Variation of the Asian Monsoon Water Vapor Budget:
Interaction with the Global-Scale Modes
Tsing-Chang Chen
(submitted: March 8, 2005)
1. Introduction
More than half of the world’s population resides in the Asian monsoon region. Because of its
importance to the well-being of such a large population, various aspects of monsoon rainfall have been the
focus of Asian monsoon studies. Thus, numerous research efforts were attracted to perform regional water
vapor budget analyses over different areas: Japan (e.g. Yoshino 1965, 1966, Akiyama 1975), East Asia (e.g.
Chen et al. 1988a), the South China Sea (e.g. Lau et al. 1998), the Tibetan Plateau (e.g. Luo and Yanai 1983;
Bollasina and Benedict 2004), Indochina Pensusula (e.g. Kiguchi and Matsumoto 2005), and other regions
within the Asian monsoon system. In order to understand the water supply through monsoon rainfall, the
regional hydrological cycles/water-vapor budgets were stressed by the monsoon research. Although these
regional hydrological cycles/water-vapor budgets in the Asian monsoon system are part of the global
hydrological cycle, we are still lacking a global perspective of the role played by the Asian monsoon in this
hydrological cycle. On the other hand, it was uncommon for the study of the global hydrological
cycle/water vapor budget analysis (e.g. Chen and Pfaendfner 1993; Ziegler et al. 2003) to explore its effect
on the regional hydrological cycle. Because the Asian summer monsoon is embedded in the global
circulation, the regional hydrological cycles/hydrological processes in Asia must be linked to the global
circulation through some special relationship. The concern with the relationship between the regional and
global hydrological cycle leads us to raise the following two issues:
(1) What is the contribution of the Asian monsoon system to the global hydrological cycle/water vapor
budget?
(2) How does the global circulation affect the hydrological cycle/water vapor budget of the Asian monsoon
region?
Issue (1) deals with the upscale interaction, while Issue (2) tackles the downscale processes.
The Asian monsoon system is characterized by the Asian continental thermal low encircled by
monsoon southerlies in South Asia and southwesterlies in Southeast and East Asia.
Within this
continental-scale monsoon circulation, the Asian monsoon system consists of several distinguishable
1
regional monsoon components: South Asia, the South China Sea and western tropical Pacific, East Asia,
and Northeast Asia. Due to differences in their response to diabatic heating, the monsoon circulations in
these regions exhibit different characteristics, including timing of onset, life cycle, and interaction with
synoptic disturbances. The South-Asian monsoon circulation includes the Indian monsoon trough, which is
modulated by the northward migration of intraseasonal monsoon trough and ridge (Sikka and Gadgil 1980,
Krishnamurti and Subrahmanyam 1982), and westward propagation of monsoon depressions (Saha et al.
1981, Chen et al. 1999).
The Southeast-Asian monsoon circulation is characterized by a
southeast-extended monsoon trough that interacts with the northward-migrating monsoon trough/ridge
(Chen and Chen 1995), the westward-propagating easterly waves (Lau and Lau 1990), and tropical
cyclones (Chen et al. 2004).
The East-Asian monsoon circulation, which is characterized by the
Meiyu-Baiu rainbelt, receives contributions of rainfall from extratropical fronts before the monsoon break
and tropical cyclones after the break (Ramage 1952, Chen et al. 2004a).
As for the contribution of the Asian monsoon to the global hydrological cycle, it may be difficult, if
not impossible, to combine water vapor budget analyses of all the individual monsoon components. Based
on the rainfall and its variance (of different timescales), the Asian monsoon covers the areas stretched from
60oE to about the International Dateline (e.g. Chen et al. 1995a). For convenience, let us designate the
hemisphere of 60oE-120oW as the Asian-monsoon hemisphere and the one of 120oW-60oE as the
extra-Asian hemisphere (which covers the North-American and West-African monsoons). Therefore, the
contrast of the water-vapor budgets of these two hemispheres will be used to illustrate the contribution of
the Asian monsoon system to the global hydrological cycle.
Numerous attempts have been made to explore the impact of the global-scale propagating modes on
the regional circulations of the Asian monsoon. However, few of these efforts gave attention to the effect of
these propagating modes on the hydrological cycle over a particular region within the Asian monsoon
system. As pointed out previously, the regional monsoon circulations over different areas of the Asian
monsoon system are characterized by different features. Of course, it may not be possible for the regional
2
hydrological cycles to respond to the global-scale propagating modes in a uniform manner. The East-Asian
monsoon will be used to illustrate the downscale processes of how the global-scale propagating mode
regulates the regional hydrological cycle.
2. Theoretical Background of the Water-Vapor Budget
The atmospheric branch of the hydrological cycle is usually depicted in terms of the water vapor
budget equation (e.g. Peixoto and Oort 1992):
W
(1)
   Q  E  P,
t
1 pS
1 pS
where W 
qdp , Q 
Vqdp , E, P, g, q, p, ps, and V are precipitable water, water vapor flux,
0
g
g 0


gravity, specific humidity, pressure, surface pressure, and the velocity vector, respectively. Water vapor,
which resides primarily in the lower troposphere, is largely transported by the low-level atmospheric
circulation. Therefore, low-level atmospheric flow is reflected through water vapor flux. Without a phase
change, water vapor behaves like a passive atmospheric constituent. According to Eq. (1), water vapor
converges (   Q  0 ) toward its sink region (E-P<0). Consequently, the residence of water vapor in the
atmosphere is maintained by the imbalance between   Q and (E-P), and represented by the water vapor
storage,
W
.
t
According to the Helomhaltz theorem, water vapor flux (Q) may be split into rotational (Q R) and
divergent (QD) components:
Q = QR + QD .
(2)
These two components of water vapor flux, (QR, QD), may be expressed in terms of the streamfunction ( Q ;
Rosen et al. 1979) and potential functions (  Q ; Chen 1985) of water vapor flux;
QR  kˆ   Q , and QD   Q .
(3)
These two water vapor flux functions can be obtained by the following two inverse transforms,
 Q   2 (kˆ    Q), and  Q   2 (  Q).
3
(4)
The water vapor budget equation can therefore be expressed as either
or
W
   QD  E  P ,
t
(1’)
W
  2  Q  E  P.
t
(1”)
Since water vapor is largely transported by low-level atmospheric flows, spatial structure of the global  Q
and  Q distributions resemble those of low-level streamfunction ( ) and velocity potential (  ). After
depictions of the global  (200mb) and  (200mb) fields were introduced by Krishnamurti (1971) and
Krishnamurti et al. (1973), the global atmospheric circulation is often portrayed by these two variables.
Thus, the water vapor transport by the global atmospheric circulation can be well delineated by  Q and
 Q . As indicated by the water vapor budget, (E-P) is only the local source/sink of water vapor. It may be
inferred from Eq. (1”) that the link between this source and sink of water vapor is established by the global
divergent circulation depicted in terms of the  Q field.
Despite their local nature, hydrological processes included in the water vapor budget can be coupled
with the global divergent circulation through the  Q field and the regional water vapor flux linked to the
global atmospheric circulation through both the  Q and  Q fields.
Using this link between local
hydrological processes and the global atmospheric circulation, the contribution of the Asian monsoon to the
global hydrological cycle will be illustrated in terms of the hemispheric water vapor budget, and the
possible impact/effect of the global atmospheric circulation on the Asian monsoon rainfall will be
demonstrated in terms of the  Q and  Q fields.
3. Contribution of the Asian Monsoon to the Global Hydrological Cycle
The Asian monsoon rainfall undergoes variations on three different time scales: diurnal, intraseasonal,
and interannual. Can the contribution of the Asian monsoon to the global hydrological cycle be reflected in
these variations? We shall illustrate them from two different perspectives: climatological and temporal
variations.
3.1 Climatology
During the northern summer, there are three major monsoon systems in the Northern Hemisphere: the
4
Asian, North American, and West African monsoons. The summer rainfall in these monsoon regions and
along the Intertropical Convergence Zone (ITCZ) is maintained by the convergence of water vapor flux
carried out by the global divergent circulation which can be inferred from the QD distribution in Fig. 1a
(Chen 1985). As revealed from summer precipitation in this figure, the amount and spatial coverage over
the Asian monsoon region and the western tropical Pacific is much larger than the other two monsoon
regions. Following Chen et al. (1995a), let us divide the globe into two hemispheres: the Asian-monsoon
hemisphere (60oE-120oW) and the extra-Asian monsoon hemisphere (120oW-60oE). For convenience, any
variable ( ) averaged over the former hemisphere is designated by ( )A, over the latter hemisphere by ( )E,
and over the globe by ( )G  12 [
 A   E ] .
Based on the water vapor budget Eq. (1), the water vapor
budgets of the two hemispheres may be written as:
WA
 (  Q) A  E A  PA ,
t
WE
and
 (  Q) E  E E  PE .
t
(5a)
(5b)
Because (  Q) G  0,
(  Q) A  (  Q) E .
(6)
The water vapor budgets of the two hemispheres are displayed in Fig. 1b. The contrasts in the water vapor
budgets between these two hemispheres may be summarized as follows:
(1) Values of PA and EA are larger than those of PE and EE, respectively. It is suggested by the contrast
between hemisphere-averaged E and P values that the intensity of the hydrological cycle in the
Asian-monsoon hemisphere is larger than that in the extra-Asian monsoon hemisphere.
(2) Rainfall in the Asian monsoon hemisphere (PA=3.2 mmday-1) is larger than evaporation (EA=2.85
mmday-1), while rainfall in the extra-Asian monsoon hemisphere (PE=2.22 mmday-1) is smaller than
evaporation (EE=2.57 mmday-1). Evidently, EA-PA<0 and EE-PE>0. In other words, the Asianmonsoon hemisphere is a water-vapor sink and the extra-Asian monsoon hemisphere is a water-vapor
source. The deficit of water vapor in the former hemisphere is supplied by the latter hemisphere, as
5
indicated by (  Q) E  0 [or (  Q) A  0] .
In view of the contrasting characteristics in the water-vapor budget between the two hemispheres
shown in Fig. 1b, several questions are raised: (1) Can the contrast of the hemispheric water vapor budget
between the two hemispheres always be maintained? (2) How is the water vapor surplus supplied in the
extra-Asian monsoon hemisphere? (3) Where does the excessive rainfall in the Asian monsoon hemisphere
go? Question (1) will be discussed later in Section 3.2, but Questions (2) and (3) are answered here. The
salinity is higher in the Atlantic ocean than the Pacific ocean (Levitus 1986; Chen et al. 1994) because
evaporation is larger over the former ocean than the latter ocean (Fig. 1b). The excessive water vapor
generated by evaporation in the Atlantic ocean converges toward the Pacific ocean (Bryan and Oort 1984;
Chen et al. 1994) to make up for its deficit of water vapor. If this water vapor supply is ceaseless, the
contrast of the water vapor budget between the two hemispheres in Fig. 1b may become sharper. Despite
interannual variations in the global and hemispheric hydrological cycle, it will be shown in Section 3.2 that
the water vapor budget contrast between the two hemispheres is always maintained. It is implied by this
water vapor budget contrast that there must be a path for water vapor to exit the Asian-monsoon hemisphere.
Compiling the interocean freshwater transport shown in Fig. 1c, Wijffels et al. (1992) showed that
freshwater is routed from the Pacific Ocean through the Bering Sea and Arctic Ocean to the Atlantic Ocean.
Based on the contrast between the QD distribution in Fig. 1a and the interocean freshwater transport in Fig.
1c, the path of water vapor exiting the Asian monsoon hemisphere is through interocean freshwater
transport shown in Fig. 1c (Chen et al. 1994). The basic role played and the contribution made by the Asian
monsoon hemisphere to the global hydrological cycle is revealed from the contrast of the water-vapor
budget between the Asian-monsoon and extra-Asian monsoon hemisphere in Fig. 1.
3.2 Temporal Variation
As previously shown in Fig. 1b, the Asian-monsoon hemisphere produces about 50% more rainfall
than the extra-Asian monsoon hemisphere. Since the global atmospheric circulation exhibits variation with
different time scales, ranging from diurnal to interannual, does the global hydrological cycle undergo
6
variations within these time scales? Can these variations of the global hydrological cycle follow those of
the Asian monsoon hemisphere? This question may be answered in terms of three different statistical
analyses. First, time series of globally and hemispherically long-term precipitation averaged every three
hours for a half month of June, daily-mean precipitation for the entire year of 1991, and summer-mean
precipitation (JJA) for the 1979-2002 period are shown in Fig. 2. The time series of hemispheric mean
divergence of water vapor flux corresponding to these precipitation time series are also shown in Fig. 3. It
is clearly revealed that variations of PG in all time scales concerned in this study follow those of PA. Second,
power spectra of the first two types of time series are displayed in Fig. 3. Signals of diurnal cycle
~
  and

intraseasonal [30~60 day ( ) and 12~24 day ( ) ] modes clearly stand out. For time series of summer
rainfall, interannual variations of summer precipitation closely follow summer SST (NINO3.4) departures.
It is of interest to note that amplitudes of rainfall variations shown in Fig. 2 are of the same order of
magnitude, regardless of their time scales. Third, in order to understand where significant signals of rainfall
variations identified above are located and to verify whether the separation of these signals into two
hemispheres is appropriate, rainfall variance distributions of these signals are shown in Fig. 4. Despite the
detailed spatial structure in their variance distributions, it is clearly indicated that these rainfall variation
modes exhibit larger values over the Asian monsoon hemisphere as inferred from their time series shown in
Fig. 2. The Asian monsoon hemisphere not only contributes more rainfall to the global hydrological cycle,
but also affects temporal variations of the global hydrological cycle more significantly than the other
hemisphere. In addition, to substantiate this inference with the hemispheric water vapor budget, we shall
also explore the hydrological implication of the possible contrast of roles played by the Asian- and
extra-Asian monsoon hemispheres in the global hydrological cycle.
a. Diurnal Variation
Several interesting features emerge from the comparison of 3-hour time series of PG, PA, and PE in
Fig.3a:
(1) PA is always larger than PE.
7
(2) Standard deviations of PG, PA, and PE diurnal variations are 0.2, 0.5, and 0.3 mm·day-1, respectively.
Amplitude of PA diurnal variation is generally larger than that of PE diurnal variation. This contrast is
consistent with variance distributions of diurnal variations of rainfall shown in Fig. 4a.
(3) Diurnal variations of PG and PA are in phase (with a correlation coefficient rPA PG  0.95 ; where rx y is
the correlation coefficient between time series of variables x and y), but are out of phase with P E (with
correlation coefficients rPA PE  0.93 and rPG PE  0.94) .
It is inferred from these statistical
relationships that the diurnal variation of global rainfall follows closely that of the Asian monsoon
hemisphere.
Both the Asian monsoon and extra-Asian monsoon hemispheres contribute to global precipitation.
Why should the diurnal variation of global precipitation follow the former hemisphere? What roles are
played by these two hemispheres in the global hydrological cycle? The hemispheric water vapor budget
analyses will be discussed further to answer these questions. Observations of evaporation are not available
and estimations of evaporation with the residual method of the water vapor budget are inaccurate.
Therefore, it is difficult, if not impossible, to perform a complete water vapor budget analysis. Because
rainfall is primarily maintained by the water vapor supply through convergence of water vapor flux,
inference of roles played by the Asian monsoon and the extra-Asian monsoon hemispheres may be made by
the following hydrological relationship:
PA~  (  Q) A  (  Q) E ~  PE
The 3-hour time series of (  Q) E are displayed at the bottom of Fig. 2a. The PA (PE) variation is in phase
(out of phase) with (  Q) E . As shown in Fig. 1b, the summer rainfall in the Asian monsoon hemisphere
is supplied by the water vapor flux diverged out of the extra-Asian monsoon hemisphere, in addition to the
in-situ evaporation.
On the other hand, the 3-hour time series of PA, PE, and (  Q) E in Fig. 2a reveal
that water vapor from the Asian-monsoon hemisphere is supplied (depleted) by (from) the extra-Asian
monsoon hemisphere through divergence (convergence) of water vapor flux to maintain the diurnal
variation of the Asian monsoon hemisphere rainfall.
8
Roles played by the two hemispheres in the diurnal variation of the global hydrological cycle may be
summarized by the composite hemispheric water vapor budget during PG maximum and minimum phases
(Fig. 5). Hemispheric evaporation is estimated by the residual, although the estimation may contain bias
from the data assimilation system and computational errors. It is of interest to note that the summer-mean
hemispheric water vapor budgets in both hemispheres are modulated by the diurnal variation, but the
contrast of the water vapor source and sink between the two hemispheres revealed from the climatological
water vapor budget shown in Fig. 1b is not changed by the diurnal variation.
At the surface, the short-wave radiative heating of the diurnal hemisphere and the long-wave radiative
cooling of the nocturnal hemisphere roughly form a wave-1 structure around the globe. The direct
atmospheric response to this east-west differential heating/cooling is reflected by the global divergent
circulation which may exhibit a low level structure similar to  Q shown in Fig. 1a. As inferred from the
spatial structure of  Q , the global divergent circulation in response to the surface radiative heating/cooling
also exhibits a wave-1 structure. Because the Earth rotates eastward, the response of the atmospheric
divergent circulation propagates westward.
First, the division of the global circulation into the
Asian-monsoon and extra-Asian monsoon hemisphere match well with the wave-1 structure of the global
divergent circulation. Second, the Asian-monsoon region may be more convectively unstable than the other
hemisphere.
The atmospheric response to the surface radiative heating is more dramatic over the
Asian-monsoon hemisphere. This hemispheric difference of rainfall in response to this diurnal variation of
heating will be illustrated in Section 4.
b. Intraseasonal Variation
Numerous research efforts have been devoted to exploring the impact and modulation of intraseasonal
modes on the onset and life cycle of monsoons in different regions within the Asian monsoon system: the
Indian monsoon by Murakami (1976), Krishnamurti and Bhalme (1976), Krishnamurti and Subrahmanyam
(1982), the South China Sea monsoon by Chen and Chen (1995), and the East-Asian monsoon by Chen and
Murakami (1988). The effect of intraseasonal modes on the hydrological cycle of the entire Asian monsoon
9
system has not been a research focus. The lack of previous research attention on this effect may be
attributed to the lack of uniform rainfall observations. Therefore, previous research efforts along this
direction were largely concentrated on regions of particular concern, e.g. East/Southeast Asia by Lau et al.
(1988), Chen et al. (1988a), Chen et al. (2004b), India by Chen et al. (1988b) and Chen and Chen (1993). It
may be difficult to obtain an overview of the possible effect of the global intraseasonal modes on the Asian
~
monsoon system by merging its impacts on several regional monsoons. As revealed from Var( P ) and
Var( P̂ ) distributions in Fig. 4, large variances of these two variables cover the entire Asian monsoon
hemisphere. In order to obtain a coherent impact of intraseasonal modes over the entire Asian monsoon and
the contribution of the Asian monsoon system to the possible intraseasonal variation of the global
hydrological cycle, we need to examine intraseasonal variation of the global and hemispheric water vapor
budget.
Time series of daily PG, PA, PE, (  Q) A , and (  Q) E are displayed in Fig. 3b. Certain interesting
features of the global and hemispheric hydrological cycles are reflected by these time series:
(1) Variation of PE is out of phase with PA, but variation of PG is coincident with PA.
Since
PG  12 ( PA  PE ) and time series of PG and PA are in phase, it is obvious that amplitude of PA is larger
than PE. The intraseasonal variation of global precipitation follows the Asian-monsoon hemispheric
rainfall. It was observed that monsoon breaks in South Asia (e.g. Krishnamurti and Subrahmanyam
1982), the South China Sea (e.g. Chen and Chen 1995), Indochina (e.g. Kihuchi and Matsumoto 2005),
and East Asia (e.g. Chen et al. 2004a) occur in late June and early July when monsoon onset takes place
in Southwest North America (e.g. Douglas et al. 1993) and West Africa (e.g. Sultan et al. 2003, Sultan
and Janicot 2003). The contrast of these monsoon rainfall variations is well revealed from the
out-of-phase variations between PA and PE during the northern summer.
(2) Regardless of intraseasonal variations in both (  Q) A , and (  Q) E , water vapor always diverges
out of the extra-Asian monsoon hemisphere [i.e. (  Q) E >0] and converges toward the Asian
monsoon hemisphere [ (  Q) A <0]. Evidently, the former hemisphere is a water vapor source, while
10
the latter hemisphere is a water vapor sink. Since precipitation is maintained by the convergence of
water vapor, it is expected that intraseasonal variations of hemispheric precipitation and convergence of
water vapor flux should be related in a coherent way. As shown in Fig. 2b, intraseasonal variation of
(  Q) E [ (  Q) A ] is in phase (out of phase) with PE. Thus, water vapor is exchanged between these
two hemispheres in such a way that an increase (decrease) of PA follows that of (  Q) E which results
in decrease (increase) of PE.
It was shown in Fig. 3 that intraseasonal variations of global and hemispheric precipitation exhibit
two pronounced modes: 30-60 day and 12-24 day. How do these two low-frequency modes contribute to
the global and hemispheric hydrological cycles? Time series of hydrological variables corresponding to Fig.
2b for these two intraseasonal modes are shown in Fig. 6. First, let us explore how significant contributions
of these two low-frequency modes are to intraseasonal variations of all hydrological variables shown in this
figure. This contribution is measured by the ratio R( ) between variance of any variable ( ) and variance of
~
^
its 30-60 day and 12-24 day modes combined ( ) + ( ) . R(PG)=70%, R(PA)=75%, R(PE)=72%, and
R[ (  Q) E ]=84%.
Apparently, intraseasonal variations of these hydrological variables are well
contributed by the two intraseasonal modes.
The life cycles of both the 30-60 day and 12-24 day modes are different. Therefore, both modes
may occasionally be in phase. We shall illustrate separately their roles in the global hydrological cycle.
~
~
~
Variation of PG ( PˆG ) is in phase with PA ( PˆA ) , but out of phase with PE ( PˆE ) .
Variation of
~
~ ˆ
~ ˆ
ˆ ] is in phase with P
(  Q) E [ (  Q)
A ( PA ) , while out of phase with PE ( PE ) . The phase relationships
E
of these hydrological variables indicate that the roles played by these two intraseasonal modes on the global
hydrological cycle behave in the same manner. During the minimum phase, the 30-60 day (12-24 day)
mode enables the extra-Asian monsoon hemisphere to diverge water vapor to the Asian monsoon
~
~
~
hemisphere in supporting maximum PA ( PˆA ) . During the PE and P̂E maximum phase, and PA and P̂A
minimum phases, the two intraseasonal modes reverse their function in the divergence of water vapor flux.
It was shown by Chen et al. (1995a, b) that the potential function of water vapor flux of these two
11
intraseasonal modes, ~Q and ̂ Q , are mainly depicted by a wave-1 pattern. The eastward propagation of
~
these two intraseasonal modes result in the seesaw exchange of water vapor flux [reflected by (  Q) E
~ ˆ
~ ˆ
~
ˆ ] and the corresponding seesaws of both P
and (  Q)
A ( PA ) and PE ( PE ) . The impacts of global  Q
E
and ̂ Q modes on the regional Asian-monsoon rainfall will be further illustrated in Section 4. In addition,
~
~
the coherent variation between PG ( PˆG ) and PA ( PˆA ) indicates that the hydrological responses of the Asian
monsoon hemisphere to the eastward-propagating intraseasonal modes are more sensitive than the
extra-Asian monsoon hemisphere.
The roles played by the two intraseasonal modes in the global hydrological cycle are concisely
summarized in the composite hemispheric water vapor budgets during their maximum and minimum
phases in Fig. 7. The major features of these water vapor budgets can be highlighted as follows:
(1) Regardless of the seesaw precipitation variation and exchange of water vapor flux between the Asian
monsoon and the extra-Asian monsoon hemisphere, a clear distinction between water vapor source and
sink is always maintained.
(2) Although evaporation is estimated, it is indicated that P[ P(max)  P(min)] is always larger than
E[ E (max)  E (min)] for both intraseasonal modes in the two hemispheres.
It suggests that
precipitation/convection is more sensitive to the eastward propagation of the intraseasonal modes,
despite the possible effect of intraseasonal modes on the latent-heat energy flux from oceans by
previous studies (e.g. Krishnamurti et al. 1988).
(3) Recall that the Asian monsoon hemisphere is a water vapor sink. It was pointed out in Section 3.1 that
the freshwater flow from the Indian and the Pacific Oceans (generated by monsoon rainfall) across the
Bering Strait and the Arctic Sea returns back to the Atlantic. Can this returning freshwater to the
Atlantic undergo the intraseasonal variations? Although it may be difficult to substantiate this
argument with existing observational data, simulations of a coupled global climate model may offer an
alternative.
c. Interannual Variation
12
The interannual variation of Indian monsoon rainfall has a long history as a research subject. The
two most compelling mechanisms causing this interannual variation are the Southern Oscillation (an
east-west surface pressure seesaw between the Indian-western Pacific and the central-eastern Pacific)
(Walker 1923, 1924) and the interannual variation of snowfall over Central Asia and the Himalayan
mountains (Blandford 1884). These two mechanisms have been extensively explored and analyzed in the
past several decades with the data generated by modern data assimilation systems. The interannual
variation of the East/Southeast Asian monsoon rainfall has also stimulated a surge of research activity in
recent decades (e.g. Wu and Wang 2000, 2002). The main focus of these research endeavors includes the
possible impact of the ENSO cycle on the East/Southeast Asian monsoon. Over Indochina, only a few
studies (e.g. Chen and Yoon 2000) explored the interannual variation of monsoon rainfall. The variance
distribution of interannual rainfall departures from their long-term summer-mean rainfall in Fig. 2d
indicates that the major global interannual variation of summer rainfall is primarily contributed by the
Asian monsoon hemisphere. A more concise perspective of the ENSO effect on the Asian monsoon system
and the contribution of the Asian monsoon system to the interannual variation of the global hydrological
cycle may be gained through the hemispheric water budget analysis.
It was suggested that the global hydrological cycle is intensified (weakened) during the ENSO warm
(cold) event (Chahine 1992, Soden 2000).
Analyzing the interannual variation of the global and
hemispheric hydrological cycle, Chen et al. (2005) recently showed that the impact of the ENSO cycle on
the global hydrological cycle is opposite to Chahine's suggestion. Does the summer global and hemispheric
water vapor budget undergo interannual variation in concert with the ENSO cycle? This concern is
clarified by time series of various hydrological variables in Fig. 2c. In order to compare these variables
with the ENSO cycle, the summer SST (NINO3.4) is also included in Fig. 2c. The summer warm and
cold events are classified based on the criterion SST (NINO3.4)  0.5oC warm and SST (NINO3.4) 
-0.5oC cold. Since SST anomalies over the western tropical Pacific and the central/eastern tropical
Pacific exhibit an interannual seesaw following the ENSO cycle, the convective activity/rainfall in the
13
western tropical Pacific has an interannual variation opposite to that in the central/eastern tropical Pacific.
Therefore, it is expected that rainfall over western tropical Pacific and the Asian monsoon region becomes
larger (smaller) during the cold (warm) summer over the eastern tropical Pacific.
The inference of the interannual variation in the hemispheric water vapor budget is confirmed by time
series of various hydrological variables shown in Fig. 2c:
(1) Interannual variation of PA is out-of-phase with summer SST (NINO3.4) anomalies, while that of PE
is in phase.
(2) Interannual variation of PG follows that of PA, because the amplitude of PA is larger than PE.
(3) Water vapor is always diverged out of the extra-Asian monsoon hemisphere [i.e. (  Q) E >0] and
converges toward the Asian monsoon hemisphere, [i.e. (  Q) A <0]. This hemispheric water vapor
exchange is always larger (smaller) during cold (warm) events.
It will be shown in Section 4 that the interannual mode of  Q exhibits a dominant wave-1 structure and
eastward propagation (as shown by Barnett 1983, 1984a, b, 1985, 1988). The convective activity/rainfall
over the Asian monsoon hemisphere is more sensitive in response to the eastward propagation of the ENSO
interannual mode.
Divergence of inter-hemispheric water vapor flux is therefore regulated by this
interannual mode in concert with the PA interannual variation. Consequently, the global rainfall variation
follows the PA variation.
The relationship between the summer SST (NINO3.4) and the global
hemispheric water vapor essentially contradicts Chahine’s (1992) suggestion.
The composite hemispheric water vapor budgets for those warm and cold summers marked on the
SST (NINO3.4) time series in Fig. 2c are displayed in Fig. 8 to summarize the interannual variation in the
global/hemispheric hydrological cycle. It was suggested that the global hydrological cycle is regulated by
the central-eastern tropical Pacific SST anomalies in such a way that it is accelerated (decelerated) during
the ENSO warm (cold) events. As revealed from the composite hemispheric water budgets, it works
opposite to this suggestion. Recall that the SST anomalies over western tropical Pacific exhibit a seesaw
with the SST (NINO3.4) anomalies. The convective activity/rainfall in the Asian monsoon hemisphere is
14
enhanced (suppressed) during the cold (warm) SST (NINO3.4) anomalies. Regardless of interannual
variations in hemispheric hydrological variables, the hydrological roles played by the extra-Asian monsoon
hemisphere as a water vapor source and the Asian monsoon hemisphere as a water vapor sink, are not
affected by the ENSO activity. Finally, it may be difficult to measure short time-scale variations in the
ocean freshwater flux. However, it may be possible to explore whether the global freshwater flux route
suggested by Wijffels et al. (1992) undergoes an interannual variation in response to that of the water vapor
flux exchange between the Asian monsoon and the extra-Asian monsoon hemisphere.
4. Modulation of the Water Vapor Transport and Rainfall in the East-Asian Monsoon Region by
Global Modes
The contribution of the Asian monsoon system to the global hydrological cycle was illustrated in
terms of the hemispheric water vapor budget of the Asian monsoon. This hemisphere contributes 50%
more than the extra-Asian monsoon hemisphere to the global rainfall. Since the rainfall of the Asian
monsoon hemisphere undergoes variations of different time scales, the global rainfall is also subjected to
the same variations. It was observed that the extra-Asian monsoon hemisphere is a water vapor source.
Therefore, the rainfall and interhemispheric divergent water vapor flux in this hemisphere vary
out-of-phase with the extra-Asian monsoon hemisphere. These east-west seesaw variations of hemispheric
hydrological variables are caused by the propagation of global-scale oscillations. In this section, the
effect/modulation of these global modes on the rainfall and water vapor transport of the regional monsoon
circulation will be illustrated by means of downscale processes.
4.1 Diurnal Variation
Because of the lack of sufficient station observations, particularly over oceans, satellite observations
around the tropical belt were often used as proxies for tropical rainfall, for example, the tropical rainfall
estimations by a threshold value of the NOAA outgoing long wave radiation (Janwiok et al. 1994), the
Tropical Rainfall Measuring Mission (TRMM) Satellite measurements (Nesbitt and Zipser 2003), and the
combination of the oceanic rainfall observed by the Special Sensor Microwave Imager (SSM/I) and the
15
TRMM measurements (Imaoka and Spencer 2000). On the other hand, Dai (2001) analyzed the diurnal
variation of global precipitation using the merged station observations over the globe with the
Comprehensive Ocean-Atmosphere Data Set (CODAS Woodruff et al. 1987). Common features of the
diurnal variation in tropical rainfall observed by these studies are: (1) Rainfall/convection tends to reach its
maximum over oceans in the early morning and over land in the evening, and (2) amplitude of the diurnal
variation of tropical rainfall/convection is much larger over land than over oceans.
The direct atmospheric response to any change in diabatic heating is the divergent circulation (Chen
and Yen 1991 a, b). Because the Earth rotates eastward, the immediate atmospheric response to the diurnal
variation of the solar heating is reflected by the westward propagation of the diurnal component of the
global divergent circulation. Most studies of diurnal variation in the global rainfall, including those referred
to above, focused primarily on precipitation in the tropical belt. It is difficult for us to infer from these
studies how the diurnal variation of tropical rainfall is coupled with that of the global atmospheric
circulation. This concern may be addressed by the following approach. Recall that   Q   2  Q .
Convergence (divergence) of water vapor flux is coupled with a positive (negative) center of the  Q field.
As shown by Eq. (1), precipitation is maintained (suppressed) by convergence (divergence) of water vapor
flux. Although these two variables are local hydrological processes, they can be linked to the global
divergent circulation through the potential function of water vapor flux. Based on this argument, it is
expected that diurnal components of both variables, (  Q , P  ) propagate westward. This inference is
substantiated by composite charts of these two variables every three hours in Fig. 10. In order to distinguish
the day and night, noon and midnight are marked by open thin-solid and thick-solid lines, respectively. The
westward propagation of solar heating is clearly indicated by these time lines.
The land-ocean contrasts of P  amplitudes (larger over lands than oceans) are clearly evident in Fig.
9. The phase of maximum P  (over land in the afternoon/evening; over oceans in early morning) is well
presented. The best example shown in Fig. 9 to support this observation would be rainfall at 12Z-15Z over
the central-western tropical Pacific and 12Z-18Z over the Asian-Maritime continent. In addition to these
16
traditional observations of tropical rainfall diurnal variation, it is also of interest to note that the spatial
coverage of P  anomalies over the Asian monsoon hemisphere is much larger than that over the
extra-Asian monsoon hemisphere.
The westward propagation of (  Q , P  ) anomalies is clearly discernable in Fig. 10. As expected,
positive (negative) P  anomalies are always coincident with a positive (negative)  Q center. It was
shown in Fig. 3a that a diurnal seesaw oscillation exists between PA and PE : PA ( PE ) gains its maximum
(minimum) at 12Z and minimum (maximum) at 00Z. This east-west diurnal seesaw oscillation of rainfall
between the two hemispheres coincides with the centers of  Q anomalies. Evidently, this diurnal seesaw
oscillation of rainfall is a result of the hemispheric hydrological cycle modulated by the westward
propagation of the global divergent circulation reflected by the global  Q anomalies. Recall that the
spatial coverage of P  over the Asian monsoon hemisphere is larger than that of the extra-Asian monsoon
hemisphere. It is not surprising to see that diurnal variation of PG follows PA . Because positive (negative)
P  anomalies coincide with positive (negative)  Q centers, diurnal variation of (  Q) E is in phase
with PA , but out of phase with PE .
The atmospheric diurnal tidal wave may be divided into migrating and non-migrating components
(Chapman and Lindzen 1970).
The migrating component displays a wave-1 structure, while the
non-migrating component exhibits the regional-scale structure. Do  Q anomalies (shown in Fig. 9)
behave like diurnal tidal waves? In addition to its westward propagation, a clear clockwise rotation of the
(  Q , P  ) anomalies appears over Asia and the Americas. To obtain a better view of the clockwise (  Q , P  )
anomalies over the Asian continent and the western Pacific, these two variables in the Asian monsoon
hemisphere are extracted out every three hours from Fig. 9 and are shown in Fig. 10. What is the cause of
such an interesting rotation of (  Q , P  ) anomalies around a major continent and its surrounding oceans?
The tropical troposphere is generally conditionally unstable (Riehl 1979), while the troposphere of middle
to high latitudes is stable. Convection can be triggered by diabatic heating faster in the warm tropics than in
the cool middle to high latitudes. On the other hand, soil and rock have smaller heat capacities than air.
17
Because the Himalayas and the Rockies are located west of the Pacific and the Atlantic, respectively, the
east-west land-ocean contrast results in an east-west differential response to the solar heating which is
enhanced by these major mountain ranges. The response of the tropical troposphere, which has a lower
land-ocean contrast and fewer major mountain ranges, is dominated by the westward-propagating wave-1
migrating component of diurnal variation.
In contrast, the tropospheric response in the
subtropics-midlatitudes is dominated by the non-migrating component, whose scale is primarily
determined by the effect of the land-ocean contrast and whose response is faster over land than ocean. The
clockwise rotations of (  Q , P  ) anomalies over Asia and North America, and their surrounding oceans are
essentially formed by the north-south difference in the tropospheric response to the westward moving solar
heating.
Analyzing the diurnal mode of the Asian summer monsoon, Krishnamurti and Kishtawal (2000)
found that the diurnal component of divergent circulation in the afternoon hours has an ascending center
over north-central India with its divergent air descending around the surrounding regions. The reverse
occurs in the morning hours. The  Q anomalies at 12Z and 00Z in Fig. 10 match their finding over South
Asia. Actually, the diurnal  Q mode [which is a reflection of   (850mb) or    (200mb)] exhibits not
only the diurnal pulsation of the continental-scale divergent circulation, but a clockwise rotation over the
Asian continent and its surrounding oceans. This rotation of (  Q , P  ) anomalies is further confirmed by
the phase (timing) of the diurnal variation of rainfall observed/compiled by the GEWEX (Global Energy
and Water Cycle Experiment) Asian Monsoon Experiment (GAME). Two approaches were adopted to
depict the phase of the diurnal rainfall variation: universal time (UTC) and local solar time (LST). The
Cressman (1959) scheme was applied to project station observations of GAME rainfall on
2.5 o  2.5 o boxes. A clear clockwise rotation in the phase of the diurnal rainfall variation emerges from
both approaches.
The diurnal variation of any meteorological variable is one of the most regular and significant climate
signals (next to annual variation). Diurnal variations of the Asian monsoon rainfall and hydrological
18
processes in maintaining this rainfall are results of complex multiple-downscale hydrological processes of
the atmosphere in response to the westward migration of solar heating. This response includes not only the
westward propagation of a global-scale migrating component, but also the clockwise rotation of a
continental-scale non-migrating component.
4.2 Intraseasonal Variation
The intraseasonal (30-60 day and 12-24 day) seesaws of P A and P E were caused by eastward
propagation of these two global intraseasonal modes (Chen et al. 1995a, b) which will be illustrated later in
this section. On the other hand, one may question how these two global intraseasonal modes affect the local
monsoon circulation and local weather systems (which are embedded in the monsoon circulation) which are
effective producers of monsoon rain. The monsoon rainfall and the activity of the rainfall-generating
synoptic systems in South Asia are generated/modulated by the coupling of the global 30-60 day mode with
the regional 30-60 day monsoon trough/ridge (Krishnamurti and Subrahmanyam 1982, Chen et al. 1988b),
and the westward propagation of 12-24 day monsoon mode (Chen and Chen 1993). How do these two
global intraseasonal modes modulate the summer monsoon rainfall in East Asia? Water vapor transport in
this region, particularly along the Meiyu rainfall belt is fluctuated by the 30-60 day mode (Chen et al.
1988b). The regional monsoons in the South China Sea and East Asia are coupled by the 30-60 day and
12-24 day monsoon modes (Chen et al. 2000). This interaction is reflected by the intraseasonal variation of
East-Asian rainfall observed by Lau et al. (1988). It is inferred from these studies that variation of the
East-Asian monsoon rainfall is profoundly affected by global intraseasonal modes through certain
downscale interactions with the regional circulation and rainfall-producing synoptic systems.
Using station measurements of rainfall in South China, Hong-Kong, Taiwan, and Japan, Ramage
(1952) pointed out that major rainfall in East Asia was generated by cold fronts before the monsoon break
and by tropical cyclones after the break. Because rainfall observations over the ocean by satellites and
reanalysis data became available, Ramage’s suggestion was confirmed by an extensive study performed by
Chen et al. (2004a). The East-Asian summer monsoon is characterized by the northward migration of the
19
Meiyu rainbelt, the monsoon life cycle (including its onset in mid-May, break in late June-early July, and
revival in late July), and the active typhoon season after the monsoon break. These characteristics are
revealed from the y-t diagram of rainfall at 120oE-125oE (Fig. 12a). In addition to the monsoon life cycle,
both the Meiyu rainbelt and the ITCZ also undergo intraseasonal variation with a time scale smaller than the
monsoon life cycle. The rainfall time series of every summer season at (22.5oN-25oN, 120oE-125oE) were
subjected to a power spectral analysis. The 30-60 day and 12-24 day intraseasonal modes emerge from the
averaged power spectra (Fig. 12b). After removing the yearly-mean, annual, and semiannual components
from the rainfall y-t diagram (Fig. 12a), the y-t diagram of rainfall departure (  P ) is shown in Fig. 12c.
Evidently, the monsoon life cycle and intraseasonal variation of shorter time scales are largely determined
by 30-60 day and 12-24 day modes. It is revealed from Fig. 12 that intraseasonal variations of the East
Asian monsoon rainfall are modulated by the global intraseasonal modes through certain downscale
processes.
In order to illustrate this downscale process, let us first understand how the global intraseasonal
modes affect intraseasonal variations of the hemispheric water-vapor budgets. This local impact of the
global intraseasonal modes is depicted by composite charts of the two global intraseasonal modes. The
composite procedure is as follows:
~
(1) Selection of 30-60 day and 12-24 day cycles: Using ( PA ) and ( PˆA ) as indices, we select these half
cycles with their amplitudes larger than 0.8 or smaller than -0.8 of the corresponding standard
deviations over the summer (May-September) season.
(2) Definition of phases: A cycle of the 30-60 day mode is divided into eight phases (Fig. 13a), while a
~
cycle of the 12-24 day mode is split into six phases (Fig. 13b). Composite charts of ( ~Q , P ) and
( ˆ Q , Pˆ ) at each phase are prepared with five and three-day averages, respectively, centered on the date
of the selected phase.
~
Composite charts of ( ~Q , P ) and ( ˆ Q , Pˆ ) over their entire life cycles are shown in Figs. 14a, and b,
respectively. Both the ~Q and ̂ Q anomalies exhibit a wave-1 structure and relatively regular eastward
20
~
propagation. PA and P̂A attain their maxima at phase 2, when positive ~Q and ̂ Q centers are located
~
over the western tropical Pacific (the Asian monsoon hemisphere). During this phase, PE and P̂E obtain
their minima, when negative ~Q and ̂ Q centers are located over northeast Brazil (the extra-Asian
~
monsoon hemisphere). Spatial phase structures of both ~Q and ̂ Q are reversed when [ PA (phase 6), and
~
P̂A (phase 5)] and [ PE (phase 2), P̂E (phase 2)] reach their minima and maxima, respectively. According to
~
~
time series of hydrological variables shown in Fig. 2, PA maxima (minima) and PE minima (maxima) are
~
~
coupled with (  Q) A minima(maxima)/ (  Q) E maxima(minima).
The same contrast is also
applicable to the 12-24 day mode. Comparing variations in time series of PA , PE , (  Q) A and
~
(  Q) E with ( ~Q , P ) and ( ˆ Q , Pˆ ) composite charts, it becomes clear that intraseasonal variations of
PA and PE are caused by eastward propagation of these two global intraseasonal modes. In addition,
hemispheric divergence of water vapor flux is also perturbed by these two global intraseasonal modes in
such a way that water vapor diverges from the hemisphere of rainfall decrease and converges toward the
hemisphere of rainfall increase.
Because rainfall and divergence of water vapor flux are local hydrological processes, how do the two
global intraseasonal modes affect these local processes? As shown in Fig. 12a, the majority of East-Asian
summer rainfall is generated along the Meiyu rainbelt before the monsoon break and tropical cyclones after
the break. It is revealed from the contrast between Figs. 12c and d that rainfall variations in East Asia are
caused by the two intraseasonal modes. The impact of the global intraseaonsal modes on the East-Asian
rainfall should be achieved through the downscale interaction of these global modes with the regional
monsoon circulation/synoptic weather systems. This downscale interaction will be illustrated in terms of
the regulation of the two intraseasonal modes on the northward migration of the Meiyu rainbelt/the frontal
~
activity and on the tropical cyclone occurrence frequency as indicated by the y-t diagrams of P and P̂ .
Note that the atmospheric circulation is well depicted by the rotational component of atmospheric flow. For
this reason, we shall use the intraseasonal modes of the streamfunction of water vapor flux to illustrate the
downscale interaction.
21
a. 30-60 day mode
In South-Southeast Asia and the western tropical Pacific, the eastward-propagating global 30-60 day
couples with the monsoon trough/ridge through the local Hadley circulation (Chen et al. 1988b; Chen and
Chen 1995). The eastward propagation of this global 30-60 day mode migrates the monsoon ridge/trough
northward from the Equator, and regulates the monsoon life cycle over this region (Krishnamurti and
Subahmanyan 1982, Lorenc 1984, Krishnamurti et al. 1985, Chen et al. 1988b). The eastern end of the
30-60 day monsoon trough/ridge generally exhibits a cyclonic/anticyclonic cell over the South China Sea
and the western tropical Pacific. This local intraseasonal circulation, which is well depicted by ~Q
anomalies in Fig. 15, migrates northward from the vicinity of the Equator to 30o-35oN, as shown by the y-t
diagram of ~Q anomalies at 120oE for the 1996 summer. The frontal activity (shaded areas) and locations
of tropical cyclones (dots) during different phases of the 30-60 day monsoon mode are superimposed on
~Q anomalies (Fig. 15a). Tropical cyclone occurrence is enhanced over the intraseasonal ~Q cyclonic
cell, while frontal activity is intensified along the northern rim of the intraseasonal ~Q anticyclonic cell,
where water vapor flux along the Meiyu rainbelt is increased.
The contrast between the tropical
cyclone/frontal activity (Fig. 15a) and rainfall anomalies (Fig. 15b) over the entire life cycle of the local
30-60 day intraseasonal mode indicates that the East-Asian monsoon rainfall is primarily generated by
these two types of synoptic disturbances. In turn, the activity and occurrence locations of these synoptic
disturbances are regulated by the regional intraseasonal mode which is coupled with the
eastward-propagating 30-60 day mode. Evidently, the intraseasonal variation of PA in East Asia is formed
by such a downscale interaction.
b. 12-24 day mode
In Fig. 12, a 12-24 signal along the Meiyu rainbelt and the ITCZ stands out in the y-t diagram of
East-Asian summer rainfall at 120oE-125oE. What are the possible synoptic disturbances generating this
signal of rainfall variations? Cold surge are the most pronounced synoptic activity in the East-Asian winter
monsoon system. The cold front associated with a cold surge is coupled with the eastward propagating
22
upper-level short-wave trough in Northeast Asia (e.g. Lau and Lau 1990, Chen et al. 2002). Following the
cold-surge front, the cold-air intrusion brings precipitation across East Asia. In late spring, the upper-level
east-coast jet becomes weaker. At this time, the passage of cold fronts over East Asia are usually slow
because the eastward propagation speed of the midlatitude short-wave trough is reduced by the weakening
of midlatitude westerlies. Therefore, the slowly moving cold fronts form the Meiyu rainbelt. This
circulation evolution will not stop until the East-Asian monsoon break. It was shown by Compo et al. (1999)
that the cold front/surge activity in East Asia exhibits a 12-24 day signal of surface pressure. This signal of
rainfall along the Meiyu rainbelt is a reflection of the frequent passage of cold fronts. On the other hand, the
rainfall variation along the ITCZ is caused by tropical cyclone activity. Chen et al. (2004b) observed that
approximately 70% of tropical cyclogeneses over the western tropical Pacific are linked to monsoon gyres
(Lander 1994). The occurrence of monsoon gyres is evolved from the perturbation of the NW-SE oriented
monsoon trough (which extends from Indochina across the northern part of the South China Sea to the
Philippine Sea) by the intraseasonal monsoon trough/ridge (Chen et al. 2005). Since the monsoon gyre
occurrence shows a preferred time scale of 12-24 days, it is not surprising that tropical cyclogenesis in the
western tropical Pacific undergoes such a temporal variation.
The mechanisms of generating the 12-24 day rainfall signal in East Asia before and after the
monsoon break in late June and early July defined by Chen et al. (2004a) differ from each other. The
modulation of the regional 12-24 day modes on the rain-producing synoptic systems and the link of these
regional modes with the global mode are illustrated in Fig. 16 in terms of composite charts of ˆ Q
anomalies for all six phases determined by the P̂A index (Fig. 13b). The rain-producing systems are cold
fronts along the Meiyu rainbelt before the monsoon break and tropical cyclones after the break. Therefore,
the cold front ( N FT ) and tropical cyclone ( N TC ) occurrence frequencies are superimposed on ˆ Q
anomalies in Fig. 16c. As indicated by the y-t diagram of ˆ Q anomalies in Fig. 16e, the cellular ˆ Q
structure migrates southeastward before the monsoon break. This propagation characteristic is a reflection
of the southeastward movement of synoptic disturbances associated with cold fronts. The SW-NE oriented
23
strip of P̂ anomalies (Fig. 16b) are coincident with the frontal occurrence frequency N FT in Fig. 16a.
Both positive (negative) values of ˆ Q and maximum (minimum) N FT appear when the P̂A index reaches
its maximum (minimum). Evidently, the East-Asian rainfall variation modulated by the regional 12-24 day
mode coincides with the hemispheric water vapor budget and the eastward propagation of the global 12-24
day mode.
After the break, the Meiyu rainbelt no longer exists in East Asia, and tropical cyclones become the
major rainfall producer. The y-t diagram of ˆ Q anomalies (Fig. 16e) exhibits a northeastward propagation
opposite to that which occurred before the break by the southeasterlies along the southwest rim of the North
Pacific anticyclone. It is of interest to note that the tropical cyclone occurrence frequency (TC locations
denoted by dots) is suppressed by a positive (anticyclonic) ˆ A cell, but enhanced by a negative (cyclonic)
ˆ A cell. Therefore, we can find in Fig. 16d that positive (negative) P̂ anomalies appear over negative
(positive) cells of ˆ Q anomalies when PA ( PE ) reaches a maximum (minimum) during phase 2(5) of the
12-24 day mode. The link between this regional 12-24 day mode activity and the hemispheric/global
hydrological cycle is revealed through composite charts of N TC , ˆ Q and P̂ anomalies over the 12-24 day
cycle.
4.3 Interannual Variation
As pointed out previously, interannual variations of rainfall in different regions of the Asian monsoon
may be caused by different mechanisms. It may seem unlikely to bring out a coherent interannual variation
of rainfall over the entire Asian monsoon system by combining these variations in different monsoon
regions. On the other hand, within the context of the ENSO cycle, the contrast of time series of
hydrological hemispheric variables and SST (NINO3.4) in the two hemispheres provides a systematic
perspective of the hydrological cycle of the Asian monsoon hemisphere: its contribution to the global
hydrological cycle, and its relationship with the hydrological cycle of the extra-Asian monsoon hemisphere.
Actually, our main concern here is how the interannual variation of rainfall in the two hemispheres is related
to the global circulation. To address this concern, the tropical x-t diagram of (  Q , P) is shown in Fig.
24
17a. The extreme summers of the ENSO cycle are defined by the following criteria: warm when
SST (NINO3.4)  0.5oC and cold when SST (NINO3.4)  -0.5 oC. The warm and cold summers are
marked by thick (thin) solid lines on the x-t diagrams of (  Q , P) .
The ENSO cycle is often depicted in the tropics by an east-west seesaw oscillation of surface pressure
between the Indo-western tropical Pacific and the central-eastern tropical Pacific, or an east-west seesaw
variation of SSTs between the two sides of the tropical Pacific. It was observed by Barnett (1983, 1984a, b,
1985, 1988) that following the ENSO cycle, tropical surface winds and pressure exhibit a clear eastward
propagation from the Indian Ocean to the central Pacific.
The interannual variation of summer
SST (NINO3.4) may not always follow that of winter SST (NINO3.4). A clear eastward propagation
emerges from the tropical x-t diagram of summer (  Q , P) anomalies. PA and PE gain their maximum
and minimum in Fig. 2c, when positive and negative  Q centers are located over the Asian- and
extra-Asian monsoon hemispheres, respectively.
In contrast, PA and PE reach their minimum and
maximum, when the reversed spatial structure of  Q occurs. It becomes clear that interannual variations
of PA and PE are modulated by the ENSO cycle through the eastward propagation of the global divergent
circulation in response to the interannual variation of the tropical Pacific SST anomalies. This eastward
propagation of summer (  Q , P) anomalies is further illustrated by the transition of the global (  Q , P)
distribution from the 1982 summer warm episode, through the 1984-85 summer cold episodes, to the 1987
summer warm episodes (Fig. 18). As shown by Chen and Yoon (2000), a positive (negative) center of
(  Q , P) is located over the Asian monsoon hemisphere during cold (warm) summers. The transition
between them is not caused by an east-west seesaw oscillation, but by the eastward propagation of
(  Q , P) anomalies. Note that PG  12 ( PA  PE ) . The in-phase interannual variations of PG and PA
indicates once again that rainfall/convection in response to the global ENSO mode is more sensitive over
the Asian monsoon hemisphere than the extra-Asian monsoon hemisphere.
Our next concern is how the global ENSO mode affects interannual variation of regional monsoon
rainfall in the Asian monsoon system. We shall illustrate its impact on the East-Asian monsoon rainfall as
25
an example of downscale interaction. Because the atmospheric circulation is largely depicted by the
rotational flow, the interannual variation of the East-Asian summer circulation to positive SST anomalies
over the western tropical Pacific is a short-wave train (emanating from this region) along the rim of the
North Pacific with an anticyclonic cell over Korea and Japan. The reversed phase anomalous short-wave
train appears during cold western tropical SST anomalies. The dipole of anomalous circulation in the
western Pacific was named by Nitta (1986, 1987) as the Pacific-Japan pattern. Later, it was shown by
Chen et al. (1998) that the southern cell of this pattern not only shifts northward, but also enhances the
tropical cyclone occurrence frequency when SST anomalies in the western tropical Pacific become warmer.
In contrast, tropical cyclone activity is weakened and migrated southward during cold SST anomalies of the
western tropical Pacific.
The transition of summer global (  Q , P) anomalies between the warm and cold SST (NINO3.4)
phases is achieved by the eastward propagation of summer (  Q , P) anomalies. What is the response of
the East-Asian summer hydrological processes to this eastward-propagating ENSO mode? The interannual
variation of the East-Asian summer monsoon is often portrayed by the Pacific-Japan pattern in accompany
with the east-west oriented anomalous rainbelts (e.g. Samel et al. 1999). Because the East-Asian monsoon
circulation is characterized by the Meiyu rainbelt and tropical cyclone activity, can the aforementioned
response of the East Asian monsoon be reflected by a north-south oscillation of the Pacific-Japan pattern
and an alternation of the accompanied anomalous rainbelts? This question is answered by the y-t diagrams
of lagged correlation coefficients between PA and u (850mb) at 140oE (Fig. 19a) and between PA
and P (140oE) (Fig. 19b). As inferred from these two y-t diagrams, the response of the East-Asian
monsoon circulation and its rainfall should be a northward migration of the Pacific-Japan pattern [indicated
by u (850mb)] with its accompanied rainfall anomalies [indicated by P ], instead of the north-south
oscillation of this anomalous circulation pattern.
The major rain producers of the East-Asian monsoon are the Meiyu rainbelt and tropical cyclones.
Because of the spatial relationship between the Meiyu rainbelt, tropical cyclone activity, and the North
26
Pacific anticyclone, it is expected that frontal and tropical cyclone activities are modulated by the
northward propagation of the Pacific-Japan pattern. Following this anomalous circulation pattern and its
accompanied rainfall anomalies, can interannual variations of cold fronts along the Meiyu rainbelt and
tropical cyclone activity also undergo a northward migration? First, let us use PA , SST (NINO3.4) and
SST (WTP) as interannual variation indices of the northern summer monsoon circulation. Four phases of
these indices are designated in Fig. 20. Composite charts of ( Q , P) , ( Q , N FT ) , and ( Q , N TC )
for these four designated phases of the warm-cold extreme summer of the ENSO cycle identified with
summer interannual variation indices are shown in Fig. 21. Salient features of these composite charts are
summarized as follows:
1) A distinct Pacific-Japan pattern and east-west oriented strips of rain anomalies emerge from composite
charts of ( Q , P) during warm and cold summers. The ( Q , P) anomalies of two transition
phases (1, 3) juxtaposed with those of the two extreme phases (2, 4) form a clear northward migration
of the ( Q , P) wave train.
2) The Meiyu rainbelt extending from east China to southern Japan is established by the eastward
propagation of cold fronts which are advected by monsoon southwesterlies.
These monsoon
southwesterlies are intensified (weakened) when the anticyclonic (cyclonic)  Q cell of the
Pacific-Japan pattern is centered over the Philippine Sea (phase 2) and east of Taiwan (phase 3).
Therefore, frontal activity and the Meiyu rainbelt can be enhanced (reduced) by intensification
(weakening) of these monsoon southwesterlies which is a reflection of the East-Asian monsoon
intensity. Frontal activity is clearly modulated by the northward propagation of the Pacific-Japan
pattern. It was observed by Nitta (1987) that Japan becomes dry/hot (wet/cool) during the warm (cold)
SST (NINO3.4) summer. The intensity modulation of monsoon southwesterlies by the Pacific-Japan
pattern also affects rainfall in South China and Taiwan.
3) The occurrence frequency of tropical cyclones is enhanced (suppressed) by the cyclonic (anticyclonic)
southern cell of the Pacific-Japan pattern (Chen et al. 1998). It is shown in Fig. 21 that tropical cyclone
27
activity is covered by the cyclonic southern cell of the Pacific-Japan pattern during phase 1 and 4 (cold).
In contrast, tropical cyclone activity is expelled from the anticyclonic southern cell during phases 2
(warm) and 3 to appear along the southern rim of this cell where strong anomalous easterlies exist.
Because tropical cyclones are major rain producers, rainfall is therefore increased (reduced) over the
cyclonic (anticyclonic) southern cell of Pacific-Japan pattern.
In summary, we conclude that the modulation of the East-Asian hydrological processes by the eastward
propagation of the global ENSO mode depicted by (  Q , P) anomalies can be accomplished through its
downscale interaction with the regional monsoon circulation reflected by the northward migration of the
Pacific-Japan pattern and the regulation of the frontal and tropical cyclone activity by this northward
propagation property.
5. Summary and Concluding Remarks
Because of the importance of summer rainfall to the monsoon region, monsoon research has long been
focused on the search for mechanisms of the onset, life cycle, and interannual variation of monsoon rainfall.
During the northern summer, major global rainfall is actually produced in the Asian monsoon-western
tropical Pacific region. This special feature of northern summer rainfall leads us to explore two issues of
the global/hemispheric hydrological cycle in the present study: (1) the role played by the Asian monsoon in
the global hydrological cycle, and (2) effects of the propagating global modes, including diurnal,
intraseasonal, and interannual modes, on hydrological processes in the Asian monsoon region. Results of
these two issues are summarized as follows:
a. The role played by the Asian monsoon in the global hydrological cycle
As indicated by the global distribution of seasonal-mean values and variance, northern summer (JJA)
rainfall occurs largely over the Asian monsoon-western tropical Pacific region. The water vapor budget
analysis of the hemisphere (60oE-120oW) covering this region (which is designated as the Asian monsoon
hemisphere) was used to illustrate the contribution of the Asian monsoon to the global hydrological cycle.
Climatologically, the Asian monsoon hemisphere contributes 50% more than the extra-Asian monsoon
28
hemisphere (120oW-60oE) to global precipitation. Therefore, part of this excessive rainfall in the former
hemisphere is supported by water vapor diverged out of the latter hemisphere and converged toward the
former hemisphere. Eventually, the excessive rainfall in this hemisphere is transported from the Pacific,
across the Bering Strait and the Arctic Ocean, into the Atlantic. Based on this hemispheric water vapor
budget, it is clear that the Asian monsoon hemisphere is a sink region of water vapor and the extra-Asian
monsoon hemisphere is a source region of water vapor. The water vapor budgets of both hemispheres
exhibit temporal variations of different time scales. Variations of three time scales, including diurnal,
intraseasonal, and interannual modes, were analyzed in this study. Several interesting features of these
variations emerge. Amplitudes of these three variation modes are in the same order of magnitude (which is
an order of magnitude smaller than that of precipitation itself). An out-of-phase, east-west seesaw of
precipitation and water vapor flux exchange is observed between the two hemispheres. The variations of
global precipitation follow those of the Asian monsoon hemisphere because it contributes more to global
precipitation than the extra-Asian monsoon hemisphere. Nevertheless, roles played by the two hemispheres
in the global hydrological cycle are not changed by the variation of any time scale at all. The upscale
interaction of the Asian monsoon with the global hydrological cycle is revealed through the water vapor
budget of the Asian monsoon hemisphere.
b. Effects of global propagating modes on Asian monsoon hydrological processes
Regional processes involved in the evolution of the Asian monsoon were the major concern of
previous monsoon research. In contrast, the coupling between the global and regional monsoon modes has
not been a focus of this research. Actually, it was shown in this study that variations of the hemispheric
water vapor budget and hydrological processes over the Asian monsoon region are modulated by the
propagating global modes.
(1) Diurnal Variation
Because the Earth rotates eastward, the immediate atmospheric response to the wave-1, east-west
differential solar heating/cooling at the surface is the westward propagation of the global divergent
29
circulation. Diurnal seesaw variations in hydrological variables of the hemispheric water vapor budgets
between the Asian monsoon and the extra-Asian monsoon hemisphere are caused by this westward
propagation of the diurnal mode of global divergent circulation. The tropical atmosphere is conditionally
unstable, while the troposphere in middle-high latitudes is climatologically stable. Thus, the atmospheric
response to the surface solar heating in triggering convection/rainfall is faster in the tropics than in
middle-high latitudes. On the other hand, the heat capacity contrast between east-west land and ocean, and
between land (mountain range) and air enables the land to be warmed up by the solar heating faster than
water and air. A collaboration of these two differential responses to the surface solar heating forms a
clockwise rotation of diurnal variation in the continental-scale rainfall.
(2) Intraseasonal variations
The intraseasonal variations of northern summer rainfall are contributed to by 30-60 day and 12-24
day modes. Both global modes exhibit a clear eastward propagation, but couplings between these two
modes and the corresponding regional monsoon modes differ from each other.

~
~
~
30-60 day mode: Maximum (Minimum) PA and minimum (maximum) PE occur at phase 2(6) of the PA
index, when the positive (negative) ~
 Q center appears in the western tropical Pacific and the negative
(positive) ~
 Q center covers the Amazon basin. Over East Asia, the regional circulation in response to
the eastward propagating 30-60 day mode is reflected by the alternate northward migration of positive
~ cells. The East-Asian monsoon rainfall produced by cold fronts along the Meiyu
and negative 
Q
rainbelt before the monsoon break and by tropical cyclones after the break is regulated by the
~ cells.
northward-migrating 
Q

12-24 day mode: P̂A reaches its maximum (minimum) and P̂E attains its minimum (maximum) at phase
2(5) of the P̂A index, when the positive (negative) ̂ Q center moves to the western tropical Pacific and
the negative (positive) ̂ Q center propagates to Central America. The response of the East Asian
monsoon to the global ( ̂ Q , P̂ ) mode before the monsoon break is a southeastward propagation of ̂ Q
cells which modulate the cold-air intrusion following the eastward propagation of cold fronts. After the
30
monsoon break, tropical cyclone activity is modulated by the northwestward propagation of a 12-24 day
̂ Q short-wave train. The East-Asian monsoon rainfall is regulated by the global 12-24 day mode
through these regional monsoon modes.
(3) Interannual variation
The east-west seesaw variation of the water vapor budgets between the Asian-monsoon and the
extra-Asian monsoon hemispheres is caused by the eastward propagation of the interannual ( Q , P) mode.
This propagating  Q mode follows the SST (NINO3.4) index in the following manner: PA gains its
maximum (minimum) and PE attains its minimum (maximum) when the positive (negative)  Q center
is located over the tropical western Pacific and the negative (positive)  Q center is situated over the
tropical central-eastern Pacific during phase 4(2) of the PA index and SST (NINO3.4)  -0.5oC (  0.5
o
C).
It was suggested by previous studies (e.g. Chahine 1992, Soden 2000) that warm (cold)
SST (NINO3.4) anomalies intensify (weaken) the tropical/global hydrological cycle. Our results show
otherwise because the convective activity/rainfall in the Asian monsoon hemisphere is more sensitive to the
eastward propagation of the interannual  Q mode. The regional response to this interannual variation
mode was often portrayed as a flip-flop oscillation between the warm and cold Pacific-Japan pattern. It was
found in this study that this response is a northward migration of the Pacific-Japan pattern. Frontal and
tropical cyclone activities are modulated by this pattern to produce the north-south juxtaposition of the
east-west oriented rainfall anomalies.
The global and regional hydrological processes presented in this study offer a new perspective of the
East-Asian monsoon hydrological cycle: (1) its relationship with the extra-Asian monsoon hemisphere and
contribution to the global hydrological cycle, (2) its role in the temporal variations of the global
hydrological cycle, and (3) the hydrological interaction between the global and regional monsoon modes
with various time scales. Numerous previous studies of the Asian monsoon focused primarily on either
special features of this monsoon’s evolution, or a given time scale, e.g. diurnal, intraseasonal, interannual,
etc. Therefore, numerical simulations of the Asian monsoon often pay special attention to either a special
31
regional phenomena (e.g. Rajendran et al. 2004) or a special time scale (e.g. Lau and Yang 1996, Lau et al.
2000). In fact, the Asian monsoon simulated by global climate models can possibly include all temporal
and spatial scales. So far, no comprehensive analyses of the Asian monsoon simulations have attempted to
cover time scales ranging from diurnal to interannual, and both the upscale and downscale interaction with
the propagating global modes. The global/regional water vapor budgets and hydrological processes
presented in this study provide a new way to validate Asian monsoon simulations with global climate
models.
Acknowledgements
The technical assistance provided by Judy Huang, Simon Wang, Paul Tsay, and Adam Clark in preparing
the manuscript is gratefully acknowledged.
32
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38
Figure Captions:
Fig. 1
The global hydrological cycle: (a) precipitation P, potential function of water vapor flux  Q , and
divergent water vapor flux QD for the northern summer (JJA), (b) the long-term mean
hemispheric water vapor budgets of the Asian monsoon hemisphere
monsoon hemisphere
 A
and the extra-Asian
 E , and (c) the long-term mean freshwater transport of the global ocean by
Wijffels et al. (1992). Units of all hydrological variables in the hemispheric water budget are
mm·day-1.
Fig. 2
~
Variance (Var) of global precipitation fluctuations: (a) diurnal cycle P  , (b) 30-60 day mode P ,
(c) 12-24 day mode P̂ , and (d) interannual variation mode P . Contour intervals of these
precipitation variances are colored in terms of scale of numerical values shown in the lower right
side of each panel.
Fig. 3
Temporal variations of hemispheric-mean P and   Q and global-mean precipitation PG : (a)
 G ,  A , and  E
diurnal variation, (b) intraseasonal variation, and (c) interannual variation.
are variables ( ) averaged over the globe (thick-solid line), Asian monsoon (thin-solid line), and
extra-Asian monsoon (short-dashed line) hemisphere, respectively. Global and hemispheric mean
values of hydrological variables are denoted by thin dashed lines.
Fig. 4
Power spectra of PG , PA , and PE time series are shown in Figs. 3a and 3b: (a) diurnal, and (b)
intraseasonal. Diurnal variation is dominated by the diurnal cycle, while intraseasonal variation is
contributed to by both the 30-60 day and 12-24 day modes.
Fig. 5
Same as Fig. 1b, except for the maximum and minimum phase of the diurnal cycle. The long-term
climatological values of all hydrological variables shown in Fig. 1b are parenthesized in this figure
for comparison.
Fig. 6
~

Intraseasonal variations of P  s and (  Q)  s : (a) 30-60 day mode ( ) , (b) 12~24 day ( ) ,
~

and (c) combination of ( ) + ( ) + ( ) , where ( ) = annual mean of ( ).
~
Fig. 7
Phase of (a) 30-60 day mode, and (b) 12-24 day mode defined in terms of PA and P̂A time series.
Fig. 8
Same as Fig. 1b, expect for 30-60 day mode [(a) maximum phase and (b) minimum phase] and
39
12-24 day mode [(c) maximum phase and (d) minimum phase]. The maximum and minimum
phases of these two intraseasonal modes are determined according to the phase schematic shown
in Fig. 7.
Fig. 9 Same as Fig. 1b, expect for the interannual ENSO mode: (a) maximum phase and (b) minimum
phase determined by PA and SST (NINO3.4) time series. An example is shown in Fig.3c.
Fig. 10 Composite charts of (  Q , P  ) (JJA) every three hours:  Q anomalies are contoured, while P 
anomalies are stippled. Local noon and midnight are marked by an open strip and thick-solid line,
respectively. The contour interval of  Q anomalies is 5x104g·kg-1m2s-1 and P  anomalies are
colored with the gray colors shown in the lower right side of the top panel of (  Q , P  ) at 00UTC.
Fig. 11 Enlarged regional composite charts of (  Q , Q , P  ) (JJA) anomalies in the Asian monsoon
hemisphere every three hours extracted out of the corresponding global composite charts in Fig.
10: P  anomalies are stippled in terms of the color scale shown in the lower right side of the
(  Q , Q , P  ) composite chart at 12 UTC. Local noon and midnight are marked by an open strip
and thick-solid line, respectively. The contour interval of  Q is 5x104g·kg-1m2s-1.
Fig. 12 Maximum phase of the P diurnal cycle expressed in terms of (a) universal time (UTC) and (b) local
solar time (LST). The maximum P phase is indicated by a stick with a dot on its root (UTC) or an
arrow (LST).
Fig. 13 y-t diagrams of various P anomalies at 120oE-125oE and power spectra of P anomalies (P
anomalies with its annual-mean, annual, and semiannual components removed): (a) P, (b) P , and
~
(c) P  Pˆ . Power spectra of P at (22.5o-25oN) of every summer (thin-solid lines) and averaged
~
power spectra (thick-solid line) are shown. Contour intervals of P, P , and P  Pˆ are shown on
the top of every panel.
~
Fig. 14 Summer (May-September) composite charts of (a) ( ~Q , P ) and ( ˆ Q , Pˆ ) anomalies based on
~
phases defined by PA and P̂A indices shown in Fig. 7. Contour intervals of ~Q and ̂ Q , and
~
scales of P and P̂ are shown on the top of the first panel in each column.
Fig. 15 Composite charts of ~Q anomalies in the East Asian region for all eight phases of the 30-60 day
40
cycle corresponding to those shown in Fig. 14a and superimposed with N FT (frontal occurrence
~
frequency) and N TC (dot; tropical cyclone locations) in (a) and with P anomalies in (b). The
northward migration of the 30-60 day ~Q mode, which may be perceived from composite charts
of ~Q over its entire life cycle (eight phases), can be summarized by (c) the y-t diagram of ~Q
(120oE) anomalies in 1996 summer. The contour interval of ~Q , numerical values of N FT , and
~
scale of P are shown around the top right corner.
Fig. 16 Same as Fig. 15, except for six phases of the 12-24 day mode of ~Q before (a, b) and after (c, d)
the Asian monsoon break, and (e) the y-t diagram of ˆ Q (125oE) in 1999 summer before the
monsoon break, and in 1999 summer after the break. The contour interval of ~Q , scale of N FT
and P̂ anomalies are shown on the top of all columns.
Fig. 17 (a) The x-t diagram of  (  Q , P ) (Eq-10oN) anomalies, and (b) time series of SST (NINO3.4)
(JJA). Warm (cold) extreme summers are defined by SST (NINO3.4)  -0.5o (  -0.5oC) and
marked by W (C). The contour interval of  Q is 5x104g·kg-1m2s-1, and the color scale of P is
shown on the top of (a).
Fig. 18 The eastward propagation of summer  (  Q , P ) anomalies is illustrated by the warm-cold-warm
ENSO cycle (1982-1987).
Fig. 19 The y-t diagram of lagged correlation coefficients (a) rP
o
A  u (850m b)140 E
u (850mb)140oE, and (b) rPA P
140 o E
between PA and
between PA and  P140oE.
Fig. 20 Composite cycle of warm and cold extreme summers (condensed from Fig. 17) determined by
PA . Summers between 1979 and 2004 that are projected on different phases of this composite
cycle are shown on the top panel. Schematic cycles of SST (NINO3.4) and SST (WTP)
matched with the PA composite cycle are shown in the bottom panel.
Fig. 21 Composite charts of (a)  ( Q , P ) , (b)  ( Q , N FT ) , and (c) ( Q , N TC / TD ) over four phases
of the extreme warm-cold summer cycle (based on summers shown in Fig. 20a). The contour
interval of  Q is 2x106g·kg-1m2s-1. Color scales of N FT and NTC / TD are shown on tops of
columns (a) and (b).
41