these waters have salinities greater than 34.5 %o while in northern Drake Passage the salinities are less than 34.1 %o. The relatively low salinity surface waters in northern Drake Passage originate in the southeastern Pacific along the southern coast of Chile (Silva and Neshyba, 1979/80), a region of heavy precipitation and continental runoff. This region may also be the source for the extensive tongue of low salinity surface water that extends westward from the South American continent between 40° and 45°S (Deacon, 1977b). The southward-flowing branch of the West Wind Drift transports the low salinity surface layer into the northern Drake Passage.
South of Australia, the front between the Subantarctic Zone and the Polar Frontal Zone can be identified by an abrupt southward decrease in both surface temperature and surface salinity (Gordon et al., 1977b). Burling (1961) called this feature the Australasian Subantarctic Front. Recognizing that this front might extend to other sectors, Houtman (1967) proposed the more general name, Subantarctic Front. In the Pacific, the property gradients across this front appear to diminish toward the east. In Drake Passage, the surface gradients are relatively weak, and, in fact, surface salinity increases toward the south (Sievers and Nowlin, 1984). Nowlin et al. (1977) selected the relatively sharp meridional density gradient at depth as a more reliable indicator of the position of the Subantarctic Front. The different salinity characteristics of the Subantarctic Front south of Australia and at Drake Passage illustrate a difficulty often encountered in attempting to identify a single objective indicator of frontal position that works equally well in all sectors. Indeed, this difficulty serves to illustrate one inherent difficulty in confirming the concept of circumpolar zonation.
In Fig. 3.6, the Subantarctic Front lies between stations 1189 and 405. Besides the intensified meridional gradients in temperature, salinity and density, the front is also the site where the subsurface salinity minimum associated with Antarctic Intermediate Water makes its rapid descent from near the surface to depths near 1,000 m. In Fig. 3.7, the location where the salinity minimum layer crosses 400 m and 600 m (about halfway through its descent) is used to bracket the position of the Subantarctic Front. The separation of these two isobaths is indicative of the sharpness of the front and the intensity of the current. In some areas, these isobaths are very widely separated. Perhaps this separation is somewhat misleading and is an artifact of using non-synoptic and widely-spaced stations. Comparing Fig. 3.7 to the map of dynamic topography shown in Fig. 3.3, it can be seen that the Subantarctic Front coincides with the northern high velocity jet which flows northward along the southeast flank of the Campbell Plateau, turns eastward at about 50°S, crosses the mid-ocean ridge system at the Eltanin Fracture Zone and then becomes somewhat diffuse in the southeastern Pacific.
Polar Frontal Zone
The transition from waters with Subantarctic characteristics to waters with
Antarctic characteristics occurs within the Polar Frontal Zone. In their description of zonation south of Australia and New Zealand, Gordon et al. (1977b) aptly referred to this as the Complex Zone. Vertical profiles of temperature within the Polar Frontal Zone exhibit multiple inversions which are accompanied by density-compensating salinity fluctuations. This interleaving of warm, salty waters from the north with cold, fresh waters from the south exhibits vertical scales between 10 m and 200 m. Inversions are most prevalent within the upper 500 m of the water column where lateral gradients of temperature and salinity along density surfaces are greatest. Apart from this interleaving, salinity generally increases monotonically with depth until the salinity maximum of Circumpolar Deep Water is reached. Weak subsurface minima in temperature between 200 m and 500 m depth can often be traced continuously southward to the intense minimum at depths of about 100 m south of the Polar Front.
In the early literature, the Polar Front, or Antarctic Convergence, was considered to be the boundary separating Subantarctic and Antarctic waters (Deacon, 1937; Sverdrup et al., 1942; Mackintosh, 1946). Many different surface and subsurface criteria were used to determine the position of this boundary (cf. Gordon, 1971). As more and better quality data became available, the zonation concept emerged, and it is currently popular to refer to the boundary between the Polar Frontal Zone and the Antarctic Zone as the Polar Front (Emery, 1977; Whitworth, 1980). So defined, the Polar Front coincides with the northern terminus of the well-defined temperature minimum layer of the Antarctic Zone. Here the temperature minimum layer begins its steep descent from depths of approximately 150 m to depths of approximately 500 m. As it deepens, the temperature minimum layer breaks up into multiple weaker minima characteristic of the Polar Frontal Zone. Consistent with the earliest descriptions of the Polar Front, this boundary is often marked by relatively sharp gradients in surface temperature, salinity and silicate (Patterson and Sievers, 1979/80; Sievers and Nowlin, 1984). In Fig. 3.6, the Polar Front is located between stations 407 and 408. The reversal in the slope of the isolines suggests the presence of a current meander between stations 405 and 408, which probably accounts for the two apparently detached temperature minima observed at station 405. The buoy trajectories shown in Fig. 3.5 confirm that meanders are common over the ridge near 135°E.
The depth of the temperature minimum core layer has been plotted by Gordon and Molinelli (1982). Its location at 200 m and 400 m is shown in Fig. 3.7 to bracket the position of the Polar Front. The break in the 400 m isobath between 140°W and 170°W implies that in this region the temperature minimum erodes away before it attains a depth of 400 m. The Polar Front coincides with the southern high velocity jet revealed in Fig. 3.3. It follows the northern flank of the Pacific-Antarctic Ridge northeastward to the Udintsev Fracture Zone where it crosses into the Southeastern Pacific Basin and then, like the Subantarctic Front, weakens until it reintensifies west of Drake Passage.
Fig. 3.8. Distribution of salinity (%c) at 100 m. Adapted from Gordon and Molinelli (1982).
The Antarctic zone is south of the Polar Front and is characterized by relatively smooth vertical profiles of temperature and salinity that exhibit a well-defined subsurface temperature minimum embedded within a halocline above 200 m. The temperature minimum marks the base of the winter mixed layer. Surface heating and melting of ice during summer warms and freshens the surface water and leaves as a residual the characteristic subsurface temperature minimum. Below the minimum, temperature and salinity increase monotonically to the maxima associated with Circumpolar Deep Water.
Within the Antarctic Zone, the isolines in the various property distributions shown in Fig. 3.6 continue to rise toward the south as far as station 409. This station is within the eastward-flowing northern limb of the Ross Sea Gyre (Fig. 3.3). South of this point the isolines become more or less horizontal across the Southeastern Pacific Basin, indicative of weak zonal flow (possibly southward flow) in this location. South of station 413 the deep isopycnals slope downward to the south. This configuration is consistent with westward flow (East Wind Drift) in the southern limb of the Ross Sea Gyre.
In the southern part of the Antarctic Zone, the temperature and salinity maxima of the Circumpolar Deep Water attenuate and there is a transition to the more vertically homogeneous water column of the continental shelf and slope. In Drake Passage, the transition is referred to as the Continental Water Boundary (Sievers and Emergy, 1978) and the waters to the south are in the Continental Zone. In Fig. 3.6, the Continental Water Boundary is between stations 22 and 5. The direction of flow at the Continental Water Boundary apparently varies with location. Eastward flow is reported in Drake Passage (Nowlin et al., 1977), at 60°E (Jacobs and Georgi, 1977), and at 132°E (Callahan, 1971). In the Weddell and Ross Seas, however, this transition occurs within the westward flow of the southern limbs of the subpolar gyres.
Flow within the Continental Zone south of the Continental Water Boundary is poorly resolved. In southern Drake Passage, direct current measurements have shown persistent subsurface flow to the west along the continental shelf and slope (Whitworth et al., 1982). This current has been studied by Nowlin and Zenk (1987) who identified the source waters as coming from the Weddell Sea. They noted that the current cannot be part of the wind-driven circulation because the mean winds throughout Drake Passage are from the west. Rather, the flow is probably a thermohaline feature: a westward extension of the Polar Slope Current of the Weddell Sea. All around the Antarctic continent, shelf waters are colder and denser than the waters offshore. While much of this shelf water is not dense enough to form Antarctic Bottom Water, Killworth (1983) has noted that it may sink to some shallower depth, turning to the west because of the earth's rotation as it sinks. Since dense shelf waters are a circumpolar feature, the implication is that a thermohaline-driven westward flow may exist in most regions within the
Continental Zone. In regions with subpolar gyres (e.g., the Weddell Gyre and the Ross Sea Gyre) such thermohaline currents may supplement the westward wind-driven flow.
As noted earlier, the similarity of zonation observed in the various sectors has led to the idea that the Subantarctic and Polar Fronts and their associated highvelocity jets exhibit circumpolar continuity. Emery (1977), in his analysis of historical expendable bathythermograph (XBT) and hydrographic station data, attempted to assess the extent of continuity across the Pacific Sector of the Southern Ocean. He concluded that the distinctive interleaving and temperature-salinity relationships which characterize the Polar Frontal Zone can be traced continuously across the Pacific, but that the bordering fronts and their associated jets seem to become diffuse and ill-defined in the southeastern Pacific. The same interpretation of flow structure in this region is suggested by dynamic topography (Fig. 3.3) and the zonally varying width of the fronts as revealed by Fig. 3.7. As noted earlier, the failure to detect high-velocity jets in this region may be an artifact of inadequate sampling, but Emery speculated that, in response to the lack of any topographic constraints in the Southeastern Pacific Basin, the jets may spread out into a broader and weaker flow.
In her analysis of the FGGE drifting buoys, Hofmann (1985) noted that in some sectors, especially south of Australia and New Zealand, the buoys drifted along preferred paths which coincide with frontal positions. These paths are evident in Fig. 3.4. The northernmost path lies between 46°S and 50°S and coincides with the Subtropical Front. Farther south the next path, which corresponds to the Subantarctic Front, crosses 140°E between 52°S and 54°S, then turns southeastward and passes around the southern flank of Campbell Plateau (cf. Fig. 3.7). The southernmost path preferred by the buoys, which corresponds to the Polar Front, crosses 140°E between 58°S and 60°S, then crosses the Southeast Indian Ridge at about 59°S, 150°E and continues eastward along the northern flank of the ridge system. Using the frontal positions defined by Clifford (1983) and surface velocities derived from the buoy trajectories, Hofmann concluded that the jets associated with the Subantarctic and Polar Fronts are continuous over large sectors of the Southern Ocean, especially in regions influenced by major bathymetric features. In the southeastern Pacific, however, this continuity could not be demonstrated.
If the jets are continuous, a dynamical explanation is lacking. It has been noted that there is consistency between the mean positions of Southern Ocean fronts and specific features of the wind-induced Ekman flow and the seasonal sea-ice distribution (Taylor et al., 1978; Deacon, 1982; Clifford, 1983). However, it is not clear how these surface forces acting alone can develop and maintain the jets which extend to the bottom. Other investigators (Thompson, 1971; Gordon, 1972a; McCartney, 1976) have related the jets in specific locations to bottom topography, but it is not clear how a zonally variable bottom configuration can produce zonally continuous jets.
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