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Anatomy of three warm-core Leeuwin Current eddies
Deep-Sea Research II 50 (2003) 2229–2243
Anatomy of three warm-core Leeuwin Current eddies Rosemary Morrowa,*, Fangxin Fanga,1, Michele Fieuxb, Robert Molcardb b
a LEGOS 18, av. Edouard Belin, 31401 Toulouse Cedex 4, France LODYC, Universit!e Pierre et Marie Curie, case 100, 4 place Jussieu, 75252 Paris Cedex 05, France
Abstract The vertical structures of three warm-core Leeuwin Current eddies are described, as well as their temporal evolution determined from altimetric data. These warm-core eddies have large positive sea-level anomalies, although their surface temperature fields are not necessarily warmer than the ambient waters. These eddies were observed some 350 km offshore from the Leeuwin Current in September 2000, around 25–30 S, and their vertical structure and temporal evolution show they were formed at the coast in May, when the Leeuwin Current is strongest. After separation from the current, the warm-core eddies drifted WNW following isopycnal contours, and were strongly steered by bathymetry. These eddies penetrated to at least 1500 m depth; the strongest eddy, Eddy B, influenced isopycnals to 2500 m. Eddy B appeared to be trapped and intensified by a Leeuwin Current squirt that directed warm, low-salinity water offshore from August to October. The average annual eddy heat and salt fluxes for waters warmer than 8 C are estimated at 0.004 PW and 3 105 kg s1, respectively. The magnitude of these eddy fluxes is 20–30% of the annual mean poleward heat and salt flux in the eastern Indian Ocean, and is within the range of values observed for Agulhas eddies entering the South Atlantic. r 2003 Elsevier Science Ltd. All rights reserved.
1. Introduction Most eastern boundary regions of the ocean are synonymous with cold upwelling, equatorward surface flow, and a large heat sink from the atmosphere (Tomczak and Godfrey, 1994). The eastern boundary of the southern Indian Ocean is the opposite: the coastal current, the Leeuwin Current, flows poleward against the prevailing equatorward winds with no coastal upwelling. More importantly, it is the only eastern boundary *Corresponding author. E-mail address: [email protected] (R. Morrow). 1 Current address: Earth Science and Engineering, Imperial College of Science, Technology and Medicine, RSM Building, Prince Consort Road, London SW7 2BP, UK.
region with a large heat transfer from the ocean to the atmosphere, with values equivalent to a western boundary current (Josey et al., 1999). The ocean–atmosphere heat flux is largest along the southwest Australian coastline where the annual heat loss exceeds 50 W m2. This coastal maxima is associated with the strong poleward Leeuwin Current, which transports around 5 Sv of warm, low-salinity waters southward from the Indonesian Throughflow region (Smith et al., 1991). However, the region of ocean heat loss is not limited to the coastal Leeuwin Current and its meanders. A large region of ocean heat loss of up to 25 W m2 extends westward from the coast, reaching 80 E at 20 S, and 95 E at 30 S (Josey et al., 1999). The spatial distribution of this heat
0967-0645/03/$ - see front matter r 2003 Elsevier Science Ltd. All rights reserved. doi:10.1016/S0967-0645(03)00054-7
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loss resembles the beta-refraction pattern associated with Rossby waves. Large sea-level anomalies propagating at Rossby wave speeds are well documented in the 20–35 S band, from altimetric data (Morrow and Birol, 1998; Potemra, 2001) and with greatly reduced amplitude in models (Birol and Morrow, 2001; Hirst and Godfrey, 1993). The propagating anomalies are maximal at annual and semi-annual frequencies, with a strong inter-annual modulation (Morrow and Birol, 1998). Wind-forcing plays a minor role, both offshore and at the coast (Birol and Morrow, 2001). Instead, the positive (warm-core) anomalies appear to detach from the coast between 20 S and 35 S in phase with downwelling coastal propagating waves, which also occur at annual and semiannual frequencies (Potemra, 2001; Birol and Morrow, 2003). Maximum warm-core anomalies occur in May and November during the monsoon transition periods, and may be forced by Pacific wind anomalies that propagate as Kelvin waves through the Indonesian Seas and Timor Passage (Potemra, 2001). Many questions remain about these westwardpropagating warm-core anomalies. Are these features Rossby waves or transient eddies propagating at Rossby wave speeds? What is their vertical structure? How are their propagation paths modified by the bathymetry and background stratification? How do they evolve in time and decay? Most importantly, what is their role in transferring heat and other tracers into the interior of the southern Indian Ocean, and in mixing and modifying the ambient water masses? To help resolve some of these questions, the Transport Indo-Pacific (TIP2000) research cruise was organised in September 2000. One of the aims was to measure the vertical hydrographic structure associated with a few of these propagating anomalies. Eddies are fully detached from the Leeuwin Current meanders after around 90 days (Pearce and Griffiths, 1991), so the September cruise allowed us to measure detached anomalies from the strong May forcing. In this paper, we examine the vertical structure, water masses, generation process and propagation of three warm-core eddies measured during TIP2000. We also estimate the heat and salt fluxes associated
with these three eddies. A companion paper includes a description of the transport and watermass properties of the Leeuwin Current and its eddies during September 2000 by Fieux et al. (2003). Fang and Morrow (2003) use altimeter data to investigate the temporal and spatial evolution of ‘‘long-lasting’’ warm-core eddies generated between 1995 and 2000, and provide details of their propagation pathways and decay rates.
2. Vertical structure of the warm-core eddies The TIP2000 cruise through the Leeuwin Current eddies took place during 1 week from 24 September 2000 to 1 October 2000. The cruise track is shown in Fig. 1, superimposed on a weekly map of positive (warm-core) sea-level anomalies (SLAs) from altimetry. The altimetric SLAs are a combination of Topex/Poseidon (T/P) and ERS2 data mapped onto a regular 0.25 grid, using an improved space/time objective interpolation scheme that eliminates long wavelength correlated noise (Le Traon et al., 1997). The SLA maps are produced by the CLS Space Oceanography Division. Fig. 1 shows that the TIP2000 cruise crossed three warm-core eddies: eddy A at 26 S, 111 E; eddy B at 30 S, 110.5 S; and eddy C at 32 S, 114 E. Their eddy statistics are given in Table 1. Also marked in Fig. 1 is an approximate westward limit for a Rossby wave forced from the eastern boundary in May 2000. The western limit of the Rossby wave position is estimated from analyses of Hovmuller diagrams at various latitudes (not shown) (see Morrow and Birol, 1998; Birol and Morrow, 2001, for Hovmuller diagrams from T/P at these latitudes from 1992 to 1999). Much of this signal is generated by annual and semi-annual coastal Kelvin waves that propagate poleward along the eastern boundary (Potemra, 2001); by linear theory the semi-annual signal can radiate offshore as Rossby waves up to the critical latitude of 25 S. Indeed, north of 25 S, these positive anomalies have longer wavelengths and resemble a coherent Rossby wave train forced from the eastern boundary. South of 25 S, the warm
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(a) 800 16
18
600
20 400
Latitude
22
24
200
26 0 28
30 200 32 400
34 100
102
104
106
108
110
112
114
116
Longitude
Fig. 1. Topex/Poseidon+ERS2 sea-level anomalies (in mm) for the cruise period 25 September–1 October 2000, with the TIP2000 cruise track (solid line) overlaid. Station numbers mark the ends of each section (39, 45, 52, 62, 72). The three warm-core eddies (A, B, C) mentioned in the text are also indicated. A dashed line shows the westward limit of Rossby waves forced from the eastern boundary in May 2000. The 3000 and 4000 m depths are contoured.
Table 1 Statistics for the three warm-core eddies observed during the TIP2000 campaign in late September 2000 Eddy
Centre
Diameter (km)
SLA (cm)
Depth (m)
Drift (cm s1)
Reference station #
A B C
26 S 111 E 30 S 110 E 32 S 114 E
180 350 160
40 80 50
750 2500 1500
6 WNW 4 WNW 1 WSW
47–48–49 56–58, 66–68 56-58, 66-68
anomalies have increasingly smaller radius, appear as more isolated features, and often ‘‘stray’’ from the Rossby wave path (Fig. 1), being locally steered or decelerated by bathymetric features and by non-linear interactions with the stratified flow (Fang and Morrow, 2003). These anomalies appear more like individual eddies, and most likely result from instabilities of the coastal Leeuwin Current at these latitudes (Pearce and Griffiths, 1991; Batteen and Butler, 1998). The arrival of a
semi-annual downwelling coastal Kelvin wave around 25 S may change the coastal density structure, leading to meanders and detached eddies with a large positive sea-level anomaly (Birol and Morrow, 2003). The vertical temperature and salinity structure associated with these warm-core eddies and their surrounding water is shown for the four TIP2000 transects between 25 S and 32 S (Fig. 2), and on a T2S diagram (Fig. 3). Stations along section 1
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were full depth to 20–100 m from the bottom; sections 2 and 4 were to 2000 m depth or 20–100 m from the bottom, whichever came first. Stations along section 3 were full depth, except for station 57 to 5150 db, station 58 to 4000 db, and stations 59–62 to 3000 db. Here, we present the water-mass characteristics in terms of their T2S structure; the full water-mass analysis, including other tracers, is given by Fieux et al. (2003). Close to the coast in Fig. 2, the poleward-flowing Leeuwin Current is evident from the downsloping isotherms and isohalines, and the characteristics of warm (T >20 C), low-salinity water (So35.4, even 35.0 around station 40). The equatorward undercurrent at depths >400 m is also present on the slope, with upsloping isotherms and isopycnals (Smith et al., 1991). The ambient water (outside the warm-core eddies) has a vertical structure similar to that defined by Warren (1981) from a hydrography section at 18 S. The surface layers are influenced by the Indonesian Throughflow Water from the north, with warm (T > 19 C) water and salinity o35.3 (Tomczak and Large, 1989) depending on the latitude; the depth of this low-density surface layer decreases from 140 m around 25 S to 70 m around 30 S. Directly under the surface layer is a 125 m thick layer of salinity maximum water (S >35.6); this water is formed from the strong evaporation in the subtropical gyre between 25 S and 35 S (Warren, 1981). Below, the linear decrease in temperature and salinity in the range 9–14 C marks Indian Central Water (Rochford, 1969; Wyrtki, 1971; You and Tomczak, 1993). This in turn overlies a thermostad of 8–9 C water that is more than 150 m thick around 600 m depth, which corresponds to Subantarctic Mode Water formed in the southeast Indian Ocean (McCartney, 1977; Karstensen and Tomczak, 1997). Antarctic Intermediate Water, with a salinity minimum o34.5 (Wyrtki, 1971; Toole and Warren, 1993; Fine, 1993), is evident from 800 to 1000 m. The salinity minimum is more clearly defined in the southern-most sections. The northern sections and those close to the coast show evidence of mixing, which may indicate some recirculation around the eastern boundary of the Perth Basin. Below 1500 m, the cold (To4 C),
saline (S >34.6) water corresponds to Indian or Circumpolar Deep Water. The three warm-core eddies have a similar vertical structure as the ambient water, but with the isopycnals shifted downwards and capped by warm, low-salinity surface water. Eddies A and B, corresponding to the stations further offshore, tend to be warmer and saltier in the thermocline and at the depths of the AAIW salinity minimum. The core of eddy A was measured at stations 43–44–46; 350 km offshore (Table 1). A lowdensity cap 30 m deep exists at station 45, with T >22 C, and So34.8; this relates to surface waters drawn in from the north around the edge of eddy A. In the centre, at station 43, the surface temperature is lower and the surface salinity higher but the warm–fresh mixed layer reaches 200 m. These temperature values are close to those observed in the coastal current in September 2000, but the low-density mixed layer is deeper than at the coast. Eddy A perturbs the ambient isothermals and isohalines to at least 850 m depth. Eddy C was measured at stations 69–70 on section 4, 150 km offshore. Its low-density cap corresponds to a 130 m thick mixed layer, and Eddy C influences the vertical density structure to B1100 m depth. Eddy B was the strongest of the three, observed at stations 60–64 at the limit of transects 3 and 4. Its low-density core had a surface mixed layer of 20 C, 35.3 water extending to 240 m, which perturbed the underlying density field to 2500 m depth. As for eddy A, this warm, low-salinity core is more deep reaching than the vertical structure of the Leeuwin Current in September.
3. Temporal evolution of the warm-core eddies The vertical structure of the offshore warm-core eddies, and particularly the depth of their lowdensity surface mixed layer, suggests that eddies A and B were formed closer to the coast at an earlier time when the coastal warm, low-salinity water extended much deeper (Smith et al., 1991, have shown that the warm current is deepest in May/ June). We have tracked the positions of these three warm-core eddies using the 10-day maps of
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Fig. 2. Vertical temperature and salinity structure associated with the four TIP2000 cross-shelf sections. Top panel: stations 39–45 centred around 25.5 S; second panel: stations 45–51 centred at 27.5 S; third panel: stations 52–62 centred around 29.5 S; bottom panel: stations 62–73 centred around 31 S. Note the change of scale at 1000 m depth. The bold line indicates the bathymetry.
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Stations 39-75 25 23
20
24
Pot. Temp. C
25
15 27 26
10
28
5
27
Eddy A (stns 42-46) Eddy B (stns 60-64) Levitus Annual 110E 29S Reference Profile N Reference Profile S 0 34
34.2
34.4
34.6
34.8
35 Salinity
35.2
35.4
35.6
35.8
36
Fig. 3. Temperature and salinity profiles on potential density surfaces for the TIP2000 stations. Eddy A (stations 42–44, 46) are marked in green, eddy B (stations 60–64) are marked in red. Other stations on the northern transect (stations 39–41, 45) are marked with a blue dotted line; on the second transect (45, 47–51) with a blue dashed line; on the third transect (52–59) with a blue dasheddotted line; and the southern transect (65–72) with a blue solid line. Reference profile N (the mean of stations 47–49) is a black solid line; reference profile S (the mean of stations 56–58, 66–68) is a black dashed line. The Levitus annual mean profile at 110 E, 29 S is marked as a pink dotted line.
T/P+ERS altimeter data. Fig. 4a shows the central position of eddies A and B over a 6-month period starting in May 2000. This starting date is chosen as the period when the eddies clearly separated from the coastal Leeuwin Current. Eddy B split in two in July 2002—the paths of its offsprings B1 and B2—are also marked. After separation from the coastal current, the eddies drift west-northwestward in the deep Perth Basin,
in paths which essentially follow constant isopycnal contours in the upper layer (Fig. 4). As a reference, we have plotted the isopycnal surfaces at 150 m depth from Levitus 1998 climatological data WOA98 (http://www.nodc.noaa.gov/ OC5/woa98.html) to show the horizontal density gradients in our region; the 150 m surface is representative of the isopycnal gradients over the top 300 m.
Fig. 4. (a) Lagrangian track of the central position of eddies A and B during the period from May to December 2000, overlaid on isopycnal contours at 150 m depth (from Levitus, 1994). Bathymetry is shaded grey. (b) Hovmuller diagrams of positive SLA (in mm) during 2000 along the line corresponding to the 26.75 isopycnal contour at 150 m depth. (c) As for (b) but for the 26.4 isopycnal at 150 m depth.
(a)
(b)
(c)
22 Nov
22 Nov
25 Oct
25 Oct
27 Sep
27 Sep
30 Aug
30 Aug
2 Aug
2 Aug
5 Jul
5 Jul
7 Jun
7 Jun
10 May
10 May
12 Apr
12 Apr
15 Mar
15 Mar
16 Feb
16 Feb
19 Jan 90
50
95
100 105 Longitude
0
110
50
19 Jan 90
50
95
100 105 Longitude
0
110
50
115
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Eddy A tends to follow the 26.4 potential density contour at 150 m depth; eddies B and C tend to follow the 26.75 contour further south. Figs. 4b and c show Hovmuller diagrams during 2000 at positions defined by the isopycnal contours at 150 m depth, which ‘‘track’’ the eddies for a longer time than for the traditional Hovmuller diagrams on constant latitude lines (e.g., Birol and Morrow, 2001). Eddy A forms at the coast, south of the Abrolhos Islands near 28.5 S in early May 2000, then moves northwest around isopycnal 26.4 (Fig. 4c). In mid-July, the eddy separates in two, one-half continues westward to be observed by altimetry in late September at 25 S, 107 E (Fig. 1). The second weaker half travels at a slower speed, and is observed during TIP2000 and by altimetry at 111 E, 26 S. Eddy C appears to form in midSeptember, and indeed its low-density surface mixed layer is close to the Leeuwin Current’s surface density structure observed in late September. Eddy B is clearly formed at the coast around 5 May 2000. In mid-July, this eddy also splits in two—eddy B1 continues along the 26.75 isopycnal surface at near constant speed, and reaches 106 E, 29.5 S by the end of September (Fig. 4a). Eddy B2 remains stuck around 110–111 E for around 3 months, maintaining and even increasing its magnitude, then continues WNW in October. What causes the extreme intensity of eddy B? Close inspection of Fig. 1 in the region northeast of Eddy B reveals a band of higher sea level extending from Eddy B to the coast between 112 E and 113 E, 28.5 S, flanked by two cyclonic eddies centred at 27 S, 112 E and 30 S, 113 E. Station 48 touches the edge of the northern cyclonic eddy, stations 55 and 56 touch the edge of the southern cold eddy. The temperature and salinity structure along transect 2 shows a warm (T >20.5 C), lowsalinity (So35.15) surface current between stations 49 and 50 flowing to the southwest, which appears to feed into Eddy B (see also Fieux et al., 2003). High-resolution satellite SST maps (courtesy of CSIRO Division of Marine Research, Australia) indicate the presence of a coastal squirt at this location, with warm surface waters from the Leeuwin Current being directed off the shelf near 27 S (Fig. 5). The time series of satellite maps shows that the two cyclonic eddies developed just
after Eddy B split in two, and that the coastal squirt was present from late August to midOctober, which may explain the blocking of Eddy B and its subsequent intensification.
Fig. 5. Map of high-resolution satellite SST in the Leeuwin Current region. The location of TIP2000 stations is also marked. Arrowheads show the magnitude and direction of the altimetric surface geostrophic currents; their scale is given in the legend. (Map courtesy of CSIRO Marine Research, Australia.)
R. Morrow et al. / Deep-Sea Research II 50 (2003) 2229–2243
4. Available heat and salt anomalies What role do these warm-core eddies have in transferring heat and salt into the interior of the southern Indian Ocean? Fig. 1 clearly demonstrates that these warm-core eddies can perturb the temperature and salinity structure to depths of 1000–2500 m. However, we are really interested in whether these eddies introduce hydrographic property changes along isopycnals. In this case, they can then influence the level of mixing, watermass modification and heat and salt fluxes along isopycnals. To answer some of these questions we have estimated the heat and salt anomalies associated with these eddies based on the methods described in van Ballegooyen et al. (1994)—hereafter VB94. The available heat anomaly (AHA) in Joules was calculated for each discrete potential density layer, sy ; within the eddy Z 2p Z R ri cp hi ðrÞ½Ts ðrÞ AHAs ¼ 0
0
Ts ðrefÞr dr dy: Within each density layer, sy ; ri is the vertically averaged density (in kg m3); cp is the vertically averaged specific heat capacity (in J kg1 C1); hi ðrÞ is the thickness of the density layer within the eddy (in m); r is the radial distance from the centre of the eddy (in m); R is the radial extent of the eddy (in m); Ts ðrÞ is the vertically averaged temperature at the distance r from the eddy centre (in C); and Ts ðrefÞ is the vertically averaged temperature ( C) in the same potential density layer at a reference station in the ambient water outside the eddy. Similarly, the available salt anomaly (ASA) in kg is calculated for each potential density layer: Z 2p Z R ASAs ¼ 0:001ri hi ðrÞ½Ss ðrÞ 0
0
Ss ðrefÞr dr dy: The factor of 0.001 converts salinity in Practical Salinity units to the salinity fraction (mass of salt per unit mass of seawater). The vertical structure of the temperature and salinity anomalies ½T0 ðrÞ T0 ðrefÞ; S0 ðrÞ S0 ðrefÞ along the four sections is
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shown in Fig. 6, now plotted along potential density surfaces. For each section, the reference profile is chosen from stations outside the warmcore eddies with a near zero altimetric SLA; i.e. their steric height value is close to the 3-year mean steric height calculated from 1993 to 1995. Since there is a considerable north–south gradient in the mean density field, and eddy A appears to propagate along a more northern trajectory, we have chosen different reference profiles for eddies A and B. The mean reference profile for eddy A is chosen as the mean of stations 47–48–49 (solid line in Fig. 3). We note that station 48 just touches the edge of a coldcore eddy. For eddy B, we use the mean of stations 56-57-58-66-67-68 (dashed line in Fig. 3). Eddies A and B tend to be warmer and saltier than the mean reference profile on the same isopycnal level. Thus, increasing the integration depth increases the eddy AHA and ASA. In the upper levels, much of the increased AHA and ASA occurs in the thermocline waters (potential density layer from 25.2 to 26.5 kg m3). The surface cap water of these offshore eddies is only slightly fresher but slightly cooler than water at the same density in the inshore reference profile. Not surprisingly, the largest heat and salinity anomalies are associated with eddy B, with upper thermocline waters being some 0.5 C warmer and 0.15 saltier than the reference station. For intermediate and deep water with densities >27.1, there is large variability at all stations, and again eddies A and B are warmer and saltier than the reference profiles. The region is one of strong bathymetric gradients, which leads to a complicated recirculation and mixing of AAIW and UCDW (Fieux et al., 2003). Both our eddies lie at the most offshore point of our sections, and the increased mixing of the AAIW minimum may be due to the interaction of the deep circulation with bathymetry, rather than due to the eddies themselves. This is particularly apparent for the reference profile for eddy A. In the following analysis, we will analyse the eddy heat and salt anomalies over the entire vertical extent of the eddy, but also down to the 10 C and 8 C isotherm (or approximately the 26.7 and 26.9 isopycnals, respectively), which eliminates the influence of this IW/DW variability in the calculation.
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Fig. 6. Vertical structure of temperature anomalies ( C) (left panels) and salinity anomalies (right panels) along isopycnal levels for the four TIP2000 cross-shelf sections. Anomalies for the two northern sections are calculated relative to the mean of station 47-48-49 on section 2. Anomalies for the two southern sections are calculated relative to the mean of stations 56–58, 66–68. The bold line indicates the surface isopycnal value.
5. Eddy heat, salt and volume transports To estimate the total eddy heat content anomaly (in J) and total eddy salt content anomaly (in kg), the anomaly values shown in Fig. 6 were
horizontally integrated from the centre of each warm-core eddy to its outside circumference 2pR; then vertically integrated over the full vertical extent of the eddy, or to a given isopycnal/ temperature level.
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Table 2 Eddy heat and salt fluxes statistics for the two warm-core eddies observed during the TIP2000 campaign in September 2000 Eddy
Depth range
Available heat anomaly (1020 J)
A
Total: 0–1500 m Surface—10 C Surface—8 C
0.69 0.23 0.24
2.79 1.34 1.40
Surface water Thermocline Mode/intermediate water
0.08 0.12 0.34
0.65 0.62 1.13
Total: 0–2500 m Surface—10 C Surface—8 C
1.38 0.64 0.66
5.24 3.41 3.47
Surface water Thermocline Mode/intermediate water Deep water
0.19 0.41 0.52 0.28
1.3 2.07 1.53 0.39
Total: 0–1500 m Surface—10 C Surface—8 C
2.36
13.10
B
Agulhas Eddy A4
Available salt anomaly (1012 kg)
Volume flux (Sv)
Heat flux (103 PW)
Salt flux (106 kg s1)
51.6 55.3
1.64 1.75
2.19 0.67 0.74
0.09 0.04 0.04
123.1 132.3
3.90 4.20
4.39 1.97 2.04
0.17 0.11 0.11
7.5
0.42
33.0 39.4
1.05 1.25
Volume (1012 m3)
Note: Eddy fluxes are calculated for the full depth range of the eddy, and for water warmer than 10 C and 8 C, and over each of the water mass groups defined in Table 3. Agulhas Eddy A4 is from van Ballegooyen et al. (1994).
Table 3 Water mass properties in the Leeuwin Current region Water mass
Potential density range
Depth range (m)
Temperature range ( C)
Salinity range
Surface water Thermocline water Mode/intermediate water Upper circumpolar deep water
24.0–26.0 26.0–26.9 26.9–27.5 27.4–28.0
0–350 200–900 700–1450 1000–3000
18–22 7–18 5–8 2–5
35.0–35.8 34.5–35.8 34.3–34.5 34.4–34.7
The results of these calculations are shown in Table 2, where the two Leeuwin Current eddies are also compared with similar results from the Agulhas region—Agulhas Eddy A4 in VB94. To aid our discussion we also calculated the AHA and ASA contributions within each different water masses of the SE Indian Ocean. This allows us to identify which water masses are most strongly affected by the eddy heat and salt anomalies. The water-mass definitions we use (Table 3) are derived from Hanawa and Talley (2001) and Sloyan and
Rintoul (2001). We note that 8 C corresponds to a potential density level of around 26.9 kg m3, and waters warmer than 8 C include the surface down to the base of the Southeast Indian SubAntarctic Mode Water (SEISAMW) layer. The volume of water >8 C in eddy A is about 55 1012 m3; in eddy B 132 1012 m3, and in the Agulhas eddy A4 39 1012 m3. If we divide by a 1-year period, we can estimate the volume transport of water >8 C as 1.75 Sv for eddy A, 4.2 Sv for eddy B and 1.25 Sv for the Agulhas Eddy A4.
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The larger volume of low-density water >8 C in the Leeuwin Current eddies is expected since we are at lower latitudes in the subtropical gyre. The extremely large volume of water >8 C transported by eddy B accounts for its large steric height and its sea-level anomaly >80 cm. Eddy B was the largest warm-core eddy observed in this region during the 6-year period from 1995 to 2000 (Fang and Morrow, 2003). The AHA for eddy A for waters warmer than 8 C is 0.24 1020 J; not surprisingly, eddy B has a larger heat content (0.66 1020 J integrated down to the 8 C isotherm). Although the surface mixed layer was deeper for eddy B, there was little anomalous heat in this surface layer. As shown in the vertical sections (Figs. 3 and 6) most of the warming occurs in the thermocline and mode water layers. The AHA for the Agulhas eddy A4 is calculated over its entire vertical extent down to 1500 m or the 3 C isotherm, with a value of 2.4 1020 J. For the Leeuwin Current eddies, A and B, the AHA calculated over the entire vertical extent is 0.7 and 1.4 1020 J, respectively, although we have reservations about the stability of our reference profile at the isopycnal levels >27.0. So being conservative, the AHA of the Agulhas eddy A4 is a factor 4–10 larger than for the Leeuwin Current eddies. The ASA follows a similar trend. Eddies A and B have an ASA integrated down to 8 C of 1.4 1012 and 3.5 1012 kg, respectively. The ASA for the Agulhas eddy A4, integrated over its entire vertical extent, is 13.1 1012 kg. Once again, the Agulhas eddies carry 4–10 times more salt into the Atlantic Ocean than the Leeuwin Current eddies carry into the Indian Ocean. The eddy heat and salt fluxes are calculated by dividing the total AHA and ASA for each eddy by the period of 1 year. Again, the Leeuwin Current eddies A and B, integrated down to the 8 C isotherm, transport, respectively, 0.00074 and 0.002 PW of heat into the Indian Ocean. When integrated over the entire vertical extent of the eddy, they contribute 0.0022 and 0.0044 PW of heat, compared with 0.0075 PW carried by the Agulhas Eddy A4 into the Atlantic. Similarly, the salt transport for the two Leeuwin Current eddies integrated over their entire vertical extent is,
respectively, 0.8 and 1.7 105 kg s1, compared with 4.2 105 kg s1 for the Agulhas eddy A4. Although we have systematically compared our results with the one, well-sampled Agulhas eddy published by VB94, there exists a wide span of heat and salt flux estimates for Agulhas eddies, as shown in the comparative study by de Ruijter et al. (1999). Our Leeuwin Current estimates lie within the range of Agulhas eddy heat fluxes values (from 0.001 to 0.025 PW in de Ruijter et al., 1999) and of eddy salt flux estimates (ranging from 0.15 to 6.3 105 kg s1).
6. Discussion One of the surprising results is that these warmcore Leeuwin Current eddies, with their associated warm–fresh cap, actually carry a considerable salt flux into the subtropical gyre at deeper levels. How is this accomplished? The Leeuwin Current is fairly shallow, and its warm, low-salinity water extends over the upper 250 m and is essentially confined to the shelfbreak, with an equatorward, saline undercurrent below 300 m (Smith et al., 1991; Fieux et al., 2003). Offshore, in the upper layers, the salinity maximum due to the large evaporation in the subtropical latitudes. So the shallow fresh cap we see in the offshore eddies is indicative of the vertical structure at the coast. Within 50 km of the shelfbreak, the subsurface structure at 200–400 m depth is dominated by the downsloping isotherms and isohalines. So the waters between 200 and 400 m depth just off the shelfbreak are both warmer and saltier than at similar depths further offshore. Eddies formed in this region will maintain their warmer, saltier structure at depth once they are transported offshore. Lateral mixing at the eddy boundary may help make these waters warmer and saltier than the reference profile along the same isopycnal surfaces, thus contributing to the maximum anomalies detected in the thermocline offshore. Coastal observations between 400 and 800 m depth show an equatorward undercurrent that transports cooler and less saline water farther north (Smith et al., 1991). This appears to reduce our offshore eddy heat and salt anomalies in the lower thermocline density range
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(25.7oso26.2). Lateral mixing also may contribute to maintaining the large heat and salt anomalies along isopycnals in the mode water range (26.5oso26.9). How sensitive are these AHA and ASA calculations to our choice of reference station? Due to the limited ship time and consequently the zig-zag cruise tracks, the strongest warm-core eddies were observed at the deepest offshore positions, and our reference profiles are limited to the stations further inshore. This means that our two reference profiles may have a geographical bias. We have two ways to test the stability of our eddy flux calculations. Firstly, we have constructed a mean reference profile from the nine stations outside the warm and cold-core eddies, and in water deeper than 1500 m (stations 47-49, 56-58, 66-68). This mean reference profile lies between the two reference profiles in Fig. 3, and increases the heat and salt fluxes for eddy A, and decreases the heat and salt fluxes for eddy B. So, for example, the eddy heat flux for waters warmer than 8 C using the nine station mean reference profile is 0.001 PW for eddy A, and 0.0015 PW for eddy B, compared with 0.00075 PW for Eddy A and 0.002 PW for Eddy B using two different reference profiles. This shows how minor adjustments in the reference profile can lead to a 30% change in our eddy heat transport magnitude, which gives an idea of the sensitivity of our transport calculations. Secondly, although we have no other hydrographic measurements offshore during the same time period, we can compare our reference profiles to the annual mean profile from the Levitus 1998 climatology, at a point located between eddies A and B at 110 E, 29 S. The Levitus T=S profile (marked in pink dots in Fig. 3) represents the mean T=S properties over the last 50 years. Fig. 3 shows that eddies A and B are considerably cooler and fresher than the Levitus profile in the surface waters (so25:5) and in the mode and intermediate waters (26.5oso27.2), and warmer and saltier in the lower thermocline (25.5oso26.5). However, these anomalies are not restricted to the warmcore eddies—all of our station profiles east of 110 E show this offset from the Levitus profile. This pattern of large scale cooling and freshening of the mode and intermediate waters, compared
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with warming, and increasing salinity in the thermocline, has also been detected over the last 25 years in repeat hydrographic sections at 32 S (Wong et al., 1999). From this we conclude that the Levitus annual climatological profile is not representative of the background reference state of the ocean in 2000,2 and so will not provide an accurate reference profile for calculating mesoscale eddy fluxes. So although we have some reservations about our chosen reference profiles, we believe that they provide the best estimate for our background mean state in September 2000. How representative are these three warm-core eddies? Fang and Morrow (2003) have analysed long-lived, anti-cyclonic eddies from 6 year of altimetric data in the southeast Indian Ocean between 20 S and 32 S (long-lived eddies are those whose SLA is more than 10 cm amplitude after 12–18 m). They find that on average, six of these long-lived warm-core eddies are generated per year—although yearly averages vary from 3 to 9 depending on inter-annual variations in the eastern boundary density structure. Eddies A and C had average amplitudes of around 30–40 cm in SLA; eddy B with an SLA of 80 cm was one of the strongest eddies measured during 1995–2000. If we assume six warm-core eddies the size of eddy A are generated each year, a rough estimate of the annual eddy heat and salt fluxes for waters warmer than 8 C would be 0.004 PW and 0.24 106 kg s1, respectively. This is a factor of 10 less than the estimates by VB94, who show an annual heat and salt flux of Agulhas eddies into the Atlantic (calculated over the entire eddy depth range) of 0.045 PW and 2.5 106 kg s1, respectively. How do these estimates compare with the mean heat and salt fluxes in the southern Indian Ocean? East of 90 E, there is a net northward mass transport of B2–3 Sv in the upper levels (Sloyan and Rintoul, 2001) which is part of the wind-driven subtropical gyre circulation. East of 90 E, inverse model estimates from the surface to B1000 m depth (Ganachaud, pers. comm.) show an upper ocean poleward heat flux of 2
The Levitus seasonal mean profile and monthly averaged September profile also show the same cooler and fresher offset at depth.
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0.01370.008 PW at 18 S and 0.2670.26 PW at 32 S. These inverse models do not resolve the narrow Leeuwin Current transport or its eddies very well, and we note that much of the poleward heat transport at 32 S comes from the equatorward transport of cold mode waters between 500 and 1000 m. However, our warm-core eddy heat flux magnitudes are 20–30% of the inverse model mean poleward heat fluxes. In this study, we have not calculated a closed heat budget: we are comparing essentially westward eddy fluxes with the mean poleward heat flux, and we have not included the contribution from cold-core eddies generates at these latitudes. However, the magnitude of the warm-core eddy fluxes suggests that Leeuwin Current eddies could contribute significantly to the upper-ocean westward heat flux in this mid-latitude band south of 20 S. The same inverse model calculation east of 90 E shows an upper-ocean poleward salt flux of 0.275.4 106 kg s1 at 18 S and 13712 106 kg s1 at 32 S. Here, the magnitude of our Leeuwin Current eddy salt fluxes is similar to this annual mean salt budget, although with large uncertainties. Despite the lower salinity flux at the surface, all of the eddies contribute a net high salinity flux from their stronger contribution at deeper levels. Talley (1999) describes how low-salinity water is modified in this region: (i) fresher surface flow from the Indonesian Throughflow is modified and exported south of 32 S as salty thermocline water; and (ii) a horizontal inflow of fresher AAIW and SAMW in the east leaves saltier in the west. Our calculations suggest that deep-reaching long-lived warm-core eddies that mix down past the thermocline could contribute to increasing the salinity in both these water masses. What is the role of mixing for these warm-core eddies? Sloyan and Rintoul (2001) have shown that in the Indian Ocean region between 18 S and 32 S, air–sea fluxes transform 7.5 Sv of surface and thermocline waters to SAMW (>26.0). Interior mixing converts a further 572 Sv of SAMW to AAIW. As a consequence of diapycnal mixing, there is an associated downward flux of heat through these levels and a downward flux of salt that increases the salinity of SAMW and AAIW. There is also an upward diapycnal salt flux
into the thermocline, which enhances the subsurface salinity maximum (Sloyan and Rintoul, 2000). The large AHA and ASA of our warm-core eddies is of the right sign to aid in the warming and increasing the salinity of the thermocline and mode water layers, particularly when combined with the strong lateral mixing from these eddies. From our simple study we cannot yet determine whether the heat and salt mixes principally along isopycnals, or whether diapycnal mixing from the strong horizontal shear at the eddy edge could also contribute. The passage of warm, low-salinity water from the Indonesian Throughflow to the Agulhas Current is a key part of the ‘‘warm route’’ of the global overturning circulation (Gordon, 1986). Most of this water is transported westward by the South Equatorial Current around 10 S towards Madagascar. The contribution from Leeuwin Current eddies has never before been addressed—mainly because they are not well resolved in hydrographic sections or global, coarse-resolution climate models. This study provides a first estimate of the heat and salt flux contributions of Leeuwin Current eddies; albeit with significant error bars in the calculation. A more detailed budget based on both warm and cold-core eddies should now be determined from a global, eddy-resolving OGCM, or a combination of altimetry with more extensive hydrographic data.
Acknowledgements Thanks to the scientists, technicians and crew aboard the R.V. Marion Dufresne who helped us make the hydrography measurements. Special thanks to Michele Fieux, for her organisational capabilities as cruise leader, for her patience as we constantly changed the cruise plans to ‘‘track’’ the eddies, and for resolving the post-cruise calibration headaches. Altimeter data were generated by the CLS Space Oceanography Division as part of the European Unions’ Environment and Climate project AGORA (ENV4-CT9560113) and DUACS (ENV4-CT96-0357), with financial support from the Centre for Earth Observation
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Programme and Midi-Pyr!en!ees regional council. Additional inverse model calculations for the eastern Indian Ocean were kindly provided by Alexandre Ganachaud. We thank l’Institut Fran@ais pour la Recherche Technique et Polaire, l’Institut National des Sciences de l’Univers, and the Programme National d’Etude Du Climat for financing the TIP2000 cruise and this analysis.
References Batteen, M.L., Butler, C.L., 1998. Modeling studies of the Leeuwin Current off western and southern Australia. Journal of Physical Oceanography 28, 2199–2221. Birol, F., Morrow, R.A., 2001. Source of the baroclinic waves in the southeastern Indian Ocean. Journal of Geophysical Research 106 (C5), 9145–9160. Birol, F., Morrow, R.A., 2003. Source of the semiannual Rossby waves in the southeastern Indian Ocean, in preparation. De Ruijter, W.P.M., Biastoch, A., Drijfhout, S.S., Lutjeharms, J.R.E., Matano, R.P., Pichevin, T., van Leeuwen, P.J., Weijer, W., 1999. Indian6Atlantic interocean exchange: dynamics, estimation and impact. Journal of Geophysical Research 104, 20885–20910. Fang, F., Morrow, R., 2002. Evolution and structure of Leeuwin Current eddies in 1995–2000. Deep-Sea Research II, this issue. Fieux, M., Molcard, R., Morrow, R., 2003. Leeuwin Current and eddies off western Australia, in preparation. Fine, R.A., 1993. Circulation of Antarctic intermediate water in the south Indian Ocean. Deep-Sea Research 40, 2021–2042. Gordon, A.L., 1986. Interocean exchange of thermocline water. Journal of Geophysical Research 91, 5037–5046. Hanawa, K., Talley, L.D., 2001. Mode waters. In: Siedler et al. (Eds.), Ocean Circulation and Climate. Academic Press, New York. Hirst, A.C., Godfrey, J.S., 1993. The role of the Indonesian Throughflow in a global ocean GCM. Journal of Physical Oceanography 23, 1057–1086. Josey, S.A., Kent, E.C., Taylor, P.K., 1999. New insights into the ocean heat budget closure problem from analysis of the SOC air–sea flux climatology. Journal of Climate 12, 2856–2880. Karstensen, J., Tomczak, M., 1997. Ventilation processes and water mass ages in the thermocline of the southeast Indian Ocean. Geophysical Research Letters 24, 2777–2780. Le Traon, P.-Y., Nadal, F., Ducet, N., 1997. An improved mapping method of multi-satellite altimeter data. Journal of Atmospheric and Oceanic Technology 15, 522–534. McCartney, M.S., 1977. Subantarctic mode water, in: a voyage of discovery. Deep-Sea Research 24, 103–119.
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Morrow, R.A., Birol, F., 1998. Variability in the south-east Indian Ocean from altimetry: forcing mechanisms for the Leeuwin Current. Journal of Geophysical Research 103, 18529–18544. Pearce, A.F., Griffiths, R.W., 1991. The mesoscale structure of the Leeuwin Current: a comparison of laboratory models and satellite imagery. Journal of Geophysical Research 96, 16739–16757. Potemra, J.T., 2001. Contribution of equatorial Pacific winds to southern tropical Indian Ocean Rossby waves. Journal of Geophysical Research 106 (C2), 2407–2422. Rochford, D.J., 1969. Seasonal variations in the Indian Ocean along 110 E, I: hydrological structure of the upper 500 m. Australian Journal of Marine and Freshwater Research 20, 1–50. Sloyan, B.M., Rintoul, S.R., 2000. Estimates of area-averaged diapycnal fluxes from basin-scale budgets. Journal of Physical Oceanography 30, 2320–2341. Sloyan, B.M., Rintoul, S.R., 2001. Circulation, renewal and modification of Antarctic Mode and Intermediate Water. Journal of Physical Oceanography 31, 1005–1030. Smith, R.L., Huyer, A., Godfrey, J.S., Church, J.A., 1991. The Leeuwin Current off western Australia 1986–87. Journal of Physical Oceanography 21, 323–345. Talley, L.D., 1999. Some aspects of ocean heat transport by the shallow, intermediate and deep overturning circulations. In: Clark, Webb, Keigwin (Eds.), Mechanisms of Global Climate Change at Millennial Time Scales. AGU Geophysical Monograph 112. Tomczak, M., Large, D.G.B., 1989. Optimal multiparameter analysis of mixing in the thermocline in the eastern Indian Ocean. Journal of Geophysical Research 94, 16141–16149. Tomczak, M., Godfrey, J.S., 1994. Regional Oceanography: an Introduction. Pergamon Press, New York. Toole, J.M., Warren, B.A., 1993. A hydrographic section across the subtropical south Indian Ocean. Deep-Sea Research 40 (10), 1973–2019. Van Ballegooyen, R.C., Grundlingh, M.L., Lutjeharms, J.R.E., 1994. Eddy fluxes of heat and salt from the southwest Indian Ocean into the southeast Atlantic Ocean. Journal of Geophysical Research 99, 14053–14070. Warren, B.A., 1981. Transindian hydrographic section at Lat. 18 S: property distributions and circulation in the south Indian Ocean. Deep-Sea Research A 28, 759–788. Wong, A.P.S., Bindoff, N.L., Church, J.A., 1999. Large-scale freshening of intermediate waters in the Pacific and Indian oceans. Nature 400, 440–443. Wyrtki, K., 1971. Oceanographic Atlas of the International Indian Ocean Expedition. Natural Science Foundation, Washington, DC, 531pp. You, Y., Tomczak, M., 1993. Thermocline circulation and ventilation of the Indian Ocean derived from water mass analysis. Deep-Sea Research 40, 13–56.
Anatomy of three warm-core Leeuwin Current eddies Rosemary Morrowa,*, Fangxin Fanga,1, Michele Fieuxb, Robert Molcardb b
a LEGOS 18, av. Edouard Belin, 31401 Toulouse Cedex 4, France LODYC, Universit!e Pierre et Marie Curie, case 100, 4 place Jussieu, 75252 Paris Cedex 05, France
Abstract The vertical structures of three warm-core Leeuwin Current eddies are described, as well as their temporal evolution determined from altimetric data. These warm-core eddies have large positive sea-level anomalies, although their surface temperature fields are not necessarily warmer than the ambient waters. These eddies were observed some 350 km offshore from the Leeuwin Current in September 2000, around 25–30 S, and their vertical structure and temporal evolution show they were formed at the coast in May, when the Leeuwin Current is strongest. After separation from the current, the warm-core eddies drifted WNW following isopycnal contours, and were strongly steered by bathymetry. These eddies penetrated to at least 1500 m depth; the strongest eddy, Eddy B, influenced isopycnals to 2500 m. Eddy B appeared to be trapped and intensified by a Leeuwin Current squirt that directed warm, low-salinity water offshore from August to October. The average annual eddy heat and salt fluxes for waters warmer than 8 C are estimated at 0.004 PW and 3 105 kg s1, respectively. The magnitude of these eddy fluxes is 20–30% of the annual mean poleward heat and salt flux in the eastern Indian Ocean, and is within the range of values observed for Agulhas eddies entering the South Atlantic. r 2003 Elsevier Science Ltd. All rights reserved.
1. Introduction Most eastern boundary regions of the ocean are synonymous with cold upwelling, equatorward surface flow, and a large heat sink from the atmosphere (Tomczak and Godfrey, 1994). The eastern boundary of the southern Indian Ocean is the opposite: the coastal current, the Leeuwin Current, flows poleward against the prevailing equatorward winds with no coastal upwelling. More importantly, it is the only eastern boundary *Corresponding author. E-mail address: [email protected] (R. Morrow). 1 Current address: Earth Science and Engineering, Imperial College of Science, Technology and Medicine, RSM Building, Prince Consort Road, London SW7 2BP, UK.
region with a large heat transfer from the ocean to the atmosphere, with values equivalent to a western boundary current (Josey et al., 1999). The ocean–atmosphere heat flux is largest along the southwest Australian coastline where the annual heat loss exceeds 50 W m2. This coastal maxima is associated with the strong poleward Leeuwin Current, which transports around 5 Sv of warm, low-salinity waters southward from the Indonesian Throughflow region (Smith et al., 1991). However, the region of ocean heat loss is not limited to the coastal Leeuwin Current and its meanders. A large region of ocean heat loss of up to 25 W m2 extends westward from the coast, reaching 80 E at 20 S, and 95 E at 30 S (Josey et al., 1999). The spatial distribution of this heat
0967-0645/03/$ - see front matter r 2003 Elsevier Science Ltd. All rights reserved. doi:10.1016/S0967-0645(03)00054-7
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loss resembles the beta-refraction pattern associated with Rossby waves. Large sea-level anomalies propagating at Rossby wave speeds are well documented in the 20–35 S band, from altimetric data (Morrow and Birol, 1998; Potemra, 2001) and with greatly reduced amplitude in models (Birol and Morrow, 2001; Hirst and Godfrey, 1993). The propagating anomalies are maximal at annual and semi-annual frequencies, with a strong inter-annual modulation (Morrow and Birol, 1998). Wind-forcing plays a minor role, both offshore and at the coast (Birol and Morrow, 2001). Instead, the positive (warm-core) anomalies appear to detach from the coast between 20 S and 35 S in phase with downwelling coastal propagating waves, which also occur at annual and semiannual frequencies (Potemra, 2001; Birol and Morrow, 2003). Maximum warm-core anomalies occur in May and November during the monsoon transition periods, and may be forced by Pacific wind anomalies that propagate as Kelvin waves through the Indonesian Seas and Timor Passage (Potemra, 2001). Many questions remain about these westwardpropagating warm-core anomalies. Are these features Rossby waves or transient eddies propagating at Rossby wave speeds? What is their vertical structure? How are their propagation paths modified by the bathymetry and background stratification? How do they evolve in time and decay? Most importantly, what is their role in transferring heat and other tracers into the interior of the southern Indian Ocean, and in mixing and modifying the ambient water masses? To help resolve some of these questions, the Transport Indo-Pacific (TIP2000) research cruise was organised in September 2000. One of the aims was to measure the vertical hydrographic structure associated with a few of these propagating anomalies. Eddies are fully detached from the Leeuwin Current meanders after around 90 days (Pearce and Griffiths, 1991), so the September cruise allowed us to measure detached anomalies from the strong May forcing. In this paper, we examine the vertical structure, water masses, generation process and propagation of three warm-core eddies measured during TIP2000. We also estimate the heat and salt fluxes associated
with these three eddies. A companion paper includes a description of the transport and watermass properties of the Leeuwin Current and its eddies during September 2000 by Fieux et al. (2003). Fang and Morrow (2003) use altimeter data to investigate the temporal and spatial evolution of ‘‘long-lasting’’ warm-core eddies generated between 1995 and 2000, and provide details of their propagation pathways and decay rates.
2. Vertical structure of the warm-core eddies The TIP2000 cruise through the Leeuwin Current eddies took place during 1 week from 24 September 2000 to 1 October 2000. The cruise track is shown in Fig. 1, superimposed on a weekly map of positive (warm-core) sea-level anomalies (SLAs) from altimetry. The altimetric SLAs are a combination of Topex/Poseidon (T/P) and ERS2 data mapped onto a regular 0.25 grid, using an improved space/time objective interpolation scheme that eliminates long wavelength correlated noise (Le Traon et al., 1997). The SLA maps are produced by the CLS Space Oceanography Division. Fig. 1 shows that the TIP2000 cruise crossed three warm-core eddies: eddy A at 26 S, 111 E; eddy B at 30 S, 110.5 S; and eddy C at 32 S, 114 E. Their eddy statistics are given in Table 1. Also marked in Fig. 1 is an approximate westward limit for a Rossby wave forced from the eastern boundary in May 2000. The western limit of the Rossby wave position is estimated from analyses of Hovmuller diagrams at various latitudes (not shown) (see Morrow and Birol, 1998; Birol and Morrow, 2001, for Hovmuller diagrams from T/P at these latitudes from 1992 to 1999). Much of this signal is generated by annual and semi-annual coastal Kelvin waves that propagate poleward along the eastern boundary (Potemra, 2001); by linear theory the semi-annual signal can radiate offshore as Rossby waves up to the critical latitude of 25 S. Indeed, north of 25 S, these positive anomalies have longer wavelengths and resemble a coherent Rossby wave train forced from the eastern boundary. South of 25 S, the warm
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(a) 800 16
18
600
20 400
Latitude
22
24
200
26 0 28
30 200 32 400
34 100
102
104
106
108
110
112
114
116
Longitude
Fig. 1. Topex/Poseidon+ERS2 sea-level anomalies (in mm) for the cruise period 25 September–1 October 2000, with the TIP2000 cruise track (solid line) overlaid. Station numbers mark the ends of each section (39, 45, 52, 62, 72). The three warm-core eddies (A, B, C) mentioned in the text are also indicated. A dashed line shows the westward limit of Rossby waves forced from the eastern boundary in May 2000. The 3000 and 4000 m depths are contoured.
Table 1 Statistics for the three warm-core eddies observed during the TIP2000 campaign in late September 2000 Eddy
Centre
Diameter (km)
SLA (cm)
Depth (m)
Drift (cm s1)
Reference station #
A B C
26 S 111 E 30 S 110 E 32 S 114 E
180 350 160
40 80 50
750 2500 1500
6 WNW 4 WNW 1 WSW
47–48–49 56–58, 66–68 56-58, 66-68
anomalies have increasingly smaller radius, appear as more isolated features, and often ‘‘stray’’ from the Rossby wave path (Fig. 1), being locally steered or decelerated by bathymetric features and by non-linear interactions with the stratified flow (Fang and Morrow, 2003). These anomalies appear more like individual eddies, and most likely result from instabilities of the coastal Leeuwin Current at these latitudes (Pearce and Griffiths, 1991; Batteen and Butler, 1998). The arrival of a
semi-annual downwelling coastal Kelvin wave around 25 S may change the coastal density structure, leading to meanders and detached eddies with a large positive sea-level anomaly (Birol and Morrow, 2003). The vertical temperature and salinity structure associated with these warm-core eddies and their surrounding water is shown for the four TIP2000 transects between 25 S and 32 S (Fig. 2), and on a T2S diagram (Fig. 3). Stations along section 1
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were full depth to 20–100 m from the bottom; sections 2 and 4 were to 2000 m depth or 20–100 m from the bottom, whichever came first. Stations along section 3 were full depth, except for station 57 to 5150 db, station 58 to 4000 db, and stations 59–62 to 3000 db. Here, we present the water-mass characteristics in terms of their T2S structure; the full water-mass analysis, including other tracers, is given by Fieux et al. (2003). Close to the coast in Fig. 2, the poleward-flowing Leeuwin Current is evident from the downsloping isotherms and isohalines, and the characteristics of warm (T >20 C), low-salinity water (So35.4, even 35.0 around station 40). The equatorward undercurrent at depths >400 m is also present on the slope, with upsloping isotherms and isopycnals (Smith et al., 1991). The ambient water (outside the warm-core eddies) has a vertical structure similar to that defined by Warren (1981) from a hydrography section at 18 S. The surface layers are influenced by the Indonesian Throughflow Water from the north, with warm (T > 19 C) water and salinity o35.3 (Tomczak and Large, 1989) depending on the latitude; the depth of this low-density surface layer decreases from 140 m around 25 S to 70 m around 30 S. Directly under the surface layer is a 125 m thick layer of salinity maximum water (S >35.6); this water is formed from the strong evaporation in the subtropical gyre between 25 S and 35 S (Warren, 1981). Below, the linear decrease in temperature and salinity in the range 9–14 C marks Indian Central Water (Rochford, 1969; Wyrtki, 1971; You and Tomczak, 1993). This in turn overlies a thermostad of 8–9 C water that is more than 150 m thick around 600 m depth, which corresponds to Subantarctic Mode Water formed in the southeast Indian Ocean (McCartney, 1977; Karstensen and Tomczak, 1997). Antarctic Intermediate Water, with a salinity minimum o34.5 (Wyrtki, 1971; Toole and Warren, 1993; Fine, 1993), is evident from 800 to 1000 m. The salinity minimum is more clearly defined in the southern-most sections. The northern sections and those close to the coast show evidence of mixing, which may indicate some recirculation around the eastern boundary of the Perth Basin. Below 1500 m, the cold (To4 C),
saline (S >34.6) water corresponds to Indian or Circumpolar Deep Water. The three warm-core eddies have a similar vertical structure as the ambient water, but with the isopycnals shifted downwards and capped by warm, low-salinity surface water. Eddies A and B, corresponding to the stations further offshore, tend to be warmer and saltier in the thermocline and at the depths of the AAIW salinity minimum. The core of eddy A was measured at stations 43–44–46; 350 km offshore (Table 1). A lowdensity cap 30 m deep exists at station 45, with T >22 C, and So34.8; this relates to surface waters drawn in from the north around the edge of eddy A. In the centre, at station 43, the surface temperature is lower and the surface salinity higher but the warm–fresh mixed layer reaches 200 m. These temperature values are close to those observed in the coastal current in September 2000, but the low-density mixed layer is deeper than at the coast. Eddy A perturbs the ambient isothermals and isohalines to at least 850 m depth. Eddy C was measured at stations 69–70 on section 4, 150 km offshore. Its low-density cap corresponds to a 130 m thick mixed layer, and Eddy C influences the vertical density structure to B1100 m depth. Eddy B was the strongest of the three, observed at stations 60–64 at the limit of transects 3 and 4. Its low-density core had a surface mixed layer of 20 C, 35.3 water extending to 240 m, which perturbed the underlying density field to 2500 m depth. As for eddy A, this warm, low-salinity core is more deep reaching than the vertical structure of the Leeuwin Current in September.
3. Temporal evolution of the warm-core eddies The vertical structure of the offshore warm-core eddies, and particularly the depth of their lowdensity surface mixed layer, suggests that eddies A and B were formed closer to the coast at an earlier time when the coastal warm, low-salinity water extended much deeper (Smith et al., 1991, have shown that the warm current is deepest in May/ June). We have tracked the positions of these three warm-core eddies using the 10-day maps of
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Fig. 2. Vertical temperature and salinity structure associated with the four TIP2000 cross-shelf sections. Top panel: stations 39–45 centred around 25.5 S; second panel: stations 45–51 centred at 27.5 S; third panel: stations 52–62 centred around 29.5 S; bottom panel: stations 62–73 centred around 31 S. Note the change of scale at 1000 m depth. The bold line indicates the bathymetry.
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Stations 39-75 25 23
20
24
Pot. Temp. C
25
15 27 26
10
28
5
27
Eddy A (stns 42-46) Eddy B (stns 60-64) Levitus Annual 110E 29S Reference Profile N Reference Profile S 0 34
34.2
34.4
34.6
34.8
35 Salinity
35.2
35.4
35.6
35.8
36
Fig. 3. Temperature and salinity profiles on potential density surfaces for the TIP2000 stations. Eddy A (stations 42–44, 46) are marked in green, eddy B (stations 60–64) are marked in red. Other stations on the northern transect (stations 39–41, 45) are marked with a blue dotted line; on the second transect (45, 47–51) with a blue dashed line; on the third transect (52–59) with a blue dasheddotted line; and the southern transect (65–72) with a blue solid line. Reference profile N (the mean of stations 47–49) is a black solid line; reference profile S (the mean of stations 56–58, 66–68) is a black dashed line. The Levitus annual mean profile at 110 E, 29 S is marked as a pink dotted line.
T/P+ERS altimeter data. Fig. 4a shows the central position of eddies A and B over a 6-month period starting in May 2000. This starting date is chosen as the period when the eddies clearly separated from the coastal Leeuwin Current. Eddy B split in two in July 2002—the paths of its offsprings B1 and B2—are also marked. After separation from the coastal current, the eddies drift west-northwestward in the deep Perth Basin,
in paths which essentially follow constant isopycnal contours in the upper layer (Fig. 4). As a reference, we have plotted the isopycnal surfaces at 150 m depth from Levitus 1998 climatological data WOA98 (http://www.nodc.noaa.gov/ OC5/woa98.html) to show the horizontal density gradients in our region; the 150 m surface is representative of the isopycnal gradients over the top 300 m.
Fig. 4. (a) Lagrangian track of the central position of eddies A and B during the period from May to December 2000, overlaid on isopycnal contours at 150 m depth (from Levitus, 1994). Bathymetry is shaded grey. (b) Hovmuller diagrams of positive SLA (in mm) during 2000 along the line corresponding to the 26.75 isopycnal contour at 150 m depth. (c) As for (b) but for the 26.4 isopycnal at 150 m depth.
(a)
(b)
(c)
22 Nov
22 Nov
25 Oct
25 Oct
27 Sep
27 Sep
30 Aug
30 Aug
2 Aug
2 Aug
5 Jul
5 Jul
7 Jun
7 Jun
10 May
10 May
12 Apr
12 Apr
15 Mar
15 Mar
16 Feb
16 Feb
19 Jan 90
50
95
100 105 Longitude
0
110
50
19 Jan 90
50
95
100 105 Longitude
0
110
50
115
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Eddy A tends to follow the 26.4 potential density contour at 150 m depth; eddies B and C tend to follow the 26.75 contour further south. Figs. 4b and c show Hovmuller diagrams during 2000 at positions defined by the isopycnal contours at 150 m depth, which ‘‘track’’ the eddies for a longer time than for the traditional Hovmuller diagrams on constant latitude lines (e.g., Birol and Morrow, 2001). Eddy A forms at the coast, south of the Abrolhos Islands near 28.5 S in early May 2000, then moves northwest around isopycnal 26.4 (Fig. 4c). In mid-July, the eddy separates in two, one-half continues westward to be observed by altimetry in late September at 25 S, 107 E (Fig. 1). The second weaker half travels at a slower speed, and is observed during TIP2000 and by altimetry at 111 E, 26 S. Eddy C appears to form in midSeptember, and indeed its low-density surface mixed layer is close to the Leeuwin Current’s surface density structure observed in late September. Eddy B is clearly formed at the coast around 5 May 2000. In mid-July, this eddy also splits in two—eddy B1 continues along the 26.75 isopycnal surface at near constant speed, and reaches 106 E, 29.5 S by the end of September (Fig. 4a). Eddy B2 remains stuck around 110–111 E for around 3 months, maintaining and even increasing its magnitude, then continues WNW in October. What causes the extreme intensity of eddy B? Close inspection of Fig. 1 in the region northeast of Eddy B reveals a band of higher sea level extending from Eddy B to the coast between 112 E and 113 E, 28.5 S, flanked by two cyclonic eddies centred at 27 S, 112 E and 30 S, 113 E. Station 48 touches the edge of the northern cyclonic eddy, stations 55 and 56 touch the edge of the southern cold eddy. The temperature and salinity structure along transect 2 shows a warm (T >20.5 C), lowsalinity (So35.15) surface current between stations 49 and 50 flowing to the southwest, which appears to feed into Eddy B (see also Fieux et al., 2003). High-resolution satellite SST maps (courtesy of CSIRO Division of Marine Research, Australia) indicate the presence of a coastal squirt at this location, with warm surface waters from the Leeuwin Current being directed off the shelf near 27 S (Fig. 5). The time series of satellite maps shows that the two cyclonic eddies developed just
after Eddy B split in two, and that the coastal squirt was present from late August to midOctober, which may explain the blocking of Eddy B and its subsequent intensification.
Fig. 5. Map of high-resolution satellite SST in the Leeuwin Current region. The location of TIP2000 stations is also marked. Arrowheads show the magnitude and direction of the altimetric surface geostrophic currents; their scale is given in the legend. (Map courtesy of CSIRO Marine Research, Australia.)
R. Morrow et al. / Deep-Sea Research II 50 (2003) 2229–2243
4. Available heat and salt anomalies What role do these warm-core eddies have in transferring heat and salt into the interior of the southern Indian Ocean? Fig. 1 clearly demonstrates that these warm-core eddies can perturb the temperature and salinity structure to depths of 1000–2500 m. However, we are really interested in whether these eddies introduce hydrographic property changes along isopycnals. In this case, they can then influence the level of mixing, watermass modification and heat and salt fluxes along isopycnals. To answer some of these questions we have estimated the heat and salt anomalies associated with these eddies based on the methods described in van Ballegooyen et al. (1994)—hereafter VB94. The available heat anomaly (AHA) in Joules was calculated for each discrete potential density layer, sy ; within the eddy Z 2p Z R ri cp hi ðrÞ½Ts ðrÞ AHAs ¼ 0
0
Ts ðrefÞr dr dy: Within each density layer, sy ; ri is the vertically averaged density (in kg m3); cp is the vertically averaged specific heat capacity (in J kg1 C1); hi ðrÞ is the thickness of the density layer within the eddy (in m); r is the radial distance from the centre of the eddy (in m); R is the radial extent of the eddy (in m); Ts ðrÞ is the vertically averaged temperature at the distance r from the eddy centre (in C); and Ts ðrefÞ is the vertically averaged temperature ( C) in the same potential density layer at a reference station in the ambient water outside the eddy. Similarly, the available salt anomaly (ASA) in kg is calculated for each potential density layer: Z 2p Z R ASAs ¼ 0:001ri hi ðrÞ½Ss ðrÞ 0
0
Ss ðrefÞr dr dy: The factor of 0.001 converts salinity in Practical Salinity units to the salinity fraction (mass of salt per unit mass of seawater). The vertical structure of the temperature and salinity anomalies ½T0 ðrÞ T0 ðrefÞ; S0 ðrÞ S0 ðrefÞ along the four sections is
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shown in Fig. 6, now plotted along potential density surfaces. For each section, the reference profile is chosen from stations outside the warmcore eddies with a near zero altimetric SLA; i.e. their steric height value is close to the 3-year mean steric height calculated from 1993 to 1995. Since there is a considerable north–south gradient in the mean density field, and eddy A appears to propagate along a more northern trajectory, we have chosen different reference profiles for eddies A and B. The mean reference profile for eddy A is chosen as the mean of stations 47–48–49 (solid line in Fig. 3). We note that station 48 just touches the edge of a coldcore eddy. For eddy B, we use the mean of stations 56-57-58-66-67-68 (dashed line in Fig. 3). Eddies A and B tend to be warmer and saltier than the mean reference profile on the same isopycnal level. Thus, increasing the integration depth increases the eddy AHA and ASA. In the upper levels, much of the increased AHA and ASA occurs in the thermocline waters (potential density layer from 25.2 to 26.5 kg m3). The surface cap water of these offshore eddies is only slightly fresher but slightly cooler than water at the same density in the inshore reference profile. Not surprisingly, the largest heat and salinity anomalies are associated with eddy B, with upper thermocline waters being some 0.5 C warmer and 0.15 saltier than the reference station. For intermediate and deep water with densities >27.1, there is large variability at all stations, and again eddies A and B are warmer and saltier than the reference profiles. The region is one of strong bathymetric gradients, which leads to a complicated recirculation and mixing of AAIW and UCDW (Fieux et al., 2003). Both our eddies lie at the most offshore point of our sections, and the increased mixing of the AAIW minimum may be due to the interaction of the deep circulation with bathymetry, rather than due to the eddies themselves. This is particularly apparent for the reference profile for eddy A. In the following analysis, we will analyse the eddy heat and salt anomalies over the entire vertical extent of the eddy, but also down to the 10 C and 8 C isotherm (or approximately the 26.7 and 26.9 isopycnals, respectively), which eliminates the influence of this IW/DW variability in the calculation.
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Fig. 6. Vertical structure of temperature anomalies ( C) (left panels) and salinity anomalies (right panels) along isopycnal levels for the four TIP2000 cross-shelf sections. Anomalies for the two northern sections are calculated relative to the mean of station 47-48-49 on section 2. Anomalies for the two southern sections are calculated relative to the mean of stations 56–58, 66–68. The bold line indicates the surface isopycnal value.
5. Eddy heat, salt and volume transports To estimate the total eddy heat content anomaly (in J) and total eddy salt content anomaly (in kg), the anomaly values shown in Fig. 6 were
horizontally integrated from the centre of each warm-core eddy to its outside circumference 2pR; then vertically integrated over the full vertical extent of the eddy, or to a given isopycnal/ temperature level.
R. Morrow et al. / Deep-Sea Research II 50 (2003) 2229–2243
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Table 2 Eddy heat and salt fluxes statistics for the two warm-core eddies observed during the TIP2000 campaign in September 2000 Eddy
Depth range
Available heat anomaly (1020 J)
A
Total: 0–1500 m Surface—10 C Surface—8 C
0.69 0.23 0.24
2.79 1.34 1.40
Surface water Thermocline Mode/intermediate water
0.08 0.12 0.34
0.65 0.62 1.13
Total: 0–2500 m Surface—10 C Surface—8 C
1.38 0.64 0.66
5.24 3.41 3.47
Surface water Thermocline Mode/intermediate water Deep water
0.19 0.41 0.52 0.28
1.3 2.07 1.53 0.39
Total: 0–1500 m Surface—10 C Surface—8 C
2.36
13.10
B
Agulhas Eddy A4
Available salt anomaly (1012 kg)
Volume flux (Sv)
Heat flux (103 PW)
Salt flux (106 kg s1)
51.6 55.3
1.64 1.75
2.19 0.67 0.74
0.09 0.04 0.04
123.1 132.3
3.90 4.20
4.39 1.97 2.04
0.17 0.11 0.11
7.5
0.42
33.0 39.4
1.05 1.25
Volume (1012 m3)
Note: Eddy fluxes are calculated for the full depth range of the eddy, and for water warmer than 10 C and 8 C, and over each of the water mass groups defined in Table 3. Agulhas Eddy A4 is from van Ballegooyen et al. (1994).
Table 3 Water mass properties in the Leeuwin Current region Water mass
Potential density range
Depth range (m)
Temperature range ( C)
Salinity range
Surface water Thermocline water Mode/intermediate water Upper circumpolar deep water
24.0–26.0 26.0–26.9 26.9–27.5 27.4–28.0
0–350 200–900 700–1450 1000–3000
18–22 7–18 5–8 2–5
35.0–35.8 34.5–35.8 34.3–34.5 34.4–34.7
The results of these calculations are shown in Table 2, where the two Leeuwin Current eddies are also compared with similar results from the Agulhas region—Agulhas Eddy A4 in VB94. To aid our discussion we also calculated the AHA and ASA contributions within each different water masses of the SE Indian Ocean. This allows us to identify which water masses are most strongly affected by the eddy heat and salt anomalies. The water-mass definitions we use (Table 3) are derived from Hanawa and Talley (2001) and Sloyan and
Rintoul (2001). We note that 8 C corresponds to a potential density level of around 26.9 kg m3, and waters warmer than 8 C include the surface down to the base of the Southeast Indian SubAntarctic Mode Water (SEISAMW) layer. The volume of water >8 C in eddy A is about 55 1012 m3; in eddy B 132 1012 m3, and in the Agulhas eddy A4 39 1012 m3. If we divide by a 1-year period, we can estimate the volume transport of water >8 C as 1.75 Sv for eddy A, 4.2 Sv for eddy B and 1.25 Sv for the Agulhas Eddy A4.
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The larger volume of low-density water >8 C in the Leeuwin Current eddies is expected since we are at lower latitudes in the subtropical gyre. The extremely large volume of water >8 C transported by eddy B accounts for its large steric height and its sea-level anomaly >80 cm. Eddy B was the largest warm-core eddy observed in this region during the 6-year period from 1995 to 2000 (Fang and Morrow, 2003). The AHA for eddy A for waters warmer than 8 C is 0.24 1020 J; not surprisingly, eddy B has a larger heat content (0.66 1020 J integrated down to the 8 C isotherm). Although the surface mixed layer was deeper for eddy B, there was little anomalous heat in this surface layer. As shown in the vertical sections (Figs. 3 and 6) most of the warming occurs in the thermocline and mode water layers. The AHA for the Agulhas eddy A4 is calculated over its entire vertical extent down to 1500 m or the 3 C isotherm, with a value of 2.4 1020 J. For the Leeuwin Current eddies, A and B, the AHA calculated over the entire vertical extent is 0.7 and 1.4 1020 J, respectively, although we have reservations about the stability of our reference profile at the isopycnal levels >27.0. So being conservative, the AHA of the Agulhas eddy A4 is a factor 4–10 larger than for the Leeuwin Current eddies. The ASA follows a similar trend. Eddies A and B have an ASA integrated down to 8 C of 1.4 1012 and 3.5 1012 kg, respectively. The ASA for the Agulhas eddy A4, integrated over its entire vertical extent, is 13.1 1012 kg. Once again, the Agulhas eddies carry 4–10 times more salt into the Atlantic Ocean than the Leeuwin Current eddies carry into the Indian Ocean. The eddy heat and salt fluxes are calculated by dividing the total AHA and ASA for each eddy by the period of 1 year. Again, the Leeuwin Current eddies A and B, integrated down to the 8 C isotherm, transport, respectively, 0.00074 and 0.002 PW of heat into the Indian Ocean. When integrated over the entire vertical extent of the eddy, they contribute 0.0022 and 0.0044 PW of heat, compared with 0.0075 PW carried by the Agulhas Eddy A4 into the Atlantic. Similarly, the salt transport for the two Leeuwin Current eddies integrated over their entire vertical extent is,
respectively, 0.8 and 1.7 105 kg s1, compared with 4.2 105 kg s1 for the Agulhas eddy A4. Although we have systematically compared our results with the one, well-sampled Agulhas eddy published by VB94, there exists a wide span of heat and salt flux estimates for Agulhas eddies, as shown in the comparative study by de Ruijter et al. (1999). Our Leeuwin Current estimates lie within the range of Agulhas eddy heat fluxes values (from 0.001 to 0.025 PW in de Ruijter et al., 1999) and of eddy salt flux estimates (ranging from 0.15 to 6.3 105 kg s1).
6. Discussion One of the surprising results is that these warmcore Leeuwin Current eddies, with their associated warm–fresh cap, actually carry a considerable salt flux into the subtropical gyre at deeper levels. How is this accomplished? The Leeuwin Current is fairly shallow, and its warm, low-salinity water extends over the upper 250 m and is essentially confined to the shelfbreak, with an equatorward, saline undercurrent below 300 m (Smith et al., 1991; Fieux et al., 2003). Offshore, in the upper layers, the salinity maximum due to the large evaporation in the subtropical latitudes. So the shallow fresh cap we see in the offshore eddies is indicative of the vertical structure at the coast. Within 50 km of the shelfbreak, the subsurface structure at 200–400 m depth is dominated by the downsloping isotherms and isohalines. So the waters between 200 and 400 m depth just off the shelfbreak are both warmer and saltier than at similar depths further offshore. Eddies formed in this region will maintain their warmer, saltier structure at depth once they are transported offshore. Lateral mixing at the eddy boundary may help make these waters warmer and saltier than the reference profile along the same isopycnal surfaces, thus contributing to the maximum anomalies detected in the thermocline offshore. Coastal observations between 400 and 800 m depth show an equatorward undercurrent that transports cooler and less saline water farther north (Smith et al., 1991). This appears to reduce our offshore eddy heat and salt anomalies in the lower thermocline density range
R. Morrow et al. / Deep-Sea Research II 50 (2003) 2229–2243
(25.7oso26.2). Lateral mixing also may contribute to maintaining the large heat and salt anomalies along isopycnals in the mode water range (26.5oso26.9). How sensitive are these AHA and ASA calculations to our choice of reference station? Due to the limited ship time and consequently the zig-zag cruise tracks, the strongest warm-core eddies were observed at the deepest offshore positions, and our reference profiles are limited to the stations further inshore. This means that our two reference profiles may have a geographical bias. We have two ways to test the stability of our eddy flux calculations. Firstly, we have constructed a mean reference profile from the nine stations outside the warm and cold-core eddies, and in water deeper than 1500 m (stations 47-49, 56-58, 66-68). This mean reference profile lies between the two reference profiles in Fig. 3, and increases the heat and salt fluxes for eddy A, and decreases the heat and salt fluxes for eddy B. So, for example, the eddy heat flux for waters warmer than 8 C using the nine station mean reference profile is 0.001 PW for eddy A, and 0.0015 PW for eddy B, compared with 0.00075 PW for Eddy A and 0.002 PW for Eddy B using two different reference profiles. This shows how minor adjustments in the reference profile can lead to a 30% change in our eddy heat transport magnitude, which gives an idea of the sensitivity of our transport calculations. Secondly, although we have no other hydrographic measurements offshore during the same time period, we can compare our reference profiles to the annual mean profile from the Levitus 1998 climatology, at a point located between eddies A and B at 110 E, 29 S. The Levitus T=S profile (marked in pink dots in Fig. 3) represents the mean T=S properties over the last 50 years. Fig. 3 shows that eddies A and B are considerably cooler and fresher than the Levitus profile in the surface waters (so25:5) and in the mode and intermediate waters (26.5oso27.2), and warmer and saltier in the lower thermocline (25.5oso26.5). However, these anomalies are not restricted to the warmcore eddies—all of our station profiles east of 110 E show this offset from the Levitus profile. This pattern of large scale cooling and freshening of the mode and intermediate waters, compared
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with warming, and increasing salinity in the thermocline, has also been detected over the last 25 years in repeat hydrographic sections at 32 S (Wong et al., 1999). From this we conclude that the Levitus annual climatological profile is not representative of the background reference state of the ocean in 2000,2 and so will not provide an accurate reference profile for calculating mesoscale eddy fluxes. So although we have some reservations about our chosen reference profiles, we believe that they provide the best estimate for our background mean state in September 2000. How representative are these three warm-core eddies? Fang and Morrow (2003) have analysed long-lived, anti-cyclonic eddies from 6 year of altimetric data in the southeast Indian Ocean between 20 S and 32 S (long-lived eddies are those whose SLA is more than 10 cm amplitude after 12–18 m). They find that on average, six of these long-lived warm-core eddies are generated per year—although yearly averages vary from 3 to 9 depending on inter-annual variations in the eastern boundary density structure. Eddies A and C had average amplitudes of around 30–40 cm in SLA; eddy B with an SLA of 80 cm was one of the strongest eddies measured during 1995–2000. If we assume six warm-core eddies the size of eddy A are generated each year, a rough estimate of the annual eddy heat and salt fluxes for waters warmer than 8 C would be 0.004 PW and 0.24 106 kg s1, respectively. This is a factor of 10 less than the estimates by VB94, who show an annual heat and salt flux of Agulhas eddies into the Atlantic (calculated over the entire eddy depth range) of 0.045 PW and 2.5 106 kg s1, respectively. How do these estimates compare with the mean heat and salt fluxes in the southern Indian Ocean? East of 90 E, there is a net northward mass transport of B2–3 Sv in the upper levels (Sloyan and Rintoul, 2001) which is part of the wind-driven subtropical gyre circulation. East of 90 E, inverse model estimates from the surface to B1000 m depth (Ganachaud, pers. comm.) show an upper ocean poleward heat flux of 2
The Levitus seasonal mean profile and monthly averaged September profile also show the same cooler and fresher offset at depth.
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R. Morrow et al. / Deep-Sea Research II 50 (2003) 2229–2243
0.01370.008 PW at 18 S and 0.2670.26 PW at 32 S. These inverse models do not resolve the narrow Leeuwin Current transport or its eddies very well, and we note that much of the poleward heat transport at 32 S comes from the equatorward transport of cold mode waters between 500 and 1000 m. However, our warm-core eddy heat flux magnitudes are 20–30% of the inverse model mean poleward heat fluxes. In this study, we have not calculated a closed heat budget: we are comparing essentially westward eddy fluxes with the mean poleward heat flux, and we have not included the contribution from cold-core eddies generates at these latitudes. However, the magnitude of the warm-core eddy fluxes suggests that Leeuwin Current eddies could contribute significantly to the upper-ocean westward heat flux in this mid-latitude band south of 20 S. The same inverse model calculation east of 90 E shows an upper-ocean poleward salt flux of 0.275.4 106 kg s1 at 18 S and 13712 106 kg s1 at 32 S. Here, the magnitude of our Leeuwin Current eddy salt fluxes is similar to this annual mean salt budget, although with large uncertainties. Despite the lower salinity flux at the surface, all of the eddies contribute a net high salinity flux from their stronger contribution at deeper levels. Talley (1999) describes how low-salinity water is modified in this region: (i) fresher surface flow from the Indonesian Throughflow is modified and exported south of 32 S as salty thermocline water; and (ii) a horizontal inflow of fresher AAIW and SAMW in the east leaves saltier in the west. Our calculations suggest that deep-reaching long-lived warm-core eddies that mix down past the thermocline could contribute to increasing the salinity in both these water masses. What is the role of mixing for these warm-core eddies? Sloyan and Rintoul (2001) have shown that in the Indian Ocean region between 18 S and 32 S, air–sea fluxes transform 7.5 Sv of surface and thermocline waters to SAMW (>26.0). Interior mixing converts a further 572 Sv of SAMW to AAIW. As a consequence of diapycnal mixing, there is an associated downward flux of heat through these levels and a downward flux of salt that increases the salinity of SAMW and AAIW. There is also an upward diapycnal salt flux
into the thermocline, which enhances the subsurface salinity maximum (Sloyan and Rintoul, 2000). The large AHA and ASA of our warm-core eddies is of the right sign to aid in the warming and increasing the salinity of the thermocline and mode water layers, particularly when combined with the strong lateral mixing from these eddies. From our simple study we cannot yet determine whether the heat and salt mixes principally along isopycnals, or whether diapycnal mixing from the strong horizontal shear at the eddy edge could also contribute. The passage of warm, low-salinity water from the Indonesian Throughflow to the Agulhas Current is a key part of the ‘‘warm route’’ of the global overturning circulation (Gordon, 1986). Most of this water is transported westward by the South Equatorial Current around 10 S towards Madagascar. The contribution from Leeuwin Current eddies has never before been addressed—mainly because they are not well resolved in hydrographic sections or global, coarse-resolution climate models. This study provides a first estimate of the heat and salt flux contributions of Leeuwin Current eddies; albeit with significant error bars in the calculation. A more detailed budget based on both warm and cold-core eddies should now be determined from a global, eddy-resolving OGCM, or a combination of altimetry with more extensive hydrographic data.
Acknowledgements Thanks to the scientists, technicians and crew aboard the R.V. Marion Dufresne who helped us make the hydrography measurements. Special thanks to Michele Fieux, for her organisational capabilities as cruise leader, for her patience as we constantly changed the cruise plans to ‘‘track’’ the eddies, and for resolving the post-cruise calibration headaches. Altimeter data were generated by the CLS Space Oceanography Division as part of the European Unions’ Environment and Climate project AGORA (ENV4-CT9560113) and DUACS (ENV4-CT96-0357), with financial support from the Centre for Earth Observation
R. Morrow et al. / Deep-Sea Research II 50 (2003) 2229–2243
Programme and Midi-Pyr!en!ees regional council. Additional inverse model calculations for the eastern Indian Ocean were kindly provided by Alexandre Ganachaud. We thank l’Institut Fran@ais pour la Recherche Technique et Polaire, l’Institut National des Sciences de l’Univers, and the Programme National d’Etude Du Climat for financing the TIP2000 cruise and this analysis.
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