Cript - Subsidence Warming in the Tropical Cyclogenesis of Cindy (2017): CPEX observations and Coupled Modeling - Hurricanes and Coupled ...
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1 Subsidence Warming in the Tropical Cyclogenesis of Cindy (2017): CPEX
2 observations and Coupled Modeling
3 Edoardo Mazza∗ and Shuyi S. Chen
4 Department of Atmospheric Sciences, University of Washington, Seattle
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Submitted to J. Atmos. Sci.
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∗ Corresponding author: emazza2@uw.edu
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ABSTRACT
6 The formation of tropical cyclones (TC) in unfavorable large-scale environments remains a chal-
7 lenge for TC forecasting. Tropical Storm (TS) Cindy (2017) formed at 1800 UTC 20 June in
8 the Gulf of Mexico despite strong vertical wind shear, low mid-tropospheric relative humidity,
9 and poorly organized convection. A key to TC genesis is the initial development of a warm
10 core within an emergent cyclonic vortex, a process which occurs on small spatial scales and is
11 often difficult to observe. TS Cindy was observed during the Convective Processes Experiment
12 (CPEX) field campaign in 2017 by the NASA DC-8 aircraft, equipped with a Doppler wind lidar,
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13 precipitation radar, and GPS dropsondes. This study combines CPEX observations and a cloud-
14 resolving, fully-coupled atmosphere-wave-ocean numerical simulation to investigate the formation
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15 of TS Cindy. Prior to TC genesis, a shallow cyclonic circulation was embedded in a deep layer
16 of west-southwesterly flow associated with an upper-level trough. Within the disturbance, a warm
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17 and dry anomaly was observed by dropsondes near the center of the cyclonic circulation, with
a maximum at about the 2.5 km level. The temperature perturbation reaches 5°C along with a
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19 dew point temperature depression of 8°C in the coupled model simulation. Backward trajectory
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20 analysis shows that subsidence is primarily associated with a thermally indirect circulation along
the western flank of the storm. Air parcels descend more than 1000 m towards the lower tropo-
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22 sphere while warming up by 9-12°C. The subsidence-induced virtual temperature perturbation in
23 the 1.5-3.5 km layer accounts for 50 % of the sea-level pressure depression. Subsidence warming
24 therefore played a key role in the genesis of TS Cindy.
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25 1. Introduction
26 The genesis of tropical cyclones is a multiscale process that involves the transformation of a
27 precursor disturbance into a warm core, low-pressure system with a closed surface circulation.
28 Precursor disturbances can be tropical waves in the tropical basins (Frank and Roundy 2006), long-
29 lasting mesoscale convective systems (MCSs) or cloud clusters (Kerns and Chen 2013), monsoon
30 lows, Central-American gyres (Papin et al. 2017) or have an extra-tropical origin (Davis and Bosart
31 2003). TC genesis is facilitated when a set of large-scale conditions are met, such as low vertical
32 wind shear, a moist mid-troposphere, a vorticity-rich low troposphere, a deep and warm ocean
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33 mixed layer (Gray 1968; McBride and Zehr 1981) and a large thermodynamic disequilibrium
34 between the tropopause and the sea surface (McTaggart-Cowan et al. 2015). TC genesis involves
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35 two fundamental processes: the amplification and organization of cyclonic vorticity, and the
36 formation of a warm core vortex. Pre-existent, low-level cyclonic vorticity can be amplified and
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37 axisymmetrized by vortical hot towers (Hendricks et al. 2004; Montgomery et al. 2006). Mid-
tropospheric vortices can instead result from diabatic heating in the stratiform region of long-lived
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39 MCSs (Chen and Frank 1993) or from evaporative cooling in the precipitating region of MCSs
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40 (Bister and Emanuel 1997). The formation of the warm core is supported by diabatic heating in
the convective and stratiform cloud region (Chen and Frank 1993; Dolling and Barnes 2012a). An
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42 observational study by Kerns and Chen (2015) showed that subsidence warming associated with
43 MCSs can contribute directly to development of warm core vortex in TC genesis.
44 About 60 % of all TC genesis events in the North Atlantic from 1948 to 2010 involved a varying
45 degree of baroclinicity (McTaggart-Cowan et al. 2013). In the Western Caribbean Basin and
46 Gulf of Mexico, TC genesis often occurs in unfavorable environments, with upper-tropospheric
47 disturbances enhancing the vertical wind shear and promoting mid-tropospheric dry air intrusions
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48 (Gray 1968). Bracken and Bosart (2000) found that a pronounced upper-tropospheric trough-ridge
49 pattern is often associated with TC genesis events in the Bahamas and Gulf of Mexico. Strong
50 vertical wind shear is considered to be unfavorable for TC genesis as observed in the Atlantic and
51 the west Pacific (McBride and Zehr 1981; Kerns and Chen 2013), while weak-to-moderate westerly
52 shear can instead assist TC genesis (Bracken and Bosart 2000; Nolan and McGauley 2012; Reasor
53 and Montgomery 2001). Wind shear induces significant structural changes in TCs, which can
54 hinder their development or intensification: idealized experiments (Jones 1995; Frank and Ritchie
55 1999, 2001) and observational studies (Rodgers et al. 1994; Black et al. 2002; Chen et al. 2006;
Corbosiero and Molinari 2002; Reasor et al. 2013) indicate that the structure of sheared TCs is
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57 highly asymmetric, as the deepest convection focuses in the downshear quadrants. Another factor
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58 that greatly affects the formation of TCs is mid-tropospheric humidity (Malkus 1958; Gray 1975;
59 McBride and Zehr 1981). TC genesis is favored by high relative humidity in the mid-levels, whereas
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60 intrusions of dry air can delay or suppress TC development (e.g. Dunion and Velden 2004; Wang
61 2012; Kerns and Chen 2013). Ventilation of dry and cool air into the developing TC disrupts its
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62 thermodynamic structure and suppresses convection by reducing the updrafts buoyancy (Simpson
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63 and Riehl 1958; Shelton and Molinari 2009; Riemer and Montgomery 2011; Tang and Emanuel
2012; Ge et al. 2013). Importantly, how TCs develop in relatively unfavorable environmental
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65 conditions over the Gulf of Mexico remains an active area of research.
66 Subsidence-driven temperature anomalies associated with oceanic convective systems have been
67 documented since the GATE field campaign (Houze 1977; Zipser 1977). Simpson et al. (1997)
68 argue that subsidence is the only viable process for maintaining a low-level warm anomaly in clear
69 air. Descent is typically observed in the form of unsaturated downdrafts underneath the anvil canopy
70 of long-lived MCSs. Chen and Frank (1993) showed the presence of a wake low in a region of
71 low-tropospheric subsidence on the edge of a simulated MCS. In mature tropical cyclones, instead,
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72 observed subsidence-induced warm anomalies are often attributed to vortex tilting (Halverson
73 et al. 2006; Heymsfield et al. 2006; Shelton and Molinari 2009). The relationship between wind
74 shear and asymmetric warm anomalies in TCs has been investigated by Tao and Zhang (2019),
75 which describe how the alignment of a mid-tropospheric and upper-tropospheric warm anomalies
76 is often seen as the storm approaches the onset of rapid intensification. A small number of studies
77 focuses on the role of subsidence during TC genesis and almost exclusively investigated the role of
78 subsidence within MCSs. Dolling and Barnes (2012b) and Dolling and Barnes (2012a) describe
79 how subsidence helped the formation of a lower-tropospheric warm core in TS Humberto (2001)
by inducing a hydrostatic pressure drop and by capping the boundary layer, allowing for a buildup
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81 of high equivalent potential temperature air, which was later ingested in the nascent eyewall.
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82 Stossmeister and Barnes (1992) document the formation of a second circulation center in TS Isabel
83 (1982) underneath a region of mesoscale subsidence. Finally, subsidence warming during TC
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84 genesis was also captured by dropsondes released in typhoons Megi and Fanapi during the Impact
85 of Typhoons on Ocean in Pacific (ITOP) field campaign (Kerns and Chen 2015).
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86 Is the presence of long-lived MCSs the only pathway to low-level subsidence warming during
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87 TC genesis? In this study, we focus on the genesis of TS Cindy (2017) to describe how subsidence-
induced warm anomalies in the lower troposphere can support TC formation in a very different
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89 dynamic and thermodynamic context. TS Cindy developed from a poorly-organized, broad cy-
90 clonic disturbance embedded in a high-shear, low mid-tropospheric relative humidity environment.
91 Airborne observations collected during the NASA Convective Processes Experiment (CPEX) field
92 campaign reveal a low-level warm anomaly formed within a shallow cyclonic circulation in a
93 cloud-free region, well removed from convective clusters and their associated anvils. Using a com-
94 bination of aircraft observations and a convection-resolving, fully-coupled atmosphere-wave-ocean
95 simulation, we address three main questions: a) what are the spatial and temporal characteristics
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96 of the subsidence-induced temperature perturbation, b) does the subsidence-induced warming pro-
97 duce a significant pressure perturbation during TC genesis? and c) what mechanisms drive the
98 descent? The thermodynamic and kinematic properties of the disturbance are investigated using
99 dropsonde and airborne wind lidar retrievals, complemented by a high-resolution model simulation
100 to overcome their limited spatial and temporal sampling. The model simulation is then employed
101 to perform backward trajectory analysis and to diagnose the driving mechanisms of subsidence.
102 The paper is organized as follows: section 2 is an overview of the meteorological evolution of TS
103 Cindy, the data and methodology employed in the study are described in section 3. The results are
presented in section 4, 5, 6, 7 and 8. A discussion of the results and their implications is included
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105 in section 9.
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106 2. Overview of Tropical Storm Cindy
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107 The National Hurricane Center (NHC) describes Cindy as a large, sprawling TS that formed
108 on 20 June. Its genesis was preceded by the interaction of two consecutive tropical waves with a
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109 Central-American Gyre (CAG). The first wave reached the Caribbean Sea on 15 June, while the
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110 second wave moved into the basin on 18 June. The second wave featured deep convection on
its eastern flank (Fig.1a) along with low-level cyclonic vorticity. The interaction with the CAG
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112 produced a cyclonic disturbance in the central Gulf of Mexico on 19 June 2017 characterized by
113 an elongated circulation, with wind speed exceeding 17 m s−1 (Berg 2018). At this stage, the
114 convective activity was primarily focused along the eastern flank of the disturbance (Fig.1b).
115 On 20 June, multiple low-level vorticity local maxima merged into a coherent center as the deep
116 convection attained a curved structure around it (Fig.1c), prompting the NHC to declare the genesis
117 of TS Cindy at 1800 UTC 20 June 2017. TS Cindy formed in an unfavorable environment: an
118 upper-level cut-off in the northwestern Gulf of Mexico advected mid-tropospheric drier air into
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119 the Gulf and enhanced the vertical wind shear in the area prior to TC genesis. As a result, TS
120 Cindy was characterized by a broad surface wind field, with an exposed low-level circulation and
121 asymmetrically-distributed convection (Fig.1c).
122 On the next day, deep convection rotated towards the NW quadrant (Fig.1d) as TS Cindy
123 intensified to reach a peak intensity of 50 kt on 21 June 0000 UTC. TS Cindy made landfall at
124 0700 UTC on 22 June just west of Cameron, LA. While inland, TS Cindy weakened to TD status
125 and finally dissipated on 24 June 0600 UTC in the mid-Atlantic states. The impact of TS Cindy
126 was primarily due to excessive rainfall: widespread accumulations in the 7-10 in. range were
measured in south-eastern Mississippi, southwestern Alabama and part of the Florida panhandle,
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128 with a maximum of 18.69 in. at Ocean Springs, MS.
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129 3. Data and methods
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130 This study uses a combination of aircraft observations from the CPEX field campaign, satellite
131 observations, and a fully coupled atmosphere-wave-ocean high resolution numerical simulation.
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132 For model comparison purposes, the track and intensity of TS Cindy are obtained by linearly inter-
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133 polating the Best-Track dataset (Landsea and Franklin 2013) to hourly interval. High-frequency,
2-minute storm center fixes from the NHC are used for the CPEX observations analysis.
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135 a. CPEX Field Campaign − Aircraft Observations
136 The CPEX field campaign took place in the North Atlantic-Gulf of Mexico-Caribbean Sea
137 region in May-June 2017. It was designed to study convective processes in the tropics using the
138 NASA DC-8 aircraft observations (Chen and Zipser 2018). Four research flights from 17-21 June
139 were conducted to capture the development of TS Cindy from its precursor tropical wave in the
140 western Caribbean Sea. On 17 June and 19 June, the flights targeted some of the convective
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141 elements embedded in the precursor disturbance. On 20 June, the NASA DC-8 airplane sampled
142 the kinematic and thermodynamic structure of TS Cindy right as the disturbance was classified as
143 TS by the NHC (Fig.2a). Its mature phase was captured by the 21 June mission (Fig.2b).
144 The wind and thermodynamic profiles analyzed in this study were obtained from the Doppler
145 Aerosol WiNd lidar (DAWN) and by Yankee Environmental System (YES) dropsondes (Black et al.
146 2017). DAWN is a coherent-detection, wind-profiling lidar system that emits a pulsed signal with
147 a 2-micron wavelength (Kavaya et al. 2014). The instrument retrieves wind speed and direction at
148 a vertical resolution of approximately 65 m. YES dropsondes measure wind and thermodynamic
variables as they descend through the atmosphere at a 2 Hz frequency. YES dropsondes have
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150 been previously employed during the Tropical Cyclone Intensity (TCI) field campaign (Doyle et al.
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151 2017). In the 20 June CPEX mission, 254 DAWN wind profiles were obtained and 16 dropsondes
152 were launched (Fig.2a). 28 dropsondes and 528 DAWN wind profiles were recorded in the 21 June
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153 CPEX mission (Fig.2b).
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154 b. UWIN-CM simulation
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155 The Unified Wave INterface-Coupled Model (UWIN-CM, Chen et al. 2013; Chen and Curcic
2016) is employed to perform a fully-coupled atmosphere-wave-ocean simulation of TS Cindy.
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157 The UWIN-CM consists of the Weather Research and Forecasting model (WRF, Skamarock et al.
158 2008), the University of Miami Wave Model (UMWM, Donelan et al. 2012) and the Hybrid
159 Coordinate Ocean Model (HYCOM, Bleck 2002). The WRF model is configured with an outer
160 domain with two nested grids with horizontal grid spacings of 12, 4, and 1.3 km, respectively.
161 There are 44 vertical levels. The outer domain covers an area of 6,468 km (E-W) x 4,320 km (N-S).
162 The inner-most 1.3-km nest (523 km x 523 km) is storm-following (Fig.3). The Kain-Fritsch
163 cumulus scheme (Kain 2004) is used in 12-km outer domain, while in both the 4- and 1.3-km nests
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164 convection is resolved explicitly. For all domains, the WRF single-moment, 6-class microphysics
165 scheme (WMS6, Hong and Lim 2006) and the YSU PBL scheme (Hong et al. 2006) are used.
166 The horizontal grid spacing of UMWM is 4 km and 36 frequency bins are used in the spectral
167 computations. HYCOM is run at a 1/25 °(approximately 4 km) horizontal grid spacing and 41
168 vertical levels. Initial and boundary conditions for WRF are provided by the ERA5 reanalysis
169 (Hersbach and Dee 2016), while the HYCOM model is initialized from HYCOM global analysis
170 fields. The simulation is initialized at 1200 UTC 19 June and terminates at 0000 UTC 23 June.
171 The 1.3-km moving nest is initialized 6 hours into the simulation at 1800 UTC 19 June. Analysis
nudging is applied every 6 hours to the wind fields in the 12-km WRF domain to improve the
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173 simulated track of TS Cindy.
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174 For the analysis of the storm structure, the UWIN-CM WRF output is linearly interpolated onto
175 constant height levels with a vertical spacing of 50 m. A storm-tracking algorithm is used to
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176 calculate the storm position and intensity using hourly model output. The algorithm locates the
177 850 hPa geopotential height minimum and calculates the storm intensity as the corresponding
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178 minimum sea-level pressure and maximum wind speed. Following the NHC official report, the
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179 time of TC genesis in the simulation is taken to be 1800 UTC 20 June 2017.
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180 c. Hydrostatic Pressure Perturbation
181 The sea-level pressure perturbation (SLP’) due to changes in the virtual temperature (Tv ) is
182 calculated using the hydrostatic equation along with the ideal gas law, in a manner consistent
183 with Stossmeister and Barnes (1992); Dolling and Barnes (2012a); Kerns and Chen (2015). SLP0
184 at each model grid point can be estimated using equation 1, where the Top Of the Atmosphere
185 (TOA) is the upper limit of integration and the overbar denotes domain-averaged quantities. The
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186 contribution from a specific atmospheric layer (e.g. 1-2 km) can be calculated by changing the
187 limits of integration in equation 1.
Z TO A
0 −g p(z) p(z)
SLP = − dz (1)
0 Rd Tv (z) Tv (z)
188 d. Trajectory Analysis
189 Backward trajectories are calculated by adapting a code developed for the Cloud
190 Model 1 (CM1, Bryan and Fritsch 2002) to work on WRF output. The origi-
nal code is available at https://github.com/tomgowan/trajectories/blob/master/
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192 trajectoriesCM12ndorder.ipynb. It is based on the work of Miltenberger et al. (2013)
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193 and uses a second-order, semi-implicit discretization in space and time. The backward trajectory
194 calculations are performed on 10-minute model output from a twin numerical simulation of TS
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195 Cindy, restarted at 0600 UTC 20 June 2017 from the parent experiment described in section 3b.
196 Three sets of parcel trajectories are initialized within the subsidence-induced warm anomaly at the
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197 elevations of 2500, 2100 and 1700 m.
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198 4. UWIN-CM Simulation of TS Cindy
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199 The storm track and intensity TS Cindy in the UWIN-CM simulation are evaluated against the
200 Best-Track dataset (Landsea and Franklin 2013). The minimum SLP is used for the comparison as
201 it is a more reliable measure of the storm position and intensity than peak winds, in particular in
202 the early stages of development. As shown in Fig.3, the observed track is sufficiently well captured
203 by the model simulation up to the NHC genesis time (1800 UTC 20 June). Later on, the simulated
204 track oscillates around the observed one, showing an anticipated and accentuated recurving to
205 the west. The average track error is 104.7 km. TS Cindy is slower in the simulation than in the
10206 observation: it makes landfall approximately 80 km east and 7 hours later than the observed storm
207 (Figs.3, 4a). The UWIN-CM simulated minimum SLP closely tracks the observation until about
208 0600 UTC 22 June when the observed Cindy made landfall, whereas the simulated storm remained
209 over the ocean because of its slower motion (Fig.4a). As a result, the simulated minimum SLP at
210 landfall is 6 hPa lower than observed. Overall, the root mean square error (RMSE) of the minimum
211 SLP from 1200 UTC 19 June to 1200 UTC 22 June is 2.9 hPa.
212 One of the most prominent large-scale environmental conditions during the development of TS
213 Cindy is strong wind shear. To assess the UWIN-CM simulation of the large-scale environment,
we compute the deep-layer (200-850 hPa) wind shear, averaged in a 200-800 km annulus around
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215 the storm and compare it to the one calculated from the ERA5 reanalysis. Prior to TC genesis
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216 (vertical dashed line), the UWIN-CM simulation correctly reproduces the observed high-wind
217 shear environment, with values well above 20 m s−1 (Fig.4b). After 00 UTC 21 June, the ERA5
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218 wind shear declines rather steadily, reaching a minimum of 11 m s−1 at 12 UTC 22 June. The
219 UWIN-CM wind shear compares well with the ERA5 one, with an overall RMSE of just 1.3 m
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220 s−1 , however there are two differences: the observed reduction in wind shear is anticipated by
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221 approximately 6 hours and the wind shear is generally weaker after TC genesis between 0000 UTC
21 June and 0000 UTC 22 June.
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223 The excessive shear reduction coincides with the storm recurving to the west and slowing down.
224 We speculate this could be associated with a more efficient rearrangement of the upper-level flow
225 in the simulation possibly due to diabatic effects. Given the similarity between the observed and
226 simulated track and intensity during the CPEX flights (Fig.3, Fig.4a) and the fact that most of the
227 analysis is performed in storm-relative coordinates, we do not expect these differences to influence
228 the results presented in this study.
11229 5. Subsidence warming
230 Subsidence in an atmospheric layer is revealed by an enhanced dew point temperature depression
231 and increased static stability, often large enough to produce a temperature inversion. Its footprint
232 on the skew T-log(p) diagram is the so-called thermodynamic “onion profile” (Zipser 1977; Houze
233 1977). In this section we examine the evidence of organized subsidence during the genesis of TS
234 Cindy in the CPEX observations and UWIN-CM simulation.
235 a. Observations from the CPEX dropsondes
The 20 June CPEX mission reveals important thermodynamic features involved in the genesis
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237 of TS Cindy. Three dropsondes were released shortly after the TS classification by the NHC at
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238 1908, 1948 and 2029 UTC 20 June (square markers in Fig.5) in an area largely cloud-free, where
239 the lower-tropospheric temperature was higher than in the surrounding environment. All these
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240 dropsondes are located within 100 km from the storm center (24, 69 and 54 km away respectively)
241 and indicate the presence of low-level subsidence warming within the developing disturbance. The
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242 thermodynamic diagram in Fig.6a shows that the subsidence is maximized around the 800-825 hPa
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243 level, where the dew point temperature depression approaches 10 °C and a cyclonic circulation
exists (wind barbs in Fig.5). Below that level, an approximately isothermal layer extends down
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245 to 900 hPa. A second inversion is also observed at 600 hPa. In the mid-to-upper troposphere,
246 the dry-air intrusion associated with the trough is revealed by dew point temperature depressions
247 exceeding 20 °C. On 20 June, 13 dropsondes were released in the near and far environment, more
248 than 100 km away from the storm in the drier environment along the western flank of the storm or
249 in proximity of precipitating clouds in the eastern flank. Their average temperature and dew-point
250 temperature profiles (Fig.6b) are in stark contrast with those near the storm center: they do not
251 show signs of organized low-level subsidence such as deep layers of temperature inversion and
12252 enhanced dew-point temperature depression. The temperature perturbation within the developing
253 TS Cindy is estimated by subtracting the mean temperature profile of the environment (Fig.6b)
254 from that of the inner disturbance ( Fig.6a). The result is shown in Fig.6c: a positive anomaly of
255 3.84 °C is collocated with the subsidence at 825 hPa, while a second anomaly (3.97 °C) is present
256 just below 600 hPa. The dropsondes data is scarce above 450 hPa but the temperature perturbation
257 profile suggests that the system does not have a well-defined warm core above 500 hPa. The
258 observed thermodynamic structure of TS Cindy indicates that near its genesis time (1900 UTC
259 20 June) a subsidence-induced positive temperature perturbation, maximized at 800-825 hPa, is
located within a developing cyclonic circulation.
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261 The thermodynamic structure of TS Cindy changed remarkably after its genesis. The dropsondes
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262 launched during the 21 June CPEX mission (Fig.7) reveal that, in its mature stage, TS Cindy is
263 characterized by a warm anomaly close to its center that extends from 890 hPa to the upper
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264 troposphere, with a maximum of 2.4 °C located at approximately 650 hPa. Some shallow inversion
265 layers can be observed in the individual dropsondes, as it can be expected next to the convection
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266 of a developing eyewall. The average thermodynamic profiles, however, do not exhibit a clear
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267 subsidence signature, suggesting that the sinking motion lacks the strength and the organization
observed on 20 June.
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269 b. UWIN-CM simulation - Temperature Perturbation
270 To assess the presence of low-level warming in the simulation of TS Cindy, we calculate the
271 temperature perturbation at each vertical level by subtracting the corresponding domain-average
272 temperature and search for its maximum value within 100 km from the storm center. The time-
273 height diagram in Fig.8a shows how the maximum temperature perturbation evolves from 0000
274 UTC 20 June to 0000 UTC to 22 June. Prior to 1200 UTC 20 June, the disturbance is characterized
13275 by a moderately positive temperature perturbation of 2-4 °C between 2000-8000 m. These warm
276 features are generally short lived and lack a coherent vertical structure. Starting at 1200 UTC 20
277 June, the maximum temperature perturbation sharply increases in the 4000 - 4500 layer and extends
278 downward to approximately 2000 m. A second pulse of lower-tropospheric warming starts before
279 0000 UTC 21 June. These warm anomalies have common characteristics: they are maximized
280 just above 2000 m, have magnitudes exceeding 5.5 °C and originate above 4000 m. In the lower
281 troposphere (around 2000 m), the maximum temperature perturbation grows by more than 3.5 °C
282 in 12 hours. The model simulation thus indicates that during the genesis of TS Cindy subsidence
produced a coherent, long-lasting warm anomaly within the developing disturbance. Later on 21
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284 June, the model portraits a structure characterized by a more elevated temperature perturbation,
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285 largely consistent with the dropsondes collected during the 21 June CPEX mission.
286 The importance of subsidence in the genesis of TS Cindy is further suggested by the time series
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287 of minimum sea-level pressure (Fig.8b): the growth of the lower tropospheric warm anomaly
288 is accompanied and followed by a 7 hPa pressure fall from 1001 hPa to 994 hPa. Once the
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289 warm anomaly dissipates, the minimum sea-level pressure stabilizes around 994 hPa and remains
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290 stationary for several hours afterwards.
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291 6. Kinematic structure
292 As discussed in section 5, both the observations and the model simulation indicate that a
293 subsidence-induced warm anomaly in the lower troposphere occurred prior and during the genesis
294 of TS Cindy between 1200 UTC 20 June and 1000 UTC 21 June. To understand the context in
295 which the subsidence occurs, we analyze the circulation associated with the system along three
296 vertical cross sections at different phases:
297 1) Pre-Subsidence (0800 UTC 20 June)
14298 2) During Subsidence (1900 UTC 20 June)
299 3) Post-Subsidence (2000 UTC 21 June)
300 To do so, we complement the CPEX DAWN wind profiles and dropsondes with the UWIN-CM
301 simulation. CPEX observations are used to validate the simulated structure of TS Cindy. The
302 model simulation provides a complete picture where CPEX observations are scarce: above the
303 flight level, in between dropsondes and in regions of strong lidar attenuation. Due to the absence
304 of airborne measurements, only model fields are presented for phase 1. For phases 2 and 3, the
CPEX flights legs (1855 - 1922 UTC 20 June and 1950-2022 UTC 21 June) are reproduced in the
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306 model output.
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307 The cyclonic circulation associated with the precursor disturbance is clearly evident in the model
308 cross section (Fig.9a, b). In phase 1, the circulation was confined in the lowest 3 km of the
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309 troposphere and featured a 120 km-wide area of low winds at its center. A southwesterly jet located
310 between 9 and 12 km approached the disturbance from the west, imposing a large wind shear
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311 gradient over the disturbance (Fig.9a).
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312 Phase 2 was characterized by persistent subsidence within the disturbance. The 20 June CPEX
flight leg intersected the storm center at 1900 UTC, shortly after the NHC classified the system
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314 as TS Cindy. The model cross section (Fig.9c, d) and CPEX observations (Fig.10a, c) show a
315 similar slanted region of low winds at the center of the circulation along with a mid-tropospheric jet
316 located at 7 km. Compared to phase 1, the low-level cyclone has strengthened. Both observations
317 and modeling suggest that the wind speed exceeds 25 m s−1 on the western flank of the circulation
318 (Fig.10d). In this stage, the upper-level jet is directly above the disturbance (Fig.10c). A notable
319 difference can be seen approximately 7 km above the disturbance: while both CPEX observations
320 and model simulation indicate the presence of a mid-tropospheric jet, its position is not entirely
15321 consistent. Airborne wind measurements show the maximum located directly above the western
322 flank of the circulation whereas in the UWIN-CM it is located above the center of the disturbance.
323 On 21 June the CPEX flight leg passed to the south of the storm center after the bulk of the
324 subsidence had occurred. Since the DAWN lidar retrievals suffered from severe attenuation in
325 the middle troposphere, the analysis of phase 3 relies more on the model simulation. Following
326 the period of sustained low-level subsidence, the model indicates that TS Cindy is now more
327 axisymmetric (Fig.9e) with a deeper cyclonic circulation whose radius of maximum wind has
328 reduced to approximately 60 km. The model cross section indicates that an eyewall-like wind
structure is also present. There are however discrepancies between the observed and the model
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330 storm structure, in particular the dropsondes indicate lower wind speeds between 4-6 km compared
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331 to the model cross-section (Fig.10d).
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332 7. Key Processese contributing to subsidence
333 To better understand the physical processes that contributed to subsidence warming in TS Cindy,
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334 we analyze the spatial and temporal evolution of the air parcels that undergo substantial warming
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335 during the genesis of TS Cindy. We do so by performing a backward trajectory analysis as discussed
in section 3d. Three sets of trajectories are initialized within the subsidence-induced warm anomaly
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337 in the lower troposphere at 1800 UTC 20 June at the elevations of 2500, 2100 and 1700 m, from
338 parcels whose temperature exceeds 18, 20, 22 °C respectively (Fig.11). The three initial levels are
339 selected to represent the behavior of parcels near the top, the center and the bottom of the warm
340 anomaly.
341 Backward trajectories reveal that intense subsidence begins for all three sets of parcels between
342 1100 UTC and 1300 UTC 20 June at an elevation between 3-3.5 km. As shown in Fig.12, prior
343 to starting their descent, the parcels originated primarily from two distinct airstreams: one rising
16344 from the lowest levels of the troposphere, and one emanating from the mid-levels. Once the parcels
345 reach the northwest quadrant of the storm, they begin subsiding towards the lower troposphere; in
346 doing so, they become increasingly warmer than their surrounding environment, reaching terminal
347 values above 6 °C (Fig.12b). The bulk of the subsidence is confined along the western flank of
348 the storm, during a 5-hour period between 1300 UTC and 1800 UTC. We can estimate some key
349 characteristics of the subsidence experienced by the air parcels. The median descent for parcels
350 initialized at 1700, 2100, and 2500 m is estimated to be 1090 m, 1130 m and 1380 m respectively,
351 with an accompanying temperature increase of 9.9 °C, 10.3 °C and 12 °C. The resulting lapse rates
along the trajectories vary between 8.7 °C km−1 and 9.1 °C km−1 , with sinking rates between -6
M
352
353 cm s−1 and -7.6 cm s−1 .
an
354 The known mechanism supporting subsidence within MCSs or squall lines relies on the evapo-
355 ration of hydrometeors in the stratiform precipitation region (e.g. Houze 1977; Zipser 1977; Chen
us
356 and Frank 1993). The trajectories show that the subsidence is focused along the western flank
357 of the circulation, downstream of the precipitating region (Figs.12, and 13). Parcels initiate their
cr
358 descent underneath the anvil clouds associated with the convection on the northern flank of the
ip
359 storm (Fig.12). In that region, the evaporation of hydrometeors can help initiate the subsidence.
Although this process may have contributed to the initial descending motion of the air parcels
t
360
361 in the stratiform region, the continued descending motion in the nearly convection-free region
362 may be forced by other processes. The prolonged descent of positively buoyant parcels suggests
363 the presence of a thermally indirect circulation. The presence of this ageostrophic circulation is
364 diagnosed via the Pettersen kinematic frontogenesis as defined in Bluestein (1993). As shown in
365 Fig.11a, a large temperature gradient is present in the lower troposphere in the northwest quad-
366 rant of the storm. Colder air wrapping around the developing cyclonic circulation gives rise to a
367 localized band of frontogenesis in the lower troposphere (Fig.14a). The position of this banded
17368 region is consistent with the convection observed in the GOES infrared imagery just few hours
369 later (Fig.5). During subsidence, the parcels are located downstream of such band, in a region
370 characterized by weak frontolysis (Fig.14a). The cross section through the frontolytical region
371 shows that it is associated with an area of subsidence (i.e. negative vertical velocity) in which the
372 parcels are embedded (Fig.14b). Such a mechanism relies on the divergence and deformation of
373 the flow, acting on the existing horizontal temperature gradient, which force subsidence to restore
374 thermal wind balance. It is worth noting that a similar mechanism has been proposed to explain the
375 descent of mid-tropospheric air in the sting jets of deep, marine extra-tropical cyclones (Schultz
and Sienkiewicz 2013; Martínez-Alvarado et al. 2014; Coronel et al. 2016).
M
376
an
377 8. Decreasing sea-level pressure from subsidence warming
378 Descent produces low level warming and drying, resulting in increased static stability and
us
379 enhanced dew point temperature depression. When the subsidence is sufficiently organized, a
380 surface meso-low can emerge (Zipser 1977; Chen and Frank 1993). As discussed in section 3c,
cr
381 a change of virtual temperature in the atmospheric column will result in a pressure perturbation.
ip
382 The presence of low-level cyclonic vorticity can determine whether the pressure perturbation is
retained or dissipated by gravity waves. In this section, we investigate the pressure perturbation
t
383
384 induced by the subsidence in the UWIN-CM simulation.
385 The lower tropospheric warming and drying within TS Cindy is shown in Fig.15. Early in the
386 genesis stage (0900 UTC), the disturbance does not display a well-organized warm anomaly. As
387 the subsidence develops within the disturbance, however, the temperature rapidly increases in the
388 inner vortex, reaching values above 18 °C at 1800 UTC 20 June. The warm anomaly is located
389 very close to the SLP minimum and is then advected along its southern and eastern flank over
390 time. This temperature perturbation is retained for several hours following TC genesis up to 0000
18391 UTC 21 June. As the dropsondes in Fig.10a suggests, subsidence-induced warming also results
392 in a consistent drying of the air mass, with low relative humidity values and an enhanced dew
393 point temperature depression. The simulated temperature perturbation is indeed associated with
394 a significant dew point temperature depression (Fig.15, middle column). At 0900 UTC 20 June,
395 the dew point temperature depression is smaller than 2 °C over most of the domain and does not
396 display a coherent spatial organization. In response to the persistent subsidence, the dew point
397 temperature depression increases near the center of the disturbance and exceeds 8 °C at 1800 UTC.
398 A primary objective of this study is to quantify the direct contribution of the subsidence warming
to the TC genesis in Cindy. To estimate how much of the total pressure perturbation is accounted
M
399
400 for by the low-level warm anomaly, we compute the hydrostatic pressure perturbation by integrating
an
401 Equation 1 for two atmospheric layers: i) the SLP pressure perturbation is computed by integrating
402 Eq.1 from the surface to the top of the atmosphere, ii) the pressure perturbation due to the low-level
us
403 warming is estimated by integrating Eq.1 from 1500 m to 3500 m. As shown in Fig.15 (right
404 column), the genesis of TS Cindy is associated with a deepening of the SLP perturbation: as the
cr
405 storm acquires a more axisymmetric look, the total SLP perturbation grows from -3 hPa at 0900
ip
406 UTC 20 June to -6 hPa at 1800 UTC 20 June and -7 hPa at 0000 UTC 21 June. This 4 hPa drop
from is consistent with the deepening of the SLP minimum displayed in Fig.10b. The pressure
t
407
408 perturbation due to the subsidence-induced warm anomaly during TC genesis is shown by the black
409 contours in figure 15. At 0900 UTC 20 June, its contribution is very limited. As the warming
410 occurs, the pressure perturbation in the 1500 - 3500 layer grows up to 3.2 hPa. At 1800 UTC 20
411 June it accounts for more than 50 % of the total perturbation. As most of the warming occurs prior
412 to 1800 UTC 20 June, the contribution of the 1500-3500 m layer does not grow significantly further
413 and remains approximately 3 hPa. Throughout the genesis, the 1500-3500 m pressure perturbation
19414 is closely colocated with the position of the SLP minimum and contributes significantly to the
415 deepening and axisymmetrization of the low-pressure system.
416 9. Summary and Conclusion
417 The presence of subsidence-induced low-level warm anomalies has often been linked to a weaken-
418 ing of tropical disturbances due to suppressed convection and mixing of lower equivalent potential
419 temperature air into the eyewall (Shelton and Molinari 2009). It is usually argued that subsidence
420 is due to the interaction between its circulation and the environmental flow (Heymsfield et al.
2006; Halverson et al. 2006) and results from the differential advection of cyclonic vorticity by the
M
421
422 environmental flow. Conversely, this study builds on Stossmeister and Barnes (1992), Dolling and
an
423 Barnes (2012b), Dolling and Barnes (2012a) and Kerns and Chen (2015) to argue that subsidence-
424 induced low-level warming can be a key process supporting tropical cyclogenesis. In this study
us
425 we investigate the genesis of TS Cindy from a broad, cyclonic circulation that moved from the
426 Caribbean Sea into the southern Gulf of Mexico (Fig.1), in an environment characterized by high
cr
427 vertical wind shear and low mid-tropospheric moisture.
ip
428 Throughout the genesis process, the structure of the disturbance changed from an initially shallow,
broad, asymmetric vortex (Fig.16a) to a vertically aligned, axisymmetric cyclone with a tighter
t
429
430 eyewall-like feature around its warm core (Fig.16c). Both the CPEX observations and UWIN-CM
431 modeling indicate that the subsidence-induced low-level warming occurred and persisted within
432 the shallow cyclonic disturbance as it organized into a TS. The temperature perturbation estimated
433 to be between 3.8 °C and 6 °C respectively and maximized in the atmospheric layer between 1500-
434 3500 m. The vertical location of the subsidence-induced warm anomaly in TS Cindy is consistent
435 with what previous studies have found in other developing TCs (Halverson et al. 2006; Heymsfield
436 et al. 2006; Dolling and Barnes 2012a). Its magnitude is also comparable to that observed by
20437 Stossmeister and Barnes (1992) in TS Isabel (1982), Heymsfield et al. (2006) in TS Chantal,
438 by Dolling and Barnes (2012a) in hurricane Humberto and by Shelton and Molinari (2009) in
439 hurricane Claudette.
440 Backward trajectories calculated from the UWIN-CM simulation indicate that subsidence is
441 focused in the western flank of the disturbance (Fig.16b). Air parcels located in the northern sector
442 of the storm at an elevation between 3-3.5 km start to descend as they exit a precipitating region and
443 move into a drier environment. The subsidence rates are calculated to be between -6 cm s−1 and
444 -7.6 cm s−1 . By integrating the hydrostatic equation we estimate that the lower tropospheric warm
anomaly accounts for approximately 3 hPa or 50% of the overall SLP depression when the storm
M
445
446 was classified as TS by the NHC. Such a pressure perturbation is also consistent with previous
an
447 studies. Stossmeister and Barnes (1992) estimated from TS Isabel that a 3-4 km deep layer, having
448 a temperature perturbation of 2 to 3 °C, would result in a pressure perturbation of 2 hPa. Similarly,
us
449 Dolling and Barnes (2012a) calculated that a 7 °C temperature perturbation would result in a 5 hPa
450 pressure perturbation.
cr
451 We describe how subsidence can produce a significant pressure perturbation in a disturbance
ip
452 vastly different from a typical long-lived MCS, one where a pre-existing, shallow and broad
cyclonic circulation is present but lacks the spatial organization required for TC classification. Due
t
453
454 to the shallow nature of the cyclonic disturbance during TC genesis, the deep-layer shear did not
455 produce an appreciable vortex tilt in TS Cindy, as confirmed by both the observed and modeled
456 storm structure, hence the subsidence cannot be interpreted as a response to vortex tilting. We
457 suggest instead that the subsidence results from two processes. The parcels initiate their descent
458 underneath the anvil clouds associated with the convection on the northern flank of the storm,
459 there the evaporation of hydrometeors can initiate the subsidence, as described by Houze (1977),
460 Zipser (1977) and Chen and Frank (1993). As the parcels move away from the precipitation, the
21461 subsidence is supported by a thermally indirect circulation linked to an area of weak frontolysis
462 along the western flank of the storm. A similar process has been proposed as the driving mechanism
463 behind sting jets in deep, marine extratropical cyclones (Schultz and Sienkiewicz 2013; Martínez-
464 Alvarado et al. 2014; Coronel et al. 2016). We acknowledge that our analysis does not rule out
465 possible concurring processes that might drive the subsidence.
466 TC genesis events, especially in the early part of the hurricane season or in the subtropics often
467 feature environments characterized by large wind shear, dry mid-tropospheric air and horizontal
468 temperature gradients in the lower troposphere. It is therefore possible that tropical or subtropical
disturbances embedded in these environments could undergo a similar physical process. This
M
469
470 study thus allows us to better understand the occurrence of subsidence during the genesis of TCs
an
471 in unfavorable environments by providing a comprehensive analysis of the subsidence-induced
472 temperature, humidity and pressure anomalies that could guide more targeted aircraft observations
us
473 during future events. Additional studies are needed to understand how frequently this process is
474 observed and its storm-to-storm variability.
cr
ip
475 Acknowledgments. We thank the CPEX science team for their support during the field campaign,
especially Dr. G. D. Emmitt and Mr. S. Greco for providing the DAWN wind data. The authors
t
476
477 are thankful to Dr. Ed Zipser and two anonymous reviewers’ whose constructive comments and
478 suggestions helped improve the manuscript. This research was supported by two NASA research
479 grants, CPEX (80NSSC18K0185) and CYGNSS (80NSSC18K0713).
480 Data availability statement. All the CPEX aircraft observations used in this study are publicly
481 available on the CPEX data repository (https://tcis.jpl.nasa.gov/data/cpex/). The
482 ERA-5 reanalysis data can be obtained from the CDS repository (DOI:10.24381/cds.bd0915c6).
22483 Access to the UWIN-CM simulation output along with the analysis codes on the Prof. Chen’s
484 group server can be easily obtained by contacting the corresponding author (emazza2@uw.edu).
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