Indonesian throughflow in a coupled general circulation model

JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 102, NO. C6, PAGES 12,341-12,358, JUNE 15, 1997

Indonesian throughflow in a coupled general circulation model

Niklas Schneider and Tim P. Barnett

Climate Research Division, Scripps Institution of Oceanography, La Jolla, California

Abstract. The Indonesian throughflow is analyzed in an extended simulation with a coupled ocean-atmosphere model. The model, developed by the Max-Planck-Institut far Meteorologie, Hamburg, Germany, combines an atmospheric general circulation model at T42 resolution (2.8 ø latitude by 2.8 ø longitude) and a primitive equation ocean model with zonal resolution of 2.8 ø and a meridional resolution of 0.5 ø in the tropics and is coupled without flux correction equatorward of a latitude of 60 ø. The onset and strength of the monsoon in the Indonesian waters agree well with climatology, and many aspects of the observed temperature fields in the eastern Indian Ocean and Timor Seas are found in simulation. Differences between simulation and observations of

temperature occur in mean and seasonal cycles in the far western Pacific. The annual cycles of sea level along the coast of Sumatra and Java are simulated satisfactorily. The simulated throughflow transports on average 13.8 Sv (106m3s -i) from the Pacific to the Indian Ocean. The vertically averaged (barotropic) component of the throughflow has a seasonal range of 13.1 Sv and is weakest in February and strongest in July. In contrast, deviations from the vertical average of the throughflow (baroclinic) are strongest in March and September. The average and seasonal cycle of the barotropic component of the throughflow are forced by winds over the Pacific and along the western coasts of Australia and South America, as described by the island rule. For closed Torres Strait, the contribution of the average bottom pressure torque is small, and friction closes the vorticity balance. For annual timescales, baroclinic flows affect the throughflow transport through the bottom pressure torque. The annual cycle of the baroclinic component of the throughflow is forced predominantly by winds over the Indonesian Seas. The throughflow exports 0.9 PW of heat from the Pacific into the Indian Ocean and is an important heat sink for the western Pacific. The throughflow is a major heat source for the Indian Ocean and is associated with reversal of the divergence of the meridional transport of heat south of 10øS that is balanced by heat fluxes from the ocean to the atmosphere.

1. Introduction

The Indonesian throughflow between the tropical Pacific and Indian Oceans is the only low-latitude transport between major ocean basins and is believed to be important for the general circulation of the Indian Ocean and for the global thermohaline circulation. The state of knowledge about its description, dynamics and impacts has been summarized in a recent issue of the Journal of Geophysical Research [Lukas et al., 1996].

The magnitude of the throughflow has been the subject of numerous observational and numerical studies. Godfrey [1996] gives estimates of the throughflow transport that vary from

'essentially nil to 20 Sv (=106 m 3 s'l), depending on data and methodology used. Several authors use the mean of this spread and take a typical value of the throughflow to be 10 Sv. The most recent, and possibly most reliable, estimate of the mean and seasonal transport [Meyers et al., 1995] is obtained from 6 years of expendable bathythermograph (XBT) observations and yielded a value of 5 Sv in the upper 400 m of the water column. A strong sensitivity to the choice of reference level and stratification below was found and 7 Sv were given as the most reasonable total transport.

Copyfight 1997 by the American Geophysical Union

Paper number 97JC00022. 0148-0227/97/97JC-00022509.00

The transport is strongest in late boreal summer and weakest in boreal winter. The exact amplitude and phase of the seasonal cycle are subject to some discussion. Wyrtki [1987] deduced from dynamic height differences between the western Pacific and eastern Indian Ocean that the throughflow is strongest during the southeast monsoon in July and August and weakest in January and February. A single year of direct current measurements in Lombok Strait [Murray and Arief, 1988]) showed maximum transports in August and smallest values in November and December. Multiyear XBT observations [Meyers et al., 1995] indicate zero transport in May-June and 12 Sv in August-September. Water mass analysis similarly suggests that the throughflow transport is weakest during northwest monsoon in northern winter [Ilahude and Gordon, 1996]. The majority throughflow waters come from the northern hemisphere, as indicated by tracer [Fine, 1985; Ffield and Gordon, 1992; Ilahude and Gordon, 1996] and drifter observations [Lukas et al., 1991].

A spectrum of models have been used to estimate and explain the magnitude of the throughflow and its subsequent impact on the Indian Ocean. Wyrtki [1961 ] assumed that the Indian to Pacific pressure gradient is balanced by friction. Godfrey [1989, 1996] estimated the throughflow from the Sverdrup balance by introducing the "Island Rule", which has been extended to include friction and shallow sills in the

Indonesian waters [Wajsowicz, 1993]. To explain the northern hemisphere origin of throughflow waters, Godfrey et al. [1993] suggested that either lateral mixing or departures from

12,341

12,342 SCHNEIDER AND BARNETt: INDONESIAN THROUGHFLOW

the Sverdrup balance close to the equator are important, and Nof [1995] suggested that advection of momentum is essential. Wind-driven numerical models in general are successful in calculating reasonable transports of the throughflow [Kindle et al., 1987; Semtner and Chervin, 1988; Kindle et al., 1989; Inoue and Welsh, 1993; Miyama et al., 1995].

The island rule [Godfrey, 1989] implies that the throughflow transport is determined by winds in the Pacific and along western and southern coasts of Australia [Godfrey, 1989; Wajsowicz, 1994] and that it is independent of winds of the open Indian Ocean. Masurnoto and Yarnagata [1996] suggest that this holds true for the seasonal cycle as well, in contradiction to the hypothesis that annual signals of the throughflow are forced by winds in the eastern Indian Ocean [Wyrtki, 1987; Kindle et al., 1987; Kindle et al., 1989; Clarke and Liu, 1993; Qu et al., 1994; Meyers et al., 1995; Yarnagata et al., 1996] and that semiannual signals in sea level along the coast of Sumatra, and thus possibly the throughflow, are forced remotely from the equatorial Indian Ocean [Clarke and Liu, 1993; Qu et al., 1994, Yarnagata et al., 1996]. We will reconcile these suggestions and show that the average and seasonal barotropic throughflow transport is governed by the island rule and is therefore independent of Indian Ocean winds. The annual cycle of the baroclinic part of the throughflow, on the other hand, is not in Sverdrup balance and is mainly forced by winds in the eastern Indian Ocean and Indonesian Seas.

The throughflow impacts the Indian Ocean circulation significantly [Godfrey, 1996]. Observed mass stream functions compare favorably with Sverdrup transports in the Indian Ocean if the throughflow is taken into account as an eastern boundary condition [Godfrey and Golding, 1981]. The implied deepening of the thermocline off western Australia is responsible for a lack of cold upwelling there, an anomaly compared to other oceans [Godfrey and Golding, 1981]. Numerical experiments with an open and a closed Indonesian archipelago confirm these findings [Hirst and Godfrey, 1993] and indicate that the throughflow heats the Indian Ocean and cools the Pacific and affects the surface temperature in remote regions where the altered thermocline structure is communicated to the surface. Similar experiments with a reduced gravity model confirmed these findings and showed that the throughflow also transfers time dependent signals from the Pacific to the Indian Ocean [Verschell et al., 1995]. More specifically, the heat transport of the throughflow was estimated by Hirst and Godfrey as 0.63 PW out of the Pacific, or one third of the total heat input into the tropical Pacific of their model.

Forcing of the Leeuwin current by the throughflow was investigated by Kundu and McCreary [ 1986], who found that vertical mixing, albeit of an unrealistic form [Godfrey and Weaver, 1991], induced a southward continuation of the throughflow in the Indian Ocean, similar to the Leeuwin current. Godfrey and Weaver [1991] and Hughes et al. [1992] showed that the supply of warm waters by the throughflow and the resulting atmospheric cooling generate the Leeuwin current. Thus the throughflow affects significantly the meridional heat transport and the surface heat fluxes in the Indian Ocean [Hughes et al., 1992].

The role of the Indian Ocean in the global thermohaline circulation is subject to some debate. Gordon [1986] suggested a "warm water return route" of thermocline waters from the Pacific returning to the North Atlantic via the

throughflow and Indian Ocean and around the southern tip of Africa. This view has been challenged by the numerical studies of Hughes et al. [1992] and Hirst and Godfrey [1993], who found little effect in the Atlantic after closing the throughflow. The inverse calculation by MacDonald [1993] showed also that the Indian Ocean exports approximately 1 PW of heat to the south across 30øS, and that induced heat fluxes in the South Atlantic are not consistent with Gordon's [1986] warm water route.

The only study reported to date that has investigated the throughflow with a coupled model is that of Latif and Barnett [1995], who used a global ocean general circulation model (GCM) coupled to a statistical atmosphere covering all the tropical oceans to study the remote impact of E1 Nifio-Southern Oscillation events on the other tropical oceans. In this study, the throughflow is investigated for the first time in a simulation with a global coupled general circulation model. An integration spanning 100 years was used to determine the transports, their seasonal variations, and underlying dynamics. Furthermore, the influence of the throughflow on the heat budget of the Indian Ocean is investigated. The advantages of a coupled simulation are that all fluxes at the surface of the ocean are calculated, rather than prescribed from observations, and the role of the throughflow in the coupled system can be assessed.

In the following, the coupled model is described (section 2) and its simulation is compared with observations (section 3). The mean and seasonal cycle of the throughflow is presented, and its dynamics are discussed (section 4). The heat transport of the throughflow is discussed in section 5 and is followed by our conclusions (section 6).

2. The Coupled Model

The coupled model (called ECHO) [Latif et al., 1994; Latif and Barnett, 1994; Schneider et al., 1996) consists of a state of the art atmospheric GCM [Roeckner et al., 1992; Deutsches Klirnarechenzentrurn, 1992] and a state of the art, primitive equation, ocean GCM. The atmospheric model has 19 levels in the vertical and was run at T42 (2.8 ø by 2.8 ø) resolution. The ocean model is global, actively forced by the atmospheric model between 60øN and 60øS without any flux correction and relaxed to climatology [cf. Levitus, 1982] at higher latitudes using a NewtonJan formulation. The horizontal resolution is variable with latitudinal spacing of 0.5 ø within 10 ø of the equator and expanding to 2.8 ø resolution poleward of 20 ø latitude. The longitudinal resolution is 2.8 ø and there are 20 levels in the vertical (see Latif et al. [1994] for additional details).

We analyze an extended simulation of the global climate spanning over 100 years. This simulation has been used to investigate decadal variability in the North Pacific [Latif and Barnett, 1994; W. Xu and T. P. Barnett, Multi-time scale heat budget in a coupled atmosphere-ocean GCM: A North Pacific study, submitted to Journal of Climate, 1995; hereinafter referred to as Xu and Barnett, submitted manuscript, 1995; W. Xu, et al., On the physics of decadal variability in the North Pacific, submitted to Journal of Climate, 1995; hereinafter referred to as Xu et al., submitted manuscript, 1995] and the warm pool [Schneider et al., 1996]. These studies established that the simulation generates all major features of the atmosphere and ocean with approximately the right strength and spatial structure. In the tropics the model suffers from a

SCHNEIDER AND BARNLWYF: INIX)NESIAN THROUGHFLOW 12,343

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Figure 1. Model bathymetry in Indonesian waters in 103m. Shaded areas show depths between coast and 1000 m. The coarse model resolution represents the complicated topography in the Indonesian waters (thin lines) by a single Indonesian Sea, connected to the Pacific via a passage at its northern end and via Torres Strait. Solid dots denote XBT lines compared to regional averages P, I, T, and J of the simulation (dotted boxes).

tendency to produce a split Intertropical Convergence Zone in the Pacific typical for coupled models and produces waters off south America that are too warm, a result of underestimating the stratus cloud cover over this region [Latif et al., 1994].

In this study we focus on the Indonesian throughflow. Because of the model resolution, Indonesian waters consist of

an Indonesian Sea and Timor Sea separated by the Lesser Sunda Islands and are connected to the Pacific by two gaps: one between New Guinea and Asia (called hereinafter "Indonesian passage"), with a sill depth of 478 m, and Torres Strait between New Guinea and Australia, with a shallow sill of 75 m

(Figure 1). The latter is usually closed in numerical models of the region [Kindle et al., 1989; Wajsowicz, 1995], as it is shallow, narrow, and full of reefs, although Hirst and Godfrey [1993] also show it open. However, as we will show in the following, the major portion of the throughflow occurs through the Indonesian Passage. The seasonal cycle of the transport through Torres Strait is, however, larger than observed and of similar magnitude as seasonal variations through the Indonesian Sea. Thus this topographic feature of the model is not of great consequence for the time averaged transport but distorts estimates of the seasonal cycle.

Because of the immense storage requirements, only surface quantities were saved as monthly averages. Mean and seasonal cycles of subsurface velocity, temperature, and salinity were reconstructed from instantaneous saves of all quantities needed to restart the model. These are typically available three to four times a year, with most restart files written in March, June, September, and December. This procedure resembles observational studies that determine long-term averages from collections of quasi-synoptic cruises. Aliasing of timescales shorter than a month does not occur, since the variance of

surface elevation from restart files equals the variance estimated from monthly averages, subsampled at months when restart files are available.

Mean and annual cycles of subsurface quantities were obtained from restart files as averages and best fits of annual and semiannual harmonics. Comparison of annual amplitudes and phases of surface elevation from restart files and monthly means validated this reconstruction. Semiannual periods are not reconstructed that well. For example, the semiannual signal propagates eastward in the equatorial Indian Ocean and has amplitudes of 1.2 cm on the coast of Sumatra,

approximately symmetric around the equator. The reconstruction of the semiannual cycle has differences in phase up to 3 months in areas of small amplitudes and underestimates the amplitude at the coast of Sumatra just north of the equator by up to 0.8 cm. Thus the reconstructed semiannual signal of subsurface quantities has to be viewed with caution.

3. Coupled Model Performance Surface fluxes and the oceanic state of this simulation have

been compared to available climatologies by Latif et al. [1994], Schneider et al. [1996], Xu et al. (submitted manuscript, 1995) and Xu and Barnett (submitted manuscript, 1995). To allow an evaluation of the relevance of the physics discussed later, the simulated seasonal cycles of winds, thermal structures, and sea level are compared with observations in the Indonesian waters and the eastern Indian Ocean. Specifically, the wind stress is compared to a recent climatology of Da Silva et al. [1994], and the simulated temperatures are juxtaposed with XBT-based climatology of Meyers et al. [1995].

3.1. Wind Stress

The simulated winds compare generally well with the observed seasonal cycle of the winds from Da Silva et al. [1994]. In February (Figure 2a), winds blow from the northwest over the Indonesian Sea and in the Timor Sea, while trades to the west of Australia blow from the southeast, creating a strong cud of the wind stress (upwelling) in the Timor Sea, while the winds blowing along the southern shore of Java transport waters onshore. These features are borne out both in the simulation and in the climatological wind stress.

In August, southeast monsoons are strongly developed and merge smoothly with southeast trades further to the west (Figure 2b) and transport waters southward from the coast of Java. The smooth transition between the monsoon and the

southeast trades is associated with a small curl, so that the Ekman pumping should be weak. The simulation agrees very well with the climatology in these aspects.

The monsoons over the China Sea are poorly simulated, probably due to the land sea flag in this region (Figure 1). However, this region is not essential for the Indonesian throughflow and should therefore not compromise its simulation by ECHO.

12,344 SCHNEIDER AND BARNEl'T: INDONESIAN THROUGHFLOW

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Figure 2. Simulated and climatological wind stress in newtons per square meter over the Indian Ocean and Indonesian waters from (top) Da Silva et al. (1994) and (bottom) ECHO for (a) February and (b) August. Stresses of 0.1 N m '2 in meridional and zonal directions are depicted over Australia.

3.2. Oceanic Temperature

Seasonal cycles of temperatures were prepared by Meyers et al. [1995] from XBT observations collected from 1983 to 1989 as bimonthly means with a vertical resolution of 20 m and horizontal spacing of 100 km along ship tracks (Figure 1). Since the model does not resolve the complicated topography of Indonesian waters, spatial averages north of the Pacific entrance of the throughflow (P, Figure 1) and in Indonesian waters (I, Figure 1) are considered. South of Java and Timor the topography is more homogenous and allows comparison of zonal averages of simulation for areas T and J (Figure 1) and XBT lines.

Annual average. The mixed layer of the simulation encompasses the top 50 m and has temperatures slightly cooler (0.3øK) in the Pacific (Figure 3) compared to Indonesian waters. Between 50- and 220-m depth, temperature differences reverse, and Indonesian waters are warmer by up to 1.8øK than waters in region P. Below 250 m, temperature differences reverse yet again, with Indonesian waters being warmer by 0.2øK. The observed thermocline (Figure 3)is less steep than the simulated, and waters in the Pacific are 0.2øK warmer than in Indonesian Seas.

In the section south of Timor the simulated thermocline

slopes downward toward the south and indicates westward flows at all latitudes (Figure 4). North of 12øS the simulation is very similar to observations. South of this latitude, observed

isotherms slope upward above 100 m. Below that depth, isotherms reach their greatest depth at 17øS and slope upward toward the coast. The observed temperature fields are associated with eastward transports south of 12.5øS [Meyers et al., 1995], opposite to the simulation.

Annual Average

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E

200

ECHO

500 ---I

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400

10 15 20 25

T/øC Figure 3. Annual averaged observed (thin lines) and simulated (thick lines) temperatures for regions P (solid lines) and I (dashed lines) (see Figure 1).

SCHNEIDER AND BARNETT: INDONESIAN THROUGHFLOW !2,345

E

T/øC, XBT, ?, annual average

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Annual average temperature in degrees Celsius Figure 4. (top) from observations and (bottom) ECHO for region T (see Figure 1).

Further to the west, comparison of simulation and observations shows similar differences (Figure 5). The simulated thermacline slopes southward all latitudes, while observations show an upward slope associated with the eastward gyre current [Meyers et al., 1995]. Only above a depth of 200 m does the simulation show a similar slope of the isotherms, indicating that the model's gyre current is trapped to the surface. The differences of simulation and observations

along these sections could be due to the unrealistically large Torres Strait in the model or subtle difference of the wind stress

in this region or due to differences in averaging. Annual harmonic. Annual cycles of the XBT data and

model are described by a least squares determination of the amplitude and phase of a sinusaid with an annual period. Missing points in observations are filled by linear interpolation in time.

In Indonesian waters, annual amplitudes are largest at the top of the thermacline at about 100-m depth both in model and observations (Figure 6). At shallower depths, simulated amplitudes are about half of observed. This is consistent with differences of simulated and observed vertical gradients if the seasonal cycle of subsurface temperature results mainly from vertical advection. Above 200 m, where amplitudes are large, temperatures are highest in February and March both in the observations and the model. Below, there are phase differences of up to 5 months, but amplitudes are small.

On the Pacific side of the throughflow, the simulation differs from XBT observations (Figure 6). Simulated amplitudes are largest below 100-m depth, while the observations are of similar magnitude above this depth. More importantly, the observed seasonal cycle in this region is in phase with Indonesian waters, while simulated signals are 6 months out of phase and lead to an overestimation of the seasonal cycle of simulated baroclinic pressure gradient. Later, we will show that this is a local failure of the model associated

with the frictional parameterization. South of Timor, simulated and observed annual phases agree

remarkably well (Figure 7). North of 12øS, the upper 250 m

attain maximum temperatures early during the year (February or March), as do waters above 100-m depth at all other latitudes. South of 12øS and below 100-m depth, temperatures are highest in July in the simulation and observations. Amplitudes show larger differences but are still surprisingly similar. For example, both simulation and observations place large amplitude close to the surface and at the poleward limits of the line, as well as centered at 14øS. There, however, simulated amplitudes show a large subsurface maximum, which is not observed. The simulation shows a band of low

amplitudes at a depth of 100 m, dominated by semiannual periods that are not found in observations. These discrepancies point to differences between simulated and observed thermoclines and vertical velocities, if subsurface temperature signals are caused by vertical advection.

Further west, gross features of the simulation are born out in observations as well, but differences exist in magnitude and details (Figure 8). Large seasonal amplitudes in the surface ocean at the southern end of the line are simulated at

approximately correct magnitude. The simulated subsurface maximum centered at 13øS is stronger than observed. Differences also exist at 20øS, where observations indicate a

maximum seasonal amplitude, but the model displays a minimum. The simulated phase of the annual signal has a very simple structure, with a line starting at 11øS extending southward at depth of 100 m delineating a half-year phase shift. Observations indicate such a structure as well, however, differences occur north of 1 løS.

In summary, many aspects of the observations are reproduced by the simulation of the oceanic temperature. Differences occur in mean and seasonal cycle on the Pacific side of the throughflow and lead to seasonally modulated baroclinic pressure gradients not documented from observations.

3.3. Sea Level

Annual and semiannual amplitudes and phases of sea level compare favorably with coastal observations of Bray et al.

100 - E

_c 200•

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400 -

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Figure 5. Annual averaged temperature in degrees Celsius from (top) observations and (bottom) ECHO in region J (see Figure 1).

12,346 SCHNEIDER AND BARNETI': INDO•I• THROUGHFLOW

C cle a Annuol , I , I

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b

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Figure 6. (a) Amplitude and (b) phase of the seasonal cycle of temperature from observations (thin lines) and simulation (thick lines) for regions P (solid lines) and I (dashed lines).

[1996]. Semiannual signals along the coast of Sumatra decrease away from the equator and reach a minimum at the eastern end of Java, in accordance with the observations. Further east, they increase again and suggest that the large signal at the coast of Timor cannot be explained by remote forcing from equatorial Indian Ocean but must be forced locally. In the model, the amplitudes are smaller than observed

by a factor of 5, likely due to sidewall friction along the coast introduced to prevent grid splitting and due to coarse resolution that resolves the coastal Kelvin waves poorly and measures their amplitude at least half a grid spacing (140 km) away from the coast.

Annual amplitudes reach 10 cm along the south coast of Java (Figure 9) during February, in good agreement with

a T, XBT, T, annual amplitude/øK b T, XBT, T, annual phose,/month 0 :•:•:•:==========================================================='-::::: .......... :::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::: 0 •"4 .. . . .• 2 •-- •.: ....... -'"'""'"•• .•.....•. :..'•j •, , , ,

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Figure 7. (a) Amplitude and (b) phase of annual cycle of temperature from (top) observations and (bottom) ECHO for region T.

SCHNEIDER AND BARNETF: [NDONESIAN THROUGHFLOW 12,347

a o

•,E 100-

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300 iiiii •.:.i-'::j•.•i½i::•:- .... •-:.s•:i:•:•:•::-.-:':.::•:.::.-'::-.-.::i.-:-:•:•:.:s•:•:?.-:..•: l '"'""J 40O

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Figure 8. (a) Amplitude and (b) phase of annual cycle of temperature from (top) observations and (bottom) ECHO for region J.

observed amplitude of 10 cm obtained in Sunda Strait during mid February [Bray et al., 1996]. The slow northwestward phase propagation (Figure 9) in the Timor Sea suggest generation off Java. The simulation also shows annual amplitudes of 10 cm in the western Timor Sea in the vicinity of 15øS and 110øE, similar to findings of Wyrtki [1962], which are largest in September and propagate westward. The near anticorrelation of sea level at the coast of Java and in the

Timor Sea causes an annual cycle of the meridional sea surface slope associated with the annual cycle of the throughflow transport.

4. The Indonesian Throughflow

4.1. Annual Average

The simulated throughflow transports on average 13.8 Sv (=10 • m3s '•) from the Pacific to the Indian Ocean (Figure 10). The major part (11.6 Sv) enters Indonesian waters via the Indonesian Sea, and only a small portion (2.2 Sv) flows through the model's Torres Strait.

Observational estimates of the transport based on a variety of methods vary from 2 Sv to 20 Sv [Godfrey, 1996], with an average of 10 Sv. Meyers et al. [1995] estimated the

ECHO annual am litude a

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1 8'%

80øE 100øE 1 20øE 140øE 1 60øE 1 80 ø

ECHO, •/, annual Rhosemonth b / I• "o .............. ........ • ' "' '•"•-','-:•:••"-.:.•ii: 6øN :'r' ' 0 ø ,

6øS

1 2øS

1 8øS

80øE 100øE 1 20øE 1 40øE 1 60øE 1 80 ø

Figure 9. (a) Annual amplitude and (b) phase of simulated sea level.

!2,348 SCHNEIDER AND BARNETT: INDON•IAN THROUGHFLOW

,,, Thrø ,ughfl, øw • • , • I •, I

20' [

15 ..................................................

10

5

t...T_o r r e s............'" '"'-.. .......... ' i , / , i , i i i

2 4 6 8 J0 J2

Month

Figure 10. Seasonal cycle of the Pacific to Indian Ocean transport in $verdrups (10•m • s4): total throughflow (solid curve), transport through Indonesian Sea (dashed curve) and transport through Torres Strait (dash dotted curve). The thin dotted line denotes the average of the total throughflow. Positive numbers correspond to transports from the Pacific into the Indian Ocean.

geostrophic transport of the Indonesian waters from XBT observations (the same ones used in section 3) and obtained a transport in the upper 400 m of 5 Sv with a stipulation that flow at and below 400 m could increase the transport to 7 S v. The coupled model simulates the transport of the Indonesian throughflow reasonably well. Transport through Torres Strait was estimated by Wyrtki [1961 ] to be zero. As expected, the unrealistic large Torres Strait in the model leads to an overestimation of this transport. Compared to the total throughflow, however, the simulated transport through Torres Strait is small.

The transport is toward the Indian Ocean at all depths and strongest at the surface, with values of 4.3x 104 m 2 s 4 (Figure 11). At a depth of 200 m, transports per unit depth in Indonesian Passage are 1.4x104 m 2 s 4 and form a local minimum. Below, flows increase to 2.1x104m2s4 at a depth of 400 m. Thus transports through the Indonesian sea occur in two vertical bands, separated at a depth of 200 m, which carry 6.2 Sv (above 200 m) and 5.4 Sv (below 200 m). This vertical structure is associated with pressure differences between Pacific and Indian Oceans due to differences in temperature (section 3.2.). Observations show a pressure gradient between the Pacific and Indian Ocean contained in the upper 200 m [Wyrtki, 1987] and indicate a surface trapping of the throughflow, even though estimates of the throughflow by transports in the Timor Sea [Meyers et al., 1995] indicate that flow below 400 m contributes about 2 Sv.

A map of the transport in the upper 500 m (Figure 12) shows the connection of the throughflow with currents in the Pacific and Indian Oceans. Most of throughflow transport is supplied by the North Equatorial Current (north of 7øN) via the western boundary current along the Philippines and from the south China Sea. On average, northward flow along the western coast of New Guinea tums eastward at the equator. As part of the seasonal cycle a small amount of water enters the

Indonesian Sea in August from the southern hemisphere via the this current.

In the Indonesian Sea, throughflow waters follow the western coast and form an anticyclonic circulation. A small portion (less than 1 Sv)drains through Sunda Strait, and the major part enters the Indian Ocean via the western tip of the Sunda Islands. Once in the Timor Sea, transports tums west and form a jet 300 to 400 km wide in a west-southwesterly direction, as expected from Sverdrup calculations [Godfrey and Golding, 1981; Godfrey, 1989]. The flow is affected by winds over the Timor Sea, as a small portion moves southward off the coast of Australia as part of the seasonal formation of an alongshore current off northwestern Australia.

Transports in Torres Strait are fed by northward boundary currents that originate at the bifurcation of the South Equatorial Current at 8øS. The majority of that flow remains in the Pacific as the South Equatorial Countercurrent, which is strongly developed in the coupled simulation [Schneider et al., 1996]. Transports through Torres Strait are only a small leakage of this flow.

4.2. Seasonal Cycle

The throughflow has a strong seasonal cycle with largest transports into the Indian Ocean of 20.2 Sv in July and smallest transport of 7.1 Sv in February (Figure 10). Transports through the Indonesian Sea and Torres Straits are in phase and contribute 7.7 Sv and 5.4 Sv, respectively, to the combined seasonal range of 13.1 Sv. Inoue and Welsh [1993] found similarly strongest transports during July to October and weakest transports in February. However, their seasonal

lOO

200

3OO

4OO

Thr, oug, flw, :' I I I

I I

I

l

I ' I ' I ' I '

1 2 3 4

Transport/10ms -' Figure 11. Time-averaged transport per unit depth from the Pacific into the Indian Ocean through Indonesian Sea (solid line) and Torres Strait (dashed line).

SCHNEIDER AND BARNEl'T: INDONESIAN THROUGHFLOW 12,349

ECHO, ov., tronsport/m=s -', O-500m

.......... .• C

1 00"E 1 20"E 1 40"E 1 O"E 1 8 0"

Figure 12. Annual average of transport in upper 500 m. Scale over Australia denotes 50 m 2 s '• in meridional and zonal directions.

range is much larger, with strongest transports of 18.1 Sv, as in this simulation, but with minimum transports of 0.5 S v from the Indian Ocean to the Pacific. The simulated seasonal

range of 5.4 Sv in Torres Strait is larger than estimates by Wyrtki [1961] of 1.4 Sv, even though phases agree.

The seasonal cycle of flow from the Pacific into the Indonesian Seas has a strong baroclinic component (Figure 13) in that the upper 200 m are out of phase with flow below. This is consistent with the out of phase relationship of seasonal cycles of temperature in the Indonesian Sea and western Pacific described in section 3.2. The baroclinic

component is strongest in March and September (Figure 13), in agreement with estimates in the upper 400 m deduced from XBT sections of nil in May-June to 12 Sv in August- September [Meyers et al., 1995]. The phase of the baroclinic signal differs from the barotropic transport which peaks in February and July and suggests that annual cycles of baroclinc and barotropic throughflow components are of different origin.

lO0

200

300

4OO

5OO

2 4 6 8 10 12 Month

Figure 13. Deviations from annual mean and vertical average of transport per unit depth from the western Pacific into the Indonesian Sea. Positive transports denote flow from the Pacific into the Indian Ocean.

Seasonal variations of baroclinic transports on the Pacific side are of the same order as the throughflow (Figure 14), while the 16øC isotherm, whose depth coincides with the zero crossing of the baroclinic throughflow and is used to describe the large-scale context of the annual cycle of the throughflow, has amplitudes of less then 5 m (Figure 15). In the Indonesian and Timor Seas, seasonal variations of the baroclinic flow are dramatic. Amplitudes of the 16øC isotherm reach 20 m off the coast of Java and more than 15 m off northern Australia. The 6-

month phase shift between these locations indicates the seasonal formation of a baroclinic jet south of Java (Figure 15), whose transport in the upper 200 m (Figure 14) leads maximum transports through the Indonesian Sea and has a range that exceeds that of the baroclinic throughflow in the top 200 m by 5 Sv and that of the entire throughflow by up to 3 Sv. This indicates that storage and local forcing in the Indonesian Sea are important aspects of the seasonal cycle of the throughflow, as suggested by Meyers et al. [1995].

BC, tra,nspo, rt, 0-187m ! ,

10 ' ' ' ' ' ' ' ........... Ekman 6 S ..,,' '...,.

5 ] ,•z •' •,.-'"" '•-... •, ,

> '.,.,.::::..,...,.,,-,q .,.-'"" I ,.," q'.q. :-.. ß ....- ,.,,,1 •.•.

-10 [ ß i ' i ß i [ i

2 4 • 8 10 12 [Vlon•h

Figure 14 Baroclinic (deviations from vertical average over entire water column) transport in the upper 200 m across sections shown in Figure 15 of the throughflow (solid curve), in the Timor Sea (dashed curve), and in the western Pacific (dash-dotted curve). Also shown is the seasonal cycle of Ekman transport across the Indonesian Sea at 6øS (dotted curve). Since Torres Strait is shallower than 200 m, its transport does not contribute to this volume budget of the baroclinic transport in the upper 200 m.

12,350 SCHNEIDER AND BARNETT: INDONESIAN THROUGHFLOW

a D

6øN

0 ø

6øS

1 2øS

1 8øS

80øE

ECHO annual am litude m

100øE 120øE 140øE 1 60øE 1 80 ø

Figure 15. (a) Annual amplitude and (b) phase of depth of the 16øC isotherm of ECHO. Phases denote month when isotherm is deepest. Dashed, dotted and dash-dotted lines in (a) denote sections used and described in Figure 14.

The annual signal of the throughflow is associated with eddies that are generated in the Timor Sea and propagate westward into the Indian Ocean with speeds of 0.1-0.2 m s '•. Their northern edges straddle the coast of the Lesser Sunda Islands and form the seasonally reversing baroclinc jet, while seasonally reversing currents are generated at its eastern edge along the coast of Australia. The eddy is anticyclonic in the second half of the year and cyclonic during the first half. The generation of the seasonally reversing currents is similar to the simulation of Woodberry et al. [1989]. Perigaud and Delecluse [1992] observe similar annual signals emanating from the eastern Indian Ocean and crossing the Indian Ocean along 12øS. In ECHO, the annual signal is attenuated by surface winds as it propagates westward.

In summary, barotropic and baroclinic components of the throughflow are not in phase. Seasonal cycles of baroclinic currents associated with the throughflow are dominated by seasonal changes in Indonesian Sea and eastern Indian Ocean. Storage of waters in the Indonesian Seas is an important aspect of the seasonal cycle of the baroclinc part of the throughflow, as waters accumulate from November to March and drain during the remainder of the year. The seasonal cycle of the throughflow is communicated into the Indian Ocean by the seasonal formation and westward propagation of an eddy in the Timor Sea.

4.3. Dynamics

Barotropic flow: The island rule. Conservation of volume dictates that throughflow transports equal northward transports in the Pacific across 44øS [Godfrey, 1996], the southernmost latitude of Australia, since Bering Strait and Bass Strait are closed in ECHO. If transports in a segment of the water column, say from a surface to a level of no motion,

are considered, the throughflow is also fed by vertical transport through this level and to the north of 44øS. The relative importance of horizontal transports above a level of no motion and upwelling through this level has been subject to speculation as no direct observations of the vertical transports in the Pacific exist. Godfrey [1996] neglected the vertical transport at 1000 m, while Gordon [1986] postulated that upwelling of deep waters is a major source of the throughflow.

In ECHO, the throughflow is supplied by northward flow below 4000 m that upwells in the low latitudes close to the equator, as evidenced by the meridional overturning stream function in the Pacific (Figure 16). This implies that an estimation of the throughflow transport has to consider the entire water column and cannot be simplified by an assumption of a level of no motion.

.E 2

5

6

40øS 20øS 0 ø 20øN 40øN 60øN

Figure 16. Meridional overturning stream function in the Pacific Ocean in Sverdrups. Values larger than the Indonesian throughflow transport of 13.5 Sv are shaded.

SCHNEIDER AND BARNETT: INDONESIAN THROUGHFLOW 12,351

ECHO, X7x•/10-"Nm

0"

\ 20øS

40"

-3

j 2 I

0.2

0.1

0.0

-0.1

-0.2

120øE 150øE 180" 1 50WV 120%V 90øW

Plate 1. Average curl of the wind stress in 10 '6 N m '3 of ECHO. Thick lines and numbers 1, 2, and 3 denote areas used in the evaluation of the island rule.

This requirement complicates the application of the island rule [Godfrey, 1989, 1996] since topography throughout the Pacific can influence the throughflow transport. The importance of topography relative to surface wind stress is quantified by a line integral along a path that surrounds Australia and the Pacific to the east of the vertically integrated horizontal momentum equation in the steady limit (i.e. the island rule [Godfrey, 1996])

• fv_.-. dl- pO -1 •[ P-H _VH. dl- pO -1 • •" dt= ,•i ,•i ,•i

F + N (1)

where P = (-v, u) is a rotated vector of zonal transport u and meridional transport v, integrated between the surface of the ocean at z =0 and its bottom at z =-H,f is the Coriolis parameter, Po is a reference oceanic density, P-H is the pressure at the ocean bottom, and x is the wind stress vector. The terms

on the right-hand side represent friction, F, and advection, N. The path 31 surrounds the island and ocean to its east and

corresponds to paths 1 and 3 for New Guinea and New Zealand and to path 2 for Australia (Plate 1). For New Guinea and New

Zealand (denote by NG and NZ), the first term on the left-hand side integrates to (f•v'-f•i)T, i = NG, NZ, wheref•v and fs denote the Coriolis parameters in the northern and southern extreme of the island and T is the meridional transport in the Pacific to the east of the island. Evaluation of this term along path 2 yields the multiple island rule [Wajsowicz, 1993] (f•v û -f•Û)TÛ - (fffz-f•ûs)7•z. If terms of the right hand-side of (1) are small, 7 •ø is the transport through the Indonesian Seas, and the difference T ̂ us - T •ø is the transport through Torres Strait.

The second term on the right hand side of (1) describes the joint effect of bottom topography and baroclinicity [Huthnance, 1984] and the third represents forcing by the wind stress. To estimate and compare the magnitude of these terms, it is convenient to rewrite them as area averages of bottom pressure torque and curl of the wind stress using Stokes' theorem. Equation (1) applied to Australia reads

p•1IldA_Vx(p_ n_VH) - pG•llaA_vx_, = F + N 2 2

(2)

ECHO, Bottom oressure toroue

20os • . o

40"S •x• -

10-"Nm -s depth 10•m - ,

120øE 150øE 180 ø 150"W 120"W 900W

1..5

1.0

05

•00

-0.5

-1.5

Plate 2. Average bottom pressure torque in 10-6N m '3 of ECHO. Superimposed are ECHO's topography in 103 m. Please note that the color scale is 5 times larger compared to the scale of Plate 1.


Table 1. Island Rule Transports in Sverdrups due to Bottom Pressure Torque, Wind Stress Curl, and Their Sum for Areas Bounded by Lines Indicated in Plate 1, and for a Combination of the Paths using the Multiple Island Rule (MIR)

Path Bottom pressure Wind Stress Curl Total torque

New Guinea 1 -25.7 -0.80 -26.4

2 6.7 15.3 22.0

New Zealand 3 7.7 47.1 54.8

Australia 2+3, MIR 8.4 26.2 34.6

Closed Torres Strait 1 +2+3+S, MIR -0.8 18.0 17.2

The last row includes the contribution of path S that fills the small gap between paths 1 and 2 and shows the island rule transport if Torres Strait is closed and the combined land mass of Australia and New Guinea is considered. Positive transports indicate flow from the Pacific into the Indian Ocean.

and similar expressions result for New Guinea and New Zealand, where the second term on the left is absent. We refer to (2) as "island rule" as it gives a complete diagnosis of transports around islands.

The bottom pressure torque and wind stress curl from the output of ECHO have been evaluated in a manner that is consistent with the model's finite difference formulation and

that satisfies Stokes' theorem. The curl of the wind stress

reflects meridional gradients of zonal wind stress, with westward trade wind and eastward winds in midlatitudes, and has

typical magnitudes of 0.1xl0'tN m '3 (Plate 1). In contrast, the bottom pressure torque is localized on the continental slopes and along undersea ridges and reaches magnitudes of 0.5xl0'tN m '3 (Plate 2), 5 times larger than the curl of the wind stress.

Transports due to wind stress and bottom pressure torque in (2) for New Guinea, New Zealand, and Australia (Table 1) show that wind stress dominates the island role transport for Australia and New Zealand and that bottom pressure torque exerts significant control. In case of New Guinea the bottom pressure torque dominates over wind stress. Combination of both forcings around Australia and New Guinea grossly overestimates transport through Torres Strait (54 Sv) and fails to capture the sign of the transport through the Indonesian waters. This is expected since in ECHO, and in reality, friction in the shallow and narrow Torres Strait redirects water to pass to the north of New Guinea. In the limit of overwhelming friction in Torres Strait, the gap between Australia and New Guinea is closed, and the island rule estimate for a path surrounding their combined land mass is appropriate. A resulting transport of 17.2 Sv is more in line with ECHO's throughflow but still too large and indicates that the remaining terms of (2), friction or advection, account for the difference. A likely candidate is friction, since ECHO employs sidewall friction with a Newtonian damping parameter of 5x10 '6 s '• that constrains flow along coasts and in the narrow strait between New Guinea and ECHO's Asian land mass.

The effect of the bottom pressure torque on the combined land mass of New Guinea and Australia is small (-0.8 Sv) and implies that bottom pressure torque does not determine the magnitude of the throughflow and that sills in ECHO's Indonesian waters do not constrain the flow. More

importantly, deep flows (Figure 16) that comprise the transport due to the bottom pressure torque do not affect the

throughflow. This is reassuring as deep overturning in ECHO is unrealistic and results from numerical diffusion [Schneider et al., 1996] as shown by a comparison with a new simulation that changed the numerical scheme, has no numerical diffusion, and lacks the deep overturning. The vertical stream function of the Pacific (Figure 16) results from a superposition of an overturning cell driven by numerical diffusion that has no effect on the throughflow and a realistic wind driven northward flow in the top 1000 m that determines the throughflow transport. In the top third of the water column these circulations cancel each other and create the impression that the throughflow is forced by flow below 3000 m in the South Pacific.

In summary, the throughflow in ECHO is driven by the surface wind stress and retarded by friction. The bottom pressure torque affects the circulation around New Zealand, Australia and New Guinea but is small when the combined land

area is considered. Deep flows associated with bottom pressure torque do not force the throughflow, a reassuring conclusion as the deep flows in ECHO result from numerical diffusion.

Barotropic flow: The seasonal cycle. The neglect of the time derivative in (1) is valid for timescales shorter than the adjustment time of the global barotropic circulation. This implies that the seasonal cycle of the throughflow is governed by the island rule and only forced by winds in the Pacific and along the coast of Australia, consistent with results of Masumoto and Yamagata [ 1996], who singled out forcing by winds along the coast of Australia. Winds in the equatorial Indian Ocean and along the shores of the Sunda Islands do not affect the island rule transport directly in contradiction to suggestions that the seasonal cycle of the throughflow is forced by upwelling in this region. If signals from the Indian Ocean affect the throughflow transport, they must do so through either friction or advective terms of the island role or through interactions with baroclinic signals via the bottom pressure torque.

The island rule transport due to direct influence of the wind stress for the combination of Australia and New Guinea and

with consideration of New Zealand (Figure 17) is in phase with ECHO's throughflow transport and has a range of 13.6 Sv. The annual cycle of transport due to the bottom pressure torque has a seasonal range of 7.4 Sv and reaches its extrema in March and September. The residue between sum of wind and bottom

SCHNEIDER AND BARNET'F: INDONESIAN THROUGHFLOW 12,353

-5

Island Rule , I , I • I I I I

... TF,• d

..... bpt , ' I i I

2 4 6 8 10 12

Month

Figure 17. Island rule transports for closed Torres Strait due to annual cycles of curl of the wind stress (solid curve) and bottom pressure torque (dash-dotted curve). The annual cycle of the throughflow from the western Pacific into the Indonesian Sea (TFin a) is shown as heavy dashed curve.

pressure torque and throughflow transport indicates the important role of friction in the annual cycle of the throughflow.

The annual amplitude of bottom pressure torque is largest on either side of the passage between the western Pacific and the Indonesian Sea. This suggests that baroclinic signals from both the Indian Ocean and the Pacific have some control on the

seasonal cycle of the throughflow via the bottom pressure torque. Wind stress contributions from the western coast of Australia dominate the response, consistent with Masumoto and Yamagata [1996].

In summary, the seasonal cycle of the barotropic transport of the throughflow is predominately forced by winds in the Pacific and along the west coast of Australia, and affected by the bottom pressure torque on either side of the sill from the Pacific into the Indonesian Sea. This indicates that baroclinic

signals from the Pacific and Indian Oceans have some effect o n the transport of the throughflow. The extent to which this result is dependent on the formulation of ECHO and its numerical diffusion remains to be investigated with other models.

Baroclinic flow. Time-averaged baroclinic flow in the upper 200 m in the Indian Ocean shows a westward jet between 9 ø and 14øS that is qualitatively consistent with a Sverdrup balance [Godfrey and Golding, 1981]. Local forcing, such as Ekman pumping, is important since transports in the upper 200 m are divergent. The southward baroclinic transport at 2øS of 1.7 Sv is smaller than the southward transport of 3.9 Sv between Java and New Guinea and a westward transport of 6.4 Sv between Java and Australia along 121øE. An exact determination of the physics of the baroclinic structure of the throughflow requires a detailed analysis of the effects of vertical mixing and vertical propagation in the ocean and goes beyond the scope of this study.

Baroclinic adjustment of the ocean takes much longer than a year and implies that the seasonal cycle of the baroclinic structure of the throughflow is not in Sverdrup balance. Therefore winds in the Indian Ocean and in the Pacific can

affect the baroclinic structure of throughflow as suggested by Wyrtki [1987] and Meyers et al. [1995].

The preponderance of forcing in Indonesian waters is suggested by the following. First, seasonal ranges of sea level (Figure 9) and displacement of the 16øC isotherm (Figure 15) in Indonesian Sea are large compared to western Pacific and are associated with large baroclinic transports in the Timor sea (Figure 14). Second, Ekman transports due to winds in the Indonesian Sea are in phase and of proper magnitude to account for the transport in the Timor Sea (Figure 14). Third, in the western Pacific, baroclinic transports are of the same order as the throughflow (Figure 14) and suggest that Ekman pumping in this region is not important.

To test the hypothesis that winds over Indonesian waters are important in forcing the seasonal cycle of the baroclinic throughflow, experiments were conducted with a one-and-a- half-layer reduced gravity model. The model solves equations of the equatorial beta plane [Moore and Philander, 1977], linearized around a state of rest and augmented by Laplacian horizontal diffusion of momentum and a sponge layer at the poleward boundaries. The model covers the Indo-Pacific from Africa to America, is bounded by walls at 30øS and 20øN, has the same resolution (2.8 longitude, 0.5 latitude) as ECHO, a land flag derived from the 200 m isobath of ECHO (this implies that Torres Strait is closed), and was forced with the annual component of the wind stress of ECHO. The depth of the upper layer is 200 m, consistent with the depth of the zero crossing of the baroclinic portion of the throughflow, and its reduced gravity is 0.05 m 2 s 'l, as suggested by the ratio of sea level (Figure 9) and displacements of the 16øC isotherm (Figure 15). Initial conditions are a state of rest and no interface displacement, and all results presented are from the last of 5 years of integration.

The seasonal cycle of interface displacement simulated by the reduced gravity model (RG) shows large amplitudes in the Timor Sea and along the southern coast of Java and small amplitudes in the western Pacific (Figure 18), consistent with results from ECHO (Figure 15). Unlike the results of ECHO, amplitudes and phase in the Indonesian Sea are similar to the western Pacific in RG, because of different frictional

characteristics of the passage between the western Pacific and the Indonesian Sea between these model. Closure of this

passage in RG, the limit of strong friction there, yields amplitudes and phases in the Indonesian Sea similar to the Timor Sea and more consistent with results from ECHO. This

suggests that unrealistic sidewall friction causes the 6 month phase shift between temperatures in the western Pacific and Indonesian Sea (Figure 6) in ECHO. Since the small phase lag of RG correspond closer to observations the influence of the Pacific on the seasonal cycle of throughflow is more realistic and smaller in RG than in ECHO.

RG's throughflow is strongest toward the Indian ocean in September (Figure 19) in agreement with the baroclinic component in the upper 200 m of the throughflow of ECHO (Figure 13). The amplitude of the response of RG depends on the fraction of wind stress projecting onto the vertical mode simulated, on frictional effects, and on losses to the barotropic flow due to the bottom pressure torque. Since the interest here is the relative importance of forcing in different areas, rather that the absolute magnitude, all results of the throughflow transport are normalized by the annual amplitude of the throughflow. For total winds applied the annual amplitude of the throughflow in RG is 12 Sv.

12,354 SCHNEIDER AND BARNETt: INDONESIAN THROUGHFLOW

6øN

o

6øS

1 2%

RO m

Ind 5 -

18øS o • 80øE 100øE 120øE 140øE 1 60øE 180 ø

Figure 18. Annual amplitude of interface displacement in meters simulated by a reduced gravity model of the Indo-Pacific within 20 ø latitude from the equator. The model was forced with the annual components of the surface wind stress of ECHO. Thin lines denote the model's topography, and thick dashed lines show borders of regions where surface wind stresses are applied in experiments Ind, Pac, and Indo.

The relative contributions to the annual cycle of the throughflow from winds over the Pacific, Indian Ocean, and Indonesian Seas are determined by applying winds over these areas only (Figure 18). The annual cycle of winds over the Indonesian Sea (region Indo in Figure 18) explains 60% of the total throughflow (Figure 19) and is the largest forcing despite its small area. Winds in the Pacific and Indian Oceans

contribute only 23% and 17%, respectively (Figure 19). These results confirm that wind stress over Indonesian waters is the

dominant forcing of the seasonal cycle of the baroclinic structure of the throughflow.

In addition, local winds generate an eddy south of the Lesser Sunda Islands that transmits the annual signal of the throughflow into the Indian Ocean in accordance with results from ECHO. The dynamics of RG confirm that this eddy is a westward propagating Rossby wave.

5. Throughflow Heat Flux

The role of the throughflow in the heat budget of the Indian Ocean is studied on both an annual average and seasonal

1.0-

o.o

RG Thro, ughrio, w I • , • , I ,

0 2 4 6 8 10 12

Month

Figure 19. Normalized throughflow simulated by the reduced gravity model for control run and for experiments where the wind stress is applied over the Pacific (Pac), Indian Ocean (Ind), and Indonesian sea only (Indo). Borders of these regions are shown as dashed lines in Figure 18.

timescale. The advective heat flux is estimated from the annual

cycles of temperature and velocity of the throughflow and in the Indian Ocean, and the net air/sea exchange over the Indian Ocean is determined. We note in advance that these estimates

from ECHO are generally in good agreement with those obtained from observations, results that will be noted where

appropriate.

5.1. Annual Heat Budget Over the Indian Ocean

For comparison with observations, the advective heat flux was integrated zonally along 29øS from Africa to Australia and referenced to 0øC. It is southward out of the Indian Ocean and

has a magnitude of 1.52 PW. Both direction and magnitude of this flux are in surprisingly good agreement with indirect estimates from bulk formulae, e.g., 1.42 by Hsiung [1985] and 0.98 by Georgi and Toole [1982], and direct estimates from hydrographic observations, e.g., 1.3 by MacDonald [1993] (Model B; see also her Table 2 for additional estimates) and 1.7 by Toole and Warren [1993]. Thus the model simulates the advective heat flux in the Indian Ocean reasonably well.

The magnitude of these estimates has to be viewed with caution. If volume transports across a section do not vanish, such as in the southern Indian Ocean due to the Indonesian

throughflow, estimates of advective heat transports depend on the scale of temperature [Hall and Bryden, 1982]. To quantify the effect of the throughflow on the advective heat flux in the Indian Ocean and to extract this sensitivity, we write the meridional heat transport as a sum of advection by a spatially averaged drift and by the remainder of the flow.

II vT = v ll T + II v' T

with v=v+v', Ildxdzv'=0 (3)

where v is the meridional velocity, T is the temperature, and integrals are taken over a zonal cross section of the Indian Ocean. The first term on the right-hand side of (3) is the advection of spatially averaged temperature and depends on the scale of temperature. We reference this term by the spatially averaged temperature of the return flow into the Pacific between Australia and Antarctica (3.4øC). The second term represents advection by baroclinic and horizontally swirling motion with zero transports and is invariant to changes of the temperature scale. In the following, these motions will be referred to as drift and baroclinic or swirling, respectively.


Direct estimate of the advective throughflow yields a baroclinic component from the western Pacific into the Indonesian sea of-0.3 PW, compared to a barotropic component, referenced to 3.4øC, of-0.4 PW. Heat transports through Torres Strait are mainly carried by the drift, with a value of-0.2 PW, while baroclinic and swirling motions in the shallow and narrow strait contribute -0.01 PW only. Rectification of seasonal cycles of advective heat fluxes is small, consistent with linear dynamics suggested above.

In the Indian Ocean the throughflow affects mainly the heat flux due to baroclinic and swirling motions and leaves a clear signature of the inflows into Indonesian Sea at 1.5øS and through Torres Strait at 10øS in meridional advection of heat (Figure 20). The flow from the western Pacific changes the heat transport by -0.65 PW, and the change associated with Torres Strait contributes -0.12 PW. The entrance of throughflow waters into the eastern Indian Ocean is accompanied by a change in slope of the meridional heat advection south of 10øS that indicates loss of heat from the ocean to the atmosphere in the southern Indian Ocean. These heat losses are anomalies for

the latitude band and are attributed to the heat flux of the

throughflow [Hirst and Godfrey, 1993; Godfrey et al., 1995]. The model's budget north of 29øS is closed by net air/sea heat flux of 0.8 PW into the Indian Ocean and a residual of 0.4 PW

due to lateral mixing, since storage of heat in the Indian Ocean north of 29øS accounts for 4xl 0 '3 PW only.

The heat transport due to the drift term in (3) is zero north of 1.5øS since there is no meridional transport of mass. South of passages to the Pacific, the transport is proportional to the sectional average of ocean temperature and shows very small changes with latitude (Figure 20). This implies that the divergence of this term is small and of secondary importance to the Indian Ocean. Referenced to 3.4øC, the heat transport due to the southward drift is on average -0.1 PW, much smaller than transports associated with swirling and overturning motions.

o

c- o

o

0.0

-0.5

-1.0

-1.5

ECHO, Indian Ocean , , , I • , , ! , , ,. I , , , • . , .

_

.

' ' ß i . . . i , , ,' i , . , i . , ß

-2O -10 0 10 2O

Latitude

Figure 20. Advective heat transports in petawatts (10 •s W) across the Indian Ocean including the Indonesian Sea. Shown are heat transports due to overturning and swirling motions that have zero meridional transport (solid curve), and due to a mean southward drift that captures meridional transports (dashed curve). For the latter, the reference temperature is 3.4øC, the temperature of the return flow to south of Australia. Latitudes of throughflow from the western Pacific and through Torres Strait are indicated by thin dashed lines.

0.0

• ß c o

-2.0

ECHO, Indian Ocean I , I , I , I , I , I

2 4 6 8 10 12

Month

Figure 21. Seasonal cycle of advective heat transport at 20øS (solid curve), and entering the Indian Ocean at 120øE from the throughflow (dashed curve). The divergence of advective fluxes and its annual average are shown by dash-dotted curve.

The net heat flux of the Indonesian throughflow compares well with results of Hirst and Godfrey [ 1993]. They estimated the Indonesian throughflow heat flux of 0.63 PW by differencing surface heat fluxes over ocean basins from integrations with open and closed throughflow.

The above results raise the question of the role of the throughflow in the maintenance of the eastern end (the Pacific portion) of the warm pool. Comparable budget studies in a box bounded by 20øN, 20øS, and 180 ø longitude show swirling and overturning motions that import from the east 2.1 PW into the west Pacific and export 1.3 PW to the north and 0.3 PW to the south. The divergence of these terms in the western Pacific is a heat gain of 0.5 PW and is balanced to large degree by the baroclinic transport into the Indonesian Sea of 0.31 PW. The heat flux divergence due to the drift through these sections, referenced to 3.4øC, is a gain of 0.1 PW, compared to heat loss by the drift of the Indonesian throughflow of 0.6 PW. The difference resulting from these advective processes of-0.3 PW is balanced by local air/sea heat exchange. Thus the throughflow is an important cooling agent for the west Pacific warm pool. Additional details of the heat budget of the west Pacific portion of the warm pool are given by Schneider et al. [1996].

5.2. Seasonal Cycle of Heat Transport of the Indian Ocean

The seasonal cycle of the advective heat transport of the throughflow has the same phase as the volume flux (Figure 21). It is close to zero (0.1PW) in February and warms the Indian Ocean at a rate of 1.4 PW in July (all heat fluxes are referenced to 3.4øC). Across 20øS the advective heat flux is almost in phase and exports between 1.0 PW in March and 1.8 PW in August. The advective divergence in the Indian Ocean has the opposite phase and cools the Indian Ocean by 0.97 PW in March and 0.19 PW in October.

If we take 15øS as a typical latitude at which to evaluate this seasonality, the greatest heat export occurs in July with a minimum occurring in approximately January. These seasonal changes are mostly located above 200-m depth, with the vast majority of the change occurring above 100-m. There is a weak

12,356 SCHNEIDER AND BARNETI': INDONESIAN THROUGHFLOW

ECHO, In,dian Oce,an, t,o 20,øS i I - . i

2.õ

2.0

0.5 C)oC) ......

-0ø5

ß ß i - i - i .i i ß i i, i

2 4 6 8 10 12 Month

2,2,, Seasonal c¾c1½ of t•e surface •eat •ux i• •etawatts •o•t• of 20•S •d east of Sumatra, Ja•a, ]20•, and Australia. A•u• a•em•e •eat gux is s•ow• as a t•i• ]i•e.

reversal of the transport with season and this occurs between about 90 and 200 m.

Advective heat transports are in contrast with the net surface heat flux in the Indian Ocean to the north of 20øS

(Figure 22). It has a strong semiannual signal, as expected from the near equatorial symmetry of the area under consideration, and is largest in March and October and smallest in January and July. The seasonal range of 3.1 PW of the surface fluxes of the Indian Ocean exceeds the seasonal

advective flux by almost a factor of 4. Thus the seasonal signal of the throughflow does not seem to be as important to the seasonal heat budget of the Indian Ocean as it is to the long- term balance.

6. Summary and Conclusion

The Indonesian throughflow is analyzed in an extended simulation with a coupled ocean atmosphere model. The model, developed by the Max-Planck-Institut fQr Meteorologie, Hamburg, Germany, combines an atmospheric GCM at T42 resolution and a primitive equation ocean model with zonal resolution of 2.8 ø and meridional resolution of 0.5 ø

in the tropics and is coupled without flux correction equatorward of a latitude of 60 ø. The global nature of the model necessitates that the complicated topography of the Indonesian archipelago is represented in a simplified manner and Pacific and Indian Oceans are connected by an Indonesian Sea and by Torres Strait between Australia and New Guinea. Following Wajsowicz [1996], details of the topography are of secondary importance dynamically but for their effects on dissipative processes, since throughflow chooses the westernmost passage available, i.e., Molucca Strait in observations [Wyrtki, 1961; Ffield and Gordon, 1992] or the coast of the Indonesian Sea in ECHO.

Within the resolution of the global model, wind stress, thermal structure, and sea level are simulated rather well by the coupled model. The onset and strength of monsoons in Indonesian waters agree well with climatology, and many aspects of observed temperature fields in the eastern Indian Ocean and Timor Seas are found in simulation. Differences

between simulation and observations occur in mean and

seasonal cycles on the Pacific side of the throughflow and suggest an average and seasonally modulated baroclinic pressure gradient in the simulation. Semiannual cycles of sea level are smaller than observations but reproduce the observed phase at the coast. The underestimation of the coastal sea level signal can be attributed to the model's coarse resolution, which ill represents coastally trapped waves. The simulated annual cycle of sea level is in good agreement with observations.

The model suffers from numerical diffusion [Schneider et al., 1996] that is responsible for unrealistic deep upwelling at the equator. The diffusion is also a likely explanation for other failures of ECHO. For example, the simulated temperature- salinity relationship in Indonesian seas does not show the observed figure eight shape [Ilahude and Gordon, 1996]. Overall, however, simulations of thermal structure and

velocities in the upper oceans are very encouraging considering that the only forcings of the coupled model are solar insulation and the surface boundary conditions poleward of 60 ø latitude.

The simulated throughflow transports on average 13.8 Sv (=106m 3 s 'l) from the Pacific to the Indian Ocean. The majority (11.6 Sv) of this transport passes through the Indonesian Sea; only a small portion (2.2 Sv)flows through Torres Strait. Waters entering the Indonesian Sea are supplied mostly by southward boundary currents in the Pacific that derive from the North Equatorial Current and the South China Sea. Little water originates from the southern hemisphere. In the Indonesian Sea, the throughflow feeds an anticyclonic circulation, which drains into the Timor Sea and forms a westward jet to the south of Java that continues into the Indian Ocean.

ECHO's throughflow has a seasonal range of 12 Sv, made up of in phase variability in the Indonesian Sea and Timor Strait, which contribute 60% and 40%, respectively. Transport is weakest in February and strongest in July. The seasonal range through the Indonesian Sea agrees in phase and roughly in magnitude with other model-derived estimates. The annual cycle in Torres Strait is overestimated compared to observations, suggesting that Torres Strait is too large in the model. The seasonal cycle of the current system in the upper 200 m associated with the throughflow is dominated by seasonal changes in the Indonesian and Timor Seas, rather than by transport and sea level variations in the western Pacific. Storage of waters in the Indonesian Sea is an important aspect of the seasonal cycle of the throughflow as suggested by Meyers et al. [ 1995].

Dynamics of barotropic and baroclinic components of the throughflow are distinct, since the barotropic flow has adjusted to forcing for seasonal and longer timescales. The barotropic throughflow is therefore governed by the island rule and forced by winds over the Pacific and along the western coasts of Australia and South America. For closed Torres Strait, forcing of the average throughflow by bottom pressure torque is small (-0.8 Sv)and implies that topography does not control the throughflow transport. For the annual cycle, transports due to the bottom pressure torque are significant and suggest that baroclinic signals interact with the barotropic throughflow transport. The extent to which this is realistic or an artifact of ECHO needs to be investigated by an estimation of terms of the island rule of other, preferably high-resolution, models. Wind stress and bottom pressure torque together overestimate the throughflow, so that advection or more likely friction is


important. The sensitivity of the throughflow to the parameterization of dissipative processes is another important question to be tackled with high-resolution models. Alternatively, precise observations of the throughflow transport and winds might be used to constrain and determine dissipative properties of the Indonesian Seas, if the bottom pressure torque can be estimated. Different dynamics of barotropic and baroclinic components of the throughflow suggest that observations of thermal structure of the ocean [Meyers et al., 1995] captures baroclinic signals of the throughflow transport, while observations of winds determine variations of throughflow transport via the island rule [Wajsowicz, 1994].

The annual cycle of the baroclinic component of the throughflow is not in Sverdrup balance and is mainly forced b y winds over Indonesian waters. This is suggested by results of ECHO and experiments with a reduced gravity model that showed that local winds in the Indonesian waters account for

over 60% of the annual cycle of the baroclinic component with the remaining signal explained in approximately equal parts by winds in the Indian Ocean and Pacific. The distinction of dynamics of barotropic and baroclinic components of the throughflow reconciles results of Masurnoto and Yarnagata [1996] and suggestion that winds along the coast of Java force the annual throughflow signal [Wyrtki, 1987; Meyers et al., 1995].

In light of the time-averaged sea level difference between the Pacific and Indian Ocean, the seasonal cycle of the baroclinc component of the throughflow can be understood in the following terms. During the southeast monsoon, Ekman transports export waters out of Indonesian waters and lower sea level. The resulting pressure gradient from the western Pacific to the Indonesian Sea increases the throughflow. This flow continues along western boundaries of the Indonesian Sea into the Indian Ocean. During the northwest monsoon, Ekman transports raise sea level in the Indonesian and Timor Sea, and the pressure gradient from the western Pacific to the Indonesian Sea is nil. The Pacific to Indian Ocean pressure differences occurs in the Timor Sea. Its higher latitudes and the absence of western boundary friction allow the pressure gradient to be balanced by rotation and result in a weak through flow.

In terms of the coupled system, the seasonal cycle of the throughflow is a forced response to the monsoons that are caused by the seasonally varying solar insolation and the distribution of the land and ocean. Thus feedbacks of the

seasonal cycle of the throughflow on the atmosphere seem be small, and the throughflow is a passive response of the ocean to the seasonal forcing by the sun.

The throughflow exports 0.9 PW of heat from the western Pacific into the Indian Ocean. The throughflow is an important heat sink for the western Pacific compared to advective heat exchanges of the western Pacific warm pool with the surrounding cooler ocean. The throughflow is major heat source for the Indian Ocean and has a marked influence on the

advective heat transport in the Indian Ocean and on the sign of the surface heat flux in the southern Indian Ocean. This result

raises the question about the role of the throughflow in the coupled ocean-atmosphere system of the Indian Ocean, which will be the subject of a future study.

The results presented in this paper show for the first time an analysis of the Indonesian throughflow in a coupled model. It goes to the credit of the developers of this model at the Max-

Planck-Institut ftir Meteorologie, Hamburg, that the quality of the coupled simulation is sufficient to allow such a study.

Acknowledgments. We wish to thank Norman Barth, Nan Bray, Sarah Gille, Michele Morris, and Janet Sprintall for many discussions. The computational expertise of Jack Ritchie, Juan Soberani, and Michael Hamilton is greatly appreciated. The integration of the coupled model was performed at Los Alamos National Laboratory under the auspices of D. Poling and with the cooperation with C. Keller. We are grateful to G. Meyers for the XBT data, A.M. da Silva for his wind stress product, J. Potemra for his reduced gravity model, and the Max- Planck-Institut for Meteorologie for sharing their coupled model with us. This work was supported by the Environmental Science Division of U.S. Department of Energy (grant DE-FG03-91ER61198) as part of the Atmospheric Radiation Measurement Program, by NSF Climate Dynamics grant ATM 93-14495, and by Scripps Institution of Oceanography.

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(Received January 17, 1996; revised October 22, 1996; accepted December 12, 1996.)

Indonesian throughflow in a coupled general circulation model

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