The 6 May 2010 Elevated Supercell During VORTEX2

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The 6 May 2010 Elevated Supercell during VORTEX2

CHRISTOPHER W. MACINTOSH AND MATTHEW D. PARKER North Carolina State University, Raleigh, North Carolina

(Manuscript received 23 August 2016, in ﬁnal form 3 April 2017)

ABSTRACT

An elevated supercell from the second Veriﬁcation of the Origins of Rotation in Tornadoes Experiment (VORTEX2)on6May2010isinvestigated. Observations show that the supercell formed over a stable inversion and was likely decoupled from the surface. Quintessential features of a supercell were present, including a hook echo (albeit bent anticyclonically) and midlevel mesocyclone, and the storm was quasi steady during the observing period. A weak surface cold pool formed, but it was apparently devoid of air originating from midlevels. Idealized modeling using near-storm soundings is employed to clarify the structure and maintenance of this supercell. The simulated storm is decoupled from the surface by the stable layer. Additionally, the reﬂectivity structure of the simulated supercell is strikingly similar to the observed storm, including its peculiar anticyclonic-curving hook echo. Air parcels above 1 km reached their LFCs as a result of the simulated supercell’s own dynamic lifting, which likely maintained the main updraft throughout its life. In contrast, low-level air in the simulation followed an ‘‘up–down’’ trajectory, being lifted dynamically within the stable layer before becoming strongly negatively buoyant and descending back to the surface. Up–down parcels originating in the lowest 100 m are shown to be a potential driver of severe surface winds. The complementary observations and simulations highlight a range of processes that may act in concert to maintain supercells in environments lacking surface-based CAPE.

1. Introduction currents and preexisting boundaries (e.g., Carbone et al. 1990), or lifting by density currents, bores, and gravity On 6 May 2010, the second Verification of the Origins waves (e.g., Rotunno et al. 1988; Koch and Clark 1999; of Rotation in Tornadoes Experiment (VORTEX2; Marsham and Parker 2006; Morcrette et al. 2006; Ziegler Wurman et al. 2012) collected what may be the first storm- et al. 2010; Marsham et al. 2011). The lifting mechanism following dataset for an elevated supercell.1 In this study, likely differs on a storm-by-storm basis and depends on convection is considered to be elevated when it receives micro-, meso-, and synoptic-scale features. The majority all of its inflow parcels from above the near-surface layer of studies on elevated convection have focused on meso- [i.e., the lowest 500 m AGL, as in Parker (2008)]. This scale convective systems, which may be due to the lack of occurs most typically in an environment without surface- elevated supercell observations from past field projects. In based CAPE, but with appreciable instability present aloft this study, we focus on the unique dynamics that may (Colman 1990a,b; Grant 1995; Moore et al. 1998; sustain a supercell over a stable layer. Thompson et al. 2003; Horgan et al. 2007; Corfidi et al. Many convective storms are maintained, at least in part, 2008). For a storm to form over a stable layer, parcels with by lifting at the edge of a cold pool (outflow). However, it is instability aloft need to be lifted to their levels of free not clear that surface cold pools occur in many elevated convection so that the elevated instability can be released. systems. It is possible that cold outflow could be produced This can be accomplished, for instance, by frontal lifting prior to stabilization (e.g., Parker 2008; Ziegler et al. 2010), (e.g., Moore et al. 1998), interactions between gravity or that in situ cooling of the stable layer could occur if the relative humidity is suitably low. Alternatively, Bryan and Weisman (2006) simulated the presence of an elevated cold 1 A supercell is a thunderstorm with a rotating updraft (e.g., pool (2–4-km layer) above a stable layer, which worked to Lemon and Doswell 1979). continually release elevated instability by lifting air along its leading edge. Bores may also provide sufficient lift to Corresponding author: Christopher W. MacIntosh, cwmacwx@ sustain elevated convection; Parker (2008) found in gmail.com his idealized simulations of MCSs in environments with

DOI: 10.1175/MWR-D-16-0329.1 Ó 2017 American Meteorological Society. For information regarding reuse of this content and general copyright information, consult the AMS Copyright Policy (www.ametsoc.org/PUBSReuseLicenses). Unauthenticated | Downloaded 09/28/21 06:44 AM UTC 2636 MONTHLY WEATHER REVIEW VOLUME 145 low-level cooling that as the convection became elevated, a Nowotarski et al. (2011) produced the first compre- ‘‘gravity-wave-like bore’’ maintained the convective up- hensive study of elevated supercells by using a numerical drafts. For supercells, the enhanced vertical perturbation model to compare surface-based supercells to those over pressure gradient acceleration has also been shown to stable boundary layers. Interestingly, most of the simu- help maintain elevated convection above stable layers lated ‘‘elevated’’ supercells in the Nowotarski et al. (2011) (Nowotarski et al. 2011; Billings and Parker 2012). It is study were still fueled by parcels from the near-surface possible that storm maintenance mechanisms vary from layer, except in the case of extreme stability [experiments case to case even for a given storm type. 500mC, 1kmA, 1kmB, and 1kmC in Nowotarski et al. Heavy precipitation (with associated flash flooding) and (2011)]. In addition, midlevel downdrafts were able to large hail are seemingly the most common hazards associ- reach the surface except in the case with the greatest ated with elevated convection (e.g., Grant 1995; Moore stability [experiment 1kmC in Nowotarski et al. (2011)]. et al. 1998; Horgan et al. 2007), but it appears that severe Within the stable layer, weaker convergence and baro- surface winds are at least possible (e.g., Grant 1995; clinic vorticity generation were found by Nowotarski et al. Schmidt and Cotton 1989; Knupp 1996; Bernardet and (2011), leading to weaker near-surface vertical vorticity. Cotton 1998; Bryan and Weisman 2006). The processes This is one possible reason for the apparently hindered behind the production of severe surface winds in elevated tornado potential in elevated supercells. Notably, the convection are somewhat mysterious, with various studies Nowotarski et al. (2011) study used statically stable concluding that different storm-scale and mesoscale pro- boundary layers; in contrast, the 6 May 2010 VORTEX2 cesses were most important. It has long been presumed that case exhibited a well-mixed boundary layer, which nev- downdrafts of elevated storms cannot penetrate through a ertheless was capped by a temperature inversion and had stable layer to the surface (e.g., Kuchera and Parker 2006). no surface-based or mixed layer CAPE. It is uncertain Parcels are thought to be slowed in their descent to the how such differences in low-level stability are likely to surface because of the reduction of negative buoyancy, affect storm evolution and severity. presumably reducing the severe threat. In contrast, Knupp Given how few other studies exist, many questions re- (1996) saw ‘‘up–down’’ trajectories associated with the main unanswered. How are parcels with appreciable in- downdrafts of a microburst-producing storm in Colorado. stability lifted to their levels of free convection? How is In such cases, parcels in the stable layer initially rise upward this lifting maintained throughout the storm’s existence? as a result of an upward-directed vertical perturbation How does the stable layer influence the structure and pressure gradient. This ascent, in turn, causes these parcels maintenance of an elevated storm? And, when observed, to become strongly negatively buoyant and descend rapidly what is the mechanism behind severe surface wind pro- back to the surface. In this case, inflow air was slightly duction in the presence of a stable layer? Observations and stabilized under the presence of an echo overhang (anvil idealized numerical modeling will be exploited to explain shading), causing somewhat negatively buoyant air to be the processes supporting the structure, maintenance, and lifted over the cold pool, become more negatively buoyant, severe surface wind production of the 6 May 2010 elevated and descend rapidly to the surface. Schmidt and Cotton supercell from VORTEX2. Section 2 presents the obser- (1989) and Bernardet and Cotton (1998) also reported up– vations from the 6 May 2010 elevated supercell, including down trajectories associated with severe downdrafts in an both the evolution of the preconvective environment and elevated squall line and an elevated derecho, respectively. the storm itself. The observations lead to several hypoth- Both studies found that the resulting surface winds ex- eses that require further testing, which is undertaken by ceeded severe criteria as a result of the acceleration of the way of idealized numerical modeling in section 3.The flow by the horizontal pressure gradient in the stable layer. dynamics behind the structure, maintenance, and severe It is generally thought that elevated storms are less surface wind production of a simulated elevated supercell likely to produce tornadoes, as reviewed in the work of Kis within the observed environment are investigated and and Straka (2010). Billings and Parker (2012) studied compared to several sensitivity tests. Section 4 summarizes several nocturnal tornado-producing storms they origi- the important conclusions from both components of this nally presumed to be elevated, but they ultimately con- study and presents avenues for future work. cluded they were likely ingesting air from the near-surface layer and were not elevated by definition. Such subtle 2. Observations of 6 May 2010 differences are potentially quite important operationally, a. VORTEX2 observational platforms and methods since base scans of elevated supercells may have compa- rable values of vertical vorticity to those in surface-based During VORTEX2, an armada of mobile observa- supercells, and are arguably operationally indistinguish- tional platforms was deployed in an effort to learn able from surface-based storms on radar. more about the dynamics of tornadoes and their parent

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FIG. 1. Data from the Goodland, KS (KGLD), WSR-88D at 0057 UTC 7 May 2010. Radar locations for DOW6, DOW7, and SMART-R2 are marked with red circles. Sounding locations are marked with black crosses. StickNets were located every 8 km along the dashed purple line. supercells (Wurman et al. 2012). Many of these plat- issues]. The staff at the National Severe Storms Labo- forms were deployed during the 6 May 2010 case. This ratory (NSSL) edited the vast majority of the SMART- case was the first storm to be sampled during the 2010 R2 data, including removing the second-trip and season of VORTEX2 and was therefore a ‘‘shake- spurious echoes and deliasing the velocity data; addi- down’’ mission for many of the instruments and mobile tional editing, including all of that for the DOW data, Doppler radars. Even so, the novelty of the target was manually performed by the authors. Because of a storm (easternmost storm in Fig. 1) makes it a dataset calibration error, DOW reflectivities were roughly worth studying. The environment near the supercell 30 dBZ lower than what was observed by surrounding was sampled with RS92 Vaisala radiosondes launched radars; the DOW reflectivities in this paper have been using the mobile GPS advanced upper-air sounding corrected by the authors. In addition, an issue with the system (MGAUS). Temperature, pressure, and hu- DOW antenna controllers resulted in the loss of some midity were measured directly and winds were calcu- data and inconsistent height readings during sampling. lated based on the GPS location of the balloon. Surface The SMART-R2 also suffered from a mechanical issue kinematic and thermodynamic observations in the causing the reflector to detach from the pedestal forward-flank, hook-echo, and inflow areas of the storm mount, which culminated in the reported elevation were made using mobile mesonets (Straka et al. 1996) angle being lower than reality. A dual-Doppler syn- and StickNets (Weiss and Schroeder 2008;locations thesis was attempted between DOW6 and SMART- shown in Fig. 1). Quality control and bias correction R2; however, these apparent errors in positioning were performed on these surface data as well as a time- prevented an accurate one from being produced. We to-space conversion to help eliminate biases associated present reflectivity and single-Doppler data from these with mobile versus stationary observations (Skinner radars subject to the uncertainties associated with the et al. 2010). aforementioned technical issues. The evolution of the supercell was sampled by a b. Observations Shared Mobile and Atmospheric Teaching Radar (SMART-R2; Biggerstaff et al. 2005) and two Doppler 1) EVOLUTION OF THE PRECONVECTIVE on Wheels radars (DOW6 and -7; Wurman et al. 1997); ENVIRONMENT the radar locations can be seen in Fig. 1. Because of their close proximity to the supercell, the mobile radars A short-wave trough was located over the western had better data quality than the conventional Doppler United States with a jet streak present over central radars nearby [the storm quickly moved out of the Colorado at 0000 UTC on 6 May 2010 (Figs. 2a,b). At range of the Goodland, Kansas, Weather Surveillance this time, a surface low was positioned over southwest- Radar–1988 Doppler (WSR-88D), while other mobile ern Kansas with a diffuse warm front stretching across radars were degraded as a result of various mechanical central and southeastern Kansas (Fig. 2c). A sounding

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2 inversion and much weaker low-level shear (45 m2 s 2 0–3-km SRH; Fig. 3b). Also of note is the somewhat limited surface moisture over much of the region, with surface dewpoints generally around 128–158C(548– 598F) in the warm sector over much of southern Kansas versus 58–98C(418–488F) north of the front in northern Kansas and southern Nebraska (Fig. 2c). This unseasonably dry surface air accounts for the absence of surface-based CAPE. Even though the environment north of the warm front was fairly dry and capped by an inversion, large-scale advection and ascent led to the development of most unstable parcel CAPE (MUCAPE) above the inversion by late af- ternoon (e.g., Fig. 4), such that storm development could occur in the presence of lift. The analyzed 0–6-km 2 shear vector magnitude exceeded 30 m s 1,further suggesting that convection could become supercellular in this environment (Fig. 3a). Convection initiation occurred north of the warm front in south- western Nebraska and northwestern Kansas around 2230 UTC. The VORTEX2 team targeted a develop- ing supercell north of McDonald, Kansas (39.88N, 101.48W; location shown in Fig. 1), at 2343 UTC and maintained operations on this storm until roughly 0140 UTC 7 May.

2) SUPERCELL OBSERVATIONS Soundings taken in the direct inﬂow (Fig. 4; see Fig. 1 for sounding locations) of the 6 May 2010 supercell re- vealed an even more favorable environment than earlier 2 in the day, with MUCAPE of 1000–1200 J kg 1. North- easterly surface winds veering to westerly aloft led to large, looping hodographs, resulting in 0–3-km SRH 2 over 400 m2 s 2. As in the 1745 UTC sounding, each near-storm proﬁle contained a robust frontal inversion near 750 hPa, and lacked surface-based CAPE, suggesting that all resulting convection was elevated. The supercell had the characteristics of the conceptual

21 model of Lemon and Doswell (1979), with a quasi- FIG. 2. NAM analyses of isotachs (kt, where 1 kt 5 0.5144 m s ; shaded), wind barbs, and heights (dam) at (a) 250 and (b) 500 hPa, steady structure that included a hook echo to the and (c) surface observations at 0000 UTC 7 May 2010. Approxi- southwest of the forward ﬂank (Fig. 5). Midlevel, cy- mate locations of the low pressure system (L), dryline (brown line), clonic rotation (inbound to the west of outbound) was and cold (blue line) and warm (red line) fronts are also shown. evident in a column to the east of the aforementioned hook echo (Fig. 6, location annotated by red arrows). north of the front (sounding locations can be seen in This suggests the presence of a midlevel mesocyclone, Fig. 1) at 1745 UTC showed a considerable inversion supporting the characterization of this storm as a su- around 750 hPa with a well-mixed boundary layer percell. Anticyclonic rotation (inbound to the east of beneath (Fig. 3a), an environment without surface- outbound) is also present to the west of the mesocyclone based CAPE. Strong low-level vertical wind shear was in the rear ﬂank (Fig. 6, locations annotated by blue present with a clockwise-looping hodograph and arrows); this will be discussed later. 2 393 m2 s 2 of 0–3-km storm-relative helicity (SRH). In the majority of observed supercells (e.g., South of the warm front, a deep, well-mixed boundary Markowski 2002), the hook echo bends counterclock- layer was in place with only a minimal capping wise because of the cyclonic rotation in the main

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FIG. 3. Skew T–logp plot of upper-air soundings and hodographs (a) north of the warm front at 1745 UTC 6 May 2010 and (b) south of the warm front at 2053 UTC 6 May 2010. Surface-based (most unstable) parcel paths are shown in red (blue). Wind barbs are in kt. updraft; however, the hook echo of the 6 May 2010 vorticity couplet that straddles the hook echo (Brandes supercell curved anticyclonically throughout the ob- 1981; Ray 1976, Ray et al. 1981; Heymsﬁeld 1978; serving period (Fig. 5). Anticyclonic-curling hook Klemp et al. 1981; Markowski 2002; Grant and van den echoes have been observed (Van Tassell 1955; Brandes Heever 2014), which is evident behind the hook echo 1981; Fujita 1981; Fujita and Wakimoto 1982)andmay at a few times (e.g., Fig. 6). Given the severe limitations be associated with the anticyclonic member of a of the observations, the evolution of vertical vorticity

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FIG. 4. (a) Skew T–logp plot of the model sounding (blue) compared to the thermal profiles of the 0039 (green) and 0117 UTC (red) 7 May 2010 soundings, and hodograph of the 0106 UTC (magenta) 7 May 2010 sounding. (b) CAPE, CIN, and delta-z vertical profiles for the model, and the 0039 and 0117 UTC 7 May 2010 soundings. Colors are the same as in (a). Delta z is the distance a parcel is required to be lifted to reach its LFC. near the hook echo will be discussed further using around 297 K, with a 2–3-K deficit underneath the cell simulations in section 3. (Fig. 7a). In contrast, equivalent potential temperature

Mobile mesonets sampling the inﬂow environment (ue) values beneath the cell closely matched the ambi- recorded ambient potential temperature (u) values ent inﬂow values (near 317 K; Fig. 7b). The 0117 UTC

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FIG. 5. Reﬂectivity (dBZ) from SMART-R2 at (a) 4.98 at 0039:51 UTC, (b) 4.98 at 0051:51 UTC, (c) 4.28 at 0100: 45 UTC, (d) 3.58 at 0109:39 UTC, (e) 3.58 at 0118:33 UTC, and (f) 2.78 at 0127:49 UTC on 7 May 2010. Snapshots were taken when the hook echo was at a height of ;2.5 km. Labeled reﬂectivity values on the color scale are 224, 211, 2, 15, 28, 41, and 54 dBZ.

7 May sounding shows that air with ue near 317 K exists experienced by VORTEX2 team members returning to only in the lowest kilometer above the ground (Fig. 7, their lodging in Hays, Kansas, with measured peak gusts 2 right). This suggests that the surface cold pool is not of 21 and 24 m s 1 occurring at Hill City (0408 UTC driven by descending midlevel air. Instead, rather 7 May; 39.48N, 99.88W) and Russell (0454 UTC 7 May; uniform ue air at the ground suggests that the surface 38.98N, 98.98W), Kansas, respectively. Unfortunately, cold pool was produced by in situ evaporation of pre- these observations occurred after the period of VOR- cipitation in the near-surface layer. This seems to be TEX2 sampling. We assess possible mechanisms for consistent with the large ambient dewpoint depressions severe wind production in this environment using sim- in the lowest 100 hPa AGL (Fig. 4a). Since equivalent ulations in section 3. potential temperature is not conserved in the presence 3) OBSERVATIONAL SUMMARY of melting, we still consider this result to be speculative and evaluate it further using parcel trajectories in During the evening hours of 6 May 2010 in north- section 3. western Kansas, a supercell formed north of a warm Despite the very weak cold pool, the storm eventually front in a strongly sheared environment without surface- 2 grew upscale and became part of a series of line seg- based CAPE. Observed MUCAPE $ 500 J kg 1 aloft ments spanning northern Kansas (not shown). Near- apparently supported the elevated supercell updraft severe2 surface winds in advance of these cells were throughout its life. Mobile mesonet observations of

surface ue suggest that midlevel downdrafts failed to penetrate the stable layer and reach the surface, sup- 2 The National Weather Service (NWS) criterion for severe porting the notion that the storm was largely decoupled 2 thunderstorm winds is $25.7 m s 1. from the air mass below the frontal inversion. Many

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21 FIG. 6. Reﬂectivity (dBZ) and radial velocity (m s ) from SMART-R2 at (a),(b) 11.48 at 0046:45 UTC; (c),(d) 10.08 at 0046:32 UTC; (e),(f) 7.88 at 0046:18 UTC; and (g),(h) 5.68 at 0045: 53 UTC 7 May 2010. Approximate locations of the midlevel mesocyclone (red) and mesoanticyclone (blue) are marked with arrows. Height of mesocyclone is labeled in each velocity image. Labeled reﬂectivity values on the lhs color scale are 224, 211, 2, 15, 28, 41, and 54 dBZ. 2 Labeled velocity values on the rhs color scale are 218, 212, 26, 0, 6, 12, and 18 m s 1.

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FIG. 7. Averaged objective analyses of reﬂectivity (dBZ; shaded) from DOW6 and DOW7, and time-to-space converted mobile mesonet and StickNet observations of storm-relative surface winds and (a) u and (b) ue in K at 0113 UTC 7 May 2010. Tracks are 3 min in length and observations taken between 0108 and 0118 UTC 7 May 2010.

Positions are relative to the location of DOW6 in km. A vertical profile of ue (K), with a green dotted line at 317 K, from the 0117 UTC 7 May 2010 sounding (x 5 60, y 5211 km), is included along the rhs for reference. Objective analyses were interpolated onto a 60 km 3 75 km 3 10 km Cartesian grid using a two-pass Barnes analysis (Barnes 1964) with horizontal grid spacing of 500 m. quintessential supercell characteristics were observed, the storm’s rear flank. Given the limitations of the including a hook echo to the southwest of the weak-echo VORTEX2 datasets (as the first case in 2010, there region and a mesocyclone in the midlevels. Addition- were a number of unexpected radar quality issues), it is ally, an unusual anticyclonically curved hook echo was difficult to delve much farther into the storm’s governing present and may be related to the mesoanticyclone in dynamics. Therefore, further study of this supercell was

Unauthenticated | Downloaded 09/28/21 06:44 AM UTC 2644 MONTHLY WEATHER REVIEW VOLUME 145 undertaken using idealized model simulations, the re- for the blended sounding was not allowed to exceed a sults of which are discussed in section 3. relative humidity of 90% at any height to prevent the presence of moist absolutely unstable layers that would spark excess precipitation. The 0106 UTC sounding 3. Numerical model simulations was employed for the model wind profile because of a. Model setup worries that the wind profiles in the 0039 and 0117 UTC soundings could be convectively contaminated farther Although unprecedented, the observations for this case aloft as those soundings drifted closer to the storm [see, are still limited in duration and scope. Therefore, an e.g., the near-supercell wind perturbations analyzed by idealized model simulation was designed to advance our Parker (2014, his Fig. 7)]. This amalgamation of near- understanding. Simulations in this study were run using inflow thermal profiles with the far-field winds allowed Cloud Model 1, version 16 (CM1; Bryan and Fritsch for a realistic inflow environment that sustained con- 2002), with the new, higher-order parcel interpolator vection in the model without convective contamination from version 17. The model domain for each simulation of the winds. was 165 km 3 165 km 3 16 km with a horizontal grid Through experimentation, a warm bubble was found spacing of 250 m. The grid moved at a constant speed to be insufficient to produce a sustained storm because throughout the simulation to keep the supercell centered. of the interactions with the environmental stable layer. A stretched vertical grid (64 levels), with spacing ranging In nature, frontal lifting likely provided ascent aloft to from 100 m in the lowest 3 km to 500 m above 9 km, was support initiation. As a proxy, convergence was ini- used to better resolve the stable layer near the surface. A tialized over the first 30min of the simulation following damping layer was put in place above the tropopause the technique of Loftus et al. (2008). However, given (11.5 km to the model top) and along each of the lateral the elevated nature of the expected convection, the boundaries (20-km width) to prevent reflection off the applied convergence was maximized near the top of the model top and sides, respectively. Clouds and pre- stable layer instead of at the surface, as Loftus et al. cipitation were represented using the NASA Goddard (2008) originally envisioned. The equations for velocity version of the Linetal.(1983)microphysics scheme (Tao potential from Loftus et al. (2008) were used but with a et al. 1989; Tao and Simpson 1993). Each simulation was 2 2 maximum divergence (set to 25.0 3 10 3 s 1), taking run out for 4 h with a large time step of 2.0 s and output place at (x , y , z ), where x and y are the horizontal every minute over the analysis times. The Coriolis force, c c max c c points at the center of the domain and z is set to surface fluxes (including drag), and radiational effects max 2000 m. With convergence maximized near the top of were neglected. the stable layer, the storm developed and survived long A horizontally homogenous environment was imple- enough for analysis of its structure and maintenance. mented at model start using a sounding representative Although this ostensibly forces the initial storm to of the observed environment (Fig. 4), created by feed primarily on elevated air, the environment is not taking a straight average, with some caveats, of two supportive of surface-based convection (zero surface- inflow soundings (0039 and 0117 UTC; green and red based CAPE; Fig. 4), so the elevated forcing for initi- thermodynamic profiles in Fig. 4, respectively) com- ation simply expedites convective development in a bined with the wind profile from another inflow realistic scenario that closely resembles the VORTEX2 sounding (0106 UTC; pink hodograph in Fig. 4).3 The observations. thermodynamic profiles were based on the 0039 and To further understand the origins and fates of parcels 0117 UTC soundings because they were taken closer to within the elevated supercell, passive tracers and tra- the storm and farther into the postfrontal air mass; the jectories were used in the model. Tracers were initiated 0106 UTC sounding appeared to be unrepresentatively at the start of the model in 500-m-deep layers from the dry aloft (not shown). Above 4 km AGL, the 0117 UTC surface up to 2.5 km. These tracers are both passively sounding was ignored as a result of its trajectory into advected and smoothed by the full scalar equations in the forward flank of the supercell, leading to a profile the model; fractional concentrations (0–1) of their that was near saturation throughout and not indicative original value reveal the displacements of air in each of of the true inflow environment. The moisture profile these layers over time. In addition, massless parcels were introduced 2 h into the simulation to give a more explicit description of specific parcel trajectories. A 3 These soundings, used individually, were not able to sustain convection in the model. As a result of the poor representation of total of 2 190 584 parcels were released near the su- the thermodynamic environment, a NAM-initialized WRF simu- percell over a 123 km 3 145 km 3 6kmgrid.Theparcel lation was also unsuccessful in simulating this storm. trajectories are integrated forward using the model

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FIG. 8. Plan views of reflectivity (shaded; dBZ), ground-relative wind vectors, and vertical velocity greater (less) 2 than 5 (25) m s 1 shown in solid (dashed) purple contours 120 min into the simulation at (from top left to bottom right) 3.0, 2.5, 2.0, 1.5, 1.0, and 0.5 km AGL. The potential temperature deficit at the surface is shown with the 0.5 km AGL reflectivity (contoured in dashed blue at 21 and 23 K). winds on the native model time step and are saved Beneath the stable inversion, there is a lack of signifi- every 30 s. cant vertical motion (Fig. 8, purple contours); however, appreciable wind perturbations are evident (Fig. 8;0.5km b. Simulation results AGL). At 1 km AGL and above, strong downdrafts are As in the observations, the simulation produced a evident in the rear flank, nearly collocated with the hook supercell storm that was quasi steady through 3 h. Because echo (dashed purple contours in Fig. 8); the dynamics and of this, for brevity, the present analysis will focus on the effects of this descent will be discussed later in this section. storm 2 h into the simulation (Fig. 8) unless otherwise In the low levels, the storm has a somewhat diffused ap- noted. At this time, the storm is representative of other pearance with reflectivity structure that is not strikingly times in the simulations and parallels the observations. supercellular. Although this may be perpetuated in the

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21 FIG. 9. Plan views of tracer concentration (shaded) and vertical velocity every 10 m s (purple) 120 min into the simulation at 5 km AGL. Tracer originates in the 0.05–0.5-, 0.5–1.0-, 1.0–1.5-, 1.5–2.0-, and 2.0–2.5-km layers with a maximum concentration of 1. The 35-dBZ contour is shown in black for reference. The dashed gray lines in the top-left frame correspond to the locations of the updraft cross sections in Fig. 17. model by diffusion, the observed storm had a similar look in the updraft (downdraft). By 3.0 km AGL, the appear- on radar at lower tilts (at altitudes below what is shown in ance of reflectivity, along with cyclonic vorticity in the Figs. 5 and 6). This may be a partial effect of the decou- updraft, conforms to the typical supercell model (Fig. 8). pling of winds from the storm beneath the inversion, such A key motivation for this study was to understand the that precipitation falling through the subinversion layer is origin of the updraft air in presumably elevated super- redistributed by a flow field that is not characteristic of the cells. At 5 km, the highest concentration of passive tracer storm aloft. At and above the stable layer, the main up- in the updraft is from above 1.5 km, with no tracer ar- draft appears to the east of a prominent anticyclonic hook riving from the lowest 1 km (Fig. 9). As noted previously, echo (Fig. 8), bearing great resemblance to the observa- no CAPE exists in the lowest 1 km, so this result makes tions (e.g., Fig. 6). The winds have a clear cyclonic (anti- physical sense. Some lower concentrations from the cyclonic) curvature to them to the east (west) of the hook 1–1.5-km layer appear west of the main updraft, sug- echo, especially at and above 2.0 km (Fig. 8), indicating gesting that low-level parcels, possibly with rather small 2 substantial cyclonic (anticyclonic) vorticity that is present instability (CAPE , 500 J kg 1), are still being lifted to

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FIG. 10. Normalized (a) origins of updraft parcels in 100-m-deep layers and (b) destinations of downdraft parcels originating above 1 km AGL 120 min into the simulation (or 30 min after model restart). Updraft parcels were 2 defined as those with w $ 10 m s 1 at 5 km and above in a box centered around the storm. Downdraft parcels were 2 defined as those with w # 23ms 1 in the lowest 2 km AGL at any point during their life and were then sorted to only include those that originated above 1 km AGL. an appreciable altitude (10.6% of analyzed updraft par- that the plotted potential temperature perturbations cels came from the 1–1.5-km layer). In many cases, this air (shadedfieldinFig. 11a) are time averaged and thus not (with substantial CIN) is strongly negatively buoyant necessarily representative of an individual parcel’s in- during ascent and is accelerated back to its original height stantaneous buoyancy at each point. Below ;1.2 km, par- or lower [i.e., up–down trajectories, as described in sec- cels follow an up–down flow branch, becoming strongly tion 1;cf.Knupp (1987)]. This regime of trajectories will negatively buoyant after being lifted to an altitude as high be discussed further momentarily. To further examine as 4 km and then moving out of the zone of lifting the origins of updraft air, parcels having acquired w $ (Fig. 11b). The zone where parcels have ascended into the 2 10 m s 1 at 5 km in a 30 km 3 40 km box around the up- warm inversion layer is clearly visible as a cold pocket draft within 2–2.5 h into the model simulation were centered at x 5 10 km, z 5 2kminFig. 11. Above 1.2 km, identified for analysis (6580 parcels in total). Placing the each averaged trajectory rises into the main updraft. strict thresholds on w and the parcel’s horizontal location To further isolate the behavior of the up–down tra- prevented the inclusion of parcels that were not repre- jectories that travel through the anticyclonic hook- sentative of the updraft. All of the parcels originate above echo region, we identified and averaged the 11 such 2 1 km in height, with the majority coming from above parcels that arrived there with z ,20.02 s 1 during a 1.5 km (Fig. 10a). This matches both the environmental 5-minperiodcenteredont 5 2 h. As a result of the CAPE profile (Fig. 4) and the lack of significant vertical impressive directional and speed shear in the wind motion below the inversion height. profile (Fig. 4), the ambient environmental parcels To look more closely at typical low-level parcel behavior, originate with appreciable streamwise horizontal vor- the aforementioned 6580 updraft parcels were averaged ticity (Fig. 12a). Streamwise horizontal vorticity will be within each 250-m-deep layer (0–250 m AGL, 250–500 m reoriented into the vertical as a result of gradients in AGL, etc.) based on their heights of origin (Fig. 11a); note vertical velocity along the parcel trajectory, as shown in

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FIG. 11. (a) Average updraft (black) and up–down (white) trajectories near the main updraft. Potential temperature deﬁcit (shaded; K) was averaged from y 5220 to 210 km at t 5 120 min. Boxes on trajectories represent the parcel’s location 120 min into the simulation. Parcel motion is from right to left. (b) Height (black; m) and 2 buoyancy 1 buoyant perturbation vertical pressure gradient acceleration (ACCB; blue, m s 2) over time (min) for the average up–down parcel in the 1.00–1.25 km AGL layer [topmost white trajectory in (a)].

Fig. 12b. In general, supercell mesocyclones are fueled these up–down parcels, limiting the amount of positive by tilting and subsequent stretching of this horizontal z produced (Fig. 13). Therefore, updraft parcels (as vorticity, yielding large values of positive z during opposed to the up–down trajectories) dominate the parcels’ period of rapid ascent (Davies-Jones 1984). center of positive z in the parent storm (i.e., the Similarly, tilting of ambient horizontal vorticity occurs mesocyclone). along the up–down trajectories, creating a small Negative buoyancy drives the up–down parcels quickly amount of positive z initially (Fig. 13). Because of their toward the surface, whereupon tilting causes the vorticity negative buoyancy, however, stretching is weak for vectors to be redirected downward (Figs. 12b and 13).

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21 FIG. 12. (a) Vertical proﬁle of the ambient streamwise vorticity (s ) from the model simulation with the origin layer for averaged parcels highlighted in the dashed green box, and (b) parcel trajectory (black) and vorticity vectors (teal) viewed from the south-southeast for an averaged up–down trajectory that descends behind the hook echo (pink trajectory in Fig. 14).

Substantial ‘‘negative’’ stretching of the air over time vertical vorticity by tilting and stretching, baroclinic results from this downward acceleration, producing large generation of horizontal vorticity is also present (e.g., values of anticyclonic z in the descending branch of these Rotunno and Klemp 1985); the parcel budgets for hori- trajectories (Fig. 13). In addition to the production of zontal vorticity do not reconcile very well, but the sign of

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(Rotunno and Klemp 1982) were calculated using the solver described by Coffer and Parker (2015, 2017)and Davenport and Parker (2015). A west–east cross section through the main updraft shows a strong dynamic low to the east of the updraft (Fig. 15a; x 5 18 km, z 5 4.5 km). This is fueling upward dynamic accelerations from near the inversion to roughly 6 km (Fig. 15,red hatching). Combined with negligible upward buoyant accelerations below 3 km, it is implied that the supercell’s own dynamic lifting plays a major role in both ﬂow ﬁelds, lifting air below the inversion into the ascending branch of the up–down trajectories, and lifting parcels above the inversion to their LFCs. In addition to the dynamic lifting, an area of negative buoyancy is present at the inversion top (Fig. 15a; x 5 12 km, z 5 2 km), which may provide additional isen-

21 tropic upglide where the potential temperature surfaces FIG. 13. Time series of (top) height (m) and vertical vorticity (s ), 2 are domed upward, much as with the elevated cool pool (middle) vertical vorticity tendency (s 2) for tilting (blue) and 2 stretching (green), and (bottom) vertical vorticity tendency (s 2)for noted by Bryan and Weisman (2006). This negative parcel tendency (orange) and tilting 1 stretching (aqua) for an av- buoyancy anomaly is associated with the zone of parcels eraged up–down trajectory that descended behind the hook echo from the stable layer that have ascended in the up–down (pink trajectory in Fig. 14)fromt 5 100 to 130 min. flow branch. At the same latitude and longitude of this buoyancy minimum, ridges and troughs are present in the baroclinicity on average produces antistreamwise the potential temperature surfaces with vertical velocity horizontal vorticity (i.e., it decreases the initially large perturbations located roughly one-quarter wavelength values of streamwise vorticity during the trajectory his- out of phase with these ripples (Fig. 16), which is in- tory; not shown). The history of vorticity vectors along dicative of gravity waves. These features are persistent the averaged trajectory (Fig. 12b) is consistent with the and continually adjacent to the simulated storm. These budget, showing that the primary process contributing to waves represent perturbations produced by the storm strong anticyclonic vorticity in the hook-echo region is that can potentially contribute to the maintenance of the the reorientation and subsequent stretching of initial updraft by causing the upglide of parcels to their LFC. ambient streamwise vorticity along the up–down trajec- In the observations (section 2, Fig. 7)wenoteda tories that flow into that part of the storm. weak surface cold pool exhibiting a lack of a significant

Interestingly, this persistent negative z along the up– ue deﬁcit. The model data reveal a similar phenome- down trajectories, acquired upon descent into the rear- non. Over the course of 120–180 min in the simulation,

flank, results in pools of negative vorticity to the west the ue deficit in the surface cold pool never exceeds of the hook echo (Fig. 14a). Trajectory analysis con- 2 K, while u deficits are typically 2–3 K (Fig. 17). Much firmed that this was the primary method of negative as in the observations, the base-state vertical profile of vorticity generation as parcels originating in the mid- ue in Fig. 17 suggests that this cold pool air could have levels were not present near the vorticity minimum only descended from below roughly 1 km AGL. Ex- (not shown). At 3 km AGL, this anticyclonic vorticity ploration of the downdrafts by way of parcels reveals 2 2 (approaching 20.05 s 1) is much greater in magnitude that downdraft trajectories (w # 23ms 1) originating than the positive vorticity present in the updraft sector above 2 km AGL indeed do not reach the surface 2 (0.01–0.015 s 1 centered at x 5 15 km; Fig. 14). De- (Fig. 10b), not even in the area of u surplus in the scending precipitation then wraps around the stronger, forward flank (x 5 15–20 km, y 529to214 km in anticyclonic zone of vorticity, resulting in the peculiar Fig. 17), where a weak heat burst is present. anticyclonic shape of the simulated hook echo that was We noted the occurrence of near-severe surface also seen in the observed storm (Fig. 14a). winds in advance of the observed cells late in the life Although we have demonstrated the origins and be- cycleofthe6May2010case(section 2). Interestingly, havior of updraft and up–down air in the low levels, the the simulated supercell produces severe surface winds persistent lifting mechanism leading to the steadiness out ahead of the surface cold pool (Fig. 18). Two par- of these flow branches is still unclear. The buoyancy cels linked to this phenomenon were singled out for field and dynamic and buoyant perturbation pressures further investigation. Each parcel follows an up–down

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21 21 FIG. 14. (a) Plan view of vertical vorticity (shaded; s ), vertical velocity #23ms (hatched), and the 35-dBZ contour of reﬂectivity (thick black line) at z 5 3 km and (b) the x–z 2 2 cross section of vertical vorticity (shaded; s 1) and vertical velocity #23ms 1 (hatched) at y 5 211 km. An averaged up–down parcel trajectory near the hook echo is shown in pink. Parcel motion is from right to left.

2 ﬂow branch, originating in the lowest 100 m AGL and severe limits, reaching values as high as 40 m s 1 ascending to greater than 400 m AGL. These parcels (Fig. 18). For both the gray and black parcels in Fig. 18, both descend in the hook echo, suggesting that evap- the air is accelerated while traveling close to the ground oration may have enhanced their negative buoyancy. because of the dynamic high that formed where the Upon descending to the ground,4 asharpincreasein downdraft meets the surface (y 529km), implying ground-relative wind speed is experienced to above continued horizontal pressure gradient forcing. The pressure gradient between a dynamic surface high in the downward branch of the up–down trajectory and a 4 These parcels actually descend below the lowest model level dynamic surface low, due to a rotorlike circulation, is (z 5 50 m). the cause of this acceleration to the east (Fig. 18b). The

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FIG. 15. Vertical cross sections through the (a) short and (b) long axes of the supercell updraft of buoyancy with storm-relative wind vectors, buoyant pressure perturbations with buoyant acceleration vectors, and dynamic 2 pressure perturbations with dynamic acceleration vectors 120 min into the simulation. Here, w $ 5ms 1 is shaded 2 2 in gray and vertical accelerations $0.05 m s 2 are hatched in red. Buoyancy is contoured every 0.10 m s 2 and the pressure perturbations are contoured every 50 Pa with negative values dashed.

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21 FIG. 16. The x–z cross sections of w (shaded; m s ), potential temperature (K), and ground-relative wind vectors 120, 140, 160, and 180 min into the simulation. Potential temperature contours are (from bottom to top) 296, 298, 300, 302, 304, 306, 308, and 310 K.

footprint and magnitude of the surface winds seem to warm, inversion layer assisted in the development of an correspond well with the observed case. elevated cold pool atop the inversion. In the descending branch of the up–down parcels, pools of negative vertical c. Simulation summary vorticity due to vortex tilting and stretching contorted the A numerical model simulation was run using a blend of tip of the hook echo into an anticyclonic bend. Up–down soundingsfromthe6May2010VORTEX2casetofur- parcels from the lowest 100 m AGL were also the driver ther investigate the mechanisms behind the structure, behind simulated severe surface winds, a result of strong maintenance, and severe surface wind production of the low-level downdrafts reaching the surface and an accel- storm. Updraft air was found to originate from above eration of the winds near the ground by a horizontal 1 km AGL and was mostly driven upward by dynamic pressure gradient. lifting from the supercell itself. Parcels originating below ;1.2 km AGL were initially lifted dynamically before 4. Conclusions and future work becoming negatively buoyant, causing the air to descend; these are known as up–down parcels. This ﬂow branch The 6 May 2010 storm observed during VORTEX2 is was the main low-level downdraft mechanism. Down- perhaps the ﬁrst coordinated dataset ever collected for an drafts originating in the midlevels were not able to pen- elevated supercell. The supercell formed to the north of a etrate through the stable layer. It therefore appears that warm front in an environment characterized by strong the surface cold pool was fueled by the evaporation of vertical wind shear and zero surface-based CAPE. Despite precipitation beneath the inversion, which is supported the lack of surface instability, MUCAPE approaching 21 by both the ambient surface ue values and the lack of 1000 J kg sustained the supercell throughout its life cy- midlevel downdraft parcels below 1 km in the simulation. cle. The storm was steady over the course of the radar Meanwhile, up–down trajectories that impinged on the sampling, moving to the east-northeast without much

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FIG. 17. (a),(c) Potential temperature and (b),(d) equivalent potential temperature deficit (shaded; K) (top) 120 and (bottom) 180 min into the simulation. The 35-dBZ contour is shown in black for reference. The vertical profile of the base-state equivalent potential temperature (K), with a green dotted line at 319 K, is shown on the right. change in structure. An anticyclonic hook echo was (to ,4 km AGL on average) until it was strongly neg- prominent with this supercell and was collocated with atively buoyant and then quickly descending to the anticyclonic rotation in the midlevels (Fig. 6). Cyclonic surface. This flow branch was the main low-level midlevel rotation existed in the column above the weak- downdraft mechanism in the storm, since parcels echo region in the single-Doppler analysis, implying the originating in the midlevels were unable to penetrate existence of a rotating updraft and justifying the classifi- the stable layer. The surface cold pool was conse- cation of this storm as a supercell. Only a weak surface quently fueled by the evaporation of precipitation be- cold pool was evident, and it was apparently devoid of air low the stable layer. Up–down parcels, halted in their originating from midlevels. ascent by the warm inversion, created an elevated cold To explore some of the relevant dynamical processes pool on top of the inversion layer. Air descending in that were difficult to observe or not directly observed, these up–down trajectories created pools of negative we simulated the storm using an idealized numerical vertical vorticity that bent the hook echo in a coun- model with a blend of near-storm soundings taken on terclockwise manner. In the lowest 100 m, the up–down 6 May 2010. The simulated supercell was similar in parcels were also the driver behind the simulated sur- structure to the observed storm, having a less well- face winds that reached severe wind speed thresholds, defined structure near the surface that became more descending to the surface rapidly and being accelerated supercellular-looking farther aloft (Fig. 8). The simu- by a dynamic perturbation pressure gradient near the lated storm also included an anticyclonic hook echo ground. This is similar to the mechanism discussed by that was the result of a pool of negative vorticity in the Schmidt and Cotton (1989) and Bernardet and Cotton rear-flank region. Only air parcels originating above (1998), both of whom found that strong surface winds the near-surface layer fueled the supercell updraft, with in a mesoscale convective system stemmed from up– most of the inflow air coming from above ;1.2 km down parcels and were enhanced by a horizontal AGL. This air was lifted dynamically by the supercell pressure gradient associated with a mesohigh–mesolow itself; buoyant accelerations were negligible in the low- couplet near the ground. levels. Below ;1.2 km AGL, the majority of air fol- Interestingly, our results differ somewhat from those of lowed an up–down trajectory, being lifted dynamically Nowotarski et al. (2011). The majority of the elevated

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21 FIG. 18. (a) Ground-relative wind magnitude (shaded; m s ) and cold pool (dashed; contoured at 21 and 23 K) at the surface with parcels associated with this wind maximum in black and gray. (b) Vertical cross section of dynamic pressure perturbations (shaded; Pa), storm- relative wind vectors, and parcel trajectories as in (a). Pressure perturbations were averaged over y 5212 to 22. Both plots are from 123 min into the simulation. supercells in their study, simulated over statically stable moist Weisman and Klemp (1982) sounding in their boundary layers with various levels of stability, were still study]. The present sounding (e.g., Fig. 4a) has zero fed by near-surface parcels and had midlevel downdrafts surface-based CAPE, such that near-ground air simply that were able to reach the surface. Only in cases of ex- cannot acquire positive buoyancy and fuel updrafts. treme stability were their simulated supercells truly de- As with any case study, it is unknown how repre- coupled from the surface. Although the reasons for the sentative this storm is of all elevated supercells (nor differences between their ﬁndings and ours are not totally how well the simulation replicates the processes of the clear, we speculate that much of the difference hinges on real-world storm). More cases would obviously be our use of an observed sounding [as opposed to the rather ideal. It is also unknown whether frontally forced

Unauthenticated | Downloaded 09/28/21 06:44 AM UTC 2656 MONTHLY WEATHER REVIEW VOLUME 145 elevated storms are fundamentally different from Bryan, G. H., and J. M. Fritsch, 2002: A benchmark simulation nocturnal elevated storms, nor how such storms might for moist nonhydrostatic numerical models. Mon. Wea. Rev., , respond to spatial or temporal changes in inversion 130, 2917–2928, doi:10.1175/1520-0493(2002)130 2917: ABSFMN.2.0.CO;2. strength. Our simulation also suggests some important ——, and M. L. Weisman, 2006: Mechanisms for the production of roles for up–down trajectories and possible gravity severe surface winds in a simulation of an elevated convective waves, but such phenomena are likely quite difficult to system. 23rd Conf. on Severe Local Storms, St. Louis, MO, observe. In terms of societal impact, the seemingly Amer. Meteor. Soc., 7.5. [Available online at https://ams. unusual mechanics of severe surface winds in our sim- confex.com/ams/23SLS/techprogram/paper_115224.htm.] Carbone, R. E., J. W. Conway, N. A. Crook, and M. W. Moncrieff, ulated elevated supercell deserve further exploration. 1990: The generation and propagation of a nocturnal squall Are there specific environmental factors, outside of the line. Part I: Observations and implications for mesoscale stable inversion, that could make an elevated supercell predictability. Mon. Wea. Rev., 118, 26–49, doi:10.1175/ environment more conducive to severe wind pro- 1520-0493(1990)118,0026:TGAPOA.2.0.CO;2. duction? And, are storms such as this truly incapable of Coffer, B. E., and M. D. Parker, 2015: Impacts of increasing low-level shear on supercells during the early evening transition. Mon. producing tornadoes? Although not discussed in detail Wea. Rev., 143, 1945–1969, doi:10.1175/MWR-D-14-00328.1. here, our simulated elevated supercell occasionally ——, and ——, 2017: Simulated supercells in nontornadic and produced short-lived near-surface vortices. We pre- tornadic VORTEX2 environments. Mon. Wea. Rev., 145, 149– sume these to be spurious, but a more advanced un- 180, doi:10.1175/MWR-D-16-0226.1. derstanding of tornadoes in supercells that appear to be Colman, B. R., 1990a: Thunderstorms above frontal surfaces in environments without positive CAPE. Part I: A climatology. Mon. elevated would be worth pursuing. The long-range goal Wea. Rev., 118, 1103–1122, doi:10.1175/1520-0493(1990)118,1103: of this line of research is to advance the knowledge of TAFSIE.2.0.CO;2. how storms respond to their environments across the ——, 1990b: Thunderstorms above frontal surfaces in environ- full range of observed lower-tropospheric stability. ments without Positive CAPE. Part II: Organization and instability mechanisms. Mon. Wea. Rev., 118, 1123–1144, , . Acknowledgments. This research was supported by doi:10.1175/1520-0493(1990)118 1123:TAFSIE 2.0.CO;2. Corfidi, S. F., S. J. Corfidi, and D. M. Schultz, 2008: Elevated convection NSF Grant AGS-1156123. We thank Sandra Yuter and and castellanus: Ambiguities, significance, and questions. Wea. Gary Lackmann, as well as the members of the Convective Forecasting, 23, 1280–1303, doi:10.1175/2008WAF2222118.1. Storms Group (CSG), at NCSU for their suggestions and Davenport, C. E., and M. D. Parker, 2015: Impact of environmental feedback. We gratefully acknowledge Casey Davenport heterogeneity on the dynamics of a dissipating supercell for her preliminary exploration and guidance on the thunderstorm. Mon. Wea. Rev., 143, 4244–4277, doi:10.1175/ MWR-D-15-0072.1. VORTEX2 observations for this case. We thank George Davies-Jones, R. P., 1984: Streamwise vorticity: The origin of up- Bryan for his tireless support of the CM1 modeling system draft rotation in supercell storms. 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