Physical Preliminaries

# Physical Preliminaries - The Arabian Sea: A Monsoon-Driven Environment

    Introduction
    Monsoon Systems
    Monsoon Variability
    Surface Layer Response
    Phytoplankton Response
    Thermohaline Circulation and Water Masses
    Fisheries of the Arabian Sea

## Introduction

The intent here is to review the northern Arabian Sea ecosystem in relationship to the region's climate and both its monsoonal, wind-driven and thermohaline ocean circulation components. An attempt is made to lay out in one place the background for an Arabian Sea U.S. GLOBEC effort and to provide an introduction to the literature. With regard to the overall literature it is worthwhile mentioning several general sources. The reader is directed to the volume edited by Fein and Stephens (1987) for an up-to-date discussion of the various aspects related to the monsoon in meteorology, literature, history, climatology and oceanography. The International Indian Ocean Expedition (IIOE) hydrographic atlas compiled by Wyrtki (1971) is an excellent source on water mass distribution and the overall ocean circulation. Qasim (1982) and Swallow (1984) provide recent reviews of the regions general oceanography. The best single sources on the biology of the Indian Ocean can be found in the volume compiled by Zeitzschel (1973) and the National Institute of Oceanography at Goa's 25th anniversary volume (Desai, 1992). The reader should also consult the U.S. JGOFS Arabian Sea Process Study document (Smith et al., 1991) for reviews of the regional processes. Finally there is a special issue of Deep-Sea Research on the northern Arabian Sea based in part on British, U.S. and German work there in 1986-87 which is in press.

This review will begin with some basic aspects of the monsoons and their climatic variability. The basic physics of the surface layers of the Arabian Sea and their relation to primary production are then treated. This is followed by a discussion of the intermediate-depth low-oxygen layer and its relationship to the circulation and biogeochemical balances.

## Monsoon Systems

Monsoon systems are planetary-scale, seasonal cycles in atmospheric circulation associated with ocean-continent thermal contrasts and typically movements of the intertropical convergence with its characteristic band of convection (Webster, 1987a; Fig. 1). The role of the land-sea temperature difference to these planetary "sea-breezes" was first appreciated by Halley (1686). Hadley (1735) pointed out the importance of the earth's rotation in determining the nature of the resulting winds. As pointed out by Webster (1987a) the only major process not covered in these two pioneering works was that of convection; i.e., the importance of evaporation and precipitation to the monsoon. The strongest monsoon system occurs between the continent of Asia and the Indian and western Pacific Oceans. These consist of the northern winter monsoon with flows off of the Asian continent and a northern-hemisphere summer monsoon with flows of moist maritime air onto the continent. A weaker system, the West African monsoon, occurs in the Gulf of Guinea (Halley, 1686; Hamilton and Archbold, 1945; Webster, 1987a).

The Asian monsoon, which can be broken into an eastern and Indian branch, consists of a build-up of high pressure over the Asian highlands in the northern-hemisphere winter. This produces the northeast Indian monsoon (NE monsoon) with near-surface winds which flow out of the northeast over the Indian subcontinent and out across the Arabian Sea. The other branch brings flow off the Chinese mainland out over the island archipelagos of the Philippines and Indonesia. Both branches cross the equator and push the intertropical convergence into the southern hemisphere. The situation reverses with the build-up of high pressure south of the equator in the Indian Ocean in April and lowering of surface pressure over the Asian highlands as summer approaches. This leads to a reversal in the winds proceeding first along East Africa and then across the center of the Arabian Sea and onto the Indian Subcontinent. The flow across the Arabian sea within this southwest monsoon takes the form of an intense low-level jet, known as the Findlater jet after Findlater (1969). As depicted in Findlater's (1969) work the jet typically bifurcates before it reaches the Indian coast into two branches which are denoted as the Somali and Split jet. The placement and time dependence of this split is important in setting the bounds of the wind stress curl induced open ocean upwelling which is a dominant feature of the Arabian Sea monsoon response. The NE and SW monsoons are separated by transition periods or monsoon lulls during which the wind speeds drop and the sea surface heats.

A major factor controlling the air-sea interaction during the monsoon and the structure of the near-surface ocean responses to this cycle is the alternate temperature rise to nearly 30 deg. C during the lulls and the cooling which occurs through air-sea heat exchange during the high-wind periods. The phenomenon known as Arabian Sea cooling (Duing and Leetmaa, 1980, McCreary and Kundu, 1989) is most intense in the SW monsoon when the northern and western Arabian Sea are dominated by upwelling. This involves coastal upwelling along both the Arabian and Somali coasts (Currie et al., 1973; Schott, 1983; Currie, 1992), and open-ocean upwelling in the mid-basin associated with the wind-stress maximum under the Findlater jet (Smith and Bottero, 1977; Swallow, 1984; Bauer et al., 1991). The upwelling and surface heat exchange creates thermal anomalies in the western Arabian Sea which vary from year to year (Brown and Evans, 1981). While it has been suggested that these in turn modify the monsoon rainfall over India (Shulka, 1975; Raman et al., 1992), this effect is not seen in all models (Washington et al., 1977). During the SW monsoon, large-scale upwelling also occurs off southwest India (Banse, 1968; Shetye, 1984).

## Monsoon Variability

The Indian monsoon is highly variable over a range of time scales (c.f. Kutzbach, 1987; Webster, 1987b). The variability is of interest to U.S. GLOBEC both in the context of climate variations and for operational reasons. Shorter-term variability within a monsoon is addressed first, followed by a discussion of the longer-term variations between monsoons.

The onset of the SW monsoon is variable but usually takes place between mid-May and mid-June. Southern regions have a larger range of onset dates than more northerly ones (cf. Krishnamurti, 1987). Onset in Bombay between 1977-1980 occurred within 7 days of June 15th (Krishnamurti, 1987). Onset over the northern Arabian Sea is essentially a tropical storm, the "onset vortex", which varies in intensity, but is strong enough to be a danger to shipping. The onset in Bombay is a good indication of the presence of the onset vortex, although Bombay is not necessarily affected in any given year. The monsoon is interrupted by lulls in rain and wind known as monsoon breaks. An example of these cycles between active monsoon periods and breaks in two normal monsoon years (see longer term variability below) is given in Fig. 2. It is not clear to what degree breaks involve wind reduction over the Arabian Sea. The monsoon breaks seem to be related to the passage of individual convective complexes (cloud bands) (Webster, 1987). At a slightly longer period there are 30-50 day oscillations in the monsoon (Krishnamurti et al., 1985a,b) which correspond to a global-scale oscillation in the tropics.

Interannual variations in monsoon strength appear from the scale of the bi-annual oscillation in the tropical atmosphere (Cadet and Diehl, 1984; Dube et al., 1990) to longer-period variations related to cycles ranging from ENSO (El Nino/Southern Oscillation) to glacial-interglacial time scales. The origin of the Southern Oscillation index stems from Walker's (1910, 1924) work on monsoon prediction. The result was a system more in tune to a Pacific basin phenomenon which became a major focus in oceanography in the 1970's (Bjerknes, 1969; Wyrtki; 1975). The ENSO phenomenon has seen a wealth of quantification over the last two decades. The connection between the ENSO and monsoons over the Arabian Sea, however, remains unclear. Barnett (1985) and Webster and Yang (1992) are but two of several attempts to consider the interaction between monsoon and ENSO. The latter authors suggest a selective coupling between the onset phase of the Asian summer monsoons and global-scale summer circulation patterns with weak global trade winds being tied to weak Asian monsoons. Although the exact nature of the latter coupling is uncertain in the sense of whether or not ENSO events start in the Indian sector, it is clear that Indian rainfall is less following ENSO events and heavy during anti-ENSO periods. A good example is the 1987 monsoon which had almost no rain over portions of India (Krishnamurti et al., 1989). It is important to note that much of the change in this monsoon involved low winds south of the equator and along the Somali coast. Winds over the central Arabian Sea were near climatological values (Bauer et al., 1991). There is little discernible trend in monsoon rainfall records over the last hundred years, but there is considerable interdecadal variation. The 1900's, 1920's, and 1960's all had lower rainfall, while the 1910's, 1930's, 1940's, 1950's, and 1970's saw higher rainfall (Fig. 3).

Finally, paleoceanographers have been able to use sediment records from the northern Arabian Sea to consider the longer-term variations tied to ice age cycles with their associated solar radiation forcing, and longer-term events such as the rise of the Tibet-Himalayas massif (Prell, 1984; Kutzbach, 1987; Prell and Kutzbach, 1987). Sediment records and modelling (Luther et al., 1990; Prell et al., 1990) suggest weaker monsoons at the height of the last glaciation (18 Kyr BP) and stronger monsoons at the interglacial maximum (9 Kyr BP).

## Surface Layer Response

The physical response of the surface layers of the northern Arabian Sea to the monsoon cycle is important to understanding the operation of the region's pelagic ecosystems. Basic descriptions of the surface forcing fields and the nature of the mixed layer can be found in Cadet and Diehl (1984), Hastenrath and Lamb (1979), and Molinari et al. (1986). Surface currents are discussed in Cutler and Swallow (1984) and Molinari et al. (1990). Here the principal features are covered along with a description of conditions along ~65 deg. E during the NE monsoon of 1986-87 and the SW monsoon of 1987 (Bauer et al., 1991). The wind patterns for the month of December and August are shown in Fig. 4.

Extremes of near-surface thermal stratification at the heights of the two monsoons are shown in Fig. 5. Note the overall shift in the slope of the thermocline; the large scale gradients of the thermocline are changed throughout the depth of sampling. Currents in the NE monsoon case are to the west (into the page). The flow reverses in the SW monsoon such that the flow is eastward and slightly stronger than in the northern winter. The SW monsoon section reveals a major mid-ocean upwelling zone separated from the coastal upwelling zone by a reversed flow. Smith and Bottero (1977) first suggested a separation of these two upwelling domains. The reversal can be interpreted as the westward end of a permanent anticyclonic circulation over the Murray Ridge (Quraishee, 1984, Bauer et al., 1991). Connections between these two upwelling systems in three dimensions remain to be fully established. This has ecological implications, particularly in respect to the coupling between coastal upwelling and mid-ocean conditions. This is one of the many cases where the Arabian Sea is a model in which to understand interactions between physics and biology in other flow-coupled pelagic ecosystems.

Mixed layers are deepest in the NE monsoon in the northwestern area (average about 125 m, Banse 1984; see also Fig. 5). During this season, they shoal to less than 20 m around 5 deg. N (Bauer et al., 1991). Conditions during the SW monsoon in 1987 show the deepest mixed layers (~60 m) around 15 deg. N with a minimum in mixed layer depth under the Findlater jet. The shallow mixed layers under the maximum winds are coincident with the upwelled dome of fluid in the SW monsoon (Fig. 5). Deepening of the mixed-layer by wind stirring is curtailed by the upwelling of the thermocline at rates of order 10-5 m s-1 (Bauer et al., 1991). The deep mixed layers to the south of the wind maximum can only be explained through advection of dense mixed layers southward in the Ekman layer and the subsequent downward motion of the thermocline associated with Ekman pumping, with maximum vertical velocities of -2.5 X 10-5 m s-1 (Bauer et al., 1991). The simple two-dimensional mixed-layer model of Bauer et al. (1991) suggests that the upwelled structure reaches an equilibrium in mixed-layer depth where the upward motion balances the turbulent entrainment rate. The deep mixed layers to the south continue to deepen in the simple model since turbulent entrainment and downwelling both act to increase mixed-layer depth.

## Phytoplankton Response

During the SW monsoon, the combination of vertical motion and entrainment causes a very abrupt transition in phytoplankton biomass and productivities along an axis from the oligotrophic deep mixed layer up across the wind maximum into the mid-ocean upwelling region. This effect is obvious in chlorophyll sections along 65 deg. E (Fig. 6). in which the coastal upwelling zone (with even higher phytoplankton biomass and productivity) is not shown. Chlorophyll biomass and carbon-14 productivities during these sections ranged from lows of 15-30 mg m-2 and 0.3-0.6 g C m-2 day-1 in the deep mixed layer regime south of 15 deg. N to extremes of 40 mg m-2 and 1.1 g C m-2 day-1 in the mid-ocean upwelling region. These values agree with estimates from the IIOE work (Banse, 1988). It is the two dimensional development of the mixed layer associated with Ekman transports that lead to this coupled bloom and oligotrophic region. See Yentsch and Phinney (1992) for a similar explanation making use of the geostrophic relation in the thermocline. Quantities are even higher in the upwelling plume off Ras al-Hadd where in the 1987 cruise Chl-a biomass was 80 mg m-2 and productivities were 2.5 g C m-2 day-1 (Hitchcock and Frazel, 1989). In contrast, the NE monsoon period (Fig. 6a) exhibits a lower Chl-a biomass over the entire region and lower productivities.

The description based on single sections through the region in single years cannot give a reasonable picture of the fields in light of the evidence discussed above for interannual variations. A powerful tool for quantifying the representative nature of limited ship based measurements is remote sensing. A series of papers (Banse and McClain, 1986; Brock and McClain, 1992; Brock et al., 1991) describe phytoplankton blooms in the Arabian Sea from CZCS data. There are blooms in both monsoons in the northern extreme of the basin (Banse, 1987), presumably caused by mixing related to cooling over the Gulf of Oman and Murray Ridge. Brock and McClain (1992) suggest interannual variability, which might be driven by the biannual oscillations in the forcing as discussed above. The reliable record, however, is very short. This should be improved with the SeaWiFS deployment in the time frame of the proposed Arabian Sea efforts.

## Thermohaline Circulation and Water Masses

In most of the world's ocean the intermediate to deep circulation driven by water mass formation is of little direct consequence to the near-surface ecosystem. This is not the case in the Arabian Sea, where a layer over a kilometer deep is depleted in oxygen (Fig. 7). It is "suboxic", not anoxic; it does not usually contain hydrogen-sulfide. It does, however, have essentially no oxygen as evidenced by the depletion of nitrate in the water column (Deuser et al., 1978; Broecker and Peng, 1983; Naqvi, 1987; Naqvi, 1991; Naqvi et al., 1992). Recognition of the influence of this low oxygen layer on the biology of the region goes back at least to Gilson (1937) and the later Soviet work of Vinogradov and Voronina (1962a) and has formed one focus of studies in the region.

The physical and biological dynamics responsible for the oxygen minimum involves a complicated balance between the advection or mixing of oxygen into the layer and the removal of oxygen in the water column by respiration. Oxygen is added to the Arabian Sea by inflow of cold oxygen rich waters from the south up either the Somali or Indian coasts or by horizontal mixing between the Arabian Basin and the equatorial zone. Some oxygen is added through the introduction of warm, low oxygen waters from the marginal basins, the Red Sea and Arabian or Persian Gulf (Banse, unpubl.; Olson et al., 1993). Alternatively, oxygen can mix into the intermediate layers from the shallower or deeper adjacent waters (Swallow, 1984; Naqvi et al., 1992). The biological uptake of oxygen to balance these inputs depends ultimately on the primary production and how it is utilized. The available data suggests that the total phytoplankton production on an annual basis is not extraordinary (Olson et al., 1993) although the intense bloom and decay cycles may enhance the ratio of new to old production. A key issue in understanding the bloom and decay cycle involves better quantification of the grazing down of the phytoplankton stocks during the monsoons and the details of the trophic dynamics during the oligotrophic monsoon lulls.

There are many indications that the low-oxygen zone may be very variable and that it has varied considerably in the past. There are at least two studies that suggest the time scales for respiration versus oxygen input are very short compared to intermediate and deep water circulation times. Using anthropogenic chloroflurocarbon F-11 as an inert tracer, high-quality oxygen measurements and the productivity work quoted above, Olson et al. (1993) estimate that the time scale of oxygen renewal in the Arabian Sea is 23 years. A similar calculation based on the nitrogen cycle rather than that of oxygen gives time scales on the order of a year for the near surface portions of the low oxygen layer (Naqvi et al., 1992). Either estimate is very short relative to the time scales estimated for chlorofluorocarbons for the geographically analogous oxygen minima in the eastern North and South Atlantic Oceans (Doney and Bullister, 1992; Warner and Weiss, 1992). These time scales suggest that the oxygen minimum layer might be sensitive to decadal variations in the monsoonal forcing of surface productivity and circulation and to changes in the input from the marginal seas.

## Fisheries of the Arabian Sea

In closing it is worthwhile to briefly explore the major pelagic fisheries of the Arabian Sea. The Indian Ocean is has one of the last under-exploited tuna fisheries (Rothschild and Yong, 1970; Sharp, 1992). As mentioned above, however, the extensive low oxygen zone apparently restricts all but some of the smaller tuna and occasional yellowfin. The southern coast of Oman supports a fishery for Indo-Pacific spanish mackerel (Scomberomorus commerson) with maximum landings in February to April (Dudley et al., 1992). Otolith analysis suggests that these fish grow up to 80 cm in their first year (Dudley et al., 1992). An old fishery centered in Salalah exploits the oil sardine (Sardinella longiceps) for camel fodder; the same species is a major contributor to the Indian fish catch. This species recruitment has been studied relative to the onset of the monsoon and coastal upwelling off southwestern India by Longhurst and Wooster (1990) but similar studies have not been carried out farther north on the Indian coast or off Oman. Their conclusion that the stocks of sardine are tied to the onset time and strength of remotely forced upwelling might also be applicable to the Omani coast. Mackerel, sardines and demersal fish (Banse, 1959) are all susceptible to upwelling of low oxygen waters into the coastal environment. The mackerel and sardine catches show marked interannual variations which are probably tied to variations in the physical forcing as suggested by the correlation between sardine catches at Cochin and sea level (Longhurst and Wooster, 1990). There is also evidence of a connection between the strength and onset of the monsoon and sardine catch (Murty and Edelman, 1971; Longhurst and Wooster, 1990) although further work with longer records is required to further quantify the connection. Fluctuations in mackerel catch in the same fishery are not correlated with deviations in sardines (Longhurst and Wooster, 1990). The variations in overall Indian mackerel catches are approximately decadal (Noble, 1992) with no long term trends. The final stock worthy of mention are the myctophid stocks in the Gulf of Oman. They are not currently being exploited although it has been suggested that they are exploitable (Gjosaeter, 1984).