Journal of the Meteorological Society of Japan, Vol. 85B, pp. 55--75, 2007

55

Role of Land-Atmosphere Interaction on Asian Monsoon Climate

Tetsuzo YASUNARI
Hydrospheric Atmospheric Research Center (HyARC), Nagoya University, Nagoya, Japan
Frontier Research Center for Global Change (FRCGC), JAMSTEC, Yokohama, Japan(Manuscript received 1 June 2007, in final form 7 June 2007)

Abstract
This paper has discussed the land-atmosphere interaction associated with the Asian monsoon climate
and its variability. The snow cover and soil moisture are internal factors of the climate system particularly in the seasonal to interannual time scales, but are suggested to play some important roles in changing large-scale surface energy and water balance, which in turn affect the monsoon circulation and precipitation. From the modeling as well as observational studies it has been suggested that the snow coveris most likely to influence atmosphere essentially through the albedo effect particularly in lower latitudes and on the Tibetan Plateau. The GCM experiments have strongly suggested that soil moistureanomalies efficiently affect the atmosphere under dry or semi-arid condition. Vegetation has also been
noted as an important variable for the formation of moist monsoon flow over the continent, in addition
to the dynamical and thermo-dynamical effect of the Tibetan Plateau.
Through the GCM and RCM experiments the anthropogenically-induced change of land cover/use including the deforestation has proved to be a great impact on regional precipitation and water cycle of themonsoon region by changing characteristics of vegetation control in energy and water cycle, which inturn affect atmospheric boundary layer (ABL) and cloud/precipitation processes in regional-scale. The
discussion has also been made on how the land surface changes could induce change of precipitation
through feedback processes of moisture convergence and in-situ evapo-transpiration processes. The critical role of moisture amount near surface and in the ABL is emphasized, to induce positive feedback tochange of precipitation over land in the monsoon region.

1.

Introduction

The monsoon is manifested as a landatmosphere-ocean interaction between continents and oceans in the seasonal cycle. The
ocean has large heat content with longer climate memory of more than a year, while theland has small heat content and its climate
memory is believed to be short (of less than aseason). Some of the recent model studies emphasized relative importance of the oceanatmosphere interaction (O-AI) compared toCorresponding author: Tetsuzo Yasunari, Hydrospheric Atmospheric Research Center (HyARC),
Nagoya University, Nagoya, Aichi 464-8601, Japan.
E-mail: yasunari@hyarc.nagoya-u.ac.jp( 2007, Meteorological Society of Japanland-atmosphere interaction (L-AI), particularly focusing on the strong impact of largescale sea surface temperature (SST) anomalies
in the tropical oceans nearby the continent. It
should be noted, however, that to specify or distinguish the roles of L-AI from those of O-AI
and vice versa is very difficult or, may be nonsense, particularly in the Asian-Australianmonsoon system, because these two processes
are strongly coupled to each other. In otherwords, the L-AI (or O-AI) in the monsoon system should, in any case, be understood as partof the full land-atmosphere-ocean interaction.
In this review I will discuss some current issues on the role of L-AI on the Asian monsoon
climate with its seasonal cycle, interannual
variability and long-term change including
anthropogenic impact, based on some recent ob-

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servational as well as climate modeling studies,by updating the discussion made in Yasunari(2006).
2.

Fundamental land surface processes
controlling Asian monsoon climate

The physical quantities of land surface which
may posses anomalous atmospheric forcing or
climate memory effect can be 1) snow cover, 2)soil moisture, and 3) vegetation. The land surface layer is generally thought to have smallheat capacity compared to the ocean surface
layer, the non-linear processes of these quantities in the seasonal and interannual variations
effectively produce relatively longer memory effect, even compatible to that of SST as shown in
Fig. 1 (Walsh et al. 1985; Shinoda and Gamo
2000). The seasonal and interannual variability of these quantities play significant rolein land-atmosphere interaction by producing
anomalous surface and atmospheric heatingor cooling, which in turn affects the monsoon
circulation and convection. One should keep in

Fig. 1. Autocorrelations
ofmonthly
anomaly of SST in the northern Pacific,
sea ice in the Arctic Sea, snow cover
over Eurasia, geopotential height at
700 hPa depicted from Walsh et al.
(1985) and NDVI in Sahel region, North
Africa obtained from Shinoda andGamo (2000). (Shinoda et al. 2003)

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mind that these quantities are interactive withthe atmospheric conditions so that detection of
simple cause-result relation is sometimes difficult. However, due to lack of the adequate observational data, and lack of understanding ofactual physical processes particularly relevantto the climate studies, large uncertainties still
exist in the relation between variabilities ofthese quantities and monsoon activity. Comprehensive and interactive observation and modeling strategy is really important for this study.
2.1 Role of continental snow cover
The snow cover has several effects on producing anomalous atmospheric conditions as follows: a) albedo effect controlling incoming solar
radiation, b) insulation of heat between theatmosphere and land surface by snow mass of
small heat conductivity, c) heat sink effect bymelting process, and d) water source by melting
process (Shinoda et al. 2003). The first threeeffects can work to change surface energy
transfer process while snow cover exists on thesurface. Only the fourth effect possibly has a
climate memory even after snowmelt, by affecting soil moisture content near the surface whichis called ‘‘snow-hydrological effect’’ as shown in
Fig. 2 (Yasunari et al. 1991).
The data sets for validating these effects ofsnow cover are still limited. Most of the observational studies have been made based on thesnow cover extent (of frequency) data from the
NOAA operational meteorological satellites.
Some studies also used ground-based snow
depth measurement for limited areas (e.g.,
the former USSR, USA, Canada etc.). The
continental-scale snow mass (water equivalent
depth) information retrieved from satellitebased microwave data has become available forthe recent one or two decades (e.g., Koike et al.
2001) though some accuracy problems still exist
due to biases from vegetation and melting processes.
Blanford (1884) first noted the relationshipbetween the Himalayan winter snow coveranomaly and the succeeding All-India Monsoon
rainfall (AIMR). Walker (1910) followed up his
study and re-confirmed this negative correlation. Since the 1960s the satellite-based snow
cover extent data became available, and somany studies have addressed this issue as a
typical example of the role of L-AI on the inter-

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Fig. 3. Year-to-year variations of the All
India Monsoon Rainfall and snow coverarea over central part of Eurasia northof 52 . (Hahn and Shukla 1976)

Fig. 2. Schematic diagrams for (1) the albedo feedback and (2) the hydrologicalfeedback of snow cover during the seasonal march from winter to summer.
(Yasunari et al. 1991)

annual variability of Asian monsoon. Hahn andShukla (1976) showed an apparent negative
correlation between the satellite-based wintersnow cover extent anomaly over Eurasia andthe following AIMR, though the data usedwere only 9 years (Fig. 3). Since then numerous
similar studies have been done using some different indices of snow cover for winter and/or
spring, and some different indices of the Indianor Asian summer monsoon activity (Dey andBhanu Kumar 1982; Dickson 1984; Bhanu
Kumar 1987, 1988; Chattopadhay and Singh
1995; Kripalani et al. 1996, 2003; Shankar-Rao
et al. 1996; Morinaga et al. 1997, 2000; Bamzai
and Shukla 1999; Kripalani and Kulkarni
1999; Robock et al. 2003). Some studies also
discussed Eurasian snow-AIMR relationship,as part of ENSO-monsoon connections (Yasunari 1987; Kahndekar 1991; Yasunari and Seki
1992; Yang 1996; Yang and Lau 1998; Kawamura 1998; Ye and Bao 2001; Liu and Yanai

2002). Some of these studies concluded thatthe snow cover-AIMR correlation might be resulted from a common large-scale atmosphericcirculation pattern that is responsible both forthe snow cover anomaly and the AIMR anomaly. One possible idea is that ENSO-related
teleconnection induces anomalous atmosphericcirculation over Eurasia, which may fundamentally be responsible for anomalous Asian summer monsoon activity (Webster et al. 1998).
The snow cover anomaly induced by the anomalous circulation may partly reinforce to produce weaker monsoon through radiation andenergy flux processes as suggested in the observational study (Yang et al. 2004) and in a GCM
experiment (Shen et al. 1998).
A central issue for the snow cover-AIMR connection may be how the climate memory ofsnow cover in winter (or spring) anomaly canbe conveyed to the following summer. Some
GCM studies attempted to this issue (Barnett
et al. 1989; Yamazaki 1989; Yasunari et al.
1991; Vernekar et al. 1995; Douville and Royer
1996). Though some of these GCM experiments
have given unrealistically large anomalous
snow cover (snow mass) as an initial condition,
both the albedo and the snow-hydrological effect of anomalous snow cover could play, tosome extent, to produce monsoon anomalies insummer. Some studies emphasized that the58

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albedo and snow-melting effect on the Tibetan
Plateau in spring showed relatively large
impact on weakening the summer monsoon
through suppressing net radiation and sensible
heat flux (Yasunari 1991; Douvelle and Royer
1996). In contrast, a recent large-ensemble
GCM experiment has suggested that in thenorthern part of the continent (e.g., Siberia),
snow-mass anomaly may be more importantfor producing atmospheric anomaly than snow
cover (albedo) forcing (Gong et al. 2004). Takata
and Kimoto (2000) pointed out the importanceof seasonal freezing (and melting) of surface
permafrost layer in the northern part of thecontinent on thermal and hydrological processes of the surface. Full inclusion of theseprocesses causes higher surface temperature(due to suppressed evaporation) and stronger
monsoon circulation in Southeast Asia. The observational study (Ohta et al. 2001) of seasonal
surface energy and water flux in East Siberia
conducted as part of GAME-Siberia partly supported the important processes noted by Takata and Kimoto (2000).
The precise observational studies on seasonal
march of temperature and circulation field overEurasia from spring to summer, however, haveshown that the influence of snow cover and related soil moisture anomaly on the temperatureand circulation anomalies in the lower troposphere is limited primarily when and where
snow cover exists seasonally as schematicallyshown in Fig. 4 (Shinoda et al. 2001; Ueda
et al. 2003; Robock et al. 2003). The long-term
surface energy flux observation on the Tibetan

Fig. 4. Latitude–time sections showing(a) mean (1966–90) seasonal evolutionof the surface temperature (degrees
Celsius) along a longitude of 30–60 E.
Thick black contour indicates surface
temperature 0 C. The dashed contour
denotes climatological snow-disappearance pentads. (b) The same as in (a),except for 850 hPa horizontal wind obtained from ECMWF reanalysis (1980–
90). The southerly wind component
greater than 1 m s1 is denoted by twotone shadings. (c) The same as in (a),but for static stability g between 850
and 700 hPa. Shading is region of gless than 5.0 K km1 . (Ueda et al. 2003)

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Plateau as part of GEWEX Asian Monsoon Experiment (GAME) in 1998 suggested that theanomalous heavy snow cover in spring on thecentral Plateau resulted in considerable decrease of sensible heating and delayed monsoon
onset over the Plateau (Tanaka et al. 2004).
2.2 Role of soil moisture
The soil moisture is a key parameter for L-AI
in the climate system. The anomaly of surface
and near-surface soil moisture is likely to havepersistence of several days to several months,
which may cause climate memory throughanomalous surface energy and moisture fluxes.
However, this quantity is very difficult to measure adequately. The long-term in-situ measurement of soil moisture is very limited bothin time and space (Robock et al. 1998, 2000).
The recent satellite-based indirect measurement using microwave sensors is a promising
method for large-scale soil moisture monitoring, but it has still been limited for soil moisture or wetness of very-near surface soil layer.
The climate memory effect of soil moisture has,
therefore, been discussed basically based uponclimate model experiments.
Delworth and Manabe (1988, 1989) firstly assessed the climate memory effect of soil moisture in GCMs using the so-call bucket model(Manabe 1969) interacting with the atmosphere through surface water balance. Their
bucket model consists of one soil layer with
15 cm water holding capacity (field capacity).
They noticed that the persistency (measuredby auto-correlation) of soil moisture anomaly issmall whereas in higher latitudes this value islarge, depending upon the ratio of P/Ep (P: Precipitation, Ep: Potential Evapo-transpiration)
as shown in Fig. 5. In the humid tropics or
monsoon regions where P is large enough tosaturate the soil layer, soil moisture anomaly
does not make sense as a climate memory. The
distribution of persistency in Fig. 5 can begenerally described in the form 1/l ¼ W  /Ep,
where l is e-folding time of exponential decayof soil moisture anomaly, and W  field capacity.
Though their definition of soil moisture and assumption of soil layer model were so simple,
this relation between soil moisture anomaly
persistency, field capacity of soil layer and Ep
was partly proved in the observation in central
Eurasia (Vinnikov and Yeserkepova 1991).

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In recent years, more realistic and sophisticated land surface models (LSMs) have beendeveloped, and many sensitivity experimentswith GCMs have been conducted using different types of LSMs under different experimental
designs. For example, Douville et al. (2001) andDouville (2002) assessed the influence of soil
moisture on seasonal and interannual variability of Asian and African monsoons. Throughsome ensemble experiments with and without
an interactive LSM (named ISBA; Interaction
between Soil Biosphere and Atmospheremodel), they found that the impact of soil moisture anomaly through the land-atmosphere interactive process is significant in relatively dry
monsoon regions (e.g., Indian subcontinent andnortheast Asia) but not in the humid monsoon
regions (e.g., Southeast Asia). They also noticed
that in the significant impact regions soil moisture anomaly primarily plays a role as an initial condition (at the monsoon onset phase)rather than the boundary condition. Kanae
et al. (2006) also performed a set of hindcast
GCM simulation (with prescribed sea surface
temperature) on boreal summer hydro-climate
for about 40 years (1951–1998), with and without realistic soil moisture anomalies, and noticed that only semi-arid regions in the peripheries of monsoon regions showed reasonably
good simulated precipitation with the observed
ones. It is interesting to note that all the GCM
experiments, though the LCM performance andsimulation design are quite different from eachother, have shown relatively high impact of soil
moisture anomaly in dry or semi-arid regions.
This general tendency of soil-precipitation feedback has comprehensively been confirmed by
the recent Global Land Atmosphere CouplingExperiment (GLACE) (Koster et al. 2004),
where twelve GCMs with different LCMs participated. For example, in the Asian monsoon
region, the sensitivity of precipitation change
to soil moisture condition is large only in relatively dry areas, e.g., Indian subcontinent andthe Yellow-River/Inner Mongolia in China as
shown in Fig. 6. In humid monsoon and tropicalregions, the impact seems to be small or insignificant primarily due to already-saturated soil
moisture condition. However, this characteristic nature of weak soil moisture memory effectin the humid tropics may, to some extent, counteract with the relatively long memory effect in60

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Fig. 5. (a) Lag one autocorrelation values of soil moisture for the months of June, July and August(JJA) for Soil Moisture Interactive GCM experiment (SMI). At each grid point, deviations ofmonthly mean soil moisture from the long-term mean for that month were correlated with datafrom the same grid point, but lagged one month. Coefficients greater than 0.16 (0.3) are significantly different from zero at the 95% (99.9%) confidence level. Values greater than 0.6 are densely
stippled, while value less than 0.4 are lightly stippled. Permanently ice-covered region are black.
(b) Potential evaporation (cm d1 ) for JJA in SMI. (Delworth and Manabe 1989)

the humid tropics (Wu and Dickinson 2004). In
addition, the soil moisture anomaly may be important as an initial condition to switch on theinteraction with precipitation in these relatively dry regions.
2.3 Role of vegetation
Another important aspect of land-atmosphere interaction in addition to soil moisture(and snow cover) may be a role of vegetation.
A general overview on the influence of vegetation on weather and climate is summarized in

Pielke et al. (1998). In the Asian monsoon
region, most of the areas are heavily coveredby vegetation, including tropical rain and monsoon forests, water-fed paddy field, grass landand boreal forest etc. The roles of vegetation inland-atmosphere interaction may be classified
as follows;1) control of radiation and energy fluxes
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Fig. 6. The land-atmosphere coupling strength diagnostic for boreal summer (the W difference,
dimensionless, describing the impact of soil moisture on precipitation), averaged across the 12
models participating in GLACE. (Insets) Area-averaged coupling strengths for the 12 individualmodels over the outlined, representative hotspot region. No signal appears in southern SouthAmerica or at the southern tip of Africa. (Koster et al. 2004)

3) control of substantial field capacity of soil
with root depth and structure
Due to lack of observations particularly in thelarge-scale vegetation and its impact on climate
and water cycle, this aspect has been noted
only in the recent years. LSMs including thesevegetation processes have also been developedand improved in the recent few decades or so(e.g., Dickinson and Henderson-Sellers 1988;
Sellers et al. 1986; Takata et al. 2003), thoughthese models need numerous number of tuning
parameters. To improve and fully utilize these
sophisticated models, however, observational
studies are also essential, including optimal determination of these parameter values. In the
Asian monsoon region, intensive field campaigns related to some international projects
such as GAME have been conducted, and haverevealed new aspects of the role of vegetationin terms of energy and water cycle processes.
In the evergreen tropical forest in Thailand,
energy and water flux measurements at several
sites with different types of vegetation wereconducted for more than two years as part of
GAME-Tropics. The evapotranspiration esti-

mated by a multi-layer model forced by theobserved radiation, leaf area index (LAI) andother realistic parameters revealed that aseasonally maximum evapo-transpiration appeared in the late dry season just before themonsoon onset (Fig. 7) when the ground surface was driest in the year (Tanaka et al.
2003). A LSM study including the effect of root
depth to pump up water in deep soil layer could
well simulate this seasonal feature when themean root depth was set at 6 m (Tanaka et al.
2004; Tanaka and Hashimoto 2006). This implies that effective field capacity must be set at
an extremely large value in this case. In thedeciduous monsoon forest in Thailand, in
contrast, the seasonal maximum of evapotranspiration appeared in late monsoon season,
associated with seasonal photosynthetic activity and water availability (Toda et al. 2002).
These dependencies of the seasonal variationof evapo-transpiration (i.e., latent heating) on
dominant vegetation types may have a great
impact on modeling of regional and large-scale
climate and water cycle, particularly in thetropics.
In the boreal forest in Siberia, similar long-

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Fig. 7. (a) Seasonal variation of simulated net radiation, latent heat and rainfall interception overan evergreen forest of northern Thailand in 1998 and 1999. Observed variations of the same parameters except latent heat are also plotted. (b) The cumulative values of the same parameters
and rainfall starting from January of 1998. Vertical bars indicate ranges of 32 simulated values
with the maximum and the minimum, respectively. (Tanaka et al. 2003)

term flux measurements were made in Lenariver basin as part of GAME-Siberia (Ohtaet al. 2001) as shown in Fig. 8. Here, surface
energy partition (sensible vs. latent energy) isstrongly controlled by phenology of the (larch)
forest coupled with seasonal melting (and freezing) condition of permafrost beneath the forest.
Namely, in May during and after the snow meltsensible heating greatly increased, which was
mostly used for melting the frozen ground. At
the beginning of June when the melted layer ofground reaches about 20 cm depth or so, the foliation of trees suddenly started which at thesame time triggered transpiration (latent heating) from the trees. During mid-summer (June,
July and the beginning of August) latent heat-ing rate was comparable or more than sensible
heating rate, and most of the annual precipitation occurs in this period. This condition prevailed for a broad area of Siberian boreal forest
zone, which contributes to water recycling aswell as maintenance of permafrost, by limiting
the melting of permafrost within the nearsurface layer. Thus, the boreal forest and permafrost of this region are likely to co-exist as a
symbiotic eco-climate system through energy
and water cycle.
3.

Interaction between land-surface, ABL
and cloud/precipitation processes

Impacts of land surface processes to the atmosphere are brought about through modifica-

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Fig. 8. Time-series for the daily net allwave radiation and the ground heatflow (a), the sensible heat flux and thelatent heat flux (b), and the Bowen ratio (c) at a site of boreal forest (Taiga)
near Yakutsk, East Siberia. The energybalance components and the Bowen ratio indicated in this figure are for a dry
canopy only. (Ohta et al. 2001)

tion of the atmospheric boundary layer (ABL),
and relevant cloud processes at or above the
ABL. Betts et al. (1998) reviewed the overall
possible interactive processes between land
surface conditions and the ABL, and emphasized the importance of vegetation and related
land cover/use differences on the ABL structure
and cloud/precipitation process. In the monsoon region, the advection or inflow of moist
flow from oceans to the land is sometimes emphasized to produce monsoon rainfall in the in
the interior of the land area. However, throughsome continental-scale experiments (CSEs)

63

under GEWEX (e.g., GAME in Asian monsoon
region, The Large Scale Biosphere-Atmosphere
Experiment in Amazonia (LBA)) moisture supply from wet land surface has also been proved
to be important particularly in the late monsoon season when the surface becomes wetter
than the earlier stage of the monsoon season.
Shinoda and Uyeda (2002) examined a possible
impact of water-fed paddy field in East Asia,
using cloud-resolving atmospheric model. They
forced the model using latent and sensible heat
fluxes over the paddy field observed in GAMEHuaihe Basin Experiment (HUBEX), and foundthat meso-scale cloud systems embedded in the
Meiyu (Baiu) front developed through moisturesupply to ABL and change of moist static stability of the lower troposphere. The change ofland-atmosphere interaction and its impact on
cloud/precipitation systems through the seasonal change of soil moisture (surface wetness)condition was also observed over the Tibetan
Plateau during the GAME-Tibet (Yamada andUyeda 2006). Yamada (2007) simulated the feature over the plateau, showing that dry elevated surface on the Tibetan Plateau is favorable for more convective rain whereas wettersurface condition is favorable for stratiform
cloud but with more rainfall amount as shownin Fig. 9. Understanding of these non-linear
feedbacks from precipitation to land surface
condition may have some clue for prediction ofseasonal as well interannual variation of precipitation in the Asian monsoon region.
4.

Large-scale feedback processes
through energy and water cycle

A key process of regional climate and water
cycle change by the land use/cover change including deforestation may be how the land surface change does induce changes of moisture
convergence ðCÞ and in-situ evapotranspiration
ðEÞ which result in change of precipitation ðPÞ.
This issue can be more generally interpreted asa fundamental energy and water cycle processof an interaction between an area-limited land
surface condition and large-scale atmosphericenvironment or circulation. This issue was discussed in terms of deforestation in the Amazon
basin by Zeng et al. (1996) and Zeng (1998).
Figure 10 shows a schematic diagram of thetwo feedback loops in the perturbed land
surface area in the atmospheric water cycle.

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Fig. 9. Diurnal sequence of (top) the standard deviation of vertical velocity with contour interval of
0.2 ms1 , (middle) hydrometer mixing ration with contour interval of 2:0  102 g kg1 , and (bottom) rain rate at the surface, averaged over the whole experimental domain in the (a) DRY and (b)WET experiments assumed in the summer monsoon season over Tibetan Plateasu. The brokenlines in the top and middle panels indicate cloud edge, defined as the mixing ratio of cloud waterand ice of 1:0  104 g kg1 . (Yamada 2007)

Change in P is controlled by two mechanisms:
C feedback (left hand side) and E feedback(right hand side). In the first mechanism, anincrease in P releases latent heat that drives
large-scale upward motion ðwÞ, which causes
more moisture convergence, leading to more P.
In the E feedback, higher P leads to a wettersurface and more evaporation, which in turncontributes to even more P. If these two positive feedbacks overlap and work at the sametime, a higher sensitivity of P to land surfacechange appears. However, if the P increasecauses decrease in radiation through, for example, cloud cover increase, the increase of C ispartially compensated by decrease of E due todecrease of radiation. An essential issue in the
C feedback may be how P increase contributes
to increase of large-scale upward motion ðwÞ.
Assuming the thermodynamic balance neglecting the horizontal temperature advection,
radiative heating approximated by Newtonian
cooling, convective latent heating proportional
to P, the moisture closure (i.e., P ¼ E þ C), andC proportional to w and humidity q in theboundary layer, the P change ðDPÞ can be approximated as a simple equation as follows(Zeng et al. 1996; Zeng 1998):
DP ¼ q@DT  /ð1  q  bÞ

ð1Þ

where q is a measure of moist static stability

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Fig. 10. The two major feedback loops inthe perturbed region: moisture convergence feedback and evaporation feedback. Here, P is precipitation, w islarge-scale upward motion, C is largescale moisture convergence, wet is surface wetness, and E is evapotranspiration. (Zeng 1998)(unstable if q > 1 ), T  is an atmospheric equilibrium temperature, bð1 Ep/P0 Þ is a ratio ofpotential evaporation Ep to a threshold precipitation rate P0 , above which soil is saturated. @is a constant relevant to radiation relaxation
and atmospheric latent heat profile.
In the tropics near the equator, this positive
C feedback roughly depends upon whether thehorizontal scale of the perturbed area is larger
than the equatorial radius of deformation(Zeng 1998). In the off-equatorial monsoon region, however, the influence of large-scale moisture advection is dominated and the feedback
loops (as shown in Fig. 10) may not be so simple. The impact of deforestation may be apparent through the two positive feedback loops
when the environmental monsoon flow is weak,
as is demonstrated in September case of Thailand (Kanae et al. 2001) as described more in
Section 7. In the relatively dry or marginalmonsoon region (e.g., the north-west part of Indian subcontinent, the Mongolia and northern
China) where the sensitivity of P to soil moisture anomaly is large as discussed in 4.2.2, thepositive feedback loops as shown in Fig. 10 may
be strong mostly due to large b under relatively

65

dry condition. We should keep in mind that relatively humid regions under the Asian summer
monsoon climate, the atmospheric boundary
layer and the lower troposphere are close to besaturated, that is, q is relatively large. Thiscondition also increases the sensitivity of P
under slight change of the surface condition(e.g., albedo), and in turn change of radiational
forcing. The large impact of water-fed paddy
field in China on development of cloud/precipitation system in the Meiyu frontal zone may
correspond to this situation.
This simple model of energy and water balance can be applied in understanding the roleof land atmosphere interaction of large-scale
Asian monsoon system. For example, when wecompare the pre-onset phase and the maturephase of the monsoon season, we notice thatDP between the two season is basically due tochange of T  derived from seasonal change ofsolar radiation, but the enhancement of the Pchange is due to change of q and b, derived
from moisture convergence and in situ evaporation, respectively (Prasanna and Yasunari2007). Without these moist processes, the large
Asian monsoon system with huge P particularly over land area may not be formed as it is.
5.

Relative role of large-scale orography
vs. soil /vegetation cover on mean
Asian monsoon climate

Recently, we have conducted a series of GCM
experiments assessing the role of continental
land surface with and without vegetation, andrelative importance of large-scale orography(e.g., Tibetan Plateau) and vegetation on theformation of Asian summer monsoon. In thiscase, various vegetation types are simply represented by difference in albedo, roughness andfield capacity of surface soil layer (Yasunari
et al. 2006). These experiments do not explicitly
include the physiological processes of vegetation in the model, but the difference in field
capacity of soil layer is expected to hold an important nature of vegetated or non-vegetatedsurface since the soil layer is basically produced
by vegetation itself. The results suggest therole of vegetation is generally very large to produce strong monsoon as it is, and relative rolesof albedo and field capacity of soil have appeared to be different from region to region. In
East Asia both the Tibetan plateau topography

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Fig. 11. Changes of seasonal precipitation for (a) South/Southeast Asia and (b) East Asia simulated
by a CCSR /NIES/FRCGC atmospheric GCM for different boundary conditions of (1) no Tibetan
Plateau, no vegetation (bared surface with rock albedo and roughness) (2) with Tibetan Plateau,
but no vegetation, (3) with Tibetan Plateau and vegetation (soil only), and (4) with Tibetan Plateau
with full vegetation (with soil and realistic vegetation albedo and roughness). (Yasunari et al.
2006)

and vegetation are important for penetration ofmoist monsoon flow and precipitation. In addition, relative importance of albedo vs. field capacity seems to be different from region toregion. In the East Asia the albedo effect of vegetation is more important, but in the South andSoutheast Asia the role of large field capacity
has proved to be more important to produce
the stronger monsoon particularly in the mid
monsoon season, as shown in Fig. 11. These numerical experiments have strongly suggested
that in the Asian monsoon region the role ofvegetation is essential for the formation of
monsoon circulation and precipitation throughstrong latent heating of the atmosphere. The
evapo-transpiration from vegetated surface
plays a crucial role in forming thick moist
boundary layer and convection for the latentheating, as will be described in the next session. The analysis of the same GCM experiments for the northern Eurasia (Saito et al.
2006) proved that the Tibetan Plateau plays
an important role to produce large amplitudeof seasonal cycle with precipitation maximum
in summer and that vegetation plays a crucialrole to enhance in-situ water recycle there to
produce summer precipitation basically from
evapotranspiration.

6.

Land-atmosphere interaction in theseasonal cycle of monsoon

The land shows strong and rapid heating(and cooling) in the seasonal cycle which inturn has a large impact on seasonal atmospheric differential heating (and cooling) processes between land and ocean. The land surface processes which modulate the seasonalheating, therefore, are likely to be responsible,
to some extent, for interannual variability ofthe monsoons.
Seasonal land surface heating and resultant
atmospheric heating over Eurasia manifest
themselves in the surface or lower troposphericpressure field. Figure 12 shows difference ofmonthly mean geopotential field at 850 hPa
from April to May. A remarkable decrease ofpressure is seen over southern Eurasia centered over India/Tibetan Plateau area. Similar
patterns of pressure change are seen over Eurasia from month to month for the period ofMarch through July. In contrast, the pressurechanges over the tropical and sub-tropicaloceans are very small and positive (increaseof pressure). The large decrease of pressurechange over Eurasia indicates seasonal heatingof land, centered over the Tibetan Plateau andJuly 2007

T. YASUNARI

67

Fig. 12. Difference of climatological mean height at 850 hPa and U, V field for May minus April byusing 22 years (1979–2000) data of NCEP/NCAR reanalysis data. Dashed lines show negativevalues. Unit of contour is gpm. Unit of UV field is shown in the right bottom. (Yasunari, 2006)

Mongolia. The decrease of pressure over North
America centered over the Rockies is verysmall, suggesting smaller heating over there
compared to the Tibetan Plateau area.
The low-pressure area over the southernEurasia (called the ‘‘monsoon trough’’) at thesurface and lower troposphere induces moist
southwesterly wind from the Indian Ocean(the southwest monsoon flow) toward India,
Southeast Asia and East Asia, and dry northerly wind from the interior of the Eurasiancontinent. The moist monsoon flow, in turn,induces convection and precipitation over thesoutheastern part of the continent, which plays
a dominant role of atmospheric latent heating
as a ‘‘moist L-AI’’ during the monsoon season.
An important issue may be how this moist
L-AI starts in the interior of the continent. The
sensible heating over the Tibetan Plateau may
be one of the important factors for triggering
this process. In other continents, dry L-AI
based on sensible heating from the land surface
cannot produce the continental-scale monsoons.

Figure 13 shows a schematic diagram of thisdry and shallow L-AI. This process is dominated over the Australian continent and theWest Africa (to the north of West African monsoon) where desert area prevails and largescale downward motion is dominated. The existences of the desert and the downward motion
are self-perpetuating through a positive feedback of surface albedo and radiation balance(Otterman 1974; Charney 1975; Webster et al.
1998). This dry L-AI cannot induce deep atmospheric heating in the interior of the continents, but can trigger shallow coastal monsoon
circulation along the periphery of the continent.
This shallow monsoon circulation, however, islikely to trigger deep convection when thecontinents are surrounded by warm tropicaloceans, e.g., the warm water pool of Arafura
and Timor seas neighboring to the AustralianContinent (Kawamura et al. 2002). The land
heating process over the Australian continentis thus important for triggering the IndonesiaAustralian monsoon.

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Vol. 85B

Fig. 13. Schematic of the circulation between (a) the desert region of North Africa and the Near
East; and, (b) the subtropical ocean regions and the precipitating (pluvial) part of the monsoon
circulation, respectively. The dominant heating or cooling terms are shown in the boxes. ‘‘SH’’ and‘‘LH’’ refer to sensible and latent heating. ‘‘RAD’’ and ‘‘EVAP’’ refer to radiational heating andevaporation, respectively. The pluses and minuses represent the sign of the processes and theirrelative intensity. (Webster et al. 1998)

Over the elevated land surface of the Tibetan
Plateau, in contrast, sensible heating plays adominant role of atmospheric diabatic heatingof the whole troposphere particularly in the onset phase of the Asian monsoon. Based upon
the GAME-reanalysis data, Ueda et al. (2003)
shows strong diabatic heating rate ðQ1Þ in
May through the whole troposphere over the
Tibetan Plateau, suggesting that both sensible
heating through deep dry convection but alsolatent heating through shallow convection in
the pre-monsoon season play an important rolein the heating process even in the pre-onset
phase. The simulation of Asian monsoon by acoupled ocean-atmosphere GCM also suggestthat without the Tibetan Plateau orography
the monsoon precipitation cannot penetrate
into the interior of the continent (Abe et al.
2003).
Diabatic heating over the Tibetan Plateau in
the pre-monsoon phase causes horizontal temperature (pressure) gradient in the upper(lower) troposphere between the surrounding
the Indian and Pacific Oceans and the interiorof the continent, which facilitate large-scaleconvective activity over the South China Sea
through the strengthening of the southwest
monsoon flow intrusion (Ueda and Yasunari1998). In addition, recent studies (e.g., Drbohlav and Wang 2004; Jiang et al. 2004; Wanget al. 2006; Waliser 2006) pointed out an important role of the existence of strong vertical east-

erly wind shear over South Asia to the south of
the Tibetan Plateau in the northward migration of convective zone with intraseasonal
time-scales (including the monsoon onset
phase) from the equatorial Indian Ocean toward the Himalayas. This may imply that theatmospheric diabatic heating over and aroundthe Tibetan Plateau (which is primarily responsible for producing the meridional thermal contrast and the vertical easterly shear over SouthAsia) is one of the essential elements, whichcharacterize the intraseasonal variability as
well as the seasonal march of the Asian summer monsoon system.
7.

Possible impact of anthropogenic land
use/cover change on Asian monsoon
climateThe Asian monsoon region is heavily populated with nearly 60% of the people on theearth are living. Therefore, the anthropogenic
land use and land cover changes have been oneof the largest in the world. These land use/coverchanges from the original vegetation has a highpossibility to have changed regional climate
and water cycle of this region. Fu (2003) investigated potential impacts of human-induced
land cover change on the East Asian monsoon,
assuming the present and the potential vegetation (Ojima 2000). Using a regional model(MM5) with a land surface scheme (BATS), he
noticed that the land cover change occurred inJuly 2007

T. YASUNARI

the history of China may have weakened theEast Asian summer monsoon and strengthened
the winter monsoon. The weakened monsoon
trough in the interior of the continent presumably induced by decreased latent heating is responsible for the weakened summer monsoon.
The potential role of large-scale vegetationand land cover on Asian monsoon has recently
been investigated also by some GCM experiments. Xue et al. (2004) showed that somecharacteristic features of the East Asian summer monsoon, e.g., abrupt northward jump of
rainfall belt in the seasonal march and lowlevel monsoon circulation, are far better simulated when they introduced more sophisticatedvegetation scheme of SSiB with realistic vegetation distribution than a simple vegetation
scheme without physiological processes (i.e.,
photosynthesis and stomatal control of transpiration). The abrupt northward jump of rainfall
could occur due to strong sensible heating in
the interior of East Asia which induced horizontal temperature and pressure gradient in
the pre-monsoon season. This probably was related to suppressed evapo-transpiration in thedry and hot season through stomatal control,
resulting in enhancement of sensible heating.
Suh and Lee (2004) also showed in the regional
climate model experiment that temperatureand rainfall biases over East Asia have improved considerably by including the biosphericprocesses of realistic vegetation over the East
Asian landmass. These experiments have suggested that physiological control of surface energy flux partitioning (i.e., Bowen ratio) greatly
affect pressure gradient over land, and in turn,circulation and rain systems. However, the reality of these biospheric processes needs to bevalidated through in-situ and satellite-based
comprehensive observations.
Xue (1996) focused on the impact of desertification in the Mongolian and the inner Mongolian grassland on East Asian monsoon climate,by changing the area of desertification in the
GCM. He found that the desertification is further intensified by altering from the grasslandtype to the desert-type land surface due to the
reduction in evaporation and convective latentheating above the surface layer. This process
seems to be contrastive to the positive feedbackof desertification in the Sahel, West Africa
where enhanced radiative cooling and associ-69

ated sinking motion induced by increased albedo is likely to be a main mechanism (Charney 1975).
In the Indo-china peninsula, a decreasing
trend of monsoon rainfall has been remarkably
seen only in September during the past several
decades. Kanae et al. (2001) tried to explain
this feature as a result of deforestation in thecentral part of the peninsula using a regional
climate model (RAMS). They noticed that theeffect of deforestation by changing albedo,
roughness and soil moisture condition significantly reduced rainfall only in September
when the monsoon westerly flow seasonally became weak considerably as shown in Fig. 14. Inother monsoon months, the effect of the deforestation was negligible because the effect ofstrong inflow and convergence of moist monsoon flow was dominated over this region.
These results suggested that the impact ofland surface condition, i.e., vegetation cover,soil moisture etc., changes depending upon theatmospheric conditions including large-scalewind field, thermal stability etc.
8.

Concluding remarks

This paper has discussed the landatmosphere interaction associated with the
Asian monsoon climate and its variability. Particular emphasis and focus has been put onwhat are essential and primary roles of ‘‘land’’of the Asian (or Eurasian) continent in the
Asian monsoon system as a huge coupled
ocean-land-atmosphere system.
As major land surface parameters controlling
the Asian monsoon variability, we have discussed roles of snow cover, soil moisture andvegetation. The snow cover and soil moisture
are internal factors of the climate system particularly in the seasonal to interannual timescales, and are likely to play some importantroles in changing large-scale surface energy
and water balance, which in turn affect themonsoon circulation and precipitation. One big
problem is that the data availability of thesetwo quantities as the internal parameters of
the climate system is still limited though the
new satellite data are considerably improving
the situation. Recently, new attempts havestarted on land-surface assimilation using satellite and in-situ observational data and land
surface models (e.g., Pathemathevamn et al.

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Journal of the Meteorological Society of Japan

Vol. 85B

Fig. 14. Observed rainfall trend during 1950–1990 in August (left, upper) and September (left, lower) and rainfall difference in August (right, upper) and September (right, lower) in Indo-China peninsula computed between vegetated condition and deforested condition in a regional atmosphericmodeling system (RAMS). Blue (red) colour shows increasing trend or positive values, and size ofcircles in the left column shows trend values. (Kanae et al. 2001)

2003; Boussetta et al. 2005), which are expected to produce better data for assessing
land-atmosphere interaction in the near future.
The quantitative estimates of the impact ofthese parameters on the monsoon-related fields
need further study, but it may be concluded
how and to what extent these parameters playa role in variability of the monsoon system depends strongly on basic climatic condition andseasonality. From the modeling as well as ob-

servational studies the snow cover is likely to
influence atmosphere essentially through thealbedo effect particularly in lower latitudes
and the Tibetan Plateau. From the GCM experiments it has been strongly suggested thatsoil moisture anomalies efficiently affect the atmosphere under dry or semi-arid condition, but
this effect needs further assessment including
vegetation effect, since soil layer usually exist
with vegetation in the real nature.

July 2007

T. YASUNARI

Another parameter, vegetation, has been
noted as an important variable for the formation of moist monsoon flow over the continent.
Vegetation has several functions for controlling
atmospheric energy and water vapor conditions, such as albedo, roughness, stomatal conductance and water-holding capacity of root/soil
structure. In fact, through the recent observations and modeling studies, it has been noticed
that the atmospheric latent heating over landis fundamentally through vegetation control ofevapo-transpiration.
The role of the Tibetan Plateau is emphasized to trigger the atmospheric heating overthe continent. At the onset or pre-monsoon
phase of the monsoon, sensible heating overthe elevated land surface plays an importantrole to form the continental-scale heat low inthe lower troposphere. This circulation system
in turn produces moist southwesterly flow penetrating into the south, southeast and East
Asia. This moist monsoon flow induces convection and precipitation over the southeasternpart of the continent, which further intensifies
atmospheric diabatic heating through latent
heat release. In addition, we emphasized thatthe moist land surface could play an importantrole for the penetration of precipitation deep interior of the continent. This moist land surface
process is essentially attributed to the existence of vegetation. In this sense, the heavily
vegetated land surface may be another key factor contributing to the huge Asian summer
monsoon system, in addition to the Tibetan
Plateau.
Through the GCM and RCM experiments
the anthropogenically-induced change of land
cover/use including the deforestation has
proved to be a great impact on regional precipitation and water cycle of the monsoon region bychanging some characteristics of vegetation
control of energy and water cycle, which inturn affect ABL and cloud/precipitation processes in regional-scale. The discussion has
also been made on how the land surfacechanges could induce change of precipitation
through feedback processes of moisture convergence and in-situ evapo-transpiration processes. The critical role of moisture amountnear surface and in the ABL is emphasized, toinduce positive feedback to change of precipitation over land in the monsoon region.

71

In the Asian monsoon system, the fundamental heating centers are located over the warm
oceans near the continent. The maintenance of
this system is coupled with the strong thermal
and thermo-dynamical effect of the Tibetan Plateau through the zonally-oriented giant circulation between the South Asia and the equatorial
Pacific region (Kitoh 2002; Abe et al. 2004). It
can be emphasized here that even the coupledatmosphere-ocean system in the tropical Pacific, which is an essential part of the Asian/Australian monsoon system, is strongly controlled by the land-atmosphere interaction overthe Eurasian continent.
Acknowledgements
I would like to thank Dr. Jun Matsumoto for
his careful reading of the manuscript and invaluable comments on this paper. Thanks areextended to Dr. Toshiki Iwasaki, editor in chiefof the special issue for his encouragement oncompleting this paper.
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