Acta Oceanol. Sin., 2014, Vol. 33, No. 1, P. 11–23

DOI: 10.1007/s13131-014-0429-2

http://www.hyxb.org.cn

E-mail: hyxbe@263.net

Interbasin exchanges and their roles in global ocean circulation:A study based on 1 400 years’ spin up of MOM4p1ZHU Yaohua1, WEI Zexun1∗ , FANG Guohong1, WANG Yonggang1, GUAN Yuping2

1 Key Laboratory of Marine Science and Numerical Modeling, First Institute of Oceanography,State Oceanic Administration, Qingdao 266061, China

2 State Key Laboratory of Tropical Oceanography, South China Sea Institute of Oceanology,Chinese Academy of Sciences, Guangzhou 510301, China

Received 16 May 2013; accepted 18 September 2013

©The Chinese Society of Oceanography and Springer-Verlag Berlin Heidelberg 2014

AbstractA global prognostic model based on MOM4p1, which is a primitive equation nonBoussinesq numericalmodel, has been integrated with 1 400 years from the state of rest based on the realistic topography tostudy the long-term pattern of combined wind-driven and thermodynamically-driven general circulation.The model is driven by monthly climatological mean forces and includes 192×189 horizontal grids and31 pressure-based vertical levels. The main objective is to investigate the mass and heat transports at inter-basin passages and their compensations and roles in the global ocean circulation under equilibrium state oflong-term spin up. The kinetic energy analysis divides the spin up process into three stages: the quasi-stablestate of wind driven current, the growing phase of thermodynamical circulation and the equilibrium state ofthermohaline circulation. It is essential to spin up over a thousand years in order to reach the thermohalineequilibrium state from a state of rest. The Arctic Throughflow from the Bering Strait to the Greenland Seaand the Indonesian Throughflow (ITF) are captured and examined with their compensations and existingdata. Analysis reveals that the slope structures of sea surface height are the dynamical driving mechanismof the Pacific-Arctic-Atlantic throughflow and ITF. The analysis denotes, in spite of O (1.4×106 m3/s) of thesouthward volume transport in the northern Atlantic, that there is still O (1 PW) of heat transported north-ward since the northward currents in the upper layer carry much higher temperature water than the south-ward flowing northern Atlantic deep water (NADW). Meridional volume and heat transports are focused onthe contributions to NADW renewals and Atlantic meridional overturning circulation (AMOC). Quantitativedescriptions of the interbasin exchanges are explained by meridional compensations and supported by pre-vious observations and numerical modeling results. Analysis indicates that the volume and heat exchangeson the interbasin passages proposed in this article manifest their hub roles in the Great Ocean ConveyorSystem.Key words: numerical modeling, global ocean, interbasin exchange, meridional transport, meridional over-

turning circulation

Citation: Zhu Yaohua, Wei Zexun, Fang Guohong, Wang Yonggang, Guan Yuping. 2014. Interbasin exchanges and their rolesin global ocean circulation: A study based on 1400 years’ spin up of MOM4p1. Acta Oceanologica Sinica, 33(1): 11–23, doi:10.1007/s13131-014-0429-2

1 IntroductionThe global-scale circulation has long been one of the

oceanography’s most challenging and exciting research top-ics. A century ago, Pillsbury (1912) pointed out that globalocean circulation transports heat poleward from the equator.But oceanographers did not focus on the heat transport rateuntil several decades ago. There has been a developing focuson the world oceanic thermohaline circulation since it is im-mediately related to the global climate change. Broecker (1987,1991) introduced the ocean conveyor belt terminology and two-layer thermohaline flow scheme to study the deep layer circu-lation and upper layer compensation currents. Schmitz (1995)summarized updated research achievements and observation-

al results and proposed a four-layer thermohaline flow schemebased on interbasin water exchange. He illustrated his thermo-haline scheme, including bottom water, deep water, intermedi-ate water and upper layer compensation water, and estimatedthe volume transport rates. Both Broecker’s two-layer schemeand Schmitz’s four-layer scheme presented canonical picturesfor global ocean thermohaline circulation.

Huisman et al. (2009) and Marotzke and Willebrand (1991)employed GFDL’s (geophysical fluid dynamic laboratory) mod-ular ocean model (MOM2) with idealized rectangle Atlantic andPacific Ocean and 4◦×4◦ coarse grid to integrate thousands ofyears to study the multiple equilibria of thermohaline circula-tion. In their studies, coarse grids and large steps were applied

Foundation item: The National Basic Research Program Grant of China under contact No. 2011CB403502; the International Cooperation ProgramGrant of China under contact No. 2010DFB23580; the International Cooperation Program of State Oceanic Administration of China under contractNo. QY0213022; project supported by the First Institute of Oceanography, the State Oceanic Administration of China under contract No. 2010G06;author Guan Yuping is supported by The National Natural Science Foundation of China under contact Nos 40976011 and 91228202.*Corresponding author, E-mail: weizx@fio.org.cn

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12 ZHU Yaohua et al. Acta Oceanol. Sin., 2014, Vol. 33, No. 1, P. 11–23

to save computational resources. Analyses were focused to findthe equilibrium states and shift criterion rather than the re-al thermohaline circulation structure. Wei et al. (2000) inte-grated MOM2 for 11 years with 1◦ global resolution and (1/3)◦high resolution nested grid to study both wind driven and ther-modynamically driven general circulation. Typical zonal andmeridional cross sections were selected to analyze climatolog-ical mass, heat and salt transport rates in the global ocean.The net northward mean through flow in the Bering Strait andthe net southward mass transport in the Atlantic, thus Pacific-Arctic-Atlantic interbasin exchange, were not indicated.

Dong et al. (2011) employed a global circulation modelfor the earth simulator (OFES), in which the model code wasbased on the MOM3 and the spin up run was 50 years from aninitial condition at rest with observed mean hydrographic da-ta. Detailed analyses were focused on the role of interocean ex-changes on decadal variations of the meridional heat transportin the South Atlantic. However, their numerical results wouldbe different if the spin up run lasted for a thousand years more.

Therefore, it becomes necessary and useful to simulatethe sufficiently developed state of the general ocean circulationwithout initial background constraint whereby under real to-pography, relatively fine resolution and realistic climatologicalforces can be applied.

Along with the development of computer technology andthanks to GFLD’s endeavor, it becomes possible to integrate therealistic topography global ocean for thousands of years withrelatively high resolution grids. This article presents an inte-gration of 1 400 years from the state of rest to the equilibri-um state in order to investigate the long-term pattern of gen-eral circulation system, based on the combined wind-drivenand thermodynamically-driven mechanism. The nonBoussi-nesq mass conserving MOM4p1 includes conservative temper-ature and pressure based vertical coordinates. The analysesin this article indicate that the interbasin exchanges are al-

ways closely related to meridional transports. The interbasinexchange results of this article reveal a Pacific-Arctic-Atlantic-Indian Ocean-Pacific circle transport pattern which composesa worldwide ocean circulation.

There are six interbasin exchange passages in the worldocean–two in the Arctic region, three in the Antarctic Circum-polar Current regime (ACCR) and one in the tropical area. TheGreenland Sea and the Norwegian Sea connect the Atlantic andArctic Ocean, therefore is an Atlantic-Arctic exchange passage(AA section). The Bering Strait is another interbasin passageconnecting Arctic and Pacific Ocean (AP section). ACCR linksthe Southern Pacific, Atlantic and Indian Ocean, therefore is aunique linkage for the great oceans to exchange bottom, deep,intermediate and upper layer water. There are three inter-basin exchange passages in ACCR, including Southern Atlantic-Indian Ocean passage (SAI), Southern Indian Ocean-Pacificpassage (SIP) and Southern Pacific-Atlantic passage (SPA). TheIndonesian Archipelago passage is the sole passage which con-nects the Pacific and Indian Ocean in the tropical region, thusis the only “express way” for mass, heat and salt transports inthe world ocean. Obviously, it is of great importance to focuson these passages’ transports because of their hub roles in theglobal ocean circulation and their relations to meridional trans-ports and Great Ocean Conveyor as well.

2 Model configurationThe MOM4p1 is one of the most recent versions of the

GFDL ocean model, featured by its generalized vertical coor-dinates. In this global spherical model, the horizontal reso-lution of 1.9◦×0.95◦ (192×189 grids) and 31 vertical levels (Ta-ble 1) are employed. Pressure-based vertical coordinate p ∗ hasbeen applied for the mass conserving, nonBoussinesq, free sur-face ocean primitive equations. Lateral boundary conditionsare cyclic in zonal direction, and solid walls in meridional di-rection.

Table 1. Mean depths and thickness of vertical levels under p ∗ coordinate

Level Depth range/104 Pa Thickness/104 Pa Central depth/104 Pa Level Depth range/104 Pa Thickness/104 Pa Central depth/104 Pa

1 0–10.1 10.1 5.1 17 631.8–744.8 113.0 684.5

2 10.1–21.9 11.8 15.2 18 744.8–876.3 131.5 805.1

3 21.9–35.6 13.7 28.7 19 876.3–1 029.2 152.9 947.4

4 35.6–51.6 16.0 42.6 20 1 029.2–1 206.9 177.7 1 110.9

5 51.6–70.1 18.5 60.5 21 1 206.9–1 413.7 206.8 1 303.0

6 70.1–91.6 21.5 79.7 22 1 413.7–1 654.2 240.5 1 524.4

7 91.6–116.6 25.0 103.6 23 1 654.2–1 933.8 279.6 1 783.9

8 116.6–145.7 29.1 129.7 24 1 933.8–2 259.0 325.2 2 083.7

9 145.7–179.5 33.8 161.7 25 2 259.0–2 637.2 378.2 2 434.3

10 179.5–218.8 39.3 197.3 26 2 637.2–3 042.9 405.7 2 840.0

11 218.8–264.5 45.7 240.4 27 3 042.9–3 448.6 405.7 3 245.8

12 264.5–317.6 53.1 288.7 28 3 448.6–3 854.3 405.7 3 651.5

13 317.6–379.4 61.8 346.6 29 3 854.3–4 260.0 405.7 4 057.2

14 379.4–451.1 71.7 412.2 30 4 260.0–4 665.7 405.7 4 462.9

15 451.1–534.6 83.5 490.1 31 4 665.7–5 071.4 405.7 4 868.6

16 534.6–631.8 97.2 579.0

The constant horizontal viscosity and diffusivity coeffi-cients are taken as Ah = 3.0× 104m2/s and Kh = 1.0× 103m2/sand the vertical ones are computed in the model with “chen”scheme. Time steps used for the internal mode (baroclinic) andexternal mode (barotropic) are 4 800 and 40 s respectively. Formore details about the model description, readers are directed

to “Elements of MOM4p1” by Griffies (2010).The topography data set is based on NOAA’s (1988)

ETOPO5 and the International bathymetric chart of the ArcticOcean (IBCAO). The surface boundary conditions are providedby NOAA National Oceanographic Data Center (NODC) WorldOcean Atlas 1994 (WOA), including monthly mean water flux,

ZHU Yaohua et al. Acta Oceanol. Sin., 2014, Vol. 33, No. 1, P. 11–23 13

heat flux, sea surface temperature (SST) and sea surface salini-ty (SSS). The spin up run is started from a state of rest and hasbeen integrated for 1 400 years to the equilibrium state. The seasurface is forced by the monthly climatological wind stress, tak-en from Hellerman and Rosenstein (1983), the water flux andthe heat flux with the surface temperature and salinity restoringtowards above SST and SSS.

Long-term fluctuations in flow patterns are an importantcomponent of ocean climate and thus are the focus of this ar-ticle. Except the kinetic energy stability analysis, the time vari-ability is not considered in this article, but not because of itslack of importance or interest. Different from Schmitz (1995),this article is focused on interbasin mass and heat transportsand their relations to meridional transports and NADW, ratherthan on vertical layer schemes and NADW renewals’ path.

Analyses in this article are based on 50 years mean com-putational result of the equilibrium state.

3 Result analysis and comparison

3.1 Three developing stages of the total kinetic energyFigure 1 shows the time variation of total kinetic energy of

the global ocean along with the years of integration. The mostsignificant feature is the three different stages of kinetic ener-gy’s developing process. The total kinetic energy approaches itsquasi-stable state within the first three months and lasts for 60years. This first stage is the quasi-stable state of wind-drivencurrent. As we have recorded, the kinetic energy series are0.478 5, 0.478 2, 0.456 7, 0.450 1, 0.447 6 (unit: 1018 J) cor-responding to the integration length of 3, 6, 9 months, 10and 60 years. The second stage is the growing phase of thethermodynamically-driven current in which the total kineticenergy has been growing dramatically until a thousand inte-grated years with total kinetic energy of 0.658 1×1018 J. Thethird stage is the equilibrium state for thermohaline circula-tion where the kinetic energy stably fluctuates in a very narrowrange.

Fig.1. Total kinetic energy variation along with inte-grated years.

Therefore, in order to obtain the equilibrium state fromthe state of rest, it is necessary to integrate more than a thou-sand years. Otherwise, the underdeveloped thermodynami-cally-driven circulation causes unreasonable vertical tempera-ture and salinity structures and underestimated transports forthe deep ocean thermohaline circulation. This is one of theimportant conclusions of the article, which is of significanceon the global ocean circulation modeling. From the integra-tion year of 1 210 to 1 400, the total kinetic energy varies from

0.659 37×1018 to 0.659 41×1018 J, with the relative deviationless than 1.0×10−4 within 200 years. Therefore, it is consid-ered as “equilibrium state”. Even though there are some crite-ria for the definition of equilibrium state, sometimes scientistsdo use their own judgment. For example, Marotzke and Wille-brand (1991) employed the basin-averaged surface heat flux un-der O(0.1) W/m2 as an equilibrium judgment.

If just for the purpose of wind driven current or shallowmarginal circulation, it could be enough to integrate the modelfor a decade or so, since it can approach equilibrium state muchquicker.

3.2 Sea surface height distribution and its slope structuresFigure 2 depicts the global sea surface height (SSH) distri-

bution. The Pacific Ocean has the highest SSH and the AtlanticOcean has the lowest one. The SSH reaches its peak in the sub-tropical convergent zone and decreases poleward. Both the Pa-cific and Atlantic have their higher SSH in the western boundarycurrent (WBC) area and lower values at the eastern ends. SSHin the Northern Pacific is higher than that in the Arctic Ocean,and the latter is higher than in the Nordic Sea. This constitutesa driving force inducing a through flow into the Arctic Oceanfrom Pacific and eventually flowing into the northern Atlantic.

Fig.2. The global ocean circulation model MOM4p1produced global SSH distribution.

Figure 3a shows the detailed SSH distribution around theArctic region. In the Bering Strait region, the SSH at the southside is approximately 0.5 m higher than the north side withinthe span of two latitudinal degrees (typical values are 0.8 m inthe south and 0.3 m in the north side of the Strait). The SSHof 0.46 m in the Bering Strait and 0.25 m at 70◦N of ChukchiSea, slightly north of Bering Strait, are significantly higher thanthe zonal mean SSH of –0.36 m at 70◦N cross section betweenGreenland and Norway. In the Arctic Ocean, the SSH on theCanada basin is higher than in Svalbard-Barents Sea side. At80◦N, the SSH peak occurs on the Canada basin with a valuebetween 0.8 and 1.0 m, while the trough occurs in Svalbard-Barents Sea side with a value between –0.8 and –1.0 m. Thisslope structure of SSH constructs the dynamical mechanism todrive a Pacific-Arctic-Atlantic interbasin transport. This Pacific-Arctic-Atlantic transport is also an important component of theGreat Ocean Conveyor. It needs to be noted that this model isnot an ocean-ice coupled model, thus the SSH and the upperlayer currents in the Arctic region could be exaggerated, sincewithout ice covering the wind stress effect is exaggerated.

14 ZHU Yaohua et al. Acta Oceanol. Sin., 2014, Vol. 33, No. 1, P. 11–23

Fig.3. The SSH distribution in the Arctic region (a) and near the Indonesian Archipelago region (b).

Figure 3b illustrates the SSH distribution around the In-donesian Archipelago region. SSH in the east side of the In-donesian Archipelago with a typical value of 1.6 m is significant-ly higher than in the west side, meanwhile the north side (westPacific region) is higher than that in the south side (south Indi-an Ocean). Particularly, 0.6 m of the typical magnitude of SSHat 30◦S Indian Ocean and 1.6 m of the SSH in the west Pacif-

ic region construct a northeast-southwestward SSH slope. It isthis SSH slope that drives ITF from northeast to southwest ofIndonesian Archipelago.

In the subantarctic frontal zone (SFZ), the SSH decreasesdramatically poleward. The Antarctic coastal area has the low-est SSH in the global ocean, with a typical magnitude of –2.5 mand an extreme magnitude of –3 m in the Weddell Sea.

Table 2. Mass, heat and salt transports through specific transoceanic sections and interbasin passages

Section Latitude Longitude Volume transport/106 m3·s−1 Heat transport/PW Salt transport/109 kg·s−1

1A 0◦N 55◦W–15◦E –1.603 0.239 –0.0430

1B 0◦N 40◦–100◦E 0.024 –0.225 0.0002

1C 0◦N 100◦E–75◦W 0.605(14.000 from 133◦E) –0.570 0.0360

2A 30◦N 90◦W–0◦E –1.585 0.687 –0.0470

2C 30◦N 120◦E–110◦W 0.953 0.474 0.0430

3A 30◦S 50◦W–17◦E –1.089 –0.109 –0.0440

3B 30◦S 30◦–120◦E –15.480 –1.737 –0.5470

3C 30◦S 150◦E–70◦W 17.190 –0.048 0.5840

AP BERING ST. 66◦N 170◦–167◦W 1.2670 0.010 0.0410

AA G & N 70◦N 25◦W–20◦E –1.422 0.247 –0.0380

4 SAI 75◦–30◦S 20◦E 292.000 2.998 10.0800

5 SIP 75◦–30◦S 135◦E 307.600 4.181 10.6300

6 SPA 73◦–53◦S 67◦W 290.800 3.509 10.0300

7 IAP 7◦S 110◦–140◦E –13.530 –0.952 –0.4660

Notes: Positive values denote eastward or northward net transport. Abbreviation are AA G&N, cross section at 70◦N between Greenland and

Norway dividing Atlantic and Arctic; AP, Bering Strait section dividing Arctic and Pacific Ocean; 4, Southern Atlantic-Indian Ocean passage; 5,

Southern Indian Ocean-Pacific passage; 6, Southern Pacific-Atlantic passage; 7, Indonesian Archipelago passage.

3.3 Volume transport and its compensationFigure 4 depicts the typical cross sections and interbasin

passages in the global ocean. The volume, heat and salt trans-

Fig.4. The specific transoceanic sections and inter-basin passages.

ports through these cross sections have been calculated andlisted in Table 2.

In the Atlantic, the volume transports through zonal sec-tion 1A (–1.603×106 m3/s ), 2A (–1.585×106 m3/s) and 3A(–1.089×106 m3/s) are relatively close since there are no signifi-cant straits and passages at both eastern and western bound-aries. This net southward volume transport comes from theArctic Ocean. Although the Norwegian Current flows north-ward, the southward East Greenland Current (EGC) is appar-ently stronger, therefore making a net –1.422×106 m3/s (south-ward) of volume transport at AA, the cross section of 70◦N be-tween Greenland and Norway. This amount of mass loss in theArctic Ocean is compensated by the northward through flow inthe Bering Strait (1.267×106 m3/s) and net P-E (precipitationminus evaporation) into the Arctic Ocean. As we have found,the net P-E is 0.155×106 m3/s in the Arctic Ocean, which exactly

ZHU Yaohua et al. Acta Oceanol. Sin., 2014, Vol. 33, No. 1, P. 11–23 15

fills the gap between 1.422×106 m3/s and 1.267×106 m3/s.Figure 5a depicts the surface currents in the Arctic re-

gion, showing clearly that there exist northward currents in theBering Strait, clockwise circulation in the Canadian basin andanticlockwise in the Barents Sea, Greenland Sea and NorwegianSea area around Svalbard. The current speed can reach 0.5 m/sin the Bering Strait, with an averaged value of 0.2 m/s in theStrait region. Compared with previous observational data andnumerical models, this article presents quite reasonable result-s on the mean current speed, volume transport and sea levelslope structure in the Bering Strait. Overland and Roach (1987)

established a two-dimensional barotropic numerical model forthe Bering Sea and Chukchi Sea, obtained 0.4 m sea level dif-ference between Bering Sea and Chukchi Sea and 1.1×106 m3/snorthward through flow in the Bering Strait. Li et al. (2005) an-alyzed current data based on the mooring stations during theSecond National Arctic Research Expedition of China in 2003and showed 20 cm/s mean northward current in the BeringStrait, which is slightly less than 24.5 cm/s of one-year mooredmeasurements from the autumn of 1990–1991 by Woodgate etal. (2005).

Fig.5. The surface currents in the Arctic region (a), and near the Indonesian Archipelago (b).

However, the volume transports among the zonal sect-ions in the Pacific and the Indian Ocean are not balanced be-cause of the existence of the Indonesian Archipelago passageand the South China Sea (SCS). Obviously there is neither vol-ume transport balance between cross section 1B and 3B in theIndian Ocean nor in the Pacific between cross section 1C and3C. In the Indian Ocean, the southward volume transport of15.48×106 m3/s at 3B comes from section 7 (13.53×106 m3/s)and the Torres Strait (2.3×106 m3/s) since only 0.024×106 m3/sof contribution from the equatorial Indian Ocean section 1B.In the Pacific Ocean, the northward volume transport at 1C(0.605×106 m3/s) is close to 2C (0.95×106 m3/s), but hugely dif-ferent from 3C (17.2×106 m3/s). But if we calculate the volumetransport through the equator from the crucial point 132.5◦E(east of Sorong, the northernmost point of the New Guinea Is-land) to the eastern end of the Pacific, then the volume trans-port is 14×106 m3/s. That means the volume transport from100◦ to 132.5◦E through the equator is –13.395×106 m3/s (a-gain, negative value indicates southward transport), which isdecomposed to –1.41×106 m3/s from the Karimata Strait and–11.99×106 m3/s from the Makassar Strait and the Maluku Sea.Thus, the northward 17.19×106 m3/s volume transport throughcross section 3C is roughly balanced by –2.3×106 m3/s from theTorres Strait, and above 14×106 m3/s through the equator, thelatter joins the North equatorial Current (NEC) and then tur-ns westward. The NEC is divided into the south branch andnorth branch before it approaches Philippines Peninsula. Thesouth branch mostly joins the Equatorial Counter Current (EC-C). Meanwhile, the north branch becomes strengthened at thewestern boundary. Some of the latter enters the Luzon Straitwestward and its majority forms the Kuroshio and flows north-ward from the area east of Taiwan.

Figure 5b shows the surface current distribution in theIndonesian Archipelago region. A part of NEC south branch,flows into the Sulawesi Sea and the Maluku Sea together withthe SCS branch from the Sulu Archipelago. All the southwardbranches from the Karimata Strait, the Makassar Strait and theMaluku Sea converge in the Banda Sea, with an average rate of13.53×106 m3/s, and further flow out toward the Indian Oceanthrough the Timor Sea together with 2.3×106 m3/s westwardcurrent from the Torres Strait.

This above 13.53×106 m3/s of the interbasin volume trans-port carried by the ITF through section 7, from the Pacific tothe Indian Ocean, is of great significance. As the sole interbasinpassage in the tropical region, the Indonesian Archipelago pas-sage acts an “express way” to fulfill the global mass, heat and salttransport balance “quickly”. Gordon (2010) calculated the aver-age strength of the ITF based on 3 years observed data from the“INSTANT” project. His 15×106 m3/s of the observational resultis in good agreement with our 13.53×106 m3/s of the numericalmodeling result. To some extent, this proves the reliability of thenumerical result of this article.

There are indeed some important water exchanges fromthe Luzon Strait, the Taiwan Strait, the Karimata Strait, the Su-lu Sea and the Lombok Strait and even other straits around theSCS, but the grid resolution of the numerical model in this ar-ticle is not fine enough to reproduce their accurate values. Inorder to integrate the global ocean over a thousand years fromthe state of rest to equilibrium state, some scientists (Huismanet al., 2009; Marotzke and Willebrand, 1991) even used 4◦×4◦grid resolution to reduce the computational time.

In the southern ocean, the eastward volume transportthrough meridional section 4 (292×106 m3/s) and the south-ward transport through the zonal section 3B (15×106 m3/s) is

16 ZHU Yaohua et al. Acta Oceanol. Sin., 2014, Vol. 33, No. 1, P. 11–23

equal to the volume transport through the meridional section5 (307×106 m3/s), meanwhile the latter is decomposed into291×106 m3/s eastward on the meridional section 6 and 17×106

m3/s northward on zonal section 3C. The 291×106 m3/s on thesection 6 and 1×106 m3/s southward transport on the section 3Amatches 292×106 m3/s on section 4. Hence, the interbasin masstransports at the Southern Atlantic-Indian Ocean, the SouthernIndian Ocean-Pacific and the Southern Pacific-Atlantic are di-rectly balanced by the Southern Ocean meridional transports.

On the north Pacific zonal section 2C, the northward vol-ume transport 0.953×106 m3/s is less than the Pacific-Arctic ex-change (1.267×106 m3/s) by 0.314×106 m3/s, which is just com-pensated by the net P-E rate. As we have calculated, the netP-E rate over the 30◦–66◦N northern Pacific Ocean is 0.31×106

m3/s. Similarly, there is a small volume transport difference be-tween the northern Atlantic sections 2A (–1.585×106 m3/s) andAA (–1.422×106 m3/s), which is mainly caused by the net P-Erate 0.12×106 m3/s into 30◦–70◦N northern Atlantic Ocean andpartly induced by the southward flow (–0.05×106 m3/s) from theBaffin Bay and the Davis Strait which joins the Labrador cur-rent afterwards. For the same reason, it is easy to understandthe variation of the volume transports at the zonal section 1A(–1.603×106 m3/s), 2A (–1.585×106 m3/s) and 3A (–1.089×106

m3/s).The annual mean P-E rate is illustrated in Fig. 6. The most

notable feature is the intensive P-E at the equatorial Pacific, upto 20 mm/d at the eastern end of the Pacific. The IndonesianArchipelago area together with its vicinity of the western Pacif-ic, the eastern Indian Ocean, gains extensive rainfall of up to 15mm/d since it’s in between the Pacific warm pool and IndianOcean warm pool. Northern Pacific, northern Atlantic, ArcticOcean and ACCR gain net P-E as well. However, the subtropicalzones at both hemispheres show negative P-E up to 5 mm/d,with a maximum value up to 10 mm/d off the Western Aus-tralian coast. As mentioned above, precipitation is explained tobe an important compensation of the mass transport balance.

Fig.6. Annual mean P-E (precipitation minus evapora-tion) rate into the ocean.

3.4 Heat transport and its compensationFigure 7 depicts the vertically integrated meridional heat

transport rate on zonal grids. The meridional heat transportrate is distributed between 0.5 PW (northward) and –0.5 PW(southward) in most of the world ocean, but is intensified in theWBC areas. It reaches 2 PW at 35◦N of Gulf Stream region, 1.5PW in the Kuroshio region and northeast of Madagascar. It ap-proaches 1 PW to the east of Tanzania, Papua New Guinea and

northeast of the Drake Passage. All of the above intensified pos-itive meridional heat transport rates are accompanied with thestrong northward WBC. Therefore it is easy to find the intensi-fied negative rates in the strong southward WBC area, especiallyin the Agulhas current, East Australia current and Brazil currentregion.

Fig.7. Vertically integrated meridional heat transporton zonal grids, which is computed from z -integral ofcpρdx v t /1015 where dx ,v and t are the zonal grid width(m), the north component of current velocity and tem-perature respectively.

In the northern Atlantic and Pacific (north of 40◦N), thepositive heat transport rate covers a majority area and thuscauses a net northward heat transport. In the subtropicalzones of the Atlantic and the Pacific (20◦ to 40◦N), the inten-sified positive meridional heat transports of the Gulf Streamand Kuroshio dominate the northward heat transport. It clear-ly shows that the Agulhas Current carries a vast amount ofheat from the subtropical zone to the ACCR to fulfill equatori-al Pacific-Indonesian Archipelago-Indian Ocean-ACCR heat cir-culation. It is worth mentioning that as strong WBCs, the BrazilCurrent carries a large amount of heat southward from 35◦ to40◦S, meanwhile the Malvinas Current (i.e., Falkland Current)carries a huge amount of heat northward from northern sec-tor of Drake Passage from 50◦ to 40◦S, which forms the Brazil-Malvinas confluence.

The remarkable features of heat transports listed in Table 2are the positive values through sections 1A (0.239 PW), 2A (0.687PW), 2C (0.474 PW) and negative values through sections 3A(–0.109 PW), 3B (–1.737 PW), 3C (–0.048 PW). They show thatthe heat is transported poleward from the tropical zone. In theAtlantic, however, it is transported northward across the equa-tor even though the mass transport is southward there.

Unlike the volume transports in the Atlantic that are alltoward the south and vary little as shown in Table 2 for crosssections 1A, 2A and 3A, the heat transports change significantly.Figure 8a indicates that the meridional heat transport is north-ward in the Atlantic and increases to its peak of 0.75 PW near35◦N, and decreases all the way back to 0 in the Arctic region.These features agree with Hall and Bryden’s (1982) result in Fig.8b, except that the model produced heat transport is southwardbetween 20◦ and 35◦S. The peak value of 0.75 PW is 20% lessthan Hall and Bryden’s result whereas the model produced re-sult is 0.2 PW higher than Hall and Bryden’s at 70◦N. As shownin Fig. 8b, the magnitudes of heat transports vary signif-

ZHU Yaohua et al. Acta Oceanol. Sin., 2014, Vol. 33, No. 1, P. 11–23 17

Fig.8. Model produced meridional heat transport in Atlantic (a), in PW and published results adopted from Hall and Bryden(1982)

derived from integrating BONKEWS (1980) air-sea exchange values(b). Single points are values obtained by the authors using the

direct method: Be, BENNETT; Br, BRYAN; BH, BRYDEN and HALL; HB, HALL and BRYDEN; R, ROEMMICH; W, WUNSCH.

icantly from models and our result in this article presents asmoother tendency.

Figure 9a is the northward heat transport across each par-allel of latitude in the global ocean, clearly showing that heat

is transported poleward from the equator, with a peak of 1.2PW near 35◦N and a trough of –2.0 PW around 30◦S. This resultis consistent with the result of Ganachaud and Wunsch (2002)(Fig. 9b) and estimates of Trenberth et al. (2001).

Fig.9. Model produced northward heat transport by the global ocean across each circle of latitude (a), and published results

adopted from Ganachaud and Wunsch (2002) (b).

Figure 10 illustrates the annual mean net heat flux into theglobal ocean. The peaks of the net heat flux into the ocean oc-cur in the equatorial area with their maximum values up to 100W/m2 at both eastern and western ends of the Pacific. The o-cean gains net heat flux in subtropical area, northern ACCR,northeast Pacific and eastern Atlantic coast. In the area between40◦ and 50◦S in Atlantic, the strong heat flux of up to 80 W/m2

according to Fig. 10, is superimposed to the Malvinas Currentregion, therefore enlarges the northward meridional heat trans-port. The troughs occur at WBC regions in the northern hemi-sphere, where the Gulf Stream and Kuroshio release heat up to200 and 150 W/m2 respectively. The ocean loses net heat fluxin the middle latitude areas (both northern and southern hemi-sphere), in the Arctic Ocean and the southern ACCR. There arestill a few other areas notably losing net heat flux, e.g., in theNorwegian Sea, the western coast of Australia and south of CapeAgulhas. Around 45◦N of western Atlantic, i.e., east of Nova Sco-

Fig.10. Annual mean surface heat flux (positive down-ward).

18 ZHU Yaohua et al. Acta Oceanol. Sin., 2014, Vol. 33, No. 1, P. 11–23

tia, seems an exception. Over there the ocean gains net heat fluxup to 80 W/m2.

Thus for the basin-scale oceanic region, the heat transportrates are not always balanced by incoming and outgoing valueson the specified cross sections due to a large amount of heat fluxinvolved. Only for certain regions with around zero total heatflux, the heat transport rates can be approximately balanced.

Figure 11 illustrates the zonally averaged heat flux shownin Fig. 10. Huge net heat flux (over 60 W/m2) is gained near theequator and decreases dramatically in the subtropical region.It seems almost symmetric at both hemispheres, with fluctua-tions under negative range.

Fig.11. Zonal mean surface heat flux into the globalocean(positive downward).

Unlike the volume transports which are roughly balancedon zonal Atlantic sections, the heat transports on 1A (0.239PW), 2A (0.687 PW) and 3A (–0.109 PW) have no way to ap-proach a balance, neither on section 5 (4.181 PW), section 6(3.509 PW) and 3C (–0.048 PW). Again, it indicates that thesurface heat flux must be considered while balancing the heattransport rates in basin-scale oceans.

Heat transport analysis denotes the heat is transportednorthward in the north Atlantic even though the volume trans-port is southward. That maintains the anomalously warm win-ter air temperatures enjoyed by northern Europe. In the globalocean, the heat is meridionally transported poleward from thetropical region. Therefore, the ocean circulation system acts a

“conveyor belt” to fulfill meridional heat transport, thus chang-ing the world climate.

It is of outstanding significance for the ITF to carry 0.952PW of heat from the Pacific Ocean to the Indian Ocean, whichis roughly equivalent to the amount of the northern Atlanticmeridional heat transport. This “express way” of the interbasinheat transport at low latitudes outlines the ITF’s important ef-fect on global climate change. The southward Agulhas Currentsystem (ACS) demonstrates an important role to transport theheat from ITF to the Southern Ocean, even to the northern At-lantic afterwards, thus keeps the “efficient” heat circulation inthe global ocean.

The following is going to explain how the MOC fulfills thepositive meridional heat transport in the northern Atlantic. Fig-ure 12 shows the vertical profile of zonal mean north velocitycomponent v and potential water temperature t respectivelyat 35◦N in the Atlantic, where the heat transport peak occurs.According to the zonal mean v , it is simply a three-layer ver-tical scheme. The relatively stronger (up to 0.005 m/s) north-ward component v occurs in the upper 1 000 m, correspond-ing to the potential temperature from 10 to 24◦C, meanwhilethe weaker (–0.002 m/s) southward component v occurs in thedepths from 1 000 m to 3 500 m, corresponding to the potentialtemperature of 3◦–10◦C. Obviously, the northward heat trans-port in the upper layer is larger than the southward heat trans-port in the deep layer, the latter performing the NADW’s south-ward movement–the lower limb of the great ocean conveyor.This is the meridional heat transport mechanism, i.e., the GulfStream carries more heat than the deep layer southward NAD-W, that causes the net northward heat transport in the northernAtlantic and thus brings warm winter air temperatures to north-ern Europe. At the depths from 4 000 m to the seabed, there is aslightly northward water movement, which is recognized as theAntarctic Bottom Water (AABW) originating from the vicinity ofthe Antarctic.

Figure 13a is a model produced SST distribution chart inthe Atlantic, simply showing how the temperature decreasespoleward from the equator, which is very consistent with Lev-itus and Boyer (1994) SST (Fig. 13b). Poleward surface currentsalways act as “warm current” whereas equatorward surface cur-rents always act as “cold current”.

3.5 AMOC and its strengthFigure 14 shows the meridional overturning circulation

Fig.12. Vertical profile of zonal mean north component velocity (a) and potential temperature (b) at 35◦N in the Atlantic Ocean.

ZHU Yaohua et al. Acta Oceanol. Sin., 2014, Vol. 33, No. 1, P. 11–23 19

Fig.13. Model produced (a) and Levitus annual mean SST (Levitus and Boyer, 1994) (b) in Atlantic.

(MOC) pattern of the global ocean. The solid line contours in-dicate positive stream functions (clockwise circulation) and thedashed line contours indicate negative stream functions (anti-clockwise). The “zero stream function” occurs in the equatorialarea from ocean surface down to 2 000 m. From the equatorto 60◦N, its depth goes down slightly from 1 500 m, and downto the bottom around 60◦N. In the northern Hemisphere, theMOC is relatively weak. The two positive peaks (correspondingto clockwise sinking) occur at 20◦N with the depth of 100 –150m, and at 50◦N with the depth of 1 000 m. The former happensin the subtropical convergent zone, whereas the latter denotesthe formation of the NADW. In the southern Hemisphere theMOC affects to a much larger extent and deeper since the ACCRlinks the major water exchange passages of all the Southern Pa-cific, Atlantic and Indian Oceans. The strongest anticlockwisecirculation (south sinking) occurs at 60◦S and down to the bot-tom. It can spread to the northern Hemisphere.

Figure 15 shows the Atlantic (a) and Pacific (b) MOC sepa-rately. The biggest difference between them is that the Atlantichas a strong north sinking in the high latitude area to form theNADW and the Pacific Ocean does not form deep water in the

Fig.14. MOC stream functions of the global ocean.

north hemisphere, but forms south sinking instead. This phe-nomenon is coincided with the existing ocean observations anddescribed as the “on” state of the Great Ocean Conveyor, dis-cussed in detail by Huisman et al. (2009) and Marotzke andWillebrand (1991).

Fig.15. MOC stream functions of the Atlantic (a) and the Pacific Ocean (b).

In the north Atlantic, when the Gulf Stream flows north-ward, it releases its heat and gets colder. While this saltier wa-ter becomes colder, it becomes heavier and starts sinking. To-gether with the even colder East Greenland current, it forms lowtemperature and high salinity NADW. It spreads between 40◦Nand 65◦N with a depth of up to 2 000 m. The Atlantic MOCstream function shows a strength up to 20×106 m3/s in this arti-

cle, just supported by Gordon’s (1986) estimation of the NADWvolume. Broecker’s (1991) revised radiocarbon-based estimateof the flux for the NADW from 23×106 m3/s to 20×106 m3/s. Heemphasized that “it is difficult to assess the error in this esti-mate but it is probably on the order of 25% (i.e.,±5×106 m3/s)”.Plenty of previous studies estimated the Atlantic MOC strengthby different approaches. For example, Roemmich and Wunsch

20 ZHU Yaohua et al. Acta Oceanol. Sin., 2014, Vol. 33, No. 1, P. 11–23

(1985) estimated 17×106 m3/s of volume transport at the 24◦NAtlantic by hydrographic section data; Talley et al. (2003) esti-mated (18±5) ×106 m3/s of the NADW formation; Ganachaudand Wunsch (2000) obtained (15±2) ×106 m3/s of the NADWoverturning in the high latitudes by the inverse model.

Underneath the NADW is the Atlantic bottom water. Itclearly comes from the Antarctic region with its MOC streamfunction strength of 15×106 m3/s in the central Atlantic. Thepeak of the stream function occurs at the depth 3 500–4 000 mand with its meridional span from 10◦N to 30◦S.

According to Fig. 12, the zonal mean northward veloci-ty component v at 35◦N is divided into three vertical layers:theupper 800 m of the northward compensating current; the in-termediate/deep layer from 800 to 4 000 m of southward deepcurrent of NADW; the northward bottom current, from 4 000m to the seabed. To calculate the mass transport rates of eachstages and compare with Gordon and Broecker’s MOC strength,50◦N is chosen as the typical latitude of NADW, with the follow-ing findings: 16.02×106 m3/s northward mass transport rate forthe upper layer (0–876 m); –18.72×106 m3/s mass transport rate(southward) for the intermediate/deep water (876–3 854 m);1.914×106 m3/s northward mass transport rate for the bottomwater (3 854–5 072 m). This is highly consistent with Schmitz’sthree layer thermohaline Great Ocean Conveyor scheme, notonly on the stage depths, but also on the mass transport ratesand the extent of NADW. The meridional mass transport rate of18.72×106 m3/s also coincides with 20×106 m3/s of the MOCstrength.

In the north Pacific, however, the Kuroshio is not ascold as the Gulf Stream when it flows northward, since thePacific-Arctic passage (the Bering Strait) is much narrower thanAtlantic-Arctic passage (the Greenland Sea and the NorwegianSea). Furthermore, in the northmost Pacific, it is not the samemechanism as in the North Atlantic where strong northerly galedrives the cold water from the Arctic Ocean convectively mix-

ing. Thus less cold water forms sinking to the depth of 1 000m and only with the stream function strength of 15×106 m3/s.This is the so-called formation of the north Pacific intermediatewater. Underneath the Pacific intermediate water is the Pacif-ic deep water and bottom water, which come from the Atlanticdeep water and bottom water. The Pacific does not form its owndeep water and bottom water.

3.6 Three renewals of NADW and their relations to the inter-

basin exchangesIn this section, the model-produced result shows more de-

tailed descriptions to explain how the NADW is related to the in-terbasin exchanges and meridional transports, thus being partof the Great Ocean Conveyor.

According to Schmitz (1995) and Gordon (1986), about10×106 m3/s of the intermediate water from ACCR flowsthrough the northern sector of Drake Passage, becomes in-volved in a Malvinas Current-Brazil Current-Subtropical gyreinteraction, and then joins the Benguela current regime (BCR)-North Equatorial Current (NEC) and the Gulf Stream and even-tually becomes the primarily renewal of the NADW. Figure 16adepicts the vertical profile of the zonal mean v (north compo-nent of current speed) at the section of east of Cape Horn. Thepositive v component indicates net northward mass transportrate from the northern sector of the Drake Passage. 14.96×106

m3/s has been integrated at 54◦S, from 65◦ to 54◦W. It providesthe flow to become the above mentioned 10×106 m3/s of theprimarily renewal of the NADW. Both Schmitz and Gordon esti-mated O (10×106 m3/s) was the magnitude to become the NAD-W renewal from this origin. Even 14.96×106 m3/s is compar-atively close to their O (10×106 m3/s) but there could be stillsome “leaking” out of it. When, where and how it leaks? Toanswer these questions, water properties need be analyzed indetail. As mentioned earlier, this article will not focus on thevertical scheme and paths of the NADW renewals.

Fig.16. Vertical profile of zonally averaged v (north) component at section 54◦S, 65◦–54◦W (east of Cape Horn) (a) and 32◦S,

8◦–18◦E (west of Cape of Good Hope) (b) respectively.

The secondary renewal originates from the O (5×1016 to6×106 m3/s) “leak” from the ITF according to Schmitz (1995)and Gordon (1986), but this amount could be up to 8×106 m3/saccording to our result. It bypasses the Cape Agulhas into theAtlantic Ocean and joins the BCR, then the NEC and the GulfStream, becoming the secondary renewal of the NADW.

Figure 17 shows the vertical profile of the meridional mean

u (east component of current speed) at the section of south ofCape Agulhas. The negative u component denotes net west-ward mass transport rate from the southern Indian Ocean tothe southern Atlantic. A transport of –40.64×106 m3/s has beenintegrated at the section of 34◦–39◦S, 25◦E. It looks much biggerthan the ITF; however, only a small part of it will turn northwardand join the BCR. Most of it will travel back to the ACCR. Figure

ZHU Yaohua et al. Acta Oceanol. Sin., 2014, Vol. 33, No. 1, P. 11–23 21

16b shows the vertical profile of the zonal mean v at the sectionof west of Cape of Good Hope. The positive v component showsthe part from the Southern Indian Ocean and turning north-ward to BCR system. Only 8.68×106 m3/s has been integratedat 32◦S, 8◦–18◦E (coast). This is somewhat close to the result ofSchmitz, which is 5×106 m3/s of the ITF “leaking” to the BCRsystem and becomes the secondary NADW renewal. However,it could be up to 8×106 m3/s according to our numerical result.It is worth emphasizing again that this part of renewal carries alarge amount of heat from the ITF, thus from the tropical area tothe northern Atlantic.

Fig.17. Vertical profile of meridionally averaged u(east) component at section 34◦–39◦S, 25◦E (South ofCape Agulhas).

The third renewal is relatively weak (1.42×106 m3/s) butit’s also an interbasin exchange from the Pacific-Arctic Ocean tothe Atlantic. This amount of the Arctic through flow is just in be-tween the result of Schmitz (1995) (1.5×106 m3/s) and Broeck-er (1991) (1×106 m3/s). It joins the Arctic circulation system

and eventually joins the EGC and becomes the cold part of theNADW renewal. One should notice that this 1.42×106 m3/s ofAtlantic-Arctic exchange rate does not mean only this amountof cold water is coming from the Arctic Ocean. It has been inte-grated at 70◦N section separately, from Greenland to Jan Mayen(10◦W) for the EGC, from Jan Mayen to Norway for the Nor-wegian Current. Therefore the 1.42×106 m3/s southward masstransport rate breaks down to 11.50×106 m3/s of the southwardEGC and 10.08×106 m3/s of the northward Norwegian Current.The cold EGC passes the Denmark Strait and flows into thenorthern Atlantic Ocean. This cold part plays an important roleto cool down the other two high salinity renewals, and alongwith the deep convection forced by atmospheric cooling there,form the cold and high salinity, thus high density water knownas the NADW. The NADW can sink to the abyss and flow south-ward with the rate of 18.72×106 m3/s at 50◦N, forming the con-veyor’s lower limb.

Figure 18a is the model-produced vertical profile of thezonal mean salinity in the Atlantic. It illustrates the upper layerhigh salinity water from the Gulf Stream sinking gradually dur-ing its convergence with the cold EGC. It highly agrees with Lev-itus (1994) result of the high salinity between 20◦ and 30◦N, withthe salinity 35 sinking to the depth of 2 500 m.

These three renewals make the total convergent rate of17.4×106 m3/s, including 1.42×106 m3/s of southward Atlantic-Arctic exchange and northward transport rate of 16×106 m3/sfrom the upper layer 0–876 m at 50◦N Atlantic, which is slight-ly different from the result of Schmitz (1995) estimation of theintermediate layer compensation (14×106 m3/s). However, ac-cording to the stream function in Fig. 15, it is easy to understandthat the magnitudes of the upper layer renewals (16.02×106

m3/s) and deep layer returning current (–18.72×106 m3/s) varyalong with the latitude, and are not always a constant number.

Fig.18. Model produced (a) and Levitus (1994, b) vertical profile of zonal mean salinity in the Atlantic.

A large transport of approximately 20×106 m3/s of AMOCin section 3.5 is in good agreement with the analysis of thethree renewal components of the NADW–O (10×106 m3/s) fromthe ACCR, O (5×106 m3/s) from the ITF (up to 8×106 m3/sin our results), 1.42×106 m3/s from Atlantic-Arctic ocean ex-change in this section. The vertical structure of the zonal meannorth component v in the north Atlantic is highly consistent

with Schmitz’s three layer thermohaline Great Ocean Convey-or scheme, not only at the layer depths but also at the masstransport rates (10×106, 5×106 and 1.5×106 m3/s of Schmitz’srenewals respectively).

4 ConclusionsA nonBoussinesq mass conserving model based on Mom-

22 ZHU Yaohua et al. Acta Oceanol. Sin., 2014, Vol. 33, No. 1, P. 11–23

4p1 has been spun up over 1 400 years from the state of rest to s-tudy the long-term pattern of the general ocean circulation andto investigate the mass and heat transports through the typicalcross-sections, especially the six interbasin exchange passages.Analysis reveals the interbasin exchanges and their relations tothe meridional transports, as well as the renewals of the NADWquantitatively.

(1) The southward net volume transports at the zonal At-lantic sections require a compensation of –1.42×106 m3/s fromthe Atlantic-Arctic interbasin exchange. Thus, the southwardEGC is 1.42×106 m3/s stronger than the northward Norwegiancurrent on volume transport. The former provides colder wa-ter to stimulate the formation of NADW along with the effect ofdeep convection forced by atmospheric cooling there.

(2) The 0.5 m higher of the SSH at the south side of theBering Strait compared with the north side induces a north-ward through flow in the Bering Strait to fulfill a Pacific-Arcticinterbasin exchange of 1.27×106 m3/s. This current joins theArctic Ocean circulation system and eventually flows out fromthe Greenland Sea and the Denmark Strait, therefore complet-ing the Pacific-Arctic-Atlantic interbasin exchange. This inter-basin through flow has been driven by the SSH slope among thenorthern Pacific, the Arctic and the northern Atlantic accord-ing to the authors. As calculated, the P-E rate in the Arctic re-gion (70◦ to 90◦N) is 0.16×106 m3/s, which exactly fills the gapof 1.42×106 m3/s of Atlantic-Arctic exchange and 1.27×106 m3/sPacific-Arctic exchange.

(3) The Indonesian Archipelago passage is the only inter-basin passage in the tropical region, thus acts an “express way”for mass, heat and salt transports from the Pacific to the Indi-an Ocean. It is of outstanding significance for the ITF to carry13.53×106 m3/s of water, 0.952 PW of heat and 0.466 kg/s of saltto fulfill the Pacific-Indian Ocean through flow. By this tropi-cal “short-cut”, the mass, heat and salt balances in the globalocean can be reached “quickly and efficiently”, thus greatly af-fect global climate change. The SSH slope structure of 1.6 min the west Pacific and 0.6 m at 30◦S Indian Ocean constructsITF’s driving force. 13.53×106 m3/s of this article is quite close toGordon’s (2010) 15×106 m3/s of the observational ITF strength,which is based on 3 years observed data from the “INSTANT”project.

(4) As a unique linkage of the southern Pacific, Atlanticand Indian Ocean passage, the ACCR acts an irreplaceable “traf-fic hub” and provides all the upper, intermediate, deep andbottom layers of water to exchange. About 300×106 m3/s ofwater, 4 PW of heat and 1.0 kg/s of salt transports take placehere. Almost all the levels of water masses originate here andreturn here. The difference of interbasin volume transports be-tween southern Atlantic-Indian Ocean passage (292×106 m3/s)and southern Indian Ocean-Pacific passage (307×106 m3/s) isbalanced by the southern Indian Ocean meridional transport(–15×106 m3/s). Similarly, the difference between southern In-dian Ocean-Pacific passage and southern Pacific-Atlantic pas-sage is balanced by the southern Pacific meridional volumetransport, meanwhile the difference between southern Pacific-Atlantic passage and southern Atlantic-Indian Ocean passageis balanced by the southern Atlantic meridional volume trans-port. If P-E rate is taken into account, the above balances wouldmatch better. However, for the heat transport balance, the sur-face heat flux needs be considered. Overall, all the interbasinexchanges are directly related to the meridional transports.

(5) The origins and strengths of the three NADW renewals

have been analyzed, specifically, O (10×106 m3/s) from the in-termediate layer of the northern sector of the Drake Passage,O (5×106 m3/s) from the Pacific (up to 8×106 m3/s in thismodel) through the Indonesian Archipelago passage (ITF) andO (1.42×106 m3/s) from the Pacific-Arctic-Atlantic interbasinexchange. The O (16×106 m3/s) of the northward upper lay-er volume transport at 50◦N zonal Atlantic section gives an ex-planatory notes on the consistency with the total-sum of abovethe NADW renewals. The three vertical layer structure of thezonal mean v component in the north Atlantic is highly con-sistent with three layer thermohaline Great Ocean Conveyorscheme (Schmitz, 1995), not only on the stage depths but al-so on the volume transports.

(6) Even with 1.585×106 m3/s southward volume transportat 30◦N Atlantic zonal cross-section, there is still 0.687 PW ofheat transported northward since the upper/intermediate layerwater has much higher temperature than the southward deepreturning current of the NADW. It is this mechanism of AtlanticMOC and meridional heat transport that causes the warm win-ter air temperatures in northern Europe.

(7) It takes only a few years for the wind-driven cur-rent or shallow marginal circulation to reach a stable state,but at least a thousand years of integration is needed for thethermodynamically-driven circulation to approach an equilib-rium state from the state of rest, otherwise the underdevel-oped thermodynamically-driven circulation causes unreason-able vertical temperature and salinity structures and underesti-mated transport rates for deep ocean thermohaline circulation.Therefore the grid resolution is limited since 1 400 years of in-tegration with a fine grid resolution is a tremendous computa-tional work. The SSH and surface current in the Arctic regioncould be exaggerated due to the exaggerated wind stress effectwithout the consideration of ocean-ice coupling.

(8) To authors’ experience, such a long time scale of inte-gration with a relative coarse resolution but without form draginduced by mesoscale eddies, could cause over- estimated zon-al transport in the ACCR unless unrealistically high coefficientsfor the horizonal viscosity are used. The horizontal viscosity co-efficient of used in this article is significantly less than of Bryan(1986) and Marotzke and Willebrand (1991). This is the mainreason causing the ACC transport rate larger than the observa-tional data. However, according to Marotzke and Willebrand’s(1991) scale analysis taking the no-slip boundary conditions in-to account, an idealized world ocean model of a design similarto his would lead one to expect an ACC transport of 200×106–400×106 m3/s.

Ac k now l e d g e m e n t sThe authors would like to thank the two anonymous re-

viewers for their valuable suggestions, and Shu Qi and SongZhenya for their help on computational environment setup.

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