Received: 1 March 2019

Revised: 30 August 2019

Accepted: 27 September 2019

Published on: 2 December 2019

DOI: 10.1002/qj.3672

RESEARCH ARTICLE

Mesoscale precipitation systems and their role in the rapid
development of a monsoon depression over the Bay of Bengal
Hatsuki Fujinami

Hidetaka Hirata

Institute for Space-Earth Environmental
Research, Nagoya University, Nagoya, Japan
Correspondence
H Fujinami, Institute for Space-EarthEnvironmental Research, Nagoya
University, Furo-cho, Chikusa-ku, Nagoya
464-8601, Japan.
Email: hatsuki@nagoya-u.jp
Present address
Hidetaka Hirata, Department ofEnvironmental Systems, Faculty of
Geo-Environmental Sciences, Rissho
University, Kumagaya, Japan

Masaya Kato

Kazuhisa Tsuboki

Abstract
The structure of mesoscale precipitation within monsoon depressions is still not as
well-known as the synoptic-scale composite cloud and precipitation structure. Here,using observational data from multi-satellite sensors and a cloud-resolving regionalmodel, we investigate the three-dimensional structure of mesoscale precipitation systems in the different stages of the life cycle of a monsoon depression. Effects of latentheating from the precipitation systems and the Bay of Bengal (BoB) on the development of the monsoon depression are also evaluated in sensitivity experiments withthe model. A typical monsoon depression developed on 17 August 2016 over the
BoB. In the rapid development phase, satellite observations reveal mesoscale convective systems with deep convective precipitation cells and stratiform precipitation

Funding information
Japan Society for the Promotion of Science,
KAKENHI Grants JP 18KK0098,
KAKENHI Grants JP 26220202, KAKENHI
Grants JP 26400465

near the head of the BoB. Extremely deep and intense convective cells appear along
a ring-like rain band when a closed cyclonic circulation becomes obvious aroundthe northernmost part of the BoB. The deep convection appears frequently, particularly along the western-to-southwestern side of the low, a convergence-prone area
between northerlies with air masses of large convective available potential energy(CAPE) that makes up the western–southwestern part of the closed cyclonic circulation, and strong monsoon westerlies with relatively stable air masses to the south ofthe low. Sensitivity experiments reveal that both cloud/precipitation processes andevaporation from the BoB are essential for the rapid development of the monsoon
depression over the BoB. Evaporation from the BoB adds a large amount of moisture to the atmospheric boundary layer near the low. Southwesterlies and southerlies
on the eastern side of the low draw in warm, humid boundary-layer air to the closed
cyclonic circulation, which maintains a large-CAPE environment in the low andenhances deep convection. A possible positive feedback process including moist convection that leads to the rapid intensification of monsoon depressions over the BoB
is discussed.
KEYWORDS
cloud-resolving regional model, latent heating, mesoscale precipitation system, monsoon depression,
spaceborne precipitation and cloud radar

Q J R Meteorol Soc. 2020;146:267–283.

wileyonlinelibrary.com/journal/qj© 2019 Royal Meteorological Society

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268

1

INTRODUCTION

Synoptic-scale monsoon low-pressure systems (LPSs) contribute a large fraction of total summer monsoon rainfall(Hurley and Boos, 2015; Praveen et al., 2015), especiallyover land in South Asia and Australia (Yoon and Chen, 2005;
Krishnamurthy and Ajayamohan, 2010; Berry et al., 2012).
Extreme precipitation and catastrophic floods in South Asia
are often associated with LPSs (Houze et al., 2011; Webster
et al., 2011; Joseph et al., 2015; Bohlinger et al., 2017). In
addition, LPSs play a role in the transport of water vapour
toward high mountain areas such as the Himalayas and the
Tibetan Plateau, affecting precipitation processes in theseregions (Lang and Barros, 2002; Ueno et al., 2008; Donget al., 2016; 2017; 2018). Thus, LPSs play an important role informing the hydroclimate around their tracks and surrounding
areas.
The study of LPSs around India has the longest history
among all monsoon areas. LPSs around India are categorized into lows, depressions, deep depressions and cyclonicstorms based on their intensity in terms of surface windspeed (Saha et al., 1981; Hurley and Boos, 2015). Monsoon
depressions have surface wind speeds of 8.5–13.5 m⋅s−1 . On
average, about 15–35% of LPSs during June–September are
categorized as monsoon depressions (Hurley and Boos, 2015;
Praveen et al., 2015; Hunt and Fletcher, 2019). Typically,
LPSs including monsoon depressions form over the head of
the Bay of Bengal (BoB) and move northwestward over the
Indian subcontinent, with a life cycle of 3–6 days and a horizontal scale of ∼1,000–2,000 km (Godbole, 1977; Sikka,
1977). The genesis location and track of LPSs were recently
reconfirmed by tracking methods using state-of-the-art atmospheric reanalysis data (Krishnamurthy and Ajayamohan,
2010; Hatsuzuka et al., 2014; Hurley and Boos, 2015; Huntet al., 2016a).
The structure of Indian monsoon depressions is characterized by a vertical tilt to the west and south, a warm-over-coldtemperature anomaly in the centre of the depression, peak horizontal winds at 700–800 hPa with cyclonic winds extending
from the surface to the upper troposphere, and precipitation/cloud maxima in the west–southwest of the depression(Godbole, 1977; Krishnamurthy and Ajayamohan, 2010).
Recently, more detailed three-dimensional composite structures of monsoon depressions have been presented, basedon more than 100 cases, using high-spatial-resolution atmospheric reanalysis and satellite datasets (Hurley and Boos,
2015; Hunt et al., 2016a). These composite structures haverevealed the robust features of monsoon depressions including the known features mentioned above and, in addition, a
deep column of high potential vorticity (PV) extending fromthe surface to the upper troposphere with PV maximum ataround 500 hPa and a weaker maximum around 700–850 hPa.
The low-level horizontal wind maximum and relevant relative

vorticity maximum are associated with the maximum around
700–850 hPa, whereas a warm core around 350 hPa, which ismainly due to latent heating from condensation, causes the
PV maximum around 500 hPa (Hunt et al., 2016a), indicating the dominance of diabatic processes (e.g. latent heating)in the growth of monsoon depressions. Murthy and Boos(2019) found that both deep convective and stratiform heatingproduce PV maxima around 500 hPa in tropical depressions.
Previous studies have proposed moist baroclinic instability
as a possible cause of depression growth (e.g. Moorthi andArakawa, 1985; Krishnamurti et al., 2013). Cohen and Boos
(2016) argued against this theory, based on a comparison of
the vertical structure of monsoon depressions, hurricanes anddiabatic Rossby waves, and indicated that the growth of monsoon depressions is more closely related to tropical depression
spin-up than to moist baroclinic instability. However, a common finding of previous studies is that diabatic heating frommoist convection plays a key role in the development of
monsoon depressions.
Recent studies have reported a favourable environment for
the genesis of LPSs around South Asia (Yanase et al., 2012;
Tsuboi and Takemi, 2014; Ditchek et al., 2016; Hatsuzuka
and Fujinami, 2017; Samanta et al., 2018). Humid and convectively unstable environments with high cyclonic vorticity
and high sea-surface temperature (SST) favour genesis overthe northwestern BoB despite relatively strong vertical wind
shear due to the vertical structure of a large-scale Asian monsoon circulation. The environment is strongly modulated byintraseasonal oscillations such as the boreal summer intraseasonal oscillation (BSISO: e.g. Wang et al., 2006; Kikuchi
et al., 2012) and quasi-biweekly oscillation (QBW: e.g. Chattergee and Goswami, 2004; Fujinami et al., 2014), clustering
the genesis points and tracks of LPSs between active andbreak phases of intraseasonal oscillations (Goswami et al.,
2003; Krishnamurthy and Ajayamohan, 2010; Yanase et al.,2012; Hatsuzuka and Fujinami, 2017).
The composite three-dimensional structure of precipitation has also been investigated using Tropical Rainfall
Measuring Mission (TRMM)–Precipitation Radar (PR) data
and the CloudSat–Cloud Profiling Radar (CPR) (Stanoet al., 2002; Hunt et al., 2016b). These studies showed
composite “synoptic-scale” features of hydrometeors inmonsoon depressions. Intense convection occurs in thesouth–southwest of monsoon depressions, with significant
precipitation for hundreds of kilometres outside the centre andwell above the mid-troposphere.
However, the composite analysis obscures mesoscale features such as rain bands and convective systems in monsoon
depressions. In addition, the mesoscale hydrometeor structure
might change in different stages of the life cycle. Therefore, in this study we investigate mesoscale precipitationand cloud features inside a synoptic-scale monsoon depression using multi-satellite sensors. We focus on the monsoon

FUJINAMI ET AL.

depression that occurred over the head of the BoB on 17
August 2016. It persisted as a monsoon depression until
20 August (e.g. India Meteorological Department cyclone
eAtlas; http://www.rmcchennaieatlas.tn.nic.in). Fortunately,
several overpasses of the Global Precipitation Measurement(GPM) satellite with a spaceborne precipitation radar captured the evolution of the monsoon depression from the development phase to the decaying phase after landfall. CloudSat
with a cloud-profiling radar and a geostationary meteorological satellite also observed the depression during thisperiod. Thus, we can capture the three-dimensional featuresof mesoscale precipitation/cloud systems of the monsoon
depression using multi-satellite sensors.
We also investigate the development of the monsoon
depression and the spatio-temporal structure of mesoscaleprecipitation systems using a cloud-resolving regional modeland observational data. Previous studies have not investigated mesoscale precipitation/convection systems of monsoon depressions using cloud-resolving models. Some previous studies have used the Weather Research and Forecasting(WRF) model to investigate the relationship between land surface feedback and landfalling monsoon depressions (Changet al., 2009; Baisya et al., 2017) and between cloud condensation nuclei and the intensity of monsoon depressions
through sensitivity experiments (Krishnamurti et al., 2013).
As demonstrated below, monsoon depressions do involve
mesoscale convective systems (MCSs) with deep convective cells and precipitating stratiform cloud regions. Thus,
non-hydrostatic and compressible cloud-resolving models are
required to better represent and understand MCSs withinmonsoon depressions and the development of monsoon
depressions.
Sensitivity experiments are also carried out to evaluate theeffects of latent heating due to condensation and evaporationfrom the BoB on the development of the monsoon depression,using the numerical model. We do not deal with the detailed
processes involved in the genesis of the monsoon depression
in the present article.

2
2.1

D A T A A N D M O D E L US E D
Data

We used precipitation data from the GPM Core Observatory that carries the first spaceborne dual-frequency phasedarray precipitation radar (DPR), operating at Ku and Ka bands(13 and 35 GHz, respectively) and a conical-scanning multichannel (10–183 GHz) Global precipitation measurement
Microwave Imager (GMI: Hou et al., 2014). The DPR canobserve the three-dimensional structure of precipitation. The
DPR L2 data version 04A (Iguchi et al., 2015) were used inthis study. The horizontal resolution of DPR is 5.0 km at nadir;the vertical resolution is 250 m. The minimum detectable

269

F I G U R E 1 Distribution of infrared brightness temperature(shading, K) from GridSat data, 850 hPa geopotential height (contours,gpm) and 850 hPa wind field (vectors) from the NCEP CFSv2 at 0000
UTC 17 August 2016. The contour interval for geopotential height is
15 gpm. The area shown (0◦ N–42.6◦ N, 55◦ E–109◦ E) corresponds tothe model domain for the numerical simulation using CReSS. The area
enclosed by the thick dashed line indicates the area for the sensitivity
experiment excluding latent-heat flux from the Bay of Bengal. See
details of the numerical simulation in the text. The thick solid line
denotes the 3,000 m topographic contour [Colour figure can be viewed
at wileyonlinelibrary.com].

rainfall rate is about 0.2 mm⋅h−1 and the swath width is about
125 km. The GMI with a swath of 885 km can observe thehorizontal structure of precipitation over a wider area than the
DPR. GMI L2 data version 04a were used.
CloudSat Cloud Profiling Radar (CPR) data were used
to investigate the vertical structure of cloud in the monsoon
depression. The radar measures at nadir only and therefore
the output swaths have effectively zero width. Radar reflectivity data from the 2B-GEOPROF R04 product were used. The
along-track resolution is 1.7 km and the vertical resolution is
250 m. We also used Gridded Satellite (GridSat) B1 data toshow the cloud distribution over the whole analysis domain(Figure 1). GridSat B1 data are a high-quality Climate Data
Record (CDR) of global infrared measurements from geostationary satellites, based on the International Satellite Cloud
Climatology Project (ISCCP) B1 dataset (Knapp et al., 2011).
The data are on a 0.07◦ latitude × 0.07◦ longitude grid with a
3-hourly time interval.
The observation frequency of the GPM and CloudSat
satellites for a single location is low because of the orbit,
and the observation width is limited. Therefore, successive
observations of precipitation/clouds in a single synoptic-scaledisturbance are unlikely. However, these satellites do provide

FUJINAMI ET AL.

270

detailed three-dimensional structures of precipitation/clouds
when data are available. In contrast, GridSat data provide
continuous brightness temperature data for a broad areafrom the infrared channel of geostationary satellites. The
complimentary use of multiple satellites can thus reveal
the mesoscale features of precipitation/clouds in monsoon
depressions.
We used the National Centers for Environmental Prediction (NCEP) Climate Forecast System version 2 (CFSv2) dataon a 0.5◦ × 0.5◦ grid to investigate the atmospheric circulation
related to the monsoon depression (Saha et al., 2010).

2.2

NEV (i.e. no evaporation from the BoB) run to evaluate theeffect of latent heating on the development of the monsoon
depression. In the NCP run, cloud microphysical processes
are switched off during the simulation and therefore condensation does not occur in the simulation domain; thus, it
behaves as a dry model without considering condensationor precipitation. In contrast, the NEV run only switches off
evaporation from the BoB (i.e. the area enclosed by the thickdashed line in Figure 1) to evaluate the effect of latent-heatflux (i.e. evaporation) from the BoB, although cloud microphysical processes are included in the experiment, as in the
CNTL run.

Model and experimental design

We used the Cloud Resolving Storm Simulator (CReSS:
Tsuboki and Sakakibara, 2002; 2007; Tsuboki, 2008), which
is a non-hydrostatic and compressible model, to simulate
the monsoon depression. CReSS has been used to examine various low-pressure systems such as typhoons, cyclones
and midlatitude cyclones, and their associated mesoscale
cloud/precipitation systems (e.g. Akter and Tsuboki, 2012;
Hirata et al., 2015; 2016; Tsuboki et al., 2015). A bulk
parametrization for cold rain (Murakami, 1990; Ikawa andSaito, 1991; Murakami et al., 1994) was applied to cloud
microphysical processes. The microphysical schemes used in
the model included both prognostic equations for the mixing ratio of water vapour, cloud water, cloud ice, rain, snow
and graupel, and for the number concentrations of cloud ice,snow and graupel. A 1.5-order turbulent kinetic energy closure scheme (Deardorff, 1980) was used for subgrid-scale
turbulence processes. The model domain was South Asiaand its surroundings, including the Himalayas and the
Tibetan Plateau (55◦ –109◦ E, 0◦ –42.6◦ N; the whole domain
of Figure 1) and the horizontal grid resolution was 0.05◦ longitude × 0.05◦ latitude. The atmospheric vertical coordinate
was a terrain-following coordinate, and 45 vertical layers were
included with the top level at a height of 22.5 km. The time
step for terms related to sound waves was 2.0 s, and for otherterms was 6.0 s. We selected 1200 UTC on 14 August 2016 as
the initial time and the integration period was 7 days. The initial and lateral boundary conditions for CReSS were derived
from the NCEP CFSv2 data with a time resolution of 6 h
and a spatial resolution of 0.5◦ longitude × 0.5◦ latitude. We
used the daily National Oceanic Atmospheric Administration(NOAA) optimal interpolation (OI) sea-surface temperature(SST) with a spatial resolution of 0.25◦ × 0.25◦ at the initialtime of the simulation. The initial conditions were specified throughout all grid cells of the CReSS domain, whileduring the rest of the integration only lateral boundary conditions were applied. With this configuration of CReSS, we
performed three different experiments: a control experiment(CNTL run) and two sensitivity experiments comprising the
NCP (i.e. no cloud microphysical processes) run and the

3
3.1

RESULTS
Hydrometeor structure

In this section, we investigate the hydrometeor structure of
the monsoon depression on both the synoptic and meso-scales
during its life cycle. Figure 1 shows the monsoon depression
at 0000 UTC 17 August 2016 when it reached minimum central pressure over the head of the BoB during its life cycle(Figure 2). Deep convection (<200 K) that reaches above
15 km develops south of its centre, especially in the southwest, while convective activity is suppressed from the northto east of the centre, showing a synoptic-scale cloud distribution that is typical of monsoon depressions (Godbole, 1977;
Sikka, 1977; Stano et al., 2002; Hunt et al., 2016b; Fujinami
et al., 2017). Active convection extends from the southwestern part of the monsoon depression to the western Tibetan
Plateau across the western Himalayas, corresponding to theconvergence zone between the prevailing monsoon westerlies
and the easterlies in the northern part of the depression in thelower troposphere.
Figure 2 shows time series of the central pressure of themonsoon depression from the CFSv2 data and the CNTL
run by CReSS from 0000 UTC 16 August to 1200 UTC
21 August. The CFSv2 data reveal that the central pressure
decreases rapidly from 0000 UTC 16 August and reaches aminimum at 0000 UTC 17 August, followed by a gradualincrease. Figure 3 shows the time sequence of the monsoon depression and its hydrometeors from 15 to 18 August.
Dates and times for the figures correspond to the four overpasses of the GPM observatory over the monsoon depression(Figure 3e–h). The observation time for infrared brightness
temperature from GridSat (Figure 3a–d) corresponds closely
to that from the GPM observations. At 0000 UTC 16 August(Figure 3a), a day before the low reached monsoon depression
intensity, the head of the BoB is largely covered by clouds
with a few contiguous cold cloud areas. Active convection isobserved in the cloud area, accompanied by high precipitation (>20 mm⋅h−1 ) (Figure 3e). A small vortex with a closed
contour appears centred on 20◦ N, 90◦ E (Figure 3a) and isFUJINAMI ET AL.

Central pressure (hPa)

271

Time
F I G U R E 2 Time series of the central pressure of a monsoon depression based on CFSv2 data (solid line with closed circles) and the CReSSCNTL run (solid line with open circles)

F I G U R E 3 Distribution of infrared brightness temperature (shading, K) from GridSat data and 850 hPa geopotential height (contours, gpm)
and wind field (vectors) from the NCEP CFSv2 at (a) 0000 UTC 16 August, (b) 1200 UTC 16 August, (c) 0000 UTC 18 August, and (d) 1200 UTC
18 August in 2016. The contour interval for geopotential height is 15 gpm. Also shown is precipitation (mm⋅h−1 ) derived from the GMI at (e) 2336
UTC 15 August, (f) 1254 UTC 16 August, (g) 2325 UTC 17 August and (h) 1245 UTC 18 August. The area enclosed by thick solid lines represents
the GMI swath [Colour figure can be viewed at wileyonlinelibrary.com].

observed from 0000 UTC 15 August over the head of the BoB(not shown).
The vortex structure of the low becomes obvious at thehead of the BoB from 1200 UTC 16 August (Figure 3b). The
BoB is largely covered by clouds, but active convection is

confined along the closed contour of the vortex (i.e. the 1,380
gpm contour in Figure 3b) away from its centre. This feature
is more remarkable in the precipitation distributions from the
GMI (Figure 3f); high precipitation occurs in a ring-like rainband at a radial distance of ∼200–300 km from the centre of

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FUJINAMI ET AL.

F I G U R E 4 (a) Spatial distribution of precipitation (shading, mm⋅h−1 ) derived from the GMI at 2336 UTC 15 August 2016; this is anenlarged view of Figure 3e around the target precipitation systems. The area enclosed by grey shading (dashed lines) indicates the GMI (GPM DPR)
swath. (b) Spatial distribution of precipitation (shading, mm⋅h−1 ) at 1,000 m above sea level from GPM DPR. The area enclosed by grey shading
indicates the GPM DPR swath. (c) Vertical–latitude section of precipitation (shading, mm⋅h−1 ) along line A–A' from GPM DPR. (d)Vertical–longitude section along line B–B'. (e,f) As in (a,b), but for 1254 UTC 16 August. (g,h) As in (c,d), but for line C–C' and line D–D'. (i,j) Asin (a,b) but for 1245 UTC 18 August. (k,l) As in (c,d), but for 1245 UTC 18 August. The observation times in (a,e,i) correspond to those in
Figure 3e,f,h for GMI, respectively [Colour figure can be viewed at wileyonlinelibrary.com].

the low (21◦ N, 90◦ E). Some MCSs are observed over northern India in front of the Himalayas where there is convergence
between the monsoon westerlies and the easterlies of the monsoon depression (Figure 3b). On 17 August the vortical low
develops into the monsoon depression (Figure 1).
The monsoon depression makes landfall at 1200 UTC
17 August and then moves northwestward (not shown).

At 0000 UTC 18 August (Figure 3c), active convection
is no longer observed over the BoB. Active convection
only appears within the monsoon depression and the areaof convergence extending northwestward from the monsoon depression over land. Mesoscale spiral rain bands
are observed within the monsoon depression after landfall(Figure 3g,h).

FUJINAMI ET AL.

(a) IR

273

0600 UTC 18 Aug(b) CloudSat CPR

A

0755 UTC 18 Aug

A
Height (km)

A'

A'

(K)

(dBZ)

F I G U R E 5 (a) Distribution of infrared brightness temperature (shading, K) from GridSat data and 850 hPa geopotential height from the
NCEP CFSv2 at 0600 UTC 18 August 2016. The contour interval for geopotential height is 15 gpm. Line A–A' indicates the CloudSat overpass at
0755 UTC 16 August 2016. (b) Radar reflectivity (dBZ) from CloudSat Cloud Profiling Radar (CPR) along line A–A' [Colour figure can be viewed
at wileyonlinelibrary.com].

Figure 4 shows the horizontal distribution and verticalcross-section of precipitation from the DPR for the same dates
and times as in Figure 3 except for 2325 UTC 17 August.
Note that the radar swaths of DPR only overlap the central
portions of the GMI swaths, as shown by the area within
the dashed lines in Figure 4a,e,i. At 2336 UTC 15 August(Figure 4a,b), the precipitation area over the BoB consists ofa mixture of vertical convective cells and stratiform rainfallwith the melting layer at a height of about 5 km (Figure 4c,d),thereby showing the features of MCSs (Maddox, 1980; Houzeet al., 2007). Some convective cells exceed 10 km in height;
most of the high precipitation signals (>20 mm⋅h−1 ) are confined below 7 km. At 1254 UTC 16 August (Figure 4e,f),
the area of high precipitation around 21◦ N, 87◦ E, which ispart of the ring-like rain band, consists of a row of vertically
upright intense convective cells that exceed 15 km in height(Figure 4g). At 1245 UTC 18 August (Figure 4i), the DPR
also detects the finer spiral structure of high precipitation(>20 mm⋅h−1 ) (Figure 4j). Almost all the precipitation cells
along the rain bands have a vertically erect convective structure (Figure 4k), but they are not as high (<10 km) as thoseon 16 August (Figure 4g).
The CloudSat overpass reveals the inner structure beneaththe cloud tops of the monsoon depression at approximately
0755 UTC 18 August (Figure 5). Dense clouds make up themonsoon depression as a whole. The mean cloud top heightis around 14 km. The layer with horizontally elongated highreflectivity around 5 km is a melting layer. There is an area of
weak reflectivity (i.e. an eye-like structure) between 23.5◦ N
and 24◦ N, close to the depression centre under high-altitude
clouds. The southern portion of the monsoon depression tends
to have vertically upright convective echoes (pink colour),whereas a horizontally elongated melting layer without deep
upright echoes exists in the northern portion.

3.2

Model experiments

Section 3.1 revealed different features of mesoscale precipitation systems such as the MCS structure with active deep

convection cells and stratiform precipitation around 0000
UTC 16 August (Figure 3e), the extremely deep convection
cells along a ring-like rain band on the southwestern side ofthe depression around 1200 UTC 16 August (Figure 3f) andthe spiral rain bands (Figure 3g,h) of high precipitation after
landfall under the synoptic-scale cloud cover (Figure 3a–d).
These different features might be associated with different
stages in the development of the monsoon depression. In thissection, we investigate the mesoscale features of precipitation
systems at different life stages of the depression using CReSS.
Then, we propose a mechanism for deep convection during
the rapid development phase of the depression.
Central pressure from the CNTL run shows a similar
temporal evolution to that in CFSv2 throughout the period(Figure 2). The SLP fields are also well reproduced in terms ofspatial distribution and temporal evolution (Figure S1). Thus,
the CNTL run simulates well the development of the monsoon
depression.
Figure 6 shows the spatial distribution of sea-level pressure (SLP) and precipitation from the CNTL run for 16 to
18 August. The UTC values in Figure 6 are not exactly thesame as in Figure 3e–h but correspond to almost the samestage in the life cycle. The horizontal distribution of precipitation from the CNTL run captures well the characteristicfeatures seen in Figure 3e–h. At 0000 UTC 16 August, MCSs
with some intense precipitation are found in the head of
the BoB with a small closed cyclonic circulation. In contrast, the low-pressure centre is obvious at 1100 UTC 16
August at the head of the BoB (Figure 6b). A ring-like rainband of high precipitation occurs away from the centre of
the monsoon depression along the closed isobar over thenorthernmost BoB. At 0000 UTC 18 August the precipitation bands are elongated in the northwestward direction, westof the monsoon depression, and the model also reproducessome rain bands and the ring-like structure of precipitation around the centre of the depression. The depression has
west–southwest precipitation maxima with some spiral-likerain bands (Figure 6d).

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FUJINAMI ET AL.

Surface precipitation (shading, mm⋅h−1 ) and sea-level pressure (contours, hPa) at (a) 0000 UTC 16 August, (b) 1100 UTC 16
August, (c) 0000 UTC 18 August and (d) 0600 UTC 18 August in 2016 simulated by the CNTL run. The contour interval for sea-level pressure is
2 hPa [Colour figure can be viewed at wileyonlinelibrary.com].

FIGURE 6

Figure 7 presents the vertical structure of the hydrometeors
in the CNTL run for the different life stages, as in Figure 6, forcomparison with the observed vertical structure from GPM
DPR (Figure 4). CReSS does not calculate precipitation intensity as a model output, aside from surface precipitation, but
provides vertical distributions of rain, graupel and snow mixing ratios. The vertical distribution of the total mixing ratioof the three hydrometeors corresponds to that of precipitation particles in the precipitation system. The cross-sections
are drawn along the meridional and zonal lines that pass oversome local maxima of simulated surface precipitation. It is
worth noting that the CNTL run successfully simulates the
DPR-observed vertical structure of the convective cells. At
0000 UTC 16 August, there is a vertically erect deep convective cell around 21◦ N and stratiform-like precipitation around19◦ N along the line A–A' (Figure 7b), indicative of MCS
structure with some convective cells at different development
phases. The MCS-like structure is also seen along line B–B'.
At 1100 UTC 16 August, convective cells along the ring-likestructure, which are located over the northernmost BoB, aredeep and intense, and reach 15 km in height along lines C–C'
and D–D' around the centre of the low (Figure 7e,f). The deep
convective cells have a large amount of graupel, indicative ofstrong updraught. After landfall of the monsoon depression,
the amount of graupel decreases and the heights of precipitation cells generally become lower (Figure 7h,i), consistent
with the GPM observations (Figure 4k,l).
An outstanding feature of the precipitation system is
extremely deep convection cells around the centre of the low,
particularly on the western-to-southwestern side during the
rapid developing phase (Figures 4g and 7e). Here, we reveal
how deep convection occurs in this area during this phasein the CNTL run. The ring-like distribution of precipitation
seen in Figure 7d corresponds to areas of local maxima of
low-level moisture flux convergence (Figure 8a). Extremely
deep convection is generated at a local maximum of moisture
flux convergence along line A–A' in the area where northerlies encounter strong low-level monsoon westerlies on thewestern side of the low; this is part of a synoptic-scale convergence zone extending northwestward up to the foothills of theHimalayas. A similar feature is found in the observed windfields (Figures 1 and 3b). The height–latitude cross-section
along line A–A' shows that low-level northerlies encounter athick layer of low-level westerlies around 21.2◦ N (Figure 8c).
The spatial distribution of equivalent potential temperature(𝜃 e ) shows that low-level monsoon westerlies from the Indian
subcontinent to the south of 21◦ N bring relatively dry air (i.e.
lower 𝜃 e ) toward the head of the BoB (Figure 8b). 𝜃 e gradually increases moving eastward between 19◦ N and 21◦ N.
Note that low-level warm, humid air of high 𝜃 e extends to thenorth of the centre of the low within a closed cyclonic circulation. Convective available potential energy (CAPE), whichis related to the potential updraught strength, is also large(>1,500 J⋅kg−1 ) in the same region. Surface-based CAPE wasused in this study. Hunt et al. (2016a) revealed large CAPE
and small convective inhibition (CIN) within the monsoon
depression from composite monsoon depressions. Air withhigh 𝜃 e is confined to below 900 hPa (i.e. the atmosphericboundary layer) to the north of 21.2◦ N, whereas dry air isfound around 700 hPa, resulting in a high degree of conditional instability with large CAPE (Figure 8d). In the area of
low-level convergence between northerlies and strong westerlies, the environmental air can be raised up to the level offree convection (LFC), which leads to deep convection due tothe release of large CAPE to the west–southwest of the centre of the low. The same mechanism is also applicable to thedeep convection at 0000 UTC 16 August shown in Figure 7b.
The height–longitude cross-section along line B–B' crosses
the centre of the low (Figure 8e), which is located around88.4◦ E. The updraught near the centre of the low is caused by
the same mechanism as that along line A–A' in Figure 8c. The
centre of the low is characterized by high 𝜃 e from the surface
to the mid-troposphere.
As seen in Figures 1 and 3, the monsoon depression develops over the head of the BoB. This means that evaporationfrom the BoB might be important for cyclogenesis. In fact,
the increase of 𝜃 e toward the leeside in the boundary layer

FUJINAMI ET AL.

275

(b)
17.5

B

B'

A

A

A'

15

15

12.5

12.5

Height (km)0000 UTC 16 Aug
A'
Height (km)

(a)

(c)
17.5

10
7.5

7.5
5

2.5

2.5

0

0

(e)

D

D'

Height (km)

C'

C

C

(f)
C'

17.5

15

15

12.5

12.5

Height (km)1100 UTC 16 Aug

(d)

10
7.5

5

2.5

2.5

0

0

0600 UTC 18 Aug

F'

E

(i)
17.5

E'

15

15

12.5

12.5

10
7.5

Height (km)F

Height (km)

E'

E

D'

7.5

(h)
(g)

D

10

5

17.5

B'

10

5

17.5

B

F

F'

10
7.5

5

5

2.5

2.5

0

0

−1

(mm h )
(g kg−1)

F I G U R E 7 (a) Surface precipitation at 0000 UTC 16 August 2016 simulated by the CNTL run; this is an enlarged view of Figure 6a aroundthe head of the Bay of Bengal. (b,c) Vertical cross-sections of simulated precipitating hydrometeors: rain (black solid contours, g⋅kg−1 ), graupel(blue dashed contours, g⋅kg−1 ) and snow (white solid contours, g⋅kg−1 ) along line A–A' and line B–B', respectively. Shading denotes areas where thesum of the rain, graupel and snow mixing ratios exceeds 0.05 g⋅kg−1 , indicating the existence of precipitating particles. The contours forhydrometeor mixing ratios are drawn at 0.05, 0.2, 0.4, 0.8, 1.2, 2.0 and 3.0 g⋅kg−1 . The thick solid contour shows the 0 ◦ C level. (d–f) As in (a–c),
but for 1100 UTC 16 August when the ring structure of precipitation was observed near the northern coast of the BoB in Figure 6b. Verticalcross-sections are for line C–C' and line D–D'. (g–i) As in (a–c), but for 0600 UTC 18 August after the monsoon depression makes landfall. Verticalsections are for lines E–E' and F–F', respectively [Colour figure can be viewed at wileyonlinelibrary.com].

FUJINAMI ET AL.

276

(a)

(b)

A'
B'

B
A

(K)

(c)

A

(d)

A'

A

(m s−1)(e)B(m s−1)

A'

(K)

(f)

B'

(m s−1)

B

B'

(K)

(a) Geopotential height (contours, gpm), water vapour flux (vectors, kg⋅kg−1 ⋅m⋅s−1 ) and its divergence (shading,
1.0 × 10 kg⋅kg ⋅s−1 ) at 975 hPa at 1100 UTC 16 August simulated by the CNTL run. The contour interval for geopotential height is 10 gpm. (b)As in (a), but for horizontal wind fields (vectors), equivalent potential temperature (shading, K) and CAPE (contours, J kg−1 ). The contour intervalof CAPE is 500 J⋅kg−1 ; contours with values of <1,000 J⋅kg−1 are omitted. (c) Height–latitude cross-section of vertical velocity (shading, m⋅s−1 ),zonal wind component (contours, m⋅s−1 ), and meridional and vertical winds (vectors; vertical wind is multiplied by 10) along line A–A' in (a). Only
vectors that exceed 1 m⋅s−1 in vertical velocity or 5 m⋅s−1 in meridional wind are plotted. The contour interval for zonal wind is 5 m⋅s−1 . (d) As in(c), but for equivalent potential temperature (shading, K) and horizontal water vapour flux divergence (contours, kg⋅kg−1 ⋅s−1 ). Only contours withvalues less than −3.0 × 10−6 kg⋅kg−1 ⋅s−1 are plotted. (e) As in (c), but for the height–longitude cross-section of the meridional wind component(contours) and the zonal and vertical winds (vectors) along line B–B' in (a). (f) As in (e), but for equivalent potential temperature (shading, K) andhorizontal water-vapour flux divergence (contours, kg⋅kg−1 ⋅s−1 ) [Colour figure can be viewed at wileyonlinelibrary.com].

FIGURE 8−6−1

FUJINAMI ET AL.

277(a) CNTL run 0000 UTC 15 Aug(b) CNTL run 0000 UTC 16 Aug(c) CNTL run 0000 UTC 17 Aug(d) CNTL run 0000 UTC 18 Aug

20 (m s−1)

100 200 300 450 (W m−2)

F I G U R E 9 Spatial distribution of sea-level pressure (contours, hPa), 10 m horizontal wind (vectors) and surface latent-heat flux (shading,W⋅m−2 ) from the CNTL run at 0000 UTC on (a) 15 August, (b) 16 August, (c) 17 August and (d) 18 August 2016. Contour interval for sea-levelpressure is 2 hPa. Wind vectors with wind speed less than 3 m⋅s−1 are omitted. Latent-heat flux less than 100W⋅m−2 is omitted [Colour figure can beviewed at wileyonlinelibrary.com].
(a) NCP run 0000 UTC 15 Aug(b) NCP run 0000 UTC 16 Aug(c) NCP run 0000 UTC 17 Aug(d) NCP run 0000 UTC 18 Aug(e) NEV run(f) NEV run(g) NEV run(h) NEV run 0000 UTC 18 Aug

0000 UTC 15 Aug

0000 UTC 16 Aug

0000 UTC 17 Aug(mm h−1)

20 (m s−1)

F I G U R E 10 Spatial distribution of sea-level pressure (contours, hPa) and 10 m wind (vectors) at 0000 UTC on (a) 15 August, (b) 16 August,(c) 17 August and (d) 18 August 2016 from the NCP run. Contour interval for sea-level pressure is 2 hPa. Wind vectors with wind speed less than
3 m⋅s−1 are omitted. (e–h) As in (a–d), but for the NEV run. Shading denotes precipitation [Colour figure can be viewed at wileyonlinelibrary.com].

suggests the intake of moisture from the ocean surface to theflowing air (Figure 8b). Figure 9 shows the time sequenceof surface latent-heat flux in the CNTL run. At 0000 UTC
15 August (Figure 9a), surface latent-heat flux is large overthe southern BoB south of 17◦ N. From 16 to 17 August(Figure 9b,c), latent-heat flux increases greatly over the headof the BoB as surface wind is enhanced due to the development of the monsoon depression. In addition, the area of highlatent-heat flux around 20◦ N, 87◦ E corresponds to that of high
SST (>29 ◦ C) during 14–16 August according to the daily OI
SST dataset (not shown).
To evaluate the respective effects of latent heating of condensation due to cloud/precipitation processes and evaporation from the BoB on the genesis and development of the monsoon depression, we performed two sensitivity experiments:

the NCP run and the NEV run. In the NCP run, cloud
microphysical processes are not employed during the simulation. Latent heating from condensation does not occur inthis run. In the NCP run (Figure 10a–d), unlike the CNTL
run (Figure 8e–h), the low does not develop at all aroundthe BoB. The NCP run revealed that latent heat of condensation from active convection (Figures 1,3 and 4) is essential tothe development of the monsoon depression, similar to othertropical disturbances such as typhoons and cyclones.
In the NEV run, cloud microphysical processes are active,
but evaporation from the BoB is eliminated. At 0000 UTC 15
August (Figure 10e), the spatial distribution of precipitationand SLP is quite similar to that in the CNTL run (Figure 8e).
After that, however, SLP does not decrease over the head of
the BoB from 16 to 17 August (Figure 10f,g). In addition,

FUJINAMI ET AL.

278(a) CNTL run 0000 UTC 15 Aug(b) CNTL run 0000 UTC 16 Aug(c) CNTL run 0000 UTC 17 Aug(d) CNTL run 0000 UTC 18 Aug(e) NEV run(f) NEV run(g) NEV run(h) NEV run 0000 UTC 18 Aug

0000 UTC 15 Aug

0000 UTC 16 Aug

0000 UTC 17 Aug(K)

20 (m s−1)

F I G U R E 11 Spatial distribution of sea-level pressure (contours, hPa), 10 m horizontal wind (vectors) and equivalent potential temperature at
975 hPa (shading, K) at 0000 UTC on (a) 15 August, (b) 16 August, (c) 17 August and (d) 18 August 2016 from the CNTL run. White shading
indicates areas above 975 hPa (∼200 m above sea level). The contour interval for sea-level pressure is 2 hPa. Wind vectors with wind speeds of<3 m⋅s−1 are omitted. (e–h) As in (a–d), but for the NEV run [Colour figure can be viewed at wileyonlinelibrary.com].

the vortex core and surrounding precipitation systems do notappear over the head of the BoB. A comparison between the
CNTL run and the NEV run demonstrates that the moisturesupply from the warm ocean surface over the BoB plays anactive role in the rapid intensification of the vortex core overthe head of the BoB.
We further investigate how the moisture supply from the
BoB provides a suitable environment for rapid development.
Figure 11 shows a time sequence of 𝜃 e in the atmosphericboundary layer for the CNTL and NEV runs. In the CNTL
run, at 0000 UTC 15 August (Figure 11a), an area of high𝜃 e extends broadly over the BoB, whereas a low-𝜃 e area (i.e.
a dry area) is observed along the eastern coast of India, leeward of the Indian subcontinent, which contains plateaus andmountains, relative to the low-level monsoon westerlies. During the rapid developing phase of the depression from 16 to
17 August (Figure 11b,c), dry air of low 𝜃 e extends eastward
between 18◦ N and 19.5◦ N over the head of the BoB as the
low develops, because of the enhanced westerlies blowing
from land. In contrast, the low/monsoon depression aroundthe northernmost part of the BoB contains air of high 𝜃 e . The
wind field during this phase is similar to that at 1100 UTC
16 August (Figure 8b). After landfall, the high-𝜃 e area almost
disappears from the northernmost part of the BoB. The NEVrun revealed that an air mass of high 𝜃 e does not form overthe BoB without evaporation from the BoB (Figure 11e–h).
This influence extends up to the Gangetic Plain. The difference between the CNTL and NEV runs suggests that surface
latent-heat flux from the BoB plays a role in maintaining

warm, humid air in the boundary layer around the depression
during the development phase.
Areas of high 𝜃 e in the boundary layer correspond to
those of large CAPE (Figure 12). On 15 August (Figure 12a),a moderate–large CAPE area extends sporadically over thehead of the BoB. From 16 to 17 August (Figure 12b,c), thelarge-CAPE area extends throughout the low-pressure area
near the northernmost part of the BoB. The observational evidence and results from model experiments suggest that deep
convection is produced by the release of the instability within
the air rising in the area of low-level convergence during the
rapid development phase of the monsoon depression. After
landfall, the large-CAPE area does not follow the depression(Figure 12d). The CAPE value is small over and around thedepression over land. Therefore, the height of the convection
cells decreases after landfall (Figure 7g–i). In the NEV run(Figure 12e–h), the lack of evaporation from the BoB leads
to a small-CAPE environment over the BoB and the Gangetic
Plain, resulting in precipitating convection being suppressedover the BoB.

4DISCUSSION

Deep convection, MCSs and a possible
4.1
feedback
MCSs with embedded deep convection appear near the low,
particularly along the western–southwestern side during the
rapid development phase of the monsoon depression over theFUJINAMI ET AL.

279(a) CNTL run 0000 UTC 15 Aug(b) CNTL run 0000 UTC 16 Aug(c) CNTL run 0000 UTC 17 Aug(d) CNTL run 0000 UTC 18 Aug(e) NEV run(f) NEV run(g) NEV run(h) NEV run 0000 UTC 18 Aug

0000 UTC 15 Aug

0000 UTC 16 Aug

0000 UTC 17 Aug(J kg−1)

F I G U R E 12 Spatial distribution of sea-level pressure (contours, hPa) and CAPE (shading, J⋅kg−1 ) at 0000 UTC on (a) 15 August, (b) 16
August, (c) 17 August, and (d) 18 August 2016 from the CNTL run. The contour interval for sea-level pressure is 2 hPa. (e–h) As in (a–d), but for
the NEV run [Colour figure can be viewed at wileyonlinelibrary.com].

head of the BoB (Figures 4a–h and 7a–f). The large CAPE
environment and release of the instability over the area of
low-level convergence probably generate MCSs with embedded deep convection (Figure 8a,b). Evaporation from the headof the BoB maintains the humid boundary layer within the
low during this phase. The maxima of surface latent-heatflux correspond to areas of strong dry air inflow from theIndian subcontinent to the south of the low (Figures 9b,c and11b,c) and SST maxima along the western coast of the head of
the BoB. These conditions might enhance evaporation fromthe head of the BoB and add a large amount of moisture to
the air flowing inward toward the low/monsoon depression(Figure 8b).
The southwesterlies/southerlies on the eastern side ofthe low draw in warm humid boundary-layer air to thelow, maintaining a large-CAPE environment and refuellingthe deep convection. Then, the warm humid air enters theboundary layer in the closed cyclonic circulation of the
low near the northernmost part of the BoB. Despite thelarge-CAPE environment, moist convection is supressed in
the northern-to-northeastern part of the low/monsoon depression (Figures 1 and 6b). In fact, weak subsidence appears in
the mid- to upper-level troposphere in the CNTL run (figure
not shown). Synoptic-scale weak subsidence was observedin monsoon depressions to the north–northeast of the low/monsoon depression (Fujinami et al., 2017). Weak subsidence
might result in the dry air (i.e. low 𝜃 e ) around 600–700 hPa
to the north of 21◦ N (Figure 8d), resulting in the high degreeof conditional instability. Then, moist boundary-layer air
is transported southward by northerlies/northeasterlies thatmake up the western part of the closed cyclonic circulation of
the low, without a large consumption of water vapour becauseof the supressed moist convection. The moist low-level
northerlies eventually encounter strong westerlies on thewestern and southwestern sides of the low. The low-level air
is raised up to the LFC and large CAPE is released, resulting
in the generation of extremely deep convection and MCSs.
Isolated deep convection and MCSs exist within the developing tropical depression/cyclone, similar to the monsoon
depression. In tropical depressions/cyclones, Houze (2010)
illustrated conceptually that the vorticity of the low-level
environment is stretched by convergence related to deep convection at the base of convective updraughts and is advected
upward. These updraughts result in centres of high positive
vorticity being extended upward. An MCS begins as one ormore such isolated instances of deep convection. The MCS
has both a convective and a stratiform component during its
mature stage. The stratiform region of the MCS also accumulates positive vorticity from the remaining portions of theconvective cells to form a mid-level mesoscale convective
vortex. The low as a whole develops rapidly through such processes related to the MCS life cycle. Murthy and Boos (2019)
also demonstrated that heating from both deep convection andstratiform precipitation systems can produce the observed vertical distribution of potential vorticity in tropical depressions,indicating the development of tropical depressions throughthe generation of positive potential vorticity from the heating.
Our study revealed that deep convection cells and stratiform
precipitation are indeed observed within the low in the rapid
development phase of the monsoon depression. Thus, theFUJINAMI ET AL.

280

development processes of tropical depressions proposed by
the previous studies may be applicable to monsoon depressions, although further detailed study is needed to clarify therespective effects of deep convective and stratiform rain systems embedded in the MCS on the rapid development of themonsoon depressions.
Further intensification of the low due to the development of deep convection and MCSs enhances the wind speed
around the low and, in turn, might promote more evaporation from the BoB and moist air inflow into the low/monsoon
depression, leading to larger CAPE and more intense moist
convection. This positive feedback might lead to the rapid
intensification of the monsoon depression over the head of
the BoB.

4.2
Genesis environment for monsoon
depressions
Recent studies have proposed that barotropic instability can
explain the structure, propagation and geographic distribution of monsoon depressions (Mishra, 2018; Diaz and Boos,2019). Intraseasonal oscillations such as the BSISO and QBWchange the location of the strong meridional cyclonic shear
zone over the BoB that strongly controls the genesis locations and tracks of the depressions (Goswami et al., 2003;
Hatsuzuka and Fujinami, 2017). When the low-level monsoon westerlies are enhanced and the core of the westerlies
appears around ∼15◦ N over the BoB due to the propagationof the BSISO and QBW, LPSs including monsoon depressions occur frequently over the head of the BoB (Hatsuzuka
and Fujinami, 2017). The monsoon depression of the present
study occurred under an atmospheric environment suitable for
cyclogenesis in low-level wind fields, similar to that shown
by Hatsuzuka and Fujinami (2017) (Figure 1). The strong
meridional cyclonic shear zone of monsoon westerlies aroundthe monsoon trough can be a favourable environment forbarotropic instability. However, it cannot explain the rapid
growth of the monsoon depression over the BoB (Diaz and
Boos, 2019). The atmospheric environment and condensational heating from the convection cells might jointly promote
the growth of monsoon depressions over the BoB. Observed
monsoon depressions tend to retain their strength in a few days
after landfall without a rapid decline, unlike a typhoon. In fact,
the monsoon depression in this study keeps its strength andvortical structure from 1800 UTC 17 August to 0000 UTC 20
August, after landfall (Figure 2). Barotropic instability might
explain why the monsoon depression can persist after landfall
in conjunction with condensation heating from convection.
The detailed mechanism of the genesis of low-pressure systems remains unclear, although latent heating is a major driverof their development. Monsoon depressions often develop
from the re-genesis of westward-propagating residual lows,which include tropical depressions and lows related to QBW,

from the South China Sea across the Indochina Peninsula(Chen and Weng, 1999). How the initial vortex is generated
in an environment of rich cyclonic vorticity related to themonsoon trough is also an important problem for future work.

5

SUM M ARY

We have investigated the hydrometeors of the monsoon
depression that occurred on 17 August 2016 over the Bay of
Bengal (BoB) using infrared brightness temperature, GPM
GMI, GPM DPR and CloudSat CPR data in the different
stages of the monsoon depression life cycle, focusing on themesoscale precipitation systems. In the rapid development
stage of the low from 0000 to 1200 UTC on 16 August, a day
before the depression reached monsoon depression intensity,
active convection with an MCS structure is observed over thenorthernmost BoB. A ring-like rain band with very deep convective cells appears at a radial distance of a few hundred
kilometres from the centre of the depression, as the vortical
low structure becomes evident. The lowest central pressure
is observed at 0000 UTC 17 August and the low becomesa monsoon depression over the northernmost BoB. Intense
convection occurs in the south–southwest region of the monsoon depression, which is typical of monsoon depressions.
After reaching minimum central pressure over the northernmost BoB, the monsoon depression moves northwestward
and makes landfall at 1200 UTC 17 August. After landfall,
the central pressure of the depression increases gradually.
The monsoon depression has mesoscale spiral rain bands. The
height of precipitation cells in the rain bands becomes lower
than in the developing stage over the BoB. Dense clouds cover
the monsoon depression with taller convective clouds in its
southern part.
We used the CReSS non-hydrostatic and compressible
cloud-resolving model to investigate the time variation ofhydrometeors observed by satellites and a mechanism fordeep convection in the rapid development phase of monsoon
depressions. CReSS performed reasonably well in reproducing the observed spatio-temporal variation of the monsoon
depression from the development stage over the BoB throughthe decaying phase after landfall. Features of the mesoscale
structure of precipitation, such as active convection cells withan MCS structure over the northernmost BoB and a subsequent ring-like rain band with deep convective cells, arealso well reproduced. The deep convective cells appear in
the western-to-southwestern part of the low during the rapid
development phase of the monsoon depression. The deep convective cells are generated by the release of large CAPE in
the area of low-level convergence between northerlies withlarge-CAPE air mass, which make up the western part of aclosed cyclonic circulation centred on the northernmost partof the BoB, and monsoon westerlies with relatively dry air
mass to the south of the cyclonic circulation.

FUJINAMI ET AL.

Two sensitivity experiments were performed to evaluate
the effects of latent heating from cloud/precipitation processes and evaporation from the BoB on the developmentof the monsoon depression. The NCP (i.e. no cloud microphysics processes) run could not reproduce the developmentof the low throughout the simulation. Thus, latent heating due to condensation from active convection is essential
for the growth of the monsoon depression, similar to tropical disturbances such as typhoons and cyclones. The effectof latent-heat supply from the BoB is evident from thecomparison between the CNTL and NEV (i.e. no evaporationfrom the BoB) runs. Evaporation from the BoB adds a large
amount of moisture to the atmospheric boundary layer near
the low over the head of the BoB. The southwesterlies/southerlies on the eastern side of the low draw in warm,
humid boundary-layer air to the closed cyclonic circulationof the low, maintaining a large-CAPE environment in the low
around the head of the BoB. Northerlies with large-CAPE
air mass encounter strong westerlies in the western/southwestern part of the low, resulting in deep convective cells
driven by the instability. Without evaporation from the BoB,
deep convection cannot develop around the BoB because of
the small-CAPE environment. Thus, both cloud/precipitation
processes and evaporation from the BoB are essential for the
rapid growth of the monsoon depression over the head of
the BoB.
Although this study examined only one typical monsoon depression, it reveals some novel findings about themesoscale structure of cloud/precipitation systems in a monsoon depression during the different stages of its life cycle. To
fully understand the common mesoscale features of cloud/precipitation in monsoon depressions, further studies areneeded of multiple monsoon depressions based on a ground
precipitation radar network that can observe the whole life
cycle of the depression with high spatio-temporal resolution.
Our results demonstrate that latent heating from the BoB
is important for the development of monsoon depressions.
Most monsoon depressions over the BoB are generated to thenorth of 15◦ N (e.g. Hatsuzuka et al., 2014; Hunt et al., 2016a),but the genesis location varies from coastal to offshore areas.
More monsoon depressions should be examined in the future
to reveal other processes related to ocean–atmosphere interactions during the rapid development phase.

ACKNOWLEDGEMENTS
We thank Nobuhiro Takahashi, Hirohiko Masunaga andHironari Kanamori for their valuable comments and suggestions. We also thank two anonymous reviewers for helping
to improve the manuscript. This study was conducted withsupport from the Japan Society for the Promotion of Science (JSPS) KAKENHI Grants JP 26220202, JP 26400465
and JP 18KK0098. Parts of this study were conducted281

under a collaborative research programme of the Institute forSpace-Earth Environmental Research, Nagoya University.
ORCID
Hatsuki Fujinami
Hidetaka Hirata

https://orcid.org/0000-0002-2400-479X
https://orcid.org/0000-0001-6078-9101

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How to cite this article: Fujinami H, Hirata H,
Kato M, Tsuboki K. Mesoscale precipitation systems
and their role in the rapid development of a monsoon
depression over the Bay of Bengal. Q J R Meteorol Soc.
2020;146:267–283. https://doi.org/10.1002/qj.3672