Relationship between synoptic forcing and polynya formation in the

JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 109, C04023, doi:10.1029/2003JC001838, 2004

Relationship between synoptic forcing and polynya formation in the Cosmonaut Sea: 2. Regional climate model simulations David A. Bailey,1,2 Amanda H. Lynch,3 and Todd E. Arbetter4 Cooperative Institute for Research in Environmental Sciences/Program in Atmospheric and Oceanic Sciences, University of Colorado, Boulder, Colorado, USA Received 26 February 2003; revised 12 November 2003; accepted 23 February 2004; published 24 April 2004.

[1] In part I of this series [Arbetter et al., 2004], a relationship was found in the

observations between synoptic atmospheric systems and polynya formation in the Cosmonaut Sea region of Antarctica. In this study, we use a regional coupled atmospheresea ice model to investigate the role of atmospheric forcing of the polynyas in this area. The model successfully simulates both latent and sensible heat polynyas which are found in the region. In particular, a deep-ocean polynya is formed coincident with the passage of an atmospheric synoptic scale system. The divergence in the sea ice is found to be highly correlated with the atmospheric divergence. We conclude that the initial formation of a polynya can be caused by the interaction of the atmospheric synoptic scale and the continental katabatic wind regimes. While we cannot directly evaluate the role of the ocean using the present model simulations, we performed experiments with different levels of prescribed basal heat flux. These experiments highlight a polynya which could be initiated by the winds and maintained by the oceanic heat. This leads to the concept of a wind-driven sensible heat polynya, not typical of other deep-ocean polynyas INDEX TERMS: 4540 Oceanography: Physical: Ice such as the Weddell polynya of the 1970s. mechanics and air/sea/ice exchange processes; 4504 Oceanography: Physical: Air/sea interactions (0312); 4207 Oceanography: General: Arctic and Antarctic oceanography; 4215 Oceanography: General: Climate and interannual variability (3309); KEYWORDS: polynya, model simulation, Antarctica Citation: Bailey, D. A., A. H. Lynch, and T. E. Arbetter (2004), Relationship between synoptic forcing and polynya formation in the Cosmonaut Sea: 2. Regional climate model simulations, J. Geophys. Res., 109, C04023, doi:10.1029/2003JC001838.

1. Introduction [2] The oceanic thermohaline circulation has a crucial role in the moderation of the climate system. The polar regions are the locations of two important branches of this global ocean circulation: formation of deep water and ventilation of deep and intermediate waters. One signal of these oceanic processes is the development of areas of open water in the sea ice, known as polynyas. Certain polynyas, such as the large Weddell Sea polynya of the 1970s, are an indication of deep convective regions where cold, highsalinity water is formed and eventually flows off the continental shelf to become Antarctic Bottom Water. Other polynyas, such as the Cosmonaut Sea polynya, involve the upwelling of relatively warm circumpolar deep water and 1

Now at School of Computational Sciences, George Mason University, Fairfax, Virginia, USA. 2 Also at Center for Ocean-Land-Atmosphere Studies, Calverton, Maryland, USA. 3 Now at School of Geography and Environmental Science, Monash University, Clayton, Victoria, Australia. 4 Now at National Snow and Ice Data Center, University of Colorado, Boulder, Colorado, USA. Copyright 2004 by the American Geophysical Union. 0148-0227/04/2003JC001838

do not seem to be locations of deep convection. In the Cosmonaut Sea, a zone of divergence between the Antarctic circumpolar current and the Antarctic coastal current allows circumpolar deep water of North Atlantic origin to be upwelled to the surface. This warmer water provides heat to melt the sea ice during the cold Antarctic winter. [3] The two main types of polynyas are named for the dominant mechanisms involved in their formation. A latent heat polynya, sometimes referred to as a coastal polynya, is formed by wind or ocean currents which remove the sea ice from an area. The rapid regrowth of sea ice causes significant release of latent heat and brine. Latent heat polynyas are well studied [Massom et al., 1998; Cotton and Michael, 1994; Bromwich et al., 1992] and will not be the focus here. A sensible heat polynya or deep-ocean polynya, on the other hand, forms through the upwelling of warm water which melts the ice through sensible heat transfer. The region usually requires a ‘‘preconditioning’’ mechanism (through atmospheric, oceanic, or ocean bottom topographic conditions) before the water can be upwelled to melt the ice. Deep-ocean sensible heat polynyas in particular are an important aspect of climate variability because of their relative size (up to 2% of the overall Antarctic winter sea ice cover) and likely impact on ocean ventilation, intermediate and bottom water formation, and atmospheric circu-

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BAILEY ET AL.: COSMONAUT SEA POLYNYA FORMATION, 2

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Figure 1. A map of the Cosmonaut Sea region with the model topography (m) and locations/features of interest. lation. For example, the exchange of heat and moisture between the atmosphere and ocean in the presence of an idealized Weddell Sea polynya has been shown in global simulations to have a large effect on the variability of the Antarctic and global climate [Glowienka-Hense, 1995]. [4] The largest puzzle of the Antarctic sea ice still remains, the formation and disappearance of the Weddell Sea polynya of the 1970s. Various modeling approaches have been tried and are ongoing to investigate the formation of the Weddell Sea polynya during the three consecutive winters of 1974– 1976. The main difficulty is determining the preconditioning mechanism and whether it was provided by the ocean or atmosphere. Martinson et al. [1981], using a two-layer ocean model, found that the initial brine rejection from ice formation would form dense cold preconditioned water, and cause a vertical overturning due to static instability. This in turn upwelled warm salty deep water which led to the formation of an ice-free region. They speculated that the region of vertical convection would be moved along with the mean current and then destroyed upon contact with the Weddell subpolar gyre. They also suggested that the polynya formation and disappearance may have also been affected by topographic conditions and fresh-water input from summer precipitation and runoff. Motoi et al. [1987] proposed that the preconditioned region for 1974 Weddell Sea polynya was formed by surface cooling alone without

ice formation, through the persistence of high salinity water from the preceding summer. Alverson and Owens [1996] presented a mechanism for deep-ocean convection through topographic preconditioning. They showed in idealized model simulations that the presence a sea mount could lead to the formation of a chimney and deep convection. They extended this to the region around Maud Rise in the Antarctic and suggested this as a possible mechanism for deep-ocean polynyas in this region. More recently, Holland [2001] emphasized the importance of ocean dynamics and proposed a mechanism that involved a divergent oceanic eddy, shed from Maud Rise. From the perspective of atmospheric preconditioning, Parkinson [1983], using an atmosphere-sea ice model, suggested that the atmospheric circulation and atmosphere-ocean interaction was intricately involved in the preconditioning and maintenance of the Weddell Sea polynya. Other studies, using a non-hydrostatic ocean model [Ka¨mpf and Backhaus, 1998] and a coupled atmosphere-sea ice-mixed layer model [Timmermann et al., 1999], have suggested that sporadic, short-lived atmospheric events have an important role in the formation of deep ocean polynyas. In coarse-resolution atmosphere-sea ice model sensitivity experiments, Wu et al. [2003] concluded that a number of major Antarctic polynyas, including the Weddell polynya, were unlikely to form and be maintained by oceanic sensible heat processes alone.

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Figure 2. SSM/I derived sea ice concentration for August (a) 6, (b) 7, and (c) 8, 1988.

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Figure 3. The specified (a) BHT barotropic ocean currents (m s 1) with stream function contours (Sv), (b) oceanic heat flux (W m 2) perturbation region, and (c) BHT surface ocean currents (m s 1) used in the model experiments.

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Figure 4. Model simulated sea level pressure (hPa) and winds (m s 1) for August 6, 1988, at (a) 0000 UT and (b) 1200 UT, and August 7 at (c), 0000 UT and (d) 1200 UT. [5] While the majority of studies have focused on the Weddell polynya of the 1970s, satellite observations from 1979 to present day have indicated the continued recurrence of deep-water and coastal polynyas in the Cosmonaut Sea region [Comiso and Gordon, 1996]. The deep-ocean Cosmonaut polynya area is significant because of its size (averaging 7.2 104 km2) and because it may supply an observable model for similar deep-ocean polynyas. The deep-ocean Cosmonaut polynya region is thought to be initiated and maintained primarily by upwelling of warmer circumpolar deep intermediate water causing ice melt and a divergence of surface water [Comiso and Gordon, 1996] as opposed to deep-ocean convection, deemed to be the main mechanism in the formation of the Weddell polynya. It should also be noted that the Cosmonaut Sea deep-ocean polynya region is actually comprised of multiple reduced ice concentration sites. These sites are sometimes distinct, defined by western and eastern polynya modes, or merged into one polynya [Comiso and Gordon, 1996]. For the remainder of the paper, we will only refer to the single merged deep-ocean polynya.

[6] The continental coastline of Antarctica, in the vicinity of the Cosmonaut Sea around 50E, has the typical steep topography (Figure 1) which causes cold dense katabatic winds to accelerate and blow off the continent. These strong katabatic winds lead to frequent latent heat polynya formation [Ishikawa et al., 1996]. Three permanent stations are located in the general vicinity, the Japanese station at Syowa, the Russian station of Molodezhnaya, and the Australian station of Mawson. The oceanography of the region is known primarily from the cruises of the Japanese Antarctic Research Expeditions (JARE) [Wakatsuchi et al., 1994; Takizawa et al., 1994; Ohshima et al., 1996]. The main barotropic (vertically averaged) oceanic features include the Antarctic Coastal Current (CC) which follows the coastline as it flows to the west. Farther to the north, the Antarctic Circumpolar Current (ACC) flows eastward. At the surface, the wind-driven circulation to the north is in the opposite sense to the ACC and the coastal circulation is in the same direction as the barotropic flow [Orsi et al., 1995; Ohshima et al., 1996]. Between these two major surface currents is the Antarctic divergence zone where the Cir-

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Figure 5. Model simulated sea ice concentration (%) for experiment 1 on August 7 at (a) 0000 UT, (b) 0600 UT, (c), 1200 UT, and (d) 1800 UT. cumpolar Deep Water (CDW) is ventilated [Orsi et al., 1995]. [7] The shear zone implied by the barotropic currents was theorized to be a mechanism leading to the formation of the deep-ocean Cosmonaut Sea polynya [Comiso and Gordon, 1996] during August and September. Their theory invoked the conservation of potential vorticity, in which a cyclonic eddy was compressed between the ACC and the CC. To conserve potential vorticity, the fluid column would have to stretch, causing enhanced upwelling of CDW, and the heat from this would be sufficient to maintain an ice free region. Comiso and Gordon [1996] also conjectured that this particular polynya would not be the site of deep convection. This idea was supported by JARE summer voyages [Takizawa et al., 1994; Wakatsuchi et al., 1994], where they found that the Cosmonaut Sea did not have the anomalously cold water typically associated with deepocean convection during the preceding winter. More recent observations in the region show the presence of Antarctic Bottom Water (AABW) at a few locations off the coast of Enderby Land [Whitworth et al., 1998]. Nonetheless, in

accordance with Comiso and Gordon [1996], the observations suggest that the AABW is not likely to be formed in this region and has been advected from farther east, in the vicinity of the Amery Ice Shelf [Orsi et al., 1999]. The ice shelf provides super-cooled water which is a crucial component in the deep water formation process [Orsi et al., 1999]. The mechanisms involved in the formation and maintenance of this particular polynya are still unclear. [8] From the analysis of the Special Sensor Microwave/ Imager (SSM/I) derived sea ice concentration, Comiso and Gordon [1996] found that the Cosmonaut Sea polynya was very active during July and August of 1988. Figure 2 displays the SSM/I derived sea ice concentration for August 6 – 8, 1988. This is a time period, during mid-winter, when the overall sea ice concentration rapidly increases to its maximum extent. The ice concentration in the center of the polynya is less than 0.3 (seen in blue) on August 7 (Figure 2b). There is an uncertainty of 10–35% in this value, especially in predominantly new ice regions, due to differences in the SSM/I algorithms [Comiso and Steffen, 2001]. As this

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Figure 6. Difference in effective ice thickness (ice area ice thickness) due to (a) dynamic transport and (b) thermodynamic growth in experiment 1 between 1200 UT and 0600 UT on August 7. Shaded areas indicate negative values. The contours are in units of centimeters with a contour interval of 2 cm. The box indicates the bullseye region used in experiments 2 – 4.

was a time of peak activity in the polynya, this period will be the focus of the model experiments in this paper. A detailed analysis using the satellite products and large-scale atmospheric analysis, in a companion paper [Arbetter et al., 2004] (hereinafter ALB) indicates an interannual link between the synoptic scale atmospheric regime and polynya occurrence in the Cosmonaut Sea. While ALB is the first in the companion sequence, the two studies arose nearly independently using two different approaches. The first part was intended to investigate the role of the atmospheric regime over several years. This work, the second part, is a first attempt to explicitly model a particular occurrence of the Cosmonaut Sea polynyas, using a coupled regional climate system model.

2. Model and Experiment Description [9] The model used for this work is the Antarctic Regional Climate System Model (AntARCSyM) [Bailey and Lynch, 2000a], which is an Antarctic implementation of its Arctic predecessor (ARCSyM) [Lynch et al., 1995; Walsh et al., 1993]. The AntARCSyM is a coupled climate system model that incorporates comprehensive treatments of the atmosphere, ocean, sea ice, and land surface for application over a limited region and contains physical parameterizations suitable specifically for highlatitude applications. The atmospheric component of the model is based on the National Center for Atmospheric Research (NCAR) Regional Climate Model Version 2 (RegCM2) as described by Giorgi et al. [1993a, 1993b]. The dynamical core of RegCM2 is based on the NCAR/Penn State Mesoscale Model Version 4 (MM4), described by Anthes et al. [1987]. It is a hydrostatic primitive equation model with a terrain-following sigma vertical coordinate (s) system. The model employs a horizontal Arakawa ‘‘B’’ grid formulation [Arakawa, 1966] and a split-explicit time integration

scheme [Giorgi et al., 1993a]. The radiative transfer schemes in ARCSyM include the CCM2 shortwave and longwave radiative models of Briegleb [1992a, 1992b], and the more recently developed longwave rapid radiative transfer model (RRTM) [Mlawer et al., 1997]. The RRTM was shown to be more accurate in polar simulations [Pinto et al., 1999]. The boundary layer physics is modeled through the explicit planetary boundary layer scheme of [Holtslag et al., 1990]. Convective cumulus cloud processes are parameterized either as a simple moisture convergence threshold [Kuo, 1974] or through a more complex scheme allowing for updrafts and downdrafts with the clouds [Grell, 1993]. Nonconvective, grid-scale moist processes can be resolved through an ‘‘implicit’’ scheme [Anthes et al., 1987] in which supersaturated water vapor is precipitated instantaneously, or an ‘‘explicit’’ scheme [Hsie et al., 1984] which includes prognostic equations for cloud water, rainwater, water vapor, and cloud ice. The atmospheric model has been coupled to a land surface model (LSM) [Bonan, 1996] that encompasses multiple soil layers and sub-grid-scale vegetation parameterizations. It has been determined, from the initial simulations [Bailey and Lynch, 2000a, 2000b], that there were inadequacies in the original elevation data set, and hence the topography of the continent was derived from the Drewry [1983] data set for the current model realizations. [10] The initial model assessment of Bailey and Lynch [2000a, 2000b] provides some confidence in the models ability to simulate the coastal climate, year-round, around Antarctica and over the Southern Ocean sea ice cover in winter. Using a coupled atmosphere-sea ice model, several experiments were performed to provide support for the upwelling mechanisms surrounding the formation of this polynya. The atmospheric-sea ice simulations in the experiments presented in this paper are very similar to the experiments performed by Bailey and Lynch [2000c],

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Figure 7. Model simulated sea ice concentration for experiment 2 for August 7 at (a) 0000 UT, (b) 0600 UT, (c), 1200 UT, and (d) 1800 UT.

which compared well to the limited station observations. The model grid covered the domain indicated in Figure 1 with a horizontal resolution of 20 km. All experiments were initialized using SSM/I derived sea ice concentration at the beginning of June 1988 to allow for 1 month of spin up for the sea ice thickness. The initial sea ice thickness was uniformly set to 50 cm. Direct ice thickness observations were not available for this region, but this value agrees with overall East Antarctic ice thickness observations [Allison et al., 1993]. The initial sea surface temperature (SST) was prescribed from climatology [Shea et al., 1992]. The atmospheric model was initialized and forced at the boundaries using the ECMWF 12-hourly operational analyses from NCAR [Trenberth, 1992]. [11] The simulations in this paper were performed using the coupled atmosphere-sea ice model with specified ocean currents and oceanic heat flux. The August monthly mean ocean barotropic or surface currents from the model simulation of Beckmann et al. [1999] (hereinafter BHT) were prescribed, for some simulations, underneath the atmo-

sphere-sea ice model. Figure 3a shows the barotropic circulation in the Cosmonaut Sea from the BHT simulation. Notice the recirculation in the domain which includes the far eastern part of the Weddell gyre. The ACC enters this domain in the upper right corner (Figure 3a) and is relatively weak at around 2 cm s 1. The CC is stronger at around 5 cm s 1. In the surface currents (Figure 3b) the CC is similar to that in the barotropic currents (Figure 3a), but the circulation to the north is counter-rotating. In the BHT model, however, the surface divergence is located farther to the north than is observed by sea ice drifters [Heil and Allison, 1999] and compared to climatology [Orsi et al., 1995]. Typically, sea ice models are forced from below using the geostrophic ocean currents. Hence, to maximize the impact of the ocean circulation on the sea ice, both the surface and barotropic currents were used in separate experiments to test the sensitivity of the model. Additionally, as the surface divergence region in the BHT currents is located farther north than suggested by observations, the ocean currents were shifted southward. Thus the surface

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Figure 8. Net difference between experiments 2 and 1 in effective ice thickness change from 0600 UT to 1200 UT on August 7, due to (a) dynamic transport and (b) thermodynamic growth. Shaded areas indicate negative values. The contours are in units of centimeters with a contour interval of 1 cm. The box indicates the bullseye region used in experiments 2 – 4.

divergence was located closer to where the deep-ocean polynya was expected to form. [12] The eddy-induced upwelling theory presented by Comiso and Gordon [1996] implies a warm water upwelling rate (we) of 2.6 10 6 m s 1 into the mixed layer. To compute an approximate upwelling flux, the expression is given by we = rcpDT, where r (1025 kg m 3) is the density of seawater, cp (3989 J kg 1 K 1) is the specific heat of seawater, and D T is the temperature difference across the thermocline. Assuming a maximum temperature difference, DT = 1.0 ( 1.8) = 2.8C, the upwelling flux would provide approximately 30 W m 2 of heat to the mixed layer. The oceanic heat flux at the base of the ice is computed using the expression Fo = rctu*DT, where ct is the turbulent transfer coefficient at the base of the ice, and DT is the temperature difference between the mixed layer and the base of the ice. The transfer coefficient is typically of the order of 10 5, and the maximum DT is about 0.5C. This would imply an ocean to ice heat flux of approximately 20 W m 2. The transfer coefficient (ct), however, is a function of the turbulent mixing at the base of the ice, and can vary up to 5 times depending on the surface forcing. Thus the instantaneous oceanic heat flux can be in excess of 100 W m 2. For the experiments testing the impact of oceanic upwelling in the region, the heat flux at the base of the ice and the ocean currents were specified based on the following experiments: (1) In experiment 1 a uniform heat flux of 15 W m 2 was used, with the barotropic ocean currents (Figure 3a) included, to test whether the atmospheric forcing could produce a polynya with minimal impact from the ocean. (2) In experiment 2 the heat flux was set uniformly to 15 W m 2, except for a patch of 30 W m 2, centered on 50E (Figure 3c) where the deep-ocean polynya was expected to form, and the ocean currents were removed. (3) Experiment 3 included the same heat flux distribution as the previous experiment, but also included the shifted surface currents from BHT (Figure 3b). (4) Finally, in

experiment 4, the ‘‘bullseye’’ heat flux was increased to 200 W m 2 as an extreme case.

3. Results [13] The atmospheric circulation for August 6– 7, 1988, from model experiment 4 is presented in Figure 4. Note that the model simulations have higher temporal and spatial resolution, in all state variables, than large-scale weatherforecast based analyses (e.g., NCEP or ECMWF). Hence the simulations capture smaller-scale system developments more readily. Despite the significant differences in sea ice concentration between the experiments (shown below), the atmospheric simulations of all the experiments were very consistent. A maximum August mean sea level pressure difference of 2 hPa was found over regions covered by ice in one experiment and open water in the other with an RMS overall difference of less than 1 hPa. Only the atmospheric results from experiment 4 are presented here. [14] On August 6 at 0000 UT (Figure 4a), there was a synoptic-scale low-pressure system along the northern boundary of the domain. The prevailing winds around the low provided a predominantly westward flow in the northern half of the region. To the south, strong katabatic winds flowed seaward off the coast turning to the west due to coriolis deflection and then were influenced by the synoptic environment. A region of divergence in the flow near the western boundary was evident. Six hours later, the lowpressure center had moved southeastward as had the region of divergence. At 1200 UT (Figure 4b), the low continued to the southeast and intensified. The trend continued 6 hours later. During this period, the minimum ice concentration in the polynya had decreased from about 45% on August 5 to less than 40% on August 6 according to the SSM/I product (Figure 2a). At 0000 UT on August 7 (Figure 4c), the lowpressure center reached a peak intensity with a central pressure of 940 hPa. There were several areas of small-

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Figure 9. Model simulated sea ice concentration for experiment 3 on August 7 at (a) 0000 UT, (b) 0600 UT, (c), 1200 UT, and (d) 1800 UT.

scale divergence in the region at this time. Further, a mesoscale disturbance was developing in the center of the domain off the coast of Cape Ann. Six hours later the cyclone weakened and moved eastward. The disturbance off the coast of Cape Ann developed a closed circulation. This sequence was consistent with typical interactions of katabatic, large-scale synoptic, and mesoscale circulations [Bromwich and Parish, 1988; Parish and Bromwich, 1998]. Note the surface divergence created by the confluence of the cyclonic circulation with the prevailing katabatic flow. By 1200 UT (Figure 4d), the synoptic low-pressure system had continued slowly to the east. Further, there was a region of divergence adjacent to the mesoscale system near Cape Ann, which will be discussed later. During this time period, the sea ice concentration minimum was reduced to near 30% and the polynya grew to its maximum size (Figure 2b). The low-pressure center continued to move eastward, the atmospheric circulation to the north weakened, and the katabatic winds returned to the prevailing direction. The polynya had shrunk in size by

August 8 (Figure 2c) and nearly disappeared by August 11 [Comiso and Gordon, 1996]. While the concurrence alluded to here does not imply cause and effect, it does suggest that atmospheric forcing played a role in the polynya development. [15] The sea ice simulation from experiment 1 (Figure 5), during the peak period of August 7 from 0000 to 1800 UT, indicated the presence of a very weak polynya (indicated by the arrow) which was distinct from the latent heat polynyas along the coast. Note that the overall ice extent simulated by the model was larger than suggested by the satellite-derived product (Figure 2). This overestimation was likely a result of insufficient oceanic heat and the simple open boundary conditions. Also, the coastal polynya region in the model was larger than that shown by the satellite-derived product. This may be due to the inability of the satellite product to resolve the coastal region sufficiently, but more likely the absence of fast-ice in the model. The difference between the deep-ocean region of low ice concentration and the coastal polynyas is seen in the total dynamic transport and

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Figure 10. Net difference between experiments 3 and 2 in effective ice thickness change from 0600 UT to 1200 UT on August 7, due to (a) dynamic transport and (b) thermodynamic growth. Shaded areas indicate negative values. The contours are in units of centimeters with a contour interval of 1 cm. The box indicates the bullseye region used in experiments 2 – 4.

thermodynamic growth from 0600 UT to 1200 UT on August 7 (Figure 6). Both the latent heat polynyas near the coast and the polynya farther away showed strong ice divergence, i.e., negative dynamic transport (Figure 6a, indicated by shading). Large ice regrowth (Figure 6b, indicated by dense contours in white) occurred in the latent heat polynyas along the coast. In the polynya to the northwest (within the box), there was negligible ice growth and even a small amount of ice melt (spatial average growth of 1 mm over the 6-hour period). Hence the processes seen in these reduced ice concentration regions were distinct and consistent with the definitions of latent and sensible heat polynyas. This experiment implied that a weak sensible heat polynya could be formed through atmospheric divergence alone, provided there was a baseline of at least 15 W m 2 of oceanic heat from below. There was, however, insufficient heat from the ocean in experiment 1 to maintain and increase the size of the polynya. In a separate experiment (not shown), with an increased uniform heat flux of 30 W m 2 and the BHT barotropic currents, the sea ice extent was smaller than that in the SSM/I product (Figure 2). This implies, in the absence of other biases, that an overall flux of 30 W m 2 was excessive for the region as a whole. [16] In experiment 2, the ocean currents were removed and a ‘‘bullseye’’ heat flux of 30 W m 2 was applied (described earlier). The sea ice concentration (Figure 7) was similar to that in experiment 1, except for slightly lower concentrations in the latent and sensible heat polynyas. The differences of the dynamic transport and thermodynamic growth between experiments 2 and 1, on August 7, are presented in Figure 8. The difference in dynamic transport, between the two experiments, in the sensible heat polynya region was relatively small (about 1 cm higher transport in experiment 2). Experiment 1 had prescribed ocean currents, while experiment 2 did not have ocean currents. In addition, the latter experiment had the bullseye heat flux (indicated

by the box). It is interesting that the removal of the relatively small ocean currents did not significantly change the sea ice transport. The thermodynamic growth difference (Figure 8b) between experiments 2 and 1 was negligible (less than 1 cm), but there was slightly more melt (more negative) in experiment 2. Overall, the results of this experiment implied that the lack of ocean currents in experiment 2 had very little effect on the polynya formation. [17] The sea ice concentration for experiment 3 (Figure 9) was similar to that in experiments 1 and 2. In this case, the surface ocean currents were prescribed to test whether the oceanic divergence would influence the polynya formation. The bullseye heat flux was also applied. The sea ice concentrations in the polynyas, from experiment 3, were slightly reduced compared to experiment 1, and almost indistinguishable from experiment 2. Again, it is difficult to see this by simple inspection. The dynamic transport and thermodynamic growth differences between experiments 3 and 2, on August 7, are presented in Figure 10. Here the differences were negligible in the vicinity of the deep-ocean polynya (western portion of the box). Dynamic transport differences of up to 4 cm were apparent in the eastern portion of the box and in the coastal region. In this case, experiment 3 had prescribed surface currents and experiment 2 had none. The currents were positioned such that a region of ocean divergence would be located in the same vicinity as the bullseye heat flux. It should be noted that the ocean currents were prescribed from a monthly mean model climatology and reached a maximum speed of 5 cm s 1. The ocean currents were very slow compared to the overlying atmospheric circulation of 30 m s 1. Hence it appears that the prescribed surface ocean currents had limited influence on the sea ice circulation. This implies that the sea ice circulation in these model simulations was strongly wind driven. Note that in the real ocean, the surface currents are also wind driven, primarily through the drag from the ice. Therefore a region of oceanic upwelling could be

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Figure 11. Model simulated sea ice concentration for experiment 4 on August 7 at (a) 0000 UT, (b) 0600 UT, (c), 1200 UT, and (d) 1800 UT. more intense and located directly under the atmospheric divergence. [18] Finally, in experiment 4, the simulation was initialized on August 1 using the results from experiment 3. As mentioned earlier, the observed oceanic heat flux can be in excess of 100 W m 2. As a test to determine whether a polynya could be sustained under oceanic forcing alone, the bullseye heat flux was increased to 200 W m 2. Note that a separate experiment (not shown) with a heat flux of 100 W m 2 in this region produced similar results to experiment 3. The sea ice results from experiment 4 show that the polynya formed sooner during this period and the concentration was significantly reduced (see Figure 11). In this case, a region of the polynya formed farther to the north (Figure 11a). The polynya did not form, however, until a synoptic-scale atmospheric system moved into the region (Figure 4) as in the previous experiments. Also, the polynya dissipated after the synoptic system had moved farther east. Hence it appears that even with substantial heat from the ocean, the polynya cannot be formed through oceanic upwelling alone.

[19] To provide more evidence for the mechanisms described here, experiment 4 produced another strong, although spurious, polynya event on August 12. (Note that the SSM/I product indicated only a weak polynya on this day.) The surface wind regime on August 12 (Figure 12c) was similar to that on August 7 (Figure 12a), but significantly weaker. The maximum wind speeds on August 7 were around 20 m s 1 and the flow was generally northward throughout the domain. The surface wind flow, on August 7 (Figure 12a), and hence sea ice drift (Figure 12b) in the deep-ocean polynya area (western part of box) was predominantly eastward during this time. However, on the eastward side of the box, the ice drift was very strong and directed northward around the lowpressure center. In contrast to the conditions on August 7, the maximum wind speed on August 12 was around 10 m s 1 (Figure 12c). The sea ice drift on August 12 (Figure 12d), on the west side of Cape Ann near the coast, was almost negligible. To the north, as on August 7, the ice drift was directed to the north around the low. There was clearly a change in the katabatic wind offshore flow, in

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Figure 12. Model simulated (a) sea level pressure (hPa) and winds (m s 1) and (b) sea ice concentration (%) and drift (cm s 1) on August 7. (c and d) Same for August 12. both cases, which was either an interaction with the synoptic-scale system [Parish and Bromwich, 1998], or a temporary cessation of the katabatic flow [Galle´ et al., 1998; Galle´ and Pettre´, 1998]. To examine this more closely, the atmospheric and sea ice divergence fields are presented in Figures 13a and 13b for August 7 at 0600 UT. The region of divergence (white) was clear in the western portion of the box, where the deep-ocean polynya formed, for both the atmosphere (Figure 13a) and sea ice (Figure 13b). In contrast, on August 12, the atmospheric and sea ice divergence (Figures 13c and 13d) were small, although generally divergent, within the box region. This allowed the bullseye heat flux to melt ice and produce a larger polynya throughout the whole region. This polynya, however, was an artifact of the specified large stationary heat flux. While this particular occurrence disappeared, other spurious polynyas formed later in August (not shown). As no spurious polynyas occurred in the first three experiments, this suggests that a stationary heat flux of 200 W m 2 was unrealistic.

[20] The time evolution of the area-averaged atmospheric/ sea ice divergence and sea ice concentration in the deepocean polynya region (the western half of the box) from experiment 2 is presented in Figure 14. The sea ice divergence was well correlated with the atmospheric divergence through the month of August. The ocean divergence (not shown) was negligible and constant (mean currents) and not significantly correlated with the sea ice divergence. On August 7, the divergence in the atmosphere and sea ice was very clear (Figure 14a). This divergence on August 7 led to a drop in the ice area approximately 6 to 12 hours later (Figure 14b) which roughly corresponds to the lag found in ALB. On August 12, the atmospheric and sea ice divergence was near zero, but increasing (Figure 14a). During this same period, a decreasing trend was seen in the sea ice concentration (Figure 14b). The spurious polynya on August 12, in experiment 4, was not present in this particular experiment (not shown). It is important to note that the correlations were not perfect. A portion of this can be explained by small-scale atmospheric variability that was not captured by the 6-hourly model sampling shown

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Figure 13. Atmospheric and (b and d) sea ice convergence/divergence (10 3 s 1) for August 7 (Figures 13a and 13b) (at 0600 UT) and August 12 (Figures 13c and 13d) (at 0600 UT). The contour intervals are 3 10 5 s 1 and 2 10 6 s 1 for the atmospheric and sea ice divergence, respectively. here. Additionally, internal sea ice mechanics will cause the sea ice to respond more slowly and not diverge/ converge at the same temporal and spatial scales as in the atmosphere. A partial explanation for the imperfect correlation of the sea ice divergence and concentration is the thermodynamic contribution. Generally, the net result of the model simulations was that a region of atmospheric and sea ice divergence occurred over the bullseye heat flux area, which allowed a polynya to form. The polynya disappeared when the atmospheric divergence moved off the enhanced heat flux region. During periods without a polynya, the sea ice was continually advected through the bullseye region and the oceanic heat was not able to melt the ice significantly.

4. Conclusions [21] In these experiments, a polynya was simulated in the deep ocean, away from the coast, in addition to recurrent coastal polynyas. While the overall ice extent was too large,

the formation of this polynya was consistent with a sensible heat polynya mechanism. The rapid ice regrowth present in the coastal areas and typical of latent heat polynya formation was not present in the deep-ocean polynya. In the model simulations presented with specified oceanic heat flux, the deep-ocean polynya was present during a period of surface wind and sea ice divergence. Despite specified heat flux (upwelling) throughout the duration of the experiments, the polynya only formed during a meso-synoptic atmospheric event. The inclusion of monthly mean barotropic or surface ocean currents did not influence the results significantly. The sequence of events for the formation of the polynya in the model simulation was (1) a change in the katabatic wind flow regime from an interaction with a synoptic-scale system or a cessation of the katabatic outflow, (2) a region of atmospheric divergence forced the ice to diverge over the enhanced heat flux region, (3) an enhanced heat flux able to melt the sea ice and create a polynya; (4) a wind regime that returned to ‘‘normal’’ and a divergence that moved away from the

Figure 14. (a) Area-averaged time series (August) of atmospheric and sea ice divergence/convergence (10 3 s 1) in deepocean polynya region (western half of box) of experiment 2. (b) Area-averaged time series of sea ice divergence along with ice concentration in the deep-ocean polynya region. 13 of 16

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Figure 14 14 of 16

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enhanced heat flux region, (5) sea ice advection, due to the change in winds, that filled in the polynya region. These simulations lead to the concept of a wind-driven sensible heat polynya. [22] It was found that a polynya would not form in the model due to ocean forcing alone, such as oceanic upwelling due to local eddy intensification, except in the experiment with the bullseye region of 200 W m 2 . The aforementioned correlation of the sea ice and atmospheric divergence suggested that the sea ice would not remain over this region long enough to be affected. This was verified in the model simulations by prescribing a large region of excessive oceanic heat. These simulations only produced a polynya when the sea ice was advected out of the region by atmospheric forcing, or remained stationary for an extended period over the surplus heat flux region. The temporal and spatial variations in the oceanic heat flux, which follow the winds and sea ice, were not present in this model. The instantaneous ocean divergence could actually be much larger, and hence the upwelling more vigorous, producing oceanic heat in excess of 100 W m 2. Note that observations of the instantaneous oceanic heat flux in the Weddell Sea were up to 1000 W m 2 [McPhee et al., 1996]. The use of a simple interactive ocean model that includes upwelling could be used to confirm this hypothesis and determine the spatial and temporal variability of the oceanic heat flux. Hence, in accordance with Comiso and Gordon [1996], we find that the atmosphere is responsible for preconditioning the region through surface divergence. We can only speculate on the role of the ocean in the formation the polynya. Certainly, during the period when the winds became nearly quiescent in experiment 4, the stationary heat flux was able to produce a spurious polynya on August 12. [23] In addition to the conclusions regarding this polynya formation event, an important result of this study is that the short-term sea ice drift in the Cosmonaut Sea is generally and primarily wind driven. In model experiments, the sea ice divergence was highly correlated with the atmospheric divergence and not with specified ocean currents, even in the absence of polynya formation. Further, the sea ice drift from the model compares well with local observations from buoys [Heil and Allison, 1999] and is an order of magnitude larger than the mean ocean currents. These results indicate that this is a true characteristic of the Cosmonaut Sea and likely in all coastal areas around Antarctica where the coastal ocean current is relatively weak. Observational estimates of the ocean currents in this region were not directly available. [24] One important question which remains is how the deep-ocean polynya is maintained. The model was unable to maintain a polynya for 6 days as seen in the work of Comiso and Gordon [1996]. Some deficiencies in the model simulation and the 1- to 2-day lag found in ALB point toward a possible oceanic mechanism. These particular model simulations could not account for oceanic divergence and upwelling or mixed-layer deepening in response to the surface wind/sea ice stresses and oceanic heat loss. Future model simulations will include an interactive ocean component to investigate this question further. [25] Acknowledgments. This work was generously supported by a grant from the NASA PORAP program to the University of Colorado and a

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CIRES graduate research fellowship. The authors thank the associate editor, R. Kwok, and two anonymous reviewers for their helpful and constructive comments. D. Bailey also thanks P. Rhines and many others at the School of Oceanography, University of Washington, for computing resources used to prepare this manuscript and many insightful discussions during his postdoctoral work.

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T. E. Arbetter, National Snow and Ice Data Center, University of Colorado, UCB 449, Boulder, CO 80309-0449, USA. (todd.arbetter@ colorado.edu) D. A. Bailey, Center for Ocean-Land-Atmosphere Studies, 4041 Powder Mill Road, Suite 302, Calverton, MD 20705-3106, USA. (bailey@cola. iges.org) A. H. Lynch, School of Geography and Environmental Science, PO Box 11A, Monash University, Clayton, Victoria 3800, Australia. (amanda. [email protected])

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