Journal of Oceanography, Vol. 60, pp. 205 to 218, 2004
⋅ Air-sea interaction,
⋅ coastal circulation,
⋅ water masses,
⋅ primary produc-
⋅ eastern Arabian
* Corresponding author. E-mail: email@example.com
Present address: National Center for Antarctic & Ocean Research,
Govt. of India, Headland Sada, Vasco-da-Gama, Goa 403804, India.
Copyright © The Oceanographic Society of Japan.
Air-Sea Interaction, Coastal Circulation and Primary
Production in the Eastern Arabian Sea: A Review
ALVARINHO J. LUIS* and HIROSHI KAWAMURA
Center for Atmospheric and Oceanic Studies, Faculty of Science, Tohoku University,
Sendai 980-8578, Japan
(Received 14 November 2002; in revised form 21 April 2003; accepted 19 May 2003)
Air-sea interaction, coastal circulation and primary production exhibit an annual
cycle in the eastern Arabian Sea (AS). During June to September, strong southwest-
erly winds (4~9 m s–1) promote sea surface cooling through surface heat loss and
vertical mixing in the central AS and force the West India Coastal Current
equatorward. Positive wind stress curl induced by the Findlater jet facilitates Ekman
pumping in the northern AS, and equatorward-directed alongshore wind stress in-
duces upwelling which lowers sea surface temperature by about 2.5°C (compared to
the offshore value) along the southwestern shelf of India and enhances phytoplankton
concentration by more than 70% as compared to that in the central AS. During win-
ter monsoon, from November to March, dry and weak northeasterly winds (2–6
m s–1) from the Indo-China continent enhance convective cooling of the upper ocean
and deepen the mixed layer by more than 80 m, thereby increasing the vertical flux of
nutrients in the photic layer which promotes wintertime phytoplankton blooms in the
northern AS. The primary production rate integrated for photic layer and surface
chlorophyll-a estimated from the Coastal Zone Color Scanner, both averaged for the
entire western India shelf, increases from winter to summer monsoon from 24 to 70
g C m–2month and from 9 to 24 mg m–2, respectively. Remotely-forced coastal Kelvin
waves from the Bay of Bengal propagate into the coastal AS, which modulate circula-
tion pattern along the western India shelf; these Kelvin waves in turn radiate Rossby
waves which reverse the circulation in the Lakshadweep Sea semiannually. This re-
view leads us to the conclusion that seasonal monsoon forcing and remotely forced
waves modulate the circulation and primary production in the eastern AS.
ing an intense low-level jet (Findlater, 1977) over the
central AS. In response to these winds a clockwise circu-
lation evolves in the AS (Wyrtki, 1971; Schott, 1983;
Swallow, 1984). The equatorward eastern boundary of this
anticyclonic circulation is known as the West India
Coastal Current (WICC) (Shetye et al., 1990). To the south
of Sri Lanka, the WICC merges with the eastward flow-
ing Southwest Monsoon Current (SMC) which bends
around the Sri Lankan coast and flows poleward into the
Bay of Bengal (Fig. 1(a)).
During the winter or northeast monsoon, generally
from November to February, the winds blow from the
northeast. These winds force the Northeast Monsoon
Current (NMC) toward the west (Fig. 1(b)). Part of this
flow bifurcates at the southwest Indian coast and flows
poleward to form the WICC (Fig. 1(b)) (Shetye et al.,
One of the highest primary productions among the
world oceans occurs in the Arabian Sea (AS) due to a
variety of physical processes introduced by the semian-
nual reversal of monsoon winds (Qasim, 1982; Banse,
1987). This basin, which forms the western flank of the
north Indian Ocean, offers a striking example of wind-
driven ocean circulation (Wyrtki, 1973). During the south-
west or summer monsoon, generally from June through
September, strong winds blow from the southwest form-
206A. J. Luis and H. Kawamura
1991). This poleward flow, which occurs in opposition to
the wind field, is facilitated by a density gradient along
the west coast of India (WCI) (Shetye and Shenoi, 1988).
Reduced solar radiation and high evaporative cooling in
the northern AS lead to high salinity water (>36 ppt), while
low salinity water occupies the region to the south of the
Indian tip (IT). This establishes a density gradient which
leads to the poleward flow of the WICC. During the
intermonsoon period, from March to April and October
to November, weak, highly variable wind regimes (2~3
m s–1) occur in the AS (Hastenrath and Lamb, 1979a) and
the basin surface circulation dissipates (Cutler and
Swallow, 1984). The southwest monsoon period is char-
acterized by sea surface cooling in the central AS
(Colborn, 1976; Düing and Leetmaa, 1980) due to sur-
face heat loss and entrainment (Hastenrath and Lamb,
1979b; Shetye, 1986; Bauer et al., 1991). While a posi-
tive wind stress curl to the north of the Findlater jet pro-
motes Ekman pumping during the summer monsoon, con-
vective cooling during the winter monsoon leads to
densification and downwelling in the northern AS.
In this review we highlight the studies carried out
along the eastern AS, extending from the Pakistan coast
(25.5°N) to south of Sri Lanka (5.5°N) (dashed line in
Fig. 1(a)), and identify “gray” areas that need further at-
tention. The purpose of emphasizing on only the eastern
AS is that recent statistics indicate that of the annual fish
catch, which ranges from 2.2 to 2.8 million tonnes (Cen-
tral Marine Fisheries Research Institute, 1995), about 73%
of the catch originates from the west coast of India due to
a high primary production (PP) in this sector of the AS.
The continental shelf, which we identify by the 500-m
contour in Fig. 1(a), is narrow (50 km) south of Karachi,
widens to ~350 km off the Gulf of Cambay and gradually
tapers towards south to ~120 km at the south Indian tip
At this point, it is relevant to note that the dynamic
processes in the eastern AS are induced by local and re-
mote forcing. During the summer monsoon, offshore di-
vergence of the alongshore wind stress component leads
to coastal upwelling and sea surface temperature (SST)
cooling (Shetye et al., 1985; Muraleedharan and Kumar,
1996; Naidu et al., 1999; Luis and Kawamura, 2002a, b).
Strong summer monsoon winds enhance evaporative cool-
ing in the central AS (McCreary and Kundu, 1989). On
the other hand, wind jets in the equatorial Indian Ocean,
between 5°S to 5°N, excite equatorial Kelvin waves
which, on reflection from the eastern boundary of the Bay
Fig. 1. (a) Geography of the northern Arabian Sea. This review focuses on the region identified by a thick dashed line. Schemat-
ics of summer-monsoon circulation are superimposed. Ekman pumping region in the northern Arabian Sea is highlighted in
yellow tone. Coastal upwelling promoted by divergence of alongshore wind stress component is indicated in green tone.
Current branches indicated are the Ras al Hadd Jet (RHJ), Lakshadweep Low (LL), West India Coastal Current (WICC),
Southwest Monsoon Current (SMC), Sri Lanka Dome (SD) and East India Coastal Current (EICC). The boxes (a), (b), (c) and
(d) are referred in Figs. 7 and 10. The Findlater Jet and wind direction are indicated by bold gray arrows. (b) As in (a), but for
winter monsoon. Convective cooling region is shown in yellow tone. Additional abbreviations shown are: Lakshadweep High
(LH) and Northeast Monsoon Current (NMC).
Air-Sea Interaction, Coastal Circulation and PP in the Eastern Arabian Sea: A Review207
of Bengal, propagate along the perimeter of this basin as
a coastal Kelvin wave and radiate Rossby waves (Yu et
al., 1991). The coastal Kelvin waves propagate along the
perimeter of the Bay of Bengal, bend around the Sri
Lankan coast and propagate poleward along the WCI. The
downwelling (upwelling) Kelvin wave radiates
downwelling (upwelling) Rossby waves which propagate
offshore and promote anticyclonic (cyclonic) circulation
in the Lakshadweep Sea during the winter (summer)
monsoon (Bruce et al., 1994; Shankar and Shetye, 1997;
Bruce et al., 1998). The anticyclonic Lakshadweep cir-
culation is also strengthened by negative wind stress curl
during the winter monsoon.
In Section 2 we describe the monsoon-ocean inter-
action. Section 3 is devoted to coastal circulation, while
PP and chlorophyll-a (Ch-a) mapping studies are high-
lighted in Section 4. The relationship between the physi-
cal dynamics and PP is also explored in Section 3. A dis-
cussion and summary are offered in Section 4.
2. Monsoon-Ocean Interaction
During the northern hemisphere summer, the solar
heating of the Indian subcontinent cause a low pressure
area over Arabia and Pakistan, establishes a north-south
pressure gradient of about 22 hPa (Tomczak and Godfrey,
1994) and drives the southwest monsoon. On the other
hand, wintertime cooling of the landmass causes a south-
ward pressure gradient of about 6 hPa that drives the win-
ter monsoon. The summer monsoon is maintained by
propagation of convective systems of synoptic (lows,
depressions, etc.) and planetary scale (tropical conver-
gence zones, TCZ) from warm tropical oceans onto the
heated subcontinent (Gadgil, 2000). The TCZ is charac-
terized by cyclonic vorticity associated with large shear
of the zonal wind in the north-south direction, at all lev-
els in the lower troposphere above the atmospheric bound-
ary. In this section we focus on seasonal forcing on the
eastern AS using ship-based measurements, satellite-
based observations and modeling studies.
2.1 Ship-based evidence
Figure 2 depicts the fields of monsoon-ocean inter-
action parameters and surface heat budget, which were
estimated using 60 years of ship observations (Hastenrath
and Lamb, 1979a, b). These are presented at 1° × 1° spa-
tial resolution. The top (bottom) row represents the con-
ditions during July (January) corresponding to the sum-
mer (winter) monsoon. During July the wind speed is high
(8~9 m s–1) along the central WCI and westerly winds
with 4–6 m s–1 speed occur south of 12°N (Fig. 2(a)).
The winter monsoon is characterized by northerly to
northeasterly winds with reduced strength compared to
that of the summer monsoon (Fig. 2(b)). Strong winds (6
m s–1) in the Gulf of Mannar (GM) are attributed to
channeling of the winds through the sea-level channel
between the elevated landmass of southern India and Sri
Lanka. These winds have been referred to as gap winds,
and their characteristic features have been highlighted and
studied elsewhere (Luis and Kawamura, 2000, 2001,
The surface atmospheric heat budget during July
exhibits a heat loss (~40 W m–2) slightly north of 10°N
and an equal gain in the south (Fig. 2(c)). In January the
net heat flux is weak and is directed into the ocean to the
south of Saurashtra (Fig. 2(d)). Heat loss in the GM is at
the expenditure of high latent heat (Hastenrath and Lamb,
Fig. 2. Spatial patterns of (a) and (b) wind speed (m/s), (c) and (d) surface net heat flux (W/m2), (e) and (f) wind stress curl
(10–8 dynes/cm3), and (g) and (h) sea surface temperature (°C). The upper (lower) panel represents summer (winter) monsoon.
The data resolution is 1° × 1°. Source: Hastenrath and Lamb (1979a, b).
208 A. J. Luis and H. Kawamura
1979b, Chart 45). The wind stress curl is favorable to
upwelling along the entire coastal region during July, with
highly positive curl induced by the Findlater jet to west
of Karachi (Fig. 2(e)). During January, a highly negative
curl (–3 × 10–8 dynes cm–3) to the southwest of the IT
(Fig. 2(f)) facilitates an anticyclonic circulation in the
Lakshadweep Sea (Bruce et al., 1994, 1998). During July,
positive wind stress curl and the divergence of the
alongshore wind stress component promote upwelling and
SST cooling along the southern WCI (Shetye, 1984). The
winter monsoon SST scenario is as follows. The Ekman
divergence at the coast promotes upwelling; the offshore
advection of this cold leads to low SST to the west of Sri
Lanka (Luis and Kawamura, 2002a). Cold, low-salinity
water from the northern Bay of Bengal (BB) is also
advected into the eastern AS by the Northeast Monsoon
Current (Fig. 1(b)), which lowers SST along the WCI
(Haugen et al., 2002). Cold SSTs that occur in the north-
ern region have been attributed to convective cooling in
the upper oceanic layers induced by dry continental winds
Based on the analysis of Fig. 2 and with inferences
from the literature, the scenario during the monsoons is
as follows. The southwest monsoon is more energetic than
the winter monsoon. The magnitude of the offshore com-
ponent of the wind during the summer monsoon is much
greater than the alongshore component along the WCI
(Shetye et al., 1985). With the alongshore wind stress
component directed toward the equator, upwelling occurs
along the Kerala coast during both monsoons (Shetye,
1984). During the winter monsoon, a high SST cooling
increases mixed layer thickness by ~100 m north of 20°N
(Rao et al., 1989; Prasad, 1996) owing to a decrease in
solar radiation and to high evaporation (Prasad, 1997)
enhanced by dry northeasterly winds (Kumar and Prasad,
2.2 Satellite-based evidence
Satellite-based monitoring offers data with repeated
periodicity, high spatial resolution and wide area cover-
age, and provides an excellent view of the oceanographic
and atmospheric features from the air-sea interfacial layer
to the top of the cloud. The satellite oceanography started
with the launch of SEASAT in 1978. Though this moni-
toring lasted for only three months, it demonstrated the
feasibility of using satellites for geophysical studies. A
few satellite-based studies on the eastern AS have been
carried out and these provide useful insights into the
physical and biological forcing. Recently, Luis and
Kawamura (2000) employed daily-mean NASA
Scatterometer (NSCAT) winds and advanced very high
resolution radiometer (AVHRR) 9-km spatial-resolution
SST and studied the role of wind on SST in the GM dur-
ing January 1997. They argued that the sea-level channel
between the elevated land topography of Sri Lanka and
south India facilitates channeling and acceleration of the
northeast trades through the GM (Fig. 3(a)); these winds
have been referred as gap winds. The study indicated that
the gap winds induce SST cooling of ~1°C due to deple-
tion of surface heat through high latent heat loss (~220
W m–2) (Fig. 3(a)).
Further analysis of the winds derived from the Spe-
cial Sensor Microwave/Imager (SSM/I), AVHRR SSTs
and meteorological variables from the National Centers
for Environmental Prediction/National Center for Atmos-
pheric Research (NCEP/NCAR) reanalysis data set
(Kalnay et al., 1996) indicated that wind stress and sur-
face heat loss were punctuated in time ranging from 15
to 30 days over the winter monsoon period. Furthermore,
the turbulence caused by wind in the surface atmospheric
boundary layer enhances the surface heat loss and lowers
the SST by 1.5°C over a 15-day time scale (Luis and
2.3 Modeling studies of monsoon-ocean interaction
Numerical modeling studies of the coastal air-sea
interaction in this sector of the AS are fragmentary. Shetye
(1986) ran a simple mixed-layer model with Kraus-Turner
thermodynamics for a 2° latitude zonal strip (9–11°N),
extending from Somalia to the Kerala coast, to understand
evolution of the seasonal cycle of SST and mixed layer.
The results indicated that momentum and heat forcing are
effective for the SST evolution during winter monsoon
over the entire strip. In addition to surface forcing, the
importance of entrainment and advection to the evolu-
Fig. 3. Monsoon-ocean interaction monitored by satellites in
the vicinity of the Indian tip during (a) January 1997 and
(b) May 1997. AVHRR sea surface temperature (gray tone)
superimposed with contours of surface net heat flux
(W m–2) and NASA Scatterometer (NSCAT) derived wind
stress vectors (N m–2). Source: Luis and Kawamura (2002a,
Air-Sea Interaction, Coastal Circulation and PP in the Eastern Arabian Sea: A Review209
tion of coastal SSTs during summer monsoon is also
For the first time, a high resolution (9 km), the three-
dimensional, sigma-coordinate, primitive-equation
Princeton Ocean Model (POM) was run to understand the
dynamics of the wintertime wind forcing in the GM (Luis
and Kawamura, 2002a). The stratified POM was fitted
with realistic bottom topography and was forced with
winds observed by NASA Scatterometer (NSCAT) (Fig.
3(a)) to gain insights into the evolution of SST during a
14-day gap-wind event during 21 January to 3 February,
1997. The model solutions revealed that coastal Ekman
dynamics, together with positive wind stress curl, lower
SST west of Sri Lanka during the first 7 days of the event.
The spatially variable and shallow shelf topography at
the IT promotes diapycnal mixing and alters the SST along
the periphery of the IT. The mixed-layer heat budget re-
vealed that the depletion of surface heat modulates SST.
Based on the satellite-based observations (Luis and
Kawamura, 2000, 2001) and with inferences from the
model results, they concluded that SST cooling in the GM
is related to the monsoon-sea-topography interaction.
Numerical study of SST cooling dynamics in the vi-
cinity of the IT prior to the summer monsoon (May) (Fig.
2(b)) also substantiated the role of physical forcing on
the ocean (Luis and Kawamura, 2002b). Using POM with
realistic bottom topography, and daily-mean NSCAT wind
forcing, these authors pointed out that wind-driven Ekman
dynamics along the southwestern India coast facilitates
upwelling-induced SST cooling. An equatorward-directed
coastal jet advects the cold water southward, where the
shallow, irregular shelf topography at the IT enhances
wind mixing, and a sharp gradient at the shelf break (~500
m) modulates upwelling strength through cross-isobath
flow from subsurface. They concluded that the southward
advection of cold water generates low SSTs near the IT,
in agreement with the AVHRR SST maps.
3. Coastal Circulation
Contrasting the eastern boundary current dynamics
of other oceans, the eastern AS show marked seasonal
changes in its eastern boundary currents. The coastal cir-
culation in the eastern AS is fully three-dimensional. We
address this scenario from various dynamical viewpoints
using hydrographic and modeling approaches.
3.1 Hydrographic approach
3.1.1 Water masses distribution
In the AS the seasonal evaporation minus precipita-
tion (E–P) influences the surface salinity. During the sum-
mer monsoon E–P lies between –1 and –1.5 m year–1 to
the east of 70°E and to the south of 20°N (Prasad, 1997);
so high precipitation leads to the formation of Equatorial
Surface Water (Darbyshire, 1967). The upper-ocean wa-
ter masses off the southwestern coast of India are charac-
terized by Indian Equatorial Water (IEW) (Sverdrup et
al., 1942; Emery and Meincke, 1986), with a contribu-
tion from less saline Bay of Bengal Water (BBW) and
more saline Arabian Sea Water (ASW) (Fig. 4; Stramma
et al., 1996). According to Darbyshire (1967), the north-
ward WICC during the summer monsoon carries the low
density ESW to the north and ASW downdrafts and deep-
ens the thermocline to about 100 m. During May to early
September, ESW retreats and ASW shoals to depths less
than 50 m and lowers SST (22°C) along the Kerala coast.
During the winter monsoon, E–P is positive and is
nearly constant (1.5 m year–1) to the north of 10°N. High
evaporation caused by dry northerly winds enhances up-
per ocean convection (Banse, 1968; Prasad, 1997) and
increases surface salinity, which leads to the formation
of the Arabian Sea High Salinity Water mass (ASHSW)
(Rochford, 1964; Kumar and Prasad, 1999). The charac-
teristics of ASHSW are: temperature, 28–24°C; salinity,
36.7–35.3 psu; and depth, 0–100 m. The ASHSW subducts
in the region north of 20°N and spreads equatorward along
the 24σt isopycnal. Below the surface is the water mass
called Red Sea Persian Gulf Intermediate Water (RSPGW;
Emery and Meincke, 1986) or North Indian Intermediate
Water (Wyrtki, 1973). This water mass is formed with
the contribution from Red Sea Water (RSW), which shows
Fig. 4. T-S distribution of conductivity-temperature-depth
(CTD) measurements along 8°N and from 68.1°E to south-
west shelf of India during 5–8 August, 1993. Stations to the
west (east) of 75°E are shown as dashed (solid) lines. Boxes
indicate the range of temperature and salinity values for
Bay of Bengal Water (BBW), Arabian Sea Water (ASW),
Indian Equatorial Water (IEW) and Red Sea Persian Gulf
Intermediate Water (RSPGW). Density surfaces of the core
of Persian Gulf Water (PGW) and Red Sea Water (RSW)
are marked by arrows. Source: Stramma et al. (1996).
210 A. J. Luis and H. Kawamura
a maximum at about 300 m depth, and a contribution from
Persian Gulf Water (PGW), which shows a salinity maxi-
mum between 550 to 800 m depth.
Figure 4 summarizes the temperature-salinity (T-S)
characteristics of different water masses. It was con-
structed from conductivity-temperature-depth (CTD)
measurements along 8°N and spanning from 68°10′ E to
the southwest Indian coast during August 1993 (Stramma
et al., 1996). The BBW is characterized by a T range of
25 to 29°C and S of <35 psu at the surface, and is located
between 74.2°E and the southern India shelf. During June–
August 1987, BBW has been identified at about 150 m
depth with S below 35 psu and T between 14 to 20°C
(Shetye et al., 1990). ASW mass is characterized by T >
29°C and S > 35 psu. It is identified to the west of 75°E.
IEW is identified to the east of 75°E and has a large range
of T from 8 to 23°C and a close S range (34.5 to 35 psu).
On the other hand, RSPGW is identified by low T (5–
14°C) and S of about 35 psu. The density surfaces of the
core of PGW and RSW are indicated by arrows in Fig. 4.
Banse (1968) summarized the upwelling events along
the WCI as follows. (1) During the southwest monsoon
cool subsurface water occupies the entire shelf from
Cochin (10°N) to Karachi (24.5°N); for the latter case
cool water persists till November. (2) Low surface tem-
peratures occur off Saurashtra and along the Pakistan coast
due to atmospheric cooling during winter monsoon.
(3) Upwelling along the southwestern coast (south of
10°N) during the southwest monsoon is due to seasonal
changes of mass distribution and to the anticyclonic cir-
culation of the Arabian Sea, as suggested by Darbyshire
(1967) and Wyrtki (1971).
During the UNDP/FAO Pelagic Fisheries Project
(1971–75) about 200 hydrographic sections were occu-
pied to the south of 17°N (Johannessen et al., 1981). They
noted shoaling of isotherms at the southern coast from
April to September. Besides the role of the local winds,
they pointed out that the anticyclonic AS circulation con-
tributes to coastal upwelling. However, Shetye and Shenoi
(1988) strongly advocated that the wind-driven circula-
tion dominates the WCI, as suggested by Wyrtki (1971).
To further substantiate their claim, hydrocast data were
collected normal to the coast during the 1987 southwest
monsoon (Shetye et al., 1990) and with follow-up obser-
vations during the 1987–88 northeast monsoon (Shetye
et al., 1991). During July–August they reported 2.5°C
surface cooling at the southern Indian coast, which was
attributed to upwelling promoted by the alongshore wind
stress component. The upwelling intensity increased
northward from the IT as the alongshore wind compo-
nent strengthened. As a signature of upwelling they iden-
tified an undercurrent with its core at 150 m depth, which
was discerned up to 20°N.
During the Indian Joint Global Ocean Flux Studies
(JGOFS), oceanographic observations along 12.5, 15 and
10°N were conducted during July–August, 1995 (Fig. 5)
(Muraleedharan and Kumar, 1996). The section along
12.5°N indicated signatures of strong upwelling, marked
by cold SST (~22°C) and low salinity (35 psu) at the coast.
As the upwelling intensity weakened, they inferred high
sea surface salinity (>36 psu). At the 10°N section, the
isotherms above 100 m exhibited shoaling at the coast.
While all of the above studies reported upwelling due to
wind-driven Ekman dynamics, numerical solutions sug-
gest that upwelling-favorable Kelvin waves from the Bay
of Bengal propagate into the eastern AS and modulate
coastal circulation and SST along the WCI (McCreary et
al., 1993; Haugen et al., 2002; Shankar et al., 2002).
3.1.3 Coastal currents
Although a comprehensive hydrographic survey of
the north Indian Ocean was conducted during the Inter-
national Indian Ocean Expedition (IIOE) (Wyrtki, 1971),
a detailed picture of the large-scale coastal circulation
emerged only in the 1990s due to intensive oceanographic
Fig. 5. Vertical distribution of temperature (a, d, g), salinity
(b, e, h) and density (c, f, i) along 15° (top panel), 12.5°
(middle panel) and 15°N (bottom panel) during July–
August, 1995. Unit: temperature, °C and salinity, psu. Re-
produced from Muraleedharan and Kumar (1996).
Air-Sea Interaction, Coastal Circulation and PP in the Eastern Arabian Sea: A Review211
surveys carried out by Indian scientists in the Exclusive
Economic Zone (EEZ) of India. Considerable insights
have been gained through the analysis of hydrocast data
collected along sections normal to the WCI during the
1987 southwest and 1987–88 northeast monsoons (Shetye
et al., 1990). During the summer monsoon they identi-
fied an equatorward shallow (between 75–100 m) surface
flow, with enhanced transport towards the south, and a
poleward flowing undercurrent with its core at 150 m or
below in close proximity to the continental slope.
The hydrographic data of December 1987 and Janu-
ary 1988 provided insights into the poleward coastal cur-
rent along the west coast in a situation where the
alongshore wind component is directed towards the equa-
tor (Fig. 1(b), Shetye et al., 1991). The presence of an
undercurrent flowing equatorward was also inferred from
vertical density sections and dynamic topography. Based
on the alongshore momentum balance equation and with
the support from the inferences from the analytic model
of McCreary et al. (1986), who simulated the intensifica-
tion of a poleward flowing Leeuwin Current and an
equatorward undercurrent off western Australia, Shetye
et al. (1991) argued that the alongshore baroclinic pres-
sure gradient was responsible for the maintenance of a
poleward surface flow during the northeast monsoon.
3.1.4 Coastal sea level
The sea level changes in the eastern AS are induced
by mass redistribution due to currents forced by monsoons
(Figs. 1(a) and (b)) and by remote forcing from the Bay
of Bengal. Locally, the seasonal sea level changes along
the coast are attributed to the astronomical tides, atmos-
pheric pressure, and thermohaline effects; the latter is
dictated by heat and salt exchange across the ocean
boundaries. Every wind system, whether stationary or
moving, will create currents, followed by redistribution
of mass (Sverdrup et al., 1942). Coastal currents forced
by monsoons (Shetye et al., 1990; Shetye et al., 1991)
also alter mass distribution through horizontal advection.
Of these four factors, astronomical tides contribute less
than 1 cm, which is an order of magnitude less than the
meteorological effects (Shankar, 2000).
The climatological seasonal cycle of sea level (cor-
rected for atmospheric pressure) exhibits an annual cy-
cle: sea surface is elevated (depressed) during northeast
(southwest) monsoon (Fig. 6). A good correspondence
between sea level and alongshore currents inferred from
ship drift at Marmagao (15°N), Mangalore (12.9°N), and
Cochin (10°N) is evident. However, a similar compari-
son with the alongshore wind exhibited discrepancies,
suggesting that remote forcing of the coastal circulation
could be important (Shankar, 2000). Shankar and Shetye
(1999), using one-and-half layer and two-and-half layer
reduced gravity models, attempted to relate salinity ef-
fect to the sea-level variability along the coast. They noted
that lower (higher) salinity at the coast leads to higher
(lower) coastal sea level. Since salinity changes are caused
by precipitation and river runoffs, as pointed out by Sastry
and D’Souza (1972), they inferred that the sea level along
the coast can be linked to rainfall in particular, and the
monsoon in general.
3.2 Modeling approach
In this section we highlight the remote forcing as-
pects on eastern AS circulation. A comprehensive numeri-
cal investigation using a two-and-half layer thermody-
namic model provided considerable insights into the lo-
cally and remotely forced dynamics of the north Indian
Ocean (McCreary et al., 1993). These authors pointed out
that the winds along the eastern boundary of the Bay of
Bengal set up coastal Kelvin waves, which bend around
Sri Lanka and propagate poleward along the WCI, forc-
ing changes in the WICC (McCreary et al., 1993; Shankar
and Shetye, 1997). The Kelvin waves originating in the
Bay of Bengal have their origins to the withdrawal and
collapse of the southwest monsoon during August–Octo-
Fig. 6. Climatological seasonal cycle of sea level anomalies
(annual mean removed, solid line) and alongshore currents
(dashed line) at seaports along the west India coast. The
sea level was corrected for atmospheric pressure. Redrawn
from Shankar (2000).
212 A. J. Luis and H. Kawamura
ber. The role of remote forcing was further emphasized
by a test run without Arabian Sea winds; the results re-
vealed that the coastal currents along the WCI were strong,
but they were weaker without Bay-of-Bengal winds.
The remote forcing mechanism is also linked to the
evolution of LH, which manifests as a large anticyclonic
eddy in the upper 300–400 m off the southwest India dur-
ing January (Bruce et al., 1994) (Fig. 1(b)). The dynamic
topographic maps indicated a dynamic topographic height
of about 15–20 cm at the eddy center, while SEASAT-
based sea-surface height anomalies relative to the 1-year
mean (August 1987 to July 1988) showed a westward
migration speed of 14–15 km day–1. Using a 3-layer, re-
duced-gravity, wind-forced tracer model, they pointed out
that the Bay-of-Bengal low-salinity water is advected into
the core of the eddy via the Northeast Monsoon Current
(NMC in Fig. 1(b)). This highlights the importance of
negative wind-stress curl near the southwest coast of In-
dia to the LH genesis. It was also noted that the Rossby
wave radiations from downwelling Kelvin waves along
the WCI also support the eddy genesis in late December
and its decay in April (Bruce et al., 1998).
Shankar and Shetye (1997), using an analytical model
and a one-and-half layer dynamically reduced-gravity
model, pointed out that the sea level changes in the
Lakshadweep Sea exhibits an annual cycle, with a low
(high) appearing during July (January). The results of the
model with an active layer of 100 m when forced with
climatological monthly-mean winds indicated that the
downwelling (upwelling) Rossby waves that are radiated
by downwelling (upwelling) and northward propagating
coastal Kelvin waves generate the anticyclonic (cyclonic)
eddy during January (July) (Figs. 1(b) and (a)). They
pointed out that only the Kelvin waves with period greater
than 40 days were effective in the eddy formation and in
interannual and intraannual variability of the LH and LL.
In an attempt to address the mechanism of the for-
mation of a warm SST pool in the Lakshadweep Sea prior
to the onset of the summer monsoon (May), Shenoi et al.
(1999) also invoked the remote forcing theory. They
pointed out that the downwelling Kelvin waves originat-
ing from the western boundary of the Bay of Bengal force
an equatorward East Indian Coastal Current that carries
low-saline Bay-of-Bengal water into the Lakshadweep
Sea (Fig. 1(b)) and generates a stable surface layer, which,
together with the downwelling Rossby waves radiated
from the IT, promote warm SSTs. This warm pool en-
hances atmospheric moisture convergence and cause deep
convection, leading to the formation of the monsoon vor-
tex during late May; this vortex forms a component of
the monsoon system (Krishnamurthi et al., 1981).
4. Primary Production
Reversal in the surface circulation during monsoons
(Figs. 1(a) and (b)), seasonality in the nutrient distribu-
tion (Banse, 1987) and irradiance (Brock et al., 1994) have
important ramifications for the PP in the north Indian
Ocean. In this section, we first delineate the seasonal dis-
tribution of nutrients using climatology (World Ocean
Atlas-1998, 1999). Then we highlight phytoplankton dis-
tribution using the ship borne measurements, model pre-
dictions and satellite-based chlorophyll-a (Ch-a) obser-
Figure 7 displays seasonal vertical profiles of sili-
cate, nitrate, phosphate and oxygen at the locations iden-
tified by box (a), (b), (c), and (d), respectively, in Fig.
1(a). We compare the coastal ocean nutrient profiles with
open ocean profile sampled at 65.5°E, 12.5°N (panel (d)
in Fig. 7). During the summer monsoon (top row) the
vertical profiles exhibit the following features. At the
northern shelf, the oxygen profile in the upper 150 m is
Fig. 7. Climatological profiles of nitrate, silicate, phosphate
and oxygen during summer (top row) and winter monsoon
(bottom row) at: (a) 23.5°N, 66.5°E, (b) 14.5°N, 73.5°E,
(c) 6.5°N, 77.5°E, and (d) 12.5°N, 65.5°E (see boxes in Fig.
1(a)). Data source: World Ocean Atlas-1998 (1999).
Air-Sea Interaction, Coastal Circulation and PP in the Eastern Arabian Sea: A Review213
characterized by a sharp decline at a rate of 0.4
ml l–1 m–1 in the upper 50 m and a maximum (20 ml l–1)
at 125 m depth. Nitrate exhibits maxima at 100 and be-
low 200 m. Silicate concentration increases below 125 m
and the profile shows maxima at 200 m, 300 m and at
deeper depths. At the central location (box (b)), the sili-
cate profile shows a maximum at 125 m, while nitrate
concentration increases above 20 µM as compared with
the northern box. The increase in the near-surface oxy-
gen is the result of upwelling. For the southern box, ni-
trate concentration exceeds 30 µM below 100 m and the
oxycline exhibits a weaker gradient of 0.03 ml l–1 m–1.
Compared to the coastal region, the open ocean profiles
show sharp contrasts. Higher vertical mixing in the cen-
tral AS and offshore advection of upwelled water from
the western AS (Kumar et al., 2000) enhances oxygen
concentration in the upper 75 m. The oxycline gradient
increases to 0.06 ml l–1 m–1.
The oxygen and nutrient profiles for winter monsoon
are shown in the lower panel of Fig. 7. The following
features are observed. At the northern box, strong con-
vection and overturning in the surface layers increases
mixed layer depth, which enhances oxygen concentration
in the upper 140 m and renders a sharp oxycline (0.05
ml l–1 m–1). Compared to the summer monsoon, the ni-
trate profile exhibits a weak maximum at 200 m. For the
central box, the oxygen concentration depletes in the up-
per 150 m and nitrate concentration peaks at 200 m. Com-
pared to the summer monsoon profile, the silicate con-
centration exceeds 20 µM below 150 m. At the southern
box, higher oxygen concentration is evident in the upper
50 m and nitrate concentration increases between 75 and
300 m. The open-ocean profiles do not show any subsur-
face maximum, but silicate concentration exceeds 30 µM
below 300 m. In conclusion, the oxygen content of the
surface mixed layer is high and enhanced nutrient con-
centration occurs just below the base of the surface mixed
layer during winter monsoon.
4.2 Primary production and phytoplankton blooms
Here the focus is on ship-based PP and phytoplankton
measurements and ecosystem models. Based on 21 cruises
in the northern AS north of 20°N, Qasim (1982) carried
out a comprehensive study of PP among other oceano-
graphic aspects. He found that surface PP is high off
Mumbai, Gulf of Kutch and Saurashtra, and off Pakistan
(Fig. 8). The annual average PP in coastal areas was re-
ported to be ~37 mg C m–3, which was three times than in
the offshore regions. The values during the summer
monsoon were one third of the value in other seasons.
Madhupratap et al. (1996) advocated that the wintertime
sea surface cooling drives convection processes that lead
to the injection of nutrients into the surface waters, lead-
ing in turn to high PP in the northeastern AS; the evi-
dence for this has been provided using Indian Ocean Color
Monitor (OCM) Ch-a composite images (Chauhan et al.,
2001). As for the column PP (integrated over the euphotic
layer), Qasim’s computations indicated 270 × 106 tonnes
C year–1, constituting more than 50% of the total produc-
tion of the northern AS. The column-integrated PP peaked
during the summer monsoon. His study also showed that
about 50% of the annual Ch-a production in the coastal
sectors of the northeastern AS, constituting ~26 mg m–3
(15 × 103 tonnes year–1). The annual-mean surface PP
based on 14C uptake showed high values (50–100
mg C m–3 day–1) along the central WCI and in the Gulf of
Kutch, while high values of column-integrated PP over
the euphotic layer (1–2 g C m–2 day–1) are identified along
the Saurashtra coast (Figs. 8(a) and (b)).
As a part of Indian JOGFS program, the seasonality
and composition of phytoplankton composition was de-
ciphered along the WCI and compared with that along
the 64°E meridian (Sawant and Madhupratap, 1996).
These authors inferred higher phytoplankton population
density during the summer and winter monsoon period,
which for the upper 120 m ranged from 1.7 to 170 and
0.3 to 10 (×108) cells m–2, respectively. Their species-
wise classification showed that diatoms constituted about
86% and cyanobacteria and dinoflagellate were 7% and
6% of the total population for all the seasons combined.
Compared to the open ocean, the cell count (cells l–1) for
the coastal population of phytoplankton species was
larger. Among the diatom population Chaetoceros sp.
topped the list with 75% higher counts in the coastal re-
gion. This study suggested that physical mechanisms,
Fig. 8. Primary production in the northern Arabian Sea.
(a) Annual mean rates of primary production (PP) based on
14C uptake at the surface (mg C/m3/day) and (b) integrated
PP for the euphotic zone (g C/m2/day). Gaps in the distri-
butions indicate no data. Source: Qasim (1982).
214 A. J. Luis and H. Kawamura
among other constraints, could be attributed to higher
phytoplankton counts in the coastal region. On the other
hand depletion in nutrients, especially nitrate, in the open
ocean waters restricts the PP (Gupta et al., 1980).
To isolate the physical mechanisms that promote the
phytoplankton activity, McCreary et al. (1996) ran a four-
component ecosystem model, with a 2 1/2 layer physical
model system with surface mixed layer embedded in the
upper layer and a biological model with advective-diffu-
sive equations in each layer that determine nitrogen con-
centrations in four compartments: nutrients,
phytoplankton, zooplankton and detritus. They compared
the model solutions with climatological Coastal Zone
Color Scanner (CZCS) data. Their model solutions show
upwelling blooms, detrainment blooms, and entrainment
blooms in the AS. The upwelling blooms are promoted
by a high underwater irradiance and a large nutrient sup-
ply to the shallow surface mixed layer through Ekman
pumping in the northern AS and through upwelling in-
duced by divergence of the alongshore wind stress com-
ponent along the southern India coast during summer
monsoon. The detrainment blooms, in initially nutrient-
rich but irradiance-constrained water, are attributed to:
(1) a gradual build up of nutrients for several months prior
to the blooming event, (2) advection of nutrients from
western sectors of the AS toward the central AS, and
(3) shoaling of the mixed layer. These blooms were iden-
tified in the northwestern and southwestern AS. The en-
trainment blooms are caused primarily by vertical entrain-
ment of nutrients and by secondary sources such as recy-
cling after the blooms onset via remineralization of de-
4.3 Satellite-based observations
Although satellite observations relate to only the
upper part of the euphotic zone, near-surface periodicity
of phytoplankton, if caused by hydrographic processes,
is bound to reflect events involving the entire mixed layer.
Based on this rationale, ocean color data from the CZCS,
OCM and Sea viewing Wide Field-of-view Sensor
(SeaWiFS) have been a central part of the following stud-
Using CSCZ Ch-a data from November 1978 through
December 1981, Lierheimer and Banse (2002) examined
seasonal and interannual variability of phytoplankton pig-
ment, between 5 and 16°N, and pointed out that
phytoplankton blooms promoted by upwelling during
summer monsoon are restricted to the shelf, and, on a
few occasions, the phytoplankton blooms do occur into
the Lakshadweep Sea and to the south of India. Based on
the analysis of time-sequence CZCS images, they ruled
out indigenous phytoplankton blooms in the nutrient-poor
Lakshadweep Sea. A similar inference was provided by
Banse and English (1994) using global maps of Ch-a
medians. They considered medians rather than temporal
means because the histograms of the monthly means were
non-Gaussian, exhibiting tails with extreme Ch-a values.
Their climatological median values indicated a highest
peak (0.9 mg m–3) during September and a secondary peak
(0.3 mg m–3) during February at 13°N, 66°E. Emphasiz-
ing on Ch-a variability in different sectors of the AS,
Banse and English (2000) examined geographical differ-
ences in seasonality of CSCZ-derived phytoplankton pig-
ment values during 1978–86. For the sector lying to the
north of 17.5°N and east of 62°E, which forms part of
our review area, Ch-a mean values varied from 0.5 to 1.5
mg m–3 with large standard deviation (indicating blooms)
on a time scale of 2–8 weeks during late winter monsoon.
With a gradual surface warming and with enhanced strati-
fication, a decrease in the Ch-a concentration to less than
0.2 mg m–3 was inferred during early June.
The seasonal Ch-a variability during 1978–86 for the
Gulf of Mannar has been described using 1-km spatial
resolution Ch-a maps derived from CSCZ (Yapa, 2000).
Yapa noted that Ch-a concentrations were higher during
the summer monsoon than the winter monsoon. Upwelling
was considered as the principle mechanism for high Ch-
a production. Recently, composite maps of SST and Ch-a
from OCM have also been employed to provide fishery
forecasts along Gujarat coast based on the mapping of
important oceanic features (Solanlki et al., 2001). Initial
results indicated that the forecasts were effective in en-
hancing the fish catch 2–3 fold. They opined that a close
coupling between physical and bio-chemical processes
could be further explored to locate potential fishing zones.
4.4 Physical mechanisms and primary production
To gain insights into the seasonal modulation of PP
and Ch-a in relation to mixed layer dynamics and light
penetration, we present the monthly progression of mixed
layer depth (MLD), photic depth, PP rate integrated for
the photic zone and surface Ch-a estimated from CSCZ
(Fig. 9). These parameters were averaged for the coastal
region along the WCI or West India Coastal Province re-
ferred to by Longhurst (1998). MLD is the depth where
the water density is lowered by 0.125σt from the surface
σt. The photic depth represents the depth where the light
intensity falls by 1% of surface irradiance for water types
I–III (Jerlov, 1964). PP rate (Pt) and Ch-a averaged for
photic zone were estimated from CSCZ data using the
algorithm of Sathyendranath et al. (1995).
MLD shoals due to coastal upwelling in response to
spin-up of the southwest monsoon and the pycnocline is
illuminated from March to October (Fig. 9, upper panel).
Vertical flux of nutrients into the surface mixed layer
enhances PP and Ch-a, which show peak values of 70
g C m–2·month and 21 mg m–2 during September and Au-
gust, respectively (Fig. 9, lower panel). A decrease in the
Air-Sea Interaction, Coastal Circulation and PP in the Eastern Arabian Sea: A Review215
Ch-a and PP after September is attributed to a buildup of
herbivore population. A weak secondary maximum in
March in PP and Ch-a is a reflection of wintertime con-
vective cooling and deepening in the MLD.
5. Discussion and Summary
The researches summarized in this review indicate
that the characteristic monsoon-ocean interaction is im-
portant in driving and modulating the strength of coastal
currents, evolution of SST, eddy formation, which even-
tually modulates the PP. With the support of the research
work outlined in the previous sections, we now attempt
to seek answers to following issues.
What is the characterization of the annual cycle of
surface forcing parameters? A temporal march of net heat
flux (Qn, negative (positive) represents sea surface cool-
ing (warming)), wind speed (W), SST, and wind stress
curl (available for only four months of the year, hence
shown by open circles), derived from Hastenrath and
Lamb (1979a, b), are shown in Fig. 10. Each point repre-
sents a 1° × 1° average for box (a), (b) and (c) identified
in Fig. 1(a). For the northern box, Qn and SST exhibit
bimodal oscillations, with peak SST lagging Qn by a
month; W shows only one peak during July (Fig. 10(a)).
The curl is highly positive (4 × 10–8 dynes cm–3) during
In the vicinity of the central WCI, just south of 15°N,
bimodal oscillations in Qn and SST are also evident, but
the amplitude is almost half of that in the northern box;
moreover, the undulations in SST profile are marginal,
with an annual mean of ~28°C. Winds are fairly weak
during the first and last fours months of the year; they are
strengthened in July (Fig. 10(b)). Moving southward,
while the Qn profile is positive throughout the year, the
wind speed displays an interesting pattern: high wind
speed (6 m s–1) occurs during June to September, with
nodes of weak magnitude (<2 m s–1) during inter-monsoon
periods (Fig. 10(c)). In brief, the annual cycle of surface
forcing reveals the following. Qn and SST show a bimo-
dal cycle along the coast, with decreasing amplitude to-
wards the south. Wind and its curl, which is positive,
weaken southward with a peak in July. The coastal re-
gions are conducive to warming during January to April
and during post-summer monsoon periods.
AS is a wind-driven basin with seasonally varying
coastal currents. Coastal upwelling is a consequence of
the alongshore wind stress component all year round along
this coast. So, what is the seasonal upwelling scenario
along the coast? With the monthly mean alongshore wind
stress component directed toward the equator (Shetye et
al., 1985), divergence of the alongshore wind-stress would
Fig. 10. Climatological seasonal march of wind speed (W),
surface net heat flux (Qn), sea surface temperature (SST),
and wind stress curl which is represented by open circles as
it was available for only January, April, July and October.
Area-mean was taken on each of the 1° × 1° boxes shown
in Fig. 1(a). The centered locations are: (a) 23.5°N, 66.5°E,
(b) 14.5°N, 73.5°E and (c) 6.5°N, 77.5°E. Data source:
Hastenrath and Lamb (1979a, b).
Fig. 9. Climatological seasonal march of mixed layer depth
(MLD), photic depth, primary production rate integrated for
photic zone (Pt) and surface chlorophyll-a (Ch-a) from
Coastal Zone Color Scanner. Source: Longhurst (1998).
216 A. J. Luis and H. Kawamura
promote upwelling throughout the year. In fact,
hydrographic data suggest that the upwelling starts in
March and continues till September along the Kerala coast
(Johannessen et al., 1981).
What is the relation between monsoon forcing and
PP? PP along the southern India coast is enhanced during
the summer monsoon due to an increase in the upwelling
intensity promoted by offshore flux of surface water in
response to equatorward alongshore wind stress compo-
nent. The northern AS supports high PP in response to
upward Ekman pumping promoted by positive wind stress
curl induced by the Findlater jet (Fig. 1(a)). The scenario
during the winter monsoon is dominated by upper-ocean
convective cooling and deepening of MLD to ~100 m
(Rao et al., 1989) and vertical influx of nutrients into the
euphotic layer, which together with an increase irradi-
ance enhance PP on an annual scale (Qasim, 1982;
Madhupratap et al., 1996). In the context of local forc-
ing, the monsoons modulate the geographical distribu-
tion of PP and Ch-a in the eastern AS.
In summary, in this review we synthesize
multidisciplinary aspects, ranging from air-sea interac-
tion to PP, based on a number of investigations published
in international and Indian journals. The intention is to
provide a condensed, illustrative summary of these in-
vestigations that would serve as a guide for future works.
It should also provide insights on what has been explored
and what still needs to be probed. Currently, there are
either no answers or only partial answers to the follow-
ing questions. (1) What is the coupling between AS SST
and rainfall over the Indian subcontinent? (2) What is the
seasonal pattern of salinity along the eastern AS? (3) What
is the seasonal cycle of Ch-a distribution in EEZ of the
Indian coast? (4) What is role of El Niño-Southern Oscil-
lation on coastal SST? (5) What is the role of shelf to-
pography on the three-dimensional circulation? A major
obstacle in addressing these issues is the lack of data of
sufficient spatial and temporal resolution. With more than
a dozen satellites in space and with the operational satel-
lite oceanography in place, and with enhanced comput-
ing systems, it would be tempting to attempt these and
many unresolved issues relating to the dynamics of this
highly productive and important sector of the north In-
This study was supported by MEXT RR2002 Project
for Sustainable Coexistence of Human, Nature and Earth
(category 7). Additional support from ADEOS-I and II
projects of NASDA, Japan, is also acknowledged.
Arabian Sea High Salinity Water
NCEP/NCAR National Centers for Environmental Prediction/
National Center for Atmospheric Research
NMC Northeast Monsoon Current
OCM Ocean Color Monitor
PGWPersian Gulf Water
POM Princeton Ocean Model
PP Primary Production
Pt Primary Production rate
Surface net heat flux
RHJ Ras al Hadd Jet
RSPGW Red Sea Persian Gulf Intermediate Water
RSWRed Sea Water
SDSri Lanka Dome
SeaWiFSSea viewing Wide Field of view Sensor
SMC Southwest Monsoon Current
SSM/I Special Sensor Microwave Imager
SST Sea Surface Temperature
TCZ Tropical Convergence Zone
UNDP/FAO United Nations Development Program/Fishery
WCI West Coast of India
WICC West India Coastal Current
Arabian Sea Water
Advanced Very High Resolution Radiometer
Bay of Bengal Water
Conductivity Temperature Depth
Coastal Zone Color Scanner
Exclusive Economic Zone
East India Coastal Current
Equatorial Surface Water
Evaporation minus precipitation
Gulf of Mannar
Indian Equatorial Water
International Indian Ocean Expedition
Joint Global Ocean Flux Studies
Mixed Layer Depth
National Aeronautics and Space Administration
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