Restriction of convective depth in the Weddell Sea

Transcription

1 GEOPHYSICAL RESEARCH LETTERS, VOL. 34, L10610, doi: /2007gl029295, 2007 Restriction of convective depth in the Weddell Sea K. Akitomo 1 Received 8 January 2007; revised 11 March 2007; accepted 25 April 2007; published 31 May [1] The depth open-ocean convection can reach around Maud Rise in the Weddell Sea is estimated based on the entrainment assumption using hydrographic data obtained in winter 1986 during which no distinct polynya was observed. While the water column is stable at all CTD stations when observed, additional cooling makes the mixed layer water heavier than the underlying water at any depth in an adiabatic sense. However, convective plume cannot reach the ocean bottom since it loses positive density anomaly due to thermobaricity by entraining the Warm Deep Water (WDW) on its downward way. The deeper convection is located over Maud Rise where the maximum temperature of the WDW, q max, is about 0.5 C while the shallower one in the warm water cell region where q max is more than 1.0 C. The mean convective depth is much shallower than the ocean depth and the ventilated depth during the Weddell Polynya years (34 km). Citation: Akitomo, K. (2007), Restriction of convective depth in the Weddell Sea, Geophys. Res. Lett., 34, L10610, doi: / 2007GL Introduction [2] A large ice-free region, called the Weddell Polynya, observed around Maud Rise during the winter seasons in the mid 1970s [e.g., Carsey, 1980], is a manifestation of openocean deep convection transporting a huge amount of heat upward and causing significant cooling in the deep layer [Martinson et al., 1981; Gordon, 1982]. The geographical feature of Maud Rise provides the conditions favorable to open-ocean deep convection, and it holds true even in usual (no-polynya) winters [e.g., Alverson and Owens, 1996]. McPhee [2000, 2003] reported that a marginal stability condition for thermobaric deep convection (TDC) was encountered around Maud Rise in winter 1994, during which a small polynya or a low ice-concentration was detected around Maud Rise [Drinkwater, 1997]. In numerical experiments, however, the depth of convective plume is at most 1.5 km over Maud Rise in usual winters [Akitomo et al., 1995; Akitomo, 2006], which is much less than the ventilated depth during the Weddell Polynya years (34km [Gordon, 1978, 1982]), and implies smaller upward heat transport (or weaker cooling at depths). [3] In the present study, the possible depth of TDC in usual winters is estimated based on the entrainment assumption [Turner, 1973] using hydrographic data, and its relation to upward heat transport is discussed. Open-ocean deep convection in the Weddell Sea which belongs to Type 2 TDC is characterized by two factors [Akitomo, 1999a, 1 Department of Geophysics, Kyoto University, Kyoto, Japan. Copyright 2007 by the American Geophysical Union /07/2007GL b]; thermobaricity of seawater, dependency of the thermal expansion rate on pressure (water depth), and a two-layered structure of the water column in which the cold fresh mixed layer (ML) overlies the warm saline deep layer. These factors play a central role in determining the depth of TDC, again. 2. Data and Methods [4] Hydrographic data is from the cruise ANT 5/2 of R/V Polarstern around Maud Rise in the Weddell Sea in austral winter 1986 [Huber et al., 1989], during which no distinct polynya was observed, i.e., no-polynya year. Figure 1 shows locations of 127 CTD stations used here, which were covered in 49 days from 18 July (Station 5) to 4 September (Station 138). CTD casts reach the ocean bottom at 59 stations and only 700 m at 68 stations. Therefore, the deeper profiles of potential temperature q and salinity S at the shallow stations are linearly interpolated using those at the adjacent deep stations. Figures 2a and 2b show the vertical section of q along the cruise route. The Warm Deep Water (WDW) occupies just below the cold ML, and some extreme warm regions of several hundreds of kilometers, called warm water cells, are found in the WDW [Gordon and Huber, 1984]. S has a similar distribution to q; lower in the ML, higher in the WDW and a peak in the warm water cell (not shown). [5] Figure 2c shows the density difference of the surrounding deep water and ML water, Dr = r sr r ml, for the observed profiles of q and S when the ML water is adiabatically lowered to depth z. Dr has a positive peak just below the ML (thick dashed line) and decreases downward (it becomes negative above Maud Rise and continental slope of Antarctica). This vertical change of Dr is because the contribution of q to density increases downward by thermobaricity. Since positive Dr below the ML acts as a barrier to convective motion (thermobaric barrier [McPhee, 2003]), the water column is on the whole stable at all stations. The barrier thickness is thinner at stations above Maud Rise than at other stations and stability of the water column is weakest [e.g., Gordon and Huber, 1990]. Particularly, it is less than 1 km at Stations (0.3 km in minimum at Station 108) surveyed during August, which is about half those at Stations surveyed about one month before (from 27 July to 2 August). [6] The depth of TDC triggered by subsequent cooling (brine ejection) is estimated as follows. First, we seek the conditions of the ML properties such as its salinity S ml and depth h ml for the water column to be marginally stable. For this, keeping the ML temperature q ml at the freezing point q f (= S ml S 3/2 ml S 2 ml ), S ml is repeatedly increased by 10 4 at a time until the barrier layer disappears. During this procedure, the L of5

2 Figure 1. Cruise ANT V/2 of R/V Polarstern around Maud Rise in the Weddell Sea during austral winter 1986, with bottom topography. Contour interval is 1 km. ML is deepened and S ml is reset to the mean value over a new ML when the salinized ML water becomes heavier than the water just below. At the marginally stable state, comparing the profiles of q and S with the observed ones, we estimate the total efflux of heat needed for keeping q ml at q f and forming sea ice (salinizing the ML). Although salinity of sea ice (or brine ejection) changes with its thickness, we assume it to be 5. This assumption may bring some uncertainties in the estimation of ice thickness (heat flux) needed for instability, but does not affect the depth of TDC. [7] Next, density change of a convective plume with depth z (positive downward) is estimated by the entrainment assumption [Turner, 1973]. The volume V, horizontal radius r, and vertical velocity w of convective plume are related by the equations, dv/dz =2aV/r, dr/dz =2a rg 0 /2w 2, and dw 2 /dz =2g 0 4aw 2 /r. a is the entrainment coefficient (0.1), and g 0 is the reduced gravity, (r pl r sr )g/r 0, where r pl (r 0 ) is the plume (reference) density. Using these relations with the initial values, r pl,0 (= r ml ), r 0, w 0, and V 0 (= p r 2 0 h ml ), we calculate the plume density r pl at depth z and then determine the convective depth H tdc where r pl becomes equal to r sr. The calculation is done at discretized vertical levels with a 10-m interval. While experiments are executed with various combinations of r 0 and w 0 since their values are somewhat arbitrary, the results will be shown for three different values of r 0, 0.25, 0.50, and 1.00 km with constant w 0 of 0.05 ms 1. Observations and model experiments suggest that convective size(2r 0 ) and velocity w 0 are at most 1 km and 0.1 ms Results [8] After the imposed cooling (brine ejection due to ice formation) removes the barrier layer at each station, the unstable (Dr < 0) region occupies from the mixed layer base to the ocean bottom (Figure 3a). The amount of cooling rate Q needed to establish this marginal state until 30 September 1986 is shown in Figure 4a. The result is similar to the result by McPhee [2003]. Only weak cooling (Q <50Wm 2 ) is needed over and to the south of Maud Rise while more than 100 Wm 2 far from the rise. The former drops in the actual range of heat flux (several tens of watts per squared meters), implying occurrence of TDC. The associated ice growth ( m) and salinization of the ML ( ) are not unrealistic despite uncertainties in the estimation (not shown). [9] The negative Dr throughout the water column at the marginal stable state means that the ML water would descend to the ocean bottom if the process were adiabatic. In reality, however, the process will proceed not adiabatically but diabatically; the descending water will entrain the surrounding water on its downward way. As seen in Figure 3b which shows Dr estimated from the entrainment assumption, H tdc is much shallower than the ocean depth. When r 0 is 0.25 km, for example, the maximum and minimum values are 1.69 km and 0.52 km, respectively, and the mean is 1.23 km. These values are not inconsistent with the model study by Akitomo [2006], and much smaller than the ocean depth and the ventilated depth during the Weddell Polynya years (34 km; Gordon, 1978, 1982). While H tdc increases with r 0 to reach about 3 km with r 0 = 1.0 km, the initial horizontal radius of convective plume is probably much smaller in the actual situation. [10] The small H tdc due to entrainment of surrounding deep warm water is not self-evident. Diabatic convective plume could reach the ocean bottom in a simple two-layered ocean with constant q and S in each layer. Thus, the profiles of q and S at depths are the key to restrict the convective depth in the actual situation. Indeed, there are significant peaks of q and S in the WDW km below the ML, and weak stratification at the deeper layer. Entraining a part of the WDW on the downward way, the sinking ML water becomes warmer and more saline. The warming reduces the thermobaric densification while the salinization densifies the sinking water. If the former overcomes the latter, the sinking water is lightened or not densified against the stratified deeper layer. As a result, the sinking water loses its density anomaly before reaching the ocean bottom. Significant negative correlation between H tdc (Figure 4b) and q max of the WDW (Figure 4c) supports this mechanism. The deeper (>1.5 km) convection appears over Maud Rise where q max is at most 0.5 C while the shallower (<0.9 km) one in the warm water cell region where q max is more than 1.0 C. The deeper convection may act to induce the larger upward heat transport, and then it can be a factor to lead to the observed reduction of ice thickness over the rise even in usual winters. [11] Maud Rise is said to have an effect to accelerate the onset of TDC by weakening stability of the water column [e.g., Alverson and Owens, 1996], and the warm water cell to delay the onset through a negative feedback process [Martinson, 1990]. Besides these effects on the onset of TDC, the present results show that the rise and warm water cell play a role in determining H tdc through thermobaricity after the onset. 4. Discussion [12] Open-ocean deep convection around Maud Rise in the Weddell Sea is largely affected by thermobaricity and a 2of5

3 Figure 2. Vertical section of q along the cruise route: (a) upper 1.5 km and (b) full depth. (c) Density difference Dr = r sr r ml between the surrounding and ML waters when the ML water is lowered adiabatically. Solid line indicates positive value; dotted line indicates negative value. Thick dashed line indicates the depth Dr has a positive peak at each station. Contour interval is 0.1 K in Figures 2a and 2b, and 0.01 kgm 3 in Figure 2c. two-layered structure with the cold fresh ML overlying the WDW (Type 2 TDC [Akitomo, 1999a, 1999b]). It is found here that these factors are also a key to determine the depth of TDC, H tdc. Entraining a part of the WDW and increasing its temperature on its downward way, the sinking water loses its density anomaly due to thermobaricity before it reaches the ocean bottom. This restriction of H tdc is a great contrast to the adiabatic scenario in which the ML water can sink to the ocean bottom once TDC starts. The predicted maximum (minimum) H tdc is found around Maud Rise (the warm water cell), showing that the structure of the WDW has a large influence on the depth of TDC. The averaged H tdc is much shallower than the ventilated depth during the Weddell Polynya years. Excessive salinity in the ML during summer may have accelerated the onset of TDC to much deeper ventilation in those years [e.g., Motoi et al., 1987]. [13] Despite the shallowness of TDC, the upward heat transport is still enough to melt out sea ice. If the water column is homogenized to the depth of 1.5 km by convective mixing and its temperature is 0.2 C, the heat transported upward from depths is estimated at Jm 2. This is about ten times as large as the heat needed to melt out sea ice with the thickness of 0.5 m, Jm 2. While this can be the reason why a small polynya has been observed in usual winters [e.g., Drinkwater, 1997], there must be other reasons why no large polynya has been observed after the Weddell Polynya years. A thin surface layer of cold but fresh water formed by ice melting may 3of5

4 Figure 3. (a) Same as in Figure 2c but for cooled until the stable barrier (positive Dr) layer vanishes. (b) Depth convection can reach H tdc, i.e., the depth on which Dr = 0, based on the entrainment assumption. Solid, dotted, and dashed lines indicate cases with r 0 = 0.25, 0.5, and 1.0 km, respectively. insulate sea ice from upwelled warm water. Intermittent TDC may fail the complete drainage of the ML water to depths within a cooling season. Akitomo et al. [1995] found a signature of the incomplete drainage. [14] Deep convection at high latitudes has a large effect on the long-term global climate change and then it is a key process needed to be properly built in an ocean general circulation model for its better performance. Comparing several parameterization schemes for vertical mixing in the Weddell Sea, Timmermann and Beckmann [2004] reported that the ocean penetrative plume scheme (OPPS [Paluszkiewicz and Romea, 1997]) based on the entrainment assumption exhibits a better performance for buoyancy-driven convection. While the OPPS represents the early stage of convective processes [Canuto et al., 2004], it may owe its performance to properties of TDC in the Weddell Sea, i.e., TDC occurs in isolated thermal plumes [e.g., Akitomo et al., 1995; Akitomo, 2006]. For universal parameterization, further investigation Figure 4. (a) Cooling rate needed for the water column to be unstable until 30 September (b) H tdc when sinking water entrains surrounding water based on the entrainment assumption with r 0 = 0.25 km. (c) Maximum potential temperature, q max, of the WDW. 4of5

5 is needed to clarify the detailed convective properties depending on the regions and the stages as well as thermobaricity, stratification, ocean-ice interaction, wind, and so on. References Akitomo, K. (1999a), Open-ocean deep convection due to thermobaricity, 1. Scaling argument, J. Geophys. Res., 104, Akitomo, K. (1999b), Open-ocean deep convection due to thermobaricity, 2. Numerical experiments, J. Geophys. Res., 104, Akitomo, K. (2006), Thermobaric deep convection, baroclinic instability, and their roles in vertical heat transport around Maud Rise in the Weddell Sea, J. Geophys. Res., 111, C09027, doi: /2005jc Akitomo, K., T. Awaji, and N. Imasato (1995), Open-ocean deep convection in the Weddell Sea: Two-dimensional numerical experiments with a nonhydrostatic model, Deep Sea Res., Part I, 42, Alverson, K., and W. B. Owens (1996), Topographic preconditioning of open-ocean deep convection, J. Phys. Oceanogr., 26, Canuto, V. M., A. Howard, P. Hogan, Y. Cheng, M. S. Dubovikov, and L. M. Montenegro (2004), Modeling ocean deep convection, Ocean Modell., 7, Carsey, F. D. (1980), Microwave observation of the Weddell Polynya, Mon. Weather Rev., 108, Drinkwater, M. R. (1997), Satellite microwave radar observations of climaterelated sea-ice anomalies, Proceedings of a Workshop on Polar Processes in Global Climate, November, 1996, pp , Am. Meteorol. Soc., Boston, Mass. Gordon, A. L. (1978), Deep Antarctic convection west of Maud Rise, J. Phys. Oceanogr., 8, Gordon, A. L. (1982), Weddell deep water variability, J. Mar. Res., 40, Gordon, A. L., and B. A. Huber (1984), Thermohaline stratification below the southern ocean sea ice, J. Geophys. Res., 89, Gordon, A. L., and B. A. Huber (1990), Southern ocean winter mixed layer, J. Geophys. Res., 95, 11,655 11,672. Huber, B. A., P. A. Mele, W. E. Haines, A. L. Gordon, J. C. Jennings, L. I. Gordon, R. F. Weiss, F. A. Van Woy, and P. K. Salameh (1989), CTD and hydrographic data from cruise ANT V/2 of R/V Polarstern, Tech. Rep. LDGO-89-3, Lamont-Doherty Geol. Obs., Palisades, N. Y. Martinson, D. G. (1990), Evolution of the Southern Ocean winter mixed layer and sea ice: Open ocean deepwater formation and ventilation, J. Geophys. Res., 95, 11,641 11,654. Martinson, D. G., P. D. Killworth, and A. L. Gordon (1981), A convective model for the Weddell Polynya, J. Phys. Oceanogr., 11, McPhee, M. G. (2000), Marginal thermobaric stability in the ice-covered upper ocean over Maud Rise, J. Phys. Oceanogr., 30, McPhee, M. G. (2003), Is thermobaricity a major factor in Southern Ocean ventilation?, Antarct. Sci., 15, Motoi, T., N. Ono, and M. Wakatsuchi (1987), A mechanism for the formation of the Weddell Polynya in 1974, J. Phys. Oceanogr., 17, Paluszkiewicz, T., and R. D. Romea (1997), A one-dimensional model for the parameterization of deep convection in the ocean, Dyn. Atmos. Oceans, 26, Timmermann, R., and A. Beckmann (2004), Parameterization of vertical mixing in the Weddell Sea, Ocean Modell., 6, Turner, J. S. (1973), Buoyancy Effects in Fluids, 367 pp., Cambridge Univ. Press, New York. K. Akitomo, Department of Geophysics, Graduate School of Science, Kyoto University, Kyoto , Japan. (akitomo@kugi.kyoto-u.ac.jp) 5of5