Physical processes contributing to the water masstransformation of the Indonesian Throughflow

Ariane Koch-Larrouy & Gurvan Madec &

Daniele Iudicone & Agus Atmadipoera & Robert Molcard

Received: 11 April 2008 /Accepted: 6 October 2008# Springer-Verlag 2008

Abstract The properties of the waters that move from thePacific to the Indian Ocean via passages in the Indonesianarchipelago are observed to vary with along-flow-pathdistance. We study an ocean model of the Indonesian Seaswith reference to the observed water property distributionsand diagnose the mechanisms and magnitude of the watermass transformations using a thermodynamical methodol-ogy. This model includes a key parameterization of mixingdue to baroclinic tidal dissipation and simulates realisticwater property distributions in all of the seas within thearchipelago. A combination of air–sea forcing and mixingis found to significantly change the character of theIndonesian Throughflow (ITF). Around 6 Sv (approxi-mately 1/3 the model net ITF transport) of the flow leavesthe Indonesian Seas with reduced density. Mixing trans-forms both the intermediate depth waters (transforming4.3 Sv to lighter density) and the surface waters (madedenser despite the buoyancy input by air–sea exchange,net transformation=2 Sv). The intermediate transforma-tion to lighter waters suggests that the Indonesiantransformation contributes significantly to the upwellingof cold water in the global conveyor belt. The mixinginduced by the wind is not driving the transformation. Incontrast, the baroclinic tides have a major role in thistransformation. In particular, they are the only source ofenergy acting on the thermocline and are responsible forcreating the homostad thermocline water, a characteristic

of the Indonesian outflow water. Furthermore, they coolthe sea surface temperature by between 0.6 and 1.5°C, andthus allow the ocean to absorb more heat from theatmosphere. The additional heat imprints its characteristicsinto the thermocline. The Indonesian Seas cannot only beseen as a region of water mass transformation (in thesense of only transforming water masses in its interior) butalso as a region of water mass formation (as it modifies theheat flux and induced more buoyancy flux). This analysisis complemented with a series of companion numericalexperiments using different representations of the mixingand advection schemes. All the different schemes diagnosea lack of significant lateral mixing in the transformation.

Keywords Indonesian Throughflow .Water masstransformation . Neutral density framework .

Thermodynamic . Tidal mixing processes

1 Introduction

The Indonesian Throughflow (ITF) is the only low latitudepassage between two oceans and thus plays a key role inthe global circulation (Hirst and Godfrey 1993; Wasjowiczand Schneider 2001). In addition, it is a region of strongtransformation within the thermocline (Ffield and Gordon1996; Hautala et al. 1996; Koch-Larrouy et al. 2007,hereafter KL07, see Gordon 2005 for a review). The ITFcools and freshens the saltier and warmer Pacific thermo-cline (Fig. 1). In the mid-thermocline, the Pacific inflowhas a salinity of about 34.9 psu and a temperature of about23°C. Whereas within the Indonesian archipelago, thesalinity decreases to 34.4 psu and the temperature reaches21.5°C (Fig. 1). This transformation is the result of strongdianeutral fluxes that modify the stratification (Ffield and

Ocean DynamicsDOI 10.1007/s10236-008-0154-5

Responsible editor: Tony Hirst

A. Koch-Larrouy (*) :G. Madec :D. Iudicone :A. Atmadipoera : R. MolcardLaboratoire d’Océanographie et du Climat: Expérimentations etApproches Numériques (LOCEAN),Paris, Francee-mail: Ariane.Koch-Larrouy@locean-ipsl.upmc.fr

Gordon 1996; Gordon 2005). The associated dianeutralmixing is mainly driven by wind and tides, which combinethe North and South Pacific waters and mix downward thebuoyancy resulting from the regional air–sea heat andfreshwater flux (Gordon 2005). The resulting relatively coldand fresh thermocline water is projected into the IndianOcean, eventually reaches the Agulhas Retroflection, and ishence perhaps of relevance to the Atlantic Ocean’s merid-ional overturning circulation (Gordon 2005).

A recent study (KL07) noted that the Indonesianarchipelago (including the Philippine Seas and the SouthChina Sea) is the only region of the world where stronginternal tide generation occurs in a semi-enclosed basin. Asa result, all of the baroclinic tidal energy remains trappedlocally inside the archipelago. KL07’s study implemented atidal parameterization, adapted to the specificities of theIndonesian archipelago. Introduced in an Oceanic GeneralCirculation Model, this parameterization allows the modelto better represent the properties of the water mass

evolution in each sub-basin, in good agreement with theobservations. In particular, the cold and fresh tongue ofIndonesian water in the Indian Ocean is well reproduced.Vertical diffusivity is large with an average of 1.5 cm2/s,which is close to the independent estimates inferred fromobservations (1–2 cm2/s, Ffield and Gordon 1992). Thissuggests that the total energy input provided by the tidalparameterization has the right order of magnitude. Further-more, one of the conclusions of the KL07 study is that theinternal tides are a key ingredient of the Indonesian watermass transformation. Nevertheless, it may not be the onlyone. In the present paper, we propose to quantify the totaltransformation and the relative contribution of all theprocesses in this transformation as well as their horizontaland vertical structure. To this end, we will use the twonumerical simulations of KL07 with and without tides andthe thermodynamic diagnostic tool of Iudicone et al.(2008a) that generalizes the classical approaches forestimating surface mass fluxes and sub-surface mixing.

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TREF SREF

SWODB

T TIDES

T WODB

STIDES

Fig. 1 Salinity (left) and tem-perature (right) on the neutraldensity surface γn=24 for theWODB and some data added byAtmadipoera at al. (2008) (toppanel), TIDES run (with tidalmixing, middle panel) and REFrun (without tidal mixing, bot-tom panel)

Ocean Dynamics

This diagnostic tool allows us to quantify the role of eachcomponent in the physics of the ITF water mass transfor-mation in a numerical model.

The paper has been divided into sections as follows. InSection 2, the model configuration and the method used toestimate of the water mass transformations are respectivelypresented. The main characteristics of the model results arethen shown in Section 3. The role of surface buoyancyfluxes and of mixing in creating/destroying water massesare presented in Section 4. Finally, Section 5 deals with theconclusions and the discussion.

2 Methodology

2.1 Model

The model configuration is the same as in Koch-Larrouyet al. (2007). It is a sub-domain of the global ocean modelORCA025 (Barnier et al. 2006). It is based on the latestversion of the Nucleus of European Model of the Ocean(Madec 2008; Madec et al. 1998). Within the ITF region,the model has a 1/4° Mercator mesh and 46 levels in thevertical, with thicknesses from 6 m at the surface to 250 mat the bottom. A partial step representation of bathymetryis used. The domain extends from 95°E to 145°E inlongitude and from 25°S to 25°N in latitude. We use theTréguier et al. (2001) open boundary condition algorithmwhich uses both a radiation condition and a relaxation toclimatology. Open boundary conditions are provided bythe reference global simulation ORCA025 at the sameresolution (Barnier et al. 2006). Also, the same climato-logical surface forcing as the global simulation is used.The solution obtained with the zoom configuration is verysimilar to the global one. In particular, the total transport isvery similar as well as its repartition from the eastern andwestern routes. In terms of the water mass structure, wefound the same T–S structure including the same bias: theSouth Pacific Subtropical Water (SPSW) salinity maxi-mum is too salty, presumably due to an excess ofevaporation at its formation site in the subtropical gyre.The purpose of our study is, using a model with referenceto the observed water property distributions, to diagnosethe local water mass transformation. This implies that theentrance water masses should be as close as possible to theobserved water masses. We choose to force the openboundaries with the observed climatological T–S structure(Levitus et al. 1998) in order to have relatively goodentrance properties and to focus on the local transforma-tion. The entrance velocity is still prescribed by the globalsimulation. Instabilities that are generated by the non-geostropic constraint at the boundaries disappear afterthree grid points and the stable solution is satisfactory,

with identical transport and good water mass properties atthe entrance.

The Indonesian archipelago is a series of complextopographic features and very sharp sills. In such a region,classical models often have problems due to the poorlyresolved or smoothed bathymetry which leads to a badestimation of the distribution of the transports and hence thesource of water masses (Gordon and Mc Clean 1999;Kamenkovich et al. 2003; Potemra et al. 1997). In thisstudy, each barrier or sill has been carefully preserved byadjusting the depth and the width of each modelledpassage.1 The first barrier encountered by Pacific waterscoming towards Makassar Strait is the 1,350-m-deepSangihe Ridge (Fig. 2), providing access to the SulawesiSea. The 680-m-deep Dewakang Sill (580 m in the model)separating the southern Makassar Strait from the Flores Seais an even more formidable barrier. Along the eastern path,Pacific water must flow over the 1,940-m barrier (1,800 min the model) of the Lifamatola Passage before passing intothe deep levels of the Seram and Banda Seas. The deepestbarrier encountered by both the western and eastern paths tothe Indian Ocean is the 1,300-m sill of the Sunda Arc nearTimor. The Savu Sea is connected to the Banda Sea downto a depth of 2,000 m (Ombai Strait, 1,800 m in the model)but is closed to the Indian Ocean at depths shallower than1,200 m. The shallowest barrier is the 350-m-deep (210 min the model) Lombok Strait. Lombok and Ombai Straitsare very narrow passages only 25 km wide and 14 km wideat depths greater than 100 m. However, the grid cell is25 km wide. Partial cells have been used to reduce thewidth of these straits (Madec 2008). The strait depth andwidth adjustments, in these key passages of the model,enable us to have the correct water masses in the differentbasins and a relatively good transport distribution.

Another key point to correctly model the water masstransformation in this region is to take into account themixing associated with internal tide dissipation (KL07).The recently developed KL07’s parameterization of tidalmixing over the semi-enclosed seas of the Indonesianarchipelago is used. It is based on the St Laurent et al.(2002) parameterization:

ktides ¼ qΓE x; yð ÞF zð ÞrN2

ð1Þ

where Γ=0.2 is the mixing efficiency, N the Brunt–Väisäläfrequency, and ρ the density.

Vertical diffusivity is parameterized as a function of theenergy transfer between the barotropic and baroclinic tidesE(x, y), and involves a distribution of this energy in thevertical F(z) and the fraction of the energy to be dissipated

1 Adjustment of the horizontal resolution to preserve the section in onegrid point strait such as Lombok.

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q. These three parameters have been adjusted to theIndonesian region. E(x, y) is provided from a tidal model(Lyard and Le Provost 2002) that ensures the right amountof energy input and its correct horizontal structure. q isequal to 1 since the Indonesian archipelago is a semi-enclosed basin that forces the internal tides to completelydissipate inside each semi-enclosed sea. F(z) is inferredfrom a 2D internal tide generation model (Gerkema et al.2004) that has been run with appropriate parameters forthe main passages of the Indonesian archipelago. F(z) isfound to scale as N2 above the core of the thermocline andas N below. Consequently, the profile of the energydissipation has a maximum in the thermocline. Theremaining vertical physics comes from the turbulentclosure scheme, forced by the wind and ocean shear(Tke parameterization, Blanke and Delecluse 1993), andthe increase of vertical mixing when N2 becomes negative(static instability). The background vertical diffusivity is0.1 cm2/s. The horizontal advection scheme is a third orderupstream biased scheme (UBS; see Madec 2008 for moredetails).

The model is forced by daily climatological wind stressfields derived from the weekly gridded European RemoteSensing satellite fields during the period 1992–2001.Surface heat fluxes and evaporation are computed withbulk formulae using climatologies from the NationalCenters for Environmental Prediction and the National

Center for Atmospheric Research reanalysis for surface airtemperature, specific humidity, and an observational clima-tology of cloud cover. The surface fresh water flux isderived from the Climate Prediction Centers MergedAnalysis of Precipitation to which is added a weakrelaxation to the surface salinity of Levitus et al. (1998).The simulation has been performed over an 11-year periodstarting from Levitus T–S at rest. After 11 years, the modeland its water masses are in equilibrium over a seasonalcycle (cf. KL07). The final year is used in this study and istaken as a climatological year.

We will analyze the same simulations as in KL07, calledTIDES and REF (with and without the tidal parameteriza-tion). A sensitivity study of the advection scheme isperformed in order to quantify its associated spuriousmixing (Section 4.2).

2.2 Approach used to diagnose the water masstransformation

The roles of the mixing and forcing in the transformation ofwater masses are quantified using the thermodynamicalmethod of Iudicone et al. (2008a). It is a generalization ofWalin’s (1982) approach to neutral surfaces which takesinto account solar radiation penetration. This method hasbeen fruitfully applied to the Southern Ocean (Iudicone etal. 2008b). It evaluates each mixing and forcing contribu-

Banda Sea

Sulawesi Sea

Flores Sea

Sulu Sea

Maluku Sea

Savu Sea

Seram Sea

Makassar Strait

Sangihe Ridge

Java Sea

Halmahera Sea

Dewakang sill 560 m

Lifamatolasill 1800m

South Savu < 1200 m

Halmaherasill 650 m

Lombok sill 210 m

Ombai eastsill 1800 m

Ombai westsill 1200 m

Libani channel

Sunda Arcsill 1300 m

Timor Sea

Fig. 2 Bathymetry of the modeland sill depths of the model

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tion in the water mass transformation from the temperatureand salinity evolution equations:

DT

Dt¼ d T

d tþ U � rT ¼ DT þ FT ð2Þ

DS

Dt¼ d S

d tþ U � rS ¼ DS þ FS ð3Þ

where T and S are the potential temperature and the salinity,U the velocity, and D and F are the mixing and the forcingterms.

Briefly, the methodology is as follows: each tendencyterm of the temperature Eq. (2) and the salinity Eq. (3) arediagnosed and then recombined using their expansion/contraction coefficients in order to obtain the associatedtendency in term of neutral density. The resulting tendencyis then projected onto neutral surfaces and summed in eachdensity layer, giving access to the associated dianeutraltransport, Ωγ:

Ωg ¼ @D

@gþ @F

@gð4Þ

where D and F are the volume-integrated effects of mixingand forcing over a volume. Ωγ is the water masstransformation in a neutral density framework whileMg ¼ �@Ωg

�@g, the convergence of Ωγ, refers to the

water mass formation at a given neutral density γ.Therefore, @D=@g and @F=@g are the transformationsdriven by ocean physics and external forcing, respectively.Mγ>0 is associated with a formation of the water mass at agiven neutral density γ, whereas Mγ<0 represents adestruction of the water mass. Let us consider a 1D fluidat equilibrium isolated from the surface forcing. Thevertical advection balances the vertical diffusion. Thediapycnal fluxes can be quantified either by advection orby diffusion. This very simple configuration illustrates wellthe link between the right hand side of Eqs. (2) or (3)(diffusion and forcing terms) and the equivalent diapycnalflux to which it can be associated (Ωγ). In this study, we areconsidering a climatological state and not the interannualvariability of the ITF. The model has reached an equili-brated state after 10 years, and thus dγ /dt is close to zeroover a seasonal cycle. The physical terms (D+F in Eq. (4))balances exactly the advection. After the projection ontodensity classes, they can be related to diapycnal fluxes.

The mixing term has been split into four parts of Eq. (4).

Ωg ¼ ΩF þΩTidesDv þΩtke

Dv þΩubsDl þΩiso

Dl ð5Þ

The first component concerns the diffusion associated withvertical tidal mixing alone Ωtides

Dv

� �; the second component

represents the remaining vertical physics (ΩtkeDv, i.e., mixing

associated with the TKE turbulent closure and convection

processes). Then, we have chosen to separate the diffusionassociated with the advection scheme used. It is a third order-biased scheme composed of two distinct parts: a non-diffusive but dispersive fourth order-centered scheme andan upstream-biased diffusive term which can be expressed asa biharmonic diffusion which coefficient is proportional tovelocity (Madec 2008). These two parts are diagnosed onlineseparately and the latter part forms the mixing decompositionΩubs

Dl

� �. Two other advection schemes are tested: a second

order centered difference scheme (referred as CEN) and asecond order flux corrected scheme (referred as TVD, Madec2008). The CEN does not have a diffusion part Ωubs

Dl ¼ 0� �

but is dispersive and creates unphysical extrema. The TVDscheme is a combination of a centered difference and a firstorder upwind scheme, which is only applied above a certaincriteria (Lévy et al. 2001). As for the UBS scheme, it isdecomposed into the centered part and the remainingdiffusing part Ωubs

Dl

� �. Finally, the last contribution to the

mixing concerns the isoneutral mixing operator ΩisoDl

� �. In

the interior of the ocean, this operator acts along isoneutralsurfaces. However, in the mixed layer, the surfaces alongwhich mixing occurs are linearly flattened to becomehorizontal at the sea surface. Therefore, the mixing operatorbecomes a Laplacian operator in vicinity of the oceansurface. We note that ΩF is the transformation induced bythe forcing, which takes into account the solar energypenetration as in Iudicone et al. (2008a). In other words, ata given location, the forcing can penetrate down to at adenser density class than the surface density itself.

Within the Indonesian Seas, a very good approximationof the neutral density is given by McDougall and Jackett(2005). The calculation of the tendency and the tracers hasbeen performed online at each modeled time step so thereare no sampling errors introduced in the computation.Nevertheless, we verify that sufficient accuracy can beachieved with an off line computation using 5-day meanoutputs only.

3 Model results

3.1 Model transport and water masses

The main route of the ITF is the western route that passesthrough the Makassar Strait and has a North Pacific origin.At the eastern entrance, a deep flow enters the IndonesianSeas from the Pacific Ocean (3 Sv) via Maluku Sea andLifamatola Strait (in good agreement with the observations1.5 Sv, van Aken et al. 1988, and the latest INSTANTmeasurements, 3 Sv, van Aken 2007), forming the easternroute. Surface and lower thermocline waters enter viaHalmahera. The transport along this route is slightlyunderestimated (0.5 Sv) compared to the value of 1 Sv

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deduced from observation (Gordon 2005). The transportthrough the three exits and their relative contributioncompares well with the estimates from observations, evenif they are 10–40% larger than observed: 2.6 vs. 1.7–2.2±0.5 Sv for the Lombok Strait (Arief and Murray 1996;Sprintall, J, Wijffels, S, Molcard R, Jaya I, 2008: Transportvariability in the exit passages of the Indonesian though-flow. J Geophys Res, to be submitted), 6.3 vs. 4.3±1.2 Svfor the Ombai Strait (Molcard et al. 2001), and 7.7 vs. 4.5–7.6±1 Sv for the Timor Passage (Molcard et al. 1996;Cresswell et al. 1993; Sprintall, J, Wijffels, S, Molcard R,Jaya I, 2008: Transport variability in the exit passages ofthe Indonesian thoughflow. J Geophys Res, to be submit-ted), respectively. Indeed, Fig. 3 shows model annual meantemperature and salinity over the top 600 m on a sectionalong the western route. The model reproduces very wellthe observed T–S evolution along this route.

Salinity erosion for both routes is in accordance with theobserved data (see Gordon 2005 for a review). Indeed, atthe entrance, the North Pacific inflow water is characterizedby the salty Subtropical Water (NPSW, 34.8 psu) between100 and 200 m and by the fresher intermediate water(NPIW, 34.4 psu) below. All along their journey, thesalinity extrema are progressively eroded (as in Ffield andGordon 1992). In Makassar Strait and Flores Sea, the veryfresh (<33.5 psu) and warm surface Java water joins thisroute and is progressively mixed. The eastern route watermass evolution is also well reproduced. In particular, somesalty SPSW enters the lower thermocline of the Banda Sea,whereas no SPSW enters the Banda Sea within the upperthermocline (see figure 3 in KL07), as observed (Gordon2005).

The model is therefore a good representation of realityboth in terms of the transport and water masses. It will be

Fig. 3 Vertical section of annu-al mean temperature (mid-panel)and salinity (bottom) along themodeled western route. Thecontours represent the isoneutralsurfaces from 21.5 to 26.5, withan interval of 0.5. The thickerblack contours represent thelimit of the water mass consid-ered (Table 1). Top panel showsthe subdivision of the total do-main (TOT) used for the calcu-lation. It also shows the sectionpath (black thick line) throughthe domain which is also iden-tified on the section distanceaxis. The green path corre-sponds to the passage throughthe Sulawesi Sea, purple to theMakassar and Western FloresSeas, and orange to the southeastern Flores Sea and southwestern Banda Sea

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very helpful in the diagnosis of the main mechanisms of theIndonesian water mass transformation.

3.2 Density decomposition

In order to study the transformation, density classes have tobe defined. Using the modeled T–S structure described(Fig. 3), the water column has been decomposed into fourdensity classes, which refer to classical water massdefinitions (Table 1): The Surface Water (SW), theThermocline Water (TW), the Intermediate Water (IW),and the Deep Water (DW). The top of the thermocline,which is usually taken as the maximum depth of the mixedlayer, corresponds in the model to γn=22 (Fig. 3).

The surface water is then defined as the water above γn=22 and contains the warm and fresh water intrusion from theJava Sea (Fig. 3) to the Flores Sea (Table 1). It concernswater above 27°C. Furthermore, the commonly used lowerlimit of the tropical thermocline water is 12°C (Wyrtki 1961;Purba and Atmadipoera 1992) which in the model corre-sponds to the neutral density γn=26 (Fig. 3). The thermo-cline is thus defined by the water between γn=22 and γn=26, which contains the North Pacific salinity maximum andvaries from 12 to 27°C. In the following, thermocline waterwill be shaded in grey for more readability. The intermediatewater, which is characterized by a salinity minimum (Fig. 3)is defined between γn=26 and γn=27.5. Deep watercorresponds to the water below γn=27.5.

The surface layers have been divided into two layersseparated by the γn=21.5 (upper black line on Fig. 3) torefine the density decomposition. The upper surface waterexists only near the Java Sea (Fig. 3). The lower surfacewater is found between 0–50 m and 0–25 m depth, and itslower limit (γn=22) only outcrops in the Northern Australianbasin (Fig. 3). In the same way, the thermocline has beendivided into three layers: the upper (22<γn<23), the mid-(23<γn<24.5), and the lower thermocline (24.5<γn<26,black lines on Fig. 3).

The mid-thermocline water, the saltiest layer, and the lowerthermocline water ends up in the Banda Sea with a constantsalinity profile (Fig. 3). Finally, the intermediate water hasbeen divided into upper, mid-, and lower intermediate water(26<26.5<27<27.5). The mid-intermediate water is the

freshest and varies between 300–350 m and 600–700 m. Inthe Indian Ocean, it encounters the saltier North West IndianIntermediate water (Fig. 3, 500 m depth at 5,000 km fromthe entrance; Fieux et al. 1994 and 1996, JADE experiment1992).

3.3 Density binned vertical transport

For these density classes, the transport distribution isplotted in Fig. 4 for the inflow and the outflow. It changesfrom the entrance to the exit. The inflow has been taken asthe sum of the transport across the Sangihe Ridge (westernentrance), through the Halmahera and Maluku entrances,and through the Java and the South China Sea entrances.Positive transport is towards the Indian Ocean. The outflowcorresponds to the westward transport at 112°E betweenJava and Australia. The inflow is maximum in the surfacelayers and in the intermediate and deep layers, and almostzero within the mid-thermocline layer. At the Indiansection, the outflow is thermocline intensified. The ther-mocline transport is equally distributed over the threethermocline layers and a large volume transport is stillconcentrated within the LIW layer in the outflow. Theupper part of the intermediate and deep water and most ofthe surface water lose their transport in favor of thethermocline layers. In the following, we will link thesediapycnal fluxes to their physical sources using Eq. (5).

4 Transformation of water masses

4.1 TIDES experiment

The dianeutral fluxes related to the physical sources in theITF are shown in Fig. 5a for the TIDES experiment, usingthe decomposition of Eq. (5) and summed over the TOTregion (blue and red colors, see Fig. 3). The main dianeutralfluxes within the thermocline are schematized on Fig. 5b asan example of how to interpret Fig. 5a. The correspondingformation rates Mγ, i.e., the convergence/divergence of thedianeutral transports taken from the derivatives of thecurves, are given in Table 3 and also in numbers on Fig. 5a.

Table 1 Definition of watermasses used in the text

Ocean Dynamics

A positive slope in Fig. 5a corresponds to a water massdestruction whereas a negative slope corresponds to a watermass formation. For example, the black curve representsthe total diapycnal fluxes. They are maximum at γn=22 andminimum at γn=26. The difference between the two

extrema corresponds to a convergence of 6.4 Sv in thethermocline Ωg 22ð Þ �Ωg 26ð Þ ¼ 6:4 Sv

� �. This conver-

gence is the result of a destruction of 2 Sv of surface waterΩg 19ð Þ �Ωg 22ð Þ ¼ �2 Sv� �

and of 4.4 Sv of intermediatewater Ωg 26ð Þ �Ωg 29ð Þ ¼ �4:4 Sv

� �.

Fig. 5 Sum of all of the tendency terms in the equation of evolutionof neutral density for the TIDES (a) and REF (c) experiments. Thedianeutral flux is the derivative of the total mixing and total forcingcurves. The right side of Eq. (5) (black continuous line) is theadvection. The left side of Eq. (5) (black dashed line) is the totalphysics. The difference between the two is the residual term(negligible). The total transformation results from a balance betweenall of the terms on the right of Eq. (5). For example, the total black

curve shows the balance between the total mixing (blue curve) and theforcing (red curve). The mixing ΩD (blue curve) is the sum of ΩTides

Dv(green), Ωtke

Dv (pink), ΩUBSDl (orange), Ωiso

Dl (cyan) as in Eq. (5). Insert(b) shows the main dianeutral fluxes in the thermocline within theTIDES simulation schematized by arrows. The intensity of eachdianeutral flux is noted. Only dianeutral fluxes larger than 1 Sv areschematized. The same color key as for (a) is used for the arrow

Fig. 4 Density partitioned transport of the inflow (yellow, calculatedon the red boundaries of the mask TOT, see Fig. 3, top panel) and theoutflow (black, calculated over the blue boundaries of the mask TOT,

see Fig. 3). The total vertically integrated transport is the same for theinflow and the outflow

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For the surface water (γn<22), the forcing tends to formlighter surface water (red curve ΩF 19ð Þ �ΩF 22ð Þ ¼8:6 Sv). It is overcompensated by a larger mixing (bluecurve, ΩD 19ð Þ �ΩD 22ð Þ ¼ �10:6 Sv), which results in thedivergence of the flux from the surface to the thermocline(black curve, Ωg 19ð Þ �Ωg 22ð Þ ¼ �2 Sv). The mixing ismainly driven by the tides (green curve), which consume5.3 Sv (green Fig. 5a and b) of lighter surface water andrepresent 50% (Table 2) of the total mixing contribution.The lateral mixing due to the diffusive part of theadvection scheme (orange Fig. 5ab) consumes 3.6 Sv oflight surface water which represents 35% of the totalmixing contribution. Horizontal mixing due to the iso-pycnal mixing (cyan) at the surface contributes 2% to thetotal mixing while the remaining part of the verticalphysics contributes 11% (pink, Table 1). In particular, thismeans that the wind makes a quite small contribution tothe transformation.

The intermediate waters (γn>26) are deep enough tonever experience the surface forcing. Mixing is the onlymechanism responsible for their transformation. It destroys4.4 Sv of intermediate and deeper water. The tides are againthe major contributor to this transformation (53%) while thecontribution of the diffusion part of the advection scheme is34%. The vertical shear creates turbulence that produces anupward 0.4 Sv (8%) dianeutral flux at the base of thethermocline and the isopycnal diffusion contributes to 5%of this transformation.

What is formed in the thermocline (22<γn<26) resultsdirectly from what has been destroyed in the surface and inthe intermediate layers. The total convergence into thethermocline is 6.4 Sv (2.2 Sv for the UTW, 2.6 Sv for theMTW and 1.6 Sv for the LTW). The forcing is still acting inthe shallower layers and tends to destroy UTW (7 Sv=8.6–1.6 Sv, Fig. 5a and b) and MTW (1.6 Sv) by lighteningthem.

The net tidal contribution over the whole thermocline is53% of the total mixing contribution, mainly distributed inthe UTW and MTW layers (3 Sv=5.3–2.3 Sv and 4.7 Sv=2.3+2.5 Sv respectively, Fig. 5b) and quasi null for thelower thermocline. On the other hand, diffusion associatedwith the advection schemes, which represents 33% of themixing transformation, is acting mainly on both sides of thethermocline and is almost zero in the mid-thermocline(Fig. 5b).

4.2 REF vs. TIDES experiment

To further understand the role played by the tides, the TIDESand REF experiments (with and without the tidal parameter-ization) are compared. For the REF experiment, i.e., withoutthe tidal mixing parameterization, the total transformation isfar less significant than for the TIDES run. Only 0.8 Sv of

light surface water is consumed and 1.8 Sv from the UIWand MIW to form 2.7 Sv within the whole thermocline(black curve, Fig. 5c, Table 2).

In the surface layer (γn<22), mixing destroys only4.9 Sv, about half the value in the TIDES experiment.The upper ocean is hence less mixed. As a consequence,the forcing is less efficient and represents only 5.4 Sv offormation of light surface water whereas for the TIDES runit was 8.6 Sv. So adding the tides has two effects at thesurface: it increases the consumption of surface water(almost five times more) and increases the heat absorbed bythe ocean. The heat flux difference is about 20 W/m2 (notshown). The T–S vertical structure is also modified. Thesurface layer is colder and saltier (Fig. 6) in the TIDES runand the thermocline is fresher and warmer (Fig. 1). Themean sea surface temperature difference is 0.5°C with localmaxima of up to 1.2°C.

The contribution of the freshwater flux vs. the heat flux inthe density forcing for both the TIDES and the REFexperiment is plotted in Fig. 7 as well as the temperaturecontribution without solar energy penetration (orange). Forthe TIDES experiment, the forcing contribution is mainlydriven by temperature (solid red), as its contribution is fourtimes stronger than the freshwater counterpart (solid blue).The solar energy penetration allows between 2.5 and 3.5 Svmore transformation by the heat flux and reaches deeperwater masses. Furthermore, adding the tidal mixing increasesthe heat contribution by 4 Sv (REF, red dashed thin,compared to TIDES, red thin). The salt contribution is alsoenhanced by 0.8 Sv. The solar energy penetration and thetidal mixing act in a similar way in deepening andintensifying the penetration of heat. Their contributionquadruples the amount of heat absorbed by the ocean thatwould be otherwise nearly equal to the salt contribution.

The transformation of the thermocline water for theREF experiment includes the formation of 1.2 Sv of UTWand 1.5 Sv of LTW (Fig. 5). These values are smaller thanfor the TIDES experiment. The major difference concernsthe MTW. In the REF run, the water diverges from thislayer leading to a destruction of 0.7 Sv instead of aformation of 2.6 Sv for the TIDES run. The tides are theonly process that allows large convergent dianeutral fluxesin the MTW, leading to far better agreement with theobservations (Fig. 1, in the density class of 24 kg m−3,MTW).

4.3 Regional transformation

This section tackles the issue of where the transformationoccurs. We quantify the relative contribution of thetransformation processes in each of the nine sub-domainswhose sum is equal to the domain previously considered(Fig. 8). The total transformation over the whole domain

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TOT (Fig. 8, top panel) is distributed as follows: 40% istransformed near the Sangihe Ridge (W1 region), 30% inthe Makassar and Flores Seas (W2+3 region), 26% in theHalmahera, Malukku, and Seram Seas (E1 region), and26% in the Ombai and Timor Straits (WE3 region).Whereas in the Banda Sea (WE1+2 region), in the JavaSea (JA region), and in the Indian Ocean (WE4 region),the transformation is in the opposite direction to the totalone (−15%, −2%, −3%, Table 3). This reverse transfor-mation concerns mainly the surface and the upperthermocline layers and is due to a larger forcing comparedto the mixing term. In these regions, the surface waters arelightened by the forcing because the tidal mixing isrelatively weak.

Diffusion due to the advection scheme is localized oververy narrow passages: the Sangihe Ridge and the Makassar,Ombai, and Timor Straits. At the entrance, this term evendominates the transformation. Everywhere else it is negli-gible. The tides are the largest contribution to mixingeverywhere except at the entrance of the western route.

This spatial decomposition shows that the Banda Sea isnot the place of strong transformation due to mixing as wasassumed from the observational analysis. The main trans-formation occurs before and after the Banda Sea. Inparticular, the Halmahera and Ceram Seas (E1) are theplace where the tidal mixing is the largest, transforming theSP waters before their entry into the Banda Sea.

5 Conclusions and discussion

The total transformation quantified in the simulation is adiapycnal convergence of 6.4 Sv into the thermocline,

resulting from a destruction of 2 Sv of surface water and4.4 Sv of intermediate and deep water. The surfacedestruction is the result of two opposite effects: theatmospheric forcing (8.6 Sv) which lightens the surfacewater and the larger total mixing (10.6 Sv). Each mixingcontribution to the diapycnal transformation has beenseparated. Firstly, vertical mixing given by the turbulentclosure scheme, forced by the wind and ocean currentshear, only represents 10% of the transformation. Itsuggests that the wind is not an essential component ofthe transformation in the Indonesian Seas. Secondly, thetides play a dominant role in the transformation andrepresent 55% of the total mixing transformation. Thisoccurs mainly at the major sills encounters by the ITFbefore and after the Banda Sea (Sangihe Ridge, Dewakang,Timor, Ombai, Halmahera, Lifamatola sills, Fig. 2).Without tidal mixing, the total transformation is far lesssignificant. Only 2.6 Sv is formed in the thermocline, out ofwhich 1.8 is transferred from intermediate waters. Themajor differences are localized in the mid-thermocline. Inthis layer, the tides are producing a strong vertical mixing,which allows the model to reproduce the observed halostadstructure and the associated cold and fresh tongue charac-teristic of the Indonesian water (see also Gordon 2005 andFig. 1 in this paper).

The tidal mixing cools the SST by 0.5 to 1°C (Fig. 6).This SST difference may influence air sea interactions incoupled climate models. The tidal mixing allows a surfaceheat flux contribution which is 3 Sv larger (representing 20–40 W/m2 more absorbed) and a freshwater flux contributionwhich is 0.8 Sv larger, and deepens their respectivecontribution. The solar energy penetration also allows anadditional heat contribution, which is 2.5 Sv larger and

Table 2 Net convergence/di-vergence (positive/negative) inthe surface, thermocline, inter-mediate, and deep waters (Table1) for the TIDES and REFexperiments (in Sv)

Table 3 Percentage of the total transformation for each region defined in Fig. 8

Red text: forcing, green: tides,pink: rest of the vertical physics,orange: lateral diffusion of theadvection scheme named UBS,blue: lateral mixing due to iso-pycnal formulation. The per-centage corresponds to thefraction of the total mixing.

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deeper. Without solar penetration or tidal mixing, the heatcontribution would only be as large as the freshwatercontribution, i.e., four times smaller. The Indonesian Seasare one of the rare places (e.g., equatorial cold tongues)where mixing is directly acting on the forcing.

Finally, spurious mixing associated with the advectionscheme produces 30% of the transformation, mainly bydestroying surface water. This only concerns lateralmixing and is limited to some narrow channels. Themodel resolution is sufficient to resolve the Rossbydeformation radius at Indonesian latitudes. However, thecomplex topography of the Indonesian straits, composed

of fine structures of a few kilometers, are constraining theflow at some critical points which are not reproduced bythe model. In particular, the very narrow passages such asLombok and Ombai Straits, the passages between theislands of Sangihe at the entrance of the western route orthe narrow channel of Makassar Strait, are resolved byonly one grid point. Thus, the model is unable to resolvethe physics inside the strait. The numerical schemebecomes strongly diffusive at the exit of these straitswhere the flow encounters the resident waters withdifferent properties. In order to resolve explicitly the flowin these passages, a resolution of 1 km would necessary.Such a resolution is very CPU costly and not yet availablefor large scale circulation studies.

Sensitivity experiments have been performed usingdifferent advection scheme which are more or lessdiffusive (see Section 2.2). We tested these schemes onboth the vertical and horizontal advection. Only changesin the horizontal advection scheme produced significantdifferences for the water masses. The total transformationranges between 4.8 and 6.4 Sv depending on the advectionscheme used (Fig. 9). Due to the dominance of otherphysical processes (such as forcing, tidal mixing, andisopycnal mixing), the remaining non-physical (i.e.,numerical) contribution is only 10% of the total mixingtransformation. This spurious numerical mixing could bereduced (e.g., using the CEN experiment) but it actuallyimproves the water mass representation (Fig. 10) andcould be interpreted as a parameterization for somemissing physics, such as a lack of lateral mixing. Theauthors propose that the unresolved barotropic tidesconstrained by topographic features could be responsible

Fig. 6 SST and SSS differencebetween the TIDES and REFexperiments. The tides cool thesurface by 0.5 to 1.5°C andincrease the salinity in theMakassar Strait and Flores Seaby 0.2 to 0.5 psu

Fig. 7 Contributions to the forcing tendency, in terms of temperature(orange and red) and salinity (blue). Orange represents the temper-ature contribution without solar energy penetration, red with solarenergy penetration. Dashed lines correspond to the REF experimentand solid lines to the TIDES experiment

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for lateral mixing in such narrow channels, by strongstirring of different properties on each side of the strait.This spurious lateral mixing may be complicated byadditional vertical mixing processes which are not repro-duced by the model such as hydraulic jumps, lee waves,the bottom boundary layer, and wind-generated nearinertial waves which penetrate the thermocline and maydissipate in it (Alford and Gregg 2001; Polton et al. 2008).One perspective to improve on this numerical limitation isto increase the resolution in order to resolve these very

narrow channels. Only a model with a few kilometerresolution would succeed in doing so.

These results renew the picture of the Indonesiantransformation for three reasons. Firstly, this study showsthat the Indonesian archipelago cannot be reduced to theonly transformation, in the sense of mixing TW, IW, andSW to form a new water mass: the Indonesian water. Indeed,the strong mixing is able to cool the SST, and thus increasesthe surface forcing. The additional forcing entering the oceanimprints its effect into the thermocline water, as seen recently

Fig. 8 Upper panel: net con-vergence (positive term, yellowboxes) or divergence (negative,blue boxes) in differentsub-domains. Convergence/divergence for the total domain(TOT) on the left, along thevertical section of the modeledwestern route, as in Fig. 3, onthe right. Bottom panel (right)shows the remaining domains,and (left) shows the legend forthe arrows indicating the de-composition of the diapycnalfluxes marked on each figure

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by repeated XBT lines (Wijffels et al. 2008) and can lead towater mass formation. Therefore, the Indonesian water is theresult of both a transformation driven by the mixing and aformation driven by both the mixing and the air sea fluxes.Both mixing and forcing, which created the new water mass,are driven by the tidal mixing. Secondly, this study concludesthat the thermocline (halostad) structure cannot be reproducedwithout strong tidal vertical mixing, which occurs before andafter the Banda Sea. This result is in contrast to the classicalvision from observational analysis that strong vertical andhorizontal mixing occurs within the Banda Sea (Hautala et al.1996; Gordon 2005). Finally, the third point concerns thetransfer of intermediate waters into the thermocline. It hasbeen quantified to be as large as 4 Sv. This suggests that theIndonesian transformation contributes significantly to theupwelling of cold water in the global conveyor belt.

This paper traces the different sources of water masstransformation using Eulerian boxes. In a companionpaper (Koch-Larrouy et al. 2008), the same transformationhas been described in a Lagrangian sense along eachdifferent path of the ITF, and gives a complementarypicture of the transformation.

Acknowledgements This work is part of the DRAKKAR projectand is supported by MERCATOR-ocean (projects 100043 and061396) and by the Marine Environment and Security for theEuropean Area project (MERSEA, SIP3 CI 2003 502885). Theocean model integrations have been performed at the Institut deDéveloppement et des Ressources en Informatique Scientifique(IDRIS, project 51140 and 1396).

Fig. 9 Total transformation for the three experiments that differ onlyby the lateral advection scheme used: CEN (blue), TVD (green), UBS(orange). The total transformation difference between UBS experi-ment (the more diffusive scheme) and CEN experiment (the lessdiffusive scheme) is 1.6 Sv, far less significant than the Ωubs

Dl ¼ 5:2 Svdiagnosed online with the stratification created by UBS scheme inFig. 5. This is due to feedbacks of the other terms: forcing term issmaller and the vertical and horizontal mixing are stronger in the CENthan in the UBS experiment (not shown). This suggests that a largepart the transformation done by the UBS diffusion is replaceable byphysical (and not numerical) mixing. The remaining part (1.6 Sv) isstill not reliable to physical processes

Fig. 10 Water masses in thethree different experiments thatdiffer only by their lateral ad-vection scheme (CEN, blue;TVD, green; UBS, red) withtidal mixing and for the REFexperiment (dashed red). Theobservations are in black

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