Jountal of the MeteorologicaL Society of Japan, Vol. 78, No. 4, pp. 461-.175, 2000

Diurnal variation of precipitation
Two-dimensional

over the Indo-China Peninsula:numerical simulation

By TakehikoSatomura

Diuis'ionof Earth and Planetary Science,GraduateSchoolof Science,Kyoto Un'iuersity,Kyoto, Japan(Manuscnpt receiued 20 September 1999, in reuisedfor"m 15 April 2000)

Abstract
The diurnai variation of precipitation over the Indo-China Peninsula is investigated using a twGdimensional, nonhydrostatic, and cloud-resolving numerical model. The model is initialized by the climatologicai monthly mean vertical profiles of zonal flow, temperature and humidity at the center of the
Indo-China Peninsula in June.
The model successfuily simulates the diurnal variation of precipitation. The simulated diurnal variation
is as follows: convection is activated at the lee-side foot of two mountainous regions located at the west
and middle of Thailand in the evening of each day; activated clouds then are organized into squall linesthat travel eastwa.rd during the night at about 5-10 m s-t. These squall lines weaken around midnight.
It is concluded that the sola.r-synchronized life cycle of the squall lines and their eastward movement
cause the nighttime maximum of the precipitation over the inland area of the Indo-China Peninsula. High
resoiution analysis of a convective activity index calculated from the GMS IR data over the Peninsula
supports this conclusion.

1. Introduction
Convective cloud systems in the tropics releasehuge amounts of latent heat and play an importantrole in driving the global circulation of the atmosphere. The temporal and spatial distributions of
the cloud systems determine the time and place ofthe heat release. In the tropics, the diurnal variation
is a prominent time scale.
Diurnal variability of precipitation over land in
the tropics and mid-latitude summer have beenstudied observationally using rain gage networks(e.g., Wallace 1975; Schwartz and Bosart 1979;
Landin and Bosart 1985; Riley et al. 1987; Oki and
Mushiake 1994) and radars (e.g., Karr and Wooten
1976; Saito and Kimura 1998). Although a rain gage
network observesprecipitation directiy, it occasionally misses local precipitation. A long time seriesof
rain gauge measurementsreduces this deficit. Using
hourly rain gauge data extending over 7-35 years,
Wallace (1975) and Riley et al. (1987) studied diurnal variations of precipitation in both warm andcold seasons in the United States. Both papers
Corresponding author: T. Satomura, Division of Eartharrti Plalci,aly Scierrc", Gr'aciua'ueSchoui oi Seierrue,i(yoto University, Kitashiraka,wa Oiwakecho, Sakyo-ku, Kyoto 606-8502, Japan.
E-mail:satomura@kugi.kyotou.ac.jp.
@)2000, Meteorological Society of Japan

found that, in the warm season, the phase of diurnal precipitation cycle has a prominent maximum
in the evening (around 1800 local time) in the leeof the Rockies. Their analysis also revealed that,as one moves eastward, a clear transition to a nocturnai maximum occured over the high plains during the warm season. Wallace (1975) suggestedthat
mountain-generated precipitation systernstended tomove eastward over the Central Plain after sunset,
but didn't identify the type and mechanism of thesystem. Using radar data in Japan, the fact that thephase of daily precipitation cycle in summer has anocturnal maximum was also observed in the northwestern part of the Kanto Piain near the mountains(Saito and Kimura 1998). Likewise, they did not
discussthe attendant mechanisms in detail.
Oki and Mushiake (1994) analyzed the seasonal and diurnal cycle of tropical precipitation(Maiaysia) in addition to precipitation in Japan
using hourly rain gage data extending more than
10 years. A pronounced peak ofrainfail at 1600 local
time in inland regions of the Maiay Peninsula and amorning peak on the windward coast for the monsoon wind were found. The eastward or westward
shift of the maxrmum phase of cirurnal preclpttatlon
was not analyzed in the inland regions. it should benoticed that the width of llalay Peninsula is about
300 km. which is much shorter than the distance462

Journal of the Meteorological Society of Japan

from the Front Range of the Rockies to the central
United States.
Infrared outgoing radiation or infrared equivalent black body temperature (?s6) is used to study
large-scalecloud activity. An index of ?66 , which
selects ?66 lower than a threshold value and averages over several degrees in iongitude and latitude, corresponds well with precipitation (e.g., Reed
and Jaffe 1981; Richard and Arkin 1981; Arkin and
Meisner 1987); it is used to study diurnal va,riabiiity of precipitation over a wide area (e.g., Reed andJaffe 1981; Albright et al. 1985; Duvel and Kandel
1985; Meisner and Arkin 1987; Janowiak et al. 1994;
Nitta and Sekine 1994; Harada et al. 1998). By
these studies, the continent/ocean contrast is now
weil known; the diurnal cycle of precipitation has
a 1800 to 0000 local time maximum over the tropical continents and a morning maximum over the
ITCZ. Using three-hourly infrared data from a geostationary satellite, Meisner and Arkin (1987) reported an ea.stwardphase shift of the diurnal cycle
near the eastern boundary of the Rocky Mountains
similar to the shift analyzed by rain gages (Wallace1975). Asai et al. (1998) analysed diurnal variabilityof cloudiness over East Asia during the warm season
and aiso showed that a systematic phase delay similar to the east of the Rocky Mountains appeared
to the east of the Tibetan Plateau. It is also possible through the analysis of Nitta and Sekine (1994)
to distinguish the analogous eastward phase shift ofthe diurnal cycle of precipitation in July over the
Indo-China Peninsuia, although this phase shift was
not discussedin their paper.
Nowadays, a numerical model is a powerful tool.
The accuracy of amount, location and time of precipitation in global numerical models is of great importance in simulating the climate of the Earth.
Randall et al. (1991) examined diurnal variability of
precipitation in a general circulation model (GCM)
and concluded that the model simuiated the afternoon maximum over the land in warm rainy seasons
and an eariy morning maximum over the oceans far
from land. Apart from the causes which they presumed (see recent arguments of Sui et al. 1998, or
Liu and Moncrieff 1998), the early morning maximum over the ocean apparently agreed with observation. The afternoon maximum over the land
discorded with the analyses using rain gages (e.g.,
Wallace 1975) and Ten (e.g., lv{eisner and Arkin
1987; Nitta and Sekine 1994), however; the reasonof this discrepancy is not understood.
A few papers on the diurnal variation of precipitation over the Indo-China Peninsula have been published. Although the analyzed area of Murakami(1983) included the lndo-Uhina Penrnsuia, he focused on the diurnai variability over the ocean andthe Tibetan Plateau. Nitta and Sekine (1994) also
analyzed the wide area included the Indo-China

Vo]. 78. No. 4

Peninsula, but focused on the diurnal variability
over the ocean. Both papers used infoared irradiance data with a resolution of 1" longitude-latitude
square mesh of comparable resolution to GCMs.
Their analyses are. therefore, unable to detect systems too small to be simulated properiy in GCMs
that may contribute the night maximum of diurnal
variation ofprecipitation over the tropical continent.
Infrared irradiance data of 0.1o resolution were analyzed over the Indo-China Peninsula by Haradaet al. (1998). They discussedthe iand/ocean contrast, spatial variation over the Peninsula, and an
intraseasonalvariation of the diurnal cycle of Tes.
They also noticed a movement of cloud systems from
north to south. The mechanism which producesnight maximum of diurnal precipitation cycle was,however, not fully discussed.
The purpose of this paper is to study the diurnal
variation of precipitation over the Indo-China Peninsula during the monsoon period using a numericalmodel. To describe cloud behavior accurately, anonhydrostatic cloud-resolving model is employedwith cloud microphysics and interactive radiation.
Since the statistical behavior of the diurnal cycle ofprecipitation is the focus herein, the initiai conditionof the model is taken from climatic data.
model and experirnental

2. Numericalsetup

The nui.nericalmodel used in this study is basedon RAMS (Regional Atmospheric Modeling System) developed by Colorado State University andthe .ASTER division of Mission Research Corporation. The detailed description of RAMS wasgiven by Pielke et al. (1992). It was configured
as a two-dimensional, nonhydrostatic and compressible equation model in our study. Surface-layer
fluxes were parameterized using a vegetation model(Avissa"r and Pielke 1989). Vegetation type was"evergteen broad leaf forest"l thatspecified as theis, one of 18 possible types based on the BATS
(Biosphere-AtmosphereTransfer Scheme;Dickinson
et al. 1986). A bulk microphysical parameterization including the ice phase was used to simulate
clouds. A radiation code adopting the DOM/adding
algorithm and a semi-random overlapping model(Nakajima et al. 1995) was used in place of the
RAMS radiation parameterizations.
The horizontal grid spacing is 935 m, whiie thevertical spacing is stretched from 25 m at the lowest
level to a maximum of 500 m. The verticai grid is aterrain-following coordinate system (Gal-Chen andSomerville 1975; Clark 1977). The coordinates in!L Lt l :l J^) ..y_J L. ^U -r r^r _\ &, - - - + \ ^ r € d e f i n e da s :
1.
)a

t **

(

rr

-

-

's

(1)

ft:n11-2.'
where Il is the height of the model top and :" is theAugust 2000

463

T. Satomura

bottom topography height. The horizontai and thevertical domains used 1600 and 65 grid points, andreached approximately 1500 km horizontally and
26 km above the sea level, respectively.
The upper boundary is a rigid horizontal surface.
Above a height of 16 km, an absorbing layer is utiiized to avoid the reflection ofvertically-propagatingwaves at the upper boundary. At the lateral boundaries, a radiative boundary condition (Klemp andWilheimson 1978) is applied to the velocity component normal to the boundary. All variables otherthan the normal velocity component have zero gradient at lateral inflow boundaries and are set values
at outflow boundaries such that the second spatiai
derivative is zero.
Topography is a highly filtered form of the actual topography transecting 95"E 14oN and 105"E
17oN through the middle of Indo-China Peninsula(Fig. 1a). This is approximately pa,rallel to the tropospheric mean wind direction over the Indo-China
Peninsula in June (see below). Figure lb compares the smoothed topography that was used in
the model to the actual topography with 30// longitude resolution. About 290 km from the left (west)
and about 200 km from the right (east) boundaries, the lower boundary of the model domain is
the sea; the Andaman Sea and the South-China
Sea, respectively. Mountains around r = 400 km
are the Dawna Range, which separate Myanmarfrom Thailand, and r x t200 km are the Annam
Range which separate Laos from Vietnam. The region between r N 800 and 1100 km is the Khorat
Plateau, north-eastern Thailand. The peaks a;roundr = 700-800 km correspond to the Dong Phraya
mountains. Because the model domain is oriented
approximately west-to-east, the right-hand direction
and the left-hand direction in the model domain are
hereafter referred to as eastward and westward, respectiveiy.
The initial thermodynamic and wind conditions
were specified by an interpolation from climatic values on June at 102.5"E 17.5oN, about the centerof the model domain. A profiie over an upstream(western) point was not employed for the initialcondition because upstream profiles over the ocean
are not representative of the Indo'China Peninsula
which is the focus herein. Figure 2 shows the ciimatic vertical profile of thermodynamic variables
and wind. The wind velocity along the model domain is depicted as [/ and the wind perpendicular
to the model domain is depicted as V. Because V
is much smaller than [/ except in the lowest 1 km,
oniy [/ is considered. The sea surface temperature is
kept constant at 302 K according to the climatic sea
surface temperature of the Andaman Sea in June.
The climatic vertical profiles in June is practicallythe same as the profiles in July. The difference of[/ is less than 1 m s-1 for the lower troposphere

(N)
,a

q)
--1

-

16

P

74

, .']

1 2

10

95

100

110

105

(E)

Long i tude
Fig. 1a. Topography of the Indo.China Peninsula. Surface elevation is contoured at
100 m, 200 m, 500 m and 1000 m. Thickline A-B is the section used in our 2-D
model.

2000
1600
Dong

ho

1200

mountains
Range

ll

800

i

KhoratPlateau

4000SouthCh ina
e a

I000

500

0

A

15 0 0(km)

X

B

Fig. 1b. Model input topography (solid line)
and unfiltered topography with 30" longitude resolution (dashedline).

(> 400 hPa) and linearly increasesto 4.5 m s-1 at
100 hPa where [/ is about 25 m s-t; the temperature
difierence is less than about 0.5 K throughout mostof the troposphere; and the difference of the mlxing
ratio is less than 3 x 10-a kg kg-t. Therefore, theclimatic vertical profiles in June are a representative vertical profile of the early-to-mature monsoon
seasonin the Indo-China Peninsula.
The experimental conditions described is hereafter referred to as CNTL. In addition, the following
three experiments were performed (Table 1):1) remove mountains and make all the land flat/--^t^---^)\rerurlcu!^^LU db\Tnltr\Tm\r\\,/1v11\a/,2) remove seasand set all the model lower boundary to be land with the same surface elevationas CNTL (referred to as NOSEA);

Vol. 78, No. l

Journal of the Meteorological Society of Japan464

3. Results

r*'' I
l

i

l

r i l

- l:tt- iI

i
I

jl

\ril
500
600
700
800
340

360

-20 - ' 1 0 0

380
(K)

10
(m/s)

Vertical profiles of Initial conFig. 2.
potential temperatureLeft:ditions.
(solid), equivalent potential temperature(dashed) and saturated equivalent potential temperature (dotted). Right: windvelocities parallel (solid) and perpendicular (dashed) to the model domain.
Table 1. Experimental

Run
CNTL
NOMNT
NOSEA
DRY

Mountain
yes
no
yes
yes

conditions.
Sea
yes
yes
no
yes

Moisture
yes
yes
yes
no

3) same as CNTL except microphysical processes
are removed (referred to as DRY).
These experiments are obviously intended to clarify the effects of mountains, the sea surface and thecondensation by comparing with CNTL.
All simulations began at 00 LST of Thailand(17 UTC) and were integrated for 3 days. An initial
random disturbance was added to the temperature
at the lowest atmospheric iayer with an amplitude
of1K.
Experiments using topography made from twoother sections through the Indo-China Peninsula
were completed; along the 16'N line and the transecting line 95'E 15oN and 105'E 18oN. Experiments using initial thermodynamic and wind conditions on June at 95"E 15"N over the Bay of Bengal
also performed. However, results of these simulations is not shown because they are essentially thesame as the results of CNTL.

3.1 CNTL simulation
3. 1.1 Dr,urnal uariatiort
On the first day of the integration, convection
scattered over the land area during daytime. On thesecond and third days, rather organized convective
systems developed over the land in the afternoon
and some remained active until midnight. The evoIution of condensateis displayed in Fig. 3 for 6-hourintervais on the third day. At midnight, weak convective systems remained over the eastern half of theland area (Fig. 3a) and disappeared in the morning(Fig. 3b). In the daytime (Fig. 3c), convective activity was still weak over the land except over the
upstream (western) slope of the Dawna Range. In
the evening, a strong convective system formed overthe downstream (eastern) foot of the Dawna Range(Fig. 3d). As on the second day, this system moved
eastward during the night.
A horizontal-time section diagram of precipitation
rate averaged on the second and the third days isshown in Fig. 4. As noticed in Fig. 3, precipitation was rarely observed in the early morning except near the eastern boundary. The lower bounda.ry for r > 1300 km is sea and the early morning or
pre-dawn maxima of precipitation over this regionis similar to the observed diurnal cycle of precipitation over the ocean (e.g.,Sui et al. 1997). Precipitation in this region is, however, not discussedfurther
because the lateral boundary will affect its cha^racteristics. In late morning, precipitation was simuIated over the upstream side of the Dawna Range(r = 400 km) . This precipitation system moved
from the west to the mountains, but did not passover the Range (1000-m high). Therefore, the precipitation continued over the upstream slope of theRange until evening. In the downstream side of the
Dawna Range, a precipitation system began developing in the afternoon at r = 450-500 km. This system moved eastward with a speed of about 7 m s-r
until midnight and reached the western siope of
the Dong Phraya mountains around r ry 700 km.
Another significant precipitation system developed
also over the eastern (downstream) side of the Dong
Phraya in the evening and moved eastward till midnight at about 10 m s-1. Over the western siopeof the Dong Phraya, precipitation occurred in theafternoon but was weak.
On inland areas that are far eastward from mountainous regions, the simulated diurnal variation of
precipitation showed a peak at night or at midnight.
Precipitation systems which started at the easternfoot of the mountainous region in the late afternoon
^^^L
calu

l^-,
u4Jr.

^-J
@uu

-^-,^J
luvvLu^^-+-,,^-,1
ucourv@ru,,,i+1"

^

c^n^l
oyLlu

^f
vr

7
t

10 m s-1, causednighttime maxima of precipitationon those iniand areas.
It should be noticed that a clear spacial gap of

465

T. Satomura

August 2000

(lsn)

Day 3

0h

LT:

10

200

400

400

LT:

( " )

0
(I'xn)

200

12 0 01400( km)

10 0 0

12 0 0

14 0 0
(lsn)

1200

1400

Day 3

12 h

800

600

LT:

(d)

10

N

400

800

600

10 0 0

Day 3

6h

LT:

(b)

200

800

600

1000

Day 3

1t h

5

0
"

0

200

400

600

800

1000

( lott)xh, (b) t:6h, (c) t : 12 h and (d)Fig. 3. Predicted potential temperature and condensatesat (a) t:0
in
cxccssof 10-a kg kg-t
ofcondcnsatcs
raiio
The
niixin8i : i8 h on ihe tirirti tiay iii case CI{TL.
(cioud boundary) is shaded. Contour interval for potential temperattrre is 4 K. Time coordinate islocal standard time in Thailand (: UTC + 7 h).

Journal of the Meteorological Society of Japan466( h nrr r)(ko)

Ave P r e c R a t e

LT:Vol. 78, No. 4

19h Day 3

6

20
D

*{
ril

l o

ri[

il

F

,.|f

1 2

4

N 3

il
500

2

1000
1

0
510 520 530 540 550 580 570 580 590 600(lm)

Fig. 4. Horizontal-timesection diagram ofprecipitation rate averaged on the second and the third days in case CNTL.
Contours of precipitation are 0.5, 1 and
10 mm h-r. Shadedarea indicatesprecipitation rate is greaterthan 1 mm h-1.
Dash-dotted line indicates moving speedof 7 m s-1. The surfacetopography isdepicted by the black sha.dingat the bottom of the figure.

----)precipitation existed at just downstream of the peakof the Dawna Range (r x 40V450 km). Anotherprecipitation gap was also found just downstream ofthe peak of the Dong Phraya (n = 750 km) although
it was narrow.
3.1.2 Structures of the mouing prec'ipitation systems
A close-up of a precipitation area moving to theright (eastward) in Fig. 5 displays two precipitation
shafts at r * 560 km and r = 570 km. Ahead of the
precipitation shafts, a new cloud was triggered by
the cold pool beneath it. The cold pool was 2-3 K
colder than the environment and its depth was about
500 m. Significant front-to-rear flow was noticed in
the lowest 1 km in front of the cold pool and in the
trailing clouds of 4-6 km height, which extended forr < 550 km. Strong upward motion in convective
clouds occurred at the front of the precipitation system and connected the lower and upper front-to-rea.r
flow. A region of strong upward motion and condensate showed down-shear tilting. Under the trailingclouds rear-to-front flow was found from 0.6 to 2 km
in height. A part of the rear-to-front flow ascended
at r x 550 km and 565 km, and the other partdescendedinto the cold pool. These structures are
sllllllalr _ __-lr r:--__ /^ ..
lrxcs
squa,lL(-) Lruprear\u.8.,

lt^_^_:^glvruuLrrc[

x

20m,ts^_ldr]uMiller 1976; Zipser 1977; Chong et al. 1987; Hauser
et al. 1988; Tao and Simpson 1989).
Figure 6 portrays a time sequenceof vertical ve-

Fig. 5. Close.up of a precipitating a.rea
at t : 19 h on the third day in case
CNTL. Shading depicts regions of mixing ratio of condensates greater than
10-4 kg kg-r. Solid lines are potential
temperature d < 300 K with 1 K interval. Winds relative to a system moving
with 7 m s-l are represented by arrows.
A 20 m s-l reference vector is shown on
the lower left. Vertical wind velocities are
exaggerated by a factor of 5.
locity and condensate for 70 min period between
18 h 10 min and 19 h 20 min on the third day. In
Fig. 6a, four cells of vertical velocity can be seen.
Near the front of the precipitation system, a newcell formed above the leading edge of the cold pool
Iocated at r : 550 km. Further rearward were decaying cells above r : 545 and 535 km, and an
older cell at tr : 527 km above an accompanying
rain shaft.
Subsequent panels. each separated by 10 min,show that the new echo first intensified and thendecayed as it moved rearward relative to the cold

pool. By 19 h 0 min (Fig. 6f), another new cellformed at the leading edge of the cold pool. The
location of the new cell relative to the cold pool was
the same as for the new ceil of Fig. 6a; this behaviorof the simulated multi-cell system allows us to define the model system's propagation speed as thatof the cold pool. An updraft always existed above
the cold pool front, which supports the advective
cell regeneration in multicell storm proposed by Lin
et al. (1998).
A density current propagation speed equation derived by Seitter (1986) is467

T. Satomura

August 2000

(I(t[)(b)

LT:18h50m Day 3

L T : 1 t h 1O m

o
4
N

2
0

540

LT:19hLT= 18hZ Orn D ay

580

Om Day4
N,
0

L T : 1 t h 1O m

L T = 1t h 3 O m

o
4
N
I

LT:1 th2Om

LT:1 8h4Om

o

4
N

0

and 19 hFig. 6. Vertical velocity and condensates presented for 10 min intervals between 18 h 10 min
5isdisplayed.
in
Fig'
as
the
domainofsection
km
90
The
same
20 min in the third day in case CNTL.
positive
Vertical velocity is contoured with 1 m s-l interval. Solid and dashed contours indicate
greaterthan
and negative velocities, respectively. Shaded region encloses mixing ratio of condensates
pool.
cold
edge
of
front
position
theofthe
10-4 k; kg-1. Open rectangle at the x axis indicates

V:k

L,P
p

(2)

where V is the vel.ocitv of the front' Ap is the difference in surface pressurebetween the density current
head and the environment, p is the density of the environment. The constant k is approximateiy unity'
In case of the system shown in Fig. 5' Ap = 50 Pa
An ambient wind in theand p = 1.16 kg *-t.
evening in this region is almost calm. The density
current speed V estimated from parameters in this
precipitation system is - 6.5 h s-1, which is close to
the travel velocity of the system (- 7 m s-11. This
result indicates that simulated precipitation systems
were not simply advected by the low level wind but
rather propagated at the speed of the gravity current.
Since structures were similar to the squall lines
and moving speeds were close to the gravity currents, precipitation systems simulated in this study
will simply be called squall Iines hereafter.

we

3.1.3 Origi,n of the squall l'ines
The squall lines were initiated over the leeward
foot of the mountains in the afternoon and propagated eastward. This initiation and propagation repeated every day, resulting in the night maxima of
precipitation in the inland region. Hence, the triggering location and time of squall lines govern thediurnal cycle of precipitation over the Indo'China
Peninsula.
Figure 7 shows a close-up time sequenceof windveiocity, condensatesand potential temperature forthe 45 min period between 16 h 00 min and 16 h
45 min on the third day around the Dawna Range.
At 16 h (Fig.7a), a deep layer of cold air began totravel on the downstream side of the mountain' but
not beyond the lee slope far from the mountain peakin the daytime. The cold air over the windward sideof the mountain was maintained by the reiativeiycoid sea surface temperature of the Andaman Sea
and also by precipitation systems over the windwardside of the Dawna Range that screenedthe solar ra-

Vol. 78, No. 4

Journal of the Meteorological Society of Japan468( lstt)LT:16h0m Day 32g-,r'V.-*'-'?y'l

t

>
J

t

.

S

-

--'*---*--

Jzz/

1:;l!^.:---.l;..::--320 340 360 380 400 4ZO 440 460 480 500 520 540(lolr)(b)

LT:16h15m Dav 3---t---l--.-r'.----

1-1---1-+-_- - \\':-r'rA'
: i

-*4+
\

320 340 360 380 400 420 440 460 480 500 520 540
(lsn)
( " )

LT:16h30m Day\\-\*r'/-_v-+''--.%J\- ----**r'r'**---r'0

3 0 0 320 340 360 380 400 420 440 460 480 500 520 540
(l@)

( d )

LT:16h45m Day 3->*zZ-+\--r'l--

/'-

S n

0

t _ 300 320 340 360 380 400?Om/ s
wfac=10.0

-

--_-z

420 440 460 480 500 520 540(lcn)

X
u o ff -0.0

Fig. 7. Close-up around the Dawna Range presented for 15 min intervals between 16 h and 16 h 45 minon the third day in case CNTL. Shading depicts regions of potential temperature 0 < 302 K with
contour line of 1 K interval. The boundarv of mixing ratio of condensates in excess of 10-a kg kg-t
(cloud boundary) is drawn by a heavy thick line. Winds are represented by arrows. A 20 m s-r and2 m s-1 reference vectors are shown on the lower left.

(hour)Ave

Pr e cRate( hour )F

Ave Prec

Rate

20

20
l o

469

T. Satomura

August 2000

t #u{rulfluffi
prflirt1fl
? 1 2

,"d'u
,i/'odl,"
Jd'
10 0 0500

500

10 0 0( tsn)

X
Fig. 8. Same as Fig. 4 except for case
NOMNT.
diation and cooled the air through evaporation of
precipitation. Subsequentpanels in this figure, eachseparated by 15 min, show that the cold air penetrated further on the lee side (Fig. 7b) and finallytriggered deep convection at r : 485 km (Fig. 7c).
This started to propagate eastward (Fig. 7d).
Although precipitation continued over the windward side of the Dawna Range during the daytime,
ciouds advected from the windward side to the iee
side dissipated and considerable precipitation did
not occur on the lee slope between 420 km < r <470 km. Downsiope wind on the lee slope prevented
the air parcei from lifting and reaching condensation
Ievel; this resulted in the precipitation gap remarked
on in Fig. 4.
3.2 NOMNT s'imulation
Diurnal variation of precipitation in caseNOMNT
is shown in Fig. 8 as an averageofthe secondand thethird days. Clearly, instead of the night maximumof precipitation found in case CNTL (Fig. 4)' afternoon maximum appeared over the land (290 km <r < 1300 km). Significant long-lived and travelingprecipitation systems did not occur over the land although rather weak systems existed. The afternoon
maximum of precipitation corresponds to the afternoon minimum of atmospheric stability due to theincreasing ground surface temperature in the afternoon. This precipitation occurred mainly during thesecond day. On the third day' only scattered small
cumuli existed with little precipitation.
These differences from CNTL indicate that themountains located in upstream of the plain govern the night maximum of precipitation by triggering squall lines at approximately the same iocal time immediateiy downstream of the mountains.
An early morning maximum arose over the ocean
near the eastern boundary of the model domain(1400 km < r). Considering that this area is nearFig. 9. Same as Fig. 4 except for caseNOSEA.
the outflow boundary, this precipitation maximum
is probably spurious.
Total precipitation amount averaged over land on
the second and the third days was 11.4 mm, about
1.5 times smaller than that in CNTL (16.8 mm)"Averaging over the inland region only (400 km <r < 1200 km), this ratio was almost the same:
11.2mm in NOMNT and 17.0 mm in CNTL. Hence,
not only did the mountainous terrain determine thephase of diurnai variation of precipitation, but it
also enhanced total precipitation amount even overthe lee-side areas. This effect of mountains on the
total precipitation amount is interesting because it
is counter-intuitive; mountains often reduce precipitation over the lee side regions, for example, in foehn
or Chinook situations.
3.3 NOSEA simulation
Diurnal variation of precipitation in NOSEA isshown in Fig. 9 as an average of the second and thethird day. Precipitation on the upstream side of the
Dawna Range (300 km ( r < 400 km) simulated incase CNTL (Fig. a) did not occure. As in CNTL,
propagating squall lines over the downstream sideof the Dawna Ridge (450 km < r < 700 km) practically disappeared in NOSEA. Both precipitation
systems were probably not supplied with sufficientvapor to maintain themselves,becausethe water vapor fed from the land surface is less than that fromthe sea surface. On the other hand, precipitation
systems over the Khorat Plateau (800 km < r <1100 km) still remained, though the precipitation
pattern was largely modified. In other words, in
NOSEA, precipitation systems disappeared over thewestern half of the Indo-China Peninsula but stayedin the eastern half of the Peninsula. This resuit indicates that the former systems were suppiied with

470

Journal of the Meteorological Society of Japan\J/kg)

LT:

1 2h

D ay

( z -r 5 0 . t u()h o u r

3

)

Ave

2500

Vol. 78. No.'ll J ( z : 11 1 . 6 m )20*.-q.'i
r'::n

2000

:tl

I O

,'jb
1500

)r

ry,

'12

,

n
O

I

10 0 0

I
0A\

500
500
0

0

500

I 000

15 0 0(krn)

XFig. 12. Horizontal-time section diagram offiltered horizontal velocity u at - 100 m
above the ground averaged on the second
and the third days in case DRY. Contours are +1, +3, *5 and *9 m s-r.
Shaded area indicates ?, < -1 m s-1.
Dash-dotted iine indicates moving speedof 3 m s-1. The surface topography rs
depicted by the black shading at the bottom of the figure.

Fig. 10.
Horizontal variation of CAPE
at the noon of the third day for case
CNTL (solid) and case NOSEA (dashed).
CAPE is passed low-pass filter with
cut-ofi length of 100 km.
( J/kg)

L T:

12h

D ay

( z ' r 5 ot u )

3

50

40

as 100-400 J kg-t for r ( 600 km and CIN wasla.rger than several tens to a hundred J kg-l.
Over
the middle of the land (600 km ( r < 1000 km),
CAPE and CIN in case NOSEA were almost thesame as those in case CNTL, which correspondedwith active convection over that area.

an

z
20

10

0

500

10 0 0

15 0 0(1o.)

Fig. 11. Sameas Fig. 10 except for CIN.
water vapor from the Andaman Sea, while the lattersystems used water lapor fed from the land surface.
Differencesof horizontal variations of CAPE (convective available potential energy) and CIN (convective inhibition) between NOSEA and CNTL wereconsistent with the difference of precipitation systems between them. In case CNTL, CAPE wasf^.---

1000

+h--

lnnn

T L--1

^.,--

+l -

t--,1

ovnanri---^^.

of high eievation (Fig. 10), and CIN was less than
20 J kg-l over the whole land (Fig. 11) at noon of
the third day. In case NOSEA, CAPE was as small

3.4 DRY s'irnulation
Figure 12 depicts horizontal velocity at the fifth
vertical level of the model (about 100 m above theground) as an average of the second and the third
day in case DRY. To examine mesoscale structuresin the horizontai velocity, a low-pass filter with cutoff wavelength 20 km was applied along each modellevel. There are two distinct features: systems moving eastward slowly during the afternoon and thenight, and stationary strong westerly flow over thelee slope of the mountains.
The stationary strong westerly wind was noticed
over the peaks and leeward slopes of the Dawna
Range (400 km < r < 500 km). Over the Dong
Phraya mountains (700 km < n < 800 km) andthe Annam Range (r = 1200 km), the stationary
strong-wind regions also occurred but were not significant. The stationary strong wind occurred overfhomnrrntqin

1-6^-,,o6+i.-L^-j-^-#^1,,,i-,1:-^-^.--^1!rua u@Lowith height up to - 1.5 km at the initial time
(Fig. 2) and also because the mountain enhances
the horizontai wind on the leeward slope (see, e.g.,August 2000

T. Satomura( lcn)

LT=

5h

Day

1

N 2
1
320 340 360 380 400 4e0 440 460 480 500 520 540
(h)
(b )

D ay 3

LT : 15h'*t"

Fig. 13. Filtered vertical velocity aroundthe Dawna Range at (a) 5 h on the firstday and (b) 15 h on the third day incaseCNTL. Negativevaluesare shaded.
Counter intervalsare (a) 0.05 m s-t and( b ) 0 . 1m s - t .
Smith 1977). The strong wind region was most significant at night but was still noticeable during daytime.
Negative (westwa,rd) horizontal wind areas
emerged in the morning on the leeward foot of the
Dawna Range and the Annam Range. A weak windarea also occurred on the leeward foot of the Dong
Phraya mountains in the morning. These negative
or weak wind areas sta.rted to move in the late morning and were traced till midnight. The travei speedof the u : 0 m s-r line was about 3 m s-r eastward, much smaller than the propagation speed of
the squall Iines in caseCNTL (7-10 m s-r).
4. Discussion
1.1 Location and timing of squall line generation
Examining Fig. 7a, it iq noticed that upward motion occurred al r x 470 km before the cold air
from the west reached this position. When the leading edge of the cold air arrived, a convective cloud
emerged (Figs. 7c and 7d). To examine what determines the initiation location of convective ciouds, alow pass filter with 30 km cut-off waveiength along
the terrain-following model coordinate (z* : const.)
was applied to the vertical velocity (Fig. 13). At
5 h on the first day when the flow was approximately adiabatic (i.e., no solar heating and no detectable cloud), mountain waves ciearly formed overthe mountain (Fig. 13a), distinguished by a phaseof positive vertical velocity approached the groundaround r : 460 km near the convection initia+::^uruu^^:-+
l/vuru

-

-

Aan

'l,n

A

"-.l

l

h'.-^

..^"-,]

r x 480 km caused the downward extension of thepositive vertical veiocity. Hence, if a mass of cold
air came down the lee slope (katabatic flow), theupward velocity forced by the leading edge of thecold pool was reinforced by the upward phase of themountain wave around r :460 km.
Even during the third day afternoon when thethermal and wind structures of the atmosphere weremodified by diabatic heating, an area of upwardvelocity over the lee slope was still evident from420 km < r <440 km at z : 3 km to 460 km <r < 480 km near the surface. Therefore, we consider that the upward velocity at the front edge ofthe cold air that penetrated down the lee side of themountain was reinforced by this area of upward velocity and triggered the squall iine each day. This
upward velocity was possibly the upward phase of
the mountain wave, though greatly modified.
Another interpretation of the reinforcement of the
upward motion is as follows: When the cold air ran
down the lee slope, the upward velocity forced at its
leading edge was not signifi.cant due to the descentof the system. When the front edge of the katabatic flow reached flatter terrain (e.g., the foot of
the mountain or the small bump at r x 480 km),the forced ascent triggered the squall line.
The moving squall lines in case CNTL is similar
to, but not the same as the moving zone of ascentover the lee side of the mountain discussedby Fujibe
et ai. (1999). I-Isinga dry numerical modei including
the solar radiation, they found that downslope wind
and valley wind converged over the lee slope near themountain top at noon. The zone of upward velocity associated with this convergence slowly moved
down on the lee slope during the afternoon. The
moving zone of the weak horizontal velocity in case
DRY seems to correspond to the moving convergence zone of Fujibe et al. (1999). O" the other
hand, in case CNTL herein, the cold air embeddedin the downslope wind piayed a major role in producing ascent. In addition, the propagation speeds
of squall lines were two-to.three times faster than
the speed of the moving zone of weak horizontal velocity in case DRY.
l. 2 Obseraational euidence
An activity index of deep convective clouds 16' isdef,ned by

I":230 - Tss : Tpn < 230K
-0
: T a B> 2 3 0 K ,(3)where ?ss is the infra.red equivalent biack bodytemperature derived from GMS. Hourly TsB datafor one month (June 1996) at 0.05o x 0.05' grid
points are utilized for the anaiysis. Although 16:
does not correspond weil with rainfall amount for
scales much smaller than several hundreds of kiloureiers (Ricirard anti Harkel i98i), iL ueverihelcss
indicates the activity of deep clouds such as squalllines.

Journal of the Meteorological Society of Japan

(h)

20
c)
tr 16
F
L2
rfl

O
J

B
+

110

i00
'

l

l

.

l

1

I Ongt r luoe

Fig. 14. Horizontal-time section diagramof amplitudes of diurnal componentsof
Ic at 15"N on June 1996 over the
Indo-China Peninsula.Only the positive
ralues are contoured with an interval of
1K.
Figure 14 shows a horizontal-time section of amplitudes of diurnal components of /6: along a latitude of 15oN over the Indo-China Peninsula. At
15oN, the Dawna Range is located around 99oE andthe Dong Phraya mountains around 101.5"E. It isclearly shown that 16: has a trend towards later maximum as increasing longitudes until 104-105'E. This
supports the simulation result that the travelingsquall line causes the diurnal variability of precipitation" It is also noticeable that the amplitudes ofdiurnal variability have locai minima on the lee sideof the Dawna Range and the Dong Phraya mountains, which harmonizes with the precipitation gap
in the simulation (r = 400-450 km and r t 750 km,
seeFig. 4).
A trend towards later maximum at decreasing
longitudes over the eastern Indo-China Peninsula(105"-108'E) is observedin Fig. 14. The cause ofthis westward propagation of cloud activity is notclear, although 16:roughly estimated from simulated
clouds also shows the similar westward propagation(not shown).

Vol. 78, No. 4

1.3 Orogenrc squall lines in midlatr,tudes
Triggering and propagation of squall lines on the
Iee siope of a mountain in the late afternoon are
not unique to the Indo-China Peninsuia. The samemechanism can work in some orogenic squail lines
in midlatitudes. Tripoli and Cotton (1989) reported
that the lee-side subsidence suppressed convectionin their 2D numerical model over the Rocky Mountains of the U. S.. Further, they suggested a mechanism similar to this study in that deep convection
was triggered when an air mass from the windwardside of the mountains reached the ascending branchof mountain waves over the lee-side slope. Tucker
and Crook (1999) aiso pointed out the importanceof cold air running down the lee slope for the generation of mesoscaleconvective systeml the coid air
in their study originated as outflow from previous
convective activity over the lee slope of the Rocky
Mountains.
Statistical analyses of diurnal variability of precipitation over the Great Plains in summer (Wallace
1975; Riley et al. 1987) show clear trends towards alater maximum eastward from the mountains. These
facts indicate that the triggering and propagationof squall lines on the lee-side of mountains are sufficiently common that statistical analyses of longterm precipitation data can detect the eastward
phase lag of the diurnal precipitation maximum.
Though Tripoli and Cotton (1989) and Tucker andCrook (1999) are specific case studies, in contrast
to this paper where climate data were used as theinitial conditions, both the numerical simuiations
and statistical analyses of rain gauge networks suggest that the same mechanisms operate in the diurnal variability of summer precipitation eastward ofmountain ranges in midlatitudes as well as in thetropics.
5. Summary
Diurnal va"riation of precipitation over the
Indo-China Peninsula was simulated by a twe
dimensional, nonhydrostatic, and cloud-resolving
numericai model. Conclusions are as follows:
1) Convection triggered at the lee-side foot ofmountains in the late afternoon forms squalllines which propagate eastward and producenight maxima of precipitation in inland areas
far eastward from the mountain.
2) The timing and location of convection initiation
are determined by the intrusion of cold air fromthe windward side into the lee side and also by
the mountain wave.
3) The Andaman Sea feeds water vapor thatcauses precipitation over the lee side of the
Dawna Range. Water vapor that maintains pre-T. Satomura

August 2000

cipitation over the eastern half of the Peninsulacomes mainly from the land surface.
4) Orogenic squall lines (CNTL) brought 1.5 times
more precipitation than cioud systems over flat
Iand (NOMNT).
5) The simulated night maxima of diurnal variation of precipitation and eastwa"rdpropagationof convectively active systems was confirmed by
the high resolution analysis of 16; .
The mechanism behind the daily triggering of convection probably works in other summer precipitation systems in the lee side of mountains. In thisstudy, climatic thermal and wind conditions wereused as the initiai data for the model. The proposed mechanism required only a general elementin addition to the environment preferable for squalllines: the cold air mass advected from the windward side of the mountain reaches a flat terrain in
the lee side where the up-branch of mountain wave
approaches the surface. Considering that neither
the initial condition of the model nor the simulated
precipitation diurnal cycle in this study are associated to one special event in the Indo-China Peninsula, the daily variation of precipitation over several
tropical regions are most probably explained by themechanism proposed herein. For example, a similarphase shift of diurnal 15: variation is detectable overBorneo as well as over the Indo-China Peninsula in
the analysis of Nitta and Sekine (1994) using low
resolution (1' x 1') data of ?66 (Fig. 4a of theirpaper). Considering that there is a high mountain
range (Iran Mountains) elongated from southwest tonortheast in Borneo, the phase shift over Borneo isarguably caused by the triggering and propagationof squall lines over the lee side of the Iran Mountains.
In contrast to the analysis of 16:, whose maximum
time was about 2 hours later than the maximum
time of precipitation over the Malay Peninsula (Okiand Mushiake 1994), the analysis of radar data directly reveals not only the maximum time but aiso
the structures of the precipitation systems. Single
radar has a limited observation range about 100200 km. Fortunately, in Thailand, a Doppler radar
network covers the whole nation. Analysis of thoseradar data during monsoon seasonsnow proceeds incooperation with the Thai Meteoroiogical Department.
Acknowledgments
The author expresses his special thanks to Dr.
S. Emori for providing subroutines for RAMS including the improved radiation parameterization.
He aiso acknowledges the contributions of T. Oki
and A. Watanabe for valuable discussions, lvI.W.

4tJ

lvloncrieff for a careful reading of the manuscriptand heipful comments.This work was supported by
the Center for Climate System Research, Universityof Tokyo and a Grant-in-Aid for Scientific Research
B-2-I1440L41 of the Ministry of Education, Science
and Culture, Japan. This research was completed asa part of GAME-Tropics project, which is a part of
The nuGEWEX Continental Scale Experiments.
merical experiment was done on the KDK system at
Radio Atmospheric Science Center and on the Data
Processing Center at Kyoto University. Figures were
Library.
drawn by the GFD-DENNOU
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August 2000

T. Satomura

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