ARTICLE IN PRESS
Deep-Sea Research I 56 (2009) 1203–1216
Contents lists available at ScienceDirect
Deep-Sea Research I
journal homepage: www.elsevier.com/locate/dsri
The deep-water motion through the Lifamatola Passage and its
contribution to the Indonesian throughflow
Hendrik M. van Aken a,, Irsan S. Brodjonegoro b, Indra Jaya c
a
b
c
NIOZ Royal Netherlands Institute for Sea Research, Texel, The Netherlands
ITB Institute of Technology Bandung, Bandung, Indonesia
IPB Bogor Agricultural University, Indonesia
a r t i c l e in fo
abstract
Article history:
Received 28 August 2008
Received in revised form
26 January 2009
Accepted 1 February 2009
Available online 7 February 2009
In order to estimate the contribution of cold Pacific deep water to the Indonesian
throughflow (ITF) and the flushing of the deep Banda Sea, a current meter mooring
has been deployed for nearly 3 years on the sill in the Lifamatola Passage as part of
the International Nusantara Stratification and Transport (INSTANT) programme. The
velocity, temperature, and salinity data, obtained from the mooring, reflect vigorous
horizontal and vertical motion in the lowest 500 m over the 2000 m deep sill, with
speeds regularly surpassing 100 cm/s. The strong residual flow over the sill in the
passage and internal, mainly diurnal, tides contribute to this bottom intensified motion.
The average volume transport of the deep throughflow from the Maluku Sea to the
Seram Sea below 1250 m is 2.5 Sv (1 Sv ¼ 106 m3/s), with a transport-weighted mean
temperature of 3.2 1C. This result considerably increases existing estimates of the inflow
of the ITF into the Indonesian seas by about 25% and lowers the total mean inflow
temperature of the ITF to below 13 1C. At shallower levels, between 1250 m and the
sea surface, the flow is directed towards the Maluku Sea, north of the passage. The
typical residual velocities in this layer are low (3 cm/s), contributing to an estimated
northward flow of 0.9–1.3 Sv. When more results from the INSTANT programme for the
other Indonesian passages become available, a strongly improved estimate of the mass
and heat budget of the ITF becomes feasible.
& 2009 Elsevier Ltd. All rights reserved.
Keywords:
Indonesian throughflow
INSTANT
Lifamatola Passage
Current measurements
1. Introduction
In his description of the warm return flow of the
thermohaline overturning circulation, Gordon (1986)
stressed that interocean exchange of thermocline waters
is an essential process to maintain this density-driven
global circulation system. One of the important exchange
processes is the Indonesian throughflow (ITF), which
transports water from the low latitude Pacific Ocean
through the Indonesian seas to the eastern Indian Ocean.
Corresponding author.
E-mail addresses: aken@nioz.nl (H.M. van Aken),
irsansb@ocean.itb.ac.id (I.S. Brodjonegoro), indrajaya@ipb.ac.id (I. Jaya).
0967-0637/$ - see front matter & 2009 Elsevier Ltd. All rights reserved.
doi:10.1016/j.dsr.2009.02.001
Gordon (1986) assumed that in the ITF about 8.5 Sv
(1 Sv ¼ 106 m3/s) flows in the upper 1000 m through
several branches in the Indonesian seas, while diapycnal
mixing in the ITF maintains a local downward air–sea heat
exchange of about 100 W/m2. Simulations with ocean
general circulation models as well as with coupled global
climate models have shown that the existence of the ITF
has far-reaching consequences for the global ocean
circulation and climate (e.g. Schneider, 1998; Wajsowicz
and Schneider, 2001; Pandey et al., 2007).
Hydrographic observations (temperature, salinity, CFCs)
have shown that this shallow inflow into the Indonesian
seas (full arrows in Fig. 1) occurs mainly through the
Makassar Strait between the islands Kalimantan and
Sulawesi (Gordon and Fine, 1996). The shallow Dewakan
ARTICLE IN PRESS
1204
H.M. van Aken et al. / Deep-Sea Research I 56 (2009) 1203–1216
10
Pacific Ocean
5
Sulawesi
Sea
0
Kalimantan
Irian
Sulawesi
-5
Banda Sea
D
Java
O
L
-10
T
Indian Ocean
Australia
-15
110
115
120
125
130
135
Fig. 1. Map of the Indonesian throughflow. The full arrows depict the main shallow throughflow route via Makassar Strait. The dashed arrows show the
deep throughflow from the Pacific Ocean via the Banda Sea to the Indian Ocean. The full line shows the 2000-m isobath, while the shaded area is
shallower than 500 m. The area in the rectangle encompasses the Lifamatola Passage and is enlarged in Fig. 2. The capitals D, L, O, and T indicate the
approximate locations of, respectively, the Dewakan sill, Lombok Strait, Ombai Strait, and Timor Passage.
sill at the southern end of Makassar Strait (D in Fig. 1) has
a sill depth of about 680 m (Gordon et al., 2003b). Current
measurements in Makassar Strait have shown that the
throughflow through Makassar Strait is about 10 Sv, with
a maximum contribution from the layer between 150
and 200 m (Gordon et al., 1999, 2003a). The mean temperature of the throughflow in Makassar Strait is nearly 15 1C
(Vranes et al., 2002; Gordon et al., 2003a). The main exits
of the ITF towards the Indian Ocean are the passages
between the lesser Sunda Islands (Nusantara), in particular Lombok Strait, Ombai Strait, and Timor Passage (L, O,
and T in Fig. 1, respectively).
Because of the limited depth of the Dewakan sill in
Makassar Strait, ventilation of the deep basins in the
Banda Sea, with depths in the Weber Deep surpassing 7000 m, requires another pathway. Van Riel (1956)
derived, from the changing temperature stratification
measured during the Snellius Expedition in 1929–1930,
that the flushing of the deep Banda Sea follows a pathway
from the Pacific Ocean, via the Lifamatola Passage east
of Sulawesi (box in Fig. 1, and Fig. 2). The sill in this
passage has a depth between 1900 and 2000 m (van Riel,
1956; Broecker et al., 1986), allowing a deep throughflow
with temperatures well below that of the throughflow in
Makassar Strait. Van Aken et al. (1988) and Gordon et al.
(2003b) have shown, from tracer distributions, that the
hydrographic stratification in the deep Banda Sea agrees
with a ventilation by deep overflow from the Pacific Ocean
via the Lifamatola Passage. Thereby the cold throughflow water descends from the Lifamatola sill along the
topography in an approximately 500 m thick layer, first
into the 5000 m deep Seram Sea, and second from the
Seram Sea over a 3500 m deep sill into the Banda Sea (Van
Aken et al., 1988). Along this path some heating of the
bottom water is observed, mainly attributed to mixing
with warmer overlying water near the sills (Van Aken
et al., 1988). Consequent deep upwelling and mixing in the
Banda Sea then brings the deep throughflow water from a
depth of 5000 to about 1000 m (Van Aken et al., 1991;
Gordon et al., 2003b). Above the latter level the water
from the deep throughflow leaves the Banda Sea towards
the Indian Ocean through the southern exits, joining the
shallow throughflow water from Makassar Strait over the
680 m deep Dewakan sill.
Direct measurements of the flow across the sill in the
Lifamatola Passage and subsequent transport estimates
are scarce. Over 70 years ago Lek (1938) carried out
observations with an Ekman current meter at an anchor
station of the RV Snellius which lasted less than 35 h.
At 1500 m he found a mean southeastward flow of about
5 cm/s. He also observed strong internal semidiurnal and
diurnal tides, with amplitudes reaching over, respectively,
ARTICLE IN PRESS
1205
H.M. van Aken et al. / Deep-Sea Research I 56 (2009) 1203–1216
1
Maluku Sea
Latitude (S)
Obimayor
Lifamatola
Mangole
2
Seram Sea
Banda Sea
3
125
126
127
128
Longitue (E)
Depth (m)
1000
1200
1400
1600
1800
2000
0
5
10
15
20
Distance (km)
25
30
35
Fig. 2. Plot of the topography near the Lifamatola Passage, between the small Lifamatola Island and the larger island Obimayor, with isobaths every
1000 m (a), based on the Smith and Sandwell (1997) 20 20 resolution ETOPO-2 topographic data set. The cross indicates the position of the current meter
mooring near the sill of the passage, while the arrow in the direction 1291 shows the observed mean current direction in the lowest 500 m. The line,
perpendicular to the vector, ending in dots, is the cross-section for which transports have been calculated. The bottom profile along this line is shown in
(b). The mooring was located at x ¼ 14.7 km.
30 and 15 cm/s at several depths. Broecker et al. (1986)
reported observations with two self-recording current
meters over the sill, each about 10 m above the bottom,
which lasted 28 days. They found a typical steady
southeastward residual velocity of about 25 cm/s, in
agreement with a deep flow of Pacific water towards the
Banda Sea. This flow was modified by weaker (10 cm/s)
diurnal and semidiurnal tides. Van Aken et al. (1988)
reported 3.5 months of current measurements with a
current meter mooring over the sill. At 60 m above the
bottom the mean velocity was 61 cm/s, directed to the
southeast, while at 200 m above the sill the mean velocity
was 40 cm/s, in the same direction. Superimposed on
the mean residual flow were tidal motions of a mixed
semidiurnal and diurnal character, leading to total
velocities in the bottom layer that regularly surpassed
1 m/s. From these current measurements the deep inflow
below 1500 m was estimated to be 1.5 Sv. Luick and
Cresswell (2001) carried out observations with a single
current meter mooring, located at 11410 N, nearly 400 km
north of the Lifamatola Passage, in a 230 km wide passage
of the Maluku Sea east of Sulawesi. Assuming a horizontally homogenous flow, they reported a southward
transport between 740 and 1500 m of 7 Sv. However, one
can question whether a single observational point can be
representative for the ocean circulation in that 230 km
wide passage, especially since the hydrographic observations (T, S, O2) indicate the presence of a cyclonic
circulation in the Maluku Sea below 1800 m (Luick and
Cresswell, 2001).
Several research programmes have aimed at the direct
measurement of the other branches of the ITF, both in
Makassar Strait (Gordon et al., 1999, 2003a) and in the
southern exit channels (Murray and Arief, 1988; Molcard
et al., 1994, 1996, 2001). However, the observations from
those experiments do not supply a contemporaneous data
set of the ITF, since the experiments were carried out
in different years. Therefore, the average structure
and magnitude of the total transport of mass, heat, and
freshwater by the ITF are not well known. This lack of
ARTICLE IN PRESS
1206
H.M. van Aken et al. / Deep-Sea Research I 56 (2009) 1203–1216
information hampers the validation of ocean general
circulation models and global climate models. In response
to this lack of knowledge the International Nusantara
Stratification and Transport (INSTANT) programme was
established to directly and simultaneously measure the
ITF, both in the northern inflow passages and in the
southern outflow passages (Sprintall et al., 2004). In this
international research programme, scientists from Indonesia, the USA, Australia, France, and the Netherlands
cooperated to determine the strength of the ITF entering
and leaving the Indonesian seas during three consecutive
years. Variations in the ITF will be related to changes in
the meteorological and oceanographic forcing (sea level,
monsoons, El Niño, Indian Ocean Dipole, etc.). This paper
deals with the observations with a current meter mooring
over the sill in Lifamatola Passage, carried out during the
INSTANT programme. It mainly focuses on the throughflow of the Lifamatola Passage and its variability, but also
describes the higher-frequency (tidal) current oscillations
in the passage.
2. The data
The sill depth in Lifamatola Passage is about 2000 m
(Fig. 2). Slightly downstream of the sill a mooring was
deployed twice for a period of about 1.5 years, leading to a
total mooring period of over 34 months (Table 1). The
deployment cruise in January 2004, the service cruise in
July 2005, and the recovery cruise in January 2006 were
all carried out with the Indonesian RV Baruna Jaya I.
During these cruises, series of CTD casts were also
recorded, with during each cruise at least one CTD cast
close to the mooring position, reaching from the sea
surface to the bottom. For the first period the mooring was
fitted with two RDI 75 kHz ADCPs (Long Ranger), covering
the upper and lower parts of the water column, upward
looking near the surface, and downward looking in the
bottom layer. At intermediate levels, three Aanderaa RCM
11 acoustic current meters were mounted. All RDI
and Aanderaa instruments were fitted with a tilt sensor
for the correction of the velocity, measured from a tilted
mooring. Each of these instruments contained a temperature sensor, while the ADCPs were also fitted with a
pressure sensor. Additionally five Sea Bird Electronics
Microcats (SBE37SM) were mounted to record pressure,
temperature, and salinity (Table 1). The mooring position
appeared to be about 30 m shallower than the deepest
part of the nearby deep channel over the sill, at a distance
of about 1 km further northeast. After recovery of this
mooring it appeared that because of a faulty but undocumented setting of these instruments, only 7 of the 80
data bins of 8 m length were recorded, strongly diminishing the information on the ITF, which is assumed to be
concentrated in the near-surface and near-bottom layers.
Moreover, it appeared that because of very strong tides
the 15-min average velocity at mid-levels regularly
surpassed 140 cm/s. This caused a serious blow-down of
Table 1
Description of the INSTANT moorings in the Lifamatola Passage.
Mooring name
Latitude
Longitude
Deployment date
Recovery date
Corrected depth (m)
LOCO-10-1
1149.10 N
126157.80 E
26-01-2004
17-07-2005
2019
Height above bottom (m)
Instrument type
Serial number
Sampling interval (min)
1535
1534
1232
1231
931
930
630
613
608
10
RDI ADCP
SBE 37 Microcat
Aanderaa RCM 11
SBE 37 Microcat
Aanderaa RCM 11
SBE 37 Microcat
Aanderaa RCM 11
SBE 37 Microcat
RDI ADCP
SBE 37 Microcat
3553
2672
243
2673
244
2674
245
2659
3714
2961
30
5
15
5
15
5
15
5
30
No data
Mooring name
Latitude
Longitude
Deployment date
Recovery date
Corrected depth (m)
LOCO-10-2
1149.10 N
126157.80 E
17-07-2005
04-12-2006
2017
Height above bottom (m)
Instrument type
Serial number
Sampling interval (min)
1233
1231
931
930
630
613
608
10
RDI ADCP
SBE 37 Microcat
Aanderaa RCM 11
SBE 37 Microcat
Aanderaa RCM 11
SBE 37 Microcat
RDI ADCP
SBE 37 Microcat
3553
2959
403
2672
243
4139
3714
2674
30
5
15
5
15
5
5
30
ARTICLE IN PRESS
1207
H.M. van Aken et al. / Deep-Sea Research I 56 (2009) 1203–1216
residual current was estimated from a harmonic analysis
with the dominant tidal frequencies (see below). For the
first deployment period this supplied us with monthly
mean residual current data from 500 to 1800 m, and for
the second period from 300 to 2000 m.
From the average current components between 1500
and 2000 m, measured during the second deployment
period, the mean deep current direction was determined
to be 1291 (arrow in Fig. 2). This direction appears to
be well aligned with the deep channel over the sill. In the
following discussion we have rotated our geographic
reference frame, naming the direction 1291 along-channel,
and the direction perpendicular to the channel, 2191,
cross-channel.
3. The character of the currents and their variability
Current data at 1500 m are available for the whole
deployment period (Fig. 3). At this level the mean velocity
is 14.7 cm/s to the southeast (1191). The magnitudes of the
temporal variation of the current components, expressed
as standard deviation, are 20.7 and 17.5 cm/s for, respectively, the east and north components. The plot of both
velocity components (Fig. 3) shows a considerable contribution from the tides, with an envelope that shows
more or less regular fortnightly variability, indicative of
the spring tide–neap tide phenomenon, as well as longerterm changes of the residual flow. At 1500 m the tides
cause a regular reversal of the current direction. This tidal
reversal of the current direction is observed even during
spring tide at 2000 m, about 20 m above the bottom.
However, during neap tide the residual current at 2000 m
is larger than the tidal contribution, maintaining a permanent southeastward bottom flow. During 0.16% of
the 1-h data records the current speed at 1500 m exceeds
1 m/s. This fraction increases to 12.6% at 2000 m. Variability
at lower frequencies can be observed in the background,
100
50
0
-50
-100
100
50
0
-50
07
06
01
/
10
/
01
/
06
01
/
06
07
/
01
/
06
04
/
01
/
01
/
05
01
/
05
10
/
01
/
05
07
/
01
/
05
04
/
01
/
01
/
04
01
/
04
10
/
01
/
07
/
04
/
01
/
01
/
01
/
01
/
04
-100
04
North component
(cm/s)
East component
(cm/s)
the mooring with an average range of 435 m, adding to the
information loss. In the second deployment period the
mooring was shortened by 300 m to reduce blow-down,
which decreased to an average range of 252 m. However,
thereby the possibility to measure velocity in the upper
300 m was lost. The faulty ADCP setting was also repaired,
so that effectively in the second period our information on
the current structure covered a larger part of the water
column, although data reached only to 300 m below the
sea surface. The bias in the horizontal velocity, induced by
the horizontal motion of the sensors due to the blowdown, was estimated to be about 2 cm/s or less, with a
periodic character. Since the focus of this paper is on the
throughflow, and since the dominant tides have an
amplitude of at least an order of magnitude larger, this
bias has been ignored.
After recovery of the moorings, the recorded directions
were corrected for the magnetic variation, and, by a
combination of low-pass filtering (filter width E1/h) and
sub-sampling every whole hour, a synchronous hourly
data set for all sensors and ADCP bins was produced,
including sensor or bin depth. A shift of the time base
of each instrument was applied in this process to correct
for the effects of different recording intervals and the
different meanings of the time stamps in the different
instruments (beginning, centre, or end of the recording
interval). From these synchronized and depth-dependent
data, time series of hourly data were calculated at fixed
depth levels by means of vertical linear interpolation,
every 100 m in the upper 1600 m and every 50 m from
1600 to 2000 m. Continuous hourly data records are
available for the interpolation depth interval from 1000
to 1500 m in the first deployment period, and for the
1000–2000 m depth interval in the second deployment
period. In order to recover information on the lowfrequency flow in the parts of the water column where
data were available only for part of the tidal period
because of the mooring blow-down, the monthly mean
Date
Fig. 3. Time series of the hourly current components at a depth of 1500 m for both deployment periods combined.
ARTICLE IN PRESS
1208
H.M. van Aken et al. / Deep-Sea Research I 56 (2009) 1203–1216
leading, for example, to a reversal of the low-pass
flow direction at 1500 m during the weeks around 1 April
2006.
The rotary spectrum (not shown) has characteristic
tidal bands centred around about N cpd (cycle per day),
N being an integer number between 1 and 12. Analysis
of the anisotropy of these tidal bands of rotary spectra at
the interpolation depths has shown that apart from the
semidiurnal tidal bands at 1500 m the variable currents
at tidal frequencies are mainly linearly polarized. The
characteristic amplitudes of the anti-clockwise (ACW) and
clockwise (CW) components for those tidal bands were all
of the same order of magnitude with ratios ACW/CW
varying between 0.7 and 1.4, on average 1.0 (70.1 stdev).
As a single exception more extreme values of the ACW/CW
ratios were found only in the semidiurnal band at 1500 m,
where the ACW/CW ratio was 2.7, indicative of the
dominance of ACW motion at this depth in this frequency
band. This suggests a large influence of the topography on
other frequencies and at deeper levels in the channel over
the sill. The tidal currents below 1500 m were all aligned
in the along-channel direction.
The layer-averaged power spectra for the alongchannel and cross-channel velocity (Fig. 4) confirm the
conclusion, presented by Van Aken et al. (1988), that
the tidal current in the Lifamatola Passage is dominated
by the diurnal (O1 and K1) and semidiurnal (M2 and S2)
frequencies, while additional peaks are found at their
higher harmonics, at N cpd (N ¼ 3, 4, 5, y). Additionally,
a spectral peak is found near the lunar fortnightly (Mf)
frequency (7.32 102 cpd). The latter is close to, but not
coincident with, the local inertial frequency of 6.37 102
cpd. The difference of the spectra for the along-channel
and cross-channel components shows that in the depth
interval from 1600 to 2000 m, completely surrounded by
the channel wall (thick lines in Fig. 4), the kinetic energy
Spectral density (cm2s-2cpd-1)
104
103
Mf
O1 K1
M2
S2
102
101
100
129° up
219° up
129° down
219° down
10-1
10-2
10-1
100
Frequency (cpd)
101
Fig. 4. Power spectrum of the velocity components in along-channel
(1291, black lines) and cross-channel (2191, grey lines) direction,
averaged over the depth intervals 1000–1500 m (up, thin lines) and
1600–2000 m (down, thick). The peaks at the dominant semidiurnal and
diurnal tidal component as well as the lunar fortnightly components are
identified. These spectra are based on successive 70-day sub-periods
during the second deployment period, with an overlap of 35 days.
of the along-channel velocity variations is on average 19
times the kinetic energy of the cross-channel variations at
diurnal and semidiurnal frequencies. At the sub-diurnal
frequencies below 0.5 cpd that ratio for the spectral
continuum is about 3, at super-semidiurnal frequencies
about 2.3. This anisotropy reflects the fact that over the
sill the variable motion takes place in a narrow channel,
8.5 km wide at the depth of the 1800 m isobath, which
hinders the cross-channel motion considerably. At the
level from 1000 to 1500 m the channel is much wider
(28 km at the 1250 m level), and consequently the
anisotropy of the variable motion is smaller, 5 at the
tidal peaks, 2 at sub-diurnal frequencies, and 1.2 at
super-semidiurnal frequencies.
To show the character of the tides in more detail, a
harmonic tidal analysis of the data has been carried out
for the fortnightly, diurnal, and semidiurnal tides, indicated in Fig. 4, as well as for higher harmonics, up to at
least 6 cpd tides. It appears that the strongest tides are,
in order of amplitude, K1, O1, M2, S2, and Mf, all five with
amplitudes at all depths between 1000 and 2000 m of
over 1 cm/s in the along-channel direction. The determination coefficient (correlation R squared) of the harmonic
analysis shows that these five tidal frequencies determine
on average 65% of the along-channel current variation
(R ¼ 0.80). All the higher harmonics have amplitudes less
than 1 cm/s. The tidal ellipses for the dominant frequencies (Fig. 5) confirm that the diurnal tides, K1 and O1, are
the strongest between 100 and 200 m above the bottom of
the sill with along-channel amplitudes of, respectively, 33
and 25 cm/s. The narrow ellipses for K1 and O1 are fairly
well aligned in the direction 1291 of the deep channel of
the Lifamatola Passage, decreasing upwards in amplitude.
The phase shift from 1800 to 1200 m was about 351. This
suggests that these tides have the character of internal
tides with an upward energy flux, generated on the slope
below the sill by interaction of the barotropic tide with
the sill topography in a stratified ocean (Gill, 1982).
The rotation sense of the ellipses of the diurnal tides
varies with depth, being anti-clockwise in the bottom
layer below 1900 m, in an 200 m thick layers around
1400 m, and above 1100 m. In the other layer the rotation
sense was clockwise. The ellipses for the semidiurnal tides
and the lunar fortnightly tide are less narrow, and their
orientation varies with depth, suggesting that these tides
are propagated across the mooring site as internal waves
from different generation areas. The rotation sense of the
ellipses of these tidal components was dominantly anticlockwise below 1200 m, turning to clockwise rotation
above this depth level. The vertically varying rotation
sense of the tidal ellipses shows that they do not behave
as free waves in an infinite ocean, but interact with lateral
walls (Gill, 1982). It is expected that the strong internal
tides near narrow straits like the Lifamatola Passage
support turbulent mixing, which contributes to the water
mass transformation of the ITF water in the Indonesian
seas (Ffield and Robertson, 2005; Koch-Larrouy et al.,
2007). Local results of intense turbulent mixing for the
temperature and oxygen distributions in the Lifamatola
Passage have been described by Van Aken et al. (1988) and
Tomczak and Godfrey (1994).
ARTICLE IN PRESS
1209
H.M. van Aken et al. / Deep-Sea Research I 56 (2009) 1203–1216
129° component (cm/s)
30
30
M2
20
20
20
10
10
10
0
0
0
-10
-10
-10
-20
-20
-20
-30
-30
-30
-20
30
129° component (cm/s)
30
O1
-10
0
10
20
-20
30
K1
20
20
10
10
0
0
-10
-10
-20
-20
-30
-30
-20
-10
0
10
20
219° component (cm/s)
-10
0
10
20
Mf
-20 -10 0
10 20
219° component (cm/s)
S2
2000 m
1800 m
1500 m
1000 m
-20
-10
0
10
20
219° component (cm/s)
Fig. 5. Plots of the tidal ellipses at depths of 2000, 1800, 1500, and 1000 m for the tidal components O1, K1, M2, S2, and Mf. These ellipses were determined
with a harmonic analysis of the depth-interpolated current data from the second deployment period.
4. The mean current profile
For the first deployment period uninterrupted interpolated hourly velocity data are available for the depth
interval of 1000–1500 m, for the second deployment
period from 1000 to 2000 m. Only for these depth
intervals can the residual currents be determined directly
by averaging of the hourly data. To extend the information
on the residual currents further in the vertical, we have
applied an alternative method. As shown above, most of
the current variability comes from only five tidal components plus sub-tidal variability with time scales from
months to years. Assuming that the along-channel current
can be described quite well as a monthly mean residual
current plus tidal contributions from the O1, K1, M2, S2,
and Mf components, harmonic analysis by means of
multiple linear regression (Emery and Thomson, 1997)
can be used to estimate the monthly mean along-channel
current components. This can also be applied to those
parts of the water column where, because of the blowdown of the moorings, the current record is not continuous, but interrupted by short periods of missing data.
This has allowed us to extend the mean current profiles,
based on monthly averages of the residual velocity, for the
first deployment period to the depth interval from 500 to
1700 m, and for the second deployment period from 300
to 2000 m.
The current profiles for the two different deployment
periods, derived from the monthly mean residual currents
(grey lines in Fig. 6), are quite similar to each other. The
resulting long-term averaged current profile (black line
with symbols in Fig. 6) shows a strong southeastward
overflow in the lowest 750 m over the sill, following the
axis of the channel. From 300 to 1200 m a small but
significant northwestward flow of about 3.2 (70.5 stderr)
cm/s is observed. In the deep overflow the average alongchannel velocity reaches a maximum of 67 (76 stderr)
cm/s at 1950 m, about 70 m above the bottom, decreasing
downwards to 51 cm/s at 2000 m.
The historic current estimates, discussed in the
introduction (crosses in Fig. 6), compare reasonably
well in order of magnitude and vertical structure with
our averaged current profile, especially if we consider
their limited accuracy and much different observational
periods. On average these estimates are 62% of our mean
currents at similar depths. The relatively low near-bottom
velocities reported by Broecker et al. (1986) agree with the
downward velocity decrease that we observe below
1950 m, especially considering that the velocities were
measured at 10 m above the bottom, while our lowest
ARTICLE IN PRESS
1210
H.M. van Aken et al. / Deep-Sea Research I 56 (2009) 1203–1216
250
Salinity
34.6
500
34.62
34.63
500
750
S
1000
1000
1250
1500
Θ
Depth (m)
Depth (m)
34.61
L
1500
1750
B
2000
0
vA
vA
20
40
60
129° Velocity Component (cm/s)
80
Fig. 6. Mean profile of the along-channel current component (1291,
black line with symbols) and estimated accuracy (twice the standard
error), derived from monthly mean values, with separate curves (grey
lines) for the first and second deployment periods. The crosses show
mean velocities from the literature, shifted vertically to the appropriate
distance from the bottom (L ¼ Lek, 1938; B ¼ Broecker et al., 1986;
vA ¼ Van Aken et al., 1988).
profile level at 2000 m was about 20 m above the bottom.
The information available from our current profile shows
that the southeastward deep overflow extends from the
bottom to a depth of about 1250 m, while Van Aken et al.
(1988), with data from only two current meters, derived
from extrapolation a current reversal at 1500 m. Their
averaged velocities were also smaller than those presented here. Therefore it can be expected that the average
transport in the deep overflow presented here is larger
than their estimate of 1.5 Sv.
5. The temperature and salinity stratification
The depth over which the temperature sensors in the
mooring moved, because of the blow-down of the
mooring, was in general of the same order of magnitude
or larger than the expected vertical tidal migration of the
isotherms of a few 100 m (Van Aken et al., 1988).
Therefore, enough observations were available in the
water column between 500 and 2000 m to determine
the long-term (3 years) mean profiles of potential
temperature and salinity (Fig. 7). These show a salinity
minimum at 800 m, where in the South Pacific Ocean a
similar minimum is found (Wyrtki, 1961). In the lowest
400 m, where the core of the overflow of deep Pacific
water into the Seram Sea/Banda Sea system is expected
(Van Aken et al., 1988) the vertical gradient of salinity
and potential temperature is larger than directly above
this layer.
While during the first deployment period the nearbottom Microcat was defective, during the second deployment period this instrument recorded continuously the
2000
2
4
6
Potential Temperature (°C)
8
Fig. 7. Mean profiles of potential temperature (thick line) and salinity
(thin line) between 500 and 2000 m, averaged over the whole
deployment period.
near-bottom temperature and salinity in the lowest 10 m.
Similar to the velocity signal, the temperature and salinity
signals show a strong tidal variability (Fig. 8), with a
clear fortnightly variation in amplitude. The temperature
and salinity variations are negatively correlated as can be
expected from the mean stratification with opposite Y
and S gradients, depicted in Fig. 7. Therefore the tidal
signals of these parameters are in opposite phase, with
temperature maxima coinciding with salinity minima
(Fig. 8). The spectra of temperature and salinity, derived
from this time series (Fig. 9), clearly show the presence of
the same dominant tides, which are also observed in the
velocity spectrum, including the higher harmonics of the
tides. The tidal character of these scalars is therefore also
mixed diurnal and semidiurnal.
At sub-tidal frequencies some variation is still present
in temperature and salinity. When we remove the tidal
variability at fortnightly and shorter time scales with a 15day running average, we can draw a time–depth diagram
of the temperature, showing the temporal variability
of the stratification at sub-tidal frequencies (Fig. 10). A
deep maximum of the sub-tidal temperature variability is
observed at 1750 m (stdev ¼ 0.14 1C), nearly twice the
variability 250 m higher or lower (stdev ¼ 0.08 1C). Below
1200 m the depths of the isotherms show more variability at low frequencies than above this level. Over short
vertical distances the correlation of the temperature in
this deep layer is significant (p ¼ 1%), being 0.79, 0.64, and
0.43 over vertical distances of, respectively, 50, 100, and
200 m. However, over larger vertical distances, e.g. 400
and 500 m, the correlation reduces to values of 0.10 and
0.02, while the correlation between the near-bottom
temperature at 2000 m and the temperature at 1300 m
ARTICLE IN PRESS
1211
H.M. van Aken et al. / Deep-Sea Research I 56 (2009) 1203–1216
34.64
S
3.0
34.62
2.8
34.60
2.6
34.58
2.4
34.56
2.2
Salinity (pss-78)
Potential Temperature (°C)
3.2
34.54
Θ
2.0
34.52
1-Apr-06
1-May-06
1-Jun-06
Date/Time (UTC)
1-Jul-06
Fig. 8. Continuous plot of the near-bottom temperature (black line) and salinity (grey line), measured with an SBE Microcat for the sub-period April until
June 2006.
10-1
Mf
-2
600
S2
7.0
Temperature
(deg. C)
6.0
800
Θ spectrum
10-3
7.5
6.5
5.5
K1
5.0
10-4
O1
S spectrum
10-5
M2
1000
4.5
S2
Depth (m)
Spectral density
10
O1 K1M2
10-6
10-7
10-2
10-1
100
Frequency (cpd)
1200
4.0
1400
101
1600
Fig. 9. Power spectra of the near-bottom temperature (black line) and
salinity (grey line), measured with an SBE Microcat in the second
deployment period.
1800
has even a negative extreme of 0.27. The significant
positive correlation is due to the apparent simultaneous
upward or downward movements of thick layers of water
with their isotherms, reflected by coincident upward or
downward peaks in the isotherms (Fig. 10). However, at
larger vertical distances of about 500 m the correlation
is reduced to near zero by additional intrusions of thick
packets of water with variable vertical temperature
gradients, while the temperature near the bottom is
partly in counter-phase with the temperature at 1300 m,
as can be expected when thermostads are a dominant
feature in the variable temperature structure. The strongest example of such a low T gradient intrusion occurs
between days 800 and 840 (10 March–19 April 2006). In
that sub-period the 2.5 and 3.0 1C isotherms move downwards, while the 3.5 and 4.0 1C isotherms move upwards, all over a vertical distance of over 100 m (Fig. 10).
Apparently a deep thermostad with a low vertical stability,
centred around 1700 m, passed Lifamatola Passage around
day 822 (1 April 2006). Such a thick thermostad may have
dynamic consequences for the deep flow, since it is
probably also connected with changing horizontal density
gradients. Evidence for this hypothesis can be seen in the
3.5
3.0
2.5
2000
200
400
600
800
Day number (1 Jan. 2004 = 1)
1000
Fig. 10. Time–depth section of the temperature measured with the
temperature sensors on the mooring. The signal was smoothed with a
low-pass filter with a width of 14 days to remove the tidal contributions
to the temporal variability.
velocity already at 1500 (Fig. 3), which shows a current
reversal around that same date.
One may question whether the thermostad is formed
locally, or whether it is advected from elsewhere. The
thermostad around 1 April 2006 lasted at least 30 days,
an earlier thermostad in October 2005 about 26 days.
During these periods no particularly strong tides or other
phenomena were observed, which may be responsible for
local mixing over the sill as the cause of the occurrence of
the thermostad. At the centre of the thermostad at a depth
of 1700 m, the passage has a width of only about 14 km.
During the 30 days of its occurrence, the mean downchannel velocity at 1700 m was 8 cm/s, suggesting an
along-channel size of about 220 km. Apart from the effects
of the convergence and divergence of the thermostad as it
passes the narrow Lifamatola sill, these rough estimates
ARTICLE IN PRESS
1212
H.M. van Aken et al. / Deep-Sea Research I 56 (2009) 1203–1216
show that the thermostad was much longer than the
width of the channel. This suggests that the thermostad
was advected from the Maluku Sea to the Lifamatola sill.
The cause of the formation of such deep thermostads
upstream of the passage remains uncertain yet. The
relatively large size of the thermostad in the along-flow
direction ensures that the change of the pressure gradient
due to the presence of a thermostad probably extends
over the whole Lifamatola sill, and thereby influences the
whole deep overflow.
6. Transports through Lifamatola Passage
With the velocity data, obtained from the moorings,
we can calculate the volume transport across a section
that runs perpendicular to the mean along-channel deep
flow in the direction of 2191. The endpoints of that section
are given in Table 2. The section runs from a ridge,
extending northwestwards from Obimayor (depth
1090 m) across the deep channel in the Lifamatola Passage
(maximum depth 2050 m) to a ‘‘shallow’’ platform
east–southeast of Lifamatola (depth 1645 m). The width
of this section is 36.2 km, while at 1750 m it is only
10.5 km wide, intersected by the bottom topography
(Fig. 2b). For the calculation of the transport we assume
that the measured velocity profile (Fig. 6) is representative
for the vertical velocity structure across the whole section.
For the levels below the deepest interpolated velocity
level, 2000 m, we apply the 2000-m velocity. Given the
narrow width of the deep channel at these levels, the
effect of this, or any other extrapolation, is very small.
The along-channel volume transport Tr over the sill
between the depth levels z0 and z1 has been computed
Table 2
Endpoints of the transport section, through which the transports are
calculated.
Latitude
Longitude
0
127105.140 E
126152.810 E
1139.90 S
1155.070 S
from the profile of the velocity in the direction of 1291 by
the following integral:
Z z1
Tr ¼
V 129 WðzÞ dz,
(1)
z0
where W(z) is the depth-varying width of the channel. We
have calculated the volume transports for the depth
intervals 450–1250 m and from 1250 to the bottom, since
the along-channel velocity profile in Fig. 6 suggests that in
these depth intervals the current is, respectively, in the
northwestern and southeastern direction. These depth
intervals are well covered by velocity estimates for the
second deployment period, but for the first deployment
period, part of the data in these intervals are missing in all
or part of the months in this period. In order to prevent a
possible bias in the transport estimate due to missing
data, we have to find a method to obtain a reliable
approximation for the transport for the first deployment.
For the second deployment period, with a complete
data coverage for both depth intervals, the deep transport
below 1250 m has been derived from the 5-day running
mean values of the interpolated along-channel velocity.
The resulting transport time series varies from 0.8 to
5.2 Sv, with an average value of 2.7 Sv (71.3 Sv stdev). This
transport is nearly twice the transport value of 1.5 Sv
derived by Van Aken et al. (1988). This is partly because
their 3.5 months of observations gave lower velocity
values (vA crosses in Fig. 6) than our 1.5-year deployment
period, and partly because their estimate of the thickness
of the overflow layer (460 m) from linear extrapolation
was smaller. The best linear correlation between the
volume transport and the velocity at a single level
(R ¼ 0.993) is found at a depth of 1500 m. The time series
at this level is also complete for the first deployment
period. We have calculated the volume transport for this
period from the velocity at 1500 m by means of the linear
regression with the along-channel velocity at 1500 m,
derived from the data in the second deployment period.
The thin full line in Fig. 11 shows that there exists a
regular fortnightly variability of the transport through the
Lifamatola Passage below 1250 m, probably connected
with the Mf tidal component. However, application of a
1.8
6
1.6
4
1.4
2
0
1.2
T1300-T2000 (°C)
Volume transport 129° (Sv)
8
-2
1.0
01/01/04
01/07/04
01/01/05
01/07/05
01/01/06
01/07/06
01/01/07
Date/Time
Fig. 11. Time series of the along-channel volume transport between 1250 m and the bottom in the direction 1291. Positive values are from the Maluku Sea
southeastwards to the Seram Sea. The thin full line shows the 5-day running mean, the thick full line the 15-day running mean, where the fortnightly Mf
tidal contribution is suppressed. The dashed line represents the 15-day running mean of the temperature difference between 1300 and 2000 dbar.
ARTICLE IN PRESS
H.M. van Aken et al. / Deep-Sea Research I 56 (2009) 1203–1216
15-day running mean (thick full line in Fig. 11) effectively
suppresses this tidal contribution. A peculiar phenomenon in the transport time series is the 22-day period around
1 April 2006, where the along-channel volume transport
below 1250 m is negative. It appears that this is caused by
the lowering of the zero-velocity interface, on average at
1250 m, to a level of nearly 1700 m, the deepest current
reversal level observed in our records. This lowering
coincides with the presence of the thermostad in the
passage, centred around the same depth of 1700 m, as
described above. A similar phenomenon, although less
extreme, occurred around 20 October 2005. Apparently
the vertical temperature gradient is related to the strength
of the deep throughflow. The temperature difference
between 1300 and 2000 m, assumed to be representative
for the vertical temperature gradient in the deep throughflow and the presence of thermostads (Fig. 11, dashed
line), shows a regular coincidence of high- or lowtransport events with high or low temperature gradients.
The along-channel transport, derived from the 15-day
running mean time series in Fig. 11, averaged over the
whole near-3-year deployment period, amounts to 2.4
(71.5 stdev) Sv, while the transport derived from the
averaged profiles in Fig. 6 amounts to 2.5 Sv. Apparently
the averaged velocity profile is not very sensitive to the
bias caused by levels with incomplete data from the first
deployment period.
In order to determine the average temperature of the
deep throughflow below 1250 m, we have to determine
the transport-weighted mean temperature. The transportweighted temperature TTW is defined as
R z1
z V 129 TðzÞWðzÞ dz
T TW ¼ 0
,
(2)
Tr
where T(z) is the depth-dependent temperature, observed
with the temperatures sensors in the mooring and Tr is
defined in (1). When we calculate the long-term mean
value of this TTW with the 5-day-averaged time series
of the along-channel velocity and temperature for the
second deployment period, we obtain a mean transportweighted temperature of 3.21 1C, while the use of mean
vertical profiles of temperature and velocity for this
period results in a mean temperature of 3.25 1C. Apparently the effects of any Reynolds terms in the total heat
(temperature) transport are negative, but small, and can
be ignored. For the total 34-month deployment period the
mean temperature of the deep throughflow, derived from
the long-term mean profiles of temperature and alongchannel velocity, is 3.22 1C. The mean salinity of the deep
throughflow, determined similarly, is 34.617, agreeing
with the (near-homogeneous) salinity in the deep Banda
Sea below 2000 m (34.615–34.620; Van Aken et al., 1988,
1991).
In the layer from 250 to 1250 m a significant mean
velocity from the Maluku Sea to the Seram Sea is observed
(Fig. 6). Given the increasing width of the Lifamatola
Passage from 1250 to the sea surface, the assumption that
the transport at these levels can be estimated from
velocity measurements at a single mooring becomes
questionable. But if we do this, the total volume transport
in the 1000 m thick layer between 250 and 1250 m
1213
amounts to 1.0 (71.1 stdev) Sv, which means a strongly
varying, but on average northwestward, transport towards
the Maluku Sea. An annual cycle of the northwestward
velocity can be discerned only at 300 and 400 m depth, in
counter-phase with the near-surface Ekman Transport at
71S in the Banda Sea (Sprintall and Liu, 2005). If we
assume that the residual velocity at 300 m extends to the
sea surface, we arrive at a small residual northwestward
transport in the upper 1250 m of 1.3 Sv. However,
estimates of the current structure in Makassar Strait have
shown that the vertical current structure in the upper
300 m of the Indonesian seas may show a considerable
seasonal variability (Gordon et al., 2003a). These estimates of the transport above 1250 m, given above, are
probably not very accurate, and should be interpreted
with care.
7. Discussion
The data obtained with the INSTANT mooring in the
Lifamatola Passage show that vigorous motion occurs in
the deep channel across the sill. That motion consists of
the residual flow, contributing to the deep ITF of cold
water from the Pacific Ocean towards the Banda Sea and
of vertically varying internal tides. Since the southern
exits of the Banda Sea are all shallower than the sill in the
Lifamatola Passage, this sill controls the deep throughflow
from the Pacific to the Indian Ocean (Gordon et al.,
2003b).
In the tidal motion the diurnal tides K1 and O1 are
dominant, although Wyrtki (1961) has reported that in
the sea areas around Sulawesi, also in the Lifamatola
Passage, the surface tide is mixed, but prevailing semidiurnal. Apparently the mooring was not intersected
directly by a semidiurnal internal tidal wave beam.
Generation of (lower-frequency) diurnal internal tides by
interaction of the barotropic tide and topography in a
stratified sea requires a smaller critical bottom slope than
the generation of the (higher-frequency) semidiurnal
internal tides, while a propagating diurnal internal wave
beam will also follow a less-steep path than the
semidiurnal tide (Gill, 1982). The steeper critical bottom
slope is probably found deeper on the sill, at a larger
distance from the mooring near the top of the sill, than the
smaller slope, where diurnal internal tides are generated.
The shorter distance to the critical slope combined with
the less-steep wave beams will favour the diurnal tide to
reach the mooring. These frequency-dependent differences in generation areas and wave propagation may
explain the observed local dominance of the diurnal
internal tides over the sill.
We can assume that the strong tidal variation of the
near-bottom temperature and salinity, measured during
the second deployment period (Fig. 8), is caused mainly by
vertical tidal motion of the isotherms and isohalines. From
the temperature signal and the mean vertical gradient
of the potential temperature we then derive the standard
deviation of the vertical excursion to be 156 m. The
standard deviation of the vertical excursion, derived from
the variations of the salinity, is 144 m, within 10% a similar
ARTICLE IN PRESS
1214
H.M. van Aken et al. / Deep-Sea Research I 56 (2009) 1203–1216
value. This value agrees with the large vertical motion
of the isotherms presented for periods of 4 days by
Van Aken et al. (1988), who show depth variations of
isothermals (peak–trough) of about 400 m (their Fig. 8).
This result also indicates that it is hard to derive the nearbottom vertical stratification from only a few CTD casts.
More detailed modelling studies on the generation
of internal tides, in combination with resulting generation
of turbulent mixing in a general circulation model,
like those carried out by Koch-Larrouy et al. (2007),
may shed light on the question where and how internal
tides contribute to the enhanced turbulent mixing in
the deep Indonesian seas, derived by Van Aken et al.
(1988, 1991).
The main transport in the Lifamatola Passage, responsible for the cold part of the ITF and the ventilation of the
deep Banda Sea, appears to be concentrated in the lower
layer of the water column below 1250 m (Fig. 6), with a
velocity maximum of over 65 cm/s at 70 m above the
bottom. The observed downward decrease of the velocity
below that level is probably related to friction effects in
the near-bottom layers. The long-term mean volume
transport in this cold ITF, estimated in different ways,
amounts to 2.4–2.5 Sv for the total 34-month deployment
period. This is nearly 60% higher than the volume
transport of 1.5 Sv, derived from a much shorter deployment period of only 3.5 months, reported by Van Aken et
al. (1988). This difference is not necessarily connected
with long-term changes. From our 34-month transport
record (Fig. 11) one can derive that on average 15% of the
3-month-average transports have a magnitude of 1.5 Sv or
less. It can be concluded that the relatively low transport,
reported by Van Aken et al. (1988), is not really a rare
event.
The time series of the cold transport, shown in Fig. 11,
shows an occasionally strong variation with a time scale of
about 1 month. It has been shown that the negativetransport event around 1 April 2006 coincided with the
presence of a deep hydrographic feature centred at
1700 m, characterized by a strongly reduced vertical
temperature gradient, a thermostad. Below 1700 m the
residual flow was then still directed towards the Seram
Sea, with near-bottom velocities at 1950 m of well over
40 cm/s. The negative correlation between the temperature at 1950 and 1300 m suggests that variations in the
vertical temperature gradient, as they occur in a thermostad, are a dominant feature below 1250 m. To analyze
whether the relation between the presence of a thermostad in the lowest 750 m and the reduction of the deep
overflow can be generalized, we have determined the
correlation between the 15-day running mean of the
volume transport below 1250 m and the temperature
difference between 1300 and 2000 m. The resulting
correlation is 0.62. Apparently, deep thermostads are
encountered regularly in the Lifamatola Passage, and they
cause a decrease in the volume transport. The source and
dynamics of these deep features are unknown yet, but the
coincidence of the strong thermostad at 1700 m with the
low-transport event, as well as the significant correlation
between transport and vertical temperature gradient,
suggests that both are related.
The long-term averaged deep transport derived from
our data, 2.5 Sv, is considerably smaller than the southward transport between 740 and 1500 m of 7 Sv, derived
by Luick and Cresswell (2001) for a section in the Maluku
Sea between Sulawesi and Halmahera. The transport over
a similar depth interval derived from our mooring data, is
0.3 Sv, directed to the northwest. This large contrast
clearly indicates, as was already suggested by Luick and
Cresswell (2001), that the assumption that the averaged
current data from a mooring at the extreme western side
of a 230 km wide section are representative for the entire
width of this passage is indeed doubtful. Representative
deep current observations are more likely obtained for our
section of only 36 km wide, although even in this case
the lateral homogeneity is still an assumption, not yet a
proven fact. At shallower levels of, e.g., 500 m the crosschannel distance between the isobaths in the Lifamatola
Passage is about one third of the section used by Luick
and Cresswell (2001), but even over such a distance it
is doubtful that the use of a single current meter mooring will produce an accurate estimate of the shallow
transport.
The inflow of cold water through the Lifamatola
Passage, in addition to the main ITF branch through the
Makassar Strait, influences not only the volume budget of
the ITF but also the heat transport. The existing transport
estimate of the strength of the warm Makassar Strait
branch based on the Arlindo Experiment (Gordon et al.,
1999) is 9.3 Sv. The deep inflow through Lifamatola
Passage, presented here, adds over 25% of cold water to
this estimate. Vranes et al. (2002) derived the transportweighted temperature in Makassar Strait by four different
methods and found a mean value of 14.6 (71.8 stdev) 1C. If
we combine these (non-contemporary!) results with the
mean transport and transport-weighted temperature from
the deep transport in Lifamatola Passage (2.5 Sv and
3.2 1C), we arrive at a total northern inflow of the ITF into
the Indonesian seas of 11.8 Sv, with a transport-weighted
mean temperature of about 12.2 1C. This temperature is
considerably lower than the temperature estimates from
the past of well over 20 1C (Gordon et al., 2003). When
considering these results, one has to be aware that these
results are obtained by putting together data collected
from different experiments, carried out over different
years. But new data from all the different contemporary
experiments of the INSTANT programme may lead to a
thorough revision of the heat budget of the ITF.
The 2.5 Sv cold deep water entering the Banda Sea
system with a mean temperature of 3.2 1C rises there to
shallower levels before it can leave for the Indian Ocean
through the passages between the lesser Sunda Islands or
around Timor. According to Gordon et al. (2003b) this
deep water leaves the Banda Sea after upwelling to depths
between 1300 and 680 m or shallower, before it leaves for
the Indian Ocean. The deeper part of this throughflow
occurs through the Timor Passage and Ombai Strait
(T and O in Fig. 1). This agrees quite well with the
available ETOPO-2 20 20 resolution topographic data set
(Smith and Sandwell, 1997). Preliminary direct transport
estimates of the throughflow entering the Indian Ocean,
derived from other parts of the INSTANT experiment,
ARTICLE IN PRESS
H.M. van Aken et al. / Deep-Sea Research I 56 (2009) 1203–1216
suggest a deep maximum in this throughflow at 1000 m
depth (Sprintall et al., 2009). At that level the Banda Sea
has a mean temperature of 5.0 1C (Van Aken et al., 1988).
The heat transported by 2.5 Sv at 5 1C is larger that by the
same volume transport in the Lifamatola Passage at 3.2 1C.
This implies that the heat flux of the deep throughflow in
the Band Sea is divergent, and extra heat has to be
supplied to warm the upwelling cold water (Munk, 1966;
Van Aken et al., 1991). This heat source is supplied by
turbulent mixing with the overlying water. Given that the
surface of the Banda Sea is 6.7 1011 m2, and the specific
heat of sea water is about 4000 J/kg/1C, the downward
turbulent heat flux in the Banda Sea at a level of 1000 m
should supply 28 W/km2 to support the divergent advective heat flux. If the cold throughflow rises up to the
shallower level of 500 m, where the mean temperature is
8 1C (Van Aken et al., 1988), a downward turbulent heat
flux of 74 W/m2 is required. Upwelling of the cold ITF to
even shallower and warmer levels needs even larger heat
inputs. This heat is supplied by the cooling of the
overlying water, and ultimately by cooling of the atmosphere in Indonesia. Apparently, maintenance of the deep
throughflow through the Lifamatola Passage may have a
considerable influence on the local climate, of the same
order of magnitude as the heat flux sustained by the
throughflow of thermocline water (100 W/m2) proposed
by Gordon (1986).
The downward turbulent heat flux in the water column
is balanced by the upward advective heat flux. Using the
vertical advection–diffusion model, proposed by Munk
(1966), and using a deep transport estimate of 1 Sv,
Van Aken et al. (1988, 1991) were able the estimate the
turbulent diffusion coefficient in the Banda Sea to be
9 10–4 m2/s. Given a known temperature stratification in
the Banda Sea, the turbulent diffusion coefficient derived
from Munk’s model is proportional to the vertical
upwelling velocity, which is in its turn proportional to
the deep inflow of cold water. With our new estimate
of an inflow of 2.5 Sv, this leads to an estimate of about
2.2 10–3 m2/s, more than a factor of 20 higher than
the generally accepted canonical value of 10–4 m2/s for
the world ocean (Gargett, 1984). Apparently, the strong
internal tides in the Banda Sea contribute strongly to the
turbulent energy budget responsible for the maintenance
of the downward turbulent heat flux and the considerable
mixing in the Banda Sea (Van Aken et al., 1988; KochLarrouy et al., 2007).
When more results from the INSTANT programme
become available, it is likely that a more reliable budget
for the ITF of mass, heat, and freshwater than currently
available can be produced. These will form a benchmark
for model simulations of the interocean exchange between the Pacific and Indian Oceans and are expected
to boost our understanding of these processes for the
oceanic climate.
Acknowledgements
We thank the captains and crew of the Baruna Jaya I of
the Indonesian Agency for Research and Application of
1215
Technology (BPPT) and the technicians and scientists from
the Agency for Marine and Fisheries Research (BRKP) for
their support during the research cruises. Many Indonesian students also participated in the three cruises. Theo
Hillebrand and Sven Ober prepared and serviced the
instruments, while Marcel Bakker, Jack Schilling, and Leon
Wuijs were responsible for the construction and handling
of the mooring, and Kees Veth replaced HMvA during the
last cruise. Funding for the mooring equipment was
received from the Netherlands Foundation for Scientific
Research (NWO) via the LOCO investment programme,
and via the IMAU of Utrecht University from the COACh
International Research School. Ship time of the RV Baruna
Jaya I was made available by the BRKP.
References
Broecker, W.S., Patzert, W.C., Toggweiler, J.R., Stuiver, M., 1986. Hydrography, chemistry, and radioisotopes in the southeast Asian waters.
Journal of Geophysical Research 91, 14,345–14,354.
Emery, W.J., Thomson, R.E., 1997. Data Analysis Methods in Physical
Oceanography. Pergamon, Oxford, p. 634.
Ffield, A., Robertson, R., 2005. Indonesian seas finestructure variability.
Oceanography 18, 108–111.
Gargett, A.E., 1984. Vertical eddy diffusivity in the ocean interior.
Dynamics of Atmospheres and Oceans 42, 359–393.
Gill, A.E., 1982. Atmosphere–Ocean Dynamics. Academic Press, London,
p. 662.
Gordon, A.L., 1986. Interocean exchange of thermocline water. Journal of
Geophysical Research 91, 5037–5046.
Gordon, A.L., Fine, R.A., 1996. Pathways of water between the Pacific and
Indian oceans in the Indonesian sea. Nature 379, 146–149.
Gordon, A.L., Susanto, R.D., Ffield, A., 1999. Throughflow within Makassar
Strait. Geophysical Research Letters 26, 3325–3328.
Gordon, A.L., Susanto, R.D., Vranes, K., 2003a. Cool Indonesian throughflow as a consequence of restricted surface layer flow. Nature 425,
824–828.
Gordon, A.L., Giulivi, C.F., Ilahude, A.G., 2003b. Deep topographic barriers
within the Indonesian Seas. Deep-Sea Research II 50, 2205–2228.
Koch-Larrouy, A., Madec, G., Bouruet-Aubertot, P., Gerkema, T., Bessières,
L., Molcard, R., 2007. On the transformation of Pacific water into
Indonesian throughflow water by internal tidal mixing. Geophysical
Research Letters 34, L04604.
Lek, L., 1938. The Snellius Expedition in the eastern part of the East
Indian Archipelago 1929–1930. Vol. II: Oceanographic results. Part 3:
Die Ergebnisse der Strom- und Seriemessungen (in German).
Luick, J.L., Cresswell, G.R., 2001. Current measurements in the Maluku
Sea. Journal of Geophysical Research 106, 13953–13958.
Molcard, R., Fieux, M., Swallow, J.C., Ilahude, A.G., Banjarnahor, J.,
1994. Low frequency variability of the currents in Indonesian
channels (Savu-Roti and Roti-Ashmore Reef). Deep-Sea Research I
41, 1643–1661.
Molcard, R., Fieux, M., Ilahude, A.G., 1996. The Indo-Pacific throughflow
in the Timor Passage. Journal of Geophysical Research 101,
12,411–12,420.
Molcard, R., Fieux, M., Syamsudin, F., 2001. The throughflow within
Ombai Strait. Deep-Sea Research I 48, 1237–1253.
Munk, W.H., 1966. Abyssal recipes. Deep-Sea Research 13, 707–730.
Murray, S.P., Arief, D., 1988. Throughflow into the Indian Ocean through
the Lombok Strait, January 1985–January 1986. Nature 333, 444–447.
Pandey, V.K., Bhatt, V., Pandey, A.C., Das, I.M.L., 2007. Impact of
Indonesian throughflow blockage on the Southern Indian Ocean.
Current Science 93, 399–406.
Schneider, N., 1998. The Indonesian throughflow and the global climate
system. Journal of Climate 11, 676–689.
Smith, W.H.F., Sandwell, D.T., 1997. Global sea floor topography
from satellite altimetry and ship depth soundings. Science 277,
1956–1962.
Sprintall, J., Wijffels, S., Gordon, A.L., Ffield, A., Molcard, R., Susanto, R.D.,
Soesilo, I., Spaheluwakan, J., Surachman, Y., van Aken, H.M., 2004.
INSTANT, a new international array to measure the Indonesian
throughflow. EOS, Transactions, American Geophysical Union 85
(39), 369–376.
ARTICLE IN PRESS
1216
H.M. van Aken et al. / Deep-Sea Research I 56 (2009) 1203–1216
Sprintall, J., Wijffels, S.E., Molcard, R., Jaya, I., 2009. Direct estimates of
the Indonesian Throughflow entering the Indian Ocean: 2004–2006.
Journal of Geophysical Research, in press.
Sprintall, J., Liu, W.T., 2005. Ekman mass and heat transport in the
Indonesian seas. Oceanography 18 (4), 88–97.
Tomczak, M., Godfrey, J.S., 1994. Regional Oceanography, and Introduction. Pergamon, Oxford, p. 422.
Van Aken, H.M., Punjanan, J., Saimima, S., 1988. Physical aspects of the
flushing of the East Indonesian basins. Netherlands Journal of Sea
Research 22, 315–339.
Van Aken, H.M., van Bennekom, A.J., Mook, W., Postma, H., 1991.
Application of Munk’s abyssal recipes to tracer distributions in the
deep waters of the South Banda Sea. Oceanologica Acta 14, 151–162.
Van Riel, P.M., 1956. The Snellius Expedition in the eastern part of the
East Indian Archipelago 1929–1930. Vol. II: Oceanographic results.
Part 5: The bottom water. Ch. II: Temperature.
Vranes, K., Gordon, A.L., Ffield, A., 2002. The heat transport of the
Indonesian Throughflow and implications for the Indian Ocean heat
budget. Deep-Sea Research II 49, 1391–1410.
Wajsowicz, R.C., Schneider, E.K., 2001. The Indonesian throughflow’s
effect on global climate determined from the COLA coupled climate
model. Journal of Climate 14, 3029–3042.
Wyrtki, K., 1961. NAGA Report. Vol. 2: Scientific results of
marine investigations of the South China Sea and the Gulf of
Thailand 1959–1961. Scripps Institution of Oceanography, La Jolla,
p. 195.