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Role of Orography, Diurnal Cycle, and Intraseasonal Oscillation in Summer Monsoon
Rainfall over the Western Ghats and Myanmar CoastSHOICHI SHIGE, YUKI NAKANO, AND MUNEHISA K. YAMAMOTO
Graduate School of Science, Kyoto University, Kyoto, Japan(Manuscript received 1 December 2016, in final form 3 August 2017)
ABSTRACT
Rainfall over the coastal regions of western India [Western Ghats (WG)] and Myanmar [Arakan Yoma
(AY)], two regions experiencing the heaviest rainfall during the Asian summer monsoon, is examinedusing a Tropical Rainfall Measuring Mission (TRMM) Precipitation Radar (PR) dataset spanning 16 years.
Rainfall maxima are identified on the upslope of the WG and the coastline of AY, in contrast to theoffshore locations observed in previous studies. Continuous rain with slight nocturnal and afternoon–
evening maxima occurs over the upslope of the WG, while an afternoon peak over the upslope and amorning peak just off the coast are found in AY, resulting in different locations of the rainfall maxima for
the WG (upslope) and AY (coastline). Large rainfall amounts with small diurnal amplitudes are observed
over the WG and AY under strong environmental flow perpendicular to the coastal mountains, andvice versa. Composite analysis of the boreal summer intraseasonal oscillation (BSISO) shows that the rainanomaly over the WG slopes lags behind the northward-propagating major rainband. The cyclonic systems
associated with the BSISO introduces a southwest wind anomaly behind the major rainband, enhancing theorographic rainfall over the WG, and resulting in the phase lag. This lag is not observed in the AY region
where more closed cyclonic circulations occur. Diurnal variations in rainfall over the WG regions are
smallest during the strongest BSISO rainfall anomaly phase.

1. Introduction
Rainfall during the Asian summer monsoon is concentrated along the western coastal regions of India[Western Ghats (WG)] and Myanmar [Arakan Yoma
(AY)] (e.g., Xie et al. 2006; Hoyos and Webster 2007;
Romatschke and Houze 2011; Kumar et al. 2014).
Rainfall maxima were previously observed offshorein satellite observations (Krishnamurti et al. 1983;
Grossman and Garcia 1990). Motivated by these observations, several numerical studies (Smith and Lin
1983; Grossman and Durran 1984; Ogura and Yoshizaki1988) investigated the rainfall maxima upstream of the
WG. More recently, Hoyos and Webster (2007)
showed localized rainfall maxima to the west of the
WG using data from the Global Precipitation Climatology Project (GPCP) (see their Fig. 2a). With
respect to the rainfall near AY, Xie et al. (2006)
showed that the rainfall maxima are located approximately 50 km offshore of Myanmar’s west coast (seeCorresponding author: Shoichi Shige, shige@kugi.kyoto-u.ac.jp

their Fig. 4), 1 using the Tropical Rainfall Measuring
Mission (TRMM) Precipitation Radar (PR) data.
While there have been efforts to explain near-coastal
rainfall maxima through numerical modeling of the flowinteracting with coastal topography, Johnson (2011)
suggested a full explanation requires consideration of thediurnal cycle. Over the Indonesian Maritime Continent(IMC), the diurnal cycle is dominant and diurnal-driven
offshore migrating systems contribute to rainfall maxima
offshore (Mori et al. 2004; Wu et al. 2009; Ichikawa andYasunari 2006). Motivated by these observations of thediurnal cycle of rainfall concentrations near tropicalcoastlines, Ogino et al. (2016) modeled the annual mean
precipitation in the tropics as a function of coastal distance and noted, ‘‘the principal and fundamental cause of
producing the coastal precipitation must be the diurnalland–sea circulation driven by the heat contrast between
them’’ (p. 1235). Romatschke and Houze (2011), however, showed a weak diurnal cycle over the WG and AY.
1This result was due to human errors (S.-P. Xie 2015, personalcommunication), as shown later.

DOI: 10.1175/JCLI-D-16-0858.1
Ó 2017 American Meteorological Society. For information regarding reuse of this content and general copyright information, consult the AMS Copyright
Policy (www.ametsoc.org/PUBSReuseLicenses).

9366

JOURNAL OF CLIMATEThe boreal summer monsoon intraseasonal oscillation(BSISO) is the dominant large-scale feature of themonsoon (see review in Webster et al. 1998), which ingeneral develops over the equatorial Indian Ocean andpropagates northward (Yasunari 1979). Hoyos andWebster (2007) showed that rainfall in AY is produced
not only by the in situ and orographic-driven convection,
but also by northward-propagating convection within
BSISOs. Transient enhanced cyclonic circulations in thenorthern region of the Bay of Bengal (referred to as Bay
of Bengal depressions) associated with BSISOs tend todrive moist air toward the mountain ranges, resulting in
rainfall. Hoyos and Webster (2007) suggested that asimilar interaction between the BSISO and orography
occurs over the WG region, but they did not examine
this hypothesis in detail.
For over 16 yr, the TRMM PR, the first spaceborne
radar, provided the most reliable precipitation productsin the tropics and subtropics through the well-developed
2A25 retrieval algorithm (Iguchi et al. 2009) (see review
by Houze et al. 2015). The sampling properties of
TRMM PR were poor due to the narrow swath of
215 (245) km before (after) the orbit boost in August
2001, but the 16-yr data accumulation allows us tocreate a high-resolution rainfall climatology. In thispaper, the role of orography, diurnal cycle, and BSISO
in determining the characteristics of summer monsoon
rainfall over the WG and AY is studied, using the highresolution rainfall climatology from the TRMM PR
observations. Section 2 describes the datasets andmethods used in this study. Section 3 revisits the positions of rainfall maxima in the WG and AY regions, andsection 4 examines the diurnal cycle in the WG and AY
regions as a function of environmental wind conditions.
Section 5 describes the interaction between the BSISO
and orography, with the focus on the WG region. Section 6 offers a summary of the findings of this paper.

2. Data and methods
In this study, data from the TRMM PR 2A25 version
7 (V7) (Iguchi et al. 2009) for a 16-yr period (1998–2013) during the monsoon months (June–August) on a
0.058 grid are used. To compare the findings presented
here with the results of Xie et al. (2006, their Fig. 1b),
the 7-yr (1998–2004) averaged June–August climatologies of the surface precipitation from version 6 (V6) of
the TRMM PR 3A25, 0.58-spatial-resolution gridded
monthly composite of instantaneous and footprintscale data (PR 2A25), are also briefly considered. The
European Centre for Medium-Range Weather Forecasts (ECMWF) interim reanalysis (ERA-Interim; Deeet al. 2011) dataset was used to describe the atmospheric

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environment, having a resolution of 0.758 3 0.758 at
6-hourly intervals.
To compare the rainfall maxima determined by the
infrared (IR) techniques used by the earlier satellite
observations, the Geostationary Operational Environmental Satellite (GOES) Precipitation Index (GPI;
Arkin and Meisner 1987) is applied to a single global(608N–608S) IR product (Janowiak et al. 2001) for a
14-yr period (2000–13) during the monsoon months(June–August). The GPI assigns a constant conditionalrain rate of 3 mm h21 to all pixels within a given region
that have a brightness temperature lower than 235 K andassigns a zero rain rate to all other pixels. The TRMM
3B42 V7 data from TRMM Multisatellite Precipitation
Analysis (TMPA; Huffman et al. 2007), the merged
microwave–IR product where the rain gauge analysis isindirectly used for the calibration over land, are also used.
In TMPA, the GPI method has been used for IR estimates.
To examine the amplitude of the diurnal cycle ofrainfall, the normalized relative amplitude (NRA) of
the diurnal cycle for the 16-yr period (1998–2013) isderived from the TRMM PR data. This is defined byaverage evening (1200–2359 LT) minus morning(0000–1159 LT) rainfall, normalized by the climatologicalmean rainfall, following Johnson (2011) and Rauniyar andWalsh (2011).
The bimodal intraseasonal oscillation index developed by Kikuchi et al. (2012) accurately represents
the state of the BSISO and Madden–Julian oscillation(MJO) modes at any particular time of year. The index isbased on the extended empirical orthogonal function
analysis of 25–90-day filtered daily outgoing longwave
radiation data with 2.58 horizontal resolution (Liebmann
and Smith 1996). For this study, the eight phases of the
BSISO mode defined by Kikuchi et al. (2012) are used.
Days on which the amplitude of the BSISO index waslarger than 1 are analyzed, resulting in 65% of all days
during the monsoon months of the 16-yr period.
Previous studies pointed out that two types of intraseasonal oscillations (10–20 and 30–60 days) are importantin the Asian monsoon region (e.g., Annamalai and Slingo2001), but the 10–20-day mode is beyond the scope of thisstudy. We focus on the interaction between the BSISO andorography over the WG region where Hartmann andMichelsen (1989) showed that only the 30–60-day mode isimportant, using a 70-yr record of daily rainfall.

3. Position of rainfall maxima
Figures 1a,b show the 16-yr averaged June–August
climatologies of the surface precipitation at 0.058 resolution estimated from the PR data and 850-hPa windvelocity from the ERA-Interim data. Rainfall on the1 DECEMBER 2017

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FIG. 1. June–August climatologies of (a) surface precipitation (mm month21) at 0.058 resolution using the TRMM PR 2A25 V7 data and(b) 850-hPa wind vectors (m s21) from ERA-Interim with orography (m) for a 16-yr period (1998–2013). The coastal regions of western
India [Western Ghats (WG)] and Myanmar [Arakan Yoma (AY)] are indicated by the black boxes in (a). (c) Cross-shore distributions ofprecipitation (black line; mm month21) and topography (gray bars; m) averaged across the AY region. (d) The 7-yr (1998–2004) averaged
June–August climatologies of surface precipitation at 0.58 resolution from the TRMM PR 2A25 V6 data.

western side of the coastal mountain ranges of western
India (WG) and Myanmar (AY), where southwesterly
flow impinges during the summer monsoon, are the regions experiencing the highest levels of rainfall. Overall,
the patterns of accumulated rain agree with those in Xie
et al. (2006, their Fig. 1b) reasonably well. Figure 1a,
however, shows that rainfall is highest along the mountain slopes over the WG region while Xie et al. (2006)
observed the rainfall maxima offshore of the WG region. Position of rainfall maxima over the AY region is
not clear. The cross-shore distribution of rainfall andorography over the AY region is shown in Fig. 1c, whichcorresponds to Fig. 4 of Xie et al. (2006). In Xie et al.
(2006), the maximum rainfall was shown to be displaced
windward of the maximum orography by as much as
50 km, but the rainfall maxima are also observed on the
coastline in Fig. 1c of the present study. Biasutti et al.
(2012) also showed that the maxima in rainfall frequency are not displaced from the upslope of AY andthe most frequent precipitation is seen on the coast (seetheir Fig. 10).
Xie et al. (2006) showed the 7-yr (1998–2004) averaged June–August climatologies of surface precipitation
at 0.58 resolution, for the PR data. Though not explicitly
noted, the data used by Xie et al. (2006) are likely from
V6 of 3A25, for which data processing started in June
2004. The 7-yr (1998–2004) averaged June–August climatologies of surface precipitation at 0.58 resolution
from V6 are shown in Fig. 1d. In this figure, we made theformat (i.e., domain, color bar, etc.) as consistent as
possible with Fig. 1b of Xie et al. (2006), which showed
the offshore displacement of rainfall maxima west of the
WG and many other locations. The most serious problem in V6 of the PR algorithm was the underestimationof rain over land, in particular, over Africa (Iguchi et al.
2009). Therefore, rainfall over the slopes of the WG
from V6 is smaller than that from V7, but the rainfall
maxima are still observed over the slopes of the WG.

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FIG. 2. June–August climatologies of surface precipitation (mm h21) based on (a) the GPI method for a 14-yr period (2000–13) and(b) TRMM 3B42 data for a 16-yr period (1998–2013). Longitude–height cross sections of the 16-yr June–August unconditional mean
precipitation rate (mm h21) from the TMM PR data and horizontal wind along with topography across the (c) WG and (d) AY regions,which the black lines indicate in (a). Black solid lines in (c) and (d) indicate surface precipitation (mm h21) based on the GPI method.

Some recent studies have shown reasonable rainfall
maximum locations in the WG and AY regions, butthey did not discuss them explicitly. For example,
Romatschke and Houze (2011, p. 24) showed rainfall
maximum locations correctly, but noted, ‘‘The coastal
regions of western India and western Myanmar are theregions of greatest monsoon precipitation, with muchof the rain falling over the ocean upwind of the coastlines (Hoyos and Webster 2007; Xie et al. 2006).’’ Thisidea might come from the influence of earlier numerical studies (Smith and Lin 1983; Grossman and Durran
1984; Ogura and Yoshizaki 1988), which investigatedoffshore maxima in rainfall near the WG. The spatial
shift of the rainfall away from the WG was determinedfrom images of high reflective clouds by the method ofKilonsky and Ramage (1976; see Fig. 2a of Grossman
and Durran 1984) or rain maps produced from infrared
radiometers (IR) and gauges by Krishnamurti et al.
(1983; see Fig. 1 of Ogura and Yoshizaki 1988).
Figure 2a shows the June–August climatology ofsurface precipitation from the GPI (Arkin and Meisner
1987), which is a popular IR technique and has beenused in the TRMM 3B42 data (Fig. 2b) from TMPA

(Huffman et al. 2007). Rainfall maxima offshore withvery weak rainfall over the upslope are found in the WG
and AY regions, as in the previous studies. To validate
the GPI rain maxima, the longitude–height cross sections of the unconditional mean precipitation rate fromthe PR over the WG and AY regions with ERA-Interim
horizontal winds are shown in Figs. 2c,d. The GPI
method determined the spatial shift of the rain maxima
away from the PR rainfall maxima over both regions.
Infrared techniques, such as the GPI method, as well as
visible sensor techniques assume that deeper, colder
clouds are more likely to produce heavy rainfall, leading
to an overestimate of rain area from the cold brightness
temperatures associated with extensive cirrus anvils(Ebert and Manton 1998) and an underestimate of theorographic rainfall associated with relatively low andwarm cloud tops (Arkin et al. 1989). Rickenbach (1999)
showed that in an environment of deep troposphericdirectional wind shear, the cold cloud shield became
spatially decoupled from the source convection oftropical oceanic squall lines, leading to the poor performance of IR techniques. Strong vertical wind shearwith low-level westerlies and upper-level easterlies is

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FIG. 3. NRA of the diurnal cycle of rainfall defined by average June–August evening (1200–
2359 LT) minus morning (0000–1159 LT) rainfall normalized by June–August mean rainfall fora 16-yr period (1998–2013) using the TRMM PR data. Gray areas indicate the occurrence ofraining pixels of ,1%.

found in Figs. 2c,d. Thus, the spatial shift of the GPI rain
maxima away from the PR rainfall maxima is attributed
to westward advection of cirrus anvils from the convection associated with relatively low and warm cloud
tops over the upslope by upper-level easterlies.
The rainfall patterns from the TRMM 3B42 data
(Fig. 2b) agree much more with the TRMM PR data
(Fig. 1a) than those estimated by the GPI method, but
the rainfall maxima occur offshore of the WG region
instead of the upslope. Over coastal mountain ranges of
the Asian monsoon region, heavy orographic rainfall
associated with low precipitation-top heights (Shige andKummerow 2016; Konwar et al. 2014) leads to conspicuous underestimation of rainfall using microwave
radiometer algorithms, which conventionally assume
that heavy rainfall is associated with high precipitationtop heights (Shige et al. 2013, 2015; Taniguchi et al. 2013;
Yamamoto and Shige 2015; Yamamoto et al. 2017).
Over land, the rain gauge analysis is indirectly used tocreate the TRMM 3B42 data, but there still remained
underestimation over the upslope of the WG region by
the microwave radiometer algorithm (Kummerow et al.
2015) used in the TMPA (see Figs. 2.2d,f of Shige et al.
2015). The GPI method used in the TMPA also contributes to the rainfall maxima offshore of the WG region. The GPI method has been also used in the GPCP
data (Adler et al. 2003; Huffman et al. 2001), resulting in
the localized rainfall maxima to the west of the WG
shown by Hoyos and Webster (2007) (see their Fig. 2a).

These data need to be used with care in the Asian summer
monsoon region with strong vertical wind shear and theorographic rainfall associated with low cloud tops.

4. Diurnal cycle
The prevailing view that the coastal precipitation in
the tropics results from daytime heating over land,
which results in land–ocean temperature gradients dueto the smaller heat capacity of land compared to that ofocean, is generally attributed to observational studies of
the IMC (Ogino et al. 2016; Yamanaka 2016). Diurnaldriven offshore migrating systems contribute to rainfall
maxima offshore of the IMC (Mori et al. 2004). In
contrast, we have shown the rainfall maxima occur overthe upslope of the WG and the coastline of the AY regions instead of offshore, in section 3. In this section, we
examine the diurnal cycle of rainfall over the WG andAY regions.
The NRA for June–August over the Asian monsoon
region (Fig. 3) shows the general preference forafternoon–evening rainfall over land and morning
rainfall over the ocean. Late night–morning rainfall overthe southern slopes of the Himalayas is distinct, which isconsistent with Hirose and Nakamura (2005). A strong
diurnal cycle is found in the IMC, which is consistent
with previous studies. In contrast, the upslope of the WG
and AY, the regions experiencing the heaviest rainfall
during the Asian summer monsoon (Fig. 1a), have very

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FIG. 4. Time–zonal sections of the composite diurnal rainfall cycle averaged across the (a) WG and (b) AY regions,which are indicated by the black boxes in Fig. 1a, where a 1–2–1 temporal filter is applied. Topography (km) is indicatedby the gray bars above the plots. Vertical dashed lines indicate the location of the coast. The mean diurnal variation of
rainfall with intervals of error bars indicating the 95% confidence interval for the mean values with the Student’s t test
for the upslope of the (c) WG and (d) AY regions, which are indicated by the arrows in (a) and (b).

small amplitudes, which is consistent with Romatschkeand Houze (2011). It should be noted that Johnson(2011) showed evening rainfall for May–June along
coastlines over the WG and AY regions using TRMM
3B42 data, cautioning using the data in the Asian summer monsoon region. On the other hand, the leeward
sides of the WG and AY have large amplitudes. Weakdiurnal cycles over the windward sides and strong cycles
over the leeward sides are evident for other coastal
mesoscale mountains in the Indochina Peninsula such as
Bilauktaung, Cardamom, and the Annam Range.
Regional variations in the diurnal rainfall cycle observed by the PR averaged over the domains perpendicular to the WG and AY (Fig. 1a) are shown in Fig. 4.
The narrow rainfall area over the upslope of the WG(Fig. 4a) indicates pure orographic rainfall over the WG.
Continuous rain with a slight nocturnal and afternoon–
evening maxima is found over the upslope of the WG(Fig. 4a), resulting in a double-peaked cycle with small

amplitudes over the upslope (Fig. 4c). On the other
hand, an afternoon peak over the upslope and a morning
peak just off the coast are found in AY (Fig. 4b), resulting in an afternoon peak over the upslope (Fig. 4d).
Romatschke and Houze (2011) showed a doublepeaked cycle for strong precipitating medium-sized
systems occurring in the WG and AY regions (seetheir Figs. 13b,d), although their selected regions arewider than those used in this study. They speculated thatthe peak in the afternoon is likely related to systems
over land, while the morning peak is likely associatedwith the behavior of systems over the ocean. This could
account for the afternoon peak over the upslope and amorning peak just off the coast in the AY region, but thispossibility seems unlikely for the WG region because
both peaks are found over the upslope.
As shown in Fig. 3, the diurnal cycles over the WG andAY regions are much weaker than over the IMC.
Nugent et al. (2014) showed ambient upstream wind

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FIG. 5. Histogram showing the frequency (number of days) of the cross-shore component of the horizontal wind(m s21) averaged over the upstream regions of (a) the WG and (b) AY. The upstream regions are indicated by theblack boxes in Fig. 6.

speed control in the transition of tropical orographic
convection over Dominica, a small mountainous islandin the eastern Caribbean Sea, from low-wind thermal tohigh-wind mechanical convection. During the summer
monsoon, southwesterly flows over the WG and AY
regions are much stronger than over the IMC (Fig. 1b).
To investigate the wind speed control of convection type(i.e., diurnal cycle amplitude) over coastal mountain
ranges, we categorized the TRMM PR data according toambient upstream wind speed.
Figure 5 shows the frequency of the cross-shore
component of upstream horizontal wind (m s21) at
850 hPa for the WG and AY regions. Three regimes aredefined as follows: weak (U # U2 s), medium (U2s , U , U1 s), and strong (U $ U1 s) regimes, where
U is the cross-shore component of upstream horizontal
wind, U is the mean of U, and s is the standard deviationof U. The mean wind speed over the WG (9.25 m s21) islarger than that for AY (8.84 m s21), while the standard
deviation for the WG (3.94 m s21) is smaller than that for
AY (4.20 m s21).
The NRA over the coastal areas of the WG and AY
regions for the weak wind regime is much larger than thatfor the strong wind regime (Fig. 6), indicating that thediurnal cycle is a function of ambient upstream wind. The
results are consistent with the fact that the thermallyforced circulation is typically observed under no or weakenvironmental wind, while the mechanically forced circulation dominates under strong environmental wind.
Banta (1990) suggested that stronger winds ventilate thesurface as it tries to heat up, and they produce mechanicalturbulence that also mixes the heat away from the surface,preventing the buildup of a strong temperature difference
that would drive thermally forced circulations.
Regional variations in the diurnal rainfall cycle averaged over domains perpendicular to the WG and AY(Fig. 1a) for weak and strong regimes are shown in Fig. 7
and Fig. 8, respectively. Rainfall for the weak wind regime over the WG and AY regions is much smaller thanthat for the strong wind regime, indicating that thecontribution of thermally driven convection to rainfall ismuch smaller than that of mechanically driven convection. For the weak wind regime, rain maxima over the
upslope in the afternoon and rain maxima over theocean at midnight and in the early morning are seen for
the WG and AY regions (Figs. 7a,b), leading to a onepeaked cycle over the upslope (Figs. 7c,d). The continuous rain with slight nocturnal and afternoon–evening
maxima for the strong wind regime over the upslope of
the WG (Figs. 8a,c) is very similar to that for all regimes
over the WG shown in Figs. 4a,c except for larger rainfall amounts. Again, mechanically driven convection is amajor contributor to rainfall, and therefore rainfall isconcentrated over the upslope of the WG regions. Over9372

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FIG. 6. NRA of the diurnal cycle of rainfall during June–August for the (left) weak and (right) strong cross-shore
component of the horizontal wind regime for (a),(b) the WG and (c),(d) the AY. The upstream regions for crossshore components of the horizontal wind are indicated by the black boxes. Gray areas indicate the occurrence ofraining pixels ,1%.

AY, morning and slight afternoon maxima over thecoast are seen for the strong wind regime (Figs. 8b,d).
According to the classical view of orographically andthermally forced winds in mountains (Lin 2007), during

the day, the mountain serves as an elevated heat source
because of the sensible heat released by mountain surface. In a quiescent atmosphere, this can induce mountain upslope flow, leading to the rain maximum over the1 DECEMBER 2017

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FIG. 7. As in Fig. 4, but for the weak wind regime.

upslope in the afternoon. At night, the opposite occurs:
surface cooling produces downslope flow. Diurnal
components of surface wind (surface divergence) arecalculated as the difference between the 6-houly surface
wind (surface divergence) and the daily mean surface
wind (surface divergence) for the weak and strong windregimes for the WG and AY regions, using the datafrom ERA-Interim for a 16-yr period (1998–2013). At
0000 UTC (0500 LT for the WG and 0600 LT for the AY),the diurnal components of the surface winds over thecoastal mountains of the WG and AY regions flow toward the oceans for weak and strong wind regimes(Fig. 9). It is likely that the diurnal components of thesurface winds in the early morning are downslope flows.
The downslope flows over the WG and AY regions forthe weak wind regime are stronger than those for thestrong wind regime, resulting in the shift of convergenceover the upslope for the strong regime to that just off thecoast for the weak regime. Therefore, the rain maximum
just off the coast at midnight and in the early morning for
the weak wind regime is attributed to the downslope flow.
Over AY, rainfall with a midday maximum is enhanced over the ocean (918–928E) for the strong windregime (Fig. 8b) compare to all regimes shown in Fig. 4b.
Romatschke and Houze (2011) showed that large systems in the Bay of Bengal region formed when cyclonic
circulations (Bay of Bengal depressions) enhanced theanomalous southwesterly flow and have a strong diurnal
cycle with a midday maximum. They speculated that theinitial convection of the mesoscale system is triggeredover the ocean during the late night or early morning,
and the nocturnally formed systems off the coast growuntil about midday, with expanding stratiform rain
areas, as in the case of Borneo (Churchill and Houze
1984; Houze et al. 1981; Williams and Houze 1987)
where widespread rainfall is likely related to the presence of the Borneo vortex (Ichikawa and Yasunari2006). The results for the strong wind regime suggest
in situ orographically driven convection (pure orographic rainfall) over the WG region and orographic
enhancement of preexisting systems over the AY region, reflecting differences in the latent heating profiles(Fig. 3 of Shige and Kummerow 2016) between the WG(bottom heavy) and AY (top heavy) and resulting indifferent locations of the rainfall maxima for the WG(upslope) and AY (coastline).

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FIG. 8. As in Fig. 4, but for the strong wind regime.

5. BSISO
Figure 10 presents the composite eight-phase BSISO
life cycle of the PR rainfall anomalies and ERA-Interim
specific humidity anomalies with 850-hPa wind anomalies.
The positive rainfall anomaly associated with cyclonicsystems forms over equatorial India in phase 1 andmoves northward over the Arabian Sea and the Bay of
Bengal. The typical northward-propagating character ofthe BSISO is consistent with previous studies (e.g.,
Wang et al. 2006; Chattopadhyay et al. 2009). Figure 10
shows the rain anomaly maximum phase band is oriented from northwest to southeast. The northwest–
southeast-tilted structure of the major rainband is also
consistent with previous studies. It is interesting to notethat the rain anomaly over the WG slope is isolated fromthe major rainband associated with the BSISO. In phase
3, there are positive specific humidity anomalies andpositive rainfall anomalies over the surrounding Arabian Sea (Fig. 10k) but negative rainfall anomalies over
the WG slope with easterly anomalies (Fig. 10c). Positive rainfall anomalies with westerly anomalies appear

over the coast of the WG region in phase 4 (Fig. 10d) andthen spread northward in phase 5 (Fig. 10e). The horizontal wind anomaly is a key factor determining the rainfall anomaly associated with BSISO. This feature isconsistent with Hartmann and Michelsen (1989) using therain gauge data, but was not reported in Wang et al. (2006,
see their Fig. 2) and Chattopadhyay et al. (2009, see theirFig. 3), since Wang et al. (2006) used only rainfall over theocean from TRMM Microwave Imager (TMI) and
Chattopadhyay et al. (2009) used TRMM 3G68 with aresolution of 0.58 3 0.58 from the PR with a 1–2–1 spatial
filter both in the zonal and meridional orientations.
Figure 11 clearly shows that the upslope rain lags
behind the offshore rain associated with the BSISO
major rainband in the WG region, in contrast to the
maxima in the rainfall anomaly over the IMC ahead of
the MJO major rainband (Rauniyar and Walsh 2011).
The cyclonic systems (Fig. 11c) associated with the
BSISO introduce southwest winds (Fig. 11a) behind the
major rainband, enhancing the orographic rainfall over
the WG, and resulting in the phase lag of rainfall between the Arabian Sea and the WG. Hartmann and1 DECEMBER 2017

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FIG. 9. Diurnal components of surface wind (m s21) and surface divergence (1025 s21) at 0000 UTC during June–
August for the (left) weak and (right) strong wind regime for (a),(b) the WG and (c),(d) the AY.

Michelsen (1989) showed simultaneous occurrence ofthe rainfall maximum in the WG region with westerly
wind anomaly, although they did not report the phaselag of rainfall between the Arabian Sea and the WG
because they used rain gauge data only over the land.

This is similar to in-phase relationship between the
ground-based radar reflectivity and zonal wind anomalies over the west coast of the Downa range in thewestern part of the Indochina Peninsula (Yokoi and
Satomura 2008). Figure 10g shows that the cyclonic

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FIG. 10. (a)–(h) PR rainfall anomaly and (i)–(p) ERA-Interim specific humidity anomaly with ERA-Interim 850-hPa wind (m s21; vectors)
during the eight BSISO phases. The L indicates the cyclonic systems referred to in the text.

system over the AY introduces a westerly wind anomaly, enhancing the orographic rainfall over the west
coast of the Downa range.
Hoyos and Webster (2007) suggested that the interaction between the BSISO and the orography are similar
over the WG and AY regions, but the phase lags among
upslope rain, offshore rain, and wind are not observed
over the AY region. This is because cyclonic circulations
over the AY region during the active phase (Fig. 10g)
exhibit smaller expansion and more closed structures
than those over the WG region during the active phase(Fig. 10d), leading to much higher relative vorticities
over the AY region (Fig. 11d) than over the WG region(Fig. 11c). Large precipitating systems prevail over the
Bay of Bengal (Hirose and Nakamura 2005; Romatschkeand Houze 2011), resulting in top-heavy heating profiles
produced by deep stratiform ice clouds (Fig. 3 of Shigeand Kummerow 2016), and enhancing the northward
propagation of the BSISO and cyclonic circulations(Chattopadhyay et al. 2009; Choudhury and Krishnan2011). Differences in precipitating system characteristics
and their associated large-scale response might explain

the differences in the interaction between the BSISO andorography between the WG and AY regions.
The averaged upstream winds throughout all phasesof BSISO shown in Fig. 11a are within the medium windregime in Fig. 5, but the weak wind regime tends to occur
in association with phases 1–2, while the strong windregime occurs in association with phases 5–6, which affect the diurnal cycle of rainfall as discussed in section 4.
Scale interactions on diurnal, intraseasonal, and seasonal time scales is important for the mean and variability of the tropical climate system (see a review by
Slingo et al. 2003). A number of studies have investigated the impact of the MJO on the diurnal cycle ofrainfall (e.g., Sui and Lau 1992; Tian et al. 2006), butmuch less attention has been paid to the impact of the
BSISO on the diurnal cycle of rainfall. It is interesting tocompare the impact of the BSISO on the diurnal cycle ofrainfall over the WG with the impact of MJO on thediurnal cycle of rainfall over the IMC.
Again, the phase lag over the WG is clearly visible(Fig. 12a). The modulation of the NRA by the BSISO isevident over the upslope of the WG (Fig. 12b). The

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FIG. 11. Upslope rainfall, offshore rainfall, and upstream wind over (a) the WG and (b) AY regions as a functionof the composite life cycle of the BSISO. A 1–2–1 temporal filter is applied. (c),(d) As in (a),(b) but with vorticity
plotted instead of wind speed. The upslope of the WG and AY regions correspond to that in Fig. 4 and the offshore
is located 28 upstream of the upslope, which is indicated in Fig. 12 for the WG.

smallest NRA over the upslope of the WG coincides
with the maxima of the BSISO rainfall anomaly duringphase 5. Strong winds amplify the mechanically driven
convection resulting in the largest BSISO rainfall
anomaly over the upslope of the WG, but reduce thethermally driven convection resulting in the smallest
NRA over this region. These results are consistent withthose in section 4. The NRA over the upslope of the WG
is greatest during phase 3, which is before the largest
BSISO rainfall anomaly during phase 5, and coincides
with the large-scale active BSISO envelope with thelargest rainfall anomaly over the surrounding Arabian
Sea.
Although there have been disputes on the impact of thediurnal cycle of rainfall by the MJO, Peatman et al. (2014)
showed that the amplitude of the diurnal cycle and thedaily mean rainfall over the IMC are greatest shortly
before the large-scale active MJO envelope arrives, with alead time of one-eighth of an MJO cycle, in agreementwith ground-based observations over the eastern Indian
Ocean off the island of Sumatra (Fujita et al. 2011;
Kamimera et al. 2012). Therefore, the change in the diurnal cycle of rainfall throughout an intraseasonal cycle isdifferent between the IMC where the diurnal cycle ofrainfall is strong and the WG where the diurnal cycle ofrainfall is weak and orographic rainfall is strong.

6. Summary
In this study, a TRMM PR dataset of 16 yr is used to
explore the physical mechanisms determining the spatial
and temporal patterns of summer monsoon rainfall in9378

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FIG. 12. Time–zonal sections of (a) rainfall and (b) NRA of the diurnal cycle of rainfall associated with thecomposite life cycle of the BSISO averaged across the WG. A 1–2–1 temporal filter is applied. Topography (m) is
indicated by the gray bars above the plots. Vertical dashed lines indicate the location of the coast. Rainfall associated with composite life cycle of the BSISO is shown by black contours in (b). The offshore and upslope areindicated by the arrows.

the WG and AY regions. Satellite observations in previous studies showed that the rainfall maxima occur
upstream of the WG and AY (Krishnamurti et al. 1983;
Grossman and Garcia 1990; Xie et al. 2006; Hoyos andWebster 2007). In contrast, the TRMM PR data showthat the rainfall maxima occur over the upslope of the
WG and on the coast of AY. Previous studies using IR
data assumed that heavy rainfall results from deep
convection, leading to an overestimate of the rain areafrom the cold brightness temperatures associated withextensive cirrus anvils and an underestimate of theorographic rainfall associated with relatively low cloud
tops. The spatial shift of the IR-derived rain maxima
away from the PR rainfall maxima is attributed to
westward advection of cirrus anvils from the convection
associated with relatively low and warm cloud tops overthe upslope by upper-level easterlies.
Over the IMC, the diurnal cycle of rainfall is dominantand diurnal-driven migrating systems contribute to therainfall maxima offshore (Mori et al. 2004; Ogino et al.
2016; Yamanaka 2016). In contrast, diurnal variations in
the rainfall are very weak in the WG and AY regions
where low-level monsoon flow is strong. Continuous
rain with a slight nocturnal and afternoon–evening
maxima occurs over the upslope of the WG, while anafternoon peak over the upslope and a morning peakjust off the coast are found in AY, resulting in different

locations of the rainfall maxima for the WG (upslope)and AY (coastline). The rainfall amount and diurnal
cycle amplitude are functions of the environmental flow
perpendicular to coastal mountains. Large rainfall
amounts but small diurnal amplitudes are observed
under strong environmental flow, and vice versa, implying that rainfall is not associated with thermally
driven convection, but rather with mechanically driven
convection. Ventilation of the surface by stronger winds
may prevent the buildup of a strong temperature difference that would drive thermally forced circulations(Banta 1990).
Consistent with past studies, the TRMM PR composites show the northwest–southeast-tilted structureand northward propagation of the major rainband associated with the BSISO. The rain anomaly over the
WG lags behind the major rainband, which was not reported by the previous studies (Wang et al. 2006;
Chattopadhyay et al. 2009). Orographic rainfall over the
WG slope is enhanced with southwesterly wind anomalies of the cyclonic system associated with the BSISO.
Hoyos and Webster (2007) suggested that an interaction
between the BSISO and orography over the WG issimilar to that over AY, but this lag is not observed in
the AY region where more closed cyclonic circulations
and more stratiform rainfall regions prevail compared to
the WG region. The amplitude of the diurnal cycle over

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the upslope of the WG is smallest during the largest
BSISO rainfall anomaly phase over the WG, and it islargest during the large-scale active BSISO phase withthe largest rainfall anomaly over the surrounding Arabian Sea, preceding the largest BSISO rainfall anomaly
over the WG.
Numerical prediction of the Asian summer monsoon,
the world’s most energetic monsoon and thus a major
component of global climate dynamics, is very important for society. The spatial and temporal nature ofsummer monsoon rainfall in the WG and AY region
described in this paper would be a critical test forclimate models.
Acknowledgments. We would like to dedicate thisarticle to the memory of the late Professor TakehikoSatomura, whose research motivated us to start thisstudy. This study is supported by JSPS KAKENHI
Grant 26282109 and the JAXA Precipitation Measuring
Mission (PMM) Project. Profs. S. Yoden, M. Satoh,
Y. N. Takayabu, and M. Hirose are thanked for useful
discussions. The authors thank Dr. S. Yokoi, and two
anonymous reviewers for their constructive commentsthat improved the clarity of the paper.
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