Meteorol Atmos Phys 89, 117–142 (2005)
DOI 10.1007/s00703-005-0125-z
1
2
National Climate Center, China Meteorological Administration, Beijing, China
Department of Physics and Materials Science, City University of Hong Kong, Hong Kong, China
The East Asian summer monsoon: an overview
Ding Yihui1 and Johnny C. L. Chan2
With 17 Figures
Received August 16, 2004; revised October 13, 2004; accepted November 7, 2004
Published online: June 20, 2005 # Springer-Verlag 2005
Summary
The present paper provides an overview of major problems
of the East Asian summer monsoon. The summer monsoon
system over East Asia (including the South China Sea (SCS))
cannot be just thought of as the eastward and northward
extension of the Indian monsoon. Numerous studies have
well documented that the huge Asian summer monsoon
system can be divided into two subsystems: the Indian and
the East Asian monsoon system which are to a greater
extent independent of each other and, at the same time,
interact with each other. In this context, the major findings
made in recent two decades are summarized below: (1) The
earliest onset of the Asian summer monsoon occurs in most
of cases in the central and southern Indochina Peninsula.
The onset is preceded by development of a BOB (Bay
of Bengal) cyclone, the rapid acceleration of low-level
westerlies and significant increase of convective activity in
both areal extent and intensity in the tropical East Indian
Ocean and the Bay of Bengal. (2) The seasonal march of
the East Asian summer monsoon displays a distinct
stepwise northward and northeastward advance, with two
abrupt northward jumps and three stationary periods. The
monsoon rain commences over the region from the
Indochina Peninsula-the SCS-Philippines during the period
from early May to mid-May, then it extends abruptly to the
Yangtze River Basin, and western and southern Japan, and
the southwestern Philippine Sea in early to mid-June and
finally penetrates to North China, Korea and part of Japan,
and the topical western West Pacific. (3) After the onset of
the Asian summer monsoon, the moisture transport coming
from Indochina Peninsula and the South China Sea plays a
crucial ‘‘switch’’ role in moisture supply for precipitation
in East Asia, thus leading to a dramatic change in climate
regime in East Asia and even more remote areas through
teleconnection. (4) The East Asian summer monsoon and
related seasonal rain belts assumes significant variability
at intraseasonal, interannual and interdecadal time scales.
Their interaction, i.e., phase locking and in-phase or outphase superimposing, can to a greater extent control the
behaviors of the East Asian summer monsoon and produce
unique rythem and singularities. (5) Two external forcing
i.e., Pacific and Indian Ocean SSTs and the snow cover in
the Eurasia and the Tibetan Plateau, are believed to be primary contributing factors to the activity of the East Asian
summer monsoon. However, the internal variability of the
atmospheric circulation is also very important. In particular,
the blocking highs in mid-and high latitudes of Eurasian
continents and the subtropical high over the western North
Pacific play a more important role which is quite different
from the condition for the South Asian monsoon. The later
is of tropical monsoon nature while the former is of hybrid
nature of tropical and subtropical monsoon with intense
impact from mid-and high latitudes.
1. Introduction
Based on studies mainly by Chinese meteorologists over many years, it has been found that many
differences exist between the monsoon circulation over India and that over East Asia. This fact
suggests that the structure and main components
of the monsoon system over East Asia is likely to
be independent of the Indian monsoon system,
even though there exist some significant interactions. In other words, the huge Asian monsoon
system can be divided into two subsystems,
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D. Yihui and J. C. L. Chan
the South Asian (or Indian) and the East Asian
monsoon systems, which are independent of each
other and, at the same time, interact with each
other (Zhu, 1934; Yeh et al, 1957; Tao and Chen,
1987). Thus, the summer monsoon over EastAsia (including the South China Sea) cannot be
just thought of as the eastward extension of the
Indian monsoon, on the one hand, and, on the
other hand, the summer monsoon over the mainland of China cannot fully be taken to be the
northward extension of the Indian monsoon. One
must take into account their own unique regional
characters. But Zhu et al (1986) emphasized the
interaction between them. They pointed out that
this interaction may be accomplished through energy exchange, the propagation of low-frequency
oscillation, and moisture transport. The recent
work made by Wang and Lin (2002) has lent a
confirmative support to the existence of the East
Asian monsoon system and further extends the
Asian summer system to incorporate the western
North Pacific region (the Asian-Pacific monsoon).
Thus, the Asian-Pacific monsoon is demarcated
into three sub-systems: the Indian summer monsoon (ISM), the western North Pacific summer
monsoon (WNPSM) and the East Asian summer
monsoon (EASM) (Fig. 1). The EASM domain
defined by them includes the region of 20 –45 N
and 110 –140 E, covering eastern China, Korea,
Japan and the adjacent marginal seas. This definition does not fully agree with the conventional
notion used by Chinese meteorologists (Tao and
Chen, 1987; Ding, 1994), who usually includes
the South China Sea (SCS) in the EASM.
Wang and Lin (2002) believe that the ISM and
WNPSM are tropical monsoons in which the low
level winds reverse from winter easterlies to
summer westerlies, whereas the EASM is a subtropical monsoon in which the low-level winds
reverse primarily from winter northerlies to
southerlies. However, if the SCS region is included in the EASM, the EASM should be a
hybrid type of tropical and subtropical monsoon.
In Fig. 1, one can also note that the ISM and
WNPSM are separated by a broad transitional
zone over Indochina Peninsula and Yun-Gui
plateau. This discontinuity provides a broad
‘‘buffer’’ zone or corridor between the ISM and
WNPSM. Over Indochina Peninsula, the rainy
season sets in late April or early May, reaches
its maximum in intensity in autumn and has double peaks occurring in May and September–
October (Matsumoto, 1997; Lau and Yang, 1997),
respectively, a characteristic that differs substantially from the rainy seasons in the adjacent ISM
and WNPSM.
The Asian monsoon region assumes the most
distinct variation of the annual cycle and the
alternation of dry and wet seasons which is in
concert with the seasonal reversal of the monsoon circulation features (Webster et al, 1998).
However, for different parts of the Asian monsoon region, the durations of dry and wet seasons
may be different, depending on their climate regions and the degree of effects of the Asian monsoon. In South Asia, the dry and wet seasons are
very well-defined while in East Asia four seasons
can be evidently perceived, although the dry and
wet seasons are main modes of annual march of
the precipitation in this region. In mid-latitude
regions of East Asia such as the central China
along the Yangtze and Huaihe River Basins and
Fig. 1. This map divides the Asian-Pacific monsoon into three subregions. The ISM and western WNPSM (see the text) are
tropical monsoon regions. A broad corridor in the Indochina Peninsula separates them. The subtropical monsoon region is
identified as the EASM. It shares a narrow borderline with the WNPSM. (Wang and Lin, 2002)
The East Asian summer monsoon
Korean Peninsula, they are not generally included in the dry and wet alternative regions because
the wet period for theses regions is shorter than
one month. These short rainy periods mainly
occur during prevalence of the summer monsoon
in these regions.
For many years, a large amount of literatures
has been contributed to the study of the Indian
summer monsoon. However, during recent two
decades, a more and more attention has been
paid to study the East Asian summer monsoon.
Recent studies on the East Asian summer monsoon have been devoted to the following aspects
(Ding, 1994; Lau et al, 2001; Ding et al, 2004;
Chang, 2004): (1) the onset of the East Asian
summer monsoon, especially in the South China
Sea (SCS); (2) the seasonal march of the East
Asian summer monsoon and associated major
seasonal rain belts; (3) the Meiyu=Baiu and associated weather disturbances; (4) multiple-scale
variability of the East Asian summer monsoon
and their effects on anomalous climate events
(droughts=floods), especially intraseaonal (ISO),
interannual (e.g., ENSO-monsoon relationship)
and interdecadal-scale variability; (5) the remote
effect of the East Asian summer monsoon through
teleconnection patterns; (6) the physical processes and mechanisms related to the East Asian
summer monsoon, and (7) the predictability and
prediction of the East Asian summer monsoon.
One major thrust for these studies is the South
China Sea Monsoon Experiment (SCSMEX,
1996–2001) and the GEWEX Asian Monsoon
Experiment (GAME, 1995-present) (Lau et al,
2001; Yasunari, 2000; Ding and Liu, 2001; Ding
et al, 2001; Ding et al, 2004). The present paper
will make a comprehensive review to highlight
major achievements and findings concerning the
above-described problems, except for the item (7).
2. Onset of the East Asian summer monsoon
The onset of the Asian summer monsoon is a key
indicator characterzing the abrupt transition from
the dry season to the rainy season and subsequent
seasonal march. Numerous investigators have
studied this problem from the regional perspectives. It is to some extent difficult to obtain a
unified and consistent picture of the climatological onset dates of the Asian summer monsoon
in different regions due to differences in data,
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monsoon indices and definitions of monsoon onset
used in these investigations. Ding (2004) has summarised the climatological dates of the onset of
the Asian summer monsoon in different monsoon
regions based on various sources, with dividing
the whole onset process into four stages: (1)
Stage 1 (late in April or early in May): the earliest onset in the continental Asia is often observed in the central Indochina Peninsula late in
April and early in May, but in some cases, the
onset may first begin in the southern part or
the western part of the Indochina Peninsula. (2)
Stage 2 (from mid to late May): this stage is
characterized by the areal extending of the summer monsoon, advancing northward up to the
Bay of Bengal and eastward down to the SCS.
(3) Stage 3 (from the first dekad to second dekad
of June): this stage is well known for the onset of
the Indian summer monsoon and the arrival of
the East Asian rainy season such as the Meiyu
over the Yangtze River Basin and the Baiu season in Japan. (4) Stage 4 (the first or second
dekad of July): the summer monsoon at this
stage can advance up to North China, the Korean
Peninsula (so-called Changma rainy season) and
even Central Japan.
Figure 2 presents an illustrative description of
this onset process (Zhang et al, 2004). During the
first pentad of May (Fig. 2a), the summer monsoon
is established only over Sumatra. In the next two
pentads (Fig. 2b, c), the tropical monsoon advances
up to the land bridge, first establishing itself
over the southwestern Indochina Peninsula and
then expanding to the entire southern peninsula.
During the pentad of May 16–20 (Fig. 2d),
the build-up of the summer monsoon is observed
over the central Indochina Peninsula. At the same
time, the onset location extends into the central
and southern SCS, accompanied by a rainfall rate
of >5 mm day 1 over the entire SCS. In the next
pentad, onset expands quickly and almost covers
the entire SCS (Fig. 2f ). On the other hand, the
Asian summer monsoon also advances northwestward to the Indian monsoon region from
the near-equatorial East Indian Ocean and the
Indochina Peninsula starting from mid-May
(Fig. 2d). Earliest onset of the Asian summer monsoon in this region may be observed over the
southern tip of the Indian subcontinent. In early
June (Fig. 2g, f ), the Asian summer monsoon
rapidly advances northwestward, arriving in the
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D. Yihui and J. C. L. Chan
Fig. 2. Climatological pentad-averaged precipitation rates (mm day 1 ) for the period from May 1–5 (a) to June 6–11 (h) in
sequence. Light and dark shadings indicate precipitation regions greater than 5 mm day 1 and 10 mm day 1, respectively. The
black dots represent the location of onset of the summer monsoon (Zhang et al, 2004)
The East Asian summer monsoon
central Indian subcontinent. Meanwhile, the onset
over the Arabian Sea and the western coast of the
Indian subcontinents is observed, due mainly to
the enhancement of the cross-equatorial airflow
off the Somali coast and the development of the
onset vortex in the central and northern Arabian
Sea (Krishnamurti et al, 1981; Ding, 1981). This
date is generally believed to be normal onset
dates for the Indian summer monsoon. So, the
onset of the East-and Southeast Asian summer
monsoon and the South Asian summer monsoon
is closely interrelated in the context of the Asian
summer monsoon system. However, the earliest
onset of the Asian summer monsoon occurs over
the Indochina Peninsula and the SCS.
The onset process over the SCS and the
Indochina Peninsula is very abrupt, with dramatic changes of large-scale circulation and rainfall
patterns occurring during a quite short time
period of about one week. After this sudden
onset, low-level easterlies and upper-level westerlies rapidly switch to westerlies and easterlies,
respectively. At the same time, the dry season
which lasts for the cold season rapidly changes
into the wet season, indicating the earliest arrival
of the summer monsoon rainy season in the
Asian–western North Pacific monsoon region.
This sudden change in rainfall is clearly illustrated in Fig. 3. Over the SCS, the major precipitation belt is steadily located in the zonal
band of 15 S–10 N before mid-May. Another
rain belt located in South China (20 –28 N) cor-
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responds to the pre-summer rainy season there.
Around mid-May the near-equatorial rain belt
suddenly moves northward and merges with the
South China rain belt. It can be seen that this
process is accomplished in a quite short time period (Fig. 3b). In contrast, over the Indian longitudes (Fig. 3a) this onset process is more or
less gradual, although a large increase in rainfall
amount in this region may be noted. This suddenness of the onset process in the SCS has been
well documented by numerous investigators with
both climatological and case studies, based on the
large-scale wind, geopotential height, rainfall and
OLR patterns (Lau and Yang, 1997; Matsumoto,
1997; Fong and Wang, 2001; Wang and Lin,
2002). From Figs. 4–5, it can be seen that a
dramatic change clearly occurs from the pentad
of May 11–15 to the pentad of May 16–20 for
these fields. The southwesterlies rapidly expand
from the equatorial East Indian Ocean region,
across the Indochina Peninsula, down to most
of the South China Sea (Fig. 4a–d). At the same
time, the OLR values significantly decrease from
240 W m 2 to values below 240 W m 2 during
this short trainsition period (Fig. 5), implying
that convective clouds and precipitation abruptly
develop over the SCS during the onset process,
heralding the end of the dry season and the arrival of the wet season in this region. The most
significant change of the low-level wind pattern
between prio-and post-onset is the acceleration
and eastward extension of tropical westerlies
Fig. 3. Latitude-time cross-sections of mean precipitation (1979–2001) along 70–80 E (a) and 110 –120 E (b). The CMAP
precipitation dataset is used here. Unit: mm day 1 (Sun, 2002)
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D. Yihui and J. C. L. Chan
Fig. 4. 21-yr (1979–1999) mean 850 hPa wind patterns:
(a) for the pentad of May 6–10, (b) for the pentad of
May 11–15, (c) for the pentad of May 16–20, (d) for the
pentad of May 21–25, and (e) the difference of mean
850 hPa wind patterns between May 21–25 and May 6–10.
Unit: m s 1 . Shading areas denote regions with wind speed
greater than 8 m s 1 (Ding and Sun, 2001)
from the tropical East Indian Ocean to the central
and southern SCS (Fig. 4e). The Somali jet upstream also undergoes a considerable intensification. From the northern part of the Bay of
Bengal to the northern SCS, a wind shear line
with two cyclonic circulations embedded is generated. This fact indicates the development of the
monsoon trough which is connected with the tailing part of mid-latitude frontal systems in the
northern SCS. Therefore, the onset of the summer monsoon in the SCS should to be considered
as a regional demonstration of the rapid seasonal
intensification of the whole Asian summer
monsoon. Correspondingly, the most significant
change in the OLR pattern is also seen in the
Arabian Sea, the tropical East Indian
Ocean and the Bay of Bengal, and the SCS
and the tropical West Pacific (Fig. 5e). These
changes reflect abrupt enhancement of cloud
and rainfall in these regions. Among them, the
change in the SCS is most marked. Another sudden change is the rapid weakening and eastward
retreat of the subtropical high over the West
Pacific from the Indochina Peninsula and the
SCS (figure not shown). At the same time, a
trough over the Bay of Bengal continuously
The East Asian summer monsoon
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Fig. 5. Same as Fig. 5, but for OLR patterns. Unit: W m 2 .
The areas with OLR magnitudes less than 230 W m 2 are
shaded (Ding and Sun, 2001)
extends southward and deepens, which greatly
favors local development of intensive convective activity as well as the acceleration and
eastward propagation of low-level westerlies in
the tropical East Indian Ocean. Now it is not
clear which one, eastward extension of lowlevel southwesterlies or the eastward retreat of
the subtropical high, is the primary cause for
leading to large-scale abrupt changes in the
above chain of events.
The most salient feature of the 200 hPa wind
patterns is the significant development and northward movement of the South Asian high over the
eastern part of the Indochina Peninsula. Before
the onset of the SCS summer monsoon, the South
Asian high is located in the southern part of the
Indochina Peninsula, and has a weaker intensity
(figure not shown). Thereafter, this high moves
toward the northwest and significantly intensifies.
The upper-level westerly jet and the easterly jet
on either flank of the high correspondinly accelerates, thus leading to intensification of upper
level divergence and convective activity in the
Indochina Peninsula and the SCS (Zhang et al,
2004). From the heating pattern during this
period, it can be known that this major outflow
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D. Yihui and J. C. L. Chan
region corresponds to an extensive area of the
heat source (Q1>0) in these regions.
Based on the above analysis, the chain of significant events during the onset of the SCS summer monsoon may be identified below:
– the development of a cross-equatorial current
in the equatorial East Indian Ocean (80 –
90 E) and off the Somali coast and the rapid
seasonal enhancement of heat sources over
the Indochina Peninsula, South China, Tibetan
Plateau, and neighboring areas;
– the acceleration of low-level westerly wind
in the tropical eastern Indian Ocean;
– the development of a monsoon depression or
cyclonic circulation and the breaking of the
continuous subtropical high belt around the
Bay of Bengal;
– the eastward expansion of tropical southwest
monsoon from the tropical East Indian Ocean;
– the arrival of the rainy season in the regions of
Bay of Bengal and Indochina Peninsula with
involvement of impacts from mid-latitudes;
– further eastward expansion of the southwesterly monsoon into the SCS region;
– the significant weakening and eastward retreat
of the main body of the subtropical high, and
eventual onset of the SCS summer monsoon
with convective clouds, rainfall, low-level
southwesterly wind and upper-level northeasterly wind suddenly developing in this region.
The case of the onset the SCS summer monsoon in 1998 has been extensively studied, because a complete dataset acquired during the
SCSMEX field phase (May–August) is available
(Ding and Li, 1999; Lau et al, 2001; Johnson and
Ciesielski, 2002; Ding et al, 2004). The onset
process in this year is in many ways similar to
climatological conditions illustrated above, but
with the earliest onset occurring over Indochina
Fig. 6. Vertically integrated (from surface to
300 hPa) moisture budgets averaged for 1990–
1999 for various monsoon regions prior to the
onset (the 1st pentad of April–the 2nd of May)
(a) and after the onset of the SCS summer monsoon (June–August) (b). Unit: 106 Kg s 1 (Ding
and Sun, 2002)
The East Asian summer monsoon
Peninsula and the northern part of the SCS concurrently. Intense cold air activity coming from
mid-latitudes induced extensive area of vigorous
meso-scale convective systems (MCSs) in the
northern SCS (Ding and Liu, 2001; Johnson and
Ciesielski, 2002). The monsoon trough or a stationary tailing part of the cold front was intensified through the feedback effect of atmospheric
convective heating caused by the subsequent development of MCSs in the trough. Thus, the
tropical southwest monsoon to the south of the
monsoon trough rapidly intensified and propagated northward, leading to the onset of the summer monsoon in this region.
The onset of the Asian summer monsoon, as a
kind of switch, plays a crucial role in heat and
moisture transport and hydrological cycle. From
Fig. 6, it can be seen that before the onset of
the SCS summer monsoon, the interhemispheric
moisture transport is rather weak and even southward. The northward moisture transport across
the northern boundaries of various regions is
generally weak, except for the regions of the
Indochina Peninsula and the SCS. The moisture
sinks occur in the regions of Bay of Bengal, the
Indochina Peninsula and South China, where
the enhanced precipitation may be observed.
After the onset the whole picture of the moisture
transport and budget rapidly changes and becomes well-organized. The cross-equatorial flow
has its maximum moisture transport in the western part of the equatorial Indian Ocean. The
second maximum moisture transport is located
in the equatorial East Indian Ocean. In the South
Asian and Southeast Asian monsoon regions, one
may see consistent eastward moisture transport,
all the way to the SCS. The moisture sinks from
the Indian Peninsula to the SCS are consistent
with the major observed precipitation regions,
with the Bay of Bengal having the maximum.
The northward moisture transport through the
northern boundaries has its maximum in the region of the Bay of Bengal. The SCS takes the
second place. But, if one combines together the
moisture transport coming from the Indochina
Peninsula and the SCS, the northern moisture
transport into the East Asian region will obviously exceeds the northward transport through the
Bay of Bengal. This fact implies the critical role
of the moisture transport from the SCS in the
precipitation in East Asia.
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3. Seasonal march of the East Asian
summer monsoon and major seasonal
rain belts
The seasonal advance and retreat of the summer
monsoon in East Asia behaves in a stepwise way,
not in continuous way. When the summer monsoon advances northward, it undergoes three
standing stages and two stages of abrupt northward shifts. In this process, as does the monsoonal airflow, the monsoon rain belt and its
associated monsoon air mass also demonstrate
a similar northward movement. These stepwise
northward jumps are closely related to seasonal
changes in the general circulation in East Asia,
mainly the seasonal evolution of the planetary
frontal zone, the westerly upper-level jet stream
and the subtropical high over the West Pacific.
Recently, Wu and Wang (2001), and Wang and
Lin (2002) have studied the large-scale onset,
peak and withdrawal of the Asian monsoon
rainy season, and have identified two phases in
the evolution process. The first phase begins with
the rainfall surges over the South China Sea
in mid-May, which establishes a planetary-scale
monsoon rainband extending from the South
Asian marginal seas (the Arabian Sea, the Bay
of Bengal, and the SCS) to the subtropical western North Pacific (WNP). The second phase of
the Asian monsoon onset is characterized by the
synchronized initiation of the Indian rainy season
Fig. 7. Latitude-time section of 5-day mean rainfall over
eastern China (110 –120 E) from April to September averaged for 1961–1990. Regions of heavy rainfall (> 50 mm)
are shaded. Unit: mm. (Ding and Sun, 2002)
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D. Yihui and J. C. L. Chan
and the Meyu=Baiu in early to mid-June. The
peak rainy seasons tend to occur primarily in three
stepwise phases, in late June over the Meiyu=
Baiu regions, the northern Bay of Bengal and
the vicinity of the Philippines; in late July over
India and northern China; and in mid-August over
the tropical WNP. The first two stepwise jumps
occurs in the East Asian region.
Based on the time-latitude cross-section of
5-day rainfall amount for eastern China (Fig. 7)
(Sun, 2002), the most conspicuous feature is the
monsoon onset between 18 and 25 N as indicated by the steep rise in precipitation starting
from the first 10-day period of May. This rainy
episode is so-called pre-summer rainy season in
South China, Hong Kong and Taiwan (e.g., Lau
et al, 1988). The first standing stage of the major
rain belt generally continues into the first 10-day
period of June, and afterwards it rapidly shifts to
the valley of the Yangtze River. This second stationary phase initiates the Meiyu rainy season in
central China. The time span of the season on the
average lasts for 20–30 days (12th June-8th July).
The wind and thermal fields in the Meiyu region are usually characterized by a low-pressure
trough (the so-called the East Asian summer
monsoon trough), a weak stationary front at
surface, significant horizontal wind shear across
the front and frequent occurrence of prolonged
heavy rainfall. The heaviest rainfall is mostly associated with eastward-moving meso- to synoptic scale disturbances along the front. The
Meiyu=Baiu and associated disturbances will be
discussed in more details in the next section.
The Baiu in Japan and Changma in Korea also
occur in a similar situation, but with a regional
difference in locations, timing and duration. As
indicated by Ninomiya and Muraki (1986), the
Baiu in Japan begins in early June when rainfall
in Okinawa reaches its peak. In the last ten days
of June, the rainfall peak moves to the western
and southern parts of Japan. Then the rainfall
peak further moves northward in the first ten
days of July. North of 40 N, no rainfall peaks
associated with the Baiu can be observed. So,
the Baiu season in Japan mainly lasts from early
June to mid-July, almost concurrently with the
occurrence of the Meiyu in China. The rainy season in Korea, the so-called Changma, accompanied with a belt-like peak rainfall zone, begins
with the influence of the quasi-stationary con-
vergence zone between the tropical maritime airmass from the south, and both continental and
maritime polar airmasses from the north (Oh
et al, 1997). Based on the precipitation peak and
lower tropospheric circulation features, the onset
date of the Northeast Asia summer monsoon or
Changma rainy season can be determined as the
period of the 37th to 39th pentad (late June–midJuly), with a significant interannual variability
(Qian and Lee, 2000). Therefore, the Changma
is a shorter monsoonal rainy season, with mean
period being 20 days long.
From mid-July, the rain belt rapidly jumps
over North China and Northeast China again, the
northernmost position of summer monsoon rainfall. This standing stage of the rain belt causes
the rainy season in the North China that generally
lasts for one month. In the early or middle part of
August the rainy season of North China comes to
end, with the major monsoon rain belt disappearing. From the end of August to early September
the monsoon rain belt quite rapidly moves back
to South China again. At this time, most of the
eastern part of China is dominated by a dry spell.
The East Asian summer monsoon assumes a
marked active-break cycle. As indicated above,
the active periods corresponds to major monsoon
rainy seasons such as the presummer rainy season in South China, and Meiyu=Baiu rainy season in the Yangtze River Basin and Japan during
May–mid-July. Afterwards, a break of the monsoonal rainy period occurs from late July to early
August in Japan (Chen et al, 2003). This break
of different spans is also observed in South
China, central China, Northeast China, Taiwan,
and Korea, but with different occurrence time.
From mid-July, the second rainy season or the revival of the rainy period (Chen et al, 2003) predominates over South China, with a gap of a time
period of about 20 days or one month between
the pre-summer rainy season and this rainy season that is mainly caused by typhoons, the movement of the ITCZ and other tropical disturbances
in the monsoonal airflow.
For other regions, after the break spell, monsoon rain resumes for a period from August
to September–October. Therefore, the monsoon
rainfall variation during the warm season in East
Asia is generally characterized by two active
rainfall periods separated by a break spell. It is
clearly seen from Fig. 8 that the Meiyu rain band,
The East Asian summer monsoon
127
Fig. 8. Latitudinal-time cross-sections of CMAP rainfall averaged over longitudinal zones of (a) 120 –125 E, (b) 125 –
130 E, and (c) 130 –140 E, and rainfall histograms of three regions: (d) Taiwan (120 –125 E, 20–25 N), (e) Korea (125 –
130 E, 35 –40 N), and (f) Japan (130 –140 E, 32.5 –40 N). Different phases of summer monsoons in three regions are
indicated by active, break and revival. The contour interval of CMAP rainfall in (a)–(c) is 1 mm day 1, while rainfall amounts
larger than 5 mm day 1 are stippled by different colors indicated by the scale shown in the lower left corner of the three upper
panels. (Chen et al, 2003)
forming in early May, progresses northward until
the end of July, and diminishes between 40 and
45 N in Northeast China and Korea, and about
40 N in Japan. The passage of the Meiyu rain
band is followed by a break spell (monsoon break)
which also propagates northward. Then, the
monsoon rainfall revival after the break is clearly
observed. Chen et al (2003) has shown that the
monsoon revival in East Asia is caused by a different mechanism associated with the development of other monsoon circulation components
including the ITCZ and weather systems in midlatitudes. The Changma break in late July is very
short, with the duration of a half month. Starting
from late August, the revival of the monsoon
rainy period is also observed in Fig. 8. The second rain spell is not long based on the study by
Chen et al (2003). But, Qian et al (2002) pointed
out that this precipitation surge can maintain
until early September, forming the autumn rainy
season in Korea.
4. The Meiyu=Baiu and associated
weather disturbances
Meiyu=Baiu is a unique rainy season in the seasonal march of the East Asian summer monsoon.
It starts nearly concurrently with the onset of
the East Asian summer monsoon onset in the
South China Sea. Then, as the summer monsoon propagates northward, the Meiyu rain belt
sequentially establishes itself in South China and
Taiwan, the Yangtze and Huaihe River Basins
and Japan, and the Korean Peninsula. As pointed
out by Chen (2004), the different terminology
has been used for this major seasonal rain belt
128
D. Yihui and J. C. L. Chan
Fig. 9. Annual mean (1975–1986) frequency distribution of 850 hPa fronts in (a) southern China and Taiwan Meiyu season
(15 May–15 June), and (b) Yangtze River Valley Meiyu season (16 June–15 July). Front frequency is counted at 12 h intervals
and analyzed at 1 lat 1 long grid intervals. Heavy dashed line indicates maximum axis (from Chen, 1988)
in different regions. In China, the term ‘‘Meiyu’’
is used for the rainy season from mid-June
to mid-July over the Yangtze River Valley (Tao
and Chen, 1987). In Japan, the term ‘‘Baiu’’ is
used both for the rainy season over Okinawa
region from early May to mid-June and over
the Japanese Main Islands from mid-June to
mid-July (Saito, 1985). In Taiwan, on the other
hand, the term ‘‘Meiyu’’ is used both for the
rainy season over Taiwan and over South China
from mid-May to mid-June (Chen, 1983; 1988;
Wang, 1970). Therefore, the ‘‘Meiyu’’ season over
South China and Taiwan discussed in this paper
corresponds to the ‘‘South China pre-summer
rainy period’’ used by many Chinese meteorologists (Tao and Chen, 1987; Ding, 1992), and
the ‘‘pre-Meiyu’’ period used by Chang et al
(2000 a, b).
Figure 9 presents the annual mean frequency
distribution of 850 hPa fronts in the Meiyu season of South China and Taiwan (mid-May to
mid-June) and of the Yangtze River Valley (midJune to mid-July) (Chen, 1988). For the former
case, the axis of maximum frequency, indicating the mean position of the Meiyu front, is
oriented approximately in an east–west direction extending from southern Japan to southern
China. The mean position shifts northward to
Japan and central China in the Meiyu season of
the Yangtze River Valley. The Meiyu front often
moves southeastward slowly in the early stage
of its lifetime and appears as a quasi-stationary
front in the late stage with an average lifetime of
8 days.
Although Meiyu in China and Baiu in Japan
both occur in the early summer rainy season in
East Asia, their structure and dynamics are not
fully same, due to different locations of the planetary frontal zone. As indicated by Chen and
Chang (1980), the structure of the eastern (near
Japan) and central (the East China Sea) resembles a typical midlatitude baroclinic front with
strong vertical filting toward a upper level cold
core and a strong horizontal temperature, whereas
the western (Southern China and the Yangtze
River Basin) section resembles a semitropical disturbance with an equivalent borotropic warm core
structure (Ding, 1992), a weak temperature gradient, and a rather strong horizontal wind shear
in the lower troposphere. Figure 10 clearly illustrates the synoptic conditions where the Baiu
in Japan and Meiyu in China form (Ninomiya,
2004). In this conceptual model the Meiyu=Baiu
Baiu cloud zone consists of a few cloud system
families, each of which consists of two parts:
a sub-synoptic scale cloud system associated with
a sub-synoptic-scale Meiyu=Baiu frontal depression (indicated by S), and a few meso--scale
cloud systems (indicated by ). The latter are
aligned along the trailing portion of the preceding sub-synoptic-scale cloud system. Cold lows
and a midlatitude blocking ridge and the Pacific
The East Asian summer monsoon
129
Fig. 10. Conceptual model of
the Meiyu-Baiu frontal cloud
zone (Ninomiya, 2004)
subtropical anticyclone all have strong influences
on the Meiyu=Baiu cloud systems, but with a
stronger effect of cold lows on Baiu (eastern
section). The subtropical and tropical monsoon
airflows have a more significant influence on
Meiyu in China. Rows of large and small arrows
in Fig. 10 indicate the 500-hPa and 850-hPa maximum wind axes, respectively. The short-wave
trough that propagates along the northern maximum wind zone becomes coupled with the shortwave trough in the Meiyu=Baiu frontal zone
under the influence of the cold low over Siberia,
Fig. 11. Climatology of the Meiyu composited for Meiyu periods based on 30-yr NCEP datasets and 740 station data in
China: (a) total rainfall amount (Unit: mm), (b) the se field at 850 hPa (Unit: K), (c) 850 hPa temperature fields (Unit: K), and
(d) the moisture transport at 850 hPa (Unit: kg(ms) 1 ). The maximum transport zone is shaded, (Ding and Liu, 2003)
130
D. Yihui and J. C. L. Chan
leading to the development of a sub-synoptic-scale
frontal depression. Subsequently, a few meso-scale cloud clusters form along the trailing portion
of the preceding sub-synoptic scale cloud system.
Figure 11 show the climatological aspects of
Meiyu over the Yangtze and Huaihe River Basins
based on the 30-yr (1971–2000) NECP datasets
and 740 surface station data in China (Ding and
Liu, 2003). It can be seen that Meiyu rainfalls
are mainly distributed over the middle and lower
valley of the Yangtze River, with the latter having the maximum rainfall amount (260 mm),
accounting for 45% of total rainfall amount for
summer (June, July and August) (Fig. 11a).
Therefore, nearly half of summer rainfalls comes
from the Meiyu season that on the average lasts
for about 25 days (from June 12 to July 8). In the
Meiyu zone, the air is very moist, with a high
specific humidity belt at low-level along the
Meiyu zone observed. Overall, the Meiyu zone
is characterized by a high se region (Fig. 11b).
An interesting feature of the low-level temperature field is its sandwich pattern, with the warmer
air to south and the north, respectively, and relatively colder air in between (Fig. 11c). This cooling in the Meiyu zone is also noted by Kato
(1987). Three reasons may be used to illustrate
the colder temperature zone along the Meiyu
precipitation region: (1) intrusion of low-level
cold air from northeast accompanied by the lowlevel northeasterlies to north of the Meiyu zone;
(2) cooling effect of precipitation evaporation
at low-level and near the surface; and (3) the
intense airmass modification over North and
Northwest China through the surface sensible
heating (Kato, 1987). This reverses meridional
thermal contrast between the Meiyu zone and the
region to its north. From the view point of wind
fields, to the south of the Meiyu zone, there are
extensive southwest and southeast monsoon at
850 hPa that merge together in the Meiyu and
Baiu zones. The strong low-level jet (LLJ) and
its vertical coupling with the upper level jet may
be observed (Chen, 2004), and the Meiyu precipitation zone is located in between. Major Meiyu
rainfalls generally occurs in the right quadrant of
entrance sector of upper-level jet which is dominated by upward motion (Cressman, 1981). The
positive vorticity to the left side of the LLJ is
also favorable for occurrence of rainfalls. A
large amount of moisture is transported into the
Meiyu=Baiu zone by the summer monsoon. The
South China Sea is a major moisture channel for
the Meiyu precipitation (Fig. 11d). Significant
moisture convergence is observed in the middle
and lower valleys of the Yangtze River and the
western Japan where the Meiyu and Baiu precipitation is highly concentrated.
Chen and Chang (1980) studied dynamics of
the Meiyu front. The vorticity budget calculated by them showed that generation of cyclonic
vorticity by horizontal convergence was counteracted by cumulus damping in the eastern section
and by boundary layer friction in the mountainous
Fig. 12. Climatologically averaged (1971–2000) Meiyu frontal
structure along 117.5 E. Solid
lines are se isolines (Unit: K)
and dashed lines are isolines of
specific humidity (Unit: g kg 1 ).
Horizontal bar at the bottom represents the averaged latitudinal
range of precipitation greater
than 200 mm (27–30 N) (Ding
and Liu, 2003)
The East Asian summer monsoon
western section. Results from theoretical, modeling and observational studies suggest that
the Meiyu frontogenetic process is initiated and
maintained by the CISK mechanism through the
interaction between the potential vorticity (PV)
anomaly and the convective latent heating (Chen
et al, 1998; Chen, 2003). The Meiyu front affecting South China and Taiwan forms in the subtropical latitude, which is a distinct area from that
for the formation of polar front in the Meiyu
season. It resembles a semitropical disturbance
with an equivalent barotropic warm core structure, a weak horizontal temperature gradient, a
rather strong horizontal wind shear, and a positive low-level PV anomaly (Chen, 2004).
Figure 12 is the mean structure of the Meiyu
front averaged for 1971–2000. An interesting
feature is the highly moist air column ahead of
the Meiyu front which very much resembles the
eye-wall region of a typical tropical cyclone. The
Meiyu rainfall intensively occurs in this region.
This implies the significant importance of convective precipitation and associated latent heat
release. Generally, the frontal structure at lowlevel or near the surface disappears or even
changes its sloping from northward tilting to
southward tilting. So, Xie (1956) previously defined the low-level part of the Meiyu front as the
equatorial front, with the relatively cold air in the
south of the Meiyu front and relatively warm air
in the north. Corresponding to the Meiyu front
shown in Fig. 12, the mean cross-front vertical
circulation is characterized by strong upward
motion throughout the entire troposphere located
in the region of Meiyu rainfalls, the southerly
component at low-level and the northerly component at upper-level in the region to the south of
the Meiyu front. Therefore, a so-called monsoon
circulation cell (anti-Hadley cell) is clearly evident. To the north of the Meiyu front, there is a
thermally direct cell.
From Fig. 12, it can be seen that the MeiyuBaiu frontal zone associated with intense convective precipitation is not characterized by the
strong convective instability, but by nearly moist
neutral stratification. This indicates the release
of the convective instability associated with the
cumulus convection. For the sustenance of the
strong convective precipitation during the Meiyu
period, some large-scale process must generate
convective instability against the stabilizing ef-
131
fect of the convective clouds. The local time
change of convective stability is due to the differential advection of e. Ninomiya (2004) has
indicated that area of negative differential advection (generation of convective instability) are
present over the Meiyu=Baiu frontal zone, which
indicates that the differential advection generates
successively convective instability against the release of the instability by the convective clouds.
As the result of these two processes, the large
precipitation and nearly moist neutral stratification are maintained within the frontal precipitation zone.
The heavy rainfalls during the Meiyu period
are mainly generated by the meso-- and meso-scale disturbances which are embedded within
and propagated along the Meiyu-Baiu cloud and
rain band or frontal zone with horizontal length
scale of several thousand kilometers (Ding, 1992).
Results of a case study of the heavy rain event in
23–25 June 1983 over the Yangtze River Valley
by Ma and Bosart (1987) revealed that a quasistationary frontal boundary, separating very warm
and moist tropical Pacific air from slightly cooler
but still moist air, served to focus the rains in a
relatively narrow latitudinal band. The meso-scale systems during the Meiyu period may
be classified into two types: the Yangtze River
Valley shear line and the low-level vortex. The
Yangtze River Valley (112–120 E, 30–35 N)
shear line is the major synoptic system, which
generates heavy rainfalls in this region (Chen,
2004). There were at least two kinds of low-level
vortices that generated heavy rains during the
Meiyu season. One was the SW (southwest) vortex.
It was generated on the lee-side of the Tibetan
Plateau and tended to be stationary if there
was no upper-level trough to steer it out of the
Sichuan Basin. It could produce heavy rainfalls
locally in Sichuan Basin. Once it is steered out
and moves eastward, it moves along the Meiyu
shear line in most cases and moves northeastward or southeastward in some cases. Another
kind of low vortex is the intermediate-scale cyclone which forms along the Meiyu front with a
horizontal scale of 1000–3000 km (Ninomiya and
Murakami, 1987; Ninomiya, 2001).
In general, the SW vortex is defined as a
700 hPa closed cyclonic circulation over southwestern China, mainly over the western part of
the Sichuan Basin. It is a low-level circulation
132
D. Yihui and J. C. L. Chan
Fig. 13. Distributions of daily geopotential height and
wind vector (unit: ms 1 ) at 850 hPa during the Meiyu period from June 29 to July 1, 1999. C3 denotes a southwest
vortex which brought about a heavy rainfall episode in the
middle and lower Yangtze River basin (Ding et al, 2001)
system, often only visible on 850 and 700 hPa
analyses. On the surface weather map, one may
often observe a negative pressure tendency during 24 hours over the low-vortex region. In this
sense, the SW vortex is also called the SW low
vortex. The SW vortex may provide strong orographic lifting to trigger convection and, consequently, a large amount of rainfalls on the steep
topography surrounding the Sichuan Basin. Many
cases may be exemplified, for example, the heavy
rainfalls in the Sichuan Basin on 1–14 July of
1981 which have been extensively studied by numerous meteorologists (Chen and Dell’Osso,
1984; Kuo, Cheng and Anthes, 1986; Wang and
Orlanski, 1987). Figure 13 is a notable example of
consecutive genesis, development and eastward
movement of a SW vortex in the 1999 Meiyu
season (Ding et al, 2001).
From the synoptic viewpoint, the genesis and
development of the SW vortex needs to meet
two requirements: (1) the existence of a vigorous
southerly airflow from the eastern slope of the
Tibetan Plateau to the Sichuan Basin. It may play
a dual role in the genesis of the SW vortex.
Dynamically, this southerly wind produces
‘‘differential frictional effects’’, a mechanism first
discussed by Newton (1956) in connection with
Colorado cyclone formation, thus leading to the
formation of a cyclonic circulation at low level.
Thermally, the southerly wind may transport
abundant warm, moist air into the eastern slope
of the Plateau and the Sichuan Basin, providing
the major moisture source for precipitation and
the release of latent heat; (2) the necessary triggerning mechanism. Most of the time, the low
pressure troughs passing over the Tibetan Plateau
may act as a triggering mechanism for the SW
vortex. Chang et al (1998) has studied the development of a low-level SW vortex which was
involved in its coupling with two upper-level disturbances. Both disturbance appeared later than
and upstream of the low-level vortex. Faster eastward movements allowed them to catch up with
the low-level vortex and led to a strong vertical
coupling and deep tropopause folding. From
the regional viewpoint, the topography of the
Tibetan Plateau is extremely important.
The development of the SW vortex is expected
to depend greatly on the effect of latent heat
release, due to the fact that this vortex is usually
accompanied by a large amount of rainfall and
convective activity. In order to document better
the effects of strong latent heat release associated
with convection, Kuo et al (1986) calculated mesoscale heat and moisture budget associated with
a SW vortex which resulted in a flood catastrophe
The East Asian summer monsoon
in the Sichuan Basin, on 11–15 July, 1981. With
weak stability at the middle levels, latent heat
release can induce strong, upward vertical motion,
which in turn enhances low-level convergence
spin-up and convective cloud development, establishing a positive feedback between the circulation of the SW vortex and the cumulus
(Chang et al, 2000). Wang et al (1993) further
indicate that the mesoscale vortex in the lee of
the Tibetan Plateau is driven diabatically.
As indicated by Chen (2004), due to the observational spatial data limitations in China, very
little work has been done on meso--scale systems. The Meiyu experiment over the middle and
lower reaches of the Yangtze River (1980–1983)
for the first time provided an opportunity for
studying this system on the horizontal scale of
25–250 km, by using the denser network of the
upper-air and surface observations. The major
findings have been summarized in the monograph by Zhang (1990). It was found that the
meso--scale systems occurred in advance of
the forward tilting minor wave trough which
was located near the Meiyu cloud and rain bands,
on the right side of the upper-level jet, and the
left side of the low-level jet. In general, this system was associated with the mesoscale shear line.
During past ten years, the availability of mesoscale observational data has been considerably
improved due to several Meiyu rainstorms experiment projects carried out in South China, Taiwan
133
and the Yangtze and Huaihe River Basins, such
as HUAMEX, TAMEX, GAME=HUBEX and the Mesoscale Rainstrom Experiment in the Yangtze River
Basin. Some new results have been achieved
in relation to meso-scale disturbances in Meiyu
fronts.
A typical example of Meiyu=Baiu frontal mesoscale disturbances is shown in Fig. 14 (Ninomiya,
2004). The Meiyu-Baiu cloud zone appears as the
chain of cloud systems on the subsynoptic-scale
and mesoscale. The wavelength of the major disturbances in Fig. 14 is estimated to be 2000 km,
which falls on the border between macro-- and
meso--scale. Therefore, these disturbances are
identified as subsynoptic-scale Meiyu=Baiu frontal
disturbances in the present report. Some authors
(Matsumoto and Nimomiya, 1971) classified them
as medium-scale disturbances.
The meso--scale cloud systems are very favorable for occurrence of meso-scale convective
systems (MCSs). The MCSs are often observed to
develop in the region of the meso--scale cloud
systems. By definition, mesoscale convective systems (MCS) are a well organized, meso--scale
(with horizontal resolution of 200–2000 km) convective system which has a nearly elliptic shape
and smooth edge. MCS includes the meso-scale
convective complex (MCC) that has been extensively studied. Activities of the MCSs are quite
frequent in China. They mainly occur in Southwest China, but are often observed in connection
Fig. 14. The longitude-time section of TBB at 32.5 N for 1991
Meiyu=Baiu period The isopleths are at 10 C intervals,
and the minus sign of TBB is
omitted (Ninomiya, 2000; 2004)
134
D. Yihui and J. C. L. Chan
with the major seasonal rain belts such as those
during the presummer rainy season in South
China and Meiyu in the Yangtze-Huaihe River
Basins. During the Baiu season in Japan MCS
are sometimes observed as an important intense
rain-producing system (Ninomiya and Murakami,
1987). The preferred locations of occurrence of
MCS are the northwestern periphery of the subtropical high over the western North Pacific
where the warm and cold air have a frequent
and vigorous interaction. Sometimes, the MCSs
also may be produced in East and South China
due to strong surface heating and local unstable
stratification.
The MCC have been intensively studied in
80’s and early 90’s. In the figure produced by
Miller and Fritsch (1991), the MCCs in China
were only observed in Southwest China which
are associated with the Southwest Vortex. But,
based on studies by Chinese meteorologists, the
gensis regions of MCCs are not only confined in
this region, they may occur over a number of other
regions. In late spring and early summer, MCCs
often occur over the southern part of China (Xiang
and Jiang, 1995) in relation to Meiyu season. Their
mean lifetime is about 18 hours, slightly longer
than that (about 10 hours) in North America.
MCCs generally generate and develop in late
afternoon and early evening, further grow into
MCC at nighttime and disspate in the morning
of the next day. Wu and Chen (1988) studied the
composite structure of environment conditions
for the 12 cases of meso--scale MCS (i.e., MCC)
over South China selected in May–June 1981–
1986 at their formation and mature stages. The
overall structure was quite similar to that for the
midlatitude MCCs in the North America as obtained by Maddox (1983). The MCCs form and
intensify in the warm sector to the south of the
Meiyu front=shear line. The strong warm advection and speed convergence (i.e., convergence due
to the downstream speed decrease) in the lowertropospheric southwesterlies, possible lifting mechanisms at the formation and intensification
stages, prevail over the area of MCCs. The MCCs
tended to form and to intensify on the cyclonic
side of the LLJ exit region. Anticyclonic circulation and diffluent flow in the upper troposphere
provided conditions favorable for the intensification of MCCs. At the genesis and development
stages, the precipitation amount is relatively
small, with severe convective weather dominating.
The heavy rainfalls mainly occur at the mature
stage, with intense rainfall rate of 30–50 mm hr 1.
Therefore, the MCCs are an important rainproducing system in the summer monsoon season
in South China and the Yangtze River Basin.
Finally, the conceptual model of the Meiyu
front in the Yangtze River Basin and South China
is presented (Fig. 15). Ahead of the Meiyu front,
a so-called monsoon vertical circulation is observed, with the upward motion in Meiyu precipitation region and downward motion in the
south. The Meiyu front at low-level evolves into
the so-called equatorial front or nearly disappears. In the Meiyu precipitation zone, the air
Fig. 15. Synoptic model of the
Meiyu season in East China (Liu
et al, 2003)
The East Asian summer monsoon
in the deep troposphere is highly moist, with high
e observed. The LLJ is observed to the south of
the Meiyu front within the lower return branch of
the secondary circulation. It is often vertically
coupled with the upper-level jet stream.
5. Intraseasonal oscillations (ISO)
and teleconnection patterns
During last two decades, a large amount of
research works have been devoted to study the
intraseasonal oscillation of the Asian monsoon.
On the intraseasonal scale, the monsoon fluctuateds mainly on two preferred time scales:
10–20-day and 30–60-day, with the latter often
referred to as the Madden-Julian Oscillation
(MJO). In the South China Sea and the East Asian
summer monsoon regions, the ISO can play three
135
fold roles: the triggering of the onset of the summer monsoon, modulation of active and break cycles of the summer monsoon and rainy seasons
and connection of summer monsoon activity of
the neighbouring regional monsoon systems of
the South Asian, the East Asian and Western
North Pacific. When the ISO can propagate or
fluctuate on an even larger-scale or the hermispheric scale, this remote connection may excite
some kind of atmospheric teleconnection patterns or Rossby wave trains.
Figure 16a is the Morlet wavelet analysis of
850 hPa zonal wind in the SCS region for
May–August of 1998 during the SCSMEX field
experiment (Xu and Zhu, 2002). Two main modes
of 30–60-day and 10–20-day low frequency
oscillations can be identified. Figure 16b has
shown that the phase of the westerly wind of the
Fig. 16a. Morlet wavelet analyses of zonal
wind at 850 hPa in the SCS. Unit: day (Xu
and Zhu, 2002). (b) Observed 850 hPa zonal
wind over the SCS region (5–20 N, 105–
120 E) in 1998 (shaded) and the temporal variations of the 30–60 day low-frequency oscillation (solid line) and corresponding kinetic
energy (dashed line). Unit: ms 1 for wind and
m2 s 2 for kinetic energy (Mu and Li, 2000)
136
D. Yihui and J. C. L. Chan
30–60-day mode occurred concurrently with bursting of the westerly monsoon at 850 hPa in this
region (Mu and Li, 2000). Also based on the data
from the SCSMEX in 1998, Chan et al (2002)
have shown that the onset and maintenance of
1998 SCS summer monsoon were controlled by
the 30–60-day oscillation and further modified
by the 10–20-day mode. Chen and Chen (1995)
previously indicated that the onset of the 1979
SCS summer monsoon occurs under the condition
of a phase-lock between the 30–60-day and the
10–20-day modes over the Northern SCS.
Recently, Mao and Chan (2004) have obtained
a more general conclusion that the 30–60-day
mode and 10–20-day mode oscillations control
the behavior of the SCS summer monsoon activities for most of years. The 30–60-day oscillation of the SCS summer monsoon exhibits a
trough-ridge seesaw over the SCS, with anomalous cyclones (anticyclones) along with enhanced
(suppressed) convection migrating northward. On
the other hand, the 10–20-day oscillation manifests as an anticyclone=cyclone system over the
western tropical Pacific with a largely zonal orientation propagating westward into the SCS.
The arrival of the ISO oscillation is not only to
be a possible triggering mechanism for the sudden onset, but also can play a crucial role in the
stepwise northward advance of the East Asian
summer monsoon and in modulating the regional
rainy seasons. Qian et al (2002) have shown that
the onset of the East Asian summer monsoon
occurs when a wet phase of the climatological
intraseasonal oscillation (ISO) arrives or develops,
and the northward propagating summer monsoon
consists of several phase-locking wet ISO. In the
East Asian summer monsoon region, the seasonal
process of the summer monsoon and the ISO
propagation are both northward and they are
interconnected at all the stages of the seasonal
march and in all the subregions of East Asia.
Wang and Xu (1997) have further identified four
cycles of statistically significant climatological
intraseasonal oscillation (CISO) from May to
October in the Asian summer monsoon regions.
The peak wet phase of these cycles corresponds
to active stage of the summer monsoon while the
dry phase corresponds to the monsoon break. It
should be pointed out that though the climatological ISO is often the primary reason for the sudden onset, the onset is paced by the seasonal
evolution of large-scale circulation and thermodynamics that determines the direction of the
onset advance. With the large-scale background
established by the seasonal evolution, the arrival
of several one-after-another ISO wet phases triggers the development of deep convection. Due to
the seasonal regulation, the ISO has a tendency
to be phase-locked with respect to the calendar
year so that the climatological onset displays
multiple stages. The stepwise march of the onset
is observed each year (Wu and Wang, 2001).
Two teleconnection patterns associated with
the Asian summer monsoon have been revealed.
Nitta (1986), and Huang and Li (1988) indicated
that heating sources caused by convective activity over the SCS and the region around the
Phillipines (over the Warm Pool) may excite a
stationary wave train, thus producing a teleconnection pattern, so-called JP pattern (JapanPacific). The immediated downstream effect of
the propagation of this wave train is exerted upon
the behavior of the subtropical high over the
western Pacific, and especially on its position.
Then, the summer rainfall will be influenced by
the anomalous behavior of the subtropical high.
Huang and Sun (1990) further analyzed the relationship between the conditions of anomalous
summer precipitation in the eastern China and
the temperature in surface and subsurface layers
of the Warm Pool at depths of between 50 and
300 m. Recently, Li and Zhang (1999), and Lau
and Wang (2002) have indicated that the thermal
forcing excited by convective activity and rainfalls in the SCS and western tropical Pacific,
through this teleconnection pattern, may affect
weather and climate not only in China, Korea
and Japan, but also possibly in North America.
Another teleconnection pattern originates from
a large amount of monsoon rainfalls and associated intense heating forcing in India, which
can exert a significant remote effect on the
general circulation on a large-scale basis. Liang
(1988) has found that the summer rainfall between India and North China has a stable and
significant positive correlation relationship, especially with a fairly consisitent occurrence of
droughts and flooding events in these two regions. Meanwhile, Guo and Wang (1988) used
a longer set of data (1951–1980) for 110 stations
in China and 31 subregions in the Indian
Peninsula to further study this problem, and have
The East Asian summer monsoon
justified the above relationship indicating that
the most significant correlative region with a
significance level of 0.95 is North China which
demonstrates a positive correlation, with their
correlation coefficient being 0.65 (the confidence
level exceeding 99.9%). In recent years, a number of investigators have paid attention to this
teleconnection patterns and have well documented its existence with significant statistical relationship and physical explanation (Hu and Nitta,
1996; Kripalani and Kulakarni, 1997; 2001). In
addition, a negative correlation between summer
rainfall variations in India and southern Japan is
further found, which reflects downward propagation of a wave-type circulation pattern over
mid-latitude Asia.
6. Physical processes and mechanisms
related to the onset and the seasonal march
of the East Asian summer monsoon
In the Asian monsoon region, the thermal contrast due to differential heating between land and
sea in the process of seasonal march of solar
radiation acts as a seasonal precondition for the
onset. However, the Asian monsoon is not only
forced by the thermal effect of land-sea contrast,
but also by the elevated heat source produced by
137
the huge massif of the Tibetan Plateau (Yeh and
Gao, 1979; Murakami and Ding, 1982; Luo and
Yanai, 1984; Ding, 1992). Based on the estimate
of heat budget made by Yeh and Gao (1979) and
others, the total energy supplied by the Tibetan
Plateau has its maximum in late spring and early
summer, with a peak occurring in May. This heat
flux from the surface to the atmosphere has its
maximum contribution from the sensible heat.
Thus, the atmosphere over the Tibetan Plateau
in May and June becomes the strongest atmospheric heat source in a year, and has abnormally
high temperature with the warmest region in July
and August found in the region of the longitudinal range of 50 –110 E. It is very interesting
that during the transition season from spring to
summer, the warming in this region occurs earlier than in other zones of the same latitude. In
March, the increase in thickness (500–300 hPa)
is also evident and attains its maximum in May
and June (Yeh and Gao, 1979), preceding the
onset of the Asian summer monsoon in timing.
All of these studies have well documented the
thermal forcing of land-sea contrast, especially
the Tibetan Plateau and its surrounding areas,
on the onset of the Asian summer monsoon.
Next, one may naturally ask why the earliest
onset occurs in the Indochina Peninsula and the
Fig. 17. Hovemoller diagrams of vertical shear of zonal wind (m s 1 ) between a 200 hPa and 850 hPa averaged over
10 –20 N, (b) temperature difference ( C) between 20 N and 10 N averaged over the 850–200 hPa layer and (c) instability
index (K=1000 hPa) averaged over 10 –20 N. Shading in (a), (b) and (c) denotes, respectively, easterly vertical shear,
positive temperature difference, and instability index over 65 K=1000 hPa. The instability index is defined as the difference
of the saturated equivalent potential temperature between 1000 hPa and 700 hPa (divided by the pressure difference) (Wu and
Wang, 2001)
138
D. Yihui and J. C. L. Chan
SCS, rather than in other locations. The study by
He et al (1987) made an initial attempt to provide
some evidence to address this important problem
by using the data of 1979. They found that a
sudden temperature increase over the eastern
Plateau and the central China plain (85 –115 E)
occurred during the period from 6 May to 15 May.
At the same time, the reversal of the meridonal temperature gradient first occurred over
the longitudes east of 85 E and then over the
longitudes west of 85 E. The two stages of the
reversal of the temperature gradient (as well as
the geopotential height gradient) coincide with
the two stages of the onset of the low-level southwesterlies and organized rains over the Bay of
Bengal and the Arabian Sea. The dominant role
played by the temperature increases over the land
areas including the plateau in this reversal has
been further documented by the works of Wu
and Wang (2001), and Zhang et al (2004).
Wu and Wang (2001) also pointed out that the
change of the wind direction or the vertical shear
(200–850 hPa) (Fig. 17a) can be explained by the
reversal of the meridional temperature gradient
(Fig. 17b). The meridional temperature gradient
averaged over the layer of 850–200 hPa reverses
first over the Indochina Peninsula because the
atmosphere heats up more quickly over the land
than over the ocean. The thermal advection of the
warm air from the Tibetan Plateau in relation to
the westerly winds at middle and upper levels
before the onset is also important. The latent heat
released by the pre-summer or spring rainfall in
South China and the Indochina Peninsula possibly make some contribution to heating of the
atmosphere. This view is supported by the development of the zone of high convective instability
(Fig. 17c). As a result, the easterly vertical shear
and the onset of the Asian summer monsoon develops first along Southeast Asian longitudes.
The arrival of the MJO oscillation is likely to
be a triggering mechanism for the sudden onset
and northward propagation of the summer monsoon. But, the MJO alone is not sufficient to trigger the onset of the summer monsoon in some
years and some regions. In such cases, the midlatitude events (troughs and ridges) may play a
substantial role in the monsoon onset (Davidson
et al, 1983; Chang and Chen, 1995; Hung and
Yanai, 2002; Liu et al, 2002). However, very
few investigators have studied the physical pro-
cesses and mechanisms of triggering the onset by
the intrusion of mid-latitude troughs or frontal
systems in detail. Ding and Liu (2001) summarized the possible triggering mechanisms in their
study on the effect of change in circulation features at mid-latitudes on the onset of the northern
SCS summer monsoon based on various previous
studies: (1) lifting effect to release the existing
convectively potential instability for occurrence
of convection and precipitation; (2) accelerating
the low-level northeasterly wind with enhancing
the meridional pressure gradient to increase the
shear vorticity and cyclonic circulation of wind
shear line; (3) enhancing the baroclinicity due to
increase of horizontal temperature gradient, thus
providing some amount of available potential
energy for development of disturbances or mesoscale systems in the frontal zone; (4) exciting the
growth of extensive convective cloud systems,
which is a favorable environment for development of meso-scale systems in the low-level wind
shear zone between northeasterly and southwesterly winds and associated low troughs which
may force the subtropical high to retreat southward and eastward through some kind of feedback process. Chan et al (2000) also emphasized
the importance of southward intrusion of cold air
from mid-latitudes to trigger the onset of the SCS
summer monsoon. Its role is to lift the warm,
moist and unstable air to release the convective
available potential energy (CAPE), when the atmospheric convective instability is already established before the onset through the heat and
moisture transport by the low-level tropical or
subtropical southwesterlies.
The impact from mid-latitudes may be observed not only for the onset of the East Asian
summer monsoon, but also for all stages of its
seasonal progress. The continuous southward intrusion of cold air and accompanying frontal systems (the so-called Meiyu=Baiu front) is excited
by the development and prevailing of blocking
highs in the mid-and high latitudes over Eurasia.
The dual blocking high situation, one located
over the Ural Mountains and another located
over the Okhotsk Sea, is the most favorable situation for prolonged Meiyu=Baiu heavy rainfall
(Ding, 1991; Zhang and Tao, 1998; Wu, 2002).
So, one of the main differences between the
Indian and East Asian summer monsoon is the
different effect of mid-latitudes events.
The East Asian summer monsoon
The East Asian summer monsoon assumes a
great interannual variability. Numerous investigators have linked this variability to changes in
Eurasian or Tibetan snow cover (Liu and Yanai,
2002) and the Pacific SST. National Climate
Center of China (1998) has identified a positive
correlative relationship between the snow cover
over the Tibetan Plateau in preceding winter
and spring and rainfalls in the following summer in the region of the Yangtze River Basin.
Recently, Zhang et al (2003) has further indicated the existence of a close relationship between the interdecadal increase of snow depth
over the Tibetan Plateau during the preceeding
spring, and the excessive summer rainfall over
Yangtze River Basin. It is proposed that the excessive snow results in decrease in heat sources
over the Tibetan Plateau, through the increased
albedo and spring snow melting, thus reducing
the land-sea thermal contrast, the driving force
of the Asian summer monsoon (Ding and Sun,
2003).
The effect of ENSO events on the East Asian
summer monsoon and related seasonal rainfalls
has been extensively studied. It has been found
that the most significant influence occurs in the
following year after the onset of El Ni~
no events
(NCCC, 1998) with above-normal rainfalls observed in the Yangtze River Basin. Under this
condition, the weak summer monsoon may be
expected. Recently, Wang et al (2000) have
found that ENSO events can affect the East
Asian climate through a Pacific-East Asian (PEA)
teleconnection, with an anomalous anticyclonic
east of the Phillipines during El Ni~
no events
often observed over West-Pacific and the southward shift of the seasonal rain belt.
The interdecadal variability of the East Asian
summer monsoon is now of considerable concern
for many investigators (Lau and Wang, 1999;
Wang et al, 1999; Chang et al, 2000; Ding and
Sun, 2003). They have linked the interdecadal
variability of the East Asian monsoon to an interdecadal change in the background state of the
coupled ocean-atmospheric system or a longterm warming tend in the tropical Indian Ocean
and Pacific. Among these contributing factors,
the Pacific Decadal Oscillation (PDO) and Indian
Ocean Dipole (IOD) may play a very important
role. Their relationship to the East Asian summer
monsoon remains to be further studied.
139
7. Conclusions
The present paper provides an overview of major
problems of the East Asian summer monsoon.
The major conclusions drawn upon this review
can be summarized below:
(1) The earliest onset of the Asian summer monsoon occurs in most of cases in the central
and southern Indochina Peninsula. The onset process over the SCS and the Indochina
Peninsula is very abrupt, with dramatic
changes of large-scale circulation and rainfall
occurring during a quite short time period of
about one week.
(2) The onset of the summer monsoon over the
Indochina Peninsula and the SCS is preceded
by development of circulation features and
convective activity in the tropical East Indian
Ocean and the Bay of Bengal that is characterized by the development of a twin cyclone
crossing the equator, the rapid acceleration of
low-level westerlies and significant increase
of convective activity in both areal extent
and intensity.
(3) The seasonal march of the East Asian summer monsoon displays a distinct stepwise
northward and northeastward advance, with
two abrupt northward jumps and three stationary periods. The monsoon rain commences over the region from the Indochina
Peninsula-the SCS-Philippines during the
period from early May to mid-May, then it
extends abruptly to the Yangtze River Basin,
and western and southern Japan, and the southwestern Philippine Sea in early to mid-June
and finally penetrates to North China, Korea
and part of Japan, and the topical western
West Pacific.
(4) After the onset of the Asian summer monsoon, the moisture transport coming from
Indochina Peninsula and the South China
Sea plays a crucial ‘‘switch’’ role in moisture
supply for precipitation in East Asia, thus
leading to a dramatic change in climate regime in East Asia and even more remote
areas through teleconnection.
(5) The East Asian summer monsoon and related
seasonal rain belts assumes significant variability at intraseasonal, interannual and interdecadal time scales. They can strongly affect
and modulate the onset, active-break cycle
140
D. Yihui and J. C. L. Chan
and propagation of the East Asian summer
monsoon. Their interaction, i.e., phase locking, and in-phase or out-phase superimposing,
can to a greater extent control the behaviors
of the East Asian summer monsoon and produce unique rythem and singularities.
(6) Tow external forcing, i.e., Pacific and Indian
Ocean SSTs and the snow cover in the
Eurasia and the Tibetan Plateau, are believed
to be primary contributing factors to physical
processes and mechanism related to the East
Asian summer monsoon. However, the internal variability of the atmospheric circulation
is also very important to affect the activity of
the East Asian summer monsoon. In particular, the blocking highs in mid-and high
latitudes of Eurasian continents and the subtropical high over the western Pacific play a
more important role which is quite different
from the condition for the South Asian monsoon. The later is of of tropical monsoon nature while the former is of hybrid nature of
tropical and subtropical monsoon with intense impact from mid-and high latitudes.
Acknowledgements
This work is jointly supported by National Climbing Project
‘‘South China Sea Monsoon Experiment (SCSMEX)’’ and
the Research Grants Council of the Hong Kong Special
Administrative Region Government of China Grant City
U 2=00C.
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Authors’ addresses: Ding Yihui, National Climate Center,
China Meteorological Administration, Beijing 100081
(E-mail: dingyh@cma.gov.cn); Johnny C. L. Chan, Department of Physics and Materials Science, City University of
Hong Kong, 83 Tat Chee Ave., Kowloon, Hong Kong, China
(E-mail: Johnny.chan@cityu.edu.hk)