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  • Ziegler, C. L., , and E. N. Rasmussen, 1998: The initiation of moist convection at the dryline: Forecasting issues from a case study perspective. Wea. Forecasting, 13, 11061131, doi:10.1175/1520-0434(1998)013<1106:TIOMCA>2.0.CO;2.

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    (a) Cyclone frequency and intensity during 2001–12 based on the JTWC data. Overlapping intensities with the same value are labeled with a lighter color to avoid confusion with the number of TCs. The specific humidity (shaded, g kg−1) and temperature anomalies (solid contours, 2-K intervals) calculated from the background area in the figure and the horizontal winds (vector, m s−1) at 850 hPa are displayed for (b) May 2007, (c) November 2007, (d) May 2010, and (e) October 2010 using the monthly mean NCEP–CFSR reanalysis data. The tracks represent the TCs that formed in 2007 and 2010. The areas of the model’s parent domain (D1) and nested domain (D2) are indicated in (b).

  • View in gallery

    The surface pressure (contours, 2-hPa intervals), wind velocity (vector, m s−1), and vorticity (shaded, ×10−5 s−1) at 850 hPa for the (a) simulations and (b) NCEP–CFSR reanalysis data at each TC genesis time (see Table 1). The black dots are the genesis points for each TC using the JTWC data.

  • View in gallery

    The surface-based CAPE (shaded, J kg−1) for the (a) simulations and (b) NCEP–CFSR reanalysis data. The black dots indicate the genesis points of each TC using the JTWC data. The display time is 2 days prior to the genesis of each TC.

  • View in gallery

    The 6-h-averaged moisture fluxes (shaded, m s−1), reflectivities (contours, 20-dBZ intervals), streamlines at 925 hPa, and the zero zonal wind shear between 850 and 200 hPa (dashed line, m s−1). The averages are calculated from the time 0600 UTC 11 May for Akash, 0600 UTC 15 May for Laila, 1800 UTC 8 Nov for Sidr, and 0600 UTC 18 Oct for Giri (i.e., 2.5 days prior to each genesis). The stars are the positions of the depressions (max wind speed of 15 kt) of each TC.

  • View in gallery

    Water vapor mixing ratio (shaded, g kg−1), reflectivities (contours, 20-dBZ intervals), temperature anomalies (dashed line, 1-K intervals) calculated from the average of domain D2, and horizontal wind (vector, m s−1) at 925 hPa at time t (time t for each case is mentioned in the text). The stars in all panels are the sounding points.

  • View in gallery

    Skew T–logp sounding and wind barbs (kt) at the location of each TC (the stars in Fig. 5). The long dashed, dot–dot–dashed, solid, and short dashed lines represent the dewpoint sounding, temperature sounding, parcel trace, and relative humidity sounding, respectively.

  • View in gallery

    Hodographs at the sounding points described in Fig. 6. The numbers represent the height in km.

  • View in gallery

    Premonsoon TCs: (a) mature-stage MCS (shaded; dBZ), 850-hPa wind velocity (m s−1), and MCS after 6 h (hatched area) for Akash (2007). The yellow line is the approximate RIJ. The vertical cross section along AB in (a) indicates the (b) vertical reflectivities (shaded; dBZ), downdrafts (thin contours; 0.2-m s−1 interval), water vapor mixing ratio (thick contours; >16 g kg−1), and cross-sectional parallel-system-relative wind with vertical velocity (vectors; m s−1). (c) Potential temperature anomalies for a 6° × 4° area for the system-wide values (shaded; K) and updrafts (thick solid contours; 2 m s−1 interval); the dashed (thin solid) contours of 1 × 10−3 s−1 represent the vertical cyclonic (anticyclonic) vorticity. (d)–(f) As in (a)–(c), but for Laila (2010). The wide arrows in (a) and (d) are the approximate MCS directions.

  • View in gallery

    As in Fig. 8, but for the postmonsoon TCs in grayscale.

  • View in gallery

    Distribution of the horizontal cyclonic vortices (contours of 1 × 10−3 s−1) for the time period from t + 1 to t + 13 h in 1-h intervals at (a) 925, (b) 500, and (c) 200 hPa for each TC. In (a)–(c) the vectors are the horizontal winds (m s−1) at t + 13. (d) The vertical cross section of the updraft (shaded; m s−1) and cyclonic vorticity (contours; 0.4 × 10−3 s−1) along the point located in (c) for each case.

  • View in gallery

    Reflectivity at 925 hPa (shaded, dBZ) and the geopotential height (shaded, m) along with the horizontal wind (vectors, m s−1) at 500 hPa for TCs in 2007.

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    Average NOAA daily-interpolated OLR (shaded, W m−2) and 850-hPa horizontal wind (vectors, m s−1) for 2 days prior to the formation of the premonsoon TCs with intensities (a) ≥64 and (b) <64 kt during 2001–12. (c),(d) As in (a),(b), but for the postmonsoon TCs.

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Mesoscale Convection and Bimodal Cyclogenesis over the Bay of Bengal

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  • 1 Department of Physics, Bangladesh University of Engineering and Technology, Dhaka, Bangladesh
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Abstract

Mesoscale convective systems (MCSs) are an essential component of cyclogenesis, and their structure and characteristics determine the intensity and severity of associated cyclones. Case studies were performed by simulating tropical cyclones that formed during the pre- and postmonsoon periods in 2007 and 2010 over the Bay of Bengal (BoB). The pre- (post) monsoon environment was characterized by the coupling of northwesterly (southwesterly) wind to the early advance southwesterly (northeasterly) monsoonal wind in the BoB. The surges of low-level warm southwesterlies with clockwise-rotating vertical shear in the premonsoon period and moderately cool northeasterlies with anticlockwise-rotating vertical shear in the postmonsoon period transported moisture and triggered MCSs within preexisting disturbances near the monsoon trough over the BoB. Mature MCSs associated with bimodal cyclone formations were quasi linear, and they featured leading-edge deep convection and a trailing stratiform precipitation region, which was very narrow in the postmonsoon cases.

In the premonsoon cases, the MCSs became severe bow echoes when intense and moist southwesterlies were imposed along the dryline convergence zone in the northern and northwestern BoB. However, the development formed a nonsevere and nonorganized linear system when the convergence zone was farther south of the dryline. In the postmonsoon cases, cyclogenesis was favored by squall-line MCSs with a north–south orientation over the BoB. All convective systems moved quickly, persisted for a long time, and contained suitable environments for developing low-level cyclonic mesovortices at their leading edges, which played an additional role in forming mesoscale convective vortices during cyclogenesis in the BoB.

Corresponding author address: Nasreen Akter, Department of Physics, Bangladesh University of Engineering and Technology, Zahir Raihan Rd., Dhaka 1000, Bangladesh. E-mail: nasreenphysics@yahoo.com

Abstract

Mesoscale convective systems (MCSs) are an essential component of cyclogenesis, and their structure and characteristics determine the intensity and severity of associated cyclones. Case studies were performed by simulating tropical cyclones that formed during the pre- and postmonsoon periods in 2007 and 2010 over the Bay of Bengal (BoB). The pre- (post) monsoon environment was characterized by the coupling of northwesterly (southwesterly) wind to the early advance southwesterly (northeasterly) monsoonal wind in the BoB. The surges of low-level warm southwesterlies with clockwise-rotating vertical shear in the premonsoon period and moderately cool northeasterlies with anticlockwise-rotating vertical shear in the postmonsoon period transported moisture and triggered MCSs within preexisting disturbances near the monsoon trough over the BoB. Mature MCSs associated with bimodal cyclone formations were quasi linear, and they featured leading-edge deep convection and a trailing stratiform precipitation region, which was very narrow in the postmonsoon cases.

In the premonsoon cases, the MCSs became severe bow echoes when intense and moist southwesterlies were imposed along the dryline convergence zone in the northern and northwestern BoB. However, the development formed a nonsevere and nonorganized linear system when the convergence zone was farther south of the dryline. In the postmonsoon cases, cyclogenesis was favored by squall-line MCSs with a north–south orientation over the BoB. All convective systems moved quickly, persisted for a long time, and contained suitable environments for developing low-level cyclonic mesovortices at their leading edges, which played an additional role in forming mesoscale convective vortices during cyclogenesis in the BoB.

Corresponding author address: Nasreen Akter, Department of Physics, Bangladesh University of Engineering and Technology, Zahir Raihan Rd., Dhaka 1000, Bangladesh. E-mail: nasreenphysics@yahoo.com

1. Introduction

The South Asian premonsoon (March–May) and postmonsoon (October–December) seasons are the transition periods between the summer (June–September) and winter (January–February) monsoons. A strong southwesterly (SW) prevailing wind transports enormous moist and warm air masses from the sea to the land in the boreal summer, while dry and cool northeasterly (NE) flows occur in the opposite direction (i.e., from the land to the sea) during the boreal winter (Ramage 1971; Das 1995). The Arabian Sea and the Bay of Bengal (BoB), which are two branches of the north Indian Ocean (NIO), play pivotal roles in the seasonal wind reversal of South Asian monsoons. The southern BoB experiences the southwest monsoon in early May (Wu and Zhang 1998; Mao and Wu 2007; Wu et al. 2012). The maximum annual sea surface temperature (SST) in the central BoB and the northward-propagating deep convection phase of the intraseasonal oscillations (ISOs) trigger an earlier monsoon onset in the BoB than in India (Jiang and Li 2011; K. Li et al. 2013; Yu et al. 2012).

The BoB is not only significant for the Asian monsoon onset but also offers a unique setting for tropical cyclone (TC) activities. The TCs over the BoB are confined within the monsoon transition periods, with a maximum frequency in October–November and a second maximum in May (McBride 1995; Harr and Chan 2005; Camargo et al. 2007; Kikuchi and Wang 2010; Akter and Tsuboki 2014, hereafter AT14). In the boreal summer, the location of the monsoon trough (MT) is well inland; the prevailing southwesterly winds and upper-level easterly winds create strong vertical shear that suppresses TC formation (Jeffries and Miller 1993; McBride 1995; Z. Li et al. 2013). Conversely, the bimodal TC activity over the BoB is modulated by the seasonal migration of MT locations that are usually in the northern and central BoB during the pre- and postmonsoon seasons, respectively (McBride 1995; AT14). Within the transition seasons, intraseasonal ISO phases often influence TC formation in the BoB (Kikuchi and Wang 2010; Kikuchi et al. 2012; Yanase et al. 2012). Both the seasonal MT position and ISO phases are associated with the formation of synoptic-scale tropical disturbances or cloud clusters that favor active cyclones (Gray 1998; Roundy and Frank 2004).

Recently, AT14 revealed that the northern position of the MT in the BoB and environmental convective inhibition (CIN) are mutually responsible for the decreased cyclone frequency in May (premonsoon), even though the BoB maintains higher SSTs that result in increased convective available potential energy (CAPE; Glickman 2000) compared with the postmonsoon season. In the premonsoon season, deep hot and dry air is advected from northwest India toward the BoB to provide environmental CIN that caps the boundary layer over the northwest BoB (Fig. 7 of AT14). Consequently, TC genesis is reduced because of suppressed convection. Therefore, seasonal environmental flow is an essential dynamical aspect that precedes TC genesis in the BoB. Ritchie and Holland (1999) and Yoshida and Ishikawa (2013) investigated five types of large-scale dynamical flow patterns associated with cyclone development: monsoon shear line (SL), monsoon confluence region (CR), monsoon gyre (GY), easterly wave (EW), and Rossby wave energy dispersion (RD) in the western North Pacific; the results demonstrated that the SL, CR, and GY patterns are related to the MT. Conversely, for the same basin, Lee et al. (2008) discussed the EW, SL, and CR patterns and three synoptic-scale flows: SW, NE, and combined NE and SW; the results showed that all of the patterns are related to the monsoon, except for the EW. Furthermore, Lee et al. (2008) noted that mesoscale convective system (MCS; Houze 2004) activities are linked with such synoptic flows.

Many studies have acknowledged that large-scale or synoptic-scale flows are not the only major contributors to TC genesis. Individual MCSs that are associated with a preexisting tropical disturbance and cyclonic mesoscale convective vortices (MCVs) that develop in the stratiform precipitation region near the middle troposphere are also fundamental precursors for cyclogenesis (e.g., Zehr 1992; Harr et al. 1996; Bister and Emanuel 1997; Ritchie and Holland 1997; Gray 1998; Dunkerton et al. 2009; Houze 2010). Two processes are hypothesized to explain lower-tropospheric vortices produced by midtropospheric MCVs: the top-down and bottom-up paradigms for cyclogenesis. The first paradigm emphasizes the downward advection of MCVs in a moist environment (Emanuel 1993; Bister and Emanuel 1997), where greater penetration occurs by the merging of individual MCVs and cyclonic low-level vorticity is further enhanced (Simpson et al. 1997; Ritchie and Holland 1997). The second paradigm incorporates cycles of deep, moist convective activity called vortical hot towers (VHTs), where the lower-tropospheric vortices are formed within the embryonic environment of MCVs (Hendricks et al. 2004; Reasor et al. 2005; Montgomery et al. 2006; Braun et al. 2010). The combined effect of the upscale growth of cyclonic vortices and their integration contributes to the development of the TC vortex.

The structure and characteristics of MCSs and the associated MCVs that form in the BoB during pre- and postmonsoon TCs are vital to the seasonal bimodal cyclogenesis process. Few studies have focused on the different types of MCSs and their contributions to precipitation in South Asia during the premonsoon and monsoon seasons (Houze et al. 2007; Romatschke et al. 2010; Romatschke and Houze 2011a,b). Previous studies have shown that the BoB experiences large systems with extremely large stratiform regions in both the premonsoon and monsoon seasons. During the premonsoon period, MCSs are primarily related to depressions and exhibit weak diurnal cycles. Moreover, no investigation on postmonsoon MCS characteristics has been conducted. Specifically, synoptic-scale flow patterns and related MCSs for TC genesis have not been identified in the BoB. Therefore, the objective of the present study is to assess how seasonal variations in the synoptic-scale flows over the BoB determine the different types of MCS formations and the relevant vorticity generation within the BoB that contributes to the seasonal bimodal distribution of tropical cyclogenesis. To achieve this objective, simulations of the structural characteristics of MCSs during cyclogenesis in the BoB are the only viable option because observational data are limited.

2. Model specifications and data used

The Advanced Hurricane Weather Research and Forecasting (WRF) Model (AHW) (version 3.3.1), which is derived from the Advanced Research version of the WRF (ARW) Model (Davis et al. 2008; Skamarock et al. 2008), is used to simulate pre- and postmonsoon cyclones for examining the MCSs that are associated with bimodal cyclogenesis. A Lambert projection map is utilized with two-way nested domains; the grid spacing is 12 km for the outer domain and 4 km for the innermost domain (Fig. 1b). Davis et al. (2008) showed that 12- and 4-km grid spacing provide the most accurate forecasts of storm position and intensity. The parent domain (D1) consists of 861 × 595 grid points, while the inner nest (D2) has 1064 × 996 grid points. The 28 terrain-following vertical levels, where the top level is 50 hPa, are used. The Kain–Fritsch cumulus parameterization, which predicts deep and shallow convection using a mass flux approach (Kain 2004), is only applied in D1 with a 5-min time step. The Ferrier microphysics scheme, which includes a prognostic mixed-phase representation of changes in water vapor and condensate and considers cloud water, rain, cloud ice, and precipitating ice (Ferrier et al. 2002), is used in both domains.

Fig. 1.
Fig. 1.

(a) Cyclone frequency and intensity during 2001–12 based on the JTWC data. Overlapping intensities with the same value are labeled with a lighter color to avoid confusion with the number of TCs. The specific humidity (shaded, g kg−1) and temperature anomalies (solid contours, 2-K intervals) calculated from the background area in the figure and the horizontal winds (vector, m s−1) at 850 hPa are displayed for (b) May 2007, (c) November 2007, (d) May 2010, and (e) October 2010 using the monthly mean NCEP–CFSR reanalysis data. The tracks represent the TCs that formed in 2007 and 2010. The areas of the model’s parent domain (D1) and nested domain (D2) are indicated in (b).

Citation: Monthly Weather Review 143, 9; 10.1175/MWR-D-14-00260.1

The Monin–Obukhov scheme (Monin and Obukhov 1954) is used for the surface layer physics; the Noah land surface model, which considers four soil temperature and moisture layers (Chen and Dudhia 2001), is also used. Moreover, the Yonsei University (YSU; Hong et al. 2006) planetary boundary layer (PBL) scheme is selected because it is a first-order closure scheme that explicitly treats entrainment processes at the top of the PBL. The asymptotic entrainment flux term is proportional to the surface flux in the inversion layer (Noh et al. 2003). The Rapid Radiative Transfer Model (RRTM; Mlawer et al. 1997) and the Dudhia scheme (Dudhia 1989) are selected for the longwave and shortwave radiation calculations, respectively.

In this study, all of the physical parameterizations are chosen according to Raju et al. (2011) and Osuri et al. (2012), who customized the parameters for simulating cyclones over the NIO using the WRF Model.

Additionally, the AHW Model includes a one-dimensional ocean mixed layer model and modified surface flux and drag formulations for high-wind conditions over the ocean. The ocean mixed layer model is based on Pollard et al. (1973), and it attempts to capture the negative feedback of the SST on TCs (i.e., the decrease in SSTs due to the passage of a TC). For this study, a 30-m ocean mixed layer depth is selected for initializing the model over the BoB; details are available in Kumar et al. (2011). Here, Donelan’s drag formulation (Donelan et al. 2004), which predicts weaker surface friction under high wind conditions compared with the Charnock formulation, is selected. The surface enthalpy flux is calculated using the formulation of Garratt (1992).

The atmospheric initial and boundary conditions are derived from National Centers for Environmental Prediction (NCEP) Final Analysis (FNL) Operational Global Analysis data at 6-h intervals with a resolution of 1° × 1°. The NCEP high-resolution real-time global sea surface temperatures (RTG_SSTs) at a 0.083° resolution are imposed for the daily sea surface temperatures. Two premonsoon TCs—Akash (2007) and Laila (2010)—and two postmonsoon TCs—Sidr (2007) and Giri (2010)—are considered in this study by running four simulations using the aforementioned experimental setup, with no bogussing, for the TC initialization. The duration of each simulation is 4 days (Akash: 0000 UTC 11 May–0000 UTC 15 May 2007, Laila: 0000 UTC 15 May–0000 UTC 19 May 2010, Sidr: 0000 UTC 8 November–0000 UTC 12 November 2007, and Giri: 0000 UTC 18 October–0000 UTC 22 October 2010). A maximum surface wind speed of 34 kt or ~17.5 m s−1 in Joint Typhoon Warning Center (JTWC) best track data is considered as the condition of its formation in this study. The maximum wind speed beyond ~17.5 m s−1 is referred to as TC intensification by self-sustaining mechanism (Zehr 1992; McBride 1995). On the basis of JTWC data, each case has taken less than 1.5 days to turn from tropical depression [maximum surface wind speed of 15 kt (1 kt = 0.5144 m s−1)] to the cyclogenesis. So the initial time for simulation, including the model spinup time, is kept at least 2.75 days prior to the commencement of cyclogenesis (Table 1). As the SST data are daily, the initial start time of each simulation is 0000 UTC and the total time from initiation of the model simulation to the formation of all TCs is 2.75 days, except for Sidr (2007), which is 12 h longer than others. The model output is acquired at a 30-min interval.

Table 1.

Pre- and postmonsoon cyclones.

Table 1.

The simulated results are verified by NCEP Climate Forecast System Reanalysis (CFSR) 6-hourly data, which have a resolution of 0.5° × 0.5°. The data have 64 hybrid sigma-pressure vertical levels, with a top pressure of ~0.266 hPa (Saha et al. 2010). In addition, monthly mean NCEP–CFSR data are used for analyzing the climatological conditions of the BoB, and JTWC data are utilized to verify the simulated position and intensity of the TCs. NCEP reanalysis and NOAA daily interpolated outgoing longwave radiation (OLR) data with a 2.5° × 2.5° resolution are also used for 2001–12.

3. Bimodal cyclones

a. Overview

Only 7% of global TC formations occur in the NIO (Neumann 1993); however, TCs in the region are devastating because of the funnel-shaped low-lying coastal areas. The pre- and postmonsoon TCs that formed in the BoB in 2007 and 2010 are considered suitable cases for resolving the seasonal differences in MCS formations during cyclogenesis. The most important criteria for selecting these years are as follows: (i) cyclones between 2001 and 2012 have the most sufficient data, including damage information; (ii) pre- and postmonsoon TCs from the same year avoid interannual variability, such as El Niño–Southern Oscillation (ENSO), between the two seasons; and (iii) very severe cyclonic storms [i.e., maximum wind speeds of at least 64 kt based on the Indian Meteorological Department (IMD) scale] occurred. According to the JTWC data from 2001 to 2012, a total of 36 cyclones formed over the BoB; 30.6% formed during the premonsoon season, and 69.4% formed during the postmonsoon season (Fig. 1a). Among these cyclones, only 6 (16.7% of the total) in the premonsoon season and 3 (8.3% of the total) in the postmonsoon season had wind speeds ≥64 kt.

The criteria mentioned above indicate that the cyclones of 2007 and 2010 are the most suitable for this study. Moreover, both years experienced La Niña, which is the cool phase of the ENSO cycle. The TC features, which were collected from the JTWC, are sequentially arranged in Table 1 (JTWC 2007, 2010), except for the casualty information. The damage estimates were obtained from Begum et al. (2013).

Akash (2007) and Laila (2010) both formed almost at the center of BoB (Figs. 1b and 1d) during May, when important monsoon evolutions occurred in the southeast BoB. Both TCs and the monsoon advancement in the BoB are influenced by ISO; however, TCs and monsoons do not occur in unison because their activities are modulated by the interannual variability (phase and intensity) of ISO (Fosu and Wang 2015). Z. Li et al. (2013) argued that April–May is the perfect premonsoon period because the oceanic and atmospheric conditions change dramatically in preparation for the monsoon onset. In connection to the greater ocean heat content, the first branch of northward-propagating ISOs is associated with the monsoon onset over the BoB, and stronger ISO intensity in April–May is a favorable environmental condition for cyclone intensification. According to a map of the southwest monsoon of IMD for 2007 (available in IMD website), the onset occurred over the east central area of the BoB on 21 May. Before that, a tropical depression formed in the same region on 11 May, which moved north-northeastward and intensified as tropical cyclone Akash on 13 May; finally, it crossed the Bangladesh coast on 15 May (Bhatia and Rajeevan 2008; Mazumdar et al. 2008). In 2010, IMD reported that the southwest monsoon set up over the southeast BoB on 17 May, 3 days prior to the normal date as a consequence of a severe cyclonic storm Laila (depression to landfall: 16–21 May 2010) over the BoB (India Meteorological Department 2010). Both Akash (2007) and Laila (2010) developed before the monsoon onset in their respective locations in BoB. Favorably, in both cases, cross-equatorial flow across the southern BoB was accelerated by the TC vortices. Thus, the cyclones in this study are representative premonsoon TCs that correspond well with the locations of MTs over the ocean in both cases (details in section 4a).

The synoptic conditions for each cyclone are also illustrated in Figs. 1b–e using the monthly mean specific humidity and temperature of the NCEP–CFSR reanalysis data. In the premonsoon season (especially May), the BoB is influenced by two different air masses: a hot, dry air forcing from northwestern India and a warm, moist air transporting from the southwest of the bay. The warm, moist air masses are wedged under deep hot and dry air that extends from 950 to 600 hPa (AT14). Conversely, in the postmonsoon (~October–November) cases, the entire environment of the BoB is nearly uniform regarding temperature and humidity. The track of each cyclone (i.e., Akash, Sidr, Laila, and Giri) that originated as a depression (i.e., a maximum wind speed of 15 kt) is also displayed in the corresponding figures.

b. Model verification

The data collected from the 12-km simulation are compared with the observed (JTWC) and NCEP–CFSR reanalysis (~55-km resolution) data for the TCs in 2007 and 2010. The surface pressure, relative vorticity, and horizontal wind at 850 hPa are illustrated in Fig. 2 at the genesis time of each TC (Table 1) to demonstrate the evolution of the TCs. The simulation results nearly accurately depict the position and intensity of the TCs. The model and reanalysis (JTWC) data suggest a minimum surface pressure of 990 and 985 (997) hPa near the vortex center for Akash (2007). Moreover, the simulated minimum surface pressures for Sidr (2007), Laila (2010), and Giri (2010) are 1003, 1005, and 1007 hPa, respectively; the reanalysis (JTWC) data suggest that the minimum surface pressure is 1006 (~996) hPa for these three TCs. The wind rotates around the TC center with a velocity greater than 16 m s−1 for all simulated cyclones, except for Giri (wind speed of ~11 m s−1). The relative vorticity is on the order of 10−4 s−1 in all simulations, and it signifies strong rotation at the TC centers. Because the data resolution is lower, the values of the vorticity are less intense in the reanalysis data than in the simulations.

Fig. 2.
Fig. 2.

The surface pressure (contours, 2-hPa intervals), wind velocity (vector, m s−1), and vorticity (shaded, ×10−5 s−1) at 850 hPa for the (a) simulations and (b) NCEP–CFSR reanalysis data at each TC genesis time (see Table 1). The black dots are the genesis points for each TC using the JTWC data.

Citation: Monthly Weather Review 143, 9; 10.1175/MWR-D-14-00260.1

Cyclogenesis depends on the formation of tropical convection or MCSs (Gray 1998; Zehr 1992), which are initiated by several important factors (e.g., the availability of CAPE and low-level convergence) (Protat and Lemaitre 2001). The surface-based environmental CAPE within the BoB is, therefore, verified with the reanalysis data two days prior to each genesis. The simulated CAPE, including the seasonal variability, over the ocean is similar to the reanalysis data (Fig. 3). The CAPE is >2500 J kg−1 during the premonsoon cases because of higher SSTs (Sasamal 2007; Alappattu and Kunhikrishnan 2009). In contrast, the CAPE is lower (<1800 J kg−1) in the postmonsoon environment. The seasonal variability in the instability is consistent with the result of AT14 using 20-yr-average seasonal data. The simulated surface wind and its convergence toward the genesis location are also analogous to the reanalysis data.

Fig. 3.
Fig. 3.

The surface-based CAPE (shaded, J kg−1) for the (a) simulations and (b) NCEP–CFSR reanalysis data. The black dots indicate the genesis points of each TC using the JTWC data. The display time is 2 days prior to the genesis of each TC.

Citation: Monthly Weather Review 143, 9; 10.1175/MWR-D-14-00260.1

4. Hierarchy of cyclogenesis in the BoB

Recent studies have suggested that synoptic-scale tropical disturbances or cloud clusters (~700 km) are preconditions for initializing TCs; these features are possibly formed by monsoonal troughs, the ITCZ, easterly waves, or equatorial waves (Gray 1998). A preexisting disturbance is characterized by cyclonic relative vorticity in the lower troposphere, where large-scale convergence (wind surge) can assist in the development of an MCS (~250 km; Gray 1998) that has deep convective cells, stratiform clouds, and precipitation (Houze 2004). Initially, an MCS may be initiated by one or more isolated VHTs (~10 km; Montgomery et al. 2006). Following the weakening of the VHT, a stratiform region may appear in the MCS over time (Houze 2010). Moreover, latent heat release in the stratiform region accelerates midtropospheric warming and evaporative cooling below. The combination of these thermodynamic profiles with updrafts above and downdrafts below can lead to the development of an MCV that is typically on the order of 100 km in diameter (Gray 1998). The mature MCS is then characterized by both cyclonic vorticity in the convective-scale VHTs and a midtropospheric MCV during the middle of the life cycle. In the later stages of the MCS, the VHT may no longer remain, but the MCV persists. If the cold pool under the MCV can be removed or warmed and the midlevel vorticity maximum structure change to be a low-level maximum, then the wind-induced surface heat exchange (WISHE; Emanuel et al. 1994) feedback mechanism may become dominant.

a. Cloud clusters

During the early genesis period, which is defined 2.5 days prior to each cyclogenesis event, the environment is determined by the 6-h-averaged moisture flux, reflectivity, and streamlines at 925 hPa from the 12-km resolution model grid (Fig. 4). The zero line of the average zonal vertical wind shear between 850 and 200 hPa indicates a low shear region or MT passing over the BoB (McBride 1995; AT14). Within the trough region, cloud clusters are organized and spread to at least ~700 km, which is consistent with the cluster length discussed by Gray (1998). Two types of synoptic flow patterns in the BoB are related to the seasonal development of TCs. In the premonsoon cases, a combination of SW and northwesterly (NW) winds appear at low levels within the BoB. The NW flow advects dry air from the arid regions of southwestern Asia and western India to the BoB up to 500 hPa; however, the equatorial SW flow from the Arabian Sea carries large amounts of moisture from the surface to 875 hPa over the BoB (also AT14). In the case of Laila (2010), NW wind turns along the boundary of the coast and flows to the north and northwest over the BoB. After entering the bay, the NW wind gradually moistens (humidity ~7 g kg−1 in the north BoB to ~15 g kg−1 in the south BoB) at the low levels (also Fig. 1) as the influx of water vapor intensifies with higher SSTs during this period (29.8°C on average, AT14); however, the total moist air masses within the NW flow are smaller than those define by the region of SW winds.

Fig. 4.
Fig. 4.

The 6-h-averaged moisture fluxes (shaded, m s−1), reflectivities (contours, 20-dBZ intervals), streamlines at 925 hPa, and the zero zonal wind shear between 850 and 200 hPa (dashed line, m s−1). The averages are calculated from the time 0600 UTC 11 May for Akash, 0600 UTC 15 May for Laila, 1800 UTC 8 Nov for Sidr, and 0600 UTC 18 Oct for Giri (i.e., 2.5 days prior to each genesis). The stars are the positions of the depressions (max wind speed of 15 kt) of each TC.

Citation: Monthly Weather Review 143, 9; 10.1175/MWR-D-14-00260.1

In the postmonsoon cases, the SW and NE winds coexist from low levels to the 600-hPa level. In this period, the SW monsoonal winds continue to retreat south of the BoB, while the onset of easterly or NE winds occurs in the northern extent of the BoB. The NE winds gradually advect more low-level moisture to the BoB relative to the surroundings; however, the average moisture is less than that in the premonsoon environment. These strong directionally distinct dynamic and thermodynamic environmental forcings in the BoB, especially the SW wind surges in the premonsoon cases and the NE wind surges in the postmonsoon cases, supply low-level moisture and support the initiation of mesoscale deep convection within cloud clusters. In the next section, the characteristics of the MCSs formed in the BoB and the associated vortices related to the active cyclones during the pre- and postmonsoon seasons are discussed using the simulation at 4-km resolution.

b. Mesoscale convection

1) Early stage convection

The initial MCS stage (widespread regions of ~35-dBZ reflectivity), which formed nearly 2.5 days (±6 h) before the cyclogenesis of each TC, is shown in Fig. 5. The time of early MCS stage is defined as “t” in each case of this study. Here, t is 0530 UTC 11 May for Akash (2007), 0800 UTC 15 May for Laila (2010), 0600 UTC 8 November for Sidr (2007), and 1130 UTC 18 October for Giri (2010). The environmental low-level horizontal wind, the water vapor mixing ratio, and the temperature anomalies at time t are also illustrated in Fig. 5. For Akash (2007), an MCS containing a small group of cells initially appears within an environment of significant temperature and humidity variations on the synoptic scale. Temperature (moisture) is gradually increasing (decreasing) from the east of MCS toward the west by the values from 23° to 28°C (21–12 g kg−1). A sharp negative horizontal moisture gradient [~2 g kg−1 (100 km)−1] associated with positively varying temperature [~1° (100 km)−1] to the west of the MCS indicates a synoptic dryline boundary (Schaefer 1986). The MCS forms in the low-level convergence when the air transported via strong SW winds (>12 m s−1) moves into relatively dry NW winds (<6 m s−1) along or near the dryline boundary. Similarly, the early MCS of Laila (2010) originates when air that is transported by intense SW winds enters the BoB and converges with air advected by weak NW winds near 4°–6°N. Because of the limitation of the domain area (D2), the SW wind is not visible in the figure; however, it is confirmed by the simulated result in D1 (also in Fig. 4). In contrast to the MCS of pre-Akash (2007), the MCS of pre-Laila (2010) emerges as a linear convective system (Houze et al. 1990) and forms in the southernmost extent of the BoB within a region of uniform temperature (22°C) and humidity (17 g kg−1) at the low level.

Fig. 5.
Fig. 5.

Water vapor mixing ratio (shaded, g kg−1), reflectivities (contours, 20-dBZ intervals), temperature anomalies (dashed line, 1-K intervals) calculated from the average of domain D2, and horizontal wind (vector, m s−1) at 925 hPa at time t (time t for each case is mentioned in the text). The stars in all panels are the sounding points.

Citation: Monthly Weather Review 143, 9; 10.1175/MWR-D-14-00260.1

Furthermore, both postmonsoon cases are characterized by a prominent north–south shear line in which relatively strong NE winds (>10 m s−1) create a horizontal wind shear line across which an abrupt change in the horizontal wind component occurs (Glickman 2000). The convective systems in these cases are tropical squall lines because they form along the confluence boundary of the shear line and contain a narrow quasi-linear band of active convection—either continuous or discontinuous—that is perpendicularly aligned with the low-level horizontal environmental shear (Barnes and Sieckman 1984; LeMone et al. 1998; Wang and Carey 2005). No remarkable synoptic-scale temperature differences are observed in the postmonsoon environment; however, air advected by the northeasterly wind is slightly cooler (<−1 K) than the surroundings.

Environmental soundings (Fig. 6) and hodographs (Fig. 7) at each MCS formation time t are obtained for the areas where moisture advection triggers MCS initiation. In all cases, the soundings indicate unstable lapse rates (Doswell et al. 1985), with lifted index values (LIs; Galway 1956) ≤−3 and high kelvin index (George 1960) values (greater than 35). These values permit the environment to develop a strong convective system with heavy rain. The level of free convection for a surface parcel is between 950- and 975-hPa level, which is very close to the lifting condensation level; deep convection is, therefore, more likely (Rasmussen and Blanchard 1998; Craven et al. 2002). More than 75% of the moisture in all cases is between the surface and ~6 km; however, dry air is noticeable in the environment of Akash (2007) at 500 hPa. Air parcels are much warmer and more buoyant than the environment in the presence of moderate to high CAPE (>1064 J kg−1) and CIN of ~0 J kg−1. Among all the cyclones, the premonsoon environment of Akash (2007) is extremely unstable (LI = −10 and CAPE = 4968 J kg−1), and there is the potential for initiating severe thunderstorms (Bluestein 1993). The severe dryline convection (Weston 1972; Ziegler and Rasmussen 1998; Hane et al. 2002; Murphey et al. 2006) during the premonsoon period may be enhanced because the MCS of Akash (2007) develops along the dryline boundary within the BoB.

Fig. 6.
Fig. 6.

Skew T–logp sounding and wind barbs (kt) at the location of each TC (the stars in Fig. 5). The long dashed, dot–dot–dashed, solid, and short dashed lines represent the dewpoint sounding, temperature sounding, parcel trace, and relative humidity sounding, respectively.

Citation: Monthly Weather Review 143, 9; 10.1175/MWR-D-14-00260.1

Fig. 7.
Fig. 7.

Hodographs at the sounding points described in Fig. 6. The numbers represent the height in km.

Citation: Monthly Weather Review 143, 9; 10.1175/MWR-D-14-00260.1

In the hodographs, the vertical wind shear from the surface to 6 km above ground level (AGL) suggests a moderate to strong total shear of 15–22 m s−1 and a bulk shear of 8–10 m s−1, which support the formation of ordinary cells or supercells (Weisman and Klemp 1982; Bunkers 2002). Differences are noted in terms of the directional shear that veers clockwise with height in the premonsoon cyclone cases and that turns counterclockwise with height (backing wind) in the postmonsoon cases. The positive (negative) values of environmental helicity, which are noted in the skew T–logp plot in Fig. 6, also provide evidence of the veering (backing) winds with height. In the premonsoon scenario, the wind veers from SW at the surface to NW at 3 km during Akash (2007) and from westerly to NW during Laila (2010). A backing wind from nearly NE at the surface to northerly aloft is observed for Sidr (2007), whereas the wind is westerly to NE for Giri (2010). The veering wind with height is an indication of warm air advection, whereas backing wind profiles are indicative of cold air advection (Bluestein 1993) and dynamical sinking air.

2) Mature stage convection

The mature MCSs that are associated with high reflectivity of at least 50 dBZ, are displayed in Figs. 8a,d for the premonsoon TCs. The hatched area in the figure shows the position and size of the MCSs after 6 h, which indicate the direction that systems are moving. Vertical cross sections approximately perpendicular to the systems are shown in Figs. 8b,c for Akash (2007) and in Figs. 8e,f for Laila (2010). These cross sections contain the reflectivities, the cross section parallel to the system-relative winds (i.e., approximately system-relative meridional wind component), the water vapor mixing ratios, the updrafts, the downdrafts, the vertical vorticities, and the potential temperature perturbations. In the mature stage, the MCS of Akash (2007) is organized as a bow echo (i.e., it presents a bowed thunderstorm outline) (Fujita 1978), and it becomes an increasingly comma-shaped asymmetric echo with a length of ~300 km (Fig. 8a). Intense convection evolves to the south of the system and is accompanied by a convective comma head to the north. The strong system-relative low-level moist (humidity >16 g kg−1) inflow from the south side of the system instigates deep convection to the south of the system and stratiform rain developing over the region to the north. The vertical height of the convection is approximately 15 km.

Fig. 8.
Fig. 8.

Premonsoon TCs: (a) mature-stage MCS (shaded; dBZ), 850-hPa wind velocity (m s−1), and MCS after 6 h (hatched area) for Akash (2007). The yellow line is the approximate RIJ. The vertical cross section along AB in (a) indicates the (b) vertical reflectivities (shaded; dBZ), downdrafts (thin contours; 0.2-m s−1 interval), water vapor mixing ratio (thick contours; >16 g kg−1), and cross-sectional parallel-system-relative wind with vertical velocity (vectors; m s−1). (c) Potential temperature anomalies for a 6° × 4° area for the system-wide values (shaded; K) and updrafts (thick solid contours; 2 m s−1 interval); the dashed (thin solid) contours of 1 × 10−3 s−1 represent the vertical cyclonic (anticyclonic) vorticity. (d)–(f) As in (a)–(c), but for Laila (2010). The wide arrows in (a) and (d) are the approximate MCS directions.

Citation: Monthly Weather Review 143, 9; 10.1175/MWR-D-14-00260.1

The most recognizable feature in a bow echo is the rear-inflow jet (RIJ; Weisman 1993), which is elevated air that descends from the rear to the front of the line of convection and supplies cool, dry midlevel air that aids in the production of the convective- and system-scale downdrafts. In this case, the 3–4-km RIJ from the north or northwest descends behind the line of convection (Fig. 8), creates a downdraft, and ultimately drives the system to the south at an average speed of ~9 m s−1. Because the midtroposphere is dry (Fig. 6), evaporative cooling creates negative buoyancy that further accelerates the downdraft toward the surface. The downward-moving air that typically spreads out in all directions after reaching the ground may produce strong, damaging winds (Fujita 1978). The near-surface reflectivity gradient in Fig. 8b indicates straight-line winds that rush downward within the core of the storm and that spread at a velocity of ~20 m s−1.

Because of the tropical moist environment, the downdraft is ≤−1.6 m s−1, which is less than that of midlatitude convection (Xu and Randall 2001). Consequently, a small low-level cold pool, with a potential temperature perturbation of ≤−4 K, is produced; however, this value is sufficient for tropical convection (COMET Program/UCAR 1999). The higher temperature associated with the updraft indicates latent heat release by the convection. The updraft is quite erect and remains at the leading edge of the convective line with a maximum intensity of ~12 m s−1. This finding may be a result of the balance between the vorticities associated with the ambient low-level vertical wind shear (18 m s−1 at 0–3 km AGL) and the low-level cold pool that forms beneath the convection (Weisman and Rotunno 2004). The intense vertical vorticity (maximum of 4 × 10−3 s−1) associated with the updraft indicates the presence of a strong mesocyclone near the apex of the comma-shaped echo. The outflow boundary that propagates ahead of the line initiates new cells downwind and helps the system persist for ~16 h (not shown).

Alternatively, a quasi-linear type of less-organized convective cells along the leading edge and a trailing stratiform rain region appear during the mature MCS stage for Laila (2010). The mature system has a height of ~15 km and a width of ~250 km. The cross-section parallel-system-relative winds reveal the southerly inflow that supplies low-level moisture to the updrafts. The updraft and downdraft are 4 and −0.6 m s−1, respectively. The horizontal vorticity generated by both the system cold pool and the ambient low-level wind shear (13 m s−1 at 0–3 km AGL) further initiate mesovortices (2 × 10−3 s−1) at low levels via tilting (Holton 1992). Moreover, the system is moving with an average velocity of ~9 m s−1 toward the northwest due to the southeasterly wind, which is the deflected SW wind from the Malay Peninsula and Sumatra. New cells form ahead of the leading line; however, these cells are not well organized. The system continues until it dissipates after ~14 h. Within a few hours, many other scattered organized and nonorganized convective systems form and merge together.

Figure 9 represents parameters similarly to those in Fig. 8 to demonstrate the postmonsoon MCSs; in both cases, the MCSs are squall lines (Houze and Betts 1981; Gamache and Houze 1982) with leading-edge convective precipitation that is trailed by a narrow region of stratiform rain. The systems are very long, with a length of ~500 km and a maximum height of 15 km. The line of convection forms and intensifies when the system-relative low-level moist westerly inflow ahead of the line converges with the easterly inflow behind the squall line. The updraft is upright within 100–300 hPa due to the proper balance between the moderate instability and the moderate to strong low-level vertical shear (>12 m s−1 at 0–3 km AGL). The storms are linearly organized along the boundary of the convergence line; the storm outflows create a weak (<−3 K) and shallow cold pool along the rear side of the system. Tropical squall lines can easily develop in low-shear and low-LFC environments, and they are triggered by weaker cold pools that are 2–4 K colder than the ambient temperature (COMET Program/UCAR 1999). The vertical cross sections in Figs. 9b,e support the concept that new cells develop along the leading edge of the system’s gust front and then propagate rearward to contribute to the growth of the trailing stratiform precipitation region. Low-level mesovortices and the updrafts are observed in Figs. 9c,f. A strong postmonsoon low-level easterly jet transports the systems at an average velocity of 8 m s−1 and promotes swelling at the midpoint of the line. The life spans of Sidr (2007) and Giri (2010) are ~18 and ~20 h, respectively.

Fig. 9.
Fig. 9.

As in Fig. 8, but for the postmonsoon TCs in grayscale.

Citation: Monthly Weather Review 143, 9; 10.1175/MWR-D-14-00260.1

c. Mesoscale convective vortices

Multiscale vortex formation and interaction occur during cyclogenesis. The distributions of vertical cyclonic vorticities on the order of 10−3 s−1 (as a function of time) are displayed at different levels (i.e., 925, 500, and 200 hPa) in Fig. 10 for both pre- and postmonsoon cases. In this study, all of the systems have quasi-linear convective systems (QLCSs) and several meso-γ-scale (2–20 km; Orlanski 1975) vortices (“mesovortices”) that form at low levels along the convergent outflow boundary (Figs. 8 and 9). The horizontal scale of the cyclonic mesovortices is 10–40 km in the premonsoon cases and only 5–10 km in the postmonsoon cases. Several studies (Weisman and Trapp 2003; Sippel et al. 2006; Wheatley and Trapp 2008; Atkins and St. Laurent 2009a) have noted similar mesovortices at low levels along the leading edge of convective systems, particularly in QLCSs (i.e., squall lines and bow echoes). The low-level mesovortices that develop in QLCSs depend on the environmental vertical wind shear within 0–2.5 km AGL (Weisman and Trapp 2003) and the presence of a positive Coriolis force (Trapp and Weisman 2003).

Fig. 10.
Fig. 10.

Distribution of the horizontal cyclonic vortices (contours of 1 × 10−3 s−1) for the time period from t + 1 to t + 13 h in 1-h intervals at (a) 925, (b) 500, and (c) 200 hPa for each TC. In (a)–(c) the vectors are the horizontal winds (m s−1) at t + 13. (d) The vertical cross section of the updraft (shaded; m s−1) and cyclonic vorticity (contours; 0.4 × 10−3 s−1) along the point located in (c) for each case.

Citation: Monthly Weather Review 143, 9; 10.1175/MWR-D-14-00260.1

For early stage MCSs in both seasons, a vortex couplet (i.e., cyclonic in the north and anticyclonic in the south) originates due to the tilting of the crosswise horizontal vorticity in the downdraft (not shown); this finding is consistent with observations of mesovortices in QLCSs (i.e., Trapp and Weisman 2003; Wheatley and Trapp 2008; Atkins and St. Laurent 2009b). These low-level cyclonic mesovortices further strengthen the circulation by merging with new vortices and stretching the planetary vorticity (Trapp 2013). The vertical structure of the vorticity in Figs. 8 and 9 indicates that cyclonic vortices intensify and extend upward during the mature stage, while anticyclonic vortices tend to weaken or diminish in the mature stage. Figure 10 also shows that the height of the low-level cyclonic vortices extends through the midtroposphere to the upper troposphere; intense vorticity occurs within the midtroposphere. A few cyclonic vortices extend to the tropopause (~15 km), and are associated with strong updrafts (>10 m s−1) between 100 and 300 hPa. These convective towers are highly vortical in nature within their core, with a maximum vertical vorticity of >1 × 10−3 s−1. Convective tower examples for each case are shown in Fig. 10d. These towers are characterized and defined as VHTs, which are fundamental coherent structures in the tropical cyclogenesis process (Hendricks et al. 2004; Montgomery et al. 2006; Houze 2010; Braun et al. 2010) and in the TC intensification process (Van Sang et al. 2008; Shin and Smith 2008; Montgomery et al. 2009). In this study, VHTs and low-level mesovortices within QLCSs that form over the BoB during cyclogenesis are essential building blocks of a cyclone vortex. However, the detailed structure and formation of VHTs and their interactions with mesovortices are not analyzed in this study.

Cyclonic vortices increase in both size and intensity with altitude as the individual convective cells embedded in the QLCS mature. Moreover, the vortices weaken as the cells decay and propagate rearward with respect to the leading edge of the line. Therefore, the environment behind the convective lines becomes enriched with vorticity over time, particularly within the midtroposphere because of the presence of the maximum vorticity. This phenomenon is signified through counterclockwise rotation of the wind at t + 13 h in Fig. 10. Several studies have also noted broadening cyclonic circulations (on the order of 100 km) in the form of a cyclonic mesovortex or line-end vortex at the end of QLCSs (Trapp and Weisman 2003; Wheatley and Trapp 2008; Atkins and St. Laurent 2009a).

In this study, the vortex circulation is more prominent in the premonsoon period than in the postmonsoon period. The dominant cyclonic vortex persists beyond the life cycle of the QLCS as a MCV, which is a deep column of cyclonic vorticity. A warm and moist boundary layer over the ocean further supports deep, moist convection near the MCV. Figure 11 shows the MCSs, the 500-hPa geopotential heights, and wind speeds for the pre- and postmonsoon TCs in 2007 at three times: the initial convection at time t, the deep moist convection that gradually organizes within the MCV, and an intensifying tropical storm that is represented by decreasing geopotential heights. Latent heat released aloft due to convection results in low pressure at levels below, which increases the magnitude of the cyclonic surface winds around the TC center.

Fig. 11.
Fig. 11.

Reflectivity at 925 hPa (shaded, dBZ) and the geopotential height (shaded, m) along with the horizontal wind (vectors, m s−1) at 500 hPa for TCs in 2007.

Citation: Monthly Weather Review 143, 9; 10.1175/MWR-D-14-00260.1

5. Discussion and summary

Cyclogenesis in the BoB is unique because the location of the MT and vertical shear primarily determine and limit cyclonic activity during the pre- and postmonsoon seasons instead of in the monsoon season (McBride 1995; AT14). The TCs that formed in the premonsoon (Akash and Laila) and postmonsoon (Sidr and Giri) environmental conditions in 2007 and 2010 were simulated using the WRF Model. Nested domains with grid resolutions of 12 and 4 km were used to examine details of MCSs and the environments associated with bimodal TC genesis. The simulations were validated with observations.

a. Synoptic features for cyclogenesis

In contrast to the different types of large-scale or synoptic-scale flow patterns observed during cyclogenesis in the western North Pacific basin (Ritchie and Holland 1999; Lee et al. 2008), only a NW–SW flow pattern was found in both premonsoon cases; the NW wind from northwestern India advected deep hot and dry air to the trough region in the BoB, while the SW wind advected shallow moist and warm air from the Arabian Sea. The SW flow during May was associated with the early advance monsoon flow in the southeast BoB and accelerated toward the preexisting low-level disturbance (i.e., low pressure) near the MT over the ocean. Notably, during the boreal summer, the MT is located over the Indian subcontinent (20°–25°N) (Wang 2006), and strong SW flow from the Arabian Sea traverses the entirety of South Asia.

However, a combination of SW and NE winds was identified in the postmonsoon cases. The SW winds retreated from the southern BoB (Fig. 4) at the end of the monsoon season. The NE winds encountered the BoB from the north and northeast and transported low-level moisture to the BoB after being blocked by the north–south-oriented mountains and converging along the east coast of the BoB, particularly near the Malay Peninsula (Chang et al. 2005; Akter and Tsuboki 2012). Therefore, the pre- (post) monsoon environment is characterized by the coupling of NW (SW) wind with the early advance SW (NE) monsoonal wind in the BoB. The dynamical forcing of relatively intense (>10 m s−1) SW (NE) wind surges in the pre- (post) monsoon period advected moisture and triggered the MCSs within the cloud clusters or within the preexisting disturbances near the MT over the BoB. This finding is consistent with the results of Gray (1998), who explained that trade wind surges or monsoon surges of >10 m s−1 are capable of initiating strong convection near tropical disturbances.

b. Mesoscale features for cyclogenesis

In the premonsoon environment, when air advected by the low-level SW winds converged toward the air advected by the NW winds, favorable regions for MCS initiation were found either along the synoptic dryline or along the convergence zone south of the dryline. Initially, a small group of cells formed along the dryline in the presence of very high CAPE (i.e., 4968 J kg−1) and strong wind shear (18 m s−1 at 0–3 km AGL) in the tropical environment (COMET Program/UCAR 1999). A strong bow echo developed, followed by a comma-shaped echo that was associated with winds of 20 m s−1 at the surface. Weisman (1993) observed similar strong and long-lived bow echo systems during the warm season for midlatitude convection that formed in an environment with at least 2000 J kg−1 of CAPE and strong shear (20 m s−1 over the lowest 2.5 km AGL). However, Johns et al. (1990) suggested that severe warm-season bow echoes are possible with very high CAPE (4500 J kg−1), which can help to maintain bow echoes even in the presence of weak forcing. Beyond the dryline and near the southern extent of the BoB, a mature-stage MCS appeared as a nonorganized linear-type QLCS, which formed in an environment with moderate CAPE (i.e., 1329 J kg−1) and veering wind shear of 13 m s−1 at 0–3 km AGL.

In the postmonsoon cases, when strong NE winds advected air over the BoB, a horizontal wind gradient or shear line formed, which assisted in the development of the squall-line MCSs that were oriented in a north–south direction. Moderate instability (1124 J kg−1 for Sidr and 1064 J kg−1 for Giri) and moderate to high wind shear at 0–3 km AGL (12 m s−1 for Sidr and 18 m s−1 for Giri) were favorable for initiating tropical squall-line MCSs. In both cases, backing winds supported cool air advection (Bluestein 1993), even though the temperatures were not sufficiently cold to form a front in the BoB.

All MCSs in this study were QLCSs that contained several mesovortices, which formed at low levels along the leading edge of the system; the intensities of the mesovortices increased through the midtroposphere. The diameters of the vortices were larger in the premonsoon cases than in the postmonsoon cases. The gradual movement of the systems and the formation of new mesovortices downwind created a midtropospheric environment with enhanced vorticity, which further assisted in the formation of an active cyclone. Some low-level vortices were very intense and resembled mesocyclones, and some exhibited characteristics analogous to VHTs and extended to the tropopause. Because of such VHTs, it may be approximated to agree with the bottom-up paradigm (Hendricks et al. 2004; Reasor et al. 2005; Montgomery et al. 2006; Braun et al. 2010) for cyclogenesis in which low-level mesovortices play an additional important role in the genesis of a cyclone vortex in the BoB. Further detailed examinations of all low-level vortices and their corresponding interactions are needed.

Table 2 provides an overall summary of the case studies, such as the basic characteristics of the seasonal MCSs and the associated environments that favor bimodal TC genesis in the BoB. However, a firm conclusion that MCSs always maintain a QLCS, such as a bow echo, a nonorganized linear system and a squall line, for all TCs or that any other types of MCSs are possible during cyclogenesis in the BoB is difficult to state. The NOAA interpolated outgoing longwave radiation (OLR) 2 days before the formation of each TC (2001–12) for intensities ≥64 and <64 kt are separately averaged and presented in Fig. 12. For intense TCs, the gradient of the OLR composites along the meridional (zonal) direction is very steep in the pre- (post) monsoon period, and it indicates possible quasi-linear deep convection in the east–west (north–south) direction within the BoB. In the case of weaker TCs, the mean OLR values do not have a significantly linear signature for QLCS, especially during the postmonsoon cases. For high-intensity TCs, a QLCS is characteristic of an MCS; in other words, the formation of a QLCS and its associated low-level mesovortices can intensify TCs. In the case of high- and low-intensity TCs, SW winds in the premonsoon period and NE winds in the postmonsoon period are significantly more intense than the ambient wind field.

Table 2.

Basic characteristics of MCSs and the environments associated with TC genesis over the BoB.

Table 2.
Fig. 12.
Fig. 12.

Average NOAA daily-interpolated OLR (shaded, W m−2) and 850-hPa horizontal wind (vectors, m s−1) for 2 days prior to the formation of the premonsoon TCs with intensities (a) ≥64 and (b) <64 kt during 2001–12. (c),(d) As in (a),(b), but for the postmonsoon TCs.

Citation: Monthly Weather Review 143, 9; 10.1175/MWR-D-14-00260.1

Overall, the various MCSs in the bimodal cyclone seasons depend on the characteristics of the low-level sheared wind surges that advect warm or cool air and on moisture and confluence within the bay ahead of the surges. For the cases here, fast-moving, long-lived QLCSs with leading-edge convection were a common characteristic of MCSs associated with cyclogenesis in the BoB. Nonetheless, more cases are needed to attain concrete conclusions regarding the pre- and postmonsoon MCS characteristics during cyclogenesis in the BoB.

Acknowledgments

The author is very grateful to the Department of Physics, Bangladesh University of Engineering and Technology, Dhaka, for providing the laboratory facilities. The JTWC data and NCEP data were downloaded from their respective websites. The Grid Analysis and Display System Software (GrADS) was used for analyzing and displaying the data.

REFERENCES

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