Journal of the Meteorological Society of Japan, Vol. 81, No. 5, pp. 909--933, 2003

909

An Evolution of the Asian Summer Monsoon Associated with Mountain
Uplift—Simulation with the MRI Atmosphere-Ocean Coupled GCM—

Manabu ABE
Doctoral program in Geoscience, University of Tsukuba, Tsukuba, Japan

Akio KITOH
Meteorological Research Institute, Tsukuba, Japanand
Tetsuzo YASUNARI1
Institute of Geoscience, University of Tsukuba, Tsukuba, Japan(Manuscript received 10 September 2002, in revised form 3 April 2003)

Abstract
Using the MRI global atmosphere-ocean coupled general circulation model, we had six simulations
with different mountain heights, i.e., 0% (M0), 20% (M2), 40% (M4), 60% (M6), 80% (M8), and 100% (M,control run) of the present global orography, respectively, to study climate changes due to progressive
mountain uplift. The changes of the Asian summer monsoon, with progressive mountain uplift is studied
in this paper.
An active convection region extends with mountain uplift to form a moist climate in South and East
Asia. Monsoon circulation such as low-level westerly, and upper-level anticyclonic circulation, is alsoenhanced with mountain uplift. The increase in precipitation, and the enhancement of southwesterly, inthe later stages of the mountain uplift, appear only over India and the south and southeastern slope of
the Tibetan Plateau. Over the coastal region of Southeast and East Asia, where the maximum precipitation appears in M0, precipitation decreases gradually with mountain uplift, and the southwesterly inthe later stages becomes weaker. In the connection with these changes, surface heat flux changes remarkably over moist Asia in the earlier stages of mountain uplift, compared with that in the later stages.
The intensity of the Indian, Southeast Asian, and East Asian monsoon was investigated with indiceswhich are defined by area mean precipitation. The Indian monsoon becomes strong gradually withmountain uplift; particularly, in the later stages, the remarkable enhancement is found. The intensity of
the South Asian monsoon is the strongest in M4. Thus, in the later stages of mountain uplift, that becomes weaker in association with the northwestward migration of the convective activity. Although theEast Asian monsoon is enhanced gradually with mountain uplift, the enhancement in the earlier stages

1

Corresponding author and present affiliation:
Manabu Abe, Hydrospheric Atmospheric Research Center, Nagoya University, Furo-cho,
Chikusa-ku, Nagoya, 464-8601, Japan.
E-mail: mabe@ihas.nagoya-u.ac.jp
Present affiliation: Hydrospheric AtmosphericResearch Center, Nagoya University, Nagoya,Japan.
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is larger than that in the later stages. In the equatorial Indian Ocean, SST also increases with mountain
uplift, resulting in the increase in precipitation. The increase in SST results from the change of the ocean
surface dynamics due to the enhanced monsoon circulation. This result could not be obtained if CGCM
was not used in this study.

1.

Introduction

Large-scale orography, such as the Tibetan
Plateau and the Rocky mountains, plays important roles to form the present global climate
system (Kutzbach et al. 1993). The large-scale
orography has dynamical and thermodynamical effects to the atmosphere, which in turninfluences the ocean circulation through theatmosphere-ocean interaction. For the Asiansummer monsoon, the Tibetan Plateau is very
important to induce and maintain the monsoon
circulation, but the effects and the role of the
Tibetan Plateau is not understood fully on theintraseasonal, interseasonal, annual, and interannual variability in present. Furthermore,
the mountain uplift occurred in the geologicalhistory must be one of the cause of the past climate change (Kutzbach et al. 1993).
Many previous studies have investigated theeffect of mountains on the Asian summer monsoon using a climate model. Most of the previous model studies conducted a comparison between mountain-run (M) and no-mountain run(NM) using an atmospheric general circulation
model (AGCM) (Manabe and Terpstra 1974;
Hahn and Manabe 1975; Tokioka and Noda
1986; Broccoli and Manabe 1992), or an atmosphere-ocean coupled GCM (CGCM) (Kitoh
1997; Kitoh 2002). However, these studies
could not reveal to what degree the change ofmountain height affect a climate and monsoon.
It can also be postulated that the effect ofmountains does not necessarily change linearlywith increasing mountain altitudes. Kutzbach
et al. (1989), by using AGCM conducted threeexperiments, i.e., no-mountain run (NM), halfmountain run (HM), and mountain run (M),to investigate how the components of monsoon
changed linearly, or non-linearly, associatedwith progressive mountain uplift. Their study
concluded that atmospheric heating rates,
mid-tropospheric vertical motion, and uppertropospheric planetary wave amplitude varied
approximately linearly with progressive mountain uplift. However, as the change from NM to

HM, and that from HM to M, were not investigated in their study, experimental design
with the three stages of mountain uplift intheir study was not sufficient to investigate thecommencement of the Asian summer monsoon
as is in present or the detailed change in a climate by progressive mountain uplift. Thus, itis necessary to study the evolution of the Asiansummer monsoon considering the changes
among the three stages set in Kutzbach et al.
(1989). Recently, in Liu and Yin (2002), 11 experiments, in which the different heights of the
Tibetan Plateau were prescribed, were conducted to examine the impact of progressive
uplift of the Tibetan Plateau on the evolution of
the East Asian monsoon using AGCM. However, their study did not take account of theocean circulation changes due to mountain uplift. The Asian summer monsoon may play anactive role in the interannual variability of theatmosphere-ocean coupled system in the tropical Pacific through the east-west circulation(Yasunari 1990). Kutzbach et al. (1993) andKitoh (1997) have made experiments using an
atmospheric general circulation model with a
50 m slab ocean model. In their studies, sea
surface temperature (SST) distribution near
the continents in mountain run was different
from that in no-mountain run. Moreover, Kitoh(2002) investigated the surface climate changes
with mountain uplift using CGCM. Kitoh(2002) showed the decrease in SST with mountain uplift, particularly over the subtropicaleastern Pacific ocean, due to less shortwave
radiation into surface by increased low-levelclouds and more evaporation by an enhanced
trade wind. In the real climate system, thepresence of mountain is important to form
SST distribution, and the oceanic circulation as
well as the atmospheric circulation. Althoughstudies have been made on the effects of mountains using GCM, there has been few studies
using CGCM that tried to reveal the effects ofmountains taking account of mountain heighton the climate system including the ocean system. The changing oceanic general circulation

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M. ABE, A. KITOH and T. YASUNARI

should be considered for fully understanding
the effects of orography on climate formation.
Thus, systematic experiments using the Meteorological Research Institute (MRI) CGCM aremade to investigate the climate changes due toprogressive mountain uplift more in detail.
For geological climate changes related tomountain uplift, it is difficult to deduce the influence of mountain uplift only from geologicalevidences. Thus, the quantitative evaluation ofclimate change related to mountain uplift usinga climate model can supply significant information and theoretical interpretations for the
past climate changes due to mountain uplift.
In this paper, we focus on the changes in the
Asian summer monsoon. Since CGCM is usedin this study, experiments in this study differ
essentially from experiments in Liu and Yin(2002) using AGCM, although the different
heights of mountain were also set in this study.
Further, our preparing paper investigates theinfluences of mountain uplift to the Pacific andIndian ocean sector. Although the global orography is changed in this study, other experiments to examine the sole effects of the Tibetan Plateau, and the Rocky mountains are also
conducted.
Section 2 describes the model and experimental design. The changes of the Asian monsoon with progressive mountain uplift are
shown in Section 3. Section 4 shows the change
in the intensity of the Asian summer monsoon
indicated by indices.
2.

Model and experimental design

The model used in this study is a global
ocean-atmosphere coupled general circulation
model which is a grid point model developed at
MRI (MRI-CGCM1, Tokioka et al. 1995). The
horizontal resolution of the atmospheric modelis 5 longitude and 4 latitude and the verticalresolution is 15 layers with an atmospheric top
at 1 hPa. The calculations of shortwave radiation are based on the parameterization of Lacis
and Hansen (1974). Rayleigh scattering andabsorption by ozone and water vapor are calculated. Longwave radiation is calculated basedon the multiparameter random model of Shibata and Aoki (1989). The model uses a parameterization of the modified Arakawa-Schubert
penetrative cumulus convection (Tokioka et al.
1988). The calculation of orographic gravity-

911

wave drag follows Palmer et al. (1986). The
land surface has 4 layers with a bottom at 10 m
depth. Ground temperature, ground wetness
and frozen soil moisture are prognostic variables calculated at each level of the land surface. The effects of vegetation are not explicitly
considered on the land surface processes. The
details of the performance of the AGCM aredescribed by Kitoh et al. (1995).
The global oceanic model is developed at MRI(Nagai et al. 1992), which contains 21 verticallevels, 11 of which are located between 300 m
depths and the oceanic surface. A realistic bottom topography is included in the model. The
ocean general circulation model (OGCM) has aresolution of 2:5 longitude and 2:0 latitude.
To resolve the equatorial waves and the narrow
equatorial upwelling region sufficiently, however, non-uniform meridional resolution from
0:5 to 2:0 is introduced within the equatorial region of 12 S–12 N. The simulation of
the oceanic mixed layer is calculated following
the Mellor-Yamada level 2 turbulence closure
scheme.
In order to produce realistic SST, and sea
surface salinity fields, the flux adjustment forsurface energy and water flux is operated. Although the original model predicts sea ice concentration and its depth, in this study, we
specified the sea ice conditions in their monthly
climatology based on the observations.
Figure 1 shows the seasonal mean distribution of precipitation and SST during June–
August in control run (M). The SST pattern isreproduced reasonably since the flux adjustment is working well. A warm pool in the western Pacific, and a cold tongue in the eastern
Pacific are well simulated. Large-scale precipitation distribution agrees with the observed
climatology. There are also the Intertropical
Convergence Zone (ITCZ), and the South Pacific Convergence Zone (SPCZ) in the Pacific.
The results of the climatological state andinterannual variability of the Asian monsoon
and atmosphere-ocean system over the Pacificin this model were investigated by Kitoh et al.
(1999). In their study, the overall features ofthe model climatology and variability of the
Asian summer monsoon in the CGCM were
found to be close to the observed one. The simulated SST and its variability in the CGCM
showed some bias compared to the observation.

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Fig. 2. Orography used in the control run(M). Unit is 1000 m and the contours
are 0.5, 1, 2, 3, and 4.

Fig. 1. (a) Seasonal mean precipitation
during June–August in M. Unit is mm
day 1 and the contour interval is 2 mm
day 1 . The values above 4 mm day 1
are shaded. (b) Seasonal mean SST
during June–August in M. Unit is  C
and the contour interval is 1 C. The
values above 28 C are shaded.

The seasonal change of wind direction and convections over the monsoon region in the model
agrees well with those of the observation. Theinterdecadal variability in the Pacific sector in
the model was found to be realistic (Yukimoto
et al. 1996).
In order to investigate the evolution of the
Asian summer monsoon with progressive
mountain uplift, six runs were performed,
which were integrated for 50 years respectively. Control run with the standard present
orography is designated as the M run. The top
elevation of the Tibetan Plateau is 4400 m in M
(Fig. 2). M0 run, M2 run, M4 run, M6 run andM8 run were conducted under the boundary

conditions of 0%, 20%, 40%, 60%, and 80% of
the global orography in M, respectively. Landsea distribution is the same in all the runs. All
the runs started from the same initial condition
with the orography height in M run. In order toavoid a calculation error due to a rapid changeof the boundary condition, the height of orography in runs, except M run, was decreased
10% of standard present orography by a month.
After the mountain height in each run reached
the designed height, the integrations were continued with the fixed mountain height. Since
the true climatological states of SST, and surface salinity in runs, except M run, are notknown, the adjustment values for flux adjustment in M were regarded as the deficientvalues for exchanges of heat energy and fresh
water flux between the atmosphere and theocean in the CGCM in this study. Thus, the flux
adjustment with the same adjustment data in
M was also operated in other runs.
Figure 3 shows the time series of annual
mean sea water temperature averaged vertically between surface and 300 m depth over theequatorial Pacific Ocean (150 E–80 W, 10 S–
10 N). The averaged temperature in runs, except M, also reaches equilibrium within about
10 years, although its interannual variability
appears for 50 years in all the runs.
M0 run has a flat surface everywhere and is
the same as NM run in Kitoh (2002). Seasonal mean data averaged for the last 30 years(21–50) integrations are used for analysis in
this study.

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M. ABE, A. KITOH and T. YASUNARI

Fig. 3. Time series of annual mean sea
water temperature averaged verticallybetween surface and 300 m depth in theequatorial Pacific (150 E–80 W, 10 S–
10 N). Unit is  C.

3.

Changes with mountain uplift

3.1 Rainfall and low-level circulation
Figure 4 shows the distribution of precipitation and wind at 850 hPa averaged from June
to August over Asia in M0, M2, M4, M6, M8,
and M. The difference of precipitation, andwind at 850 hPa between M2 and M0, M6 andM4, and, M and M8, are shown in Fig. 5. Precipitation in M0 (with land-sea distribution,
but without mountains) is shown in Fig. 4a.
Although all mountains are removed, active
convections are found from west Africa throughthe Indochina Peninsula to the East China Sea
in M0, which can be identified as summer
monsoon rainfall. A convection center over Asia
in M0 is located over the Bay of Bengal, theIndochina Peninsula and the Arabian Sea with
precipitation above 10 mm day 1 . Monsoon
westerly centered at 10 N also appears overSouth Asia in M0. The maximum zonal windvelocity is located over the Bay of Bengal in M0,
which coincides with the center of active convection in M0. Further southwesterly fromSouth Asia blows through Japan to the North
Pacific. Over the equatorial Indian Ocean,cross-equatorial wind from the southern hemisphere to the northern hemisphere appears.
However, the Somali Jet is not found clearly,
but the maximum meridional wind velocity is
located over the eastern Indian Ocean. Further,
the small amount of precipitation is seen innorthern India, where northerly blows to converge with southwesterly at about 10 N to form
monsoon westerly in M0. Besides, there is relatively strong mid-latitude westerly over 40 N–

913

60 N of Eurasia in M0. Precipitation over theland within 20 N–40 N, where the Tibetan
Plateau and a dry climate region are located in
the present, is also small. In the region over
and around the Caspian Sea, however, the climate is more rainy compared to that south ofthere.
In M2 (Fig. 4b), some differences from M0
appear. Precipitation in M2 increases in eastern China, to the east of 70 E, and Mongolia,
compared with that in M0 (Fig. 5a). One canfind that the area of precipitation above 4 mm
day 1 extends into the interior in East Asia,
as Fig. 5a shows positive values above 1 mm
day 1 in the whole of East Asia. Furthermore,
the increase in precipitation is found over the
Arabian Sea, and the Bay of Bengal and Myanmar region in Fig. 5a. Southwesterly over
the Arabian Sea and cross-equator flow over
the western Indian Ocean become strong, andsouthwesterly over northwestern India is alsostronger in M2 than in M0. The changes of flow
at low-level enhance water vapor flux into theland area over South and East Asia to increase
precipitation. The precipitation increase, andthe circulation changes from M0 to M4, also
become more obvious in regions where the increase in precipitation is found in M2. The
migration of active convections from the Bayof Bengal to northern India is also foundover northwest India from M0 to M4. Over thewestern Tibetan Plateau, northwesterly in M0changes to northerly in M4, which flows into
northern India as northwesterly. Further,moist southwesterly also intrudes into northern India and combines with dry northwesterly over India, which may be related to the increase in precipitation over India.
In M6 (Fig. 4d), a significant difference from
M4 appears over the Bay of Bengal and the
Indochina Peninsula, where precipitation decreases (Fig. 5b). In northern India, the areaof precipitation above 4 mm day 1 continues toexpand from southeast to northwest to result inincreased precipitation (Fig. 5b).
In M8 (Fig. 4e) and M (Fig. 4f ), precipitation
decreases further over the Bay of Bengal andEast Asia (Fig. 5c). In addition, the decreaseof precipitation is found over the Arabian Sea,and the South China Sea from M8 to M. On theother hand, the northwestward migration ofactive convection continues in northwestern

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Fig. 4. Seasonal mean precipitation and wind at 850 hPa during June–August in each run. Units ofprecipitation are mm day 1 and the contour interval is 4 mm day 1 . The values above 4 mm day 1
are shaded. Units of wind are m sec 1 .

India in M8 and M. Thus, an active convection
region extends northward in India and the Bayof Bengal region with mountain uplift. From
M8 to M, precipitation increases in the dry re-gion of northwestern India, and the west of the
Tibetan Plateau (Fig. 5c). Monsoon westerly
over South Asia becomes stronger in M6 compared with that in M4 (Fig. 5b). The maximum

October 2003

M. ABE, A. KITOH and T. YASUNARI

Fig. 5. Distribution of difference of seasonal mean precipitation and wind at
850 hPa during June–August between(a) M2 and M0, (b) M6 and M4, and (c)M and M8. Units of precipitation are
mm day 1 and the contour interval is
1 mm day 1 . The positive values (increased precipitation) are shaded. Unitsof wind are m sec 1 .

wind speed of the monsoon westerly appears
over the Arabian Sea in M6, while it is locatedover the Bay of Bengal in M0 (Fig. 6). Over the
Bay of Bengal and the Indochina Peninsula,monsoon westerly becomes weak from M6 to M,
which coincides with the decrease in precipitation, while it remains stronger and moves
northward over the Arabian Sea. Figure 5cclearly shows westward wind anomalies fromthe Bay of Bengal through the Indochina Peninsula to the western Pacific. These wind
anomalies indicate that monsoon westerly becomes weaker over the Bay of Bengal and the
Indochina Peninsula, and that trade wind over
the western Pacific is enhanced with mountain
uplift. However, as shown in Fig. 5a and 5b, thewesterly is enhanced over the north of the Bayof Bengal and the Indochina Peninsula from

Fig. 6. Longitudinal distribution of seasonal mean zonal wind speed at
850 hPa during June–August averaged
at 6 N–18 N in each run. Units is
m s 1.

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Fig. 7. Distribution of seasonal mean sea level pressure during June–August. Units are hPa. The
values above 1020 hPa are shaded. For easy comparison, constant value is added for each experiment so that global mean values become the same as in M.

M0 to M8. These changes indicate the centerof monsoon southwesterly moves northwestward into the interior with mountain uplift inSoutheast and East Asia, coinciding with thenorthwestward expansion of an active convection region over Southeast and East Asia.
Moreover, the flow over the Bay of Bengal has
a strong northward component between M6
and M than that in the earlier mountain upliftstages. In this connection, the southwesterly
in Northeast Asia moves into the interior, andflows from the Pacific through Japan become
obvious between M6 and M, as the easterly

wind anomaly from the Pacific exists overJapan as shown in Fig. 5c. These results overSoutheast and East Asia may be related to an
enhanced subtropical high over the Pacific.
Figure 7 shows the seasonal mean sea levelpressure during June–August in each run. Although the atmospheric mass was kept theconstant value in all the runs, a reduction to
the mean sea level pressure results in different
global mean sea level pressure values. For easycomparison, a constant value is added to sea
level pressure in each run so that the global
mean values become the same as in M in theOctober 2003

M. ABE, A. KITOH and T. YASUNARI

figures. It is found obviously that the subtropical high over the Pacific is enhanced withmountain uplift. However, this change of thesubtropical high is not affected only by Asian
mountains, but also by mountains in other regions, e.g., the Rocky mountains (Hoskins 1996;
Rodwell and Hoskins 2001). The enhanced subtropical high with mountain uplift causes astrong easterly wind from the western Pacific,
particularly in M (Fig. 6). The enhanced subtropical high over the north Pacific results ina stronger trade wind over the Pacific and enhanced southeasterly from the western Pacificover Southeast Asia. The northwestward movement of monsoon southwesterly over East Asia
may be also affected by the enhanced subtropical high with mountain uplift. Although theSomali Jet in the model intrudes into the East
African continent, the enhancement of it withmountain uplift is found in Fig. 4. In brief,
while the monsoon westerly is enhanced overSouth Asia from M0 to M4, the westerly from
M6 to M is enhanced only over the Arabian Sea
and the southern slope of the Tibetan Plateau,
and weakened over the Bay of Bengal and the
Indochina Peninsula.
Incidentally, weak westerly disappears overthe land in the mid-latitude between M6 and
M. The gradual extension of a small-precipitation region is also found in the north of Central
Asia with mountain uplift. This result is consistent with a result by Broccoli and Manabe(1992), although precipitation increases from
M8 to M in a part of Central Asia. In the equatorial Indian Ocean, precipitation increases
with mountain uplift as shown by Fig. 5.
The northwestward extension of an active
convection region is examined more clearly.
Figure 8 shows the meridional distribution of
precipitation averaged from 75 E to 85 E (lineA-B drawn in Fig. 8a) in each experiment, andthat in precipitation along line C-D drawn in
Fig. 8a, which was drawn from Central Asia to
Southeast Asia. In Fig. 8b, extremely heavy
precipitation at 10 N in M0 is found. While theprecipitation at 10 N decreases gradually withmountain uplift, precipitation at broad latitudeof 14 N–26 N increases significantly. Further,
the latitude at which the extreme increase of
precipitation occurs among each consecutive
mountain uplift, moves north within 14 N–
26 N with progressive mountain uplift. In the917

later stages of mountain uplift, precipitation
also increases particularly in India. In Fig. 8c,
the maximum precipitation in M0 is locatedat 18 N. Precipitation over the north of the Indochina Peninsula and the southern slope of
the Tibetan Plateau increases with mountain
uplift. The precipitation over the Indochina
Peninsula increases as whole from M0 to M4,
particularly over the north of the peninsula.
However, while the precipitation over the northof the peninsula and the slope of the Tibetan
Plateau increases gradually from M6 to M, thatover the Indochina Peninsula decreases, whichindicates an active convection region intrudes
northwestward into the interior in the laterstages of mountain uplift.
3.2

Expansion of moist Asian climate toward
the interior
Figure 9 shows the difference of mean sea
level pressure (SLP) between the successive
two stages (M2-M0, M4-M2, M6-M4, M8-M6,
and M-M8). As stated above, since the atmospheric mass was kept the constant value in all
the runs, a reduction to the SLP results in different global SLP values. The difference is calculated using the value which is modified byadding the constant value to SLP in each run sothat its global mean value becomes the same asin M. In Fig. 9a and 9b, the decrease of SLP
occurs over the Tibetan Plateau, and the westof the Tibetan Plateau including the Arabian
Peninsula. On the other hand, over the east of
the Tibetan Plateau, the small increase in SLP
from M0 to M4 is found. From M6 to M, while
SLP over the Tibetan Plateau and the west of
the Plateau decreases gradually, the increase
in SLP appears over India and the coastal region of East Asia. These SLP changes are consistent with the northwestward migration ofsouthwesterly at 850 hPa over Southeast andEast Asia with mountain uplift. While thesouthwesterly blowing from India to East Asia
is enhanced over the slope of the Tibetan Plateau, the flow becomes weaker with mountain
uplift in the southern side of the southwesterly.
Therefore, precipitation over the southern partof the Bay of Bengal and the Indochina Peninsula, which is larger in M0, decreases withmountain uplift. Further, the southwesterly
over the Arabian Sea is enhanced over thecoastal region due to the gradual decrease in918

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Fig. 8. The meridional distribution of seasonal mean precipitation during June–August averaged
from 75 E to 85 E (A-B) in each run and that along C-D (from Central to Southeast Asia). The abscissa (ordinate) indicates precipitation (latitude). Units of precipitation are mm day 1 . In (c), thevalues on the line C-D, which are averaged for three grids (cross) on normal line to the line C-D,
are plotted.

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Fig. 9. (a) Seasonal mean sea level pressure (SLP) in M0 and the difference of seasonal mean SLP
during June–August between (b) M2 and M0, (c) M4 and M2, (d) M6 and M4, (e) M8 and M6 and(f ) M and M8. Units are hPa. The positive values are shaded. The difference is calculated using thevalue which is modified by adding the constant value to SLP in each run so that its global meanvalue becomes the same as in M.

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Journal of the Meteorological Society of JapanSLP over the Arabian Peninsula and the IranAfghanistan region. The northward migrationof the southwesterly in low-level also transports water vapor into India, resulting in theincrease in precipitation over India.

Vol. 81, No. 5

The increase in SLP is found over East andSouth Asia (Fig. 9), coinciding with a region
where surface air temperature decreases, andground wetness increases over East and South
Asia (Fig. 10). The large increase in groundFig. 10. (a) Seasonal mean ground wetness in M0 and the difference of seasonal mean ground wetness during June–August between (b) M2 and M0, (c) M4 and M2, (d) M6 and M4, (e) M8 and M6,and (f ) M and M8. Units are %. The positive values are shaded.

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wetness over Northeast and East Asia from M0
to M4, and that over North India from M8 to M
are induced by the increase in precipitation.
Therefore, the increase in SLP over moist Asia
may be associated with reduced surface temperature by the increase in precipitation. In
addition, the increase in SLP over the coastal
region of East Asia may be affected by the enhanced subtropical high over the Pacific.
Figure 11 shows the difference of seasonalmean latent heat flux and sensible heat flux at

921

surface between M2 and M0, M6 and M4, and
M and M8, respectively. The changes of both
fluxes with mountain uplift show the opposite
sign in each other. Latent heat flux at surface
increases over Northeast Asia and East Asia,
particularly in Mongolia, while sensible heat
flux at surface decreases about two third of thelatent heat flux over the same region. Both heat
flux changes also coincide with the increase ofground wetness. Further, the change of bothheat flux in other stages of mountain upliftFig. 11. Difference of seasonal mean latent and sensible heat flux at surface during June–August
between (a) (d) M2 and M0, (b) (e) M6 and M4, and (c) (f ) M and M8. Units are W m 2 . The positive
values are shaded.

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Fig. 12. Difference of seasonal mean shortwave radiation flux at surface and cloud amount during
June–August between (a) (d) M2 and M0, (b) (e) M6 and M4, and (c) (f ) M and M8. Units of shortwave radiation are W m 2 . The positive values are shaded.

agrees with the increase in ground wetness.
From M6 to M, latent (sensible) heat flux
over India increases (decreases), although thesevalues are smaller than those in the earlier
stages of mountain uplift. Further, shortwave
radiation flux at surface over the moist Asia
region decreases due to cloud increased withmountain uplift, as shown in Fig. 12. The decrease in shortwave radiation in East Asia isremarkable from M0 to M2 (Fig. 12a). In Central Asia, however, shortwave radiation in-creases with mountain uplift due to the decrease in cloud amount, particularly, in thelater stages of the mountain uplift. In northern
India, the large decrease in shortwave radiation appears from M8 to M (Fig. 12c).
Figure 13 shows the change of latent heat,sensible heat, shortwave radiation, and longwave radiation flux at surface with mountain
uplift in India (75 E–80 E, 10 N–30 N), Southeast Asia (95 E–105 E, 10 N–30 N), East Asia
(105 E–120 E, 30 N–45 N), and Central Asia

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Fig. 13. The change of latent heat, sensible heat, shortwave radiation, and longwave radiation flux
at surface with mountain uplift in (a) India (75 E–80 E, 10 N–30 N), (b) Southeast Asia (95 E–
105 E, 10 N–30 N), (c) East Asia (105 E–120 E, 30 N–45 N), and (d) Central Asia (40 E–65 E,
40 N–50 N). Units are W m 2 . Open circle (square) indicates latent (sensible) heat flux at surface,and closed circle (square) indicates shortwave (longwave) radiation flux at surface. The positive
values of shortwave radiation flux at surface indicate a downward direction, and that of otherfluxes an upward direction.

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Journal of the Meteorological Society of Japan(40 E–65 E, 40 N–50 N), respectively. Thechange of shortwave radiation flux with mountain uplift is consistent with that of cloudamount. In connection with the migration of
the active convection region, latent heat flux
over India increases gradually with mountain
uplift, while sensible heat, shortwave radiation,and longwave radiation decrease. Shortwave
radiation flux over Southeast and East Asia
decreases with mountain uplift, particularly inthe earlier stages of mountain uplift (Fig. 13b
and Fig. 13c). Sensible heat and longwave radiation flux also decrease largely over Southeast and East Asia in the earlier uplift stages.
These changes are consistent with the changes
in the earlier uplift stages that ground wetness
over Southeast and East Asia increases largely(Fig. 10), and that surface temperature decreases. The increase in shortwave radiation,
longwave radiation and sensible heat flux overCentral Asia begins to appear from M4 to M6(Fig. 13d). These changes are outstanding in
the later stages of mountain uplift, as well as
the decrease in those fluxes over South andEast Asia. Further, these changes are consistent with the change of ground (surface air)temperature and ground wetness.
These results stated above indicate surface
air (ground) temperature is affected significantly by ground wetness and surface heat fluxchange induced by the increased precipitation.
With mountain uplift, the decrease (increase)in shortwave radiation flux into surface (cloudamount) coincides with the increase of precipitation, as well as the increase in latent heat
flux. Therefore, ground surface temperature
over moist Asia decreases with mountain uplift
due to the decrease in shortwave radiation into
surface by cloud effect, and the wetter ground,in addition to the rise of ground surface.
3.3

SST change over the tropical Indian
Ocean
Figure 14 shows SST in the Indian Ocean
averaged within June–August in M0, M4, andM8 and its difference in M2, M6, and M8 fromthat in M0, respectively. In M0 (Fig. 14a), SST
in the northern hemisphere is larger than 28 C.
From M0 to M2 (Fig. 14b), SST over the Arabian Sea and the Bay of Bengal decreases dueto the decrease in shortwave radiation flux into
surface by cloud effect and the increase in la-Vol. 81, No. 5

tent heat flux (Fig. 12a). In M6 (Fig. 14d), SST
in the Bay of Bengal also increases due to decreased shortwave radiation flux associatedwith the change of cloud amount. On the other
hand, SST in the equatorial Indian Ocean increases with mountain uplift; particularly large
increase is found in the eastern equatorial Indian Ocean. Although southerly wind at lowlevel across the equator in M0 is stronger overthe eastern Indian Ocean, the southerly windover the eastern equatorial Indian Ocean becomes weaker, and southerly wind over thewestern becomes stronger with mountain uplift. Thus, these changes are related to theweakness of the upwelling in the eastern Indian Ocean. The enhanced southerly wind over
the western Indian Ocean causes the enhancement of upwelling in the western Indian Ocean,
and a western part of the Arabian Sea. Further,
the westerly component of sea surface windstresses in the equatorial Indian Ocean becomes larger with mountain uplift. This change
also causes the weakness of a sea surface current toward poleward, related to the increase in
SST in the equatorial Indian Ocean. The increase in SST also results in the increase of
precipitation in the equatorial Indian Ocean, as
shown in Fig. 5. Both shortwave radiation flux,
and latent heat flux in the equatorial Indian
Ocean, decrease with mountain uplift. Therefore, the increase in SST in the equatorial Indian Ocean results from the change of theocean dynamics. Using CGCM in this study, we
can describe the increase of precipitation consistent with the increase in SST over the equatorial Indian Ocean, as well as the change in
SST. In Chandrasekar and Kitoh (1998), thepresence of positive SST anomalies south of theequator over the Indian Ocean resulted in thedecrease in precipitation over Indian Ocean,and the weakening of the mean monsoon circulation. Therefore, the increase in SST in theequatorial Indian Ocean with mountain upliftmust affect the change of the monsoon precipitation and circulation over South and East
Asia. If using AGCM with observed SST,
warmer SST in the equatorial Indian Ocean
may affect the monsoon precipitation in M0.
This result indicates that summer SST distribution in the Indian Ocean is affected significantly by the evolution of the Asian summer
monsoon. The effect of mountain uplift should

October 2003

M. ABE, A. KITOH and T. YASUNARI

Fig. 14. Seasonal mean SST and surface wind stress in (a) M0, (c) M4 and (e) M8, and its difference
during June–August between (b) M2 and M0, (d) M6 and M4, and (f ) M and M8. Units of SST areC and those of surface wind stress are N m 2 . Shaded area indicates the values above 28 C in (a),(c), and (e), and the positive values in (b), (d), and (f ).

925

926

Journal of the Meteorological Society of Japan

Vol. 81, No. 5

Fig. 15. (a) (c) (e) Seasonal mean distribution of wind at 200 hPa during June–August in M0, M4,
and M8, and (b) (d) (f ) the difference of wind at 200 hPa between M2 and M0, M6 and M4, and M
and M8. Units are m s 1 . The contour shows mountain height at 1500 and 3000 m.

be investigated with coupled GCM including
full ocean processes.
3.4 Upper-level atmospheric circulation
Wind distribution at 200 hPa in M0, M4, andM8, and its difference between M2 and M0, M6
and M4, and, M and M8 are shown in Fig. 15.
In M0 (Fig. 15a), westerly is found zonally inthe whole north of 30 N; particularly, an ex-treme flow called the subtropical jet exists at
30 N over Asia. There is also weak anticyclonic
circulation linked with convective activity overSouth Asia. Further, over the equatorial Indian
Ocean, northeasterly crossing the equator isfound. The important change with mountain
uplift for the circulation at 200 hPa is that anticyclonic circulation is enhanced, and becomesconspicuous over South Asia.

October 2003

M. ABE, A. KITOH and T. YASUNARI

927

Fig. 16. Seasonal mean distribution of mean temperature between 500 hPa and 200 hPa during
June–August in each run. Units are  C and the contour interval is 2 C. Temperature above 26 C
are shaded. The thick contour shows mountain height at 1500 and 3000 m.

As shown by the wind difference between M2
and M0 in Fig. 15b, easterly is enhanced overSouth Asia and the equatorial Indian Ocean.
As anticyclonic circulation anomalies are foundover East Asia, and the South Indian Ocean,
the anticyclonic circulation centered in Myanmar becomes more outstanding in M2. In M4,

anticyclonic circulation over the whole of South
Asia is enhanced. Thus, the subtropical jet isdiverted north over East Asia, and the northeasterly over the equatorial Indian Ocean isstronger than that in M0.
Figure 16 shows the seasonal mean distribution of mean temperature between 500 hPa and928

Journal of the Meteorological Society of Japan

200 hPa during June–August, averaged for thelast 30 years. In M0 (Fig. 16a), the temperature
over the low latitude region is above 28 C. A
warm area above 26 C in M0, also exits overSouth Asia and the North Indian Ocean, but
the reversal of the meridional temperature
gradient between the continents and the Indian
Ocean is not distinguishable (Fig. 16a). Although the area above 26 C is smaller in M2(Fig. 16b) and M4 (Fig. 16c) than in M0, the reversal of the temperature gradient from theequatorial Indian Ocean to South Asia, appears
over South Asia. Figure 17 shows the change ofthe north-south and the east-west land-sea
contrast of mean temperature between 500 hPa
and 200 hPa, with mountain uplift. Figure 17aFig. 17. Change of (a) the north-southand (b) west-east land-sea contrast ofmean temperature between 500 hPa
and 200 hPa with mountain uplift. (a)shows the difference of the temperature
between 30 N and the equator, and (b)the difference between 90 E and 130 E.
Units are  C.

Vol. 81, No. 5

shows the difference of the temperature between 30 N and the equator, and Fig. 17b the
difference between 90 E and 130 E. The northsouth contrast over South Asia is negative in
M0. The reversal of the contrast from negative
to positive appears over the eastern continent
in the earlier stages, the reversal is also foundover the entire South Asia region in the laterstages. The maximum north-south contrast exists in the earlier stages at about 110 E, butthat in M8 and in M, it changes to 95 E. Particularly, the north-south contrast between theplateau and sea increases largely in the laterstages. The east-west contrast over East Asia
affects significantly East Asian monsoon. In M0,
the east-west contrast of the temperature issmaller over East Asia, while the contrast overSoutheast Asia is positive. In the earlier stages,
the east-west contrast over Southeast Asia isenhanced. While a small change of the contrast
is found from M4 to M6, that over East Asia becomes significantly stronger in the later stages.
The center of the anticyclonic circulation
moves from about (100 E, 20 N) in M0 to about(90 E, 26 N) in M6. While the difference for thecenter of the anticyclonic circulation between
M0 and M4 is not noticeable, that between theearlier stages (M0, M2, M4) and the laterstages (M6, M8, M) is remarkable. Easterly
anomalies over South and East Asia, shown in
Fig. 15d and 15f, indicate that the anticyclonic
circulation moves to the northwest. In the laterstages of mountain uplift, therefore, the centerof the anticyclonic circulation moves graduallyto the top of the Tibetan Plateau. Moreover, as
the anticyclonic circulation extends to Northeast Asia, the northerly is dominant over East
Asia particularly in M8 and M. Over the ocean
to the east of Japan, a westerly wind component also decreases from M6 and M, as shownin Fig. 15d and 15f.
In M8 (Fig. 16e), an area above 28 C appears only over South Asia, and then temperature above 24 C is found over the Tibetan
Plateau. Therefore, the temperature gradientfrom the ocean to the land is outstanding overthe Indian Ocean/South Asia compared withthat in the earlier stages (M0-M4). The eastwest temperature gradient is also found overEast Asia and the northwestern Pacific. In M,
the warmer area above 26 C is found centered over the Tibetan Plateau. Further, theOctober 2003

M. ABE, A. KITOH and T. YASUNARI

temperature gradient from the ocean to thewarm center over the Tibetan Plateau, is noticeable in M compared with that in the earlier
stages of mountain uplift.
Incidentally, the subtropical jet is diverted
and becomes weaker by the development of the
anticyclonic circulation with mountain uplift.
Flows crossing the equator over the Indian
Ocean are enhanced in consistent with the enhancement of the anticyclonic circulation in theearlier stages (Fig. 15b), while a remarkablechange of the circulation in M6 and M8 is notfound (Fig. 15d and 15f ). However, the northeasterly becomes weaker over the eastern Indian Ocean, and the maritime continent in M.
4.

Intensity of the Asian summer
monsoonAs shown by figures used in the previous
section, the various local changes of some meteorological components are found with progressive mountain uplift. In this section,change in the intensity of the Asian summer
monsoon is focused on, using four indices
shown in Fig. 18. They are: the seasonal mean
precipitation during June–August averaged
in India (75 E–80 E, 10 N–30 N) (IMR), inSoutheast Asia (95 E–105 E, 10 N–30 N)(SEA), and in East Asia (105 E–120 E, 30 N–
45 N) (EA), and the seasonal mean zonal wind
shear between 850 hPa and 200 hPa during
June–August averaged in an area (40 E–
100 E, 2 N–18 N) as defined in Webster andYang (1992) (WYI).
The abscissa in this figure indicates the stageof mountain uplift. The intensity of the Asiansummer monsoon indicated by IMR increases
from M0 to M. Moreover, the increase rate from
M0 to M6 is a half of that from M6 to M. While
the migration of the active convection region
over India is slower in the earlier stages ofmountain uplift, that is faster in the laterstages. The standard deviation of the interannual variability of IMR also increases withmountain uplift. SEA index indicates the enhancement in the earlier stages of mountain
uplift and the weakening in the later stages.
Also, the EA index shows weak enhancementin the later stages of mountain uplift, compared with enhancement in the earlier stages.
From M0 to M4, the precipitation area extends

929

largely in the land of South and East Asia andthen precipitation increases in the ArabianSea, the Bay of Bengal and India, as stated in
the previous section (Fig. 4). While the increasein precipitation over the slope of the Tibetan
Plateau and India occurs from M6 to M, precipitation decreases in the southern part of the
Bay of Bengal and the Indochina Peninsula.
The large increase in precipitation are also restricted in a smaller area over the northernpart of the Arabian Sea, India and the slopeof the Tibetan Plateau in the later stages ofmountain uplift. Therefore, the enhancementof SEA and EA index in the earlier stages ofmountain uplift is consistent with the increase
and the areal expansion of large precipitation
over Southeast and East Asia. Further, the decrease in precipitation over the coastal regionof Southeast East Asia is associated with theweakening of SEA and the weaker enhancement of EA in the later stages. However, theincrease in the standard deviation of their interannual variability is small compared with
IMR and WYI.
The intensity of the Asian summer monsoon
indicated by the WYI is shown in Fig. 18d. The
Asian summer monsoon becomes strong from
M0 to M8 as well as that indicated by EA. Also,
the standard deviation of its interannual variability increases with mountain uplift. The
large enhancement of WYI occurs from M0 to
M6, compared with that from M6 to M8, which
differs from the enhancement indicated by
IMR. From M0 to M6, westerly wind speed at
850 hPa increases over the Arabian Sea, India
and the Bay of Bengal. However, weakened
westerly at 850 hPa is found over the Bay of
Bengal and the Indochina Peninsula from M6
to M in association with the northward migration of southwesterly, as stated in the previous
section. At 200 hPa, the easterly is enhanced
over the whole of South Asia and the North Indian Ocean, with progressive mountain uplift.
Thus, the weak enhancement of the Asiansummer monsoon in the later stages of mountain uplift results from the weakened westerly
at low-level over the Bay of Bengal and the Indochina Peninsula.
The changes of the Asian summer monsoon,with progressive mountain uplift, based upon
the changes of four indices. The Indian summer
monsoon becomes strong with mountain uplift.

930

Journal of the Meteorological Society of Japan

Vol. 81, No. 5

Fig. 18. Indices of the Asian summer monsoon: the seasonal mean precipitation during June–
August averaged in (a) India (75 E–80 E, 10 N–30 N) (IMR), in (b) Southeast Asia (95 E–
105 E, 10 N–30 N) (SEA), and in (c) East Asia (105 E–120 E, 30 N–45 N) (EA). (d) The seasonalmean zonal wind shear between 850 hPa and 200 hPa during June–August averaged in an area(40 E–100 E, 2 N–18 N) as defined in Webster and Yang (1992) (WYI). The abscissa in this figure
indicates the stage of the mountain uplifts. The dot-dot-dash line indicates the standard deviationof the interannual variability in each stage.

October 2003

M. ABE, A. KITOH and T. YASUNARI

Also, the Indian sumer monsoon in the laterstages of mountain uplift shows great enhancement, and the migration into northern India ofan active convection region remarkably occurs.
On the contrary, the monsoon circulation indicated by WYI, becomes even weaker from M8 to
M, as decreased precipitation and weakened
monsoon westerly at low level appear in Southand Southeast Asia. The enhancement of theEast Asian monsoon with mountain uplift issimilar to that of the large-scale monsoon circulation defined by WYI index. In Southeast
Asia, however, the monsoon is the strongest in
about a half mountain stage, unlike the Indianand East Asia monsoon. In Liu and Yin (2002),
the change of the intensity of the Indian,Southeast Asian, and Northeast Asian monsoon with mountain uplift were investigatedusing the indices based on the seasonal changeof an averaged surface wind direction and velocity between summer and winter. The changeof the Southeast Asian monsoon with mountain
uplift in their study is similar to that in thisstudy. In their study, however, the enhancement of the Indian monsoon was not found.
5.

Concluding remarks

Rainfall, and low-level westerly representing
the Asian summer monsoon, appear in M0
which has land-sea distribution, but without
orography. Weak anticyclonic circulation is also
found at the upper-level over the north Indian
Ocean in M0. There is northward expansion ofan active convection region, and the monsoon
westerly over South Asia, clearly in the earlier
stages of mountain uplift. Thus, the moist climate region appears over South and East Asia,
even in the lower, than a half of the present
mountain case, while a dry climate region isfound over Iran-Afghanistan region, as is in thepresent. In M0, precipitation over South Asia
is concentrated at about 10 N and quite large.
In the earlier stages of mountain uplift, precipitation over South, Southeast and East Asia
increases quantitatively. Its region migrates
northward spatially, with mountain uplift. In
the later stages of mountain uplift, precipitation migrates largely into the south and southeastern slope of the Tibetan Plateau and North
India, but decreases particularly in the Bay of
Bengal and the Indochina Peninsula. In association with the quantitative and spatial

931

change in precipitation, ground wetness increases largely in the earlier stages of mountain uplift, in which the dry ground surface inM0 becomes wet ground surface over South andEast Asia. These changes change surface heat
flux in the land over Southeast and East Asia
significantly in the earlier stages. However, inthe later stages of mountain uplift, groundwetness, and surface heat flux have smallerchanges, because the wetter ground surface already has presented over Southeast and East
Asia, under the condition with the mountainheight lower than the half. Ground temperature also decreases further due to the wetter
ground surface and the decrease in shortwave
radiation into surface by cloud effect, in addition to the rise of ground surface. Further, SST
in the Arabian Sea and the Bay of Bengal, decreases with mountain uplift, due to the reduction of shortwave radiation by cloud effect andthe increase in latent heat flux. These changes
with mountain uplift also contribute to suppress the fall of pressure, or to raise pressure
over South Asia and East Asia. Therefore,
while westerly is accelerated further in the inland side of South and Southeast Asia, that is
weakened with mountain uplift in the oceanside, consistent with the weekend monsoon
westerly over the Bay of Bengal and the Indochina Peninsula. As a result, southwesterly inSoutheast and East Asia migrates northwestward into the interior with mountain uplift,
and precipitation decreases over the coastal
region in Southeast and East Asia. The northwestward movement of southwesterly in Southeast and East Asia may also be affected by theenhanced subtropical high with mountain uplift.
The north-south temperature contrast at
middle-upper level between the land and thesea over South Asia is smaller in the earlier
stages of mountain uplift. In the later stages,
the contrast, particularly, over an area including the Tibetan Plateau, becomes large.
The west-east temperature contrast at middleupper level over Southeast Asia appears in theearlier stages. However, the contrast over East
Asia becomes large with mountain uplift. In M8
and M, the maximum contrast is also found at
30 N.
The intensity of the Indian, Southeast Asian,and East Asian monsoon was investigated with

932

Journal of the Meteorological Society of Japan

indices which were defined by area mean precipitation. Indian monsoon becomes strong
gradually with mountain uplift; particularly inthe later stages its remarkable enhancement isfound. This result differs from the result in Liu
and Yin (2002). The intensity of Southeast
Asian monsoon is the strongest in M4. Thus, inthe later stages of mountain uplift, that becomes weaker in association with the northward migration of convective activity. Although
East Asian monsoon is enhanced graduallywith mountain uplift, the enhancement in theearlier stages is larger than that in the laterstages. Results in Liu and Yin (2002) showedthat Southeast Asia monsoon has intensified
during 0%–70% of the present mountain, but
weakened during 70%–100%. Their result issimilar to our result. However, since their indices were defined by the seasonal change of
surface wind between summer and winter, ourindices are different from theirs. Further, theirexperiments using AGCM is essentially different from ours using CGCM. In the equatorial
Indian Ocean, SST also increases with mountain uplift, resulting in the increase in precipitation. The increase in SST results from thechange of the ocean surface dynamics due tothe enhanced monsoon circulation, and that
could not be obtained if CGCM was not used inthis study.
In this study, the height of global orographychanged in each run. The results in this study,
therefore, also include the effect of mountains
besides the Tibetan Plateau. As the high terrain in the eastern African continent is related
to the evolution of the Somali jet, the jet may beenhanced not only by the Tibetan Plateau uplift, but also by the high terrain. The enhanced
subtropical high with global mountain uplift
also influences the Asian monsoon circulation.
In particular, the Rocky mountains also enhance the subtropical high over the North Pacific (Hoskins 1996; Rodwell and Hoskins 2001).
To investigate the sole effect of Tibetan Plateau
and the Rocky mountains separately, we are
performing other experiments.
The quantitative and spatial changes of aclimate by mountain uplift may depend significantly on physical processes (e.g., the parameterization of cumulus convection) in the model.
Precipitation, for example, shows large variety
among GCMs in the world (Gadgil and SajaniVol. 81, No. 5

1998; Kang et al. 2002), depending upon different parameterization of cloud/precipitation
processes. Therefore, experiments using different models and/or schemes may be required inorder to understand more comprehensively theclimate changes associated with mountain uplift.
Acknowledgments
The authors are grateful to Dr. F. Kimura,
Dr. H.L. Tanaka, Dr. H. Ueda, and graduate
students in the climatology and meteorology
group in the University of Tsukuba for theirvaluable advice. We also thank two anonymous
reviewers for their valuable comments. This
work was done as a co-operative study between
the Institute of Geoscience, University of Tsukuba and the Meteorological Research Institute. It was also partially supported by Large
Scale Numerical Simulation Project of Science
Information Processing Center, University ofTsukuba.
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