MARCH 1999 VAVRUS 873

The Response of the Coupled Arctic Sea Ice±Atmosphere System to Orbital Forcing and Ice Motion at 6 kyr and 115 kyr BP

STEPHEN J. VAVRUS Center for Climatic Research, Institute for Environmental Studies, University of WisconsinÐMadison, Madison, Wisconsin

(Manuscript received 2 February 1998, in ®nal form 11 May 1998)

ABSTRACT A coupled atmosphere±mixed layer ocean GCM (GENESIS2) is forced with altered orbital boundary conditions for paleoclimates warmer than modern (6 kyr BP) and colder than modern (115 kyr BP) in the high-latitude Northern Hemisphere. A pair of experiments is run for each paleoclimate, one with sea-ice dynamics and one without, to determine the climatic effect of ice motion and to estimate the climatic changes at these times. At 6 kyr BP the central Arctic ice pack thins by about 0.5 m and the atmosphere warms by 0.7 K in the experiment with dynamic ice. At 115 kyr BP the central Arctic sea ice in the dynamical version thickens by 2±3 m, accompanied bya2Kcooling. The magnitude of these mean-annual simulated changes is smaller than that implied by paleoenvironmental evidence, suggesting that changes in other earth system components are needed to produce realistic simulations. Contrary to previous simulations without atmospheric feedbacks, the sign of the dynamic sea-ice feedback is complicated and depends on the region, the climatic variable, and the sign of the forcing perturbation. Within the central Arctic, sea-ice motion signi®cantly reduces the amount of ice thickening at 115 kyr BP and thinning at 6 kyr BP, thus serving as a strong negative feedback in both pairs of simulations. Ice motion causes the near- surface air to cool in both sets of experiments, however, thus representing a positive feedback at 115 kyr BP and a negative feedback at 6 kyr BP. The excess cooling with ice motion at 115 kyr BP is caused by the enhanced, advective spreading of the ice pack into the North Atlantic dominating over the warming tendency from the thinner central Arctic sea ice. The reduced atmospheric warming due to ice dynamics at 6 kyr BP is caused by sea-ice ridging, a thickening process that partially counteracts the orbitally induced atmospheric warming per- turbation.

1. Introduction Sea-ice motion is one process that has often been Polar regions have been identi®ed by modeling stud- omitted from climate models, even though it is regarded ies as the areas expected to experience the most dramatic as the most likely candidate to counter the positive feed- future climate changes from increased levels of atmo- backs inherent in the thermodynamic-only sea-ice codes embedded in most atmospheric GCMs (WCRP 1994). spheric CO2. One of the strongest positive feedbacks cited in these analyses is associated with the retreat and The negative feedbacks effected by ice dynamics are thinning of sea ice in high latitudes, particularly the believed to stem from two primary features that are Arctic (IPCC 1990). Unfortunately polar regions are absent in a stationary ice cover: ice advection and local also among the least understood climatically, due to the ice thickness variations, which include leads (Hibler logistical dif®culties they pose for obtaining data and 1984). Ice motion causes regions of convergence and the added complexities inherent in modeling cryo- divergence within the pack, giving rise to regional ice spheric components. This paper will test the robustness thickness variations. In some areas, such as compressive of enhanced polar climate sensitivity simulated in nu- regions along coasts, this deformation results in the spa- merous GCMs by incorporating a frequently neglected tial pattern of sea-ice thickness being dominated by at- process (sea-ice transport) believed to provide negative mospheric wind forcing rather than thermodynamic feedbacks and by performing data-model comparisons forcing. In these regions, one would expect that exter- of past climate changes in high latitudes. nally forced atmospheric thermal perturbations would have a relatively small bearing on the thickness of the ice cover. The second distinguishing aspect of a dynamic ice pack is the ridging caused by small-scale motion Corresponding author address: Dr. Stephen J. Vavrus, Center for that results in a range of ice thicknesses locally and the Climatic Research, Institute for Environmental Studies, University of WisconsinÐMadison, 1225 West Dayton St., Madison, WI 53706- creation of open water. This variable thickness distri- 1695. bution is important because sea ice grows at a rate in- E-mail: [email protected] versely proportional to its thickness; thus, the overall

᭧ 1999 American Meteorological Society

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FIG. 1. Mean-annual sea-ice drift and sea level pressure (mb) [from R. Colony, University of Washington Polar Science Center, in Barry et al. (1993)]. growth rate of the pack can be dominated by that of the thin ice portion of the spectrum (Maykut 1982). There- fore, if mechanical forcing produces a region of thin ice, then the relatively rapid thermodynamic growth helps to restore the ice toward its original thickness. Likewise, if thermodynamic forcing causes a thinner ice pack, the relative ease with which thin ice can ridge will favor a mechanically induced thickness restoration (Hibler 1984). The importance of ice motion in shaping the modern Arctic sea-ice cover is apparent from a consideration of FIG. 2. Arctic sea-ice draft (m) during (a) summer and (b) winter its average circulation characteristics, as described by [from Bourke and Garrett (1987), in Wadhams (1994)]. Barry et al. (1993) (Fig. 1). The motion of the interior ice pack strongly resembles the implied surface wind stress, consisting of a clockwise circulation around the Ice motion is necessary to reproduce the observed pileup Beaufort Gyre and a drift of ice away from the Siberian of thick ice along the North American coast and the ice coast. After sea ice crosses the North Pole some of it tongue along Greenland (Hibler and Walsh 1982; Pol- converges along the North American coast; the remain- lard and Thompson 1994). Even small amounts of leads, der exits the basin through the Fram Strait and into the which are created by moving ice, cause substantial East Greenland Current. The presence of these two path- warming at the Arctic Ocean surface (Ledley 1991) and ways is strikingly apparent in ice thickness ®elds (Fig. in the lower troposphere of both polar regions (Vavrus 2), which show a buildup of sea ice north of Greenland 1995; Simmonds and Budd 1991). The inclusion of and the Canadian Archipelago and a tongue of ice ex- leads also improves the simulation of Antarctic sea-ice tending along the east coast of Greenland. extent (Washington et al. 1976), whose seasonal cycle Model simulations of the modern Arctic and Antarctic is dif®cult to reproduce without the inclusion of ice ice packs are also strongly affected by ice dynamics. advection (Hibler and Walsh 1982). Furthermore, ice

Unauthenticated | Downloaded 10/01/21 10:31 PM UTC MARCH 1999 VAVRUS 875 transport from the Arctic Ocean prevents spurious en- change in Arctic sea-ice characteristics upon the inclu- hancement of the pycnocline strength (Ranelli and Hib- sion of ice transport in a doubled CO2 experiment, but ler 1991), and ice export from the Arctic and Antarctic the response of Antarctic sea ice was signi®cantly mut- Oceans results in considerable thinning of sea ice in ed. both hemispheres (Ledley 1991). Ice dynamics also ap- As a means of clarifying the role of dynamic sea ice pear to have important consequences for climate vari- in shaping high-latitude climate changes, this paper will ability; Hibler and Zhang (1994) report that the absence assess the effect of ice motion on Arctic climate sen- of Arctic ice advection results in a spurious poleward sitivity and will evaluate the ability of a GCM to sim- displacement of the Marginal Ice Zone in the North ulate two time periods (6 kyr and 115 kyr BP) whose Atlantic. climates were strongly affected by alterations of Earth's The importance of ice motion in affecting the sen- orbital con®gurations. This pair of paleoclimates are sitivity of sea ice to climate perturbations has been high- noteworthy because they represent the response of the lighted in simulations ranging in sophistication from high-latitude Northern Hemisphere to a signi®cant in- stand-alone one- and two-dimensional sea-ice models solation surplus (6 kyr BP) and de®cit (115 kyr BP) to three-dimensional regional and global GCMs. Al- relative to present. The approach used here allows for though the consensus view that has emerged from these a number of improvements over previous studies. First, studies is that ice dynamics have a potent stabilizing this methodology provides an opportunity to test the effect on sea-ice variations, the sensitivity tests have in¯uence of ice motion over a range of external forc- usually focused on warming perturbations; very little ings, from the warm perturbation at 6 kyr BP to the consideration has been given to the in¯uence of mobile cold anomaly at 115 kyr BP. Global snapshots of these ice in cooling scenarios. Early simulations by Hibler paleoclimates have been simulated before but without (1984) of the Weddell Sea found that a dynamic-ther- dynamic sea ice and usually with less realistic surface modynamic sea-ice model showed less sensitivity than boundary conditions (e.g., prescribed SSTs and sea-ice a thermodynamic-only model to atmospheric warming extent). Second, because the model is driven by actual perturbations. Later simulations over the same domain insolation perturbations known to have occurred, this by Lemke (1987) showed that ice dynamics mitigate the approach allows for comparison with paleoenvironmen- response of sea ice to snowfall variations, a conclusion tal evidence and is more realistic than sea-ice sensitivity corroborated for Arctic sea ice by Holland et al. (1993). studies that have simply increased the solar radiation Curry et al. (1995) demonstrated that an ice thickness ¯ux by a prescribed fraction. Finally, the dynamics sea- distribution can substantially reduce the sensitivity of ice code used in this study [the cavitating ¯uid model the equilibrium Arctic sea-ice thickness to downwelling of Flato and Hibler (1990)] is more sophisticated than atmospheric radiation anomalies. Bitts et al. (1996) those included in previous climate models, many of found that the monthly variance of mean Arctic sea-ice which fail to conserve momentum and apply arbitrary thickness decreased by 50% when an ice export term thickness thresholds at which ice motion ceases. It will was included, suggesting that ice advection regulates be shown that blanket statements regarding the climatic thermodynamically induced ¯uctuations. In a doubled feedback produced by ice dynamics are inappropriate;

CO2 experiment Pollard and Thompson (1994) reported the effect of mobile sea ice depends on the climatic that ice dynamics damp the decrease in sea-ice extent variable, the season, and the sign of the atmospheric slightly in the Arctic and by 25% in the Antarctic. thermal perturbation. In general, however, ice transport A closer examination of modeling simulations with acts as a cooling mechanism by spreading the ice pack dynamic ice, however, reveals numerous contradictions, equatorward. which muddle efforts to pinpoint the role of sea-ice motion in modifying climate. Hibler (1984) showed that 2. Model description the inclusion of an ice thickness distribution has little effect on the sensitivity of sea ice area or volume to The National Center for Atmospheric Research's atmospheric warming, and Flato (1997) found that such (NCAR) Global Environmental and Ecological Simu- a distribution actually enhances the ice thickness sen- lation of Interactive Systems, version 2.0 (GENESIS2) sitivity to downwelling longwave radiation increases. is a coupled atmosphere±mixed layer ocean global cli- Semtner's (1987) Arctic Ocean GCM with sea-ice dy- mate model that originated from the NCAR Community namics simulated a summertime melt off of sea ice when Climate Model version 1. Because a detailed account the air temperature was increased by only 2 K, exceed- of GENESIS2 is given by Thompson and Pollard ing the response of thermodynamic-only models to (1997), only a brief summary of the GCM's features is warmer anomalies (Parkinson and Kellogg 1979; Kutz- given here. bach and Gallimore 1988). A simulation by Bjork GENESIS2 consists of three primary components: an (1992) showed that ice motion may reduce or enhance atmospheric model, a static mixed layer ocean, and a the sensitivity, depending on how much of the absorbed land surface package that contains a sea-ice code and solar energy within the mixed layer is transferred to the prescribed vegetation. The atmospheric model employs ice bottom. Pollard and Thompson (1994) found little a horizontal spectral truncation of T31 (approximately

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3.75Њϫ3.75Њ) and contains 18 vertical levels in a hybrid overlying snow cover becomes heavy enough to depress coordinate system, which tend from sigma coordinates the snow±ice interface below the ocean surface, then near the ground to normalized pressure coordinates at some of the snow is converted to ``white ice.'' the top of the atmosphere. Clouds are predicted using Sea-ice motion is included using the cavitating ¯uid prognostic three-dimensional water cloud amounts (Se- model of Flato and Hibler (1990), in which ice resists nior and Mitchell 1993) and resolved into three separate compression but offers no resistance to divergence or types: cirrus, stratus, and convective. The physical ef- shear. Ice is transported in a momentum-conserving fects of vegetation are accounted for by the land surface manner by wind and ocean stresses. The wind stress is transfer model, which exchanges energy, mass, and mo- determined directly from the atmospheric GCM's mentum between the atmosphere and vegetation. Be- (AGCM) prognostic surface wind velocities, whereas cause the off-line results of a global vegetation model the ocean stress depends on the relative velocity of sea are used to prescribe the distribution and physical at- ice with respect to the underlying ocean currents, whose tributes of the modern terrestrial biosphere, there are no prescribed values are obtained from an off-line 2Њϫ2Њ interactive climate±vegetation feedbacks. The GCM's ocean GCM for the modern climate (Thompson and ocean component is a mixed layer of ®xed 50-m depth Pollard 1997). Thus the paleoclimatic sea-ice velocities with a prognostic meridional heat transport that depends are somewhat constrained to present conditions, but it on latitude, the land±ocean fraction, and the latitudinal appears that the modern ocean circulation pattern of ice SST gradient. As the sea-ice fraction in a grid box ap- discharge through the Fram Strait has been remarkably proaches 1 the heat ¯ux convergence is linearly reduced stable over the past 180 kyr (Hebbeln and Wefer 1997). toward zero in the Northern Hemisphere and 6 W mϪ2 in the Southern Hemisphere. An enhanced wintertime 3. Paleoclimate simulations heat ¯ux is applied to the Norwegian Sea when the SST falls below 274.2 K to simulate the buffering effect of The Arctic climates of 6 kyr BP and 115 kyr BP advection by warm ocean currents and a deepening provide models with targets to hindcast environmental mixed layer during winter. This extra heat source pre- conditions known to have been signi®cantly different vents sea-ice formation over this domain. The mixed than present, yet with most boundary conditions similar layer depth in Hudson Bay is reduced to 25 m to permit (e.g., continental con®guration, sea level, ice sheet cov- the observed modern freeze-up during December. erage) or well established (insolation). Around 6000± Sea ice is represented by both thermodynamic and 9000 yr ago the earth's tilt was larger than today, and dynamic components. The thermodynamic portion re- modern boreal wintertime perihelion occurred instead solves ice and snow cover into three layers, whose melt- during summer. 115 kyr ago a larger eccentricity, small- ing and freezing formulations follow Semtner (1976). er tilt, and boreal wintertime perihelion coincided to Heat diffuses linearly through ice and snow, and the reduce summertime insolation in the Northern Hemi- thickness can change by melting at the surface and by sphere substantially below modern values. Because po- freezing or melting at the ice base. The bottom melt lar regions receive virtually all of their sunlight during rate is proportional to the underlying mixed layer tem- the summer half-year, these anomalous orbital con®g- perature, whereas the lateral melt rate is a function of urations produced considerable mean-annual Arctic ra- the temperature of the adjacent lead. During melting diation anomalies, especially at 115 kyr BP,which equal conditions a subgrid-scale ice thickness distribution is or exceed those from doubled atmospheric CO2 (IPCC applied, in which the probability density function of sea- 1990). ice thickness is assumed to be triangularly shaped be- These insolation perturbations appear to have forced tween zero and twice the mean thickness of the grid signi®cant changes in the climate of the high-latitude box. This kind of distribution agrees with observed Northern Hemisphere. Interpretations of various proxy thickness distributions from the Arctic (Tucker et al. climatic data for the mid-Holocene (around 6 kyr BP) 1992) and the Antarctic (Wadhams et al. 1987). Mini- depict an Arctic with considerably warmer air temper- mum lead fractions of 0.002 and 0.10 are imposed in atures (2±4 K), warmer SSTs, and reduced sea ice the Northern and Southern Hemisphere, respectively, (Stewart and England 1983; Frenzel et al. 1992; Koc et with an additional exponential reduction as ice thickens al. 1993). Quantitative comparisons for 115 kyr BP are above 1 m. Ice forms at a temporary thickness of 0.2 less comprehensive, but this time period marks the be- m, then combines with any existing ice in the grid box ginning of the most recent glacial inception, when a to produce a single ice thickness. The respective sea- major period of global ice buildup began (Shackleton ice albedos for visible and infrared wavelengths are 1987). Various estimates suggest that temperatures in ®xed at 0.8 and 0.7 under cold conditions (surface tem- the ice sheet source regions would have needed to be perature below 268 K) and decrease linearly toward 0.7 5±12 K colder than present (Rind et al. 1989). A com- and 0.4 as the surface temperature approaches the melt- plete data-model comparison for both of these paleo- ing point. The sea-ice code does not explicitly include climates is presented in section 6. surface melt ponds but does account for enhanced ther- These substantial environmental changes have pro- mal inertia from brine pockets within the ice. If the vided an impetus for GCMs to hindcast the early±mid-

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Holocene, including 9 kyr BP, when Arctic insolation house gases were required to achieve widespread ter- anomalies were even higher than at 6 kyr BP. Kutzbach restrial ice cover. and Gallimore (1988) simulated signi®cantly reduced This paper describes the results of a conservative ap- Arctic sea-ice area and thickness at 9 kyr BP, along with proach toward simulating the high-latitude climates of warmer temperatures in every season poleward of 60ЊN. the Northern Hemisphere at 6 kyr and 115 kyr BP, in Similar ®ndings were reported in the 9 kyr BP simu- that insolation and sea-ice motion are the only variables lation of Mitchell et al. (1988), who found that the re- that are altered. Combined changes in CO2, SSTs, and sidual Laurentide ice sheet caused signi®cant cooling vegetation would probably amplify the responses sim- downstream. Because this cooling effect interferes with ulated here, but the purpose of this procedure is to iso- attempts to isolate the orbital forcing signal, many stud- late the effects of orbital forcing and sea-ice dynamics ies have opted to concentrate instead on the mid-Ho- in driving climate changes in the Arctic. As such, a pair locene, when the impact of the Laurentide had become of paleoclimate experiments were performed for both 6 minimal. Simulations by Kutzbach et al. (1991) showed kyr and 115 kyr BP: one with ice transport (6KDI, signi®cant decreases in Arctic sea-ice cover (more than 115KDI) and one without (6KTI, 115KTI). The simu- 1 m annually in the central basin) and warmer air tem- lations were run for 20 yr, consisting of a 10-yr spinup peratures poleward of 60ЊN during summer (1±4 K) and and a 10-yr equilibrium interval. The ®nal decade of annually (1±3 K). Foley et al. (1994) simulated a similar the four simulations was used in the analysis, during annual temperature change (1.8 K) for the orbital anom- which time Arctic climatic variables showed no trend. aly alone and ampli®ed warming (3.4 K) when imposed The results of these paleoclimate runs were compared vegetation feedbacks were included. Liao et al. (1994) with 15-yr long modern simulations with sea-ice motion also report high-latitude warming, but only show results (MODDI) and without (MODTI), the ®nal 5 yr of which for summertime. All of the aforementioned modeling were available for analysis (the small interannual var- studies used coupled atmosphere±mixed layer ocean iability in the model permits the use of such a short models with interactive but nondynamic sea ice. The sampling interval). results from other recent 6 kyr BP simulations (Hewitt and Mitchell 1996; Hall and Valdes 1997) are not pre- 4. Modern simulation sented here, because their use of prescribed SSTs and sea ice short-circuit important feedbacks that could a. Comparison with observations cause much stronger high-latitude climate changes. Many of the characteristics of GENESIS2's simula- Numerous modeling efforts have also been directed tion of the modern Arctic climate with ice transport are toward simulating glacial inception around 115 kyr BP. described by Maslanik et al. (1996), so only an overview Using insolation as the only altered boundary condition, is given here. In general the model does a credible job Royer et al. (1983) simulated a mean annual cooling of of reproducing observed Arctic ®elds and shows sig- over 2 K and increased precipitation across the hypoth- ni®cant improvement over most other GCM simula- esized region of ice sheet formation in Canada, although tions, which traditionally have had dif®culty in simu- they were unable to generate permanent snow cover. lating this region (Walsh and Crane 1992). The follow- Rind et al. (1989) altered several boundary conditions ing are the major errors in GENESIS2's simulated Arctic (insolation, atmospheric CO2, sea-ice extent, and SSTs) climate: SLP gradients are too weak over the Arctic in an attempt to initiate ice sheets. Only when all of Ocean and pressures are too high during summer, sea these cooling mechanisms were invoked did the model ice is too thin, sea-ice concentrations are too low during simulate permanent snow cover, and even then it oc- summer, and ice pack margins are too diffuse (Maslanik curred only over a restricted region of Baf®n Island. et al. 1996). Syktus et al. (1994) also reported dif®culty in simulating The pattern of simulated mean-annual air temperature glacial inception; orbital changes had to coincide with (Fig. 3) matches reasonably well with measured values, lower CO2 in order for July snow cover to greatly ex- as the model captures the coldest overall conditions pand. Dong and Valdes (1995) were able to simulate (Ϫ16Њ to Ϫ18ЊC) north of Greenland and the Canadian perennial snow cover in the Canadian Archipelago at Archipelago and also simulates the depressed temper- 115 kyr BP with only insolation reductions, and the atures off the east Greenland coast, where sea ice im- model was able to support a permanent snowpack in parts a local chilling effect. Large cold model biases are various high-latitude areas when CO2 and SSTs were evident in northern Eurasia, whereas smaller ones ap- lowered. Gallimore and Kutzbach (1996) used exag- pear in the Norwegian Sea and over most of the Arctic gerated insolation forcing for this time period and found Ocean. that glacial inception could be achieved only with re- The sea-ice coverage at the end of the melt season duced CO2 and a prescribed expansion of tundra veg- (Fig. 3) is adequately simulated in the convergent ice etation. Schlesinger and Verbitsky (1996) found that region north of Greenland and the Canadian Archipel- orbital forcing alone was insuf®cient to initiate North ago, as well as off the east coast of Greenland. The American ice sheets; reduced concentrations of green- model's marginal ice zone around Svalbard and Franz

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FIG. 3. (a) Mean-annual surface air temperature (ЊC) and (b) late-summer sea-ice fraction observed (left) and simulated (right) [observed air temperatures are from NCEP reanalysis (Kalnay et al. 1996); observed sea ice fraction is from NASA (1984)].

Josef Land is somewhat too far poleward, however, and The pattern of sea-ice drift is largely responsible for there is too much open water off the northern coasts of the seasonal maximum and minimum ice thickness dis- Alaska, northwestern Canada, and Siberia. tributions (Figs. 2 and 5), which feature a buildup of The model does a remarkably good job of reproducing ice north of Greenland and the Canadian Archipelago, the observed sea ice drift (Figs. 1 and 4), as it correctly as well as a tongue off eastern Greenland containing 1± simulates the clockwise circulation in the Beaufort Gyre, 2-m-thick sea ice in both the simulations and obser- the general ¯ow of ice from northern Eurasia toward vations. The model reproduces the ice margin in the North America, and the ice export through Fram Strait. northern Greenland Sea during winter, but shifts it too Maximum mean annual ice velocity in the Fram Strait far poleward during late summer. Except in the con- is 0.12 m sϪ1, in good agreement with measurements vergent ice region north of North America, the modeled (Barry et al. 1993). sea-ice thicknesses are noticeably underestimated at

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FIG. 4. Mean-annual simulated sea-ice drift (m sϪ1) for modern conditions. both times of the year. The mean annual ice thickness ness, and SLP. The ice transport parameterization causes in the central Arctic (80Њ±90ЊN) is 2.2 m, compared to reduced ice coverage from the Laptev Sea westward to conventional estimates of around 3 m (Bourke and Gar- Svalbard, due to the mean offshore ice drift along the rett 1987). Eurasian shelf. Maximum open water anomalies occur The distribution of sea level pressure (SLP) is sim- in the Svalbard±Franz Josef Land region, due to strong ulated quite well (Fig. 6). The model captures the ice divergence caused by the higher ice speeds in the strength and location of the Beaufort high, while re- East Greenland Current (Fig. 4). Much greater ice cov- producing the tongue of low pressure that penetrates the erage appears just east of Greenland, where only a min- Arctic from the Norwegian Sea. Not only is this realistic imal amount of sea ice grows thermodynamically, and SLP pattern a necessary ingredient for a satisfactory these higher ice concentrations extend equatorward climatology, but the resulting surface wind ®eld is large- along the southern Greenland coasts. ly responsible for producing credible sea-ice velocities. The ice coverage differences strongly affect the sur- face air temperatures (Fig. 7). Much warmer conditions prevail in the vicinity of reduced ice concentrations, b. Comparison with and without ice motion from Scandinavia northeastward into the Arctic Ocean, A detailed comparison of the modern simulation with and cover most of the Arctic Basin. Conversely, the and without ice transport is given by Maslanik (1997), excessive ice coverage near the Greenland coast forces so only a brief summary is given here. Figure 7 shows colder readings on- and offshore of southern Greenland. a composite of difference maps (MODDI Ϫ MODTI) The differences in air temperature and ice drift between for mean-annual values of ice fraction, surface air tem- MODDI and MODTI cause generally thinner ice over perature, midtropospheric geopotential height, ice thick- most of the interior Arctic Ocean and thicker ice north

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FIG. 5. Simulated Arctic sea-ice draft (cm) during (a) summer and (b) winter. and east of Greenland. The surface temperature differ- ences propagate aloft into the middle troposphere, caus- ing differences in geopotential height almost directly FIG. 6. Mean-annual sea level pressure (mb) over the Arctic (a) above the core surface anomalies. These anomalous observed and (b) simulated [observations are from Walsh et al. thickness patterns aloft appear to force SLP anomalies, (1996)]. resulting in lower pressure to the east of the anomalous trough around Greenland and higher pressure down- (60Њ±80ЊN), and lower Arctic land (60Њ±80ЊN) to de- stream of the anomalous ridge over the Barents and Kara termine whether ice motion signi®cantly changes the Seas. climatic response over the range of 6±115 kyr BP forc- ings. The variables chosen, which are deemed repre- 5. Climate sensitivity in paleoclimate simulations sentative of the ice±atmosphere system in the high-lat- itude Northern Hemisphere, include sea-ice descriptors a. Sensitivity over the range of 6 kyr to 115 kyr BP (ice thickness and concentration) and atmospheric de- Statistical sensitivity tests are performed separately scriptors (surface and lower-tropospheric air tempera- on regionally averaged climatic variables over the cen- tures, SLP,precipitation, snow depth, cloud fraction, and tral Arctic Ocean (80Њ±90ЊN), the lower Arctic Ocean incident solar energy at the surface). More rigorous anal-

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FIG. 7. Mean-annual differences between the modern simulation with and without dynamic sea ice (MODDI Ϫ MODTI). Variables are (a) sea-ice fraction, (b) surface air temperature (K), (c) sea-ice thickness (cm), (d) geopotential height (m) at the ␴ ϭ 0.695 model level, and (e) SLP (mb).

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TABLE 1. Statistically signi®cant (95% level) seasonal [Mar±May (MAM), etc.] and annual (ANN) changes across the range of thermal forcing between the versions with and without ice motion [(6KDI± 115KDI) minus (6KTI±115KTI)]. A positive (negative) sign indicates signi®cantly enhanced (reduced) sensitivity with dynamic sea ice. MAM JJA SON DJF ANN Central Arctic Ocean (80Њ±90ЊN) Ice thickness Ϫ Ϫ Ϫ Ϫ Ϫ Ice concentration ϩ Ϫ Ϫ Ϫ Air temperature Ϫ Ϫ Ϫ Ϫ 970-mb temperature Ϫ 929-mb temperature Precipitation Snow depth Cloud fraction Incident solar energy Ϫ Sea level pressure Lower Arctic Ocean (60Њ±80ЊN) Ice thickness Ϫ Ϫ Ϫ Ϫ Ϫ Ice concentration ϩ Ϫ Ϫ ϩ Air temperature 970-mb temperature 929-mb temperature Precipitation Snow depth Ϫ Cloud fraction Incident solar energy Ϫ Sea level pressure

volume, whereas the 6 kyr BP simulations show a win- tertime drop of up to 25% and late-summer decreases of more than 50%. Table 1 summarizes the prevalence and type of sta- tistically signi®cant (95% con®dence level) differences produced by ice motion, as given by a Student's t-test of the difference between the magnitude of response from 6 kyr to 115 kyr BP with ice dynamics included and without (i.e., |6KDI Ϫ 115KDI| Ϫ |6KTI Ϫ 115KTI|). Positive (negative) values of this difference imply that ice motion signi®cantly enhances (reduces) sensitivity and thus acts as a positive (negative) net FIG. 8. Annual cycle of simulated Northern Hemisphere sea-ice (a) area and (b) volume in the modern, 6 kyr BP, and 115 kyr BP ex- feedback over this range of thermal forcing, although periments with dynamic sea ice. the sign and magnitude of the feedback may differ be- tween the 6 kyr and 115 kyr BP departures from modern. Regionally averaged, sea-ice dynamics act as a predom- ysis is performed on those variables that showed sig- inantly negative feedback in both the central Arctic and ni®cant changes, in order to isolate the sign of the feed- lower Arctic basins and largely con®ne their impact to back and whether sea-ice dynamics are more important the ice variables; only near-surface air temperatures in in the warm or the cold scenario. the central Arctic show a noteworthy sensitivity. The The insolation deviations at 6 kyr and 115 kyr BP only positive feedbacks are associated with ice concen- are big enough to induce large changes in the cryo- tration, and even they are limited to spring and winter. sphere, as illustrated by the annual cycles of hemispheric Despite their infrequency, however, these positive feed- sea-ice area and volume in the experiments with dy- backs will be shown below to initiate numerous im- namic sea ice (Fig. 8). Changes in wintertime ice area portant atmospheric anomalies, particularly in the cold from modern are rather modest, but the late-summer scenario (115 kyr BP). Virtually no signi®cant changes minimum is over 40% larger at 115 kyr BP and is re- were induced by ice motion over lower Arctic land. duced by about 25% at 6 kyr BP. Sea-ice volume de- For those variables that showed signi®cant sensitivity partures from modern are considerably more pro- to ice motion, three climate change parameters (␭R, ␭ϩ, nounced: at 115 kyr BP wintertime increases are 50%± and ␭Ϫ) are calculated. Here ␭R indicates whether the 100% and there is nearly a tripling of late-summer ice sensitivity is relatively greater for the cold or warm

Unauthenticated | Downloaded 10/01/21 10:31 PM UTC MARCH 1999 VAVRUS 883 forcing, ␭ϩ gives the sign and strength of the feedback from their modern values. Because identical radiative due to ice dynamics for the warm forcing, and ␭Ϫ ex- perturbations were used in both model versionsÐthat presses the sign and magnitude of the feedback for the is, ⌬Q 6KDI ϭ⌬Q 6KTI and ⌬Q115KDI ϭ⌬Q115KTIÐthe feed- cold scenario. Here ␭R is de®ned as back parameters derived here are valid over any tem- poral and spatial domain. (V Ϫ V ) Ϫ (V Ϫ V ) ␭R ϭ 6KDI MODDI 6KTI MODTI , (1) Sea-ice motion acts almost exclusively as a negative (V 115KDIϪ V MODDI) Ϫ (V 115KTIϪ V MODTI) feedback with respect to ice thickness in the interior and where V is the average of the variable under consid- periphery of the Arctic Ocean in both the 6 kyr and 115 ϩ Ϫ eration and the subscripts refer to the particular exper- kyr BP simulations (␭ and ␭ Ͻ 1), with a larger impact R iment. The terms from the modern forcing are included in the central Arctic in the cold scenario (|␭ | Ͻ 1) and R to account for any linear model biases that exist between farther south in the warm scenario (|␭ | Ͼ 1) (Table 2). the MODDI and MODTI simulations. If the averages These results are probably caused by a combination of are identical in these two control runs, then (1) reduces effects: the net divergence in a moving ice pack causes to the ratio of the change from the DI to the TI version a mean equatorward ¯ow and thus thins the interior sea with 6 kyr BP forcing applied to the change with 115 ice, the residence time in the rapid ice growth region kyr BP forcing applied. Absolute values of ␭R greater of the central Arctic is limited with ice transport, and (less) than one imply that the variable is relatively more the easier ridging of thin ice helps to counteract the (less) sensitive to dynamic sea ice when the warm orbital thermodynamic melting. The ®rst two of these processes forcing anomaly is applied. The feedback parameters, appear to dominate in the 115 kyr BP experiment, ␭ϩ and ␭Ϫ, are used to identify whether ice dynamics whereas the third seems to be most important at 6 kyr induce a feedback whose sign and strength depend on BP. the sign of the external perturbation (orbitally forced). The effect of ice transport on ice concentration is The feedback parameters were derived from the in- more complicated than its in¯uence on ice thickness verse of the standard climate feedback parameter, ␭, (Table 2). In the warm scenario dynamic sea ice acts as which is rede®ned in its inverted form as a negative feedback during the warmest two seasons (summer and autumn), but as a slightly negative or pos- ⌬T itive feedback during the colder seasons. With a cold ␭ ϭ , (2) ⌬Q orbital forcing, however, ice motion produces an en- hanced response year-round in the central Arctic but where ⌬T is the temperature departure and ⌬Q is the only during the cold season in the lower Arctic. There initial forcing perturbation at the troposphere, such as is also a pronounced seasonal variation in the magnitude an orbitally induced change. For this analysis the meth- of the feedback between the two scenarios, as the ice od used to determine the sign of the feedback parameters concentration sensitivity over the entire orbital range is ϩ Ϫ ␭ and ␭ is similar to the technique of Cess et al. dominated by the positive 115 kyr BP feedback in spring (1991). They used the ratio of two feedback parameters and winter but by the negative 6 kyr BP feedback an- to represent various simulated responses to a SST per- nually and during summer and autumn. As explained turbation with (␭) and without (␭S) an interactive snow below, these temporally and spatially varying responses cover scheme in the models. A ␭/␭S ratio greater (less) are due to the competing effects of ice dynamics in than one indicates a positive (negative) snow feedback. extending the ice margin equatorward while generating Likewise, in this analysis the ratio of the feedback pa- leads within the pack. rameters obtained from the DI and TI model versions The nature of the feedbacks produced by ice dynamics ϩ Ϫ are used to de®ne ␭ and ␭ as is clari®ed by examining the annual cycle of the dif- ference in sea-ice volume and area between the TI and V Ϫ V 6KDI MODDI DI versions in the warm and cold scenarios (Fig. 9). ΂΃⌬Q6KDI Both the hemispheric and central Arctic time series of V 6KDIϪ V MODDI ␭ϩ ϭϭ sea-ice volume consistently show increases (decreases) V 6KTIϪ V MODTI V 6KTIϪ V MODTI in ice volume at 6 kyr (115 kyr) BP upon the inclusion ΂΃⌬Q6KTI of ice motion, corresponding to a persistent negative feedback. The effect of ice transport on sea-ice area, on

V 115KDIϪ V MODDI the other hand, is seasonally dependent and varies be- tween the warm and cold scenarios. During winter and ΂΃⌬Q115KDI V Ϫ V ␭Ϫ ϭϭ115KDI MODDI , (3) spring, ice dynamics extend the hemispheric coverage V 115KTIϪ V MODTI of sea ice under both warming and cooling perturba- V 115KTIϪ V MODTI tions, whereas the absence of changes in the central ⌬Q ΂΃115KTI Arctic indicates that this hemispheric expansion is al- where V has been generalized to be the average of any most exclusively due to an advective, equatorward ad- variable, not just temperature, and ⌬Q is the insolation vance of the ice margin. During the summer and autumn, perturbation induced by the change in orbital parameters however, the pattern generally reverses and ice motion

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TABLE 2. Seasonal and annual values in each experiment of variables showing signi®cant sensitivity to ice motion across the range of thermal forcing from 6 kyr to 115 kyr BP. Also shown are the climate sensitivity parameters. MAM JJA SON DJF ANN Central Arctic Ocean (80Њ±90ЊN) Ice thickness (m) 6KDI 2.22 1.79 1.02 1.56 1.65 115KDI 4.40 4.52 4.06 4.14 4.28 6KTI 1.79 1.01 0.41 1.18 1.11 115KTI 6.10 6.06 5.86 5.90 5.98 MODDI 2.63 2.39 1.70 2.12 2.21 MODTI 3.07 2.71 2.25 2.66 2.67 ␭ϩ 0.32 0.35 0.37 0.38 0.36 ␭Ϫ 0.58 0.64 0.65 0.62 0.63 ␭R Ϫ0.69 Ϫ0.90 Ϫ0.93 Ϫ0.75 Ϫ0.81 Ice concentration 6KDI 0.9820 0.8080 0.6520 0.9720 0.8570 115KDI 0.9860 0.8820 0.8950 0.9950 0.9380 6KTI 0.9980 0.8810 0.6610 0.9790 0.8830 115KTI 0.9996 0.9996 0.9996 0.9996 0.9996 MODDI 0.9835 0.8408 0.7941 0.9887 0.9016 MODTI 0.9989 0.9885 0.9580 0.9986 0.9860 ␭ϩ 1.67 0.31 0.48 0.85 0.43 ␭Ϫ 3.57 3.71 2.42 6.30 2.68 ␭R Ϫ0.33 2.48 2.61 0.55 2.56 Air temperature (ЊC) 6KDI Ϫ20.28 0.88 Ϫ9.11 Ϫ27.07 Ϫ14.20 115KDI Ϫ21.25 Ϫ0.88 Ϫ15.91 Ϫ31.75 Ϫ17.02 6KTI Ϫ20.89 0.75 Ϫ9.97 Ϫ27.76 Ϫ14.80 115KTI Ϫ23.67 Ϫ1.00 Ϫ18.31 Ϫ34.60 Ϫ18.93 MODDI Ϫ19.37 0.12 Ϫ11.97 Ϫ28.54 Ϫ14.88 MODTI Ϫ22.29 0.11 Ϫ15.05 Ϫ31.75 Ϫ17.17 ␭ϩ Ϫ0.65 1.19 0.56 0.37 0.29 ␭Ϫ 1.36 0.90 1.21 1.13 1.22 ␭R 4.63 1.09 3.26 7.00 4.45 Lower Arctic Ocean (60Њ±80ЊN) Ice thickness (m) 6KDI 1.24 0.71 0.32 0.78 0.77 115KDI 2.07 1.70 1.37 1.72 1.71 6KTI 1.03 0.45 0.11 0.59 0.55 115KTI 2.01 1.71 1.46 1.73 1.73 MODDI 1.31 0.81 0.47 0.90 0.87 MODTI 1.25 0.75 0.44 0.87 0.83 ␭ϩ 0.32 0.33 0.45 0.43 0.36 ␭Ϫ 1.00 0.93 0.88 0.95 0.93 ␭R ϱ Ϫ2.86 Ϫ1.50 Ϫ4.00 Ϫ3.00 Ice concentration 6KDI 0.659 0.366 0.217 0.556 0.454 115KDI 0.703 0.470 0.423 0.639 0.556 KTI 0.641 0.403 0.190 0.538 0.447 115KTI 0.655 0.550 0.471 0.590 0.566 MODDI 0.670 0.407 0.306 0.589 0.493 MODTI 0.652 0.473 0.336 0.565 0.507 ␭ϩ 1.00 0.59 0.61 1.22 0.65 ␭Ϫ 11.00 0.82 0.87 2.00 1.07 ␭R 0.00 Ϫ2.07 Ϫ3.17 Ϫ0.24 5.25

usually causes reduced ice cover in the interior Arctic fectively ridging the thinning ice into a thicker and more and across the hemisphere, suggesting that the process compact ice cover. of open-water formation through the creation of leads A summary of the integrated seasonal climatic var- within the ice pack is relatively more important during iations induced by mobile sea ice is provided in Fig. this time of year. One important exception to this pattern 10, which shows mean annual anomalies in the central is the positive late-summer anomaly at 6 kyr BP, which Arctic Ocean of variables that displayed signi®cant sen- is probably a result of net equatorward ice ¯ow ®lling sitivity to ice motion (similar but less pronounced pat- in thermodynamically created open-water areas and ef- terns of climate sensitivity occur in the lower Arctic).

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FIG. 9. Annual cycle of the difference in the simulated sea-ice (a) volume and (b) area between the paleoclimate experiments with and without ice dynamics for the Northern Hemisphere (left) and central Arctic (right).

The simulated decrease in ice thickness at 6 kyr BP is alies simply thin or thicken the pack, while contracting about two-thirds less with dynamic ice than with the or expanding the ice margin. These same thermody- thermodynamic-only version. With the 115 kyr BP forc- namic tendencies also occur in a dynamic ice pack, but ing, ice motion reduces the gain by more than one-third. the superimposed mechanical effects can reinforce or A similarly strong negative feedback is evident under mitigate these tendencies. Because sea ice resists com- the warm forcing with respect to ice concentration (or pression more than divergence, even in the absence of area), as the inclusion of ice transport cuts the decrease a net vorticity forcing by the atmosphere an ice pack by over half, whereas ice motion more than doubles the will tend to spread equatorward (Hibler 1984), causing increased coverage under the cold orbital forcing. These thinning in the interior. If the ice is thin enough, how- sea-ice differences translate into changes in surface air ever, this mechanically induced thinning can be coun- temperature, which follow more closely with ice con- teracted by easier ridging of the ¯oes. In addition, ther- centration anomalies and are particularly striking at 6 modynamically forced opening and closing of leads can kyr BP. Mean annual surface air temperatures in the be enhanced or reduced by ice motion within the pack, central Arctic rise by more than 2 K with the thermo- depending on the particular pattern of ice velocity within dynamic-only model but by only 0.7 K when a mobile the interior. ice pack is added. The extra ice concentration in the 115 kyr BP scenario with dynamic sea ice enhances the b. Spatial patterns cooling by almost half a kelvin. The physical mechanisms responsible for the different In addition to the statistical analyses that consider climate sensitivities between a thermodynamic-only ice changes over distinct geographic domains, it is also in- pack and a mobile one are presented schematically in structive to examine the variations among the experi- Fig. 11. Without ice motion, atmospheric thermal anom- ments in a two-dimensional context to pinpoint locally

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within the Laptev and east Siberian Seas (Fig. 12). This increase in open water is large enough to warm the overlying air in these areas by more than 1 K, with much greater warmings in Greenland, presumably due to its exposure to warmer, moister regional conditions. There is considerable thinning of ice only along northern Greenland and the Canadian Archipelago, where the ice in the modern simulation is the thickest. The surface air temperature anomalies are transmitted into heating anomalies aloft, causing in¯ated midtropospheric heights around Greenland. The SLP distribution is also modi®ed, in the form of a weaker Aleutian low and anomalous low pressure over northern Eurasia, both of which are statistically signi®cant (90% level). As expected, the anomaly maps for the 115 kyr BP scenario show substantial cooling and enhanced sea-ice coverage (Fig. 13). A dramatic change is the increased sea-ice cover around the east coast of Greenland and its expansion east of Iceland. The ice pack might have spread even further if not for the parameterization that effectively prevents the Norwegian Sea from reaching the freezing point (section 2). The sea-ice margin from Greenland to Novaya Zemlya also shifts southward and in combination with the aforementioned ice advance causes enhanced cooling over the Greenland and Iceland Seas and from northern Greenland to northern Russia. Corresponding to the large-scale cooling across the en- tire Arctic, sea-ice thickens up to several meters near the Canadian Archipelago, whereas midtropospheric heights drop dramatically, especially in the vicinity of the largest ice concentration anomalies. There are also noteworthy changes in the SLP ®elds, consisting of gen- eral pressure rises over the Arctic Ocean, a weakening of the Icelandic low, and decreased pressure over north- eastern Europe. The causes of these surface pressure changes (signi®cant at the 90% level) are dif®cult to diagnose. The increased atmospheric pressure over the Arctic Ocean is probably a hydrostatic effect caused by FIG. 10. Departures from modern in the paleoclimate simulations of central Arctic (a) sea-ice thickness, (b) sea-ice concentration, and surface cooling, whereas the dipole high±low pattern (c) surface air temperature. from Greenland to northern Russia is probably not, since it is reminiscent of the 6 kyr BP anomaly pattern, even though the ice coverage and surface temperature anom- alies were nearly opposite between 6 kyr and 115 kyr sensitive regions and forcing mechanisms. Figures 12± BP. This discrepancy suggests that upper-air dynamical 15 show annually averaged difference maps of several forcing may be strongly in¯uencing the surface pressure variables (sea-ice concentration and thickness, surface anomalies. Observational studies by LeDrew (1983, air temperature, midtropospheric geopotential height, 1985) demonstrate that surface cyclones and anticy- and SLP) for the following four comparisons: 6KDI Ϫ clones over polar regions are more strongly affected by MODDI, 115KDI Ϫ MODDI, (6 KDI Ϫ MODDI) Ϫ large-scale vorticity and thickness advection than by the (6KTI Ϫ MODTI), and (115KDI Ϫ MODDI) Ϫ hydrostatic response to diabatic heating at the surface. (115KTI Ϫ MODTI). The ®rst two sets of maps isolate Comparisons between the MODDI and 6KDI or the impact of orbital forcing with the improved sea-ice 115KDI simulations capture the effect of altered orbital parameterization, whereas the last two sets of maps iso- parameters in conjunction with more complete model late the effect of ice dynamics under both the warm and physics, but they fail to isolate the impact of dynamic cold atmospheric forcing scenarios. sea ice. One can, however, extract the importance of ice Under the warm orbital forcing (6KDI Ϫ MODDI) motion in each paleoclimate experiment by examining there is a pronounced reduction in ice cover along the the simulated departures from modern with ice transport eastern and southern coasts of Greenland, as well as compared to those without. Such a comparison for the

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FIG. 11. Schematic diagram in a y±z plane illustrating the differences between a thermodynamic-only sea-ice pack (left) and one that includes ice dynamics (right). Solid black arrows indicate the direction of ice response to warm (W) and cold (C) perturbations; open arrows show the effect of ice dynamics (D). The pack with ice dynamics is thinner and more extensive due to net equatorward motion and contains openings in the pack (leads). Warm (cold) perturbations in either pack cause thinning (thickening) of ice accompanied by a contraction (expansion) of the ice margin while widening (narrowing) the leads. Ice motion opposes the ice margin contraction (``1'') in a warming scenario (negative feed- back) but reinforces the expansion in a cooling scenario (positive feedback). Dynamical effects can either enhance or retard the thermal forcing on ice thickness (``2''), depending on how thick the ice is: thin ice is more prone to ridging and thus ice motion opposes the thermal perturbation, whereas a thermally induced thickening is offset by equatorward spreading of the pack. Motion within the pack can either shrink or expand leads (``3''), the magnitude of which depends on how thick and compact the pack is: a thin, open pack diverges and converges easily; a thick, compact pack is more cohesive. Thus the mechanical forcing of openings within the pack is relatively more important for thin, diffuse ice packs.

6 kyr BP simulations is illustrated in Fig. 14, which region. The midtropospheric height expression of the shows the difference between 6KDI Ϫ MODDI Ϫ 6KTI surface temperature anomalies consists of enhanced Ϫ MODTI (thus, if the patterns were identical in the ridging in the southern Greenland±Labrador Sea area two control runs, then these maps simply show 6KDI and widespread enhanced troughing from northeastern Ϫ 6KTI). The inclusion of ice dynamics enhances melt- Greenland into northern Eurasia. These height anoma- off along the southern coasts of Greenland but mitigates lies aloft are consistent with upper-air forcing of SLP the meltback north of Svalbard, along the northern half departures, whose major features are a weakening of the of Greenland's eastern coast, and within the Beaufort Icelandic low and anomalous low pressure over most Sea. This result can be understood to be a result both of northern Eurasia. of differences between the modern control simulations The difference maps of (115KDI Ϫ MODDI) Ϫ (MODDI and MODTI) and feedbacks introduced by the (115KTI Ϫ MODTI) elucidate the in¯uence of dynamic ice motion. For example, because more sea ice exists sea ice under a negative atmospheric thermal anomaly along the southern coast of Greenland in the MODDI (Fig. 15). The most prominent effect of ice transport in simulation, there is more to melt off when a warm at- this scenario is to produce excessive sea ice around the mospheric perturbation is applied. On the other hand, Denmark Strait and Iceland Sea, as discussed earlier. In the tendency of ice motion to spread the pack equator- addition, relatively less ice coverage occurs in the Fram ward counteracts the thermally induced poleward con- Strait and along the northern half of the east coast of traction and results in a much more modest meltback Greenland, apparently due to higher lead fractions in- along the ice pack margins. These differences in ice duced by the dynamic ice pack. The more important coverage anomalies between the model versions with difference is the excessive ice cover equatorward, which and without ice motion cause corresponding signi®cant forces colder near-surface conditions in situ and down- differences in the surface temperature ®elds, creating stream over the northeastern Atlantic and into northern relatively cooler conditions over most of the Arctic, Europe. The effect of ice transport on the ice thickness especially from northeastern Greenland to northern Rus- ®elds is striking; one can clearly surmise the mean ice sia, but somewhat warmer readings around southern drift from the dipole pattern of anomalously thin ice in Greenland and the Labrador Sea. The ridging effect of the divergent sea-ice zone of the Arctic Ocean and rel- ice transport in mitigating thermodynamic vertical melt- atively thick ice in the convergent portion. As a result ing is also clearly seen, as the ablation in the convergent of the surface heating anomalies there is an eastward sea ice zone is offset by more than 1 m in much of the displacement of the midtropospheric trough, which fa-

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FIG. 12. Mean-annual differences between the 6 kyr BP and modern simulation with dynamic sea ice (6KDI Ϫ MODDI). Variables are as in Fig. 7.

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FIG. 13. Mean-annual differences between the 115 kyr BP and modern simulation with dynamic sea ice (115KDI Ϫ MODDI). Variables are as in Fig. 7.

Unauthenticated | Downloaded 10/01/21 10:31 PM UTC 890 JOURNAL OF CLIMATE VOLUME 12 vors a weakened Icelandic low and lower pressure over whereas the model generates a much more modest northern Russia due to the upper-air forcing described warming of generally less than 1 K (Fig. 12), especially earlier. over the Arctic Ocean, where the negative feedbacks from the dynamic sea-ice code have the most impact. The only area with a simulated temperature increase of 6. Paleodata-model comparisons more than 2 K is Greenland. The performance of the model can be evaluated by The simulated sea-ice changes are also smaller than its ability to reproduce the climatic conditions of 6 kyr what is suggested by the paleodata. The 6 kyr BP sea- and 115 kyr BP inferred from paleoenvironmental data. ice reconstructions by A. deVernal (1997, personal com- As might be expected, more records are available for munication) show 4±5 months less ice cover (Ն50% the mid-Holocene period, although there are data from areal coverage) than present in Hudson Bay and 3 the earlier time that indicate a substantially cooler Arc- months less in the Labrador Sea. Conversely, the model tic. A description of the records and their climatic im- simulates only a 1-month decrease in Hudson Bay and plications is provided in the text, tables, and maps in a 2-month reduction in Labrador Sea. Similarly, the this section. model produces no statistically signi®cant decrease in summertime ice concentration around the Canadian Ar- chipelago, nor along the coasts of Svalbard, despite ev- a. 6 kyr BP (data) idence to the contrary. The climatic estimates for this time period are based on surface air temperature and sea-ice coverage inferred c. 115 kyr BP (data) from proxy ®eld data from a number of independent sources described below. Comprehensive reconstruc- 115 kyr BP is generally regarded as the beginning of tions of temperature appear in Frenzel et al. (1992), who the most recent glacial period, based primarily on co- used palynological transfer functions to derive mean- inciding oxygen±isotope changes in marine sediment annual temperature departures from present. These maps cores. A period of global ice sheet buildup is believed indicate warming of up to a few degrees over most of to have taken place around this time, when the Northern the Arctic, in agreement with other pollen data indi- Hemisphere's ice volume increased to about half its peak cating a mid-Holocene warm period in northern Canada, glacial maximum value from 120 to 110 kyr BP (Shack- Greenland, Siberia, and Scandinavia (Funder 1989; leton 1987). Pollen records from northern Europe sug- Guiot et al. 1993; Koshkarova 1995; Williams et al. gest that air temperatures dropped rapidly around 115 1995). Other proxy data generally consistent with these kyr BP, accompanied by signi®cant vegetation changes reconstructions include isotope records from Arctic ice (Woillard 1978). Data from the Greenland ice cores in- cores, lake and coastal microfossils, and peat deposits dicate temperatures 10 K colder than present on the (e.g., Edlund 1986; Funder and Weidick 1991; Polyak summit at 113 kyr BP (Johnson et al. 1995). and Solheim 1994; Johnson et al. 1995). Records from the North Atlantic Ocean also suggest In addition to these temperature proxies, there are widespread cooling at this time. Evidence from coc- ®eld data from various sources that imply reduced mid- coliths and planktonic foraminifera indicate a rapid Holocene sea-ice cover over a wide region of the North- southward shift of the polar front (Ruddiman and ern Hemisphere (Table 3). Unfortunately, many of these McIntyre 1977; Larsen and Sejrup 1990), in conjunction proxies are qualitative and therefore cannot be used di- with a sudden SST drop of about6KintheNorwegian rectly for data±model comparisons. However, taken to- Sea and Denmark Strait and an end to ice-free conditions gether and in concert with the temperature reconstruc- in the Greenland±Norwegian Seas (Kellogg 1980; Lar- tions, these data strongly suggest that there were sig- sen and Sejrup 1990). High-latitude penetration of tem- ni®cant reductions in Arctic sea ice cover. Work is under perate Atlantic water appears to have ceased, as the way at quantifying past sea-ice changes over a larger North Atlantic drift stopped reaching the Norwegian and domain, including the central Arctic basin (A. deVernal Greenland Seas (Kellogg 1977) and the seasonally ice- 1997, personal communication). free Fram Strait shifted into a perennially ice-covered state (Hebbeln and Wefer 1997). Circumstantial evi- dence for the widespread advance of sea-ice cover is b. 6 kyr BP (model output) indicated from the slow accumulation along the east Consistent with paleoenvironmental evidence for a Greenland ice sheet starting around 115 kyr BP, sug- mid-Holocene Arctic warming, the model simulates sta- gestive of a reduced moisture source region (Funder tistically signi®cant climate changes that include warm- 1989). er conditions with reduced sea-ice cover and thickness at 6 kyr BP. The magnitude of these climate changes, d. 115 kyr BP (model output) however, is considerably smaller than what the ®eld data suggest. The data indicate mean-annual temperature The model results presented in section 5 are quali- anomalies of ϩ2toϩ4 K over most of the Arctic, tatively in agreement with the evidence for dramatic

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FIG. 14. Mean-annual differences in the 6 kyr BP departures from modern between the simulations with and without ice dynamics [(6KDI Ϫ MODDI) Ϫ (6KTI Ϫ MODTI)]. Variables are as in Fig. 7. cooling at 115 kyr BP (Fig. 13). Sea ice thickened sig- earlier, the temperature drops and sea-ice advance may ni®cantly and advanced equatorward in the DI version, have been greater had the Norwegian Sea been allowed accompanied by considerable mean-annual cooling of to freeze over). The colder Arctic, especially the 3±4- up to a few degrees in parts of the Arctic (as noted K summertime temperature decreases over northern

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FIG. 15. As in Fig. 14 but for 115 kyr BP.

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TABLE 3. Description of paleoenvironmental evidence for reduced ice cover during the mid-Holocene (around 6 kyr BP). Site Location Evidence Interpretation Reference 1 Ellesmere Island Algae Warmer with thinner lake ice around 5700 yr Blake (1989) BP 2 Ellesmere Island Driftwood More open water from 4500 to 6500 yr BP Blake (1972) 3 Ellesmere Island Molluscs Less sea ice (sub-Arctic mollusc species lived Stewart and England (1983) far north of present location at 6400 yr BP) 4 Ellesmere Island Whale tusk Receded ice shelves (narwhal tusk dated 6830 Evans (1989) yr BP found behind present ice shelf) 5 Ellesmere Island Driftwood Receded ice shelves (driftwood dated 6900± Crary (1960) 3000 yr BP found behind present ice shelves) 6 Ellesmere Island Raised beaches Reduced sea-ice cover along northern coast Evans (1988) 7 Queen Elizabeth Islands Whale and walrus More open water (bones dated 6500±7500 yr Harrington (1975) bones BP) 8 Arctic Ocean Ostracodes Assemblages suggest seasonal ice cover until Cronin et al. (1995) 4±5 kyr BP 9 Hudson Bay Dino¯agellate cysts Sea-ice cover reduced by 4±5 months yrϪ1 A. deVernal, pers. comm. (1997) 10 Labrador Sea Dino¯agellate cysts Sea-ice cover reduced by 3 months yrϪ1 A. deVernal, pers. comm. (1997) 11 Canadian Archipelago Driftwood More open water (driftwood peak from 6000 Dyke and Morris (1990) to 3500 BP) 12 Svalbard Driftwood More open water around 5±7 kyr BP Haggblom (1982) 13 East Greenland coast Diatoms Retreat of sea ice along east Greenland coast Koc et al. (1993) around 5±7 kyr BP 14 Ellesmere Island Driftwood Reduced summertime sea-ice cover from Stewart and England (1983) 6000 to 4200 BP 15 Greenland Ice Sheet Glacial deposits Receded ice sheet margins Funder (1989)

Canada, create reduced snow-free conditions (not to paleoenvironmental reconstructions, which show shown) that are favorable for glacial inception, yet no mean-annual Arctic temperatures 2±4 K warmer than terrestrial snow cover survives the summers. In addition, present and considerably less sea ice. This discrepancy nothing like a 10-K cooling over central Greenland oc- for 6 kyr BP, in tandem with the absence of a simulated curs in the model; mean-annual surface air temperatures glacial onset at 115 kyr BP, suggests that orbital forcing across the interior ice sheet are only about 1 K cooler alone is not suf®cient to initiate the observed climate than in the modern simulation. Thus, although the sim- anomalies at these times, even with a more realistic ulation supports the notion of a much colder Arctic cli- treatment of sea ice. Changes in other earth system com- mate at this time, the change in the orbital boundary ponents, such as vegetation, atmospheric CO2, and condition alone is unable to account for the intensity of ocean circulation, may be needed to produce realistic cooling suggested by the paleoenvironmental evidence. simulations. A recent modeling study (Foley et al. 1994) found that the presumed high-latitude vegetation chang- es during the mid-Holocene exerted strong climatic le- 7. Conclusions verage in the Arctic, inducing positive feedbacks that A coupled atmosphere±mixed layer ocean GCM is were comparable in magnitude to the insolation anom- used to estimate the effect of orbital forcing and sea- aly. Most GCMs have been unable to reproduce wide- ice dynamics on the Arctic climate. When the model spread glaciation at 115 kyr BP with orbital forcing version with ice motion is perturbed with positive (6 alone, without simultaneously invoking lower CO2 kyr BP) and negative (115 kyr BP) insolation, signi®cant (Syktus et al. 1994) and/or vegetation changes (Galli- climate anomalies ensue that are in the same sense as more and Kutzbach 1996). the orbital forcing changes: the 6 kyr BP (115 kyr BP) The inclusion of sea-ice dynamics in¯uences the sim- simulation produces warmer (colder) temperatures and ulations considerably. Ice motion acts as a predomi- reduced (enhanced) sea ice. In the cold scenario in the nantly negative climatic feedback across the range of central Arctic sea ice thickens by 2 m and air temper- thermal forcing from 6 kyr to 115 kyr BP, but this com- atures drop by2Kintheannual average. Signi®cant posite effect is due to strong negative feedbacks under cooling also occurs in northern Canada, but no per- warm orbital forcing exceeding moderate positive feed- manent snow cover is produced, despite evidence for backs under cold orbital forcing. Under both kinds of rapid polar ice sheet growth at this time. In the warm thermal anomalies ice transport causes negative feed- scenario, the central Arctic experiences a sea ice thin- backs with respect to sea-ice thickness in every season. ning of about 0.5 m and less than a 1-K warming. The By enhancing the spread of ice equatorward through modest 6 kyr BP changes simulated here are in contrast advection, ice motion generally serves as a cooling

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been driven by concerns of anticipated future anthro- pogenic climate change, it is also important to consider the role of ice dynamics under colder climates than pre- sent, since the impact may be quite different under a different background atmospheric state. Considerably colder conditions in boreal high latitudes have occurred over the past few centuries, in conjunction with large increases in sea-ice coverage along the coast of Iceland (Lamb 1995). The use of altered orbital parameters as inputs to the GCM used here allows insight into the physical forcing mechanisms associated with ice motion FIG. 16. Hypothesized feedback loop illustrating the interaction under both warm and cold climates, the effects of which among ice dynamics and the atmosphere. would probably be similar for other types of warming and cooling scenarios. mechanism, corresponding to a positive feedback with Acknowledgments. This work was supported by a respect to ice coverage in the cold scenario but as a NASA Earth System Science Fellowship (NGT-30346) negative feedback in the warm scenario. Because the and NSF Grant ATM-93-18973. Computer time was surface air temperature responds more strongly to provided at the National Center for Atmospheric Re- changes in ice concentration than ice thickness, ice search. The author is very grateful to Jim Maslanik, transport serves as a negative feedback on air temper- Jeremy Dunn, and Dave Pollard for providing the output ature at 6 kyr BP but as a positive feedback at 115 kyr for the modern control simulations and to John Kutz- BP. bach for his thoughtful comments and guidance in the A dynamic ice pack creates important climatic chang- development of this paper. es by initiating ice coverage anomalies that alter lower- tropospheric heating patterns at the surface. 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