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: [email protected]
Interbasin exchanges and their roles in global ocean circulation:
A study based on 1 400 years’ spin up of MOM4p1
ZHU 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
Abstract
A global prognostic model based on MOM4p1, which is a primitive equation nonBoussinesq numerical
model, has been integrated with 1 400 years from the state of rest based on the realistic topography to
study 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 and
31 pressure-based vertical levels. The main objective is to investigate the mass and heat transports at interbasin passages and their compensations and roles in the global ocean circulation under equilibrium state of
long-term spin up. The kinetic energy analysis divides the spin up process into three stages: the quasi-stable
state of wind driven current, the growing phase of thermodynamical circulation and the equilibrium state of
thermohaline circulation. It is essential to spin up over a thousand years in order to reach the thermohaline
equilibrium state from a state of rest. The Arctic Throughflow from the Bering Strait to the Greenland Sea
and the Indonesian Throughflow (ITF) are captured and examined with their compensations and existing
data. Analysis reveals that the slope structures of sea surface height are the dynamical driving mechanism
of the Pacific-Arctic-Atlantic throughflow and ITF. The analysis denotes, in spite of O (1.4×106 m3 /s) of the
southward volume transport in the northern Atlantic, that there is still O (1 PW) of heat transported northward since the northward currents in the upper layer carry much higher temperature water than the southward flowing northern Atlantic deep water (NADW). Meridional volume and heat transports are focused on
the contributions to NADW renewals and Atlantic meridional overturning circulation (AMOC). Quantitative
descriptions of the interbasin exchanges are explained by meridional compensations and supported by previous observations and numerical modeling results. Analysis indicates that the volume and heat exchanges
on the interbasin passages proposed in this article manifest their hub roles in the Great Ocean Conveyor
System.
Key words: numerical modeling, global ocean, interbasin exchange, meridional transport, meridional overturning circulation
Citation: Zhu Yaohua, Wei Zexun, Fang Guohong, Wang Yonggang, Guan Yuping. 2014. Interbasin exchanges and their roles
in 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
al results and proposed a four-layer thermohaline flow scheme
based on interbasin water exchange. He illustrated his thermohaline scheme, including bottom water, deep water, intermediate water and upper layer compensation water, and estimated
the volume transport rates. Both Broecker’s two-layer scheme
and Schmitz’s four-layer scheme presented canonical pictures
for global ocean thermohaline circulation.
Huisman et al. (2009) and Marotzke and Willebrand (1991)
employed GFDL’s (geophysical fluid dynamic laboratory) modular ocean model (MOM2) with idealized rectangle Atlantic and
Pacific Ocean and 4◦ ×4◦ coarse grid to integrate thousands of
years to study the multiple equilibria of thermohaline circulation. In their studies, coarse grids and large steps were applied
1 Introduction
The global-scale circulation has long been one of the
oceanography’s most challenging and exciting research topics. A century ago, Pillsbury (1912) pointed out that global
ocean circulation transports heat poleward from the equator.
But oceanographers did not focus on the heat transport rate
until several decades ago. There has been a developing focus
on the world oceanic thermohaline circulation since it is immediately related to the global climate change. Broecker (1987,
1991) introduced the ocean conveyor belt terminology and twolayer thermohaline flow scheme to study the deep layer circulation and upper layer compensation currents. Schmitz (1995)
summarized updated research achievements and observation-
Foundation item: The National Basic Research Program Grant of China under contact No. 2011CB403502; the International Cooperation Program
Grant of China under contact No. 2010DFB23580; the International Cooperation Program of State Oceanic Administration of China under contract
No. 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: [email protected]
1
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ZHU Yaohua et al. Acta Oceanol. Sin., 2014, Vol. 33, No. 1, P. 11–23
to save computational resources. Analyses were focused to find
the equilibrium states and shift criterion rather than the real thermohaline circulation structure. Wei et al. (2000) integrated MOM2 for 11 years with 1◦ global resolution and (1 /3 )◦
high resolution nested grid to study both wind driven and thermodynamically driven general circulation. Typical zonal and
meridional cross sections were selected to analyze climatological mass, heat and salt transport rates in the global ocean.
The net northward mean through flow in the Bering Strait and
the net southward mass transport in the Atlantic, thus PacificArctic-Atlantic interbasin exchange, were not indicated.
Dong et al. (2011) employed a global circulation model
for the earth simulator (OFES), in which the model code was
based on the MOM3 and the spin up run was 50 years from an
initial condition at rest with observed mean hydrographic data. Detailed analyses were focused on the role of interocean exchanges on decadal variations of the meridional heat transport
in the South Atlantic. However, their numerical results would
be different if the spin up run lasted for a thousand years more.
Therefore, it becomes necessary and useful to simulate
the sufficiently developed state of the general ocean circulation
without initial background constraint whereby under real topography, relatively fine resolution and realistic climatological
forces can be applied.
Along with the development of computer technology and
thanks to GFLD’s endeavor, it becomes possible to integrate the
realistic topography global ocean for thousands of years with
relatively high resolution grids. This article presents an integration of 1 400 years from the state of rest to the equilibrium state in order to investigate the long-term pattern of general circulation system, based on the combined wind-driven
and thermodynamically-driven mechanism. The nonBoussinesq mass conserving MOM4p1 includes conservative temperature and pressure based vertical coordinates. The analyses
in this article indicate that the interbasin exchanges are al-
ways closely related to meridional transports. The interbasin
exchange results of this article reveal a Pacific-Arctic-AtlanticIndian Ocean-Pacific circle transport pattern which composes
a worldwide ocean circulation.
There are six interbasin exchange passages in the world
ocean–two in the Arctic region, three in the Antarctic Circumpolar Current regime (ACCR) and one in the tropical area. The
Greenland Sea and the Norwegian Sea connect the Atlantic and
Arctic Ocean, therefore is an Atlantic-Arctic exchange passage
(AA section). The Bering Strait is another interbasin passage
connecting Arctic and Pacific Ocean (AP section). ACCR links
the Southern Pacific, Atlantic and Indian Ocean, therefore is a
unique linkage for the great oceans to exchange bottom, deep,
intermediate and upper layer water. There are three interbasin exchange passages in ACCR, including Southern AtlanticIndian Ocean passage (SAI), Southern Indian Ocean-Pacific
passage (SIP) and Southern Pacific-Atlantic passage (SPA). The
Indonesian Archipelago passage is the sole passage which connects the Pacific and Indian Ocean in the tropical region, thus
is the only “express way” for mass, heat and salt transports in
the world ocean. Obviously, it is of great importance to focus
on these passages’ transports because of their hub roles in the
global ocean circulation and their relations to meridional transports and Great Ocean Conveyor as well.
2 Model configuration
The MOM4p1 is one of the most recent versions of the
GFDL ocean model, featured by its generalized vertical coordinates. In this global spherical model, the horizontal resolution of 1.9◦ ×0.95◦ (192×189 grids) and 31 vertical levels (Table 1) are employed. Pressure-based vertical coordinate p ∗ has
been applied for the mass conserving, nonBoussinesq, free surface ocean primitive equations. Lateral boundary conditions
are cyclic in zonal direction, and solid walls in meridional direction.
Table 1. Mean depths and thickness of vertical levels under p ∗ coordinate
Level Depth range/104 Pa Thickness/104 Pa Central depth/104 Pa
1
2
3
4
5
6
7
8
9
10
11
12
13
14
15
16
0–10.1
10.1–21.9
21.9–35.6
35.6–51.6
51.6–70.1
70.1–91.6
91.6–116.6
116.6–145.7
145.7–179.5
179.5–218.8
218.8–264.5
264.5–317.6
317.6–379.4
379.4–451.1
451.1–534.6
534.6–631.8
10.1
11.8
13.7
16.0
18.5
21.5
25.0
29.1
33.8
39.3
45.7
53.1
61.8
71.7
83.5
97.2
5.1
15.2
28.7
42.6
60.5
79.7
103.6
129.7
161.7
197.3
240.4
288.7
346.6
412.2
490.1
579.0
The constant horizontal viscosity and diffusivity coefficients are taken as A h = 3.0 × 104 m2 /s and K h = 1.0 × 103 m2 /s
and the vertical ones are computed in the model with “chen”
scheme. Time steps used for the internal mode (baroclinic) and
external mode (barotropic) are 4 800 and 40 s respectively. For
more details about the model description, readers are directed
Level Depth range/104 Pa Thickness/104 Pa Central depth/104 Pa
17
18
19
20
21
22
23
24
25
26
27
28
29
30
31
631.8–744.8
744.8–876.3
876.3–1 029.2
1 029.2–1 206.9
1 206.9–1 413.7
1 413.7–1 654.2
1 654.2–1 933.8
1 933.8–2 259.0
2 259.0–2 637.2
2 637.2–3 042.9
3 042.9–3 448.6
3 448.6–3 854.3
3 854.3–4 260.0
4 260.0–4 665.7
4 665.7–5 071.4
113.0
131.5
152.9
177.7
206.8
240.5
279.6
325.2
378.2
405.7
405.7
405.7
405.7
405.7
405.7
684.5
805.1
947.4
1 110.9
1 303.0
1 524.4
1 783.9
2 083.7
2 434.3
2 840.0
3 245.8
3 651.5
4 057.2
4 462.9
4 868.6
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 Arctic
Ocean (IBCAO). The surface boundary conditions are provided
by NOAA National Oceanographic Data Center (NODC) World
Ocean Atlas 1994 (WOA), including monthly mean water flux,
ZHU Yaohua et al. Acta Oceanol. Sin., 2014, Vol. 33, No. 1, P. 11–23
heat flux, sea surface temperature (SST) and sea surface salinity (SSS). The spin up run is started from a state of rest and has
been integrated for 1 400 years to the equilibrium state. The sea
surface is forced by the monthly climatological wind stress, taken from Hellerman and Rosenstein (1983), the water flux and
the heat flux with the surface temperature and salinity restoring
towards above SST and SSS.
Long-term fluctuations in flow patterns are an important
component of ocean climate and thus are the focus of this article. Except the kinetic energy stability analysis, the time variability is not considered in this article, but not because of its
lack of importance or interest. Different from Schmitz (1995),
this article is focused on interbasin mass and heat transports
and their relations to meridional transports and NADW, rather
than on vertical layer schemes and NADW renewals’ path.
Analyses in this article are based on 50 years mean computational result of the equilibrium state.
3 Result analysis and comparison
3.1 Three developing stages of the total kinetic energy
Figure 1 shows the time variation of total kinetic energy of
the global ocean along with the years of integration. The most
significant feature is the three different stages of kinetic energy’s developing process. The total kinetic energy approaches its
quasi-stable state within the first three months and lasts for 60
years. This first stage is the quasi-stable state of wind-driven
current. As we have recorded, the kinetic energy series are
0.478 5, 0.478 2, 0.456 7, 0.450 1, 0.447 6 (unit: 1018 J) corresponding to the integration length of 3, 6, 9 months, 10
and 60 years. The second stage is the growing phase of the
thermodynamically-driven current in which the total kinetic
energy has been growing dramatically until a thousand integrated years with total kinetic energy of 0.658 1×1018 J. The
third stage is the equilibrium state for thermohaline circulation where the kinetic energy stably fluctuates in a very narrow
range.
13
0.659 37×1018 to 0.659 41×1018 J, with the relative deviation
less than 1.0×10−4 within 200 years. Therefore, it is considered as “equilibrium state”. Even though there are some criteria for the definition of equilibrium state, sometimes scientists
do use their own judgment. For example, Marotzke and Willebrand (1991) employed the basin-averaged surface heat flux under O(0.1) W/m2 as an equilibrium judgment.
If just for the purpose of wind driven current or shallow
marginal circulation, it could be enough to integrate the model
for a decade or so, since it can approach equilibrium state much
quicker.
3.2 Sea surface height distribution and its slope structures
Figure 2 depicts the global sea surface height (SSH) distribution. The Pacific Ocean has the highest SSH and the Atlantic
Ocean has the lowest one. The SSH reaches its peak in the subtropical convergent zone and decreases poleward. Both the Pacific and Atlantic have their higher SSH in the western boundary
current (WBC) area and lower values at the eastern ends. SSH
in the Northern Pacific is higher than that in the Arctic Ocean,
and the latter is higher than in the Nordic Sea. This constitutes
a driving force inducing a through flow into the Arctic Ocean
from Pacific and eventually flowing into the northern Atlantic.
Fig.2.
The global ocean circulation model MOM4p1
produced global SSH distribution.
Fig.1.
Total kinetic energy variation along with integrated years.
Therefore, in order to obtain the equilibrium state from
the state of rest, it is necessary to integrate more than a thousand years. Otherwise, the underdeveloped thermodynamically-driven circulation causes unreasonable vertical temperature and salinity structures and underestimated transports for
the deep ocean thermohaline circulation. This is one of the
important conclusions of the article, which is of significance
on the global ocean circulation modeling. From the integration year of 1 210 to 1 400, the total kinetic energy varies from
Figure 3a shows the detailed SSH distribution around the
Arctic region. In the Bering Strait region, the SSH at the south
side is approximately 0.5 m higher than the north side within
the span of two latitudinal degrees (typical values are 0.8 m in
the south and 0.3 m in the north side of the Strait). The SSH
of 0.46 m in the Bering Strait and 0.25 m at 70◦ N of Chukchi
Sea, slightly north of Bering Strait, are significantly higher than
the zonal mean SSH of –0.36 m at 70◦ N cross section between
Greenland and Norway. In the Arctic Ocean, the SSH on the
Canada basin is higher than in Svalbard-Barents Sea side. At
80◦ N, the SSH peak occurs on the Canada basin with a value
between 0.8 and 1.0 m, while the trough occurs in SvalbardBarents Sea side with a value between –0.8 and –1.0 m. This
slope structure of SSH constructs the dynamical mechanism to
drive a Pacific-Arctic-Atlantic interbasin transport. This PacificArctic-Atlantic transport is also an important component of the
Great Ocean Conveyor. It needs to be noted that this model is
not an ocean-ice coupled model, thus the SSH and the upper
layer currents in the Arctic region could be exaggerated, since
without ice covering the wind stress effect is exaggerated.
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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 Indonesian Archipelago region. SSH in the east side of the Indonesian Archipelago with a typical value of 1.6 m is significantly higher than in the west side, meanwhile the north side (west
Pacific region) is higher than that in the south side (south Indian Ocean). Particularly, 0.6 m of the typical magnitude of SSH
at 30◦ S Indian Ocean and 1.6 m of the SSH in the west Pacif-
ic region construct a northeast-southwestward SSH slope. It is
this SSH slope that drives ITF from northeast to southwest of
Indonesian Archipelago.
In the subantarctic frontal zone (SFZ), the SSH decreases
dramatically poleward. The Antarctic coastal area has the lowest SSH in the global ocean, with a typical magnitude of –2.5 m
and 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
1A
1B
1C
2A
2C
3A
3B
3C
AP BERING ST.
AA G & N
4 SAI
5 SIP
6 SPA
7 IAP
0◦ N
0◦ N
0◦ N
30◦ N
30◦ N
30◦ S
30◦ S
30◦ S
66◦ N
70◦ N
75◦ –30◦ S
75◦ –30◦ S
73◦ –53◦ S
7◦ S
55◦ W–15◦ E
40◦ –100◦ E
100◦ E–75◦ W
90◦ W–0◦ E
120◦ E–110◦ W
50◦ W–17◦ E
30◦ –120◦ E
150◦ E–70◦ W
170◦ –167◦ W
25◦ W–20◦ E
20◦ E
135◦ E
67◦ W
110◦ –140◦ E
Volume transport/106 m3 ·s−1
Heat transport/PW
Salt transport/109 kg·s−1
–1.603
0.024
0.605(14.000 from 133◦ E)
–1.585
0.953
–1.089
–15.480
17.190
1.2670
–1.422
292.000
307.600
290.800
–13.530
0.239
–0.225
–0.570
0.687
0.474
–0.109
–1.737
–0.048
0.010
0.247
2.998
4.181
3.509
–0.952
–0.0430
0.0002
0.0360
–0.0470
0.0430
–0.0440
–0.5470
0.5840
0.0410
–0.0380
10.0800
10.6300
10.0300
–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 compensation
Figure 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 interbasin passages.
ports through these cross sections have been calculated and
listed in Table 2.
In the Atlantic, the volume transports through zonal section 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 significant straits and passages at both eastern and western boundaries. This net southward volume transport comes from the
Arctic Ocean. Although the Norwegian Current flows northward, the southward East Greenland Current (EGC) is apparently stronger, therefore making a net –1.422×106 m3 /s (southward) of volume transport at AA, the cross section of 70◦ N between Greenland and Norway. This amount of mass loss in the
Arctic Ocean is compensated by the northward through flow in
the Bering Strait (1.267×106 m3 /s) and net P-E (precipitation
minus 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
fills the gap between 1.422×106 m3 /s and 1.267×106 m3 /s.
Figure 5a depicts the surface currents in the Arctic region, showing clearly that there exist northward currents in the
Bering Strait, clockwise circulation in the Canadian basin and
anticlockwise in the Barents Sea, Greenland Sea and Norwegian
Sea area around Svalbard. The current speed can reach 0.5 m/s
in the Bering Strait, with an averaged value of 0.2 m/s in the
Strait region. Compared with previous observational data and
numerical models, this article presents quite reasonable results on the mean current speed, volume transport and sea level
slope structure in the Bering Strait. Overland and Roach (1987)
15
established a two-dimensional barotropic numerical model for
the Bering Sea and Chukchi Sea, obtained 0.4 m sea level difference between Bering Sea and Chukchi Sea and 1.1×106 m3 /s
northward through flow in the Bering Strait. Li et al. (2005) analyzed current data based on the mooring stations during the
Second National Arctic Research Expedition of China in 2003
and showed 20 cm/s mean northward current in the Bering
Strait, which is slightly less than 24.5 cm/s of one-year moored
measurements from the autumn of 1990–1991 by Woodgate et
al. (2005).
Fig.5. The surface currents in the Arctic region (a), and near the Indonesian Archipelago (b).
However, the volume transports among the zonal sections in the Pacific and the Indian Ocean are not balanced because of the existence of the Indonesian Archipelago passage
and the South China Sea (SCS). Obviously there is neither volume transport balance between cross section 1B and 3B in the
Indian Ocean nor in the Pacific between cross section 1C and
3C. In the Indian Ocean, the southward volume transport of
15.48×106 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 /s
of 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 different from 3C (17.2×106 m3 /s). But if we calculate the volume
transport through the equator from the crucial point 132.5◦ E
(east of Sorong, the northernmost point of the New Guinea Island) to the eastern end of the Pacific, then the volume transport is 14×106 m3 /s. That means the volume transport from
100◦ to 132.5◦ E through the equator is –13.395×106 m3 /s (again, negative value indicates southward transport), which is
decomposed 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 through
cross section 3C is roughly balanced by –2.3×106 m3 /s from the
Torres Strait, and above 14×106 m3 /s through the equator, the
latter joins the North equatorial Current (NEC) and then turns westward. The NEC is divided into the south branch and
north branch before it approaches Philippines Peninsula. The
south branch mostly joins the Equatorial Counter Current (ECC). Meanwhile, the north branch becomes strengthened at the
western boundary. Some of the latter enters the Luzon Strait
westward and its majority forms the Kuroshio and flows northward from the area east of Taiwan.
Figure 5b shows the surface current distribution in the
Indonesian Archipelago region. A part of NEC south branch,
flows into the Sulawesi Sea and the Maluku Sea together with
the SCS branch from the Sulu Archipelago. All the southward
branches from the Karimata Strait, the Makassar Strait and the
Maluku Sea converge in the Banda Sea, with an average rate of
13.53×106 m3 /s, and further flow out toward the Indian Ocean
through the Timor Sea together with 2.3×106 m3 /s westward
current from the Torres Strait.
This above 13.53×106 m3 /s of the interbasin volume transport carried by the ITF through section 7, from the Pacific to
the Indian Ocean, is of great significance. As the sole interbasin
passage in the tropical region, the Indonesian Archipelago passage acts an “express way” to fulfill the global mass, heat and salt
transport balance “quickly”. Gordon (2010) calculated the average strength of the ITF based on 3 years observed data from the
“INSTANT” project. His 15×106 m3 /s of the observational result
is in good agreement with our 13.53×106 m3 /s of the numerical
modeling result. To some extent, this proves the reliability of the
numerical result of this article.
There are indeed some important water exchanges from
the Luzon Strait, the Taiwan Strait, the Karimata Strait, the Sulu Sea and the Lombok Strait and even other straits around the
SCS, but the grid resolution of the numerical model in this article is not fine enough to reproduce their accurate values. In
order to integrate the global ocean over a thousand years from
the state of rest to equilibrium state, some scientists (Huisman
et al., 2009; Marotzke and Willebrand, 1991) even used 4◦ ×4◦
grid resolution to reduce the computational time.
In the southern ocean, the eastward volume transport
through meridional section 4 (292×106 m3 /s) and the southward 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 section
5 (307×106 m3 /s), meanwhile the latter is decomposed into
291×106 m3 /s eastward on the meridional section 6 and 17×106
m3 /s northward on zonal section 3C. The 291×106 m3 /s on the
section 6 and 1×106 m3 /s southward transport on the section 3A
matches 292×106 m3 /s on section 5. Hence, the interbasin mass
transports at the Southern Atlantic-Indian Ocean, the Southern
Indian Ocean-Pacific and the Southern Pacific-Atlantic are directly balanced by the Southern Ocean meridional transports.
On the north Pacific zonal section 2C, the northward volume transport 0.953×106 m3 /s is less than the Pacific-Arctic exchange (1.267×106 m3 /s) by 0.314×106 m3 /s, which is just compensated by the net P-E rate. As we have calculated, the net
P-E rate over the 30◦ –66◦ N northern Pacific Ocean is 0.31×106
m3 /s. Similarly, there is a small volume transport difference between the northern Atlantic sections 2A (–1.585×106 m3 /s) and
AA (–1.422×106 m3 /s), which is mainly caused by the net P-E
rate 0.12×106 m3 /s into 30◦ –70◦ N northern Atlantic Ocean and
partly induced by the southward flow (–0.05×106 m3 /s) from the
Baffin Bay and the Davis Strait which joins the Labrador current afterwards. For the same reason, it is easy to understand
the 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, up
to 20 mm/d at the eastern end of the Pacific. The Indonesian
Archipelago area together with its vicinity of the western Pacific, the eastern Indian Ocean, gains extensive rainfall of up to 15
mm/d since it’s in between the Pacific warm pool and Indian
Ocean warm pool. Northern Pacific, northern Atlantic, Arctic
Ocean and ACCR gain net P-E as well. However, the subtropical
zones at both hemispheres show negative P-E up to 5 mm/d,
with a maximum value up to 10 mm/d off the Western Australian coast. As mentioned above, precipitation is explained to
be an important compensation of the mass transport balance.
Fig.6. Annual mean P-E (precipitation minus evaporation) rate into the ocean.
3.4 Heat transport and its compensation
Figure 7 depicts the vertically integrated meridional heat
transport rate on zonal grids. The meridional heat transport
rate is distributed between 0.5 PW (northward) and –0.5 PW
(southward) in most of the world ocean, but is intensified in the
WBC areas. It reaches 2 PW at 35◦ N of Gulf Stream region, 1.5
PW in the Kuroshio region and northeast of Madagascar. It approaches 1 PW to the east of Tanzania, Papua New Guinea and
northeast of the Drake Passage. All of the above intensified positive meridional heat transport rates are accompanied with the
strong northward WBC. Therefore it is easy to find the intensified negative rates in the strong southward WBC area, especially
in the Agulhas current, East Australia current and Brazil current
region.
Fig.7. Vertically integrated meridional heat transport
on zonal grids, which is computed from z -integral of
c p ρdx v t /1015 where dx , v and t are the zonal grid width
(m), the north component of current velocity and temperature respectively.
In the northern Atlantic and Pacific (north of 40◦ N), the
positive heat transport rate covers a majority area and thus
causes a net northward heat transport. In the subtropical
zones of the Atlantic and the Pacific (20◦ to 40◦ N), the intensified positive meridional heat transports of the Gulf Stream
and Kuroshio dominate the northward heat transport. It clearly shows that the Agulhas Current carries a vast amount of
heat from the subtropical zone to the ACCR to fulfill equatorial Pacific-Indonesian Archipelago-Indian Ocean-ACCR heat circulation. It is worth mentioning that as strong WBCs, the Brazil
Current carries a large amount of heat southward from 35◦ to
40◦ S, meanwhile the Malvinas Current (i.e., Falkland Current)
carries a huge amount of heat northward from northern sector of Drake Passage from 50◦ to 40◦ S, which forms the BrazilMalvinas confluence.
The remarkable features of heat transports listed in Table 2
are the positive values through sections 1A (0.239 PW), 2A (0.687
PW), 2C (0.474 PW) and negative values through sections 3A
(–0.109 PW), 3B (–1.737 PW), 3C (–0.048 PW). They show that
the heat is transported poleward from the tropical zone. In the
Atlantic, however, it is transported northward across the equator even though the mass transport is southward there.
Unlike the volume transports in the Atlantic that are all
toward the south and vary little as shown in Table 2 for cross
sections 1A, 2A and 3A, the heat transports change significantly.
Figure 8a indicates that the meridional heat transport is northward in the Atlantic and increases to its peak of 0.75 PW near
35◦ 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 southward
between 20◦ and 35◦ S. The peak value of 0.75 PW is 20% less
than Hall and Bryden’s result whereas the model produced result is 0.2 PW higher than Hall and Bryden’s at 70◦ N. As shown
in 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 a
smoother tendency.
Figure 9a is the northward heat transport across each parallel of latitude in the global ocean, clearly showing that heat
is transported poleward from the equator, with a peak of 1.2
PW near 35◦ N and a trough of –2.0 PW around 30◦ S. This result
is 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 the
global ocean. The peaks of the net heat flux into the ocean occur in the equatorial area with their maximum values up to 100
W/m2 at both eastern and western ends of the Pacific. The ocean gains net heat flux in subtropical area, northern ACCR,
northeast Pacific and eastern Atlantic coast. In the area between
40◦ and 50◦ S in Atlantic, the strong heat flux of up to 80 W/m2
according to Fig. 10, is superimposed to the Malvinas Current
region, therefore enlarges the northward meridional heat transport. The troughs occur at WBC regions in the northern hemisphere, where the Gulf Stream and Kuroshio release heat up to
200 and 150 W/m2 respectively. The ocean loses net heat flux
in the middle latitude areas (both northern and southern hemisphere), in the Arctic Ocean and the southern ACCR. There are
still a few other areas notably losing net heat flux, e.g., in the
Norwegian Sea, the western coast of Australia and south of Cape
Agulhas. Around 45◦ N of western Atlantic, i.e., east of Nova Sco-
Fig.10. Annual mean surface heat flux (positive downward).
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 flux
up to 80 W/m2 .
Thus for the basin-scale oceanic region, the heat transport
rates are not always balanced by incoming and outgoing values
on the specified cross sections due to a large amount of heat flux
involved. Only for certain regions with around zero total heat
flux, the heat transport rates can be approximately balanced.
Figure 11 illustrates the zonally averaged heat flux shown
in Fig. 10. Huge net heat flux (over 60 W/m2 ) is gained near the
equator and decreases dramatically in the subtropical region.
It seems almost symmetric at both hemispheres, with fluctuations under negative range.
Fig.11.
Zonal mean surface heat flux into the global
ocean(positive downward).
Unlike the volume transports which are roughly balanced
on zonal Atlantic sections, the heat transports on 1A (0.239
PW), 2A (0.687 PW) and 3A (–0.109 PW) have no way to approach a balance, neither on section 5 (4.181 PW), section 6
(3.509 PW) and 3C (–0.048 PW). Again, it indicates that the
surface heat flux must be considered while balancing the heat
transport rates in basin-scale oceans.
Heat transport analysis denotes the heat is transported
northward in the north Atlantic even though the volume transport is southward. That maintains the anomalously warm winter air temperatures enjoyed by northern Europe. In the global
ocean, the heat is meridionally transported poleward from the
tropical region. Therefore, the ocean circulation system acts a
“conveyor belt” to fulfill meridional heat transport, thus changing the world climate.
It is of outstanding significance for the ITF to carry 0.952
PW of heat from the Pacific Ocean to the Indian Ocean, which
is roughly equivalent to the amount of the northern Atlantic
meridional heat transport. This “express way” of the interbasin
heat transport at low latitudes outlines the ITF’s important effect on global climate change. The southward Agulhas Current
system (ACS) demonstrates an important role to transport the
heat from ITF to the Southern Ocean, even to the northern Atlantic afterwards, thus keeps the “efficient” heat circulation in
the global ocean.
The following is going to explain how the MOC fulfills the
positive meridional heat transport in the northern Atlantic. Figure 12 shows the vertical profile of zonal mean north velocity
component v and potential water temperature t respectively
at 35◦ N in the Atlantic, where the heat transport peak occurs.
According to the zonal mean v , it is simply a three-layer vertical scheme. The relatively stronger (up to 0.005 m/s) northward component v occurs in the upper 1 000 m, corresponding to the potential temperature from 10 to 24◦ C, meanwhile
the weaker (–0.002 m/s) southward component v occurs in the
depths from 1 000 m to 3 500 m, corresponding to the potential
temperature of 3◦ –10◦ C. Obviously, the northward heat transport in the upper layer is larger than the southward heat transport in the deep layer, the latter performing the NADW’s southward movement–the lower limb of the great ocean conveyor.
This is the meridional heat transport mechanism, i.e., the Gulf
Stream carries more heat than the deep layer southward NADW, that causes the net northward heat transport in the northern
Atlantic and thus brings warm winter air temperatures to northern Europe. At the depths from 4 000 m to the seabed, there is a
slightly northward water movement, which is recognized as the
Antarctic Bottom Water (AABW) originating from the vicinity of
the Antarctic.
Figure 13a is a model produced SST distribution chart in
the Atlantic, simply showing how the temperature decreases
poleward from the equator, which is very consistent with Levitus and Boyer (1994) SST (Fig. 13b). Poleward surface currents
always act as “warm current” whereas equatorward surface currents always act as “cold current”.
3.5 AMOC and its strength
Figure 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 indicate positive stream functions (clockwise circulation) and the
dashed line contours indicate negative stream functions (anticlockwise). The “zero stream function” occurs in the equatorial
area from ocean surface down to 2 000 m. From the equator
to 60◦ N, its depth goes down slightly from 1 500 m, and down
to the bottom around 60◦ N. In the northern Hemisphere, the
MOC is relatively weak. The two positive peaks (corresponding
to clockwise sinking) occur at 20◦ N with the depth of 100 –150
m, and at 50◦ N with the depth of 1 000 m. The former happens
in the subtropical convergent zone, whereas the latter denotes
the formation of the NADW. In the southern Hemisphere the
MOC affects to a much larger extent and deeper since the ACCR
links the major water exchange passages of all the Southern Pacific, Atlantic and Indian Oceans. The strongest anticlockwise
circulation (south sinking) occurs at 60◦ S and down to the bottom. It can spread to the northern Hemisphere.
Figure 15 shows the Atlantic (a) and Pacific (b) MOC separately. The biggest difference between them is that the Atlantic
has a strong north sinking in the high latitude area to form the
NADW 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 phenomenon is coincided with the existing ocean observations and
described as the “on” state of the Great Ocean Conveyor, discussed in detail by Huisman et al. (2009) and Marotzke and
Willebrand (1991).
Fig.15. MOC stream functions of the Atlantic (a) and the Pacific Ocean (b).
In the north Atlantic, when the Gulf Stream flows northward, it releases its heat and gets colder. While this saltier water becomes colder, it becomes heavier and starts sinking. Together with the even colder East Greenland current, it forms low
temperature and high salinity NADW. It spreads between 40◦ N
and 65◦ N with a depth of up to 2 000 m. The Atlantic MOC
stream function shows a strength up to 20×106 m3 /s in this arti-
cle, just supported by Gordon’s (1986) estimation of the NADW
volume. Broecker’s (1991) revised radiocarbon-based estimate
of the flux for the NADW from 23×106 m3 /s to 20×106 m3 /s. He
emphasized that “it is difficult to assess the error in this estimate but it is probably on the order of 25% (i.e., ±5×106 m3 /s)”.
Plenty of previous studies estimated the Atlantic MOC strength
by different approaches. For example, Roemmich and Wunsch
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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◦ N
Atlantic by hydrographic section data; Talley et al. (2003) estimated (18±5) ×106 m3 /s of the NADW formation; Ganachaud
and Wunsch (2000) obtained (15±2) ×106 m3 /s of the NADW
overturning in the high latitudes by the inverse model.
Underneath the NADW is the Atlantic bottom water. It
clearly comes from the Antarctic region with its MOC stream
function strength of 15×106 m3 /s in the central Atlantic. The
peak of the stream function occurs at the depth 3 500–4 000 m
and with its meridional span from 10◦ N to 30◦ S.
According to Fig. 12, the zonal mean northward velocity component v at 35◦ N is divided into three vertical layers:the
upper 800 m of the northward compensating current; the intermediate/deep layer from 800 to 4 000 m of southward deep
current of NADW; the northward bottom current, from 4 000
m to the seabed. To calculate the mass transport rates of each
stages and compare with Gordon and Broecker’s MOC strength,
50◦ N is chosen as the typical latitude of NADW, with the following findings: 16.02×106 m3 /s northward mass transport rate for
the 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 bottom
water (3 854–5 072 m). This is highly consistent with Schmitz’s
three layer thermohaline Great Ocean Conveyor scheme, not
only on the stage depths, but also on the mass transport rates
and the extent of NADW. The meridional mass transport rate of
18.72×106 m3 /s also coincides with 20×106 m3 /s of the MOC
strength.
In the north Pacific, however, the Kuroshio is not as
cold as the Gulf Stream when it flows northward, since the
Pacific-Arctic passage (the Bering Strait) is much narrower than
Atlantic-Arctic passage (the Greenland Sea and the Norwegian
Sea). Furthermore, in the northmost Pacific, it is not the same
mechanism as in the North Atlantic where strong northerly gale
drives the cold water from the Arctic Ocean convectively mix-
ing. Thus less cold water forms sinking to the depth of 1 000
m and only with the stream function strength of 15×106 m3 /s.
This is the so-called formation of the north Pacific intermediate
water. Underneath the Pacific intermediate water is the Pacific deep water and bottom water, which come from the Atlantic
deep water and bottom water. The Pacific does not form its own
deep water and bottom water.
3.6 Three renewals of NADW and their relations to the interbasin exchanges
In this section, the model-produced result shows more detailed descriptions to explain how the NADW is related to the interbasin exchanges and meridional transports, thus being part
of the Great Ocean Conveyor.
According to Schmitz (1995) and Gordon (1986), about
10×106 m3 /s of the intermediate water from ACCR flows
through the northern sector of Drake Passage, becomes involved in a Malvinas Current-Brazil Current-Subtropical gyre
interaction, and then joins the Benguela current regime (BCR)North Equatorial Current (NEC) and the Gulf Stream and eventually becomes the primarily renewal of the NADW. Figure 16a
depicts the vertical profile of the zonal mean v (north component of current speed) at the section of east of Cape Horn. The
positive v component indicates net northward mass transport
rate 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 provides
the flow to become the above mentioned 10×106 m3 /s of the
primarily renewal of the NADW. Both Schmitz and Gordon estimated O (10×106 m3 /s) was the magnitude to become the NADW renewal from this origin. Even 14.96×106 m3 /s is comparatively close to their O (10×106 m3 /s) but there could be still
some “leaking” out of it. When, where and how it leaks? To
answer these questions, water properties need be analyzed in
detail. As mentioned earlier, this article will not focus on the
vertical scheme and paths of the NADW renewals.
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.
Fig.16.
The secondary renewal originates from the O (5×1016 to
m3 /s) “leak” from the ITF according to Schmitz (1995)
and Gordon (1986), but this amount could be up to 8×106 m3 /s
according to our result. It bypasses the Cape Agulhas into the
Atlantic Ocean and joins the BCR, then the NEC and the Gulf
Stream, becoming the secondary renewal of the NADW.
Figure 17 shows the vertical profile of the meridional mean
6×106
u (east component of current speed) at the section of south of
Cape Agulhas. The negative u component denotes net westward mass transport rate from the southern Indian Ocean to
the southern Atlantic. A transport of –40.64×106 m3 /s has been
integrated at the section of 34◦ –39◦ S, 25◦ E. It looks much bigger
than the ITF; however, only a small part of it will turn northward
and 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
16b shows the vertical profile of the zonal mean v at the section
of west of Cape of Good Hope. The positive v component shows
the part from the Southern Indian Ocean and turning northward to BCR system. Only 8.68×106 m3 /s has been integrated
at 32◦ S, 8◦ –18◦ E (coast). This is somewhat close to the result of
Schmitz, which is 5×106 m3 /s of the ITF “leaking” to the BCR
system 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 a
large amount of heat from the ITF, thus from the tropical area to
the northern Atlantic.
Fig.17.
Vertical profile of meridionally averaged u
(east) component at section 34◦ –39◦ S, 25◦ E (South of
Cape Agulhas).
The third renewal is relatively weak (1.42×106 m3 /s) but
it’s also an interbasin exchange from the Pacific-Arctic Ocean to
the Atlantic. This amount of the Arctic through flow is just in between the result of Schmitz (1995) (1.5×106 m3 /s) and Broecker (1991) (1×106 m3 /s). It joins the Arctic circulation system
21
and eventually joins the EGC and becomes the cold part of the
NADW renewal. One should notice that this 1.42×106 m3 /s of
Atlantic-Arctic exchange rate does not mean only this amount
of cold water is coming from the Arctic Ocean. It has been integrated at 70◦ N section separately, from Greenland to Jan Mayen
(10◦ W) for the EGC, from Jan Mayen to Norway for the Norwegian Current. Therefore the 1.42×106 m3 /s southward mass
transport rate breaks down to 11.50×106 m3 /s of the southward
EGC and 10.08×106 m3 /s of the northward Norwegian Current.
The cold EGC passes the Denmark Strait and flows into the
northern Atlantic Ocean. This cold part plays an important role
to cool down the other two high salinity renewals, and along
with the deep convection forced by atmospheric cooling there,
form the cold and high salinity, thus high density water known
as the NADW. The NADW can sink to the abyss and flow southward with the rate of 18.72×106 m3 /s at 50◦ N, forming the conveyor’s lower limb.
Figure 18a is the model-produced vertical profile of the
zonal mean salinity in the Atlantic. It illustrates the upper layer
high salinity water from the Gulf Stream sinking gradually during its convergence with the cold EGC. It highly agrees with Levitus (1994) result of the high salinity between 20◦ and 30◦ N, with
the salinity 35 sinking to the depth of 2 500 m.
These three renewals make the total convergent rate of
17.4×106 m3 /s, including 1.42×106 m3 /s of southward AtlanticArctic exchange and northward transport rate of 16×106 m3 /s
from the upper layer 0–876 m at 50◦ N Atlantic, which is slightly different from the result of Schmitz (1995) estimation of the
intermediate layer compensation (14×106 m3 /s). However, according to the stream function in Fig. 15, it is easy to understand
that the magnitudes of the upper layer renewals (16.02×106
m3 /s) and deep layer returning current (–18.72×106 m3 /s) vary
along 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 AMOC
in section 3.5 is in good agreement with the analysis of the
three renewal components of the NADW–O (10×106 m3 /s) from
the ACCR, O (5×106 m3 /s) from the ITF (up to 8×106 m3 /s
in our results), 1.42×106 m3 /s from Atlantic-Arctic ocean exchange in this section. The vertical structure of the zonal mean
north component v in the north Atlantic is highly consistent
with Schmitz’s three layer thermohaline Great Ocean Conveyor scheme, not only at the layer depths but also at the mass
transport rates (10×106 , 5×106 and 1.5×106 m3 /s of Schmitz’s
renewals respectively).
4 Conclusions
A nonBoussinesq mass conserving model based on Mom-
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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 study the long-term pattern of the general ocean circulation and
to investigate the mass and heat transports through the typical
cross-sections, especially the six interbasin exchange passages.
Analysis reveals the interbasin exchanges and their relations to
the meridional transports, as well as the renewals of the NADW
quantitatively.
(1) The southward net volume transports at the zonal Atlantic sections require a compensation of –1.42×106 m3 /s from
the Atlantic-Arctic interbasin exchange. Thus, the southward
EGC is 1.42×106 m3 /s stronger than the northward Norwegian
current on volume transport. The former provides colder water to stimulate the formation of NADW along with the effect of
deep convection forced by atmospheric cooling there.
(2) The 0.5 m higher of the SSH at the south side of the
Bering Strait compared with the north side induces a northward through flow in the Bering Strait to fulfill a Pacific-Arctic
interbasin exchange of 1.27×106 m3 /s. This current joins the
Arctic Ocean circulation system and eventually flows out from
the Greenland Sea and the Denmark Strait, therefore completing the Pacific-Arctic-Atlantic interbasin exchange. This interbasin through flow has been driven by the SSH slope among the
northern Pacific, the Arctic and the northern Atlantic according to the authors. As calculated, the P-E rate in the Arctic region (70◦ to 90◦ N) is 0.16×106 m3 /s, which exactly fills the gap
of 1.42×106 m3 /s of Atlantic-Arctic exchange and 1.27×106 m3 /s
Pacific-Arctic exchange.
(3) The Indonesian Archipelago passage is the only interbasin passage in the tropical region, thus acts an “express way”
for mass, heat and salt transports from the Pacific to the Indian Ocean. It is of outstanding significance for the ITF to carry
13.53×106 m3 /s of water, 0.952 PW of heat and 0.466 kg/s of salt
to fulfill the Pacific-Indian Ocean through flow. By this tropical “short-cut”, the mass, heat and salt balances in the global
ocean can be reached “quickly and efficiently”, thus greatly affect global climate change. The SSH slope structure of 1.6 m
in the west Pacific and 0.6 m at 30◦ S Indian Ocean constructs
ITF’s driving force. 13.53×106 m3 /s of this article is quite close to
Gordon’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, Atlantic
and Indian Ocean passage, the ACCR acts an irreplaceable “traffic hub” and provides all the upper, intermediate, deep and
bottom layers of water to exchange. About 300×106 m3 /s of
water, 4 PW of heat and 1.0 kg/s of salt transports take place
here. Almost all the levels of water masses originate here and
return here. The difference of interbasin volume transports between southern Atlantic-Indian Ocean passage (292×106 m3 /s)
and southern Indian Ocean-Pacific passage (307×106 m3 /s) is
balanced by the southern Indian Ocean meridional transport
(–15×106 m3 /s). Similarly, the difference between southern Indian Ocean-Pacific passage and southern Pacific-Atlantic passage is balanced by the southern Pacific meridional volume
transport, meanwhile the difference between southern PacificAtlantic passage and southern Atlantic-Indian Ocean passage
is balanced by the southern Atlantic meridional volume transport. If P-E rate is taken into account, the above balances would
match better. However, for the heat transport balance, the surface heat flux needs be considered. Overall, all the interbasin
exchanges 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 intermediate layer of the northern sector of the Drake Passage,
O (5×106 m3 /s) from the Pacific (up to 8×106 m3 /s in this
model) through the Indonesian Archipelago passage (ITF) and
O (1.42×106 m3 /s) from the Pacific-Arctic-Atlantic interbasin
exchange. The O (16×106 m3 /s) of the northward upper layer volume transport at 50◦ N zonal Atlantic section gives an explanatory notes on the consistency with the total-sum of above
the NADW renewals. The three vertical layer structure of the
zonal mean v component in the north Atlantic is highly consistent with three layer thermohaline Great Ocean Conveyor
scheme (Schmitz, 1995), not only on the stage depths but also on the volume transports.
(6) Even with 1.585×106 m3 /s southward volume transport
at 30◦ N Atlantic zonal cross-section, there is still 0.687 PW of
heat transported northward since the upper/intermediate layer
water has much higher temperature than the southward deep
returning current of the NADW. It is this mechanism of Atlantic
MOC and meridional heat transport that causes the warm winter air temperatures in northern Europe.
(7) It takes only a few years for the wind-driven current or shallow marginal circulation to reach a stable state,
but at least a thousand years of integration is needed for the
thermodynamically-driven circulation to approach an equilibrium state from the state of rest, otherwise the underdeveloped thermodynamically-driven circulation causes unreasonable vertical temperature and salinity structures and underestimated transport rates for deep ocean thermohaline circulation.
Therefore the grid resolution is limited since 1 400 years of integration with a fine grid resolution is a tremendous computational work. The SSH and surface current in the Arctic region
could be exaggerated due to the exaggerated wind stress effect
without the consideration of ocean-ice coupling.
(8) To authors’ experience, such a long time scale of integration with a relative coarse resolution but without form drag
induced by mesoscale eddies, could cause over- estimated zonal transport in the ACCR unless unrealistically high coefficients
for the horizonal viscosity are used. The horizontal viscosity coefficient of used in this article is significantly less than of Bryan
(1986) and Marotzke and Willebrand (1991). This is the main
reason causing the ACC transport rate larger than the observational data. However, according to Marotzke and Willebrand’s
(1991) scale analysis taking the no-slip boundary conditions into account, an idealized world ocean model of a design similar
to 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 s
The authors would like to thank the two anonymous reviewers for their valuable suggestions, and Shu Qi and Song
Zhenya for their help on computational environment setup.
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