Indian Ocean Dipole: Processes and Impacts

INDIAN OCEAN DIPOLE: PROCESSES AND IMPACTS

P. N. Vinayachandran,1 P. A. Francis2 and S. A. Rao3

Equatorial Indian Ocean is warmer in the east, pendently. Coupled models have been able to re- has a deeper thermocline and mixed layer, and produce IOD events and process experiments by supports a more convective atmosphere than in such models — switching ENSO on and off — the west. During certain years, the eastern In- support the hypothesis based on observations that dian Ocean becomes unusually cold, anomalous IOD events develop either in the presence or ab- winds blow from east to west along the equator sence of ENSO. There is a general consensus among and southeastward off the coast of Sumatra, ther- different coupled models as well as analysis of data mocline and mixed layer lift up and the atmo- that IOD events co-occurring during the ENSO are spheric convection gets suppressed. At the same forced by a zonal shift in the descending branch time, western Indian Ocean becomes warmer and over to the eastern Indian Ocean. Processes that enhances atmospheric convection. This coupled- initiate the IOD in the absence of ENSO are not ocean atmospheric phenomenon in which convec- clear, although several studies suggest that anoma- tion, winds, sea surface temperature (SST) and lies of Hadley circulation is the most probable forc- thermocline take part actively is known as the In- ing function. Impact of the IOD is felt in the vicin- dian Ocean Dipole (IOD). Propagation of baro- ity of Indian Ocean as well as in remote regions. clinic Kelvin and Rossby waves, excited by anoma- During IOD events, biological productivity of the lous winds, play an important role in the de- eastern Indian Ocean increases and this in turn velopment of SST anomalies associated with the leads to death of corals over a large area. More- IOD. Since mean thermocline in the Indian Ocean over, the IOD affects rainfall over the maritime is deep compared to the Pacific, it was believed continent, Indian subcontinent, Australia and east- for a long time that the Indian Ocean is passive ern Africa. The maritime continent and Australia and merely responds to the atmospheric forcing. suffer from deficit rainfall whereas India and east Discovery of the IOD and studies that followed Africa receive excess. Despite the successful hind- demonstrate that the Indian Ocean can sustain its cast of the 2006 IOD by a coupled model, forecast- own intrinsic coupled ocean-atmosphere processes. ing IOD events and their implications to rainfall About 50% percent of the IOD events in the past variability remains a major challenge as is under- 100 years have co-occurred with El Nino South- standing reasons behind an increase in frequency ern Oscillation (ENSO) and the other half inde- of IOD events in recent decades.

1. INTRODUCTION

Monsoons dominate the climate of the North Indian Ocean and the neighbouring land mass. Winds over the Indian Ocean (Fig. 1) blow from the southwest during May–September and north- 1Centre for Atmospheric and Oceanic Sciences, east during November–February and drive a cir- Indian Institute of Science, Bangalore 560 012, culation that reverses its direction completely, un- India. e-mail: [email protected] seen in any other part of the world oceans (Schott 2 Indian National Centre for Ocean Information and McCreary, 2001; Shankar et al., 2002). Be- Services, P. B. No. 21, Hyderabad 500 055, India. cause of monsoons, the spatial distribution of sea e-mail: [email protected] surface temperature (SST) in the Indian Ocean is 3Indian Institute of Tropical Meterology, Pashan, Pune 411 008, India.e-mail: [email protected] characterised by warm water on the eastern side Citation: P. N. Vinayachandran, P. A. Francis, S. and cooler on the west which is in contrast to Pa- A. Rao, Indian Ocean dipole: processes and cific and Atlantic Oceans that are warmer on the impacts, Current Trends in Science, N. Mukunda west. In fact, all characteristic properties of the (ed), Indian Academy of Sciences, Bangalore, 2009 ocean show marked east-west assymetry in Indian (in presss). Ocean. Freshwater input into the ocean, in the Figure 1. Climatology of winds (vectors in m/s), SST (oC, color shaded) and outgoing longwave radiaion (OLR, contours in Wm−2) over the Indian Ocean. form of rain and drainage from the land, is greater exceeding 1 ms−1, known as Wyrtki (Wyrtki, 1973) on eastern side causing the eastern part of the basin jets. Wyrtki jets transport warmer upper layer to have lower salinity than the west. A major up- water towards the east and accumulates near the welling system is located off the western boundary boundary causing the thermocline in the east to and consequently biological productivity is higher be deeper than in the west. The thermocline slope in the west than in the east. The thermocline, becomes greater during spring and fall and the vol- ume of warm water and the heat content of the which separates warmer water of the oceanic mixed Indian Ocean is larger on the eastern side than layer from the colder water underneath, is located in the west. (When the thermocline is shallow, deeper in the east than in the west. Asymmetry in winds have to do lesser work to bring cooler water the SST also influences the overlying atmosphere: into the mixed layer than when it is deep. That convection is higher over the eastern part of the is, for the same wind-strength, a shallower ther- ocean and lesser over the west (Fig. 1). mocline can facilitate cooling of the mixed layer The equatorial Indian Ocean, however, experi- and SST, whereas a deeper thermolcine may not.) ences a somewhat different seasonal wind forcing Thus, Wyrtki jets dictate the shape of the thermo- and hence has a different response than the Ara- cline in the equatorial Indian Ocean. Climatologi- bian Sea and Bay of Bengal. Winds over the cally, since the eastern equatorial Indian Ocean is equatorial Indian Ocean, particularly their zonal warmer, it supports a more convective atmosphere than in the west (Figure 1). The departure of the component, are weak during monsoons. Relatively ocean-atmosphere system from this mean state oc- strong westerly winds, however, appear during the curring during certain years, characterized by op- transition between monsoons, first during April– posite sign of anomalies in the east and west, is May (spring) and then again during October– known as the Indian Ocean Dipole (IOD). November (fall). These winds drive strong east- Possibilities of air-sea coupled processes over the ward currents along the equator, attaining speeds Indian Ocean has been suggested by several stud- Figure 2. Mean September–October composite anomaly patterns of (a) SST (oC, shaded) and depth of 20oC isotherm (m, contours) and (b) OLR (Wm−2, shaded) and surface winds (ms−1, vectors) for the positive IOD years during 1979–2008 for which satellite data of OLR is available. The composite is based on positive IOD years (Meyers et al., 2007) of 1982, 1991, 1994, 1997 and 2006. SST data used is HadISST available at (http://badc.nerc.ac.uk); depth of 20oC isotherm is derived from Simple Ocean Data As- similiation (SODA,(Carton et al., 2000) ) available at (http://iridl.ldeo.columbia.edu); NOAA AVHRR OLR data and NCEP Reanalysis surface wind data are obtained from (http://www.cdc.noaa.gov). ies in the past. Reverdin et al. (1986) analyzed erly winds. Saji et al. (1999) recognised anomalous interannual variations of convection and SST in conditions such as in 1961, 1994 and 1997 in the the equatorial Indian Ocean using ship reports. tropical Indian Ocean as a dipole mode character- They suggested that coupled air-sea dynamics over ized by low SST off Sumatra and high SST in the the Indian Ocean should be considered in order western equatorial Indian Ocean accompanied by to understand the interannual variability. This wind and rainfall anomalies. They showed that the suggestion was prompted by anomalous conditions IOD is a coupled ocean atmosphere phenomenon, in 1961, characterized by significantly cold SST it is independent of ENSO and about 12% of the anomalies in the east (between 80oE–90oE) and SST variability in the Indian Ocean was associated warm SST anomalies in the west. From limited with the dipole mode. Highlighting the 1997–98 ship observations, Reverdin et al. (1986) concluded event, Webster et al. (1999) also have suggested that the SSTA affects cloudiness, rainfall and con- that anomalous conditions present during this pe- sequently, causes westward wind anomalies along riod was due to an internal mode of the climate the equator. Anomalies during 1961 were not con- system of the Indian Ocean. A coupled ocean at- current with an ENSO which pointed to the role mosphere process in which oceanic Rossby waves cause a deepening of the thermocline was proposed of air-sea interaction processes controlled by In- to be a necessary feature that led to the sequence dian Ocean. Hastenrath et al. (1993) showed that of events culminating in anomalies that contrast zonal structure of the equatorial Indian Ocean re- the in east-west direction Webster et al. (1999). sponds to changes in the strength of westerlies oc- The ocean-atmosphere coupling during an IOD in- curring during the transition from summer to win- volves atmospheric convection, winds, SST and up- ter monsoon (October–November). Stronger west- per ocean dynamics. The discovery of the IOD led erlies drive a faster equatorial jet causing deeper to the recognition that the Indian Ocean can sup- mixed layer and warmer SSTs in the east and en- port its own coupled ocean-atmosphere variations tails cold water upwelling in the western equatorial (Schott et al., 2009) and is not merely a slave to the Indian Ocean. Changes in the ocean state affects events happening over the Pacific Ocean in connec- the atmosphere, particularly rainfall over eastern tion with ENSO. Africa. SST variability in the eastern Indian Ocean The Indian Ocean circulation and its impact on during 1994 was observed to have strong cou- regional climate has been a topic of great interest pling to the shallowing of the thermocline (Mey- recently and several reviews have appeared in the ers, 1996). The shallow thermocline during 1994 literature. The monsoon circulation of the Indian was due to unusually weak equatorial jets (Vinay- Ocean has been reviewed by Schott and McCreary achandran et al., 1999) caused by anomalous east- (2001). Yamagata et al. (2004) and Annamalai and Figure 3. Same as in 2, but for negative IOD years since 1979 listed in Meyers et al. (2007), ie., 1980, 1981, 1985 and 1992.

−1

−2 1950 1955 1960 1965 1970 1975 1980 1985 1990 1995 2000 2005 Figure 4. The dipole mode index (DMI) deﬁned (Saji et al., 1999) as the diﬀerence in SST anomalies between westen (50oE–70oE) and eastern (10oS–10oN) Indian Ocean. The DMI is normalized by its standard deviation and smoothed by a 5-month running mean before plotting.

Murtugudde (2004) have reviewed the interannual is discussed in section IV, followed by mechanisms variability in the Indian Ocean and its impact on that can trigger IOD in section V. Section VI sum- climate variability. The role of tropical Atlantic, marizes the impact of IOD on regional climate and Pacific and Indian Oceans on climate variability outstanding issues are listed in section VII. in the tropics has been reviewed by Chang et al. (2006). In a recent review, Schott et al. (2009) 2. INDIAN OCEAN DIPOLE have summarized the role of Indian Ocean circulation on climate variability. It is well recognised The Indian Ocean Dipole refers to an anoma- that the dominant signal of coupled interannual lous state of the ocean-atmosphere system (Saji variability in the Indian Ocean is the IOD. The et al., 1999; Webster et al., 1999; Murtugudde objective of this review is to summarize our un- et al., 2000). During the mature phase of an IOD, which happens during September–October, derstanding of the IOD, particularly the processes eastern equatorial Indian Ocean becomes unusu- associated with IOS events and their implications, ally cold and the western equatorial Indian Ocean onsidering that there has been rapid development unusually warm (Fig. 2). Cold SST anomalies in our understanding of these. A description of (SSTA) suppress atmospheric convection in the the IOD phenomenon is given in the next section east whereas warm SSTA enhances convection in followed by a section devoted to processes related the west. Winds blow westward over the equato- to IOD. The relationship between ENSO and IOD rial Indian Ocean and from the southwest off the Figure 5. Monthly mean composite anomaly patterns of SST and depth of 20oC for the pIOD years. Alternate months starting from March in the pIOD year are plotted. The top right panel corresponds to January of the year following pIOD. coast of Sumatra, the latter being favourable for nals are strongly coupled to surface signals (Rao coastal upwelling. Equatorial jets become weak, et al., 2002). Due to the vertical movement of reducing the eastward transport of warm water en- isotherms, magnitude of temperature anomalies at tailing a shallower than usual thermocline in the any given depth is much greater than at the sur- east. Sea level decreases in the eastern equatorial face. Anomalous state of the ocean-atmosphere Indian Ocean and rises in the central part. Ther- system described above is referred to as a positive mocline rises in the east and deepens in the cen- IOD (pIOD). The reverse, negative IOD (nIOD), tral and western equatorial Indian Ocean. Tem- also occurs, characterized by warmer SST anoma- perature anomalies associated with IOD are also lies, enhanced convection, higher sea level, deeper seen in sub-surface ocean (Vinayachandran et al., thermocline in the east and cooler SSTA, lower sea 2002; Horii et al., 2008) and the sub-surface sig- level, shallower thermocline and suppressed con- Figure 6. Monthly mean composite anomaly patterns of OLR (Wm−2, shaded) and surface wind (ms−1, vectors)) for the pIOD years. Alternate months starting from March in the pIOD year are plotted. The top right panel corresponds to January of the following year of pIOD. vection in the west (Fig. 3). Negative IOD can ern equatorial Indian Ocean (Fig. 4). Averages of be considered as an intensification of the normal anomalies are calculated for a box in the western o o o state whereas positive IOD represents conditions Indian Ocean bounded by 50 E–70 E and 10 S– o o o nearly opposite to the normal. Therefore, major 10 N and in the east by 10 S–Equator and 90 E– 110oE. In general, the DMI is expected to be higher focus of attention has been to understand pIOD. than one standard deviation and should remain This article concerns mostly with pIOD events and so for 3 to 4 months, in an IOD year. The west for simplicity referred to as IOD hereafter. minus east anomalies are positive during a pIOD An IOD event can be detected using Dipole year and vice versa. During the 1870–1999 period, Mode Index (DMI) which is defined as the differ- 27 pIOD events occurred and recent IOD events ence in SST anomalies between western and east- are 1961, 1963, 1967, 1972, 1982, 1994, and 1997 (Yamagata et al., 2002; Vinayachandran et al., phyll concentration in the western Indian Ocean 2007). During 2003, a pIOD event was initiated and along the coast of India shows a decrease dur- but it was terminated midway (Rao and Yamagata, ing the IOD years of 1997 and 2006 (Wiggert et al., 2004). Most recently, an IOD event occurred dur- 2006). Enhanced biological productivity off the ing 2006 (Vinayachandran et al., 2007). Twenty coast of Indonesia during 1997 led to the death nIOD events have also occurred during the 1870– of coral reef in a large area due to the asphyxia 1999 period (Meyers et al., 2007). caused by the oxidation of organic matter in the Evolution of the IOD (Fig. 5) is strongly water column (Abram et al., 2003). locked to the seasonal cycle due to the thermo- Coral records have provided evidences for the dynamic air-sea feedback between an atmospheric presence of IOD in the present as well as in the anticyclone located to the east of Sumatra (Anna- past because upwelling and cooling taking place malai et al., 2003) and the ocean underneath being dependent on the seasonal cycle of winds (Li in the eastern ocean during IOD events leave their et al., 2003). Typically, an event begins to appear signature in the marine environment which are pre- during late spring/early summer, matures during served in coral records. Anomalies of Sr/Ca indi- September–November and most of the anomalies cate temperature anomalies in the ocean whereas 18 disappear by January of the following year. East- δO indicates influence of both temperature and erly equatorial wind anomalies and warm SST salinity. Such records from Mentawai Islands, off anomalies in the central part of the Indian Ocean Indonesia showed that several IODs occurred in the begin to appear during spring. SST anomalies as- past, including an event in 1877 which was stronger sociated with the IOD peak during September– than that during 1997 (Abram et al., 2003). On November. For certain events, however, there is the western side, coral density bands of Malindi a clear indication that the atmospheric compo- Marine Park, Kenya also present evidences of the nent appears much earlier in spring in the form IOD (Kayanne et al., 2006). Here, anomalies of of easterly wind anomalies and southeasterlies off δO18 were correlated with the increase in rainfall Sumatra (Vinayachandran et al., 1999 2007). Even during IOD events. The Sr/Ca ratio and δO18 though the general features of every event is simi- (after removing temperature signals) from coral lar, there are differences in the location and timing records carry distinct signals of temperature and of the peak anomalies. For example, the peak SST salinity anomalies respectively, and therefore, can anomalies during the 1994 event occurred about 2 be used for reconstructing past records of IOD months earlier compared to the 1997 event. The (Abram et al., 2008). Such records show that the eastern pole is located more or less in the same IOD during middle Holocene were characterized by region for every event because coastal upwelling is an important cooling mechanism. In the west, longer duration of surface cooling caused by en- however, the region covered by positive anomalies hanced cross equatorial winds. vary considerably between events. During the 2006 event, for example, positive SSTA were located 3. PROCESSES to the west of the commonly considered IOD box bounded by 50oE–70oE; 10oS–10oN(Vinayachan- Evolution of SST anomalies involves cou- dran et al., 2007). pled ocean-atmosphere processes (Vinayachandran Copious upwelling takes place along the east- et al., 1999; Murtugudde and Busalacchi, 1999; ern boundary of the EIO during an IOD which Vinayachandran et al., 2002). Oceanic processes leads to enhanced biological productivity and af- that are responsible for cooling SST anomalies fects the bio-geochemistry of the region. The satel- are: (1) shoaling of thermocline and (2) varia- lite data coverage of chlorophyll concentration is tions in mixed layer depth. The first process de- restricted to the events of 1997-98 and 2006 (Wig- pends chiefly on subsurface ocean dynamics (i.e, gert et al., 2009; Iskandar et al., 2009). During the coastal and open ocean upwelling), while the sec- strong IOD event of 1997–1998, the eastern equatorial Indian ocean, which is not usually very pro- ond one helps in re-distributing the heat energy in ductive, experienced phytoplankton blooms (Mur- the top few meters of the ocean. Similarly, atmo- tugudde and Busalacchi, 1999). This event was spheric processes that are involved in evolution of also associated with an increase in chlorophyll a SST are (1) atmosphere-ocean heat fluxes and (2) concentration in the southeastern Bay of Bengal, large-scale atmospheric circulations. In order to and a decrease in the southwestern part of the bay understand the role of oceans and atmosphere in (Vinayachandran and Mathew, 2003). In the Ara- evolution of SST anomalies during IOD events, we bian Sea, influence of IOD led to a decrease in describe oceanic and atmospheric processes sepa- chlorophyll a concentration, primary productivity rately in this section, focussing on their major con- and sea to air CO2 flux (Sarma, 2006). Chloro- tributions to SST anomalies. I I I I I I I I I I I

Figure 7. First EOF mode of sea level anomalies from SODA (upper panel). Positive Indian Ocean Dipole events are marked as I in the corresponding time series (lower panel).

3.1. Oceanic processes development(Vinayachandran et al., 2007). The Ocean general circulation models have provided western warming is mostly controlled by air-sea useful insights into processes that lead to the devel- fluxes and advection by ocean currents. Anoma- opment of SSTA during IOD events (Murtugudde lous downwelling Rossby waves deepens the ther- et al., 2000; Vinayachandran et al., 2002). Forced mocline which prevents cooling of the SST by en- by winds and surface fluxes, they have been able trainment (Murtugudde et al., 2000). The role of to simulate the evolution of IOD extremely well waves is elaborated further below. (Vinayachandran et al., 2007). Such models have 3.2. Role of baroclinic waves illustrated that upwelling, forced either by winds The semiannual zonal winds over the central or through the propagation of baroclinic waves, equatorial Indian Ocean excite Kelvin waves which is an important process responsible for SST cool- remotely affect sea level variations and ther- ing in the east. The semiannual zonal winds over mal structure along the eastern equatorial Indian the central equatorial Indian Ocean drive eastward Ocean (Clarke and Liu, 1993; Yamagata et al., equatorial jets. Anomalies in zonal winds weaken 1996; Meyers, 1996). Easterly wind anomalies or even reverse the Wyrtki jet, which casues a de- along the equator drive an anomalous upwelling crease in eastward transport resulting in a shallow Kelvin wave which radiates into the eastern In- thermocline in the east (Hastenrath et al., 1993; dian Ocean. After reaching the eastern basin they Vinayachandran et al., 1999). Alonghore compo- reflect into coastally trapped Kelvin and Rossby nent of winds effectively upwells cooler water to the waves. This upwelling Kelvin wave lifts up the mixed layer when the thermocline is shallow. Air- thermocline in the eastern equatorial Indian Ocean sea fluxes also play a significant role in the SST thereby triggering upwelling in the southeastern cooling, partcularly during initial stages of IOD tropical Indian Ocean (SETIO) (Vinayachandran et al., 1999; Murtugudde et al., 2000; Vinayachan- equatorial Indian Ocean, erodes (forms) during dran et al., 2002; Rao et al., 2002). Local along- positive dipole years as a response to decreased (en- shore winds, which are southeasterly during de- hanced) precipitation (Murtugudde et al., 2000). velopment and peak phase of the Indian Ocean This feeds back to the SST in the form of enhanced Dipole, also drive the coastal upwelling in the SE- cooling (warming) due to strong vertical mixing TIO (Vinayachandran et al., 1999; Murtugudde (shallow mixed layer) in the eastern (western) In- et al., 2000). Due to the strong easterly anoma- dian Ocean. Masson et al. (2004) examined the lies along equator and southeasterly winds along role of salinity on equatorial currents. The anoma- Sumatra/Java coasts, strong anomalous anticy- lous easterlies over the equator drive south equa- clonic wind stress curl forms in the southern trop- torial current at surface and an equatorial under ical Indian Ocean (Xie et al., 2002). This anoma- current at subsurface. The south equatorial cur- lous wind stress curl forces downwelling Rossby rent transports the normal fresh pool westward, waves in the southern tropical Indian Ocean dur- increases stratification and creats a barrier layer in ing developing phase of the IOD (Rao and Behera, the western equatorial Indian Ocean. As a result of 2005). Similarly, downwelling Rossby waves are this shallow stratification, wind mixing is restricted formed in the northern tropical Indian Ocean. In and amplitude of the momentum increases in the normal years open-ocean upwelling also occurs in surface layer, resulting in strengthening of south the southwestern tropical Indian Ocean forced by equatorial current and continuation of equatorial negative wind stress curl (Masumoto and Meyers, under current (Masson et al., 2004). These results 1998; Xie et al., 2002). Downwelling Rossby waves suggest that there is a positive feedback between radiated from the southeastern Indian Ocean prop- salinity and SST in the tropical Indian Ocean dur- agate into this upwelling region, deepen the ther- ing a positive IOD event. mocline and warm the SST in the western tropi- 3.4. Atmospheric processes cal Indian Ocean (Murtugudde et al., 2000; Vinay- Compared to the southeastern coasts of tropi- achandran et al., 2002). These features are brought cal Pacific and tropical Atlantic Oceans character- out clearly in a simple empirical orthogonal func- ized by cold SST tongue and low stratus clouds, tion (EOF) analysis of sea level anomalies obtained the SETIO (off Sumtra) is a region of warm pool from SODA (Carton et al., 2000). The first EOF (Vinayachandran and Shetye, 1991) with deep con- mode of sea level anomalies clearly show the struc- vection (Gadgil et al., 2004; Li et al., 2003). Over ture of Kelvin and Rossby waves in the tropical a warm ocean such as the tropical Indian Ocean, Indian Ocean. During IOD years, the anomalous a modest SST anomaly may induce deep convec- upwelling Kelvin waves propagate eastward, re- tion in situ, so that SST-cloud relation is in phase flect and then propagate as coastal trapped Kelvin (Li et al., 2003). Since, a major part of the tropi- waves along Sumatra/Java and the Bay of Bengal cal Indian Ocean is covered by a warm pool (SST (Fig. 7). This results in low sea level anomalies in greater than 28 oC) (Vinayachandran and Shetye, the eastern equatorial Indian Ocean and the Bay of 1991), the SST-cloud relation is in phase. On the Bengal. The anomalous downwelling Rossby waves other hand, in tropical Pacific, significant phase forced by anti-cyclonic wind stress curl during IOD differences are observed; maximum SST anomalies events are seen as high sea level anomalies with are observed in the eastern tropical Pacific during northwest-southeast slope (Rao and Behera, 2005). a warm ENSO, but maximum anomaly in convection associated with ENSO appears in the central 3.3. Role of Salinity Pacific. In-phase SST-cloud relation in tropical Interannual sea surface salinity variation in the Indian Ocean leads to negative feedback between Indian Ocean are caused both by freshwater forc- the atmosphere and ocean. As cloud amounts in- ing and ocean dynamics (Murtugudde and Busalac- crease in the tropical Indian Ocean, in response to chi, 1998; Perigaud et al., 2003; Yu and McCreary, the increase in SST, the net downward shortwave 2004; Vinayachandran and Nanjundiah, 2009). In- radiation decreases and cools the SST (Li et al., terestingly, a coupled model simulation show that 2003). Another thermodynamic air-sea feedback significant interannual salinity variations in the In- in the SETIO is the Bjerknes feedback mechanism. dian Ocean occur only during IOD years (Vinay- For example, a drop in SST in the SETIO implies a achandran and Nanjundiah, 2009). Scanty obser- corresponding drop in cloud amount in situ. A de- vations of salinity in the tropical Indian Ocean pre- crease in atmospheric convection leads to develop- vent the identification of the role of salinity on ment of descending Rossby waves to the west (Gill, IOD. However, using OGCMs, a few researchers 1980), resulting in an anomalous anticyclonic cir- have examined the role of salinity on SST during culation to the west of decreased convection cen- IOD event. The barrier layer that is present (ab- ter. Since, winds along SETIO are southeasterly sent) during normal years in the eastern (western) during the development phase of IOD, the anti- Warm Upwelling Kelvin Wave SSTA EQ Cool Sumatra SSTA Downwelling Rossby Wave Coastal Upwelling Anomalous Anticyclone

Mean Southerly

Figure 8. Schematic of coupled positive feedback in th southeastern equatorial Indian Ocean. Modified after (Li et al., 2003). cyclonic circulation enhances wind speeds in the Sumatra and Java, contribute further to IOD ter- SETIO. The enhanced winds, in addition to driv- mination. ing coastal upwelling and vertical mixing, induce 3.5. Intraseasonal Oscillation strong evaporative cooling of the sea surface, lead- Madden and Julian (1972) first reported vari- ing to more cooling of the SETIO (Fig. 8). This ability at time scales of 30–60 day in zonal winds positive feedback acts only during boreal summer, and convection, popularly known as MJO. Anoma- as mean winds along SETIO reverse in the fall. lous convection first develops in western tropical During strong IOD years (such as 1994 and 1997), Indian Ocean, propagates eastward across the In- seasonal southeasterlies associated with the Asian dian Ocean and maritime continent, and then de- cays east of 180oE. Associated with the propagat- summer monsoon reinforce coastal upwelling asso- ing anomalous convection, baroclinic circulation ciated with the IOD near Sumatra in boreal sum- anomalies in the wind field are excited; easter- mer and fall, assisting in initiation and strength- lies (westerlies) to the east (west) of the anoma- ening of these events (Halkides et al., 2006). On lous convection in the lower (upper) troposphere. the other hand, boreal fall equatorial westerlies In the suppressed convection phase of MJO, baro- deepen the thermocline in the eastern basin, re- clinic circulation anomalies also occur in the at- ducing the cold SSTA in the eastern equatorial mosphere with opposite directions. The MJO activity becomes weak during IOD years in response region and near Sumatra and Java. Such a pro- to the strong subsidence in the SETIO (Rao and cess contributed to the early termination of the Yamagata, 2004). During pure IOD events, MJO 1994 IOD. Seasonal reversal of monsoon winds dur- appears before the termination of IOD and the as- ing boreal winter reduces coastal upwelling along sociated westerlies anomalous downwelling Kelvin waves and deepen the thermocline in the SETIO, musson and Carpenter, 1982; Huang and Kinter, arresting air-sea coupling. Suppressed convection 2002; Hendon, 2003; Krishnamurthy and Kirtman, phase of MJO, associated with equatorial easterly 2003). wind anomalies, can trigger a IOD event by excit- Simultaneous correlations between SSTA in the ing anomalous upwelling Kelvin waves (Han et al., eastern Indian Ocean and ENSO indices tend to 2006; Rao et al., 2009). Their model simulations suggest that IOD events are forced by ENSO (Al- suggested that before the onset of the strong 1997 lan et al., 2001; Bequero-Bernal et al., 2002; Xie dipole, basin-wide wind-driven oceanic resonance et al., 2002; Hastenrath, 2002; Krishnamurthy and with a period near 90 days, involving the propaga- Kirtman, 2003; Annamalai et al., 2003). The tion of equatorial Kelvin and first-meridional-mode strong IOD event of the recent times that oc- Rossby waves across the basin, was excited. As- curred during 1997 coincided with a very strong sociated with the resonance was a deepened ther- ENSO. However, IODs have occurred together mocline in the eastern basin during August and with ENSO as well as independently, and SSTA early September, which reduced the upwelling in caused by ENSO lack the dipole pattern as dur- the eastern antinode of the IOD, thereby delaying ing the IOD events. Moreover, anomalies associated with IOD are phase-locked to the sea- the reversal of the equatorial zonal SST gradient sonal cycle, they appear during boreal summer and — an important indicator of a strong IOD — by fall, whereas, basin-wide anomalies associated with over a month compared to 1994 IOD event. ENSO appear mostly during winter and spring (Tozuka et al., 2008). EOF of SST anomalies show 4. DEPENDENCE OF IOD ON EL NINO a basin-wide pattern as the first mode due to the effects of ENSO. The second mode, on the other 4.1. Observations hand, shows remarkable sea-saw between the east- The mean thermocline in the Indian Ocean is ern and western Indian Ocean which is due to the deep, compared to the Pacific. Therefore, ini- IOD and this mode is independent of ENSO (Ya- tiation of a coupled phenomenon that involves magata et al., 2004). About 50 percent of the shallowing of the thermocline (so that the sub- IOD events have occurred together with ENSO surface ocean can interact with the SST through and the rest independently (Saji and Yamagata, wind mixing) is considered to be an uphill task. 2003; Meyers et al., 2007). Clearly, the ocean- This conviction led to the tenet that the Indian atmosphere coupling intrinsic to the Indian Ocean Ocean is passive and merely responds to the at- can lead to the development of an IOD event (Saji mosphere (Kiladis and Diaz, 1989; Venzke et al., et al., 1999; Yamagata et al., 2003 2004; Behera 2000; Alexander et al., 2002), and interannual vari- et al., 2006; Chang et al., 2006; Schott et al., 2009). ability of the Indian Ocean SST are forced mainly by ENSO, through an atmospheric bridge (Klein 4.2. Coupled models et al., 1999; Alexander et al., 2002). The bridge Coupled ocean-atmosphere general circulation is through Walker circulation which, climatologi- models have been successful in reproducing the cally, has ascending motion over the western Pa- IOD events, as illustrated first by (Iizuka et al., cific and maritime continent, and descending mo- 2000) and characteristics of IODs simulated by coupled models are comparable to that seen in tion over eastern Pacific. This normal Walker observations (Iizuka et al., 2000; Bequero-Bernal cell reverses during ENSO, with strong descend- et al., 2002; Lau and Nath, 2004; Cai et al., 2005; ing motion over western Pacific and equatorial In- Behera et al., 2006; Song et al., 2007). The major dian Ocean. About 4 months after large-scale SST issue that has been addressed using such models is anomalies are observed in the equatorial Pacific, to test whether IOD is generated by processes in- the whole of the Indian Ocean is covered with SST trinsic to the Indian Ocean or by externally forcing anomalies of the same sign. The warming (cooling) from the Pacific. of the Indian Ocean during warm (cool) phase of In a coupled simulation for 50 years (Iizuka ENSO is usually basin-wide and mainly due to sur- et al., 2000), the model generated 8 IODs com- face heat fluxes (Klein et al., 1999; Venzke et al., pared to 20 IOD events in a 100 year simulation 2000). The 4 month delay for the Indian Ocean using CCSM 2 (Vinayachandran and Nanjundiah, to warm in response to ENSO is due to the ther- 2009). Iizuka et al. (2000) found poor correlation mal inertia of the ocean surface layer. The In- between DMI and SSTA in the NINO3 region and dian Ocean warming associated with ENSO, how- concluded that the IOD in this model is indepen- ever, is not basin-wide at all stages of ENSO. In dent of ENSO. Bequero-Bernal et al. (2002) car- the early stages of development of an ENSO, there ried out two simulations using a coupled model, are cold SSTA in the eastern equatorial Indian first with effects of El Nino included and then ex- Ocean but they disappear as El Nino matures (Ras- cluded. They concluded that dipole like variability in the model is forced primarily by ENSO but they cline in coupled models show considerable varia- also found that the model additionally generates a tion between models (Saji et al., 2006) and the dipole structure forced by atmospheric fluxes, in- frequency of occurrence of IOD in these models dependent of ENSO. The SINTEX model simula- is crucially dependent on the shape of the thermo- tion had 20 IODs in its 80 year model integration cline. In ocean atmosphere-coupled models, dy- (Gualdi et al., 2003). The peak model SSTA as- namical coupling takes place through wind stress sociated with the IOD occurred about 3 months and thermodynamic coupling through SST and air- before the SSTA peak in the NINO3 region sug- sea heat flux. Thus generation of ENSO in a cou- gesting that IOD evolution in the model is gen- pled model can be suppressed by not allowing the erally independent of ENSO. However, the model atmosphere to interact with the model SST. This results suggested that sea level pressure anomalies is typically achieved by replacing model SST with associated with ENSO might produce atmospheric a climatology (Elliott et al., 2001; Bequero-Bernal conditions that are favourable for the generation of et al., 2002; Behera et al., 2005). The dynamical IOD. In the GFDL coupled model simulations (Lau coupling through wind stress is present in such ex- and Nath, 2004), IOD pattern was caused primar- periments which can produce a Bjerknes type of ily by ENSO; changes in surface wind field over the feed back in the presence of a shallow thermocline. Indian Ocean associated with ENSO modify both Consequently, oceanic variability associated with air-sea fluxes and oceanic process leading to dipole ENSO is not well suppressed and influences oceanic SST pattern. The model also produced strong IOD regions in the SETIO (Fischer et al., 2005). A events in the absence of ENSO and these were second method also has been used (Fischer et al., caused by an annular mode associated with sea 2005), in which, the dynamical coupling is sup- level pressure anomalies south of Australia. Using pressed by prescribing climatological wind stress the SINTEX model, Fischer et al. (2005) carried in the tropical Pacific. This constraint removes out a coupled model experiment in which the ef- the atmospheric signals associated with ENSO and fect of ENSO was suppressed by constraining wind reduces the oceanic signals dramatically (Fischer stress over tropical Pacific. The IOD in this simu- et al., 2005). Considering that oceanic processes lation was identical to the run that included ENSO. associated with the IOD can evolve even in the ab- Modifying the SST field to suppress ENSO (Behera sence of ITF (Vinayachandran et al., 2007), the et al., 2006) in the same model yielded a similar latter method could be considered as the preferred result. The IOD events in the GFDL CM2.1 cou- mode of suppressing ENSO in coupled models. pled model are preceded by anomalous conditions in the west Pacific warm pool, marked by an east- 5. TRIGGERING ward shift in convection, westerly winds and SST (Song et al., 2007) suggesting that variations in Evolution of the 1997–98 event together with an Indo-Pacific warm pool are crucial for generating El Nino propmted the hypothesis that ENSO can IOD conditions. From process experiments, they induce contrasting wind anomalies in the equato- concluded that IOD forms as a result of processes rial Indian Ocean which can mature into an IOD intrinsic to the Indian Ocean. They also noted that (Ueda and Matsumoto, 2001). This mechanism, IODs can also be caused by ENSO even though the however, is not capable of explaining all IODs, number of cases of the latter are relatively smaller. particularly those formed during non-ENSO years. The CSIRO Mark 3 coupled model produced 74 Signatures of IOD events begin to appear in spring ENSO and 34 IOD events out of 240 years of (March–April) and therefore, several studies have model simulation (Cai et al., 2005). Among these, explored a mechanism that could induce easterly there were only occasional coincidence of IOD and wind anomalies along the equator and southeast- ENSO whereas 23 IOD events occurred one year erlies off Sumatra. Spring rainfall in the eastern after ENSO. Unrealistic interactions between the equatorial Indian Ocean (EEIO) is simultaneously Indian and Pacific Oceans lead to development of correlated with the west-central Pacific SSTA and such IOD events (Cai et al., 2005) suggesting that unrelated to local SSTA (Annamalai et al., 2003). mechanism that leads to the formation of IOD like Based on this, Annamalai et al. (2003) have pro- events in coupled models need further investiga- posed that initiation of the IOD takes place in bo- tion. real spring which is time window through which In summary, both observations and models agree external forcing from the Pacific can influence the that it is not necessary to have forcing by ENSO Indian Ocean. In this mechanism, warm SSTA in in order to generate an IOD. That is, the In- west-central Pacific modifies the Walker circulation dian Ocean can support its own air-sea interac- and leads to subsidence over the EEIO and the tion process. A crucial role in the evolution of IOD grows by the Bjerknes feedback. However, the IOD is played by the interaction of the ther- their experiments with an AGCM showed rather mocline with the SST. The slope of the thermo- weak response of the EEIO rainfall to west-central Pacific SSTA and no attempt was made to demon- A key parameter involved in the coupled ocean strate that SST and thermocline depth anomalies atmosphere interaction culminating in the IOD is evolve into conditions favourable for the IOD. Ac- depth of the thermocline (Li et al., 2003). A shal- cording to Kajikawa et al. (2001), intensification low thermocline in the east is a necessary condi- of the Hadley cell over the western Pacific can en- tion for upwelling to work efficiently to cool the hance the southeasterly in the EEIO, SST cooling SST (Vinayachandran et al., 1999 2002 2007; Mur- and suppression of convection, triggering an IOD. tugudde et al., 2000). Therefore, preconditioning The intensification takes place during summer and of thermocline provides a favourable background, can be either due to ENSO or monsoon. However, and, probably, a necessary condition for the for- wind anomalies associated with the IODs appear mation of IOD (Annamalai et al., 2005). The pre- during spring (Vinayachandran et al., 1999 2007) conditioning of the Indian Ocean can be caused which the above mechanism fails to explain. by Pacific decadal variability which can influence the Indian Ocean either through atmosphere or Fischer et al. (2005) examined triggering mech- through the ITF (Annamalai and Murtugudde, anisms of IOD in the presence as well as absence 2004). The former is effected through zonal wind of El Nino using a coupled model. In the presence anomalies along the equator and an anticyclone to of ENSO, as the convection shifts eastward in the the east of Sumatra, and the latter by the influence Pacific during the development of ENSO, easterly of Indonesian throughflow (ITF) on the thermo- wind anomalies occur in the eastern Indian Ocean cline (Annamalai et al., 2005). Consequently, IOD and the thermocline becomes shallow. This is fol- events are stronger in decades during which Pacific lowed by development of upwelling favorable winds decadal variabiltiy favors a shallow thermocline in and reduced precipitation triggering the growth of the SETIO such as the 1960s and 1990s. an IOD. In the absence of ENSO, the trigger orig- Being a coupled ocean-atmospheric process, trig- inates from an anomaly in Hadley circulation dur- gering of an IOD can originate either in the ocean ing April–May featuring southerly cross equatorial or in atmosphere. All the studies mentioned above, winds and reduced rainfall in the SETIO, com- except Rao et al. (2009), point to a trigger from pletely independent of atmospheric and oceanic the atmosphere, either by a modification of the processes in the Pacific. Interestingly, these are Walker circulation or of the Hadley cell. This preceded by anomalies of SSTA in March–April is consistent with evolution of IOD events in the suggesting that the trigger originates in the ocean. last decade; observations suggest that first signs of In quite a deviation from all other theories, Fran- the IOD appears as anomalous easterlies along the cis et al. (2007) proposed that cyclones over the equator and anomalous southeasterlies off Sumatra Bay of Bengal during spring season can trigger IOD (e.g. the 2006 IOD event (Vinayachandran et al., events. Initiation of suppressed convection in the 2007)). Shallowness of the thermocline is crucial SETIO is caused by meridional pressure gradients for further development of the IOD as it is a crucial created by severe cyclones over the bay. The de- parameter for the Bjerknes feedback. There is a velopment of this anomalous convection into IOD near-perfect sea-saw relationship in convection be- is effected by the fact that suppression of convec- tween the eastern and westen boxes of IOD; when- tion in the east is associated with enhancement of ever convection is enhanced in the eastern side, it is suppressed in the west and vice versa (Gadgil convection in the west (Gadgil et al., 2003 2004), the former leading to southeasterlies off Sumatra and easterlies over EIO. IOD events are preceded by a downwelling Rossby wave that originates from the eastern boundary (Vinayachandran et al., 2002). These waves precondition the western equatorial Indian Ocean by deepening the theremocline. Such wave propagations during 1993 and 1996 provided a favourable background for the evolution of the IOD events during 1994 and 1997 respectively. Rossby waves can also excite a Kelvin wave which pre- conditions the thermocline in the eastern Indian Ocean as in the case of 2003 and 2006 (Rao et al., 2009). Intensification of the northward portion Figure 9. Correlation between DMI and precipitation of the ITCZ in March or a suppressed convection over the Asian monsoon region. DMI is computed using phase of an MJO can also restrict convection, trig- HadISST available at (http://badc.nerc.ac.uk) and pre- gering an IOD (Rao et al., 2009). cipitation used is CPC merged analysis of precipitation (CMAP) available at http://iridl.ldeo.columbia.edu. et al., 2003 2004). Thus, once the thermocline is IOD (Saha, 1970; Reverdin et al., 1986; Goddard preconditioned, enhancement of convection either and Graham, 1999; Nicholls, 1995; Nicholson and in the west or in the east can trigger a positive Kim, 1997). Many of these studies focussed on feedback. the anomalous convection and rainfall in the east African countries in 1961/62 and attributed it to the anomalously warm western equatorial Indian 6. INFLUENCE OF THE INDIAN OCEAN DIPOLE IN REGIONAL CLIMATE SYSTEM Ocean. In general, the influence of Indian Ocean SST tends to dominate in the rainfall variability 6.1. IOD and East African Rainfall of Africa in the warm phase of ENSO and the At- The variability of east African rainfall has pro- lantic Ocean controls the rainfall of Africa in the found impact on the livelihood of millions of peo- cold phase (Nicholson and Kim, 1997). Experi- ple in the developing countries in this region, who ments with atmospheric general circulation models mainly depend on the rain-fed agriculture and to quantify the relative influence of the Indian and regional fisheries. Relationhip between the SST Pacific Oceans suggest that even though the Pa- in the tropical Indian Ocean and the climate in cific SST anomalies influence the rainfall activity the surrounding regions has been explored by sev- in the African region, atmospheric response to the eral studies even prior to the discovery of the Indian Ocean variability is essential for the simulation of correct rainfall response over central, southern and eastern Africa, particularly the simulation of observed central-eastern/southern African dipole pattern in rainfall (Goddard and Graham, 1999). Interestingly, there are some studies, which suggest that the teleconnection between the east African rainfall and ENSO is a manifestation of a link between ENSO and the IOD (Black et al., 2003). After discovery of the IOD (Saji et al., 1999; Webster et al., 1999), several studies have inves- tigated the role of IOD on regulating regional climate. There is a tendency for increased rainfall in the tropical eastern Africa and drought in Indone- sia in IOD years (Saji et al., 1999). The correla- Figure 10. Rainfall over Africa in a) November 1961 tion between the DMI and precipitation over the and b) October 1997. Rainfall data used is global grid- Asian Monsoon regime confirms this (Fig. 9). The ded station rainfall data (Chen et al., 2002), available at extreme floods in East African countries in 1961– ftp://ftp.cpc.ncep.noaa.gov/precip/50yr/gauge/0.5deg. 62 and 1997–1998 (Fig. 10) were associated with anomalous conditions in the Indian Ocean (Saji 2 61 et al., 1999; Webster et al., 1999; Saji and Yam- Corln : 0.53 56 agata, 2003; Birkett et al., 1999). Extreme flood 83 75 88 conditions in the east-African region in 1997–1998 1 94 59 70 can not be attributed to an exaggerated response to the El Nino as the correlation between the east 97 African rainfall and the central Pacific SST is only 0 0.24 (Webster et al., 1999; M. et al., 1999). Fur- ther, a similar flood in 1961–62 did not co-occurr −1 85

ISMR Anomaly with an El Nino. Enhanced rain in the east African 4 68 74 82 8651 66 region is associated with increased moisture con- 65 79 87 vergence due to strong westerlies over the central −2 2 72 Africa and easterlies from the equatorial Indian −3 −2 −1 0 1 2 Ocean as a response of warm western pole of the Nino 3.4 index IOD (Ummenhofer et al., 2009). SST anomaly in western pole of the IOD have more impact on the east African short rains compared to that of the Figure 11. Scatter diagram of normalized ISMR east pole (Ummenhofer et al., 2009), these SST anomaly versus Nino 3.4 SST anomaly. Both anomalies anomalies result from the strong sub-surface in- normalized by their respective standard deviations. fluence due to propagating Rossby waves related ISMR data is obtained from http://tropmet.res.in to the IOD events (Rao and Behera, 2005). This and NINO 3.4 index is obtained from slowly propagating air-sea coupled mode could be http://www.cpc.noaa.gov/data/indices/sstoi.indices an indicator for the prediction of the IOD induced ISMR is weak as seen from (Fig. 12). This strong short rain events well in advance. Atmospheric positive relation between the IOD and ISMR dur- and coupled ocean-atmosphere general circulation ing 1960’s and 1990’s could be one of the rea- models have been able to simulate the observed as- sons for excess monsoon seasons in 1961, 1994 and sociation between the DMI and east African rain- near normal monsoon in 1997. However, neither fall reasonably well (Conway et al., 2006; Behera the ENSO nor the IOD can explain all the ex- et al., 2003). Hence coupled models also may be cess/droughts in the Indian summer monsoon. For used to predict East African short rains. example, drought years of 1985 and 1974 are not 6.2. IOD and Indian Summer Monsoon associated with negative IOD or El Nino. The Rainfall droughts during 1 1966, 1968, and 1979 are associ- A major advance in our understanding of the in- ated with very small SST anomalies in the central terannual variation of the Indian summer monsoon Pacific and no IOD signal in the Indian Ocean. rainfall (ISMR) occurred during the 1980s, with Hence, even though a few modelling studies with the discovery of a strong link with El Nino and AGCMs and coupled OAGCMs also suggest that Southern Oscillation (ENSO), the dominant signal of interannual variation of the coupled ocean- the positive IOD events can intensify the Indian atmosphere system over the Pacific (Sikka, 1980; summer monsoon rainfall (Ashok et al., 2001 2004; Pant and Parthasarathy, 1981; Rasmusson and Guan and Yamagata, 2003; Guan et al., 2003), we Carpenter, 1983). There is a high propensity of deficit monsoon during warm ENSO events, with deficit in 21 out of 25 such events during 1875– 1979, which included 9 of 11 seasons with largest anomalies (Rasmusson and Carpenter, 1983). It can be seen from Fig. 11 that this true in general. However, while majority of the droughts are associated with a positive Nino 3.4 index (El Nino condition), there are droughts of 1974 and 1985 which are associated with negative SST anomaly in the Nino 3.4 region. The NINO 3.4 index is very small in the drought years such as 1966, 1968 and 1979. Again, while in majority of the excess monsoon seasons, SST anomaly in the Nino 3.4 region is negative (La Nina condition), it is positive in 1983 and 1994. Further, in 1997, during the strongest El Nino year in the century, the ISMR was above nor- Figure 12. 41-month sliding correlations between mal. Consequently, Kumar et al. (1999) suggested NINO3 index and ISMR (dashed curve, multiplied by -1) that the relationship between the Indian monsoon and DMI and ISMR (solid curve). Correlation coefficient and ENSO had weakened in the recent decades. of 0.38 is significant at 90%. (after Ashok et al. (2001)) The correlation between DMI and ISMR is rather weak (Fig. 9). However, if we consider case by case, out of the 11 intense (anomalies in DMI more than one standard deviation) positive IOD events that occurred during 1958–1997, eight events (1961, 1963, 1967, 1977, 1983, 1994, 1993 and 1997; ie., 73% of the positive IOD events during this period) are associated with positive anomalies of the concurrent ISMR (Ashok et al., 2001). Similarly, out of the three negative IOD events during this 1958–1997, two events (1960 and 1992) correspond to negative anomalies of the ISMR. Interestingly, the sliding correlation between the ISMR and ENSO has an opposite tendency to that between ISMR and IOD; ie., when the ISMR-ENSO (negative) correlation is strong, ISMR-IOD (positive) correlation is weak. Simi- Figure 13. NOAA AVHRR OLR (WM−2 anomaly larly, when the ISMR has a strong positive corre- patterns for a) July 2002, b) July 2003, c) August lation with IOD, the relation between ENSO and 1986 and d) August 1994. Data is obtained from http://www.cdc.noaa.gov still need to understand the role of IOD in the in- convection) over the EEIO. The pattern was just terannual variation of the ISMR. opposite in 2003. These are not isolated events. In fact, the OLR anomaly patterns in the monsoon 6.3. ISMR and Equatorial Indian Ocean months of 1986 and 1994 are quite similar to that Oscillation The extreme drought in 2002 and the very in 2002 and 2003 respectively (Fig. 13c&d). Gadgil fact that none of the predictions (irrespective of et al. (2004) termed this oscillation in the convec- whether it is using complex GCMs or empirical tion over the WEIO and EEIO as the equatorial models) could anticipate this drought, triggered Indian Ocean Oscillation (EQUINOO). They sug- debates among researchers around the world on gested that the winds along the equatorial Indian the factors influencing the ISMR variability. This Ocean could be considered as measure of the inten- drought was neither associated with a very strong sity of the EQUINOO as it responds to the pres- El Nino or a negative IOD event (Gadgil et al., sure gradient associated with differential heating in 2002). Subsequent to this monsoon failure, in 2003, the atmosphere due to the difference in convection. the ISMR was very close to its long term mean. EQUINOO index (EQWIN) is defined as the neg- Interestingly, the July mean outgoing longwave ra- ative of zonal wind anomaly over central equato- o o o diation (OLR) (which may be used as a proxy for rial Indian Ocean (averaged over 60 E–90 E,2.5 S– o tropical convection) pattern over the Indian region 2.5 N). The IOD is a coupled phenomenon and and the equatorial Indian Ocean in 2003 was ex- its signatures are seen in convection and surface actly opposite to that in 2002 (Fig. 13a&b). In wind over the equatorial Indian Ocean also (Saji 2002, there was a high OLR anomaly over the et al., 1999). Interestengily, there are occasions on western part of the Indian region and the west- which such convection and surface wind anomalies ern equatorial Indian Ocean (WEIO). At the same are present without an SST dipole signal(Francis, time OLR anomaly was negative (indicating more 2006). Hence, even though all the IOD events are associated with EQUINOO, the reverse is not nec- essarily true. 2 a) b) 2 When the relative influence of EQUINOO and 1 1 IOD on ISMR is considered, the ISMR is better

0 0 correlated to the EQUINOO than IOD (Gadgil ISMR ISMR et al., 2004). The relation of ISMR with EQWIN −1 −1 and DMI is shown in Fig. 14a&b. The DMI- −2 Corln. : 0.45 Corln. : 0.05 −2 ISMR correlation is only 0.05. On the other hand, −2 −1 0 1 2 3 −2 −1 0 1 2 3 EQWIN DMI EQWIN-ISMR correlation is signiﬁcantly high

Figure 14. Scatter plot of ISMR versus a) EQWIN and JUN−SEP b) DMI. EQWIN is computed from the monthly mean 2 88 surface wind data from NCEP Renalysis, available at 75 http://www.cdc.noaa.gov. 70 1

85 74 59

61 0 83 79 66 68 86 94 4

ENSO INDEX (JJAS) −1

2 72 65 82 −2

−3 −2 −1 0 1 2 3 EQWIN (JJAS)

Figure 16. Extremes of the ISMR is represented in the phase-plane of EQWIN and ENSO index (both indices are normalized by their respective standard deviations) in the period 1958–2008. Color code: dark red, red, Figure 15. Correlation between ISMR and OLR over blue and dark blue closed circles represent ISMR anomaly the Indian monsoon region. Since lower OLR values cor- larger than -1.5 standard deviation, between -1.5 and -1 respond to deeper convection, negative correlation im- standard deviation, between 1 and 1.5 standard devia- plies enhanced (suppressed) convection associated with tion and larger than 1.5 standard deviation respectively. above normal (below normal) ISMR. Adapted from Gadgil et al. (2004) (0.45). While the ISMR is negatively correlated to vere droughts in Australia. The negative corre- the OLR over the WEIO, it is positively correlated lation between the Australian rainfall and Indian to OLR over the EEIO (Fig. 15). Hence, an excess Ocean variability has been successfully simulated (drought) monsoon is associated with enhanced by AGCMs (Ashok et al., 2003; Ummenhofer et al., (suppressed) convection over the WEIO and sup- 2008). Cold SST anomalies prevailing west of the pressed (enhanced) convection over the EEIO. In Indonesian archipelago during the IOD events in- other words, a positive EQUINOO is favourable troduce an anomalous anticyclonic circulation at for Indian monsoon and a negative EQUINOO lower levels over Australia, which in turn reduce highly unfavourable. All the excess/drought mon- the rainfall activity over this region (Ashok et al., soon seasons in the period 1958–2003 are associ- 2003). ated with the favourable/unfavourable phase of either ENSO or EQUINOO or both (Gadgil et al., 6.5. IOD and Intraseasonal Oscillations 2004). There is a clear separation between excess An active MJO is characterized by slow eastward and drought seasons when represented in the phase movement of stronger than average precipitation plane of the EQWIN and ENSO indices (which and westerly winds (Madden and Julian, 1972). In is the negative of normalized SST anomaly in the summer monsoon months (June-September), rain- NINO3.4 region) as shown in Fig. 16. The linear fall and wind anomalies also propagate northward reconstruction of ISMR on the basis of multiple on intraseasonal time scale over India, southeast regression from NINO3 index and the zonal wind and east Asia and the adjacent oceans (Sikka and over the equatorial Indian Ocean better specifies Gadgil, 1980; Yasunari, 1980). Both these propa- the ISMR than the regression with only NINO3 gations are strongly modulated by the IOD (Shin- index (Ihara et al., 2007). Clearly, monitoring and oda and Han, 2005; Ajayamohan et al., 2008 2009). predicting the EQUINOO and IOD is essential for During the active phase of IOD no significant the prediction of the ISMR. Even though the asso- intraseasonal oscillations occur in Indian Ocean ciation between the EQUINOO and the extremes equatorial zonal winds (Rao and Yamagata, 2004) of ISMR is now established, the mechanism by and 6–90 day oscillations in winds and OLR tend which the EQUINOO influences the Indian mon- to weaken (Shinoda and Han, 2005). During nga- soon is yet to be understood. One explanation for tive dipole years, submonthly (6–30 day) convec- the increased rainfall in activity over the Indian tion generated in the southeast Indian Ocean is region during the positive IOD event is through associated with a cyclonic circulation. These sub- the enhanced cross-equatorial flow to the Bay of monthly disturbances propagate southwestward, Bengal due to the strong meridional SST gradient which generate equatorial westerly winds in the across the EEIO (Guan et al., 2003). However, eastern and central Indian Ocean and northwester- during most of IOD events, the largest precipita- lies near the coast of Sumatra. During large posi- tion and convection anomalies are observed in the tive dipole years, submonthly convective activity is western part of the country (Fig. 13). More fo- largely suppressed in the southeast Indian Ocean cussed research is required to to address this prob- and thus no equatorial westerly is generated.The lem. poleward propagation of intraseasonal oscillations during summer monsoon is coherent (incoherent) 6.4. IOD and Australian Rainfall from 5S to 25N during negative (positive) IOD The years of extensive Australian drought were years (Fig. 17). The mean structure of moisture associated with generally low SST in the low lati- convergence and meridional specific humidity dis- tudes of eastern Indian Ocean (Streten, 1981 1983). tribution undergoes radical changes in contrasting The rainfall over much of Australia tends to be IOD years, which in turn influences the meridional below average during El Nino events (Ropelewski propagation of boreal summer intraseasonal oscil- and S., 1987). However, in some El Nino events lations (Ajayamohan and Rao, 2008; Ajayamohan (such as 1986/87), Australian rainfall anomalies et al., 2009). are weak or of limited extent while in others (such as 1982/83), the anomalies are major and 6.6. IOD and other global teleconnections widespread (Nicholls, 1989). The first pattern The influence of IOD on other global climate sys- of the principal component analysis of Australian tems is far less explored. One major difficulty is winter rainfall shows a broad-band stretching from that many of the positive IOD events in the recent the northwest to southeast corners. This is best re- years are associated with strong El Nino events lated to the difference in SST between the Indone- and it is difficult to separate the influence of IOD sian region and the central Indian Ocean (Nicholls, from El Nino. Nevertheless, when the influence 1989), which is infact very strong during the IOD of ENSO is removed, in the tropical regions sur- years. This suggests that the El Nino events that rounding the Indian Ocean, there is a clear pat- co-occur with IOD are generally associated se- tern of warm temperature anomalies over land re-

Figure 17. (a) Regressed filtered anomalies of CMAP precipitation (mm day1) averaged over 7095E as a function of latitude and time lag during the 19802004 period. As in (a), but for (b) negative and (c) positive IOD years. Contour interval is 0.6. Only statistically significant (0.1 significance level using a t test) anomalies are plotted (taken from Ajayamohan et al. (2008)) gions to the west and cool anomalies over regions ter, 1983). Thus, the influence of IOD and ENSO to the east of the Indian Ocean (Saji and Yama- on the Sri Lankan rainfall is inter-linked. How- gata, 2003). Over the extratropics, IOD is asso- ever, the influence of IOD is highly significant even ciated with warm temperature anomalies, reduced when the ENSO influence is statistically removed rainfall and positive geopotential height anomalies. (Lareef et al., 2003). Some of the recent studies The interannual variation of rainfall over central suggest that the evolution of the ENSO itself is in- Brazil and subtropical La Plata Basin during aus- fluenced by the SST variations in the Indian Ocean tral spring are found to be influenced by the IOD, (Yu et al., 2002), in particular the IOD events (Be- with a positive IOD associated with increased rain- hera and Yamagata, 2003; Annamalai et al., 2005). fall over the subtropical La Plata Basin, while it Thus, the influence of the IOD events are not re- is associated with decreased rainfall over central stricted on the regional climate. The mechanisms Brazil (Chan et al., 2008). Another impact of the of the global teleconnections of the IOD events are IOD is on the ’Maha’ rainfall in Sri Lanka dur- yet to be understood clearly. ing September to December. Warm SST anomalies in the WEIO associated with the IOD event lead 7. CONCLUDING REMARKS to enhanced convergence in the lower troposphere over the WEIO, which in fact, extends up to Sri ENSO, which is the dominant signal of tropi- Lanka (Lareef et al., 2003). This leads to increased cal interannual variability has significant influence rainfall activity over Sri Lanka. In general, El Nino on the Indian Ocean. Research in the last decade events also lead to increase in the ’Maha’ rainfall clearly demonstrate that the IOD is a major signal (Ropelewski and S., 1987; Rasmusson and Carpen- of interannual climate variations in the region. The question whether IOD is independent of ENSO or role in taking IOD events to maturity through its not has been debated long and it turns out that the control over atmospheric convection. Indian Ocean can breed and nurture its own IOD Impact of IOD on the climate of Asia, Africa by coupled air-sea processes, without being forced and Australia is significant and, therefore, it is ab- externally by the Pacific Ocean. Both analysis of solutely necessary that the scientific community be observations and model simulations converge to the able to forecast its evolution well in advance. Con- fact that IOD can be generated either by ENSO sidering that this is a fully coupled phenomenon, a or otherwise. There is a general consensus that coupled model is a basic requirement. The strong the modification of Walker cell during ENSO is as- IOD event of 2006 (Vinayachandran et al., 2007) sociated with subsidence over SETIO which sup- was successfully predicted by a coupled model but presses convection and can trigger a positive feed it had limited success in forecasting the evolution of back leading to the development of an IOD. Mech- the weak IOD event of 2007 (Luo et al., 2008). Ex- tremes of the Indian Summer monsoon are found anisms that cause IOD formation in the absence to coincide with the positive and negative IOD of ENSO is still unclear, although there are indi- (Gadgil et al., 2003 2004) and therefore the urgent cations that anomalies in Hadley circulation holds need for the forecast arises from the Indian sub- the key. Observations suggest that subsurface tem- continent. Result of forecasting experiments (Luo perature anomalies, accompanied by strong west- et al., 2008) are encouraging and the usefulness of ward current anomalies appear about 3 months such a system for Indian Summer monsoon as well before the SSTA (Horii et al., 2008; Nagura and as other regions where IOD affects rainfall is yet to McPhaden, 2008). Whether the preconditioning be tested and put into practice. of the ocean, with a lifted thermocline in the east, is a necessary condition is yet to be established. ACKNOWLEDGMENTS. This study was funded by the Indian Ocean Modeling Program (INDOMOD), Indian This aspect is of particular relevance to the recent National Centre for Ocean Information Services(INCOIS), intensification of IOD activity and monsoon IOD Ministry of Earth Sciences, Govt. of India. feedback (Abram et al., 2008). Following the availability of high quality sea surface wind data, ocean models have been able to REFERENCES reproduce the IOD events remarkably well (Mur- tugudde et al., 2000; Vinayachandran et al., 2007). Abram N J, Gagan M K, McCulloch M T, Wahyoe J C, and Coupled-models, on the other hand. have sev- Hantoro S 2003 Coral Reef Death during the 1997 Indian Ocean Dipole linked to Indonesian Wildfires; Science 301 eral limitations. 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