Effects of surface freshwater flux induced by sea ice transport on the global thermohaline circulation

Transcription

1 JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 108, NO. C2, 3047, doi: /2002jc001476, 2003 Effects of surface freshwater flux induced by sea ice transport on the global thermohaline circulation Y. Komuro and H. Hasumi Center for Climate System Research, University of Tokyo, Tokyo, Japan Received 14 May 2002; revised 11 September 2002; accepted 9 October 2002; published 25 February [1] Effects of surface freshwater flux induced by sea ice formation and melting on the thermohaline circulation are investigated by using a sea-ice-ocean coupled general circulation model forced by monthly climatology. Restoring to the observed sea surface salinity is not employed in order to evaluate precisely the sea ice effects. Because of the model s improper representation of interaction between sea ice and the Northern Hemisphere deep convection, the discussion is focused on the effect of sea ice in the Southern Hemisphere. In the control case, deep water formation around Antarctica occurs under compact sea ice cover, where positive annual mean sea ice production is an essential factor to induce deep convection. When sea ice motion is turned off, deep water formation is maintained primarily by thermal destabilization of water columns, as annual mean sea ice production is almost zero everywhere. Consequently, the deep ocean around Antarctica is warmer and less saline compared with that for the control case, and the Atlantic bottom circulation is weakened by 16%. Northward salt transport from the Southern Ocean in the bottom layer also decreases. The amount of the decrease is greater by 1 order of magnitude than that of the surface salt input associated with sea ice. In a case where the freshwater flux at the ice-ocean interface is turned off, results are similar to the case without sea ice transport, although surface heat and momentum fluxes significantly differ between them. This suggests that the influence of sea ice on freshwater flux is more important than on heat and momentum fluxes in affecting the global thermohaline circulation. INDEX TERMS: 4540 Oceanography: Physical: Ice mechanics and air/sea/ ice exchange processes; 4255 Oceanography: General: Numerical modeling; 4532 Oceanography: Physical: General circulation; 4215 Oceanography: General: Climate and interannual variability (3309); KEYWORDS: ice-ocean coupled model, deepwater formation, thermohaline circulation, freshwater budget, Southern Ocean Citation: Komuro, Y., and H. Hasumi, Effects of surface freshwater flux induced by sea ice transport on the global thermohaline circulation, J. Geophys. Res., 108(C2), 3047, doi: /2002jc001476, Introduction [2] The ocean has enormous mass and heat capacity. The oceanic thermohaline circulation accounts for a significant part of the meridional heat transport in the present state of the climate, and thus is considered to be one of the controlling processes for the global climatic features. The circulation also transports salt and other dissolved materials. Since the thermohaline circulation has very long timescale, it could control the long-term mean state of the ocean and the climate. [3] The thermohaline circulation is driven by the spatial inhomogeneity in surface fluxes, consisting of heat and freshwater fluxes. Its intensity and flow pattern depends strongly on mixing in the ocean interior [Munk and Wunsch, 1998] and where and how deep waters are formed. Generally, heat flux tends to induce deep water formation at high latitudes, while freshwater flux does at low latitudes. Deep water formation of the present climate state takes Copyright 2003 by the American Geophysical Union /03/2002JC place at high latitudes, where freshwater flux has a tendency to prevent the deep water formation. Previous numerical studies [e.g., Manabe and Stouffer, 1997] point out that additional freshwater input to the deep water formation area could significantly weaken the thermohaline circulation. Some studies show that freshwater flux at low latitudes is also important for the strength of the thermohaline circulation [Zaucker et al., 1994; Hasumi, 2002a]. These results suggest that the atmospheric hydrological cycle and consequent freshwater flux at the sea surface are important for the thermohaline circulation. [4] Freshwater budget at the sea surface is composed of precipitation, evaporation, river runoff, and formation and melting of sea ice. Sea ice affects sea surface freshwater flux by brine rejection and freshwater release. In addition, transport of sea ice separates the region of sea ice formation and sea ice melting [e.g., Hasumi and Suginohara, 1995]. The contribution of sea ice to the freshwater budget is very important in the polar oceans where precipitation and evaporation are relatively small [Aagaard and Carmack, 1989; Toggweiler and Samuels, 1995]. Sea ice especially plays a major role in the deep water formation at high 29-1

2 29-2 KOMURO AND HASUMI: SEA ICE EFFECTS ON THE GLOBAL CIRCULATION latitudes [Killworth, 1983; Gordon and Huber, 1990]. For example, the increase of sea ice export from the Fram Strait hindered the deep water formation in the Greenland Sea when Great Salinity Anomaly [Dickson et al., 1988] occurred [Walsh and Chapman, 1990; Häkkinen, 1995]. Thus, transport of sea ice, which is an important part of the global hydrological cycle, should affect the global thermohaline circulation. From such a point of view, this study intends to investigate how sea ice affects the global thermohaline circulation and salt transport in the deep ocean by using a sea-ice-ocean coupled model. [5] There have been several modeling studies that examine the role of sea ice in the global thermohaline circulation and salinity of the deep ocean. England [1993] simulates the present-day world ocean in an ocean general circulation model (OGCM) without sea ice. He employs values higher than the observed as the surface salinity boundary condition in the Southern Ocean to represent the effects of sea ice, and shows that this modification makes salinity of the deep water and the deep water circulation more realistic, although Toggweiler and Samuels [1995] argue that this sort of modification might distort the deep water formation process. However, models that explicitly represent sea ice are necessary to accurately evaluate its effects. In recent years, there have been a few studies by use of sea-ice-ocean coupled models to investigate relation between sea ice and the global thermohaline circulation. Stössel et al. [1998] use a free surface OGCM coupled to a sea ice model, and examine the roles of the Southern Ocean sea ice in affecting the deep water formation and hence the global thermohaline circulation. They deal with the sensitivity of the global thermohaline circulation to sea ice from the aspects of freshwater, heat, and momentum budgets. However, they restore surface salinity to observed data when and where sea ice does not exist, even in the polar oceans. This restoring might lead to underestimate of the effects of sea ice on salinity since it tends to reduce the difference in salinity between the cases with and without sea ice transport or salt and freshwater fluxes at the ice-ocean interface. Goosse and Fichefet [1999] use a sea-ice-ocean coupled model with a slope convection scheme and show that salt and freshwater fluxes induced by sea ice as well as cooling of surface water by ice-ocean heat exchange activate the deep water formation in the Southern Ocean, thus affecting the world ocean. However, they perform no experiment for studying the effect of sea ice transport. In addition, they also restore surface salinity to the observed, although their restoring is weak and limited to the middle and low latitudes where there is no ice. [6] The focus of this study is on estimating the effect of freshwater budget associated with sea ice on the global thermohaline circulation. For this purpose, we conduct numerical experiments with an ice ocean coupled general circulation model by artificially changing sea ice physics. Then, it is a prerequisite that the remaining freshwater components (i.e., precipitation, evaporation, and river runoff) are realistic and identical among all the experiments. If restoring to the observed surface salinity is employed, the restoring flux is not identical among the experiments unless the sea surface salinity (SSS) is identical. Consequently, it is difficult to distinguish the effect of sea ice induced flux changes from those caused by the change in restoring flux. Therefore, we drive the sea-ice-ocean coupled model without surface salinity restoring. [7] The paper is organized as follows. The model and the boundary conditions used in this study are outlined in section 2. The control case representing the present state of the ocean is described in section 3. The settings and results of sensitivity experiments are presented in section 4. Finally, summary and discussions are presented in section Model and Forcing 2.1. Ocean Model [8] The OGCM used in this study is CCSR Ocean Component Model (COCO) version 3. Configuration of the model is almost identical to that used by Hasumi [2002a]. Only a brief description is shown here. COCO version 3 is a z coordinate, free surface model. Sea surface height is explicitly predicted by the method of Killworth et al. [1991]. The model includes the uniformly third-order polynomial interpolation algorithm (UTOPIA) [Leonard et al., 1993] for tracer advection, and isopycnal diffusion [Cox, 1987]. The horizontal resolution is about 2.8 both in the zonal and the meridional direction. There are 40 vertical levels, whose grid width varies from 50 m for the top level to 200 m for the bottom level. The deepest level is at 5500 m depth. The model domain is global, although the northernmost two grids are treated as land. The isopycnal diffusion coefficient is m 2 s 1. The background Laplacian horizontal diffusion coefficient is m 2 s 1. The vertical diffusion coefficient is prescribed as a function of depth, varying from m 2 s 1 at the surface to m 2 s 1 at the bottom (case III of Tsujino et al. [2000]). Coefficients for the vertical and the horizontal eddy viscosity are m 2 s 1 and m 2 s 1, respectively. Convection is represented by a simple convective adjustment scheme Sea Ice Model [9] The sea ice model has both the thermodynamic and dynamic components. The thermodynamic part is the zero layer model of Semtner [1976]. In the dynamic part, the momentum equation and the equations for mass and concentration are taken from Mellor and Kantha [1989], although harmonic and biharmonic diffusion terms are adopted in the latter two equations, whose coefficients are m 2 s 1 and m 4 s 1, respectively. Treatment for lateral melting/freezing also follows the method of Mellor and Kantha [1989], where sea ice concentration changes with an empirical factor under a predicted change of sea ice mass due to melting or freezing. Internal ice stress is formulated by the elastic-visco-plastic rheology [Hunke and Dukowicz, 1997]. Leads are formed explicitly by sea ice motion. Any parameterization for lead formation is not applied. Since we use monthly mean climatological wind forcing, the formation of leads may be underestimated. Sea surface temperature remains at the freezing point where sea ice exists. Salinity of sea ice is fixed at 5 psu. We assume the constant water turning angle [Hibler, 1979] of 25 in the Northern Hemisphere and 25 in the Southern Hemisphere for the drag between sea-ice and the ocean.

3 KOMURO AND HASUMI: SEA ICE EFFECTS ON THE GLOBAL CIRCULATION Surface Boundary Conditions [10] The model is driven by the surface heat, freshwater, and momentum fluxes, given as monthly mean data. All these fluxes for the boundary conditions are obtained from NCEP reanalysis daily data [Kalnay et al., 1996] averaged over 40 years from 1961 to Flato and Hibler [1992] point out that sea ice motion becomes slower when the wind forcing is monthly averaged with viscous-plastic rheology. Nevertheless, sea ice transport in this study is realistic especially in the Southern Hemisphere, as will be discussed later. [11] The momentum flux is directly taken from the NCEP wind stress data set. The heat flux F H is calculated after the method of Haney [1971], F H ¼ Q 2 T A * T S ; ð1þ where T S is the temperature at the top of sea ice or the sea surface temperature. The coefficient Q 2 and the apparent air temperature T* A are calculated using the NCEP skin temperature, air temperature at 2 m, wind speed at 10 m, wind stress, net shortwave and longwave radiative fluxes at surface, specific humidity at 2 m, relative humidity at sigma 0.995, and sea level pressure. In the Atlantic to the north of 49 N, T * A is artificially lowered when and where sea ice is not present and the ocean is losing heat in the NCEP data, in order to reproduce realistic NADW transport in the control experiment. The decrease in annual mean T* A averaged zonally in the Atlantic is 0.3 C at 49 N and gradually increasing up to 4.9 C at 67 N. The zonally averaged, annual mean meridional profile of T* A is depicted in Figure 1. For the freshwater flux, evaporation E and precipitation P are prescribed by the NCEP data. Rain and snow are not distinguished in the original data, thus precipitation is regarded as snowfall when 0.65 T A T skin <0 C, where T A is the air temperature at 2 m height and T skin is the skin temperature of the NCEP data (this condition is described in a note about the NCEP model at ftp:// wesley.wwb.noaa.gov/pub/reanal/random_notes/model). Runoff R is diagnosed from the NCEP land runoff data by using the river routing subprogram of CCSR/NIES AGCM [Numaguti et al., 1997]. The horizontal resolution of the Figure 1. Zonnaly averaged, annual mean meridional profile of T *. Figure 2. Zonally averaged annual mean distribution of freshwater flux given as the boundary condition. The solid line is for all the basins, dashed line is for the Atlantic Ocean, dotted line is for the Pacific Ocean, and dash-dotted line is for the Indian Ocean. river routing map used here is about 5.6. The amount of R calculated in this way is 0.32 Sv (1 Sv = 10 6 m 3 s 1 )inthe Arctic Ocean and 0.35 Sv in the Southern Ocean. These values are too large compared to the observed river runoff in the Arctic Ocean (0.08 Sv by Baumgartner and Reichel [1975] and 0.10 Sv by Aagaard and Carmack [1989]) and the observed runoff water including calving icebergs from Antarctica (0.08 Sv by Jacobs et al. [1992]). Since the global thermohaline circulation is not realistically reproduced with R diagnosed from the NCEP data, the above mentioned observed values by Baumgartner and Reichel and Jacobs et al. are used in the Arctic Ocean and the Southern Ocean, respectively. R is also larger than the observed value of Baumgartner and Reichel by about 70% in the Indian Ocean, although R in the Atlantic and the Pacific is in good agreement with the same data. However, R in the Indian Ocean is not modified since P E is much larger than R. Finally, freshwater flux F W = P E + R is adjusted so that the globally integrated F W becomes zero. This freshwater flux is applied to all the experiments, and no salinity restoring is employed. The annual mean meridional profiles of F W for each basin and all the basins are shown in Figure Control Experiment [12] An experiment is performed in order to reproduce the present-day state of the ocean and sea ice. This control case is hereafter referred to as CTRL. CTRL is initiated from a state of rest and constant temperature (5 C) and salinity (34.7 psu), and integrated for 3100 years. Note that the globally averaged salinity remains unchanged over the whole period of the integration as the amounts of water and salt in the model are preserved. After the integration, no long-term trend is found in physical quantities in the deep ocean. The results discussed below are based on the average over the last 100 years of the integration. [13] Figure 3 illustrates the zonally integrated overturning circulation in the Atlantic and the Pacific. In the Atlantic,

4 29-4 KOMURO AND HASUMI: SEA ICE EFFECTS ON THE GLOBAL CIRCULATION Figure 4. Annual mean sea ice thickness and extent for CTRL (solid lines) and extent from the observational estimates of Nomura [1993] (dashed line) in the Northern Hemisphere. The ice edge is defined as the sea ice concentration of 15%. Contour interval for solid lines is 0.5 m. Figure 3. Stream function of zonally integrated meridional circulation for CTRL in (a) the Atlantic and (b) the Pacific. Contour interval is 1 Sv. there are two deep circulation cells, as in the real ocean: the upper one is associated with the formation of North Atlantic Deep Water (NADW) and the lower one is associated with the formation of Antarctic Bottom Water (AABW). The transport across 33 S is 14.9 Sv for NADW and 4.9 Sv for AABW. In an observational estimate, these values are 14 Sv for NADW and 4 Sv for AABW [Schmitz, 1995]. Thus, it can be said that the overturning circulation in the Atlantic in this model is realistically simulated. In the Pacific, the circulation originating in the Circumpolar Deep Water (CDW) formation occupies the deep layer. The transport of CDW across 11 S is 9.5 Sv. This value is close to the observations of about 10 Sv through the Samoan Passage at 10 S [Roemmich et al., 1996; Rudnick, 1997]. The modeled transport at the southernmost latitude of the Pacific is 11.5 Sv, which is within the range of the observational estimation of Sv [Schmitz, 1995]. [14] Figures 4 and 5 show the annual mean sea ice thickness and extent in the Northern and the Southern Hemisphere, respectively. The sea ice extent is quite reasonable, partly because of the thermal boundary condition that implicitly contains the sea ice information. That is, in the original NCEP data surface air temperature is very low and net downward shortwave radiative flux is small where sea ice exists. Since T* A is derived from such data, T* A tends to be very low where sea ice should exist in reality. There are some differences, however, in sea ice extent between the observation and the model result. For example, the sea ice in the Greenland-Iceland-Norwegian Seas (GIN Seas) and the Labrador Sea is too extensive, especially in winter (the figure is not shown). This may be because the Norwegian Current, which flows northward along the Norwegian coast, is too weak in the model. The reason and influence of these differences on results will be discussed later. [15] In this study, the amount of sea ice transport is more important than the amount of sea ice itself. Aargaard and Carmack [1989] estimate that annual mean southward sea ice transport through the Denmark Strait is Sv. The corresponding value in the model is Sv, which is larger than their estimation. On the other hand, annual mean net export from the Weddell Sea is Sv in the model. This value is in good agreement with the observational Same as Figure 4 but in the Southern Hemi- Figure 5. sphere.

5 KOMURO AND HASUMI: SEA ICE EFFECTS ON THE GLOBAL CIRCULATION 29-5 estimation of 0.05 Sv [Harms et al., 2001]. An observational estimate of annual mean net export from the Ross Sea continental shelf is 0.95 m yr 1 [Jacobs et al., 1985]. The Ross Sea continental shelf is not well resolved in the model, but modeled annual mean export from the whole Ross Sea (to the south of 72 S in the Pacific sector) is 1.00 m yr 1 (0.041 Sv). Therefore, the amount of the sea ice transport in the Southern Hemisphere is quite reasonable compared to the observations. [16] Figures 6a and 6b show the winter sea ice concentration and deep convection sites around Antarctica, respectively. Figure 6c depicts effective surface salt flux F S, which is defined as F S ¼ S s F W 1 S I S s P I ; ð2þ where S s is the modeled SSS, S I is the salinity of sea ice, and P I is the net production rate of sea ice. Positive F S works to raise SSS. The deep convection in the Atlantic sector mainly occurs under compact sea ice cover. Zonally averaged temperature in the Atlantic sector around 70 S is near the freezing point from the top to the bottom (Figure 7a). These results indicate that salinity is the destabilizing factor for vertical stratification to induce deep convection in this region, and that the sea ice there does not melt away as the deep convection transports little heat to the surface layer. On the other hand, in the Indo-Pacific sector, low sea ice concentration is found where deep convection occurs. F S at the deep convection sites freshens the surface water there. Zonally averaged temperature in the corresponding sector (Figure 8a) is significantly higher than the freezing point in the deep layer. These results show that the deep convection there is induced by thermal destabilization of the water column, and that the deep convection transports more heat to the surface than that in the Atlantic sector, which eventually melts sea ice. [17] In the regions where deep convection takes place, buoyancy is lost at the sea surface. Generally, buoyancy is lost near Antarctica and gained at lower latitudes. Here, the total buoyancy loss is calculated by integrating the surface buoyancy flux from the southernmost grids to the latitudes where the flux changes the sign. It is kg s 1 for the Atlantic sector (68 W 34 E, integrated to the south of 66 S) and kg s 1 for the Indo-Pacific sector (the remaining part, integrated to the south of 69 S). In this sense, the deep convection in the Atlantic sector and that in the Indo-Pacific sector make a comparable contribution to the deep water formation in the Southern Ocean. [18] In the real ocean, dense shelf water is formed on the continental shelf in the Weddell Sea where the temperature is near the freezing point at all the depths. Then the shelf water flows down on the continental slope, entraining the overlaying warmer waters, and finally becomes AABW [Price and Baringer, 1994]. In the model s Atlantic sector, the deep convection occurs in the open ocean, where the temperature is near the freezing point from the surface to the bottom as on the continental shelf in the real ocean. Then the formed deep water is mixed with the warmer water at lower latitudes. On the other hand, in the Indo-Pacific sector, the deep convection directly mixes the surface cold water with the deep warmer water. Because of the low sea ice concentration, more heat is released from the surface in Figure 6. (a) Sea ice concentration (contours are for 0.15, 0.50, and 0.85; the shaded area is under 0.15) and (b) the maximum depth of convection (over 1000 m) for CTRL in the Southern Hemisphere in winter (from June to September). (c) Annual mean surface salt flux F S (the area where F S is positive is shaded, contour interval is kg m 2 s 1 except for the contour of ) for CTRL in the Southern Hemisphere. this region, resulting in unrealistic distribution of the surface heat flux. However, the amount of the deep water formed in the whole Southern Ocean is realistic, as mentioned above, which means that the total buoyancy loss in the Southern

6 29-6 KOMURO AND HASUMI: SEA ICE EFFECTS ON THE GLOBAL CIRCULATION formation in winter locally balances melting in summer, and annual mean freshwater flux by sea ice is almost zero everywhere. Heat and momentum fluxes in these cases are expected to significantly differ from those in CTRL, as the lack of motion changes ice ocean drag, sea ice extent, and sea ice thickness. In NF-S, NF-N, and NF-G, no freshwater and salt fluxes at the ice-ocean interface exist. Therefore, melting or freezing of sea ice have no effect on SSS. The difference in heat and momentum fluxes between these cases and CTRL is expected to be small. Differences between the cases without sea ice motion and those without freshwater and salt fluxes at the ice-ocean interface show the influence of sea ice on heat and momentum fluxes. [21] For all the experiments, the other settings are identical to those for CTRL. The period of integration is 3100 years, and the data averaged over the last 100 years are used Sensitivity to Southern Hemisphere Sea Ice Effect [22] Figure 9 shows effective surface salt flux F S around Antarctica for NM-S and NF-S. In NF-S, NF-N, and NF-G, F S is defined by F S = S s F W in the region where freshwater and salt fluxes at the ice-ocean interface are turned off (the Figure 7. Zonally averaged annual mean (a) potential temperature (contour interval is 0.5 C) and (b) salinity (contour interval is 0.1 psu) in the Atlantic sector for CTRL. Ocean is realistic. Thus, the simulated heat loss in the whole Southern Ocean is reasonable as the haline part of the density flux (the surface freshwater flux and the effects of sea ice) is realistic, although the distribution of heat flux is not realistically reproduced. Since we argue mainly about the effect of sea ice on F S, sensitivity experiments carried out below seem to make sense. 4. Sensitivity Experiments [19] Various factors such as brine rejection, freshwater release, and spatial separation of regions for net sea ice production and melting induced by sea ice motion affect surface salinity. We perform sensitivity experiments to examine the role of these effects in the deep water formation and hence the global thermohaline circulation. [20] Seven experiments, including CTRL, are carried out (Table 1). In NM-S, NM-N, and NM-G, there is no sea ice motion except for weak diffusion (the coefficients are 2 and 4 order smaller in the harmonic and biharmonic terms, respectively), which is adopted to stabilize the calculation. Freshwater flux induced by sea ice still exists and affects SSS. However, when a steady state is achieved, sea ice Figure 8. for CTRL. Same as Figure 7 but in the Indo-Pacific sector

7 KOMURO AND HASUMI: SEA ICE EFFECTS ON THE GLOBAL CIRCULATION 29-7 Southern Hemisphere for NF-S, the Northern Hemisphere for NF-N, and the world ocean for NF-G). The differences of F S among the cases are caused only by the sea ice induced freshwater flux. The freshwater flux acts to freshen the surface water in the most of the area of the Southern Ocean (Figure 9b). This is because precipitation exceeds evaporation in this region, and there are icebergs and meltwater of the ice sheet from Antarctica, which is treated as river runoff in the model. However, there are regions where F S is positive (salinifying the surface water) near Antarctica for CTRL (Figure 6c). In the modeled Southern Ocean, sea ice is transported toward lower latitudes, as is the case in the observation [Emery et al., 1997], and there is net production of sea ice near Antarctica. This is consistent with the freshwater budget analysis on the Ross Sea continental shelf [Jacobs et al., 1985]. The modeled sea ice transport in the Weddell Sea is also in good agreement with the observation, as described in the previous section. For these reasons, the salinifying effect of sea ice near Antarctica seems to be realistic. On the other hand, net sea ice production is very small in NM-S because of the lack of sea ice transport, so the spatial distribution of F S in NM-S (Figure 9a) resembles that in NF-S (Figure 9b). [23] The salt flux has a large influence on convection occurring in the sea below. Brine rejection under sea ice destabilizes the stratification of the water column underneath and leads to the deep convection in the Atlantic sector for CTRL. There is a large area of positive F S in the sector, that is, the surface salinity is increased by the salt flux there. In contrast, in the Atlantic sector for NM-S, deep convection occurs where sea ice concentration is small (Figure 10), and the zonal mean temperature at the corresponding latitudes is significantly higher than the freezing point (Figure 11a). These results show that the deep convection is caused by the thermal destabilization of the water column. Hereafter, we call the kind of the convection occurring in CTRL as salinity-driven convection and in NM-S as temperature-driven convection. [24] The difference between CTRL and NM-S is interpreted as follows. Salinity-driven convection brings up less saline deep water to the surface. If such convection takes place every winter, the surface water must be salinified by another source in order to maintain the convection. The transport of sea ice is responsible for the salinity increase in CTRL. However, there is no sea ice transport in the Southern Ocean for NM-S, so salinity-driven convection is not maintained. Instead, temperature-driven convection occurs. In the real ocean, the convection on the continental shelf is Table 1. List of Experiments Abbreviation Description CTRL control case NM-S no sea ice motion in the Southern Hemisphere NM-N no sea ice motion in the Northern Hemisphere NM-G no sea ice motion in both the hemispheres NF-S no freshwater/salt flux at the ice-ocean interface in the Southern Hemisphere NF-N no freshwater/salt flux at the ice-ocean interface in the Northern Hemisphere NF-G no freshwater/salt flux at the ice-ocean interface in both the hemispheres Figure 9. Annual mean surface salt flux F S in the Southern Hemisphere for (a) NM-S and (b) NF-S. The area where F S is positive is shaded. Contour interval is kg m 2 s 1. salinity-driven [Killworth, 1983]. Therefore, it would not occur if sea ice were not exported. Thus, the response in the model to the lack of sea ice transport is realistic in that sense although the deep convection occurs not on the continental shelf but in the open ocean. [25] Corresponding to the change of the deep water formation, the temperature is higher and the salinity is lower in the Atlantic sector of the Southern Ocean for NM-S than for CTRL (Figures 7 and 11). These changes make the Atlantic bottom circulation weaker than in CTRL by 16% (Table 2). Note that the bottom layer salinity in the North Atlantic for NM-S is higher than that for CTRL. This is because the globally averaged salinity in the model ocean is the same for all the cases. The bottom water in the North Atlantic is affected strongly by NADW. The salinity is lower for NM-S than for CTRL in the Southern Ocean and the Indo-Pacific sector (Figure 12b), so the salinity of NADW must be higher than for CTRL. Therefore, the high salinity at the bottom of the North Atlantic is reasonable. [26] In the Pacific sector, the deep convection occurs mainly in the region where the sea ice concentration is small in both CTRL and NM-S. However, the zonal mean temperature in the Indo-Pacific sector of the Southern Ocean is higher for NM-S than for CTRL except the surface

8 29-8 KOMURO AND HASUMI: SEA ICE EFFECTS ON THE GLOBAL CIRCULATION kg s 1 at 33 S (the southernmost latitude at which the Atlantic Ocean is zonally enclosed by the continents) in the layer deeper than 3500 m depth (where zonally integrated salt transport is northward at 33 S in CTRL) and by kg s 1 at 33 N. These values are 7% and 12%, respectively, of the salt transport in CTRL. This shows that the change in F S by the sea ice transport in the Southern Ocean also increases the salt transport from the Southern Ocean to the global deep ocean. Moreover, the increase in the bottom layer is larger by an order of magnitude than that at the surface. The salt transport in the bottom layer is accompanied by the Atlantic bottom circulation and the Pacific deep circulation. The existence of sea ice transport not only increases salinity of the deep waters but also strengthens the circulations, resulting in amplification of salt transport. [29] In NF-S, features similar to those found in NM-S are recognized. The deep convection around Antarctica is temperature-driven. The Atlantic bottom circulation is decreased by 16% in NF-S compared with for CTRL (Table 2). The reduction rate is close to that for NM-S. The salt transport at 33 S and 33 N below 3500 m increases by kg s 1 and kg s 1, respectively. These values are also comparable with for NM-S. Since sea ice transport Figure 10. Same as Figures 6a and 6b but for NM-S. layer (Figures 8a and 12a). The larger vertical temperature gradient in NM-S indicates that thermal factor is more important in NM-S for the destabilization of the water column, although the deep convection in the Indo-Pacific sector is temperature-driven convection for NM-S as well as CTRL. In the deep layer of the Indo-Pacific sector higher temperature and lower salinity for NM-S than for CTRL are also realized (Figures 8 and 12). These changes in the deep convection and the deep water are similar to the change of convection from salinity-driven for CTRL to temperaturedriven for NM-S in the Atlantic sector, although they are not so clear as in the Atlantic sector. [27] Here we analyze salt transport in the ocean to quantitatively evaluate the changes in the deep circulation and salinity. It envisages how the sea ice induced salt flux affects the global salt transport in the ocean. Note that the globally integrated F S is not exactly zero, since F S depends on SSS. [28] Figure 13 depicts the difference in the salt transport between CTRL and NM-S. In CTRL, the northward transport of sea ice yields salt input of kg s 1 to the south of 63 S (where zonally integrated net sea ice production is positive in CTRL). Transported sea ice melts to the north of there, and roughly the same amount of salt is removed. This salt budget enhances the meridional haline circulation in the Southern Ocean [Hasumi and Suginohara, 1995]. In addition, the northward salt transport increases by Figure 11. NM-S. Same as Figure 7 but in the Atlantic sector for

9 KOMURO AND HASUMI: SEA ICE EFFECTS ON THE GLOBAL CIRCULATION 29-9 Table 2. Cross-Equatorial Flow of NADW in the Atlantic, AABW in the Atlantic, and CDW in the Pacific in Each Case (in Sverdrups) Case NADW AABW CDW CTRL NM-S NM-N NM-G NF-S NF-N NF-G is taken into account for NF-S, heat and momentum fluxes for NF-S are not so different from those for CTRL. On the other hand, no sea ice transport in NM-S leads to these fluxes quite different from those in CTRL. In addition, sea ice formation and melting cause seasonal variation of F S in NM-S, but not in NF-S. Therefore, the small difference in the global thermohaline circulation and the global salt transport between NM-S and NF-S suggests that sea ice affects the global thermohaline circulation primarily via changes in surface freshwater flux, and seasonal variation in it is of minor importance. Figure 12. for NM-S. Same as Figure 7 but in the Indo-Pacific sector Figure 13. Difference in zonally integrated salt flux between CTRL and NM-S ((CTRL) (NM-S)). Values are salt flux difference through each section ( positive means northward/upward). [30] In both NM-S and NF-S, while the AABW transport decreases, the cross-equatorial NADW transport increases by 2 4% in these cases compared with that in CTRL (Table 2). These changes are consistent with antiphase behaviors of the Atlantic deep and bottom circulations shown in many observations and modeling studies [Stocker, 1998] Sensitivity to Northern Hemisphere Sea Ice Effect [31] In NM-N and NF-N, where the sea ice s influence on salinity is turned off only in the Northern Hemisphere, the transport of the Atlantic deep circulation is approximately the same as in CTRL (Table 2). Nevertheless, Northern Hemisphere wintertime deep convection for these cases takes place in the Labrador Sea as well as the regions of the deep convection for CTRL (Figure 14). In the model, the Labrador Sea is the area where sea ice is transported from north and melts. The annual mean F S in NM-N and NF-N is larger than that of CTRL there, although the value is negative in most part (Figure 15). Thus, it is suggested that this salt flux activates the deep convection there. [32] In the northern North Atlantic and the GIN Seas, deep convection takes place in all the three cases. However, the convection in the GIN Seas makes little contribution to NADW since there is a strong front at around 64 N inthe Atlantic (Figure 7) and the Atlantic deep circulation is almost closed to the south of this latitude (Figure 3a). This is because outflow form the GIN Seas over sills is poorly reproduced in this model, although it is a common shortcoming in typical coarse resolution OGCMs [e.g., Hirst and McDougall, 1996]. The lack of outflow also causes the North Atlantic Current, which is surface inflow to the GIN Seas, too weak. The latitude of the front coincides with the wintertime edge of sea ice in CTRL (Figure 4). Consequently, there is little effect of sea ice on the NADW formation. [33] In contrast to the present model, deep convection occurs in the Arctic Ocean, the GIN Seas, and the Labrador Sea [Killworth, 1983]. Sea ice exists at least seasonally in these regions and plays an important role in the deep water formation there [Aagaard and Carmack, 1994]. Thus, our

10 29-10 KOMURO AND HASUMI: SEA ICE EFFECTS ON THE GLOBAL CIRCULATION NM-N and NF-N if the NADW formation becomes more realistic. Salinity-driven convection might occur in these cases, since F S is positive in some parts of the GIN Seas in these cases, although Häkkinen [1995] points out that brine rejection is not so important as heat loss for the deep convection in the Greenland Sea Comparison With Previous Studies [34] In NM-G and NF-G, sea ice dynamics or ice ocean salt flux is globally turned off. The resulting changes in Figure 14. The maximum depth of convection (over 1000 m) in the Northern Hemisphere in winter (from December to March) for (a) CTRL, (b) NM-N, and (c) NF-N. model is inadequate for arguing the sea ice effects on the NADW formation. Flows over sills in the northern North Atlantic is necessary to represent more realistic NADW formation process. Since there is net melting of sea ice in the GIN Seas, as confirmed by observations [Aagaard and Carmack, 1989], the NADW transport might be stronger in Figure 15. Annual mean surface salt flux F S in the Northern Hemisphere for (a) CTRL, (b) NM-N, and (c) NF-N. Contour interval is kg m 2 s 1.

11 KOMURO AND HASUMI: SEA ICE EFFECTS ON THE GLOBAL CIRCULATION strength of the Atlantic deep and bottom circulations from CTRL are qualitatively similar to those for NM-S and NF-S (Table 2). [35] Stössel et al. [1998] perform an experiment corresponding to NM-G (named NID in their study). In that case, little change is found in the strength of SH cell, which is defined by the deep minimum value of the stream function of global meridional overturning circulation and used as an index for the strength of the deep circulation originating in the Southern Ocean. In contrast, the strength of SH cell in our model decreases by about 10%, from 21.2 Sv for CTRL to 19.1 Sv for NM-G. This difference may be caused by SSS restoring applied by Stössel et al. Although their SSS restoring is limited where sea ice does not exist, it would reduce the difference between the cases since most of the sea ice cover in the Southern Ocean melts in summer. Therefore, they might underestimate the effects of sea ice transport on the global thermohaline circulation. [36] Stössel et al. [1998] perform another experiment corresponding to NF-G (named SSI in their study). Goosse and Fichefet [1999] also perform an experiment corresponding to NF-G (named NSF in their study). The circulation originating from the Southern Ocean is weakened in both cases compared to respective control cases. The strength of SH cell also decreases in our NF-G to 19.5 Sv. However, salinity in the deep layer of the Southern Ocean increases as given by Stössel et al. [1998] but decreases as given by Goosse and Fichefet [1999]. Goosse and Fichefet point out this attributes to different processes for the deep water formation, i.e., the deep water is formed only by open ocean convection by Stössel et al. [1998], while the deep water formation also occurs on the continental shelf in a control case of Goosse and Fichefet [1999]. The deep water formation on the continental shelf is suppressed in the case NSF, resulting in the opposite response in salinity between those two studies. In this study, the deep water is formed only by open ocean convection but the salinity in the deep layer of the Southern Ocean decreases in NF-G. The freshening in our result is caused by the change in the type of convection, from salinitydriven to temperature-driven. The convection type is controlled by the vertical stratification there. Thus, the opposite change in the deep layer salinity between the case SSI in Stössel et al. [1998] and our NF-G may be because of the difference in the stratification in the Southern Ocean between them. In the control case of Goosse and Fichefet [1999], the deep water is formed by convection on the continental shelf, which is salinity-driven, and by open ocean convection, which is temperature-driven. In this regard, the suppression of the convection on the continental shelf in the case NSF of Goosse and Fichefet [1999] corresponds to the change in the convection type from salinity-driven to temperature-driven. What is essential to the response of deep ocean salinity is the type of convection rather than the region of convection, that is, whether convection takes place on the continental shelf or in the open ocean. 5. Summary and Discussion [37] In this study, we have evaluated the effects of sea ice on the deep water formation and hence the global thermohaline circulation by using a sea-ice-ocean coupled model. The fixed freshwater flux is employed for the boundary condition and SSS restoring is not applied. [38] In NM-S, equatorward transport of sea ice in the Southern Ocean and associated salinification around Antarctica is turned off. Consequently, salinity-driven convection for CTRL is changed to temperature-driven for NM-S, resulting in 16% weakening of the Atlantic bottom circulation for NM-S. The evaluation of the global salt transport shows that the surface salt flux induced by the sea ice transport has an amplifying effect on the global salt transport. The intensification of the thermohaline circulation originating from the Southern Ocean is responsible for this amplification. For NF-S, the distribution of heat and momentum fluxes differs from that in NM-S, while the thermohaline circulation and global salt transport are qualitatively and quantitatively similar to those in NM-S. The change in sea-ice motion alters all of the three forcing fluxes to the ocean, namely, heat, freshwater, and momentum fluxes, while ignoring sea ice induced freshwater flux does not much affect the other two fluxes. The similarity in results thus supports the idea that the influence of sea ice on freshwater flux is more important than heat and momentum fluxes in affecting the global thermohaline circulation. Furthermore, the result also suggests that seasonal variation in freshwater flux has minor importance. [39] In the northern North Atlantic, the deep water formation occurs mainly to the south of the sea ice edge. Thus, the influence of sea ice on the deep water formation is weak except in the Labrador Sea, where the sea ice transport tends to suppress the deep convection in CTRL. It is difficult to discuss about the sea ice effects on the NADW formation by the present results since the model fails to reproduce its source water formation under sea ice cover. [40] If surface salinity is restored to the observed values, the flux induced by the restoring tends to diminish the difference of the freshwater flux generated by sea ice transport, especially at high latitudes. The restoring also causes uncontrollable differences in surface freshwater flux when sensitivity experiments are performed. Goosse and Fichefet [1999] apply SSS restoring only at middle and low latitudes where no sea ice appears throughout the year. If the SSS in NF-G is restored to the values obtained in CTRL with the same time constant as theirs (300 days for 50 m depth), the resulting salt input integrated between 54 S and 54 N is kg s 1. This value is about four times larger than the salt input induced by sea ice transport around Antarctica. Therefore, we think the restoring boundary condition should not be used even if it is applied only in low latitudes in order to perform reliable experiments on effects of sea ice induced freshwater flux on the global ocean. [41] The fixed freshwater flux, on the other hand, causes some shortcomings in the present study, for example, poor reproductivity of SSS and vertical haline stratification. The realistic field of physical quantities is desirable to perform sensitivity experiments. Since numerical models are still imperfect and observations of surface freshwater flux are not adequate, the restoring boundary condition may be preferable in experiments in which freshwater flux is not mainly concerned, as in the experiments investigating the effects of heat and momentum fluxes by Stössel et al.

12 29-12 KOMURO AND HASUMI: SEA ICE EFFECTS ON THE GLOBAL CIRCULATION [1998] and Goosse and Fichefet [1999]. In these experiments, Stössel et al. show that an effect of sea ice (and snow on it) on insulating heat flux and an interaction of sea ice and wind variation are also important, and Goosse and Fichefet indicate that a cooling of surface water by contacting with sea ice is important. We suggest the relative importance of the freshwater flux compared with the other two fluxes in the present study. Considering that the reproductivity of our model is not sufficient, however, more investigation is needed for understanding what effect of sea ice is dominant, or relatively important, on the global thermohaline circulation. [42] Another shortcoming in this model is in the processes of the deep water formation. Convection on continental shelves is the main source of the deep waters around Antarctica, though the deep convection at sensible heat polynyas in the open ocean also makes some contribution [Killworth, 1983]. In our model, by contrast, deep water formation occurs only in the open ocean. The convection on the continental shelves occurring in the real ocean is salinity-driven. It would not be maintained without the sea ice export. In this sense, the salinity-driven convection in our results corresponds to that on continental shelves in the real ocean. Thus, our result of NM-S and NF-S seems qualitatively correct from a viewpoint of the effect of sea ice transport. However, it might affect the quantitative aspect. For example, when dense water formed on the continental shelves becomes deep water, it flows down on the continental slope and entrains the surrounding water. This process is not represented in this model. If it is reproduced, the properties and formation rate of the deep water may be affected. In addition, sensible heat polynyas in the open ocean are not formed in CTRL. According to a numerical experiment that reproduces sensible heat polynyas [Goosse and Fichefet, 2001], both transport of sea ice from polynya regions and heat supply from deep layer are needed to form and maintain polynyas. Since salinity-driven convection occurs in CTRL, supply of the heat from the deep layer is not sufficient and polynyas are not formed. This may result in an overestimate of the effects of sea ice transport on the global thermohaline circulation. In the Northern Hemisphere, on the other hand, the modeled deep water formation is unrealistically located at lower latitudes. The simple convective adjustment scheme might also affect the formation processes. Stössel et al. [2002] show that the use of a subgrid-scale plume convection parameterization reduces open ocean convection in the Southern Ocean, and increases sensitivities to sea ice effects. Thus a change in the convection scheme to a plume convection parameterization might affect deep water formation and change sensitivities in our model. [43] To realistically represent the deep water formation processes in OGCMs, downslope flow and flow over a sill must be reproduced well. Explicit representation of such flows requires very high resolution in both vertical and horizontal [Hasumi, 2002b]. Thus, a parameterization for bottom boundary processes [e.g., Nakano and Suginohara, 2002] is needed. Realistic reproduction of water masses in the source regions of the deep waters is also necessary for that. In this sense, the modeling of polar oceans, especially the GIN Seas and the Arctic Ocean, should be improved for the sake of global ocean modeling. [44] Acknowledgments. 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