Mesoscale Convection and Bimodal Cyclogenesis Over the Bay of Bengal

Home , Cyclogenesis, Cyclone Akash, Dry line, Mesoscale meteorology, Tropical wave

SEPTEMBER 2015 A K T E R 3495

Mesoscale Convection and Bimodal Cyclogenesis over the Bay of Bengal

NASREEN AKTER Department of Physics, Bangladesh University of Engineering and Technology, Dhaka, Bangladesh

(Manuscript received 8 August 2014, in ﬁnal form 11 May 2015)

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.

1. Introduction and the northward-propagating deep convection phase of the intraseasonal oscillations (ISOs) trigger an earlier The South Asian premonsoon (March–May) and postmonsoon onset in the BoB than in India (Jiang and Li monsoon (October–December) seasons are the transition 2011; K. Li et al. 2013; Yu et al. 2012). periods between the summer (June–September) and The BoB is not only significant for the Asian monsoon winter (January–February) monsoons. A strong south- onset but also offers a unique setting for tropical cyclone westerly (SW) prevailing wind transports enormous (TC) activities. The TCs over the BoB are confined moist and warm air masses from the sea to the land in the within the monsoon transition periods, with a maximum boreal summer, while dry and cool northeasterly (NE) frequency in October–November and a second maxi- flows occur in the opposite direction (i.e., from the land mum in May (McBride 1995; Harr and Chan 2005; to the sea) during the boreal winter (Ramage 1971; Das Camargo et al. 2007; Kikuchi and Wang 2010; Akter and 1995). The Arabian Sea and the Bay of Bengal (BoB), Tsuboki 2014, hereafter AT14). In the boreal summer, which are two branches of the north Indian Ocean the location of the monsoon trough (MT) is well inland; (NIO), play pivotal roles in the seasonal wind reversal of the prevailing southwesterly winds and upper-level South Asian monsoons. The southern BoB experiences easterly winds create strong vertical shear that sup- the southwest monsoon in early May (Wu and Zhang presses TC formation (Jeffries and Miller 1993; McBride 1998; Mao and Wu 2007; Wu et al. 2012). The maximum 1995; Z. Li et al. 2013). Conversely, the bimodal TC annual sea surface temperature (SST) in the central BoB activity over the BoB is modulated by the seasonal mi- gration of MT locations that are usually in the northern and central BoB during the pre- and postmonsoon sea- Corresponding author address: Nasreen Akter, Department of Physics, Bangladesh University of Engineering and Technology, sons, respectively (McBride 1995; AT14). Within the Zahir Raihan Rd., Dhaka 1000, Bangladesh. transition seasons, intraseasonal ISO phases often in- E-mail: [email protected] fluence TC formation in the BoB (Kikuchi and Wang

DOI: 10.1175/MWR-D-14-00260.1

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2010; Kikuchi et al. 2012; Yanase et al. 2012). Both the cycles of deep, moist convective activity called vortical seasonal MT position and ISO phases are associated hot towers (VHTs), where the lower-tropospheric vor- with the formation of synoptic-scale tropical distur- tices are formed within the embryonic environment of bances or cloud clusters that favor active cyclones (Gray MCVs (Hendricks et al. 2004; Reasor et al. 2005; 1998; Roundy and Frank 2004). Montgomery et al. 2006; Braun et al. 2010). The com- Recently, AT14 revealed that the northern position of bined effect of the upscale growth of cyclonic vortices the MT in the BoB and environmental convective in- and their integration contributes to the development of hibition (CIN) are mutually responsible for the de- the TC vortex. creased cyclone frequency in May (premonsoon), even The structure and characteristics of MCSs and the as- though the BoB maintains higher SSTs that result in sociated MCVs that form in the BoB during pre- and increased convective available potential energy (CAPE; postmonsoon TCs are vital to the seasonal bimodal Glickman 2000) compared with the postmonsoon sea- cyclogenesis process. Few studies have focused on the son. In the premonsoon season, deep hot and dry air is different types of MCSs and their contributions to pre- advected from northwest India toward the BoB to pro- cipitation in South Asia during the premonsoon and vide environmental CIN that caps the boundary layer monsoon seasons (Houze et al. 2007; Romatschke et al. over the northwest BoB (Fig. 7 of AT14). Consequently, 2010; Romatschke and Houze 2011a,b). Previous studies TC genesis is reduced because of suppressed convection. have shown that the BoB experiences large systems with Therefore, seasonal environmental flow is an essential extremely large stratiform regions in both the pre- dynamical aspect that precedes TC genesis in the BoB. monsoon and monsoon seasons. During the premonsoon Ritchie and Holland (1999) and Yoshida and Ishikawa period, MCSs are primarily related to depressions and (2013) investigated five types of large-scale dynamical exhibit weak diurnal cycles. Moreover, no investigation flow patterns associated with cyclone development: on postmonsoon MCS characteristics has been conducted. monsoon shear line (SL), monsoon confluence region Specifically, synoptic-scale flow patterns and related (CR), monsoon gyre (GY), easterly wave (EW), and MCSs for TC genesis have not been identified in the BoB. Rossby wave energy dispersion (RD) in the western Therefore, the objective of the present study is to assess North Pacific; the results demonstrated that the SL, CR, how seasonal variations in the synoptic-scale flows over and GY patterns are related to the MT. Conversely, for the BoB determine the different types of MCS formations the same basin, Lee et al. (2008) discussed the EW, SL, and the relevant vorticity generation within the BoB that and CR patterns and three synoptic-scale flows: SW, contributes to the seasonal bimodal distribution of tropi- NE, and combined NE and SW; the results showed that cal cyclogenesis. To achieve this objective, simulations of all of the patterns are related to the monsoon, except for the structural characteristics of MCSs during cyclogenesis the EW. Furthermore, Lee et al. (2008) noted that me- in the BoB are the only viable option because observa- soscale convective system (MCS; Houze 2004) activities tional data are limited. are linked with such synoptic flows. Many studies have acknowledged that large-scale or 2. Model specifications and data used synoptic-scale flows are not the only major contributors to TC genesis. Individual MCSs that are associated The Advanced Hurricane Weather Research and Fore- with a preexisting tropical disturbance and cyclonic casting (WRF) Model (AHW) (version 3.3.1), which is mesoscale convective vortices (MCVs) that develop in derived from the Advanced Research version of the the stratiform precipitation region near the middle tro- WRF (ARW) Model (Davis et al. 2008; Skamarock et al. posphere are also fundamental precursors for cyclo- 2008), is used to simulate pre- and postmonsoon cy- genesis (e.g., Zehr 1992; Harr et al. 1996; Bister and clones for examining the MCSs that are associated with Emanuel 1997; Ritchie and Holland 1997; Gray 1998; bimodal cyclogenesis. A Lambert projection map is Dunkerton et al. 2009; Houze 2010). Two processes are utilized with two-way nested domains; the grid spacing is hypothesized to explain lower-tropospheric vortices 12 km for the outer domain and 4 km for the innermost produced by midtropospheric MCVs: the top-down and domain (Fig. 1b). Davis et al. (2008) showed that 12- and bottom-up paradigms for cyclogenesis. The first para- 4-km grid spacing provide the most accurate forecasts of digm emphasizes the downward advection of MCVs in a storm position and intensity. The parent domain (D1) moist environment (Emanuel 1993; Bister and Emanuel consists of 861 3 595 grid points, while the inner nest 1997), where greater penetration occurs by the merging (D2) has 1064 3 996 grid points. The 28 terrain- of individual MCVs and cyclonic low-level vorticity is following vertical levels, where the top level is 50 hPa, further enhanced (Simpson et al. 1997; Ritchie and are used. The Kain–Fritsch cumulus parameterization, Holland 1997). The second paradigm incorporates which predicts deep and shallow convection using a mass

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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 2 (shaded, g kg 1) and temperature anomalies (solid contours, 2-K intervals) calculated from the background area in 2 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).

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TABLE 1. Pre- and postmonsoon cyclones.

Casualty Cyclone Formation time Formation Max wind Damage Year name and day location (kt) Categorya Landfall coast (millions USD) Deaths 2007 Akash 1800 UTC 13 May 16.38N, 91.38E 65 1 Bangladesh–Myanmar $982 14 border Sidr 0600 UTC 11 Nov 10.18N, 92.28E 140 5 Bangladesh $1700 3447 2010 Laila 1800 UTC 17 May 11.78N, 86.98E 65 1 India $118 65 Giri 1800 UTC 20 Oct 17.18N, 91.68E 135 4 Myanmar $359 157 a Safﬁr–Simpson hurricane scale.

flux approach (Kain 2004), is only applied in D1 with a (NCEP) Final Analysis (FNL) Operational Global 5-min time step. The Ferrier microphysics scheme, Analysis data at 6-h intervals with a resolution of 18318. which includes a prognostic mixed-phase representation The NCEP high-resolution real-time global sea surface of changes in water vapor and condensate and considers temperatures (RTG_SSTs) at a 0.0838 resolution are im- cloud water, rain, cloud ice, and precipitating ice posed for the daily sea surface temperatures. Two pre- (Ferrier et al. 2002), is used in both domains. monsoon TCs—Akash (2007) and Laila (2010)—and two The Monin–Obukhov scheme (Monin and Obukhov postmonsoon TCs—Sidr (2007) and Giri (2010)—are 1954) is used for the surface layer physics; the Noah land considered in this study by running four simulations using surface model, which considers four soil temperature the aforementioned experimental setup, with no bogussing, and moisture layers (Chen and Dudhia 2001), is also for the TC initialization. The duration of each simulation is used. Moreover, the Yonsei University (YSU; Hong 4 days (Akash: 0000 UTC 11 May–0000 UTC 15 May 2007, et al. 2006) planetary boundary layer (PBL) scheme is Laila: 0000 UTC 15 May–0000 UTC 19 May 2010, Sidr: selected because it is a first-order closure scheme that 0000 UTC 8 November–0000 UTC 12 November 2007, explicitly treats entrainment processes at the top of the and Giri: 0000 UTC 18 October–0000 UTC 22 October PBL. The asymptotic entrainment flux term is proportional 2010). A maximum surface wind speed of 34 kt or 2 to the surface flux in the inversion layer (Noh et al. 2003). ;17.5 m s 1 in Joint Typhoon Warning Center (JTWC) The Rapid Radiative Transfer Model (RRTM; Mlawer best track data is considered as the condition of its for- et al. 1997) and the Dudhia scheme (Dudhia 1989)are mation in this study. The maximum wind speed beyond 2 selected for the longwave and shortwave radiation calcu- ;17.5 m s 1 is referred to as TC intensification by self- lations, respectively. sustaining mechanism (Zehr 1992; McBride 1995). On In this study, all of the physical parameterizations are the basis of JTWC data, each case has taken less than chosen according to Raju et al. (2011) and Osuri et al. 1.5 days to turn from tropical depression [maximum 2 (2012), who customized the parameters for simulating surface wind speed of 15 kt (1 kt 5 0.5144 m s 1)] to the cyclones over the NIO using the WRF Model. cyclogenesis. So the initial time for simulation, including Additionally, the AHW Model includes a one- the model spinup time, is kept at least 2.75 days prior to dimensional ocean mixed layer model and modified the commencement of cyclogenesis (Table 1). As the surface flux and drag formulations for high-wind con- SST data are daily, the initial start time of each simu- ditions over the ocean. The ocean mixed layer model is lation is 0000 UTC and the total time from initiation of based on Pollard et al. (1973), and it attempts to capture the model simulation to the formation of all TCs is the negative feedback of the SST on TCs (i.e., the de- 2.75 days, except for Sidr (2007), which is 12 h longer crease in SSTs due to the passage of a TC). For this than others. The model output is acquired at a 30-min study, a 30-m ocean mixed layer depth is selected for interval. initializing the model over the BoB; details are available The simulated results are verified by NCEP Climate in Kumar et al. (2011). Here, Donelan’s drag formula- Forecast System Reanalysis (CFSR) 6-hourly data, tion (Donelan et al. 2004), which predicts weaker sur- which have a resolution of 0.5830.58. The data have 64 face friction under high wind conditions compared with hybrid sigma-pressure vertical levels, with a top pressure the Charnock formulation, is selected. The surface en- of ;0.266 hPa (Saha et al. 2010). In addition, monthly thalpy flux is calculated using the formulation of mean NCEP–CFSR data are used for analyzing the cli- Garratt (1992). matological conditions of the BoB, and JTWC data are The atmospheric initial and boundary conditions are de- utilized to verify the simulated position and intensity of rived from National Centers for Environmental Prediction the TCs. NCEP reanalysis and NOAA daily interpolated

Unauthenticated | Downloaded 10/05/21 07:25 AM UTC SEPTEMBER 2015 A K T E R 3499 outgoing longwave radiation (OLR) data with a 2.583 the southwest monsoon of IMD for 2007 (available in 2.58 resolution are also used for 2001–12. 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- 3. Bimodal cyclones northeastward and intensified as tropical cyclone Akash on 13 May; finally, it crossed the Bangladesh coast on a. Overview 15 May (Bhatia and Rajeevan 2008; Mazumdar et al. 2008). Only 7% of global TC formations occur in the NIO In 2010, IMD reported that the southwest monsoon set (Neumann 1993); however, TCs in the region are dev- up over the southeast BoB on 17 May, 3 days prior to the astating because of the funnel-shaped low-lying coastal normal date as a consequence of a severe cyclonic storm areas. The pre- and postmonsoon TCs that formed in the Laila (depression to landfall: 16–21 May 2010) over the BoB BoB in 2007 and 2010 are considered suitable cases for (India Meteorological Department 2010). Both Akash resolving the seasonal differences in MCS formations (2007) and Laila (2010) developed before the monsoon during cyclogenesis. The most important criteria for onset in their respective locations in BoB. Favorably, in selecting these years are as follows: (i) cyclones between both cases, cross-equatorial flow across the southern 2001 and 2012 have the most sufficient data, including BoB was accelerated by the TC vortices. Thus, the cy- damage information; (ii) pre- and postmonsoon TCs from clones in this study are representative premonsoon TCs the same year avoid interannual variability, such as El that correspond well with the locations of MTs over the Niño–Southern Oscillation (ENSO), between the two ocean in both cases (details in section 4a). seasons; and (iii) very severe cyclonic storms [i.e., maxi- The synoptic conditions for each cyclone are also il- mum wind speeds of at least 64 kt based on the Indian lustrated in Figs. 1b–e using the monthly mean specific Meteorological Department (IMD) scale] occurred. Ac- humidity and temperature of the NCEP–CFSR re- cording to the JTWC data from 2001 to 2012, a total of analysis data. In the premonsoon season (especially 36 cyclones formed over the BoB; 30.6% formed during May), the BoB is influenced by two different air masses: the premonsoon season, and 69.4% formed during the a hot, dry air forcing from northwestern India and a postmonsoon season (Fig. 1a). Among these cyclones, warm, moist air transporting from the southwest of the only 6 (16.7% of the total) in the premonsoon season and bay. The warm, moist air masses are wedged under deep 3 (8.3% of the total) in the postmonsoon season had wind hot and dry air that extends from 950 to 600 hPa (AT14). speeds $64 kt. Conversely, in the postmonsoon (;October–November) The criteria mentioned above indicate that the cy- cases, the entire environment of the BoB is nearly uni- clones of 2007 and 2010 are the most suitable for this form regarding temperature and humidity. The track of study. Moreover, both years experienced La Niña, which each cyclone (i.e., Akash, Sidr, Laila, and Giri) that is the cool phase of the ENSO cycle. The TC features, originated as a depression (i.e., a maximum wind speed of which were collected from the JTWC, are sequentially 15 kt) is also displayed in the corresponding figures. arranged in Table 1 (JTWC 2007, 2010), except for the b. Model verification casualty information. The damage estimates were obtained from Begum et al. (2013). The data collected from the 12-km simulation are Akash (2007) and Laila (2010) both formed almost at compared with the observed (JTWC) and NCEP–CFSR the center of BoB (Figs. 1b and 1d) during May, when reanalysis (;55-km resolution) data for the TCs in 2007 important monsoon evolutions occurred in the southeast and 2010. The surface pressure, relative vorticity, and BoB. Both TCs and the monsoon advancement in the horizontal wind at 850 hPa are illustrated in Fig. 2 at the BoB are influenced by ISO; however, TCs and monsoons genesis time of each TC (Table 1) to demonstrate the do not occur in unison because their activities are mod- evolution of the TCs. The simulation results nearly ac- ulated by the interannual variability (phase and intensity) curately depict the position and intensity of the TCs. The of ISO (Fosu and Wang 2015). Z. Li et al. (2013) argued model and reanalysis (JTWC) data suggest a minimum that April–May is the perfect premonsoon period be- surface pressure of 990 and 985 (997) hPa near the cause the oceanic and atmospheric conditions change vortex center for Akash (2007). Moreover, the simu- dramatically in preparation for the monsoon onset. In lated minimum surface pressures for Sidr (2007), Laila connection to the greater ocean heat content, the first (2010), and Giri (2010) are 1003, 1005, and 1007 hPa, branch of northward-propagating ISOs is associated with respectively; the reanalysis (JTWC) data suggest that the monsoon onset over the BoB, and stronger ISO in- the minimum surface pressure is 1006 (;996) hPa for tensity in April–May is a favorable environmental con- these three TCs. The wind rotates around the TC center 2 dition for cyclone intensification. According to a map of with a velocity greater than 16 m s 1 for all simulated

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21 FIG. 2. The surface pressure (contours, 2-hPa intervals), wind velocity (vector, m s ), and vor- 2 2 ticity (shaded, 310 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.

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2 cyclones, except for Giri (wind speed of ;11 m s 1). The midlevel vorticity maximum structure change to be a 2 2 relative vorticity is on the order of 10 4 s 1 in all simu- low-level maximum, then the wind-induced surface heat lations, and it signifies strong rotation at the TC centers. exchange (WISHE; Emanuel et al. 1994) feedback Because the data resolution is lower, the values of the mechanism may become dominant. vorticity are less intense in the reanalysis data than in the a. Cloud clusters simulations. Cyclogenesis depends on the formation of tropical During the early genesis period, which is defined convection or MCSs (Gray 1998; Zehr 1992), which are 2.5 days prior to each cyclogenesis event, the environ- initiated by several important factors (e.g., the avail- ment is determined by the 6-h-averaged moisture flux, ability of CAPE and low-level convergence) (Protat and reflectivity, and streamlines at 925 hPa from the 12-km Lemaitre 2001). The surface-based environmental resolution model grid (Fig. 4). The zero line of the av- CAPE within the BoB is, therefore, verified with the erage zonal vertical wind shear between 850 and 200 hPa reanalysis data two days prior to each genesis. The indicates a low shear region or MT passing over the BoB simulated CAPE, including the seasonal variability, (McBride 1995; AT14). Within the trough region, cloud over the ocean is similar to the reanalysis data (Fig. 3). clusters are organized and spread to at least ;700 km, 2 The CAPE is .2500 J kg 1 during the premonsoon ca- which is consistent with the cluster length discussed by ses because of higher SSTs (Sasamal 2007; Alappattu Gray (1998). Two types of synoptic flow patterns in the and Kunhikrishnan 2009). In contrast, the CAPE is BoB are related to the seasonal development of TCs. 2 lower (,1800 J kg 1) in the postmonsoon environment. In the premonsoon cases, a combination of SW and The seasonal variability in the instability is consistent northwesterly (NW) winds appear at low levels within with the result of AT14 using 20-yr-average seasonal the BoB. The NW flow advects dry air from the arid data. The simulated surface wind and its convergence regions of southwestern Asia and western India to the toward the genesis location are also analogous to the BoB up to 500 hPa; however, the equatorial SW flow reanalysis data. 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 4. Hierarchy of cyclogenesis in the BoB boundary of the coast and flows to the north and Recent studies have suggested that synoptic-scale northwest over the BoB. After entering the bay, the NW 2 tropical disturbances or cloud clusters (;700 km) are wind gradually moistens (humidity ;7gkg 1 in the 2 preconditions for initializing TCs; these features are north BoB to ;15 g kg 1 in the south BoB) at the low possibly formed by monsoonal troughs, the ITCZ, east- levels (also Fig. 1) as the influx of water vapor intensifies erly waves, or equatorial waves (Gray 1998). A preex- with higher SSTs during this period (29.88C on average, isting disturbance is characterized by cyclonic relative AT14); however, the total moist air masses within the vorticity in the lower troposphere, where large-scale NW flow are smaller than those define by the region of convergence (wind surge) can assist in the develop- SW winds. ment of an MCS (;250 km; Gray 1998) that has deep In the postmonsoon cases, the SW and NE winds co- convective cells, stratiform clouds, and precipitation exist from low levels to the 600-hPa level. In this period, (Houze 2004). Initially, an MCS may be initiated by one the SW monsoonal winds continue to retreat south of or more isolated VHTs (;10 km; Montgomery et al. the BoB, while the onset of easterly or NE winds occurs 2006). Following the weakening of the VHT, a strati- in the northern extent of the BoB. The NE winds form region may appear in the MCS over time (Houze gradually advect more low-level moisture to the BoB 2010). Moreover, latent heat release in the stratiform relative to the surroundings; however, the average region accelerates midtropospheric warming and evap- moisture is less than that in the premonsoon environ- orative cooling below. The combination of these ther- ment. These strong directionally distinct dynamic and modynamic profiles with updrafts above and downdrafts thermodynamic environmental forcings in the BoB, es- below can lead to the development of an MCV that is pecially the SW wind surges in the premonsoon cases typically on the order of 100 km in diameter (Gray and the NE wind surges in the postmonsoon cases, 1998). The mature MCS is then characterized by both supply low-level moisture and support the initiation of cyclonic vorticity in the convective-scale VHTs and a mesoscale deep convection within cloud clusters. In the midtropospheric MCV during the middle of the life cy- next section, the characteristics of the MCSs formed in cle. In the later stages of the MCS, the VHT may no the BoB and the associated vortices related to the active longer remain, but the MCV persists. If the cold pool cyclones during the pre- and postmonsoon seasons are under the MCV can be removed or warmed and the discussed using the simulation at 4-km resolution.

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21 FIG. 3. The surface-based CAPE (shaded, J kg ) 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.

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21 FIG. 4. The 6-h-averaged moisture ﬂuxes (shaded, m s ), reﬂectivities (contours, 20-dBZ intervals), streamlines 2 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.

2 2 b. Mesoscale convection horizontal moisture gradient [;2gkg 1 (100 km) 1]asso- 2 ciated with positively varying temperature [;18 (100 km) 1] 1) EARLY STAGE CONVECTION to the west of the MCS indicates a synoptic dryline boundary (Schaefer 1986). The MCS forms in the low- The initial MCS stage (widespread regions of ;35-dBZ level convergence when the air transported via strong 2 reflectivity), which formed nearly 2.5 days (66 h) before SW winds (.12 m s 1) moves into relatively dry NW 2 the cyclogenesis of each TC, is shown in Fig. 5. The time winds (,6ms 1) along or near the dryline boundary. of early MCS stage is defined as ‘‘t’’ in each case of this Similarly, the early MCS of Laila (2010) originates when study. Here, t is 0530 UTC 11 May for Akash (2007), air that is transported by intense SW winds enters the 0800 UTC 15 May for Laila (2010), 0600 UTC 8 No- BoB and converges with air advected by weak NW winds vember for Sidr (2007), and 1130 UTC 18 October for near 48–68N. Because of the limitation of the domain area Giri (2010). The environmental low-level horizontal (D2), the SW wind is not visible in the figure; however, it wind, the water vapor mixing ratio, and the temperature is confirmed by the simulated result in D1 (also in Fig. 4). anomalies at time t are also illustrated in Fig. 5. For In contrast to the MCS of pre-Akash (2007), the MCS of Akash (2007), an MCS containing a small group of cells pre-Laila (2010) emerges as a linear convective system initially appears within an environment of significant tem- (Houze et al. 1990) and forms in the southernmost extent perature and humidity variations on the synoptic scale. of the BoB within a region of uniform temperature (228C) 2 Temperature (moisture) is gradually increasing (de- and humidity (17 g kg 1) at the low level. creasing) from the east of MCS toward the west by the Furthermore, both postmonsoon cases are character- 2 values from 238 to 288C (21–12 g kg 1). A sharp negative ized by a prominent north–south shear line in which

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21 FIG. 5. Water vapor mixing ratio (shaded, g kg ), reﬂectivities (contours, 20-dBZ intervals), temperature anomalies (dashed line, 1-K intervals) calculated from the average of domain D2, and horizontal wind (vector, 2 ms 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.

2 relatively strong NE winds (.10 m s 1) create a hori- horizontal environmental shear (Barnes and Sieckman zontal wind shear line across which an abrupt change in 1984; LeMone et al. 1998; Wang and Carey 2005). No the horizontal wind component occurs (Glickman 2000). remarkable synoptic-scale temperature differences are The convective systems in these cases are tropical squall observed in the postmonsoon environment; however, air lines because they form along the conﬂuence boundary advected by the northeasterly wind is slightly cooler of the shear line and contain a narrow quasi-linear band (,21 K) than the surroundings. of active convection—either continuous or discontinuous— Environmental soundings (Fig. 6) and hodographs that is perpendicularly aligned with the low-level (Fig. 7) at each MCS formation time t are obtained for

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FIG.6.SkewT–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.

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enhanced because the MCS of Akash (2007) develops along the dryline boundary within the BoB. In the hodographs, the vertical wind shear from the surface to 6 km above ground level (AGL) suggests a 2 moderate to strong total shear of 15–22 m s 1 and a bulk 2 shear of 8–10 m s 1, which support the formation of or- dinary 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 in- dication 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 FIG. 7. Hodographs at the sounding points described in Fig. 6. The premonsoon TCs. The hatched area in the figure shows numbers represent the height in km. the position and size of the MCSs after 6 h, which indicate the direction that systems are moving. Vertical the areas where moisture advection triggers MCS initi- cross sections approximately perpendicular to the sys- ation. In all cases, the soundings indicate unstable lapse tems are shown in Figs. 8b,c for Akash (2007) and in rates (Doswell et al. 1985), with lifted index values (LIs; Figs. 8e,f for Laila (2010). These cross sections contain Galway 1956) #23 and high kelvin index (George 1960) the reflectivities, the cross section parallel to the system- values (greater than 35). These values permit the envi- relative winds (i.e., approximately system-relative me- ronment to develop a strong convective system with ridional wind component), the water vapor mixing ratios, heavy rain. The level of free convection for a surface the updrafts, the downdrafts, the vertical vorticities, and parcel is between 950- and 975-hPa level, which is very the potential temperature perturbations. In the mature close to the lifting condensation level; deep convection stage, the MCS of Akash (2007) is organized as a bow is, therefore, more likely (Rasmussen and Blanchard echo (i.e., it presents a bowed thunderstorm outline) 1998; Craven et al. 2002). More than 75% of the mois- (Fujita 1978), and it becomes an increasingly comma- ture in all cases is between the surface and ;6 km; shaped asymmetric echo with a length of ;300 km however, dry air is noticeable in the environment of (Fig. 8a). Intense convection evolves to the south of the Akash (2007) at 500 hPa. Air parcels are much warmer system and is accompanied by a convective comma head and more buoyant than the environment in the presence to the north. The strong system-relative low-level moist 2 2 of moderate to high CAPE (.1064 J kg 1) and CIN of (humidity .16 g kg 1) inflow from the south side of the 2 ;0Jkg 1. Among all the cyclones, the premonsoon system instigates deep convection to the south of the environment of Akash (2007) is extremely unstable system and stratiform rain developing over the region to 2 (LI 5210 and CAPE 5 4968 J kg 1), and there is the the north. The vertical height of the convection is potential for initiating severe thunderstorms (Bluestein approximately 15 km. 1993). The severe dryline convection (Weston 1972; The most recognizable feature in a bow echo is the Ziegler and Rasmussen 1998; Hane et al. 2002; Murphey rear-inflow jet (RIJ; Weisman 1993), which is elevated et al. 2006) during the premonsoon period may be air that descends from the rear to the front of the line of

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21 FIG. 8. Premonsoon TCs: (a) mature-stage MCS (shaded; dBZ), 850-hPa wind velocity (m s ), and MCS after 6 h (hatched area) for Akash (2007). The yellow line is the approximate RIJ. The vertical cross section along AB in 2 (a) indicates the (b) vertical reﬂectivities (shaded; dBZ), downdrafts (thin contours; 0.2-m s 1 interval), water vapor 2 mixing ratio (thick contours; .16 g kg 1), and cross-sectional parallel-system-relative wind with vertical velocity 2 (vectors; m s 1). (c) Potential temperature anomalies for a 68348 area for the system-wide values (shaded; K) and 2 2 2 updrafts (thick solid contours; 2 m s 1 interval); the dashed (thin solid) contours of 1 3 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.

Unauthenticated | Downloaded 10/05/21 07:25 AM UTC 3508 MONTHLY WEATHER REVIEW VOLUME 143 convection and supplies cool, dry midlevel air that aids organized. The system continues until it dissipates after in the production of the convective- and system-scale ;14 h. Within a few hours, many other scattered orga- downdrafts. In this case, the 3–4-km RIJ from the north nized and nonorganized convective systems form and or northwest descends behind the line of convection merge together. (Fig. 8), creates a downdraft, and ultimately drives the Figure 9 represents parameters similarly to those in 2 system to the south at an average speed of ;9ms 1. Fig. 8 to demonstrate the postmonsoon MCSs; in both Because the midtroposphere is dry (Fig. 6), evaporative cases, the MCSs are squall lines (Houze and Betts 1981; cooling creates negative buoyancy that further acceler- Gamache and Houze 1982) with leading-edge convective ates the downdraft toward the surface. The downward- precipitation that is trailed by a narrow region of strati- moving air that typically spreads out in all directions form rain. The systems are very long, with a length of after reaching the ground may produce strong, damag- ;500 km and a maximum height of 15 km. The line of ing winds (Fujita 1978). The near-surface reflectivity convection forms and intensifies when the system-relative gradient in Fig. 8b indicates straight-line winds that rush low-level moist westerly inflow ahead of the line con- downward within the core of the storm and that spread verges with the easterly inflow behind the squall line. 2 at a velocity of ;20 m s 1. The updraft is upright within 100–300 hPa due to the Because of the tropical moist environment, the proper balance between the moderate instability and the 2 2 downdraft is #21.6 m s 1, which is less than that of moderate to strong low-level vertical shear (.12 m s 1 midlatitude convection (Xu and Randall 2001). at 0–3 km AGL). The storms are linearly organized Consequently, a small low-level cold pool, with a po- along the boundary of the convergence line; the storm tential temperature perturbation of #24 K, is produced; outflows create a weak (,23 K) and shallow cold pool however, this value is sufficient for tropical convection along the rear side of the system. Tropical squall lines (COMET Program/UCAR 1999). The higher tempera- can easily develop in low-shear and low-LFC environ- ture associated with the updraft indicates latent heat ments, and they are triggered by weaker cold pools release by the convection. The updraft is quite erect and that are 2–4 K colder than the ambient temperature remains at the leading edge of the convective line with a (COMET Program/UCAR 1999). The vertical cross 2 maximum intensity of ;12 m s 1. This finding may be a sections in Figs. 9b,e support the concept that new cells result of the balance between the vorticities associated develop along the leading edge of the system’s gust front 2 with the ambient low-level vertical wind shear (18 m s 1 and then propagate rearward to contribute to the growth at 0–3 km AGL) and the low-level cold pool that forms of the trailing stratiform precipitation region. Low-level beneath the convection (Weisman and Rotunno 2004). mesovortices and the updrafts are observed in Figs. 9c,f. 2 2 The intense vertical vorticity (maximum of 4 3 10 3 s 1) A strong postmonsoon low-level easterly jet transports 2 associated with the updraft indicates the presence of a the systems at an average velocity of 8 m s 1 and pro- strong mesocyclone near the apex of the comma-shaped motes swelling at the midpoint of the line. The life echo. The outflow boundary that propagates ahead of spans of Sidr (2007) and Giri (2010) are ;18 and ;20 h, the line initiates new cells downwind and helps the sys- respectively. tem persist for ;16 h (not shown). c. Mesoscale convective vortices Alternatively, a quasi-linear type of less-organized convective cells along the leading edge and a trailing Multiscale vortex formation and interaction occur stratiform rain region appear during the mature MCS during cyclogenesis. The distributions of vertical cyclonic 2 2 stage for Laila (2010). The mature system has a height of vorticities on the order of 10 3 s 1 (as a function of time) ;15 km and a width of ;250 km. The cross-section are displayed at different levels (i.e., 925, 500, and parallel-system-relative winds reveal the southerly in- 200 hPa) in Fig. 10 for both pre- and postmonsoon cases. flow that supplies low-level moisture to the updrafts. In this study, all of the systems have quasi-linear con- 2 The updraft and downdraft are 4 and 20.6 m s 1, re- vective systems (QLCSs) and several meso-g-scale spectively. The horizontal vorticity generated by both (2–20 km; Orlanski 1975) vortices (‘‘mesovortices’’) that the system cold pool and the ambient low-level wind form at low levels along the convergent outflow 2 shear (13 m s 1 at 0–3 km AGL) further initiate meso- boundary (Figs. 8 and 9). The horizontal scale of the 2 2 vortices (2 3 10 3 s 1) at low levels via tilting (Holton cyclonic mesovortices is 10–40 km in the premonsoon 1992). Moreover, the system is moving with an average cases and only 5–10 km in the postmonsoon cases. Sev- 2 velocity of ;9ms 1 toward the northwest due to the eral studies (Weisman and Trapp 2003; Sippel et al. southeasterly wind, which is the deflected SW wind from 2006; Wheatley and Trapp 2008; Atkins and St. Laurent the Malay Peninsula and Sumatra. New cells form ahead 2009a) have noted similar mesovortices at low levels of the leading line; however, these cells are not well along the leading edge of convective systems, particularly

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FIG.9.AsinFig. 8, but for the postmonsoon TCs in grayscale.

in QLCSs (i.e., squall lines and bow echoes). The low- originates due to the tilting of the crosswise horizontal level mesovortices that develop in QLCSs depend on the vorticity in the downdraft (not shown); this ﬁnding is environmental vertical wind shear within 0–2.5 km AGL consistent with observations of mesovortices in QLCSs (Weisman and Trapp 2003) and the presence of a posi- (i.e., Trapp and Weisman 2003; Wheatley and Trapp tive Coriolis force (Trapp and Weisman 2003). 2008; Atkins and St. Laurent 2009b). These low-level For early stage MCSs in both seasons, a vortex couplet cyclonic mesovortices further strengthen the circulation (i.e., cyclonic in the north and anticyclonic in the south) by merging with new vortices and stretching the

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23 21 FIG. 10. Distribution of the horizontal cyclonic vortices (contours of 1 3 10 s )forthetimeperiodfromt 1 1tot 1 13 h in 1-h intervals 2 at (a) 925, (b) 500, and (c) 200 hPa for each TC. In (a)–(c) the vectors are the horizontal winds (m s 1)att 1 13. (d) The vertical cross section 2 2 2 of the updraft (shaded; m s 1) and cyclonic vorticity (contours; 0.4 3 10 3 s 1) along the point located in (c) for each case.

Unauthenticated | Downloaded 10/05/21 07:25 AM UTC SEPTEMBER 2015 A K T E R 3511 planetary vorticity (Trapp 2013). The vertical structure represented by decreasing geopotential heights. La- of the vorticity in Figs. 8 and 9 indicates that cyclonic tent heat released aloft due to convection results in vortices intensify and extend upward during the mature low pressure at levels below, which increases the stage, while anticyclonic vortices tend to weaken or di- magnitude of the cyclonic surface winds around the minish in the mature stage. Figure 10 also shows that the TC center. height of the low-level cyclonic vortices extends through the midtroposphere to the upper troposphere; intense 5. Discussion and summary vorticity occurs within the midtroposphere. A few cyclonic vortices extend to the tropopause (;15 km), and Cyclogenesis in the BoB is unique because the loca- 2 are associated with strong updrafts (.10 m s 1) between tion of the MT and vertical shear primarily determine 100 and 300 hPa. These convective towers are highly and limit cyclonic activity during the pre- and post- vortical in nature within their core, with a maximum monsoon seasons instead of in the monsoon season 2 2 vertical vorticity of .1 3 10 3 s 1. Convective tower (McBride 1995; AT14). The TCs that formed in the examples for each case are shown in Fig. 10d. These premonsoon (Akash and Laila) and postmonsoon (Sidr towers are characterized and defined as VHTs, which and Giri) environmental conditions in 2007 and 2010 are fundamental coherent structures in the tropical cy- were simulated using the WRF Model. Nested domains clogenesis process (Hendricks et al. 2004; Montgomery with grid resolutions of 12 and 4 km were used to ex- et al. 2006; Houze 2010; Braun et al. 2010) and in the TC amine details of MCSs and the environments associated intensification process (Van Sang et al. 2008; Shin and with bimodal TC genesis. The simulations were vali- Smith 2008; Montgomery et al. 2009). In this study, dated with observations. VHTs and low-level mesovortices within QLCSs that a. Synoptic features for cyclogenesis form over the BoB during cyclogenesis are essential building blocks of a cyclone vortex. However, the de- In contrast to the different types of large-scale or tailed structure and formation of VHTs and their in- synoptic-scale flow patterns observed during cyclogen- teractions with mesovortices are not analyzed in esis in the western North Pacific basin (Ritchie and this study. Holland 1999; Lee et al. 2008), only a NW–SW flow Cyclonic vortices increase in both size and intensity pattern was found in both premonsoon cases; the NW with altitude as the individual convective cells embed- wind from northwestern India advected deep hot and ded in the QLCS mature. Moreover, the vortices dry air to the trough region in the BoB, while the SW weaken as the cells decay and propagate rearward with wind advected shallow moist and warm air from the respect to the leading edge of the line. Therefore, the Arabian Sea. The SW flow during May was associated environment behind the convective lines becomes with the early advance monsoon flow in the southeast enriched with vorticity over time, particularly within the BoB and accelerated toward the preexisting low-level midtroposphere because of the presence of the maxi- disturbance (i.e., low pressure) near the MT over the mum vorticity. This phenomenon is signified through ocean. Notably, during the boreal summer, the MT is counterclockwise rotation of the wind at t 1 13 h in located over the Indian subcontinent (208–258N) (Wang Fig. 10. Several studies have also noted broadening cy- 2006), and strong SW flow from the Arabian Sea tra- clonic circulations (on the order of 100 km) in the form verses the entirety of South Asia. of a cyclonic mesovortex or line-end vortex at the end of However, a combination of SW and NE winds was QLCSs (Trapp and Weisman 2003; Wheatley and Trapp identified in the postmonsoon cases. The SW winds 2008; Atkins and St. Laurent 2009a). retreated from the southern BoB (Fig. 4) at the end of In this study, the vortex circulation is more prom- the monsoon season. The NE winds encountered the inent in the premonsoon period than in the post- BoB from the north and northeast and transported low- monsoon period. The dominant cyclonic vortex level moisture to the BoB after being blocked by the persists beyond the life cycle of the QLCS as a MCV, north–south-oriented mountains and converging along which is a deep column of cyclonic vorticity. A warm the east coast of the BoB, particularly near the Malay and moist boundary layer over the ocean further sup- Peninsula (Chang et al. 2005; Akter and Tsuboki 2012). ports deep, moist convection near the MCV. Figure 11 Therefore, the pre- (post) monsoon environment is shows the MCSs, the 500-hPa geopotential heights, characterized by the coupling of NW (SW) wind with the and wind speeds for the pre- and postmonsoon TCs in early advance SW (NE) monsoonal wind in the BoB. 2 2007 at three times: the initial convection at time t,the The dynamical forcing of relatively intense (.10 m s 1) deep moist convection that gradually organizes within SW (NE) wind surges in the pre- (post) monsoon pe- the MCV, and an intensifying tropical storm that is riod advected moisture and triggered the MCSs within

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21 FIG. 11. Reﬂectivity at 925 hPa (shaded, dBZ) and the geopotential height (shaded, m) along with the horizontal wind (vectors, m s )at 500 hPa for TCs in 2007. the cloud clusters or within the preexisting distur- Program/UCAR 1999). A strong bow echo de- bances near the MT over the BoB. This ﬁnding is veloped, followed by a comma-shaped echo that was 2 consistent with the results of Gray (1998),whoex- associated with winds of 20 m s 1 at the surface. plained that trade wind surges or monsoon surges Weisman (1993) observed similar strong and long- 2 of .10 m s 1 are capable of initiating strong convection lived bow echo systems during the warm season for near tropical disturbances. midlatitude convection that formed in an environ- 2 ment with at least 2000 J kg 1 of CAPE and strong b. Mesoscale features for cyclogenesis 2 shear (20 m s 1 over the lowest 2.5 km AGL). How- In the premonsoon environment, when air advec- ever, Johns et al. (1990) suggested that severe warm- ted by the low-level SW winds converged toward the season bow echoes are possible with very high CAPE 2 air advected by the NW winds, favorable regions for (4500 J kg 1), which can help to maintain bow echoes MCS initiation were found either along the synoptic even in the presence of weak forcing. Beyond the dryline or along the convergence zone south of the dryline and near the southern extent of the BoB, a dryline. Initially, a small group of cells formed along mature-stage MCS appeared as a nonorganized the dryline in the presence of very high CAPE (i.e., linear-type QLCS, which formed in an environment 2 2 2 4968 J kg 1) and strong wind shear (18 m s 1 at 0– with moderate CAPE (i.e., 1329 J kg 1) and veering 2 3 km AGL) in the tropical environment (COMET wind shear of 13 m s 1 at 0–3 km AGL.

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TABLE 2. Basic characteristics of MCSs and the environments associated with TC genesis over the BoB.

Characteristics Premonsoon Postmonsoon Formation location Within monsoon trough Within monsoon trough Environmental forcing Warm and extremely moist low-level Slightly cool and moderately moist southwesterly winds low-level northeasterly winds 0–3-km total shear Moderate to strong Moderate to strong Wind veering with height Wind backing with height CAPE High or moderate Moderate Mature-stage MCSs Type (i) Severe bow echoes when the conﬂuence Squall line with a narrow trailing zone is along the dryline region of stratiform rain (ii) Nonorganized quasi-linear leading-edge convective systems when the conﬂuence zone is outside the dryline Length ;300 km, east–west oriented ;500 km, north–south orientation Height 15 km or higher 15 km or lower Avg speed Fast-moving, meridional direction Fast-moving, zonal direction Lifetime Long lived Longer than premonsoon MCSs Mesoscale vortices 10–40 km 5–10 km diameter

In the postmonsoon cases, when strong NE winds seasonal MCSs and the associated environments that advected air over the BoB, a horizontal wind gradient or favor bimodal TC genesis in the BoB. However, a firm shear line formed, which assisted in the development of conclusion that MCSs always maintain a QLCS, such the squall-line MCSs that were oriented in a north–south as a bow echo, a nonorganized linear system and a 2 direction. Moderate instability (1124 J kg 1 for Sidr and squall line, for all TCs or that any other types of MCSs 2 1064 J kg 1 for Giri) and moderate to high wind shear at are possible during cyclogenesis in the BoB is difficult 2 2 0–3 km AGL (12 m s 1 for Sidr and 18 m s 1 for Giri) to state. The NOAA interpolated outgoing longwave were favorable for initiating tropical squall-line MCSs. radiation (OLR) 2 days before the formation of each In both cases, backing winds supported cool air advec- TC (2001–12) for intensities $64 and ,64 kt are tion (Bluestein 1993), even though the temperatures separately averaged and presented in Fig. 12.For were not sufficiently cold to form a front in the BoB. intense TCs, the gradient of the OLR composites All MCSs in this study were QLCSs that contained along the meridional (zonal) direction is very steep several mesovortices, which formed at low levels along in the pre- (post) monsoon period, and it indicates the leading edge of the system; the intensities of the possible quasi-linear deep convection in the east– mesovortices increased through the midtroposphere. west (north–south) direction within the BoB. In the The diameters of the vortices were larger in the pre- case of weaker TCs, the mean OLR values do not monsoon cases than in the postmonsoon cases. The have a significantly linear signature for QLCS, gradual movement of the systems and the formation of especially during the postmonsoon cases. For high- new mesovortices downwind created a midtropospheric intensity TCs, a QLCS is characteristic of an MCS; in environment with enhanced vorticity, which further other words, the formation of a QLCS and its asso- assisted in the formation of an active cyclone. Some low- ciated low-level mesovortices can intensify TCs. In level vortices were very intense and resembled meso- the case of high- and low-intensity TCs, SW winds in cyclones, and some exhibited characteristics analogous the premonsoon period and NE winds in the post- to VHTs and extended to the tropopause. Because of monsoon period are significantly more intense than such VHTs, it may be approximated to agree with the the ambient wind field. bottom-up paradigm (Hendricks et al. 2004; Reasor Overall, the various MCSs in the bimodal cyclone et al. 2005; Montgomery et al. 2006; Braun et al. 2010) seasons depend on the characteristics of the low-level for cyclogenesis in which low-level mesovortices play an sheared wind surges that advect warm or cool air and additional important role in the genesis of a cyclone on moisture and confluence within the bay ahead of vortex in the BoB. Further detailed examinations of all the surges. For the cases here, fast-moving, long-lived low-level vortices and their corresponding interactions QLCSs with leading-edge convection were a common are needed. characteristic of MCSs associated with cyclogenesis in Table 2 provides an overall summary of the case the BoB. Nonetheless, more cases are needed to studies, such as the basic characteristics of the attain concrete conclusions regarding the pre- and

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22 21 FIG. 12. Average NOAA daily-interpolated OLR (shaded, W m ) and 850-hPa horizontal wind (vectors, m s ) 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. postmonsoon MCS characteristics during cyclogenesis ——, and ——, 2009b: Bow echo mesovortices. Part II: Their in the BoB. genesis. Mon. Wea. Rev., 137, 1514–1532, doi:10.1175/ 2008MWR2650.1. Barnes, G. M., and K. Sieckman, 1984: The environment of fast- Acknowledgments. The author is very grateful to the and slow-moving tropical mesoscale convective cloud lines. Mon. Department of Physics, Bangladesh University of En- Wea. Rev., 112, 1782–1794, doi:10.1175/1520-0493(1984)112,1782: . gineering and Technology, Dhaka, for providing the TEOFAS 2.0.CO;2. Begum, S., M. S. Alam, and M. S. Ali, 2013: Cyclone monitoring laboratory facilities. The JTWC data and NCEP data towards disaster management of Bangladesh. Open J. Comput. were downloaded from their respective websites. The Sci., 1, 1–10. Grid Analysis and Display System Software (GrADS) Bhatia, R. C., and M. Rajeevan, 2008: Monsoon 2007—A report. was used for analyzing and displaying the data. Synoptic Meteorology Rep. 6/2008, National Climate Centre, India Meteorological Department, 135 pp. Bister, M., and K. Emanuel, 1997: The genesis of Hurricane REFERENCES Guillermo: TEXMEX analyses and a modeling study. Mon. Wea. Rev., 125, 2662–2682, doi:10.1175/1520-0493(1997)125,2662: Akter, N., and K. Tsuboki, 2012: Numerical simulation of Cyclone TGOHGT.2.0.CO;2. Sidr using a cloud-resolving model: Characteristics and for- Bluestein, H. B., 1993: Synoptic-Dynamic Meteorology in Mid- mation process of an outer rainband. Mon. Wea. Rev., 140, latitudes. Vol. 2. Oxford University Press, 594 pp. 789–810, doi:10.1175/2011MWR3643.1. Braun, S. A., M. T. Montgomery, K. J. Mallen, and P. D. Reasor, ——, and ——, 2014: Role of synoptic-scale forcing in cyclogenesis 2010: Simulation and interpretation of the genesis of Tropical over the Bay of Bengal. Climate Dyn., 43, 2651–2662, doi:10.1007/ Storm Gert (2005) as part of the NASA tropical cloud systems and s00382-014-2077-9. processes experiment. J. Atmos. Sci., 67, 999–1025, doi:10.1175/ Alappattu, D. P., and P. K. Kunhikrishnan, 2009: Premonsoon es- 2009JAS3140.1. timates of convective available potential energy over the oce- Bunkers, M. J., 2002: Vertical wind shear associated with left- anic region surrounding the Indian subcontinent. J. Geophys. moving supercells. Wea. Forecasting, 17, 845–855, doi:10.1175/ Res., 114, D08108, doi:10.1029/2008JD011521. 1520-0434(2002)017,0845:VWSAWL.2.0.CO;2. Atkins, N. T., and M. St. Laurent, 2009a: Bow echo mesovortices. Camargo, S. J., K. A. Emanuel, and A. H. Sobel, 2007: Use of a genesis Part I: Processes that inﬂuence their damaging potential. Mon. potential index to diagnose ENSO effects on tropical cyclone Wea. Rev., 137, 1497–1513, doi:10.1175/2008MWR2649.1. genesis. J. Climate, 20, 4819–4834, doi:10.1175/JCLI4282.1.

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