JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 116, D13111, doi:10.1029/2010JD015554, 2011
Interannual relationships between the tropical sea surface
temperature and summertime subtropical anticyclone
over the western North Pacific
Pei‐Hsuan Chung,1 Chung‐Hsiung Sui,2 and Tim Li3
Received 27 December 2010; revised 8 March 2011; accepted 6 April 2011; published 13 July 2011.
[1] The interannual variability of the Western North Pacific Subtropical High (WNPSH)
in boreal summer is investigated with the use of the NCEP/NCAR Reanalysis Data. The
most significant change of the 500 hPa geopotential height field appears at the western
edge of the WNPSH, with dominant 2–3 year and 3–5 year power spectrum peaks.
The 2–3 year oscillation of the WNPSH and associated circulation and sea surface
temperature (SST) patterns possess a coherent eastward propagating feature, with a warm
SST anomaly (SSTA) and anomalous ascending motion migrating from the tropical
Indian Ocean in the preceding autumn to the maritime continent in the concurrent summer
of a strong WNPSH. A strong WNPSH is characterized by anomalous anticyclonic
circulation and maximum subsidence in the western North Pacific (WNP). The anomalous
WNPSH circulation has an equivalent barotropic vertical structure and resides in the
sinking branch of the local Hadley circulation, triggered by enhanced convection over the
maritime continent. A heat budget analysis reveals that the WNPSH is maintained by
radiative cooling. The 3–5 year oscillation of the WNPSH exhibits a quasi‐stationary
feature, with a warm SSTA (anomalous ascending motion) located in the equatorial
central eastern Pacific and Indian Ocean and a cold SSTA (anomalous descending motion)
located in the western Pacific. The anomaly pattern persists from the preceding winter
to the concurrent summer of a high WNPSH. The greatest descent is located to the
southeast of the anomalous anticyclone center, where a baroclinic vertical structure is
identified. The zonal phase difference and the baroclinic vertical structure suggest that the
anomalous anticyclone on this timescale is a Rossby wave response to a negative latent
heating associated with the persistent local cold SSTA. ECHAM4 model experiments
further confirm that the 2–3 year mode is driven by the SSTA forcing over the maritime
continent, while the 3–5 year mode is driven by the local SSTA in the WNP.
Citation: Chung, P.‐H., C.‐H. Sui, and T. Li (2011), Interannual relationships between the tropical sea surface temperature and
summertime subtropical anticyclone over the western North Pacific, J. Geophys. Res., 116, D13111,
doi:10.1029/2010JD015554.
1. Introduction
[2] The subtropical high over the North Pacific is a dominant component of atmospheric general circulation that
experiences a distinct seasonal cycle with the maximum
strength and northernmost position occurring in the boreal
summer. The evolution of the subtropical high over the
western North Pacific (hereafter WNPSH) in summer is
associated with the onset and withdrawal of the Asian sum1
Department of History and Geography, Taipei Municipal University of
Education, Taipei, Taiwan.
2
Department of Atmospheric Sciences, National Taiwan University,
Taipei, Taiwan.
3
International Pacific Research Center, University of Hawai’i at Mānoa,
Honolulu, Hawaii, USA.
Copyright 2011 by the American Geophysical Union.
0148‐0227/11/2010JD015554
mer monsoon. The zonal shift of the western part of the
WNPSH dominates the position of large‐scale frontal zones
in East Asia [Tao and Chen, 1987]. A southwestward
extension of the WNPSH would result in more water vapor
convergence and excessive rainfall along the Yangtze River
valley [Chang et al., 2000a, 2000b; Zhou and Yu, 2005]. The
precipitation characters in Japan and Korea were also greatly
influenced by the movement of the WNPSH through changing the transport of water vapor into the Baiu or Changma
frontal area [e.g., Tomita et al., 2004; Lee et al., 2005].
[3] Traditionally, subtropical highs are regarded as the
descending branch of the Hadley circulation maintained by
radiative cooling. This idealized, zonally symmetric Hadley
circulation model, however, would lead to a weaker
(stronger) subtropical high in the summer (winter) hemisphere, which is contrary to observations [Hoskins, 1996].
Moreover, the observed summertime subtropical highs are
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in the form of high‐pressure cells rather than a zonally
uniform high‐pressure belt. The dynamics of the summertime subtropical high is better understood in the context of
planetary waves, as suggested by Ting [1994], who found
that summertime subtropical stationary wave patterns are
maintained by the global diabatic heating.
[ 4 ] Hoskins [1996] and Rodwell and Hoskins [2001]
applied the monsoon‐desert mechanism proposed by
Rodwell and Hoskins [1996] to argue that the remote diabatic
heating over the North American monsoon region could
induce a Rossby wave pattern to the west and contribute to
the lower tropospheric subtropical anticyclone in the North
Pacific. Rodwell and Hoskins [2001] further examined the
dynamics of the summertime subtropical high in a global
nonlinear model and concluded that the combined effect of
the topography and the monsoonal heating is of primary
importance for the generation of the surface subtropical
highs and subsidence aloft over the eastern Pacific. This
differs from Chen et al. [2001], who argued that summertime
subtropical highs can be maintained by downstream energy
propagation of stationary Rossby waves forced by convection heating over the Asian monsoon region in a linear quasi‐
geostrophic model. Their experiments, however, could not
produce the distinct meridional dependence of the observed
vertical structures of the summertime planetary wave, possibly due to the lack of other heating sources such as sensible
heating over the western portion of the continent and radiation cooling off the west coast of the continent. Liu et al.
[2004] performed atmospheric GCM experiments to show
the importance of the global surface sensible heating and
longwave radiative cooling in the formation of the subtropical anticyclones. Miyasaka and Nakamura [2005] demonstrated that the summertime subtropical high can be
accurately reproduced by the near‐surface thermal contrast
between a cooler ocean and a hotter land to the east. The role
of local air‐sea interactions in generating the subtropical
anticyclone was investigated numerically by Seager et al.
[2003]. All studies above suggest the important role of the
land‐ocean‐atmosphere feedback in maintaining the subtropical high.
[5] In addition to the annual cycle, the subtropical high in
the boreal summer also exhibits remarkable interannual
variations. Convective heating over the western Pacific
warm pool is regarded as an important forcing that affects
the extent of the subtropical high in the boreal summer.
According to Nitta [1987], convective activity in the tropical western Pacific during summer can influence the
Northern Hemispheric circulation; that is, Rossby waves
generated by the anomalous tropical heat source can propagate into higher latitudes. This leads to a north–south
dipole between subtropics and middle latitude East Asia,
known as the Pacific‐Japan (PJ) Pattern. Lu [2001] found
that stronger convection over the warm pool is associated
with more eastward and northward migration of the subtropical high. Lu and Dong [2001] performed numerical
model experiments and showed that suppressed convection
caused by a cold SSTA in the warm pool results in anomalous anticyclonic circulation in the lower troposphere over
the subtropical western North Pacific (WNP), providing a
favorable condition for the westward extent of the subtropical high.
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[6] In addition to the warm pool heating, the El Nino‐
Southern Oscillation (ENSO) may also influence the East
Asian climate through the modulation of low‐level winds
over the WNP [e.g., Wang et al., 2000; Chang et al., 2000a,
2000b; Lau and Nath, 2000; Kawamura et al., 2001; Lau
and Wu, 2001; Chou, 2004]. An anomalous low‐level
anticyclone dominates the WNP during the mature phase of
an El Nino, whereas an anomalous cyclone occurs during
the La Nina mature winter. The anomalous anticyclone
persists through the following spring and summer, leading
to a weak WNP monsoon in the ENSO decaying summer
[e.g., Wang et al., 2000, 2003; Chou et al., 2003]. The
anticyclone in early summer favors the northward movement of the Mei‐yu/Baiu front and thus enhances (weakens)
the East Asia (WNP) summer monsoon rainfall [Chang
et al., 2000a, 2000b].
[7] The wintertime anomalous low‐level anticyclone in
WNP may be regarded as a Rossby wave response to the
suppressed convection over the western Pacific, which is
possibly caused both by the remote warm SSTA over the
eastern Pacific [Wang et al., 2000; Chou, 2004; Chen et al.,
2007] and the local cold SSTA [Lau and Nath, 2000; Wang
et al., 2000]. The maintenance of the WNP anticyclone from
the El Nino mature winter to its decaying early summer is
attributed to a season‐dependent thermodynamic air‐sea
feedback [Wang et al., 2000, 2003; Kawamura et al., 2001;
Chou, 2004; Li and Wang, 2005]. During the El Nino
decaying summer, the SSTA in the eastern Pacific has
diminished, yet the low‐level anticyclone still exists over the
WNP. It is likely that both the cold SSTA in the WNP in late
spring/early summer [Sui et al., 2007; Wu et al., 2010] and
the basin‐wide warming in the IO in summer [Yang et al.,
2007; Xie et al., 2009; Wu et al., 2009a, 2009b] are
responsible for maintaining the WNPSH during the summer.
[8] The Asian summer monsoon shows a strong quasi‐
biennial variability [e.g., Shen and Lau, 1995; Barnett,
1991; Chang and Li, 2000; Li et al., 2001; Li and Zhang,
2002], which is termed as the tropospheric biennial oscillation (TBO) in the Asian‐Australian monsoon region
[Meehl, 1994, 1997]. The tropical Pacific and southeast IO
(SEIO) SSTA associated with the Asian summer monsoon
also exhibits a biennial characteristic [Chang et al., 2000a;
Wang et al., 2001]. A hybrid coupled GCM simulation
shows that the TBO may exist even in the absence of ENSO
[Li et al., 2006]. On the other hand, ENSO may be linked to
the monsoon TBO through the following three branches of
teleconnection: (1) WNP monsoon‐SEIO SSTA teleconnection or the Philippine‐Sumatra pattern [Li et al.,
2002; Li et al., 2006; Wu et al., 2009b]; (2) IO convection‐Central Eastern Pacific (CEP) SSTA teleconnection
[Meehl, 1987; Chang and Li, 2000; Yu et al., 2002;
Watanabe and Jin, 2002; Li et al., 2003]; and (3) Pacific‐
East Asia teleconnection [Wang et al., 2000, 2003].
[9] Because the WNPSH is regarded as one of the
important components of the Asian summer monsoon system, the objective of this study is to conduct an analysis of
the interannual evolution of the WNPSH and investigate
possible mechanisms that link to the variability. The rest
of the paper is organized as follows. Section 2 describes the
data used in this study. In section 3 a WNPSH index is
defined and two distinctive spectrum peaks on quasi‐biennial
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(2–3 year) and lower‐frequency (3–5 year) periods are
identified. A regression analysis is subsequently performed
to obtain the large‐scale circulation, SST, outgoing longwave radiation, and heat and moisture budgets associated
with the two oscillation periods. By synthesizing all features,
possible mechanisms responsible for the two oscillations of
the WNPSH are proposed. The idealized experiments of the
ECHAM4 AGCM are also examined in section 4. In section
5 the seasonal lead‐lag patterns of regressed SSTA and circulation fields associated with the 2–3 year and 3–5 year
oscillations of the WNPSH are discussed. Section 6 comprises a summary and discussions.
2. Data and Model Description
2.1. Reanalysis Data
[10] The primary data set used in this study is the National
Centers for Environmental Prediction–National Center for
Atmospheric Research (NCEP/NCAR) Reanalysis [Kalnay
et al., 1996] that includes wind, geopotential height,
stream function, vertical p velocity, sea level pressure,
velocity potential, and outgoing longwave radiation flux on
a 2.5 × 2.5 grid for 48 years (1958–2005). This is the longest reanalysis product available. Another set of data used in
the study is the NOAA Extend Reconstructed Sea Surface
Temperature version 2 (ERSST v2) developed on a 2° × 2°
grid for the 48 years (1958–2005). The extended reconstruction of historical SST was using improved statistical
methods from 1854 to the present. A full description of the
ERSST v2 is available in the work of Smith and Reynolds
[2004].
[11] Further, this study used the European Centre for
Medium‐Range Weather Forecasts (ECMWF) Reanalysis
(ERA), which also contains the above variables at the same
resolution from 1957 to 2002. A full description of the ERA
is available in the work of Gibson et al. [1996, 1997]. The
analysis results from this study, based on NCEP/NCAR
data, were compared against those from ERA and the major
features were found to be essentially the same. Most of the
results discussed in this study are based on NCEP/NCAR
reanalysis, except the following.
[12] In order to discuss the dynamic features of the climate oscillations in this study, the heat and moisture budgets were computed, the so‐called apparent heat source (Q1)
and apparent moisture sink (Q2) [e.g., Yanai et al., 1973],
using ERA. This decision is made due to the fact that the
overall simulation of the diabatic heating (and therefore
the moisture) field over the monsoon region in ERA is quite
close to reality [Annamalai et al., 1999]. In addition, a
comparison of the heat and moisture budgets based on
ERA, NCEP/NCAR R‐1 Reanalysis, and NCEP/DOE
AMIP‐II (R‐2) Reanalysis [Kanamitsu et al., 2002] was
completed. The results show that budgets computed based
on both ERA and R‐2 reanalysis data are quite similar to the
observed rainfall variability in the tropical and East Asian
monsoon regions. However, the R‐2 reanalysis data only
cover the satellite period from 1979 to present. Therefore
the Q1 and Q2 budgets were computed at each vertical level
(23 pressure levels) using the 6‐hourly ERA data from 1958
to 2001.
[13] According to Yanai et al. [1973], the large‐scale heat
and moisture budget residuals, Q1 and Q2, respectively, are
defined and computed by
Q1 cp
k p
@
@
þ~
r þ !
p0
@t
@p
ð1Þ
and
@q
@q
þ~
rq þ !
Q2 L
@t
@p
ð2Þ
where is the potential temperature, q is the water vapor
mixing ratio, ~
is the horizontal velocity, w is the vertical p
velocity. L is the latent heat of vaporization, k = R/cp with R
the gas constant and cp the specific heat capacity at constant
pressure of dry air, and p0 = 1000 hPa.
, , and q denote large‐scale
[14] In equations (1) and (2), ~
fields that are mathematically defined as Reynolds averaged
quantities (normally with an overbar, being neglected here)
and are customarily assumed resolvable by reanalysis data.
So Q1 and Q2 are computed by 6‐hourly data at each grid
point from ERA products. The 6‐hourly estimates of Q1
and Q2 are averaged into monthly mean values and multiplied by c−1
p to be expressed in terms of equivalent warming/
cooling rate (K/day).
[15] The Q1 and Q2 can be interpreted by:
Q1 ¼ QR þ Lðc eÞ r s′ v′ @s′ !′
@p
ð3Þ
and
Q2 ¼ Lðc eÞ þ Lr q′ v′ þ L
@q′ !′
@p
ð4Þ
where QR is radiative heating rate, c and e are condensation and evaporation per unit mass of air respectively, s is
the dry static energy, and the prime denotes deviation from
the Reynolds average. The primed variables denote subgrid
scale fluctuations that are assumed to be well separated from
the resolvable large‐scale variable. Equation (3) represents
the total effect of radiative heating, latent heat released by
net condensation L(c − e), and the horizontal and vertical convergence of sensible heat fluxes due to subgrid‐scale eddies
such as cumulus convection and turbulence. Equation (4),
on the other hand, represents the total effect of net condensation and divergence of eddy moisture fluxes due to
subgrid‐scale eddies. The contributions of the eddy horizontal transport terms r · s′v′ in equation (3) and r · q′v′ in
equation (4) are generally small and may be negligible compared to the horizontal transports by the large‐scale motion
[Yanai and Johnson, 1993].
2.2. Atmospheric General Circulation Model
[16] The Atmospheric General Circulation Model (AGCM)
used in the study is ECHAM version 4.6 (hereafter ECHAM4),
which was developed by the Max Planck Institute for Meteorology [Roeckner et al., 1996]. The model was run at a
horizontal resolution of spectral triangular 42 (T42), with 2.8°
grid and 19 vertical levels in a hybrid sigma pressure coordinate system. The climatological monthly SST constructed
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Figure 1. The standard deviation of 2–7 year band‐pass filtered summer‐mean geopotential height at
500 hPa (Z500) normalized by the corresponding zonal mean Z500 (shaded) and 5870‐m contour line
for a strong (red), normal (green), and weak (blue) WNPSH composite.
by phase II of the Atmospheric Model Intercomparison
Project (AMIP II) was used as the model low boundary
forcing. The ECHAM4 model is one of the Top‐3 AMIP II
models in simulating the interannual variability modes of the
Asian‐Australian monsoon [Zhou et al., 2009a]. It shows
excellent performance in simulating the El Nino‐related tropospheric temperature anomalies [Zhou and Zhang, 2011].
Thus the model is used in the study for investigating the
interannual variations of the summertime circulation in WNP.
3. Characteristics Associated With
the Anomalous Subtropical High
3.1. WNPSH Index
[17] In this study the interannual variation of the summer
subtropical high over the WNP is measured using the 500 hPa
geopotential height field (Z500). Although the summer
Pacific subtropical high can be detected in the lower troposphere, the WNPSH is most pronounced in the middle troposphere. The center of the subtropical high in the lower
troposphere is located in the eastern Pacific, while the center
of the subtropical high in the middle troposphere resides over
the western North Pacific [Liu et al., 2001; Liu and Wu,
2004]. For this reason, the variations of the WNPSH in
this study are measured by the Z500 variability, following
many previous studies [e.g., Tao and Chen, 1987; Zhou and
Li, 2002]. Figure 1 shows the standard deviation (~
Z) of the
summer (JJA) mean Z500 during the period of 1958–2005. It
is calculated from 2 to 7 year band‐pass filtered data and has
been normalized by its zonal mean value to illustrate the
zonally asymmetric character. Figure 1 shows that the largest
variability occurs in the northwest Pacific region, southwest
of the climatological mean subtropical high as depicted by
the contour line of Z500 = 5870 m. A WNPSH index is
defined as the area mean value of Z500 anomalies in JJA in
the region of (120°–140°E, 10°–30°N). This index is used to
measure the strength of the interannual variability of the
WNPSH. For comparison, two additional indices are defined
based on 850 hPa stream function (y850) and geopotential
height anomalies (Z850) in JJA averaged over the region of
(120–150°E, 15–25°N), where the strongest interannual
variation of y850 and Z850 occurs (hereafter termed as y850
and Z850 indices). The y850 and Z850 fields were used by Lee
et al. [2005] and Lu [2001] to study the subtropical high
variability over the WNP.
[18] The power spectrums of the Z500, y850, and Z850
indices are examined. The Z500 index shows two dominant
periods at 2–3 years and 3–5 years, both of which pass the
95% confidence level (Figure 2a). The y850 and Z850 index
also presents two similar peaks although with the 3–5 year
band (2–3 year band) being above (slightly below or at) the
95% confidence level (figures not shown). Additional power
spectrum analyses with 44‐year ERA40 Z500 data (figure not
shown) show a similar result. Thus in the following, the Z500
index will be used to represent the WNPSH quasi‐biennial
(2–3 year) and lower‐frequency (3–5 year) variability.
[19] Through a Fourier series decomposition of the
WNPSH index series, two band‐pass filtered time series
were constructed by summing Fourier components within
the 2–3 year and 3–5 year periods (Figure 2b). Hereafter, the
two time series are referred to as the 2–3 year and 3–5 year
WNPSH indices. Positive indices represent the strong and
westward extending phase of the subtropical high, while
negative indices represent the opposite. In the following
analysis, the circulation and SST fields are regressed against
the two WNPSH indices to investigate spatial and temporal
structures associated with the 2–3 year and 3–5 year oscillations of the WNPSH.
3.2. Circulation Patterns Associated With the 2–3 Year
and 3–5 Year Oscillations
[20] The structures related to 2–3 year and 3–5 year
oscillations of the WNPSH were investigated by examining
the horizontal patterns of regressed fields of summer mean
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Figure 2. (a) Power spectrums of the summer‐mean geopotential height at 500 hPa (WNPSH index) and (b) decomposed band‐pass filtered time series of the (top) 2–3 years
and (bottom) 3–5 years WNPSH index. The pink curve in
Figure 2a denotes the 95% confidence level.
ERSST, outgoing longwave radiation (OLR), Z500, 500 hPa
vertical p velocity (w500), 200 hPa and 850 hPa stream
function (y200 and y850), 200 hPa and 850 hPa velocity
potential (c200 and c850), and wind fields at 850 hPa and
500 hPa (V850 and V500). Since the derived results of regression can be applied to both the positive phase (represented by
a strengthened and westward WNPSH) and the negative
phase (represented by a weakened WNPSH), the interpreta-
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tion and discussion of physical mechanisms afterward will be
merely based on the features of the positive phase.
[21] The regressed Z500, V500, y850, and y200 are shown in
Figure 3. The WNPSH high anomaly of the 2–3 year
oscillation is associated with a meridional standing wave
pattern with a zonally elongated positive Z500 anomaly
and anticyclonic circulation in the WNP, a negative Z500
anomaly over the northern East Asia, and a positive
anomaly over Eurasia (Figure 3b). The horizontal distribution of y850 in Figure 3c shows a pair of anticyclonic circulations symmetric about 10 N. One is located in the WNP
(overlapping with a positive Z500 anomaly center) and the
other is located south of the Equator. The circulation pattern
of y200 (Figure 3a) shows that an anticyclonic circulation
extends from southern to southeastern China and the WNP.
[22] In the 3–5 year oscillation, an anticyclone and a
positive Z500 anomaly center (Figure 3e) are located near
Taiwan when the WNPSH extends westward. The positive
Z500 anomaly extends further to the tropical central eastern
Pacific than that in 2–3 year oscillation. The positive
anomaly in Z500 and y850 fields (Figures 3e and 3f) acts as
the source of a Rossby wave train propagating from the
WNP to North America. The distribution of y200 (Figure 3d)
exhibits a cyclonic circulation pattern over China. In the
WNP, y200 is in general out of phase with y850.
[23] Figure 4 shows the regressed SST patterns for the 2–
3 year and 3–5 year modes. The SST field has been normalized at each grid point based on the climatological
standard deviation to capture the obvious SSTA variations.
The SST distribution associated with the westward extending
WNPSH in the 2–3 year oscillation (Figure 4a) features a
warm SSTA extending from the northern IO southeastward
to South China Sea, maritime continent and northeastern
Australia, and a cold SSTA in the central eastern equatorial
and northeastern and southeastern Pacific. For the 3–5 year
oscillation (Figure 4b), however, a warm SSTA appears in
the tropical IO and the central eastern equatorial Pacific. The
SSTA in the central eastern Pacific is surrounded by cold
SST anomalies to its north and south and in the maritime
continent.
[24] In the 2–3 year oscillation, the large‐scale divergent
circulation in upper and lower levels are shown by regressed
c200 and c850 fields in Figures 5a and 5b. A broad overturning circulation occurs in the summer of a strong
WNPSH over the western Pacific, with the low‐level convergent (divergent) flow toward Java and the eastern IO (out
of the tropical central eastern Pacific region) (Figure 5b).
The regressed OLR field in Figure 5c shows enhanced
convection over the maritime continent and suppressed
convection over the Philippine Sea and the equatorial eastern Pacific, consistent with the divergent wind in Figures 5a
and 5b. A zonal band of enhanced convection is also
observed to the north of the WNPSH. The anomalous OLR
pattern in the western Pacific resembles a negative phase of
the PJ pattern [Nitta, 1987].
[25] For the 3–5 year oscillation, the divergent circulation
pattern differs from that of the 2–3 year oscillation. The
regressed c200 and c850 fields (shown in Figures 5d and 5e)
reveal two anomalous Walker cells in the tropics, with two
upper‐tropospheric divergence centers located in the tropical
central‐eastern Pacific and IO, respectively, and a convergence center located in the region extending from southern
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Figure 3. Regressed summer‐mean fields of stream function at (a, d) 200 hPa and (c, f) 850 hPa and
(b, e) geopotential height with significant wind fields at 500 hPa for the 2–3 year (Figures 3a, 3c, and 3e)
and 3–5 year (Figures 3b, 3d, and 3f) oscillations of WNPSH. Shaded (vectors) areas pass the 90% confidence level.
Philippines to the maritime continent and eastern Australia
(Figure 5d). The regressed OLR field (Figure 5f) shows the
regions of suppressed convection in the maritime continent
and subtropical North Pacific and the regions of enhanced
convection in the equatorial eastern Pacific and western
tropical IO, consistent with the features of the divergent
flows shown in the c200 and c850 fields (Figures 5d and 5e).
[26] The meridional overturning circulation of the 2–
3 year oscillation over the western Pacific is further revealed
by the regressed zonally averaged vertical p velocity (w) at
each pressure level and latitude‐vertical divergent wind
profile (v, w) along 110°E–140°E (Figure 6a). Consistent
with the regressed OLR and c fields, the flow is upward
over the maritime continent (0°–10°S) and East Asia near
Japan (30°–40°N), with the former being stronger than the
latter. The southern branch of upward motion causes upper‐
tropospheric divergent flows toward the Philippine Sea (at
10°–20°N) where the downward motion is pronounced. The
low‐level southward divergent flow to the south of the
WNPSH may bring the moisture into the equatorial region
and enhance convection in the maritime continent. The
enhanced convection further reinforces the WNPSH subsidence via a positive feedback process.
[27] For the 3–5 year oscillation, the regressed vertical
overturning circulation along 110°E–150°E (Figure 6b)
shows two descending branches at 0°–10°S and 15°–25°N,
coinciding with the two suppressed convective centers in
the OLR field. There is no direct linkage between the WNP
and equatorial regions (Figure 5f). It can be seen from c
analysis in Figures 5d and 5e that the equatorial branch of
the descending motion within the western Pacific is a result
of the anomalous large‐scale Walker circulation. In the c
field, however, the convergent flow over the WNP is not as
strong as that in the maritime continent/eastern Australia
(Figure 5d). The result suggests that unlike the 2–3 year
oscillation, the subsidence in the WNP is not directly
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Figure 4. Regressed summer‐mean SSTA fields in (a) the 2–3 year and (b) 3–5 year oscillations of
WNPSH, normalized by climatological standard deviations at each grid point. Shaded areas pass the
90% confidence level.
attributed to the equatorial forcing in the 3–5 year oscillation. Rather, it may be forced through a Rossby wave
response to anomalous heating associated with the local
cold SSTA in Figure 4b [Sui et al., 2007; Wu et al., 2010].
3.3. Vertical Structures of the WNPSH Anomalies
[28] The analysis above indicates that the large‐scale
circulations associated with an enhanced summertime
WNPSH in the 2–3 year and 3–5 year oscillations are distinctly different. The 2–3 year oscillation is accompanied by
an anomalous local Hadley circulation that connects the
maritime continent and the WNP, while the 3–5 year
oscillation is characterized by the anomalous Walker circulation and a cold SSTA to the southeast of the anomalous
WNPSH. To reveal the specific processes that affect
WNPSH, the horizontal and vertical structures of various
dynamic and thermodynamic variables associated with the
two oscillations are further compared.
[29] First, the relative zonal position between the anomalous anticyclone and descending motion centers during a
strong WNPSH phase are examined by depicting the regressed y850 (contours) and w500 (blue shades) fields in
Figure 7. Only the stronger (at the 95% significant level)
portion of the regressed positive w500 field is shown in
Figure 7. While the subsidence center (suppressed convection in Figure 5c) in the 2–3 year oscillation is zonally in
phase with the anticyclone center, the maximum subsidence
(suppressed convection in Figure 5f) in the 3–5 year oscillation is located to the southeast of the anomalous anticyclone center. Such a phase difference suggests that the
anomalous WNPSH in the former is possibly a direct
response to the local Hadley circulation anomaly, whereas in
the latter it is a baroclinic Rossby wave response to a negative heating anomaly induced by the local cold SSTA
(Figure 4b).
[30] To reveal the vertical structure of the anomalous
WNPSH, the domains of the strongest downward motion
and anomalous subtropical high center on the 2–3 year and
3–5 year timescales respectively were chosen. The vertical
profiles of the domain‐averaged regressed fields of vertical
p velocity, relative vorticity, Q1 and Q2 are computed at
each pressure level.
[31] Figure 8a shows the regressed vertical p velocity
profiles over the area of strongest downward motion in the
2–3 year and 3–5 year oscillations. The strongest subsidence
is located in the upper troposphere (300 hPa) in the 2–3 year
oscillation but in the middle troposphere (400–500 hPa) in
the 3–5 year oscillation. The magnitude of the subsidence in
the 3–5 year oscillation is weaker than that in the 2–3 year
oscillation. Figure 8b shows the vertical profiles of the
regressed relative vorticity fields over the anomalous anticyclone center associated with the 2–3 year and 3–5 year
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Figure 5. Regressed summer‐mean velocity potential and divergent wind at (a, d) 200 hPa and (b, e)
850 hPa and (c, f) upward longwave radiation flux (similar to outgoing longwave radiation) for the
2–3 year (Figures 5a, 5c, and 5e) and 3–5 year (Figures 5b, 5d, and 5f) oscillations of WNPSH. The
red (blue) contours denote positive (negative) velocity potential anomalies or enhanced (suppressed) convection. Shaded regions pass the 90% confidence level.
oscillations. While it exhibits a clear equivalent barotropic
structure in the 2–3 year oscillation, the vorticity has a
baroclinic structure in the 3–5 year oscillation, changing
from negative anomalies in the lower troposphere to a
positive anomaly in the upper troposphere (above 250 hPa).
This result is consistent with circulation patterns over the
WNP area shown in Figures 3a, 3c, 3d, and 3f.
[32] The vertical distributions of regressed Q1 and Q2
anomalies averaged over the strongest descent region are
shown in Figures 8c and 8d. Note that the amplitude of the
negative Q1 anomaly in the middle‐upper troposphere is
greater in the 2–3 year oscillation than in the 3–5 year
oscillation (Figure 8c), while the amplitude of the negative Q2
anomaly in the lower troposphere is greater in the 3–5 year
oscillation than in the 2–3 year oscillation (Figure 8d). This
implies that enhanced longwave radiative cooling due to
the reduction of the vertically integrated moisture caused by
the descent‐induced dry advection plays an important role
in the establishment of the anomalous WNPSH on the 2–
3 year timescale. On the other hand, a peak of anomalous
Q2 at 850 hPa in the 3–5 year oscillation implies strong
negative condensation heating at the top of the boundary
layer. The negative condensation heating is induced by
the local cold SSTA over the region of (150°–170°E, 15°–
20°N) that is to the southeast of the anomalous anticyclone
and is responsible for a baroclinic vertical structure.
[33] In both the oscillations the warm SSTA under the
anomalous anticyclone may be regarded as a passive
response to the atmospheric forcing [Wang et al., 2005; Wu
and Kirtman, 2005; Zhou et al., 2008]. The anomalous
anticyclone leads to the suppression of convection and thus
the increase of downward solar radiation into the ocean
surface. Meanwhile, less evaporation due to the weakening
wind speed may also increase SST.
[34] The distinctive horizontal and vertical structures
above suggest that the physical mechanisms responsible
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Figure 6. Regressed vertical velocity (contours) and divergent wind fields (a) averaged over 110°E–140°E
for 2–3 year and (b) averaged over 110°–150°E for 3–5 year oscillations of WNPSH. Shaded regions pass
the 90% confidence level.
for the anomalous WNPSH may differ in the 2–3 year and
3–5 year oscillations. According to the Gill [1980] model, a
Rossby wave response is located to the west of a heating
source. This corresponds well to the 3–5 year oscillation
case, in which there is a phase shift between the anomalous
anticyclone and the anomalous heating (or vertical motion).
Further, the Rossby wave response to the anomalous heating
is baroclinic. These features are accurately observed in the
3–5 year oscillation. Therefore the diagnosis result in this
study suggests that the variation of the WNPSH on the 3–
5 year timescale is primarily caused by an atmospheric
Rossby wave response to a negative heating associated with
a cold SSTA to the southeast of the anticyclone center. In
contrast, in the 2–3 year oscillation, the anomalous WNPSH
center is zonally colocated with the vertical motion, and its
vertical structure is approximately barotropic. This implies
that the Rossby wave response is unlikely and the downward motion is a major driving mechanism. That is, the
strengthening of the WNPSH on the 2–3 year timescale
is primarily attributed to the radiative cooling due to the
descent‐induced dry advection; the less clouds and water
vapor in the atmosphere, the more longwave radiation is
emitted into space. The change of local vertical motion and
circulation in the WNP is induced by the remote forcing via
anomalous local Hadley circulation associated with enhanced
convective heating in the maritime continent.
[35] The regressed results above are consistent with the
summertime subtropical LOSECOD diabatic heating pattern
identified by Wu and Liu [2003] and Wu et al. [2009], who
showed that the dominant diabatic heating sources over the
northwestern Pacific region are convective heating and
longwave radiation cooling. The corresponding diagnoses of
Figures 5, 7, and 8 indeed demonstrate that in the 2–3 year
oscillation, the longwave radiation cooling associated with
the anomalous descent by the local Hadley circulation prevails over the negative convective heating; while in the 3–
5 year mode, the negative convective heating due to the cold
SSTA prevails over the longwave radiation cooling. The
negative convective heating will induce an atmospheric
Rossby wave response to the west as the anomalous
WNPSH. As a consequence, the subtropical anticyclone
anomaly of the 2–3 year oscillation possesses a barotropic
structure, whereas in the 3–5 year oscillation possesses a
baroclinic structure as presented in Figure 8b. This is
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Figure 7. Regressed summer‐mean stream function at 850 hPa (contour) and positive p velocity (only
most significant descending region is indicated here by shading) at 500 hPa for (a) the 2–3 year and (b) 3–
5 year oscillations of WNPSH.
because a convective heating in the subtropics can generate
a Sverdrup balance stream function pattern with anticyclone
(cyclone) to the east and cyclone (anticyclone) to the west
of the heating (negative heating) in the lower troposphere
and the opposite in the upper troposphere, as discussed by
Wu and Liu [2003], Liu et al. [2004], and Wu et al. [2009].
Therefore the distinct difference between the two oscillation
modes should be attributed to the different dominant diabatic heating over the northwest Pacific.
4. Model Experiment Results
[36] To test the hypothesis that the WNPSH variability on
the 3–5 year timescale is primarily forced by local SSTA in
the WNP while the WNPSH variability on the 2–3 year
timescale is attributed to the forcing over the maritime
continent, the following control and sensitivity experiments
have been developed. In the control run (CTRL), ECHAM4
was integrated for 20 years, forced by monthly climatological SST (based on the period of 1955–2000). In the
WNP run, the model was forced by the climatological SST
plus a uniform SSTA of −1°C over a WNP domain (150°E–
160°W, 10°–25°N) (blue box in Figure 9a). In the MC run,
the same simulation strategy was used as in the WNP run
except with a +1°C SSTA specified in the maritime continent region (90°–150°E, 20°S–5°N) (red box in Figure 9b).
The amplitude of this specified SSTA in the WNP and MC
regions is 2–3 times larger than the composite SSTA field
associated with the observed high WNPSH indices, in order
to obtain a robust response. The WNP and MC runs each
consists of seven ensemble members which have initial
conditions starting from 27 April 0000 UTC to 30 April
0000 UTC, separating by 12 h apart. All experiments were
integrated from May to August. The atmospheric response
to the local and remote SSTA forcing is estimated based on
the difference between the ensemble mean of the WNP/MC
run and the CTRL run (i.e., WNP‐CTRL and MC‐CTRL).
[37] The anomalous summertime (JJA mean) precipitation
and y850 field responses to the WNP SSTA are shown in
Figure 9a. The imposed local cold SSTA causes the decrease
of local precipitation to the southwest and an increase of
precipitation in the tropical central western Pacific along
5°N. An anomalous low‐level anticyclone forms to the east of
Taiwan in response to the cold SSTA forcing. The simulated
patterns are qualitatively consistent with the observations
from the reanalysis data sets (Figures 3f, 5f, and 7b), but
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Figure 8. Vertical profiles of regressed (a) vertical p velocity, (c) Q1, and (d) Q2 averaged over the
region of strongest descending motion and (b) relative vorticity averaged over the region of anomalous
WNPSH center for the 2–3 year (red) and 3–5 year (blue) oscillations.
the amplitude of the simulated precipitation and anticyclonic
circulation is too strong, due to the greater SSTA specified.
In the observation, the SSTA in the WNP decreases gradually from June to August [Wu et al., 2010].
[38] The strong positive SSTA forcing in the MC induces
a positive precipitation anomaly over the northern part of the
SSTA domain and a negative precipitation anomaly in the
northern Philippines along 15°N (Figure 9b). An anomalous
anticyclone is simulated near Taiwan and is located to the
north of the negative precipitation anomaly center, similar to
the observed anticyclone‐vertical motion phase relation. In
the tropical western Pacific, anomalous easterly flows in the
southern part of the anticyclone extend to the north IO and
converge toward the MC region. The simulated patterns in
the MC run are quite similar to those observed shown in
Figure 3c, 5c, and 7a but with less zonal extent and a
slightly northward shift due to the neglect of tropical central
Pacific cold SSTA forcing in Figure 4a.
[39] It is seen from Figures 9a and 9b that the spatial
phase patterns between the anomalous precipitation center
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and the anticyclone center in both the WNP and MC runs
are similar to those observed. In addition, the vertical profiles of the vertical velocity and relative vorticity fields
averaged over the maximum precipitation and anticyclonic
vorticity center in both the WNP and MC runs were evaluated. In Figure 9c the anomalous descending motion in the
MC run is larger than that in the WNP run. The strongest
descending motion is located in the middle troposphere
around 400–600 hPa in the MC run but at 500–700 hPa in
the WNP run. In the lower troposphere below 850 hPa,
the amplitude of the descent in the WNP and MC run is
slightly reversed as in the observation of NCEP/NCAR R‐1
(Figure 8a). In addition to the p velocity, the vertical distributions of relative vorticity over the anticyclone center in
the WNP and MC run also shows a distinct difference,
similar to the observation. The anomalous anticyclone in the
MC run (Figure 9d) is associated with the negative vorticity
throughout the troposphere that resembles the observed
barotropic structure in the 2–3 year oscillation (Figure 8b).
In the WNP run (Figure 9d), the negative vorticity in the
lower troposphere changes sign to a positive one at 400 hPa,
consistent with the observed baroclinic structure in the 3–
5 year oscillation (Figure 8b), although the baroclinicity in
the model is a little stronger due to neglecting other minor
remote SSTA forcing in the simulation. The vertical distribution of the specific humidity (figure not shown) also
resembles similar result in observed Q2 (Figure 8d) that the
amplitude of the negative specific humidity anomaly in the
lower troposphere is greater in the WNP run than in the MC
run.
[40] The horizontal and vertical structures of the anomalous circulation in the idealized numerical experiments
accurately capture the observed characteristics in the 2–
3 year and 3–5 year oscillations. Numerical experiments in
the study demonstrate the role of SSTA forcing in the WNP
and MC regions in generating the anomalous anticyclone
associated with the westward extension of the WNPSH
in the two oscillations. The baroclinic vertical structure in
the WNP run confirms that a Rossby wave response to the
local cold SSTA is a key process involved in the 3–5 year
oscillation. Since the cold SSTA in the WNP weakens from
early to late summer when El Nino is decaying in the 3–
5 year oscillation, the SSTA forcing in the tropical IO may
also contribute to the westward movement of the subtropical
high [Xie et al., 2009; Wu et al., 2010]. The numerical
experiment of Zhou et al. [2009b] has also shown that a
warmer tropical IO may contribute to the westward extension WNPSH in the interdecadal time scale. In an additional
idealized experiment with a +1°C SSTA specified in the
tropical IO basin (20°S–20°N, figure not shown), an
anomalous low‐level anticyclonic circulation is simulated in
the WNP, with a maximum center located slightly southwestward over the South China Sea. Since the simulated
anticyclone differs from the observed WNPSH in its location (a southwestward shift from the observed center) and
vertical structure (the simulated baroclinic structure is much
weaker), the warm SSTA in the tropical IO may not be the
major forcing in causing the westward shift of the WNPSH
in the 3–5 year oscillation. In the MC run, the simulated
horizontal pattern and the vertical structure resemble the
observed in the 2–3 year oscillation. This confirms the role
of the remote warm SSTA forcing in the MC in inducing the
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anomalous WNPSH through an anomalous local Hadley
circulation as discussed in section 3.
5. Temporal Evolution Features
[41] To better understand the interannual variability of the
WNPSH, the seasonal evolution of anomalous SST and
circulation fields of w500, y850, and V850 associated with the
WNPSH 2–3 year and 3–5 year oscillations are examined.
More than one full cycle (from MAM(0) to JJA(2)) and
nearly one full cycle (from MAM(0) to JJA(3)) of regressed
fields are presented for 2–3 year and 3–5 year oscillations.
Seasons of MAM(1), JJA(1), SON(1) denote the concurrent
spring, summer and autumn of the westward WNPSH,
whereas the numbers 0, 2, and 3 in parentheses denote the
lead and lag years relative to the current year (1).
5.1. SSTA and Vertical Velocity
[42] The evolution of regressed SSTA for the 2–3 year
oscillation (Figure 10a) shows the following features.
Starting from the preceding spring MAM(0), cold SST
anomalies dominated the northern IO, MC, and northern
Australia, while a weak warm SSTA appeared in the equatorial central Pacific. The warm SSTA in the central Pacific
moved eastward and intensified from JJA(0) to SON(0),
weakened afterward through MAM(1), and became negative by the current JJA(1). After JJA(1), the cold SSTA
strengthened until D(1)JF(2) and then decayed from MAM(2)
to JJA(2) when the cold SSTA turned into a weak warm
SSTA. Accompanying the above evolution, a weak warm
SSTA in the central IO in the previous summer JJA(0) grew
through SON(0). By D(0)JF(1), the warm SSTA had spread
widely to the northern IO, MC, and northern Australia. The
broad‐scale warm SSTA further extended southeastward
from D(0)JF(1) to MAM(1) and eastward from MAM(1) to
SON(1). By the following summer JJA(2), the SSTA had
evolved to a pattern similar to that in MAM(0). The above
evolution feature resembles the quasi‐biennial oscillation of
the tropical SSTA signal found in previous studies [e.g.,
Barnett, 1991; Li et al., 2006].
[43] Contrary to the 2–3 year oscillation, the evolution
of the 3–5 year oscillation shows a standing oscillation,
with the SSTA persisting from MAM(0) to SON(1) and
reversing its sign from D(1)JF(2) to JJA(3) (Figure 10b). A
warm SSTA developed in the central eastern Pacific from
MAM(0), reached its warmest phase in D(0)JF(1), and then
weakened in the following three seasons to D(1)JF(2) when
a weak cold SSTA grew. The cold SSTA became stronger
from MAM(2) to D(2)JF(3), weakened in MAM(3), and
continued to last for two seasons until SON(3) (figure not
shown). Accompanying the warm SSTA in the equatorial
Pacific, the surrounding cold SSTA to the northwest and
southwest also persisted from JJA(0) to SON(1). In D(1)JF(2),
the areas surrounding the equatorial Pacific were replaced
by a warm SSTA that can persist to JJA(3). In addition, warm
SST anomalies developed in the IO from D(0)JF(1) strengthened and extended to the WNP and marginal seas along the
East Asia coast through MAM(1) to SON(1). In the following
seasons, weak cold SST anomalies appeared in the IO from
MAM(2), extended to the north and south from JJA(2) to D(2)
JF(3), and reached the strongest intensity over the whole
ocean basin in JJA(3). The evolution of the regressed SSTA
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Figure 9. Anomalous precipitation (shaded, only areas passing the 90% confidence level plotted) and
850 hPa stream function in (a) the WNP run and (b) the MC run, and vertical profiles of (c) anomalous
vertical p velocity over the strongest descending region and (d) anomalous relative vorticity over the
anticyclone center in the WNP (blue line) and the MC (red line) runs. The thick blue and red box in
Figures 9a and 9b shows the region where uniform (either −1°C or +1°C) SSTA is specified.
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patterns described above resembles a typical El Nino evolution with a long life cycle [e.g., Wang et al., 2000, 2003].
The 3–5 year oscillation of the WNPSH may be modulated
by the remote low frequency El Nino SSTA forcing [Wu and
Zhou, 2008] via local air‐sea coupling in WNP [Wang et al.,
2000; Sui et al., 2007].
[44] The time‐longitude section of the regressed SST and
w500 fields along the Equator from MAM(0) to JJA(2) (to
JJA(3) in 3–5 year oscillation) is shown in Figure 11. For
the 2–3 year oscillations (Figures 11a and 11c), there is an
obvious eastward propagation of the warm SSTA and
upward motion from the IO to the western Pacific during the
period from JJA(0) to JJA(2). The warm SSTA propagated
to the tropical western Pacific in D(0)JF(1) where the
upward motion developed two seasons later in JJA(1). In the
central eastern Pacific, warm SSTA (upward motion) existed
from MAM(0) to MAM(1) and was followed by a cold
SSTA (downward motion) from JJA(1) to JJA(2) with a
slight eastward propagation. The center of the upward
motion within (170°W–150°W) is located to the west of the
warm SSTA in the region (150°W–120°W) over the tropical
central eastern Pacific. However, over the tropical western
Pacific, the two fields do not show a significant phase
difference.
[45] The regressed fields for the 3–5 year oscillation
(Figures 11b and 11d) exhibit a standing sandwich pattern
characterized by a warm (cold) SSTA and upward (downward) motion in the IO, a cold (warm) SSTA and downward
(upward) motion in the western Pacific, and a warm (cold)
SSTA with upward (downward) motion in the central eastern Pacific from MAM(0) to D(1)JF(2) [from MAM(2) to
SON(3)]. Although warm (cold) SST anomalies are generally accompanied by upward (downward) motion in the
ocean basins, this feature is not so clear in the IO where a
significant SSTA is only associated with weak vertical
motion. In the tropical western Pacific, the strongest descending (ascending) motion lagged the strongest cold
(warm) SSTA by 1 month, developing in JJA(0) (MAM(2)),
reaching the strongest intensity in D(0)JF(1) [D(2)JF(3)],
and then persisting to JJA(1) (JJA(3)).
[46] The spatial and temporal evolutions of the SSTA in
the 2–3 year and 3–5 year oscillations are consistent with the
study by Barnett [1991], who found that the near‐equatorial
characteristics of the quasi‐biennial SST/SLP mode are
those of a quasi‐progressive wave while the 3–7 year mode
is closer to a standing wave.
5.2. Stream Function and Wind
[47] The examination of temporal evolution of the
regressed c200 (figures not shown) and w500 (Figure 11c)
from JJA(0) to SON(1) indicates that in the 2–3 year oscillation, a large‐scale upper‐level divergence (ascending) was
located in the eastern Pacific and northern IO from JJA(0)
to MAM(1), while a large scale upper‐level convergence
(descending) appeared over the western Pacific. The pattern
was nearly out of phase in JJA(1), with a dominant upper‐
level divergence near Java and a large‐scale convergence
over the tropical central eastern Pacific (Figure 5a). It should
be noted that the c200 pattern has no significant signal in the
transition season MAM(1). This implies that the signal from
the preceding boreal summer JJA(0) to winter D(0)JF(1) may
have difficulty to pass through the spring season MAM(1)
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on the 2–3 year oscillation. This differs from the 3–5 year
oscillation, in which the large‐scale upper‐layer convergent
flow over the tropical western Pacific persisted from the
preceding summer JJA(0) to the current summer JJA(1).
[48] The evolutions of the regressed y850 and V850 fields
(Figure 12) are, in general, consistent with the c200 field. In
the 2–3 year oscillation (Figure 12a), a strong anomalous
anticyclone existed over southern Asia in SON(0) and
moved to the Philippine Sea in D(0)JF(1). The anomalous
anticyclone weakened and disappeared in MAM(1) but
reemerged over the WNP in JJA(1), again indicating a spring
barrier for the Philippine Sea anticyclone to pass through the
boreal spring on this timescale. The local Hadley circulation
in JJA(1), therefore, is essential for the development of the
summer anticyclonic circulation over the WNP on this
timescale. The evolution is also consistent with the vertical
velocity fields that descending motion within the tropical
western Pacific weakens in MAM(1), where the ascending
motion becomes stronger and statistically significant in JJA
(1) within the green box in Figures 11a and 11c.
[49] In the 3–5 year oscillation (Figure 12b), however,
the anomalous anticyclone over the Philippine Sea and
low‐level northeasterly wind over WNP persisted from
D(0)JF(1) to JJA(1), consistent with the c200, SSTA and
vertical velocity pattern in Figure 10b and Figures 11b and
11d. A positive thermodynamic air‐sea feedback can
explain how the anomalous anticyclone is maintained from
the preceding winter to the current summer [Wang et al.,
2000, 2003; Sui et al., 2007]. Here the anomalous anticyclone, initiated by a cold SSTA as a Rossby wave response
to a negative heat source in the boreal winter, may further
reinforce the cold SSTA through the wind‐evaporation‐
SST feedback under prevailing climatological northeasterly
trade winds to the spring. Although the seasonal change
of prevailing winds leads to a negative air‐sea feedback
and thus a weakening of the local cold SSTA in the boreal
summer, the negative SSTA may still have a significant
impact on the WNP monsoon, particularly in early summer.
Thus this local cold SSTA may have a large impact on the
summertime WNPSH variability.
6. Conclusion and Discussion
[50] The western edge of the summertime Pacific subtropical high in the Northern Hemisphere (WNPSH) exhibits
a significant zonal shift on the interannual timescale. The
largest variability is found in the western North Pacific. This
variability is represented in this study by an index defined by
geopotential height anomalies at 500 hPa averaged in the
region where the largest WNPSH variability occurs. A power
spectrum analysis of the WNPSH index indicates two
dominant peaks at 2–3 year and 3–5 year periods. The
oscillation characteristics of the WNPSH at the two periods
are investigated through a regression analysis of atmospheric
and oceanic variables against the two WNPSH indices at the
2–3 year and 3–5 year periods. The idealized numerical
experiments with the ECHAM4 model are also carried out to
examine the dominated SSTA and associated mechanisms
for the WNPSH variability on the two timescales. Since the
location and strength of the subtropical anticyclone greatly
change from June to August, the conclusion obtained from
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Figure 10. (a) Evolutions of the SSTA field from the preceding spring (MAM(0)) to the summer of the
following year (JJA(2)) regressed against the WNPSH index for the 2–3 year oscillation. (b) Evolutions of
the SSTA field from the preceding spring (MAM(0)) to the summer of the year after next year (JJA(3))
regressed against the WNPSH index for the 3–5 year oscillation.
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Figure 11. (a, b) Longitude‐time section of regressed SSTA and (c, d) 500 hPa vertical p velocity
anomaly fields averaged between 15°S and 15°N from MAM(0) to JJA(2) for 2–3 year oscillation
(Figures 11a and 11c) and from MAM(0) to JJA(3) for 3–5 year oscillation (Figures 11b and 11d).
Contours in warm (cold) colors indicate warm (cold) SSTA or downward (upward) motion. Shaded
regions pass the 90% confidence level. Green boxes denote the current summer of a strengthened
WNPSH.
this study may not be applicable to individual summer
months.
[51] The 2–3 year oscillation of the WNPSH and the
associated SST and circulation possess a coherent eastward
propagating feature. The primary signals include the
southeastward spread of a warm SSTA and ascending
motion from the Indian Ocean in the preceding autumn,
SON(0), to the maritime continent in the current summer
JJA(1) of the high WNPSH index. In the central and eastern
Pacific, the SSTA changes its sign from positive in the
preceding summer through winter to negative in JJA(1),
analogous to the La Nina development condition. In the
current summer, the WNPSH is characterized by an anomalous anticyclonic circulation and descending motion north
of the Philippines. The descent is a part of the anomalous
local Hadley circulation, with enhanced convection and a
warm SSTA appearing in the maritime continent. A heat
budget analysis reveals that the descent and anticyclone are
maintained primarily by radiative cooling, and the corre-
sponding vertical structure is equivalently barotropic. The
anticyclonic circulation in the WNP is accompanied by an
anomalous cyclonic circulation to its north. As a result,
the overall structure appears like a meridionally oriented PJ
type pattern. The role of the remote SSTA forcing in the
maritime continent in affecting the WNPSH variability is
verified by the idealized ECHAM4 model experiment.
[52] The 3–5 year oscillation of the WNPSH exhibits a
standing wave feature, that the warm SSTA in the equatorial
central eastern Pacific, its surrounding cold SSTA to the
northwest and southwest, and a warm SSTA in the Indian
Ocean and along the East Asian marginal seas, persisting
from the preceding winter to the current summer of a high
WNPSH index. Along with the evolution above, the
ascending motion over the tropical eastern central Pacific
and Indian Ocean and the descending motion over the
western Pacific also persist from the preceding autumn to
the current summer. These features resemble the tripole
SSTA pattern and associated overturning Walker circulation
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Figure 12. Regressed 850 hPa stream function and wind fields from the preceding summer (JJA(0)) to
the current autumn (SON(1)) for (a) the 2–3 year and (b) 3–5 year oscillations of WNPSH. Shadings and
vectors denote regions passing the 90% confidence level.
pattern frequently observed during the decaying phase of
El Nino episodes [e.g., Wang et al., 2000, 2003]. For the 3–
5 year oscillation, the maximum descent of the WNP is
slightly east of the anomalous anticyclone center. The
greatest descent appears in the lower troposphere, accompanied by a negative latent heating and a cold SSTA to the
southeast. The corresponding vertical structure of relative
vorticity in the descent region bears a baroclinic nature,
and the zonal phase difference between the anomalous
anticyclone center and the descent implies a Rossby wave
response to a negative heating anomaly in association with
the local cold SSTA. The local SSTA forcing mechanism is
confirmed by the idealized ECHAM4 model experiment.
[53] Our findings of the 2–3 year and 3–5 year oscillations
of the WNPSH are consistent with previous studies of
Barnett [1991] and Wang et al. [2001] who found that there
are quasi‐biennial and low‐frequency variabilities in the
tropical SST/SLP fields and in Asian monsoon circulation
fields. Our findings are also consistent with the observational fact that ENSO has both the quasi‐biennial and lower‐
frequency spectral peaks. The consistency above indicates
that the variability of the WNPSH connected with the Asian
Monsoon system might be closely linked to the tropical
SSTA forcing in the interannual time scale. The possible
mechanisms that connect the subtropical high and tropical
SST include the monsoon‐warm ocean interaction associated with the TBO [Li et al., 2006], ENSO‐East Asia teleconnection [Wang et al., 2000; Sui et al., 2007], and a
season‐dependent Indian Ocean impact on the WNP monsoon [Wu et al., 2009b].
[54] In addition to the tropical forcing, midlatitude processes may also play a role in the interannual variability of
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the WNPSH [e.g., Enomoto, 2004]. Since the present study
focus on the relationship between the WNPSH and tropical
forcing, the influence of the midlatitude processes have not
been discussed here. Our ECHAM4 model simulation
confirms that the 2–3 year mode is driven by the SSTA
forcing over the maritime continent, while the 3–5 year
mode is driven by the local SSTA in the WNP. Despite the
model results in support of the hypothesized physical
mechanisms in the 2–3 year and 3–5 year oscillations, the
role of the air‐sea interaction in the WNP region, where the
atmospheric feedback to the SSTA is regarded as an
important process, have not been illustrated. Further observational analyses and coupled GCM modeling studies are
needed to investigate the relative importance of these factors
and their causality in causing the interannual variability of
the WNPSH from dynamic and thermodynamic discussions.
[55] Acknowledgments. We appreciate very much for the constructive comments from three anonymous reviewers. This research
was supported by NSC 98‐2111‐M‐133‐002 and NSC 99‐2111‐M‐
133‐001. C.H.S. was supported by NSC 97‐2111‐M‐008‐010‐MY2
and NSC 99‐2745‐M‐002‐001‐ASP. T.L. was supported by ONR grants
N000140810256 and N000141010774 and the International Pacific
Research Center that is sponsored by the Japan Agency for Marine‐
Earth Science and Technology (JAMSTEC), NASA (NNX07AG53G),
and NOAA (NA17RJ1230). This is SOEST contribution 8127 and IPRC
publication 773.
References
Annamalai, H., J. M. Slingo, K. R. Sperber, and K. Hodges (1999), The mean
evolution and variability of the Asian summer monsoon: Comparison
of ECMWF and NCEP–NCAR reanalyses, Mon. Weather Rev., 127,
1157–1186, doi:10.1175/1520-0493(1999)127<1157:TMEAVO>2.0.CO;2.
Barnett, T. P. (1991), The interaction of multiple time scales in the tropical
climate system, J. Clim., 4, 269–285, doi:10.1175/1520-0442(1991)
004<0269:TIOMTS>2.0.CO;2.
Chang, C.‐P., and T. Li (2000), A theory for the tropical tropospheric biennial oscillation, J. Atmos. Sci., 57, 2209–2224, doi:10.1175/1520-0469
(2000)057<2209:ATFTTT>2.0.CO;2.
Chang, C.‐P., Y. Zhang, and T. Li (2000a), Interannual and interdecadal
variations of the East Asian summer monsoon and tropical Pacific SSTs.
Part 1: Role of the subtropical ridge, J. Clim., 13, 4310–4325,
doi:10.1175/1520-0442(2000)013<4310:IAIVOT>2.0.CO;2.
Chang, C.‐P., Y. Zhang, and T. Li (2000b), Interannual and interdecadal
variations of the East Asian summer monsoon and tropical Pacific SSTs.
Part 2: Southeast China rainfall and meridional structure, J. Clim., 13,
4326–4340, doi:10.1175/1520-0442(2000)013<4326:IAIVOT>2.0.
CO;2.
Chen, P., M. P. Hoerling, and R. M. Dole (2001), The origin of the subtropical anticyclones, J. Atmos. Sci., 58, 1827–1835, doi:10.1175/1520-0469
(2001)058<1827:TOOTSA>2.0.CO;2.
Chen, J.‐M., T. Li, and J. Shih (2007), Fall persistence barrier of sea surface temperature in the South China Sea associated with ENSO, J. Clim.,
20, 158–172, doi:10.1175/JCLI4000.1.
Chou, C. (2004), Establishment of the low‐level wind anomalies over
the western North Pacific during ENSO development, J. Clim., 17,
2195–2212, doi:10.1175/1520-0442(2004)017<2195:EOTLWA>2.0.
CO;2.
Chou, C., J. Y. Tu, and J. Y. Yu (2003), Interannual variability of the western North Pacific summer monsoon: Differences between ENSO and
non‐ENSO years, J. Clim., 16, 2275–2287, doi:10.1175/2761.1.
Enomoto, T. (2004), Interannual variability of the Bonin High associated
with the propagation of Rossby wave along the Asian jet, J. Meteorol.
Soc. Jpn., 82, 1019–1034, doi:10.2151/jmsj.2004.1019.
Gibson, J. K., P. Kallberg, and S. Uppala (1996), The ECMWF Reanalysis
(ERA) Project, ECMWF Newsl., 73, 7–17.
Gibson, J. K., A. Hernandez, A. Nomura, and E. Serrano (1997), Reanal.
Proj. Rep. Ser. 1, 72 pp., Eur. Cent. for Medium Range Weather Forecast.,
Reading, U. K.
Gill, A. E. (1980), Some simple solutions for heat‐induced tropical circulation, Q. J. R. Meteorol. Soc., 106, 447–462, doi:10.1002/qj.49710644905.
D13111
Hoskins, B. J. (1996), On the existence and strength of the summer subtropical anticyclones, Bull. Am. Meteorol. Soc., 77, 1287–1292.
Kalnay, E., et al. (1996), The NCEP/NCAR 40‐Year Reanalysis Project,
Bull. Am. Meteorol. Soc., 77, 437–471, doi:10.1175/1520-0477(1996)
077<0437:TNYRP>2.0.CO;2.
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.
Am. Meteorol. Soc., 83, 1631–1643, doi:10.1175/BAMS-83-11-1631.
Kawamura, R., T. Matsuura, and S. Iisuka (2001), Interannual atmosphere‐
ocean variations in the tropical western North Pacific relevant to the
Asian summer monsoon‐ENSO coupling, J. Meteorol. Soc. Jpn., 79,
883–898, doi:10.2151/jmsj.79.883.
Lau, N.‐C., and M. J. Nath (2000), Impact of ENSO on the variability of
the Asian‐Australian monsoons as simulated in GCM experiments,
J. Clim., 13, 4287–4309, doi:10.1175/1520-0442(2000)013<4287:
IOEOTV>2.0.CO;2.
Lau, K.‐M., and H. T. Wu (2001), Principal modes of rainfall‐SST variability of the Asian summer monsoon: A reassessment of the monsoon‐
ENSO relationship, J. Clim., 14, 2880–2895, doi:10.1175/1520-0442
(2001)014<2880:PMORSV>2.0.CO;2.
Lee, E. J., J. G. Jhun, and C. K. Park (2005), Remote connection of the
Northeast Asian summer rainfall variation revealed by a newly defined
monsoon index, J. Clim., 18, 4381–4393, doi:10.1175/JCLI3545.1.
Li, T., and B. Wang (2005), A review on the western North Pacific monsoon: Synoptic‐to‐interannual variabilities, Terr. Atmos. Ocean. Sci., 16,
285–314.
Li, T., and Y. S. Zhang (2002), Processes that determine the quasibiennial
and lower‐frequency variability of the south Asian monsoon, J. Meteorol.
Soc. Jpn., 80, 1149–1163, doi:10.2151/jmsj.80.1149.
Li, T., Y. S. Zhang, C.‐P. Chang, and B. Wang (2001), On the relationship
between Indian Ocean SST and Asian summer monsoon, Geophys. Res.
Lett., 28, 2843–2846, doi:10.1029/2000GL011847.
Li, T., Y. S. Zhang, E. Lu, and D. Wang (2002), Relative role of dynamic
and thermodynamic processes in the development of the Indian Ocean
dipole, Geophys. Res. Lett., 29(23), 2110, doi:10.1029/2002GL015789.
Li, T., B. Wang, C.‐P. Chang, and Y. Zhang (2003), A theory for the Indian
Ocean dipole‐zonal mode, J. Atmos. Sci., 60, 2119–2135, doi:10.1175/
1520-0469(2003)060<2119:ATFTIO>2.0.CO;2.
Li, T., P. Liu, X. Fu, B. Wang, and G. A. Meehl (2006), Tempo‐spatial
structures and mechanisms of the tropospheric biennial oscillation in
the Indo‐Pacific warm ocean region, J. Clim., 19, 3070–3087,
doi:10.1175/JCLI3736.1.
Liu, Y., and G. Wu (2004), Progress in the study on the formation of the
summertime subtropical anticyclone, Adv. Atmos. Sci., 21, 322–342,
doi:10.1007/BF02915562.
Liu, Y. M., G. X. Wu, H. Liu, and P. Liu (2001), Condensation heating
of the Asian summer monsoon and the subtropical anticyclone in
the Eastern Hemisphere, Clim. Dyn., 17, 327–338, doi:10.1007/
s003820000117.
Liu, Y. M., G. X. Wu, and R. C. Ren (2004), Relationship between the subtropical anticyclone and diabatic heating, J. Clim., 17, 682–698,
doi:10.1175/1520-0442(2004)017<0682:RBTSAA>2.0.CO;2.
Lu, R. (2001), Interannual variability of the summertime North Pacific subtropical high and its relation to atmospheric convection over warm pool,
J. Meteorol. Soc. Jpn., 79, 771–783, doi:10.2151/jmsj.79.771.
Lu, R., and B. Dong (2001), Westward extension of North Pacific subtropical high in summer, J. Meteorol. Soc. Jpn., 79, 1229–1241, doi:10.2151/
jmsj.79.1229.
Meehl, G. A. (1987), The annual cycle and interannual variability in
the tropical Pacific and Indian Ocean region, Mon. Weather Rev., 115,
27–50, doi:10.1175/1520-0493(1987)115<0027:TACAIV>2.0.CO;2.
Meehl, G. A. (1994), Coupled land‐ocean‐atmosphere processes and
south Asian monsoon variability, Science, 266, 263–267, doi:10.1126/
science.266.5183.263.
Meehl, G. A. (1997), The south Asian monsoon and the tropospheric biennial oscillation, J. Clim., 10, 1921–1943, doi:10.1175/1520-0442(1997)
010<1921:TSAMAT>2.0.CO;2.
Miyasaka, T., and H. Nakamura (2005), Structure and formation mechanisms of the Northern Hemisphere summertime subtropical highs, J. Clim.,
18, 5046–5065, doi:10.1175/JCLI3599.1.
Nitta, T. (1987), Convective activities in the tropical western Pacific and
their impact on the Northern Hemisphere summer circulation, J. Meteorol.
Soc. Jpn., 65, 373–389.
Rodwell, M. J., and B. J. Hoskins (1996), Monsoons and the dynamics of
deserts, Q. J. R. Meteorol. Soc., 122, 1385–1404, doi:10.1002/
qj.49712253408.
Rodwell, M. J., and B. J. Hoskins (2001), Subtropical anticyclones and
summer monsoon, J. Clim., 14, 3192–3211, doi:10.1175/1520-0442
(2001)014<3192:SAASM>2.0.CO;2.
18 of 19
D13111
CHUNG ET AL.: TROPICAL SSTA AND SUBTROPICAL HIGH
Roeckner, E., et al. (1996), The atmospheric general circulation model
ECHAM‐4: Model description and simulation of present‐day climate,
Rep. 218, 90 pp., Max‐Planck‐Inst. for Meteorol., Hamburg, Germany.
Seager, R., R. Murtugudde, N. Naik, A. Clement, N. Gordon, and J. Miller
(2003), Air‐sea interaction and the seasonal cycle of the subtropical
anticyclones, J. Clim., 16, 1948–1966, doi:10.1175/1520-0442(2003)
016<1948:AIATSC>2.0.CO;2.
Shen, S., and K.‐M. Lau (1995), Biennial oscillation associated with the
East Asian summer monsoon and tropical Pacific sea surface temperatures, J. Meteorol. Soc. Jpn., 73, 105–124.
Smith, T. M., and R. W. Reynolds (2004), Improved extended reconstruction of SST (1854–1997), J. Clim., 17, 2466–2477, doi:10.1175/15200442(2004)017<2466:IEROS>2.0.CO;2.
Sui, C.‐H., P. H. Chung, and T. Li (2007), Interannual and interdecadal variability of the summertime western North Pacific subtropical high, Geophys.
Res. Lett., 34, L11701, doi:10.1029/2006GL029204.
Tao, S., and L. Chen (1987), A review of recent research on the East Asian
summer monsoon in China, in Monsoon Meteorology, edited by C.‐P.
Chang and T. N. Krishuamurti, pp. 60–92, Oxford Univ. Press, New York.
Ting, M. (1994), Maintenance of northern summer stationary waves in a
GCM, J. Atmos. Sci., 51, 3286–3308, doi:10.1175/1520-0469(1994)
051<3286:MONSSW>2.0.CO;2.
Tomita, T., T. Yoshikane, and T. Yasunari (2004), Biennial and lower‐
frequency variability observed in the early summer climate in the western
North Pacific, J. Clim., 17, 4254–4266, doi:10.1175/JCLI3200.1.
Wang, B., R. Wu, and X. Fu (2000), Pacific‐East Asian teleconnection:
How does ENSO affect East Asian climate?, J. Clim., 13, 1517–1536,
doi:10.1175/1520-0442(2000)013<1517:PEATHD>2.0.CO;2.
Wang, B., R. Wu, and K.‐M. Lau (2001), Interannual variability of the
Asian summer monsoon: Contrasts between the Indian and the western
North Pacific–East Asian monsoons, J. Clim., 14, 4073–4090,
doi:10.1175/1520-0442(2001)014<4073:IVOTAS>2.0.CO;2.
Wang, B., R. Wu, and T. Li (2003), Atmosphere‐warm ocean interaction
and its impact on Asian‐Australian monsoon variation, J. Clim., 16,
1195–1211, doi:10.1175/1520-0442(2003)16<1195:AOIAII>2.0.CO;2.
Wang, B., Q. Ding, X. Fu, I.‐S. Kang, K. Jin, J. Shukla, and F. Doblas‐
Reyes (2005), Fundamental challenge in simulation and prediction of
summer monsoon rainfall, Geophys. Res. Lett., 32, L15711,
doi:10.1029/2005GL022734.
Watanabe, M., and F. F. Jin (2002), Role of Indian Ocean warming in the
development of Philippine Sea anticyclone during ENSO, Geophys. Res.
Lett., 29(10), 1478, doi:10.1029/2001GL014318.
Wu, R., and B. P. Kirtman (2005), Roles of Indian and Pacific Ocean air‐
sea coupling in tropical atmospheric variability, Clim. Dyn., 25, 155–170,
doi:10.1007/s00382-005-0003-x.
Wu, G. X., and Y. M. Liu (2003), Summertime quadruplet heating pattern
in the subtropics and the associated atmospheric circulation, Geophys.
Res. Lett., 30(5), 1201, doi:10.1029/2002GL016209.
Wu, B., and T. 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.
Wu, G. X., Y. Liu, X. Zhu, W. Li, R. Ren, A. Duan, and X. Liang (2009),
Multi‐scale forcing and the formation of subtropical desert and monsoon,
Ann. Geophys., 27, 3631–3644, doi:10.5194/angeo-27-3631-2009.
D13111
Wu, B., T. Zhou, and T. Li (2009a), Contrast of rainfall‐SST relationships
in the western North Pacific between the ENSO developing and decaying
summers, J. Clim., 22, 4398–4405, doi:10.1175/2009JCLI2648.1.
Wu, B., T. Zhou, and T. Li (2009b), Seasonally evolving dominant interannual variability mode over the East Asia, J. Clim., 22, 2992–3005,
doi:10.1175/2008JCLI2710.1.
Wu, B., T. Li, and T. Zhou (2010), Relative contributions of the Indian
Ocean and local SST anomalies to the maintenance of the western North
Pacific anomalous anticyclone during the El Nino decaying summer,
J. Clim., 23, 2974–2986, doi:10.1175/2010JCLI3300.1.
Xie, S. P., H. Jan, T. Hiroki, D. Yan, T. Sampe, K. Hu, and G. Huang
(2009), Indian Ocean capacitor effect on Indo–Western Pacific climate
during the summer following El Niño, J. Clim., 22, 730–747,
doi:10.1175/2008JCLI2544.1.
Yanai, M., and R. H. Johnson (1993), Impact of cumulus convection on
thermodynamic fields, in The Representation of Cumulus Convection
in Numerical Models, Meteorol. Monogr., vol. 46, edited by K. Emanuel
and D. J. Raymond, pp. 36–62, Am. Meteorol. Soc., Boston, Mass.
Yanai, M., S. Esbensen, and J. H. Chu (1973), Determination of bulk properties of tropical cloud clusters from large‐scale heat and moisture budgets, J. Atmos. Sci., 30, 611–627, doi:10.1175/1520-0469(1973)
030<0611:DOBPOT>2.0.CO;2.
Yang, J. L., Q. Y. Li, 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, J.‐Y., C. R. Mechoso, A. Arakawa, and J. C. Williams (2002), Impacts
of the Indian Ocean on the ENSO cycle, Geophys. Res. Lett., 29(8), 1204,
doi:10.1029/2001GL014098.
Zhou, T. J., and Z. Li (2002), Simulation of the East Asian summer monsoon by using a variable resolution atmospheric GCM, Clim. Dyn., 19,
167–180, doi:10.1007/s00382-001-0214-8.
Zhou, T. J., and R. Yu (2005), Atmospheric water vapor transport associated
with typical anomalous summer rainfall patterns in China, J. Geophys.
Res., 110, D08104, doi:10.1029/2004JD005413.
Zhou, T., and J. Zhang (2011), The vertical structures of atmospheric temperature anomalies associated with two flavors of El Nino simulated by
AMIP II models, J. Clim., 24, 1053–1070, doi:10.1175/2010JCLI3504.1.
Zhou, T. J., R. Yu, H. Li, and B. Wang (2008), Ocean forcing to changes in
global monsoon precipitation over the recent half‐century, J. Clim., 21,
3833–3852, doi:10.1175/2008JCLI2067.1.
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. Clim., 22, 1159–1173,
doi:10.1175/2008JCLI2245.1.
Zhou, T., et al. (2009b), Why the western Pacific subtropical high has
extended westward since the late 1970s, J. Clim., 22, 2199–2215,
doi:10.1175/2008JCLI2527.1.
P.‐H. Chung, Department of History and Geography, Taipei Municipal
University of Education, Taipei 10042, Taiwan. ([email protected])
T. Li, International Pacific Research Center, University of Hawai’i at
Mānoa, Honolulu, HI 96822, USA.
C.‐H. Sui, Department of Atmospheric Sciences, National Taiwan
University, Taipei 10617, Taiwan.
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