Meteorology
and AtmosphericPhysics

Meteoro|. Atmos. Phys. 44, 29-41 (1990)

9 by Springer-Verlag 1990

551.515.5:551.513
Institute of Geoscience, University of Tsukuba, Tsukuba, Japan

Impact of Indian Monsoon
on the Coupled Atmosphere]Ocean System in the Tropical PacificT. YasunariWith 9 Figures
Received May 19, 1989
Revised September 25, 1989

SummaryThis study addresses the relationship between the Indian summer monsoon (ISM) and the coupled atmosphere/ocean system in the tropical Pacific on the interannual time scales.
High positive correlations are found between ISM rainfall
and both mixed layer sea water temperature (SWT) and sea
surface temperature (SST) anomalies of the tropical western
Pacific in the following winter. Negative correlations between
ISM rainfall and SST in the central/eastern Pacific also appear to be most significant in the following winter. These
parameters are correlated with each other mainly on a biennial time scale. Lag-correlations between the zonal wind
and SST along the the equatorial Pacific show that the westerly (easterly) surface wind stress anomalies over the central/western Pacific are greatly responsible for the formation ofnegative (positive) SST/SWT anomalies in the western Pacificand positive (negative) SST/SWT anomalies in the central/eastern Pacific. Furthermore, it is evidenced that these lagcorrelations are physically based on the anomalies in thelarge-scale convection over the Asian monsoon region andthe associated east-west circulation over the tropical Pacific,
which first appear during the Ihdian summer monsoon season
and evolve during the following autunm and winter.
These results strongly suggest that the Asian summer
monsoon may have an active, rather than a passive, role on
the interannual variability, including the ENSO events, ofthe coupled atmosphere/ocean system over the tropical Pacific.
1. Introduction
Since the early p a r t o f this c e n t u r y ( W a l k e r a n d
Bliss, 1932), m a n y studies h a v e been d o n e o n ther e l a t i o n s h i p b e t w e e n the I n d i a n s u m m e r m o n s o o n( I S M ) a n d the E1 N i f i o / S o u t h e r n Oscillation( E N S O ) p h e n o m e n a . In recent years, E N S O has
b e c o m e a m a t t e r o f great interest as an a t m o s p h e r e / o c e a n i n t e r a c t i o n which affects the intera n n u a l variability o f global climate. In particular,some w o r k s have stressed the role o f E N S O in thei n t e r a n n u a l variability o f the I n d i a n m o n s o o n . Nicholls (1984), for example, a d o p t e d S S T a n o m a l i e s
in the I n d o n e s i a n - n o r t h A u s t r a l i a n area as a pred i c t o r o f the I n d i a n m o n s o o n rainfall. R a s m u s s o n
a n d C a r p e n t e r (1983) suggested t h a t SST a n o m alies in the eastern Pacific can be a p o t e n t i a l pred i c t o r for the I n d i a n m o n s o o n rainfall. Angell(1981), o n the o t h e r h a n d , n o t e d t h a t S S T a n o m alies in the e q u a t o r i a l eastern Pacific d u r i n g falla n d w i n t e r are highly c o r r e l a t e d with the p r e c e d i n g
I n d i a n m o n s o o n rainfall. S h u k l a a n d P a o l i n o(1983) also stressed the influence o f the I n d i a n
m o n s o o n o n E N S O , p o i n t i n g out t h a t I n d i a n
m o n s o o n rainfall a n o m a l i e s are highly c o r r e l a t e dwith sea-level pressure a n o m a l i e s at D a r w i n (asan index o f the S o u t h e r n Oscillation) in 1Lhe following a u t u m n a n d winter. G l o b a l analyses o f sea
level pressure (Barnett, 1985) a n d a t m o s p h e r i c circ u l a t i o n fields ( Y a s u n a r i , 1987) also suggested therole o f the I n d i a n m o n s o o n o n the e v o l u t i o n o f
the E N S O cycle. These studies s h o w e d distinct
signatures o v e r the I n d i a n O c e a n a n d the E u r a s i a n

30

T. Yasunaricontinent which preceded the ENSO events in theequatorial eastern Pacific.
On the other hand, oceanographical evidence(Wyrtki, 1975, etc.) suggests the importance of the
thick and warm oceanic mixed layer in the equatorial western Pacific as an origin of the E1 Nifio
events in the eastern Pacific. Recent oceanic modelstudies with observed wind forcing (White et al.,
1987; Miyakoda et al., 1987) also showed that an
anomalously high heat content in the surface layerof the western tropical Pacific is formed one year
prior to a mature ENSO event by strong easterly
wind stress over the tropical north Pacific. Veryrecently, Yamagata and Masumoto (1989) demonstrated, by use of a simple atmosphere/ocean
coupled model, that an ENSO event evolves automatically on the condition that warm SST andpositive anomaly of oceanic heat content (OHC)
covers all the equatorial Pacific, as was pointed
out by Zebiak and Cane (1987). They have alsosuggested, however, that this condition is prepared
only when the atmospheric westerly bursts andpositive OHC anomaly in the western Pacific arecoupled. Pazan et al. (1986) noted the role of offequatorial Rossby waves, which propagate westward from the central Pacific, on the anomalously
high dynamic height in the western Pacific. In viewof this, the important question may be what actually controls the heat content and dynamic
height anomalies of the oceanic mixed layer in thetropical western Pacific, and how the Asian monsoon activity is related to these anomalies. This
may give a clue as to what the mechanism of the
ENSO cycle is.
The present paper attempts to solve these problems based on observational data. Some atmospheric and oceanic key parameters will be systematically examined by lag-correlation and composite techniques, particularly relevant to Indian
monsoon activity. Some apparent evidences will
be presented that the variability of the Indian summer monsoon and the associated convective activity over the Asian monsoon region play a crucialrole in the ENSO cycle.
2. DataMonthly, gridded 15 year (1970-1984) SST datacompiled at N O A A are used in the present study.
The 2 ~x 2 ~ lat./lon, grid data for the Pacific basin
were converted into areal mean data of 4 ~ x 4" lat./lon. blocks. January and July SWT profile data
along 137 ~ for the 19 years (1967-1985), regularly observed by the research vessel Ryofu-Maruof the Japan Meteorological Agency (JMA), arealso utilized. This data set is invaluable for examining the depth and temperature of the oceanic
mixed layer in the tropical western Pacific. ISM
rainfall (June to September) index (1871-1986) is
adopted from Parthasarathy (1987), which was
originally computed from the monthly rainfall
data of about 3000 raingages by weighting theareas of the different meteorological subdivisionsof India (Parthasarathy and Mooley, 1978).
Monthly mean zonal wind anomalies at 700 mb
and 200rob for the 17 years (1968-1984) are
adopted from the National Meteorological Center(NMC)'s operational tropical wind field analysis.
Since the wind data at 850 mb and the surface do
not adequately cover the analysis period, we utilized the wind field at 700 mb as representative ofthe lower troposphere. The validity of this assumption is discussed in Yasunari (1985). Monthly
mean outgoing longwave radiation (OLR) datacompiled at N M C for the period from 1974 to1984 are also used. The anomalies of each parameters described here are computed simply by subtracting the mean monthly value, which is obtained by averaging the all available data.
3. Indian Monsoon and SST in the Pacific Basin

Figure 1 shows the lag-correlations between the
ISM rainfall and the SST anomalies for wider
areas in the equatorial western and eastern Pacific.
To deduce the relationships for seasonal means,
the SST data are smoothed by three-months moving average. Y(0) denotes the year of the summer
monsoon referenced. It is noteworthy that the correlations of SST to ISM rainfall gradually increase
from the summer of Y(0) to the following winter,and reach the maximum (minimum) in January
or February of Y ( + 1) in the western (eastern)
Pacific. Particularly in the western Pacific, themaximum value of the correlation coefficientreaches + 0.79, which is far above the 99% confidence level. Relatively high values with opposite
signs to Y(0) are also found in the summer andfall months (June to November) of the year before(Y( - 1)), though these values are not as significant
as those of the Y(0)/Y( + 1) winter. This secondary
maximum (or minimum) of correlation in Y ( - 1)

hnpact of Indian Monsoon on the Coupled Atmosphere/Ocean System
t.0

i

II

08

I

/,,

Feb.

Apr. ,June Aug.Oct. Dec

Feb. Apr. ,June Aug Oct. Dec. Feb. Apr. June Aug Oct, Nov,

seems to be due to the strong biennial nature ofboth ISM rainfall (Bhalme and Jadhav, 1984) and
SST in the equatorial Pacific (Yoshino and Kawamura, 1987; Meehl, 1987).
Figure 2 shows the spatial pattern of lag-correlations between the ISM rainfall and the SST
anomalies for the (a) winter of Y(-1)/Y(0), (b)summer of Y(0) and (c) winter of Y(0)/Y(+ 1).
Systematic changes in correlation patterns can beseen from the preceding winter to the following
winter. The change of sign between t h e westernand the eastern Pacific is noticeable particularly
along the equatorial belt. The overall pattern ismost significant in the winter of Y(0)/Y(+ 1), andis very similar to the SST anomaly pattern for the
typical E1 Nifio phase (e.g., Rasmusson and Carpenter, 1982). However, it should also be notedthat the positive correlations in the western Pacificare as significant as the negative correlations in
the central through the eastern Pacific, which is
not obvious from the real SST anomaly pattern.
The pattern for the summer of Y(0) is similar tothat for the winter of Y(0)/Y( + 1), but the values
as a whole are not as significant. Thus, the SST
anomalies in the tropical Pacific, particularly inthe western part are closely associated with Indianm o n s o o n activities in such manner that SST follows the m o n s o o n with a lag of one to two seasons.
In the subtropical western Pacific and along theeast coast of the Asian continent, a narrow areaof negative correlations is found in the winter ofY(0)/Y(+ 1), which is not seen in the summer ofY(0) (Fig. 2 (b)). This anomalous pattern seems

31

Fig. 1. Lag-correlationsbetweenIndian monsoon rainfall and sea surface temperature in the western (0~8 ~ 130 ~~ and the eastern(0~176 170~176
Pacific.
The reference monsoon season isshown with thick black bar. Y(0)denotes the year of referencemonsoon and Y(- 1) (Y(+ 1)) denotes
the year before (after) the reference
monsoon year

to be closely related to the effect of anomalous
winter m o n s o o n surges from Siberia to areas of
SST anomalies in the northwest Pacific, as discussed by Kawamura (1986) and Hanawa etal.
(1988). If this is the case, then there should be asignificant association between the Asian summer
monsoon activity and the following Asian winterm o n s o o n activity.
4. Indian Monsoon and SWTin the Western Pacific

The SWT anomalies may represent more adequately the measure of heat content anomalies in
the oceanic mixed layer. This is true particularlyin the western Pacific, where variances of SST
anomalies are small compared to those in the eastern Pacific. Figure 3 shows the vertical time sectionof SWT anomalies from the surface to 800 m depth
averaged from 2 ~ through 10 ~ along the longitude line of 137 ~ A remarkable feature is thatthe anomalies show a nearly in-phase fluctuation
through the whole mixed layer. The maximum
variance occurs at 100m through 200m depth,
which corresponds to the main thermocline. Another feature to be noted in this diagram is thedistinct biennial nature of the anomalies.
The lagged-correlations between the ISM rainfall and SWT anomalies averaged for the surface
mixed layer (0-30 m) are shown in Table 1. The
most significant correlation appears in January ofY ( + 1) with a value of 0.83, whereas relatively
high negative correlations are also seen in the pre-

32

T. Yasunari
WINTER

OF Y(-t)/Y(O)

~"

/"

f"

.-~---._-__-o.~

IOOE

120

t40

160

SUMMER

OF

Y(O)

(

160

140

180

?.

1

/

~.

,

120

140

100

80

60W

t00

80

60W

....

3

..IS.-"

IOOE

I20

I60

180

160

140

120

(Y(0)) plus January (Y(+ 1)) of each depth isshown in Fig. 4. High positive correlations exceeding the 99% confidence level are found in thelayer above 300 m depth, which roughly corresponds to the bottom of main thermocline of thislatitude zone (2 ~~ (Kurihara, 1985). The
maximum correlation (nearly 0.80) appears at 10
to 20m depth, e.g., the core layer of the surface
mixed layer, while the second maximum exists atthe core layer of the main thermocline (about
150 m), where the variance of the SWT anomalies
appears to be maximum (see Fig. 3).
The time series of ISM rainfall anomalies arecompared to the SWT anomalies at 20m and100 m in the following January, as shown in Fig. 5.
It is obvious from this figure that the strong (weak)summer monsoons are followed by the positive(negative) SWT anomalies. It is also evident thatthese two time series show the quasi-biennial oscillation, which may be responsible for the secondary maxima in the correlations between the
ISM rainfall and the SST and SWT in the preceding summer through winter (Fig. 1 andTable 1).
5. Indian Monsoon
and Tropical East-West Circulation

WINTER
60,

20

-

-

,

OF

Y(O)/Y(+I)

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~40

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~OOE ~20~ 9 \ ~ . ~ . ~S

~80

~60

~40

9

i
9

~20

9

~00

....

80

60W

Fig. 2. Spatial distribution of lag correlation between the
I n d i a n m o n s o o n rainfall of Y(0) and sea surface temperature
anomalies for (a) preceding winter (DJF), (b) contemporary
summer (JJA) and (c) following winter. Positive (negative)
values of more (less) than 0.6 ( - 0 . 6 ) are shaded (stippled)ceding seasons (July of Y ( - 1 ) and January ofY(0)) with a minimum value of -0.75. These results are consistent with those from the SST data
(Fig. 1).
The vertical profile of the correlations between
the ISM rainfall and the SWT anomalies of July

The results of lag correlations between ISM rainFall and SST and SWT in the tropical Pacific suggest that the strong convective activities over south
Asia through the central Pacific associated withthe active Indian summer monsoon may cause the"piling up" of warm surface waters over the tropical western Pacific to the east of the Philippine
islands, via a strong tropical east-west circulation(Krishnamurti, 1971). The oceanic process for thepiling up of warm water by strong surface easterly
wind stress is discussed by White et al. (1987) andothers. What is crucial for the evolution of SST
and SWT anomalies over the Pacific, as shown in
Fig. 1 and Fig. 2, is that the strong easterly windstress over the western Pacific should be maintained or further intensified throughout the seaonsfrom the norther summer to the following winter.
This feature is clearly shown in the lag correlation
diagram between [SM rainfall and zonal wind
anomalies in the upper (200 rob) and the lower(700rob) troposphere along the equatorial belt
(5 ON-j oS) as shown in Fig. 6. At 700rob (Fig.
6 (a)), negative correlations are significantly highImpact of Indian Monsoon on the Coupled Atmosphere'Ocean Systemo1020305075100~125v 15032 200-:,o, ;o

V

,?.o

33

(Ol

~-- 2 5 @O_
LU 300f'~ 400500600700
800900 j~p

T

iJan. July IJao. July ,)an. july ,Jan. ,July

"8t

I

"82

'83

'84

I

Ja'/SJ;Y

Fig. 3. Vertical time section of January and July sea water temperature anomalies fl-om the surface to 800 m depth averaged
for the 137 ~ line from 2 ~ to 10 ~Table 1. Lag-Correlations Between Indian Monsoon Rainfall

and Sea Water Temperature Anomalies Averagedfor Surface
Mixing Layer (0 to 30 m) of the Western Pacific

Y(-1)

Y(O)

hI

/
/.)"

Y(+t)
200

Jan.

Jul.

Jan.

Jul.

Jan.

Jul.

-0.49

-0.61

-0.75

-0.22

-0.83

-0.10

i/

"I
l~ 400(more than 0.7), particularly between 150 ~ and180 ~ from the summer m o n s o o n season of Y(0)through the following winter. Over the eastern
Indian Ocean through the Indonesian maritime
continent (90 ~ to 130 ~in contrast, positive
correlations are seen. A similar correlation pattern, but with the opposite sign, can be seen in
the summer of Y ( - 1 ) through the winter ofY ( - 1 ) / Y ( 0 ) associated with the QBO nature of
the tropical atmosphere, though the pattern is lesssignificant. At 200 mb (Fig. 6 (b)), the overall correlation pattern is very similar to, or even more
systematic than, that of 700 mb except with the
opposite sign. Again, the m a x i m u m east-west contrast of correlation appears in the winter of Y(0)/Y ( + 1), when the absolute coefficient values exceeds 0.7 both in the western and the eastern Pacific. These timelag-correlation diagrams also
show the eastward propagation of correlation pattern from the Indian Ocean side toward the eastern
Pacific from Y(0) through Y ( + 1), which corresponds to the eastward propagation of the tropical

!
i

u.I

600

/

r

9%

e
800~

...
99%

900Lo l~'~l
-0.5

-0,3

J

l I

-0.1 0 0.1

0.3

(15

0.7 0.8

Fig. 4. Vertical profile of correlations between Indian monsoon rainfall and sea water temperature anomalies averaged
for July of Y(0) and January of ( + 1)

anomalous east-west circulation in the QBO timescale (Yasunari, 1985).
The anomalous east-west circulations suggested
above should be associated with the anomalous
large-scale convection over the Asian monsoon
region through the equatorial Pacific in the seasonal cycle starting from the Indian m o n s o o n season. This features is demonstrated in Fig. 7, where

34

T. Yasunari: Impact of Indian Monsoon on the Coupled Atmosphere/Ocean System
251

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is b

l""

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l

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20

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YEAR

Fig. 5. Time series of Indian monsoon rainfall anomalies (thick solid line) and sea water temperature anomalies at 20 m (thickdashed line) and 100 m (thin dashed line) depth averaged for the 137 ~ line (2 ~~ in the succeding January

the anomalous OLR and 700 mb wind fields areseasonally composited in reference to the strong
monsoon years. Since the OLR data is available
only after 1974, only 4 strong monsoon years(1975, 1978, 1981, 1983) are composited in thiscase. However, the anomalous wind field composited for 6 strong monsoon years (not shown)
appeared to be substantially the same pattern as
Fig. 7. In March-April-May (MAM) of the strong
summer monsoon year (Fig. 7 (a)), convection isrelatively strong in the central through the eastern
Pacific, whereas it is still weaker than normal in
the western Pacific. The anomalous wind fieldalong the equator is consistent with this anomalous OLR field. Thus, the large westerly anomalies
with anomalous convergence are noticeable in theeastern Pacific, whereas the anomalous divergenceis apparent near the positive anomalies of OLR
in the western Pacific. It is also of interest to notethat the OLR anomalies are positive along theSouth Pacific Convergence Zone (SPCZ).
In June-July-August (JJA) (Fig. 7 (b)), the OLR
field drastically changes from the previous season.
The negative anomalies (stronger convection) appear over India through southeast Asia, and thepositive anomalies (weaker convection) in thewestern Pacific become less significant and shift
their position more eastward. The area of negative
anomalies in the eastern Pacific is considerably
diminished. A notable change in the wind field isthe appearance of easterly anomalies over the central and the eastern Pacific. The anomalous south-

westerlies over the Arabian sea are manifested asstronger monsoon flows.
In September-October-November (SON) of thestrong ISM year (Fig. 7 (c)), the negative OLR
anomalies further develop and extend the areas
over southeast Asia throuth the equatorial western
Pacific, while the positive anomalies over the central/eastern Pacific also develop. Associated withthis east-west contrast of OLR anomalies is thefurther intensification of easterly anomalies along
the equatorial Pacific.
In December-January-February (DJF) (Fig.
7 (d)), the overall feature in the anomalous OLR
and wind field is very similar to that in SON, oris further developed. An area of negative OLR
anomalies is also developed over SPCZ and SouthAmerica, wshich suggest the intensified tropicaleast-west circulation along the Indian Ocean
through the South American sector. It should be
noted that at this stage the SST (SWT) anomalies
reach their maxima in the western Pacific while
they reach their minima in the eastern Pacific (Table 1 and Fig. 2). These anomalous patterns in
OLR and wind field established in SON throughDJF of strong ISM year continue through MAMof the post strong ISM year (Fig. 7 (e)), though
the anomalies are to some extent diminished as awhole. These patterns are, however, completely
collapsed or changed to the opposite pattern by
the start of the next Indian summer monsoon as
suggested in the change from Fig. 7 (a) to Fig.
7 (b). In the composites of weak ISM year, onlya

LAG-R BETHEEN [NOIA HONSOON RAINFALL AND 700M8 U
OW,..,

180

(IF

Y(-1)

u

y(+l)

0

60E

b

120E

180

120W

60W

LAG-R BETMEEN [NOIA HONSOON RAINFALL ANO 200it8 b

OE

180

OM,&N

JN

-_D_c
~,N

LIN

Y(o)

EC
a,N

JN

Y(+I)

EC
0

60E

1201-

180

120W

60W

0

Fig. 6. Time l a g - longitude section of correlation coefficients between Indian monsoon rainfall and zonal wind at (a) 700 mb
and (b) 200 mb along the equatorial belt (5 ~~ Units are 0.2. Negative values are shown with dashed lines. Positive(negative) values of more (less) than 0.4 ( - 0 . 4 ) (0.6 and - 0 . 6 for (b)) are dotted (shaded). Positive (negative) values of more(less) than 0.6 ( - 0 . 6 ) in (a) are dark shaded

va

STRONG

ISM

TEAR

MAN-O

(UVTOO)
5-OO

~-~:*

~,-'

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T SM

J

TEAR

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">~

I"

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oq~~., . .,._. .~(UVTO0)"%"- 7"--"~

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37

T. Yasunari: Impact of Indian Monsoon on the Coupled Atmosphere/Ocean System
2 . 0 m s -~$TRON6 l'O RE^g MONSOON

)

130

140E

150

160E

170

180

170

160W 150

140W

130

120W

110

lOOW

90

Fig. 8. Longitude-time section of anomalous SST and zonal wind at 700 mb along the equator composited for two years fromstrong ISM year to weak ISM year. Positive (negative) values in SST field are shown with solid (dashed) lines (unit: 2.0C)
and negative values are stippled

mirror images (i.e., very similar patterns but with
opposite signs both in O L R and wind field) of
Fig. 7 are found (not shown).
Thus, a stronger (weaker) than normal Indianm o n s o o n is most probably followed by stronger(weaker) easterly wind over the central Pacific andstronger (weaker) westerly wind to the west of130~ i.e., stronger (weaker) low-level convergence over the equatorial western Pacific during
the succeding a u t u m n and winter. This means thatthe anomalies of some atmospheric parameters
first appear over the Indian m o n s o o n region andmove towards the equatorial central Pacific,
through the seasonal migration of convection center from the norther summer to the northern winter and spring. This characteristic nature of theinterannual variability in the tropics was also substantiated by Meehl (1987).

6. SST/Zonal Wind Relations
Along the Equatorial PacificIn the previous sections, the evidence was shownthat the SWT anomalies in the western Pacific and
the SST anomalies in the whole tropical Pacificare well correlated with the preceding ISM activities. This correlation was also suggested to bephysically based on the characteristic nature ofanomalies in convection and associated circulation
field over the Asian m o n s o o n region through thetropical Pacific, which start from the northernsummer and persist up to the end of the succeding
winter or spring.
To further examine how the atmosphere andocean are coupled to each other along the equatorial Pacific associated with the ISM and the seasonal cycle, the zonal wind anomalies at 700 nab

Fig. 7. Composites of anomalous OLR and anomalous wind at 700 mb for (a) March-April-May, (b) June-July-August, (c)September-October-November, (d) December-January-February of strong ISM year and (e) March-April-May of the year
next. Positive (negative) values in OLR field are shown with solid (dashed) contours (unit: 2.0 W m-1) and negative values
less than -4.0 are stippled

38

T. Yasunari
140~ SST VS U700/t

/ ~ 3 ""
\ -O.5~-0.3--..
}

(i
I

I

r

I

(a)

~

-o.a.
1%.

i

I1

I

]

I

I

l

i

-7

170~ SST VS U700 ( b )
O

5
I

I

I

I

I

I

-7
-5

I

r

I

12o~

I

i

SST

I

VS

i

I

UZ00

(C)

x

3I
1730" 110 ~ 1;0 ~ 1;0 ~ 1;O~ 180 ~ 1;0~(~----0,3~ 1;0 ~ 140 ~ 130 ~ 120 ~ 110 ~ 1;0 ~ ;0 ~ 80"
8and SST anomalies are composited in reference
to the ISM activity. The important parameters forthe atmosphere/ocean interaction in the equatorial
Pacific are zonal wind and OHC of the surface
layers (Cane et al., 1986; Wyrtki, 1985; Yamagata
and Masumoto, 1989; etc.). Strictly speaking, themean SWT of the surface layer, instead of SST,
should be used as a measure of OHC in the oceanic
mixed layer. In the eastern Pacific, where themixed layer depth is far smaller and its variability
is far larger than those in the western Pacific, SST
anomaly is supposed to be a good measure of OHC
anomaly. The correlations between the observed
mean mixed layer temperature at 160~ and100 ~ along the equatorial belt adopted from
Meyers et al. (1988) and the SST at the same points
used here are certainly proved to be very high (e.g.,
0.89 and 0.93, with number of freedom of about
60). In the western Pacific, where the large OHC
is sustained by the combined effect of radiation
balance of the surface, ~orizontal advection anddown- or up-welling, it may generally be supposedthat SST anomaly does not necessarily represent
OHC anomaly. However, the SWT temperature
anomalies, which show nearly in-phase fluctuationfrom the surface to the bottom of the mixed layer

Fig. 9. Lag-correlations between
700 mb zonal wind anomaliesalong
the equator and SST anomalies at(a) 140~ (b) 170 ~(c) 120 ~at the equator. Contours of more(less) than 0.3 (-0.3) are shownwith solid (dashed) lines (unit: 0.1).
Negativelagsimplythat zonal windleads SST

as shown in Fig. 3, strongly suggest that even in
the western Pacific the SST anomaly may wellreflect the OHC anomaly as far as a proper spatial
and time average is applied. We have tentatively
concluded, therefore, that SST anomaly adopted
here is a good measure of OHC anomaly in thesurface mixed layer.
Figure 8 shows a longitude-time section of thesetwo parameters composited for two years, by using
6 cases starting from the strong ISM years (i.e.,
1971, 1973, 1975, 1978, 1981 and 1983). Note thatin these cases the second years all correspond to
weak ISM years including ENSO years. Therefore,
the biennial cycles are particularly strong. At
nearly the same time when the strong ISM starts,
easterly anomalies develop over the whole Pacificbasin, particularly in the western Pacific to thewest of the dateline. The development of easterly
anomalies in the western Pacific from an ISM
season to the succeding autumn (September
through November) is followed by large negative
SST anomalies in the central/eastern Pacific andlarge positive SST anomalies in the western Pacificin the succeding winter. On the other hand, thesenegative SST anomalies in the central/eastern Pacific in October through January are followed byImpact of Indian Monsoon on the Coupled Atmosphere/OceanSystem
the easterly wind maxima in late winter throughthe following spring (January through April). In
the second year the change from easterly to westerly anomalies seems to be associated with the startof weak ISM. The westerly anomalies are significant again over the longitudes 150 ~ to 160 ~in the western Pacific. The composites startingfrom the weak ISM years (not reproduced here)
show just the opposite features to Fig. 8, as iseasily supposed from the anomalies of the second
year (i.e., weak monsoon year) in Fig. 8.
The relations between zonal wind anomalies
and SST evolutions along the equator are moreclearly demonstrated in the lag-correlation diagrams between zonal wind anomaly of each longitude in the Pacific basin and SST anomalies ofsome specified longitudes, as shown in Fig. 9. It
is noteworthy to state that not only in the western
Pacific, but also in the central/eastern Pacific, SST
is significantly correlated to the zonal wind between 150 ~ and 160 ~ in the western Pacific witha few or several months lag. The signs of correlations are opposite to each other. That meansthat strong easterly anomalies over these longitudes are largely responsible for producing thepositive (negative) SST anomalies to the west
(east) several months later, and vice versa. It
should be noted here that the correlation of zonal
wind with ISM is highest at these longitudes, as
shown in Fig. 6. In other words, the anomalous
ISM activity is closely associated with zonal wind
anomalies over this region possibly through thetropical east-west circulation, which may, in turn,produce SST anomalies with some time lag. We
may emphasize here that the wind field of theselongitudes leads the SST field over the whole equatorial Pacific.
In the central/eastern Pacific (Fig. 9 (b) and9 (c)), the secondary maxima in lag-correlations
are also noted, showing that SST leads the zonal
wind with a few months. This suggests that thepositive feedback process that the SST anomalies,
once produced by the wind anomalies, further intensify the anomalous wind field has an additionalrole on the evolution of the coupled atmosphere/
ocean system, particularly during the maturephase of warm or cold events there (Fig. 8).
7. Summary and Remarks
Lag correlation analyses between the ISM rainfall
and SST or SWT anomaly in the tropical Pacific

39

show that the interannual fluctuation of the ISM
activity with a QBO nature strongly controls the
OHC anomaly of the oceanic mixed layer in thetropical Pacific, particularly in the western Pacific.
Lag correlations between ISM rainfall anomaly
and zonal wind anomaly in the lower troposphere,
and that between zonal wind anomaly and SST
anomaly along the equator strongly suggest that
the OHC anomalies are formed by the anomalous
easterly or westerly surface wind stress particularlyin the western Pacific (150 ~ to 160 ~ Very recently, Kutsuwada (1988) also found that the QBO
mode in the surface zonal wind stress field is particularly strong over the tropical western Pacific,
even compared to the ENSO mode with about 4year period.
Generally speaking, lag correlations do not necessarily imply immediate cause and effect relations, particularly in case the correlations show
cyclic nature as shown in Fig. 1. However, all thestatistical results integrated here can be interpreted
consistently with each other, as manifestations oftime sequences of the some physical processes. For
example, both SST and zonal wind at the equatorial Pacific show remarkably large lagged correlations, compared to lead correlations, with the
ISM rainfall, as shown in Fig. 2 and Fig. 6. The
time lags between the extreme phases of these twoparameters suggested from these correlations are
consistent with the systematic lags as shown in
Fig. 9. In such a case, it may be reasonable topostulate that in reality the SST and the SWT
anomalies in winter may be more directly (or physically) linked to the ISM rainfall anomaly in thepreceding summer, rather than in the succedingsummer, by means of the persistent anomalous
wind field. The seasonal evolutions of anomalous
wind and OLR field composited for strong (weak)
ISM years (Fig. 7) further substantiate the evidence that these wind anomalies over the tropical
Pacific are linked with the anomalous convection
over the western Pacific which starts from ISM
season and persists through the following winterand spring.
The scenario of the ENSO cycle derived here
may be very different from that deduced fromsome coupled atmosphere/ocean models such as
Zebiak and Cane (1987), Battisti (1988) among
others, which are based solely on the physics ofatmosphere/ocean interaction over the Pacific basin. In these models the ENSO cycle depends on40

T. Yasunari

flowing back of warm (or cold) water westward
along the equatorial or off-equatorial latitudes
through the western boundary reflection of
Rossby waves during the cold (or warm) phase(Cane et al., 1986), where the OHC anomaly fundamentally leads the wind stress anomaly in theequatorial Pacific. These models also suggest thatthe spreading of positive OHC anomaly in thewhole equatorial Pacific basin is important for thepreconditioning of the ENSO event, or in otherwords, the triggering of the unstable ocean/atmosphere coupled mode. In the real system, however, the wind stress anomaly in the western Pacificseems to be essential for producing the OHC
anomaly later on in the whole equatorial Pacific,
as shown in Fig. 8 and Fig. 9. By using a coupledatmosphere/ocean model, Yamagata and Masumoto (1989) demonstrated that once this condition (or preconditioning) is established the ENSO
event evolves automatically. They also suggested
that the predictability skill of the ENSO shownby Cane etal. (1986) is essentially based on the
instability of this preconditioned state,
We should comment here that although individual ENSO years appear every 3 to 6 years, they
are all phase-locked with the phases of the QBO
in the SWT anomalies in the western Pacific. For
example, as shown in Fig. 3, the first half of eachcalender year of ENSO events (1986, 1972, 1976,1982) corresponds to the transition phase fromthe maximum to the minimum in the SWT and
SST anomalies in the western Pacific. In 1974 and1979, however, although the same transitions occurred in the western Pacific (Fig. 3), E1 Nifio
events in the eastern Pacific did not occur. In otherwords, the variation of OHC anomaly in the western tropical Pacific during winter and spring,which seems to be modulated by the preceding
Asian monsoon activity via the anomalous surface
wind field, may provide a necessary or at least avery favorable dynamical condition for triggering
ENSO events in the eastern Pacific. Masumoto
and Yamagata (1989) have recently shown in the
coupled model that both the OHC anomaly andthe strong westerly wind anomaly in the western
Pacific are very essential to produce the preconditioning mentioned above.
In fact, the basic time scale of ENSO cycle may
be decided by the heat storage of the whole equatorial Pacific basin and the accumulation rate ofwarm water in the western Pacific as noted by

Wyrtki (1985). The present results strongly suggestthat the accumulation rate of warm water and also
the timing of the preconditioning are modulated
fundamentally by the anomalous surface windfield over the central/western Pacific. This windfield may be part of the coupled land/atmosphere/
ocean system involving the Asian monsoon system, rather than of the coupled atmosphere/oceansystem over the equatorial Pacific.
An important problem now reserved for futureis the mechanism of the QBO in this coupled land/
atmosphere/ocean system, which seems to be "intrinsic" to the tropics. The amplitude modulationof this oscillation may actually controls the frequency and intensity of so-called ,,westerly burst"
over the western/central Pacific, which seems to
be a key factor for triggering the unstable ocean/
atmosphere coupled mode. Further analysis forthese problems is currently in progress using longterm global data sets.
Acknowledgements

The author would like to thank Dr. T. Yamagata for his
helpful commentsand discussion. He is also grateful to Dr.
K. Kurihara of JMA for providing the SWT data, He is also
acknowledgedto Dr. R. Kawamura and Mr. S.-F. Tien forcomputing assistance. Dr. S. Ozaki drafted the figures. This
researchwas financiallysupported by Grant-In-Aidfrom the
Ministry of Education for 1987, 1988 (No. 62540298) and1989 (No. 01460055).
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Authors address: Tetsuzo Yasunari, Institute of Geoscience,
University of Tsukuba, Tsukuba, Ibaraki 305, Japan.