Modelling the Asian summer monsoon using CCAM

Modelling the Asian summer monsoon using CCAM

Kim Chi Nguyen Æ John L. McGregor

Received: 4 July 2007 / Accepted: 6 November 2008 / Published online: 27 November 2008

� Springer-Verlag 2008

Abstract A ten-year mean (1989–1998) climatology of

the Asian summer monsoon is studied using the CSIRO

Conformal-Cubic Atmospheric Model (CCAM) to down-

scale NCEP reanalyses. The aim of the current study is to

validate the model results against previous work on this

topic, in order to identify model strengths and weaknesses

in simulating the Asian summer monsoon. The model

results are compared with available observations and are

presented in two parts. In the first part, the mean summer

rainfall, maximum and minimum temperatures and winds

are compared with the observations. The second part

focuses on validation of the monsoon onset. The model

captures the mean characteristics such as the cross-equa-

torial flow of low-level winds over the Indian Ocean and

near the Somali coast, rainfall patterns, onset indices,

northward movements, active-break and revival periods.

Keywords Regional climate modelling �Asia summer monsoon � Dynamical downscaling �Variable resolution model

1 Introduction

The Asian summer monsoon (ASM) plays a vital role in

China (Ding and Liu 2001), Thailand, Vietnam, Cambodia,

India (Fein and Stephens 1987) and other nearby countries.

The Asian region contains more than 60% of the world

population. Its agricultural and industrial productivity are

heavily dependent on the monsoon seasons (e.g., Fein and

Stephens 1987). However, the monsoon may also bring

destruction to the region and cause thousands of lives to be

lost or displaced, as occurred during the 1998 and 2000

floods in China. It is therefore important to predict the

monsoon seasons to help the affected communities to better

prepare for such destruction, and also to better utilize the

benefits that the monsoon may bring.

Many researchers have described the main dynamical

seasonal features of the ASM: the low-level cross-equato-

rial southwesterly flow over the Indian Ocean towards

South Asia and east China, starting in May (Krishnamurti

1985); the easterly trade winds in the tropical Pacific; the

cyclonic vortex over the south Indian Ocean and a high

pressure system in the western Pacific (Li and Yanai 1996).

Also, in the upper troposphere at 200 hPa, there is a huge

anticyclonic circulation (the Tibetan High) located on the

south side of the Tibetan Plateau (Krishnamurti 1985; Li

and Yanai 1996), with an associated westerly jet to its

north and an easterly jet to its south.

The ASM is affected by the surrounding orography,

especially the Tibetan Plateau (He et al. 1987; Luo and

Yanai 1983, 1984). Due to orography intruding up to the

600 hPa level of the troposphere, the Plateau has a strong

influence on the atmospheric general circulation over the

region, both thermally and dynamically. The Plateau exerts

dynamical blocking of the air motion, resulting in strong

atmospheric perturbations (Liu and Yin 2001) and induces

vertical circulations (He et al. 1987). Sensible heat fluxes

from the elevated ground surface of the Tibetan Plateau in

spring lead to a reversal of the meridional temperature

gradient in the upper troposphere south of the Plateau,

which occurs at the same time as the onset of the ASM (Li

and Yanai 1996). The region is also surrounded by open

oceans, the Indian and the Pacific Ocean, which are the two

major moisture sources contributing to the ASM. Lim et al.

K. C. Nguyen (&) � J. L. McGregor

CSIRO Marine and Atmospheric Research,

PB1, Aspendale, VIC 3195, Australia

e-mail: [email protected]

123

Clim Dyn (2009) 32:219–236

DOI 10.1007/s00382-008-0492-5

(2002) showed that the sea level pressure anomaly over

these oceans is the governing mechanism for the monsoon.

As reported by Kang et al. (2002), general circulation

models (GCMs) have great difficulty in simulating the

correct location and amount of rainfall for the Asian

summer monsoon. Using a limited-area model with a large

domain, Gao et al. (2006) found that the circulation and

rainfall patterns during the mid- to late-monsoon months

improved with increasing model resolution, and they

advocated a minimum grid resolution of 60 km. It has also

been found that increasing model horizontal resolution can

substantially improve the simulation of Meiyu-Baiu frontal

structures and the associated rain (Kusunoki et al. 2006;

Kitoh and Kusunoki 2007). Other researchers have also

used limited-area models to study extreme flood events

with short climate simulations of a few months (Leung

et al. 1999; Wang et al. 2003; Hsu et al. 2004).

The current study uses the CSIRO Conformal Cubic

Atmospheric Model (CCAM) to study the ASM. The

purpose of this research is to verify whether the model can

reproduce significant features reported by Ding (2004) and

other authors. In this way we aim to identify model

strengths and weaknesses in simulating the monsoon in this

region. The west Pacific summer monsoon has close links

to the East Asian summer monsoon and both are influenced

by ENSO; during warm ENSO years, the anomalous anti-

cyclones in the western north Pacific strongly affect the

East Asian monsoon (Wang and Li 2004). The decaying

phase of ENSO has pronounced effects on anomalous

summer rainfall in northern China and Japan (Wu and

Wang 2002). The meridional shift of the ridge of the

western Pacific subtropical high and its zonal migration

influence the locations of convection associated with the

East Asian summer monsoon (Tao and Chen 1987).

The layout of the paper is as follows. The model

description and simulation design are provided in Sect. 2.

The model climatic results over Asia are investigated in

Sect. 3. Section 4 concentrates on further aspects of the

simulation of the Asian summer monsoon. Conclusions are

drawn in Sect. 5.

2 Model description and simulation design

2.1 Model description

CCAM is formulated on a quasi-uniform grid, derived by

projecting the panels of a cube onto the surface of the

Earth. The conformal-cubic grid geometry was devised by

Rancic et al. (1996). Using the Schmidt (1977) transfor-

mation, CCAM can be run in stretched-grid mode to

provide high resolution over any selected region. Com-

pared to the more traditional nested limited-area modelling

approach, CCAM provides great flexibility for dynamical

downscaling from any global climate model, requiring only

sea surface temperatures (SSTs) and (optionally) some

form of nudging from the host model (McGregor and Dix

2001). It avoids a number of problems that may occur with

limited-area regional climate models, such as reflections at

lateral boundaries, as discussed by Wang et al. (2004).

The dynamical formulation of CCAM includes a num-

ber of distinctive features. The model is hydrostatic, with

two-time-level semi-implicit time differencing. It employs

semi-Lagrangian advection with bi-cubic horizontal inter-

polation (McGregor 1993, 1996), together with total-

variation-diminishing vertical advection. The grid is un-

staggered, but the winds are transformed reversibly to/from

C-staggered locations before/after the gravity wave calcu-

lations, providing improved dispersion characteristics

(McGregor 2005b). Three-dimensional Cartesian repre-

sentation is used during the calculation of departure points,

and also for the advection or diffusion of vector quantities.

As with most semi-Lagrangian models, the time differ-

encing is made weakly implicit by off-centering (Rivest

et al. 1994) in order to avoid resonances near steep orog-

raphy for large Courant numbers. Special care is also taken

in the advection of temperature and surface pressure near

terrain. Further details of the model dynamical formulation

are provided by McGregor and Dix (2001) and McGregor

(2005a).

CCAM includes a fairly comprehensive set of physical

parameterizations. The GFDL parameterization for long-

wave and shortwave radiation (Schwarzkopf and Fels

1991) is employed, with interactive cloud distributions

determined by the liquid and ice-water scheme of Rotstayn

(1997). The gravity wave drag scheme of Chouinard et al.

(1986) is included. The model employs a stability-depen-

dent boundary layer scheme based on Monin–Obukhov

similarity theory (McGregor et al. 1993), together with

non-local vertical mixing (Holtslag and Boville 1993) and

also enhanced mixing of cloudy boundary layer air (Smith

1990). A canopy scheme is included, as described by

Kowalczyk et al. (1994), having six layers for soil tem-

peratures, six layers for soil moisture (solving Richard’s

equation), and three layers for snow. CCAM also includes

a simple parameterization to enhance sea surface tempera-

tures under conditions of low wind speed and large

downward solar radiation, affecting the calculation of

ocean surface fluxes.

CCAM employs a mass-flux cumulus convection

scheme (McGregor 2003), which includes both downdrafts

and detrainment. Cloud base is determined by proceeding

downwards to find the first moist adiabatically unstable

layer. It is assumed that the mixing ratio of the air parcel at

the potential cloud base is enhanced by 5% over sea and

15% over land; the latter value is intended to represent the

220 K. C. Nguyen, J. L. McGregor: Modelling the Asian summer monsoon using CCAM

123

greater spatial variability expected in the boundary layer

over land. Convection is permitted up to the uppermost

model layer for which the parcel is moist adiabatically

unstable. The closure is that convection is calculated to be

exhausted during a 20-min convective time-scale, the same

as the model time step for these simulations. Exhaustion

occurs for the smallest possible mass flux such that the

modified environment no longer provides the given cloud

base or supports the convective plume. A simple detrain-

ment scheme is used for deep convection, in that a fraction

of the condensed moisture is transferred to liquid and ice-

water in the environment surrounding the upper part of the

convective tower; a fairly large fraction of 30% is pre-

scribed, which acts to transfer a significant fraction of the

precipitation to the resolved scale. In any time step, three

passes of the convection scheme are performed to avoid the

possibility that the cloud-base or cloud-top layers are only

marginally satisfying the convective stability criteria.

Although the cloud base condition is rather simple; it

permits an estimate of the new cloud-base conditions as a

result of the environment modification, and hence permits a

simple and natural cumulus closure.

2.2 Simulation design

The present CCAM regional climate simulation over Asia

uses a C63 global grid (six panels each having 63 9 63

grid points), with 18 levels in the vertical extending up to

4.5 hPa, with the lowest level at about 38 m. A Schmidt

stretching factor of 0.37 was used, giving about 60 km

resolution over the region of interest, which includes 75–

140E and 5–60N. The dominant orography in this region is

the Himalayas, including the Tibetan Plateau, as shown in

Fig. 1. Figure 1 also displays the conformal-cubic grid

cells, shown at one-third resolution. The vegetation types

over the Asian region are specified as mainly broadleaf

evergreen trees (tropical forest) and agriculture or grass-

land. The soil in east China is specified as fine clay to

coarse-medium sandy-loam. West China is specified as

silty-clay-loam, with the same soil type being used over

India. The Indochina Peninsula is specified as silty-clay to

coarse sandy-clay-loam.

The CCAM model can be run in stand-alone mode,

requiring only initial conditions and prescribed fields for

SST and sea-ice. For the simulation described in this

paper, global nudging of winds above 500 hPa from the

large-scale fields is also employed, with an e-folding time,

sefold, of 24 h, to help the jet stream locations of the

modelled synoptic systems follow those of the reanalysis.

This technique is appropriate when comparisons are being

made with specific years (such as when the forcing is

being provided by reanalysis fields). The equations are

described here for the u component, as an example, when

nudged by the host model (or reanalysis) field, uhost.

Applied in a time-split manner, the nudging tendency

equation is

Fig. 1 Orography (m) used for

the simulation. The conformal-

cubic grid cells are also shown,

at one-third resolution

K. C. Nguyen, J. L. McGregor: Modelling the Asian summer monsoon using CCAM 221

123

ou

ot¼ uhost � u

sefold

: ð1Þ

If the value of u before the nudging correction is denoted

by u0, the analytic solution of (1) is

usþ1 ¼ uhost þ exp�Dt

sefold

� �ðu0 � uhostÞ: ð2Þ

From (2), it can be seen that if the nudging is applied for a

time interval equal to sefold, then the difference of u from

uhost will be reduced during that time interval by a factor

exp(1). In practice, the following simple finite difference

solution to (1) is used

usþ1 ¼ u0 þ Dtðuhost � u0Þ=sefold: ð3Þ

It may be noted that for some other stretched regional

climate simulations using CCAM, a far-field wind nudging

technique has been employed (McGregor et al. 2002),

where the far-field model resolution may be considered too

coarse to provide a reliable climatology in the coarse

region. Note that the wind-nudging techniques avoid the

lateral boundary problems experienced by limited-area

regional climate models, which typically use a one-way

nesting technique.

A 10.5-year simulation from July 1988 to December

1998 was carried out with initial conditions, sea surface

temperatures (SSTs), sea-ice and upper level nudging of

winds provided by the NCEP-1 reanalyses (Kalnay et al.

1996). The 6 months of 1988 are regarded as a spin-up

period for the simulation. The model output was saved

once per day at 00 GMT, except rainfall was saved every

3 h. The simulation design is very similar to that used for

the CCAM simulation submitted to the Regional Model

Intercomparison Project (Fu et al. 2005), the main differ-

ence being the use of a more recent model version.

3 Mean states of JJA rainfall, temperature and wind

The 10-year (1989–1998) monthly-mean CCAM rainfall

was averaged to produce seasonal averages for June, July

and August (JJA). These averages are compared in Fig. 2

against the observed seasonal rainfall provided by the

Climate Prediction Center Merged Analysis of Precipita-

tion (CMAP) and the Climate Research Unit (CRU), and

also against rainfall from the host NCEP reanalyses. The

horizontal resolution is 2.5� for NCEP and CMAP and

about 0.5� for CRU (which thus should be most compa-

rable to the CCAM rainfall).

In general, CCAM captures well the rainfall patterns

over the land compared to CRU, including high-rainfall

coastal regions such as the long thin strip along the west

coast of India, the east coast of the Bay of Bengal and the

east coast of the Gulf of Thailand. These high-rainfall

strips are related to land–sea contrasts (land roughness and

diurnal temperature effects) and advection of humid air

onto a rising orography. The coastal rainfall peaks along

the west coast of India can be related to deceleration of the

low-level Somali jet by the Indian subcontinent, which

produces surface wind convergence with associated

upward flux of moisture in the eastern Arabian Sea

(Halpern and Woicesshyn 1999); in support of this rela-

tionship, the model produces about 95% correlation (not

shown) of the area-averaged 850 hPa zonal wind compo-

nent over the Arabian Sea (50–65E, 5–15N) with the

rainfall over the land along the west coast (70–77E, 10–20N).

The simulated rainfall peaks along the west coast of India

occur just after midday (not displayed). Over the Bay of

Bengal, the model reproduces the heavy rainfall shown by

CMAP, although the peak is displaced a little westward.

Along the east coast of the Bay of Bengal, simulated

rainfall occurs throughout the day but is heaviest during the

morning with a peak 1 h before midday (not displayed).

CCAM reproduces, but more excessively, the rain sha-

dow (dry) regions off the eastern coasts of India, Sri Lanka,

Malaysia and Vietnam. The dry regions are located over

the water, and on the leeside of the land masses. Over the

oceans, the high rainfall in the western Pacific and north of

the Philippines is reproduced by CCAM, except it is a little

excessive, more east–west oriented and displaced a little

north. On the other hand, NCEP rainfall exceeds CMAP

over east China, Thailand and north Vietnam, but appears

under-estimated over the sea, in particular near the Phil-

ippines and also over the Bay of Bengal. The CCAM

rainfall patterns have much improved detail over the land

areas, compared to the NCEP patterns.

In the foothills of the Himalayas, CCAM produces high

rainfall exceeding 16 mm day-1, somewhat more than the

CRU and CMAP values of 4–6 mm day-1. Note that there

are few observations in this region, and that neither CRU

nor CMAP include corrections for orographic effects or

gauge-undercatch for strong winds. The modelled rainfall

peak is also displaced south of the peak given by the

observation stations, located around latitude 28.4N (see

Fig. 1 of Barros et al. 2000). The CCAM rainfall peak over

the foothills of the Himalayas is produced in the early

morning hours around 0400 LT (not displayed). This

modelled nocturnal rainfall occurs about 2–3 h later than

the observations documented by Barros et al. (2000) and

by Barros and Lang (2003); a more recent description of

the observations is provided at http://www.hyarc.nagoyau.

ac.jp/game/6thconf/html/abs_html/pdfs/T8APB19Oct04100

318.pdf. According to Barros and Lang (2003), there is a

nocturnal rainfall peak during the summer monsoon, and

this is primarily due to an interaction of the monsoon flow

with the south slope of the Himalayas, modulated by the


123

http://www.hyarc.nagoyau.ac.jp/game/6thconf/html/abs_html/pdfs/T8APB19Oct04100318.pdf.



diurnal variability of the atmospheric state. Note also that

the average rainfall rate of low and high altitudes for 1999

and 2000 (their Fig. 2) from June to September for all

stations, is roughly 45 mm day-1, providing better agree-

ment with CCAM than the CRU and CMAP values.

Maximum and minimum surface-air temperatures over

land are plotted for both CRU and CCAM in Figs. 3 and 4.

CCAM reproduces the maximum except in India, eastern

China and Thailand where the model is about 2�C cooler

than the observations. The minimum temperatures are

reasonably reproduced, although a cold bias is produced

over central Indochina and a warm bias over northern

China (both biases about 2�C). Minimum and maximum

temperatures have a warm bias of about 6�C over Mon-

golia, and a warm bias of about 4�C over Afghanistan.

Winds and geopotential heights at three levels are

compared to NCEP in Fig. 5. The main features at 850 hPa

are the Mascarene High (55E, 32S), the Australian High

(135E, 35S), the Subtropical High (155E, 35N) and the

Somali jet. It is seen that CCAM reproduces these features.

However, the CCAM Somali jet is stronger and extends

further east into the South China Sea, in association with an

eastward displacement of the western Pacific high. Asso-

ciated with this, CCAM is unable to simulate the correct

location of the confluent area established by southeasterly

flows from the high and the subtropical westerly flows in

the south of Japan. As a result, less moisture from the

Pacific is transported into Korea and Japan (see Fig. 13).

Excessive rainfall in CCAM over the Bay of Bengal and

the western Pacific, as mentioned above, is consistent with

the 850 hPa low-level jet being too strong (Fig. 5), trans-

porting excessive moisture from the southern Indian Ocean

to these regions (as will be seen in Fig. 7). A band of strong

winds stretches from the Arabian Sea to the Philippine Sea

(a) (b)

(d)(c)

Fig. 2 Rainfall rate (mm day-1) for JJA for (a) CMAP, (b) CRU, (c) NCEP and (d) CCAM, shaded for values exceeding 2 mm day-1


123

(40–140E, 5–20N), and appears to correlate with the

maximum rainfall, north of the Philippines, described

above (Fig. 2). Lau et al. (2004a) report a similar dis-

placement of the monsoon trough towards the east in their

GCM simulation.

The main features at 500 hPa are the Subtropical High

over the western Pacific, the Iranian High in the western

domain and a weak cyclonic circulation over the northern

Bay of Bengal. CCAM experiences great difficulty in

simulating the Subtropical High at this level with its centre

displaced eastwards. The reason for the displacement is

unclear, but it is conceivably related to the model treatment

of flow over the Tibetan Plateau. At 200 hPa, the westerly

jet stream agrees with NCEP having a maximum speed

over the Tibetan Plateau, and stretching out into the Pacific.

In the equatorial part of the domain, there is easterly flow

with maximum wind speed over the Arabian Sea and south

of India. There is a huge anti-cyclonic centre around 25N.

CCAM also has a secondary anti-cyclonic centre located

south of Japan, which is not seen in NCEP.

As mentioned above, the monsoon rainfall has its

moisture source from the equatorial Indian Ocean. To

illustrate this, evaporation minus precipitation is plotted in

Fig. 6. The moisture-sink regions (net precipitating) are

mainly in the northern hemisphere and located east of 70E;

intense sinks are seen along the coastal regions, south of

the Tibetan Plateau and the western north Pacific, whilst

the moisture sources are seen to come from the Southern

Hemispheric Indian Ocean. CCAM and NCEP agree

approximately on the locations of the intense moisture-sink

regions. In general CCAM captures the main patterns, but

produces stronger moisture sinks than NCEP, and the

Fig. 3 Average maximum surface-air temperatures (�C) for JJA over land for CRU (left) and CCAM (right), where temperatures above 20�C are

shaded

Fig. 4 Average minimum surface-air temperatures (�C) for JJA over land for CRU (left) and CCAM (right), where temperatures above 10�C are

shaded


123

orientation of the moisture-sink band is more east–west.

This is consistent with Fig. 2d, which shows that CCAM

produces more rain than NCEP.

The vertically integrated (1,000–300 hPa) moisture

transport from the CCAM simulation is shown in Fig. 7

(see Ding 2004, Fig. 26, p. 37). This diagram confirms that

moisture is advected into the Bay of Bengal and the South

China Sea, and later transported northwards into east Asia,

and that much of the moisture originates south of the

equator from the Indian Ocean. This feature agrees with the

description by Ding (2004).

4 Onset and progress of the Asian summer monsoon

In this section, some unique features of the Asian mon-

soon will be discussed, such as abrupt changes, northward

propagation with a sudden jump of rainfall distributions,

Fig. 5 JJA geopotential height contours (m) and wind vectors for NCEP (left) and CCAM (right), at 200, 500 and 850 hPa, with wind speeds

shaded above 5 m s-1 (but shaded above 20 m s-1 for 200 hPa)


123

and active-break-revival characteristics of the rainfall

for selected regions in southeast Asia. Criteria for the

onset of the East Asian summer monsoon were devised by

Qian and Lee (2000) and were reproduced in Fig. 3 of

Ding (2004): the low-level wind at 850 hPa becomes

westerly; the outgoing longwave radiation (OLR) is less

than 230 W m-2; the brightness temperatures are below

244 K; the daily rainfall rate exceeds 6 mm day-1 over

the South China Sea (110E–120E, 10N–20N). Figure 8

shows 5-day averages (pentads) from CCAM averaged for

10 years over the South China Sea for the OLR, the zonal

wind and the rainfall rate; here pentad one is the average

of January 1 to January 5 inclusive; Fig. 8 may be

compared to Fig. 3 of Ding (2004). It is seen that the

zonal wind becomes westerly (positive) and the OLR falls

below 220 W m-2 around 21–25 May (pentad 28).

However, CCAM rainfall for the South China Sea

increases gradually from approximately 3 mm day-1

(before onset) to 5 mm day-1 (after onset), whereas the

observations shown by Ding (2004) exhibit a jump

from about 2 mm day-1 to almost 8 mm day-1. Based on

this result (OLR and zonal westerly wind), the average

onset date over the South China Sea as produced by

CCAM is around 21–25 May (pentad 28), which is a little

later than the date (16–20 May) reported by Zhang et al.

(2004). In the following section, we include extra con-

ditions which produce closer agreement for the average

onset date.

Fig. 6 Moisture sources (positive, white) and sinks (negative, shaded) from NCEP (left) and CCAM (right); units are 10-5 kg m-2 s-1

Fig. 7 Vertically integrated

(1,000–300 hPa) moisture

transport vectors from CCAM

for JJA, with shading for

magnitudes exceeding 100

kg m-1s-1


123

4.1 Northward propagation of the monsoon and abrupt

changes

To provide a picture of the entire Asian region, showing

northward propagation of the monsoon onset, a map of the

modelled mean onset index is plotted in Fig. 9 for suc-

cessive pentads from early May, together with rainfall

(shaded). The onset index used here is adopted from Zhang

et al. (2004). It is a combination of westerly low-level

(850 hPa) winds, easterly upper-level (200 hPa) winds and

OLR below 240 W m-2 (note that an additional condition

of 200 hPa zonal easterly wind is imposed in Fig. 9 com-

pared to Fig. 8). Ding and Liu (2001) used a threshold

OLR of 230 W m-2, stating that the threshold is somewhat

arbitrary and that the rapid change of prevailing wind

direction for a prolonged period is a more essential crite-

rion to define an active monsoon. During May 1–5, the

onset index is south of latitude 13N, over the Indian Ocean

near the southern tip of the Indian subcontinent, the Bay of

Bengal and the Indochina Peninsula. Even though onset is

still in the south of India, heavy rainfall can already be seen

over the Tibetan Plateau and over China, along the Yangtze

River, and is thought to have its origin in the mid-latitudes

or as a combination of tropical and frontal rainfall (Ding

2004).

By May 6–10, the onset location arrives at 12N, but

extends over the entire southern peninsula and eastward to

the southern South China Sea with increased rainfall over

these regions. At the same time, rainfall over the Arabian

Sea and the southern coast of India is widespread, due to

development of a prominent anticyclonic circulation in

the northern Arabian Sea and enhancement of the cross-

equatorial flow off the Somali coast (Krishnamurti et al.

1981). By May 11–15 the onset advances northward in

the Bay of Bengal and Indochina. During May 16–20,

there is significant rainfall over the southern Arabian Sea.

At the same time, the onset location advances rapidly

northeastward into the South China Sea and the Philip-

pines. During this period, rainfall over the Tibetan Plateau

is increasing. The rainfall is also widespread over south-

east China.

During May 21–25, the onset rapidly spreads north and

northwestward from the Bay of Bengal to reach the east

coast of India. Over the South China Sea, the maximum

rainfall moves north to southeast China and the onset

advances to 18N, into the south of China. There is heavy

rainfall (exceeding 10 mm day-1) over east China, the

northern Bay of Bengal and the Arabian Sea. Rainfall over

the South China Sea also reaches a maximum and the onset

advances eastward into the Philippine Sea, but there is slow

progress over India. During the May pentads, the rain belt

over east China is displaced too far north, towards the

Yangtze River, compared to the observed rainfall shown by

Ding (2004) (his Fig. 10, p. 18). By the end of June (not

shown), the onset advances well into southern China and

also covers most of India, except for its northwestern

corner. Heavy rainfall is also seen in the vicinity of Cal-

cutta. The rain belt over China is pushed northward, across

the Yangtze River.

From the model onset index, it is seen that the onset first

occurs over the southern tip of India, the Bay of Bengal and

the Indochina Peninsula. It then spreads northeastward to

the South China Sea and southeast China, and last to the

Indian subcontinent. Also, the results show that the mon-

soon onset advances northwards in a discontinuous

stepwise manner, in agreement with Ding (2004); there is

little change between May 11–15 and May 16–20 over

Indochina and the South China Sea where the onset arrives

around 16N, but during May 21–25 the onset spreads to

about 18N, then there is little change for the next two

pentads but it advances to 22N during June 1–6. In the

current study, the onset dates as derived by averaging for

each day over 10 years of data, for the Bay of Bengal, the

South China Sea, southeast China and the centre of India

are respectively early May, mid May, early June and late

June. These dates appear to agree with the three stages of

the Asian summer monsoon onset suggested by Wu and

Zhang (1998) and supported by Fong and Wang (2001) and

Mao et al. (2004). In general, the model onset dates over

land also agree with Zhang et al. (2004), as reproduced by

Ding (2004). However, for onset over the Arabian Sea and

the eastern side of the Philippines, the CCAM onset occurs

rather earlier. It may be that the precipitation scheme is

producing excessive rainfall and therefore cloud cover

which acts to decrease the OLR compared to NCEP, even

though CCAM captures the general pattern of the OLR (not

shown).

160

180

200

220

240

260

280

5 10 15 20 25 30 35 40 45 50 55 60 65 70-10

-5

0

5

10

15

pentad numberOLR

rainfallzonal wind

Fig. 8 Ten-year averages over the South China Sea of CCAM

rainfall (mm day-1), OLR (W m-2) and 850 hPa zonal wind speeds

(m s-1) plotted against pentad number. The scale on the left is for

OLR and the scale on the right is for rainfall and wind speeds


123

Fig. 9 CCAM precipitation rate (shaded above 5 mm day-1) and onset index (black dots) showing the progression of the summer monsoon,

displayed for successive pentads from early May


123

The abrupt changes of the onset are illustrated in

Fig. 10, which shows pentad 28 (May 21–25, after onset in

the South China Sea) minus pentad 25 (May 6–10, before

onset in the South China Sea) for 200 hPa winds, 500 hPa

geopotential height, 850 hPa winds and OLR, for both

NCEP and CCAM. Figure 10c shows the intensification

during monsoon onset of the 850 hPa equatorial westerly,

the Somali jet over the Arabian Sea, the Bay of Bengal and

the South China Sea. CCAM produces a similar pattern,

but with stronger wind changes than those of NCEP, pre-

sumably related to larger latent heat fluxes and convective

rainfall of CCAM in these regions. CCAM also shows

clearer cyclonic flows than NCEP over the Gulf of Oman,

the Bay of Bengal and Japan, although a little displaced.

However, it does not fully capture the changes over central

and southern China and the South China Sea. Centres of

deepening of the 500 hPa geopotential height (Fig. 10b)

over the Bay of Bengal and Japan, and a weakening of the

subtropical high over the South China Sea, are in general

agreement between CCAM and NCEP. However, CCAM

does not accurately reproduce the changes over Japan and

the South China Sea. The deepening of the 500 hPa trough

over the Bay of Bengal and Japan (Fig. 10c) is closely

related to the development of the 850 hPa cyclonic circu-

lations (Ding 2004). Note that although the CCAM low-

level westerly flows are too intense over the Indian Ocean,

they are too weak over southeast Asia and west North

Pacific; these weak westerly flows lead to some deficien-

cies in the development of the monsoon over East Asia and

the associated summer precipitation. CCAM also shows a

centre of deepening over the Arabian Sea which is very

weak in NCEP.

There is agreement between NCEP and CCAM for the

pronounced westward acceleration region of the upper-

level 200 hPa winds (Fig. 10a, shaded, 40–120E, 20–25N)

during the monsoon onset. The two anticyclonic vortices

located east and west of the Tibetan Plateau and to the

north of the accelerated region are seen in both NCEP and

CCAM. However, the two anticyclones in CCAM are not

well separated due to the CCAM wind changes being

weaker and more westerly between 70–110E and 30–40N

(Fig. 10a). The CCAM wind changes are much stronger

than NCEP in the Indian Ocean southward of 10N and

from 40E to 80E, where the NCEP and CCAM wind

regimes show greater differences. Regions of deep con-

vection during monsoon onset can be seen in Fig. 10d

where there is a significant decrease of OLR in both data

sets over the Bay of Bengal, the South China Sea and the

Arabian sea. Over the Arabian Sea, CCAM produces much

larger decreases of OLR than NCEP; this appears to be

correlated with an overestimation of the change of low-

level jet speed in this region as seen in Fig. 10c. The

location of simulated convection (Fig. 10d) over the

western Pacific is displaced further south. This may be due

to CCAM failing to simulate the correct location of the

confluent region as described above (Fig. 5).

The onset dates of the East Asian summer monsoon are

greatly affected by the western Pacific subtropical high

(Lau et al. 2004b; Lim et al. 2002), whose movement can

bring flood or drought to certain locations. For example,

flood over southern China, the Korean peninsula and Japan

in 1998 is believed to be due to the persistent and strong

low-level anticyclonic anomaly in the subtropical western

Pacific (Shen et al. 2001). The progress with time of the

western Pacific subtropical high during May is shown in

Fig. 11, as plots of 850 hPa wind vectors and geopotential

height. It is noticeable that there are strong westerly winds

in the northern periphery of the western Pacific subtropical

high, with correspondingly strong gradients of geopoten-

tial. For May 1–5 (Fig. 11a), there are low-level westerly

winds over the Indian Ocean near the equator. Over the

western Pacific, there is a southeasterly wind of similar

strength located to the south of the subtropical high. Over

the Bay of Bengal, the wind becomes southwesterly around

6–10 May (not shown) and strengthens around mid-May

(not shown), in agreement with He et al. (1987). The

westerly and southwesterly winds bring some rain to the

southern tip of India and the Bay of Bengal (as seen in

Fig. 9). As time progresses, there is development of cross-

equatorial flow originating from the southern Indian Ocean,

turning into westerlies off the east coast of Africa

(Fig. 11b). This westerly flow intensifies and accelerates

over the Bay of Bengal and extends into the South China

Sea; it replaces the southeasterly flow, as the subtropical

high retreats further into the Pacific. According to Fong and

Wang (2001), the withdrawal of the subtropical high out of

the South China Sea is probably due to convective feed-

back processes and mesoscale activities over the Indochina

Peninsula and the South China Sea. By the end of June (not

shown), the influence of the subtropical high has greatly

weakened in the region. The CCAM subtropical high and

wind ahead of the high are weaker than NCEP (Fig. 11b).

4.2 The sudden jump of monsoon rain

Another characteristic of the Asian summer monsoon is its

sudden jump of rainfall towards high latitudes. This feature

is generally captured by CCAM, as shown in Fig. 12 for

monthly sequences of longitudinally-averaged strips. The

various plots show rainfall for India (70–85E), the Bay of

Bengal (85–95E), the Indochina Peninsula (95–110E), the

South China Sea (110–120E) and the western Pacific (120–

140E).

Rain starts to fall on the southern tip of India (around

latitude 10N) in May and reaches its peak in June–July–

August. The Indian monsoon season usually occurs


123

(a)

(b)

(c)

(d)


123

between June and September. A similar rainfall structure is

reported by Wu et al. (1999) using observed rainfall from

version 1 of the GEOS-DAS (Schubert et al. 1993) from

1980 to 1993, although these authors used a narrower

longitude strip, 75–85E. CCAM captures this rainfall peak,

but produces an extra peak located between 25 and 30N

which covers northern India and the western part of Nepal.

For the Bay of Bengal strip, heavy rainfall reaches

Burma, Thailand, Calcutta and Bangladesh by May; heavy

rain arrives much earlier in April south of the Tibetan

Plateau (25–30N), which includes the eastern part of

Nepal. Both CMAP and CCAM show a sharp gradient of

the rainfall pattern near 30N.

For the Indochina strip, CMAP shows a rainfall peak

from July to October, centred around 10N and covering the

southern tip of Thailand/Malaysia, the Gulf of Thailand

and the southernmost part of Vietnam. This peak is dis-

placed northwards and centred around 15N in CCAM. The

effect of the Tibetan Plateau is still seen in this region; in

CCAM the peak rainfall is southeast of the Plateau,

between 25 and 30N. Over the equatorial part of the

Indochina strip, CCAM slightly underestimates the CMAP

rainfall for the April–October period. Also CCAM is

unable to simulate rainfall intensity for November and

December.

Over the South China Sea strip, heavy rainfall from

January to February is confined between 10S and 5S, which

is mainly the Indonesia region. There is little rain in the

South China Sea (10–20N) during January to April, with

sudden heavy rain over the South China Sea from June to

September around 10–20N, and moving back to the Borneo

region with easing intensity from mid-September. The

monsoon rainfall arrives at northern China (35–40N) about

mid-June or July. Over east China (25–30N), there is some

pre-monsoon rain, due to moisture transport into east China

from the western Pacific, north of the subtropical high

(Ding 2004). According to CMAP, heavy rainfall arrives to

these latitudes after April with a short peak in June. CCAM

simulates heavy rainfall over this region much earlier. In

general, Fig. 12 shows that CCAM underestimates rainfall

for the SCS region. Unlike CMAP rainfall, CCAM rainfall

is fragmented and noisy. The monsoon rainfall peak occurs

around mid-June and the rain band gradually moves back

Fig. 11 Plots of wind vectors (m s-1) and 850 hPa heights (m) for the first (top) and fifth (bottom) May pentads for NCEP (left) and CCAM

(right)

Fig. 10 Changes occurring during monsoon onset (May 21–25 minus

May 6–10) for (a) 200 hPa winds (with speeds shaded above 10 m

s-1), (b) 500 hPa geopotential height (m), (c) 850 hPa winds (with

speeds shaded above 2 m s-1) and (d) OLR (W m-2), with NCEP

(left) and CCAM (right)

b


123

Fig. 12 Ten-year mean rainfall

(mm day-1) for CMAP (left)and CCAM (right) plotted for

latitude versus month for five

longitudinally-averaged strips

representing (top to bottom):

India , Bay of Bengal,

Indochina, South China Sea,

western Pacific, shaded above

4 mm day-1


123

southward around August, which is earlier than CMAP,

with intensity remaining unchanged throughout subsequent

months.

Over the western Pacific strip, CMAP shows a similar

structure as in SCS with some heavy rainfall during

December–February, located in the Southern Hemisphere

between 15 and 5S, which is the northern tip of Australia,

Arafura Sea and New Guinea and its surrounding waters.

Sudden heavy rainfall is seen in the Northern Hemisphere

around mid-May, lasting until mid-October, with its peak

between June-mid September, first occurring between 5

and 10N, which is the southern Philippines and surround-

ing water. The heavy rain arrives in the East China Sea

(20–25N) around July and remains there until September

when its starts to move back to the south. On the other

hand, a more sudden jump of this rain band is seen in

CCAM, with its peak confined between 15 and 25N.

CCAM is unable to capture the narrow rain band between

10S and 5N from March to May; instead its rainfall

intensity is confined between 5S and the equator during

these months.

In summary, the above results show that between Jan-

uary and April, heavy rainfall is confined between 10S and

10N for all strips. From May to September, the heavy rain

band is seen north of 10N. These characteristics show a

sudden and northward jump of the monsoon rainfall. In

general, CCAM captures the main features, but it tends to

produce a dry patch around the equator for all regions from

April to October; the cause of this dry patch is unclear.

During the monsoon period, CCAM shows secondary

peaks which are sometimes not seen in CMAP, between 25

and 30N for the India and Indochina strips, possibly related

to CCAM better resolving the orography of the Tibetan

Plateau. CCAM shows less clearly than CMAP the rainfall

at 5–10N which commences around mid-May over all

regions.

0

5

10

15

6 12 18 24 30 36 42 48 54 60 66 72

SCSCMAP

CCAM

0

5

10

6 12 18 24 30 36 42 48 54 60 66 72

SECCMAP

CCAM

0

5

10

6 12 18 24 30 36 42 48 54 60 66 72

IndiaCMAP

CCAM

0

5

10

15

6 12 18 24 30 36 42 48 54 60 66 72

JapanCMAP

CCAM

0

5

10

6 12 18 24 30 36 42 48 54 60 66 72

KoreaCMAP

CCAM

Fig. 13 Rainfall rates

(mm day-1) derived from 10-

year averages for CMAP and

CCAM displayed for each

pentad of the year for the

following regions: South China

Sea (110–115E, 15–20N, both

land and sea), southeast China

(105–130E, 20–31N) , India

(75–80E, 15–25N), Japan (130–

140E, 32.5–37.5N) and Korea

(125–135E, 35–40N)


123

4.3 Monsoon breaks

During the monsoon season, there is a period of heavy

rainfall (active, exceeding 5 mm day-1) and also a dry

(break) period when there is little (below 5 mm day-1) or

no rainfall, or a drop in rainfall rate. The summer break

over this region is highly variable from year to year. For

example, over India, Ramamurthy (1969) investigated an

80-year period, 1888–1967, and found that there are years

with no break or up to four breaks. The break length also

varies depending on normal or dry years (Krishnamurti and

Bhalme 1976). The current study shows the climatological

break from CCAM for a 10-year period 1989–1998, whilst

the observations from CMAP are from 1979 to 1999. The

characteristic monsoon active-break-revival (re-occurrence

of heavy rainfall), described by Chen et al. (2004), is seen

from the CCAM results for three of the following five

regions: the South China Sea (110–120E, 15–20N),

southeast China (105–130E, 20–30N, including Taiwan),

India (75–80E, 15–25N), Japan (130–140E, 32.5–37.5N)

and Korea (125–130E, 35–40N). For those five regions, the

time series of area-averaged rainfall rate for both CMAP

and CCAM is plotted for the entire year against pentad

number in Fig. 13.

Over the South China Sea, CCAM rainfall starts

increasing steadily around mid-March (pentad 15). A break

is seen in June (pentads 31–34) and a second break is seen

in July (pentads 37–38); the rainfall attains its peak in late

October (pentad 59), and the rainfall declines thereafter.

Compared to CMAP, CCAM over-predicts the rainfall in

the early and late months of the year. In the observations, a

jump from 3 to 8 mm/day occurs in mid-May (pentad 28);

the rainfall maintains this intensity until early July (pentad

37–38) when a short break is observed. This short break

coincides with the second break in CCAM rainfall.

Over the southeast China region, CCAM rainfall follows

the CMAP rainfall reasonably, although it under-estimates

the rainfall a little. Both plots show rainfall over this region

increasing with time and attaining its peak during June.

CCAM rainfall shows an earlier and longer break, which

starts mid-June (pentad 33) and lasts until mid-July (pentad

40), whereas CMAP shows a break starting around pentad

38, but also reviving around pentad 40. Over India, CCAM

captures the rainfall pattern and the break, although CCAM

slightly over-estimates the rainfall.

Over Japan, CCAM rainfall follows CMAP reasonably

well until mid-May (pentad 26) when CMAP shows its first

peak, which is not captured by CCAM. CCAM predicts the

break period (pentads 41–51) reasonably, but is unable to

capture the heavy rainfall period just before this break,

except for a peak at pentad 34, which is one pentad earlier

than CMAP. In the revival period, CCAM under-predicts

the rainfall.

Over Korea, CCAM rainfall follows CMAP until mid-

June (pentad 32) when CCAM rainfall peaks early; a break

is seen from mid-July till early August (pentads 40–44).

CMAP rainfall peaks a little later (pentad 36) and lasts

longer (till pentad 49), followed by a short break. CCAM

under-predicts rainfall for the second half of the year. It is

seen that CCAM produces much less rain than CMAP over

Korea, as discussed earlier. The under-estimation of rain-

fall over Japan and Korea, in particular for the revival

stage, may be related to incorrectly simulating the conflu-

ent region which is associated with the westward extension

of the western Pacific high being too far into the Pacific

(Fig. 5).

According to Chen et al. (2004) and other researchers,

the active and break periods described for the above

regions, except India and the South China Sea, are caused

by the northward movement of both the Meiyu rain band

and the western Pacific subtropical high. However, the

revivals in southeast China and Taiwan are caused by

northward progression of the ITCZ embedded in the

monsoon trough over the South China Sea (Ramage 1952).

The monsoon revivals over Korea and Japan are caused by

the northward movement of the monsoon rainfall which is

linked to the northward progression of the ITCZ (Chen

et al. 2004). In summary, CCAM generally captures the

rainfall distribution and breaks over India, SEC and Japan.

5 Conclusions

A 10-year simulation has been performed, using CCAM to

downscale the NCEP reanalysis to a resolution of about

60 km over the Asian region. Weak wind nudging was

used above 500 hPa to assist the modelled storm tracks to

approximately follow those of the reanalysis. The simula-

tion captures the mean state of the Asian JJA summer

monsoon season. There is good agreement of rainfall

compared to CRU over land. CCAM appears to systemat-

ically produce excessive dryness downwind of large land

masses. CCAM captures the timing of nocturnal rainfall

over Nepal, although the peak is a little weaker and dis-

placed southward.

Maximum surface-air temperatures have a cold bias (but

less than 2�C) over eastern China, Laos and Thailand.

Minimum surface-air temperatures are reasonably simu-

lated, although a cold bias is produced over central

Indochina and a warm bias over northern China. A large

warm bias in both minimum and maximum temperatures is

seen over Mongolia (6�C) and Afghanistan (4�C). The

modelled low-level Somali jet for JJA is too strong and

intrudes too far eastwards into the western Pacific. This

may contribute to excessive rainfall over the Bay of Bengal

and the western Pacific. CCAM experiences difficulties in


123

simulating the locations of the western Pacific high at both

850 and 500 hPa. As a result, CCAM simulates insufficient

rainfall over Japan and Korea in the second half of the year.

The timing of the climatological onset of the monsoon

season is reasonably captured, although the rainfall does

not increase as rapidly as the observations. The changes of

wind, OLR and geopotential height before and after mon-

soon onset are generally captured. However, after monsoon

onset CCAM produces much stronger wind changes than

NCEP over the Indian Ocean southwards of 10N, both at

850 and 200 hPa. CCAM experiences some difficulties in

capturing these changes over Japan, central and southern

China and the South China Sea (Fig. 10b, c). The simula-

tion generally captures the characteristic northward jump

of rainfall of the Asian summer monsoon over regions

spanning 70–140E. It also skillfully reproduces the active-

break-revival behaviour for the southeast China region,

India and probably Japan. However, it has some difficulty

in capturing the rainfall distribution over the South China

Sea during early and late months. Also, CCAM produces

insufficient rainfall along the equator during the monsoon

season (Fig. 12).

Acknowledgments The authors are grateful to Dr. Huqiang Zhang,

Prof. Murari Lal and the anonymous reviewers for their constructive

comments. The observed rainfall was provided by ftp://ftp.cpc.ncep.

noaa.gov/precip/cmap/pentad and ftp://ftp.cpc.ncep.noaa.gov/precip/

cmap/monthly.

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Modelling the Asian summer monsoon using CCAM