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Relative Roles of Large-Scale Orography and Land Surface Processes in the Global
Hydroclimate. Part I: Impacts on Monsoon Systems and the Tropics
TETSUZO YASUNARI
Frontier Research Center for Global Change, Japan Agency for Marine-Earth Science and Technology, Yokohama, Kanagawa, andHydrospheric Atmospheric Research Center, Nagoya University, Nagoya, Aichi, JapanKAZUYUKI SAITO AND KUMIKO TAKATA
Frontier Research Center for Global Change, Japan Agency for Marine-Earth Science and Technology, Yokohama, Kanagawa, Japan(Manuscript received 12 May 2005, in final form 5 January 2006)
ABSTRACTA series of numerical simulations by an atmospheric general circulation model (AGCM) was conducted
to evaluate the relative roles of Tibetan Plateau (TP) and continental-scale land surface processes on the
Asian monsoon and hydroclimates in other regions in the Tropics. Four boundary conditions were used todefine experiments: a flat and bare-rock (nonvegetated) surface (NMR), a realistic TP with bare-rocksurface (MR), MR conditions but with a water-holding soil layer (with 20 cm of field capacity) (MS), andTP with a vegetated surface with specified albedo and roughness of the current vegetation and the soil layer(MVS).
Systematic increases of precipitation (P) over the Asian monsoon region and other land areas in the
Tropics were noticed from NMR, MR, MS, to MVS runs. Precipitation over the warm tropical western
Pacific, in contrast, decreased from NMR to MVS runs although the same SST climatology was prescribedin all the experiments. The effect of TP orography produces wet (dry) climate regions to the west (east) of
TP as shown in other GCM studies. In the Asian monsoon region both land surface effect (LSE) and TP
effect (TE) contribute nearly equally to increase P, but in the West African monsoon region, TE contributes
to decrease P while LSE plays a major role to increase P. LSE also enhances the subtropical anticyclone
over the North Pacific through enhanced land–ocean thermal contrast.
In most of the monsoon regions and Tropics, the albedo (and roughness) effect of vegetation substantially
increases available radiation, which in turn increases both evaporation (E ) and atmospheric moisture
convergence (C), and, as a result, significantly increases P. This positive feedback between E and C is likely
to be a characteristic nature when and where vegetation albedo effect is significant under sufficient radiative
energy, suggesting a prerequisite for Charney’s original hypothesis.

1. Introduction
Monsoons are large-scale circulation systems induced
by seasonal differential heating between land andoceans. Variabilities in monsoon systems can therefore
be attributed to changing characteristics in energy storage and transfer processes between land, atmosphere,
and oceans. Many previous studies have examined theinfluence of oceans on monsoon systems, focusing particularly on the impact of sea surface temperature

Corresponding author address: Tetsuzo Yasunari, Frontier Research Center for Global Change, JAMSTEC, Yokohama, Kanagawa 236-0001, Japan.
E-mail: yasunari@hyarc.nagoya-u.ac.jp© 2006 American Meteorological Society

JHM515(SST) changes, including El Niño–Southern Oscillation(ENSO), regulating large-scale circulation and moisture supplies relevant to the monsoon system.
This study explored how various land surface processes impact the monsoon systems through modulation of land–ocean heating contrast. Many previous
modeling studies have considered the effect of largescale orographies (e.g., the Tibet–Himalayan mountains) (e.g., Manabe and Terpstra 1974; Hahn andManabe 1975; Nigam et al. 1986, 1988; Broccoli andManabe 1992) by removing or changing them. Some
recent studies using coupled ocean–atmosphere generalcirculation models (GCMs) have also suggested thatchanges in the Tibetan Plateau greatly impact the Asianmonsoon system and the atmosphere–ocean coupledAUGUST 2006

YASUNARI ET AL.

system in the equatorial Pacific (Abe et al. 2003, 2004;
Kitoh 2002). Surface vegetation and soil conditions,such as prescribed albedo and field capacity, were constant in these studies.
Other studies that used prescribed orography haveconsidered the effect of vegetation by varying parameters related to vegetation (e.g., albedo, roughness parameters, stomatal conductance) (Sato et al. 1989; Xueet al. 2004). These studies showed that changing surface
vegetation conditions also strongly influence the intensity and time–space structure of the Asian monsoon.
Recent studies have further examined how changing
soil moisture conditions, which are determined by prescribing realistic soil moisture conditions or by changing soil layer characteristics such as the water-holding
field capacity (Manabe et al. 1965) of the surface soil
layer affect rainfall in the monsoon region (e.g., Milly
and Dunne 1994; Dirmeyer and Shukla 1996; Kanae etal. 2006).
However, no attempt has yet been made to evaluate
the relative importance or role of both large-scale orography and land surface processes (including soil, vegetation skin, and subgrid-scale biophysics) on the formation of the monsoon or on the hydroclimate over thecontinent and surrounding oceans. Soil and vegetation
conditions, in reality, constitute a combined land surface system; the soil layer has formed through interactions with vegetation (Jenny 1941) as well as throughphysical and chemical processes such as weathering.
The consideration of relative roles of these two major
land processes/conditions (dynamical effects of largescale orography and physical effects of vegetation) may
also be instrumental for assessing the role of continents
in the evolution of a climate system in the long historyof the earth.
To deduce the effects of different land surface conditions on the monsoon climate system, a simulation
with a “bare rock” surface, representing the nonvegetated “original continent,” should first occur. That
simulation can be followed by simulations that separately consider three effects: orography, soil under thesurface, and a vegetated surface. These three factorsthat control land surface conditions represent the basicstages in the evolution of continental soil and surface
during the earth’s climatic history. Simulations that include present realistic continental surfaces are an integration of all these factors (Buol et al. 1997). Comparisons of results from different numerical settings canprovide quantitative and qualitative estimates of theessential and key role of land as part of the fundamental processes in the real climate system. At this stage ofthe research, we decided to use a medium horizontal

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resolution (ca. 2.85° gridding) and a simple land surface
scheme with bucket-type hydrology with simple bulk
stomatal resistance, and with vegetation albedo androughness. This simple setting may partly distort themodel outputs due to a lack of detailed vegetation biophysical processes and to a coarse resolution. Nonetheless, we consider that the result of this preliminary
study is useful to draw a simplistic picture of the basic
mechanisms on atmospheric energy process and hydrological cycle over land, derived from the same configuration in the land surface scheme with different valuesof relevant parameters (e.g., surface albedo, field capacity, and altitude), and from relatively long integration (50 yr). With this framework, relative effects oforography, soil, and vegetation skin (by specified vegetation albedo and roughness) are investigated, in contrast to Xue et al. (2004) where biophysical effects werestudied. The first part of this paper focuses on the monsoon system through modulation of heating contrast
between land and ocean. The second part discusses seasonal climatology and hydroclimate evolution overEurasia (Saito et al. 2006, hereafter Part II).
This study discusses results from a series of numericalexperiments using an atmospheric GCM. The experimental design is based on the speculation on large-scaleland surface conditions stated above, and is describedin the next section. Section 3 describes the results, focusing on global or hemispheric atmospheric changes;
section 4 details regional aspects of hydroclimatechanges over some monsoon regions and the Tropics. A
discussion and conclusions follow in section 5.

2. Model and experimental design
The model used in this study is version 5.7b of theCenter for Climate System Research/National Institutefor Environmental Studies/Frontier Research Center
for Global Change atmospheric general circulation
model (CCSR/NIES/FRCGC AGCM). The model is athree-dimensional hydrostatic primitive equation(Haltiner and Williams 1980) model in spherical sigma
coordinates. Overall features and performance of theatmospheric part of the model are described in Numaguti (1999). The model included triangular truncationof up to 42 wavenumbers (T42 horizontal resolution);
this truncation is equivalent to a horizontal grid size ofca. 2.5° in longitude and latitude. The vertical axis is asigma coordinate and has 20 levels, up to about 10 hPa
in the lower stratosphere; vertical discretization follows
Arakawa and Suarez (1983). A semi-implicit leapfrog
time integration is used. Ocean surface conditions (seasurface temperatures and sea ice condensation) are pre-

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scribed by monthly climatology for 1979–1989 adopted
from the Atmospheric Model Intercomparison Project(AMIP) (Gates 1992).
Radiation in the model is parameterized using thetwo-stream k-distribution scheme (Nakajima et al. 2000).
Cumulus parameterization is a simplified Arakawa–
Schubert scheme (Arakawa and Schubert 1974; Moorthi and Suarez 1992). Large-scale condensation prognostic cloud water is parameterized as in Le Treut andLi (1991). The second-order Mellor–Yamada level-2
closure scheme (Mellor and Yamada 1974) with simple
cloud effects parameterizes vertical diffusion by turbulence. Orographic gravity wave drag is parameterized
following McFarlane (1987).
The land surface scheme used in this study is basedon bucket-type hydrology as stated in the introduction.
The explicitly controlled factors of the “surface vegetation” processes are vegetation albedo and roughness.
The scheme, however, does not include subgrid-scale,biophysical parameterizations in comparison with “vegetation” models such as the Simple Biosphere Model(SiB2; Sellers et al. 1996), and therefore the assessment
made in this study should be considered as preliminary.
Vegetation types and albedo values (invariant with seasons) are adopted from Matthews (1983). The roughness is specified for each surface vegetation type after
Kondo (1994) and is modified in accordance with snow
amount. The fixed values for daytime and nighttime
conditions are assigned to stomatal resistance.
The land surface scheme has one hydrological layerof 1-m depth and three thermal layers to the depth of 2
m (Numaguti 1999). The water-holding capacity (saturation soil water storage) of the soil is parameterized bya predefined value of field capacity (FC). The groundwetness is prognostically determined by the balance of
precipitation, evapotranspiration, snowmelt, and runoff. Water in the soil layer that exceeds the FC is removed from the system as runoff. The heat conduction
equation determines prognostically the ground temperature profile so that it reproduces diurnal and seasonal variations. Vertical turbulent fluxes at the surface
are estimated by bulk formulas. Aerodynamic resistance is diagnostically determined. The evaporation efficiency is estimated from soil wetness, aerodynamic
resistance, and stomatal resistance (specified as in theprevious paragraph), so that latent heat flux indirectly
includes the effect of transpiration.
Four simulations with different surface boundary
conditions were considered in this study. The bare-rockland surface was considered as a nonvegetated surface,
and therefore the effects of soil/vegetation were divided
into two parts: above the surface and below the surface.
Factors above the surface include albedo and roughness

VOLUME 7

length (Z0). A vegetated surface generally has a lower
surface albedo that affects incoming radiation at thesurface. Similarly, roughness length directly affects
fluxes of momentum, heat, and water vapor between
the boundary layer and the free atmosphere. Factors
below the surface include water-holding capacity (orFC), which represents the most essential function ofvegetation and soil below the surface. The FC directly
controls hydrological flux and indirectly controls heat
flux across the surface through apparent heat capacity.
Last, surface altitude was included to gauge the effectof orography.
The first experiment was the run with bare-rock flat
surface of the Eurasian continent, by removing thelarge-scale orography east of 30°E over Eurasia including the Arabian Peninsula and the Tibetan Plateau. To
mimic bare-rock (i.e., nonvegetated) conditions, FC
was set to 0.01 m, and surface albedo was set to desert
condition (⫽0.3). This was the NMR run. The second
experiment was the same as the first experiment butwith the orography over the Eurasian continent (MRrun). Differences between the MR and NMR runs suggest marginal effects linked to the presence of largescale orography over eastern Eurasia where the land
surface shows desert conditions. The third experiment
was the run with the same conditions as those of the
MR run except the soil surface condition (with FC ⫽0.2 m). This was MS run. Differences between the MS
and MR runs suggest effects related to soil-layer surfaces.
The fourth experiment (MVS run) was the run withthe same conditions as those of the MS run, except thevegetated surface conditions. Albedo and roughness
length at each grid point are given values that correspond to the prescribed realistic vegetation type at thegrid point. Albedo varied from 0.11 to 0.28, and roughness length varied from 0.003 to 2.0 m for different
vegetation types. At all land grid points FC was set to
0.2 m. This run was, therefore, closest to the reality interms of surface boundary conditions.
Figure 1 shows experimental designs for the NMR,
MR, MS, and MVS runs, and Table 1 summarizes theparameter values of each experiment. Figure 2 shows
spatial distribution of albedo and orography for the two
typical cases, that is, (a) MVS run and (b) NMR runs.
The models in each of the four experiments were integrated for 50 yr following a 5-yr spinup. Monthly orseasonal “mean climate” values were computed as averages of the last 40 yr; pentad “mean climate” values
represent averages of the last 25 yr. Mean climates foreach experiment were used to deduce the impacts ofland surface conditions.

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FIG. 1. Schematic illustration of vegetation, soil, and Asian orography for four
experimental settings.

3. Changes in the planetary-scale monsoon system
Changes in land surface conditions affect planetaryscale monsoon systems during the Northern Hemisphere summer. Figure 3 shows precipitation for (a)NMR, (b) MR, (c) MVS, and (d) the observation takenfrom the Climate Prediction Center (CPC) Merged
Analysis of Precipitation (CMAP) data (Xie and Arkin1997). Figure 4 shows the difference map of June–
August (JJA) precipitation and wind field at 850 hPa
for (a) MR minus NMR and (b) MVS minus MR. Thezone of precipitation between 10° and 15°N was
considered an intertropical convergence zone (ITCZ)
in NMR (Fig. 3a). In this ITCZ, more precipitation(⬎400 mm month⫺1) fell over oceans than over land.
Heaviest amounts fell over the Pacific, with somewhat
lighter amounts occurring over the South China Sea
and the Bay of Bengal in the Asian monsoon region.
Precipitation was light (less than 100 mm month⫺1)
over most of the continent. When there is no Tibetan
Plateau (TP), no vegetated surface with high albedo,
and near-zero FC over land, overall monsoon effects
are very weak (Fig. 3a). The MR run (Fig. 3b) included
a TP but no vegetated surface. Precipitation in thissimulation increased over the Bay of Bengal and a zoneof light precipitation (around 200 mm month⫺1) appeared over land, centered south and southeast of the
TABLE 1. Summary of the land surface conditions in thesimulations.

MVS
MS
MR
NMR

Vegetation
type
at grids

AlbedoField
AsianRoughness capacity(m)(m m⫺1) mountains

SpecifiedDesertDesertDesert

0.11–0.28
0.3
0.3
0.3

0.003–2.0
0.02
0.02
0.02

0.2
0.2
0.01
0.01

With
With
With
Without

TP as shown in Fig. 4a. Orographic effects of the TP areimportant in the formation of the Asian summer monsoon. Note, however, that precipitation amounts particularly over land in this MR simulation were lighter
than observed values. The MVS run (Fig. 3c) showedthat heaviest precipitation appears over the Bay of
Bengal, but the ITCZ over the Pacific Ocean was drastically weakened. A broad area of south, southeast, andeast Asia received heavy precipitation (exceeding 300
mm month⫺1). This distribution over land is very similar to that observed in the summer monsoon season(Fig. 3d).
Figure 5 shows global streamfunction and wind field
except south of 30°S in the lower troposphere (850 hPa)
for the three runs (NMR, MR, and MVS). The difference fields of these parameters between MR and NMR
and between MVS and MR are also shown in Fig. 6. In
the NMR run (Fig. 5a) although cross-equatorial flow
and monsoon-like westerly flow is dominated over thenorthern tropical Indian Ocean, a center of convergence associated with precipitation maximum (see Fig.
3a) is located over the western Pacific. A cyclonic circulation corresponding to the monsoon trough over the
Indian subcontinent through the Indo-China Peninsula
is so broad and no southwesterly appears over southeast Asia and China. In the MR run (Fig. 5b) the southwesterly monsoon flow is now dominated over the Arabian Sea, and southwesterly appears over the IndoChina and to the southeast of TP, and the southerly
component, though weak, appears over China. The difference field between MR and NMR (Fig. 6a) clearly
demonstrates anomalous anticyclonic circulation overnortheastern Eurasia, which is formed primarily
through deflection of midlatitude westerly flow by thelarge-scale orography over there. Another remarkablechange between MR and NMR is the appearance of
cyclonic circulation to the west of TP as clearly shown

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FIG. 2. (a) Albedo maps for vegetated land surfaces (shades) and the Asian orography
including Tibetan Plateau (contours) used for MVS run. Contour interval is 500 m. (b) As in(a), but for a desert land surface without the Asian orography used for NMR run.

in Fig. 6a. This may be due to a stationary Rossby wave
response to dynamical and thermodynamical effect of
TP as Rodwell and Hoskins (1996) discussed. In the
MVS run (Fig. 5c) the overall feature as shown in Fig.
5b is further intensified. Southwesterly monsoon flow is
now dominated to the southeast of TP and east Asia,
and westerly flow over the western Pacific is weakened.
Of particular interest may be a further intensification of
the subtropical high over the North Pacific, as is clearlyshown in the difference field between MVS and MR(Fig. 6b). This may be closely related to the intensification of surface and atmospheric heating (as suggested
in Fig. 4b, and as shown in Fig. 3 of Part II) over thevegetated Eurasian continent, presumably through the

Rossby wave response originating to the west of themonsoon heating (Rodwell and Hoskins 2001).
Figure 7 shows global streamfunction, except southof 30°S, in the upper troposphere (200 hPa) for thethree runs. In the NMR run (Fig. 7a), a subtropical
anticyclone dominated over the western Pacific, a region that corresponded to the center of convection andheating (Fig. 4a). A minor anticyclone also developed
over Mexico. In the MR run (Fig. 7b), features over andaround Asia were dramatically different from those in
the NMR run. A very large subtropical anticyclone
formed over the subtropics of Eurasia, centered over
the TP. This anticyclone corresponds to the “Tibetanhigh” in the present climate system. The differenceAUGUST 2006

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also stronger in MR than in NMR even though orographic changes occurred only over Eurasia (mainly
TP). Tibetan Plateau orography may force a planetaryscale stationary wave that also intensifies the anticyclonic circulation over the Mexican highlands. This effect is also shown in the difference field between MR
and NMR in Fig. 8a.
In the MVS run (Fig. 7c), a broader Tibetan anticyclone over Asia and deeper trough over the central
North Pacific further enhanced the overall circulation
pattern. Difference fields between MVS and MR runs
(Fig. 8b) show overall northward shift of the upper anticyclonic circulation particularly over the eastern Eurasia, showing further enhancement of the planetaryscale monsoon circulation.

4. Changes in regional-scale hydroclimates andsurface budgets
a. Changes in regional-scale surface energy budgets

FIG. 3. Map of precipitation (contours) and wind at 850 hPa(arrows) for (a) NMR, (b) MR, and (c) MVS runs, and (d) observation from CMAP for JJA. The contour interval is 50 mm
month⫺1.

field (Fig. 8a) shows enhancement of anticyclonic circulation to the west TP, suggesting that dynamical andthermodynamical effect of TP enhances the Tibetan
high through the stationary Rossby wave response(Rodwell and Hoskins 1996).
The anticyclonic circulation over North America was

Regional energy and water budgets change as orography and land surface conditions (i.e., NMR, MR, MS,
and MVS) change. This study focused on changes over
six land areas in the Asian monsoon region and othertropical regions shown in Fig. 9: south/Southeast Asia(SA), east Asia (EA), West Africa (WA), the Indonesian Maritime Continent (MC), Amazonia (Am), andSiberia (Sb).
Table 2 shows that the overall effect of vegetation
albedo and roughness on the energy budget (as inferredfrom changes from MR to MS, MS to MVS) is to reduce
sensible heating (H ) and increase latent heating (LE).
In south/Southeast Asia and east Asia, the ratios of LE
to net radiation (R), LE/R, increase from NMR to MR
due to moisture supply increase by invasion of monsoon flow. No significant changes appear in MC and in
Sb. The increase of FC from MR to MS dramaticallyincreases LE/R in all the regions through increase ofevaporation efficiency. In the monsoon regions increase of FC seems to be more effective in EA (LE/R
increases from 0.58 to 0.82) than in SA (LE/R increases
from 0.77 to 0.89) in forcing evaporation because of
more effective soil moisture memory under relatively
dry conditions. From MS to MVS increases of R arecommon in all the regions, though the ratios LE/R do
not change significantly. The overall effects of vegetation albedo and roughness (as inferred from thechanges of LE and LE/R from MR to MVS) are significant particularly in the tropical rain forest regions(MC and Am), and in mid-/high latitudes of Asia (EA
and Sb), where the increasing rate of LE and LE/R
range from 180% to 230%. The rate in the tropicalmonsoon regions (SA and WA) is relatively small

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FIG. 4. (a) Difference map of JJA precipitation (contours) and 850-hPa wind (arrows)
between (a) MR and NMR and (b) MVS and MR. The contour interval is 40 mm month⫺1.
Zero contours are suppressed. Only wind differences with speeds exceeding 1 m s⫺1 areshown.

(around 130%), since the original LE values are already high in MR.

b. Changes in regional-scale water budgets
Figure 10 shows changes in the seasonal course of
precipitation in each area. The observed precipitation
climatology is also shown in each region, adopted from
CMAP. Table 3 shows changes in atmospheric water
budgets, during the same seasons as in Table 2.
The overall features of MVS precipitation in thetropical monsoon regions (SA and WA) very wellmatch the present precipitation climatology both inquantities and in seasonal patterns. Seasonal march of

EA in MVS is also similar to that observed, though theamount in rainy season is overestimated to a small extent by about 20%. We may assume, therefore, the
MVS corresponds to the most realistic simulation of thecurrent Asian monsoon system. Precipitation in NMR
in the Asian monsoon regions (SA and EA) shows
about a half or less of that in MVS. This suggests that aland–ocean heating contrast without TP and land surface processes explain only a half or less of the precipitation in the realistic Asian monsoon system. Once TP
topography is introduced, precipitation over these
monsoon regions increased to about two-thirds of thatin MVS (as referred from differences between MVSAUGUST 2006

YASUNARI ET AL.

FIG. 5. JJA streamfunction and wind at 850 hPa for (a) NMR, (b) MR, and (c) MVS.
Contours are every 10 ⫻ 106 m2 s⫺1.

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FIG. 6. As in Fig. 5 but for streamfunction and wind at 850 hPa. Contours are every 2 ⫻ 106
m2 s⫺1. Zero contours are suppressed. Only wind differences with speeds exceeding 1 m s⫺1
are shown.

and MR). Atmospheric heating over and around TP
plays an essential role in producing the monsoon flow,by enhancing cyclonic circulation (monsoon trough) as
shown in Fig. 6a.
A distinct difference between SA and EA arises incomparisons of MR, MS, and MVS runs. As is also
discussed in section 4a, in EA the increase of FC increases E more effectively (as inferred from the difference between MS and MR) compared to SA. That is,
the amount of E in EA is comparable to that in SA for
the MS and MVS run, but it is only two-thirds of thatfor the MR run. The ratio of E to convergence (C), E/C,
for the MVS run exceeds 2 over EA, but it is only 1.1
over SA. Thus, the soil/vegetation layer over EA is

more likely to increase precipitation efficiently by enhancing E from the wet soil/vegetation layer.
On the other hand, the change from MR to MS condition (i.e., increase of FC) may induce a negative feedback between E and C over EA, while a positive feedback tends to occur over SA. This positive feedback
between E and C may be a characteristic nature ofhumid Tropics where increase of E could enhancelarge-scale convection, which, in turn, enhances convergence (Dirmeyer and Shukla 1996; Zeng 1998).
Differences of precipitation over WA between thefour runs differ considerably from differences in the
Asian monsoon regions. Interestingly, precipitation in
MR (with TP) is lighter than that in NMR (without TP).

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FIG. 7. JJA streamfunction at 200 hPa for (a) NMR, (b) MR, and (c) MVS. Contours are
every 10 ⫻ 106 m2 s⫺1.

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FIG. 8. As in Fig. 7 but for 200 hPa. Contours are every 3 ⫻ 106 m2 s⫺1.

The TP orography plays an essential role in enhancing
downward motion and dry climate conditions overwestern Africa. The cause is likely an enhanced east–west circulation between Asia and Africa, and interaction with the stationary Rossby wave to the west of
monsoon heating, as suggested in previous modelingstudies (e.g., Manabe and Terpstra 1974; Rodwell and
Hoskins 1996).
Over MC, the effect of TP orography seems to be
negligible (from the difference between NMR and MR
shown in Fig. 10d), and a systematic increase of precipitation from MR, MS, to MVS is noticeable, mainly
due to increase of E (Table 3). The land surface effect(FC, vegetation albedo, and roughness) included in the
MVS run suggests the positive feedback between E andC, similar to the other tropical regions. However, allfour runs have failed to simulate observed seasonal
cycle with a maximum precipitation in northern winter.
This systematic error may, at least partly, be related tolack of spatial resolution in the complex geography ofthis region that winter monsoon rainfall over this region
cannot be fully reproduced in the relatively coarse
model resolution. Over Am in the equatorial SouthAmerica, the overall changes of precipitation from MR,
MS, to MVS (Fig. 10e) are similar to those over MC,
indicating the dominant role of vegetation in increasing
precipitation. The MVS run well simulated precipitation with its seasonal cycle except early for summer
months (November and December).
Changes in precipitation, energy, and water balances
over Sb are also shown in Fig. 10f and Tables 2 and 3,
respectively, to compare with hydroclimate responses

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FIG. 9. Definition of six land areas used for analysis.

in the Tropics. The seasonal cycle is too much enhanced, particularly with a large maximum precipitation in northern summer when the TP orography isintroduced. In addition, summer precipitation increases
dramatically when vegetation albedo and roughness are
introduced (MS and MVS runs). In midsummer (July toAugust) the FC effect is important for increasing precipitation (because of enhanced evaporation). In spring
to early summer (March to June), albedo effects aremost important, presumably due to soil moisture supplied by increased snowmelt. However, caution may beneeded to discuss the large precipitation and evaporation in summer, because these seem to be extremely
large compared to the observations as shown in Fig. 10f.
These systematic biases in summer water balance may
be, at least partly, related to the exclusion of some important effect of vegetation (e.g., full stomatal resis-

tance and/or leaf area index) in the land surface processes. Sato et al. (1989) and Xue et al. (2004) showed,
for example, that evapotranspiration and associatedprecipitation over the forested area in northern continents were considerably suppressed by adopting the
SiB scheme in the GCM simulation. In addition, the“warm bias” in surface temperature over the highlatitude continental area in the GCM simulation (cf.
section 3 in Part II) may also be attributed to larger
evaporation than the observed. Further details on energy and water balance changes in the cold regions of
the continent will be discussed in Part II.

5. Discussions and summary
A series of AGCM simulations were conducted toevaluate the relative roles of large-scale orography (i.e.,

TABLE 2. Surface energy budget for different regions. Averages are taken over the rainy season for each region: MC (annual), SA[Jun–Sep (JJAS)], EA (JJAS), Am (annual), WA [Jul–Sep (JAS)], and Sb [Jun–Aug (JJA)]; R, LE, and H denote net radiation andlatent and sensible heat flux, respectively, in W m⫺2. Net upward flux is shown as a positive value.
MVS

MC
SA
EA
Am
WA
Sb

MS

MR

NMR

R

LE

H

R

LE

H

R

LE

H

R

LE

H

⫺148
⫺146
⫺146
⫺148
⫺137
⫺135

116
121
117
96
98
96

31
26
22
52
41
23

⫺118
⫺123
⫺119
⫺112
⫺118
⫺101

101
109
98
67
83
71

16
15
15
44
38
13

⫺109
⫺119
⫺108
⫺103
⫺117
⫺98

68
92
63
42
72
41

41
28
38
60
46
37

⫺114
⫺114
⫺102
⫺104
⫺121
⫺98

64
70
45
46
87
41

50
44
53
57
34
40

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FIG. 10. Seasonal changes in precipitation for different regions from January to December. Thick solid, dashed, dotted, and thinlines denote MVS, MS, MR, and NMR runs, respectively. Thin dashed–dotted line is observation from CMAP.

the Tibetan Plateau) and land surface processes (represented by vegetation albedo/roughness and soil layer)on the hydroclimate of monsoon regions and the Tropics. Oceanic influences were set to be similar in eachrun, by using the monthly climatological SSTs in eachrun.
Hemispheric hydroclimate and circulations, and regional-scale energy and water balance, were comparedAUGUST 2006

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TABLE 3. Atmospheric water budget for different regions. Averages are for the same period as in Table 2 for each region; C, E,
and P denote convergence, precipitation, and evaporation, respectively, in mm month⫺1.
MVS

MC
SA
EA
Am
WA
Sb

MS

MR

NMR

C

E

P

C

E

P

C

E

P

C

E

P

62
115
51
40
51
5

121
125
121
99
102
100

182
241
172
139
152
105

39
112
32
14
35
14

105
113
101
70
86
74

145
225
133
84
121
88

43
100
55
19
24
12

70
95
66
44
74
43

114
196
120
63
98
54

25
60
29
22
34
10

66
72
46
48
90
42

91
133
76
70
125
53

in the numerical experiments with the following four
land surface conditions: a flat bare-rock surface(NMR), a bare-rock surface with orography (MR), a
desert surface with soil and orography (MS), and a surface with realistic albedo/roughness with soil and orography (MVS). Major results are summarized as follows:
1) In the NMR run, the spatial distribution of precipitation during the northern summer monsoon was
basically characterized by the ITCZ with a maximum over the warm western North Pacific (Fig. 3a).
In the MR run (Fig. 3b), a more monsoonal pattern
appeared, with maxima over the Bay of Bengal,
along the southern and eastern slope of TP, andover eastern Asia. Precipitation over the warm western Pacific decreased compared to the NMR run(Fig. 4a). The overall spatial pattern of precipitation
was similar to the current climate, but the amounts
of precipitation in the monsoon region were about
60%–80% smaller than those in the MVS run andobservations. Adding the albedo and roughness associated with the vegetation in the MVS run (Fig.
3c) further increased precipitation, particularly
south and east of the TP, over the whole tropicalland areas, and over boreal forests in northern Eurasia. Simultaneously, precipitation decreased further over the warm northwestern Pacific (Fig. 4b).
2) The TP generally enhanced components of the
Asian monsoon circulation, deepening the monsoon
trough over India (Fig. 5b) and the Tibetan high in
the upper troposphere, with the northwest shift ofthe center position (Fig. 7b). These changes in thecirculation (Figs. 6a and 8a) were basically interpreted as a stationary Rossby wave response to localized monsoonal heating over Asia (Rodwell and
Hoskins 1996). In contrast, adding a vegetated surface with soil layer to Eurasia (the MVS run) resulted in a more selective enhancement of the subtropical anticyclone over the North Pacific and thenortheastward extension of the Tibetan high in theupper troposphere (Figs. 6b and 8b). A continental
vegetated surface is, therefore, more likely respon-sible for the large-scale heating contrast between thecontinents and the oceans, which may be a primaryfactor for formation of a subtropical anticyclone
over the oceans (e.g., Rodwell and Hoskins 2001;
Abe et al. 2004).
3) Regionally, the effects of TP on water budgets, as
deduced by comparing the MR and NMR runs
(Table 3), resulted in enhancement of precipitationin the Asian monsoon region (SA and EA), andreduction of precipitation in the West African monsoon region (WA). Enhanced downward motion
over western Africa was associated with an enhanced large-scale, east–west circulation induced by
TP orography, as noted by the previous GCM experiments (e.g., Broccoli and Manabe 1992; Rodwelland Hoskins 1996). Also, TP increased precipitationin the Maritime Continent, by enhancing convergence. Basically TP effect was very small in the water budget of northern Eurasia.
4) The sole effects of soil layer (inclusion of FC with 20cm), as deduced from differences between the MS
and MR runs (Table 3), appeared to be the increaseof precipitation in all the regions from the Tropics tohigh latitudes. The increase of precipitation was primarily attributed to increase of evaporation, except
for the tropical monsoon regions (SA and WA)
where convergence also increased to some extent.
The increase of FC has two effects: the surface temperature decreases because of an increase in heat
capacity in the near-surface layer, and evaporation
increases as available soil moisture increases. The
net effect of increased FC is to increase net radiationand to increase evaporation (decrease the Bowen
ratio, ratio of sensible heat flux to latent heat flux)
as shown in Table 2. The decrease of surface temperature may tend to stabilize the atmosphere andreduce convergence in most of the regions.
5) The effects of vegetated surface (with albedo androughness) are deduced by comparing the MVS andMS runs. A fundamental component of this effect is
to increase available radiation as seen in the in-

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JOURNAL OF HYDROMETEOROLOGY

crease of R (Table 2), and increase of E (under thelarge FC of the MVS run), which in turn increases
precipitation (Table 3). It is interesting to note thatin most of the regions in the Tropics and midlatitudes where available solar energy are large enough(as suggested from R in Table 2), both convergenceand evaporation increase, and as a result, precipitation increases considerably. This positive feedback
between evaporation and convergence (Dirmeyer
and Shukla 1996; Zeng 1998) is likely to occur easily
when and where radiative energy is sufficient andvegetation change is significant. This may imply that
Charney’s original hypothesis (Charney 1975; Charney et al. 1977) that the albedo significantly changes
circulation and precipitation may be valid or effective only under this condition. In higher latitudes
like Siberia (Sb), the increase of available energy
solely increases evaporation, which is a primarysource of precipitation there.
6) It may be worthwhile to roughly examine the relative impact of land surface processes (MVS run minus MR run) and that of TP topography (MR run
minus NMR run). Note that the comparison is only
preliminary and marginal, for the full evaluation
may need to consider the nonlinearity and several
feedbacks in the climate system. We presume that,as compared to the TP orography effect (TE), thesurface vegetation (vegetation albedo and roughness) and soil layer (large FC) together should constitute a combined effect [land surface process effect(LSE)], since in reality vegetation and soil form an
integrated surface system on land as mentioned in
the introduction. Thus, the differences in precipitation between the MVS run and MR run, and thosebetween the MR run and the NMR run could be aproxy measure of the LSE and TE, respectively.
In the Asian monsoon region (SA and EA), both
LSE and TE values are around 50 (mm month⫺1),suggesting that both land surface processes and TP
contribute nearly equally to increase precipitation inthis region. In MC and Am, LSE (⬃70) far exceeds TE(⬃0 to 20), suggesting LSE is more fundamental in thetropical rain forest regions. In Amazonia, the TE of
the Andes, in fact, may play an additional role for precipitation, though this effect is implicitly included in
the experiments here. In WA, LSE (⫽54) is as large asin the Asian monsoon region, though TE (⫽⫺27)
considerably decreases precipitation. In Sb, precipitation in summer is basically attributed to LSE (⫽51),
where TE is negligible but TE substantially induces
large seasonal cycle with precipitation maximum in

VOLUME 7

summer. Further description and discussion on changes
in hydrology and hydroclimate in high latitudes will bemade in Part II.
In this study, the present-day climatological SSTs
were used, and because of warm oceans around thecontinents about a half of observed precipitation fellover most of the tropical and subtropical land regions
even without TP and vegetation. However, some recent
studies (e.g., Kitoh 2002; Abe et al. 2004) using coupledatmosphere–ocean (AO) GCMs have shown that thechanges in land surface conditions, such as TP orography, change global SST distribution through radiative,
thermodynamic, and dynamic feedbacks between land,
atmosphere, and oceans (e.g., Kitoh 2002; Abe et al.
2004). Abe et al. (2004), for example, showed that without TP the warm water pool in the western Pacific (andthe Maritime Continent) could not be formed. It couldbe, therefore, that weaker heating conditions in simulations with no TP and/or less vegetation may further
suppress precipitation over land by cooling the warm
pool region.
Another issue that warrants further study will be anevaluation with more realistic vegetation schemes, including stomatal control of transpiration. Over thetropical rain forest areas (MC and Am), for example,
the simulated precipitation shows some systematic errors in the seasonal cycle of precipitation comparedwith observations (Figs. 10d and 10e). Over Sb, theamount of precipitation in summer season in the MVS
run (Fig. 10f) is too large compared to the observed
amount. Such discrepancies between the model results
and observations may, at least partly, be related to systematic errors in the simple surface soil/vegetation
scheme used here that does not allow realistic physiological control of energy and water budgets of thesurface. Sophisticated vegetation biophysical schemes[e.g., SiB, biosphere–atmosphere transfer scheme(BATS), etc.] will provide skillful evaluation on thedetailed variations of evapotranspiration [e.g., its components (transpiration/interception evaporation/
ground surface evaporation), differences with vegetation types, short-term variations] with the photosynthetic processes, the canopy effects, and the multilayer
treatment of soil moisture.
These remaining issues suggest that we pursue further investigation using a coupled AOGCM that incorporates a more sophisticated biosphere scheme, toevaluate the integrated effects of large-scale orography
and full vegetation effects.
Acknowledgments. The authors would like to thankthe anonymous reviewers for constructive suggestions
and comments.

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YASUNARI ET AL.
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