent and sensible heat fluxes aredisplayed in Figs. 2c and 2d, respectively. Each of thesecomponents of the surface energy balance have maximathat occur in July and indicate a decrease in the upwardlatent and sensible heat fluxes from the surface. As withthe solar radiation absorbed at the surface, the samecharacteristic patterns are observed in Figs. 2c and 2dfor the surface fluxes. The decreases in latent surfaceheat fluxes from the Arctic land in summer in the CBexperiment are approximately 40 W m22. Of the threeexperiments, those with colder controls display thegreatest changes (not shown). The decrease in sensibleheat flux from the surface at 116 kyr BP is similar incharacter to the change observed in latent heat flux buthas a summer decrease in the upward flux of 15 W m22

(Fig. 4d).At the lower boundary of the atmosphere, the trans-

ports of heat and water vapor across the earth’s surfaceare proportional to the vertical gradients in temperatureand specific humidity, respectively. Under snow-cov-ered conditions, the latent heat flux is solely determinedby the potential evapotranspiration. Therefore, at 116kyr BP the summer evaporation changes observed inthe Arctic Circle are by and large determined by thevertical specific humidity gradient at the surface sincethe snow-covered surface is considered saturated. Incontrast to this, the modern evaporation during summerwill also be a function of the moisture availability atthe surface and the evapotranspiration slope factor. Aswas seen in a study of the influence of changes in soilmoisture parameterization under mid-Holocene bound-ary conditions (Vettoretti et al. 2000a), the nature of theparameterizations must be considered when evaluatingthe behavior of the model under significantly differentboundary conditions.

The July changes in surface heat fluxes do not comeclose to balancing the decreases in absorbed shortwaveradiation. For example, the CB experiment has a deficitof approximately 20–25 W m22 that is not accountedfor in July at 758N. A decrease of 20 W m22 in theemitted thermal radiation is, however, observed 2months later in the CB annual cycle and may simplyindicate that there is less energy available at the surface

for snowmelt during the summer season at 116 kyr BP.A plot of the sum of the four surface flux anomalies(not shown) indicates that the time evolution of the sur-face temperature is changed at 116 kyr BP. This changein energy balance would act to either maintain snowcover or melt snow cover depending upon the time ofyear. The 116 kyr BP energy balance anomalies indicatethat the land is warming less quickly in summer andcooling less quickly in fall. This would in turn corre-spond to less snowmelt in summer and possibly de-creased snow formation in fall at 116 kyr BP.

The 116 kyr BP zonal land surface temperature anom-alies in the Arctic Circle (608–908N) for the entire yearare displayed in Fig. 3a for the CB model. The primarydecrease in polar cap surface temperature in each of theexperiments occurs from June through December withthe minimum in September. The magnitudes of the dif-ferences are very similar in each of the experiments with116 kyr BP zonal land surface temperature anomaliesdisplaying minima of 2108 to 2128C. The differencebetween each set of experiments, from warm to coldbiased lies in the southward latitudinal position of thecold anomalies, with the CB experiment displaying min-ima at 758N latitude (Fig. 3a) while the other two ex-periments (WB and UB) have minima at 858N (notshown). This southward movement of the region of



strongest thermal gradient is likely to strengthen thebaroclinicity along the polar front in each of the ex-periments. The spatial distribution of the summer polartemperature anomalies (not shown) is quite similar ineach experiment and all three experiments produce Arc-tic land temperature anomalies at 116 kyr BP that aresignificant at the 99% confidence level according to aunivariate t test. The maximum summer temperaturedecreases are in eastern and central Siberia and over theCanadian Arctic archipelago. Scandinavia and Alaskaalso experience decreases in temperature but of ap-proximately half the magnitude as the other glaciallysensitive regions.

These decreases in surface temperature are, as onewould expect, strongly correlated with changes in surfacealbedo (Fig. 3b). Increases in surface albedo progres-sively increase in southward areal extent from the WBexperiment to the CB experiment (not shown). The CBexperiment has zonal average albedo changes of greaterthan 0.4 at 758N latitude. This distribution of temperatureand surface albedo changes at 116 kyr BP is a goodexample of the positive ice–snow–albedo temperaturefeedback, which strengthens as glacial inception pro-gresses and extends southward over the millennia. Thesechanges in surface albedo are a major component influ-encing the radiative balance in the land surface schemeemployed within this model. The changes in surface tem-perature and surface albedo have minima and maximathat lag the decreases in insolation at 116 kyr BP byapproximately 2–3 months (see Fig. 1 of Part I).

Time series of the precipitation and evaporation northof the Arctic Circle (Figs. 3c and 3d) reveal the temporalevolution of this mechanism at 116 kyr BP. The pre-cipitation and evaporation are both dependent on theatmospheric moisture content, which is nonlinearly re-lated to the changes in atmospheric temperature throughthe Clausius–Clapeyron relation. The 116 kyr BP anom-alies in precipitation (Fig. 3c) and evaporation (Fig. 3d)are negative in summer in the CB experiment. All threeglacial inception experiments display the same behavior.The changes in precipitation have a minimum of 20.6mm day21 in July and August. The evaporation anom-alies at glacial inception vary up to 21.6 mm day21 inJuly at 708–758N latitude. The changes in evaporationare, of course, directly proportional to the decreases inlatent heat flux from the surface (Fig. 2c) and are wellcorrelated with the summer reduction of TOA solar in-solation. The precipitation anomalies occur more evenlyover summer and fall in each of the three experimentswith the changes in summer being most intense. Theexcess decrease of evaporation over precipitation insummer in the 116 kyr BP experiments leads to an in-crease of precipitation minus evaporation (P 2 E) atthis time and is central to the discussion that follows.

b. Feedbacks associated with the dynamics of theatmosphere

A recent study of last glacial maximum (LGM) cli-mate using the GCMII model (Vettoretti et al. 2000b)

has demonstrated that the changes in the hydrologicalcycle under cold climate perturbations have a significantinfluence on snow accumulation. The enhanced reduc-tion in evaporation relative to precipitation was foundto lead to a net increase in the precipitation minus evap-oration (P 2 E) anomaly over the Laurentide ice sheetat LGM. Thus, even though the vigor of the atmosphericcomponents of the hydrological cycle are significantlydecreased in cold climate conditions, there can never-theless be a positive impact on the terrestrial storagecomponent of the hydrological cycle. This phenomenonis precisely what is occurring under the cold climateperturbation at 116 kyr BP, except that the changes areconfined to the north polar region.

Changes in the July hydrological components aver-aged over land areas north of 608N latitude for the WB,UB, and CB experiments are displayed in Fig. 4. The116 kyr BP July anomalies in precipitation and evap-oration for each of the three experiments demonstratethat reductions in precipitation are about 50% less thanthose in evaporation (Fig. 4a). The changes in all threeexperiments, therefore, have approximately the same in-creases in P 2 E at 116 kyr BP. Contrary to this, thechanges in snowfall and sublimation increase at 116 kyrBP from the warm-biased to cold-biased experiments(Fig. 4b). The snowfall is increasing much more rapidlythan the sublimation as the control climate becomescolder. The differences between total P 2 E and solidP 2 E anomalies (Fig. 4c) illustrate that as the climatebecomes colder the total arctic P 2 E is undergoing atransition from the liquid to solid state. One would as-sume that these two polar hydrological curves shouldintersect as Northern Hemisphere glaciation progresses.This increased summer P 2 E over polar landmassesat 116 kyr BP is a strong candidate for a positive hy-drological cycle feedback mechanism on ice sheet ex-pansion over millennial timescales.

The action of this feedback mechanism over land insummer may be investigated in greater detail throughexamination of the atmospheric moisture balance equa-tion:

]W1 = · q 5 E 2 P. (1)

]t

In this equation the excess of evaporation over precip-itation, [E(u, f, t) 2 P(u, f, t)], at the earth’s surfaceis balanced by the local time rate of change of verticallyintegrated specific humidity in pressure coordinates,

po

21W 5 g q(u, f, p, t) dp, (2)E0

and by the divergence of the total atmospheric watervapor transport,

po

21q 5 g q(u, f, p, t)u(u, f, p, t) dp. (3)E0



FIG. 4. Arctic hydrological 116 kyr BP anomalies in Jul over landin the WB, UB, and CB experiments. (a) Changes between precip-itation and evaporation. (b) Changes between snowfall and subli-mation. (c) Total differences between (a) and (b).

FIG. 5. Schematic diagram of the polar water vapor feedback mech-anism involving the total Arctic precipitation minus evaporation (P2 E ), the flux of vapor (Fy) into the Arctic Circle, and the rate ofchange in Arctic precipitable water (]W/]t).

The vertical integrals are from the surface, p0, to thetop of the atmosphere, while u is the latitude coordinateand f is the longitude coordinate.

In the atmospheric branch of the hydrological cycle,the time rate of change of precipitable water, ]W/]t, isusually very small compared with the divergence andsource terms over global scales and on annual average(Peixoto and Oort 1992). However, on a high-latitudeseasonal basis this is no longer the case. Following Peix-oto and Oort, Eq. (1) can be reformulated in terms ofthe temporal and spatial average over a region boundedby a conceptual vertical wall at 608N latitude that resultsin a relation for the inflow and outflow of atmosphericwater vapor contained within the Arctic region:

]W5 F 1 (E 2 P). (4)y]t

In this equation the local rate of change of total pre-cipitable water ]W/]t contained within the atmospherenorth of the Arctic Circle is balanced by the polar excessof evaporation over precipitation, E 2 P, and by

p 2po p p21F 5 g y , f, p, t q , f, p, t df dp,y E E 1 2 1 23 30 0

(5)

the northward meridional flux of water vapor across thevertical wall at 608N. The balance between these com-ponents is illustrated in Fig. 5.

By considering the atmospheric branch of the polarhydrological cycle as described by Eq. (4), we can in-vestigate the nature of positive or negative feedbacksoccurring as a result of the strong reduction in polarinsolation at 116 kyr BP. Figure 6 displays the threequantities in Eq. (4) for the region north of 608N latitudefor the three modern controls and three 116 kyr BPexperiments. The annual cycle of the area averaged rateof change in total precipitable water and total E 2 Pare displayed in Figs. 6a and 6b, respectively, while theannual cycle of the total northward transport of watervapor into the Arctic Circle is displayed in Fig. 6c. Therate of change in precipitable water has a maximum inlate spring/early summer and a minimum in late sum-mer/early fall. The maxima and minima in the moderncontrols range from 665 to 695 kg m22 yr21 (5 665to 695 mm yr21) owing to the increased moisture con-tent of the atmosphere in summer in the cold to warm-biased experiments. The 116 kyr BP experiments rangefrom 635 to 660 mm yr21. The observed maximumand minimum rates of change in precipitable water areslightly more than 670 mm yr21 (Peixoto and Oort



FIG. 6. Time series of (a) the rate of change in total Arctic pre-cipitable water, (b) Arctic E 2 P, and (c) northward transport of watervapor into the Arctic Circle for the three modern control and three116 kyr BP experiments. Observed values are indicated in black.Arrows indicate the direction of increasing divergence and conver-gence. Units are kg m22 yr21.

FIG. 7. Time series of the 116 kyr BP anomalies for (a) the rateof change in total Arctic precipitable water, (b) Arctic E 2 P, and(c) northward transport of water vapor into the Arctic Circle. Arrowsindicate the direction of increasing divergence and convergenceanomalies. Units are kg m22 yr21.

1992). In each of the modern controls and 116 kyr BPexperiments there is also an excess of precipitation overevaporation all year round in the Arctic region with amaximum of 500 mm yr21 in winter and 300 mm yr21

in summer. The values from the modern control exper-iments are substantially more than the values obtainedfrom observations (Legates and Willmott 1990), whichsuggest that model polar P 2 E is excessive by ap-proximately 100–200 mm yr21 depending upon the timeof year (Fig. 7b). The total meridional transport of mois-ture by all motions across the Arctic Circle, not sur-prisingly, has the same characteristic annual cycle as P2 E with a winter maximum of between 80 and 120mm yr21 and a summer minimum of between 40 and80 mm yr21 in the modern controls and 116 kyr BP

experiments. Compared with the meridional transportsderived from observed specific humidity and wind ve-locities (Mitchell et al. 1993, 1996), the model controlsagree well with the general behavior of the observedtransport but again flow is too strong into the polarregions, most notably in the fall (Fig. 6c). It is notsurprising that the model quantities do not balance,which likely results from a combination of the procedureused for eliminating negative specific humidities (spec-tral hole filling) that is employed in the model (Mc-Farlane et al. 1992) and the limited resolution of themodel in the polar cap. The observed values do notbalance either, suggesting a lack of spatial and temporalcoverage at high polar latitudes. Taking into account thedeficiencies of both the model and the observationaldata, this comparison may nevertheless reveal a great



deal of information about the dynamics of the climatechange that is occurring in the polar region under thesignificantly modified solar forcing at 116 kyr BP.

The differences between 116 kyr BP and modern forthe three quantities in Eq. (4) are displayed in Fig. 7.The 116 kyr BP anomalies in the rate of change in totalprecipitable water in the polar cap display uniformchanges between the WB, UB, and CB experiments sug-gesting that the total amount of polar precipitable wateris linearly correlated with the polar mean temperaturein each of the simulations (Fig. 6a). The total polarprecipitable water has a bell-shaped annual cycle witha maximum value in summer and a minimum value inwinter (time integral of Fig. 6a). The differences in therate of change of precipitable water between the threemodern and three 116 kyr BP simulations are almostidentical (Fig. 7a). This indicates a reduction in the rateof increase of polar precipitable water content in latespring of 230 mm yr21 and an enhanced rate of decreaseof polar precipitable water content in late summer of 25mm yr21 between 116 kyr BP and modern. The simi-larities in the anomalies of the three experiments indi-cates that the 116 kyr BP anomalies in the rate of changein polar precipitable water are not dependent on themodern control simulation. Changes in E 2 P between116 kyr BP and modern (Fig. 7b) indicate increasedanomalies of total polar P 2 E in summer over landand ocean of approximately 25–40 mm yr21 in all threeexperiments. There is also a decrease in P 2 E overthe same region in fall, which indicates there is a re-duced excess of precipitation over evaporation in thisseason at 116 kyr BP. Thus, in summer, there is enhancedconvergence from the P 2 E component in the polarcap at 116 kyr BP considering ]W/]t . 0. It was alsodemonstrated in Fig. 4b that there are increases in solidP 2 E over much of the polar land areas north of theArctic Circle in summer in the 116 kyr BP simulations.It is interesting that there is a reduction in overall sum-mer precipitation in this region yet an increase in sum-mer snowfall.

The fall increase in evaporation over precipitation inall three simulations arises in part due to the weakeningin the rate at which the content of polar precipitablewater is decreasing from the atmosphere in fall. Thismay also explain the decrease in the surface net upwardlongwave (Fig. 2b) over land at this time of the yearthrough the evaporative cooling mechanism. From Eq.(1), the decrease in summer evaporation over summerprecipitation at 116 kyr BP (Fig. 4a) implies a localmoisture convergence anomaly over land since ]W/]t .0 between July and August in all the simulations (Fig.6a). The 116 kyr BP anomaly in northward transport ofmoisture in summer of approximately 20–40 mm yr21

in all three experiments indicates convergence of mois-ture into the Arctic polar cap region at this time of theyear (Fig. 7c). This suggests a balance with the increasedsummer P 2 E component as ]W/]t . 0 in midsummer.The fact that each biased experiment displays the same

behavior in summer but not necessarily in winter impliesthat this climate mechanism is exerting a strong influ-ence when a strong reduction in TOA insolation at 116kyr BP is implemented in the model. This cold pertur-bation positive feedback mechanism is a result of theincreased northward moisture transport into the polarcap and compensating increased P 2 E while the tem-poral change in moisture content of the summer at-mosphere remains nearly the same. Although the polaratmosphere is drier at 116 kyr BP, the increased moistureconvergence over land that feeds summer snowfall maybe considered as the end product of a stronger polar‘‘cryospheric moisture pump.’’ This mechanism is ul-timately a product of the large drop in the summer in-solation at these high latitudes.

The modern simulated vertically integrated annualmean total water vapor transport reveals the circumpolarnature of the transport of moisture about the Arctic Cir-cle. The modern simulated transport for the UB exper-iment is displayed in Fig. 8a. The other (WB and CB)simulations have total water vapor transports that arevery similar in structure to the UB simulation (notshown). The general circulation of this quantity agreesfairly well with the observed circulation (e.g., see Vet-toretti et al. 2000b) but with an intensity that is increasedby approximately 35% in magnitude along the midlat-itude jet streams owing to the overly intense simulationof the hydrological cycle in this model. The simulatedand observed transports are in much better agreementat high latitudes owing to the much drier conditionspresent in the atmosphere in this region. The atmo-spheric water vapor transport at high latitudes is influ-enced by the orography and the midlatitude jet streammaxima over the eastern Atlantic and eastern PacificOceans. In the western Pacific, just south of the ArcticCircle, moisture is advected northward over Alaska be-cause of the influence of the planetary wave forcing inthe atmosphere by the western cordillera. In the modernsimulation some of the moisture transport over Alaskais split into two branches, one directed over easternSiberia and the other much stronger branch flowing overthe Canadian Arctic archipelago. A similar atmosphericflow pattern occurs over the Scandinavian region wheremoisture is again forced northward across the ArcticCircle and across central Siberia through stationarywave forcing. Considering the circumpolar vortex struc-ture of the total atmospheric water vapor transport with-in the Arctic Circle and the regions of perennial snowcover at 116 kyr BP, it will be observed that the nucle-ation regions (see Fig. 7 of Part I) lie in the arid regionsjust to the north of the outflow branch of the total watervapor transport across the Arctic Circle. It is also in-teresting to note that the regions that did not nucleateperennial snow fields, namely Alaska and Scandinavia,lie in the inflow branches of the atmospheric water vaportransport across the Arctic Circle. The fact that Alaskadid not experience significant glaciation in the periodsubsequent to the marine oxygen isotope stage (MOIS)



FIG. 8. (a) Contour plot of the modern simulated annual average total water vapor transport for the UB experiment.The 116 kyr BP anomalies in total water vapor transport for the (b) WB, (c) UB, and (d) CB experiments. Thecontour interval is 10 kg m21 s21 in (a) and 5 kg m21 s21 in (b), (c), and (d).

5e/5d transition is clearly in excellent accord with thisresult. The fact that Fennoscandia did become glaciated,however, suggests that the glaciation of Canada mayhave been required before this could occur, the resultof which could well have been a repositioning of theinflow branch of the atmospheric water vapor transportacross the Arctic Circle away from the Scandinavianregion. This might be an expected consequence of apartial collapse of the process of North Atlantic DeepWater formation.

The 116 kyr BP total water vapor transport anomaliesin the three glacial inception experiments, that range

from warm to cold biased, are displayed in Figs. 8b–d.A common feature characteristic of all three flow anom-alies is that the transport vectors are in a direction thatis opposite to the circumpolar moisture current in themodern control simulation. Either the circumpolar cur-rent is reduced in magnitude in the cold post-Eemianexperiments or there is a southward migration of themain branches of the midlatitude moisture transport ora combination of both. The WB and UB anomalies havefeatures that are very similar to one another. One com-mon feature is the decrease in outflow of moisture fromthe Canadian Arctic archipelago as indicated by the



tongue-like anomaly in this region. This feature is alsoseen in the CB experiment. This may be one of thereasons why there is an increase in moisture conver-gence over the Arctic landmasses simulated during 116kyr BP summer conditions. The same phenomenon isoccurring in the easternmost portion of Siberia wherethe outflow is reduced in magnitude by approximately20% in all three glacial inception experiments. It is alsointeresting to note that there is an increase in the flowover Scandinavia in the WB and UB experiments butnot in the CB experiment. In fact, the changes in theCB experiment indicate a very strong decrease in theatmospheric transport over Alaska, Scandinavia, andwestern Europe due to the significant depression in sum-mer temperatures. The significant glaciation seen in cen-tral and eastern Siberia in the CB experiment may be aresult of this significant realignment of the atmosphericinflow and outflow of water vapor into this region re-sulting in increased moisture convergence over land ar-eas and areas of perennial snow cover.

The energetics of the atmosphere, as described by themain intrinsic forms of energy, namely the internal,gravitational-potential, kinetic, and latent heat energyconstitute essential components in the atmospheric heatengine. A complete analysis of the changes in energybalance in the 116 kyr BP atmosphere is beyond thescope of this study and is part of an ongoing investi-gation. However, it is useful to briefly explore the natureof the changes in the transient eddy activity of the at-mosphere in the north polar region to illustrate somedifferences in the high-latitude atmospheric dynamicsof the post-Eemian period. The meridional transport oflatent and sensible heat by transient eddies is arguablythe most important component of the energy exchangein the general circulation of the midlatitude NorthernHemisphere atmosphere, especially in the vicinity of thehighly baroclinic polar front. The moisture (latent heat)transport by transient eddies, predominantly in the mid-latitudes, is responsible for a much greater fraction ofthe total moisture transport in the meridional than in thezonal direction. The modern simulated annual meannorthward transport by transient eddies is displayed inFig. 9a for the UB experiment. The modern simulatedtransport for the other two experiments is much the sameas in Fig. 9a and agrees well with the magnitude andbehavior of the observed atmospheric eddy moisturetransport (not shown) (Mitchell et al. 1993, 1996; Peix-oto and Oort 1992). Much of the northward moisturetransport into the Arctic Circle by transient wave activ-ity is dominated by maxima that lie just east of the twomain continental landmasses, namely, North Americaand Asia. These maxima, which are associated with theGulf Stream and Kuroshio western ocean boundary cur-rents, are the source of a large portion of the northwardtransport of moisture in these strongly baroclinic re-gions. Although the maxima occur at approximately358N latitude in the annual mean, the eddy activity isstill significant even across the Arctic Circle in eastern

Canada and eastern Siberia. Two other maxima in north-ward eddy moisture transport occur to the southwest ofAlaska and to the southeast of Iceland. These maxima,which are of approximately half the magnitude, occurat the tail ends of the Gulf Stream and Kuroshio Currentsalong the polar front where warm subtropical air inter-acts with cold polar air. Also of note are two weakminima of southward moisture transport by transienteddies that are situated symmetrically, 1808 in longitudeapart, over Greenland and the Russian Arctic Ocean.

The 116 kyr BP annual anomalies in the northwardmoisture (latent heat) transport by transient eddy activityfor the three experiments are displayed in Figs. 9b to9d. The changes in summer transport are approximatelytwice the magnitude as those in winter, with each seasondisplaying a similar spatial distribution (not shown). Afairly consistent pattern emerges in the 116 kyr BP mois-ture transport anomalies with the reduction in summerinsolation. The two regions of southward transport overGreenland and the Russian Arctic Ocean are signifi-cantly reduced while the transport over the Canadianwest coast and Alaska are increased. Also of note is thedecrease in northward transport over the Canadian Arc-tic archipelago as well as over eastern Siberia in allthree experiments. Over eastern Europe and Scandinaviathe northward transport becomes successively strongerwith increasing cold bias in each of the three experi-ments. The CB experiment delivers a strong northwardtransport in eastern Europe.

The modern annual mean northward sensible heattransport by transient eddies shown in Fig. 10a for theUB experiment is very similar to the moisture or latentheat transport in Fig. 9a. The northward transport ofsensible heat by transient eddies plays a dominant rolein the exchange of energy between warm and cold airmasses in regions of strong baroclinicity, especially inthe winter hemisphere. Therefore, it is not surprising toexpect increased summer transient wave activity duringthe post-Eemian epoch. The 116 kyr BP anomalies forthe annual mean meridional sensible heat transport bytransient eddies for the three experiments (Figs. 10b–d) are also characteristically similar to the eddy moisturetransport displayed in Figs. 9b–d, but are of approxi-mately twice the magnitude. The 116 kyr BP eddy heattransport anomalies are decreased substantially overeastern Siberia and the Canadian Arctic archipelago.The southward eddy heat transport over Greenland andthe Russian Arctic Ocean are again significantly de-creased. Likewise, the northward heat transport increas-es over eastern Europe and Scandinavia as the experi-ments go from warm to cold biased, with the CB ex-periment showing the greatest northward eddy transportof heat in this region.

The locations of perennial snow cover in the Arcticregion (see Fig. 7 of Part I) and the location of themaxima and minima in the 116 kyr BP anomalies ofnorthward transport of latent and sensible heat (Figs. 9and 10) by transient eddy activity are strongly corre-



FIG. 9. (a) Contour plot of the annual mean northward transport of latent heat by transient eddies for the UBexperiment. The 116 kyr BP anomalies in northward transient eddy latent heat transport for the (b) WB, (c) UB, and(d) CB experiments. The contour interval is 2 3 107 W m21 in (a) and 1 3 107 W m21 in (b), (c), and (d). To convertto latent heat transport (W m21), the moisture transport (kg m21 s21) is multiplied by the latent heat of vaporization:Le 5 2.5 3 106 J kg21.

lated. The absence of perennial snow cover formationin Alaska and Scandinavia in any of our experiments isassociated with the fact that increases of northward heatand moisture transport by transient eddies occurs in boththese regions. This increased activity at 116 kyr BP is

likely the crucial factor that maintains sufficient snow-melt during the summer season. The regions of nucle-ation over the Canadian Arctic archipelago and overcentral and eastern Siberia, on the contrary, coincidewith regions of decreased northward latent and sensible



FIG. 10. (a) Contour plot of the annual mean northward transport of sensible heat by transient eddies for the UBexperiment. The 116 kyr BP anomalies in northward transient eddy sensible heat transport for the (b) WB, (c) UB,and (d) CB experiments. The contour interval is 5 3 107 W m21 in (a) and 2 3 107 W m21 in (b), (c), and (d). Toconvert to sensible heat transport (W m21), the temperature flux (8C m s21) is multiplied by the specific heat capacityof air at constant pressure, the surface pressure, and the inverse of gravity: cpp0g21 . 1 3 107 J m22 8C21.

heat transport by transient eddies in the UB and CBexperiments. It seems entirely plausible that the zonalspatial heterogeneity of the Arctic nucleation zones isin part a result of increases and decreases in the north-ward transient eddy transport of latent and sensible heat

into the Arctic Circle at 116 kyr BP. These relationshipsmay actually be further amplified as the process of gla-cial inception proceeds and continental ice sheets growand advance in this polar region [see the LGM eddymoisture transport in Vettoretti et al. (2000b, their Fig.



17)] and will be the focus of further study to be de-scribed elsewhere.

3. Summary and conclusions

An investigation of glacial inception conducted usingthe Canadian Centre for Climate Modelling and Anal-ysis AGCM has demonstrated the operation of severalclimate feedbacks that may be instrumental in the de-termination of climate state in the post-Eemian period.The three experiments consisting of three modern con-trols and three 116 kyr BP experiments were performedusing mixed layer ocean and sea ice modules to assessthe sensitivity of the model to the summertime insola-tion minimum that occurred at this time. The three con-trol experiments, which were warm biased, unbiased,and cold biased with respect to the Atmospheric ModelIntercomparison Project 2 SSTs, provide a means oftesting the sensitivity of the ability of the model to de-liver perennial snow cover at high latitudes in responseto a reduction in incoming summer solar radiation.

The absence of nucleation over Alaska and Scandi-navia in the two 116 kyr BP experiments that do exhibitperennial snow cover pose a number of interesting ques-tions concerning the role of these regions in the evo-lution of the Northern Hemisphere into a glacial state.The three sets of sensitivity experiments, each withvarying degrees of glacial inception at 116 kyr BP, wereemployed to investigate several climate feedback mech-anisms that might be instrumental in further stimulatingperennial snow cover and may help explain the absenceof nucleation in regions such as the St. Elias mountainrange in Alaska.

In particular, the following points are of note.

1) The effect of the spectral topography on the massbalance of the polar regions at 116 kyr BP must betaken into consideration when assessing the resultsof a numerical simulation. The anomalies in surfaceelevation in the model indicate that lapse rate cor-rection to the mass balance would affect the distri-bution and accumulation of perennial snow cover inthe post-Eemian simulations. Significant changes inaccumulation and ablation would occur in the BaffinIsland region and along the western cordillera. Thesoutheastern portion of Baffin Island and the St. EliasMountains would have more positive mass balanceif lapse rate corrections were calculated. The spectralanomalies also suggest that Alaska would have amore negative mass balance in accord with geologicevidence that this region remained free of perennialsnow cover at this time.

2) The Arctic anomalies in the land surface energy bal-ance, dictated by the balance of absorbed shortwaveradiation, emitted longwave radiation, and the ex-change of latent and sensible heat fluxes from thesurface to the atmosphere are very well correlatedwith the reductions in summer polar insolation at

116 kyr BP. Reductions in the 116 kyr BP net short-wave absorbed anomalies at the surface are well be-low the reduction in insolation with minima occur-ring in July in the three sets of experiments. Thelargest decreases in absorbed shortwave radiationreach 80 W m22 in July in the cold-biased experi-ment. Decreases in the net longwave radiation at thesurface relative to the shortwave radiation minimumare delayed by approximately 2 months in each ofthe experiments, with the largest changes again oc-curring in the cold-biased experiment in September.Decreases in net longwave are virtually nonexistentduring the July insolation minimum. Decreases inupward latent and sensible heat fluxes have maximathat also occur in July but with sensible heat fluxanomalies of about half the magnitude of the de-creases in latent heat flux. In midsummer the ex-periments have an imbalance between the decreasein the upward latent and sensible heat fluxes and thenet solar radiation absorbed. Thus, at 116 kyr BPthere is excess heat flux being emitted from the sur-face and that may be an important mechanism fordecreasing summer snowmelt and maintaining pe-rennial snowfall. In the fall, the 2-month delay inthe net longwave radiation anomaly at the surfacemay simply indicate that the surface is cooling moreslowly than it does under modern climate conditions.The combined effect is that the snow cover is main-tained for much longer durations of time at thesehigh polar latitudes at 116 kyr BP.

3) The enhanced reduction of summer evaporation oversummer precipitation in the Arctic Circle over landresults in a net increase in precipitation minus evap-oration (P 2 E) at 116 kyr BP. This is the firstindication that there may be a climate mechanismthat is acting to create a convergence of moistureover Arctic land at 116 kyr BP. A subsequent detailedanalysis of the atmospheric moisture balance equa-tion for the Arctic polar cap north of 608N indicateda change in behavior of the polar climate at 116 kyrBP. In the polar climate system the rate of changein precipitable water is balanced by the meridionalflux of water vapor into the Arctic Circle and thetotal polar excess of evaporation over precipitation.The moisture being pumped into the polar region bythe general circulation is much stronger in winterthan in summer. With the reduction in insolation at116 kyr BP, this mechanism becomes more vigorousin summer. At 116 kyr BP the decrease in evapo-ration over precipitation balances the increased in-flux of moisture transport into the polar region. Withthe rate of change in precipitable water being fairlyconstant in summer the change in the rate of changeof precipitable water at 116 kyr BP is negligible.Thus, we find an atmospheric–cryospheric feedbackmechanism that is pumping more moisture into theArctic Circle along with enhanced convergence.Coupled with the fact that there are increased snow-



fall anomalies over land in the 116 kyr BP simula-tions, this mechanism constitutes a positive feedbackdue to the reduction in insolation at 116 kyr BP andwould enhance ice sheet growth over the ensuingmillennia.

4) Changes in the northward transport of latent and sen-sible heat by transient eddy activity display similarbehavior in the Arctic region in each of the three116 kyr BP experiments. The main features consti-tute a decrease in the northward heat transport intothe Canadian Arctic archipelago and over easternSiberia. Over eastern Europe and Scandinavia thetransport becomes successively stronger as the tem-perature of the modern control simulation decreases.Also of note is the consistent increase of latent andsensible heat transport over Alaska in all three 116kyr BP experiments. The eddy sensible heat transportanomalies are found to be approximately twice themagnitude of the eddy latent heat flux anomalies. Ingeneral it is found that the lack of formation of pe-rennial snow cover in Alaska and Scandinavia in anyof the experiments coincides with increases of north-ward heat transport by transient eddy activity. Thezonal spatial heterogeneity of the Arctic nucleationzones appears to be in part a result of increases anddecreases in the northward transient eddy transportof heat and moisture into the Arctic Circle at 116kyr BP. This effect is likely to increase with de-creasing temperature and increased snow cover overtime. Of particular note, over Alaska, the increasein heat and moisture transport is the likely expla-nation for the large summer snowmelt observed inthe 116 kyr BP experiments and the cause of theabsence of nucleation in this region.

This investigation suggests that a number of climatemechanisms are acting to modify the 116 kyr BP climatesuch as to make conditions for perennial snow covermore favorable. These mechanisms seem to constitutea set of positive feedback mechanisms that couple withthe main Milankovitch ice–albedo positive feedback re-sulting from the decrease in summer insolation. A moredetailed investigation of the energy balance in the polaratmosphere may provide greater insight into alterationsin the atmospheric energy cycle that may have occurredat 116 kyr BP. Downscaling the representation of thetopography in GCMII (e.g., Glover 1999) in order toachieve a more appropriate mass balance in simulatingglacial inception conditions may also provide a newdirection for research in the modeling of this process.A transient simulation of post-Eemian glacial inceptionwith a coupled AGCM–ISM spanning approximately10 000 years may also reveal whether some of the can-didate mechanisms suggested to constitute positivefeedback mechanisms in this study are actually furtheramplified. The use of a coupled AOGCM and ISM willultimately allow a more detailed investigation of thetransient evolution of continental ice sheet expansion in

the Northern Hemisphere and insight into the role thatthe Atlantic and Pacific Oceans might play in the de-termination of ice advance in these regions. The absenceof perennial snow cover in the Scandinavian region maybe a result of inadequate resolution of the topographyin this region. However, it is also possible that for Scan-dinavian glaciation to occur, the Northern American re-gion needs to be in an advanced state of glaciation suchas would have been characteristic of prevailing condi-tions at the time of the later insolation minimum thatoccurred at 70 kyr BP. The strength of the thermohalinecirculation would likely have been somewhat reducedat this time, which would also have served to furthercool the Scandinavian region as this lies downstream ofthe regions where deep water currently forms and theoverlying atmosphere is thereby significantly warmed.As paleoclimate models become more sophisticated andpaleoclimate data more abundant, a greater understand-ing of past climates will ensue, undoubtedly leading toa greater understanding of climate change processes ingeneral, including those that will govern our future.

Acknowledgments. The authors would like to thanktwo anonymous reviewers for their helpful commentsand suggestions. Support for this work has been pro-vided by the Climate System History and DynamicsResearch Network that is funded by the Natural Sciencesand Engineering Research Council of Canada and bythe Meteorological Service of Canada.

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Post-Eemian Glacial Inception. Part II: Elements of a