Journal of the Meteorological Society of Japan, Vol. 82, No. 6, pp. 1545--1564, 2004

1545

Submonthly Variability of Convection and Circulation over and aroundthe Tibetan Plateau during the Boreal Summer

Hatsuki FUJINAMI
Core Research for Evolutional Science and Technology, Japan Science and Technology Agency, Kawaguchi,
Japan

and
Tetsuzo YASUNARI
Hydrospheric Atmospheric Research Center, Nagoya University, Nagoya, Japan(Manuscript received 29 September 2003, in final form 31 August 2004)

Abstract
Convective variability at submonthly timescales (7–20 days) over the Tibetan Plateau and the associated large-scale atmospheric circulation and convection were examined over regions affected by the
Asian Monsoon. The mature phase of the Asian summer monsoon (July–August) was analyzed for those
years (1986, 93, 98) in which convective variability on timescales of 14 days was notable over the Tibetan
Plateau.
Composite analyses of OLR, based on the filtered Tbb time series over the southern Tibetan Plateau,show that significant convective signals rotate clockwise around 28 N, 90 E, affecting the Tibetan Plateau, Indochina, the Bay of Bengal, and India. Significant signals also appear around the Philippines
and the South China Sea. A well-developed wave train extending from North Africa to far-east Asia
along the Asian subtropical jet is associated with convective fluctuations over the plateau. The waves are
quasi-stationary and have a Rossby wave-like downward wave train with wavenumber 7.
The waves control convective fluctuations over the plateau. During the transition to active (inactive)
convection, an upper-level trough (ridge) develops west of the plateau. Simultaneously, cyclonic anomalies strengthen over India between the lower and middle troposphere. The development of the twotroughs induces a southerly flow of moist air toward the plateau. Moistening of the lower atmospherecreates favorable conditions for subsequent active moist convection.
Possible processes for forming the wave train over the subtropical jet and a link for convective signals
between midlatitudes and the Asian monsoon are discussed.

1.

Introduction

Intraseasonal oscillations (ISOs) of convective activity and the associated atmospheric
Corresponding author: Hatsuki Fujinami, Hydrospheric Atmospheric Research Center, Nagoya,464-8601, Japan.
E-mail: hatsuki@hyarc.nagoya-u.ac.jp( 2004, Meteorological Society of Japan

circulations are prominent phenomena associated with the Asian summer monsoon. Many
studies have investigated 30–60-day oscillations, the so-called Madden-Julian Oscillation(MJO) (Madden and Julian 1972), and the relationships between the MJO and the active/
break periods of the Asian summer monsoon(e.g., Yasunari 1979, 1980, 1981; Murakami
1984; Annamalai and Slingo 2001). ISOs at

1546

Journal of the Meteorological Society of Japansubmonthly (6–25 days) timescales (e.g., Vincent et al. 1998) also occur over the region affected by the Asian summer monsoon (e.g.,
Krishnamurti and Bhalme 1976; Murakami
1976; Yasunari 1979; Krishnamurti and Ardanuy 1980; Chen and Chen 1993; Fukutomi andYasunari 1999, 2002; Annamalai and Slingo2001). These studies suggest that the effects ofthe submonthly oscillations on monsoon activity can be as important as the effects of 30–
60-day oscillations.
Common features of the 10–20-day modeover the Asian summer monsoon region havebeen reported. Murakami (1976) used Indian
Daily Weather Reports and upper-air observations from the summer of 1962 to show that aquasi-biweekly oscillation influences the active/
break cycle of the monsoon over India. Krishnamurti and Bhalme (1976) analyzed nine
monsoon elements and revealed a biweekly
signal. Later, Krishnamurti and Ardanuy(1980) showed that monsoon breaks over India
coincide with the arrival of a ridge in the 10–
20-day westward propagating mode. Chen andChen (1993) investigated the horizontal structure of 10–20-day ISOs in the Asian summer
monsoon for the boreal summer of 1979. They
found that the 10–20-day monsoon mode hasa double-cell structure: one cell is centered between 15 and 20 N, and the second cell is overthe equator. The two linked cells propagate
westward along the Indian monsoon troughand along the equator, respectively. Fukutomi
and Yasunari (1999) found prominent 10–25day oscillations over the South China Sea
during the early summer. Enhanced (suppressed) convective anomalies are associatedwith cyclonic (anticyclonic) circulation anomalies in the lower troposphere. These circulation
anomalies propagate westward from east of thePhilippines to the Arabian Sea.
Analyses of precipitation data yield conflicting results. Annamalai and Slingo (2001) used
the ECMWF 15-year reanalysis and OLR datafor 1979 to 1995 to produce a generalized viewof the submonthly variability in summer (Junethrough September), based on daily All-India
monsoon rainfall (AIR). The power spectrum of
the 17-year average daily AIR series has statistically significant peaks at both 12 and 16
days. The 10–20-day mode originates over theequatorial western Pacific, in association with

Vol. 82, No. 6

westward-propagating Rossby waves. By contrast, Hartmann and Michelsen (1989) analyzed daily precipitation from Indian stations
between 1901–70, and a 10–20-day mode is notstatistically significant in the 200-day time series beginning on 10 April except over parts ofnorthern India. They did show statistically significant peaks at 5–7-day and 40–50-day periods in the precipitation records. The data andanalysis periods in the two studies differ, so a
simple comparison is impossible. However, theinconsistency suggests that activity with submonthly timescales over India has variability
at large interannual or longer timescales. Annamalai and Slingo (2001) show that the 30–
60-day mode is strong during monsoon onset(early June) over India, whereas the 10–20-day
mode is more pronounced after the monsoon isestablished (July–August). Careful attention is
warranted for interpreting power spectra forlong-term time series. Our study regards the
10–20-day variation as a robust feature of the
Asian summer monsoon.
The southern part of the Tibetan Plateau isa major part of the Asian summer monsoon
domain (Murakami and Matsumoto 1994). Remarkable diurnal variation in convection occurs over the plateau during the summer (e.g.,
Murakami 1983; Nitta and Sekine 1994; Yanai
and Li 1994; Fujinami and Yasunari 2001).
The convection depends strongly on thermallydriven regional-scale circulations induced byplateau topography (Ueno 1998; Kuwagata etal. 2001; Kurosaki and Kimura 2002). Nevertheless, few authors have investigated convective variability on submonthly timescales over
the Tibetan Plateau. Nitta (1983) described
intraseasonal variation with periods of 10–15
and @30 days in total heating, precipitation,and relative vorticity in a case study for thesummer of 1979 for the upper troposphere
over the eastern Tibetan Plateau. Endo et al.
(1994) used upper-air soundings from 1993 andshowed that quasi-biweekly fluctuations in meteorological elements dominate over the central
plateau after the onset of the summer monsoon.
Fujinami and Yasunari (2001, hereafter FY01;
2002) showed fluctuations in the diurnal amplitude of summertime convection with periodsof about 14 days and about 30 days over theplateau. ISOs greatly influence diurnal variability over the plateau.

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H. FUJINAMI and T. YASUNARI

ISOs also occur in the extratropics (e.g.,
Krishnamurti and Gadgil 1985; Hsu and Lin
1992; Kiladis and Weickmann 1992; Ambrizzi
et al. 1995; Terao 1998, 1999). Quasi-stationary
Rossby waves occur in the Asian subtropicaljet during northern summer, with well-confinedzonal wavenumbers 5–7 (Terao 1998, 1999).
Spectral peaks have periods of 14 days and 30–
45-days. Propagation routes correspond to awave-guide for Rossby waves. FY01 discussed
a possible relationship between ISOs in convection over the plateau and ISOs in theupper-level wave-like disturbances. Murakami(1981) investigated 12–20-day oscillations on
the Asian subtropical jet over and aroundthe Tibetan Plateau during northern winter.
Plateau topography profoundly influences oscillation formation, propagation, and dissipation.
Although submonthly ISOs (SISOs) exist
over midlatitudes and the region influenced by
the Asian summer monsoon, SISOs near the
Tibetan Plateau during the boreal summer
have not been described precisely. Therefore,
the objectives of this paper were: 1) to detailthe evolution of convective activity and circulation fields in the 7–20-day band over andaround the Tibetan Plateau during the summerand 2) to reveal a link between midlatitudes
and the Asian monsoon, especially over India
and the Bay of Bengal, during the transition
from inactive to active convection over the plateau. The slightly broader temporal band usedin this study is based on spectral analysis results.
Section 2 describes the data and analysis
procedures, and includes results of time series
analyses of equivalent blackbody temperature(Tbb) over the Tibetan Plateau. Section 3 describes large-scale mean fields of circulationand convection for years with enhanced submonthly convective variability over the plateau. Section 4 details composite relationships
between 7–20-day convection over the plateau
and large-scale OLR and circulation anomalies,
and reveals a link in circulation changes between midlatitudes and the Asian monsoon region. Section 5 discusses the formation mechanisms for the wavetrain over the subtropicaljet that links the convective signals between
midlatitudes and the Asian monsoon. Section 6
contains concluding remarks.

2.

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Data and methods of analysis

2.1 Data
This study uses 3-hourly IR-Tbb data derived
from a geostationary meteorological satellite(GMS). Tbb data are defined at each 1  1grid point. This relatively fine spatial and temporal resolution describes the convection overthe plateau. Unfortunately, IR-Tbb data do notcover areas west of 80 E, so daily averaged interpolated outgoing longwave radiation (OLR)
data produced by the National Oceanic andAtmospheric Administration (NOAA) Climatic
Diagnostics Center (CDC) are also used as aproxy for large-scale convection (Liebmann andSmith 1996). Daily averaged National Center
for Environmental Prediction (NCEP)/National
Center for Atmospheric Research (NCAR) reanalysis data represent the large-scale circulation (Kalnay et al. 1996) from 1985 through1998. Both OLR and reanalysis data aredefined on global 2:5  2:5 grids. Streamfunction ðcÞ and velocity potential ðwÞ were calculated from the wind data using a spectral
transform method with spherical harmonic R25
truncation.
2.2 Time-series analysis of Tbb
Figure 1 shows the horizontal distributionof the 14-year mean convective index ðIcÞ for
July–August. The Ic is defined as Ic ¼ 250–Tbb(K) if 250–Tbb is positive; Ic ¼ 0 if 250–Tbb
is negative. FY01 also used the 250 K thresh-

Fig. 1. 14-year averaged (1985–98) Ic for
July–August (thin solid line). The contour interval for Ic is 1. The topographic
contour for 3,000 m is also shown (thicksolid line).

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the study region and the red-noise spectrum
and its 95% confidence level based on the rednoise model of Gilman et al. (1963) and Mitchell (1966). There is a pronounced peak at aperiod of about 14 days and the peak exceeds
the 95% confidence level. By contrast, Fig. 3-bshows the mean for the other years (1985–98(a) Tbb
Power spectrum (K2day)

old to identify convective activity. Convective
activity is most active during July–August, asshown in FY01. The region (28 –34 N, 85 –94 E) over the southern part of the plateau,
where Ic shows a local maximum, is used to
study convective fluctuations.
A spectral analysis using a fast Fouriertransform (FFT) was used to detect the optimal
frequency band of the fluctuations. The seasonal cycle was removed from the daily averaged Tbb data before applying the FFT by subtracting the first three harmonics of the annual
cycle (about 120 days) at each grid point. Cosine tapering was applied to 10% of the timeseries at either end. FFT analyses used the detrended data averaged over the study region
between June 1 and September 30 (122 days).
A 3-point running mean in the frequency domain was applied to the raw spectra to reduce
estimate errors.
Attention is focused on the SISO centered on
the 14-day period. Figure 2 shows the interannual variation of the 11.6–16-day variance ofTbb, which helps to determine when a prominent SISO occurs over the plateau. There is alarge quasi-biennial variation in the varianceat this timescale. We use 1986, 1993, and 1998
to show typical SISO features near the Tibetan Plateau, because the variance is largest in
those three years. Variance is also large in
1985, but because the 14-day period variability
only occurred intermittently with large amplitude during July–August, that year was not
used.
Figure 3-a shows both the 3-year averaged
power spectrum of the daily Tbb anomalies for

900
800
700
600
500
400
300
200
100
0

102

JJAS

area: 28-34N, 85-94E
mean(1986,1993,1998)

101
period (days)(b) Tbb
Power spectrum (K2day)

1548

900
800
700
600
500
400
300
200
100
0

102

JJAS
area: 28-34N, 85-94E
mean(1985-98 except 1986,93,98)

101
period (days)

Fig. 2. The interannual variability of thetotal power spectrum of Tbb for timescales in the 11.64–16-day range overthe base region (28 –34 N, 85 –94 E).

Fig. 3. Ensemble power spectra of areaaveraged (base region: 28 –34 N, 85 –94 E) Tbb daily time series from June to
September (122 days). (a) 3-year mean
(1986, 1993, 1998); (b) 11-year mean
(1985–1998 except 1986, 1993, 1998). A
red-noise spectrum (dashed curve) andits 95% level of significance (solid curve)
based on a lag-1 autocorrelation are
also shown.

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H. FUJINAMI and T. YASUNARI

except 1986, 1993 and 1998); the 14-day variance blends into a red background spectrum.
There is a peak at a period of about 30 days,
but the peak fails to reach the confidence
level. Therefore, this important period in intraseasonal convective variability has large interannual variability. Figure 3-a shows distinct
spectral gaps at periods of about 7 and 20 days,
so a 7–20-day band was chosen as the submonthly timescale for this study. The data
were then temporally filtered using a Lanczos
filter (Duchon 1979) to pass only variability inthe selected period range. This band accounts
for more than 45% of the total Tbb varianceover the study region during July–August for

Fig. 4. Time series of detrended Tbb
anomalies (dashed line) and 7–20-day
filtered anomalies (solid line) averaged
over the base region (28 –34 N, 85 –94 E). Closed circles and crosses denoteselected active phases and inactive
phases of convective activity for a composite analysis, respectively.

1549

the 3 years 1986, 1993, and 1998 (not shown) if
seasonal cycles and diurnal variability are excluded. The region contains an action centerof Tbb fluctuation at these timescales, and thestudy region is treated as a base region. A similar procedure was followed for other variables
used in this study.
2.3 Composite procedure
Temporal variation in the circulation andlarge-scale convection associated with submonthly convective fluctuations over the
Tibetan Plateau was examined using a timecomposite technique. Figure 4 shows the detrended daily Tbb anomalies (dashed line) andthe corresponding 7–20-day filtered anomalies(solid line) over the base region for the 3 years
1986, 1993, and 1998. Large amplitude 7–20day oscillations appear continuously from Julyto August in each year. Dots and crosses respectively denote active and inactive convective events that exceed 1.0 standard deviation(5.6 K) except for two cases. Thirteen cycles
were selected for the composite.
Eight phases comprise each cycle. Figure 5
shows the composite Tbb anomalies over thebase region. Phases 3 and 7 correspond to theinactive and active extremes of convection.

Fig. 5. Composite variability of the 7–
20-day filtered Tbb anomalies averaged
over the base region (28 –34 N, 85 –94 E).

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Since the cycle period is about 14 days, themean interval between phases is about 1.8
days. Composite fields were made for eachphase. Each composite anomaly map wastested for statistical significance at the 95%
level. Before the samples were averaged, composite maps for OLR and streamfunction
anomalies for each year were made to verify
similar spatial structures in each year.
3.

Large-scale mean conditions

Figure 6 shows 3-year (1986, 1993, 1998)
means for July–August of streamfunction andwind fields at; (a) 200 hPa, (b) 500 hPa, and (c)OLR and wind fields at 700 hPa. In Figs. 6-aand b, shading denotes positive values for thedifference in zonal wind between the 3-year
mean and the other years (1985–98 except
1986, 1993, and 1998) mean. Shading in Fig. 6c denotes negative OLR differences. Closed circles denote that the difference is significant
at the 95% confidence level. There is a large
200 hPa anticyclone (the Tibetan high) at 82 Ewith a ridge axis along 28 N (Fig. 6-a). The
westerly Asian subtropical jet is at 40 N. Easterlies prevail south of 28 N. The zonal winddifference has a zonally tripole structure. There
are positive anomalies along 30 N and negative
anomalies along 15 N and 50 N from 40 to150 E. Zonally-elongated positive (negative)
anomalies at 30 N (50 N) suggest that the subtropical jet is at lower latitudes in the 3-year
mean. Significant positive signals are also observed over the jet entrance region (around45 N, 30 E). Enhanced submonthly variability
over the Tibetan Plateau is strongly affected bychanges in the atmospheric basic state over the
midlatitudes. There are two anticyclones along25 N at 500 hPa (Fig. 6-b). One is over Africa
and the other is over the Pacific. The Tibetan
Plateau is between these two anticyclones. A
cyclonic circulation is centered near 20 N,
82 E over India and the Bay of Bengal. At
700 hPa, a monsoon trough is centered overnorthern India (Fig. 6-c). Enhanced convection(OLR < 220 W m2 ) occurred over India, theBay of Bengal, the South China Sea, and the
Tibetan Plateau. Active convective anomalies(shown as shading) spread zonally along 40 N;
suppressed convective signals are along 20 N.
Significant negative OLR anomalies extend
from China to Japan and correspond to en-

Fig. 6. Spatial distributions of the 3-year
mean (1986, 1993, 1998) fields for
July–August. (a) Streamfunction ðcÞ
and wind fields at 200 hPa. The contour
interval is 6:0  10 6 m 2 s1 . The difference between the 3-year averaged
zonal wind and that averaged for theother years (1985–98 except 1986,
1993, 1998) is also shown. Positive
anomalies are shaded. Closed circles
indicate a significant difference at the
95% confidence level. The topographic
contour for 3,000 m is also shown (thicksolid line). (b) As in (a) but for 500 hPa.
The contour interval is 3:0  10 6 m 2 s1 .
(c) OLR and wind fields at 700 hPa.
The contour interval is 20 W m2 .
Shading denotes areas with negative
OLR anomalies. The topographic contour for 1,500 m is also shown (thicksolid line). The reference arrow is (a)40, (b) 20, and (c) 30 m s1 .

December 2004

H. FUJINAMI and T. YASUNARI

Fig. 7. (a) 3-year averaged (1986, 1993,
1998) 7–20-day filtered 200-hPa meridional wind variance for July–August.
The contour interval is 10 m 2 s2 .
Shading represents the percent of thetotal variance (with the seasonal cycle
removed) explained by the 7–20-day
band. The thick solid line is the 3,000 m
topographic contour. (b) As in (a) but for
the years 1985–98, except 1986, 1993,
and 1998. (c) Interannual variation of
7–20-day filtered 200-hPa meridional
wind variance for July–August averaged over 5 –15 E, 40 –60 N (marked‘A’ in (a), solid line with cross marks),
40 –50 E, 40 –60 N (marked ‘B’ in (a),solid line with open circles) and 70 –80 E, 30 –50 N (marked ‘C’ in (a), solidline with closed squares).

hanced convection in the subtropical (Baiu)
frontal zone.
Convective fluctuations over the Tibetan Plateau with timescales of 7–20 days are strongly
affected by waves with the same timescales in

1551

the Asian subtropical jet. Figure 7 shows 200hPa meridional wind variance with timescalesof 7–20 days during July–August. Shading
denotes the percent of the total variance explained by the 7–20-day band (without the seasonal cycle). In the 3-year mean (Fig. 7-a), thevariance field shows large values originating
upstream from the plateau. Variance peaks occur at roughly regular intervals from 10 to130 E, which suggests that the waves are zonally phase-locked. The 7–20-day variance accounts for more than half the total variancein the 3-year mean field over the plateau. By
contrast (Fig. 7-b), in the other years, large
variance occurs downstream from the plateau.
A large variance is present over the three regions marked ‘A’, ‘B’ and ‘C’ in Fig. 7-a. The
variation between the three regions has anearly in-phase relationship (Fig. 7-c), suggesting dynamic linkage between the regions. Variance over region C strongly influences convective variability over the plateau; large values
occur especially in 1986 and 1993. In 1998, although the values are not as large as in 1986
and 1993, the variance peak is over the same
region.
The power spectrum of meridional windtime series over region C is shown in Fig. 8
to confirm the dominant period of the midlatitude waves. In the 3-year mean (Fig. 8-a),the spectrum has a significant peak at 16 days.
Submonthly variability is pronounced in theupper-level midlatitude waves. Unlike the Tbb
spectrum in Fig. 3-b, the meridional mean windspectrum for the remaining years shows aspectral peak at 14 days (Fig. 8-b), although thepeak fails to reach the confidence level by anarrow margin. The variance near 42 N, 90 E
in Fig. 7-b, which is about 1,000 km east of
Region C, has a significant peak with a 14day period (not shown). The Asian subtropicaljet and Tibetan high have vigorous submonthly intraseasonal variability during theboreal summer.
4.

Spatial structure and temporal
evolution

4.1 Convective activity
Figure 9 shows successive composite maps ofthe 7–20-day OLR anomalies from Phases 1 to
8. Open circles denote locally statistically significant grids at the 95% confidence level ac-

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Journal of the Meteorological Society of Japan

Power spectrum (m2s-2day)(a) vwnd
1100
1000
900
800
700
600
500
400
300
200
100
0

102

JJAS

area: 30-50N, 70-80E
mean(1986,1993,1998)

101
period (days)

Power spectrum (m2s-2day)(b) vwnd
1100
1000
900
800
700
600
500
400
300
200
100
0

102

JJAS
area: 30-50N, 70-80E
mean(1985-98 except 1986,93,98)

101
period (days)

Fig. 8. As in Fig. 3, except for the areaaveraged (70 –80 E, 30 –50 N) 200hPa meridional wind time series.

cording to a standard t-test. Negative (positive)
OLR anomalies indicate active (inactive) convective phases.
Phase 1 is characterized by an asymmetric
east-west pattern around the plateau. Significant negative OLR anomalies occur over theeastern part of the Tibetan Plateau and aroundthe Philippines, while positive anomalies extend from the westernmost region of theplateau to over India. Nitta (1983) noted the
14-day period in the out-of-phase precipitation
relationship between central India and theeastern Tibetan Plateau. This relationshipVol. 82, No. 6

is also present in the OLR anomalies. The
asymmetric east-west pattern changes into an
asymmetric north-south (or dipole) pattern between the Tibetan Plateau and the Bay of Bengal as Phase 1 evolves into Phase 3.
The suppressed area of convection over theplateau moves eastward from Phase 3 to 6.
During Phase 4, negative anomalies blossom
over the western flank of the plateau, and thenegative OLR anomalies over the Bay of Bengal migrate westward to the Indian subcontinent. By Phase 5, the asymmetric north-south
anomaly pattern again becomes an asymmetric
east-west pattern, but its sign is opposite thatin Phase 1. Phase 7 is nearly opposite to Phase
3. From Phase 7 to 8, negative OLR anomalies
spread eastward over the plateau, while there
are positive anomalies over the Bay of Bengal
and western part of the plateau. Then, thestructure returns to Phase 1.
This Phase sequence describes a clockwise
rotation of OLR anomalies around 28 N, 90 E
over and around the Tibetan Plateau, including
Indochina, the Bay of Bengal, and India. Yasunari (1979) showed a phase-lag relation for a
10–20-day mode of cloudiness with three reference areas, which are located over central India, Malaysia, and the eastern Tibetan Plateau.
The phase differences in our study agree withthat result.
Significant signals also occur near the Philippines, which have an out-of-phase relation to
the Bay of Bengal, India, and the western Tibetan Plateau, and are in phase with the eastern Tibetan Plateau. Negative OLR anomalies
develop around the Philippines and then move
westward, decaying slightly as they move. The
negative anomalies expand rapidly as theyreach the head of the Bay of Bengal. These
anomalies then propagate westward to India.
These active convective signals are accompanied by lower tropospheric cyclonic anomalies (not shown). Krishnamurti and Ardanuy(1980) and Chen and Chen (1993) show a similar westward propagating mode. Heavy precipitation over India during the Asian summer
monsoon accompanies monsoon lows or monsoon depressions that move westward fromthe Bay of Bengal. Yasunari (1981) noted thatthese monsoon depressions appear in phasewith negative geopotential anomalies withquasi-biweekly periods as well as 40-day

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1553

Fig. 9. Spatial distribution of composite OLR anomalies in the 7–20-day band from Phase 1 (a)through Phase 8 (h). The contour interval is 5 W m2 . The zero contour is not plotted. OLR
anomalies less than 5 W m2 are shaded. Open circles indicate locally statistically significant
grids. The 1,500 m topographic contour is shown as a thick solid line.

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Fig. 10. Spatial distribution of composite streamfunction ðc200 Þ and locally statistically significant
wind vector anomalies in the 7–20-day band at 200 hPa from Phase 1 (a) through Phase 8 (h). The
contour interval is 1:5  10 6 m 2 s1 . The unit vector is 10 m s1 . Shading denotes the Tibetan
Plateau.

periods. The southern portion of the clockwise
rotation includes the monsoon depressions.
Other major components of the clockwise rotation extend over the Tibetan Plateau. The next
focus is on formation processes for OLR
anomalies along the northern part of the clockwise rotation.

4.2 Upper-level atmospheric circulation
Figure 10 shows the evolution of streamfunction ðc200 Þ and wind vector anomalies at
200 hPa from Phase 1 to 8. Wind vectors areplotted where either the u or v component exceeds the 95% local confidence level.
A well-developed wave train extends from

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H. FUJINAMI and T. YASUNARI

North Africa to far-east Asia associated with 7–
20-day convection over the plateau. Anomalies
also overspread nearly the entire Tibetan high.
An anomalous anticyclone (positive c200 anomaly), centered near 42 N, 65 E, develops to thewest of the plateau from Phase 1 to Phase 2 aspart of the wave train. Anomalous northerlies
lie along the eastern fringe of this anticyclone,accompanied by suppressed convection. Thereis an anomalous cyclone (negative c200 anomaly) over the plateau. This cyclone shows aslight southwest-northeast tilt with a southwesterly flow on its eastward flank, where
there is also active convection. There arealso significant easterly anomalies near 18 N
from 60 to 100 E associated with strengthened easterlies. In Phase 3, northerly anomalies dominate over the base region only. Southof 30 N, an anticyclonic anomaly over western
Tibet and easterly anomalies to its south bothmove westward. The westward-propagating
anticyclonic anomaly and a second anticyclonicanomaly centered near 50 N, 30 E merge intoa single north-south elongated anticyclonicanomaly along 35 E by Phase 4.
Structures change dramatically upstream
from the wave train from Phase 3 to 4. c200
anomalies acquire a sign opposite to those at
Phase 1 and 2. A cyclonic anomaly develops
rapidly near 43 N, 60 E, west of the plateau,
and active convection strengthens ahead of thistrough. From India to Indochina along 20 N,
wind anomalies reverse and westerly anomalies become prominent. At Phase 5, the c200
anomalies are the inverse of those in Phase 1
as the spatial relationship between the Tibetan
Plateau and anomalous trough and ridge reverses. Figs. 9-a and e show that the east-west
asymmetric pattern of OLR anomalies has the
opposite sign between the two phases. In Phase6, local strengthening of the Tibetan high is
manifest as a positive c200 anomaly over theplateau maximizing before the peak in convection over the plateau. Convective activity peaks
after the ridge strength peaks. At submonthly
timescales, waves do not amplify in response toheating by convection over the plateau. Rather,
midlatitude waves force the convective activity
over the plateau.
Figure 11 shows a composite phase-longitude
cross-section of the c200 anomalies averaged
between 30 and 45 N to show the charac-

1555

Fig. 11. Composite phase-longitude section of the 7–20-day streamfunction
anomalies ðc200 Þ at 200 hPa averaged
between 30 and 45 N. The contour interval is 1  10 6 m 2 s1 . c200 anomalies
less than 2  10 6 m 2 s1 are shaded.
See the text for the explanation of thearrows.

teristics of the waves more clearly. A quasistanding wave pattern is obvious, with a zonal
wavenumber of about 7. These waves slowly
move eastward with an average eastward
phase speed of about 4 m s1 , as indicated by
the dashed arrow. By contrast, wavelike structures and successive downstream developmentof new centers of action occur, as denoted by
the solid arrow. This phenomenon is generallyinterpreted as Rossby wave energy dispersion
on the sphere (Blackmon et al. 1984) and isclearly distinct from the phase propagation associated with the higher-frequency transients(baroclinic waves). Ambrizzi et al. (1995) andTerao (1998) describe similar downstream developments along the Asian jet during northernsummer, which they interpreted as evidenceof Rossby wave dispersion along a wave-guidefor Rossby waves. The wave characteristics inthose studies are similar to wave characteristics in this study, so an interpretation similar
to the previous studies is applicable here. A
composite group velocity of about 22 m s1
can be estimated from the slope of the solid arrow. Murakami (1981) has shown that in thenorthern winter, 200-hPa 12–20-day vorticity
perturbations have prominent standing oscillations superimposed on modes propagating
slowly eastward (at about 4 m s1 ) in the subtropical jet (45 m s1 ) south of the Tibetan Plateau. FY01 shows a similar standing oscillationin spatial structures around the plateau during March–April (their Fig. 10). The standing

1556

Journal of the Meteorological Society of Japan

Fig. 12. Composite phase-longitude sections of the 7–20-day anomalies between 30 and 35 N. (a) Velocity potential ðw200 Þ anomalies at 200 hPa. The
contour interval is 3  10 5 m 2 s1 ; negative (divergent) anomalies are shaded.
(b) OLR anomalies. The contour interval is 4 W m2 . Anomalies less than
4 W m2 are shaded. A map with adifferent aspect ratio is shown in thebottom figure.

oscillation is associated with marked zonal
wind anomalies over the plateau in winter andspring and meridional wind anomalies in summer.
Upper-level divergence is associated withthe submonthly convective variation and fluctuates with timescales of 7–20-days over andaround the plateau. Figure 12-a shows a phaselongitude section of composite velocity potential
ðw200 Þ anomalies at 200 hPa averaged between
30 and 35 N. The velocity potential represents
only large-scale divergence fields because theinverse Laplacian operation smoothes smallerscale features. Negative anomalies represent
divergence and regions of rising motion. Figure12-b shows the same region as Fig. 12-a, butshows OLR anomalies. For convenience, thecorresponding geographic map is shown at theVol. 82, No. 6

bottom. Regions of upper-air divergence (convergence) agree with negative (positive) OLR
anomalies to the east of 65 E. Significant OLR
anomalies do not appear west of 65 E, suggesting that the relationship between midlatitude
waves and convective activity is made visible
by a divergent circulation to the east of 65 E.
From Phase 8 to 2, upper-air convergence occurs where convection is suppressed over thewestern flank of the plateau. Upper-air divergence and active convection occur over theeastern flank of the plateau. The upper-level
convergence moves eastward from Phase 2 to 4,
and becomes established over the eastern flankof the plateau in Phase 5. In Phase 3, convergence implies large-scale subsidence over theplateau. From Phase 4 to 6, divergence develops over the western flank of the plateau near72 E, as a cyclonic anomaly approaches fromthe west and strengthens near 65 E. The
divergence then moves eastward from Phase
6 to 8 with an active convection signal. Notethat the divergent anomalies are already established over the western Tibetan Plateau andinfluence the base region before the active convection over the base region. Upper-level divergent anomalies probably induce subsequent
active convective anomalies.
Figure 13 shows composites of the total circulation field (left) and velocity potential anddivergent wind anomalies (right) at 200 hPa
from Phase 2 to 4. Positive and negative c200
anomalies in Fig. 10-b, c, d correspond to theridge and trough, respectively, in the total field.
The center of the Tibetan high during Phase 2 isnear 28 N, 70 E (Fig. 13-a). There is an asymmetric east-west pattern of w200 anomalies centered near 90 E between 10 and 50 N (Fig. 13d) that corresponds to the OLR anomalies inFig. 9-b. Around India, easterly divergent wind
anomalies contribute to the easterly anomalies
present in Fig. 10-b. Convergent fields dominate Phase 3 over the plateau (Fig. 13-e), reflecting subsidence that suppresses convection
behind the deep trough. As enhanced convection migrates westward to the head of the Bayof Bengal, an anomalous ascending branch develops there. An asymmetric east-west patternof w200 anomalies occurs at Phase 4 (Fig. 13-f ).
Divergence becomes prominent over the western flank of the plateau where southwesterlies
cross over the plateau ahead of a deepening

December 2004

H. FUJINAMI and T. YASUNARI

1557

Fig. 13. Composite of (a–c) total 200-hPa streamfunction ðc200T Þ and wind fields; (d–f ) 200-hPa
velocity potential ðw200 Þ and locally statistically divergent wind anomalies in the 7–20-day bandfrom Phase 2 to 4. The contour interval is (a–c) 6:0  10 6 m 2 s1 and (d–f ) 2:0  10 5 m 2 s1 . The
unit vector is (a–c) 30 m s1 and (d–f ) 1 m s1 . Shading denotes the Tibetan Plateau.

trough (Fig. 13-c). Negative OLR anomalies
over the western flank of the plateau is inducedby forced ascent over the plateau barrier. Convergence over the eastern Tibetan Plateau
forces subsidence. These sequences suggestthat the midlatitudes and Asian monsoon
region around the plateau interact throughanomalous vertical circulations. Murakami(1981) has revealed that in winter, velocity potential fields at 200 hPa are greatly enhanced
by topographic effects of the plateau, whichtends to reduce the eastward vorticity advection by the strong westerly flow. A plateauinduced dynamic effect also influences thecharacteristics of the circulation fields aroundthe plateau in summer.

1558

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Vol. 82, No. 6

Fig. 14. Composite of (a–c) the 500-hPa streamfunction ðc500 Þ and locally statistically significant
moisture flux vector anomalies; (d–f ) potential instability ðye /dpÞ anomalies from Phase 2 to 4.
Potential instability is defined by the difference in ye between 400 and 500 hPa. The contour
interval is (a–c) 5:0  10 5 m 2 s1 and (d–f ) 0.5 K  100 hPa1 . (a–c) A unit vector is 0.01 g g1 m s1 .
Shading denotes the Tibetan Plateau. (d–f ) Positive (unstable) anomalies are shaded. Open circles
indicate locally statistically significant grids.

4.3

A link between the midlatitudes and the
Asian monsoon region
Circulation fields in the lower troposphereinfluence convection. Figure 14 shows streamfunction ðc500 Þ and the statistically significant
moisture flux vector anomalies at 500 hPa (left)
and the vertical gradient of equivalent poten-

tial temperature (ye , calculated as in Bolton(1980)) between 400 and 500 hPa (right) for
Phase 2 through Phase 4. 500 hPa is the lowest
NCEP/NCAR standard reanalysis pressurelevel above the plateau.
Phase 2 c500 anomalies have an asymmetric
east-west pattern centered along 90 E between

December 2004

H. FUJINAMI and T. YASUNARI

10 and 50 N (Fig. 14-a). There is an anticyclonic circulation near 40 N, 65 E, similar tothe location at 200 hPa (Fig. 10-b). Therefore,
to the north of 30 N, except just over the plateau, the waves are nearly barotropic. A second
positive c500 center is over India. There arenortherly wind anomalies over the plateau.
Northerlies also dominate over the plateau
from 500 through 100 hPa in the total windfields (not shown), moving a dry, cold air massover the plateau. The vertical stratificationis the most stable over the plateau in thisphase (Fig. 14-d) and cloud development is inhibited by entraining or mixing of dry environmental air. Consequently, convective activity
decreases. In Phase 3 (Fig. 14-b), a positive c500
center is over the plateau and comparison of canomalies at 200 hPa (Fig. 10-c) suggests thata baroclinic structure is more prominent overthe plateau. The anticyclonic anomalies suggest a strengthening of divergent flow at thislevel. The dryness, upper-level subsidence, anddivergent flow in the lower troposphere combine to greatly suppress convection over theplateau.
The Phase 4 circulation pattern has approximately reversed Phase 2. The negative c500
anomaly at 20 N, 120 E in Phase 2 moves
westward and reaches the Indian subcontinent
in Phase 4 with active convection anomalies.
Simultaneously, a second negative c500 anomaly develops west of the plateau, corresponding
to a deepening of the upper-level trough (Fig.
10-d). This east-west pattern of c500 anomalies
induces anomalous southerlies towards theplateau, which transports warm, humid air inthe lower atmosphere. The lower atmospherebecomes convectively unstable over the plateau during this phase (Fig. 14-f ). During
Phases 5 and 6, the east-west pressure gradientstrengthens. The most unstable conditions formoist convection appear over the base region
during Phase 6 (not shown).
From Phase 2 to 4, negative c700 anomalies
propagate westward from the Philippines to
India in a manner similar to Figs. 14-a, b and c(not shown), and this propagation corresponds
to a deepening of the monsoon trough over India in the total field.
Figure 15 shows composites of the total
streamfunction at (a) 200, (b) 500, and (c)700 hPa in Phase 5 to highlight the enhance-

1559

Fig. 15. (a) Total streamfunction andwind vectors at 200 hPa for Phase 5. Asin (a) except for the moisture flux vectorat (b) 500 and (c) 700 hPa. The contour
interval is (a) 6:0  10 6 m 2 s1 and (b),(c) 3:0  10 6 m 2 s1 . The unit vector is(a) 40 m s1 , (b) 0.04 g g1 m s1 , and(c) 0.1 g g1 m s1 . Shading denotes the
Tibetan Plateau.

ment of the southerly moisture flow toward theplateau. Figure 15-a shows wind vectors andFigs. 15-b and c show moisture flux at eachlevel. The upper-level trough deepens to the1560

Journal of the Meteorological Society of Japan

west of the plateau, and the ridge strengthensover the plateau. There is a deep trough at
500 hPa from 50 N, 65 E to 10 N, 80 E, andmoisture flow toward the plateau is clearlyobserved to the east of it; this flow moistens
the lower atmosphere over the plateau. At
700 hPa, as the monsoon trough deepens, moisture transport toward the plateau strengthens
to the east of the trough. There is significant
moisture transport toward the plateau between
the lower and middle troposphere.
The Phase 7 circulation features are nearly
identical to those in Figs. 12-a and 13-a in
FY01, which are composites of circulation fields
at 200 and 500 hPa for the active convective
phase over the plateau in summer. Strengthening of the low-level cyclonic circulation enhances low-level convergence, although thesoutherly moisture inflow is reduced over thebase region. The specific humidity at 500 hPa
reaches its maximum because of the increase inprecipitation. The antecedent moistening andconvergence at low levels induces a convective
maximum over the plateau.
5.

Discussion

Convective variability of about 14-day period
over the base region has large interannual
variability (Fig. 2). The variance peak northwest of the plateau (around Region C in Fig. 7a) reflects an alternation between southerly
and northerly winds that import moist and dry
air masses, respectively, over the base region
and cause the large amplitude convective fluctuation in the 3 years (1986, 1993, 1998). By
contrast, the variance peak around 42 N, 87 E
in the other years is north of the plateau (Fig.
7-b) and is about 1,000 km east compared to
Region C. In 1989, 1991, and 1997, @14-day
convective fluctuations are weak over the baseregion (Fig. 2). Although the dominant meridional wind period is @14 days around 42 N,
87 E in these years, convection fluctuates with
9-, 7-, and 6-day periods over the base region(not shown). In 1997, the 40-day period was
also dominant. These periods correspond to
the timescales of the northerly (or dry air) intrusions into the base region. In 1987, 1990,
and 1994, the @14-day convective fluctuations
are also weak because wave activity with 14day timescales is suppressed along the jet near
the plateau (not shown). In any case, the pre-Vol. 82, No. 6

dominant periods in convection over the plateau are closely related to the variability of the
Asian subtropical jet. An interannual change
in the spatial relationship between the waves
and plateau may vary the dominant periods ofconvective variability because of a shift in themeridional wind variance maximum.
Interannual variation in the basic atmospheric state is also likely to control the variation in the dominant periods for convective
fluctuations. The Asian subtropical jet is displaced to lower latitudes in the 3-year mean
fields (Fig. 6-a), indicating that the circulation
around the plateau is sensitive to the influenceof the midlatitude waves. Note that the zonal
wavenumber 7 scale (about 5,000 km in wave
length at 40 N) is almost the same zonal width
as the Tibetan Plateau. The 7–20-day wave
fluctuations may be amplified by wave-plateau
interaction through a resonance-like phenomenon.
This diagnostic study revealed that a welldeveloped wave train extending from North
Africa to far-east Asia along the Asian subtropical jet is associated with convective fluctuations over the Tibetan Plateau. Fukutomi andYasunari (1999) showed that a wave train
grows in both the lower and upper level atmosphere associated with submonthly convective
fluctuations over the South China Sea. The
wave growth is a Rossby mode response toanomalous heating (or cooling). However, significant submonthly-scale convective anomalies
upstream from the midlatitude waves are notobserved. This implies that the wave variability
in Fig. 10 is induced by an internal unstable
dynamic mode. Using a numerical simulationof a large-amplitude anticyclone forced by alocalized mass source on a Beta plane, Hsu
and Plumb (2000) showed that anticyclones areshed westward periodically due to instability.
Popovic and Plumb (2001) used NCEP/NCAR
reanalysis data for the period from 1987 to1990 and confirmed the presence of westwardmigrating anticyclones that detach from themain Tibetan High a few times each summer.
Figure 10 for Phase 3 to 4 shows an anticyclonic anomaly south of 30 N over western
Tibet moving westward. The anticyclone approached 40 E at Phase 4. This wave variation
may be induced dynamically by an unstablemode in the large-scale circulation fields. Re-

December 2004

H. FUJINAMI and T. YASUNARI

cently, Enomoto et al. (2003) described a quasistationary Rossby wave along the Asian jet inAugust and called the waves ‘‘the Silk Road
pattern’’. They proposed that localized midtropospheric descent over the eastern Mediterranean Sea and the Aral Sea is enhanced
through the monsoon-desert mechanism, andassociated upper-tropospheric convergence acts
as a wave source near the jet entrance.
The clockwise rotation of convective signals
implies that dynamic interactions exist between the midlatitudes and Asian summer
monsoon. This study did not investigate thedynamic aspects of the interactions in detail;nevertheless, it is possible to show linkages inatmospheric circulation changes between thetwo regions. That is, simultaneous changes tothe circulation fields west of the plateau andover India enhance convective activity overthe plateau. A second linkage occurs north of
the Philippines. Figure 10-g shows cyclonic
upper air anomalies around 35 N, 125 E andextending to the Philippines in Phase 7. This
cyclonic circulation is above mid- and lowertropospheric cyclonic anomalies (not shown),
which suggests barotropic enhancement of thecyclonic circulation. The cyclonic anomalies denote a weakening of the easternmost region of
the Tibetan high in the upper troposphere, andthe westernmost region of the subtropical highin the middle and lower troposphere. Subsequently, upper-air cyclonic anomalies weaken
as cyclonic anomalies develop around the Philippines in the lower and middle troposphere
coupled with an active convection signal in
Phase 8. This sequence suggests that the SISO
of the midlatitude waves initiates a SISO in thecirculation and convection north of the Philippines. Enomoto et al. (2003) hypothesized thatthe equivalent barotropic structure of the Bonin (Ogasawara) high near Japan in summer
results from the propagation of stationary
Rossby waves along the Asian subtropical jet. A
similar process may influence changes in circulation anomalies around the Philippines.
Figure 16 shows the difference in 7–20-day
OLR variance between the 3-year mean andthe other years mean during July–August.
Positive anomalies extend over the base regionand over other Asian monsoon regions as well.
Submonthly variability increased over midlatitudes and the Asian monsoon regions in the1561

Fig. 16. Difference in the 7–20-day OLR
variance between the 3-year (1986,
1993, 1998) and 11-year means (1985–
1998 except 1986, 1993, 1998) for July–
August. Positive areas are shaded.
Open circles indicate locally statistically significant grids in the difference. The 1500-m topographic contour
is shown as a thick solid line.

3 years. The upper-level Asian subtropical jet isdisplaced to lower latitudes in the 3-year mean
fields. Further statistical and dynamic studies
will determine how this displacement facilitates extratropical-subtropical interaction on
submonthly timescales.
6.

Concluding remarks

This paper examined the submonthly (7–20day timescales) intraseasonal oscillation (SISO)
of convective activity over the Tibetan Plateau
and the associated large-scale convective activity and atmospheric circulations over andaround the plateau during the mature phaseof the Asian summer monsoon (July–August).
A diagnostic study examined the three years
when prominent submonthly convective variability was observed over the plateau. Relationships between midlatitudes and the Asiansummer monsoon region were identified andare summarized as follows.
1) The variation in convection over thesouthern part of the plateau (the base region)
with @14-day periods shows large interannual
variation. The variance is prominent especiallyin 1986, 1993, and 1998. Convective fluctuations with a timescale of 7–20 days account for
more than 45% of the total variance during

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Journal of the Meteorological Society of Japan

July–August in the 3-year mean field over thebase region.
2) Climatologically, areas of the Asian subtropical jet and Tibetan high near the plateau
have vigorous submonthly intraseasonal variability during the boreal summer. In 1986,
1993, and 1998, the 7–20-day meridional windvariance fields have large values upstream
from the plateau with a peak northwest of theplateau. In addition, the upper-level Asiansubtropical jet is displaced to lower latitudes.
3) The submonthly convective anomalies
show a clockwise rotation around 28 N, 90 E
over and around the Tibetan Plateau, including
Indochina, the Bay of Bengal, and India. Othersignificant anomalies appear near the Philippines and South China Sea. The southern trackof the clockwise rotation includes westward
moving monsoon depressions, and the northern
track is induced by waves on the Asian subtropical jet.
4) A well-developed wave train extends from
North Africa to far-east Asia along the Asiansubtropical jet. The wave train is associatedwith submonthly convective fluctuations overthe plateau. The waves have quasi-stationary
behavior and a Rossby wave-like structure witha wavenumber of approximately 7. The waves
also have a slow eastward phase speed. Circulation anomalies are observed over all of the
Tibetan high.
5) Before convection minimizes over the plateau, an upper-air ridge develops to the west of
the plateau as part of the wave train. The ridge
induces northeasterly flow that moves dry, cold
air over the plateau. Suppressed convection
over the plateau results from antecedent drying in the troposphere, strengthening of upperlevel subsidence, and low-level divergence.
6) Before convection maximizes over the plateau, an upper-level trough develops and deepens to the west of the plateau. Southwesterly
flow ahead of the trough induces forced ascent(descent) over the western (eastern) flank of
the plateau with active (suppressed) convection, causing an asymmetric east-west patternin convection.
7) As the trough deepens to the west of theplateau, a cyclonic circulation also forms overIndia in the lower through middle troposphere.
The enhanced trough forces a southerly moist
flow in the lower atmosphere over the plateau.

Vol. 82, No. 6

The moistening helps create a convectively unstable condition. Enhanced low-level convergence supports a convective maximum over theplateau.
The SISO of convection over the plateau
mainly consists of convection with strong diurnal variation. The area of active convection
occurs within an anticyclonic circulation withrelatively weak synoptic-scale winds. Further
studies will be needed to solve the relationshipbetween synoptic conditions and the local circulation over the plateau. The origin of thelarge-scale atmospheric circulation with a period of @14 days, which modulates convective
variability at the same timescale over the plateau, should also be investigated.
Acknowledgments
We thank Dr. H. Ueda of the University ofTsukuba for his thoughtful discussions andcomments. Valuable comments and suggestions
provided by Prof. F. Kimura and Prof. H.L. Tanaka of the University of Tsukuba and Prof. A.
Kitoh of the Meteorological Research Institute(MRI) are gratefully acknowledged. Thanks
are also due to Dr. N. Endo of the Frontier Research Center for Global Change (FRCGC) for
helpful suggestions. The authors express their
thanks to Dr. Y. Fukutomi of FRCGC and Mr.
M. Hayasaki of the National Institute forEnvironmental Studies (NIES) for their helpfulcomments and providing some computational
programs. GMS-IR data were provided by Dr.
N. Yamazaki of MRI. The Generic Mapping
Tools (GMT) graphics package developed by
Wessel and Smith (1995) was used in preparing
a part of the figures. This work was partially
supported by a Grant-in-Aid for Scientific Research (A) (1), No. 14204044 from the Ministry of Education, Culture, Sports, Science andTechnology, and by the Japan Science andTechnology Agency (JST).
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