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MONTHLY WEATHER REVIEWVOLUME 137

The Effects of Midlatitude Waves over and around the Tibetan Plateau on Submonthly
Variability of the East Asian Summer Monsoon
HATSUKI FUJINAMI AND TETSUZO YASUNARI
Hydrospheric Atmospheric Research Center, Nagoya University, Nagoya, Japan(Manuscript received 16 October 2008, in final form 17 January 2009)
ABSTRACT
Convective variability at submonthly time scales (7–25 days) over the Yangtze and Huaihe River basins(YHRBs) and associated large-scale atmospheric circulation during the mei-yu season were examined using
interpolated outgoing longwave radiation (OLR) and NCEP–NCAR reanalysis data for 12 yr having active
submonthly convective fluctuation over the YHRBs within the period 1979–2004. Correlations between
convection anomalies over the YHRBs and upper-level streamfunction anomalies at every grid point show
two contrasting patterns. One pattern exhibits high correlations along the northern to eastern peripheries of
the Tibetan Plateau (defined as the NET pattern), whereas the other has high correlations across the Tibetan
Plateau (defined as the AT pattern). Composite analysis of the NET pattern shows slow southward migrationof convection anomalies from the northeastern periphery of the Tibetan Plateau to southern China, in
relation to southward migration of the mei-yu front caused by simultaneous amplification of upper- and lowlevel waves north of the YHRBs. In the AT pattern, convection anomalies migrate eastward from thewestern Tibetan Plateau to the YHRBs. A low-level vortex is created at the lee of the plateau by eastwardmoving upper-level wave packets and associated convection from the plateau. Rossby wave trains along the
Asian jet characterize the upper-level circulation anomalies in the two patterns. The basic state of the Asianjet during the mei-yu season differs between the two patterns, especially around the Tibetan Plateau. The
Asian jet has a northward arclike structure in NET years, while a zonal jet dominates in AT years. These
differences could alter the Rossby wave train propagation route. Furthermore, the larger zonal wavenumberof AT waves (;7–8) than of NET waves (;6) means faster zonal phase speed relative to the ground in the AT
pattern than in the NET pattern. These differences likely explain the meridional amplification of waves northof the YHRBs in the NET pattern and the eastward wave movement across the plateau in the AT pattern.

1. Introduction
Climatologically, the rainy season in China has large
seasonal variation, with three high-rainfall periods andlocations: a presummer rainy season in mid-May overSouth China, the mei-yu season over the Yangtze andHuaihe River basins (YHRBs) in central China from
mid-June to mid-July, and a summer rainy season in
North/Northeast China after mid-July (e.g., Ding 1992;
Tian and Yasunari 1998; Ding and Chan 2005). The seasonal variation shows abrupt northward shifts among thethree regions, associated with abrupt seasonal changes inlarge-scale atmospheric circulation over East Asia (e.g.,
Ueda and Yasunari 1996).

Corresponding author address: Dr. Hatsuki Fujinami, Hydrospheric Atmospheric Research Center, Nagoya University, Furo-cho,
Chikusa-ku, Nagoya 464-8601, Japan.
E-mail: hatsuki@hyarc.nagoya-u.ac.jp
DOI: 10.1175/2009MWR2826.1
Ó 2009 American Meteorological Society

In the mei-yu season, a quasi-stationary front (called
the mei-yu front in China and the baiu front in Japan)
is frequently observed in the synoptic weather map,
extending from the southeastern edge of the Tibetan
Plateau to the east of Japan. Once the front appears
over East Asia, the location of the east–west elongatedcloud zone (i.e., mei-yu–baiu front) changes very slowly,
although mesoscale precipitation systems often move
eastward along it.
The mei-yu–baiu front is an important cloud andprecipitation system of the East Asian monsoon, providing large amounts of precipitation in early summer.
The front forms as a subtropical front under the influence of atmospheric circulation in both the midlatitudes
and subtropical–tropical monsoon regions (Ninomiya
1984). The frontal zone corresponds to areas of a large
meridional moisture gradient, and moisture transported
from the South Asian and western North Pacific monsoon regions plays an important role in maintaining the

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FUJINAMI AND YASUNARI

mei-yu–baiu front. Although the mei-yu and baiu systems have common features, their associated atmospheric structures are not identical. Over the YHRBs,
the atmospheric structure shows a weak meridionaltemperature gradient and strong horizontal wind shear
in the lower troposphere, whereas areas around the East
China Sea and Japan are characterized by large meridional temperature gradients (Ninomiya and Muraki
1986; Ding 1992; Ding and Chan 2005).
Heavy rainfalls during the mei-yu period are mainly
generated by meso-a- and meso-b-scale disturbancesthat are embedded on, and propagate along, the mei-yu
front. Meso-a-scale convective systems of the mei-yu
period may be classified broadly into two types: theYangtze River Valley shear line and the low-level vortex(Chen 2004; Ding and Chan 2005). A low-level vortex often occurs over the eastern flank of the Tibetan Plateau.
A low-level vortex originating over the Tibetan Plateau
sometimes moves eastward and triggers the vortex when
it moves off the plateau (Tao and Ding 1981; Wang
1987; Wang and Orlanski 1987; Yasunari and Miwa
2006). The vortex develops and is maintained mainly
by latent heating due to strong convection (Wang and
Orlanski 1987; Wang et al. 1993).
The convection activity over the YHRBs exhibits substantial intraseasonal variability during the mei-yu period.
Intraseasonal oscillations (ISOs) on so-called submonthly(;6–25 days; e.g., Vincent et al. 1998) and Madden–
Julian oscillation (MJO; 30–60 days; e.g., Madden andJulian 1972, 1994) time scales strongly control the activityof convections over the YHRBs and South China Sea
during the summer monsoon. Previous studies havenoted a relationship between ISOs in convection over the
YHRBs and those over the South China Sea (Fukutomi
and Yasunari 1999; Fukutomi and Yasunari 2002; Chenet al. 2000; Zhu et al. 2003; Mao and Chan 2005). Chenet al. (2000) reported that both 12–24-day and 30–60-day
ISOs of convection tend to fluctuate out of phasebetween the two regions from May to August. When
convection is active (suppressed) over the YHRBs andsuppressed (active) over the South China Sea, an anticyclonic (cyclonic) circulation anomaly appears in thelower troposphere over the South China Sea. The anticyclonic (cyclonic) anomaly accompanies westerly–
southwesterly (easterly–northeasterly) wind anomalies to
its north, indicating that an increase (decrease) in moisture supply to the YHRBs from the south is associatedwith active (suppressed) YHRB convection. Submonthly-scale convection anomalies originate from thewestern North Pacific and propagate westward over theSouth China Sea in the early summer, accompanied by
low-level circulation anomalies (Fukutomi and Yasunari
1999; Fukutomi and Yasunari 2002; Mao and Chan 2005).

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During the northern summer, ISOs also dominate in
the extratropics (e.g., Ambrizzi et al. 1995; Terao 1998;
Fujinami and Yasunari 2004, hereafter FY04). Quasistationary Rossby waves occur in the Asian subtropical jetwith periods of ;14 and 30–45 days (Terao 1998). Thesewaves, called the ‘‘Silk Road pattern,’’ follow the waveguide for Rossby waves (Terao 1998; Enomoto et al. 2003).
FY04 revealed that 7–20-day convective fluctuation over
and around the Tibetan Plateau during July and August is
strongly controlled by Rossby waves along the Asian jet.
The mei-yu front includes both tropical and extratropical variations because of its geographical location.
Therefore, tropical and extratropical interactions are an
inherent property of the variation (Hsu 2005). Mao andWu (2006) examined the ISO of rainfall over the
YHRBs during the mei-yu season and related atmospheric circulation for the summer of 1991. An ISO
of 15–35 days was dominant in 1991; however, thisISO time scale is not typical over the YHRBs. The lowlevel troposphere has an anomalous anticyclone (cyclone) appearing from the northern South China Sea
to the Philippine Sea in the active (break) phase,whereas the upper-level troposphere is characterized byan anomalous anticyclone (cyclone) over eastern China
and an anomalous cyclone (anticyclone) over the northern Tibetan Plateau. The coupled anomalous circulation
pattern between the lower and upper troposphere favors
large-scale ascending (descending) motion, and hence
enhanced (reduced) rainfall over the YHRBs.
However, previous studies have not examined the detailed spatial and temporal evolution of large-scale convection and the structure of atmospheric circulation
associated with ISOs of convection over the YHRBs
during the mei-yu season. The effects of midlatitude
waves (e.g., Rossby waves) along the Asian westerly jet
and of the Tibetan Plateau on mei-yu activity also remain
poorly understood. During boreal summer, midlatitude
systems that vary on submonthly time scales favorably
propagate along the Asian westerly jet (e.g., FY04).
Thus, a close connection between midlatitude systems
and mei-yu variability in early summer could be expected. Therefore, based on the submonthly convective
fluctuations over the YHRBs during the mei-yu season,
we examined the spatiotemporal evolution of convection
and circulation over East Asia, focusing on the relationship between upper-level atmospheric circulation in
the midlatitudes and convection over the YHRBs.
In section 2, we describe the data and data processing.
In section 3, we present the general characteristics ofmean convective activity and atmospheric circulation
during the mei-yu period. Time series analysis is also
included to demonstrate the interannual variability of
submonthly-scale ISO (SISO) of convection over the2288

MONTHLY WEATHER REVIEWVOLUME 137

YHRBs and to reveal dominant periods of active SISO
years based on wavelet analysis. In section 4, we present
composite analysis results for active years of SISO toreveal the space–time evolution of the large-scale convection and circulation associated with submonthlyscale convective fluctuation over the YHRBs. Casestudies are also shown to clarify how the midlatitude
waves affect real weather systems on submonthly timescales over the YHRBs. In section 5, we discuss effectsof background time-mean flows on the propagationroute of midlatitude waves and the role of the Tibetan
Plateau in submonthly-scale convection over the YHRBs.
Finally, in section 6, we summarize the results.

2. Data and processing
Interpolated outgoing longwave radiation (OLR) data
produced by the National Oceanic and Atmospheric
Administration/Climatic Diagnostics Center (NOAA/
CDC) are used as a proxy for convection (Liebmann
and Smith 1996). National Centers for Environmental
Prediction–National Center for Atmospheric Research(NCEP–NCAR) reanalysis data represent large-scale
circulation and include zonal and meridional windcomponents (u, y) (Kalnay et al. 1996). Both types ofdata include daily averages for the 26-yr period (1979–2004) at global coverage on a 2.58 latitude–longitudegrid. Streamfunction (c) and divergence were calculated from the wind data using a spectral transform
method. Gauge-based analysis data of daily precipitation on a 0.58 latitude–longitude grid over East Asia
were also used for a case study (Xie et al. 2007).
Climatologically, circulation fields change abruptly
before and after the mei-yu period (10 June–10 July;
Ding 1992; Ding and Chan 2005). The different mean
basic state probably affects the development and structure of disturbances over the YHRBs and associatedlarge-scale circulation (e.g., Matthews and Kiladis
1999a,b; Fukutomi and Yasunari 2002). Therefore, data
during the climatological mei-yu period were extracted
to examine the features of SISO over the YHRBs.
Daily anomalies of OLR and the reanalysis data werecomputed by subtracting the first three harmonics of theannual cycle (about 120 days) at each grid point for eachyear. Submonthly (7–25 day) perturbations were thencomputed by applying a Lanczos filter (Duchon 1979) tothe detrended anomalies.

3. Submonthly convective variability over the
Yangtze and Huaihe River basinsa. Large-scale mean conditions
Figure 1 shows the climatological structure of convective activity and atmospheric circulation, based on

FIG. 1. (a) Mean OLR (shading), 700-hPa streamfunction (c),and wind vectors for the mei-yu period (10 June–10 July) 1979–
2004. The contour interval for c is 2 3 106 m2 s21. The topographic
contour line for 1500 m is also shown (thick dashed line). The
key convective variation region is enclosed by the 258–358N,
1108–1208E grid box. (b) Same as (a), but for the 200-hPa streamfunction (c) and wind vectors. Two rectangular domains, labeled N(358–408N, 1108–1208E) and S (208–258N, 1108–1208E), are indicatedby dotted lines (as explained in the text). The contour interval is1 3 107 m2 s21. Shading denotes areas higher than 1500 m in surface
altitude.

the 26-yr average during the mei-yu period. A zonallyelongated, banded active convection zone (OLR , 230
W m22) is observable from East China to Japan, corresponding to the mei-yu–baiu frontal zone (Fig. 1a).
The frontal zone is accompanied by strong low-levelsouthwesterlies. Deep convections (OLR , 200 W m22)
extend over the Bay of Bengal, Indochina Peninsula,and to the west of the Philippines. Active convection
also appears over the southeastern part of the Tibetan
Plateau. A subtropical anticyclone broadly covers thewestern North Pacific, where convection is suppressed.
In the upper troposphere, an upper-level anticyclone

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FIG. 2. Interannual variation of 7–25-day OLR variance overthe key region (258–358N, 1108–1208E) during the mei-yu period(10 June–10 July). The dashed line denotes the 26-yr mean of thevariance (508.7 W2 m24).

(Tibetan high) is evident south of the Tibetan Plateau(Fig. 1b). To the north of the Tibetan high, the Asian
westerly jet lies along the jet axis at approximately 37.58N.

b. Characteristics of submonthly-scale convective
variability during the mei-yu period
The boxed region (258–358N, 1108–1208E) in Fig. 1a istreated as a key region for the detailed examination of
submonthly-scale convective fluctuation over the YHRB.
The region includes active convection over East China
and is characterized as a convergence zone along the
low-level strong southwesterly flow and a shallow upperlevel trough (Figs. 1a,b). Figure 2 shows the interannual
variation of 7–25-day OLR variance over the key region
during the mei-yu period. The variance exhibits large
interannual variation from 1979 to 1999, whereas large
variance persists after 2000 with no clear interannual

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variation. In years of suppressed SISO, we found ISOsof approximately 30 days (e.g., 1988, 1991, 1994, and1998) and less remarkable ISOs of both submonthly andlonger time scales (e.g., 1980, 1983, 1987, 1996; data
not shown). Hereafter, we scrutinize active SISO years(12 yr: 1981, 1985, 1986, 1989, 1992, 1993, 1995, 2000,
2001, 2002, 2003, and 2004), characterized by variance
exceeding the 26-yr mean value (508.7 W2 m24; dashedline in Fig. 2), to show typical SISO features over the
YHRBs.
We applied wavelet analysis (Torrence and Compo
1998) to the detrended OLR time series, averaged overthe key region for the active SISO years. To avoid redundancy, the wavelet spectra are presented only for1992 and 1993, the years with the minimum and maximum submonthly variance, respectively, among the
12 yr (Fig. 2). In both years, submonthly variations are
dominant during the climatological mei-yu period, reflecting the appearance of the mei-yu front. The 1992
OLR anomalies have quasi-periodic oscillation withthree active convection phases during the mei-yu period(Fig. 3a). The active phases include large and continuous precipitation (;10–20 mm day21). The wavelet
spectrum shows a statistically significant peak with aperiod of approximately 7–12 days during the mei-yu
period (Fig. 3b). In 1993, the detrended OLR anomalies
also exhibit remarkable intraseasonal variation, butlarger amplitude than in 1992. The wavelet spectrum
has a significant peak in an 8–25-day period range during the mei-yu period (Fig. 3d). In the other years, except for 2002, statistically significant (at the 95% level)
peaks are evident in a submonthly (7–25 day) range

FIG. 3. (a) Time series of detrended OLR anomalies (dashed line), 7–25-day filtered anomalies (solid line), andunfiltered precipitation (black bar) averaged over the key region for 1992. (b) The local wavelet power spectrum of
the detrended 1992 OLR anomalies, normalized by 1/s2 [s2 5 556.4 (W m22)2]. The contours are at 1, 2, 3, 4, 8, and12. Heavy shading denotes areas with power values .4, whereas light shading denotes areas between 2 and 4. The
thick solid (dotted) contour line encloses regions of .95% (90%) confidence for a red noise process with a lag
1 coefficient. The dashed line indicates the cone of influence outside of which edge effects become important. (c),(d)As in (a),(b), respectively, but for 1993. The power is normalized by 1/s2 [s2 5 331.6 (W m22)2].

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MONTHLY WEATHER REVIEW

FIG. 4. The 7–25-day OLR variance (contours) and the percentage it composes of total variance (shading) during the mei-yu
period (10 June–10 July) for active SISO years (e.g., 1981, 1985,
1986, 1989, 1992, 1993, 1995, 2000, 2001, 2002, 2003, and 2004). The
contour interval is 100 (W m22)2. The thick dashed line indicates
the topographic contour line for 1500 m.

within the mei-yu period (not shown). Thus, submonthly oscillation was the dominant mode during themei-yu period in these years.
Figure 4 presents the 12-yr mean structure in varianceand the percentage variance of the 7–25-day filtered
OLR. A zonal distribution of high variance (.600
W2 m24), extending from South China to the westernpart of Japan, corresponds to the mei-yu–baiu frontal
activity (Fig. 1a). There is a local maximum of variance
exceeding 700 W2 m24 within the key region. The key
region overlaps the percentage of total variance exceeding 50%–60%. High values up to 900 W2 m24 arelocated over the Bay of Bengal, the South China Sea,and the Philippine Sea. Areas exceeding 50% of thepercentage variance are also seen over the southern
Tibetan Plateau, the northeastern part of the Indochina
Peninsula, the north South China Sea, and the Philippine
Sea, suggesting a possible relationship in submonthly convective variability between the YHRBs and these regions.

4. Spatial structure and temporal evolution
a. Two contrasting relationships between convection
and upper-level circulation
In this section, we examine the temporal variation inconvection and circulation associated with submonthly

VOLUME 137

convective fluctuations over the YHRBs, based on atime composite analysis. Before the composite analysis,
the relationship between convection over the YHRBs
and upper-level atmospheric circulation was investigated. In the upper-level atmosphere, there is a large
meridional pressure gradient between regions to thenorth (N region) and south (S region) of the key region,
as shown in Fig. 1b. Active convection associatedwith mei-yu frontal activity generally occurs along thesouthern side of the upper-level westerly jet, where
upper-level divergence dominates (Figs. 1a,b). It is expected that midlatitude waves propagating along the jet
could change upper-level wind fields by varying theupper-level meridional pressure gradient and divergence fields over the YHRBs. This, in turn, could influence the mei-yu frontal activity (e.g., FY04). Therefore, we focus on the relationship between convection
over the YHRBs and upper-level atmospheric circulation in the N and S regions. Figure 5a is a scatter diagram of the correlation coefficients between 7–25-day
OLR anomalies averaged over the key region and 7–25day 200-hPa streamfunction anomalies in the N and S
regions for each of the 12 yr. The largest absolute correlation coefficient within each region (15 grids) is
plotted for each year. In the S region, OLR anomalies
correlate negatively with the 200-hPa streamfunction(c200) anomalies, except in 2000 and 2002. In contrast, in
the N region, the signs of the correlation coefficientsvary from year to year. Based on the sign of the correlation in the N region, we roughly classified theseyears (except 2000 and 2002) into two types: positivecorrelation years (e.g., 1981, 1985, 1989, 1992, and 2004;
closed circles in Fig. 5a) and negative-correlation years(e.g., 1986, 1993, 1995, 2001, and 2003; open circles inFig. 5a). The negative-correlation years include weakcorrelation years in the N region (e.g., 1986). However,a remarkable difference between the positive- andnegative-correlation years is whether positive correlation exists in the N region. Therefore, we used the year
to examine the general features of positive- and negativecorrelation years. This does not alter the following results for negative-correlation years. Figures 5b,c show
the spatial distribution of the correlation coefficients
between the YHRB OLR anomalies and c200 anomalies
in every grid for the two types. We assembled the filtered OLR anomalies averaged over the key region andc200 anomalies in every grid during the mei-yu period in
the positive- and negative-correlation years, respectively, following the sequential order of the years tomake a new dataset (155 days). The correlation coefficient patterns in Figs. 5b,c were constructed using the
newly formed time series. In the positive-correlation
years (Fig. 5b), large values extend from the northern toJULY 2009

FUJINAMI AND YASUNARI

FIG. 5. Scatter diagram of the correlation coefficient between
7–25-day OLR anomalies over the key region and 7–25-day 200hPa streamfunction (c) anomalies in the N and S regions for theactive SISO years. Filled and open circles denote years of NET andAT patterns, respectively. (b) Correlation coefficients between
OLR anomalies over the key region and 200-hPa c anomalies inevery grid in the 7–25-day bands for the NET pattern. The contour
interval is 0.1. The zero contour line is not plotted. Areas that havecorrelation coefficients ,20.3 are shaded. (c) As in (b), but for theAT pattern. (d) As in (b),(c), but for 700-hPa c anomalies for boththe NET and AT patterns. See the text for an explanation of thecorrelation coefficient calculation.

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the eastern peripheries of the Tibetan Plateau (hereafter defined as the NET pattern); there is a high positive
correlation (more than 0.6) around 408N, 1158E, but anegative correlation (less than 20.6) centered on 258N,
1208E. Correlations larger than ;0.2 are significant
at the 99% level in Figs. 5b–d. Active (suppressed)
convections over the YHRBs are associated with ananomalous trough (ridge) to the north of the YHRBs
and anomalous ridge (trough) to the south of theYHRBs in the upper troposphere. However, in thenegative-correlation years (Fig. 5c), high correlation
coefficients appear along approximately 308N across the
Tibetan Plateau with a wave train–like distribution(hereafter defined as the AT pattern); the entire key
region has negative correlations. Active (suppressed)
convections over the YHRBs are associated with theupper-level anomalous ridge (trough). The two types ofcorrelation coefficient patterns suggest that the dominant processes inducing SISO in mei-yu activity are
different even though they have the same submonthly
time scales. Therefore, we performed composite analysis separately for years with NET and AT patterns.
For 2000, the correlation pattern was similar to that for
the AT pattern, but the overall pattern was shifted
westward by about 108 relative to that for the AT pattern (not shown). In 2002, the correlation distribution
showed a wavelike structure from the head of the Bay of
Bengal to the east of Japan (not shown), and positive
correlations extended broadly from northern China toeast of Taiwan, including the YHRBs.
Figure 5d is the same as Figs. 5a,b, but for 700-hPa
streamfunction (c700) anomalies. The data for both
NET and AT pattern years are assembled together (310days) because overall correlation distributions in bothpattern years resemble each other in the key region.
Correlation coefficients tend to show opposite signs
between the YHRBs and areas to their south including
the Indochina Peninsula, the South China Sea, and thePhilippine Sea. Active (suppressed) convection over the
YHRBs is associated with low-level cyclonic (anticyclonic) circulation over the YHRBs and anticyclonic(cyclonic) anomalies from the Indochina Peninsula to
the north of the Philippines, indicating the enhancementof southwesterly (northeasterly) anomalies over the
YHRBs. One difference between the two patterns isthat significant positive correlations (;0.2) in NET
years spread more northward (to 458N) than those in
Fig. 5d, centered on 32.58N, 112.58E. In AT years, significant positive correlations are confined between 258
and 358N and centered on 308N, 1108E. South of theYHRBs, the NET years show a zonal distribution ofsignificant negative correlations from 908 to 1308E witha single center on the South China Sea, whereas the AT

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MONTHLY WEATHER REVIEWVOLUME 137

TABLE 1. Dates of the selected maximum (phase 3) and minimum (phase 7) OLR indices in the key region (258–358N, 1108–1208E) in
NET and AT years. SF and L indicate a stationary front and a low over the key region, respectively, according to a surface weather chartfor day 2 1 and day 1 1 of the date of the minimum OLR indices. SF1L indicates that a low is embedded on a stationary front.
NET
Yr

Max OLR (phase 3)

1981
1981
1981
1985
1989
1989
1989
1992
1992
1992
1992
2004
2004
2004

—
15 Jun5 Jul
28 Jun—23 Jun
5 Jul—
18 Jun28 Jun
7 Jul
10 Jun19 Jun28 JunAT
Min OLR (phase 7)
10 Jun26 Jun10 Jul
3 Jul
18 Jun28 Jun10 Jul
14 Jun23 Jun
3 Jul—
14 Jun23 Jun
—

L
SF1L
SF
SF
SF1L
SF
SF
SF1L
SF1L
SF1L
—
SF
SF1L
—

years show a more zonally elongated structure of negative correlations with local minima over the western
Tibetan Plateau, India, the Indochina Peninsula, and
the Philippine Sea.

b. Composite procedure
We used time composite techniques to establish therelationships between convections over the YHRBs andlarge-scale atmospheric conditions in years of NET andAT patterns. To create the composites, we selected 12
active and 11 break peaks from the 7–25-day OLR timeseries for the NET pattern and 12 active and 12 breakpeaks for the AT pattern (Table 1). According to asurface weather chart, 22 active peaks (total of 24 samples) were accompanied by a stationary front (i.e., themei-yu front) over the key region. In addition, 17 of the
22 active cases had a low embedded on the front. Thus,
composite variations based on the selected samples
represent mei-yu frontal activities. As described in the
Glossary of Meteorology published by the American
Meteorological Society (see online at http://amsglossary.
allenpress.com/glossary; Glickman 2000), a front thatis moving at less than approximately 5 kt (2.6 m s21) isgenerally considered quasi-stationary. In synoptic weather
analysis, a quasi-stationary front is defined as one thathas not moved appreciably from its position on the last
previous synoptic chart (e.g., from 3 or 6 h previously).
Thus, a ‘‘quasi-stationary front’’ means that the frontchanges little in position as a snapshot. Here we define
stationary or quasi-stationary to mean little change inposition compared to the previous time or phase.
Hence, quasi-stationary phenomena (i.e., fronts, wave
trains) may vary on submonthly time scales.

Yr

Max OLR (phase 3)

1986
1986
1986
1993
1993
1993
1993
1995
1995
1995
2001
2001
2003
2003

15 Jun25 Jun
8 Jul—
15 Jun25 Jun10 Jul11 Jun19 Jun27 Jun—15 Jun17 Jun2 Jul

Min OLR (phase 7)
20 Jun
3 Jul—
11 Jun20 Jun30 Jun—15 Jun23 Jun
3 Jul
10 Jun20 Jun24 Jun
9 Jul

SF1L
SF1L
—
L
SF1L
SF1L
—
SF1L
SF1L
SF1L
SF1L
SF1L
SF1L
SF1L

For all the samples, standard deviations exceed themean of the active SISO years (May–August) by 0.8(11.5 W2 m24), except for three break samples. The
standard deviations of composite active and breakpeaks exceed 1.5 (21.6 W2 m24) for both NET and AT
patterns. Each cycle was divided into eight different
phases, as done by FY04 and Mao and Chan (2005).
Phases 3 and 7 correspond to the inactive and active
extremes of the YHRBs convection. The mean timespan between two subsequent phases was ;1.5 days inboth the NET and AT patterns, because the composite
cycle was ;12 days. Composite fields were createdfor each phase. The statistical significance of the composite results at each grid was estimated using Student’s
t tests.

c. NET patternFigure 6 illustrates the relationships of convections
over the YHRBs, atmospheric circulation, and convective activity from phases 5 to 8 of the NET pattern using
composite maps of 7–25-day filtered OLR, wind vectors,
and streamfunction (c) anomalies. The anomaly structures of phases 1–4 are nearly the same as those fromphases 5 to 8, but of opposite sign. Therefore, only
phases 5–8 are illustrated, including the transition frominactive to active convection over the YHRBs.
The left-hand panels (Figs. 6a–d) show composites of
convection and 700-hPa wind vector and streamfunction(c700) anomalies. A notable feature of the phase sequence is the southward movement of active convection
anomalies (dark shading) from areas to the northeastof the Tibetan Plateau (around 388N, 1158E) to South
China with a mean speed of 2.6 m s21, having a zonal

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FIG. 6. (a)–(d) Composites of 7–25-day OLR, 700-hPa wind vectors, and streamfunction (c)anomalies from phases 5 to 8 for the NET pattern. OLR anomalies less (greater) than 25 (5) W m22
are represented by dark (light) shading. Open circles indicate locally statistically significant grids at
the 95% confidence level for the OLR anomalies. Solid (dashed) contours represent positive(negative) c values. Zero contours are omitted. The contour interval is 3.0 3 105 m2 s21. Only 95%(90%) statistically significant vectors are indicated as thick (thin) solid arrows. The thick dashed line
indicates the topographic contour line for 1500 m. (e)–(h) As in (a)–(d), respectively, but for 200
hPa. The contour interval for the streamfunction anomalies is 1.0 3 106 m2 s21.

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southwest–northeast tilt. In phase 5 (Fig. 6a), an active
convection anomaly appears northeast of the Tibetan
Plateau. A closed cyclonic circulation anomaly is centered on 458N, 1258E, and an anticyclonic circulation
anomaly extends from southern China to the northern Philippines. Northwesterly wind anomalies in thesouthwestern flank of the cyclonic anomaly and preexisting southwesterly wind anomalies in the northwesternpart of the anticyclonic circulation anomalies converge
to the northeast of the Tibetan Plateau, initiating convections there. A similar initiation process was proposed by Takahashi (2003), who examined the southward migration of the mei-yu front in 1992. By phase 6(Fig. 6b), the active convection signal has moved
southward and developed over the YHRBs. The main
cyclonic circulation anomaly remains around 458N,
1258E but spreads southward locally into the key region.
Associated with the southwestward extension of thecyclonic circulation anomaly, the anticyclonic circulation anomaly over southern China in the previous phasehas moved southward and developed over the northern
Philippines. This circulation structure induces enhanced
southwesterly wind anomalies over the southern part of
the key region, enhancing the transport of abundantmoisture into the active convection region. In phase 7(Fig. 6c), YHRBs convections reach a maximum andtilt zonally to the southwest–northeast, accompanying strong southwesterly wind anomalies to the south.
The anticyclonic circulation anomalies centered on thePhilippines during phases 6–7 are associated with suppressed convection anomalies over the Philippine Sea
and the Indochina Peninsula. Circulation anomalies
over the YHRBs have an opposite sign from those over
the South China Sea and the Philippine Sea, as noted byChen et al. (2000). From phases 7 to 8, the OLR
anomaly pattern exhibits a zonally tilted tripole structure with positive anomalies in areas northeast of the
Tibetan Plateau and from the Indochina Peninsula tothe south of Japan, with negative anomalies over the
YHRBs. Circulation structures change dramatically inthe midlatitudes in phase 8 (Fig. 6d), and c700 anomalies
acquire the sign opposite that in phase 5. An anticyclonic anomaly develops rapidly, centered on 508N,
1208E, with easterly wind anomalies to its south.
Northeasterly wind anomalies become dominant overthe northern part of the YHRBs because of a closed
cyclonic circulation anomaly over the YHRBs. The
circulation structure provides a low-level divergencearound 388N, 1108E, northeast of the Tibetan Plateau,
where convection is suppressed. Subsequently, thesuppressed convection moves southward to South China
from phases 8 to 4, accompanied by a low-level anticyclonic circulation anomaly.

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The right-hand panels (Figs. 6e–h) show the same
composite, but for 200-hPa wind vector and streamfunction (c200) anomalies. Upper-level circulation anomalies associated with submonthly convection signals overthe YHRBs spread to the north of 208N. In addition, as
anticipated from Fig. 5a, circulation anomalies of large
amplitude appear from the northern to eastern peripheries of the Tibetan Plateau. The composite sequenceshows that a cyclonic circulation anomaly north of the
key region has quasi-stationary behavior from phases
5 to 8, developing and moving very slowly southsoutheastward. In phase 5 (Fig. 6e), a zonally banded
cyclonic circulation anomaly extends along 488N from908 to 1308E, whereas an anticyclonic circulation anomaly is observable to its south, east of the Tibetan Plateau. The vertical structure of the circulation anomalies
is characterized by nearly barotropic structure to thenorth of 408N. The simultaneous enhancement of cyclonic circulation in the low- and upper-level troposphere north of the key region triggers subsequent active convection over the key region. In phase 6 (Fig. 6f),
a cyclonic circulation anomaly develops rapidly around
558N, 658E. The development affects the behavior of
circulation anomalies on the downstream side in thesubsequent phases. From phases 6 to 8, an anticyclonic
circulation anomaly develops from the southern toeastern sides of the upstream cyclonic circulation anomaly. Associated with this, the cyclonic circulation anomaly northeast of the Tibetan Plateau seen in phase 6
moves slightly southeastward and gains strength. Thus,
circulation anomalies appear to grow downstream, implying Rossby wave energy propagation to the east. In
phase 7 (Fig. 6g), the anomaly structure around the
Tibetan Plateau is quite similar to the correlation pattern (Fig. 5b), but with opposite sign. Active convection
anomalies over the YHRBs lie along the southeastern
flank of the upper-level anomalous cyclone, where
southwesterly winds dominate. The anomaly structure isfairly similar to that shown by Mao and Wu (2006, theirFig. 6h), although the dominant periods of the ISOs are
not similar. By phase 2, the circulation anomaly over thewestern Tibetan Plateau acquires a sign opposite to thatin phases 7 and 8 (not shown). A cyclonic circulation
anomaly develops rapidly near 408N, 758E, with thedownstream development of anticyclonic circulation
centered on 408N, 1108E.
Figure 7 illustrates the composites of total OLR andcirculation fields of inactive and active extremes of
YHRB convection. In phase 3, suppressed convection
extends broadly from northern to southern China,
where a low-level anticyclone dominates (Fig. 7a). This
anticyclone is not part of the subtropical anticyclonethat brings fine weather after the mei-yu period. ActiveJULY 2009

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FIG. 7. Composites of (a) total OLR (shading), 700-hPa streamfunction (c), and wind vectors at phase 3 in the NET
pattern. The contour interval for c is 2 3 106 m2 s21. The thick dashed line indicates the topographic contour line for1500 m. (b) As in (a), but for the total 200-hPa streamfunction (c) and wind vectors. Shading denotes the total 200-hPa
divergence (1026 s21). The contour for (c) is 6 3 106 m2 s21. (c),(d) As in (a) and (b), respectively, but for phase 7.

convections (frontal zone) appear from the Philippine
Sea to Japan along the western to northern peripheriesof the subtropical anticyclone. Thus, the low-level
moisture inflow into the YHRBs from the south is significantly reduced in the total field. The upper-leveltroposphere shows a north–south wavy flow in the
midlatitude westerlies, including the Asian jet (Fig. 7b).
An enhanced anticyclonic ridge extends from 308N,
1008E to 608N, 1108E and a deep trough can be seen to
the east of the ridge, east of China. Convergent fields
dominate over the YHRBs, reflecting subsidence thatsuppresses convection behind the deep trough. In phase 7,
zonally elongated active convection (i.e., the mei-yu/
baiu frontal zone) extends from the eastern Tibetan
Plateau to Japan (Fig. 7c). Strong horizontal wind shear
is observed over the key region. A subtropical anticyclone extends into the northern South China Sea, inducing strong southwesterly flow in the active convection band. This circulation structure causes convergenceand maintains active convection over the YHRBs. In
the upper-level troposphere (Fig. 7d), a shallow andwide trough appears from the eastern Tibetan Plateau
to the east of Japan, with upper-level divergence to its
southern flank, where convection is active.
Figure 8 shows the total convection and circulation
series during 14–18 June 1989. Our aim is to understandhow the midlatitude waves affect real weather systems
on submonthly time scales over the YHRBs. In 1989,
the NET pattern is highly dominant (Fig. 5a) and three
active convection periods appear during the mei-yu
period (Table 1). Each active period is associated with asimilar atmospheric circulation field (not shown). Therefore, the convection and circulation series during 14–18
June 1989 can be regarded as a typical example of YHRB
convection that is enhanced by the NET pattern. In
Fig. 6, 14 June and 17–18 June correspond to phases
5 and 7, respectively. A band of convection moves
southward from northeast of the Tibetan Plateau tosouthern China during this period (Figs. 8a–e). The southward movement of banded convection corresponds tothat of a surface front over China, which is stationary,
except on 15 June (Figs. 8f–j). Low-level horizontal

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FIG. 8. (a)–(e) Spatiotemporal evolution of the 200-hPa total streamfunction (contours) and OLR (shadings)
during 14–18 Jun 1989. The contour interval for streamfunction is 6.0 3 106 m2 s21. OLR values less than 230 W m22
are represented by shading. The contour interval for OLR (white contours) is 30 W m22. The thick solid line indicates
the topographic contour line for 1500 m. (f)–(j) As in (a)–(e), but for 700 hPa, with the plots magnified only aroundthe key region. The 700-hPa wind vector is also shown, and the position of a surface front is drawn from the surface
weather chart of 1200 UTC for each day. The contour interval for streamfunction is 1.5 3 106 m2 s21. The contour line
for OLR is not drawn.

convergence (or wind shear) dominates along thebanded active convection zone. In the upper-level troposphere, an enhanced ridge appears over the westerly
jet around 408N, 508E on 14 June. At this time, aTibetan high also develops to the west of the Tibetan
Plateau (Fig. 8a). From 908 to 1408E, the westerly jetshows a nearly zonal structure. There is a trough around
508N, 1108E, at a similar location to that in the lower
troposphere, demonstrating a barotropic structure. The

upper-level trough is located around 1008E on 8 June(not shown) and moves eastward very slowly. By
18 June, the westerly jet changes into the wavy state;
there is clear downstream development of the wave
around 408N from 408 to 1608E, without marked phasepropagation. Thus, the wave is quasi-stationary andhas a zonal wavenumber of ;6. Associated with thedownstream development, the upper-level trough develops around 458N, 1308E on 15 June and deepens

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FIG. 9. As in Fig. 6, but for the AT pattern.

southward by 18 June. Thus, the southward amplification of the barotropic trough, due to the downstream
development of the Rossby waves, causes the southward movement of the mei-yu front around the YHRBs
in the NET pattern.

d. AT pattern
The development of OLR and atmospheric circulationin the AT pattern (Fig. 9) differs greatly from that in
the NET pattern. From phases 5 to 7, eastward-moving

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active OLR anomalies from the Tibetan Plateau to the
YHRBs are evident (Figs. 9a–d, dark shading) at amean speed of 3.7 m s21 at 308N. In phase 5 (Fig. 9a),an anticyclonic circulation anomaly is located over the
YHRBs, with suppressed convection. In phase 6 (Fig. 9b),when the active convection anomalies reach the eastern
flank of the Tibetan Plateau, a cyclonic anomaly develops there at 700 hPa, accompanied by significant
southerly wind anomalies to its east. In phase 7 (Fig. 9c),a closed cyclonic anomaly is established over the key
region and has the most active convection. By phase 8(Fig. 9d), active convections show a northeastward tilt
with a banded structure from the Indochina Peninsula
to the YHRBs. From phase 7, suppressed convection
anomalies appear over the western Tibetan Plateau andmove eastward in the subsequent phases.
In the upper-level troposphere (right panels), a wave
train extends from west of the Tibetan Plateau to Japan,
crossing the plateau. In phase 5 (Fig. 9e), cyclonic circulation anomalies develop over the western Tibetan
Plateau, centered on 388N, 758E. At 500 hPa, whichcorresponds to low-level atmosphere over the Tibetan
Plateau, a center of cyclonic circulation is formed over
the western Tibetan Plateau, similar to the location ofsuch circulation at 200 hPa, but extending farther eastward to 1008E (not shown). The development of thetrough induces a southerly flow of moist air toward theplateau, enhancing convection over it (e.g., FY04). By
phase 6, the overall wave pattern moves slightly eastward at a mean eastward phase speed of 3.5 m s21 at
358N. Wave anomalies also grow downstream. Associated with this, the cyclonic circulation over the western
Tibetan Plateau decays slightly, and anticyclonic circulation is strengthened over the eastern periphery of
the plateau. In phase 7 (Fig. 9g), the anomalous anticyclone still remains over the YHRBs. Comparison of
the anomaly signs between the low-level (Fig. 9c) andupper-level troposphere reveals that they are completely out of phase over the YHRBs. Active moist
convection over the YHRBs induces upper-level divergence. To compensate, low-level convergence is also
required. The mass compensation could be primarily
responsible for the enhancement of low-level cyclonic
circulation in the eastern part of the Tibetan Plateau,although the upper-level anticyclonic circulation above
the cyclonic circulation is mainly due to the eastward
movement of the upper-level wave train. The anomalous cyclone over the western Tibetan Plateau decays
significantly, whereas an anticyclonic circulation anomaly develops rapidly over the western fringe of the
Tibetan Plateau. By phase 8 (Fig. 9h), the signs of circulation anomalies near the Tibetan Plateau are completely reversed compared to those in phase 5 (Fig. 9e).

VOLUME 137

The anomalous anticyclone is further strengthened
over the western Tibetan Plateau, and an anomalous cyclone is established rapidly at the downstream side of
the anomalous anticyclone. Southwest–northeast-banded
convection lies to the east of the anomalous trough.
In the total field, in phase 3, convection is suppressedover the YHRBs, even though low-level southwesterlies
flow into the southern regions of the basins (Fig. 10a).
The mei-yu–baiu front covers an area from South China
to Japan, with the most active region over Japan. Thus,
the suppressed convection is not influenced by northwestern intrusion of the subtropical high, as in the NET
pattern. In the upper-level circulation (Fig. 10b), largescale convergence behind an upper-level trough is evident over a large area of the YHRBs, suggesting thatsubsidence suppresses convection. In phase 7 (Fig. 10c),the distribution of low-level southwesterlies flowing
into the south of the key region is nearly the same as inphase 3, but the wind speed is greater. Cyclonic vorticity
is dominant over the YHRBs, reflecting low-level convergence. The difference in the low-level flow patternaround the South China Sea between phases 3 and 7 issmall compared to that in the NET pattern (Fig. 7). In
the upper-level troposphere, anticyclonic circulation
dominates, with divergence over the YHRBs (Fig. 10d).
Figure 11 shows a series of a typical AT pattern observed during 18–22 June 1993. The year 1993 is characterized by the highest variance of submonthly-scalevariability in convection over the YHRBs (Fig. 2) andthe predominance of the AT pattern (Fig. 5a). The three
active periods of convection during the 1993 mei-yu
period (Fig. 3c and Table 1) were induced by atmospheric circulations with a similar feature in each case.
Therefore, the period 18–22 June was chosen as a typical example showing that the YHRBs convection isenhanced by the AT pattern. Maximum convection(i.e., phase 7) is found on 21 June over the YHRBs. On
18 June there is a banded active convection zone oversouth China (Fig. 11a), associated with a surface stationary front (Fig. 11f). There is also active convection
over the Tibetan Plateau. From 19 to 20 June, the active
convection over the plateau moves eastward to the
YHRBs (Figs. 11b,c), where a surface low is generated
at 308N, 1158E along the stationary front (Figs. 11g,h).
On 21 June, the surface low begins to move eastward
with active convection, gaining in strength. The surface
front over the YHRBs becomes cold on 21 June. It is
located along a low-level convergence zone ahead of aclosed cyclonic circulation at 700 hPa. The upper-level
circulation contains a wavy jet at approximately 408N on
18 June (Fig. 11a). The structure starts developing on
16 June (not shown). Ridges and troughs on the jet remain at almost the same position from 17 to 19 June,JULY 2009

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FIG. 10. As in Fig. 7, but for the AT pattern.

showing a characteristic of Rossby waves with zonal
wavenumbers approximately 7–8. The vertical structureof the wave is barotropic to the north of 358N by
19 June, but a vertical phase reversal becomes prominent over the YHRBs after 20 June, when the convection becomes active there. After 19 June, the overall
wave structure at 408N moves eastward at approximately 5.0 m s21. Simultaneously, a ridge appears and isstrengthened over the YHRBs, accompanied by active
convection. The wavy jet breaks into the relatively zonal
jet structure on 21 June. The generation of a low-levelvortex (or low) on the eastern flank of the Tibetan
Plateau, probably due to the eastward propagation of
Rossby waves with active convection over the plateau, isassociated with the AT pattern.

5. Discussion
a. Effects of the background time-mean flow and
the Tibetan Plateau on the two contrasting
propagation patterns
The NET and AT propagation patterns are likely
affected by the contrasting structures of the time-mean
flow in the upper-level troposphere for each pattern.

Figures 12a,b plot the time-averaged streamlines for the
NET and AT patterns, respectively. The 5-yr averages
of a 25-day low-pass-filtered u and y during the mei-yu
period were used as the background flow for each pattern. A low-pass filter was applied to reduce the effects
of high-frequency transients, including submonthly timescales, on the structure of the basic flow. The basic flows
for the NET and AT patterns show statistically significant differences in the meridional wind around the Tibetan Plateau. In the NET pattern, which was associated with enhancement of the Tibetan high around thesouthern Tibetan Plateau, the axis of the westerly jetmeanders northward in an arc from 808 to 1208E to thenorth of the plateau, showing a wavy-jet structure. The
wavy background flow likely supports wave propagation
along the northern flank of the Tibetan Plateau in NET
pattern years. In contrast, the background westerly flowin the AT pattern shows a more zonal structure aroundthe Tibetan Plateau, associated with the enhancementof the Tibetan high to the west of the plateau (Fig. 12b).
The zonal jet provides favorable conditions for waves topropagate downstream across the Tibetan Plateau. The
zonal wind speed is significantly larger than that in the
NET pattern in upstream areas of the plateau. This

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FIG. 11. As in Fig. 8, but for 18–22 Jun 1993.

might contribute to eastward movement of wave trainsfrom the west of the plateau in the AT pattern. Thus,
the contrasting structures of the background basic
flow could cause the different propagation patterns of
Rossby waves along the Asian jet around the plateau. In
addition to the difference in the background basic flow
around the plateau, the topography of the plateau may
influence the alternation of the two wave propagation
patterns. Topography might affect the northward arching time-mean flow around the plateau, creating amplification northward/northeastward through the dynamical action of the plateau. Further study is required
to understand why the appearance of these wave trains
is so closely correlated with the geography.

b. Characteristics of submonthly-scale midlatitude
waves in the NET and AT patterns
As mentioned in the previous section, the midlatitude
waves that affect submonthly convective variation overthe YHRBs have Rossby wave characteristics. They arenearly barotropic to the north of the Tibetan Plateau inboth patterns, although the vertical phase reversal becomes prominent to the east of the plateau in the AT
pattern as convection intensifies over the YHRBs. The
waves are also barotropic from 308 to 1508E to the northof the plateau (not shown). The waves in the NET
pattern have an approximate wavenumber of 6, with an
eastward phase speed of 0–2.5 m s21 along 408N (Figs.

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FIG. 12. Charts of the 200-hPa streamline constructed with 25-day low-pass-filtered windduring the mei-yu period (a) for the NET pattern and (b) for the AT pattern. Shadings denoteareas more than 10 m s21 in the zonal wind component. The thick dashed line denotes the 1500-m
topographic contour line. Closed (open) circles denote grids of statistically significant negative(positive) difference at the 95% significance level in the low-pass-filtered meridional windcomponent between the NET and AT patterns.

6e–h and 8a–e). The waves in the AT pattern havewavenumbers of 7–8, with phase speeds of 3.5–5 m s21(Figs. 9e–h and 11a–e). The different phase speeds areprobably the reason for the different directions of convection motion over the YHRBs, as well as for thedifference in the mean basic state of the Asian jet. Herewe qualitatively compare the theoretical phase velocity
of a barotropic Rossby wave (e.g., Holton 1992) withobserved results. The zonal phase speed relative to thebackground zonal wind speed is
c

u5

b
2

k 1 l2,(1)

where c is the zonal phase speed; u is the backgroundzonal wind speed; k and l are the zonal and meridionalwavenumbers, respectively; and b is the meridional
gradient of the Coriolis parameter. We can assume thatthe zonal and meridional wavenumbers are of similar
scale (k ’ l) for midlatitude synoptic-scale disturbances(Holton 1992). The value of b is 1.75 3 10211 m21 s21 at
408N. In the NET pattern, the zonal wavenumber is 6.
Hence, the Rossby wave speed relative to the background flow calculated from (1) is 25.6 m s21, whereasthat in the AT pattern with wavenumber 7 is 24.1 m s21.
Here, the background zonal wind for the barotropic
Rossby waves is defined as the vertical mean 25-day
low-pass-filtered zonal wind (from the lowest standard
pressure level above the ground to 100 hPa), averaged
over 308–508N, 808–1208E, where the Rossby waves areobserved. The background mean zonal flows are then
7.8 and 8.5 m s21 in the NET and AT patterns, respectively. Hence, the phase speeds of the waves relative tothe ground are 2.2 and 4.4 m s21 in the NET and AT
patterns, respectively. Although these values are roughestimates only, they agree well with the observed eastward phase speeds of the waves in the NET and AT
patterns, respectively.
During the mei-yu period, the dominant period ofISOs in convection exhibits large interannual variation

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over the YHRBs (see Fig. 2). This interannual variation
likely depends on the upper-level waves around theYHRBs. It would be interesting in future research to
examine the processes determining the dominant periodof ISOs in midlatitude waves along the Asian jet during
the mei-yu period.

c. Effects of the Tibetan Plateau on submonthly
convective variability over the YHRBs
This study reveals that upper-level wave activities
over the midlatitude westerly zone around the Tibetan
Plateau have a significant impact on convective fluctuations at submonthly time scales over the YHRBs (Figs.
6–12). The Tibetan Plateau also plays an active role inthe submonthly convective variability over the YHRBs.
In the NET pattern, the topography of the plateau isimportant for convection. Low-level convergence between northerlies and southerlies is limited to the eastof the plateau due to the Himalaya–Tibet massif. In
the NET pattern, because the atmospheric circulation
changes drastically from the north of the YHRBs, thelow-level convergence starts in regions to the northeastof the plateau. Thus, the change in atmospheric circulation due to the midlatitude waves and the plateau
topography creates favorable conditions for low-level
convergence to the east of the Tibetan Plateau. In the
AT pattern, both the plateau topography and the eastward movement of active convection and associatedlow-level atmospheric circulation over the plateau areimportant, before convection is enhanced over the
YHRBs (e.g., Figs. 9a and 11a,b). Although we did not
examine the detailed enhancement process of the lowlevel vortex to the east of the plateau, some previous
studies presented plausible hypotheses. Yasunari and
Miwa (2006) showed that, when the active convection
and associated convergence zone reaches over theeastern edge of the plateau, cyclogenesis is induced
by vertical stretching of the vortex due to downslope
movement on the eastern edge of the plateau and tolatent heating due to convection. The submonthly-scale
midlatitude waves do affect the location of the active
convection and associated convergence zone over the
Tibetan Plateau (e.g., Figs. 9e–h and 11a–e). The development of the upper-level trough over the western
Tibetan Plateau (Figs. 9e and 11a,b) causes the enhancement of convection from the central to easternparts of the plateau because low-level humid southerlies toward the plateau are strengthened ahead of
the trough (cf. FY04). Thus, the slow eastward movement of the midlatitude waves and the enhancement of
convections over the Tibetan Plateau are probably essential for the genesis of low-level vortices in the AT
pattern.

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6. Summary
We examined convective variability at submonthly
time scales (7–25 days) over the YHRBs and the associated large-scale atmospheric circulation during themei-yu season (10 June–10 July), using OLR and NCEP–
NCAR reanalysis data for the period from 1979 to 2004.
The 7–25-day filtered OLR anomalies averaged over the
key region (258–358N, 1108–1208E) were used as an index
to study submonthly variation over the YHRBs. The
fluctuation of OLR anomalies was closely associatedwith the activities of the quasi-stationary (mei-yu) front
and embedded mesoscale vortex, or low. We found twocontrasting relationships between convection in the
YHRBs and the midlatitude waves along the Asian jet.
Results are summarized as follows:
1) The variance in submonthly-scale convective fluctuation over the YHRBs shows large interannual variation. Submonthly-scale convective variability is notable during the mei-yu season in 1981, 1985, 1986, 1989,
1992, 1993, 1995, 2000, 2001, 2002, 2003, and 2004.
2) Correlations between the convection index and upper-level circulation during the mei-yu season reveals two contrasting patterns: one with positive(negative) correlations to the north (south) of the
key region (e.g., 1981, 1985, 1989, 1992, and 2004),
and the other with negative correlations over theentire key region (e.g., 1986, 1993, 1995, 2001, and2003). In the former pattern, high correlations extend along the northern to eastern peripheries of the
Tibetan Plateau with an arclike structure (defined as‘‘NET’’). In the latter pattern, high correlations appear across the Tibetan Plateau, with a wave train–
like structure (defined as ‘‘AT’’).
3) Composite analyses indicate quite different spatiotemporal structures in the NET and AT patterns,
even though YHRBs convection fluctuations are thesame on submonthly time scales. Composites of the
NET pattern show that convection anomalies migratesouthward slowly from the northeastern peripheryof the Tibetan Plateau to South China and have azonally oriented southwest–northeast tilt, which corresponds to the southward migration of a surface
quasi-stationary front (mei-yu front). The initiationand development of the active (inactive) convection
anomalies is induced by the development of cyclonic(anticyclonic) circulation anomalies to the north of
the key region in both the upper and lower troposphere. The southward migration of the cyclonic(anticyclonic) circulation anomalies from the northcauses active (inactive) convection anomalies.
4) In contrast, the composite sequence for the AT patternshows that convection anomalies migrate eastward

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FUJINAMI AND YASUNARI

from the western Tibetan Plateau to the YHRBs.
When the active (inactive) convection anomalies
reach the eastern flank of the Tibetan Plateau, a cyclonic (anticyclonic) anomaly rapidly develops thereat 700 hPa, and upper-level circulation shows an anticyclonic (cyclonic) anomaly over the key region. The
atmospheric structure between the Tibetan Plateau
and the YHRBs in the AT pattern provides favorable
conditions for the genesis of a low-level mesoscalevortex or low at the eastern edge of the plateau. The
eastward migration of convection anomalies is due tothat of upper-level circulation anomalies along the
Asian jet.
5) In the total fields, during the active extreme of
YHRB convection in both the NET and AT patterns, a strong low-level southerly-southwesterly
wind carries abundant moisture into the YHRBs andits convergence yields active convection over the
YHRBs. Although upper-level circulation featurescontrast in the NET and AT patterns, upper-level
divergence dominates over the active convection
regions in both patterns. In the inactive phase, tropospheric subsidence associated with upper-level
convergence behind an upper-level trough is a major
suppressor of YHRB convection in both patterns. In
the lower troposphere, southerly-southwesterly flow
is significantly weakened in the NET pattern, whereas
in the AT pattern, southerly-southwesterly flow dominates over the southern part of the YHRBs.
6) In both the patterns, the midlatitude wave trains playa crucial role in regulating the YHRB convection on
submonthly time scales. The waves are Rossby-like,with the NET pattern having an approximate wavenumber of 6 and 0–2.5 m s21 eastward phase speed
relative to the earth’s surface and the AT patternhaving a wavenumber of 7–8 and eastward phasespeed of 3.5 to 5.0 m s21. In addition, the mean basic
state of the Asian jet is significantly different aroundthe Tibetan Plateau between the two patterns. The
Asian jet has wavy characteristics in the NET pattern, with a northward arclike structure aroundthe plateau. This structure may help develop largeamplitude waves to the north-northeast of the plateau. The zonal jet structure is more dominant in the
AT pattern, creating a condition in which midlatitude waves pass eastward over the plateau. Thus,
these differences probably lead to the differences inconvection anomaly behavior over the YHRBs between the two patterns.
Acknowledgments. We thank Prof. K. Nakamura andProf. H. Uyeda of Nagoya University and other members of the Lower Atmosphere and Precipitation Study(LAPS)/Core Research for Evolutional Science andTechnology (CREST) project for their many constructive comments and suggestions. Thanks are also extended to Dr. K. Tuboki of Nagoya University, Dr.
T. Terao of Kagawa University, and Dr. Y. Fukutomiof the Frontier Research Center for Global Change(FRCGC) for their many insightful comments. This
study was partially supported by the CREST program of
the Japan Science and Technology Agency (JST) and a
Grant-in-Aid for Young Scientists (B-20740268) fromthe Japanese Ministry of Education, Culture, Sports,
Science, and Technology.
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