15 APRIL 2006

TAKAHASHI AND YASUNARI

1545

A Climatological Monsoon Break in Rainfall over Indochina—A Singularity in the
Seasonal March of the Asian Summer Monsoon
HIROSHI G. TAKAHASHI
Graduate School of Environmental Studies, Nagoya University, Nagoya, JapanTETSUZO YASUNARI
Hydrospheric Atmospheric Research Center, Nagoya University, Nagoya, Japan(Manuscript received 21 February 2005, in final form 20 October 2005)
ABSTRACTThis study investigated the climatological pentad mean annual cycle of rainfall in Thailand and theassociated atmospheric circulation fields. The data used included two different data of rainfall: rain gauge
data for Thailand from the Thai Meteorological Department and satellite-derived rainfall data from the
Climate Prediction Center (CPC) Merged Analysis of Precipitation (CMAP).
Climatological mean pentad values of rainfall taken over 50 yr clearly show a distinct climatologicalmonsoon break (CMB) occurring over Thailand in late June. The occurrence of the CMB coincides with adrastic change of large-scale monsoon circulation in the seasonal march. The CMB is a significant singularity
in the seasonal march of the Southeast Asia monsoon, which divides the rainy season into the earlymonsoon and the later monsoon over the Indochina Peninsula.
A quasi-stationary ridge dynamically induced by the north–south-oriented mountain range in Indochina
is likely to cause the CMB. The formation of the strong ridge over the mountain ranges of Indochina ispreceded by a sudden enhancement (northward expansion) of the upstream monsoon westerlies along alatitudinal band between 15° and 20°N in late June. The CMB also has an impact downstream. The
orographically induced stationary Rossby waves enhance the cyclonic circulation to the lee of Indochina,
and over the South China Sea. The enhancement of cyclonic circulation may be responsible for the summer
monsoon rains peaking in late June over the South China Sea and the western North Pacific, and in the baiu
front.

1. Introduction
The last decade has seen a surge in studies of themonsoon over the Indochina Peninsula (IP). The summer monsoon over the IP has many peculiar featuresthat distinguish it from the adjacent Indian and western
North Pacific [WNP; including the South China Sea(SCS)] monsoons (Matsumoto 1997; Wang and LinHo
2002).
Matsumoto (1997) noted that there are many significant rain events over Thailand, referred to as premonsoon rains, before the monsoon onset in mid-May.
Rainfall increases gradually from the premonsoon rainsto the summer monsoon onset. The monsoon onset

Corresponding author address: Hiroshi G. Takahashi, Graduate
School of Environmental Studies, Nagoya University, Nagoya,
Aichi, 464-8601, Japan.
E-mail: s030111d@mbox.nagoya-u.ac.jp© 2006 American Meteorological Society

over Thailand is a more gradual process than the abruptchange from very dry conditions at onset over India or
the SCS. Wang and LinHo (2002) analyzed rainy season characteristics and noted that the monsoon over the
IP differs in many essential aspects from the adjacentIndian and WNP summer monsoons. They suggested
that the IP might serve as a transition zone, or boundary, between the two tropical monsoon subsystems: theIndian monsoon and the WNP monsoon. Interesting
features in monsoon rainfall over the IP include an
early onset, late peak, and withdrawal.
Matsumoto (1997) found that the rainy season in
Bangkok starts in mid-May. Zhang et al. (2002) foundthat rainy season onset over all Thailand is around 9
May. Wang and LinHo (2002) mapped the onset of therainy season for the entire Asian–Pacific monsoon andshowed that the summer rainy season starts in the IP inearly May (P26–27: “pentad” will be referred to as “P”
in this study), which is the earliest onset over land in

1546

JOURNAL OF CLIMATE

Asia. The rainy season peaks over the IP in September(P52–56), which is the latest peak over land in Asia
(Wang and LinHo 2002).
Another unusual feature of the IP monsoon is its
significant bimodality in the seasonal march. There is astrong peak in rainfall in September and a weaker peakin late May to early June (Fig. 2). A notable break inmonsoon rainfall occurs over Thailand (Fig. 2) in lateJune. This midsummer break is called the climatological monsoon break (CMB). The cause of the CMB overThailand has not been studied. The primary purpose ofthis study is to document circulation changes before,
during, and after the CMB and to uncover the fundamental processes responsible for the CMB.
It has long been recognized that the annual evolutionof the rainy season and summer monsoon over Asia is
not smooth. Stepwise evolution, subseasons, and sudden changes have been studied extensively for many
years (e.g., Yoshino 1965; Kato 1985; Ninomiya andMuraki 1986; He et al. 1987; Tao and Chen 1987; Lau etal. 1988; Matsumoto 1992; Nakazawa 1992; Murakami
and Matsumoto 1994; Ueda et al. 1995; Ueda andYasunari 1996; Wang and Xu 1997; Kawamura andMurakami 1998; Kang et al. 1999; Wu and Wang 2000,
2001; LinHo and Wang 2002). Wang and Xu (1997)
used the term climatological intraseasonal oscillations(CISOs) to refer to intraseasonal components of outgoing longwave radiation (OLR) in climatological timeseries. They identified four subseasonal cycles in the
Asian–Pacific summer monsoon. Each CISO event was
statistically significant, had a coherent dynamic structure, and propagated in a systematic way; consequently,
they nearly qualified as a physical mode on climatological maps. They showed that extremely wet and dryphases of the CISO represent monsoon singularities,
that is, wet and dry events that occur with regularity
during the fixed pentads. LinHo and Wang (2002)
found rapid variation (on an intraseasonal time scale)in the climatological time series and named this the fastannual cycle (FAC). The FAC over the IP is much
weaker than over the WNP or South Asia (however,Fig. 2 shows that the CMB time scale is intraseasonal).
Exploring how the CMB is related to intraseasonalvariation (ISV) is warranted.
Monsoon breaks over India occur when the monsoon
trough is close to the foot of the Himalayas. At thistime, there is a striking decrease in rainfall over most ofIndia, although rainfall increases along the Himalayas
and over southern India (Rao 1976). Many observational studies during the past two decades have sought
to explain active/break monsoon spells from the perspective of ISV in the summer monsoon, especiallyover south Asia. Yasunari (1979, 1980, 1981) found

VOLUME 19

cloud bands to propagate northward over India on atime scale of 30–50 days and related the bands to monsoon active/break spells over India. Krishnamurti andArdanuy (1980) related monsoon breaks over India tohigh pressure anomalies in a westward-propagating
mode with a period of 10 to 20 days. Monsoon breaks
over Asia are now thought to be equivalent to breakspells in the ISV. However, this dose not explain the
CMB over Thailand. Wang and Xu (1997) and LinHo
and Wang (2002) showed that climatological ISV components of OLR over the IP is much weaker than theother monsoon regions. However the CMB over Thailand is prominent (Fig. 2). Therefore, this study addresses why the transient ISV over the IP is much
weaker than over India and the WNP, but the CMB
over Thailand is obvious. Understanding the causes of
monsoon singularities in other regions of the Asian–Pacific monsoon will be enhanced as causes of the CMB
are revealed in this study.
Section 2 documents the data used in this study. Section 3 describes the characteristics of the annual rainfall
over Thailand. Spatial structures in rainfall and atmospheric circulation fields are investigated in section 4.
Possible dynamics behind the CMB are discussed in
section 5, and conclusions are given in section 6.

2. Data
Daily rainfall data were taken from 32 stations inThailand (Fig. 1) and compiled by the Thai Meteorological Department (TMD). Data used are from 1951 to
2000. In Fig. 1, the stations are almost uniformly distributed in Thailand. These stations have 50-yr daily
rainfall records. Missing data from each station are limited to less than 1%.
Climate Prediction Center (CPC) Merged Analysis
of Precipitation (CMAP; Xie and Arkin 1997) data
were used to analyze the large-scale rainfall distribution. CMAP is a global dataset on a 2.5 ⫻ 2.5 grid from1979 to 2000.
The National Centers for Environmental Prediction–
National Center for Atmospheric Research (NCEP–
NCAR) reanalysis data (Kalnay et al. 1996) yielded
zonal and meridional winds (u, ␷) and geopotential
heights (z) at 850 and 700 hPa on a 2.5 ⫻ 2.5 global grid.
Also, to calculate vertically integrated moisture flux,
horizontal winds, geopotential heights, relative humidity, and air temperature from the surface to 500 hPa as
well as surface pressure were used.
Pentad data calculated from daily data for each field
describe the climatological subseasonal variability.

15 APRIL 2006

TAKAHASHI AND YASUNARI

1547

FIG. 2. Climatological pentad mean time series of rainfall index
for 50 yr in units of mm day⫺1. The index is averaged rain gauge
observations at 32 stations. To illustrate atmospheric circulation
associated with this rainfall index, the rainy season was divided
into eight periods.

FIG. 1. Location of the 32 stations in Thailand for which daily
rainfall records are available for the period 1951–2000.

3. A singularity in Thailand rainfall
Matsumoto (1997) described two rainfall peaks in thesummer monsoon in Thailand. However, his rainfall
data were limited to Bangkok, Thailand, and he focused on the onset and withdrawal of the summer monsoon in Thailand. This study uses area-averaged pentad
mean rainfall for 50 yr (1951–2000) for 32 stations overThailand, as shown in Fig. 2, to examine the characteristics of the climatological seasonal march of rainfall
over Thailand.
The summer monsoon (rainy season) over the IP occurs from May to October, while the winter monsoon(dry season) is from November to April. Seasonal rainfall variation shows two significant peaks in the climatological annual cycle over Thailand: the first rainy
peak (FRP) occurs from P28 to P30; and the second
rainy peak (SRP) occurs from P49 to P51. A monsoon
break occurs in P34 to P36 between the FRP and SRP,
and this break is defined as the CMB. The rainfall timeseries is averaged over 50 yr at 32 stations, but the CMB

is still obvious. Climatological seasonal variation in therainfall at each station shows that the CMB occurs at
almost all stations during the same period (not shown).
The CMB occurs over all Thailand at nearly the sametime each year.
The CMB is statistically meaningful for climatology
because it has statistical significance over 50 yr. We
confirmed that the CMB is a true monsoon singularity
rather than an artifact owing to extreme events in someyears by examining the occurrence frequency of eachpentad such that each pentad’s rainfall was more than athreshold value over 50 yr. The threshold value was themean rainfall during the rainy season, defined as theaverage from P26 to P60 based on previous studies(Matsumoto 1997; Zhang et al. 2002). The rainy seasonal mean was 6.7 mm day⫺1. Figure 3 shows the frequency of occurrence of precipitation more than thethreshold value. Low frequencies occur during the dry
season and around the CMB, which indicates that rainfall during the CMB is less than the rainy season mean
in most years. A small number of extreme events do notmake up the CMB. Instead, the CMB is a real monsoon
singularity that is fixed to the seasonal march. A periodof high frequency also appears during the SRP. The
SRP is an interesting phenomenon because the Indian
monsoon and WNP monsoon are weak in this period,
when monsoon rainfall over Thailand is highest. By
contrast, monsoon precipitation is high in all three regions during the FRP. Monsoon rains over Thailand
have unique characteristics that distinguish them fromthe Indian and WNP monsoons. The CMB divides the
monsoon season over Thailand into the FRP and SRP;
surrounding large-scale atmospheric circulation fea-

1548

JOURNAL OF CLIMATEVOLUME 19

a. The first increasing phase (periods 1 to 4)

FIG. 3. The occurrence frequency for each pentad when eachpentad rainfall is more than a threshold value. The thresholdvalue was determined as the rainy seasonal mean. The rainy seasonal mean averaged was at 6.7 mm day⫺1.

tures before and after the CMB also differ. The CMB
may not be a local phenomenon, but rather a significant
monsoon singularity associated with the seasonal marchof the Asian–Pacific monsoon system.
This study defines eight periods in the summer monsoon season to facilitate an examination of atmosphericcirculations associated with seasonal variation in rainfall over Thailand. The definitions are shown in Fig. 2.
The first and second periods occur during times in theannual cycle after the dry season when rainfall totals
are increasing. The third period is the FRP. During thefourth period, rainfall totals are at a peak and start to
decline. The fifth period is the CMB. During the sixthand seventh periods, rainfall totals again increase until
the eighth period, which is the SRP.

4. Atmospheric circulation associated with rainfall
over Thailand
The previous section presented the characteristics of
the CMB, namely, a monsoon singularity with a considerable drop in rainfall. Figure 4 shows a sequence of
total CMAP rainfall distribution to illustrate the largescale seasonal march of rainfall. Figure 5 shows theassociated atmospheric circulations, namely the 850hPa geopotential height and wind fields. The eight periods shown in Figs. 4 and 5 are as defined in section 3.
Figures 6 and 7 show difference fields between one
period and the previous period for rainfall and atmospheric circulation, respectively. The focus is on the
CMB as an important singularity in the seasonal marchof the Asian–Pacific monsoon.

In period 1 (P22–24), the climatological threepentad mean of the 850-hPa horizontal wind (Fig. 5a)shows no monsoon westerlies over the south and Southeast Asia monsoon regions. The largest rainfall iscentered at 5°N, 90°E over the eastern Indian Ocean(Fig. 4a).
During period 2 (P25–27), a trough appears over the
Bay of Bengal (BOB). At the same time, weak monsoon southwesterlies appear over the eastern Indian
Ocean and reach the IP (Fig. 5b). Simultaneously, highprecipitation is observed over the eastern Indian Oceanand southern IP (Fig. 4b).
During period 3 (P28–30), a large amount of rainfall
is observed over the IP (Fig. 4c). The strong monsoon
rainfall over the BOB and IP precedes the other tropical monsoon regions. The monsoon southwesterlies are
established over the BOB, as the trough deepens overthe BOB (Fig. 5c). Over the IP, the wind direction issouthwesterly. The difference fields in rainfall showthat the increase of rainfall is observed over the BOB,
IP, northern SCS, and the south of Japan (Fig. 6a). The
monsoon westerlies are enhanced along the 10°N latitude (Fig. 7a), which is consistent with the increase in
rainfall. Thus, the southwesterlies from the trough overthe BOB is likely to induce the increase in rainfall overthe IP.
The monsoon westerlies over south Asia are drastically stronger during period 4 (P31–33), which corresponds with the monsoon onset in India (Fig. 5d). The
intensified monsoon is consistent with the abrupt increase in rainfall over India (Fig. 4d).

b. During the CMB
During the CMB (P34–36), the strongest band of
monsoon rainfall is oriented northwest–southeast from
north India to the WNP (Fig. 4e). An area of low rainfall amount is locally observed over the western IP,
which is consistent with area-averaged rain gauge rainfall. The monsoon rainband migrates northward over
the Arabian Sea, Indian subcontinent, and BOB (Fig.
5e). Over the IP, the monsoon wind changes direction
from southwesterly to westerly as the trough over the
BOB weakens. A strong ridge is formed over western
Indochina, which is consistent with the area of low rainfall amount there. Simultaneously, a trough over the
SCS, where the monsoon westerlies meet trade winds,
deepens, as rainfall increases, and the monsoon westerlies enlarge eastward to the Philippines. At the sametime, the subtropical high over the south of Japan intensifies, which strengthens moisture transport that

15 APRIL 2006

TAKAHASHI AND YASUNARI

1549

FIG. 4. Spatial distributions of rainfall by CMAP in each period based on rainfall index. Light (dark) shaded region denotes therainfall in excess of 6 mm day⫺1 (10 mm day⫺1).

supports increased rainfall over Japan during the onsetof the baiu season there.
The differences between periods 4 and 5 show that
decreased rainfall is observed over the western and central IP (Fig. 6b). The area of decrease in rainfall isconfined to Thailand, while the monsoon rainfall increases over most of the Asian monsoon regions, suchas the northern Indian subcontinent, SCS, WNP, andthe baiu region, except for western IP. Over India andthe BOB, the northward expansion of the stronger
monsoon westerlies is noticeable, which is consistent
with the northward migration of rainfall (Figs. 6b and7b). When the low-level monsoon westerlies strengthen
between 15° and 20°N upstream from the monsoon(over the BOB), a strong ridge appears over the western IP, which is clearly anchored at the Indochina

1550

JOURNAL OF CLIMATEVOLUME 19

FIG. 5. Each period mean distribution of horizontal wind vector and geopotential height at 850 hPa. Unit vector is 20 m s⫺1 asdenoted at upper right. Contour interval of geopotential height is 10 gpm.

mountain ranges (the Arakan, Dawna, and Tenasserim
Mountains). This ridge formation is likely to be responsible for the reduction of rainfall to form the CMB overthe IP. The formation dynamics of the strong ridge will
be discussed in the next section. Note also that a clear
cyclonic circulation appears downstream from the monsoon westerlies in association with the deepening of thetrough over the eastern IP and SCS. Another cyclonic
circulation also appears over the WNP. Both the cyclonic circulations are consistent with the rainfall increase there. The processes of rainfall increase over the
SCS and WNP are discussed in the next section. In
addition, the subtropical high centered at 20°N, 135°E
over the WNP intensified, and this may be a response of
the greater amounts of rainfall and convection over the
SCS and WNP (Nitta 1987).

c. The second phase of increasing rainfall (periods
6 to 8)During period 6 (P37–41) to period 7 (P42–46), theamount of rainfall is gradually decreased over the SCS
and WNP regions. Rainfall also decreases over theJapanese Archipelago, which corresponds with the endof baiu (Figs. 4f,g). Simultaneously, the subtropicalhigh retreats northward over the WNP. The reduction
in rainfall over the Japanese Archipelago and SCS isassociated with weakening of moisture transport along
the periphery of the subtropical high, which is causedby the retreat of the subtropical high.
During periods 7 and 8 (P42–46 and P47–51), themonsoon westerlies reach the WNP. The trough deepens over the WNP. During the SRP, the IP has a high

15 APRIL 2006

TAKAHASHI AND YASUNARI

1551

FIG. 5. (Continued)
FIG. 7. As in Fig. 6 except for in geopotential height and windat 850 hPa.

amount of rainfall, which is consistent with the rain
gauge rainfall data. The SRP is the late rainy peak over
the whole tropical Asian monsoon regions. The strong
trough is deepened over the SCS and the WNP, which
is clearly consistent with the frequent tracks of tropical
cyclones in this period.
To examine the enhancement of cyclonic activity
over the Asian monsoon regions, a Hovmöller diagramof vorticity at 700 hPa along the latitude band between
10° and 20°N is produced as shown in Fig. 8. Positive
vorticity over the SCS and WNP at this latitude band
increases with time between the CMB and SRP, thendecreases after the SRP. The cyclonic vorticity isstrongest from August to mid-September, expanding
westward to the IP. The cyclonic vorticity likely induces
moisture convergence and heavy rainfall over the IP
during the SRP. In addition, each wave packet of positive vorticity shows systematic westward phase propagation after the CMB, starting from the SCS and WNP,
which corresponds to the frequent tracks of tropical
cyclones.

d. Different rainfall systems
FIG. 6. Difference in rainfall (CMAP) between one period andthe previous period, based on the rainfall index. Light (dark)
shaded denotes rainfall increased (decreased). The periods are (a)periods 3 ⫺ 2 and (b) periods 5 ⫺ 4.

In this subsection, the two different rainfall systems
before and after the CMB are discussed. Figure 9 shows
related series of atmospheric circulation changes.

1552

JOURNAL OF CLIMATEVOLUME 19

FIG. 8. Hovmöller diagram of daily-averaged vorticity in thelower troposphere (700 hPa) along the latitudinal band averaged
within 10°–20°N.

The upper panel shows a trough over the BOB whenmonsoon westerlies are enhanced along a latitudinalband within 5°–10°N. The monsoon southwesterlies
from the trough over the BOB are associated with increased rainfall over the IP during the FRP.
During the CMB, a trough appears over the SCS and
WNP. During this period, the rainfall center shifts eastward from south Asia to the WNP, which is consistent
with vorticity fields in Fig. 8.
In the bottom panel, the trough is formed over the
SCS and WNP, which has intensified since the CMB.
After the trough has developed over the SCS and
WNP, transient waves with positive vorticity (cyclones)
display systematic westward phase propagation (Fig. 8).
Harr and Elsberry (1991, 1995) showed that variability
in the low-level zonal wind is related to the motion oftropical depressions over the SCS and WNP. Periods 7
and 8 contain atmospheric structures that favor westward-propagating tropical depressions. During the
SRP, these tropical depressions are associated with theincreased rainfall over the IP.
The CMB is a transition period to change rainfall
systems from the monsoon southwesterlies to tropicaldepressions. Thus, the dynamics that cause rain during
the FRP differ from those of the SRP.

FIG. 9. Atmospheric change associated with rainfall over Thailand during (a) the former increasing rainfall phase (FRP; period3), (b) the CMB (period 5), and (c) the latter increasing phase(SRP; period 8). The solid (dashed) lines denote ridges (troughs).

5. Discussions
a. Topographic effects
A strong ridge anchored at the Indochina mountain
ranges is intensified during the CMB, which is associated with the decrease in rainfall, when the monsoon
westerlies are enhanced along 15°–20°N. This suggests
an enhanced interaction between the monsoon westerlies and topography. This subsection investigates amechanism of the formation of the ridge over Indochina, in terms of topographic effects.
Gadgil (1977) explained the observed trough in thesouthern BOB and the ridge to the east over Burma as
stationary Rossby waves generated downstream fromthe Western Ghats Mountains on the Indian subcontinent. The IP is quite similar topographically to the Indian subcontinent. The latitudinal mountain ranges on

15 APRIL 2006

TAKAHASHI AND YASUNARI

the IP, the Arakan, Dawna, and Tenasserim Mountains, all have an average height exceeding 1 km, which
is higher than the Western Ghats. The strong monsoon
westerlies extend up to 600 hPa (not shown) and thestrongest westerlies are around 850 and 700 hPa. Monsoon westerlies are thicker than the mountain ranges
over Indochina, so stationary Rossby wave theory canbe applied to the IP. Gadgil (1977) discussed featureson the seasonal time scale. Figure 7b clearly shows evidence of topographic effects during three pentads. In
terms of the conservation of potential vorticity, therefore, a strong ridge forms over the Indochina mountain
ranges and a cyclonic circulation forms downstream of
the IP as a stationary Rossby wave. In this case, thewavelength of the stationary Rossby wave ␭ is given by冑␭ ⫽ 2␲

U,␤

共1兲

where ␤ is the planetary vorticity gradient and U is theaverage mean flow. When ␤ ⫽ 2.2 ⫻ 10⫺11 at 15°N, andwhen U ⫽ 10 m s⫺1 over the BOB at 850 hPa, thewavelength is about 4200 km. The half-wavelength(2100 km) is nearly equivalent to the zonal extent of thecyclonic circulation over the SCS (from the eastern IP
to the Philippines).

b. The mechanism of increase in rainfall over the
SCS and WNP during the CMB
The CMB over the IP appears simultaneously withan increase in rainfall over the SCS and WNP, which isconsistent with a strong ridge over the IP and a deep
trough over the SCS and WNP (Figs. 6b, 7b). In thissubsection, the processes of rainfall increase over the
SCS and WNP during the CMB are discussed.
The previous studies noticed a drastic increase inconvections and rainfall over the SCS, WNP, and baiu
region around late June (e.g., Wang and LinHo 2002).
Wang and Xu (1997) noted that the wet event of CISO
in mid-June corresponds with the drastic increase of
convections over the WNP as a monsoon singularity.
Wu and Wang (2001) pointed out the importance for adramatic increase in rainfall of the slow seasonal evolution of large-scale circulation as well as the CISO
over the SCS and WNP. LinHo and Wang (2002) foundthat the negative OLR anomalies on an intraseasonal
time scale propagate eastward rapidly from south Asia
to the SCS and WNP around the same time. Rainfall
shows a simultaneous increase over the SCS, WNP, andbaiu region during the CMB of the IP. Ueda et al.
(1995) suggested that SST warming is not a sufficientcondition, although it is certainly one important ingredient, to allow such a dramatic increase in convections

1553

FIG. 10. Hovmöller diagram of pentad-averaged moisture convergence along the latitudinal band averaged within 5°–15°N,
which is integrated from the surface to 500 hPa (shaded and thincontours). The averaged latitude is consistent with the rainfall
increase over the WNP. The thick contours denote the latitudinal(5°–10°N) averaged zonal wind at 850 hPa from ⫺2 to 2 m s⫺1.
The contour interval is 2 m s⫺1.

over the WNP. Warm SSTs support unstable conditions
over the SCS and WNP. An additional condition, suchas a rapid atmospheric trigger, is necessary for a drastic
increase in rainfall. The deep trough over the SCS appears, which is associated with the stationary Rossby
waves generated over Indochina mountain ranges inthe strong monsoon westerlies. The appearance of thetrough over the SCS could trigger the increase in rainfall over the SCS.
Figures 6b and 7b shows that rainfall also increaseover the WNP, and another cyclonic circulation is located there.
To reveal the processes of rainfall increase over the
WNP, the Hovmöllor diagram of convergence of water
vapor flux integrated from the surface to 500 hPa alonga latitudinal band between 5° and 15°N was produced(Fig. 10). In this latitude band, the rainfall increases
greatly during the CMB. To confirm the relationshipthe rainfall increases and the expansion of the monsoon
westerlies, the zonal wind along 5°–15° was added in
the figure. This latitude is the southern part of the cyclonic circulation over the WNP where the westerlies
are enhanced (Fig. 7b). In this figure, the moisture convergence over the WNP is drastically intensified during
the CMB. The intensification of moisture convergenceseems to be a trigger for causing the drastic increase in
rainfall over the WNP in the CMB. At P34, in addition,
the monsoon westerlies rapidly expand eastward andreach the WNP, which corresponds with the intensification of moisture convergence over the WNP. Thissuggests that the enhancement of monsoon westerlies,1554

JOURNAL OF CLIMATE

FIG. 11. The time–latitude cross section of climatological zonal
wind (a) averaged over 82.5°–87.5°E and (b) its year-to-year std
dev.

which is associated with the CMB, presumably makes
the moisture convergence intensify over the WNP.

c. Seasonal phase lock of the CMBThis section discusses mechanisms behind the strong
tendency of the CMB to occur at nearly the same timeeach year. The northward expansion of monsoon westerlies over the Indian monsoon region is likely responsible for the CMB, and how this drives the CMB atnearly the same time each year will be addressed. Figure 11 shows a time–latitude cross section of the climatological mean zonal wind over 50 yr (Fig. 11a) and its
year-to-year standard deviation (Fig. 11b) along 85°E
(the BOB longitude). In early and mid-June, strong
zonal winds exist between the equator and 10°N. In lateJune, the area of the strong zonal wind expands dramatically northward. The standard deviation in early
June between 5° and 10°N and in late June between 15°
and 20°N is smaller than at any other period and latitude in Fig. 11b. Enhanced monsoon westerlies appear
over the BOB at about the same time each year, inearly June between 5° and 10°N and in late June be-

VOLUME 19

tween 15° and 20°N. This determines CMB timing.
However, the origins and speed of the northward expansion of the monsoon westerlies may differ from year
to year. The timing of the northward expansion of themonsoon westerly is consistent with the first seasonal
phase lock of ISV of the monsoon westerlies (Nakazawa 1992). In addition, the area of expanded monsoon
westerlies matches the area of increased rain, whichindicates that northward migration of the active rainfall
area accompanies the expansion of the monsoon westerlies. The northward migration of rainfall is likely to
be associated with the northward propagation of ISV(Yasunari 1979, 1980, 1981). The first northward propagation of ISV over south and Southeast Asia tends tooccur at nearly the same time almost every year, so thefirst northward propagation of the ISV over south Asia
in June could be referred to as a climatological ISV
over the South Asian monsoon region. Examination of
the 850-hPa zonal winds for each year shows that thenorthward propagation occurs in June, although theorigins and northward speeds of the ISV vary. Therefore, the CMB is associated with climatological ISV
over the BOB. The ISV of OLR is weak (Wang and Xu
1997; LinHo and Wang 2002) over the IP, whereas the
CMB in rainfall over Thailand is obvious. Because thedecrease in rainfall during the CMB is spatially limited
and complicated, OLR could not show the decrease in
rainfall.
What physical mechanism causes the seasonal phaselocking of the CMB to the annual cycle? The timing of
the CMB may be significant because the summer solstice is coincident with the timing of the northward expansion of the monsoon westerlies and the heavy rainsin the Indian monsoon, and this expansion is associatedwith the CMB. Jiang et al. (2004) noted that easterly
vertical shear is important for the northward propagation of the ISV over the BOB. Easterly vertical shear ischaracterized by low-level monsoon westerlies and upper-level monsoon easterlies that are part of the Tibetan high. Intensification of monsoon easterlies in theupper troposphere as part of the Tibetan high is essential for the northward-propagating ISV. The Tibetan
high is intensified by heating over and around the Tibetan Plateau during the onset of the Indian summer
monsoon, so the timing of the peak in solar radiation,around the summer solstice, is closely related to thenorthward expansion of the low-level monsoon westerlies. In late July, low-level monsoon westerlies reintensify (Fig. 11a). However, the standard deviation in late
July exceeds that in late June (Fig. 11b). The Asianmonsoon may fluctuate because of internal dynamics
rather than pure external forcing after the CMB.

15 APRIL 2006

1555

TAKAHASHI AND YASUNARI

6. Conclusions
This study investigated the climatological pentad
mean annual cycle of rainfall in Thailand and the associated atmospheric circulation fields. The data used included rain gauge data for Thailand from the Thai Meteorological Department, satellite-derived rainfall datafrom CMAP, and atmospheric fields from the NCEP–
NCAR reanalysis.
Climatological mean pentad values of rainfall takenover 50 yr clearly show a distinct CMB occurring overThailand in late June. The atmospheric circulation
changes drastically during the CMB not only over the
IP but also over the Asian–Pacific monsoon region.
Thus the CMB is a singularity in the Southeast Asianmonsoon.
The decrease in rainfall is associated with the formation of a strong ridge over the mountain ranges of Indochina. The strong ridge anchored at the mountain
ranges forms when monsoon westerlies strengthenalong the latitudinal band of 15°–20°N. Enhancementof the monsoon westerlies over the BOB accompanies
the northward expansion of the monsoon westerlies
over south Asia as a part of the climatological ISV. The
strong ridge in the lower troposphere over the mountain ranges of Indochina associated with the CMB ispart of a stationary Rossby wave train that is generated
by the strong monsoon westerlies and north–southoriented topography. Therefore, small displacements ofthe enhanced monsoon westerlies have a significant
nonlinear effect because of the effects of the mountain
ranges of Indochina. The CMB also has an impact
downstream. The stationary Rossby waves propagate
downwind and a trough is enhanced over the SCS. This
trough likely triggers a drastic increase in rainfall over
the SCS. The trough over the SCS was deepened, whichmakes the monsoon westerlies reach the Philippines.
The eastward-expanded monsoon westerlies intensify
the moisture convergence over the WNP, which induces rainfall increase there.
Two types of rainfall systems were identified over
Indochina. Rainfall associated with the first type peaks
in late May to early June and is caused by an intensification of the monsoon southwesterlies along 10°N.
Rainfall associated with the second type peaks in early
September and is related to active tropical depressions
originating in the SCS and WNP. The CMB occurs between the two peaks.
Thus, the CMB, namely the formation of the wave
train on the strong monsoon westerlies, is a fundamental component the seasonal march of the Southeast
Asian monsoon. It serves as a bridge that links monsoon rainfall from south Asia to the SCS and WNP. The

CMB is, therefore, not solely a local phenomenon, but
is an important singularity in the seasonal march of themonsoon over the Asian–Pacific region. Further study
using regional climate models may help verify CMB
forcing mechanisms.
Acknowledgments. Rainfall data were kindly provided by the Thai Meteorological Department. We
thank Dr. H. Fujinami of Nagoya University for histhoughtful discussions and comments. We thank Prof.
B. Wang of the University of Hawaii for valuable comments and suggestions. We are grateful to Prof. J. Matsumoto of the University of Tokyo for valuable comments. Some figures were made with the Generic Mapping Tools (GMT) graphics system developed by
Wessel and Smith in 1995. This study was supported by
Dynamics of the Sun-Earth-Life Interactive System,
No. G-4, of the 21st Century COE program.
REFERENCES
Gadgil, S., 1977: Orographic effects on the southwest monsoon: A
review. Pure Appl. Geophys., 115, 1413–1430.
Harr, P. A., and R. L. Elsberry, 1991: Tropical cyclone track characteristics as a function of large-scale circulation anomalies.
Mon. Wea. Rev., 119, 1448–1468.
——, and ——, 1995: Large-scale circulation variation over thetropical western North Pacific. Part I: Spatial patterns andtropical cyclone characteristics. Mon. Wea. Rev., 123, 1225–
1246.
He, H., W. McGinnis, Z. Song, and M. Yanai, 1987: Onset of the
Asian summer monsoon in 1979 and the effect of the Tibetan
Plateau. Mon. Wea. Rev., 115, 1966–1995.
Jiang, X., T. Li, and B. Wang, 2004: Structures and mechanisms ofthe northward propagating boreal summer intraseasonal oscillation. J. Climate, 17, 1022–1039.
Kalnay, E., and Coauthors, 1996: The NCEP/NCAR 40-Year Reanalysis Project. Bull. Amer. Meteor. Soc., 77, 437–471.
Kang, I.-S., C.-H. Ho, Y.-K. Lim, and K.-M. Lau, 1999: Principal
modes of climatological, seasonal, and intraseasonal variations of the Asian summer monsoon. Mon. Wea. Rev., 127,
322–340.
Kato, K., 1985: On the abrupt change in the structure of the baiu
front over the China Continent in late May of 1979. J. Meteor.
Soc. Japan, 63, 737–750.
Kawamura, R., and T. Murakami, 1998: Baiu near Japan and its
relation to summer monsoons over Southeast Asia and thewestern North Pacific. J. Meteor. Soc. Japan, 76, 619–639.
Krishnamurti, T. N., and P. Ardanuy, 1980: The 10 to 20-day westward propagation mode and “breaks in the monsoons.” Tellus, 32, 15–26.
Lau, K.-M., G. J. Yang, and S. H. Shen, 1988: Seasonal and intraseasonal climatology of summer monsoon rainfall overEast China. Mon. Wea. Rev., 116, 18–37.
LinHo, and B. Wang, 2002: The time–space structure of the
Asian–Pacific summer monsoon: A fast annual cycle view. J.
Climate, 15, 2001–2019.
Matsumoto, J., 1992: The seasonal changes in Asian and Australian monsoon regions. J. Meteor. Soc. Japan, 70, 257–273.
——, 1997: Seasonal transition of summer rainy season over In-

1556

JOURNAL OF CLIMATE

dochina and adjacent monsoon region. Adv. Atmos. Sci., 14,
231–245.
Murakami, T., and J. Matsumoto, 1994: Summer monsoon over
the Asian Continent and Western North Pacific. J. Meteor.
Soc. Japan, 72, 719–745.
Nakazawa, T., 1992: Seasonal phase lock of intraseasonal variation during the Asian summer monsoon. J. Meteor. Soc. Japan, 70, 597–611.
Ninomiya, K., and H. Muraki, 1986: Large-scale circulation overEast Asia during Baiu period 1979. J. Meteor. Soc. Japan, 64,
409–429.
Nitta, T., 1987: Convective activities in the tropical western Pacificand their impacts on the Northern Hemisphere summer circulation. J. Meteor. Soc. Japan, 65, 373–390.
Rao, Y., 1976: Southwest Monsoons. Meteor. Monogr., No. 1, Indian Meteorological Department, 367 pp.
Tao, S., and L. Chen, 1987: A review of recent research on the
East Asian summer monsoon in China. Monsoon Meteorology, C.-P. Chan and T. N. Krishnamurti, Eds., Oxford University Press, 60–92.
Ueda, H., and T. Yasunari, 1996: Maturing process of the summer
monsoon over the Western North Pacific—A coupled ocean/
atmosphere system. J. Meteor. Soc. Japan, 74, 493–508.
——, ——, and R. Kawamura, 1995: Abrupt seasonal change oflarge-scale convection activity over the western Pacific innorthern summer. J. Meteor. Soc. Japan, 73, 795–809.
Wang, B., and X. Xu, 1997: Northern Hemisphere summer mon-VOLUME 19

soon singularities and climatological intraseasonal oscillation.
J. Climate, 10, 1071–1085.
——, and LinHo, 2002: Rainy season of the Asian–Pacific summer
monsoon. J. Climate, 15, 386–398.
Wu, R., and B. Wang, 2000: Interannual variability of summer
monsoon onset over the western North Pacific and the underlying processes. J. Climate, 13, 2483–2501.
——, and ——, 2001: Multi-stage onset of the summer monsoon
over the western North Pacific. Climate Dyn., 17, 277–289.
Xie, P., and P. Arkin, 1997: Global precipitation: A 17-year
monthly analysis based on gauge observations, satellite estimates, and numerical model outputs. Bull. Amer. Meteor.
Soc., 78, 2539–2558.
Yasunari, T., 1979: Cloudiness fluctuations associated with thenorthern hemisphere summer monsoon. J. Meteor. Soc. Japan, 57, 227–242.
——, 1980: A quasi-stationary appearance of 30- to 40-day period
in the cloudiness fluctuations during the summer monsoon
over India. J. Meteor. Soc. Japan, 58, 225–229.
——, 1981: Structure of the Indian monsoon system with around40-day period. J. Meteor. Soc. Japan, 59, 336–354.
Yoshino, M., 1965: Four stages of the rainy season in early summer over East Asia (part 1). J. Meteor. Soc. Japan, 43, 231–
245.
Zhang, Y., T. Li, B. Wang, and G. Wu, 2002: Onset of the summer
monsoon over the Indochina Peninsula: Climatology and interannual variations. J. Climate, 15, 3206–3221.