Waworuntu2000

Recipe for Banda Sea water

by Jorina M. Waworuntu1, Rana A. Fine1, Donald B. Olson1 and Arnold L. Gordon2

ABSTRACTWater from the western Paci c ows through the Indonesian Seas following different pathways

and is modi ed by various processes to form the uniquely characterized isohaline Banda Sea Water.The processes contributing to the isohaline structure are studied using data from three ARLINDOcruises in 1993, 1994, and 1996.An inverse-modelanalysisusing salinity and CFC-11 data is appliedto a vertical section along the main path of ow, from theMakassar Strait to the Flores Sea and BandaSea. The model reproduces the seasonal and interannual variability of the throughow and showsreversals of ow in the vertical structure.The model solutions suggest strong baroclinic ows duringthe southeast monsoon of 1993 and 1996 and a small, more barotropic ow during the northwestmonsoon of 1994. The isohaline structure can be accounted for by isopycnal mixing of differentsource waters and by vertical exchanges, which are signi cant in this region. A downward uxequivalent to a downwelling velocity of 5 3 10 2 7 m/s is estimated for the section. The total balancealso suggests that seasonally and possibly interannually variable back ushing of water from theBanda Sea into the Straits contribute to the isohaline structure of Banda Sea water.

1. Introduction

The Indonesian Seas are the tropical link between the Paci c Ocean and the IndianOcean and are part of the global thermohaline circulation (Gordon, 1986). The circulationand transport through the Indonesian Seas have a large annual variation due to the strongmonsoonal change in the wind pattern in this area. The through ow has the maximumtransport during the south monsoon (northern summer), and the ow reaches its minimumin the north monsoon (northern winter) (e.g., Wyrtki, 1987; Murray and Arief, 1988;Meyers et al., 1996; Molcard et al., 1996). The change in the circulation pattern andvolume of the through ow must affect the different processes that occur in the IndonesianSeas. Dynamics of interaction with coastlines and shelves invigorate processes likeupwelling, downwelling, and internal wave action.

Changes in the ow pattern and transport variability in turn produce changes in watermasses and properties. During different monsoon seasons, waters from different sources ow to the Indonesian Seas, and the result can be seen in the variation of salinity,temperature, and other tracers from the surface down to the lower thermocline. Local

1. Rosenstiel School of Marine andAtmospheric Science, University of Miami, 4600 Rickenbacker Causeway,Miami, Florida, 33149,U.S.A. email: [email protected]

2. Lamont Doherty Earth Observatory, Columbia University, Palisades, New York, 10964,U.S.A.

Journal ofMarine Research, 58, 547569, 2000

547

processes such as precipitation, tidal mixing and coastal upwelling also contribute to thesechanges. Historical and recent data show the seasonal as well as interannual changes in thewater properties in the Indonesian Seas (e.g., Wyrtki, 1961; Boely et al., 1990; Zijlstra etal., 1990; Gordon and Fine, 1996; Ilahude and Gordon, 1996).The source of the through ow waters depends on the nonlinear properties of the

Mindanao Current and South Equatorial Current retro ection at the Paci c entrance to theIndonesian Seas and the tropical Paci c wind elds (Nof, 1996;Wajsowicz, 2000) (Fig. 1).There are two main pathways of the through ow in the Indonesian Seas, the western routeand the eastern route. The primary is water that ows through the western route which ismostly from the North Paci c. It ows throughMakassar Strait, Flores Sea, and Banda Seabefore exiting to the Indian Ocean via Timor Sea and Ombai Strait (Fine, 1985; F eld andGordon, 1992; Fieux et al., 1994, 1996; Gordon et al., 1994; Bingham and Lukas, 1995;Ilahude and Gordon, 1996; Gordon and Fine, 1996). Part of this water also exits to theIndian Ocean through Lombok Strait (Murray andArief, 1988). Prior to entering the BandaSea, water that takes the eastern route (east of Sulawesi Island) is mostly from the SouthPaci c (VanAken et al., 1988;Gordon, 1995; Ilahude and Gordon, 1996;Gordon and Fine,1996; Hautala et al., 1996). It ows via Halmahera and Seram Sea to the Banda Sea and

Figure 1. Station locations duringArlindoMixing 93 cruise (6August12 September 1993),ArlindoMixing 94 cruise (26 January28 February 1994), and Arlindo Circulation cruise (18 Novem-ber15 December 1996).

548 Journal of Marine Research [58, 4

Timor Sea. There is probably also a mixture of water coming from the North Paci cthrough the eastern route via the Maluku Sea. Most studies neglect the through ow fromthe western Paci c through the South China Sea, Karimata Strait and Java Sea because ofthe shallowness of the shelf (mean depth of Java Sea is around 40 m). There is also someevidence of ow from the Indian Ocean into the Banda Sea through the Straits of LesserSunda Islands. This ow from the Indian Ocean is mostly con ned to the surface layerduring boreal winter (Michida and Yoritaka, 1996).

As the water ows through the Indonesian sills and straits, the extremely complextopography and the combination of active monsoonal changes and higher frequencyforcing in the low latitude region contribute to momentum transfer and water masstransformation throughout the water column. Different studies show a large range ofvertical diffusivity estimates for the different basins in Indonesia. Gordon (1986) uses ahigh value of vertical diffusivity (Kz . 3 3 10 2 4 m/s) within the Indonesian thermoclineto account for the modi cation of the North Paci c thermocline temperature-salinitystructure. F eld and Gordon (1992), using a simple advection-diffusionmodel for averagebasin pro les, estimate a lower limit of vertical diffusivity, Kz at 1 3 10 2 4 m/s, fortransformation of Paci c waters into Indonesian water. These vertical diffusivities aregreater than those typically estimated for the interior ocean (e.g., Gargett, 1984; LedwellandWatson, 1991; Toole et al., 1994).

Hautala et al. (1996) examine isopycnal advection and mixing between North and SouthPaci c low latitude western boundary current sources to form Banda Sea water. Threeregimes are apparent in their analysis. The surface and upper thermocline layers, down tos u 5 25.8, required vertical mixing of surface precipitation and runoff. Vertical mixing isalso apparent below s u 5 27.0. In the lower thermoclinepurely isopycnal spreading gives asimple variation in the ratio of sources toward a larger South Paci c contribution withdepth.

Besides a large range of transport and diffusivity estimates, little is known about verticaldistribution of ow and mixing in the Indonesian Seas. Changes in the direction of thecurrents suggest that the ow is highly baroclinic. Change in direction and strength of thewind force the change in the surface Ekman ows, moving surface water toward oppositedirections. The build up of water along the boundaries change the pressure gradient acrossthe shallow sills surrounding this area and presumably force the baroclinic ow throughoutthe whole region.Vertical motion could also play an important role in the water distributionand transformation in this region. Lack of sufficient direct and long-term measurementsmake it difficult to get a complete picture of circulation and processes in the region,especially the vertical distribution of the circulation.

In this paper we will discuss some new observations from theArlindo Mixing cruises in1993 and 1994 (AM93, AM94) onboardKAL Baruna Jaya I andArlindo Circulation 1996cruise (AC96) onboard KAL Baruna Jaya IV. We will describe an inverse analysis forquantifying the mixing and advection parameters along the pathways of the through ow

2000] 549Waworuntu et al.: Recipe for Banda Sea water

that form Banda Sea water. This analysis makes use of Arlindo temperature, salinity andCFC-11 data.As more tracer data become available, different techniques are developed to quantify the

velocity eld and mixing parameters from the tracer distribution. One of the rst attemptsto obtain velocities from tracer distributionswas done by Fiadeiro and Veronis (1984). Leeand Veronis (1989) show that velocities and mixing coefficients can be obtained from aninversion of perfect tracer data, but inversion of imperfect data leads to substantial errors.Hogg (1987) obtains the ow elds and diffusivities on two isopycnal surfaces fromtemperature, oxygen and potential vorticity distribution of Levitus Atlas data (Levitus,1982).One way to solve this problem is by using different tracer distributions,with at least one

transient tracer. Wunch (1987, 1988) studies the transient tracer inverse as a problem incontrol theory and the regularization aspect of using transient tracers. Jenkins (1998)combines 3H-3He age with salinity, oxygen, and geostrophy to compute isopycnal diffusivi-ties, oxygen consumption rates, and absolute velocities for the North Atlantic subductionarea. Mc Intosh and Veronis (1993) show that an underdetermined tracer inverse problemcan be solved by spatial smoothing and cross validation.Here an inverse method similar tothat of McIntosh and Veronis (1993) is used to quantify the advective and diffusiveprocesses in the eastern Indonesian Seas.

2. Observations

a. Data

A total of 103 hydrographic stations were occupied duringAM93 cruise from 6 Augustto 12 September 1993 in the Southeast Monsoon (SEM). The AM94 cruise from 26January to 28 February 1994 during the Northwest Monsoon (NWM) has a similar stationdistribution (station 104209). Cruise AC96, including 37 hydrographic stations into theeastern Banda Sea were occupied from 18 November to 15 December 1996 (Fig. 1).The CTD data were collected by Lamont Doherty Earth Observatory (LDEO) group and

results from AM93 and AM94 have been reported before (Gordon and Fine, 1996; Ilahudeand Gordon, 1996; Mele et al., 1996). Chloro ourocarbon-11 (CFC-11) was measuredduring the three cruises by the University of Miami. The CFC analyses were done via acustom built extraction system and gas chromatograph similar to Bullister and Weiss(1988). The average analytical blanks for CFC-11 for theAM93,AM94, andAC96 cruisesare 0.001, 0.029, and 0.001 pmol/kg, respectively. The high blank levels especially at thebeginning of AM94 are due to high CFC concentrations in the laboratory air where thesamples were analyzed. The average bottle blanks for CFC-11 for the AM93, AM94, andAC96 cruises are 0.009, 0.060, and 0.002 pmol/kg, respectively. Improvements in theanalytical system and lower concentrations of CFC in the laboratory air during AC96cruise increased the quality of CFC measurements signi cantly. The precision of the CFCanalyses was estimated by drawing duplicate samples from the same Niskin bottle. For


measurement of CFC-11 greater than 0.1 pmol/kg, the averaged concentration-differencebetween duplicates in the eight to twenty replicates are 3.48%, 3.32%, and 0.32% for theAM93,AM94, andAC96 cruises, respectively.

b. Observed properties and circulation based on AM93, AM94, and AC96 data

The Banda Sea water is unique in its almost isohalinewater column at around 34.5 fromthe thermocline down to the bottom (see for example Gordon, 1986; F eld and Gordon,1992;Gordon et al., 1994;Gordon and Fine, 1996; Ilahude and Gordon, 1996; Fieux et al.,1996). Observations of water masses in the Banda Sea during the three Arlindo cruisesshow the same thermocline isohaline structure in the Banda Sea (Fig. 2). Lower salinitywater is observed at the Banda surface especially during AM93 in the SEM. At theintermediate depths the temperature-salinity (T/S) characteristics match the AntarcticIntermediateWater and so must come from the South Paci c.

Figure 2. (a) T/S diagram for stations in Sulawesi Sea, Northern Makassar Strait, Seram Sea, andBanda Sea duringAM93. Shown also are the NCEP/NCAR Reanalysis daily average surfacewindduring AM93, provided through the NOAA Climate Diagnostic Center (ftp.cdc.noaa.gov). (b)Same as (a) for AM94, with additional T/S diagram for stations in southern Makassar Strait andHalmahera Sea. (c). Station groups used for the T/S diagrams in (a) and (b).


i. Upper layer.All around Indonesia, the upper layer tracer distributionchanges seasonallyand interannually (Figs. 2, 3 and 5). DuringAM93 (SEM), North Paci c SubtropicalWater(NPSW) salinity maximum of 34.75 is observed at about 150 m depth in the MakassarStrait ( s u 5 24.5). This maximum is observed into the Flores Sea with diminishedintensity. It is not apparent in the Banda Sea. During AM94 (NWM), very low surfacesalinity is found in the southern Makassar Strait. The low salinity signal penetrates to over100 m depth. At this time, a very weak salinity maximum (34.6) is observed in Flores Seaat 150 m.The T/S curves in theMakassar Strait between s u 5 23 and s u 5 25.5 duringNWM lack

the NPSW salinity maxima and are similar to the Banda Sea T/S curves during SEM. Thereare four possible explanations for this lack of salinity maxima. (1) Qiu et al. (2000) statethat the salinity maxima input to the Indonesian Seas displays strong variability atintraseasonal time scales, resulting from eddy formations at the Mindanao retro ection.Gordon and Fine (1996) suggest that relaxation of the through ow during NWM allowsmore attenuation of the salinity maximum signal in the Makassar Strait. (2) Ilahude andGordon (1996) suspect a southward drift of Sulu Sea water wipes out the NPSW salinitymaxima causing a localized salinity minimum. (3) Upwelling in the frontal boundarybetween the out ow of Makassar Strait and current from the Flores Sea can contribute tothe properties of water in the south Makassar Strait as well (Ilahude, 1970; Hughes, 1982).(4) Here a back ushing of modi ed Banda Sea Water into the Makassar Strait is offered asanother possible explanation. The term back ushing as used here does not distinguishbetween horizontal advection and diffusion processes.The back ushing possibility is best addressed by using a transient tracer that can

differentiate temporal changes. In 1993 (SEM), a subsurface CFC-11 maximum isobserved just above the salinity maximum core at about 100 m ( s u 5 24.5) (Fig. 5).CFC-11 concentration is up to 1.8 pmol/kg at this depth, north of Dewakang sill in theMakassar Strait and decreases to about 1.4 pmol/kg in the Banda Sea at the same level.During NWM, the CFC-11 distribution follows the trend of salinity in the upper 150 m, thesubsurface maxima in the Makassar Strait is not observed; however, a weak CFC-11maximum with concentration of 1.7 pmol/kg is observed in the Flores Sea at 100 m depth.By examining the salinity distribution alone, one may conclude that the salinity maxima inthe thermocline is destroyed by vertical mixing of fresh water at the surface. LocalizedCFC mimima above the Flores maximum suggest that it cannot be a result of mixing downof surface water that is high in CFC. The strong strati cation suppresses such mixing.The salinity distributions during AC96 are similar to AM93 with the exception of a

lower surface salinity in the beginning of NWM, and a deeper thermocline by about 50 mcompared withAM93 andAM94 (Fig. 5). The latter may be a clue to interannual variations(F eld et al., 2000). In 1996, a more continuous subsurface CFC maximum is observedfrom the northernmost station in the Makassar Strait to Flores Sea at about 150 m depth.The concentration is 1.9 pmol/kg in Makassar Strait and decreases to 1.7 pmol/kg in thewestern Banda Sea. Higher CFC-11 concentrations compared to the previous cruises are


expected in the subsurface waters because of the increase in the atmospheric inputfunction. CFC-11 concentrations in the atmosphere increased until year 1993 and decreaseslightly after 1993 (Walker et al., 2000). Water below 50 m in this region, during AM93,AM94, and AC96, has been ventilated prior to 1993 and, therefore, increases in CFC-11concentration with time. Changes in the circulation pattern and ventilation rate due toseasonal and interannual variability can also alter the CFC concentration.Later, the inverseanalysis results will show that part of the increase is due to the change in the circulation.

Fresh water in ux at the surface is an important contributor to the T/S structure. Freshwater increases the stability of the upper layers. The major sources of fresh water are riverrunoff from islands plus precipitation over the ocean. Maximum fresh water in ow occursduring the NWM, when the wind brings the warm moist air from the South China Sea. Inthe NWM, there is a sea-surface salinity minimum in the Southern Makassar Strait,possibly advected from the Java Sea. The surface salinity minimum in Banda Sea duringthe month of AM93 cruise is suspected to be advected to the east from the previous NWMseason. Other possible sources of fresh water in the upper Banda Sea are river out owsfrom Irian Jaya (New Guinea; Wyrtki, 1961; Zijlstra et al., 1990). In the SEM, the surfacetemperature of the Banda Sea is about 4C cooler than in the NWM. The surface cooling isexpected to be the result of upwelling and local evaporative cooling processes when thewind is blowing from the southeast bringing dry air.

ii. Mid and lower thermocline. The mid and lower thermocline layers have a less variabletracer distribution. The North Paci c Intermediate Water (NPIW) salinity minimum isclearly observed in the Makassar Strait at s u 5 26.5 with salinity 34.45 during all threecruises (Figs. 2 and 5). The NPIW salinity minimum around 300 m is a permanent featurein both monsoon seasons. This suggests that the mid-depth circulation eld is moreconstant. The salinity minimum can be traced downstream to Flores Sea and all the way tothe Banda Sea with diminished intensity.

The increase in salinity at the NPIW level along the path from Makassar Strait to theFlores and Banda Seas can alternatively occur from vertical mixing or as a product ofisopycnalmixing between the low salinityNPIWwater with a saltier water of South Paci corigin. The latter conclusion is drawn from the Hautala et al. (1996) mixing analysis, wherepurely isopycnal mixing between North and South Paci c thermocline water can producethe observed Banda Sea water. Here CFCs are used to constrain the problem.

In the upper layer, CFC-11 partial pressure age is younger in the northernMakassar, agesincrease in the Flores and Banda Seas (Fig. 3). Along the section in the middle of the BandaSea, above 500 m, CFC-11 ages increase from south to north during AM93. CFC ages inthe Southern Banda Sea suggest that it is more recently ventilated and that the centralBanda Sea main thermocline is ventilated by ow from the Flores Sea. However, the trendchanges below 400 m, where younger ages of CFC-11 are observed in the Seram Sea andnorthern Banda Sea, as compared with the Flores Sea and Makassar Strait. Figure 4 shows


the CFC-11 pro les in the northern Makassar Strait, Banda Sea, and Seram Sea duringAC96. Below the subsurface maximum, the CFC-11 concentrations in the northernMakassar Strait decrease rapidly to near blank levels. In the Banda Sea the CFC-11 pro lehas no subsurface maxima, and the concentration decreases less with depth.Banda Sea lower thermocline and intermediate water is in uenced by South Paci c

water entering above the sills at the eastern entrance to the Banda Sea. The deepest sillconnecting the Paci c and the Banda Sea is the Lifamatola sill (1950 m) at 126 578 E,1 498 S, between Sulu Archipelago and Obi Island (Fig. 1). The Halmahera Sea sill depthis around 500 m. Seram Sea water is characterized by salty lower thermocline water thatcan only be derived from the South Paci c. In 1993 the salinity maximum is 34.7 at 250 mdepth (s u 5 26.0) at station 41. A more pronounced intrusion of South Paci c SubtropicalWater is found during the AM94 in the Halmahera and Seram Seas. The intrusion ischaracterized by a salinity maximum located on s u 5 26.0, with salinity up to 34.9(Ilahude and Gordon, 1996). CFC-11 pro les in the Banda Sea are similar to the pro les inSeram Sea. The scatter in the data at the low concentrationsmakes it difficult to determinethe trend of the CFC-11 age along the eastern route (Fig. 3). It is possible that somerecirculation occurs between the Seram Sea and the Banda Sea at this level.

Figure 3. CFC-11 derived ages on s u surfaces along different sections in the IndonesianSeas duringAM93,AM94, andAC96.


iii. Intermediate water. Below the Dewakang sill depth (550 m), at the end of MakassarStrait (118 488 E, 5 248 S), the CFC-11 concentrationsare low, but still show a clear trendbetween basins. In Makassar Strait, north of the sill, CFC-11 concentrations are near blanklevel at 1000 m and below. The presence of this old water suggests a blocked ow. Atdepths around 1000 m, the Banda Sea has de nitely higher concentrations than theMakassar water at the same depth.At this depth the Seram Sea has slightly higher CFC-11

Figure 4. CFC-11 pro les in the Makassar Strait (station 1-6), Seram Sea (station 22, 2426) andBanda Sea (station 2737) duringAC96 cruise.


concentrations as compared to the Banda Sea (Fig. 4). At these intermediate depths theconcentrations decrease from Banda Sea to the west (Flores Sea). The relatively highCFC-11 below 500 m in the Flores Sea in theAM94 data set have been discountedhere dueto blank problems at the beginning of the cruise.Both T/S and CFC pro les suggest the possibility of some back ushing of water from

the Banda Sea to Seram Sea, Flores Sea and Makassar Strait. There is also exchange withthe Java Sea for the top 50 m. These observations described in this section agree with thepattern of circulation illustrated byWyrtki (1961) and Van Bennekom (1988).

3. An inverse model for the passage to the Banda Sea

The tracer distribution at a point is a result of the advection and mixing history prior tothe observation. The result is also subject to the different boundary conditions for eachindividual tracer. The observed tracer distributions re ect the dynamics integrated overtime. If a system is stationary and there are enough linearly independent tracers, it ispossible to obtain an optimal solution for the circulation eld and mixing parameters thatcan produce the observed tracer eld. In most cases, there are not enough observations, orthe information in the tracer elds are not linearly independent, leading to an underdeter-mined problem. The Arlindo program provides sufficient data for a multi-tracer inverseanalysis for the region.The domain of interest is a two-dimensional vertical section along the ow path from

Makassar Strait to Flores Sea and Banda Sea (Fig. 5), for 1993, 1994, and 1996. It isassumed that these Straits and Seas act as a channel and the ow and tracer elds areuniform across the channel. The 2-D channel assumption does not ignore the input for theBanda Sea from the eastern channel. In fact, the model solutions show that there is owfrom the east in the lower thermocline during AM93 and AC96. The sections are griddedinto 200 km bins in horizontal direction and by 50 m in vertical direction. By excluding thetop 50 m of the water column, the Java Sea in uence and the surface exchange isminimized. The region below 650 m is not evaluated because the tracer eld is almostuniform there. When tracer gradients are parallel in a region, or gradients disappear, theproblem becomes degenerate and solutions are indeterminate (McIntosh and Veronis,1993). The model uses open boundary conditions. The top, bottom, northern and southernboundary conditions allow varying streamfunctions, so that exchanges with regionsoutside the domain are determined by the solutions of the model only.A quasi-steady stateis assumed for the tracer distributions for each cruise in turn. The tracers used in thecalculations are salinity and CFC-11 age. The information contained in the potentialtemperatures is used in the CFC-11 age calculation.The steady advective-diffusive equation for conservative tracer C can be written as

2 v. = C 1 = .(K = C ) 5 0. (1)


A steady advective-diffusive equation for an ideal tracer age T becomes

2 v. = T 1 = .(K = T ) 1 1 5 0 (2)

where factor one comes from a linear increase in age with time. Here the calculation isposed in terms of the two-dimensional streamfunction w , where u 5 2 w / z and w 5 w / x. Vertical diffusivity Kz is assumed uniform throughout the channel. There is nohorizontal diffusivity in the model. The terms chosen to balance the advective-diffusiveequation are similar to the ones used by F eld and Gordon (1992), where advection isbalanced by vertical diffusion. F eld and Gordon (1992) use a moving reference frame, sothe advection term appears as the local rate of change in tracer concentration.

For an m 3 n grid (see Fig. 6), there are (m 1 1)(n 1 1) unknowns for w . For k tracersthere will be k 3 m 3 n equations, in this case k 5 2 (salinity and CFC-11 age). For thecase where vertical diffusivity is equal to zero, the inhomogeneous terms needed to solvethe inverse problems come from the transient tracer, in this case the CFC-11 age.

In each grid, tracer concentrationsare evaluated at the center of the grid and streamfunc-tion w at the corners. The discrete form of the advective-diffusive equations in the inverse

Figure 5. Model input tracer distribution. The contour lines for salinity are every 0.05, for u every1C, and for CFC-11 every 0.1 pmol/kg, but the color shadings are irregular. The white dashedlines are sigma-theta contours. The map shows the location of the vertical section along theMakassar Strait, Flores Sea, to the Banda Sea.


calculation is:

12[(Ci 1 1, j 2 Ci, j 1 1)C ij 1 (Ci, j 2 1 2 Ci 1 1, j) C i, j 2 1

(Ci, j 1 1 2 Ci 2 1, j) C i 2 1, j 1 (Ci 2 1, j 2 Ci, j 2 1)C i 2 1, j 2 1] (3)

5 (Ci, j 1 1 1 Ci, j 2 1 2 2Ci, j)K* 1 tij

where

C 5 w /(D x D z)

K* 5 Kz /(D z)2

and tij equals zero for conservative tracers and minus one for age equations. Over the entiregrid, the difference equation can be written in matrix form as

Ax 5 b (4)

whereA is a (mnk 1 1) by (m 1 1)(n 1 1) coefficientmatrix, x is the unknownvector of Cand b is a vector of zeros and minus ones plus the diffusive terms. The one additionalequation augmented on the set of the tracer equations is for the arbitrary constant for thestreamfunction. This system of equations contains two kinds of errors, errors in thecoefficients that come from measurement error, and errors in the model.To solve this system of linear equations, the technique similar to that of McIntosh and

Veronis (1993) is used. This inverse method seeks the optimal solution for the advective-diffusive tracer problem by spatial smoothing and cross validation.

Figure 6. Grid used in the inverse calculation.Ci,j is the tracer, w is the velocity streamfunction.


a. Smoothing

This technique solves the system of equations Ax 5 b by balancing the solutionsmoothness, based on spatial derivatives of the solution, and the residual in the advective-diffusive equation. A penalty function, P, is formed as a weighted linear combination ofroughness and the residual equation e 5 Ax 2 b.

P 5 x Wxx 1 e Wee. (5)

Wx is a roughness matrix containing semi-norm operators and weights for the streamfunc-tion.

Wx 5 W w . (6)

W w is a derivative operator. When operated on the solution eld it gives the measure ofroughness. A smooth solution is obtained by minimizing the second derivatives (secondorder seminorm) for streamfunction.We is the error covariancematrix. In the calculation, anoncorrelated, uniformly weighted covariance matrix is used. It is a diagonal matrix withdiagonal elements that are inversely proportional to the relative measurement error forsalinity (ws) and CFC-11 age (wa). The numerical values of ws and wa are unimportant, itis only the size ofWe relative toW w that is important, and the relative size is controlled bythe smoothing parameter . The CFC measurements during AC96 are improved signi -cantly compared to the AM93 and AM94. In the calculation, different error weights, wa,are applied. For AC96 wa 5 1 year, and wa 5 10 years for AC93 and AC94. For salinityws 5 0.001 for all three observations.These weights are relative weights to the error in thestreamfunctions. The optimal solution (smooth solution consistent with observation) isobtained by minimizing P as a function of parameters.

b. Cross validation

To estimate , a parameter estimation technique known as the generalized cross-validation technique is used. The idea is to withhold one equation, solve theM 2 1 problem,and compare the withheld datum with the value predicted from the reduced problem(McIntosh and Veronis, 1993).A cross validation function V is formed.

V() 5M2 1e Wee

[M2 1trace(I 2 Rb)]2(7)

where Rb 5 AA 2 g and

A 2 g 5 [A WeA 1 Wx] 2 1A We

Rb is called the data resolutionmatrix.A 2 g is the generalizedmatrix inverse and x 5 A 2 gb.The entire minimization process is done by minimizing this single function V. The

minimization procedure used is the simplex minimization routine fmins in MATLAB.


c. Model results

The inverse calculationswere carried out over a range of vertical diffusivity coefficients,Kz. Figure 7 shows the ratio between the horizontal velocity, u, at distance 5 100 km anddepth 5 300 m to the input vertical diffusivity. Although this gure only shows the valuefrom one point in the section, the pattern is the same throughout the entire grid. A linearrelation between velocity and vertical diffusivity is obtained when Kz is larger than102 5 m2/s. When Kz is smaller than 10 2 5 m2/s, u/Kz is inversely proportional to Kz. Thisbehavior can be explained as follows. When Kz is large, the factor one in the age equationbecomes negligible. The balance between the terms in the advective-diffusive equation isthen between the advective term and the vertical diffusion term. When Kz is small, thebalance is between the horizontal advection and the vertical advection. A velocity-diffusivity ratio (u/Kz) that is inversely proportional to Kz, shows that u stays constant eventhough Kz decreases. This gives the limit of advection in the absence of diffusion. In otherwords, the solution for u remains the same for any diffusivity below Kz 5 10 2 5 m2/s.Figures 8 and 9 show the solutions of the inverse model for two extreme cases. The rst

one is for a no diffusivity case (Kz 5 0), and the second gure shows the solution for a largevertical diffusivity case (Kz 5 10 2 3 m2/s). Table 1 summarizes the parameters used and thesmoothingparameter of the optimal solution for each case. The no diffusivity case gives anunrealistic small advection and is shown mainly for comparison. The vertical diffusivity102 3 m2/s is chosen to give the solution in the range of observed horizontal velocities. Thesurface currents based on ship drift data in theMakassar Strait are of the order of a few tenscm/s (Mariano et al., 1995). Other diffusivities larger than 102 5 m2/s will give very similarvelocity pro les, scaled with the mixing coefficient used, as explained above. The dashedlines on the velocity pro les are the residuals between the optimal solution and theobserved tracer coefficient multiplied by the generalized inverse matrix. The solution x 5A 2 gb is equal to a vector of streamfunction on every grid. The streamfunction residual is

Figure 7. Ratio of the output horizontal velocity u at distance 5 100 km and depth 5 300 m to theinput vertical diffusivityKz.


calculated as xe 5 A 2 ge where e 5 Ax 2 b is the residual. Horizontal and vertical velocityresiduals are then obtained from xe. Except for the no diffusion case in AC96, all theresiduals are small compared with the optimal solution.

The solution of the inverse calculation is limited somewhat because of the nature of theminimization process. The model minimizes the residuals, and the residuals are minimumwhen Kz is small. Therefore the transport tends to be small for a given value of diffusivity.In a different experiment where Kz is taken as an unknown variable, the optimal solutiongives unrealistically small u and Kz.

4. Discussion

Results from the inverse calculation are limited by the accuracy in the data and theassumptions of the model. The model assumes steady state conditions, while the data aretaken from three synoptic cruises. This condition may apply assuming that the circulationadjustment time to monsoon change is short, which is generally the case in monsoon

Figure 8. (a) Horizontal velocity u for AM93, AM94, and AC96 sections from Makassar Strait toFlores Sea. Output from inverse calculationwith no vertical diffusivity (Kz 5 0). The dashed linesare the corresponding residual horizontal velocity. (b) Vertical velocity w for AM93, AM94 andAC96 sections from Makassar Strait to Flores Sea. Output from inverse calculation with novertical diffusivity (Kz 5 0). The dashed lines are the corresponding residual vertical velocity(positive upward).


regions in the world (Ray, 1984; Savidge et al., 1990; Liu et al., 1992). The ushing time ofthe Makassar Strait depends on the advective velocity. Examples of advective velocities inthe Makassar Strait are given by surface drifters. The surface drifter data are compiled bythe Drifting BuoyDataAssembly Center (DAC), Atlantic Oceanographic andMeteorologyLaboratory (AOML), NOAA, Miami. Four out of ve surface drifters found in theMakassar Strait entered from the Sulawesi Sea between February and May, but none ofthem entered during theArlindo cruises. These drifter velocities reach up to 1 m/s. At thesevelocities, it takes only 16 days to carry water and tracers as far as 1400 km. The velocity isexpected to be lower in the intermediate water. At velocities on the order of a few tens ofcentimeters per second, the ushing time can increase to on order two to three months.Because the ushing time is less than the semi annual period, the quasi steady stateassumption is still, however, applicable.The model does not allow Kz to vary spatially. The impacts of this assumption are

lessened because the model domain does not include the top mixed layer or the bottomfriction layer. This suppresses the effect of varying Kz due to turbulence in the surface and

Figure 9. (a) Horizontal velocity u for AM93, AM94, and AC96 sections from Makassar Strait toFlores Sea. Output from inverse calculation with uniform vertical diffusivity coefficient Kz 510 2 3 m2/s. The dashed lines are the corresponding residual horizontal velocity. (b) Verticalvelocity w for AM93, AM94 andAC96 sections from Makassar Strait to Flores Sea. Output frominverse calculation with uniform vertical diffusivity coefficientKz 5 102 3 m2/s. The dashed linesare the correspondingresidual vertical velocity (positive upward).


bottom layer. Vertical mixing above rough bottoms and sills indeed can be larger (e.g.,Polzin et al., 1996, 1997; Van Bennekom et al., 1988). Adding these spatially varying Kzterms in the model will change the solutions, but havingmore unknown terms in the systemof equations will make the system more underdetermined, and a solution might not beobtained at all.

Some dynamical terms like horizontal diffusion are not represented in the model. Sincehorizontal advection is the only horizontal process represented in the model, it has to beinterpreted as a combination of all processes that carry or distribute tracers horizontally.Likewise, in the case where Kz 5 0, vertical advection is the only way to distribute thetracers vertically. Assuming a channel-like geometry means that the velocity eld in themodel solutions also has to compensate for ows across channel that might occur in reality,such as the Lombok out ow and Java Sea exchange.

The errors in the data are accounted for in the error covariance matrix that contains theweights for measurement errors for CFC-11 and salinity.With these weights, the size of theresidual equation turned out to be relatively small compared with the optimal solution.Thisshows that, despite the assumptions in the model, the simple dynamics chosen canreproduce the observed tracer elds with relatively small error.

The optimal solutions of the inverse calculation suggest that strong baroclinic owsoccur duringAM93 and AC96. On the other hand, a small and more barotropic ow is theoptimal solution obtained for the tracer distribution in AM94. The general ow issouthward and eastward, and sometimes accompanied by a reversal in ow. These owpatterns in the thermocline have different directions as compared with the surface currentsthat usually follow the local wind direction (Figs. 2 and 10). During SEM, the surfacecurrent in the Flores Sea is westward, but the solutions for AM93 andAC96 show a strongeastward ow in the upper thermocline. During NWM, the surface current is eastward,while theAM94 model gives a weak eastward ow in the upper thermocline.

There are two possible explanations for the difference between the observed surfacecirculation and the solutions given by the model for subsurface ows. First, tracer

Table 1. Input parameters and smoothingparameters 5 em for no diffusionand large diffusion case. is obtained by minimizing the generalized cross validation functionV. Shown also are the meanoptimal u and w for each case.ws and wa are the weights for error in salinity and age coefficients,respectively.

Kv(m2/s) ws wa m* V ()

Mean u(102 2 m/s)

Mean w(102 7 m/s)

AM93 0 102 3 10 1.929 1.144 0.046 2 3.55AM94 0 102 3 10 2.615 1.297 2 0.019 2 2.90AC96 0 102 3 1 2.304 4.699 2 0.006 2 5.28AM93 102 3 102 3 10 2 0.350 5709 7.91 57.9AM94 102 3 102 3 10 2.480 10841 1.25 74.4AC96 102 3 102 3 1 0.037 23536 6.62 32.5

*( 5 em)


distributions re ect the circulation prior to observation. The model result may re ect thecirculation during the opposite season of the measurement. This could explain the opposite ow in the thermocline as compared with the surface current and wind direction.A more likely explanation for the difference between observed surface circulation and

model subsurface ows is that the ows in the deeper thermocline are not directly drivenby the local wind. During SEM, upwelling occurs in theArafura and the eastern Banda Sea.The westward- owing surface current induces coastal upwelling along the eastern bound-ary. The upwelled water is supplied by the thermocline, and so strengthens the eastward ow in the thermocline, as observed in the solutions for AM93 and AC96. During NWM,the surface circulation in the eastern Banda Sea is downwelling favorable, pushing thewater westward in the lower thermocline. A westward ow from the Banda Sea to theFlores Sea is required for the AM94 solution (Fig. 10). Recent current meter data in theMakassar Strait indicates that for the boreal winter the surface layer is moving opposite tothe thermocline ow (Gordon et al., 1999). It is interesting to note that the solutions for theNWM vertical velocities show weakest downwelling in the no diffusivity case or strongestupwelling in the large diffusivity case (Table 1), supporting the type of circulationdescribed above.Flow to the west is also part of the solution for AC96 in the lower thermocline and is

more pronounced when the vertical mixing is small. A secondary downwelling cell mayoccur below the upwelling cell during SEM. The idea that the ow to the east in the upperthermocline is driving a secondary cell with a reversal in ow in the lower thermocline issuggested in Gordon et al. (1994). Here it is shown that the reverse ow can extend to theupper thermocline, like in theAM94 case, and vary seasonally and interannually.The back ushing of water from the Banda Sea to the Makassar Strait through the Flores

Sea could contribute to the similarity observed in the T/S curve in the Makassar Straitduring AM94 as compared to the Banda Sea. The salinity minimum in the Flores Sea isclearly coming from the North Paci c Intermediate Water, which ows through theMakassar Strait. The water below NPIW could have more variable sources. Lowerthermocline water in the Flores Sea is affected by sources from the east, through thepassages east of Sulawesi. Higher CFC-11 concentrations in the deep Banda Sea than whatare found at the comparable depths in the Makassar Strait suggest a South Paci ccomponent. The exact route and ratio of North and South Paci c sources is undetermined.Considering the longer residence time of water in the Banda Sea (F eld and Gordon,1992), a modi ed Banda Sea water, that is a mixture of North Paci c, South Paci c andBanda Sea water from previous seasons, appears to be ushed back into the MakassarStrait from time to time. Recent observations of current meter data in the Makassar Straitfrom December 1996 to July 1998 show some evidence of reversal of the ow in thecurrent meter at 350 m and deeper (Gordon et al., 1998; Gordon and Susanto, 1999). Thecurrent meter at 240 m does not show a ow reversal. A shorter record of ADCP data (1December 19979 March 1997) from the same location shows a maximum southward ownear 110 m depth, and a variable northward and southward ow in the top 60 m (Gordon et


al., 1999). Thus, although from the property distributions and model an advective ordiffusive back ushing cannot be distinguished, the current meter data are consistent withan advective back ushing process.

The changes in the ow character are caused by the changes in the pressure eld alongthe channel. The pressure changes because of the change in the local wind direction andintensity, and the changes in the remote pressure eld. The solutions suggest that thechanges vary with depth. The exact mechanism of how and why these changes occurcannot be determined by an inverse model.A diagnostic or process study is needed to infer

Figure 10. Schematic diagram of the circulation in the Flores Sea and Banda Sea during NWM andSEM. The vertical cross-section(right panel) runs fromwest to east in the Java Sea to Flores Sea tothe eastern Banda Sea. The Makassar Strait input is shown as a cross or dot symbolizing ow intoor out of the Strait. The velocity pro le (left panel) is a pro le approximately in the Flores Sea orwestern Banda Sea. This would correspond to the velocities at the end of the model domain. Themodel results indicate that the absolute velocities depend on the diffusivity chosen, and for thisreason the u axis origin (vertical line on the velocity pro le) does not necessarily lie at the zerovelocity.The scale of the diagram is exaggerated to illustrate the seasonal circulationat the surfaceand in the thermocline.


the forcing and their responses. The seasonal ow variability given by the model solutionsupports the observation that the through ow is minimum during the northwest monsoonseason. In the large diffusivity case, the horizontal velocity for theAM94 (NWM) is about15% to 20% of the velocity in the AM93 (SEM) and AC96 (end of SEM). Interannualvariability may play an important role in the variability of the ow. The three observationswere in different ENSO phases, the 1993 and 1994 cruises occur during a prolonged ElNino period, and the 1996 cruise was in a La Nina period. The solutions show a relativelystronger ow to the east in the upper thermocline during AC96 as compared with AM93.This supports the argument that during a La Nina phase, the build up of the warmpool inthe western Paci c will strengthen the through ow (Gordon et al., 1999).While the patternof the circulation in the Indonesian Seas is strongly dependent on the local forcing and thedistributionof boundaries, the strength of the through ow seems to be more closely relatedto the pressure difference between the remote Paci c and the Indian Ocean.The model solutions show that the upper thermocline eastward ow in 1996 is also

accompanied by a reverse ow to the west in the lower thermocline. The strengthening ofthe pressure gradient in the lower thermocline between the Paci c Ocean and theIndonesian Seas during a La Nina phase could increase the ow in the deeper thermoclinethrough the eastern passages. Since the sill depths in the eastern passages are much deeperthan the Makassar Strait sill depth, an increase in the pressure gradient in the lowerthermocline will increase the through ow from the eastern passages.Mean vertical velocities of the solutions range from 2 5 3 10 2 7 m/s to 74 3 10 2 7 m/s.

Although the values are not unrealistic, the pattern depends on the size of the verticaldiffusivity chosen. In the case where vertical diffusivity is small, the downward verticaladvection is needed to satisfy the advective-diffusive balance, and when Kz is large, theoptimal vertical velocity is mostly upward in the upper thermocline. Whichever processactually occurs, whether it is large vertical mixing or downward advection, the modelresults suggest a generally downward ux in this area. Vertical exchange is an importantprocess that carries the heat and tracers downward. Downward movement in one placemust be accompanied by upwelling somewhere else. Upwelling brings cooler andnutrient-rich water to the surface. Amean downward velocity of 0.5 3 102 6 m/s will carryabout 42 W/m2 of heat over a 20C temperature gradient from the surface to the lowerthermocline.This study shows a marked change in the through ow volume and character seasonally

and interannually. The inverse model is also able to estimate the horizontal and verticalexchange processes that produce the observed water masses and are important to the localweather patterns and productivity.

Acknowledgments. JMW would like to thank Drs. A. Mariano and C. Thacker for valuablediscussions on the inverse method. RAF acknowledges support by National Science Foundationgrants OCE 9302936 and 9529847. The research of ALG is funded by NSF: OCE 9529648 and theOffice of Naval Research: N00014-98-1-0270.DBO is supported by JPL/NASA: 960606 and NSF:OCE 9314447. We acknowledge Kevin Sullivan and Kevin Maillet of RSMAS, and Muswerry


Muchtar and Suseno of LIPI for the CFC analysis. Gratitude is extended to Dr. A. Gani Ilahude ofLIPI and Dr. Indroyono Soesilo of BPPT for organizing and promoting the ARLINDO joint project.We thank the captain and crews of Baruna Jaya I and IV for providing professional support for ourscienti c mission.

REFERENCES

Bingham, F. M. and R. Lukas. 1995. The distribution of intermediatewater in the western equatorialPaci c during January-February1986. Deep-Sea Res., 42, 15451573.

Boely, T., J. P. Gastellu-Etchegorry, M. Potier and S. Nurhakim. 1990. Seasonal and interannualvariations of the sea surface temperatures (SST) in the Banda and Arafura Sea area. Neth. J. SeaRes., 25, 425429.

Bullister, J. L. and R. F. Weiss. 1988. Determination of CCl3F and CCl2F2 in seawater and air.Deep-Sea Res., 35, 839853.

F eld,A. andA. L. Gordon. 1992.Vertical mixing in the Indonesian thermocline.J. Phys. Oceanogr.,22, 184195.

F eld, A., K. Vranes, A. L. Gordon, R. D. Susanto and S. L. Garzoli. 2000. Temperature variabilitywithin Makassar Strait. Geophys. Res. Lett., 27, 237240.

Fiadeiro, M. E. and G. Veronis. 1984. Obtaining velocities from tracer distributions. J. Phys.Oceanogr., 14, 17341746.

Fieux, M., C. Andrie, E. Charriaud, A. G. Ilahude, N. Metzl, R. Molcard and J. C. Swallow. 1996.Hydrological and chloro uoromethanemeasurements of the Indonesian throughow entering theIndianOcean. J. Geophys. Res., 101, 1243312545.

Fieux, M., C. Andrie, P. Delecluse, A. G. Ilahude, A. Kartavtseff, F. Mantisi, R. Molcard and J. C.Swallow. 1994. Measurements within the Paci c-Indian oceans throughow region. Deep-SeaRes., 41, 10911130.

Fine, R. A. 1985. Direct evidence using tritium data for throughow from the Paci c into the IndianOcean. Nature, 315, 478480.

Gargett,A. E. 1984.Vertical eddy diffusivity in the ocean interior. J. Mar. Res., 42, 359393.Gordon,A. L. 1986. Interoceanexchangeof thermoclinewater. J. Geophys. Res., 91, 50375046. 1995.When is appearance reality?A comment on why does the Indonesian throughow appear

to originate from the North Paci c. J. Phys. Oceanogr., 25, 15601567.Gordon, A. L., A. F eld and A. G. Ilahude. 1994. Thermocline of the Flores and Banda Sea. J.

Geophys. Res., 99, 1823518242.Gordon,A. L. and R. A. Fine. 1996. Pathways of water between the Paci c and Indian oceans in the

Indonesian seas. Nature, 379, 146149.Gordon,A. L. and R. D. Susanto. 1999. Makassar Strait transport: Initial estimate based on Arlindo

results. Marine TechnologySociety, 32, 3445.Gordon, A. L., R. D. Susanto and A. F eld. 1999. Throughow within Makassar Strait. Geophys.

Res. Lett., 26, 33253328.Gordon,A. L., R. D. Susanto,A. F eld and D. Pillsbury. 1998.Makassar Strait transport: Preliminary

Arlindo results fromMAK-1 and MAK-2. InternationalWOCE Newsletter, 33, 3032.Hautala, S. L., J. L. Reid and N. Bray. 1996. The distribution and mixing of Paci c water masses in

the IndonesianSeas, J. Geophys. Res., 101, 1237512389.Hogg, N. G. 1987. A least-squares t of the advective-diffusive equations to Levitus Atlas data. J.

Mar. Res., 45, 347375.Hughes, R. L. 1982. On a front between two rotating ows with application to the Flores Sea. Dyn.

Atmos. Ocean, 6, 153176.


Ilahude, A. G. 1970. On the occurrence of upwelling in the southern Makassar Strait. Mar. Res. inIndonesia,10, 353.

Ilahude, A. G. and A. L. Gordon. 1996. Thermocline strati cation within the Indonesian Seas. J.Geophys. Res., 101, 1240112409.

Jenkins, W. J. 1998. Studying subtropical thermocline ventilation and circulation using tritium and3He. J. Geophys. Res., 103, 1581715831.

Ledwell, J. R. and A. J. Watson. 1991. The Santa Monica Basin tracer experiment: A study ofdiapycnal and isopycnalmixing. J. Geophys. Res., 96, 86958718.

Lee, J. H. and G. Veronis. 1989.Determiningvelocities and mixing coefficients from tracers. J. Phys.Oceanogr., 19, 487500.

Levitus, S. 1982. ClimatologicalAtlas of the World Ocean, NOAA Professional Paper 13, NationalOceanic andAtmosphericAdministration,Rockville,Maryland, 173 1 xv pp.

Liu, K. K., G. C. Gong, C. Z. Shyu, S. C. Pai, C. L. Wei and S. Y. Chao. 1992. Response of Kuroshioupwelling to the onset of the northeast monsoon in the sea north of Taiwan: Observations and anumerical simulation. J. Geophys. Res., 97, 1251112526.

Mariano,A. J., E. H. Ryan, B. D. Perkins and S. Smithers. 1995. The Mariano global surface velocityanalysis 1.0. U.S. Coast Guard Technical Report, CG-D-34-95.

McIntosh, P. C. and G. Veronis. 1993. Solving underdetermined tracer inverse problems by spatialsmoothing and cross validation. J. Phys. Oceanogr., 23, 716730.

Mele, P. A., B. A. Huber, A. L. Gordon, A. G. Ilahude, K. Sullivan and J. Marra. 1996. ArlindoMixing: CTD and hydrographic data from the August 1993 and January 1994 cruises, TechnicalReport: LDEO-96-6, LDEO, Palisades, NewYork.

Meyers, G. 1996. Variation of Indonesian throughow and the El NinoSouthern Oscillation. J.Geophys. Res., 101, 1225512263.

Michida,Y. and H.Yoritaka. 1996. Surface currents in the area of the Indo-Paci c throughow and inthe tropical IndianOcean observedwith surface drifters. J. Geophys. Res., 101, 1247512482.

Molcard, R., M. Fieux andA. G. Ilahude. 1996. The Indo-Paci c throughow in the Timor Passage.J. Geophys. Res., 101, 1241112420.

Murray, S. P. and D. Arief. 1988. Through ow into the Indian Ocean through the Lombok Strait,January 1985January 1986. Nature, 333, 444447.

Nof, D. 1996. What controls the origin of the Indonesian throughow? J. Geophys. Res., 101,1230112314.

Polzin, K. L., K. G. Speer, J. M. Toole and R. W. Schmitt. 1996. Intense mixing of Antartic BottomWater in the equatorialAtlantic Ocean. Nature, 380, 5457.

Polzin, K. L., J. M. Toole, J. R. Ledwell and R. W. Schmitt. 1997. Spatial variability of turbulentmixing in the abyssal ocean. Science, 276, 9396.

Qiu, B., M. Mao andY. Kashino. 1999. Intraseasonalvariability in the Indo-Paci c Through ow andthe regions surrounding the IndonesianSeas. J. Phys. Oceanogr., 29, 15991618.

Ray, T. K. 1984. Response of the upper layers of the Arabian Sea and the Bay of Bengal to thesouthwest monsoon.Mausam, 37, 103106.

Savidge, G., J. Lennon andA. J. Matthews. 1990.A shore-based survey of upwelling along the coastof Dhofar region, southernOman. Cont. Shelf Res., 10, 259275.

Toole, J. M., K. L. Polzin and R. W. Schmitt. 1994. Estimates of diapycnal mixing in the abyssalocean. Science, 264, 11201123.

Van Aken, H. M. Pujanan and S. Saimima. 1988. Physical aspects of the ushing of the EastIndonesianBasin. Neth. J. Sea Res., 22, 315339.

Van Bennekom, A. J., Kastoro and W. C. Patzert. 1988. Recirculation in the Banda throughow,traced with dissolved silica. Neth. J. Sea Res., 22, 355359.


Wajsowicz,R. C. 1999.Variations in gyre closure at the water-mass crossroadsofWesternEquatorialPaci c Ocean, J. Phys. Oceanogr., 29, 30023024.

Walker, S. J., R. F. Weiss and P. K. Salameth. 2000. Reconstructed histories of the annual meanatmospheric mole fractions for the halocarbons CFC-1, CFC-12, CFC-113 and carbon tetrachlo-ride. J. Geophys. Res., 105, 1428514296.

Wunsch, C. 1987. Using transient tracers: the regularizationproblem.Tellus, 39B, 477492. 1988. Transient tracers as a problem in control theory. J. Geophys. Res., 93, 80998110.Wyrtki, K. 1961. Physical oceanography of the South East Asian Waters. NAGA Report Vol 2,

Scripps Institute of Oceanography,195 pp. 1987. Indonesian throughow and the associated pressure gradient. J. Geophys. Res., 92,

1294112946.Zijlstra, J. J., M.A. Baars, S. B. Tijssen, F. J. Wetsteyn, J. I. J.Witte,A. G. Ilahudeand Hadikusumah.

1990. Monsoonal effects on the hydrography of the upper waters (, 300 m) of the eastern BandaSea and northern Arafura Sea, with special reference to vertical transport processes. Neth. J. SeaRes., 25, 431447.

Received: 3 July, 1999; revised: 3 May, 2000.


Waworuntu2000

Documents

Transcript of Waworuntu2000