relative contributions of the indian ocean and local...

Relative Contributions of the Indian Ocean and Local SST Anomalies tothe Maintenance of the Western North Pacific Anomalous Anticyclone

during the El Nino Decaying Summer*

BO WU

LASG, Institute of Atmospheric Physics, Chinese Academy of Sciences, and Graduate University of Chinese

Academy of Sciences, Beijing, China

TIM LI

IPRC, and Department of Meteorology, University of Hawaii at Manoa, Honolulu, Hawaii

TIANJUN ZHOU

LASG, Institute of Atmospheric Physics, Chinese Academy of Sciences, Beijing, China

(Manuscript received 19 June 2009, in final form 15 January 2010)

ABSTRACT

To investigate the relative role of the cold SST anomaly (SSTA) in the western North Pacific (WNP) or

Indian Ocean basin mode (IOBM) in maintaining an anomalous anticyclone over the western North Pacific

(WNPAC) during the El Nino decaying summer, a suite of numerical experiments is performed using an

atmospheric general circulation model, ECHAM4. In sensitive experiments, the El Nino composite SSTA is

specified in either the WNP or the tropical Indian Ocean, while the climatological SST is specified elsewhere.

The results indicate that the WNPAC is maintained by the combined effects of the local forcing of the

negative SSTA in the WNP and the remote forcing from the IOBM. The former (latter) contribution grad-

ually weakens (enhances) from June to August. The negative SSTA in the WNP is crucial for the maintenance

of the WNPAC in early summer. However, because of a negative air–sea feedback, the negative SSTA

gradually decays, as does the local forcing effect. Enhanced local convection associated with the IOBM

stimulates atmospheric Kelvin waves over the equatorial western Pacific. The impact of the Kelvin waves on

the WNP circulation depends on the formation of the climatological WNP monsoon trough, which does not

fully establish until late summer. Therefore, the IOBM plays a crucial role in late summer via the Kelvin wave

induced anticyclonic shear and boundary layer divergence.

1. Introduction

The most pronounced low-level circulation anomalies

over the tropical Pacific during El Nino decaying sum-

mer is an anomalous anticyclone over the western North

Pacific (WNPAC; see review papers by Li and Wang 2005;

Lau and Wang 2006). The WNPAC, persisting from the

El Nino mature winter to the subsequent summer, plays

a crucial role in the El Nino–East Asian summer monsoon

teleconnection (Zhang et al. 1996; Wang et al. 2000; Chang

et al. 2000a,b; Lau and Nath 2000; Wang and Zhang 2002;

Wu et al. 2003; Li et al. 2007; Zhou et al. 2009a).

During the El Nino mature winter and the following

spring, the WNPAC would enhance the precipitation

over southeastern China through anomalous moisture

transport (Zhang and Sumi 2002). During the El Nino

decaying summer, the WNPAC would enhance the mei-

yu–baiu precipitation by modulating the western Pacific

subtropical high (Chang et al. 2000a). In addition, the

WNPAC forced by negative precipitation anomalies

over the Philippine Sea is a part of meridional wave train

propagating northward from the Philippine Sea, which

* School of Ocean and Earth Science and Technology Contri-

bution Number 7908 and International Pacific Research Center

Contribution Number 684.

Corresponding author address: Dr. Tim Li, International Pacific

Research Center and Department of Meteorology, School of

Ocean and Earth Science and Technology, University of Hawaii at

Manoa, Honolulu, HI 96822.

E-mail: [email protected]

2974 J O U R N A L O F C L I M A T E VOLUME 23

DOI: 10.1175/2010JCLI3300.1

� 2010 American Meteorological Society

resembles Pacific–Japan (PJ) teleconnection pattern (Nitta

1987, Huang and Sun 1992; Kosaka and Nakamura 2006;

Wu et al. 2009b).

The maintenance of the WNPAC from the El Nino

mature winter to the subsequent spring is attributed to

a positive thermodynamic air–sea feedback (Wang et al.

2003). A cold sea surface temperature anomaly (SSTA)

in the western North Pacific (WNP) suppresses local

convection (Su et al. 2000), which stimulates a low-level

atmospheric anticyclone anomaly to its west. The north-

easterly anomalies to the eastern flank of the WNPAC

enhance the mean trade wind and cool the in situ SST

through enhanced upward surface latent heat flux (Wang

et al. 2000).

With the onset of the WNP summer monsoon, the mean

southwesterly monsoon circulation replaces the northeast-

erly trade wind. As a result, a negative evaporation–wind–

SST feedback appears in the region. The negative air–sea

feedback weakens the negative SSTA in the WNP dur-

ing the El Nino decaying summer (Chou et al. 2009). In

contrast to the weakening of the local negative SSTA,

the WNPAC is strengthened from the preceding spring

to the summer (Wu et al. 2009b).

A possible cause of strengthening of the WNPAC dur-

ing the El Nino decaying summer is attributed to the re-

mote SSTA forcing from the tropical Indian Ocean (TIO;

see Yang et al. 2007; Li et al. 2008; Wu et al. 2009b; Xie

et al. 2009). An Indian Ocean basin mode (IOBM) es-

tablishes after the ENSO mature winter. The basinwide

warming in the TIO is possibly caused by anomalous

surface heat flux associated with the descending branch

of the anomalous Walker circulation (Klein et al. 1999;

Venzke et al. 2000; Lau and Nath 2003) or the change of

tropical tropospheric temperature (Chiang and Sobel

2002). In addition, ocean dynamic processes also play

a role (Li et al. 2002, 2003). The IOBM may strengthen

the WNPAC in boreal summer through the following

two processes: First, through enhanced convection over

the TIO, the IOBM may stimulate a Kelvin wave–type

easterly response in the western Pacific. The anticyclonic

shear vorticity associated with the easterly anomalies

may weaken the WNP summer monsoon through Ekman

pumping divergences (Wu et al. 2009b; Xie et al. 2009).

Second, the IOBM may increase the surface moisture

and thus enhance the low-level moisture transport and

convection over the Maritime Continent, the latter of

which may further induce the subsidence over the WNP

through an anomalous Hadley circulation (Chang and Li

2000; Li et al. 2001; Sui et al. 2007; Wu and Zhou 2008;

Wu et al. 2009b).

Both the local forcing of the negative SSTA in the

WNP and the remote SSTA forcing from the TIO may

impact the WNPAC during the El Nino decaying summer.

However, their relative contributions are still not clear.

The clarification of this issue may help us to identify better

predictors for climate prediction. Through just an obser-

vational analysis, it is difficult to indentify the relative

contributions of the local and remote forcing. In the study,

our strategy is to examine the response of an atmospheric

general circulation model (AGCM) to specified regional

SSTA forcing.

The rest of the paper is organized as follow. A de-

scription of the model and data is provided in section 2.

Section 3 reveals the essential characteristics of the at-

mospheric circulation, precipitation, and surface flux

anomalies during the El Nino decaying summer through

observational diagnosis. The results of numerical exper-

iments are analyzed in section 4, in which we focus on the

relative contribution of the TIO and WNP in maintaining

the WNPAC. Section 5 summarizes our findings.

2. Model and data description

The AGCM used in the study is ECHAM version 4.6

(hereafter ECHAM4), which was developed by the Max

Planck Institute for Meteorology (MPI; Roeckner et al.

1996). The model was run at a horizontal resolution

of spectral triangular 42 (T42), roughly equivalent to

2.8 latitude 3 2.8 longitude, with 19 vertical levels in

a hybrid sigma–pressure coordinate system extending

from surface to 10 hPa. The mass flux scheme of Tiedtke

(1989) is used to represent the deep, shallow, and middle

convection. The original moisture convergence closure

has been replaced by a convective available potential

energy (CAPE)-based closure scheme (Nordeng 1994).

Previous evaluation shows that ECHAM4 is one of the

best AGCMs in the simulation of the Asian–Australian

monsoon (Zhou et al. 2009a), and it thereby has been

used in many modeling studies of monsoon variability

(Fu et al. 2008; Li et al. 2008; Zhou et al. 2009b).

The SST data constructed for phase II of the Atmo-

spheric Model Intercomparison Project (AMIP II) are

used as the model lower boundary condition and for the

observation analysis (Hurrell et al. 2008). The dataset

is a merged product based on the Hadley Centre sea

ice and SST dataset version 1 (Rayner et al. 2003) and

version 2 of the National Oceanic and Atmospheric Ad-

ministration (NOAA) optimum interpolation SST analy-

sis (Reynolds and Smith 1994).

Other observational datasets used in the study include

1) the observational precipitation field derived from the

Climate Prediction Center (CPC) Merged Analysis of

Precipitation (CMAP) for the period of 1979–2007 (Xie

and Arkin 1997); 2) National Centers for Environment

Prediction (NCEP)/Department of Energy (DOE) AMIP

II Reanalysis (Kanamitsu et al. 2002) for the period of

1 JUNE 2010 W U E T A L . 2975

1979–2007; 3) objectively analyzed air–sea fluxes (OAFlux)

for global oceans (Yu et al. 2008), with latent heat flux

from 1982 to 2004 and shortwave radiative flux from

1984 to 2004. To focus on the interannual time scale,

variations longer than 8 yr are filtered out from the orig-

inal datasets with a Lanczos filter (Duchon 1979).

In the paper, all of the observational analysis and

numerical experiments are based on five strong El Nino

events (i.e., the 1982, 1991, 1994, 1997, and 2004 events)

selected from the period of 1979–2008. In the period, the

rest of two strong events, the 1986 and 2002 events, are

excluded for following reasons. The 1986 event persists

for 2 yr, which is much longer than the normal El Nino

life cycle. During its decaying summer (the 1988 sum-

mer), a cold SSTA, rather than a warm SSTA, covers

large area of the TIO, so that the in situ precipitation is

suppressed, rather than enhanced as conventional El Nino

events (figure not shown). For the 2002 events, a signifi-

cant Indian Ocean dipole pattern, rather than an IOBM,

is seen in the TIO (figure not shown) during its decaying

summer. To enhance the Indian Ocean and local SSTA

forcing signals, we exclude the two ‘‘unconventional’’

El Nino events.

3. Observational analysis

The composite SSTA evolution patterns during the

El Nino decaying summer are shown in Fig. 1. The TIO

is controlled by the IOBM throughout the summer. The

largest warm SSTA appears in the southeastern Indian

Ocean and the tropical western Indian Ocean. The am-

plitude of the warm SSTA in the TIO slightly weakens

from June to August. Over the WNP, a significant neg-

ative SSTA occurs in June, and it gradually attenuates

from west to east and weakens with time (Fig. 1).

Corresponding to the IOBM during the El Nino

decaying summer, convection over the TIO is enhanced

(Figs. 2a–c). The maximum precipitation anomalies are

located in the northern and southeastern Indian Ocean.

In contrast, negative precipitation anomalies persistently

control the WNP throughout the summer. The negative

precipitation center slowly moves eastward from June to

July. As a Rossby wave response to the negative heating,

a large-scale anomalous anticyclone (WNPAC) extending

from 1108 to 1708E maintains over the WNP throughout

the summer (Figs. 2d–f). Although the WNPAC persists

throughout the summer, its impact on the East Asian

summer monsoon is the most significant in June. A posi-

tive anomalous rainband extends from central China to

southern Japan (Fig. 2a), indicating that the mei-yu–baiu

rainband is significantly enhanced. The temporal depen-

dence of the WNPAC impact is associated with the north-

ward movement of the East Asian climatological rainfall

band (Tao and Chen 1987).

The WNPAC and associated negative precipitation

anomalies have a strong feedback to the underlying SST

(Chou et al. 2009; Wu et al. 2009a). Figure 3 shows the

June–August mean downward shortwave radiative and

downward latent heat flux anomalies for the El Nino

composite. They indicate that the decay of the negative

SSTA in the WNP is caused by the combined effects of

increased downward shortwave radiative and downward

latent heat flux anomalies. The increased shortwave

radiation is closely associated with the negative in situ

precipitation anomalies for their analogous spatial pat-

terns throughout the summer (Figs. 3a,b). The increase

of the downward latent heat flux is caused by the east-

erly anomalies in the southern flank of the WNPAC,

which decrease the southwesterly wind speed of the WNP

monsoon. Because the monsoon southwesterly flow is es-

tablished from west to east, the positive downward latent

FIG. 1. Observed monthly mean SST anomalies (8C) from June

to August for the El Nino composite. The solid lines represent the

5% significant level.


heat flux anomalies in the Philippine Sea gradually in-

tensify and expand eastward (Figs. 3d–f). It is worthy to

note that even in the presence of the negative thermo-

dynamic air–sea feedback, the amplitude of the negative

SSTA in the WNP is still quite large in early summer. It

is expected that such a strong SSTA may have a signifi-

cant impact on the local circulation.

4. Numerical experiments

a. Experiment design

To reveal the relative contribution from the TIO and

WNP SSTA forcing, we design idealized numerical ex-

periments in which we separate the anomalous SST

forcing from the TIO and WNP. Figure 4 shows the model

geographic locations for the TIO and WNP domains.

Table 1 lists four sets of designed numerical experiments.

The first is a control run, in which ECAHM4 was in-

tegrated for 20 yr, forced by monthly climatological SST.

The climatological SST is based on the period of 1979–

2004. This experiment is referred to as CTRL run. To

investigate the relative contributions of the remote TIO

forcing and the local WNP forcing to the WNPAC during

the El Nino decaying summer, three sets of sensitivity

experiments are designed, in which the composite SSTA

shown in Fig. 1 was added to the climatological SST in

the global ocean (hereafter the GB run), the TIO only

(hereafter the TIO run), or the WNP only (hereafter the

WNP run). Note that in the TIO (WNP) run the SST cli-

matology was prescribed in the regions outside of the TIO

(WNP). Each experiment was integrated from April to

August, with 20 ensemble members. The ensemble mean

results of June–August are analyzed.

FIG. 2. Observed composite patterns of (left) monthly mean precipitation anomalies (mm day21) and (right)

850-hPa wind anomalies (m s21) from June to August during the El Nino decaying summer. The contour values

for the precipitation anomaly are 61, 62.5, and 64. The shading denotes the 5% significant level for the (left)

precipitation and (right) zonal wind.

1 JUNE 2010 W U E T A L . 2977

The difference between the GB and CTRL runs pro-

vides an evaluation on the performance of the model in

reproducing the WNPAC. The difference between the

TIO and CTRL runs examines the sole contribution of

the remote IOBM forcing to the WNPAC, and the dif-

ference between the WNP and CTRL runs examines the

sole contribution of the local WNP negative SSTA forc-

ing. A Student’s t test is used to examine the statistical

significance of model response.

b. Model climatology

Because of the close relationship between the WNP

circulation anomaly and the climatological circulation

(Chou et al. 2009; Wu et al. 2009b), it is necessary to

evaluate the model’s performance in the climatology sim-

ulation. The simulations of June–August precipitation and

850-hPa wind fields derived from the CTRL run are shown

in Fig. 5, along with the observed counterparts. In boreal

summer, the mean precipitation centers are located in the

southeastern Indian Ocean, the Arabian Sea, the Bay of

Bengal, the South China Sea, and the WNP, respectively.

Correspondingly, strong low-level winds blow from the

Southern Hemisphere to the Northern Hemisphere and

from the Arabian Sea toward the Philippine Sea. In June,

the westerlies are confined to the west of the South China

Sea, and then they gradually extend eastward as the

summer progresses. The rapid eastward expansion of the

westerlies in late summer is accompanied by an enhanced

convection associated with the onset of the WNP mon-

soon. Over the East Asia, the most remarkable feature is

a mei-yu–baiu rainband extending from central China to

south of Japan in June.

The pattern correlations between the simulated and

observed climatological precipitation fields in the region

FIG. 3. Observed composite patterns of (left) monthly mean downward shortwave radiative flux anomalies (W m22)

and (right) downward latent heat flux anomalies (W m22) from June to August during the El Nino decaying summer.

Because of the limitation of data length, only (left) 1991, 1994, and 1997 events and (right) 1982, 1991, 1994, and 1997

events are used.


of 208S–408N, 408–1808E are calculated. The correlation

coefficients for June, July, and August reach 0.65, 0.68, and

0.69, respectively. ECHAM4 reproduces the major fea-

tures of the Asian summer monsoon (Figs. 5d–f), such as

the precipitation centers over the Arabian Sea and the Bay

of Bengal, the Somalia cross-equatorial jet, and westerlies

blowing from the Arabian Sea to the Bay of Bengal. The

model also reasonably reproduces the WNP monsoon

trough and its buildup and eastward advance from June to

August. The main deficiency is that the simulated mon-

soon trough shifts northward by about 58 latitude and

extends eastward excessively. The simulated WNP rain-

band is stronger than the observation. The model also fails

to simulate the mei-yu–baiu rainband. The discrepancies

of the simulated WNP monsoon trough and mei-yu–baiu

rainband have been common problems for many AGCMs

(Lau et al. 1996; Kang et al. 2002).

c. Response of the WNPAC to remote and local SSTAforcing

The precipitation and 850-hPa wind anomalies simu-

lated by the GB run are shown in Fig. 6. The GB run

realistically reproduces low-level circulation anomalies

in the WNP from June to August. Some key character-

istics of the stimulated WNPAC, such as zonal extension,

pronounced easterly anomalies in its southern flank, and

eastward migration of the anticyclone center from June to

August, resemble the observation. The associated negative

precipitation anomalies in the WNP are also reasonably

simulated. The major difference between the GB run

and the observation is that the simulated WNPAC and

negative precipitation anomalies are slightly stronger than

those in the observation. Compared with the observation,

the negative precipitation anomalies shift northward

slightly, and unrealistic positive precipitation anomalies

appear in the equatorial WNP. The discrepancies may

be associated with the northward shift of the climato-

logical rainband over the WNP.

In the GB run, the TIO is dominated by positive pre-

cipitation anomalies over the tropical southeastern and

northwestern Indian Ocean. The amplitude of the posi-

tive precipitation anomalies gradually intensifies from

June to August, consistent with the observation. Nega-

tive precipitation anomalies and biased atmospheric cir-

culation anomalies appear in some part of the TIO in

June and August. However, they are statistically insignif-

icant. The net effect of the anomalous heating over the

TIO is enhanced convection as anomalous low-level winds

to its east converge into the positive heating region.

Although the simulated WNPAC resembles the obser-

vation, the simulated mei-yu–baiu rainfall anomalies are

much weaker than those observed (Fig. 6a). The discrep-

ancy is associated with the model climatology, that is, the

model fails to reproduce the climatological mei-yu–baiu

rainband (Fig. 5).

Corresponding to the reasonable precipitation and

low-level wind anomalies over the WNP, the model re-

produces the major characteristics of the surface flux

anomalies (Fig. 7). The positive downward shortwave

radiative flux anomaly caused by the negative precipi-

tation anomaly is dominant throughout the summer, al-

though its meridional extent is narrower and its strength

is greater than that observed (Figs. 7a–c). Meanwhile, the

model reproduces the eastward extension of the positive

downward latent heat flux anomalies associated with the

gradually establishment of the WNP summer monsoon

trough (Figs. 7d–f).

FIG. 4. Geographic domains for the tropical Indian Ocean

and western North Pacific, where the SST anomalies are specified

as the model lower boundary condition for the WNP and TIO

runs.

TABLE 1. List of numerical experiments.

Experiment SST forcing field Integration

Control run (CTRL) Global climatological SST 20 yr

Global SSTA forcing (GB) Add the composite SSTA to climatological

SST in the global ocean

20 realizations

Tropical Indian Ocean

forcing (TIO)

Add the composite SSTA to climatological

SST in the tropical Indian Ocean only

20 realizations

Western North Pacific

forcing (WNP)

Add the composite SSTA to climatological

SST in the western North Pacific only

20 realizations

1 JUNE 2010 W U E T A L . 2979

The reasonable responses of the GB run in repro-

ducing the WNPAC give us confidence to further in-

vestigate the relative roles of the remote IOBM forcing

and the local WNP forcing through comparing the re-

sults of the WNP and TIO runs with that of the GB run.

The precipitation and 850-hPa wind anomalies simu-

lated by the WNP run are shown in Fig. 8. The WNPAC

and local negative precipitation anomalies are repro-

duced throughout the summer. In June, the patterns of

the WNPAC and negative precipitation anomalies in

the WNP run resemble those in the GB run (Figs. 6a,d),

with weaker amplitudes. In July, the WNPAC and the

local negative precipitation anomalies further weaken.

In August, a positive precipitation anomaly appears over

the Philippine Sea in response to the underlying warm

SSTA. This causes the eastward shift of the WNPAC by

more than 108 longitude, relative to that in the GB run.

For all 3 months, easterly anomalies in the southern

flank of the WNPAC simulated by the GB run extend to

the west of the South China Sea (cf. Fig. 6), whereas the

easterly anomalies in the WNP run are confined to the

east of the Philippine Sea. The difference between the WNP

and GB runs is mainly attributed to the lack of the re-

mote forcing from the TIO.

The precipitation and 850-hPa wind anomalies in the

TIO run are shown in Fig. 9. The WNPAC is reproduced

throughout the summer. Compared with the WNP run,

the WNPAC and the associated negative precipitation

anomalies in the TIO run are clearly shifted westward

and trapped in the South China Sea and Philippine Sea,

particular in June and July. The intensity of the WNPAC

gradually intensifies from June to August. In June, east-

erly anomalies appear in the tropical Pacific, but the an-

ticyclonic vorticity is weak and confined to the South

China Sea. In July and August, the WNPAC significantly

strengthens and expands eastward. The eastward expan-

sion of the WNPAC is accompanied by the establishment

of the WNP monsoon trough (Fig. 5).

The mechanism through which the IOBM impacts the

WNP circulation anomaly during the El Nino decaying

summer was discussed in Wu et al. (2009b). Enhanced

convective heating anomalies in the TIO stimulate Kelvin

FIG. 5. Climatological June–August precipitation (mm day21) and 850-hPa wind (m s21) fields derived from the

(left) observation and (right) ECHAM4 CTRL run.


wave–type easterly anomalies in the western Pacific. The

anticyclonic shear of the easterly anomalies causes a bound-

ary layer divergence in the off-equatorial WNP through

an Ekman pumping processes. The extent to which the

boundary layer divergence affects the local precipitation

anomaly depends on its influences on the mean state in

the WNP. In June, the WNP monsoon trough has not

fully established. As a result, the divergence has a limited

impact on the precipitation and wind anomalies over the

WNP. In contrast, in July and August, the WNP monsoon

trough fully establishes. The Kelvin wave–induced bound-

ary layer divergence suppresses the mean convection,

leading to a significant negative heating anomaly. The

negative heating anomaly further stimulates an atmo-

spheric Rossby wave response, and thus a low-level

anomalous anticyclone in the WNP.

The above idealized numerical experiments demon-

strate that both the remote IOBM forcing and the local

WNP SSTA forcing contribute to the maintenance of

the WNPAC during the El Nino decaying summer. To

reveal the relative contribution of the remote and local

forcing, we calculate the area-averaged vorticity anom-

aly over the region of 108–358N, 1208–1608E from June to

August, simulated by the WNP and TIO runs. The re-

sults are shown in the Fig. 10. While the seasonally av-

eraged voricity intensities are comparable in the two

runs, the subseasonal evolutions are different. The local

SSTA forcing leads to the strongest WNPAC response

in June, and the response weakens in July and even turns

to a cyclone anomaly in August. The remote TIO forc-

ing, on the other hand, is relatively weak in early sum-

mer and strengthens from June to August. While the

FIG. 6. Difference between the (left) precipitation (mm day21) and (right) 850-hPa wind (m s21) between the

GB and CTRL runs. The contour values at left panel are 61, 62.5, and 64. Solid (dashed) lines denote posi-

tive (negative) values. The shading represents the 5% significant level for the (left) precipitation and (right) zonal

wind.

1 JUNE 2010 W U E T A L . 2981

weakening of the local response is primarily attributed

to the weakening of the local SSTA because of the neg-

ative thermodynamic air–sea feedback, the strengthening

of the remote response is caused by the evolution of the

mean flow, even though the IOBM weakens slightly from

early to late summer.

5. Conclusions and discussion

a. Conclusions

Some recent studies (e.g., Xie et al. 2009) stressed the

effect of the remote forcing from the tropical Indian

Ocean in maintaining the WNPAC during the El Nino

decaying summer, by arguing that the seasonal mean

SSTA is negligibly small in the WNP resulting from the

negative in situ evaporation–wind–SST feedback. Here,

with the aid of a set of numerical experiments, we dem-

onstrate that both the negative SSTA in the WNP and the

IOBM contribute to the maintenance of the WNPAC.

When solely forced by the SSTA in the western North

Pacific (WNP run) or the tropical Indian Ocean (TIO run),

the WNPAC is somewhat reproduced, but its amplitude is

weaker than that forced by the global SSTA (GB run).

The WNPAC that is simulated by the WNP and TIO

runs shows distinctive evolution characteristics. In the

WNP run, the WNPAC is forced by the local SSTA. A

significant negative SSTA appears in the WNP prior to

the El Nino decaying summer because of a positive

thermodynamic air–sea feedback (Wang et al. 2003). As

the summer comes, the reversal of the mean wind leads

to a negative air–sea feedback. Despite of the negative

feedback, the cold SSTA in the WNP is still quite sig-

nificant in June and July (Fig. 1) and impacts the local in

situ circulation anomaly. The decay of the local SSTA

forcing leads to the weakening of the simulated WNPAC

from June to August.

Different from the WNP run, the evolution of the

WNPAC in the TIO run is associated with the change of

FIG. 7. Same as Fig. 6, but for the (left) downward shortwave radiative flux and (right) downward latent heat flux. It is

worth noting that the contour values of the model results are larger than that in the observation (Fig. 3).


the background mean circulation. Whereas the ampli-

tude of the IOBM weakens slightly from June to August,

the simulated WNPAC significantly intensifies and ex-

pands eastward. The intensification and expansion are

consistent with the establishment of the climatological

WNP monsoon trough. As a Kelvin wave response to

positive heating anomalies in the TIO, easterly anoma-

lies in the western Pacific cause an anticyclonic shear,

and thus a boundary layer divergence over the WNP

through Ekman pumping (Wu et al. 2009b). The bound-

ary layer divergence leads to a greater negative pre-

cipitation anomaly response in late summer (July and

August), when the monsoon trough is fully established.

This explains why the remote Indian Ocean forcing effect

is strengthened from June to August even though the

IOBM weakens.

In summary, the WNPAC during the El Nino decay-

ing summer is attributed to both the remote and local

SSTA forcing. In early summer, the WNPAC is primarily

supported by the local negative SSTA, whereas in late

summer, the IOBM plays a more important role in

maintaining the WNPAC.

b. Discussion

Although the results above are robust based on the

ECHAM4 numerical experiments, the quantitative esti-

mation of relative contributions of the local and remote

SSTA effects on the WNPAC might be model dependent.

It is worth noting that the strengths of the WNP cold

SSTA and the IOBM differ in each El Nino event. The

contributions shown in the ideal experiments may be

model dependent, especially under the condition of biased

model climatology. Also, in reality there are two-way

interactions between the WNP and TIO SST/circulation

anomalies; as a result, it is difficult to assess the pure ef-

fect of the TIO SSTA on the WNP circulation anomaly.

Although both the IOBM and the WNP cold SSTA

contribute to the WNPAC, from a prediction point of

FIG. 8. Same as Fig. 6, but for the difference between the WNP and CTRL runs.

1 JUNE 2010 W U E T A L . 2983

view, the latter may be a better predictor for the East

Asian summer monsoon. This is because the WNPAC

exerts a dominant impact on the East Asian summer

monsoon in June, when the climatological mei-yu–baiu

rainband is the strongest. Therefore, the WNP negative

SSTA, which dominates the WNPAC in June, may de-

serve more attention for the seasonal predication of the

East Asian summer monsoon.

As noted in the introduction, the PJ pattern is pri-

marily excited by the convection anomalies over the

Philippine Sea. Because the convection anomalies re-

sult from both the local SSTA and the TIO remote

forcing during the El Nino decaying summer, it is in-

teresting to investigate their relative roles. Figure 11

shows the 200-hPa voritcity anomalies in the observa-

tion and the GB, TIO, and WNP simulations. It is clear

that the GB run can simulate the well-organized me-

ridional wave train propagating from the Philippine Sea,

which is analogous to that in the observation. However,

the voriticy anomalies in the TIO and WNP runs are

weaker and loosely organized, suggesting that the forma-

tion and maintenance of the PJ patter may rely on both the

remote and local forcing during the El Nino decaying

summer.

FIG. 9. Same as Fig. 6, but for the difference between TIO and CTRL runs.

FIG. 10. Temporal evolutions of area-averaged vorticity anom-

alies (1026 m2 s21) over the region of 108–358N, 1208–1608E for the

WNP (solid line) and TIO (dashed lines) runs.


Acknowledgments. This work was supported by NSFC

Grants 40628006 and 40821092, the National Basic

Research Program of China (2006CB403603), and the

China Meteorological Administration under Grants

GYHY200706010. TL was supported by ONR grants

N000140710145 and N000140810256 and by the Interna-

tional Pacific Research Center, which is sponsored by the

Japan Agency for Marine-Earth Science and Technology

(JAMSTEC), NASA (NNX07AG53G), and NOAA

(NA17RJ1230).

REFERENCES

Chang, C.-P., and T. Li, 2000: A theory of the tropical tropospheric

biennial oscillation. J. Atmos. Sci., 57, 2209–2224.

——, Y. S. Zhang, and T. Li, 2000a: Interannual and interdecadal

variations of the East Asian summer monsoon and tropical

Pacific SSTs. Part I: Roles of the subtropical ridge. J. Climate,

13, 4310–4325.

——, ——, and ——, 2000b: Interannual and interdecadal varia-

tions of the East Asian summer monsoon and tropical Pacific

SSTs. Part II: Meridional structure of the monsoon. J. Climate,

13, 4326–4340.

Chiang, J. C. H., and A. H. Sobel, 2002: Tropical tropospheric

temperature variations caused by ENSO and their influence

on the remote tropical climate. J. Climate, 15, 2616–2631.

Chou, C., L.-F. Huang, J.-Y. Tu, L. Tseng, and Y.-C. Hsueh, 2009:

El Nino impacts on precipitation in the western North Pacific–

East Asian sector. J. Climate, 22, 2039–2057.

Duchon, C., 1979: Lanczos filtering in one and two dimensions.

J. Appl. Meteor., 18, 1016–1022.

Fu, X., B. Yang, Q. Bao, and B. Wang, 2008: Sea surface temper-

ature feedback extends the predictability of tropical intra-

seasonal oscillation. Mon. Wea. Rev., 136, 577–597.

Huang, R., and F. Sun, 1992: Impacts of the tropical western Pacific

on the East Asian summer monsoon. J. Meteor. Soc. Japan, 70,

243–256.

Hurrell, J. W., J. J. Hack, and D. Shea, 2008: A new sea surface

temperature and sea ice boundary dataset for the community

atmosphere model. J. Climate, 21, 5145–5153.

Kanamitsu, M., W. Ebisuzaki, J. Woollen, S.-K. Yang, J. J. Hnilo,

M. Fiorino, and G. L. Potter, 2002: NCEP–DOE AMIP-II

Reanalysis (R-2). Bull. Amer. Meteor. Soc., 83, 1631–1643.

Kang, I. S., and Coauthors, 2002: Intercomparison of the climato-

logical variations of Asian summer monsoon precipitation

simulated by 10 GCMs. Climate Dyn., 19, 383–395.

Klein, S. A., B. J. Soden, and N. C. Lau, 1999: Remote sea surface

temperature variations during ENSO: Evidence for a tropical

atmospheric bridge. J. Climate, 12, 917–932.

Kosaka, Y., and H. Nakamura, 2006: Structure and dynamics of the

summertime Pacific-Japan (PJ) teleconnection pattern. Quart.

J. Roy. Meteor. Soc., 132, 2009–2030.

Lau, K. M., J. H. Kim, and Y. Sud, 1996: Intercomparison of hy-

drologic processes in AMIP GCMs. Bull. Amer. Meteor. Soc.,

77, 2209–2227.

Lau, N.-C., and M. J. Nath, 2000: Impact of ENSO on the variability

of the Asian-Australian monsoons as simulated in GCM ex-

periments. J. Climate, 13, 4287–4309.

——, and ——, 2003: Atmosphere–ocean variations in the

Indo-Pacific sector during ENSO episodes. J. Climate, 16,3–20.

FIG. 11. (a) Observed composite pattern of June–August mean 200-hPa vorticity anomalies during the El Nino

decaying summer. (b) The difference of June–August mean 200-hPha vorticity between (b) the GB, (c) the TIO run,

and (d) the WNP run and the CTRL runs.

1 JUNE 2010 W U E T A L . 2985

Lau, N.-G., and B. Wang, 2006: Interaction between the Asian

Monsoon and El Nino/Southern Oscillation. The Asian Mon-

soon, B. Wang, Ed., Praxis, 479–512.

Li, S. L., J. Lu, G. Huang, and K. M. Hu, 2008: Tropical Indian

Ocean basin warming and East Asian summer monsoon: A

multiple AGCM study. J. Climate, 21, 6080–6088.

Li, T., and B. Wang, 2005: A review on the western North Pacific

monsoon: Synoptic-to-interannual variabilities. Terr. Atmos.

Oceanic Sci., 16, 285–314.

——, Y. S. Zhang, C.-P. Chang, and B. Wang, 2001: On the re-

lationship between Indian Ocean SST and Asian summer

monsoon. Geophys. Res. Lett., 28, 2843–2846.

——, ——, E. Lu, and D. Wang, 2002: Relative role of dynamic and

thermodynamic processes in the development of the Indian

Ocean dipole: An OGCM diagnosis. Geophys. Res. Lett., 29,2110, doi:10.1029/2002GL015789.

——, B. Wang, C. P. Chang, and Y. S. Zhang, 2003: A theory for the

Indian Ocean dipole–zonal mode. J. Atmos. Sci., 60, 2119–

2135.

Li, Y., R. Y. Lu, and B. W. Dong, 2007: The ENSO–Asian monsoon

interaction in a coupled ocean–atmosphere GCM. J. Climate,

20, 5164–5177.

Nitta, T., 1987: Convection activities in the tropical western Pacific

and their impact on the Northern Hemisphere circulation.

J. Meteor. Soc. Japan, 67, 375–383.

Nordeng, T. E., 1994: Extended versions of the convective pa-

rameterization scheme at ECWMF and their impact on the

mean and transient activity of the model in the tropics. Eu-

ropean Centre for Medium-Range Weather Forecasts Re-

search Dept. Tech. Memo. 206, 41 pp.

Rayner, N. A., D. E. Parker, E. B. Horton, C. K. Folland,

L. V. Alexander, D. P. Rowell, E. C. Kent, and A. Kaplan,

2003: Global analyses of sea surface temperature, sea ice, and

night marine air temperature since the late nineteenth cen-

tury. J. Geophys. Res., 108, 4407, doi:10.1029/2002JD002670.

Reynolds, R. W., and T. M. Smith, 1994: Improved global sea

surface temperature analyses using optimum interpolation.

J. Climate, 7, 929–948.

Roeckner, E., and Coauthors, 1996: The atmospheric general cir-

culation model ECHAM-4: Model description and simulation

of present-day climate. Max-Planck-Institute for Meteorology

Rep. 218, 90 pp.

Su, H., J. D. Neelin, and C. Chou, 2000: Tropical teleconnection

and local response to SST anomalies during the 1997–1998

El Nino. J. Geophys. Res., 106, 20 025–20 043.

Sui, C.-H., P.-H. Chung, and T. Li, 2007: Interannual and inter-

decadal variability of the summertime western North Pacific

subtropical high. Geophys. Res. Lett., 34, L11701, doi:10.1029/

2006GL029204.

Tao, S. Y., and L. X. Chen, 1987: A review of recent research of the

East Asian summer monsoon in China. Monsoon Meteorol-

ogy, C.-P. Chang and T. N. Krishnamurti, Eds., Oxford Uni-

versity Press, 60–92.

Tiedtke, M., 1989: A comprehensive max flux scheme for cumulus

parameterization in large-scale models. Mon. Wea. Rev., 117,

1779–1800.

Venzke, S., M. Latif, and A. Villwock, 2000: The coupled GCM

ECHO-2. Part II: Indian Ocean response to ENSO. J. Climate,

13, 1371–1383.

Wang, B., and Q. Zhang, 2002: Pacific–east Asian teleconnection.

Part II: How the Philippine Sea anomalous anticyclone is es-

tablished during El Nino development. J. Climate, 15, 3252–

3265.

——, R. G. Wu, and X. H. Fu, 2000: Pacific–East Asian tele-

connection: How does ENSO affect East Asian climate?

J. Climate, 13, 1517–1536.

——, ——, and T. Li, 2003: Atmosphere–warm ocean interac-

tion and its impacts on Asian–Australian monsoon variation.

J. Climate, 16, 1195–1211.

Wu, B., and T. J. Zhou, 2008: Oceanic origin of the interannual and

interdecadal variability of the summertime western Pacific

subtropical high. Geophys. Res. Lett., 35, L13701, doi:10.1029/

2008GL034584.

——, ——, and T. Li, 2009a: Contrast of rainfall–SST relationships

in the western North Pacific between the ENSO developing

and decaying summers. J. Climate, 22, 4398–4405.

——, ——, and ——, 2009b: Seasonally evolving dominant in-

terannual variability modes of East Asian climate. J. Climate,

22, 2992–3005.

Wu, R., Z.-Z. Hu, and B. P. Kirtman, 2003: Evolution of ENSO-

related rainfall anomalies in East Asia. J. Climate, 16, 3742–3758.

Xie, P., and P. A. Arkin, 1997: Global precipitation: A 17-year

monthly analysis based on gauge observation, satellite esti-

mate, and numerical model outputs. Bull. Amer. Meteor. Soc.,

78, 2539–2558.

Xie, S.-P., K. Hu, J. Hafner, H. Tokinaga, Y. Du, G. Huang, and

T. Sampe, 2009: Indian Ocean capacitor effect on Indo-

western Pacific climate during the summer following El Nino.

J. Climate, 22, 730–747.

Yang, J. L., Q. Y. Liu, S.-P. Xie, Z. Y. Liu, and L. X. Wu, 2007:

Impact of the Indian Ocean SST basin mode on the Asian

summer monsoon. Geophys. Res. Lett., 34, L02708, doi:10.1029/

2006GL028571.

Yu, L., X. Jin, and R. A. Weller, 2008: Multidecade Global Flux

Datasets from the Objectively Analyzed Air-sea Fluxes (OAFlux)

Project: Latent and sensible heat fluxes, ocean evaporation,

and related surface meteorological variables. Woods Hole

Oceanographic Institution, OAFlux Project Tech. Rep. OA-

2008-01, 64 pp.

Zhang, R. H., and A. Sumi, 2002: Moisture circulation over East

Asia during El Nino episode in northern winter, spring and

autumn. J. Meteor. Soc. Japan, 80, 213–227.

——, ——, and M. Kimoto, 1996: Impact of El Nino on the East

Asian monsoon: A diagnostic study of the ’86/87 and ’91/92

events. J. Meteor. Soc. Japan, 74, 49–62.

Zhou, T. J., B. Wu, and B. Wang, 2009a: How well do atmospheric

general circulation models capture the leading modes of the

interannual variability of the Asian–Australian monsoon?

J. Climate, 22, 1159–1173.

——, and Coauthors, 2009b: Why the western Pacific subtropical

high has extended westward since the late 1970s. J. Climate, 22,

2199–2215.


relative contributions of the indian ocean and local...

Documents