A change of mean physical quantity in the monsoon and pre-monsoon.
A. Watanabe1, Y. Tachibana2, S. Ogino３, Y. Shibagaki７, T. Ohsawa３,Y., M. Yamanaka3, T. Oki６, T. Satomura4,A. Sumi５, K. Mushiake１(1Fukushima University, 2Tokai University, 3Kobe University 4Kyoto University, ５CCSR University ofTokyo, ６University of Tokyo, ７Osaka Electro-Communication University, )Fukushima University, Japan,960-1296, Kanayagawa, Fukushima, Japan. e-mail:may@sss.fukushima-u.ac.jp

Abstract
It is very important to understand a difference of the atmospheric circulation between the pre-monsoon and the monsoon periods.
We are able to divide into 3 classifications on Apr. 15 to May 13, May 14 to Jun. 15, and Aug. 15 to Sep. 15, to understand thefluctuation of the monsoon. And we calculated the mean field of the meteorological element every 3 classification. The pressure
decreases about 2hPa under the 5km in the monsoon than the pre-monsoon. But, the pressure increases about 3hPa over the 5km in
the monsoon than the pre-monsoon. Furthermore, the temperature decreases about 2℃ under the 3km, and the temperature
increases about 4℃ over the 3km in monsoon than the pre-monsoon. Accordingly, we are unable to explain the pressure decreaseof the lower layer with the temperature change. We are conceivable that this mean field was made with a convective activity. It
became clear that decreasing of pressure of the lower layer strengthens the south wind, and convective instability is strengthenedby advection of water vapor by the southerly wind.

Keyword: monsoon onset, rawinsonde observation, convective activity, Thailand

1.Introduction
We carried out rawinsonde observation tounderstand the atmospheric circulation in thepre-monsoon and the monsoon periods over Thailand.
We observed the upper atmospheric structure by
rawinsonde four times a day from Apr. 15 to Jun.15
and Aug.15 to Sep.15 1998 at Chiang Mai, Ubon
Ratchathani and Bangkok in Thailand. It is very
important to understand a difference atmosphericcirculation between the pre-monsoon and the monsoon
periods. The reason is that we can understand themonsoon onset mechanism due to the difference of theatmospheric circulation. In particular, we clarify thedifference of the mean field of meteorological element
to compare the pre-monsoon period with the monsoon
period.
Many investigators have reported the trigger ofmonsoon onset. Orgill (1967) and Eguchi (1996) haveindicated that many of the monsoons have started by
the northing of onset vortex. And, Wu and Zhang(1998) pointed out that the trigger of monsoon onset isoverlapping of phase of Madden-Julian oscillation andthree-week oscillation. Further more, Xu and Chan
(2001) have indicated the trigger of monsoon in 1998
is the northing of cyclone. Although the trigger of themonsoon is pointed out various as above there is little
research of observation. We calculated the mean value
with Bellamy method by using the rawinsonde
observation data that we observed for every 6 hours at
three stations in Thailand.
Our purpose is to indicate the difference of eachmeteorological element between the pre-monsoon andthe monsoon periods, and to understand the trigger ofmonsoon onset in the Indochina Peninsula.

2. Analytical method
We obtained the meteorological data of upper
atmosphere by rawinsonde four times a day from Apr.
15 to Jul. 15, and from Aug. 15 to Sep 15 in 1998 at
Chiang Mai, Ubon Ratchathani and Bangkok inThailand. We have calculated the average ofrawinsonde data every 100m from surface to 15km
above the sea level. And we obtain time series
meteorological data every 100m during each period.
Each missing data are calculated by Kriging method.
As we calculate the average every 6 hours at eachlevel, we can understand the situation of the mean

field during each period. On the other hand, it ispossible that we calculate the physical quantity using
average data of the 3 stations.
Here, u and v are eastward and northward
component of wind at the centroid of three stations.
We can express the wind component of eastward andnorthward by following equations.

ui = axi + byi + c
Vi = dxi + eyi + f(1)

Here, indicate each station. We can calculate theconstant (a, b, c, d, e, f) of equation (1) by substituting
observation data of three stations. Then, we can obtain
the wind component (u, v) at the centroid bysubstituting coordinate x, y of the centroid.

u = ax + by + c
V = dx + ey + f(2)

And it is possible to calculate the physical quantity bydifferentiating equations (2) in x or y. Similarly, it ispossible that we calculate the zenith component of thedivergence (D) and vorticity ( ς ) as followingequations.
∂u∂v
D =+(3)
∂y∂x

ς =∂v
∂u−
∂x
∂y(4)

In addition, it is possible to calculate the verticalvelocity (w) using the continuity equation as followingequations.

w = −∫ (∂u ∂v+ )dz∂x ∂y(5)Here, we calculated the vertical velocity as w=0 with
z=0, and vertical velocity is integrated every 100m.

3. Analytical resultWe have calculated the each physical quantity, andwe consider the difference of mean field ofmeteorological elements between the pre-monsoon
and monsoon periods.

Figure 1 shows the time-lag correlation ofmeteorological elements at Chiang Mai that wascalculated by Watanabe (2002). We can find that thelag-correlation of the meteorological elements shows
the negative value from May 14. Accordingly, We areable to divide into 3 classifications on Apr. 15 to May
13 (pre-monsoon), May 14 to Jun. 15 (monsoon), andAug. 15 to Sep. 15 (mature-monsoon). We averageeach meteorological element every 3 classification. Weconsider the trigger of the monsoon onset from thecharacteristic of the mean value.
E q u iv a le n t P ote n tial T e m p e ra tu re

M ixin g R atio

P re ssu re

T e m p e ra tu re

E astw ard C om p on e n t of W in d

N orth w ard C om p on e n t of W in d

1

in the monsoon and mature-monsoon is lower than thepre-monsoon. And, the temperature in the higher layer
than 3km in the monsoon and mature-monsoon ishigher than the pre-monsoon. The temperature doesn’trise when it enters into the monsoon. The solar
radiation is interrupted by the cloud in the monsoon.
Figure 3 shows the vertical profile of pressure
difference of Chiang Mai and Ubon Ratchathani fromApr. 15 to Jun. 15, 1998. We can find that the pressure

0 .8
0 .6

Correlation

0 .4
0 .2
0- 0 .2- 0 .4- 0 .6- 0 .8-1
16

Fig.3 The vertical profile of pressure difference (hPa) ofChiang Mai and Ubon Ratchathani from Apr. 15 to Jun. 15,1998.

11 16 21 26 31 36 41 46 51 56 61
T im e (d ay fro m A p ril 1 5 , 1 9 9 8 )Fig.1 Time-lag correlation of Meteorological elements at
Chiang Mai. The coefficient of lag correlation is calculatedwith the vertical distribution of the deviation from eachheight mean value.

Height（ｋｍ）

Figure 2 shows the vertical profile of temperature at
the centroid. The temperature in three periods is in
151413
12
1110
98765
pre-monsoon
4monsoon3
mature-monsoon
210-80 -70 -60 -50 -40 -30 -20 -10
Temperature(℃）

010203040

Fig.2 Vertical profiles of temperature at the centroid in theeach period.
accord with about 10℃ at 3km above the sea level. It
is the highest temperature in the lower layer than 3km
at the pre-monsoon in the three periods. And, it is thelowest temperature in the higher layer than 3km at thepre-monsoon in the three periods. Namely, the cloud
does not develop in the pre-monsoon; therefore thetemperature is high relatively near the surface.
However, the temperature in the lower layer than 3km

difference of the two stations don’t change largely inthe lower layer than 2km. However, the pressure
difference becomes conspicuously small in the upperlayer than 10km after the monsoon onset. It is dependon the pressure rise at Chiang Mai after monsoon
onset. Xu and Chan (2001) indicated that the reversionof the temperature gradient has not occurred in the
Indian Ocean and Tibetan highland of the upper part
troposphere. But, it is clearly that the pressure
difference of the north and south decrease after
monsoon onset. This is corresponding well with therise of the temperature of the upper layer. It is clearthat the decrease of the air pressure gradient in theupper layer is the trigger of a monsoon onset.
The vertical profile of the eastward component ofwind is shown in Figure 4, and Figure 5 shows thevertical profile of northward component of wind. The
southwesterly wind is excelling with the lower layer
than 5.5km, and the northwesterly wind is excelling
between 5.5km and 13km in the pre-monsoon.
However, the southwesterly wind is excelling fromsurface to 10km above the sea level in the monsoon.
The wind of the lower layer does not change largely
with the monsoon and pre-monsoon. The wind ofupper layer over 5.5 km is conspicuously different.
The decrease of the pressure difference of the south
and the north in the upper layer are gradually
weakening the west component of wind in the upper
layer clearly. The northeasterly wind excels in thelayer over 7 km in the mature-monsoon and thepressure difference of the south and the north mustreverse. The vertical distribution of the horizontal
advection of water vapor in each period is shown in
Figure 6. We consider the horizontal water vaporpre-monsoon

Height (km)1514131211109876543210-25

mature-monsoon

monsoon-20-15-10-50
5Eastwardof Component of Wind(m/s)

1015

20Fig.4 Vertical profiles of eastward component of wind ineach period.
pre-monsoon

mature-monsoon

monsoon1514
13
121110
9876
54
3
2
10

find that the horizontal water vapor advection in thelayer of lower than 1km is obvious in the all period. Itis corresponding with the strength of the southerly
wind of the layer. However, the horizontal water vapor
advection between 1 km and 3 km shows the value ofnegative and the correspondence with the direction of
the wind is not necessarily evident. We find that thehorizontal advection of water vapor is very small,
although the value of negative is shown in the upper
layer than 3 km. It is the positive horizontal advectionof water vapor in the southwest wind range of thelower layer than 5 km, and it is negative horizontal
advection of water vapor in the northeast wind rangeof the upper layer than 5 km in the mature monsoon.
As for the distribution of the vapor of the lower layer
the monsoon period is largest and the surface
temperature is low.
Accordingly it can dosolidification easily with a little convergence in thelower layer than 1km. We are conceivable that thedried air advection between 1km and 3km is doing
the work that continues solidification.
The distribution of such wind influences vorticity
production largely. Figure 7 shows the verticalprofile of vorticity. The vorticity with the entire layer
is negative in the pre-monsoon. Also, the vorticity is-5-4-3-2-1
01
NorthwardComponent of Wind (m/s)

234Fig.5 Vertical profiles of northward component of wind ineach period.

advection difference by the difference of the directionof the wind.
Figure 6 shows the vertical profile of horizontal
advection of water vapor mixing ratio in 12 hours. We
pre-monsoon
monsoonmature-monsoon

Height (km)

15
14
13
12
11
10
9
8
7
65
4
3
2
10-0.5

00.5

11.5

2

AdvectionofMixingRatio(g/kg12hour)Fig.6 Vertical profiles of horizontal advection of watervapor mixing ratio in 12 hours

15
14
13
12
11
10
98
7
65
43
2
10-3monsoon

mature-monsoon

Height (km)Height (km)pre-monsoon-2.5-2-1.5-1-0.50Vorticity(X10-5sec-1)

0.5

1

1.5

Fig.7 Vertical profiles of mean vorticity in each period.
positive in the lower layer than 4km and the vorticity
is negative in the upper layer than 4km in the monsoon.
Furthermore, it is the positive vorticity with all thelayers in the mature-monsoon. The northwesterly windof the pre-monsoon is strong in the north of theIndochina Peninsula, be weak the south. The
southwest wind of the mature-monsoon is strong in thesouth of the Indochina Peninsula, be weak with north.
The distribution of such wind is deciding the vorticity.
Figure 8 shows the vertical profiles of horizontal
temperature advection. The cold air advection isobvious with a lower layer than 1 km in thepre-monsoon. However, the temperature advection in
the monsoon is showing the value of positive with allof lower layers than 12 km. Also, the horizontal
temperature advection of the mature-monsoon
indicates the value of negative with the upper layer
than 5km, it of positive with a lower layer than 5 km.
The negative temperature advection region correspondwith the region of the southward component of wind,and positive temperature advection region correspondwith the region of the northward component of windexcept for the region of the negative temperature.
pre-monsoon

monsoonmature-monsoon

15141312
11

Height (km)

10
98
7
6
5
4
3
2
10-4-3.5-3-2.5-2-1.5-1-0.5

00.5

11.5

2

Advection Temperature (K/12hour)Fig.8Vertical profiles of mean horizontal temperatureadvection in each period.
Figure 9 shows the profiles of mean vertical speed ineach period. A relatively strong convergence region isappearing with the lower layer than 1km and thisregion become an ascent flow in the pre-monsoon andmature-monsoon. The descent flow appears between
1km and 3 km and weak ascent flow appears with aupper layer than the 3 km in the pre-monsoon.
pre-monsoon

monsoonmature161412Height (km)10864

4. Summary and discussion
Our purpose is to indicate the difference of eachmeteorological element between the pre-monsoon andthe monsoon periods, and to understand the trigger ofmonsoon onset in the Indochina Peninsula.
We calculated the mean value from the data ofrawinsonde observation in each period. We haveclarified the characteristics of the mean field of themeteorological element in each period by using themean value. And, we tried to clarify the cause ofmonsoon onset from the difference of the mean field.
As a result the following became clear. The excellenceof the northwest wind that makes a detour theHimalayas mountainous region generates the negative
vorticity and be causing to the dry air advection in thepre-monsoon. However, the water vapor advection islarge with the lower layer than 1 km, the southwest
wind is excelling in the layer. Accordingly, theatmosphere instability is largest in the pre-monsoon,
the convective activity is expected to become active inthe period. But, the convective activity has notoccurred in the period. It is reason that the drying air issubsiding with the layer between 1 km and 3 km.
The stable layer is formed in the upper layer. The
trigger of monsoon onset is the disappearance of thisstable layer. A stable layer sometimes disappears
with the irruption of cyclone, as Orgill (1967) andEguchi (1996) are pointing out it. However, it is abasis that the intrusion of the drying air that makes adetour the Himalayas mountainous region isdisappeared over the Indochina Peninsula. The
reversion of the pressure gradient is necessary in thesouth and the north of the Indochina Peninsula by thetemperature rising in the northern region. The
temperature rise in the northern region, the irruption of
the southwest wind strengthens and the irruption of
humid air is become active over the region. This is thebeginning of monsoon when was analyzed from themean field. The north and south temperature gradientincreases 2 ℃
in the monsoon period andpre-monsoon period in the upper layer than 2km,
although it decreases with the lower layer than 2 km.
The pressure gradient of the north and south decreases
with the lower layer than 2 km, it is increases 1.5 hPa
in the upper layer than 2km.

20-0.15

References
-0.10

-0.05

0.00

0.05

0.10

0.15

Vertical velocity (cm/sec)

Fig.9Vertical profiles of mean vertical speed in eachperiod.

However, the convergence is weak with the lowerlayer than 1km, and it is divergence in the upper layer
than 1km in the monsoon. Accordingly, although only
the ascent flow appears in the lower layer as for theothers, the descent flow is appearing in the monsoon.
The mean quantity of convergence is not corresponding
directly with the convection activity over the Indochina
Peninsula. However, it is conceivable that the condition of
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