circulation and water masses of the southern ocean: a review€¦ · water (aaiw) above and...

Developments in Earth & Environmental Sciences, 8F. Florindo and M. Siegert (Editors)r 2009 Elsevier B.V. All rights reservedDOI 10.1016/S1571-9197(08)00004-9

Chapter 4

Circulation and Water Masses of theSouthern Ocean: A Review

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

1Antarctic Research Centre, Victoria University, P.O. Box 600, Wellington,New Zealand2Department of Earth Sciences, University of Cambridge, Downing Street, CambridgeCB2 3EQ, UK3National 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. Itscirculation is complex, with a zonal Antarctic Circumpolar Current (ACC)interacting with a meridional thermohaline circulation. The ACC is a highlyvariable, deep-reaching eastward flow driven mainly by the westerly winds. It isthe longest (24,000 km), largest (transport 137–147.106m3 s�1) and only currentto connect the major oceans. The Ekman component of the westerly winds alsodrives surface waters north. Near the ACC’s northern limit, these waters sink toform Subantarctic Mode and Antarctic Intermediate waters, which continue northat depthsoB1,400m. Interacting with the ACC is the density-forced thermohalinecirculation. 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 basinsaround Antarctica. Less dense water is entrained by the ACC and mixed withdeep water moving south from the Atlantic, Indian and Pacific oceans. Theresultant Lower Circumpolar Deep Water is tapped off by deep western boundarycurrents that enter the three oceans at depths WB2,000m. These northwardinflows, with a total volume transport of B55.106m3 s�1, disperse Antarctic and

E-mail: [email protected] (L. Carter).

�Corresponding author. Tel.: þ64 4 463 6475; Fax: þ64 4 463 5186;

86 L. Carter et al.

northern-sourced waters throughout the world ocean. Other circulation elementsare the deep-reaching, cyclonic Weddell, Ross and unnamed gyres located south ofthe ACC. Further south again are the westward Antarctic Slope and Coastalcurrents that pass along the Antarctic continental margin under easterly polarwinds.

4.1. Introduction

The Southern Ocean has a profound influence on the world’s ocean andclimate. Cold, dense water sinks to abyssal depths around the margins ofAntarctica and migrates northwards into the Atlantic, Indian and Pacificoceans via deep western boundary currents (Fig. 4.1; Stommel, 1958;Warren, 1981). As succinctly noted by Warren (1971), ‘y water from theAntarctic 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 upperocean. There, they contribute to the warm surface circulation that extendswest from the Pacific and Indian Oceans into the Atlantic where the warm,saline water moves north. Approaching high northern latitudes, the watercools and sinks to form North Atlantic Deep Water (NADW), whichmigrates south, sandwiched between northbound Antarctic IntermediateWater (AAIW) above and Antarctic Bottom Water (AABW)/LowerCircumpolar Deep Water (LCDW) below (Fig. 4.2). En route, NADWmixes with other waters and eventually rises at the Antarctic continentalmargin. Thus, one cycle of the global thermohaline circulation (THC) – amajor regulator of Earth’s ocean and climate – is completed and anothercycle 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 andmeteorologically forced circulations, and (ii) their direct access to the majorocean 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 aflavour of over 70 years of oceanographic research in the Southern Ocean.Thus, we refer the reader to the reference list for a more detailed insightinto the workings of this region. We present the basic elements undertwo sections: (1) Section 4.2 examines the main water masses, focusing ontheir properties and the mechanisms that control their distribution, and(2) Section 4.3 reviews the structure and dynamics of the world’s largestocean current, the ACC, together with that of the subpolar gyres and

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edG

.

SW Indian R.

Mid A

tlantic R.

Chatham Rise

PacificAntarctic R.

F. Pl.

S. Tasman Rise

Main DWBCinflow

40°S

Figure 4.1: The main Oceanographic elements of the Southern Oceanincluding: (i) the ACC contained by the Subantarctic Front (SAF) andsouthern limit of UCDW or southern boundary (SB); (ii) the Ross, Weddelland unnamed subpolar gyres; and (iii) the main exit points of deep westernboundary currents from the Southern Ocean (blue arrows). The general pathfor 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).

Circulation and Water Masses of Southern Ocean 87

currents residing south of the ACC, and the deep THC. The chapter endswith a discussion in Section 4.4 of the present debate regarding the SouthernOcean’s response to a rapidly warming climate (e.g. Gille, 2002; Curry et al.,2003; Jacobs, 2004; IPCC, 2007).

ACC

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0

LATITUDE

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PT

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)

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Figure 4.2: Schematic section of the main water masses and their meridionaltransport as compiled from Whitworth (1988); Orsi et al. (1995); Speer et al.(2000) and Rintoul et al. (2001). Water masses are SAMW, SubantarcticMode Water; AAIW, Antarctic Intermediate Water; UCDW, UpperCircumpolar Deep Water; LCDW, Lower Circumpolar Deep Water;NADW, North Atlantic Deep Water; AABW, ‘true’ Antarctic BottomWater (gnW28.27 kgm�3). 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.

88 L. Carter et al.

4.2. Water Mass Formation and Dispersal

4.2.1. Surface Ocean

A series of ocean fronts – narrow, variable bands defined by abrupt changesin water properties, in particular, temperature and salinity – divide thesurface waters of the Southern Ocean into several zones (e.g. Gordon, 1975;Deacon, 1982; Whitworth, 1988). Early studies identified (from south tonorth) the Polar, Subantarctic and Subtropical fronts (Fig. 4.2). More recenthydrographic transects, especially those carried out during the World OceanCirculation Experiment (WOCE), have revealed additional boundarieslocated south of the Polar Front, and termed the ‘southern’ and ‘southern


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

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E. Pacific Rise

Kerguelen

Plateau

Cape Basin

TasmanBasin

EnderbyBasin

Madg.Basin

Mozb.Basin

P14

S2

Figure 4.3: Location of the principal ocean frontal systems in the SouthernOcean (based on Orsi et al., 1995, but modified for the New Zealand regionaccording to Carter et al., 1998 and Morris et al., 2001). Repeatedhydrographic transects, satellite observations and drifting floats reveal thefrontal systems as dynamic features with marked temporal and spatialvariability but generally within the constraints imposed by the ocean floortopography (Moore et al., 1999). P14 and S2 are locations of WOCEhydrographic transects portrayed in Figs. 4.4 and 4.5. Madg., MadagascarBasin; 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 bytheir potential temperature (y), salinity (S) and oxygen content (seehydrographic charts in Orsi and Whitworth, 2005). (1) Near-freezing andrelatively fresh Antarctic Surface Water (AASW) forms a layer about 100mthick that extends from the Antarctic continental shelf to the Polar Front,commonly defined as the northernmost extent of the subsurface temperatureminimum (Belkin and Gordon, 1996; Figs. 4.4 and 4.5). AASW temperaturesare 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 andWeddell seas, whereas AASW elsewhere around Antarctica commonly hasSo34.0 psu. (2) Between the Polar and Subantarctic fronts resides surfacewater that is transitional between AASW and Subantarctic Surface Water(SASW). The structure is complex in response to mixing and interleaving ofAASW 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 andy is 3–81C (Fig. 4.5). (3) SASW occurs north of the Subantarctic Front andencompasses water that usually warms northwards from B61C to 121C.Salinity is typically W34.3 psu except in the SE Pacific and Drake Passagewhere 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 theSubtropical Front where temperatures increase sharply by 4–51C and salinityby 0.5 psu (Fig. 4.3; Deacon, 1982). Subtropical surface water prevails northof the Subtropical Front.

4.2.2. Subantarctic Mode Water and Antarctic Intermediate Water

Isopycnals – surfaces of constant density – of near-surface to deep waters inthe Southern Ocean rise up in a step-like profile towards Antarctica. Close toocean fronts, isopycnals may outcrop indicating either the rise of deep watersto the ocean surface or the descent of surface waters (Figs. 4.2, 4.4 and 4.5). Inthe case of the latter, winter cooling and mixing of the surface waters on thenorthern side of the Subantarctic Front forms Subantarctic Mode Water(SAMW) (McCartney, 1977; Morris et al., 2001; Rintoul et al., 2001). Thiswell mixed, ventilated water descends northwards to B500m depth alongmuch of the front (Fig. 4.2). AAIW also descends from the surface, passing

Figure 4.4: Sections of potential temperature (A), salinity (B) and neutraldensity (C) from WOCE Line P14 across a major constriction in the ACCbetween the Ross Sea and New Zealand (see Fig. 4.3 for location). Isolines atthe Antarctic margin indicate descent of dense shelf waters, which may bemixed with NADW-influenced, LCDW as suggested by the salinity field. Theresultant AABW (cf. Fig. 4.5) is contained within the SW Pacific Basin.At the surface, north of the polar front, low salinity AAIW descendsnorthwards (Fig. 4.4B). Hydrographic profiles were derived from the WOCESouthern Ocean Atlas at http://woceatlas.tamu.edu/. Names of fronts are

given in Fig. 4.2.

Distance (km)

De

pth

(m)

° ° °°[C]

NEUTRAL DENSITY

AABW

LCDW

Figure 4.4: (Continued).

92 L. Carter et al.

under SAMW to reach a maximum depth of B1,400m (Figs. 4.2 and 4.4).AAIW is identified by a salinity minimum (34.3–34.5 psu) and temperatures ofB3–71C. However, the processes driving AAIW formation are unclear.Formation may be related to wind-forced or density-driven sinking of coldAASW and indeed there appears to be continuity between the AASW andAAIW salinity fields (Figs. 4.4 and 4.5). However, McCartney (1977)suggested AAIW may evolve at least in part from dense SAMW. Whateverthe origin, compared to the widespread formation of SAMW, new AAIWpresently appears to form mainly in the SW Atlantic and SE Pacific. Fromthese sites AAIW circulates the oceans in anticyclonic subtropical gyres thatextend north towards and locally beyond the equator before returning southas ‘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 CircumpolarDeep Water (CDW). It extends from B1,400m to W3,500m depth, but it

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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. Ofnote is (i) the step-wise ascent of saline, NADW-rich LCDW towardsAntarctica where it is capped by fresh, cold AASW in the south andsubducting AAIW in the north and (ii) the containment of a large volume ofclassic Antarctic Bottom Water (AABW; gnW28.27 kgm�3) within theWeddell Basin (cf. Fig. 4.4). Hydrographic profiles are derived from theWOCE 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 CircumpolarDeep Water (UCDW) is identified by the oxygen minimum and elevatednutrient concentrations, and has an open ocean depth range of B1,400–2,500m, and (2) Lower Circumpolar Deep Water (LCDW) whose signatureis 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 theSouthern Ocean (Orsi et al., 1995; Rintoul et al., 2001). Upon meeting theACC, NADW is entrained and transported east around the Antarcticcontinent, all the while mixing with waters from the Indian and Pacificoceans plus dense waters from Antarctica to form LCDW. Despite thevigorous mixing, the high-salinity signature of NADW is retained (Reid andLynn, 1971; Gordon, 1975). At several locations around Antarctica, LCDWrises at the continental slope where mixing with super-cold shelf water renewsnot only the deep circulation of LCDW but also generates Antarctic BottomWater (AABW), the deepest water mass in the Southern Ocean (Figs. 4.2, 4.4and 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 westernboundary currents (see Section 4.3.3 and Schmitz, 1995; Hogg, 2001).According to Rintoul et al. (2001) mixing with fresher waters, together withthe biological depletion of oxygen, slowly modify LCDW into a less dense,lower oxygenated variant that returns south as UCDW. In some WOCEsections, such as P15 in the Indian Ocean, oxygen depletion of UCDW maybe influenced by the direct injection of nutrient-enriched deep waters fromthe North Indian and North Pacific oceans.

4.2.4. Antarctic Bottom Water

In their analysis of abyssal water masses, Mantyla and Reid (1983) drewattention to the often inappropriate use of the term Antarctic Bottom Water,which has tended to be used generically for any southern-sourced bottomwater. They demonstrated that true AABW did not extend far fromAntarctica before mixing with other waters (see also Orsi et al., 1999). This isparticularly true for the deep western boundary currents in which AABW ismixed 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 kgm�3.Such dense waters are confined mainly to the deep (down to B6,000m),


circum-Antarctic Basins that include the Argentine and Brazil basins (SWAtlantic), Mozambique, Crozet and Australian-Antarctic basins (Indian)and the SE Pacific Basin (Figs. 4.1, 4.4–4.6). Less dense, southern bottomwater with 28.18ogno28.27 kgm�3, is not confined to the circum-Antarcticbasins, but instead spreads out from the deep levels of the northern edge ofthe 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; yr�11C) bottomwater occurs in the Weddell Sea, whereas the SE Pacific Basin has the mostsaline and least cold (SZ34.72 psu; �0.6oyo�0.31C) bottom water. TheAustralian-Antarctic Basin contains water with properties intermediatebetween the two end members. Traditionally, the Weddell Sea was regardedas the prime source of AABW, but recently two other sources have cometo the fore. The Weddell Sea’s contribution is now regarded as B50%,with the Wilkes Land margin including the Adelie coast (Rintoul, 1998)contributingB30%mainly to the Indian Ocean, and the Ross Sea producingB20% 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 shelfwater (HSSW) by brine rejection from sea ice. However, salt-driven increasesin density may also be influenced by intrusions of NADW-bearing LCDWonto the upper slope and shelf (Toggweiler and Samuels, 1995). Super cold,Ice Shelf Water (ISW), formed by freezing and melting below ice shelves, canmix with HSSW and reach the outer shelf before flowing down slope (Bainesand Condie, 1998). Alternatively, the simple mixing of cold AASW and high-salinity LCDW, with or without ISW, may produce negatively buoyantwaters at the shelf edge (e.g. Jacobs, 2004). Finally, dense waters may sinkvia convection chimneys and polynyas such as the well-documented butshort-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 106m3 s�1). Anthropogenic tracers, inparticular chlorofluorocarbons, record a flux of 8.1 Sv for AABW descend-ing at the 2,500m isobath off Antarctica (Orsi et al., 2001). This compares to7.6 Sv of lower NADW flowing out of the Nordic and Labrador seas at theN Atlantic source. Because AABW is colder than its northern counterpart,Antarctic overturning probably plays the dominant role in cooling the deepocean. The time at which deep water circulates through the ocean is oftenquoted to be a millennium or more, for example 1,000 years between theN Atlantic and Southern Oceans and a further 1,000 years from the Southern

Figure 4.6: Extent of dense Antarctic Bottom Water as identified by theneutral 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 viathe deep Argentine and Brazil basins, the remaining AABW is captured incircum-Antarctic basins where further northwards dispersal is inhibited byoceanic ridge systems shown in white. Basins annotated as Arg. B., ArgentineBasin, 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/.

96 L. Carter et al.


Ocean to the North Pacific. However, in a re-analysis of radiocarbon datedocean waters, Matsumoto (2007) indicates much shorter circulation agesthereby supporting but refining earlier radiocarbon-based studies (Stuiveret al., 1983). Thus, the circulation age for the Southern Ocean below 1,500mis B300 14C years with a similar age for the Atlantic. For the Pacific, thebasin 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 systemthat is rightly described by superlatives. It is the only current to connect themajor ocean basins and hence plays a prominent role in the globaldistribution of heat, salt and gases (Fig. 4.1). It is the longest current withan estimated pathway of B24,000 km (Whitworth, 1988). Finally, the ACCis the largest major current in terms of volume transport with a mean of136.777.8 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 (Orsiet al., 1995; Whitworth et al., 1998). However, use of the term wind driftmasks the role played by the buoyancy-driven component of the circulation(see Rintoul et al., 2001). Complexities aside, the net result is an eastwardcurrent system that extends from the ocean surface to bottom, its pathguided 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 ridgesystems where the current shifts poleward (Fig. 4.1). Large submarineplateaux also exert an influence. The ACC widens to the north and south asit passes around the Kerguelen Plateau, whereas the Campbell and Falklandplateaux form constrictions (Fig. 4.1; Whitworth and Peterson, 1985; Morriset al., 2001). Interestingly, this interaction with the ocean floor was inferredas early as the 1950s. Estimates for a purely wind-driven ACC yieldedtransports that were excessive compared to observations. Thus, it wasconcluded that the wind stress was partly balanced by bottom stress (seeWhitworth, 1988; Rintoul et al., 2001). The passage of westerly winds overthe ACC also induces an Ekman drift to the north – a process that probablyplays a role in the subduction and transport of mode and intermediatewaters. 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 thecircumpolar fronts (see Section 4.2 Surface Ocean) with the northern andsouthern boundaries of the ACC defined, respectively, by the Subantarcticand Southern Boundary fronts (Figs. 4.1–4.3). Both eddy-resolving modelsand 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 isconstricted as off Campbell Plateau and within Drake Passage (Nowlin andKlinck, 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 widerange of values for the entire ACC. Nevertheless, closely spaced and long-term monitoring sites have improved estimates of transport. In DrakePassage the mean transport is 136.777.8 Sv (Cunningham et al., 2003)compared to 147710 Sv between Australia and Antarctica (Rintoul andSokolov, 2001). Much of the transport in the Australasian reach of the ACCoccurs within the Subantarctic Front, which has a mean of 10577 Sv offAustralia (Rintoul and Sokolov, 2001) and B90 Sv off southern NewZealand (Morris et al., 2001). However, ACC transport can be highlyvariable. Time series from Drake Passage record variations at several timescales (Whitworth and Peterson, 1985). Short-term fluctuations, related to14-day solar and lunar tides, are superimposed on longer-period fluctuationsof B1 year and longer that can lead to changes in transport of B30–40 Svwithin weeks.The interaction of the ACC with the topography and southerly extensions

of western boundary currents generates eddies that play important roles inthe transfer of heat and momentum (Bryden and Heath, 1985; Morrow et al.,1992; Rintoul et al., 2001). Off SE New Zealand, for example interception ofthe 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 anticyclonicfeatures have been observed from current meter and satellite data, whichsuggest a frequency of occurrence of B9 eddies annually (Stanton andMorris, 2004). Modelled eddy kinetic energy, verified by current meter dataand ocean floor sedimentary evidence, attest to the power of theseperturbations, which are likely to be the cause of abyssal benthic storms(Hollister and McCave, 1984; Carter and Wilkin, 1999). Seabed topographyalso encourages intense mixing within the ACC. The Scotia Sea and


potentially other areas of marked seabed relief below the ACC, are zones ofthe highest turbulent mixing in the ocean and result in rapid upwelling thatmay locally short-circuit the classic meridional overturning as portrayed inFig. 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 Antarcticcontinental margin (Fig. 4.1). The better known are the Weddell and Rossgyres 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 wassuggested to occur south of Kerguelen Plateau (Bindoff et al., 2000), and hasnow been confirmed by McCartney and Donohue (2007). As documented forthe Weddell Gyre (Orsi et al., 1993), it is likely that all three gyres transportsalt and heat from the ACC to the Antarctic continental margin where deepand bottom waters are produced (e.g. Jacobs, 2004). Furthermore,McCartney and Donohue (2007) suggest a strong connectivity between thethree gyres with a westward flow along their southern limbs and an eastwardflow joining their northern limbs. This latter flow is just south of anothereastward flow, this time associated with an anticyclonic supergyre coveringmost of the S Pacific and S Indian oceans (Ridgway and Dunn, 2007). Thiseastward limb of the supergyre appears to reside between the Subtropicaland Subantarctic fronts. As a result, it may entrain SAMW, formed in thevicinity of the Subantarctic Front and help distribute it through the oceanbasins 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 surfaceit has the form of a NE–SW aligned, elongated gyre whereas at depth itcomprises two cyclonic cells located to the east and west of 151W. Basically,the Weddell Gyre occupies the W4,000m deep re-entrant formed byAntarctica and the ridge systems that extend eastward from near the tip ofthe Antarctic Peninsula (Fig. 4.1). While such a location implies contain-ment, the northern limb of the gyre overlaps the ridges to interact with theACC. CDW entrained from the ACC is moved within the gyre and caneventually 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 tothe W4,000m deep western reach of the SE Pacific Basin (Fig. 4.1). Like itsWeddell Sea counterpart, the Ross Gyre is a deep-reaching feature thatentrains 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, mostof which is W4,000m deep (Fig. 4.1). Indeed, the volume transport field iscompressed against the Antarctic continental margin and the associatedwestward slope current (Bindoff et al., 2000; McCartney and Donohue,2007). Transport along the eastward-flowing northern limb of the gyre isestimated at 35 Sv. However, the amount of transport along the west-moving, southern limb is unclear due to merging with the slope current, thecombined 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 isbathed by two westward currents. One is associated with the Antarctic SlopeFront that constitutes the boundary between fresh, cold Antarctic shelfwaters and less cold, saline CDW (Jacobs, 1991; Whitworth et al., 1998).Deacon (1937) regarded the frontal flow as a consequence of the prevailingeasterly winds and coined the name, East Wind Drift. On the basis ofclassical theory, he reasoned that polar easterly winds produced an onshoreEkman transport with the resultant generation of a westward geostrophiccurrent below the wind-mixed layer (Whitworth et al., 1998; Bindoff et al.,2000). The second significant feature is the Antarctic Coastal Current, whichforms a narrow, rapid flow across broad sections of the continental shelf, forexample, in the SW Weddell Sea. However, where the shelf is narrowthe coastal current is difficult to distinguish from flows associated with theAntarctic Slope Front and the southern limbs of the subpolar gyres wherethey 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 watercooled (or heated) by contact with cold (or warm) air, or modified by sourcesand sinks of cold water. May also include flows whose characteristics are


significantly altered by upwelling and/or mixing. Water sinking at highlatitudes tends to return equatorwards in relatively strong, narrow currentscalled DWBC’. The high latitude sinks noted are primarily the NorthAtlantic 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 continentalslope. Descent may take place in several ways according to Baines andCondie (1998). For prominent deep-water formation areas, dense water maydescend as broad sheets or plumes. For weaker sources, the resultantoutflows may geostrophically descend the slope to a level where the layerthins and viscous drainage prevails. Descending flows may be disrupted byeddies into discrete parcels that overall move down slope. Further complex-ities are introduced by the slope topography, for example, submarinecanyons and channels can contain and steer flows whose density mayincrease through entrainment of sediment. Whatever the mode of descent,the initial dispersal of AABW is westward as outlined by the distribution ofchlorofluorocarbons (Orsi et al., 1999). Those tracers show that the densestAABW usually follows a cyclonic path presumably under the influence of thesubpolar gyres and basin topography (Fig. 4.1). However, the Weddell–Enderby Basin experiences limited outflow at its northern rim where leastdense AABW (gnB28.28 kgm�3) passes north into adjoining basins of theAtlantic and Indian oceans (Fig. 4.6; Orsi et al., 1999). In contrast, the maininflows 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 BottomWater, a local variety of AABW, together with WeddellSea Deep Water and CDW from the Drake Passage, are carried by anorthbound DWBC into the Atlantic Ocean below southward-movingNADW (Warren, 1981; Schmitz, 1995). The DWBC is typically found inwater depths exceeding 3,500–4,000m along the western boundary presentedby the continental margin off South America. However, the current pathwayis interrupted by a succession of deep basins including (from south to north)Georgia, Argentine and Brazil basins. As noted earlier, densest AABW iscaptured within the deep basins leaving less dense waters to move north viagaps and channels through the inter-basin ridges. Even so, not all the deepwater escapes and some recirculates within the basins themselves (Hogg andJohns, 1995; Hogg, 2001). For example, around 6.9 Sv flows from the


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 andMozambique basins (the latter being a dead end), which largely constrainAABW to south of 301S and 341S, respectively (Orsi et al., 1999). LCDWfrom the Crozet Basin leaks northwards through fractures zones in the SWIndian Ridge (Warren, 1974; Johnson et al., 1991; Mantyla and Reid, 1995).Deep transport is via a northbound DWBC, which like its Atlanticcounterpart, has a complex pathway dictated by multiple basins and ridges(Warren, 1981; McCave et al., 2005). Basins encourage recirculation of deepwaters and together with mixing, account for a northward dissipation of theDWBC. For instance, LCDW below B3,800m has a northwards transportof 3.8 Sv through Madagascar Basin, but only 1.7 Sv exits north into theSomali 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 connectedto the Southern Ocean thereby allowing the northward intrusion of twoother boundary currents; one along the eastern side of SE Indian Ridge andthe other along the eastern flank of Ninetyeast Ridge (Warren, 1981; Tooleand Warren 1993; Reid, 2005). Again, the deep water carried north isLCDW, with the denser AABW retained in Australian-Antarctic Basin (Orsiet al., 1999). When extended to the 2,000 dbar reference level, the combinednorthward 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 theclassic model of the global abyssal circulation by Stommel (1958) andStommel and Arons (1960), and was later confirmed by the hydrographicsections of Warren (1971, 1973). Because the Tasman Basin is essentiallyclosed at its northern end, the main inflow is off southernmost New Zealandwhere the DWBC enters in concert with the ACC (Fig. 4.7; Carter andMcCave, 1997; Carter and Wilkin, 1999). Initially, the combined inflowintercepts Macquarie Ridge to form meanders and eddies although somecurrent filaments pass through narrow gaps in the ridge (Boyer and Guala,

SINKING

SINKING

UPWELLING

UPWELLIN

G

A.C.C.A.C.C.

UPW

ELLING

Cold deep current

Cool shallow current

Warm shallow current

Thermohaline Circulation

SINKING

SINKING

Figure 4.7: A schematic portrayal of the THC, which is a series of looselylinked, recirculation systems that transport, heat, salt, nutrients andventilating gases through the world ocean. Only the main elements of theTHC are shown, for example the Indian Ocean has three, deep northwardinflows that include from W to E, the margin off eastern Africa (shown), andthe 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 northeastalong the 3,000–3,500m high flanks of Campbell Plateau to around 491Swhere the ACC veers east leaving the DWBC to continue northwards intothe Pacific Ocean (Fig. 4.7). It eventually departs the Southern Ocean offChatham Rise between 441S and 421S (McCave and Carter, 1997). There,Warren (1973) observed a volume transport of B20 Sv, the largest for asingle DWBC (Schmitz, 1995).

4.4. Oceanographic Variability and Change

The present phase of climate change has drawn considerable attention tothe 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


hindered by; (i) sampling bias, (ii) the short history of observations, whichtypically only encompass the last 50–60 years, (iii) a multiplicity of forcingmechanisms, some with and some without clear connections to a warmingclimate 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). Ofspecial note is the Southern Annular Mode (SAM), which appears todominate inter-annual to centennial variability in the Southern Ocean (Halland Visbeck, 2002; Lovenduski and Gruber, 2005; Lovenduski et al., 2007).SAM is essentially a zone of climate variability that encircles the South Poleand strongly influences zonal winds, sea ice formation and oceaniccirculation. When in a positive phase, SAM is typically associated withenhanced westerly winds over the ACC and weakened winds further north.This favours a strengthening of the ACC, a reinforcement of the northwardEkman drift and subduction of surface waters, and a northward expansionof sea ice. To compensate, there is enhanced poleward transport and risingof deep water at the Antarctic continental margin. Under a negative SAM,windiness and storminess appear to migrate to mid-latitudes thusweakening both zonal and meridional ocean transport.Despite uncertainties associated with the aforementioned limitations, it is

nonetheless important to examine and evaluate recent changes in Antarcticaand the Southern Ocean in light of their actual or potential influence on theglobal 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 resultsemphasise 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 ofþ0.561C/decade from 1951 to 2000. The next largest warming trend outsideof the Antarctic Peninsula is in the western Ross Sea (Scott Base) with a riseof þ0.291C/decade, although the rise is not statistically significant.Elsewhere, significant trends are unclear. Annual temperature trends forcoastal and interior sites on the East Antarctic Ice Sheet suggest a slightcooling. 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 pronouncedatmospheric warming has been accompanied by an equally marked warmingof the surface ocean with summer temperatures increasing by 1.21C over thelatter half of the twentieth century (Meredith and King, 2005). At waterdepths of 700–1,100m, Gille (2002) reported an average 0.171C rise since1950, which accounts for about two-thirds of the total increase in heat


content in the ocean from 0 to 3,000m depth (IPCC, 2007). Much of thewarming is concentrated within the Subantarctic Front of the ACC. Thisraises the possibility that the warming at depth may result from the sinkingof atmospherically warmed SAMW. Gille (2002) further suggests thatwarming may also be related to a general southward displacement of theACC. This would be consistent with a suite of model simulations that suggestwarming of the Southern Ocean will be accompanied by a southward shift ofzonal westerly winds together with a narrowing and intensification of theACC (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 tosubtropical and tropical regions, which display increased salinity withthe exception of the central Pacific Ocean (Curry et al., 2003; Boyer et al.,2005). Such freshening of the upper Southern Ocean may be responsiblefor a reduction in the salinity of AAIW (Wong et al., 1999; Curry et al.,2003). Reduced salinity implies increased freshwater input that may resultfrom one or more of the following causes: (i) greater net precipitation;(ii) changes in the extent of sea ice – a major contributor to winter salinitythrough brine rejection; (iii) increased melting of ice shelves, ice sheetsand 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, hydrographicrecords spanning over 40 years show local variability in salinity trends.The upper 50m of the ocean, west of the Antarctic Peninsula, hasbecome more saline although underlying waters have freshened slightlyin line with the regional trend. While the more saline surface conditionsappear 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 salineconditions are consistent with the reduced sea ice cover and the timing ofthe hydrographic measurements. With less sea ice production there isless freshening of the ocean in the summer when most of the hydrographicmeasurements are made. On the opposite side of the continent, at LawDome, ice core records spanning a century and longer, identify a 20% lossof sea ice since 1950 although this trend is strongly overprinted withcyclical variations with an B11 year frequency (Curran et al., 2003).Reduced sea ice along with increased precipitation and melt waterfrom the West Antarctic Ice Sheet have been cited as contributing to theobserved freshening of surface waters associated with the Ross Sea Gyre(Jacobs et al., 2002). But like Law Dome, the trend is obscured by cycles, thistime by 5–6- and 9-year oscillations in HSSW formation (Assmann andTimmerman, 2005).


Salinity and temperature changes have the potential to affect bottom waterproduction and the THC. Consequently, these changes have receivedconsiderable attention (IPCC, 2007 and references therein). At glacial-interglacial time scales, palaeoceanographic evidence reveals markedvariations in the position and degree of convective overturning of the NAtlantic sector of the THC (Rahmstorf, 2002 and references therein). Duringinterglacial periods, overturning is most active and reaches its most northerlyextent. In contrast, glacial periods are likely to witness a southward shift andpossible slow-down in overturning. Any slow-down may be compensated bya greater production of bottom water from Antarctica as suggested by grainsize (Hall et al., 2001), magnetic properties (Venuti et al., 2007) and diatomproxies (Stickley et al., 2001). However, such conclusions are sometimes atodds with geochemical tracers such as d13C (e.g. Moy et al., 2006) that pointto little change in the passage of NADW through the Indian and Pacificsectors at least over recent glacial-interglacial cycles. At millennial timescales, abrupt changes such as those associated with Heinrich events, maystop N Atlantic overturning altogether as the density of the surface ocean isreduced by rapid influxes of melt water (Rahmstorf, 2002). Again, cessationof N Atlantic production may be compensated by enhanced Antarcticproduction. However, responses to the latest phase of climate/oceanwarming are unclear. In the N Atlantic, which is the best observed deep-water source, long-term trends are equivocal due to decadal variability, apaucity of long-term observations and other factors (IPCC, 2007). A similarsituation applies to Antarctica where estimations of bottom-water produc-tion are inconsistent in response to: (i) natural cycles; (ii) differences in thedefinitions and techniques to estimate production rates, and (iii) a biastowards summer observations (Jacobs, 2004). On the basis of chlorofluor-ocarbons and 14C data, which allow water masses to be traced at decadal tocentury scales respectively, Orsi et al. (2001) revealed no declinein bottom water production over the twentieth century as indicated earlierby Broecker et al. (1999). Nevertheless, the changes recorded in recenthistorical times cannot be ignored. The freshening of the Ross Gyre over thelast 50 years (Jacobs et al., 2002) and an accompanying downstreamfreshening of AABW in the adjacent Australia-Antarctic Basin (Aoki et al.,2005) are consistent with increased freshwater input. On a larger scale, thehistorical salinity data of Curry et al. (2003) also reveals a freshening of deepand 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 efficientenergy technologies with reliance on a range of sources) point to an average25% reduction in N Atlantic overturning by the year 2100 (IPCC, 2007).


None of the runs point to a shut-down; rather they favour reductionsin overturning of up to 50%. While the Southern Ocean sector has receivedless attention from modellers, simulations based on a warmer or fresherocean may enhance or stabilise N Atlantic overturning (Saenko et al., 2003;Weaver et al., 2003). To further emphasise the complexity of north–southrelationships, the projected strengthening of the Southern Hemispherewesterly winds will increase the northward Ekman transport of upper ocean.To compensate, the poleward flow of deep water below 2,000–2,500m depth,could be expected to strengthen and possibly stimulate the southward flow ofNADW (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 oftemporal and spatial scales, it is critical to improve our knowledge of thisocean/climate system. To re-emphasise the introduction to this chapter, theSouthern Ocean has a profound influence of the distribution of salt, heat andventilating gases throughout global seas. At the same time it is alsoundergoing some of the most rapid environmental changes on Earthhighlighted by the warming, glacial retreat and ice shelf collapse around theAntarctic Peninsula. Thus, to address the inevitable questions relating toimpacts of a rapidly changing climate on the Southern Ocean we require astrong modelling effort, supported by multi-seasonal oceanographic andremotely sensed observations and high-resolution palaeoceanographicrecords of past warm extremes. While this may seem to be a well-wornmessage, it is still appropriate at a time of certain change and uncertainconsequences.

ACKNOWLEDGEMENTS

We are indebted to the World Ocean Circulation Experiment (WOCE) forpermission to use their data for Figures 4.1, 4.3 to 4.6, which are attributedto Orsi, A. H., T. Whitworth III, Hydrographic Atlas of the World OceanCirculation 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 fromthe critiques by the external reviewers, Will Howard and Alejandro Orsi andtheir input is appreciated. Funding for L. Carter was provided by theFoundation for Research Science and Technology, contracts CO50410 andVICX0704.


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