1. Introduction
Tropical cyclogenesis is a process by which a preexisting tropical disturbance develops into a warm-core, nonfrontal tropical cyclone (TC genesis hereafter). The large-scale environmental conditions favorable for TC genesis have been known since the pioneering work by Gray (1968). These conditions include adequately warm sea surface temperature, high relative humidity in the lower to middle troposphere, relatively low vertical wind shear, and sufficient convective instability. However, the detailed physical processes responsible for TC genesis remain elusive (Emanuel 2018; Tang et al. 2020). A tropical disturbance occurring under a favorable large-scale environment may but does not necessarily develop into a TC. Partially due to the incomplete understanding, it has been a lasting challenge for both regional and global models to accurately predict TC genesis (Majumdar and Torn 2014).
Central to the dynamics of TC genesis is the scale interactions between the preexisting tropical disturbance and the embedded deep moist convection. Ritchie and Holland (1997) and Simpson et al. (1997) hypothesized that one or more mesoscale convective vortices (MCVs) associated with mesoscale convective systems (MCSs) within the preexisting tropical disturbance feed into the large-scale cyclonic circulation and lead to TC genesis. Bister and Emanuel (1997) suggested a “top-down” dynamical pathway of TC genesis: the cooling effects of evaporation and melting below the MCSs’ stratiform clouds produce strong subsidence, which extends the midlevel MCVs downward to the boundary layer. However, observations showed that evaporation driven downward motions can rarely reach the boundary layer in developing tropical disturbances (e.g., Zipser and Gautier 1978; Houze et al. 2009). Alternatively, Hendricks et al. (2004) and Montgomery et al. (2006) proposed a vortical hot tower route to TC genesis. They pointed out that bursts of small-scale deep convection within a vorticity-rich mesoscale environment accelerate the large-scale near-surface circulation and result in TC genesis. This “bottom-up” dynamical pathway of TC genesis has been validated in many recent observation and modeling studies (e.g., Braun et al. 2010; Wang et al. 2010b,a; Wang 2012; Gjorgjievska and Raymond 2014; Bell and Montgomery 2019). For example, through examining the convective activity during the formation of Hurricane Ophelia (2005), Houze et al. (2009) noted that a developing tropical disturbance may include a mix of deep convective cells and MCSs in various life stages. Deep convective cells generate vorticity in the lower to middle troposphere, while the vorticity of midlevel MCVs were mainly derived from older deep convective cells. Hence, TC genesis can be summarized as a “bottom-up” process as deep convective cells are the main vorticity source.
Recent studies also provided abundant evidence that a deep and moist preexisting cyclonic disturbance is a necessary condition for the rapid enhancement of deep convective cells during tropical cyclogenesis (e.g., Nolan 2007; Dunkerton et al. 2009; Montgomery et al. 2010; Wang et al. 2010a). Dunkerton et al. (2009) indicated that a deep and moist preexisting cyclonic disturbance (i.e., a wave pouch) can protect the embedded deep convection from lateral dry-air intrusion and deformation by the large-scale horizontal and vertical wind shears. Repeated deep convection occurring within the wave pouch can gradually moisten and enhance the parent vortex and lead to TC genesis. Consistent with the “marsupial pouch paradigm” proposed by Dunkerton et al. (2009), Bell and Montgomery (2019) found that organized deep convection and the resulting stratiform clouds that occurred prior to the cyclogenesis of Hurricane Karl (2010) alternately enhanced and moistened the preexisting wave pouch and provided a favorable environment for the genesis of Karl. Most interestingly, the deep convective activity occurring a few days prior to the genesis of Karl showed a distinct diurnal cycle with the most intense deep convection occurring during the mid- to late mornings (Davis and Ahijevych 2012). Through sensitivity experiments, Melhauser and Zhang (2014) pointed out that radiational modulations on the convective inhibition and relative humidity of the preexisting wave pouch are possible physical processes leading to the diurnal cycle of pregenesis deep moist convection of Karl. In fact, many recent statistical (e.g., Chang et al. 2017; Zhang et al. 2022) and case studies (e.g., Molinari et al. 2004; Davis and Ahijevych 2012) have noted that quasi-diurnal cycle of organized deep convection is a common feature that occurs a few days prior to the rapid TC genesis. In addition, pregenesis deep convection does not always peak at the same local time (e.g., mid- to late morning) during different days prior to the genesis (e.g., Chang et al. 2017) and/or for different TCs (e.g., Davis and Ahijevych 2012). This implies that radiation might not be the only physical process contributing to the pregenesis diurnal cycle of organized deep convection. In this study, using Tropical Cyclone Mora (2017) as an example, we show that convectively forced diurnal gravity wave may also play a role in modulating the diurnal pregenesis deep convection and influencing the TC genesis.
In what follows, section 2 describes the observed genesis and development of Tropical Cyclone Mora. Section 3 introduces the high-resolution regional reanalysis dataset used in this study. Sections 4 and 5 describe the roles and physical mechanisms of the diurnal MCSs prior to the cyclogenesis of Mora, respectively. Conclusions and further discussion are given in section 6.
2. Tropical Cyclone Mora (2017)
Based on the Saffir–Simpson hurricane wind scale, TC Mora was a category 1 TC that formed over the Bay of Bengal (BOB) during the 2017 north Indian Ocean cyclone season. Figure 1 shows the track and intensity of Mora from the Joint Typhoon Warning Center (JTWC) best track database. Mora’s best track record begins at 0600 UTC 27 May (the local noon is 0600 UTC), when it was still a tropical depression with a maximum surface wind speed around 12.9 m s−1. From 0600 UTC 27 May to 0000 UTC 28 May, Mora was located over the central BOB and slowly moved east-northeastward (Fig. 1a). It intensified into a tropical storm at 1800 UTC 27 May, with its maximum surface wind speed reaching ∼18 m s−1 (Fig. 1b). After 0000 UTC 28 May, Mora moved to the north and became a category 1 TC at 1800 UTC 29 May. It reached its peak intensity at 0000 UTC 30 May with a maximum surface wind speed of around 41 m s−1 (Fig. 1b). Although its peak intensity was moderate, Mora was one of the deadliest TCs during 2017. It caused widespread devastation and severe flooding with 135 total fatalities when it made landfall over Bangladesh on 30 May.
(a) Track and (b) intensity of TC Mora (2017) from the JTWC best track database. The JTWC best track record of Mora starts from 0600 UTC 26 May 2017. The black dots in (a) show the 12-hourly positions of Mora from 1200 UTC 26 May to 1200 UTC 30 May.
Citation: Journal of the Atmospheric Sciences 80, 6; 10.1175/JAS-D-22-0203.1
TC Mora developed from a cyclonic tropical disturbance associated with an active Madden–Julian oscillation (MJO) event over the equatorial Indian Ocean. Emmanuel et al. (2021, E21 hereafter) provided detailed examination on the large-scale environment of TC Mora. Using the wavenumber–frequency filtering analysis, they found that the initial tropical cyclonic disturbance that became Mora was closely related to an equatorial Rossby wave associated with the MJO, while the MJO provided abundant lower- to midlevel moisture for the development of deep moist convection. Since the large-scale environment has been thoroughly discussed in E21, here we will only provide a brief description of the formation and development of TC Mora using satellite observations. Interested readers are referred to E21 for more detailed information.
Figure 2 shows the observed infrared brightness temperature (Tb) from the National Oceanic and Atmospheric Administration (NOAA)’s global merged geostationary satellite 11.5-μm infrared Tb product (Janowiak et al. 2017). Vectors are the 700-hPa horizontal winds from the ERA5 global dataset (Hoffmann et al. 2019). On 21 May, a westerly wind burst associated with the MJO event can be found over the equatorial Indian Ocean. The initial tropical cyclonic disturbance that became Mora was located to the north of the equator, with its lower-tropospheric circulation centered around 7°N, 92°E (E21). The tropical cyclonic disturbance moved to the west during 22–23 May and started moving to the northeast after 23 May due to the large-scale northeastward steering flow (not shown). Convection embedded within the cyclonic disturbance was weak and scattered during 24 May. An organized MCS developed over the western quadrant of tropical disturbance in the early morning (∼0000 UTC or ∼0600 LST) of 25 May. The convection weakened and was advected downstream and inward to the circulation center during the afternoon and evening (1800 UTC 25 May). In the early morning (0000 UTC) of 26 May, another MCS initiated and organized over the western quadrant of the tropical disturbance and then weakened during the daytime. A similar early morning MCS developed again on 27 May. With the reoccurrences of the early morning MCSs, the circulation of the tropical disturbance was gradually enhanced from 24 to 27 May. The rapid intensification of TC Mora occurred after 27 May, and the TC made landfall over the west coast of Bangladesh on 30 May at its peak intensity (Fig. 1b). Mora’s tropical cyclogenesis time is defined as 0600 UTC 27 May in this paper. This definition is based on two main reasons: 1) the JTWC best track of TC Mora starts from 0600 UTC 27 May, at which time Mora became a tropical depression with a well-defined surface circulation; 2) the vortex and convective structures of Mora showed noticeable changes after the tropical cyclogenesis time (to be discussed later). The physical mechanisms contributing to the diurnal MCSs prior to the genesis of TC Mora and the potential roles of the diurnal MCSs in the formation of Mora will be the focus of this paper.
Observed infrared brightness temperatures (color shading) from 1800 UTC 21 May to 0000 UTC 30 May 2017. Only the brightness temperatures lower than 250 K are shown. Vectors show the 700-hPa horizontal winds from ERA5.
Citation: Journal of the Atmospheric Sciences 80, 6; 10.1175/JAS-D-22-0203.1
3. High-resolution TMeCSR
A novel high-resolution tropical mesoscale convective system reanalysis (TMeCSR) is used in this study to investigate the diurnal MCSs prior to the cyclogenesis of TC Mora. TMeCSR is produced by assimilating all-sky infrared radiances and atmospheric motion vectors (AMVs) from geostationary satellites and other conventional observations into an MCS-resolving (9-km grid spacing) regional model using ensemble Kalman filter (Chan et al. 2022, C22 hereafter). Due to the assimilation of satellite all-sky infrared radiances, AMVs and the high model resolution, TMeCSR provides a better analysis of tropical MCSs and their surrounding environments compared to the global reanalyses (e.g., ERA5). The diurnal cycle, spatial distribution, intensity, and structure of tropical MCSs are better represented in TMeCSR compared to ERA5 (C22). Most importantly, C22 showed that the analyzed dynamical and thermodynamical fields in TMeCSR are dynamically consistent with the cloud fields. As a result, forecasts initialized from TMeCSR outperform the forecasts initialized from ERA5 for forecast lead times of more than 10 h. Interested readers are referred to C22 for detailed information on the configurations and validations of TMeCSR.
Figure 3 shows the simulated infrared Tb and 700-hPa horizontal winds from TMeCSR. Although an underestimation of cloud area is noticeable compared to the observation (Fig. 2), TMeCSR realistically captures the evolution of the circulation and convection of the pregenesis tropical disturbance (Fig. 3). In particular, the reoccurring early morning MCSs prior to the TC genesis are well simulated in TMeCSR. It indicates that TMeCSR could be a useful dataset to investigate the physical mechanisms of the pregenesis diurnal MCSs. Nevertheless, the limitations of using TMeCSR will be discussed in section 6.
As in Fig. 1, but for simulated infrared brightness temperatures (color shading) and 700-hPa horizontal winds (vectors) from TMeCSR.
Citation: Journal of the Atmospheric Sciences 80, 6; 10.1175/JAS-D-22-0203.1
4. Evolution of the wave pouch and vortex
Following some previous studies (e.g., Wang et al. 2010a; Davis and Ahijevych 2012), the diurnal MCSs prior to the genesis of TC Mora are analyzed in a pouch comoving frame. To get an approximate propagation speed of the large-scale tropical disturbance, we first performed a low-pass filter on the 700-hPa wind fields from TMeCSR to filter out perturbations with spatial scales smaller than 1000 km. The propagation speed of the large-scale disturbance is then estimated by tracking the filtered 700-hPa circulation center. The pouch-relative winds are calculated by subtracting the propagation speed of the large-scale tropical disturbance from the TMeCSR original 700-hPa winds. Then, the pouch center at each hour is determined by the intersection of the wave critical surface (where the pouch-relative zonal velocity is zero) and the trough axis (where the pouch-relative meridional velocity is zero) following Dunkerton et al. (2009). We visually checked the estimated pouch centers at each hour and found them well tracking the cyclonic circulation center from the early tropical disturbance stage to the mature TC stage (Fig. 4). We also varied the filtering length of the low-pass filter from 500 to 1500 km and found the tracking results are insensitive to the small changes of the filtering length (not shown).
Pouch-relative horizontal winds (vectors) and Okubo–Weiss (OW) values (shading) at 700 hPa.
Citation: Journal of the Atmospheric Sciences 80, 6; 10.1175/JAS-D-22-0203.1
The temporal evolutions of the meso-α- and meso-β-scale vortices of the wave pouch during 23–30 May are shown in Fig. 5. During this period, the preexisting tropical disturbance moved northward across the BOB and developed into a TC (Figs. 2 and 3). Following Wang (2012), the intensities of the meso-α- and meso-β-scale vortices are estimated as the mean relative vorticity averaged within 300 and 100 km from the pouch center, respectively. During both the pregenesis and postgenesis periods, the meso-α- and meso-β-scale vortices are roughly aligned with the pouch center, especially after 0000 UTC 25 May (not shown). Note that while the mean relative vorticity averaged within 300 and 100 km from the pouch center are mostly contributed by the local mesoscale vorticity, the larger scales also have noticeable contributions.
Temporal evolutions of relative vorticity averaged over a distance of (a) 300 and (b) 100 km from the pouch center. Black crosses denote missing values. Black dashed line indicates 0600 UTC 27 May. Ticks of the x axis indicate 0000 UTC.
Citation: Journal of the Atmospheric Sciences 80, 6; 10.1175/JAS-D-22-0203.1
The meso-α-scale vortex was gradually enhanced in the first few days (23–26 May), and its intensification became more rapid after 0600 UTC 27 May (i.e., the tropical cyclogenesis time, Fig. 5a). In addition, a distinct diurnal cycle of the meso-α-scale vortex can be found during 25–27 May, with the vortex started to intensify from early morning (∼0000 UTC) and reached its daily maximum intensity around early afternoon (∼0600 UTC). At the meso-β scale, the vortex also showed rapid intensification after 0600 UTC 27 May, but no clear enhancement can be found before 27 May (Fig. 5b). In addition, the enhanced meso-β-scale vortex showed a bottom-heavy structure with the cyclonic circulation peaked around 850-hPa after 0600 UTC 27 May. As the low-level vorticity continued to increase, the meso-β-scale cyclonic circulation also extended upward to the upper troposphere during 29–30 May (Fig. 5b). The evolutions of the meso-α and meso-β vortices are quite similar with that described in Wang (2012), especially during the postgenesis phase after 0600 UTC 27 May. Based on the vortex evolutions, we roughly separated the development of the wave pouch into two phases: the pregenesis (from 0000 UTC 23 May to 0600 UTC 27 May) and the postgenesis (from 0600 UTC 27 May to 0000 UTC 30 May) phases.
A gradual enhancement of the meso-α-scale vortex during the pregenesis phase (before 0600 UTC 27 May) can be clearly seen from the cumulative vorticity budget analysis (Fig. 6a), and a much faster intensification and deepening of the meso-α-scale vortex occurred after 0600 UTC 27 May during the postgenesis phase. In addition, a diurnal cycle of the meso-α-scale vortex can be found after 25 May, especially during 25–27 May. These features are consistent with Fig. 5a. The rapid intensification of the meso-β-scale vortex started after 0600 UTC 27 May (Fig. 6b), which was around the same time when the meso-α-scale vortex began to rapidly develop (Fig. 6a). This coenhancement of the meso-α- and meso-β-scale vortices during the TC formation has also been found in previous studies (e.g., Wang et al. 2010a; Wang 2012). The enhancement of the meso-β-scale vortex started from the surface and lower troposphere during 27–28 May, and then extended to the whole troposphere during 28–30 May. This TC formation process is consistent with the bottom-up formation theory (Hendricks et al. 2004; Montgomery et al. 2006). Previous studies suggested that a strong and moist preexisting meso-α-scale vortex is usually required for such a rapid bottom-up TC formation to happen (e.g., Nolan 2007; Bell and Montgomery 2019).
Meso-α-scale (a) relative vorticity tendency [
Citation: Journal of the Atmospheric Sciences 80, 6; 10.1175/JAS-D-22-0203.1
Term 1 was mostly responsible for the enhancement of the meso-α-scale vortex during both the pregenesis and postgenesis phases (Fig. 6c). In addition, term 1 showed a clear diurnal cycle after 25 May, similar to the temporal evolution of the vorticity (Fig. 6a). Later, we will show that the vortex stretching effect associated with the pregenesis early morning MCSs was the main process leading to the vortex enhancement. Term 2 of the meso-α-scale vortex was dominated by the diffusion and subgrid processes, which tended to slow down the lower-level cyclonic circulation (Fig. 5e). A similar frictionally induced slowdown process has also been noted by Wang et al. (2010a) and Wang (2012). One should note the diffusion and subgrid processes are calculated as the residual term of the budget as in Wang (2012). The residual term may also be partially contributed by the inadequate temporal resolution of reanalysis output (hourly) used to calculate the budgets and the linear updates of the ensemble Kalman filter.
At the meso-β scale, term 1 greatly contributed to the lower-level rapid enhancement of the vortex (Fig. 6d), mainly through the convergence of cyclonic vorticity (or the stretching effect). The diffusion and subgrid processes slowed down the cyclonic circulation in the lower level, while the tilting effect contributed to the enhancement of the meso-β-scale vortex in the middle to upper troposphere during 28–30 May (Fig. 6f). These processes are quite consistent with the vortical hot tower pathway to TC genesis proposed in Montgomery et al. (2006).
The burst of deep convection and the associated enhancement of the meso-β-scale vortex at the center (or sweet spot) of a wave pouch during TC formation has been extensively investigated in the previous studies. The open questions we want to explore in this study are as follows: What process contributed to the enhancement of the preexisting meso-α-scale vortex that provided a vorticity-rich environment for the following genesis of Mora? Why was there a distinct diurnal cycle of the meso-α-scale vortex evolution? To understand the physical mechanisms, we further examine the original meso-α-scale vorticity budget [Eq. (1)] from 1200 UTC 24 May to 0000 UTC 28 May (Fig. 7). A diurnal cycle of vorticity tendency can be found during 25–27 May (Fig. 7a). Positive vorticity tendencies started from early morning (0000 UTC). They first peaked in the lower troposphere, and then gradually shifted to the middle troposphere from morning to early afternoon. From afternoon to evening, a positive vorticity tendency can be found in the middle troposphere while a negative vorticity tendency can be found in the lower troposphere (Fig. 7a). The diurnal variations of the vorticity tendencies were mostly contributed by term 1 (Fig. 7b), while term 2 tended to slow down the circulation through the diffusion and subgrid processes (Fig. 7c). By further separating term 1 into the stretching effect (Fig. 7d) and the horizontal advection (Fig. 7e), we found that the diurnal cycle and gradual enhancement of the meso-α-scale vortex were mainly controlled by the vortex stretching effect, which was closely related to the latent heating produced by the pregenesis early morning MCSs (Figs. 2 and 7f). During the early life cycle stage of the early morning MCSs (∼0000 UTC), they showed a deep convective latent heating structure (heating between 700 and 300 hPa, Fig. 7f), which generated vorticity in the lower to middle troposphere (Figs. 7a,d). After stratiform became more dominant in their later life stage (after 0300 UTC), these MCSs transitioned to a top-heavier latent heating structure (heating above 500 hPa, Fig. 7f) with the positive vorticity tendency mostly confined within the middle troposphere (Figs. 7a,d). To better show the latent heating structure changes, the mean 0000 UTC (i.e., 0600 LST) and 0600 UTC (i.e., 1200 LST) latent heating profiles averaged during 25–27 May are further compared in Fig. 7g. We can see that the latent heating profile showed a deep convective latent heating structure between 700 and 300 hPa during the early morning (0000 UTC, red line in Fig. 7g), and then transitioned to a top-heavier latent heating structure at noon (0600 UTC, blue line in Fig. 7g). Note that the stretching effect also produced positive vorticity tendencies near the surface and negative vorticity tendencies in the lower troposphere (∼600–700 hPa) during the evening, which cannot be explained by the stratiform precipitation (Figs. 7d). The vorticity tendency pattern was likely due to the mix of different precipitation processes during the evening, including cumulus congestus (Wang 2014). In summary, the diurnal MCSs played an important role in the gradual enhancement of the pregenesis meso-α-scale vortex, which provided a vorticity-rich mesoscale environment for the rapid development of inner core deep convection during the TC formation.
Temporal evolutions of the meso-α-scale (a) relative vorticity tendency, (b) term 1, and (c) term 2 of isobaric vorticity budget. (d),(e) The stretching [
Citation: Journal of the Atmospheric Sciences 80, 6; 10.1175/JAS-D-22-0203.1
5. Diurnal MCSs and gravity waves
The physical mechanisms responsible for the diurnal cycle of the pregenesis MCSs over the western quadrant of the wave pouch are investigated in this section. Figure 8 shows the evolution of the column-integrated cloud water and ice (blue contours), pouch-relative 700-hPa winds (vectors), and infrared Tb (color shading). The column-integrated cloud water and ice are used to indicate the location and intensity of moist convection (Melhauser and Zhang 2014). Green vectors show the environmental vertical wind shear, which is defined as the difference between the mean wind vectors at the 200- and 850-hPa levels over an outer region extending from the radius of 200–800 km around the pouch center following Chen et al. (2006).
Infrared brightness temperature (color shading), column-integrated cloud water and ice (blue contours; interval: 0.3 kg m−2 starting from 0.3 kg m−2), environmental vertical wind shear (green vectors) and 700-hPa translated horizontal winds (black vectors) from 1800 UTC 24 May to 1200 UTC 27 May 2017.
Citation: Journal of the Atmospheric Sciences 80, 6; 10.1175/JAS-D-22-0203.1
The gradual enhancement of the meso-α-scale vortex during the pregenesis phase occurred under a moderate to strong (∼12 m s−1) vertical wind shear. In addition, the vertical wind shear played an important role in modulating the asymmetric initiation and development of moist convection. During nighttime, scattered moist convection can be found over the downshear quadrant (from −45° to +45° relative to the vertical wind shear direction) of the pouch (first column of Fig. 8). In early morning, moist convection over the downshear quadrant became organized, and mostly located at the leading edge of mesoscale cold clouds (second column of Fig. 8). During 25 and 26 May, the organized moist convection became scattered, and was advected to the downstream quadrants from morning to noon (third column of Fig. 8). Some moist convection was also advected to the center of the wave pouch but did not last long due to the strong vertical wind shear and relatively weak pouch vortex. On 27 May, with the meso-α-scale vortex being gradually enhanced, the moist convection organized during the early morning was able to persist for a longer time (last two panels).
The impacts of the vertical wind shear on the diurnal organized moist convection (or MCSs) can be clearly seen from the time-azimuth diagram (Fig. 9a). Moist convection was initiated and organized over the downshear quadrant (from −45° to +45° relative to the vertical wind shear direction, the vertical wind shear direction is shown by the magenta line) of the pouch during the early mornings of 25–27 May, and then advected downstream by the cyclonic circulation. Some early morning moist convection was also advected inward to the pouch center (Fig. 9b). During the postgenesis phase (after 27 May), bursts of deep convection occurred within 100 km from the pouch (Fig. 9b), which is consistent with the vorticity budget analysis shown in the last section. Previous studies have shown that the shear induced asymmetric convection distribution is a common feature of TCs or pregenesis wave pouches (e.g., Marks et al. 1992; Frank and Ritchie 2001; Molinari et al. 2004; Chen et al. 2006; Molinari et al. 2006; Rappin and Nolan 2012; Uddin et al. 2021), which might be related to quasi-steady lifting associated with a balanced mesoscale vortex in shear (Raymond and Jiang 1990; Trier et al. 2000). Observation studies also showed that environmental wind shear can induce an asymmetric distribution of moist instability, with a larger CAPE value over the downshear quadrant of a cyclonic circulation (e.g., Molinari and Vollaro 2010). We compared the daily mean CAPE averaged over the downshear and upshear quadrants of the wave pouch (averaged within 300 km from the pouch center). The results show that the downshear quadrant generally had a higher CAPE than the upshear quadrant during the pregenesis phase (Fig. 9c). The daily mean CIN averaged over the downshear and upshear quadrants of the wave pouch were both small (lower than 15 J kg−1), while the CIN in the downshear was slightly greater than that in the upshear quadrant (not shown). The quasi-balanced lifting arising from the interaction of the ambient vertical shear with the wave pouch (Raymond and Jiang 1990; Trier et al. 2000) and the favorable thermodynamic conditions over the downshear quadrant likely led to the asymmetric distribution of convective initiation.
(a) Time–azimuth diagram of column-integrated cloud water and ice averaged between 100 and 300 km from the pouch center. Solid magenta line shows the temporal evolution of vertical wind shear direction (i.e., downshear direction). 0°, 90°, 180°, and 270° indicate the north, east, south, and west directions. Dashed magenta lines show the downshear quadrant (from −45° to +45° relative to the vertical wind shear direction). (b) Time–radius diagram of averaged cloud water and ice over the downshear quadrant. (c) Daily mean CAPE averaged over the downshear (blue) and upshear (brown) quadrants.
Citation: Journal of the Atmospheric Sciences 80, 6; 10.1175/JAS-D-22-0203.1
Then, the next open question to answer is why a distinct diurnal cycle of MCSs existed in the downshear quadrant of the pregenesis wave pouch. Analyses show that the diurnal MCSs were closely related to the diurnal cycle of tropospheric temperatures. Figure 10a shows the time–radius diagram of the mean column-integrated saturation specific humidity averaged within the pouch’s downshear quadrant. The mean column-integrated saturation specific humidity increased from afternoon to evening and peaked at ∼1700 UTC (2300 LST), which indicates a warming of the troposphere. Organized moist convection occurred during early to late mornings when the tropospheric temperatures were reduced (indicated as smaller values of the column saturation specific humidity). The warming of the troposphere also contributed to the reduction of the column saturation fraction (CSF hereafter, which is defined as the ratio between the column-integrated specific humidity and the column-integrated saturation specific humidity) within the pouch’s downshear quadrant from morning to evening (Fig. 10a). For instance, the mean downshear CSF averaged between 100 and 300 km from the pouch center decreased by ∼4% (from ∼0.852 to ∼0.818) from 0000 to 1700 UTC 25 May. This change was dominated by the increase of the column-integrated saturation specific humidity (increased by ∼5%), while the column-integrated specific humidity increased by ∼0.6% during the same period of time. Past studies showed that the formation of tropical MCSs and their precipitation are closely related to the environmental CSF (e.g., Bretherton et al. 2004; Ahmed and Schumacher 2015; Wolding et al. 2020; Chen et al. 2022b). High CSF limits the negative effects of dry-air entrainment on the plume buoyancy, and facilitates the transition to deep convection and organization of tropical MCSs (e.g., Raymond 2000; Neelin et al. 2009; Wang 2012; Ahmed and Neelin 2018; Schiro and Neelin 2019). In addition, Nolan (2007) and Bell and Montgomery (2019) indicated that high CSF is a necessary condition for TC genesis. With the occurrence of the diurnal MCSs, the CSF within 300 km from the pouch center increased gradually during 24–27 May (Fig. 10c). The moist and vorticity-rich pouch after 0000 UTC 27 May provided a favorable environment for the rapid development of deep convection during the postgenesis phase (Fig. 6). The evolutions of convection and moisture over the inner wave pouch region (Figs. 9b and 10b) were similar to that in Wang (2014), which indicated the transition from equatorial wave to TC was marked by a dramatic increase in CSF and vorticity generation by deep convection in the inner wave pouch region. Besides the CSF, the diurnal cycle of tropospheric temperatures also modified the atmospheric stability within the pouch. Figure 10c shows the 2-day-mean temperature difference between 1700 and 0000 UTC during 24–26 May [i.e., (24/17–25/00 + 25/17–26/00)/2.0]. The temperature difference is averaged within the downshear quadrant and within 300 km from the pouch center. It can be found that the late evening warming of the troposphere peaked in the middle troposphere (∼500 hPa). As a result, the lower- to midtropospheric stability increased from afternoon to evening. The diurnal variations of the stability might be another thermodynamical condition leading to the diurnal cycle of pregenesis MCSs over the downshear quadrant of the wave pouch.
(a) Time–radius diagram of averaged saturation specific humidity (color shading) and column-integrated cloud water and ice (black contours; interval: 0.3 kg m−2 starting from 0.3 kg m−2) over the downshear quadrant. (b) As in (a), except color shading shows the azimuthally averaged CSF over the downshear quadrant. (c) Mean temperature difference between 1700 and 0100 UTC during 24–26 May.
Citation: Journal of the Atmospheric Sciences 80, 6; 10.1175/JAS-D-22-0203.1
The diurnal tropospheric warming within the pouch was unlikely to be caused by the diurnal radiation. The tropospheric temperature peaked in the late evening (∼1700 UTC), while the column-integrated radiative heating peaked during the early afternoon (∼0700–0800 UTC, not shown). To understand the possible physical mechanisms, we further examine the diurnal variations of the 500-hPa temperature prior to the genesis of Mora. Figure 11 shows the diurnal variation during 25 May as an example. A noticeable interesting feature is the southward propagation of the warm 500-hPa temperature from the northern coast of the BOB to the wave pouch during the afternoon to evening period. Deep convection (using IR Tb as proxies, black contours in Fig. 11) developed and the 500-hPa temperature became warmer over the inland side of the northern BOB coast during the afternoon (0800–1100 UTC). The warm midlevel temperature gradually propagated to the BOB from afternoon to evening (1100–1700 UTC) and arrived in the pouch region at ∼1700 UTC. After 1700 UTC, the midlevel temperature over the pouch region was gradually reduced (2000–0500 UTC).
Spatial distributions of 500-hPa temperature (color shading) and horizontal winds (vectors) from 0800 UTC 25 May to 0500 UTC 26 May. Black dots show the pouch center location. Black circles are 300 km from the pouch center. Black contours show IR Tb starting from 250 K with a −10-K interval. To clearly show the temperature propagation over the ocean, only IR Tb over the northern BOB inland area is shown. Narrow blue boxes are regions used in analysis shown in Fig. 12.
Citation: Journal of the Atmospheric Sciences 80, 6; 10.1175/JAS-D-22-0203.1
This diurnally southward propagating warm temperature occurred almost every day prior to the genesis of Mora (Fig. 12a). Since the northward movement of the wave pouch is slow prior to the TC genesis (gray line in Fig. 12a), the diurnal midlevel warm temperature always arrived in the pouch region at ∼1600–1800 UTC during 24–26 May. Figure 12b shows the 2-day-mean diurnal perturbation of the 500-hPa temperature from 0000 UTC 24 May to 2300 UTC 25 May. Diurnal perturbations are defined relative to the daily mean temperature in each latitude belt. The mean latitude of the pouch center during these two days is shown by the gray line in Fig. 12b. The 500-hPa temperature over the inland side of the northern coast started to increase from ∼0700 UTC. A southward propagation of the positive temperature diurnal perturbation can be found from afternoon to early morning. The diurnal cycle of the 500-hPa temperature over the pouch region is largely influenced by this southward propagation, with a maximum midlevel temperature occurring at ∼1700 UTC, which is consistent with Fig. 10a.
(a) Hovmöller diagram of 500-hPa temperature zonally averaged along the meridian of the blue boxes in Fig. 10. The solid black line shows the location of the coastline. The solid gray line represents the latitude of the pouch center. (b) Hovmöller diagram of the 2-day-mean 500-hPa temperature diurnal perturbations during 24–25 May. Diurnal perturbations are defined relative to the daily mean temperature in each latitude belt. The diurnal cycle is repeated twice for clarity.
Citation: Journal of the Atmospheric Sciences 80, 6; 10.1175/JAS-D-22-0203.1
The diurnal southward propagation of the 500-hPa warm temperature was not caused by the horizontal advection of the midlevel prevailing flow. The 500-hPa cyclonic circulation center of the pouch was located to the west of the 700-hPa circulation center due to the environmental vertical wind shear (Fig. 11). The mean 500-hPa winds were weak southerly over the northern Bay of Bengal (BOB, >18°N) and were easterly over the central BOB during 0800–1700 UTC 25 May (Fig. 11), while the estimated propagation of the warm temperature was southward at a speed of ∼14.8 m s−1 (Fig. 12b). To explore the possible physical mechanisms, we further examine the cross sections of the 2-day-mean temperature diurnal perturbations during 24–25 May along the meridian of the blue box in Fig. 11. As seen in Fig. 13, in the early morning (0100–0300 UTC), the inland area of the northern BOB coast showed a negative temperature diurnal perturbation due to the radiative cooling effect. During noon to early afternoon, the land heated up and a deep mixed layer developed (0500–0700 UTC). Inland convection initiated and developed in the afternoon (0700–1500 UTC). Convective latent heating excited positive temperature diurnal perturbations which peaked in the middle troposphere. Similar convectively forced diurnal temperature anomaly has also been noticed in other tropical coastal or island regions (e.g., Mapes et al. 2003; Ruppert et al. 2020a). The positive temperature perturbation gradually propagated from the coastal region to the BOB from afternoon to evening, and arrived in the pouch region around 1700 UTC, which is consistent with the analysis of Figs. 9c and 11b. The positive temperature perturbation gradually decayed in the next few hours and MCSs began to initiate and organize over the pouch region during early morning. The diurnal propagating pattern is similar to diurnal gravity waves that have been found over the west coast of South America (Mapes et al. 2003). The propagation of the temperature perturbation is also consistent with theory. The propagation speed of a linear hydrostatic gravity wave for an atmosphere at rest is equal to the vertical wavelength divided by the Brunt–Väisälä period. Based on Fig. 13, the vertical wavelength during 1100–1500 UTC is around 9 km. Assuming the Brunt–Väisälä period is around 10 min in the tropical lower to middle troposphere (Mapes et al. 2003), the resulting theoretical propagation speed of the gravity wave should be ∼15 m s−1. The result is quite consistent with the estimated propagation speed in Fig. 11b.
Cross sections of 2-day-mean temperature diurnal perturbations during 24–25 May zonally averaged along the meridian of the blue boxes in Fig. 11. The solid gray line represents the latitude of the mean pouch center location. Black shading denotes topography.
Citation: Journal of the Atmospheric Sciences 80, 6; 10.1175/JAS-D-22-0203.1
To further verify if the temperature diurnal perturbations were associated with gravity waves, the 2-day-mean meridional wind diurnal perturbations during 24–25 May are also analyzed. The meridional wind diurnal perturbations show a consistent southward propagation as the temperature diurnal perturbations (Figs. 13 and 14). In addition, the meridional wind diurnal perturbations are approximately in quadrature with the temperature diurnal perturbations. These characteristics are in agreement with the theoretical gravity wave structures (e.g., Kiladis et al. 2009; Ruppert et al. 2020a). One should also note that the midtropospheric warm perturbations over the pouch region transitioned to cold perturbations in the early morning (Fig. 12b). The ascending motion promoted by the diurnal gravity waves might also have contributed to the initiation of early morning convection within the wave pouch (Ruppert et al. 2022). While the roles of convectively forced diurnal gravity waves in modulating offshore moist convection has been documented in previous studies (e.g., Mapes et al. 2003), it is interesting to see that they also play a potential role in regulating the diurnal MCSs prior to the TC genesis.
As in Fig. 13, but for 2-day-mean meridional wind diurnal perturbations during 24–25 May.
Citation: Journal of the Atmospheric Sciences 80, 6; 10.1175/JAS-D-22-0203.1
6. Summary and discussion
In this study, diurnal MCSs that occurred a few days prior to the genesis of Mora are studied using a high-resolution regional reanalysis (TMeCSR). We found that the spatiotemporal evolution of the pregenesis MCSs is influenced by the environmental vertical wind shear and convectively forced diurnal gravity waves. The early morning MCSs persistently occurred over the downshear quadrant of the preexisting wave pouch. Inland afternoon deep convection over the northern BOB excited gravity waves, which propagated southward toward the ocean. The warm temperature perturbations associated with the diurnal gravity waves reached the downshear quadrant of the preexisting wave pouch during afternoon to evening. The tropospheric warming induced by the diurnal gravity waves reduced the column saturation fraction and increased the static stability within the wave pouch. We hypothesize that these environmental changes hindered the initiation and organization of deep convection during afternoon to evening. As a result, a distinct diurnal cycle of convective activity can be found within the wave pouch and most MCSs were organized during early morning. The vorticity budget shows that the pregenesis diurnal MCSs enhanced the meso-α-scale vortex and the column saturation fraction of the pouch, therefore provided a favorable mesoscale environment for the development of deep convection and the meso-β-scale vortex during the TC formation.
While it is interesting to see that the convectively forced gravity waves may play a role in the diurnal cycle of pregenesis convection, one should note that the relationship between the diurnal gravity waves and the pregenesis convection is inferred from a regional reanalysis data. Future modeling studies examining how exactly the convective initiation and organization within the preexisting tropical disturbance respond to the diurnal gravity waves are needed to conclusively demonstrate the impacts of diurnal gravity waves on the genesis of Mora. More importantly, TC genesis (or intensification) usually occurs through a fairly steady progression, which can be punctuated by a diurnal modulation. This study explores the potential physical mechanisms responsible for the diurnal variations of pregenesis convection but does not examine whether the diurnal convection plays a major role in determining the occurrence of TC genesis. This is an important open question to be answered. At 9-km resolution, TMeCSR might be able to capture the primary characteristics of tropical MCSs and TCs (e.g., Chan et al. 2020; Ruppert and Chen 2020; Chen et al. 2021b,a, 2022a), but even higher resolutions are needed to understand the detailed physical processes at the storm scale. To better understand the roles of pregenesis diurnal convection in tropical cyclogenesis, future investigations will require systematic high-resolution sensitivity experiments (e.g., by modulating diurnal radiation or convective activity), which are beyond the scope of this study.
In addition, Mora is a TC formed over a bay area enclosed by coastlines. Strong afternoon convection frequently occurred along the coastlines of the BOB during the summer season. This might be a reason why the convectively forced diurnal gravity waves played an important role in modulating the pregenesis convection of Mora. While the diurnal cycle of pregenesis convection was in phase with the gravity waves, we cannot completely rule out the potential impacts of the diurnal radiation cycle. Recent studies have shown that cloud–radiative feedbacks play a crucial role in tropical cyclone development (Ruppert et al. 2020b). The diurnal radiation cycle can also lead to a nocturnal enhancement of pregenesis convection through the modulations on atmospheric stability and convective inhibition (Randall et al. 1991; Melhauser and Zhang 2014). Future studies that systematically investigate the relative roles of diurnal gravity waves and cloud–radiative feedbacks in modulating pregenesis convection and TC formation are also important.
Acknowledgments.
We thank Dr. Rich Rotunno for insightful discussions on this work. We thank Editor Dr. Zhuo Wang and three anonymous reviewers for their helpful comments, which helped us strengthen the science of our study and improve the quality and clarity of the manuscript. This study is supported by the Office of Science of the U.S. Department of Energy Biological and Environmental Research as part of the Regional and Global Model Analysis program area, the Office of Naval Research (ONR) Grant N00014-18-1-2517, and the National Aeronautics and Space Administration (NASA) Grant 80NSSC22K0613. The computations were carried out using the computing resources at the National Energy Research Scientific Computing Center (NERSC) and the Texas Advanced Computing Center (TACC). PNNL is operated for the Department of Energy by Battelle Memorial Institute under Contract DE-AC05-76RL01830.
Data availability statement.
The ERA5 dataset is available at http://apps.ecmwf.int/datasets/. The Global Merged IR data are obtained from the NASA Goddard Earth Sciences Data and Information Services Center: https://doi.org/10.5067/P4HZB9N27EKU/. The TMeCSR dataset is available at https://www.datacommons.psu.edu/commonswizard/MetadataDisplay.aspx?Dataset=6298.
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