Seasonal-to-Interannual Time-Scale Dynamics of the EquatorialUndercurrent in the Indian Ocean*

GENGXIN CHEN

State Key Laboratory of Tropical Oceanography, South China Sea Institute of Oceanology, Chinese Academy

of Sciences, Guangzhou, China

WEIQING HAN AND YUANLONG LI

Department of Atmospheric and Oceanic Sciences, University of Colorado, Boulder, Boulder, Colorado

DONGXIAO WANG

State Key Laboratory of Tropical Oceanography, South China Sea Institute of Oceanology, Chinese Academy

of Sciences, Guangzhou, China

MICHAEL J. MCPHADEN

Pacific Marine Environmental Laboratory, National Oceanic and Atmospheric Administration, Seattle, Washington

(Manuscript received 1 November 2014, in final form 24 February 2015)

ABSTRACT

This paper investigates the structure and dynamics of the Equatorial Undercurrent (EUC) of the Indian

Ocean by analyzing in situ observations and reanalysis data and performing ocean model experiments using

an ocean general circulationmodel and a linear continuously stratified oceanmodel. The results show that the

EUC regularly occurs in each boreal winter and spring, particularly during February and April, consistent

with existing studies. The EUC generally has a core depth near the 208C isotherm and can be present across

the equatorial basin. The EUC reappears during summer–fall of most years, with core depth located at dif-

ferent longitudes and depths. In the western basin, the EUC results primarily from equatorial Kelvin and

Rossbywaves directly forced by equatorial easterly winds. In the central and eastern basin, however, reflected

Rossby waves from the eastern boundary play a crucial role. While the first two baroclinic modes make the

largest contribution, intermediate modes 3–8 are also important. The summer–fall EUC tends to occur in the

western basin but exhibits obvious interannual variability in the eastern basin. During positive Indian Ocean

dipole (IOD) years, the eastern basin EUC results largely from Rossby waves reflected from the eastern

boundary, with directly forced Kelvin and Rossby waves also having significant contributions. However, the

eastern basin EUC disappears during negative IOD and normal years because westerly wind anomalies

force a westward pressure gradient force and thus westward subsurface current, which cancels the eastward

subsurface flow induced by eastern boundary–reflected Rossby waves. Interannual variability of zonal

equatorial wind that drives the EUC variability is dominated by the zonal sea surface temperature (SST)

gradients associated with IOD and is much less influenced by equatorial wind associatedwith Indianmonsoon

rainfall strength.

1. Introduction

In the Pacific and Atlantic Oceans, the Equatorial

Undercurrent (EUC) is quasi permanent in nature (e.g.,

McPhaden 1986;Metcalf and Stalcup 1967; Izumo 2005);

it flows eastward near the top of the thermocline under a

typically westward-flowing surface current, driven by

the prevailing easterly trade winds. By contrast, in the

* Pacific Marine Environmental Laboratory Contribution

Number 4292.

Corresponding author address: Dongxiao Wang, State Key

Laboratory of Tropical Oceanography, South China Sea Institute

of Oceanology, Chinese Academy of Sciences, 164 West Xingang

Road, Guangzhou, China.

E-mail: dxwang@scsio.ac.cn

1532 JOURNAL OF PHYS ICAL OCEANOGRAPHY VOLUME 45

DOI: 10.1175/JPO-D-14-0225.1

� 2015 American Meteorological Society

Indian Ocean the EUC is a transient feature. It is asso-

ciated with equatorial waves driven by the strong, sea-

sonally varying component of surface wind (Fig. 1;

Schott and McCreary 2001). Early observations sug-

gested that the EUCwas consistently detected from year

to year only during boreal winter and boreal spring,

typically from February to June. Based on current meter

measurements, Knauss and Taft (1964) observed an

EUC, with maximum speeds increasing from 0.27 to

0.81m s21 from the west to the east during March and

April 1963, but they did not detect an eastward EUC in

August 1962. Swallow (1964, 1967) reported a strong

EUC from March to June 1964 near 588 and 67.38E,with a maximum speed exceeding 1.20m s21, a mag-

nitude comparable to the EUC of the Pacific Ocean.

Indeed, the existence of an eastward EUC during

February–June has been further confirmed by later ob-

servations at various locations along the equator be-

tween 538 and 918E (e.g., Taft 1967; Knox 1974, 1976;

Rao and Jayaraman 1968; Schott et al. 1997; Reppin

et al. 1999; Sengupta et al. 2007). Meanders of the EUC

with its core being displaced from 38S to 28N along

section 55.58E was also reported for February–June of

1975 and 1976 (Leetmaa and Stommel 1980). In this

paper, seasons refer to those of the Northern Hemi-

sphere. The EUC is defined as an eastward zonal flow in

the equatorial Indian Ocean with its core located in the

thermocline above 300-m depth and beneath a westward-

or weaker eastward-flowing surface current, and it lasts

for at least 1 month.

While the Indian Ocean EUC regularly occurs only

during winter and spring based on earlier observations at

different locations, reappearance of the EUC in late

summer and early fall has also been detected for some

years. Bruce (1973) observed an EUC in the western

Indian Ocean during late August of 1964, with a core

depth of near 75m. Reppin et al. (1999) reported a re-

appearance of EUC in August 1994 at 80.58E with a core

near 150m, following the winter–spring EUC observed

from February toMay. The magnitude of the reappeared

EUC exceeds 0.40ms21, and it flows under a westward

surface current (Reppin et al. 1999). Note that 1994 is a

year of positive Indian Ocean dipole (IOD) event (e.g.,

Saji et al. 1999), as is 2006 when the presence of EUC in

August is also observed (Iskandar et al. 2009). Iskandar

et al. (2009) analyzed 6yr (December 2000–November

2006) of mooring data of the eastern equatorial Indian

Ocean at 08, 908E and showed that the subsurface zonal

current between 90- and 170-m depths generally exhibits

apparent semiannual variability, flowing eastward during

winter and summer except for 2003, when the summer-

time eastward current vanishes. They suggested that up-

welling Kelvin waves driven by wintertime easterly winds

is the primary cause for the eastward subsurface flow in

boreal winter–spring. During summer, they suggested

that both wind-driven Kelvin waves and eastern

boundary–reflected Rossby waves contribute.

Numerical ocean models are able to simulate the ob-

served EUC over the Indian Ocean. Using an ocean

general circulation model (OGCM), Anderson and

FIG. 1. The monthly climatology of CCMP surface wind stress (arrows; Nm22) and AVISO sea surface height (cm) in January, March,

May, July, September, and November averaged for 2001–11. Red (black) arrows show wind stress with easterly (westerly) wind com-

ponents. The scale in upper-left corner represents 0.2Nm22.

JUNE 2015 CHEN ET AL . 1533

Carrington (1993) reproduced the winter–spring EUC

and the strong semiannual variability of surface and

subsurface currents. Han et al. (2004) successfully sim-

ulated the observed winter–spring EUC and late sum-

mer reappearance in 1994. They pointed out that

easterly wind anomalies associated with the positive

IOD events (referred to as Indian Ocean zonal mode in

Han et al. 2004) during 1994 and 1997 drive a westward

surface current and an eastward EUC from August to

December, a situation that resembles the EUC in the

Pacific and Atlantic. They further argued that during

positive IOD years, the EUC forms in August and lasts

throughDecember. The EUC shoals in the eastern basin

because of a shallower thermocline, which is associated

with upwelling in the eastern equatorial Indian Ocean

during IOD events. Swapna and Krishnan (2008) and

Krishnan and Swapna (2009) examined the effects of

IOD and Indian monsoon winds on setting up the

summertime (June–September) EUC. They concluded

that intensification of summer monsoon winds during

positive IOD events produces nonlinear amplification of

easterly wind stress anomalies to the south of the equator

and enhances upwelling in the eastern Indian Ocean off

Sumatra–Java, so that the thermocline shoaling provides

a zonal pressure gradient and drives anomalous EUC in

the subsurface. Recently, Nyadjro and McPhaden (2014)

demonstrated how surface wind stress, zonal pressure

gradient, and zonal flow in the thermocline vary consis-

tently with the phase of the IOD, using 52yr (1960–2011)

of the ECMWF Ocean Reanalysis System, version 4

(ORAS4), ocean reanalysis product.

Even though progress has beenmade in observing and

understanding the Indian Ocean EUC, its spatial struc-

ture, seasonal-to-interannual variability, and the funda-

mental dynamics that govern the EUC and its variability

remain unclear and have not yet been systematically in-

vestigated. Equatorial waves have been suggested to

be involved in the EUC dynamics (e.g., Schott and

McCreary 2001; Zhang et al. 2014), and they appear to

be important in generating semiannual variability of

subsurface zonal flow (Iskandar et al. 2009) as well as

interannual variability of surface zonal current associated

with IOD events (Nagura and McPhaden 2010b; Zhang

et al. 2014; Nyadjro and McPhaden 2014). However,

quantitative assessments of the roles played by the

equatorial waves in determining the EUC and its vari-

ability have not yet been done. The goal of this study is to

provide a systematic investigation on the structure and

dynamics of the Indian Ocean EUC, including its sea-

sonal and interannual variability. Our approach is to

combine data analyses from moorings with modeling

experiments using a hierarchy of ocean models. The rest

of the paper is organized as follows: Section 2 describes

data and models. Section 3 discusses the structure and

variability of the observed EUC. Section 4 investigates

the dynamics of the EUC, and section 5 provides a sum-

mary and discussion.

2. Data and ocean models

a. Ocean data

The current measurements used in this study are

obtained from two moorings of the Research Moored

Array for African–Asian–AustralianMonsoon Analysis

and Prediction (RAMA; see McPhaden et al. 2009) lo-

cated at the Indian Ocean equator. One mooring is de-

ployed at 08, 908E and provides data from 14 November

2000 to 7 June 2012. The other is deployed at 08, 80.58Eand provides data from 27 October 2004 to 17 August

2012. The two moorings cover depth ranges from 40 to

410m and 25 to 350m with every 10- and 5-m interval,

respectively. These resolutions are reasonable for re-

solving the vertical structure of the EUC. Daily data

from these moorings are used for our analysis.

In addition, monthly currents from ORAS4

(Balmaseda et al. 2013) from 1958 to 2011 are used to

examine temporal variability and spatial structure of the

EUC.ORAS4 assimilates temperature and salinity data,

spanning the period 1958 to the present with a horizontal

resolution of 18 3 18 and 42 vertical levels. The National

Oceanic and Atmospheric Administration (NOAA)

weekly sea surface temperature (SST) data (Reynolds

et al. 2002) from 1982 are used to calculate the dipole

mode index (DMI; http://stateoftheocean.osmc.noaa.

gov/sur/ind/dmi.php), which is defined as the SST

anomaly difference between the western node (108S–108N, 508–708E) and eastern node (108S–08, 908–1108E)regions (Saji et al. 1999). Homogeneous Indian monthly

rainfall datasets (1871–2012) from the Indian Institute

of Tropical Meteorology (http://www.tropmet.res.in/

Data%20Archival-51-Page) are used to produce the

monsoon index. The data diagnostics also include sea

surface height averaged for the 2001–11 period from the

Archiving, Validation, and Interpretation of Satellite

Oceanographic Data (AVISO) project (http://www.

aviso.altimetry.fr/en/data/products/sea-surface-height-

products.html), surface winds for the period 1948–2012

from the National Centers for Environmental Prediction

(NCEP) reanalysis (Kistler et al. 2001), and the cross-

calibrated multiplatform (CCMP) satellite ocean surface

wind vectors available during July 1987–December 2011

(Atlas et al. 2008).

b. The OGCM and experiments

TheOGCMused in this study is a recent version of the

Hybrid Coordinate OceanModel (HYCOM; e.g., Bleck

1534 JOURNAL OF PHYS ICAL OCEANOGRAPHY VOLUME 45

2002; Wallcraft et al. 2009) configured to the Indian

Ocean basin (508S–308N, 308–122.58E) with a horizontal

resolution of 0.258 3 0.258 and 26 vertical layers. The

surface forcing fields include the 0.258 3 0.258 CCMP

ocean surface winds available from July 1987 to

December 2011; precipitation from the 0.258 3 0.258Tropical Rainfall Measuring Mission (TRMM) Multi-

satellite Precipitation Analysis (TMPA) level 3B42 prod-

uct (Kummerow et al. 1998) available for 1998–2011; 2-m

air temperature and humidity from the ECMWF interim

reanalysis (ERA-Interim) products (Dee et al. 2011)

available since 1979; and surface shortwave and long-

wave radiation from Clouds and the Earth’s Radiant

Energy System (CERES;Wielicki et al. 1996; Loeb et al.

2001) available since 2000. Surface latent and sensible

heat fluxes are calculated from wind speed, air temper-

ature, specific humidity and model sea surface temper-

ature using the Coupled Ocean–Atmosphere Response

Experiment, version 3 (COARE 3.0), algorithm (Kara

et al. 2005). Further details about the model and forcing

fields can be found in Li et al. (2014). Earlier versions of

HYCOM have been successfully used to understand

wave dynamics in the equatorial Indian (e.g., Han et al.

2011) and Atlantic Oceans (e.g., Han et al. 2008).

Taking theWorld Ocean Atlas 2009 (WOA09) annual

climatology of temperature and salinity as initial con-

ditions, the model is spun up from a state of rest for

30 yr. Restarting from the spun-up solution, HYCOM is

integrated forward in time from 1 March 2000 to

30 November 2011 with the daily forcing fields described

above. To exclude the transient effect of the model

simulation in 2000, we analyze the 3-day-averaged model

outputs from 2001 to 2011.

c. The linear ocean model and experiments

To help elucidate the wind-driven equatorial wave

dynamics, a continuously stratified linear ocean model

(LOM) is also used. The model is described in detail in

McCreary (1980, 1981) and is applied to several Indian

Ocean studies (e.g., Shankar et al. 1996; McCreary et al.

1996; Yuan and Han 2006). Although nonlinearity is

essential for understanding the mean equatorial zonal

flow near the surface, it is relatively weak in the thermo-

cline on seasonal time scales (Nagura and McPhaden

2014). Consequently, linear dynamics can be a good ap-

proximation for seasonal and interannual variability of

the transient EUC in the Indian Ocean. Here, we only

introduce aspects that are essential for our discussion. The

equations of motion are linearized about a background

state of rest with stratification represented by Brünt–Väisälä frequency, and the ocean bottom is assumed

flat at 4000m. Under these restrictions, solutions for

the zonal velocity u, meridional velocity y, and pressure

p can be expanded in the vertical normal modes with

eigenfunctions cn (z):

u5 �N

n50

uncn , (1a)

y5 �N

n50

yncn, and (1b)

p5 �N

n50

pncn , (1c)

whereN is the total mode number. The terms un, yn, and

pn are expansion coefficients that satisfy the following

equations:

�›t 1

A

c2n

�un 2 f yn 1

1

rpnx5 txZn/(rHn)1 y2=

2un ,

(2a)�›t 1

A

c2n

�yn 1 fun1

1

rpny5 tyZn/(rHn)1 y2=

2yn, and

(2b)�›t 1

A

c2n

�pnrc2n

1 unx 1 yny 5 0. (2c)

In the above, cn represents the equatorial Kelvin wave

speed for vertical mode number n. The cn values for the

first eight baroclinic modes (n5 1, 2, . . . , 8) are 264, 167,

105, 75, 60, 49, 42, and 37m s21 with observed stratifi-

cation within the Indian Ocean (e.g., McCreary et al.

1996; Han et al. 2004). The factors Zn 5Ð 02D Z(z)cn dz

and Hn 5Ð 02D c2

n dz determine how strongly the driving

wind couples to each mode; f 5by is the Coriolis pa-

rameter under equatorial beta-plane approximation;

and Z(z) is the vertical profile of wind that is introduced

as a body force, whereZ(z) is constant in the upper 50m

and linearly decreases to zero from 50- to 100-m depth.

The terms associated with A/c2n represent vertical fric-

tion, with A 5 0.000 13 cm2 s23. Density r 5 1 g cm23

is a typical density value of seawater. More details can

be found in McCreary et al. (1996).

The LOM is configured for the tropical Indian Ocean

north of 298S with a horizontal resolution of 25 km 325 km, which is close to the 0.258 3 0.258HYCOM grids.

The total solution is the sum of the first 25 modes [N 525 in Eqs. (1a)–(1c)], and the solutions are well con-

verged with this choice. The linear model is first spun up

for 20 yr and then integrated forward in time for the

period of 1988–2011 using monthly CCMP wind stress

forcing. This solution is referred to as LOM main run

(LOM_MR). To isolate the effects of eastern boundary–

reflected Rossby waves, a damper in the eastern equa-

torial ocean is applied [see McCreary et al. (1996) for

JUNE 2015 CHEN ET AL . 1535

detailed descriptions of the damper]. The damper effi-

ciently absorbs the energy of incoming equatorial Kelvin

waves, and thus no Rossby waves are reflected back

into the ocean interior from the eastern boundary. We

refer to this solution as LOM_DAMP. The difference

(LOM_MR 2 LOM_DAMP) isolates the reflected

Rossby wave effects.

Note that the damper is not effective at the western

boundary even for a straight north–south meridional

wall. Therefore, we will not assess the effect of western

boundary reflection, which is more complex than the

eastern boundary because eastward-propagating short

Rossby waves are also involved in addition to the Kelvin

wave, and strong mixing near the western boundary can

quickly damp the short Rossby waves. This is consistent

with Le Blanc and Boulanger (2001), who suggested that

at the western boundary of the equatorial Indian Ocean,

reflection of the equatorial Rossby wave into the Kelvin

wave is subject to large seasonal and interannual vari-

ability. The reflection efficiency at the western boundary

is much lower than that of the eastern boundary, where

reflection efficiency is 85% in terms of amplitude for

equatorial Kelvin waves reflecting into first meridional

mode Rossby waves. Nagura and McPhaden (2010a,b)

studied the equatorial wave effect on seasonal to in-

terannual variability of the Wyrtki jet by assuming re-

flection efficiencies of 85% at both eastern and western

boundaries. However, as Nagura and McPhaden (2014b)

showed, velocity variability in the interior of the Indian

Ocean is relatively insensitive to western boundary

reflections. Thus, LOM_DAMP primarily measures

the directly forced Kelvin and Rossby waves, with

western boundary reflection playing a minor role.

3. Observed EUC: Temporal variability and spatialstructure

To emphasize the seasonal and interannual variability

of EUC, we first apply a 31-day running mean to the

daily RAMA data and calculate monthly mean values

from the 3-day HYCOM output. Monthly climatology

of zonal currents is obtained for the 2001–11 period,

using the values of each year (rather than just the strong

EUC years) to calculate the climatological mean. Then

the seasonal variability of EUC is identified from the

monthly climatology, and interannual variability is

identified using the monthly data of each year.

Consistent with previous studies, the 31-day running-

mean zonal currents from the mooring deployed at 08,908E for the 2001–12 period show that the EUC in the

Indian Ocean is a transient phenomenon; it consistently

appears for each year during boreal winter–spring, par-

ticularly from February to April (Fig. 2), based on its

definition given in section 1. The depth of the winter–

spring EUC core is near the depth of 208C isotherm (D20)

within the thermocline, with the exception of 2007 when

the EUC core is located near D23, which is in the upper

thermocline (Fig. 2), where the vertical temperature gra-

dient often reaches the maximum amplitude (not shown).

The maximum speed of the EUC during February–April

can reach 0.75ms21. From April to May, the eastward

spring Wyrtki jet (Wyrtki 1973) appears, and the EUC

weakens and cannot be identified inMay. Consistent with

previous observational studies (e.g., McPhaden 1982;

Reppin et al. 1999), there is a clear upward phase

propagation, which suggests downward propagation of

equatorial wave energy.

In agreement with Iskandar et al. (2009), subsurface

zonal flow between 90 and 170m in the eastern equatorial

Indian Ocean exhibits significant semiannual variability,

flowing eastward again from late summer to early fall in

most years (Fig. 2). However, not all eastward-flowing

subsurface currents can be defined as EUC; only the ones

with subsurface maxima flowing under westward or

weakly eastward surface currents are defined as EUC.

From August–October, the EUC reappears in several

years during the 2001–11 period, but it is completely ab-

sent in 2003, 2008, and 2010 at this mooring location

(Fig. 2). At the mooring location of the central equatorial

Indian Ocean (08, 80.58E), the observed subsurface cur-

rents from 2004 to 2012 also show significant semiannual

variability (Fig. 2). There are, however, apparent differ-

ences between the two locations. For example, an EUC

with its core velocity of 0.58ms21 is observed in July

(with missing data in August) 2008 at 80.58E, but it isabsent at 908E. The obvious discrepancy between the twomoorings suggests that theEUC—particularly its summer

and fall reappearance—is a less robust feature of the

circulation than the boreal winter–spring EUC.

To reveal the spatial structure of the EUC, we ex-

amine the time evolution of the zonal current from

ORAS4 data, since the long current time series from

RAMA moorings are only available at 908 and 80.58E.ORAS4 data agree reasonably well with the RAMA

observations, even though the EUC magnitudes in

ORAS4 currents are systematically weaker than the

mooring observations (Fig. 3) as also found by Nyadjro

and McPhaden (2014), likely because of the 18 3 18 areamean for ORAS4 data compared to RAMA data for a

specific location. The 2001–11 mean monthly climatol-

ogy of the ORAS4 zonal current (Fig. 4) shows that the

winter–spring EUC can exist across the equatorial basin

and obtains its maximum amplitude (;0.38m s21) in the

central and western basins from March to April. The

EUC reappears in August–September with a much

weaker magnitude (,0.21ms21). By analyzing vertical

1536 JOURNAL OF PHYS ICAL OCEANOGRAPHY VOLUME 45

sections of zonal velocity along the equator for in-

dividual years (figure not shown), we find that the EUC

exists from summer to early fall of all years during 2001–

11 except for 2010 (a negative IODyear), when the EUC

disappears across the basin (see section 4c). While the

winter–spring EUC generally has a basin-scale structure

(section 4b), the summer–fall EUC has large in-

terannual variability regarding its location and depth,

which explains its weak amplitude in the 2001–11 11-yr

average. Indeed, for all the 54-yr ORAS4 products

from 1958 to 2011, the EUC essentially occurs for each

spring, and it reappears in the summer–early fall of most

FIG. 2. (top three) The 31-day running-mean daily zonal current observed by the RAMA

mooring at (08, 908E) from2001 to 2012.White and thin black lines represent the 0 and 0.2m s21

contours, respectively. The red (thick black) lines represent D23 and D20 obtained from

RAMA observations (ORAS4 data). Gray contours show temperature from the RAMA

mooring. RAMA and ORAS4 D23 and D20 agree very well. (bottom two) As in (top three),

but for zonal current observed by the RAMA mooring at 08, 80.58E from 2005 to 2012.

JUNE 2015 CHEN ET AL . 1537

years at different regions of the equatorial Indian

Ocean.

4. Model-simulated EUC and dynamics

a. Model/data comparison

To verify the HYCOM model performance, we first

compare the model solutions with RAMA data and

ORAS4 products. The simulatedmonthly zonal currents

at the RAMA location (08, 908E) from HYCOM gen-

erally show consistent patterns of the EUCs and Wyrtki

jets with the RAMA observations, even though the

model sometimes overestimates the current magnitudes

(Fig. 5). To further quantify the model/data comparison

and estimate the simulation errors, the 60–140-m depth-

averaged zonal currents, which span the EUC core

depths of both winter–spring and summer–fall, are

shown in Fig. 6. The phase of the monthly zonal currents

from the model agrees well with the observations, with a

correlation coefficient of 0.83 (above the 95% confi-

dence level). The observed and simulated zonal current

standard deviations (STDs) are 0.21 and 0.23ms21,

with a root-mean-square error (RMSE) of the model

relative to the observations of 0.12m s21. At 08, 80.58E,HYCOM subsurface currents also agree well with the

mooring data, with a correlation coefficient of 0.84 from

October 2004 to December 2011, observed and simu-

lated STDs of 0.21 and 0.25ms21, and RMSE of

0.13m s21. Note that the HYCOM results during pe-

riods when RAMA data are missing are not used to

estimate model errors. Note that not all eastward cur-

rents during summer–fall shown in Fig. 6 can be defined

as EUC. This point will be further discussed in sections

4b and 4c below.

HYCOMalso reasonably simulates the climatological

features of the EUC, including its semiannual variability

in the central-western basin (cf. Figs. 4 and 7). The

correlation coefficient of the area-averaged zonal ve-

locities for 0.58S–0.58N, 508–958E and 60–140m between

HYCOM and ORAS4 reaches 0.93 (Fig. 8).

The LOM also reasonably captures the subsurface

current amplitude and variability (Figs. 6, 8), but it often

systematically overestimates the current variability am-

plitude. This is because nonlinearity in the ocean (both

observations and HYCOM) tends to limit the linear

growth that occurs in theLOM.Time series for 60–140-m-

averaged zonal currents from LOM and RAMA have

reasonable agreement, with correlation coefficients of

0.51 and 0.54 at 908 and 80.58E, respectively (Fig. 6). TheSTDs are 0.25 and 0.28ms21 at 908 and 80.58E for LOM,

which are larger than the 0.21ms21 from RAMA data.

The area-mean subsurface zonal currents across the In-

dian Ocean equator for 0.58S–0.58N, 508–958E and 60–

140m from LOM agree well with those from ORAS4 and

HYCOM (Fig. 8). The correlation coefficients between

LOMandHYCOM (ORAS4) are 0.75 (0.74). These good

agreements demonstrate that the LOM can be used to

shed light on the roles played by the wind-driven equato-

rial waves in determining the Indian Ocean EUC and

variability.

To examine the sensitivity of model/data agreement

to the depth selection, we also analyzed the zonal cur-

rents averaged over the 100–140-m depths and obtained

similar agreement with those shown in Figs. 6 and 8 (not

shown). Because the 100–140-m depth is below the

summer–fall EUC core for the positive IOD years (e.g.,

2006, 2007, 2008, and 2011 of Fig. 5; see section 4c), we

choose to show currents for 60–140-m depth, which

FIG. 3. Zonal currents for (a)monthly averagedRAMAdata at 08, 908Eand (b)monthlyORAS4 data averaged for 0.58S–0.58N, 89.58–90.58E.White and black lines represent the 0 and 0.2m s21 (0.15m s21 for ORAS4) contours, respectively.

1538 JOURNAL OF PHYS ICAL OCEANOGRAPHY VOLUME 45

cover the EUC core for both the winter–spring and

summer–fall seasons.

b. Dynamics: Winter–spring EUC

Since the basinwide EUC regularly occurs each year

during winter–spring when equatorial easterly wind

prevails, we first examine the time evolution of wind and

EUC from January to May. Given that HYCOM solu-

tions agree with RAMA moorings as well as ORAS4

data, we will primarily present HYCOM results here-

after, together with LOM solutions to elaborate the

winter–spring EUC dynamics. The spatial patterns of

subsurface currents from LOM_MR have overall

agreement with those of theHYCOMandORAS4 data,

even though significant differences exist regarding de-

tailed structures and magnitudes (cf. Figs. 9b–d). These

differences reflect the effects of nonlinearity in the

oceanic system, which are included in ORAS4 data and

HYCOM solution but excluded from LOM_MR. Note

that from February to April, subsurface zonal currents

averaged over 60–140m of each year (Fig. 9) can be

defined as EUC in most regions of the Indian Ocean as

discussed above (also see red solid and red dashed lines

of Fig. 10). Therefore, we generally refer to the 60–

140-m-averaged current shown in Fig. 9 as the EUC.

From January to February, easterlywind occupiesmost

regions of the equatorial Indian Ocean, particularly the

western and central basin (Fig. 9a). From March to May,

the easterlies are gradually replaced by westerlies, which

drive the spring Wyrtki jet at the surface (Wyrtki 1973;

Han et al. 1999; Nagura and McPhaden 2010a). Corre-

sponding to the easterly winds, the winter–spring east-

ward EUC from HYCOM is set up along the equator

after;1 month (Fig. 9b; see also Nyadjro andMcPhaden

2014). The correlation coefficient between zonal wind

stress averaged over 58S–58N, 508–908E and zonal current

from HYCOM averaged over 0.58S–0.58N, 508–908E, andfrom 60–140m reaches 20.44 (20.47) above 95% sig-

nificance when the wind leads by 2 (1) months for the

entire (January–May) 2001–11 period. During 2001, 2008,

and 2011, boreal wintertime easterlies are confined to the

western basin west of 708E (Fig. 9a). Corresponding to

FIG. 4. The monthly climatology of vertical sections of 0.58S–0.58N averaged zonal current along the Indian Ocean equator from ORA-S4

data for 2001–11. Data of each year (rather than just the strong EUC years) is used to obtain the climatological mean. White and black lines

represent the 0 and 0.15m s21 contours, respectively. The grey lines are the temperature contours and the red ones represent D23 and D20.

JUNE 2015 CHEN ET AL . 1539

FIG. 5. Monthly zonal current at 08, 908E during (a) February–April from the mooring, (b) February–April fromHYCOM, (c) August–

October from the mooring, and (d) August–October from HYCOM. White and black lines represent the 0 and 0.2m s21 contours,

respectively.

1540 JOURNAL OF PHYS ICAL OCEANOGRAPHY VOLUME 45

these wind anomalies, the EUCs are evidently weaker

and appear primarily in the western basin (Fig. 9b).

Currents from the ORAS4 reanalysis product are very

similar to those from HYCOM (Figs. 9b,c).

To assess the roles played by the wind-driven equatorial

wave dynamics, we examine the solutions to the LOM.

Overall, Rossby andKelvin waves directly forced by winds

are themajor cause for generating the EUC in the western

basin, where the EUC often obtains its maximum (cf.

Figs. 9e and 9d). Reflected Rossby waves from the eastern

boundary (Fig. 9f), however, also have significant contri-

butions in the western basin and play a role in generating

the EUC in the eastern and central basin. For example, the

2002EUC,which obtains itsmaximum in the central basin,

results primarily from the reflected Rossby waves. The

2006 and 2010 EUCs also receive large contributions from

the reflected Rossby waves (cf. Figs. 9d–f). To further

quantify the above arguments, we obtain the time series of

EUC core values from LOM_MR and contributions from

directly forced and reflected Rossby waves (Fig. 10). As

shownbyFig. 10 (red solid anddashed lines), all subsurface

eastward currents from February to April are defined as

EUC. The STD of the EUC core speed is 0.12ms21 for

LOM_MR, 0.15ms21 for LOM_DAMP, and 0.11ms21

for (LOM_MR2 LOM_DAMP) from 2001 to 2011. The

large STD value from LOM_DAMP suggests the

importance of directly forced Rossby and Kelvin waves

to the EUC core, which is located in the western basin for

most years (Figs. 9b–d). The correlation between the total

EUC-core strength (Fig. 10, line with squares) and di-

rectly forced waves (line with circles) is 0.69, which is

much larger than the correlation coefficient of 0.12

between the total EUC and reflected Rossby waves (line

with triangles), even though the STD value of reflected

Rossby waves (0.11ms21) is comparable to the STD of

LOM_MR (0.12ms21). This is because most EUC cores

are located in the western basin, which is dominated by

the directly forced response and thus the high correlation.

Even though the EUC varies from year to year, it

occurs every winter–spring and often extends across

the equatorial basin. Thus, we further explore the

fundamental dynamics of winter–spring EUC below

using the monthly mean climatology from 2001 to

2011. To reveal the spatial structure of the equatorial

waves, we show horizontal maps of pressure and cur-

rents averaged over the EUC depth (Fig. 11). To iso-

late the long-wave dynamics associated with the

EUC, we consider the situation with only zonal wind

stress forcing, under long wavelength approximation

f[›t 1 (A/c2n)]yn 5 0 in Eq. (2b)g, and neglecting hori-

zontal mixing. Under these assumptions and consid-

ering the fact that wind is exerted near the ocean

surface and drives the subsurface EUC by setting up

pressure gradient, and that y 5 0 at the equator,

Eqs. (2a) and (2b) for EUC depth yield�›t 1

A

c2n

�un 1

1

rpnx5 0, and (3a)

un 521

rb

›2pn›y2

. (3b)

Note that we took the y derivative of Eq. (2b) to ob-

tain Eq. (3b). In this simplified system, only long Rossby

and Kelvin waves are retained, and their associated

FIG. 6. The 60–140-m-averaged zonal velocities fromRAMAmoorings, HYCOM, and LOM

at (a) 908E and (b) 80.58E. February–April and August–October are shaded. The mean values

of each data are removed.

JUNE 2015 CHEN ET AL . 1541

subsurface zonal flow obeys equatorial ‘‘geostrophy,’’ as

shown by Eq. (3b). Indeed, the 2(1/rb)(›2p/›y2) line

agrees very well with the subsurface zonal current at the

EUC depth range for all seasons (cf. the thick solid lines

of gray and black of Fig. 8), suggesting the validity of

long-wave approximation and the importance of long

Rossby and Kelvin waves in affecting the EUC. Equa-

tion (3a) describes the relationship between subsurface

zonal flow and zonal pressure gradient associated with

the damped long waves.

During January, easterly wind components prevail in

the western and central equatorial Indian Ocean

FIG. 7. As in Fig. 4, but for the HYCOM model for 2001–11.

FIG. 8. Area-averaged monthly zonal velocities for 0.58S–0.58N, 508–958E and 60–140m from HYCOM (thin black line), ORAS4

(dashed line), and LOM (thick black line). Gray line is the second derivative of pressure term (PGD),2(1/rb)(›2p/›y2), which equals u in

Eq. (3b) under the long-wave approximation. The 2001–11 mean value (including all months) of the data is removed.

1542 JOURNAL OF PHYS ICAL OCEANOGRAPHY VOLUME 45

(Fig. 1), which excites eastward-propagating equatorial

Kelvin waves and westward-propagating Rossby waves

[Eq. (2a)]. After these waves radiate out, pressure in-

creases in the west and the eastward pressure gradient

force is set up, which tends to balance the westward

(easterly) wind near the surface [Eq. (2a)]. Note that we

neglect horizontal mixing and vertical friction and con-

sider f 5 0 on the equator in the above discussion. The

setup of the eastward pressure gradient force takes

about 1 month (February panel of Fig. 11), a time scale

consistent with the 1-month lead of wind discussed

earlier (Figs. 9a,b). The eastward pressure gradient force

drives an eastward EUC [Eq. (3a); February panel of

Fig. 11]. During February, the easterly wind component

increases (Fig. 9a) and the eastward pressure gradient

force, together with the EUC, strengthen and extend

eastward during March (Fig. 11). Meanwhile, in the

eastern equatorial basin, high pressure occurs in March,

which corresponds to the relaxation of the equatorial

easterly wind component (Figs. 1, 9a). Reflected Rossby

waves from the eastern boundary are the primary cause

for the eastward subsurface current in the eastern basin

and have significant contributions to the EUC in the

central and western basin from February to April

(Figs. 9d–f). From April to May, equatorial westerlies

prevail, which increase the subsurface pressure in the

eastern basin and weaken the eastward EUC there.

Previous studies suggest that equatorial Kelvin and

Rossby waves associated with the first and second

baroclinic modes primarily determine the seasonal

FIG. 9. (a) Longitude–time plot of 58S–58N averaged monthly CCMP zonal wind stress (Nm22) from January to May for 2001–11, and

the line contours show zero value. (b) As in (a), but for 0.58S–0.58N, 60–140m averagedmonthly HYCOM zonal current (m s21). (c) As in

(b), but for ORAS4. (d) As in (b), but for LOM_MR from the sum of the first 25 baroclinic modes. (e) As in (b), but for LOM_DAMP.

(f) Difference between (d) and (e) (LOM_MR 2 LOM_DAMP), which measures Rossby waves reflected from the eastern ocean

boundary. Note that color bars in (a) and (b) are different from those of (c)–(f). Color bars range from20.5 to 0.7m s21 for (b) but from

20.5 to 0.5m s21 for (c)–(f).

JUNE 2015 CHEN ET AL . 1543

variability of zonal surface currents in the equatorial

Indian Ocean (Han et al. 1999; Nagura and McPhaden

2010a). Consistent with these studies, the first two baro-

clinic modes also dominate the surface Wyrtki jets in our

LOM_MR (not shown). However, the situation is notably

different in the subsurface layer. Currents at 60–140m

along the equator are well represented by the sum of the

first eight modes (Figs. 12a,b). The low-order modes (e.g.,

modes n 5 1, 2) have larger cn (section 2b) and thus

weaker friction becauseA/c2n is small [Eq. (3a)]. They are

approximately the response of inviscid Kelvin andRossby

waves and are very important for generating the winter–

spring EUC (Figs. 12c–e). The second baroclinic mode is

the dominant mode of semiannual variability of equato-

rial zonal current, as is also shown by previous studies

(e.g., Han et al. 1999; Nagura andMcPhaden 2010a; Yuan

and Han 2006). The intermediate modes (n5 3, 4, . . . , 8)

have relatively lower cn and therefore larger friction. They

are equatorial waves with significant damping and are also

important for the winter–spring EUC (Figs. 12f,g), ac-

counting for;40% of the climatological EUC amplitude

near the EUC core during February–March. The high-

order modes (modes 9–25) have small cn values and thus

large friction. They are essentially in ‘‘pseudo Ekman

balance,’’ which is the balance among friction, Coriolis

force, and surface wind stress and have little contribution

to the pressure gradient force and EUC.

The lag correlation between wind stress and EUC

shown above suggests that equatorial waves need 1–

2 months to set up the pressure gradient force in the

equatorial basin. For the first baroclinic mode, it takes

the equatorial Kelvin wave ;1 month and first meridi-

onal mode Rossby wave ;3 months to cross the entire

equatorial Indian Ocean basin. For the second baroclinic

mode, it takes the Kelvin wave ;1.5 months and first

meridional mode Rossby wave 4.5 months to cross the

equatorial basin. Given that strong easterly winds occur

primarily in the western basin and that directly forced

waves are the major cause for the winter–spring EUC

core, the time scale for the directly forced waves to affect

the western basin EUC is much shorter than the time to

cross the entire basin. More importantly the first eight

modes contribute to the total EUC velocity (Figs. 12a,b),

and the wind–EUC correlation reflects this relationship.

As shown in Figs. 12a and 12b, from February to April, it

takes the EUC signal (the sum of modes 1–8) less than a

month to propagate from the central basin (;708E) to theeastern boundary, somewhatmore than amonth from the

central basin to the western boundary, and;2 months to

cross the entire equatorial basin (Fig. 12b). This explains

the 1-;2-month time lag betweenEUCand surfacewind

correlations.

c. Dynamics: Summer–fall EUC reappearance

During summer and fall, eastward subsurface flow

occurs every year along the Indian Ocean equator

(Fig. 13). The EUC (with a subsurface maximum),

however, disappears in 2010 but reappears in all other

years (section 3). Recall that not all eastward subsurface

currents in the depth range 60–140m can be defined as

EUC, since some flow at shallow depths in this range

results from the downward extension of surface Wyrtki

jets (Figs. 2–4). The eastward subsurface flow generally

has a good correspondence with the equatorial easterly

wind that appears in the eastern and western basins

except for 2010, when westerly wind prevails in this

FIG. 10. Time series of EUC strength from the LOM_MR (line with squares), which is de-

fined as the averaged value of the EUC from 2108 to 108 longitude of the EUC core shown in

Fig. 9d; the line with filled circles represents the contribution from directly forced response

(LOM_DAMP) and that with triangles shows eastern boundary reflection (LOM_MR 2LOM_DAMP). To demonstrate that this current can be defined as EUC, the time series of the

subsurface current from ORAS4 data averaged for the same region as the LOM_MR line is

shown (red). The surface currents averaged for (0–50m above the EUC) is also shown (red

dashed line). Because the surface current is either westward or weakly eastward compared to

subsurface flow, the subsurface current is EUC.

1544 JOURNAL OF PHYS ICAL OCEANOGRAPHY VOLUME 45

negative IOD year (Figs. 13a,b, 14a). Warm water ac-

cumulation in the eastern basin and the development of

zonal temperature gradient (warmer in the east com-

pared to the west) may be a key element that contrib-

uted to the strengthening of westerly anomalies

(Krishnan et al. 2006). The eastward subsurface current

simulated by the LOM has an overall agreement with

the HYCOM solution, even though significant differ-

ences exist between the two (cf. Figs. 13b,d). Directly

forced Kelvin and Rossby waves dominate the currents

in the western basin for all years, and they also have

significant contributions in the central-eastern basin in

2006, 2007, and 2011 (cf. Figs. 13d and 13e) when posi-

tive IOD events occur (Fig. 14a). Strong easterly wind

anomalies associated with the IOD events have been

suggested to be important in causing the summer–fall

reappearance of EUC (section 1). Different from the

winter–springEUC that often reaches itsmaximum in the

western basin (Fig. 9), summer–fall eastward subsurface

flow often reaches its maximum in the central-eastern

basin. Evidently, reflected Rossby waves from the

eastern boundary make larger contribution to current

maxima in this region than the directly forced response

(Figs. 13d–f).

Here, we investigate further the idea that equatorial

winds associated with the IOD affect the EUC during

summer and fall. The September–November (SON)

mean DMI suggests that the positive IOD events with

larger than 0.5 STD occur in 2002, 2006, 2007, and 2011

(Fig. 14a), which is verified by the spatial distributions of

SST anomaly (SSTA) for individual years (not shown).

For instance, in 2002 obvious positive and negative

SSTAs appear in September in the western node and

eastern node regions, respectively. These SSTAs cor-

respond to easterly wind anomalies in the central-

eastern equatorial Indian Ocean. The positive SSTA in

the western node becomes weaker in October, whereas

the negative SSTA becomes stronger corresponding

to the stronger easterly wind anomaly in the eastern

equatorial Indian Ocean (also in November). Even

FIG. 11. The spatial structures of 60–140-m-averagedmonthly climatology of pressure (3980 g cm s22; color) and currents (vectors) for the

2001–11 period from LOM_MR, which are the sums of the first 25 baroclinic modes.

JUNE 2015 CHEN ET AL . 1545

though the ‘‘negative SSTA’’ obtains its maximum

magnitude from 58 to 108S and never reaches the equa-

tor, the area-averaged SSTA in the eastern node sub-

tracted from the SSTA in the western node is;1 STD of

the DMI, and thus we identify 2002 as a ‘‘weak IOD’’

year. The 2011 case is somewhat different. It is charac-

terized by strong easterly wind anomalies along the

equator with weak southeasterly wind anomalies in the

tropical southeastern basin. Cai et al. (2009) suggested

that 2008 is also a positive IOD event. Considering that

large easterly wind anomalies occur during June–August

but sharply reduce afterward, we identify 2008 as an

aborted IOD event (Fig. 14b). Atmospheric intraseasonal

disturbances may lead to termination of the IOD event

(Rao and Yamagata 2004), which is beyond the scope of

this study. A strong negative IOD event occurred in 2010

and a weaker one occurred in 2005.

To reveal the different characteristics of the summer–

fall EUC associated with the IOD, we perform com-

posite analyses for zonal currents, SSTA, and wind

anomaly during August–October for positive IOD years

(2002, 2006, 2007, 2008, and 2011), negative IOD years

(2005 and 2010), and normal years (2001, 2003, 2004, and

2009). Note that the aborted IOD of 2008 is included in

the positive IOD composite.

A common feature of EUC for each panel (left column

of Fig. 15) is the existence of a maximum in the western

basin. A striking difference exists between positive IOD

and negative IOD years (Figs. 15a,b). During a positive

IOD, the EUC extends across the equatorial basin

(Fig. 15a). An eastern basin maximum occurs at a shal-

lower core depth than that of the western basin maxi-

mum, and the EUC flows below the westward surface

current (Fig. 15a) due to forcing by strong easterly wind

anomalies, which shoal the thermocline and thus EUC

core depth, drive an eastward pressure gradient force and

EUC, and force a negative SSTA in the eastern equatorial

basin (Fig. 15e). In contrast, during a negative IOD, the

EUC disappears in the eastern basin with a strong east-

ward flow near the surface (Fig. 15b), which is driven by

the strong westerly wind anomalies that also induce pos-

itive SSTA in the eastern basin (Fig. 15f). The normal

years share similar characters with the negative IOD

years but with a relatively weak eastward flow near the

surface (cf. Figs. 15b and 15c). By performing statistical

Monte Carlo tests at each grid of zonal current sections,

we find that the mean difference in zonal velocities in the

eastern basin (758–958E) between the positive IOD years

and other years exceeds the 90% significance level.

However, zonal currents in the western basin present less

difference during these years. Monte Carlo tests also

suggest that no obvious difference in zonal currents exists

between the negative IOD and normal years. Rao and

Yamagata (2004) considered 2003 as an aborted positive

IODevent.By examining thewind andEUC for this year,

we found that westerly wind anomalies dominate the

equatorial Indian Ocean during August–September, and

theEUCpresents similar features as 2001, 2004, and 2009.

In addition to the IOD, equatorial wind variability

associated with Indian monsoon rainfall has also been

FIG. 12. (a) Time–longitude plot of the 2001–11 mean zonal velocities along the Indian Ocean equator (0.58S–0.58N average) from

LOM_MR for the 60–140-m layer from the sum of modes 1–25. (b)–(f) As in (a), but from the contributions of modes 1–8: (b) mode 1,

(c) mode 2, (d) modes 1–2, (e) modes 3–4, and (f) modes 5–8. White lines represent the 0m s21 contours.

1546 JOURNAL OF PHYS ICAL OCEANOGRAPHY VOLUME 45

suggested to affect the summer–fall EUC of IOD years

(Swapna and Krishnan 2008; Krishnan and Swapna

2009). For the period of 1988–2011 when both CCMP

winds and NOAA SST data are available, the correla-

tion coefficient is 20.81 between SON-mean DMI and

JJA-mean equatorial zonal wind anomalies and 20.91

between SON-mean DMI and SON-mean equatorial

zonal wind anomalies averaged over 58S–58N, 508–958E,with both correlations exceeding 95% significance.

When all months are included, the correlation drops

to20.56 using monthly DMI andmonthly CCMPwinds.

As a comparison, the correlation between summer

monsoon rainfall and equatorial zonal wind anomalies

is20.34 during June–September and20.44 (above 95%

significance) for June–August. These results suggest

that equatorial easterly wind anomalies associated with

strong (weak) Indian summer monsoon could help to

intensify (weaken) the zonal wind anomalies and

therefore summer–fall EUC associated with IOD events,

a result consistent with existing studies (Swapna and

Krishnan 2008; Krishnan and Swapna 2009). Interannual

variability of equatorial zonal wind over the Indian

Ocean, however, is predominantly determined by zonal

SST gradients associated with IOD events as shown by

their high correlations, with Indian monsoon rainfall

playing a less important role.

Following Swapna and Krishnan (2008), we use all-

India rainfall anomalies to define strong or weak mon-

soons. Excess (deficit) monsoons are defined when the

June–September all-India rainfall is more (less) than

10% of the long-term climatological normal, and those

within 10% of the long-term norm are defined as ‘‘nor-

mal monsoon years.’’ For our period of interest from

2001 to 2011, only 2007 is a year of positive IOD and

strong monsoon (Fig. 14), which does correspond to a

fairly strong eastern basin EUC during summer and fall.

FIG. 13. As in Fig. 9, but for July–December. The core areas of the EUC in (b) are marked by the word EUC. The year 2010 without EUC

mark shows just subsurface eastward flow.

JUNE 2015 CHEN ET AL . 1547

Year 2002 is a positive IOD and weak monsoon year

(Fig. 14). The EUC, however, is comparable to that of

2007 with a maximum occurring in the eastern basin

(Fig. 13). It shares similar EUC features with the posi-

tive IOD and normal-to-strong monsoon years (2006,

2007, 2008, and 2011), except for a weaker but thicker

and deeper core in the eastern basin (not shown). This

result suggests that the situation for weak monsoon

years can be rather complex.

Does an EUC exist in weak monsoon years? Com-

posite analyses of the three weak monsoon years 2002,

2004, and 2009 show that the EUC is not weaker than

that of normal years (Figs. 15c,d). The western basin is

dominated by easterly wind anomalies associated with

weak monsoon (Fig. 15h), which can drive a westward

surface flow and an eastward EUC. The EUC can ex-

tend to the central-eastern basin, albeit with weaker

magnitude compared to positive IOD years.

As the situation for winter–spring EUC, the summer–

fall EUC in theLOM isweaker than that in theHYCOM.

However, the LOM can reproduce the EUC in the

western basin and interannual variability of the EUC in

the eastern basin (Fig. 16). For all cases, the western basin

EUCmaximum results primarily from the directly forced

FIG. 14. (a) DMI averaged in SON from 1982 to 2011; only results from 2001 to 2011 are

shown. The SON season is the peak of the IODevents. (b) DMI for eachmonth. (c) Zonal wind

stress anomalies (with climatological mean removed) averaged for 58S–58N, 508–958E.(d) June–September accumulated all-India rainfall anomalies, which are the departures from

the long-term climatological normal of ;853mm (Swapna and Krishnan 2008). September–

November in (b) andAugust–October in (c) are shaded; the two horizontal dashed lines in each

panel show 60.5 STD.

1548 JOURNAL OF PHYS ICAL OCEANOGRAPHY VOLUME 45

Kelvin andRossbywaves (cf. the solid and dashed lines of

Fig. 16), whereas the eastern basin EUC is dominated by

reflected Rossby waves from the eastern ocean boundary

(dashed–dotted lines). It is only during the positive IOD

years that direct wind forcing also has significant contri-

bution to the eastern basin EUC (Fig. 16a).

5. Summary and discussion

The EUC is a transient feature in the Indian Ocean,

which was thought to regularly occur during spring but

reappear in summer–fall only in some years—the years

when both positive IOD event and strong Indian mon-

soon occur. Because of the time and space limitations of

observational data, EUC spatial structure and temporal

variability are unclear and have not yet been systematically

investigated. Combining data from moorings and ocean

reanalysis with modeling experiments using HYCOM and

LOM, this study provides a systematic investigation on the

structure and dynamics of the Indian Ocean EUC, in-

cluding its seasonal and interannual variability, for the

period of 2001–11. We also examined the EUC structures

for both spring and summer–fall using ORAS4 reanalysis

products for the entire 1958–2011 period.

Consistent with existing studies, the Indian Ocean

EUC indeed regularly occurs in each winter–spring,

particularly from February to April, and it can exist

in the western, central, and eastern equatorial basin.

The EUC generally extends from about 60m to more

than 200m, with a core depth near D20 within the

FIG. 15. Composite analyses of August–October mean zonal velocities from the HYCOM for (a) positive IOD

years (2002, 2006, 2007, 2008, and 2011), (b) negative IOD years (2005 and 2010), and (c) normal years (2001, 2003,

2004, and 2009). (d) Years 2002, 2004, and 2009 are weak monsoon (or monsoon deficit). (e)–(h) The composites for

SSTA (color) andwind stress anomaly (vector) corresponding to the years shown in (a)–(d), respectively.Wind stress

vectors with westward components are red and eastward components are black. The scale in upper-left corner of

(e)–(h) represents 0.05Nm22.

JUNE 2015 CHEN ET AL . 1549

thermocline. Rossby and Kelvin waves directly forced

by winds are the major cause for generating the EUC in

western basin, where theEUC core is located formost of

the years we studied. Reflected Rossby waves from the

eastern boundary, however, also have a significant

contribution in the western basin and play a de-

terministic role in producing the EUC in the eastern and

central basin (Figs. 9, 10). The EUC associated with the

equatorial Kelvin and long Rossby waves satisfies

equatorial geostrophy, with u ’ 2(1/rb)(›2p/›y2)

[Fig. 8; Eq. (3b) of section 4b]. Easterly wind excites

equatorial Kelvin and Rossby waves and increases the

subsurface pressure in the western equatorial basin,

which drives an EUC in the western basin [Fig. 11; Eq.

(3a) of section 4b]. The EUC strengthens and extends

eastward subsequently and finally weakens by the high

subsurface pressure in the eastern basin forced by the

equatorial westerly winds during April–May. Different

from the situation near the surface, where the current is

controlled by the first two baroclinic modes, the current

in the subsurface layer is affected not only by the low-

order modes 1–2, which essentially represent inviscid

Kelvin and Rossby waves’ effects, but also by the in-

termediatemodes 3–8. Equatorial waves associated with

these intermediate modes are subjected to significant

damping, and they contribute ;40% of the winter–

spring EUC core amplitude (Fig. 12).

The EUC often reappears during summer–fall, par-

ticularly from August to October, and can also be ob-

served in July (e.g., 2008) and December (e.g., 2006) in

some years. Indeed, the summer–fall EUC reappears in

all years from 1958 to 2011, except for 2010 in ORAS4

data. In 2010, the summer–fall EUC completely dis-

appeared. By analyzing vertical sections of ORAS4

zonal currents along the equator for individual years, we

find that the summer–fall EUC reappears at different

regions of the equatorial Indian Ocean. It usually occurs

in the western basin but also is present in the central and

FIG. 16. The averaged zonal currents from LOM_MR

(solid lines), LOM_DAMP (dashed lines), and the difference

between LOM_MR and LOM_DAMP (dotted–dashed lines) for

(a) positive IOD years, (b) negative IOD years, and (c) normal

years. Because of the weaker subsurface currents in the LOM,

the 80–140-m data are used here to examine the summer–

fall EUC.

FIG. 17. (a) Composite analyses of NCEP surface wind anomaly

(m s21; vectors) for weak monsoon years of 1951, 1965, 1966, 1968,

1972, 1974, 1979, 1982, 1985, 1986, 1987, 2002, 2004, and 2009.

(b) As in (a), but for NCEP surface wind anomaly and NOAA SST

anomaly (color) for 1982, 1985, 1986, 1987, 2002, 2004, and 2009.

1550 JOURNAL OF PHYS ICAL OCEANOGRAPHY VOLUME 45

eastern basin during positive IOD years (Fig. 15). Dur-

ing the negative IOD and normal years, the EUC does

not exist in the eastern basin and only occurs in the

western basin. During all years, the western basin EUC

is caused primarily by direct, wind-driven Rossby and

Kelvin waves. By contrast, the eastern basin EUC dur-

ing positive IOD years results largely from reflected

Rossby waves from the eastern boundary, with directly

wind-forced waves also having significant contributions.

During negative IOD and normal years, eastward sub-

surface flow induced by eastern boundary–reflected

Rossby waves is canceled by the westward flow di-

rectly driven by winds (Fig. 16).

Zonal wind anomalies in the equatorial Indian Ocean

associated with strong Indian monsoon rainfall may help

to intensify the zonal wind anomalies and thus summer–

fall EUC associated with positive IOD events, consistent

with existing studies. Note, however, that interannual

variability of equatorial zonal wind during summer–fall is

predominantly determined by the zonal SST gradients

associated with IOD events, with monsoon rainfall vari-

ability playing a less important role. The situation for

weak monsoon years appears to be more complex.

Composite analysis shows that the summer–fall EUC

during weak monsoon years can be stronger than that of

normal years. This is because the easterly wind anomalies

in the western equatorial basin associated with the weak

monsoon can drive an eastward pressure gradient force

and thus EUC, which can extend eastward to the central

and eastern basin (Fig. 15).

From 1948 to 2012, there were 14 weak monsoon

years: 1951, 1965, 1966, 1968, 1972, 1974, 1979, 1982,

1985, 1986, 1987, 2002, 2004, and 2009. The NCEP

wind anomalies based on the 14 yr show easterlies

extending across the Indian Ocean, especially in the

western basin (west of 658E; Fig. 17a). By analyzing

54-yr ORAS4 products from 1958 to 2011, we do not

find that the EUC weakens significantly or disappears

during the last 13 weak monsoon years. The available

NOAA SST anomalies for 1982, 1985, 1986, 1987,

2002, 2004, and 2009 suggest that the SST tends to be

higher in the western basin during weak monsoon

years (Fig. 17b), which could aid the formation of

the stronger easterly wind anomaly in the western

basin by creating a westward pressure gradient force

through pressure gradient and vertical diffusion

mechanisms.

Acknowledgments. This work is supported by National

Basic Research Program of China (Grant 2011CB403503),

the ‘‘Strategic Priority Research Program’’ of the Chinese

Academy of Sciences (Grant XDA11010103), the Exter-

nal Cooperation Program of BIC, Chinese Academy of

Sciences (Grant GJHZ201319), National Natural Science

Foundation ofChina (Grants 41476011 and41206008), and

Pearl River S&T Nova Program of Guangzhuou

(2013J2200087).W. Han is supported by NSF CAREER

Award OCE 0847605. M. McPhaden is supported

by NOAA.

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