relative contributions of the indian ocean and local...
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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.
2976 J O U R N A L O F C L I M A T E VOLUME 23
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.
2978 J O U R N A L O F C L I M A T E VOLUME 23
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.
2980 J O U R N A L O F C L I M A T E VOLUME 23
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).
2982 J O U R N A L O F C L I M A T E VOLUME 23
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.
2984 J O U R N A L O F C L I M A T E VOLUME 23
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).
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