Circulation and Water Masses of the Southern Ocean: a Review

Home , Antarctic Bottom Water, Ross Gyre, Water mass, Weddell Gyre

Chapter 4 Circulation and Water Masses of the Southern Ocean: A Review

Lionel Carter1,Ã, I. N. McCave2 and Michael J. M. Williams3

1Antarctic Research Centre, Victoria University, P.O. Box 600, Wellington, New Zealand 2Department of Earth Sciences, University of Cambridge, Downing Street, Cambridge CB2 3EQ, UK 3National Institute of Water and Atmospheric Research, P.O. Box 14901, Wellington, New Zealand

ABSTRACT

The Southern Ocean is a major component of Earth’s ocean and climate. Its circulation is complex, with a zonal Antarctic Circumpolar Current (ACC) interacting with a meridional thermohaline circulation. The ACC is a highly variable, deep-reaching eastward flow driven mainly by the westerly winds. It is the longest (24,000 km), largest (transport 137–147.106 m3 s1) and only current to connect the major oceans. The Ekman component of the westerly winds also drives surface waters north. Near the ACC’s northern limit, these waters sink to form Subantarctic Mode and Antarctic Intermediate waters, which continue north at depths oB1,400 m. Interacting with the ACC is the density-forced thermohaline circulation. Super cooling and increased salinity of shelf waters off the Weddell, Wilkes Land and Ross coasts cause these waters to sink and flow equatorwards. The densest component, Antarctic Bottom Water, is captured in deep basins around Antarctica. Less dense water is entrained by the ACC and mixed with deep water moving south from the Atlantic, Indian and Pacific oceans. The resultant Lower Circumpolar Deep Water is tapped off by deep western boundary currents that enter the three oceans at depths WB2,000 m. These northward inflows, with a total volume transport of B55.106 m3 s1, disperse Antarctic and

ÃCorresponding author. Tel.: þ64 4 463 6475; Fax: þ64 4 463 5186; E-mail: [email protected] (L. Carter). 86 L. Carter et al.

northern-sourced waters throughout the world ocean. Other circulation elements are the deep-reaching, cyclonic Weddell, Ross and unnamed gyres located south of the ACC. Further south again are the westward Antarctic Slope and Coastal currents that pass along the Antarctic continental margin under easterly polar winds.

4.1. Introduction

The Southern Ocean has a profound influence on the world’s ocean and climate. Cold, dense water sinks to abyssal depths around the margins of Antarctica and migrates northwards into the Atlantic, Indian and Pacific oceans via deep western boundary currents (Fig. 4.1; Stommel, 1958; Warren, 1981). As succinctly noted by Warren (1971),‘y water from the Antarctic is largely responsible for keeping the rest of the deep sea cold’. Through a process of slow upwelling, these deep cold waters rise to the upper ocean. There, they contribute to the warm surface circulation that extends west from the Pacific and Indian Oceans into the Atlantic where the warm, saline water moves north. Approaching high northern latitudes, the water cools and sinks to form North Atlantic Deep Water (NADW), which migrates south, sandwiched between northbound Antarctic Intermediate Water (AAIW) above and Antarctic Bottom Water (AABW)/Lower Circumpolar Deep Water (LCDW) below (Fig. 4.2). En route, NADW mixes with other waters and eventually rises at the Antarctic continental margin. Thus, one cycle of the global thermohaline circulation (THC) – a major regulator of Earth’s ocean and climate – is completed and another cycle begins (e.g. Broecker, 1991; Schmitz, 1995; Rahmstorf, 2002). This powerful and far-reaching influence of Antarctica and the surround- ing Southern Ocean largely reflects; (i) the strong buoyancy-driven and meteorologically forced circulations, and (ii) their direct access to the major ocean basins via the Antarctic Circumpolar Current (ACC) and its offshoots, the deep western boundary currents (Fig. 4.1; Moore et al., 1999; Orsi et al., 1999; Rintoul et al., 2001). In this brief synopsis we can only provide a flavour of over 70 years of oceanographic research in the Southern Ocean. Thus, we refer the reader to the reference list for a more detailed insight into the workings of this region. We present the basic elements under two sections: (1) Section 4.2 examines the main water masses, focusing on their properties and the mechanisms that control their distribution, and (2) Section 4.3 reviews the structure and dynamics of the world’s largest ocean current, the ACC, together with that of the subpolar gyres and Circulation and Water Masses of Southern Ocean 87

180°

150°E 150°W New Zealand Chatham Rise

Campbell Plateau

S. Tasman Rise ° 120°E 120 W Pacific Antarctic R.

Ross G. E. Pacific Rise

. G d e m a n - ° n 90°W 90 E U

Drake Passage 70˚S 60°S F. Pl. ell G. 50°S Wedd 40°S

° 60°W 60 E SW Indian R. SB

Mid Atlantic R. SAF Legend

Depths <3500 m 30°E 30°W ACC 0° Main DWBC inflow Figure 4.1: The main Oceanographic elements of the Southern Ocean including: (i) the ACC contained by the Subantarctic Front (SAF) and southern limit of UCDW or southern boundary (SB); (ii) the Ross, Weddell and unnamed subpolar gyres; and (iii) the main exit points of deep western boundary currents from the Southern Ocean (blue arrows). The general path for the ACC is from Orsi et al. (1995) with modiﬁcations based on Heath (1985) and Morris et al. (2001). Bathymetric elevations are Annotated as R., ridge; K. Pl., Kerguelen Plateau; and F. Pl., Falkland Plateau. The base chart is Modiﬁed from Orsi and Whitworth (2005). currents residing south of the ACC, and the deep THC. The chapter ends with a discussion in Section 4.4 of the present debate regarding the Southern Ocean’s response to a rapidly warming climate (e.g. Gille, 2002; Curry et al., 2003; Jacobs, 2004; IPCC, 2007). 88 L. Carter et al.

ACC ASF SB SF PF SAF STF 0 SAMW AAIW 1000 UCDW 2000

3000 NADW AABW

DEPTH(m) Antarctic Mid-Ocean 4000 Margin Ridge LCDW

5000

70°S 60°S 50°S 40°S LATITUDE Figure 4.2: Schematic section of the main water masses and their meridional transport as compiled from Whitworth (1988); Orsi et al. (1995); Speer et al. (2000) and Rintoul et al. (2001). Water masses are SAMW, Subantarctic Mode Water; AAIW, Antarctic Intermediate Water; UCDW, Upper Circumpolar Deep Water; LCDW, Lower Circumpolar Deep Water; NADW, North Atlantic Deep Water; AABW, ‘true’ Antarctic Bottom Water (gnW28.27 kg m3). Frontal systems are ASF, Antarctic Slope Front; SB, Southern Boundary of the ACC; SF, Southern Front; PF, Polar Front (formerly the Antarctic Convergence), SAF, Subantarctic Front; STF, Subtropical Front (formerly Subtropical Convergence). The ﬂow of ACC is directed towards the reader.

4.2. Water Mass Formation and Dispersal

4.2.1. Surface Ocean

A series of ocean fronts – narrow, variable bands deﬁned by abrupt changes in water properties, in particular, temperature and salinity – divide the surface waters of the Southern Ocean into several zones (e.g. Gordon, 1975; Deacon, 1982; Whitworth, 1988). Early studies identiﬁed (from south to north) the Polar, Subantarctic and Subtropical fronts (Fig. 4.2). More recent hydrographic transects, especially those carried out during the World Ocean Circulation Experiment (WOCE), have revealed additional boundaries located south of the Polar Front, and termed the ‘southern’ and ‘southern Circulation and Water Masses of Southern Ocean 89

boundary’ fronts (Figs. 4.2 and 4.3; Orsi et al., 1995; Orsi and Whitworth, 2005). Furthermore, these detailed and sometimes repeated transects, along with satellite-borne observations of ocean height, temperature and drifter tracks, reveal the complex and dynamic character of the frontal systems

180°

150°E 150°W

40°S Tasman STF Basin ° SW Pacific 50 S Basin ° 120°E SAF 120 W P14 60°S

in s a ° e B 70 S SF Ris c cific ti . Pa c E r ta n -A n a SE Pacific li a Basin tr s u A PF 90°E 90°W n ele gu er au K ate Pl Enderby l Basin el Crozet dd e n Basin W si Ba

60°E SB 60°W Madg. Basin Argentine Mozb. Basin Basin S2

Legend 30°E Cape Basin 30°W Depths <3500 m

0° Figure 4.3: Location of the principal ocean frontal systems in the Southern Ocean (based on Orsi et al., 1995, but modified for the New Zealand region according to Carter et al., 1998 and Morris et al., 2001). Repeated hydrographic transects, satellite observations and drifting floats reveal the frontal systems as dynamic features with marked temporal and spatial variability but generally within the constraints imposed by the ocean floor topography (Moore et al., 1999). P14 and S2 are locations of WOCE hydrographic transects portrayed in Figs. 4.4 and 4.5. Madg., Madagascar Basin; Mozb., Mozambique Basin. Names of fronts are given in Fig. 4.2. The base chart is modified from Orsi and Whitworth (2005). 90 L. Carter et al.

(Hofmann, 1985; Davis, 1998; Moore et al., 1999; Rintoul et al., 2001; Kostianoy et al., 2004; Sokolov and Rintoul, 2007). While cognizant of these complexities, the main fronts can still be used to define the distribution of three major surface waters characterised mainly by their potential temperature (y), salinity (S) and oxygen content (see hydrographic charts in Orsi and Whitworth, 2005). (1) Near-freezing and relatively fresh Antarctic Surface Water (AASW) forms a layer about 100 m thick that extends from the Antarctic continental shelf to the Polar Front, commonly defined as the northernmost extent of the subsurface temperature minimum (Belkin and Gordon, 1996; Figs. 4.4 and 4.5). AASW temperatures are typically o01C, but may rise to 2.51C near the Front or where warm, deep water approaches the surface (Gordon, 1975; Deacon, 1982). Salinity (S) varies regionally with highest values of SW34.3 psu found in the Ross and Weddell seas, whereas AASW elsewhere around Antarctica commonly has So34.0 psu. (2) Between the Polar and Subantarctic fronts resides surface water that is transitional between AASW and Subantarctic Surface Water (SASW). The structure is complex in response to mixing and interleaving of AASW as it sinks near the Polar Front (e.g. Gordon, 1975; Rintoul et al., 2001). Thus, properties are variable, but generally S is B34.0–34.4 psu and y is 3–81C(Fig. 4.5). (3) SASW occurs north of the Subantarctic Front and encompasses water that usually warms northwards from B61Cto121C. Salinity is typically W34.3 psu except in the SE Pacific and Drake Passage where values decline to o34.16 psu. Like its more southern counterpart, SASW may be affected by vertical mixing as surface waters subduct and mix (e.g. Morris et al., 2001). The northern limit of subantarctic waters is the Subtropical Front where temperatures increase sharply by 4–51C and salinity by 0.5 psu (Fig. 4.3; Deacon, 1982). Subtropical surface water prevails north of the Subtropical Front.

4.2.2. Subantarctic Mode Water and Antarctic Intermediate Water

Isopycnals – surfaces of constant density – of near-surface to deep waters in the Southern Ocean rise up in a step-like profile towards Antarctica. Close to ocean fronts, isopycnals may outcrop indicating either the rise of deep waters to the ocean surface or the descent of surface waters (Figs. 4.2, 4.4 and 4.5). In the case of the latter, winter cooling and mixing of the surface waters on the northern side of the Subantarctic Front forms Subantarctic Mode Water (SAMW) (McCartney, 1977; Morris et al., 2001; Rintoul et al., 2001). This well mixed, ventilated water descends northwards to B500 m depth along much of the front (Fig. 4.2). AAIW also descends from the surface, passing Figure 4.4: Sections of potential temperature (A), salinity (B) and neutral density (C) from WOCE Line P14 across a major constriction in the ACC between the Ross Sea and New Zealand (see Fig. 4.3 for location). Isolines at the Antarctic margin indicate descent of dense shelf waters, which may be mixed with NADW-influenced, LCDW as suggested by the salinity field. The resultant AABW (cf. Fig. 4.5) is contained within the SW Pacific Basin. At the surface, north of the polar front, low salinity AAIW descends northwards (Fig. 4.4B). Hydrographic profiles were derived from the WOCE Southern Ocean Atlas at http://woceatlas.tamu.edu/. Names of fronts are given in Fig. 4.2. 92 L. Carter et al.

[C] ° ° ° ° ) m ( AABW Depth LC DW

NEUTRAL DENSITY

Distance (km) Figure 4.4: (Continued). under SAMW to reach a maximum depth of B1,400 m (Figs. 4.2 and 4.4). AAIW is identified by a salinity minimum (34.3–34.5 psu) and temperatures of B3–71C. However, the processes driving AAIW formation are unclear. Formation may be related to wind-forced or density-driven sinking of cold AASW and indeed there appears to be continuity between the AASW and AAIW salinity fields (Figs. 4.4 and 4.5). However, McCartney (1977) suggested AAIW may evolve at least in part from dense SAMW. Whatever the origin, compared to the widespread formation of SAMW, new AAIW presently appears to form mainly in the SW Atlantic and SE Pacific. From these sites AAIW circulates the oceans in anticyclonic subtropical gyres that extend north towards and locally beyond the equator before returning south as ‘old’ AAIW, which is transported within western boundary currents (Rintoul et al., 2001; Ridgway and Dunn, 2007).

4.2.3. Circumpolar Deep Water

The most voluminous water mass in the Southern Ocean is Circumpolar Deep Water (CDW). It extends from B1,400 m to W3,500 m depth, but it F

[A] T

S 60.0° S 50.0° S 40.0° S – – – – – – – – — – – – – – – – – – — – – – – – – – – – — – – – – – 0 — – 34.5 34.4 34.2 34.7 AASW A 34.3 AIW 34.68 34.4 34.5 34.3 1000 34.4 34.6 34.66 34.71 34.74 34.76 2000 34.78 N 34.66 AD W-LC (m) DW 3000 AABW 34.72 34.8 34.71

Depth 34.73 34.68 34.74

34.78 n i

34.76 g

4000 r

LCDW a

34.73 n

i r 5000 f Cape Basin A

Weddell Basin 34.68 SALINITY 6000 0 1000 2000 3000 4000 ° ° ° [B] 60.0 S 50.0 S 40.0 S

0 – – – – – – – – — – – – – – – – – – — – – – – – – – – – — – – – – – 28

28.12 27

27.2 27.4 28.2 1000 28.2 27.6 27.8

28.26 28 2000

28.32 AABW (m) 3000 28.36 28.12 Depth

28.2 4000

5000 28.4

28.4 28.26 NEUTRAL DENSITY 6000 0 1000 2000 3000 4000 Distance (km) Figure 4.5: Hydrographic sections of the salinity (A) and neutral density (B) fields from WOCE Line S2 in the Atlantic sector of the Southern Ocean. Of note is (i) the step-wise ascent of saline, NADW-rich LCDW towards Antarctica where it is capped by fresh, cold AASW in the south and subducting AAIW in the north and (ii) the containment of a large volume of classic Antarctic Bottom Water (AABW; gnW28.27 kg m3) within the Weddell Basin (cf. Fig. 4.4). Hydrographic profiles are derived from the WOCE Southern Ocean Atlas at http://woceatlas.tamu.edu/. Names of fronts are given in Fig. 4.2. 94 L. Carter et al. rises to meet AASW or even outcrop along the Antarctic continental margin (Figs. 4.2, 4.4 and 4.5). CDW has two basic types: (1) Upper Circumpolar Deep Water (UCDW) is identified by the oxygen minimum and elevated nutrient concentrations, and has an open ocean depth range of B1,400– 2,500 m, and (2) Lower Circumpolar Deep Water (LCDW) whose signature is the salinity maximum (34.70–34.75 psu) (Gordon, 1975; Orsi et al., 1995). This maximum reflects the input of NADW that has migrated into the Southern Ocean (Orsi et al., 1995; Rintoul et al., 2001). Upon meeting the ACC, NADW is entrained and transported east around the Antarctic continent, all the while mixing with waters from the Indian and Pacific oceans plus dense waters from Antarctica to form LCDW. Despite the vigorous mixing, the high-salinity signature of NADW is retained (Reid and Lynn, 1971; Gordon, 1975). At several locations around Antarctica, LCDW rises at the continental slope where mixing with super-cold shelf water renews not only the deep circulation of LCDW but also generates Antarctic Bottom Water (AABW), the deepest water mass in the Southern Ocean (Figs. 4.2, 4.4 and 4.5; Foster and Carmack, 1976; Jacobs et al., 1985; Orsi et al., 1999). LCDW is carried equatorwards in all three major oceans by deep western boundary currents (see Section 4.3.3 and Schmitz, 1995; Hogg, 2001). According to Rintoul et al. (2001) mixing with fresher waters, together with the biological depletion of oxygen, slowly modify LCDW into a less dense, lower oxygenated variant that returns south as UCDW. In some WOCE sections, such as P15 in the Indian Ocean, oxygen depletion of UCDW may be influenced by the direct injection of nutrient-enriched deep waters from the North Indian and North Pacific oceans.

4.2.4. Antarctic Bottom Water

In their analysis of abyssal water masses, Mantyla and Reid (1983) drew attention to the often inappropriate use of the term Antarctic Bottom Water, which has tended to be used generically for any southern-sourced bottom water. They demonstrated that true AABW did not extend far from Antarctica before mixing with other waters (see also Orsi et al., 1999). This is particularly true for the deep western boundary currents in which AABW is mixed with CDW derived from the ACC. Thus, at 301N in the NW Atlantic, Amos et al. (1971) recorded o20% AABW near the seabed. To better characterise AABW and thereby improve assessments of its dispersal and contribution to bottom waters worldwide, Orsi et al. (1999) defined AABW by its neutral density (gn) whereby gnW28.27 kg m3. Such dense waters are confined mainly to the deep (down to B6,000 m), Circulation and Water Masses of Southern Ocean 95 circum-Antarctic Basins that include the Argentine and Brazil basins (SW Atlantic), Mozambique, Crozet and Australian-Antarctic basins (Indian) and the SE Pacific Basin (Figs. 4.1, 4.4–4.6). Less dense, southern bottom water with 28.18ogno28.27 kg m3, is not confined to the circum-Antarctic basins, but instead spreads out from the deep levels of the northern edge of the ACC into all major oceans (Orsi et al., 1999; also see Section 4.3.3). AABW density is determined by a combination of different sources that produce regionally distinct waters (Mantyla and Reid, 1983; Orsi et al., 1999; Jacobs, 2004). The freshest and coldest (Sr34.64 psu; yr11C) bottom water occurs in the Weddell Sea, whereas the SE Pacific Basin has the most saline and least cold (SZ34.72 psu; 0.6oyo0.31C) bottom water. The Australian-Antarctic Basin contains water with properties intermediate between the two end members. Traditionally, the Weddell Sea was regarded as the prime source of AABW, but recently two other sources have come to the fore. The Weddell Sea’s contribution is now regarded as B50%, with the Wilkes Land margin including the Ade´ lie coast (Rintoul, 1998) contributing B30% mainly to the Indian Ocean, and the Ross Sea producing B20% AABW that is destined primarily for the SE Pacific Basin (Jacobs, 2004). Equally varied are the modes of bottom water formation (Jacobs, 2004). Foster and Carmack (1976) invoked the formation of highly saline shelf water (HSSW) by brine rejection from sea ice. However, salt-driven increases in density may also be influenced by intrusions of NADW-bearing LCDW onto the upper slope and shelf (Toggweiler and Samuels, 1995). Super cold, Ice Shelf Water (ISW), formed by freezing and melting below ice shelves, can mix with HSSW and reach the outer shelf before flowing down slope (Baines and Condie, 1998). Alternatively, the simple mixing of cold AASW and high- salinity LCDW, with or without ISW, may produce negatively buoyant waters at the shelf edge (e.g. Jacobs, 2004). Finally, dense waters may sink via convection chimneys and polynyas such as the well-documented but short-lived Weddell Sea polynya (Gordon, 1982). Rates of AABW formation, as estimated from hydrographic data, usually fall within a range of 5–15 Sv (1 Sv ¼ 106 m3 s1). Anthropogenic tracers, in particular chlorofluorocarbons, record a flux of 8.1 Sv for AABW descending at the 2,500 m isobath off Antarctica (Orsi et al., 2001). This compares to 7.6 Sv of lower NADW flowing out of the Nordic and Labrador seas at the N Atlantic source. Because AABW is colder than its northern counterpart, Antarctic overturning probably plays the dominant role in cooling the deep ocean. The time at which deep water circulates through the ocean is often quoted to be a millennium or more, for example 1,000 years between the N Atlantic and Southern Oceans and a further 1,000 years from the Southern 96 L. Carter et al.

Figure 4.6: Extent of dense Antarctic Bottom Water as identified by the neutral density field at 3500 m (Orsi et al., 1999; Orsi and Whitworth, 2005). With the exception of the SE Atlantic, where AABW extends well north via the deep Argentine and Brazil basins, the remaining AABW is captured in circum-Antarctic basins where further northwards dispersal is inhibited by oceanic ridge systems shown in white. Basins annotated as Arg. B., Argentine Basin, which extends north into the Brazil basin (not shown on chart); W.B., Weddell Basin; E.B., Enderby Basin; C.B., Crozet Basin; Aust. A.B., Australian-Antarctic Basin, and SE P.B., SE Pacific Basin. Chart is generated from the WOCE Southern Ocean Atlas at http://woceatlas.tamu.edu/. Circulation and Water Masses of Southern Ocean 97

Ocean to the North Pacific. However, in a re-analysis of radiocarbon dated ocean waters, Matsumoto (2007) indicates much shorter circulation ages thereby supporting but refining earlier radiocarbon-based studies (Stuiver et al., 1983). Thus, the circulation age for the Southern Ocean below 1,500 m is B300 14C years with a similar age for the Atlantic. For the Pacific, the basin circulation age is B900 14C years.

4.3. Ocean Circulation

4.3.1. Antarctic Circumpolar Current

The Southern Ocean circulation is dominated by the ACC, a current system that is rightly described by superlatives. It is the only current to connect the major ocean basins and hence plays a prominent role in the global distribution of heat, salt and gases (Fig. 4.1). It is the longest current with an estimated pathway of B24,000 km (Whitworth, 1988). Finally, the ACC is the largest major current in terms of volume transport with a mean of 136.777.8 Sv as measured across Drake Passage (Cunningham et al., 2003). The ACC was originally termed the West Wind Drift (Deacon, 1937)in recognition of its forcing by middle-to-high latitude westerly winds (Orsi et al., 1995; Whitworth et al., 1998). However, use of the term wind drift masks the role played by the buoyancy-driven component of the circulation (see Rintoul et al., 2001). Complexities aside, the net result is an eastward current system that extends from the ocean surface to bottom, its path guided by submarine topography (Fig. 4.1; Gordon et al., 1978; Orsi et al., 1995). For much of its passage, the ACC flows along the flanks of mid- oceanic ridges except within major gaps in the Pacific and Indian ridge systems where the current shifts poleward (Fig. 4.1). Large submarine plateaux also exert an influence. The ACC widens to the north and south as it passes around the Kerguelen Plateau, whereas the Campbell and Falkland plateaux form constrictions (Fig. 4.1; Whitworth and Peterson, 1985; Morris et al., 2001). Interestingly, this interaction with the ocean floor was inferred as early as the 1950s. Estimates for a purely wind-driven ACC yielded transports that were excessive compared to observations. Thus, it was concluded that the wind stress was partly balanced by bottom stress (see Whitworth, 1988; Rintoul et al., 2001). The passage of westerly winds over the ACC also induces an Ekman drift to the north – a process that probably plays a role in the subduction and transport of mode and intermediate waters. This equatorward flow is compensated at depth by the southward 98 L. Carter et al. transport and eventual upwelling of CDW thus contributing to the THC (Wyrtki, 1961; Toggweiler and Samuels, 1995). Rather than a uniform flow, the ACC is a system of deep-reaching zonal jets that separates zones of relatively quiet water. The jets are marked by the circumpolar fronts (see Section 4.2 Surface Ocean) with the northern and southern boundaries of the ACC defined, respectively, by the Subantarctic and Southern Boundary fronts (Figs. 4.1–4.3). Both eddy-resolving models and satellite observations highlight the complex flow of the frontal jets (Morrow et al., 1992; Gille, 1994; Carter and Wilkin, 1999). Meanders, eddies and intricate branches are well shown especially where the flow is constricted as off Campbell Plateau and within Drake Passage (Nowlin and Klinck, 1986; Morris et al., 2001; Cunningham et al., 2003). Most ACC transport takes place within the fronts, but their complex flow patterns and different criteria for estimating transport have led to a wide range of values for the entire ACC. Nevertheless, closely spaced and long- term monitoring sites have improved estimates of transport. In Drake Passage the mean transport is 136.777.8 Sv (Cunningham et al., 2003) compared to 147710 Sv between Australia and Antarctica (Rintoul and Sokolov, 2001). Much of the transport in the Australasian reach of the ACC occurs within the Subantarctic Front, which has a mean of 10577 Sv off Australia (Rintoul and Sokolov, 2001) and B90 Sv off southern New Zealand (Morris et al., 2001). However, ACC transport can be highly variable. Time series from Drake Passage record variations at several time scales (Whitworth and Peterson, 1985). Short-term fluctuations, related to 14-day solar and lunar tides, are superimposed on longer-period fluctuations of B1 year and longer that can lead to changes in transport of B30–40 Sv within weeks. The interaction of the ACC with the topography and southerly extensions of western boundary currents generates eddies that play important roles in the transfer of heat and momentum (Bryden and Heath, 1985; Morrow et al., 1992; Rintoul et al., 2001). Off SE New Zealand, for example interception of the ACC by the South Tasman Rise and Macquarie Ridge spawns bottom- reaching eddies that migrate NE along the steep margin of Campbell Plateau (Boyer and Guala, 1972; Gordon, 1972). Both cyclonic and anticyclonic features have been observed from current meter and satellite data, which suggest a frequency of occurrence of B9 eddies annually (Stanton and Morris, 2004). Modelled eddy kinetic energy, verified by current meter data and ocean floor sedimentary evidence, attest to the power of these perturbations, which are likely to be the cause of abyssal benthic storms (Hollister and McCave, 1984; Carter and Wilkin, 1999). Seabed topography also encourages intense mixing within the ACC. The Scotia Sea and Circulation and Water Masses of Southern Ocean 99 potentially other areas of marked seabed relief below the ACC, are zones of the highest turbulent mixing in the ocean and result in rapid upwelling that may locally short-circuit the classic meridional overturning as portrayed in Fig. 4.2 (Garabato et al., 2004, 2007).

4.3.2. Subpolar Circulation

4.3.2.1. Major gyres

A key element of the circulation, south of the ACC, is three large, deep- reaching cyclonic gyres that extend from the ACC to the Antarctic continental margin (Fig. 4.1). The better known are the Weddell and Ross gyres that occupy ridge-bounded sectors of the Weddell and Ross seas (e.g. Orsi et al., 1993; Jacobs et al., 2002). A third and as yet unnamed gyre was suggested to occur south of Kerguelen Plateau (Bindoff et al., 2000), and has now been confirmed by McCartney and Donohue (2007). As documented for the Weddell Gyre (Orsi et al., 1993), it is likely that all three gyres transport salt and heat from the ACC to the Antarctic continental margin where deep and bottom waters are produced (e.g. Jacobs, 2004). Furthermore, McCartney and Donohue (2007) suggest a strong connectivity between the three gyres with a westward flow along their southern limbs and an eastward flow joining their northern limbs. This latter flow is just south of another eastward flow, this time associated with an anticyclonic supergyre covering most of the S Pacific and S Indian oceans (Ridgway and Dunn, 2007). This eastward limb of the supergyre appears to reside between the Subtropical and Subantarctic fronts. As a result, it may entrain SAMW, formed in the vicinity of the Subantarctic Front and help distribute it through the ocean basins as suggested by Rintoul et al. (2001). Of these cyclonic circulations, the Weddell Gyre is the largest, extending from B501W to between 201–301E(Fig. 4.1; Orsi et al., 1993). At the surface it has the form of a NE–SW aligned, elongated gyre whereas at depth it comprises two cyclonic cells located to the east and west of 151W. Basically, the Weddell Gyre occupies the W4,000 m deep re-entrant formed by Antarctica and the ridge systems that extend eastward from near the tip of the Antarctic Peninsula (Fig. 4.1). While such a location implies containment, the northern limb of the gyre overlaps the ridges to interact with the ACC. CDW entrained from the ACC is moved within the gyre and can eventually mix with cold shelf waters to form Weddell Sea Bottom Water, the local variety of AABW and the densest water in the Southern Ocean (Foster and Carmack, 1976). 100 L. Carter et al.

The Ross Gyre extends from 1601E to 1401W, and is largely confined to the W4,000 m deep western reach of the SE Pacific Basin (Fig. 4.1). Like its Weddell Sea counterpart, the Ross Gyre is a deep-reaching feature that entrains CDW to make it available for mixing with shelf and slope waters. Situated to the south and southwest of Australia, the unnamed gyre appears to favour the southern part of the Australian-Antarctic Basin, most of which is W4,000 m deep (Fig. 4.1). Indeed, the volume transport field is compressed against the Antarctic continental margin and the associated westward slope current (Bindoff et al., 2000; McCartney and Donohue, 2007). Transport along the eastward-flowing northern limb of the gyre is estimated at 35 Sv. However, the amount of transport along the west- moving, southern limb is unclear due to merging with the slope current, the combined flows reaching 76 Sv (McCartney and Donohue, 2007).

4.3.2.2. Antarctic slope and coastal currents

As summarised by Heywood et al. (2004) much of the Antarctic margin is bathed by two westward currents. One is associated with the Antarctic Slope Front that constitutes the boundary between fresh, cold Antarctic shelf waters and less cold, saline CDW (Jacobs, 1991; Whitworth et al., 1998). Deacon (1937) regarded the frontal flow as a consequence of the prevailing easterly winds and coined the name, East Wind Drift. On the basis of classical theory, he reasoned that polar easterly winds produced an onshore Ekman transport with the resultant generation of a westward geostrophic current below the wind-mixed layer (Whitworth et al., 1998; Bindoff et al., 2000). The second significant feature is the Antarctic Coastal Current, which forms a narrow, rapid flow across broad sections of the continental shelf, for example, in the SW Weddell Sea. However, where the shelf is narrow the coastal current is difficult to distinguish from flows associated with the Antarctic Slope Front and the southern limbs of the subpolar gyres where they approach the continental margin.

4.3.3. Thermohaline Circulation

Wunsch (2002) drew attention to the imprecise meaning of the term, THC. Following Jacobs (2004), we prefer the broad definition of Schmitz (1995) whereby the THC is the ‘y buoyancy-driven flow field associated with water cooled (or heated) by contact with cold (or warm) air, or modified by sources and sinks of cold water. May also include flows whose characteristics are Circulation and Water Masses of Southern Ocean 101 significantly altered by upwelling and/or mixing. Water sinking at high latitudes tends to return equatorwards in relatively strong, narrow currents called DWBC’. The high latitude sinks noted are primarily the North Atlantic and Antarctica (Stommel, 1958; Warren, 1981). As described earlier, dense water formed mainly over shelf areas of the Weddell and Ross seas, and along the Wilkes Land coast, sinks and descends down the continental slope. Descent may take place in several ways according to Baines and Condie (1998). For prominent deep-water formation areas, dense water may descend as broad sheets or plumes. For weaker sources, the resultant outflows may geostrophically descend the slope to a level where the layer thins and viscous drainage prevails. Descending flows may be disrupted by eddies into discrete parcels that overall move down slope. Further complexities are introduced by the slope topography, for example, submarine canyons and channels can contain and steer flows whose density may increase through entrainment of sediment. Whatever the mode of descent, the initial dispersal of AABW is westward as outlined by the distribution of chlorofluorocarbons (Orsi et al., 1999). Those tracers show that the densest AABW usually follows a cyclonic path presumably under the influence of the subpolar gyres and basin topography (Fig. 4.1). However, the Weddell– Enderby Basin experiences limited outflow at its northern rim where least dense AABW (gnB28.28 kg m3) passes north into adjoining basins of the Atlantic and Indian oceans (Fig. 4.6; Orsi et al., 1999). In contrast, the main inflows to the adjoining oceans are via deep western boundary currents, which carry mainly LCDW from the northern boundaries of the ACC (Mantyla and Reid, 1983; Schmitz, 1995).

4.3.3.1. Atlantic DWBC inﬂow

Weddell Sea Bottom Water, a local variety of AABW, together with Weddell Sea Deep Water and CDW from the Drake Passage, are carried by a northbound DWBC into the Atlantic Ocean below southward-moving NADW (Warren, 1981; Schmitz, 1995). The DWBC is typically found in water depths exceeding 3,500–4,000 m along the western boundary presented by the continental margin off South America. However, the current pathway is interrupted by a succession of deep basins including (from south to north) Georgia, Argentine and Brazil basins. As noted earlier, densest AABW is captured within the deep basins leaving less dense waters to move north via gaps and channels through the inter-basin ridges. Even so, not all the deep water escapes and some recirculates within the basins themselves (Hogg and Johns, 1995; Hogg, 2001). For example, around 6.9 Sv ﬂows from the 102 L. Carter et al.

Argentine Basin into the Brazil Basin through Vema and Hunter channels, but only 3.2 Sv leaves the Brazil Basin, leaving 3.7 Sv to recirculate.

4.3.3.2. Indian DWBC inﬂow

The inﬂow of AABW into the western Indian Ocean is via Crozet and Mozambique basins (the latter being a dead end), which largely constrain AABW to south of 301S and 341S, respectively (Orsi et al., 1999). LCDW from the Crozet Basin leaks northwards through fractures zones in the SW Indian Ridge (Warren, 1974; Johnson et al., 1991; Mantyla and Reid, 1995). Deep transport is via a northbound DWBC, which like its Atlantic counterpart, has a complex pathway dictated by multiple basins and ridges (Warren, 1981; McCave et al., 2005). Basins encourage recirculation of deep waters and together with mixing, account for a northward dissipation of the DWBC. For instance, LCDW below B3,800 m has a northwards transport of 3.8 Sv through Madagascar Basin, but only 1.7 Sv exits north into the Somali Basin (Johnson et al., 1998). Crozet Basin, in the western Indian Ocean, is not the only gateway for DWBCs into the Indian Ocean. The eastern Indian Ocean is also connected to the Southern Ocean thereby allowing the northward intrusion of two other boundary currents; one along the eastern side of SE Indian Ridge and the other along the eastern ﬂank of Ninetyeast Ridge (Warren, 1981; Toole and Warren 1993; Reid, 2005). Again, the deep water carried north is LCDW, with the denser AABW retained in Australian-Antarctic Basin (Orsi et al., 1999). When extended to the 2,000 dbar reference level, the combined northward transport of the three DWBCs into the Indian Ocean is B27 Sv (Toole and Warren, 1993).

4.3.3.3. Paciﬁc DWBC inﬂow

The general pathway of the DWBC into the Pacific was first outlined in the classic model of the global abyssal circulation by Stommel (1958) and Stommel and Arons (1960), and was later confirmed by the hydrographic sections of Warren (1971, 1973). Because the Tasman Basin is essentially closed at its northern end, the main inflow is off southernmost New Zealand where the DWBC enters in concert with the ACC (Fig. 4.7; Carter and McCave, 1997; Carter and Wilkin, 1999). Initially, the combined inflow intercepts Macquarie Ridge to form meanders and eddies although some current filaments pass through narrow gaps in the ridge (Boyer and Guala, Circulation and Water Masses of Southern Ocean 103

SINKING Thermohaline Circulation

Cold deep current Cool shallow current

Warm shallow current

WELLIN UP G U U P PW W E E L L L L I N I N

G G

A.C .C. .C. A.C SINKING SINKING SINKING Figure 4.7: A schematic portrayal of the THC, which is a series of loosely linked, recirculation systems that transport, heat, salt, nutrients and ventilating gases through the world ocean. Only the main elements of the THC are shown, for example the Indian Ocean has three, deep northward inﬂows that include from W to E, the margin off eastern Africa (shown), and the eastern sides of the SE Indian Ridge and Ninetyeast Ridge (not shown) (image modiﬁed from Manighetti, 2001).

1972; Gordon, 1972). This perturbed combined flow continues northeast along the 3,000–3,500 m high flanks of Campbell Plateau to around 491S where the ACC veers east leaving the DWBC to continue northwards into the Pacific Ocean (Fig. 4.7). It eventually departs the Southern Ocean off Chatham Rise between 441S and 421S(McCave and Carter, 1997). There, Warren (1973) observed a volume transport of B20 Sv, the largest for a single DWBC (Schmitz, 1995).

4.4. Oceanographic Variability and Change

The present phase of climate change has drawn considerable attention to the behaviour of Antarctica and the Southern Ocean (e.g. Gille, 2002; Jacobs et al., 2002; Walther et al., 2002; Cook et al., 2005). However, confident identification of any effect, especially in the oceans, has been 104 L. Carter et al. hindered by; (i) sampling bias, (ii) the short history of observations, which typically only encompass the last 50–60 years, (iii) a multiplicity of forcing mechanisms, some with and some without clear connections to a warming climate and (iv) marked variability at a range of temporal and spatial scales (e.g. Jacobs and Giulivi, 1998; Orsi et al., 2001; Vaughan et al., 2001). Of special note is the Southern Annular Mode (SAM), which appears to dominate inter-annual to centennial variability in the Southern Ocean (Hall and Visbeck, 2002; Lovenduski and Gruber, 2005; Lovenduski et al., 2007). SAM is essentially a zone of climate variability that encircles the South Pole and strongly influences zonal winds, sea ice formation and oceanic circulation. When in a positive phase, SAM is typically associated with enhanced westerly winds over the ACC and weakened winds further north. This favours a strengthening of the ACC, a reinforcement of the northward Ekman drift and subduction of surface waters, and a northward expansion of sea ice. To compensate, there is enhanced poleward transport and rising of deep water at the Antarctic continental margin. Under a negative SAM, windiness and storminess appear to migrate to mid-latitudes thus weakening both zonal and meridional ocean transport. Despite uncertainties associated with the aforementioned limitations, it is nonetheless important to examine and evaluate recent changes in Antarctica and the Southern Ocean in light of their actual or potential influence on the global ocean and climate. Climatic trends for the past 50 years have been identified by Turner et al. (2005) using records from 19 Antarctic meteorological stations. The results emphasise the marked geographic variability of the continental climate (e.g. Vaughan et al., 2001) as well as its temporal variability at interdecadal scales. The Antarctic Peninsula has warmed at a statistically significant rate of þ0.561C/decade from 1951 to 2000. The next largest warming trend outside of the Antarctic Peninsula is in the western Ross Sea (Scott Base) with a rise of þ0.291C/decade, although the rise is not statistically significant. Elsewhere, significant trends are unclear. Annual temperature trends for coastal and interior sites on the East Antarctic Ice Sheet suggest a slight cooling. However, all but one site exhibited warmer winters. Notwithstanding its temporal and spatial variability, the upper Southern Ocean has warmed between 1955 and 2003 in concert with the world ocean (Levitus et al., 2005). Off the Antarctic Peninsula, the pronounced atmospheric warming has been accompanied by an equally marked warming of the surface ocean with summer temperatures increasing by 1.21C over the latter half of the twentieth century (Meredith and King, 2005). At water depths of 700–1,100 m, Gille (2002) reported an average 0.171C rise since 1950, which accounts for about two-thirds of the total increase in heat Circulation and Water Masses of Southern Ocean 105 content in the ocean from 0 to 3,000 m depth (IPCC, 2007). Much of the warming is concentrated within the Subantarctic Front of the ACC. This raises the possibility that the warming at depth may result from the sinking of atmospherically warmed SAMW. Gille (2002) further suggests that warming may also be related to a general southward displacement of the ACC. This would be consistent with a suite of model simulations that suggest warming of the Southern Ocean will be accompanied by a southward shift of zonal westerly winds together with a narrowing and intensification of the ACC (IPCC, 2007 and references therein). Coherent historical trends are also evident for salinity. Subpolar regions have generally become fresher between 1955 and 1998 in contrast to subtropical and tropical regions, which display increased salinity with the exception of the central Pacific Ocean (Curry et al., 2003; Boyer et al., 2005). Such freshening of the upper Southern Ocean may be responsible for a reduction in the salinity of AAIW (Wong et al., 1999; Curry et al., 2003). Reduced salinity implies increased freshwater input that may result from one or more of the following causes: (i) greater net precipitation; (ii) changes in the extent of sea ice – a major contributor to winter salinity through brine rejection; (iii) increased melting of ice shelves, ice sheets and glaciers (e.g. Cook et al., 2005), and (iv) changes in the oceanography, especially any reduced upwelling of saline LCDW (Wong et al., 1999; Jacobs et al., 2002; Curry et al., 2003). Nearer Antarctica, hydrographic records spanning over 40 years show local variability in salinity trends. The upper 50 m of the ocean, west of the Antarctic Peninsula, has become more saline although underlying waters have freshened slightly in line with the regional trend. While the more saline surface conditions appear to be out-of-step with the strong glacier retreat on the Peninsula (e.g. Cook et al., 2005), Meredith and King (2005) suggest more saline conditions are consistent with the reduced sea ice cover and the timing of the hydrographic measurements. With less sea ice production there is less freshening of the ocean in the summer when most of the hydrographic measurements are made. On the opposite side of the continent, at Law Dome, ice core records spanning a century and longer, identify a 20% loss of sea ice since 1950 although this trend is strongly overprinted with cyclical variations with an B11 year frequency (Curran et al., 2003). Reduced sea ice along with increased precipitation and melt water from the West Antarctic Ice Sheet have been cited as contributing to the observed freshening of surface waters associated with the Ross Sea Gyre (Jacobs et al., 2002). But like Law Dome, the trend is obscured by cycles, this time by 5–6- and 9-year oscillations in HSSW formation (Assmann and Timmerman, 2005). 106 L. Carter et al.

Salinity and temperature changes have the potential to affect bottom water production and the THC. Consequently, these changes have received considerable attention (IPCC, 2007 and references therein). At glacial- interglacial time scales, palaeoceanographic evidence reveals marked variations in the position and degree of convective overturning of the N Atlantic sector of the THC (Rahmstorf, 2002 and references therein). During interglacial periods, overturning is most active and reaches its most northerly extent. In contrast, glacial periods are likely to witness a southward shift and possible slow-down in overturning. Any slow-down may be compensated by a greater production of bottom water from Antarctica as suggested by grain size (Hall et al., 2001), magnetic properties (Venuti et al., 2007) and diatom proxies (Stickley et al., 2001). However, such conclusions are sometimes at odds with geochemical tracers such as d13C (e.g. Moy et al., 2006) that point to little change in the passage of NADW through the Indian and Pacific sectors at least over recent glacial-interglacial cycles. At millennial time scales, abrupt changes such as those associated with Heinrich events, may stop N Atlantic overturning altogether as the density of the surface ocean is reduced by rapid influxes of melt water (Rahmstorf, 2002). Again, cessation of N Atlantic production may be compensated by enhanced Antarctic production. However, responses to the latest phase of climate/ocean warming are unclear. In the N Atlantic, which is the best observed deep- water source, long-term trends are equivocal due to decadal variability, a paucity of long-term observations and other factors (IPCC, 2007). A similar situation applies to Antarctica where estimations of bottom-water production are inconsistent in response to: (i) natural cycles; (ii) differences in the definitions and techniques to estimate production rates, and (iii) a bias towards summer observations (Jacobs, 2004). On the basis of chlorofluorocarbons and 14C data, which allow water masses to be traced at decadal to century scales respectively, Orsi et al. (2001) revealed no decline in bottom water production over the twentieth century as indicated earlier by Broecker et al. (1999). Nevertheless, the changes recorded in recent historical times cannot be ignored. The freshening of the Ross Gyre over the last 50 years (Jacobs et al., 2002) and an accompanying downstream freshening of AABW in the adjacent Australia-Antarctic Basin (Aoki et al., 2005) are consistent with increased freshwater input. On a larger scale, the historical salinity data of Curry et al. (2003) also reveals a freshening of deep and bottom water at Antarctic and N Atlantic sources. Simulations by 19 model runs under IPCC greenhouse gas scenario, A1B (rapid economic growth, world population peaks mid-century, new and efficient energy technologies with reliance on a range of sources) point to an average 25% reduction in N Atlantic overturning by the year 2100 (IPCC, 2007). Circulation and Water Masses of Southern Ocean 107

None of the runs point to a shut-down; rather they favour reductions in overturning of up to 50%. While the Southern Ocean sector has received less attention from modellers, simulations based on a warmer or fresher ocean may enhance or stabilise N Atlantic overturning (Saenko et al., 2003; Weaver et al., 2003). To further emphasise the complexity of north–south relationships, the projected strengthening of the Southern Hemisphere westerly winds will increase the northward Ekman transport of upper ocean. To compensate, the poleward flow of deep water below 2,000–2,500 m depth, could be expected to strengthen and possibly stimulate the southward flow of NADW (e.g. Toggweiler and Samuels, 1995; Toggweiler et al., 2006). Again, such a trend is overprinted with marked inter-annual variability. Because of the importance, size and complexity of the Southern Ocean, the incompleteness of observations, and its high variability at a range of temporal and spatial scales, it is critical to improve our knowledge of this ocean/climate system. To re-emphasise the introduction to this chapter, the Southern Ocean has a profound influence of the distribution of salt, heat and ventilating gases throughout global seas. At the same time it is also undergoing some of the most rapid environmental changes on Earth highlighted by the warming, glacial retreat and ice shelf collapse around the Antarctic Peninsula. Thus, to address the inevitable questions relating to impacts of a rapidly changing climate on the Southern Ocean we require a strong modelling effort, supported by multi-seasonal oceanographic and remotely sensed observations and high-resolution palaeoceanographic records of past warm extremes. While this may seem to be a well-worn message, it is still appropriate at a time of certain change and uncertain consequences.

ACKNOWLEDGEMENTS

We are indebted to the World Ocean Circulation Experiment (WOCE) for permission to use their data for Figures 4.1, 4.3 to 4.6, which are attributed to Orsi, A. H., T. Whitworth III, Hydrographic Atlas of the World Ocean Circulation Experiment (WOCE). Volume 1: Southern Ocean (eds. M. Sparrow, P. Chapman and J. Gould), International WOCE Project Ofﬁce, Southampton, U.K., ISBN 0-904175-49-9, 2005. The chapter beneﬁted from the critiques by the external reviewers, Will Howard and Alejandro Orsi and their input is appreciated. Funding for L. Carter was provided by the Foundation for Research Science and Technology, contracts CO50410 and VICX0704. 108 L. Carter et al.

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