Climate Dynamics (2003) 21: 11–25
DOI 10.1007/s00382-003-0315-7
C. Chou
Land–sea heating contrast in an idealized Asian summer monsoon
Received: 18 January 2002 / Accepted: 30 December 2002 / Published online: 4 March 2003
Springer-Verlag 2003
Abstract Mechanisms determining the tropospheric
temperature gradient that is related to the intensity of
the Asian summer monsoon are examined in an intermediate atmospheric model coupled with a mixed-layer
ocean and a simple land surface model with an idealized
Afro–Eurasian continent and no physical topography.
These include processes involving in the influence of
the Eurasian continent, thermal effects of the Tibetan
Plateau and effects of sea surface temperature. The
mechanical effect on the large-scale flow induced by the
Plateau is not included in this study. The idealized land–
sea geometry without topography induces a positive
meridional tropospheric temperature gradient thus a
weak Asian summer monsoon circulation. Higher prescribed heating and weaker surface albedo over Eurasia
and the Tibetan Plateau, which mimic effects of different
land surface processes and the thermal effect of the uplift
of the Tibetan Plateau, strengthens the meridional
temperature gradient, and so as cold tropical SST
anomalies. The strengthened meridional temperature
gradient enhances the Asian summer monsoon circulation and favors the strong convection. The corresponding monsoon rainbelt extends northward and
northeastward and creates variations of the monsoon
rainfall anomalies in different subregions. The surface
albedo over the Tibetan Plateau has a relatively weak
inverse relation with the intensity of the Asian summer
monsoon. The longitudinal gradient of ENSO-like SST
anomalies induces a more complicated pattern of the
tropospheric temperature anomalies. First, the positive
(negative) longitudinal gradient induced by the El Niño
(La Niña)-like SST anomalies weakens (strengthens) the
Walker circulation and the circulation between South
Asia and northern Africa and therefore the intensity of
the Asian summer monsoon, while the corresponding
C. Chou
Environmental Change Research Project,
Institute of Earth Sciences, Academia Sinica,
Taipei, 115, Taiwan
E-mail: [email protected]
monsoon rainbelt extends northward (southward). The
El Niño (La Niña)-like SST anomalies also induces
colder (warmer) tropospheric temperature over Eurasia
and warmer (colder) tropospheric temperature over the
Indian Ocean. The associated negative (positive) meridional gradient of the tropospheric temperature
anomalies is consistent with the existence of the weak
(strong) Asian summer monsoon.
1 Introduction
The Asian summer monsoon has a significant impact on
the regional climate and it also plays an important role
in the global climate system. It has been thought that
land–sea heating contrast is a fundamental mechanism
driving the summer monsoon circulation (e.g. Webster
1987; Young 1987). The relation of the land–sea heating
contrast and the Asian summer monsoon has been discussed in Fu and Fletcher (1985), Meehl (1994a) and Li
and Yanai (1996, LY96 hereafter). They found that the
intensity of the Asian summer monsoon is related to the
strength of the temperature gradient due to the land–sea
heating contrast. The greater the temperature gradient,
the more intense the Asian summer monsoon. Liu and
Yanai (2001) showed that the summer all-India rainfall
is positively correlated to the mean tropospheric temperature at 200–700 hPa averaged over Eurasia in
summer, while the highest correlation appears in western
Eurasia. On the seasonal time-scale, the onset of the
Asian summer monsoon is coincident with the reversal
of meridional upper tropospheric temperature gradient
(LY96).
Among the monsoons of the world, the Asian summer monsoon is the most intense and also the most
complicated because of the existence of the Himalaya–
Tibetan Plateau. Many studies (e.g. Flohn 1968; Hahn
and Manabe 1975; He et al. 1987; Murakami 1987;
Yanai et al. 1992; Ye 1981) show the importance of the
12
Chou: Land–sea heating contrast in an idealized Asian summer monsoon
Tibetan Plateau as an elevated heat source in the Asian
monsoon circulation. Luo and Yanai (1984) and Yanai
and Li (1994a) showed that dry convection associated
with sensible heat flux from the surface is an important
process in heating up the atmosphere above the Plateau
before the onset of the Asian summer monsoon. After
the onset, the latent heat dominates over the eastern
Plateau (Nitta 1983; Luo and Yanai 1984). Wu and
Zhang (1998) showed that the Tibetan Plateau plays a
crucial role in organizing the three consequential stages
of the Asian summer monsoon onset: the Bay of Bengal,
the South China Sea and India. Some studies discuss the
impacts of the uplift of the Tibetan Plateau on the
evolution of the Asian monsoon (e.g. Ramstein et al.
1997; An et al. 2001). As the uplift of the Tibetan Plateau continues, the intensity of the Asian summer
monsoon enhances and central Asia tends to be drier.
Broccoli and Manabe (1992) showed that the existence
of the Tibetan Plateau induces the dryness of interior
Eurasia in summer and reduces the moisture transport
across the continent.
Another component affecting the seasonal land–sea
heating contrast is the variation of snow cover over
Eurasia (e.g. Bamzai and Shukla 1999; Dickson 1984;
Douville and Royer 1996; Hahn and Shukla 1976; Kripalani et al. 1996; Liu and Yanai 2002; Yang and Xu
1994). The Eurasian snow cover has an inverse relation
with the succeeding Asian summer monsoon rainfall.
Heavy snow in the previous winter and spring will cool
and delay the heating over Eurasia in summer because of
the high albedo of snow and consumption of heat for
snow melting. It results in a weaker Asian summer
monsoon and a delay of the monsoon onset (Shukla
1984; Shukla and Mooley 1987). However, the inverse
relationship between the Eurasian snow cover and the
monsoon rainfall tends to be disrupted by El Niño
events (Yang 1996). Snow cover also plays an active role
in the tropospheric biennial oscillation (TBO) in cooperation with the mid-latitude atmospheric circulation
(Iwasaki 1991; Meehl 1994b; Nicholls 1978; Yasunari
and Seki 1992). Yanai and Li (1994b) showed that the
snow cover leads the Asian summer monsoon intensity
in the quasi-biennial range.
In addition to the land surface conditions and topography discussed already, sea surface temperature
(SST) plays a crucial role in a complicated process to
determine the horizontal temperature gradient associated with the Asian summer monsoon circulation. The
SST anomalies of ENSO, for instance, induce a longitudinal gradient of tropospheric temperature anomalies
and modify the Walker circulation. The positive (negative) gradient of SST anomalies for the warm (cold)
phase of ENSO reduces (enhances) the Walker circulation which is one of the three principal circulations of
the Asian summer monsoon (Webster et al. 1998). The
ENSO-like SST anomalies are also accompanied by the
variation of the tropospheric temperature over the
region of the Eurasian continent (LY96). The El Niño
(La Niña)-like SST anomalies is associated with a cold
(warm) tropospheric temperature over Eurasia and a
negative (positive) meridional gradient of the tropospheric temperature anomalies. Yang and Lau (1998)
showed that a premonsoon moistening and cooling of
the Eurasian landmass associated with the El Niño SST
anomalies weakens the intensity of the following Asian
summer monsoon. Other studies, such as Meehl (1987),
Rasmusson and Carpenter (1983) and Webster and
Yang (1992), also discussed the relationship between the
weak and strong Asian summer monsoon and ENSO.
The present work aims to explore mechanisms that
influence the horizontal temperature gradient in a simple
Afro–Eurasian geometry. The emphasis is on understanding the effects of surface conditions of Eurasia, the
Tibetan Plateau and the SST of the surrounding oceans.
The relationship between the horizontal temperature
gradient and the intensity of the Asian summer monsoon
is also examined. A brief review of the model and a
detailed description of experiments design is given in
Sect. 2. The impacts due to the variations of the land
surface albedo and the heat source over Eurasia are
examined and discussed in Sect. 3. This is followed by a
discussion in Sect. 4 on the thermal effect of the Tibetan
Plateau regarding the uplift of the Plateau and snow
cover. The influence of SST are complicated and experiments associated with meridional and longitudinal
gradients of SST anomalies are presented in Sect. 5. The
latter is relevant to ENSO-like SST anomalies. The last
section summarizes my finding.
2 The model and experiment design
To examine mechanisms determining the horizontal temperature
gradient associated with the Asian summer monsoon circulation, a
coupled ocean–atmosphere–land model of intermediate complexity
(Neelin and Zeng 2000; Zeng et al. 2000, ZNC hereafter) with
prescribed divergence of ocean heat transport (Q-flux) is used.
Based on the analytical solutions derived from the Betts-Miller
moist convective adjustment scheme (Betts and Miller 1993),
typical vertical structures of temperature, moisture and winds for
deep convection are used as leading basis functions for a Galerkin
expansion (Neelin and Yu 1994; Yu and Neelin 1994). The atmospheric model constrains the flow by quasi-equilibrium thermodynamic closures and is referred to as QTCM1 (quasi-equilibrium
tropical circulation model with a single vertical structure of temperature and moisture for deep convection). In QTCM1, the vertical structure of horizontal winds are divided into the baroclinic
and barotropic components (see Fig. 1 of ZNC). The vertical
profile of temperature is associated with the moist adiabatic process, while the moisture profile is estimated from a typical profile of
moisture in the tropical deep convective region (see Fig. 1 of ZNC).
The convection in QTCM1 is driven by a measure of the projected
convective available potential energy (CAPE) and a positive CAPE
indicates the conditional instability at the subgrid scale. Because
the basis functions are based on vertical structures associated with
convective regions, these regions are expected to be well represented and similar to the GCM with the Betts-Miller moist convective adjustment scheme. Far from deep convective regions,
QTCM1 is a highly truncated Galerkin representation equivalent to
a two-layer model.
A cloud-radiation scheme (Chou and Neelin 1996; ZNC) simplified from the full radiation schemes (Harshvardhan et al. 1987;
Fu and Liou 1993) is included. In the scheme, deep and cirrocumulus/cirrostratus cloud fraction is estimated by an empirical
Chou: Land–sea heating contrast in an idealized Asian summer monsoon
13
Table 1 List of experiment types. The motivation for each experiment is discussed in the text
Experiment name
Description
Control
Surface albedo 0.25, Qmax 50 Wm–2,
and no prescribed heating
Maximum/minimum values of
prescribed heating varied: 50/–50
and 100/–100 Wm–2 over Eurasia;
100/–100, 200/–100, and 300/–
100 Wm–2 over the Tibetan Plateau
Surface albedo varied: 0.2 and 0.15
over Eurasia with no prescribed
heating; 0.2 and 0.15 over the Tibetan
Plateau with maximum/minimum
prescribed heating 300/–100 Wm–2
Q-flux varied: 70, 90, and 110 Wm–2 in
Qmax for the meridional direction; 30,
–30, and –60 Wm–2 in maximum
Q-flux anomaly at the dateline for
the longitudinal direction
Prescribed heating
Fig. 1 An idealized land–sea geometry for Eurasia, Africa, the
Indian subcontinent, the Indochina peninsula and the region of the
prescribed Tibetan Plateau-like heating (see text)
parametrization (Chou and Neelin 1999). An intermediate landsurface model (ZNC) is used to simulate interaction between the
atmosphere and land-surface. This model simulates processes such
as evapotranspiration and surface hydrology in a single land surface layer for calculating energy and water budgets. Soil moisture is
balanced by precipitation, evaporation, surface run-off and ground
run-off. To examine interaction between the atmosphere and the
ocean, a mixed-layer ocean with prescribed Q-flux is used (Chou
et al. 2001). The Q-flux excluding seasonal variation is specified to
be symmetric about the equator:
Q-flux ¼ cosðu 3:0Þ Qmax e3juj
ð1Þ
with the maximum value (Qmax) at the equator and latitude / in
radians. The magnitude of Q-flux is roughly estimated from the
zonal average of the observed annual mean. An idealized continental configuration (Fig. 1) is used to represent Eurasia, Africa,
the Indochina peninsula, and the Indian subcontinent. South
America, North America and Australia are not included in this
study. A constant surface albedo, 0.25, forest surface stype and
Qmax = 50 Wm–2 are used in the control run.
A prescribed seasonally varied heat source over the Eurasian
continent is used to examine the sensitivity of the meridional
heating contrast created by the Eurasian continent and the surrounding oceans. Two values of maximum heating (cooling) are
used: 50 (–50) and 100 (–100) Wm–2. The maximum heating appears in July, while the maximum cooling appears in January.
Besides the prescribed heating source over Eurasia, the Eurasian
surface albedo can also modify the meridional heating gradient. In
contrast to the surface albedo, 0.25 in the control run, 0.2 and 0.15
are used to examine an effect equivalent to, for instance, change of
the snow cover over Eurasia.
The version of the model used here does not include topography. Thus to simulate the seasonal thermal effect of the Tibetan
Plateau, a prescribed Tibetan Plateau-like heating is specified (see
Fig. 1) with the maximum heating in July and the maximum
cooling in January. The elevation of the Tibetan Plateau is assumed
to be zero in the model. Because of the specific treatment for the
vertical profile in the model, the prescribed heating has been vertically integrated over the entire troposphere. The vertical structure
of the prescribed heating is similar to a typical deep convective
heating (e.g. Q1 in Yanai et al. 1973). Fennessy et al. (1994) found
an improvement of the simulated Indian monsoon rainfall with a
mean orography rather than an enhanced silhouette orography. I
thus use a constant value of the heating over the entire Tibetan
Plateau. Three experiments with the maximum values of the heating, 100, 200, and 300 Wm–2 are conducted and the maximum
value of the cooling is –50 Wm–2 for all three experiments. These
experiments mimic the thermal effect of the uplift of the Tibetan
Plateau. However, the lack of mechanical forcing due to the topography (Murakami 1958; Riehl 1954; Staff Members 1958a, b;
Trenberth and Chen 1988; Wu and Zhang 1998; Yin 1949) implies
caveats on the discussion of the effect of the Tibetan Plateau. Caveats on the climatology of the Asian summer monsoon simulated
by this model, such as a less distinctive maximum precipitation over
Bay of Bengal, have been discussed in Chou and Neelin (2002).
Nevertheless, we should be able to read some insight into the
Surface albedo
Q-flux
mechanisms of a Tibetan Plateau-like heating that possibly affect
the Asian summer monsoon. Experiments with different surface
albedos (0.25, 0.2 and 0.15) over the Tibetan Plateau are used to
examine effects due to variations of the snow cover over the Plateau
and the thermal effect of the Tibetan Plateau is simulated by a
prescribed heating with the maximum value, 300 Wm–2.
In the mixed-layer ocean model, SST is determined by the
surface energy balance and Q-flux. To examine the effects of SST
on the horizontal temperature gradient and therefore the intensity
of the Asian summer monsoon, Q-flux is modified. First for the
meridional direction, three values of Qmax, 70, 90, and 110 Wm–2
are used for comparison to the control run with the value of Qmax,
50 Wm–2. For the longitudinal temperature gradient, Q-flux
anomalies are added on to the original Q-flux in the control run:
Q-flux0 ¼ sinð/ pÞ Q0max ej/pj e3juj
ð2Þ
with the maximum value (Q¢max) at the dateline and longitude / in
radians. These create ENSO-like SST anomalies. The values of
Q¢max, are –30 and –60 Wm–2 for the El Niño cases and +30 Wm–2
for the La Niña case.
The results of all experiments are seasonal climatology from a
10 year average. In addition to the control run, 15 different experiments are designed to examine mechanisms determining the
horizontal temperature gradient and the intensity of the Asian
summer monsoon. A summary is given in Table 1.
3 Thermal influence of the Eurasian continent
Figure 2a shows July precipitation and 850 hPa wind
for the control run in which a constant albedo is used
everywhere and there is no topography. The monsoon
rainbelt starts from tropical Africa and extends along
the south coast of Asia to the east coast of Asia and the
western North Pacific. Most regions of the rainbelt is
found over ocean surface and the rainbelt slightly tilts
from the southwest to the northeast. The east–west
asymmetry of the rainbelt has been discussed in an
idealized monsoon study (Chou et al. 2001). Two
mechanisms are responsible for the east–west asymmetry: the interactive Rodwell-Hoskins (IRH) mechanism
and ventilation. In the IRH mechanism, Rossby-wave
descent induced by the monsoon diabatic heating in the
east interacts with the convergence zone. The ventilation
14
Chou: Land–sea heating contrast in an idealized Asian summer monsoon
introduces the asymmetry by importing low moist static
energy air from ocean regions into the continent. Both
effects reduce convection over the western part of the
continent and favor convection over the eastern part of
the continent. In Fig. 2a, the cross equatorial flow (at
850 hPa) along the east coast of Africa becomes the
southwesterly monsoon flow and merges with the trade
winds in the western Pacific. This concurs with the
monsoon rain zone. The southwesterly flow characterizes a special feature of the Asian summer monsoon. The
trade winds over the western Pacific have a strong influence on the intensity of the Asian summer monsoon
(Webster and Yang 1992).
Fig. 2 a July precipitation and winds at 850 hPa from the control
run. Differences of July precipitation and winds at 850 hPa for the
prescribed heating experiments over Eurasia between b Q (the
prescribed heating) = 50 Wm–2 and the control run, and c Q =
100 Wm–2 and Q = 50 Wm–2. The contour interval is 4 mm day–1
for a and b and 8 mm day–1 for c. Wind vectors are in m s–1
Both patterns of the precipitation and 850 hPa winds
in the control run capture the major features of a typical
Asian summer monsoon even without topography in the
experiment. However, the monsoon rainfall is relatively
weak and the rainbelt does not extend as far north as
shown in most observations in which the rainbelt extends to 30–40N at East Asia and the western North
Pacific. Examining the corresponding mean temperature
averaged over 200–500 hPa in the control run (Fig. 3a),
the positive meridional temperature gradient is relatively
weak. According to the studies in Meehl (1994a) and
LY96, the weak meridional temperature gradient is associated with a weak Asian summer monsoon and a low
monsoon rainfall. To enhance the meridional temperature gradient and examine the sensitivity of the Asian
summer monsoon to the temperature gradient, two different prescribed heat sources, 50 and 100 Wm–2 with
shape of Eurasia are used. Figure 2b, c shows the differences between the two experiments with the prescribed heating over Eurasia and the control run. As
the prescribed heating increases from 0 to 100 Wm–2, the
monsoon precipitation becomes more intense and the
Fig. 3a–c As in Fig. 2 but for the upper tropospheric (200–
500 hPa) temperature. The contour interval is 2 C for a and
1 C for b and c
Chou: Land–sea heating contrast in an idealized Asian summer monsoon
rainbelt extends farther northward. The angle of the
northeast–southwest tilt of the rainbelt also increases
due to the stronger IRH mechanism and ventilation
(Chou et al. 2001). To the south of the positive precipitation anomalies, the precipitation decreases. The value
of the negative precipitation anomalies is less than that
of the corresponding positive anomalies in the north.
Local maximum precipitation anomalies are found over
the Indian subcontinent and the Indochina peninsula in
the prescribed heating experiments. They resemble the
Fig. 4 Differences of July precipitation and winds at 850 hPa for
the surface albedo experiments over Eurasia between a Rs (surface
albedo) = 0.2 and Rs = 0.25 (the control run), and b Rs = 0.15
and Rs = 0.2. The contour interval is 1 mm day–1. Wind vectors
are in m s–1
Table 2 Indices of the Asian
summer monsoon intensity for
experiments listed in Table 1.
WYI is the vertical shear of
zonal wind between 850 and
200 mb averaged over the
region (0–20N, 40–110E)
(Webster and Yang 1992). CI is
the mean OLR averaged over
(10–25N, 70–120E) (Li and
Yanai 1996)
15
heavy rainfall along the west coast of India and the Bay
of Bengal in observations.
The 850 hPa wind anomalies in Fig. 2b, c behave
similarly to the corresponding mean winds. The increase
of the prescribed heating is equivalent to the enhancement of the land–sea heating contrast (Fig. 3) which
induces the stronger cyclonic circulation that associates
with the enhanced Asian summer monsoon. Thus, the
corresponding low-level moisture convergence associated with the monsoon rainfall moves northward and
enhances the monsoon rainfall. The differences of the
temperature averaged over 200–500 hPa are shown in
Fig. 3b, c. Both meridional and longitudinal temperature gradients are increased, while the change of the
meridional gradient is more than the longitudinal gradient. The maximum temperature tends to shift northeastward to East Asia and the western North Pacific. To
measure the intensity of the Asian summer monsoon,
two monsoon indices are used here. The first monsoon
index (WYI) is based on a vertical shear of zonal wind
between 850 hPa and 200 hPa averaged over the region
of 0–20N, 40–110E (Webster and Yang 1992). The
second index (CI) is a mean OLR averaged over 10–
25N, 70–120E (LY96). The first three rows in Table 2
are the monsoon indices for the experiments shown in
Fig. 2. The intensity of the Asian summer monsoon indeed increases as the meridional temperature gradient is
enhanced by the increase of the prescribed Eurasian
continental heating. This result is consistent with the
studies in Meehl (1994a) and LY96.
To examine the surface albedo effect a series of experiments with different surface albedo over the Eurasian continent are conducted and the results are shown
in Figs. 4 and 5. When the Eurasian surface albedo decreases, the upper tropospheric temperature gradient
increases. The increase of the meridional temperature
gradient is relatively stronger than the longitudinal
temperature gradient (Fig. 5). The corresponding summer monsoons become more intense (see Table 2) and
extend farther northward. The northward extension of
the summer monsoon rainfall creates a north–south
Experiment
Control
Prescibed heating
Over Eurasia
Surface albedo
Over Eurasia
Prescribed heating
Over the Tibetan Plateau
Surface albedo
Over the Tibetan Plateau
Q-flux
Meridional
Q-flux
Longitudinal
50/–50 Wm–2
100/–100 Wm–2
0.20
0.15
100/–100 Wm–2
200/–200 Wm–2
300/–300 Wm–2
0.20
0.15
70 Wm–2
90 Wm–2
110 Wm–2
30 Wm–2
–30 Wm–2
–60 Wm–2
WYI
CI
13.7
28.5
42.2
19.9
25.6
18.8
24.1
29.0
29.6
30.6
21.0
27.1
32.6
15.2
12.3
10.8
266.8
256.8
247.6
264.4
261.6
268.6
270.1
270.5
271.0
271.5
252.8
246.6
246.6
270.8
262.4
257.3
16
Chou: Land–sea heating contrast in an idealized Asian summer monsoon
Fig. 5a, b As in Fig. 4 but for the upper tropospheric (200–
500 hPa) temperature. The contour interval is 0.5 C
monsoon study. On the other hand, the drier landmassinduced high surface temperature enhances the monsoon
intensity, so both effects associated with the drier landmass cancel each other out. Xue (1996) showed that the
effect of the reduction in evaporation dominates the
change of the surface temperature in a desertification
experiment in central Asia, so the Asian monsoon is
weakened. Another effect associated with a drier landmass is surface albedo: a drier land surface has greater
surface albedo. The greater surface albedo associated
with a drier landmass reduces convection and weakens
the corresponding monsoon circulation. In a region with
strong surface albedo, such as northern Africa, the
monsoon convection cannot occur even with saturated
soil moisture (Chou and Neelin 2002). The effect of snow
cover is different to the effect of drier landmass. When
the snow cover increases, the surface temperature decreases due to the absorption of solar radiation for
melting the snow, while the surface albedo is increased
due to the greater albedo of the snow. Both effects of the
high snow albedo and the snow-induced low surface
temperature weaken the summer monsoon intensity.
4 Effects of the Tibetan Plateau
dipole pattern of the monsoon precipitation anomalies
similar to the results in Fig. 2. Meanwhile, in contrast to
the positive anomalies at the southern and southeastern
part of Asia, negative precipitation anomalies at the
northwestern Indian Ocean and Africa create an east–
west dipole pattern. This east–west dipole indicates the
increase of the tilting angle of the east–west asymmetry
of the monsoon rainbelt. When the convection in the
east increases, the enhancement of the subsidence associated with Rossby-wave in the west and the drier air
imported from the north both contribute the negative
anomalies of the precipitation in the west. The surface
albedo experiments over Eurasia can mimic the effect of
the surface albedo due to the variations of snow cover
and vegetation. The strong surface albedo induced by
more snow cover and desert area, for instance, reduces
the intensity of the Asian summer monsoon and extends
the monsoon rainbelt southward. This is consistent with
the results of Meehl (1994a) for the effect of land surface
albedo.
The prescribed heating experiment over Eurasia are
associated with a higher land surface temperature compared to the control run, so these experiments can mimic
the mechanisms associated with a drier landmass and
reduced snow cover. The drier landmass and reduced
snow cover over Eurasia is favorable to a strong Asian
summer monsoon. However, the local hydrological
process plays a different role in determining the location
of the monsoon rainfall and the intensity of the summer
monsoon. For instance, the drier Eurasian landmass has
less evaporation to maintain the monsoon rainfall, so
the Asian summer monsoon rainfall decreases. Chou
et al. (2001) showed the importance of soil moisture
in determining the monsoon rainfall in an idealized
The major difference between Asia and other continents
is the existence of the Tibetan Plateau which averages
over 4000 m in height. The air above the Tibetan Plateau in summer is warmer than the surrounding atmosphere (Flohn 1957, 1960; Yanai et al. 1992). It is
generally believed that the Plateau acts as an elevated
heat source (Yeh et al. 1957; Yeh and Gao 1979; Luo
and Yanai 1984; Yanai and Li 1994a). Enhanced and
vertically integrated heating prescribed over the Tibetan
Plateau shown in Fig. 1 is used to mimic the heating
effect of the elevated Plateau surface. Figure 6 shows
differences of the precipitation and the 850 hPa winds
between the prescribed heating experiments over the
Tibetan Plateau with different values of the heating, 100,
200, and 300 Wm–2 and the control experiment
(Fig. 2a). The Asian summer monsoon rainbelt extends
northeastward as the heating over the Tibetan Plateau
increases (Fig. 6). The north–south and east–west dipole
patterns of the precipitation anomalies are also found in
Fig. 6. The corresponding 850 hPa wind anomalies
show a low-level moisture convergence along the rainbelt anomaly. The Pacific trade wind anomalies are
stronger than the southwesterly monsoon flow anomalies. It implies that the easterly trade winds from the
Pacific are more dominant than the southwesterly
monsoon flow from the Indian Ocean in modifying the
intensity of the Asian summer monsoon when the
heating of the Tibetan Plateau increases. A strong
southward anomaly associated with the Rossby wave
transports dry air into the western part of the Indian
subcontinent and the ocean in the west of the Indian
subcontinent and therefore the tilting angle of the east–
west asymmetry of the monsoon rainbelt increases.
Chou: Land–sea heating contrast in an idealized Asian summer monsoon
Fig. 6 Differences of July precipitation and winds at 850 hPa for
the prescribed heating experiments over the Tibetan Plateau
between a Q = 100 Wm–2 and the control run, b Q = 200 Wm–2
and Q = 100 Wm–2, and c Q = 300 Wm–2 and Q = 200 Wm–2.
The contour interval is 1 mm day–1 for a and b and 2 mm day–1 for
c. Wind vectors are in m s–1
However, the negative precipitation anomalies may be
overestimated because of the lack of the Iran–Afghanistan-western Plateau which plays a role in the Indian
summer monsoon (Yanai et al. 1992).
To examine the relation between the monsoon intensity and the prescribed heating, Table 2 shows two
monsoon indices with respect to the different prescribed
heat sources. When the prescribed heating over the Tibetan Plateau increases, the WYI shows an increase of
the monsoon intensity, while the corresponding CI has
little change. The monsoon circulation associated with
winds becomes intense, but the strength of the corresponding convection in to the region for calculating CI
does not change much. Examining Fig. 6 carefully, the
17
monsoon convection does increase, but the corresponding positive rainfall anomalies are not confined in
to the region (10–25N, 70–120E) used for calculating
CI. This implies that the monsoon index is sensitive to
the region which is chosen to calculate the index.
The increase of the prescribed heating represents an
equivalent thermal effect of the uplift of the Tibetan
Plateau, but it does not include the mechanical effect of
the Tibetan Plateau. The simulations in Fig. 6 imply that
the uplift of the Tibetan Plateau enhances the Asian
monsoon, which is consistent with the studies in An
et al. (2001) and Ramstein et al. (1997). Figure 6 also
shows a northeastward extent of the monsoon rainbelt.
When considering regional Asian monsoon, the impact
of the uplift of the Tibetan Plateau on the summer
monsoon rainfall is different. As the prescribed heating
over the Tibetan Plateau increases, the Indian summer
monsoon rainfall has a relatively small change, while the
rainfall over the Indochina peninsula, the southeastern
and eastern parts of China and the western North Pacific
increases significantly. The rainfall over the ocean
southeast of China is smaller. This implies that the East
and Southeast Asian monsoons are more sensitive to the
uplift of the Tibetan Plateau than the South Asian
summer monsoon (Liu and Yin submitted 2002).
Figure 7 shows the upper tropospheric temperature
for the different prescribed heating experiments over the
Tibetan Plateau in Fig. 6. The patterns of the positive
temperature anomalies in Fig. 7 are narrow in latitude
compared to the prescribed heating experiments over
Eurasia because the heating area of the Tibetan Plateau
is smaller than Eurasian. However, the temperature
anomalies spread much wider than the area of the prescribed heating, especially in longitude. This implies that
the impact of small regional heating cannot be distinguished from the large-scale heating in the upper tropospheric temperature. Therefore, the temperature
gradient presents a relation with a broader scale of the
summer monsoon circulation rather than a subregional
scale of the detailed monsoon structure, such as the
Indian monsoon and the Southeast Asian monsoon. As
in Fig. 3, the maximum temperature is found in South
Asia. In comparison to the observations (e.g. Fig. 3 in
LY96), the prescribed heating, 300 Wm–2 over the Tibetan Plateau (Fig. 9a) induces reasonable magnitude of
the meridional temperature gradient, so an experiment
with the prescribed heating, 300 Wm–2 is used to simulate the effect of the Tibetan Plateau in the surface albedo experiments over the Tibetan Plateau. However,
300 Wm–2 is larger than the observed heat source shown
in Yanai et al. (1992), Yanai and Li (1994a) and Yanai
and Tomita (1998). This is because the model must
warm up the air from the surface around 1000 hPa to
the tropopause, compared to the air above the Tibetan
Plateau which is from 600 hPa to the tropopause.
When considering effects of snow cover, surface albedo experiments over the Tibetan Plateau (Fig. 8) are
conducted. As the surface albedo over the Tibetan Plateau becomes smaller due to less snow cover, the Asian
18
Chou: Land–sea heating contrast in an idealized Asian summer monsoon
Fig. 7a–c As in Fig. 6 but for the upper tropospheric (200–
500 hPa) temperature. The contour interval is 1 C
summer monsoon intensifies (see WYI in Table 2) and
moves farther northeastward. Dickson (1984) showed a
similar negative correlation between the Himalaya snow
cover and the subsequent Indian summer monsoon
rainfall. The north–south dipole pattern of the precipitation anomalies is found, while the east–west dipole
pattern of the precipitation anomalies is weakened.
Since the positive and negative precipitation anomalies
cancel each other out, the CI of the surface albedo experiments over the Tibetan Plateau has little change. The
tropospheric temperature in Fig. 9 shows only small
differences between the surface albedo experiments over
the Tibetan Plateau and the control run. The rainfall
anomalies in Fig. 8 are also relatively small compared to
the other experiments discussed before (e.g. the experiments in Fig. 6), especially the Indian summer monsoon
rainfall. Bamzai and Shukla (1999) showed no significant relation between the Himalaya snow cover and the
Indian monsoon rainfall in the following summer. In
East Asia and Southeast Asia, the impacts of the snow
cover over the Tibetan Plateau are different. The East
Asian summer rainfall is negatively correlated to the
snow cover over the Tibetan Plateau associated with the
Fig. 8 a July precipitation and winds at 850 hPa for an experiment
with the prescribed heating, 300 Wm–2 over the Tibetan Plateau.
Differences of July precipitation and winds at 850 hPa for the
surface albedo experiments over the Tibetan Plateau between b Rs
= 0.2 and Rs = 0.25, and c Rs = 0.15 and Rs = 0.2. The contour
interval is 4 mm day–1 for a and 0.5 mm day–1 for b and c. Wind
vectors are in m s–1. Note the wind vector in a is different from b and
c but the same as in Fig. 2
surface albedo, but the Southeast Asian summer rainfall
is positively correlated to the snow cover over the Tibetan Plateau. The snow cover of the Tibetan Plateau
shows a relationship with the regional Asian summer
rainfall, such as the East and Southeast Asian summer
rainfall, but not with the Asian rainfall as a whole.
5 Impacts of sea surface temperature
The relation between SST and the Asian summer monsoon rainfall is complicated (e.g. Arpe et al. 1998; LY96;
Chou: Land–sea heating contrast in an idealized Asian summer monsoon
Fig. 9a–c As in Fig. 8 but for the upper tropospheric (200–
500 hPa) temperature. The contour interval is 2 C for a and
0.2 C for b and c
Rasmusson and Carpenter 1983; Shen and Lau 1995;
Wang et al. 2000; Wang et al. 2001; Webster and Yang
1992; Weng et al. 1999; Yang and Lau 1998). To study
the effect of SST directly involving in the tropospheric
temperature on the Asian summer monsoon, two
idealized Q-flux anomalies are specified: a zonally
symmetric anomaly to the equator and an ENSO-like
anomaly in the tropics. The zonally symmetric anomaly
associated with varied Qmax induces a meridional gradient of SST anomalies and the ENSO-like anomaly
induces a longitudinal gradient of SST anomalies. For
the meridionally Q-flux experiments, three values of
Qmax, 70, 90, and 110 Wm–2 are used for comparison to
the control run with Qmax, 50 Wm–2. Figure 10 shows
differences of the corresponding SST (or land surface
temperature, Ts) between the meridional Q-flux experiments and the control run. Because the ocean transports
energy out of the tropics more efficiently, the tropical
SST decreases as Qmax increases. The SST anomalies are
symmetric to the equator. Over land, clouds associated
with the monsoon rainbelt reflect more solar radiation
and the land surface temperature in convective regions
decreases. Outside the convective regions, the decrease
19
Fig. 10 Differences of July sea (land) surface temperature for the
meridional Q-flux experiments between a Qmax = 70 Wm–2 and
Qmax = 50 Wm–2 (the control run), b Qmax = 90 Wm–2 and Qmax
= 70 Wm–2, and c Qmax = 110 Wm–2 and Qmax = 90 Wm–2. The
contour interval is 2 C
of the land surface temperature is caused by changes in
surface heat fluxes which is induced by the cold air
temperature (Fig. 11).
Figure 11 shows the difference of the upper tropospheric temperature between the meridional Q-flux experiments and the control run. The upper tropospheric
temperature anomalies are symmetric to the equator
which is consistent with the SST anomalies in Fig. 10. A
similar symmetric pattern of the upper tropospheric
temperature anomalies is also found in the observation
and is related to the existence of the Asian summer
monsoon (Liu and Yanai 2001). The tropospheric temperature is colder than the control run and the maximum
anomaly is at the equator. The distribution of the
tropospheric temperature anomalies is smooth and does
not resemble the detailed variation of the land surface
temperature (Fig. 10). The meridional gradient of the
tropospheric temperature anomalies is enhanced, while
the longitudinal gradient of the tropospheric temperature anomalies has little change. Because of the wave
20
Chou: Land–sea heating contrast in an idealized Asian summer monsoon
Fig. 11a–c As in Fig. 10 but for the upper tropospheric (200–
500 hPa) temperature. The contour interval is 1 C
Fig. 12a–c As in Fig. 10 but for July precipitation. The contour
interval is 3 mm day–1. Wind vectors are in m s–1
propagation as discussed in Su et al. (2001) and Wallace
et al. (1998), the change of the meridional temperature
gradient extends into the atmosphere over land surface.
The precipitation differences between the meridional
Q-flux experiments and the control run are shown in
Fig. 12. The rainbelt extends farther northward as the
upper tropospheric temperature gradient increases. The
north–south dipole pattern of the precipitation anomalies is found, but there is no clear east–west dipole pattern. The northward extension of the monsoon rainbelt
is found not only in Asia, but also in Africa. The positive
rainfall anomalies in Africa and Asia appear to connect
to each other and the Asian summer monsoon extends
farther northward than the African summer monsoon
(Liu and Yanai 2001). This southwest–northeast tilt of
the positive rainfall anomalies is induced by similar
mechanisms to the idealized monsoon study in Chou
et al. (2001). The meridional gradient of the tropospheric temperature anomalies is zonally symmetric to
the equator in a global scale, but the northward extent of
the convection zones occurs only over land surface. It
is evident that the existence of land is important in
the monsoon circulation. In the meridional Q-flux
experiments, the tropospheric temperature is colder than
that in the control run, but the monsoon rainfall does
not fall as the temperature becomes colder. The horizontal temperature gradient is more dominant than the
tropospheric temperature itself in determining the summer monsoon rainfall.
As the positive monsoon rainfall anomalies extend
into the African continent, the southwesterly monsoon
flow and the easterly trade wind merge over tropical
Africa instead of South Asia in the previous experiments. The cross-equatorial flow over the Indian Ocean
along the east coast of Africa is significantly weakened.
The easterly trade winds extend farther west and into the
African continent. The cyclonic circulation associated
with the Rossby wave subsidence also extends farther
west and limits the northward extent of the African
convection zone. The WYI of these experiments shows
an increase in the monsoon intensity as the tropical SST
becomes lower. The monsoon intensity is largely modified by the strong variations of the upper tropospheric
winds induced by the global meridional temperature
Chou: Land–sea heating contrast in an idealized Asian summer monsoon
21
Fig. 13 Differences of the July SST for the longitudinal Q-flux
experiments between a Q¢max = 30 and the control run, b Q¢max =
–30 and the control run, and c Q¢max = –60 and the control run.
The contour interval is 0.2 C
Fig. 14 Differences of the July upper tropospheric (200–500 hPa)
temperature for the longitudinal Q-flux experiments between a
Q¢max = 30 and the control run, b Q¢max = –30 and the control run,
and c Q¢max = –60 and Q¢max = –30. The contour interval is 0.2 C
gradient. The corresponding CI does not show any increase of the monsoon convection because most rainfall
anomalies occur outside of the region for calculating CI.
Next, the impacts of the longitudinal SST gradient on
the variations of the Asian summer monsoon are
examined. In the longitudinal Q-flux experiments,
La Niña-like SST anomalies (Fig. 13a), weak El Niñolike anomalies (Fig. 13b) and strong El Niño-like SST
anomalies (Figs. 13c) are specified. For the La Niña (El
Niño) SST anomalies, strong cold (warm) SST anomalies in the central and eastern Pacific are associated with
weak warm (cold) SST anomalies in the western Pacific.
The corresponding upper tropospheric temperature differences are shown in Fig. 14. Two longitudinal gradients of the temperature anomalies are found: western
Pacific to eastern Pacific and western Pacific to Africa.
These two longitudinal gradients of the temperature
anomalies modify the transverse circulations (Webster
et al. 1998): the Walker circulation between South Asia
to the eastern Pacific and the circulation between
northern Africa and South Asia associated with Rossby-
wave dynamics. For the La Niña case (Fig. 14a), the two
transverse circulations are enhanced by the longitudinal
gradients of the temperature anomalies, so a strong
monsoon is expected (see WYI in Table 2). Over Eurasia, a clear positive meridional gradient of the tropospheric temperature anomalies associated with a strong
summer monsoon is also found. For the two El Niño
cases, the transverse circulations are reduced and the
corresponding monsoons are weakened (see WYI in
Table 2). A negative meridional gradient of the tropospheric temperature anomalies is also found. The variations of the longitudinal SST gradient change not only
the longitudinal upper tropospheric temperature gradient, but also the meridional upper tropospheric temperature gradient. The upper tropospheric temperature
anomalies associated with ENSO are consistent with
observations (e.g. Fig. 12 of LY96).
The meridional gradient of the tropospheric temperature anomalies associated with ENSO is caused by the
SST anomaly induced Rossby wave and Kelvin wave.
For instance, the cold tropospheric temperature over
Eurasia in Fig. 14c is induced by a Rossby wave in
22
Chou: Land–sea heating contrast in an idealized Asian summer monsoon
response to the cold SST anomalies in the western Pacific in Fig. 13c. The warm tropospheric temperature
over the Indian Ocean is induced by a Kelvin wave in
response to the warm SST anomalies in the central and
eastern Pacific. The tropospheric temperature anomalies
over Eurasia are farther modified by coupling with land
hydrology. Yang and Lau (1998) showed that wetter and
colder surface condition, due to more snow cover over
Eurasia is induced by the El Niño SST anomalies. The
tropospheric temperature over Eurasia in Fig. 14c will
be farther reduced by interaction between the atmosphere and land surface.
Figure 15 shows the corresponding precipitation
difference. For the La Niña case (Fig. 15a), a dipole
pattern of the rainfall anomalies is found with positive
anomalies in the south and negative anomalies in the
north. The summer monsoon rainbelt moves southward,
while the intensity of the monsoon is enhanced
(Table 2). The observation (e.g. Fig. 14 of LY96) also
shows a similar pattern for the strong monsoon case.
The patterns of rainfall anomalies in Fig. 15 is a little
further south compared to the observations, so the CI is
Fig. 15a–c As in Fig. 14 but for the precipitation. The contour
interval is 1 mm day–1. Wind vectors are in m s–1
inconsistent with the WYI. For the El Niño case
(Fig. 15b, c), the intensity of the summer monsoon is
reduced and the monsoon rainbelt extends northward.
The northward extension of the summer monsoon
rainbelt associated with El Niño is due to an anomalous
anticyclone induced by local cold SST anomalies and
Rossby wave-induced subsidence in response to the
warm SST anomalies in the central and eastern Pacific
(Chang et al. 2000a, b; Wang et al. 2000). The anticyclonic circulation is confined to the lower troposphere.
The anomalous anticyclonic circulation provides lowlevel moisture convergence from the western Pacific and
extends the corresponding monsoon rainbelt northward.
Figure 15b, c shows a similar anticyclonic circulation
over the ocean southeast of Southeast Asia.
6 Summary and conclusions
Motivated by the close relationship between the intensity of the Asian summer monsoon and the horizontal
temperature gradient (LY96), a series of numerical experiments are conducted in an idealized Afro–Eurasian
geometry. The focus of the study is to examine the
mechanisms that influence the horizontal temperature
gradient induced by land–sea heating contrast and
modified by other processes involving in the surface
conditions of Eurasia, the thermal influence of the Tibetan Plateau and the SST distribution over the surrounding oceans. The relation between the intensity of
the Asian summer monsoon and the horizontal temperature gradient induced by the different processes are
also discussed.
The meridional temperature gradient can be intensified by colder tropical SST anomalies and higher prescribed heating sources and weaker surface albedo over
Eurasia and the Tibetan Plateau. These mimic effects of
different land surface processes and the thermal effect of
the uplift of the Tibetan Plateau. Caveats are underlined
on these experiments based on the prescribed heat
source and albedo. Without the physical topography in
the model, the mechanical effect of the Tibetan Plateau
has been neglected and the prescribed heating is larger
than the observation in order to warm up the thicker air
column. Varied surface albedo can roughly estimate the
effect of snow cover, but its hydrological effect is neglected. Nevertheless, the intensified meridional temperature gradient enhances the Asian summer monsoon
circulation and the corresponding rainfall. The Asian
summer monsoon rainbelt also extends northeastward
and creates different regional rainfall responses, particularly over East Asia. Thus, the meridional temperature
gradient can represent a broader scale of the Asian
summer monsoon with regional differences in rainfall
anomalies.
The prescribed heating experiments over Eurasia
show that the higher prescribed heating warms up the
atmosphere above Eurasia and the Eurasian landmass,
so the meridional tropospheric temperature gradient
Chou: Land–sea heating contrast in an idealized Asian summer monsoon
increases. The positive gradient of the meridional temperature anomalies enhances the intensity of the Asian
summer monsoon and extends the corresponding monsoon rainbelt northeastward. In the surface albedo experiments over Eurasia, the greater surface albedo
results in lower surface and tropospheric temperatures
over Eurasia and the meridional temperature gradient
decreases. The Asian summer monsoon weakens and the
monsoon rainbelt moves southward. The prescribed
heating experiments over the Tibetan Plateau indicates
that the higher prescribed heating favors a stronger
Asian summer monsoon and the corresponding monsoon rainbelt extends farther northeastward. The tilting
angle of the monsoon rainbelt also increases. When the
heating reaches a critical value, the East Asian monsoon
continues to extend northward, while the South Asian
monsoon has little change in position. This implies that
the East Asian monsoon is more sensitive to the uplift of
the Tibetan Plateau than the South Asian monsoon (Liu
and Yin submitted 2002). In the surface albedo experiments, the enhancement of the Asian summer monsoon
is relatively weak and it is confined to Southeast Asia
and East Asia. The Southeast Asian summer monsoon
rainfall has a positive relation with the surface albedo
over the Tibetan Plateau, while the East Asian summer
monsoon rainfall has a negative relation with the surface
albedo.
To discuss the impact of SST anomalies, two types of
Q-flux experiments are conducted: meridional Q-flux
experiments and longitudinal Q-flux experiments. The
positive meridional gradient of SST anomalies enhances
the meridional tropospheric temperature gradient and
therefore the intensity of the Asian summer monsoon
increases. The effect of the longitudinal gradient of SST
anomalies is much more complicated than the effect of
the meridional gradient of SST anomalies. The El Niñolike SST anomalies (Fig. 13b, c) weaken the Walker
circulation and induce a positive longitudinal gradient of
the tropospheric temperature anomalies across the entire
Pacific, so the intensity of the Asian summer monsoon
reduces. The circulation between South Asia and
northern Africa is also weakened by the negative longitudinal gradient of the tropospheric temperature
anomalies in the regions. Likewise, the La Niña-like SST
anomalies (Fig. 13a) are associated with a stronger
Asian summer monsoon. However, the corresponding
monsoon rainbelt extends northward, while the intensity
of the Asian summer monsoon reduces. It is due to the
anticyclonic circulation over the south and southeast of
South Asia which transports more moisture from the
western Pacific into the monsoon rainbelt.
The ENSO-like SST anomalies also induce meridional tropospheric temperature anomalies (Fig. 14). The
variations of the tropospheric temperature over Eurasia
are associated with a Rossby wave in response to the
SST anomalies over the equatorial western Pacific. The
variations of the tropospheric temperature over the Indian Ocean result from a Kelvin wave in response to the
warmer SST over the central and eastern Pacific. In the
23
El Niño-like SST anomaly case, the colder tropospheric
temperature over Eurasia and the warmer tropospheric
temperature over the Indian Ocean induce a negative
meridional gradient of the tropospheric temperature
anomalies. This further weakens the Asian summer
monsoon that has already been reduced by the longitudinal gradient of the tropospheric temperature anomalies. Land hydrology associated with snow cover might
play a role in further reducing the atmospheric temperature above Eurasia (Yang and Lau 1998).
Acknowledgements The author thanks Prof. M. Yanai for his
valuable suggestions. Comments from Prof. W. L. Gates and an
anonymous reviewer are greatly appreciated. This work was supported under National Science Council grant 90-2119-M-001-020.
The author thanks M.-J. Yang and J.-Y. Yu for provision of their
computing facilities.
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