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    Map of the locations of the instrumented tower, sounding (Edmond), radar sites (Norman and Cimarron), and town of Arcadia, Oklahoma. The Norman radar is located at the origin. The tornado occurred south of Arcadia between approximately 1700 and 1710 CST 17 May 1981. Contours of the reflectivity factor (dBZ) measured by the Cimarron radar at 1.0 km AGL at 1704 CST are also shown. The crescent-shaped region indicates where the dual-Doppler between-beam angle is greater than 45° (and less than 135°).

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    Sounding taken from Edmond, Oklahoma, at 1430 CST 17 May 1981. (a) Skew T–logp diagram. The solid and dashed lines indicate temperature (°C) and dewpoint (°C), respectively. Full (half) wind barbs represent 5 m s−1 (2.5 m s−1); flags represent 25 m s−1. Winds are shown at heights indicated in km AGL. Tower observations at 444 m AGL at 1638 CST (within updraft) are plotted as closed dots. (b) Hodograph. Each circle represents 5 m s−1 of wind speed. The heights are in kilometers AGL. The small triangle indicates observed storm motion (u = 10 m s−1, υ = 6 m s−1).

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    Cimarron radar reflectivity factor (dBZ) and horizontal storm-relative winds at 5.0 km AGL. The x (east–west) and y (north–south) distances (km) are relative to the Norman radar. Local maxima in cyclonic and anticyclonic vertical vorticity are marked “C” and “A,” respectively. Updrafts are hatched. (a) 1634 CST, 0.5 h before tornado, w > 15 m s−1 shaded; (b) 1647 CST, 0.25 h before tornado, w > 20 m s−1 shaded; and (c) 1722 CST, just after tornado, w > 11 m s−1 shaded.

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    The maximum vertical velocity vs height within the updraft of the Arcadia storm for various times.

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    Cimarron radar reflectivity factor (dBZ) and horizontal storm-relative winds at 0.5 km AGL; x and y as in Fig. 3. Gust front positions are marked by a dashed line: (a) 1630 CST; (b) 1643 CST; (c) 1651 CST; (d) 1704 CST, TVS location denoted by a dot and damage path indicated by a heavy line; (e) 1710 CST, TVS location denoted by a dot and damage path indicated by a heavy line; and (f) 1717 CST.

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    (Continued)

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    Reflectivity factor scan by the Cimarron radar at 0.7° at 1658 CST. Range is given in kilometers.

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    Time series of the height of the Arcadia storm top, as indicated by the 40-dBZ contour in raw Cimarron radar scans.

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    The maximum vertical vorticity (10−2 s−1) vs height within the mesocyclone of the Arcadia storm for various times.

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    Horizontal storm-relative winds and vertical velocity (m s−1) at 1.0 km AGL at 1704 CST; x and y as in Fig. 3. Contours for negative values are dashed. The TVS location is marked with a dot.

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    Horizontal storm-relative winds and vertical vorticity (10−2 s−1) at 0.5 km AGL for the high-resolution (500-m grid spacing) analyses; x and y as in Fig. 3. Contours for negative values are dashed. The locations of the TVS (closed dot) and tornado damage path (thin solid line segment) are marked: (a) 1704 CST and (b) 1710 CST.

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    Low-level horizontal storm-relative winds and Cimarron radar reflectivity factor (dBZ) at 1638 CST at 0.5 km AGL; x and y as in Fig. 3. (a) Dual-Doppler analysis winds at 0.5 km AGL at 1638 CST. (b) Time-to-space conversion of winds measured by the instrumented tower at 444 m AGL over a 30-min period centered at 1638 CST. (c) Vector difference between the winds in (a) and those in (b).

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    Vertical shear of the low-level horizontal winds and Cimarron radar reflectivity factor (dBZ) at 1638 CST at 0.5 km AGL; x and y as in Fig. 3. (a) Vertical shear in the 500–1000-m AGL layer in the dual-Doppler analysis winds at 1638 CST. (b) Time-to-space conversion of vertical shear between 7 and 444 m AGL measured by the instrumented tower over a 30-min period centered at 1638 CST.

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    Vertical velocities measured by the instrumented tower at 444 m AGL (heavy line) and synthesized in the dual-Doppler analysis at 0.5 km AGL at 1638 CST (contours, m s−1); x and y as in Fig. 3. Contours for negative values are dashed. The time-to-space conversion of tower data covers the period within 15 min of the center time (1638 CST). The heavy straight line indicates the direction of storm motion. The displacement of the curve above (updraft) or below (downdraft) this line indicates the vertical velocity. The maximum and minimum vertical velocities are 6.1 m s−1 and −6.8 m s−1, respectively.

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    Vertical cross sections constructed from a time-to-space conversion of instrumented tower data between 1635 and 1645 CST. Equivalent distance for storm motion of 11.7 m s−1 is indicated. (a) Equivalent potential temperature (K) and storm-relative winds. (b) Virtual potential temperature (K) and storm-relative winds. (c) Storm-relative winds and contours of the component of wind normal to the cross section (m s−1). Solid (dashed) contours indicate winds into (out of) the page.

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    (Continued)

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    Vertical profile of equivalent potential temperature (K) in the sounding taken from Edmond, Oklahoma, at 1430 CST (cf. Fig. 2).

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    Demonstration of how horizontal vorticity associated with vertical shear can be concentrated into a vortex at low levels. The analysis is carried out in terms of a material circuit in the fluid; the process involves tilting of the circuit into the horizontal by a downdraft, followed by convergence of the circuit as it is passed into an updraft. The effects of deformation acting on the material curve are neglected for ease of interpretation. The environmental wind profile is indicated at the left.

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    Material curves traced backward (using storm-relative trajectories) from a ring with a 2.5-km diameter centered on the mesocyclone at 0.5 km AGL; x and y as in Fig. 3. Velocities are computed by linear interpolations (in time) between the dual-Doppler volume at the time given in the figure and the volume at a previous time. Heights (km AGL) of points along the curve are shown. Contours of vertical vorticity (10−2 s−1) are also shown: (a) followed back 400 s from 1638 CST, and (b) followed back 400 s from 1704 CST.

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    Hodograph for low-level winds measured by the instrumented tower at 1530 (about 50 km ahead of the updraft) and 1634 CST (2 km ahead of the updraft). Each circle represents 5 m s−1 of wind speed.

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    Schematic of features at 1.0 km AGL. Downdrafts (w ≤ −1 m s−1) are shaded lightly, updrafts (w ≥ 3 m s−1) are shaded heavily, and mesocyclones (ζ ≥ 0.6 × 10−2 s−1) are denoted by solid curves. The TVS location at 1704 CST is marked.

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    Perturbation pressure (mb) retrieved from dual-Doppler analyses at 1704 and 1710 CST; x and y as in Fig. 3. Contours for negative values are dashed: (a) vertical cross section along x = 12 km [see (b)], and (b) horizontal cross section at 2.5 km AGL.

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The Arcadia, Oklahoma, Storm of 17 May 1981: Analysis of a Supercell during Tornadogenesis

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  • 1 School of Meteorology, University of Oklahoma, Norman, Oklahoma
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Abstract

On 17 May 1981, an extensive dataset was collected for a supercell thunderstorm that produced an F2 tornado near Arcadia in central Oklahoma. Coordinated dual-Doppler scans of the storm by 10-cm research radars were collected at approximately 5-min intervals from 30 min before the tornado touched down until 15 min after the tornado had dissipated. The Arcadia storm was also well sampled by a 444-m-tall instrumented tower. The low-level inflow, updraft, mesocyclone, and rear precipitation core of the supercell all passed across the tower.

A comparison of the instrumented tower measurements with a dual-Doppler synthesis reveals that the latter qualitatively resolved the low-level flow. However, the magnitudes of the low-level horizontal winds and updraft speed were underestimated. In addition, the vertical shear of the horizontal wind in the lowest kilometer was unresolved in the Doppler winds.

In the storm environment, horizontal vorticity was strong (∼1.5 × 10−2 s−1) and approximately streamwise over the depth of the instrumented tower. Just upstream (northeast) of the updraft, the magnitude of horizontal vorticity was nearly twice this value and had likely been enhanced by baroclinic generation of horizontal vorticity and/or stretching of horizontal vorticity. Tilting of the resulting horizontal vorticity into the vertical produced the pretornadic low-level mesocyclone. Low-level mesocyclone inflow was primarily from the east, but during the tornadic stage, parcels approaching from the north and west were also drawn into the circulation.

The tornado formed southeast of the mesocyclone center and near the tip of the reflectivity hook echo while low-level mesocyclone vorticity was increasing. Tornadogenesis occurred near the nose of the rear downdraft within a region of horizontal shear between southeasterly inflow into the storm and westerly outflow from the rear downdraft. Pressure retrievals suggest the rear downdraft south of the mesocyclone center was associated with a downward-directed perturbation pressure gradient force. The tornado and the parent storm dissipated as outflow surged eastward ahead of the updraft.

This case study is the first to include a comparison of independent measurements of the wind field in and near the low-level mesocyclone of a supercell. The wind analysis is also complemented by the instrumented tower thermodynamic measurements.

Corresponding author address: David C. Dowell, School of Meteorology, University of Oklahoma, 100 East Boyd, Room 1310, Norman, OK 73019-0628.

Email: ddowell@ou.edu

Abstract

On 17 May 1981, an extensive dataset was collected for a supercell thunderstorm that produced an F2 tornado near Arcadia in central Oklahoma. Coordinated dual-Doppler scans of the storm by 10-cm research radars were collected at approximately 5-min intervals from 30 min before the tornado touched down until 15 min after the tornado had dissipated. The Arcadia storm was also well sampled by a 444-m-tall instrumented tower. The low-level inflow, updraft, mesocyclone, and rear precipitation core of the supercell all passed across the tower.

A comparison of the instrumented tower measurements with a dual-Doppler synthesis reveals that the latter qualitatively resolved the low-level flow. However, the magnitudes of the low-level horizontal winds and updraft speed were underestimated. In addition, the vertical shear of the horizontal wind in the lowest kilometer was unresolved in the Doppler winds.

In the storm environment, horizontal vorticity was strong (∼1.5 × 10−2 s−1) and approximately streamwise over the depth of the instrumented tower. Just upstream (northeast) of the updraft, the magnitude of horizontal vorticity was nearly twice this value and had likely been enhanced by baroclinic generation of horizontal vorticity and/or stretching of horizontal vorticity. Tilting of the resulting horizontal vorticity into the vertical produced the pretornadic low-level mesocyclone. Low-level mesocyclone inflow was primarily from the east, but during the tornadic stage, parcels approaching from the north and west were also drawn into the circulation.

The tornado formed southeast of the mesocyclone center and near the tip of the reflectivity hook echo while low-level mesocyclone vorticity was increasing. Tornadogenesis occurred near the nose of the rear downdraft within a region of horizontal shear between southeasterly inflow into the storm and westerly outflow from the rear downdraft. Pressure retrievals suggest the rear downdraft south of the mesocyclone center was associated with a downward-directed perturbation pressure gradient force. The tornado and the parent storm dissipated as outflow surged eastward ahead of the updraft.

This case study is the first to include a comparison of independent measurements of the wind field in and near the low-level mesocyclone of a supercell. The wind analysis is also complemented by the instrumented tower thermodynamic measurements.

Corresponding author address: David C. Dowell, School of Meteorology, University of Oklahoma, 100 East Boyd, Room 1310, Norman, OK 73019-0628.

Email: ddowell@ou.edu

1. Introduction

Recent progress in numerical simulations and advances in observational capabilities have sparked renewed interest in verification of the details of how a supercell thunderstorm (Browning 1964) produces a tornado. Computer simulations of supercells are becoming more and more successful at resolving tornadolike vortices (Wicker and Wilhelmson 1995; Grasso and Cotton 1995). With the nearly nationwide coverage provided by the WSR-88D network, tornadic storms are being sampled relatively frequently by Doppler radar (Guerrero and Read 1993; Magsig and Burgess 1996). During 1994–95, scientists collaborated on the Verification of the Origins of Rotation in Tornadoes Experiment (VORTEX) in the southern and central plains of the United States (Rasmussen et al. 1994). Radiosondes, mobile mesonet instrument packages, portable Doppler radar, airborne Doppler radar, and other implements were all directed toward the broad goal of understanding the formation of tornadoes.

Tornadogenesis relies on horizontal convergence within the boundary layer to amplify vertical vorticity to magnitudes characteristic of tornadoes (Ward 1972; Lewellen 1993). In some cases, thunderstorms concentrate vertical vorticity already present in the storm environment (Wilson 1986; Brady and Szoke 1989; Wakimoto and Wilson 1989). In contrast, supercell thunderstorms appear to act upon low-level vertical vorticity produced by the storm itself (Barnes 1970; Brandes 1984b; Rotunno and Klemp 1985).

Tilting in the storm updraft of horizontal vorticity associated with the environmental vertical wind shear is generally believed to be the source of rotation for the midlevel mesocyclone of a supercell (Barnes 1970; Brandes 1984b; Rotunno and Klemp 1985). On the other hand, the source of low-level rotation is more controversial. Numerical simulations by Rotunno and Klemp (1985) suggest that the development of low-level rotation awaits the presence of evaporatively cooled air near the surface. Once the storm’s cold pool is established, horizontal vorticity produced baroclinically along the cool air boundary upstream of the updraft is tilted into the vertical, and the vertical vorticity is amplified by stretching within the updraft.

In contrast to the work of Rotunno and Klemp (1985), the simulations by Walko (1993) demonstrate that environmental horizontal vorticity can be the source for low-level rotation. (Although environmental horizontal vorticity may itself have been generated by baroclinic effects, it is distinguished from the type of baroclinically generated horizontal vorticity described by Rotunno and Klemp by its existence prior to storm formation.) Wicker (1996) stresses the importance of phasing of horizontal vorticity of both types in the development of the low-level mesocyclone.

An issue that remains even more ambiguous is how the low-level mesocyclone and tornado are related. Tornadic vortex signatures (TVSs) in Doppler radar data (Brown et al. 1978) provide observational evidence of the vastly different magnitudes of size and vorticity in mesocyclones and tornadoes. Furthermore, the behavior of the TVS within the mesocyclone varies from storm to storm; it may appear first aloft, first at low levels, or simultaneously over a large depth (Brown et al. 1978; Trapp and Mitchell 1995). The diagnosis of how mesocyclones and tornadoes are related continues to stretch the limits of observational capabilities. Another issue that remains unclear is how much the processes that produce low-level rotation vary from case to case; the variety of observed storm types (Doswell and Burgess 1993) hints that not all tornadogenesis mechanisms are alike. Investigation of such issues was a major motivation for VORTEX (Rasmussen et al. 1994).

In this paper, we examine a dataset that has awaited analysis many years but that relates directly to the goals of VORTEX. On 17 May 1981, an extensive dual-Doppler dataset of an isolated, tornadic supercell was collected with the National Severe Storms Laboratory (NSSL) 10-cm research radars (Fig. 1). Dual-Doppler volumes span a period beginning 30 min before an F2 tornado formed near Arcadia, Oklahoma, and ending 15 min after the tornado had dissipated. Previous studies of tornadic supercells using ground-based multiple-Doppler radar have not benefited from such a complete coverage of the supercell life cycle surrounding and including the tornadic stage (Brandes 1978, 1981; Ray et al. 1981; Brandes et al. 1988). In addition, the mean interval between consecutive dual-Doppler volumes (5 min) affords better time resolution than is typically achieved in airborne Doppler radar studies (Dowell et al. 1997; Wakimoto et al. 1996).

A 444-m-tall instrumented tower sampled the Arcadia supercell as the storm moved over the north side of Oklahoma City. A high-resolution cross section of the wind and temperature structure both within the pretornadic storm (including the low-level updraft and mesocyclone) and its environment were obtained, affording a unique opportunity for intercomparison of multiplatform observations. The tower measurements document details of the boundary layer flow that are relevant to the development of low-level rotation but that cannot be resolved by distant Doppler radars.

The main purposes of the research described in this paper are twofold: to compare independent measurements of the wind field within the 17 May 1981 supercell and to look for clues about the processes of low-level mesocyclogenesis and tornadogenesis.

2. Description of the 17 May 1981 dataset

a. The Arcadia storm and its environment

Conditions on 17 May 1981 were characteristic of a classic tornado outbreak in Oklahoma (Taylor 1982). A short-wave trough at 500 mb over the southern Rockies during the morning moved into the Plains during the day. Ahead of the trough, a warm front at the surface raced northward through Oklahoma, and a dryline advanced eastward into west-central Oklahoma. Afternoon temperatures near 30°C and dewpoints over 20°C yielded a potentially unstable environment over central Oklahoma.

Thunderstorm development began near the intersection of the warm front and dryline in northern Oklahoma around 1400 CST (all times in CST); later convective initiation occurred progressively farther south along the dryline (Brewster 1984). First echoes of what was to become the Arcadia storm appeared on the NSSL Norman radar display before 1500.

A sounding taken from Edmond, Oklahoma (less than 10 km west of where the tornado later occurred), at 1430 (Figs. 1 and 2) indicates that the environment was characterized by substantial conditional instability and strong vertical shear of the horizontal wind. Since in the environment the convective available potential energy (CAPE) was 2250 J kg−1 (2500 J kg−1 if computed using virtual temperature), the 0–6-km shear near 5 × 10−3 s−1, and the bulk Richardson number 22, the environment was supportive of supercell formation (Weisman and Klemp 1984).

Despite the strong vertical wind shear, the vertically integrated storm-relative environmental helicity from 0 to 3 km AGL was only 110 m2 s−2, that is, below the threshold value for tornadoes of 160 m2 s−2 suggested by the work of Davies-Jones et al. (1990). Owing to counterclockwise curvature in the hodograph between 2 and 4 km AGL (Fig. 2b), the 2–3-km layer contributes no net helicity. Davies and Johns (1993) have suggested that the height of the level of free convection (LFC) (nearly 2 km AGL in this case) (Fig. 2a) may be a more appropriate storm inflow depth to use in the helicity integration. The 110 m2 s−2 helicity from 0 to 2 km for the Edmond wind profile has the same helicity density as a 0–3-km layer with the threshold 160 m2 s−2 helicity.

Between approximately 1700 and 1710 on 17 May 1981, an F2 tornado carved a path 5-km long south of the town of Arcadia, Oklahoma (Taylor 1982), destroying a mobile home, downing power poles, and causing heavy damage to timber (D. Burgess 1993, personal communication). Twelve dual-Doppler volumes of the Arcadia storm were collected by the Norman and Cimarron 10-cm research Doppler radars in central Oklahoma (Fig. 1) over a period of nearly an hour beginning at 1628. Thus, a relatively detailed documentation is available of the storm before, during, and after the tornado, except for a 10-min gap between dual-Doppler volumes during the pretornadic stage, when the Norman radar was briefly focused on another supercell in the vicinity.

Throughout the period covered by the Doppler dataset, the geometry was favorable for dual-Doppler analysis; the between-beam angle exceeded 45° for the mesocyclone region of the storm (Fig. 1). During the tornadic stage, the mesocyclone was approximately 45 km from each radar; at this range, the low-elevation beams were 400–500 m above the ground, which was below cloud base. When Doppler data collection ended at 1725, the Arcadia storm was losing supercell characteristics.

b. Dual-Doppler analysis methodology

The Norman (Cimarron) radar collected data at 1.0° (0.6°) azimuthal increments; typical elevation angle increments between sweeps were 0.8°–2.0°. Each radar has a beamwidth of 0.8°. Reflectivity and radial velocity data from each radar were interpolated to Cartesian grids with horizontal and vertical grid spacings of 0.8 km and 0.5 km, respectively. Given the 1.0° azimuthal increment between beams for the Norman radar, the 0.8-km spacing corresponds to the distance between adjacent gates during the tornadic stage of the supercell, when the storm was 45 km away from the radar. The vertical grid structure was chosen such that the lowest level of the grid was at ground level (a physical boundary), and the next lowest level was near the height of the low-elevation radar sweeps. A spherical influence region of radius 1.2 km was used for the Cressman (1959) interpolation to the grid. The radius of influence was large enough to account for the vertical spacing between consecutive sweeps and to produce a relatively smooth wind field. Objectively analyzed data at ground level were extrapolated downward from the height of the lowest sweeps (400–500 m AGL). Thus, implicit in the Doppler analysis is the assumption that velocities (and divergence) are constant in the lowest 500 m.

Ground clutter, range-folding contamination, sidelobe contamination, and other noisy data were removed before the analysis. Range-folding contamination was significant for this dataset owing to distant severe storms north of the Norman radar; however, this contamination primarily affected the light precipitation region in the central and northeast portions of the storm. Radial velocities were dealiased manually. We used an empirical relationship between the radar reflectivity factor and raindrop terminal fall speed to subtract the component of the latter from the interpolated radial velocities.

To account for motion of the storm while data were being collected, the data were advected horizontally to locations corresponding to a central time of each data volume. The beginning of each volume was 1–2 min before the center time. Storm motion was estimated by visually correlating reflectivity features on the southwest side of the storm between successive volumes. The analyses confirmed that this storm motion estimate matched that of the low-level mesocyclone during the period surrounding and including the tornado.

The three-dimensional wind field was obtained from the dual-equation system using iterative downward integration of the anelastic mass continuity equation subject to the condition w = 0 at the storm top; downward integration was chosen because it tends to damp the vertical propagation of errors during the integration (Ray et al. 1980). Due to inexact upper boundary conditions and accumulation of divergence errors, this process typically yielded nonzero vertical velocities at the lower boundary. Therefore, a constant horizontal divergence correction was added to each column that would, upon reintegration of the continuity equation, require the vertical velocities to be zero at both the ground and the storm top; this procedure is commonly termed the linear O’Brien (1970) adjustment.

Since the objective analysis scheme produced a qualitatively smooth horizontal wind field, no further smoothing was applied to the horizontal velocities. Vertical velocities were noisier in appearance; a single-pass Leise (1981) filter was applied in order to eliminate features with wavelengths shorter than 2 km. Quantities such as vorticity were computed using centered finite differences. Doviak et al. (1976) and Ray et al. (1980) provide a discussion of the errors involved in dual-Doppler wind synthesis. Mesocyclone-scale (several kilometers), but not tornado-scale (several hundred meters), features will be resolved in the dual-Doppler analyses.

c. Characteristics of the instrumented tower

The low-level inflow region, mesocyclone, updraft, and rear precipitation core of the pretornadic Arcadia storm all passed across the 444-m-tall instrumented tower (Carter 1970) on the north side of Oklahoma City (Fig. 1). Horizontal wind, vertical velocity, temperature, wet-bulb temperature, and other variables were measured at 1.5-s intervals at up to seven levels between the surface and 444 m. In the analysis, instrumented tower data were smoothed using a 20-s running mean. Since pressure measurements made on the tower appear to be erroneous, they were not used in the analysis.

3. Storm morphology and evolution

The Arcadia storm was a mature supercell by the time it was observed simultaneously by the Norman and Cimarron radars, as evidenced by a well-defined mesocyclone–mesoanticyclone couplet in the early scans at midlevels (Fig. 3a). (For the purposes of this paper, we refer to a relatively long-lived storm with updraft rotation on scales of several kilometers as a “supercell.”) The storm had already produced, and was continuing to produce, golfball-sized hail (Taylor 1982). For the horizontal scales resolved in the analysis, mesoanticyclone vorticity at 5.0 km AGL tended to exceed mesocyclone vorticity in magnitude (e.g., −0.019 s−1 vs +0.012 s−1 at 1647 CST) during the pretornadic stage (Fig. 3b).

Trajectories (not shown) are consistent with vortex couplet production by the tilting of horizontal vorticity associated with environmental vertical shear. The air on the northwest side of the mesoanticyclone and the far south side of the mesocyclone originated at midlevels, while the air within the southeast part of the mesoanticyclone and within much of the mesocyclone ascended from low levels. This finding agrees with conventional theory, which states that midlevel mesovortices in supercells arise from tilting of environmental horizontal vorticity by the primary storm updraft (Barnes 1970; Brandes 1984b; Rotunno and Klemp 1985). Horizontal vortex tubes (normal to the vertical shear) associated with the mean shear between the surface and 4 km AGL (Fig. 2b) were oriented from southeast to northwest; tilting of such vortex tubes by an updraft would produce a mesocyclone–anticyclone configuration like that observed in the pretornadic stage of the storm (Figs. 3a,b).

During the time period roughly coinciding with the tornadic stage, the midlevel vortex couplet gradually shifted from a northwest-to-southeast orientation (Fig. 3b) to a northeast-to-southwest configuration (Fig. 3c). This may be related to weakening updrafts at 5 km at this time (Fig. 4). As the updraft dwindled, the mesoanticyclone was advected downstream to the northeast (Fig. 3c) relative to the mesocyclone; in contrast, the mesocyclone tended to remain collocated with the updraft (e.g., Fig. 3a), even while the updraft was weakening. Investigation of the magnitudes of the terms in the vorticity equation for the mesoanticyclone (not shown) confirms that the relative importance of horizontal advection was increasing compared to the other terms (stretching, tilting, and vertical advection) as the midlevel updraft was weakening.

Close inspection of the supercell storm reveals that it consisted of multiple cells (Dowell et al. 1997) during the Doppler radar observation period. The lobe of relatively high reflectivity at 1634 CST around x = −10, y = 45 is associated with both a weak vorticity couplet and an updraft (Fig. 3a). This small cell moved rapidly northeastward away from the primary cell. It appears that storm splitting (Klemp and Wilhelmson 1978) was in progress when Doppler scanning began. The primary cell had much stronger updrafts (Fig. 3a, around x = −6, y = 37) and higher reflectivities than the smaller left-split cell; the primary cell eventually became tornadic.

While the tornado was forming, a third cell was developing on the south flank of the storm east–southeast of the tornadic cell. By 10 min after tornado dissipation, reflectivity in the posttornadic core had decreased to less than 50 dBZ (around x = 17, y = 55 in Fig. 3c) and had reached 60 dBZ in the new core (around x = 25, y = 50 in Fig. 3c). This new cell may have been invigorated by the low-level convergence associated with the surging outflow (Fig. 5e) from the tornadic cell. However, the new cell failed to become tornadic (Taylor 1982). Later single-Doppler scans indicated a weakening, multicellular, disorganized structure to the storm.

The most notable change in the low-level pretornadic horizontal winds between 1630 and 1643 was the development of strong storm-relative northerlies on the southwest side of the storm, which occurred coincident with a growing appendage in reflectivity (Fig. 5). At 1630 there were strong inflow southerlies (and southeasterlies) along the entire southern flank of the storm. At 1643 air still entered from the south to the east of the mesocyclone, but the wind field southwest of the mesocyclone was dominated by a surge of outflow northerlies.

By 1651, the storm-relative low-level winds on the southwest side of the mesocyclone had backed to northwesterly behind the advancing gust front (Fig. 5c). Then, with the development of westerlies south of the vorticity maximum, the low-level storm-relative flow became circular (Figs. 5d,e). The onset of such storm-relative westerlies within the hook often signifies the transition to the tornadic stage (Brandes 1978, 1981). A sharp increase in Cimarron radar Doppler velocities (up to 30 m s−1 ground relative) in the southern (outbound) member of the inbound–outbound couplet immediately preceded tornado touchdown. The tornado eventually dissipated as the outflow surged rapidly eastward (Figs. 5e,f), leaving the low-level mesocyclone behind.

Immediately before the tornado touched down, a pronounced reflectivity hook developed on the southwest side of the isolated supercell (Fig. 6). The Arcadia tornado was associated with a local horizontal shear anomaly in the single-Doppler radial velocities (Brown et al. 1978) at the tip of the hook (Brandes 1981). This feature shows up most clearly in the low-level Cimarron radial velocities at 1704 and the low-level Norman Doppler velocities at 1710 (not shown). No clearly defined shear anomaly was present prior to the tornadic stage. In these respects, this case is similar to the 20 May 1977 Del City storm (Brandes 1981). We will refer to the shear anomaly as a TVS, although it lacks the time continuity of a classic TVS (Brown et al. 1978). Gate-to-gate shears at 1704 (not shown) were 4.0 to 6.8 × 10−2 s−1 in the lowest six elevation scans by the Cimarron radar, which is on par with what was observed in the Del City storm (Brandes 1981).

Tornadogenesis coincided with the time when the storm top was collapsing (Fig. 7), reflectivity was generally decreasing, and midlevel updrafts were weakening (Fig. 4) (Fujita et al. 1976; Burgess 1982). However, the dual-Doppler measurements suggest the low-level updrafts were stronger during the tornadic and immediate pretornadic stages (Fig. 4). Tornadogenesis also corresponded to the time when low-level mesocyclone vorticity was increasing rapidly (Fig. 8). At 1643, midlevel vorticity exceeded that at low levels; during tornadogenesis (1651 to 1704) midlevel vorticity was increasing in magnitude, but low-level vorticity was increasing even more rapidly, as is common in tornadic supercells (Brandes 1984a). After the tornado had dissipated, the vertical gradient of vorticity was reversed; low-level vorticity exceeded midlevel vorticity after 1715, as is also commonly observed (Brandes 1984a).

The vertical velocity field at 1704 at 1.0 km AGL (Fig. 9) exhibits classic tornadic structure (Lemon and Doswell 1979; Brandes 1984b); an arc-shaped updraft region surrounds a downdraft at the rear of the storm. The tornado formed within the mesocyclone in a zone of strong vertical velocity gradient (but within the updraft) near the nose of the rear downdraft. Further details of tornadogenesis will be discussed in section 5.

To see if any of the flow features close in size to the minimum resolvable wavelength (∼2 km) could be better illustrated, we reanalyzed the Doppler volumes at 1704 and 1710 on a finer grid (500-m horizontal grid spacing with a 500-m radius of influence). With such a small radius of influence, numerous gaps between data in the vertical were produced in the grid, which precluded the computation of vertical velocities on the fine grid. However, since the radial velocities at elevation angles near zero are relatively unaffected by vertical particle motion, we are able to estimate well the horizontal velocities at low grid levels where there are radial velocity data from both radars. The resulting analyses show more detail in the flow (Fig. 10).

The TVS at 0.5 km AGL was within 500 m of the estimated damage path of the tornado (Fig. 10). The tornado formed southeast of the center of the low-level mesocyclone and was associated with a local vorticity maximum within the larger-scale cyclonic flow.

4. Instrumented tower data

The low-level mesocyclone and updraft of the Arcadia storm passed the instrumented tower on the north side of Oklahoma City (Fig. 1) around 1638. We will compare the dual-Doppler analysis centered at this time to the tower data through the use of time-to-space conversion of the latter (Barnes 1973; Goff 1976; Johnson et al. 1987). For a storm that is relatively steady state, the time cross section through the storm at low levels can be thought of as a spatial cross section. A detailed quantitative comparison of tower versus Doppler wind measurements is possible but is beyond the scope of this paper. Instead, we will focus on a few key similarities and differences in the analyses of the low-level flow.

Storm-relative dual-Doppler winds at 1638 at 500 m AGL (Fig. 11a) capture many of the features in the flow sampled directly by the tower (Fig. 11b). Both indicate strong east–southeasterly inflow into the storm and strong northerlies within the growing reflectivity appendage on the southwest side of the storm.

Although the independent measurements of wind directions within the storm agree, the dual-Doppler wind speeds tend to be underestimates of the tower wind speeds (Fig. 11c). Some of the differences in wind speed may be the result of storm evolution. At the time of the dual-Doppler analysis (1638), the strength of the storm-relative inflow was decreasing, while the strength of the storm-relative outflow was increasing (Figs. 5a,b). At the edges of Fig. 11c, where corresponding measurements were made approximately 10 min apart, it would thus appear to be improper to compare the values.

For the measurements made within 4 min of 1638, the mean difference in magnitude between dual-Doppler and tower winds is 6 m s−1. Sources of disagreement may include (a) slightly different heights of the measurements (444 m for the tower, 500 m for the Doppler analysis); (b) differences in measurement type (point vs volumetric samples); (c) differences in filtering of the raw data; and (d) ground clutter contamination of the Doppler velocities.

A comparison of the vertical shear of the horizontal wind measured by the tower between the “surface” (7 m AGL) and 444 m AGL with that in the Doppler analysis over the 500–1000-m layer revealed little correlation between the two (Fig. 12). Differences in the southwestern portion of the storm may be partly affected by the heights over which the shear is calculated, relative to the depth of the storm outflow. However, this may not explain all dissimilarities within the storm. Strong vertical shear (1.5 to 2.5 × 10−2 s−1 over the depth of the tower) was typical of the storm environment, including the inflow portion of the storm (Fig. 12b). The environmental sounding (cf. Fig. 2) indicates strong shear (approaching 1 × 10−2 s−1) also over the 500–1000-m layer. However, the Doppler analysis resolves very little shear in the storm inflow (Fig. 12a). We believe this highlights a limitation of the Doppler data; horizontal vorticity associated with the vertical shear of the horizontal wind is not resolved at low levels. We conclude that a detailed vorticity budget associated with tornadogenesis near the surface is not possible. Lack of resolution at low levels has been a problem in general for multiple-Doppler datasets of tornadic storms (Brandes 1984b, 1988).

Both the in situ and remotely sensed measurements indicate a strong updraft within the cyclonically curved low-level flow (Figs. 11a, 13). However, the analyzed updraft centers are displaced 1 km from each other, which results in a slight difference in the analyses of the phase relationship between the updraft and the region of maximum cyclonic curvature in the horizontal winds. The displacement may be the result of a southwestward tilt with height of the low-level updraft (Fig. 14); the Doppler radars did not sample the vertical variation in convergence in the lowest 500 m necessary to resolve the updraft slope.

The maximum vertical velocity measured by the tower within the updraft is roughly a linear function of height in the lowest 266 m (not shown). However, above 266 m the linear relationship breaks down; the maximum updraft speed at 444 m is only slightly greater than at 266 m. Since the convergence was mostly concentrated below the height of the lowest Doppler-radar beams, it is not surprising that the peak updraft speed (Fig. 13) in the Doppler analysis (nearly 3 m s−1) is only half that measured by the tower (6 m s−1). The Doppler data are limited by the radar horizon and by smoothing inherent in both the remote data collection and the postanalysis. Overall, the comparison of the dual-Doppler analysis and instrumented tower data suggests that the former is able to resolve features in the low-level winds with horizontal scales of around 2 km or longer but that the magnitude of the wind speeds in such features may be underestimated.

The tower observations and the Edmond sounding (Fig. 2a) indicate that the updraft air at 444 m is neither saturated nor buoyant; a virtual temperature deficit in the updraft of approximately 1 K is suggested relative to the environmental air sampled by the instrumented tower 1 h previously (and by the Edmond sounding 2 h previously). The rising motion must have been associated with an upward-directed perturbation pressure gradient force (Brandes 1984a; Wicker and Wilhelmson 1995). With additional lift of around 1 km without entrainment, the air would have become buoyant (Fig. 2a).

We applied a pressure retrieval to the Doppler analyses to determine whether an upward pressure gradient force could be resolved within the low-level updraft. Perturbation pressure and virtual temperature were retrieved using the technique of Hane and Ray (1985), except that the velocity time tendency terms were also incorporated into the current study. Each retrieval is based upon a pair of consecutive dual-Doppler volumes. Velocity time derivatives (estimated by finite differences) and mean velocities are computed in a storm-relative reference frame.

The retrievals around 1638 (not shown) indicate a slight downward-directed perturbation pressure gradient force within the low-level (1.0 km AGL) updraft, which is inconsistent with the observation that nonbuoyant air was being lifted in the updraft. Furthermore, the thermodynamic retrieval (not shown) failed to resolve the observed temperature gradient of 3 K km−1 across the mesocyclone (Fig. 14b). The radar horizon problem, ground clutter problems, and smoothing in the Doppler analysis limit the ability of the thermodynamic retrieval to resolve the low-level temperature field. The thermodynamic retrieval was further hindered by the presence of hail in the storm. Since the Arcadia storm was a prolific producer of golfball-sized hail (Taylor 1982) and contained radar reflectivity factors as high as 70 dBZ, there is considerable uncertainty in the retrieval owing to difficulty in relating the reflectivity factor to precipitation mixing ratio. Additionally, there may be error owing to the general difficulties in applying a thermodynamic retrieval to a volume with incomplete data (Hane and Ray 1985).

A vertical cross section within the low-level mesocyclone (Figs. 14a,b) reveals two features that are reminiscent of a density current (Goff 1976): a sloped baroclinic zone and a relatively narrow updraft that precedes it. However, the horizontal winds indicate that no gust front surge along the direction of storm motion (east-northeast) has yet developed. At the time of this cross section (1635–1645), the outflow was surging southward in a storm-relative sense (Fig. 5b). Shear in the wind component normal to the cross section is the dominant feature in the horizontal winds (Fig. 14c).

Equivalent potential temperature θe in the inflow and updraft air was as high as 349 K (Fig. 14a). In the rainy downdraft air on the west side of the mesocyclone, θe was as low as 335 K; the Edmond sounding (Fig. 15) suggests that this air had descended from above 750 mb. The secondary θe maximum farther west (t = 1642 to 1643 in Fig. 14a) is associated with high dewpoints within the heavy rain and area of strongest descending motion.

5. Low-level mesocyclogenesis and tornadogenesis

a. Theory from numerical simulations and laboratory experiments

The frictionless Boussinesq vorticity equations are given by
i1520-0493-125-10-2562-e1
where D/Dt represents the total derivative, ζ is the vertical vorticity, ω is the 3D vorticity, h subscripts denote horizontal components, B is buoyancy, and the other quantities have their usual definitions. The first term on the right-hand side of (1) represents tilting of horizontal vorticity into the vertical; the second is the stretching of vertical vorticity. The first term on the right-hand side of (2) includes stretching of horizontal vorticity and tilting of vortex lines; the second represents baroclinic generation of horizontal vorticity.

Significant vertical vorticity is already present in some thunderstorm environments and can be readily amplified by stretching at the base of a thunderstorm updraft (Wilson 1986; Brady and Szoke 1989; Wakimoto and Wilson 1989). We will focus here on how mesocyclogenesis and tornadogenesis can occur when such vorticity is not already available.

According to conventional theory, supercell thunderstorms derive their rotation (about a vertical axis) from vorticity that is initially purely horizontal (Barnes 1970; Brandes 1984b; Rotunno and Klemp 1985). Tilting of horizontal vorticity into the vertical by the primary storm updraft, followed by stretching of vertical vorticity in the updraft, is the traditional conceptual model for the formation of a mesocyclone (vertical vorticity ≳ 0.01 s−1). However, recent work suggests that this mechanism is insufficient to explain tornado-intensity rotation (i.e., vorticity greater than 1 s−1) in the lowest few hundred meters of the atmosphere (Davies-Jones and Brooks 1993; Walko 1993). By the time tilting and stretching in the updraft have given an air parcel tornadic vertical vorticity, the parcel has risen far away from the surface. Instead, Walko (1993) and Davies-Jones and Brooks (1993) hypothesize that a downdraft is also required in the genesis of the tornado’s parent circulation.

An idealized numerical experiment by Walko (1993) demonstrates how a downdraft aids in low-level vortexgenesis. This process is investigated in terms of the circulation (Rotunno and Klemp 1985) about a material curve (i.e., the area integral of the component of vorticity normal to a surface outlined by the curve) in the fluid. Circulation is defined by
i1520-0493-125-10-2562-e3
where v is the 3D velocity, l is the space coordinate, and the integration is taken along a closed curve. The circulation about a curve moving with the fluid changes only owing to baroclinic and friction effects.

In the Walko (1993) experiment, the circulation of a low-level vortex is derived from the environmental vertical shear of the horizontal wind; a simplified model of how this mechanism works is depicted in Fig. 16. There is a circulation about the initially tilted material curve owing to the vertical wind shear. This circulation is conserved as the downdraft tips the material circuit into a horizontal plane just above the surface. Equivalently, weak vertical vorticity is produced via tilting of horizontal vorticity over a region many radii larger than the eventual vortex core (Walko 1993). When the material circuit then enters the base of the updraft, the low-level vortex spins up as the curve converges. The situation is more complicated inside a thunderstorm, where momentum transported by a downdraft could be representative of both the environment and a perturbation produced by the storm.

Davies-Jones and Brooks (1993), while also stressing the importance of the downdraft in tilting low-level horizontal vorticity, diagnosed the development of a low-level rotation that originated as baroclinically generated horizontal vorticity. In their numerical thunderstorm simulation, horizontal vorticity produced at the edge of the rain-cooled air north of the updraft was tilted into the vertical by the vertical velocity gradient within descending air. The vertical vorticity was then amplified as it entered the region of low-level convergence associated with the updraft. A similar mechanism was noted in a more idealized numerical experiment by Trapp and Fiedler (1995).

In the Davies-Jones and Brooks (1993) process, the material curve that eventually encloses the low-level vortex initially has little circulation about it. However, circulation about the tilted fluid circuit builds up owing to baroclinicity in the inflow (i.e., a temperature gradient pointing in the +y direction in Fig. 16). Otherwise, the subsequent flattening and shrinking of the material curve is similar to that depicted in Fig. 16.

We now return to the Arcadia storm to compare what was observed in nature to the theories of tornadogenesis deduced from numerical simulations and laboratory experiments.

b. Observations of the Arcadia storm

Analysis of the trajectories of material curves most clearly demonstrates the difference between the pretornadic and tornadic low-level mesocyclones. Unfortunately, it is difficult to obtain quantitative results from a trajectory and circulation analysis applied to real data, owing to limited spatial resolution and noise in the data. For the Arcadia storm, no radar scans of the mesocyclone are available below 400 m AGL. Data at the lowest level of the Cartesian grid have been extrapolated downward from the height of the lowest radar scans; in reality, the tower measurements confirm significant vertical shear of the horizontal winds in this layer (Fig. 12b). We will focus on features in the flow at 500 m AGL. However, since many trajectories ascend to this level, the results must be interpreted with caution. We anticipate that the trajectory analysis will still provide a sense of parcel origins. The trajectories depicted herein represent an improvement over those in a preliminary analysis (Dowell and Bluestein 1996) in that velocity time derivatives have been incorporated into the current calculations.

In the pretornadic Arcadia storm (Figs. 5b, 17a), all air parcels entering the mesocyclone at 500 m AGL originated east of the mesocyclone. The west side of the material curve was tipped upward by the updraft, and thus low-level horizontal vorticity was tilted into the vertical.

During the tornadic stage (Fig. 17b), some air parcels also approached from the east and rose into the low-level mesocyclone. However, during tornadogenesis air parcels also approached the mesocyclone from the north and west (Davies-Jones and Brooks 1993); air parcels were converging from a broader area during tornadogenesis than earlier.

Instrumented tower data provide insight into the nature of the vorticity source for the low-level mesocyclone; the low-level air was rich in horizontal vorticity. Tower measurements far from the storm indicate a vertical shear of approximately 1.5 × 10−2 s−1 in the lowest 440 m of the atmosphere (Fig. 18). Assuming that vertical motions were relatively weak far from the storm, we can infer that the horizontal vorticity (which is equal in magnitude, but perpendicular, to the vertical shear) was directed toward the west-southwest. As the storm approached, the magnitude of the vertical shear gradually increased to 2.8 × 10−2 s−1 at 1634 (Fig. 18) and then peaked at over 3.0 × 10−2 s−1 within the updraft (not shown). Possible explanations for the increased low-level vertical shear from the far storm to the near storm environment include stretching of horizontal vorticity in the accelerating storm inflow and baroclinic generation of horizontal vorticity.

A gradual temperature drop of 5°C (2°C) at the surface (at 444 m AGL) was experienced by the instrumented tower during a period of over 1 h between the sunny environment outside the storm and the shaded region with light precipitation immediately upstream of the updraft (not shown). Preliminary analyses of supercells during VORTEX (Rasmussen et al. 1994) have shown forward-flank storm baroclinicity to be rather diffuse (Rasmussen and Straka 1996; P. Markowski 1997, personal communication). The instrumented-tower temperature trace of the Arcadia storm also hints that the temperature gradient was spread over a broad region, but we lack measurements perpendicular to the tower cross section necessary to confirm this. Although baroclinicity may have been weak, air parcels residing in the broad baroclinic zone for long periods could have still acquired significant horizontal vorticity (P. Markowski 1997, personal communication).

The west-southwest to east-northeast orientation of the reflectivity contours (Figs. 5, 11) beneath the downstream anvil of the storm suggests that the temperature contours were oriented similarly (i.e., cool air within the precipitation region, warm air outside the storm). This implies baroclinic generation of horizontal vorticity directed toward the west-southwest, which is the same direction as the preexisting environmental horizontal vorticity and the same direction as the observed increase in vorticity as the storm approached (Figs. 12b, 18). Thus, low-level mesocyclogenesis may have been enhanced by the addition of baroclinic effects to the preexisting horizontal vorticity (Wicker 1996).

Since the orientation of the tower cross section through the pretornadic low-level updraft is approximately parallel to the low-level horizontal vorticity (Figs. 12b, 13), we are able to compute a rough estimate of the rate of vertical vorticity generation by tilting on the inflow (east) side of the updraft. Tower measurements at 444 m indicate a vertical velocity gradient of 6 m s−1 over a distance of 1 km (Fig. 13). If we assume that half this magnitude of gradient is representative of the overall depth of the tower and that horizontal vorticity is of magnitude 2.5 × 10−2 s−1, then this implies a rate of tilting of horizontal vorticity into the vertical of
i1520-0493-125-10-2562-e4
With an updraft-relative flow of about 15 m s−1 (Fig. 11b), parcels could pass through the vertical velocity gradient in 70 s, long enough for a vertical vorticity of approximately 0.5 × 10−2 s−1 to be produced by tilting. This vorticity could then be amplified within the updraft by convergence (Brandes 1984b).

The tornado formed southeast of the mesocyclone center at 0.5 km AGL in a region of horizontally sheared flow between westerlies on the rear downdraft side of the mesocyclone and storm-inflow southeasterlies (Figs. 9 and 10). Storm intercept teams have long observed that the formation of the rear downdraft precedes the formation of significant supercell tornadoes (Lemon and Doswell 1979); the rear downdraft is manifest visually by a clear slot, which is a relatively bright region within the cloud on the right side (with respect to storm motion) of the mesocyclone. With the limitations of the radar data, it is difficult to diagnose accurately the role the rear downdraft plays in tornadogenesis. However, the proximity of the tornado to the nose where the rear downdraft penetrates into the mesocyclone (Fig. 9) suggests a dynamical relationship between the two features, although the nature of the relationship is unclear. Westerly winds that were a part of the horizontally sheared flow in which the tornado was embedded were found just downstream of the rear downdraft, suggesting that westerly momentum brought downward in the downdraft may have played an important role in the development of vertical vorticity near the ground, as in the Walko (1993) simulation (Fig. 16).

In the Arcadia storm, the rear downdraft appeared to separate from the rainy downdraft north of the updraft and with time progress cyclonically around the updraft (Fig. 19). Numerical simulations of supercells indicate that the primary storm-scale rear downdraft (north and northwest of the updraft in the Arcadia storm) is driven by the cooling associated with evaporation of precipitation (Klemp and Rotunno 1983; Brooks et al. 1994). As the mesocyclone intensifies, it wraps precipitation around the updraft (Brooks et al. 1994), which promotes downdraft development farther behind the updraft (west of the updraft in the Arcadia storm).

Dynamic effects associated with the intensifying mesocyclone may also lead to downdraft formation. In the simulation of Klemp and Rotunno (1983), the “occlusion downdraft” (that portion of the rear flank downdraft within the low-level mesocyclone and immediately behind the convergence line of the occluded gust front) was driven dynamically by the strong circulation. Large vorticity near the surface with weaker rotation aloft was associated with a downward-directed perturbation pressure gradient force.

Pressure retrievals (Hane and Ray 1985) for the tornadic stage of the Arcadia storm do suggest a downward perturbation pressure gradient force south and west of the mesocyclone at low levels (e.g., near y = 45 in Fig. 20a), which could be related to the vertical velocity minimum that wraps into the southeast side of the mesocyclone (Fig. 19). This region of downward-directed pressure forces is apparent only in the retrievals from 1651 through 1713 (i.e., from tornadogenesis to tornado demise).

In the numerically simulated storm of Klemp and Rotunno (1983), downward acceleration owing to dynamic effects was primarily driven by strong vorticity near the surface. The region of positive ∂p/∂z in the Arcadia storm is related both to low pressure near the surface and to high pressure south and west of the mesocyclone aloft (at 2–3 km AGL) (Fig. 20). Relatively high pressure coincided with diffluence in the flow west-southwest of the mesocyclone and confluent flow south and southeast of the mesocyclone (Fig. 20b); both of these flow features developed as the mesocyclone at 2–3 km was intensifying (Fig. 8) just prior to tornado formation. However, as discussed in section 4, the analyses must be interpreted with caution owing to the many difficulties in retrieving pressure from dual-Doppler measurements (Hane and Ray 1985). Further diagnosis of rear downdraft formation and its role in tornadogenesis will likely require the use of numerical simulations.

The traditional explanation for tornado demise is that the outflow surges ahead of the primary updraft and mesocyclone, cutting off the low-level inflow into the updraft and, in general, disrupting the balance of mesocyclone processes (Lemon and Doswell 1979; Brooks et al. 1994). This is consistent with the analysis of the Arcadia storm at least in that the Doppler analysis shows westerly outflow winds farther and farther east of the updraft and mesocyclone in successive volumes around the time of tornado dissipation (cf. Fig. 5).

6. Conclusions

In many ways, the Arcadia supercell was remarkably similar to storms that have been documented in previous dual-Doppler studies, including the Del City storm (Brandes 1981). Hopefully the common behavior of the storms indicates that there is indeed a common mechanism acting to produce many of the tornadoes within isolated supercells.

When radar observations began, the Arcadia storm had a strong updraft and mesocyclone–anticyclone couplet aloft but only a weak mesocyclone at low levels. With time, low-level northerly outflow increased on the rear side of the storm, a reflectivity appendage formed, and the rear downdraft wrapped cyclonically around the updraft.

Tornadogenesis occurred while low-level mesocyclone vorticity was rapidly increasing. The tornado itself was located southeast of the mesocyclone center, at the tip of the reflectivity hook. During tornadogenesis, convergence was increasing within the low-level mesocyclone.

The tornado dissipated while low-level outflow surged well ahead of the primary updraft and mesocyclone. After tornado demise, mesocyclone vorticity at low levels exceeded that aloft, and the Arcadia cell soon lost supercell characteristics. Another cell developed near the leading edge of the outflow, but it proved to be short lived.

Before the time of tornadogenesis, the low-level mesocyclone of the Arcadia storm consisted of a rising stream of environmental air that entered the storm from the east (Fig. 17a). With time, a rear downdraft formed southwest of the updraft, and air began to converge from a broader region into the mesocyclone (Figs. 17b, 19).

The development of the rear downdraft in the Arcadia storm occurred simultaneously with approximately a doubling of vorticity in the low-level mesocyclone. The intensification of the two features may result from a synergetic interaction. As the rear downdraft intensifies, it may aid in strengthening the low-level mesocyclone through downward momentum transport and/or increased convergence at the edge of where the downdraft air spreads out at the surface. The intensification of the mesocyclone may in turn lead to strengthening of the rear downdraft by dynamic effects and/or wrapping of precipitation around the updraft. Pressure retrievals for the Arcadia storm suggest that the low-level rear downdraft south and southwest of the mesocyclone center during the tornadic stage was in a region with a downward-directed perturbation pressure gradient force associated with relatively high pressure aloft southwest of the intensifying mesocyclone and low pressure near the surface.

Perhaps more importantly, a downdraft may be necessary to produce the strong, localized, near-surface vorticity characteristic of tornadoes, as hypothesized by Walko (1993) and Davies-Jones and Brooks (1993). The Arcadia tornado formed near the nose of the rear downdraft (Fig. 9) in a region of horizontally sheared flow between southeasterly storm inflow and westerly outflow downstream of the rear downdraft (Fig. 10). The evidence, although circumstantial, suggests an active role of the rear downdraft in tornadogenesis.

The most novel aspect of this study is that it is the first to synthesize independent measurements of the 3D low-level flow in and near the mesocyclone of a supercell. Doppler analyses provide an overall perspective of the wind field within the Arcadia storm and instrumented tower measurements resolve some of the finer-scale details of the low-level flow and thermodynamic structure.

Tower data show that the low-level environment of the Arcadia storm was rich in horizontal vorticity; this vorticity entered the pretornadic storm from the east in an approximately streamwise fashion and was tilted upward into the low-level mesocyclone. Instrumented tower measurements indicate low-level vertical shear immediately upstream of the updraft almost twice as great as observed farther away in the environment. This increase may have resulted from enhancement of the low-level horizontal vorticity by baroclinic generation and/or stretching of horizontal vorticity in the storm inflow.

Acknowledgments

The Doppler radar data were analyzed using the following software developed at the National Center for Atmospheric Research: rdss (data editing), reorder (objective analysis), and cedric (wind synthesis). NCAR is supported by the National Science Foundation. Graphics were generated using zxplot, developed at the University of Oklahoma by Ming Xue.

We are indebted to Don Burgess, who suggested that we take a look at this case. Jerry Wardius provided a copy of the radar and instrumented tower data. David Stensrud made upper-air and surface data available to us. Carl Hane assisted with the pressure retrieval and Jeff Trapp provided helpful comments on some of the topics in this paper.

This work was funded by NSF Grant ATM-9302379.

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Fig. 1.
Fig. 1.

Map of the locations of the instrumented tower, sounding (Edmond), radar sites (Norman and Cimarron), and town of Arcadia, Oklahoma. The Norman radar is located at the origin. The tornado occurred south of Arcadia between approximately 1700 and 1710 CST 17 May 1981. Contours of the reflectivity factor (dBZ) measured by the Cimarron radar at 1.0 km AGL at 1704 CST are also shown. The crescent-shaped region indicates where the dual-Doppler between-beam angle is greater than 45° (and less than 135°).

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2562:TAOSOM>2.0.CO;2

Fig. 2.
Fig. 2.

Sounding taken from Edmond, Oklahoma, at 1430 CST 17 May 1981. (a) Skew T–logp diagram. The solid and dashed lines indicate temperature (°C) and dewpoint (°C), respectively. Full (half) wind barbs represent 5 m s−1 (2.5 m s−1); flags represent 25 m s−1. Winds are shown at heights indicated in km AGL. Tower observations at 444 m AGL at 1638 CST (within updraft) are plotted as closed dots. (b) Hodograph. Each circle represents 5 m s−1 of wind speed. The heights are in kilometers AGL. The small triangle indicates observed storm motion (u = 10 m s−1, υ = 6 m s−1).

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2562:TAOSOM>2.0.CO;2

Fig. 3.
Fig. 3.

Cimarron radar reflectivity factor (dBZ) and horizontal storm-relative winds at 5.0 km AGL. The x (east–west) and y (north–south) distances (km) are relative to the Norman radar. Local maxima in cyclonic and anticyclonic vertical vorticity are marked “C” and “A,” respectively. Updrafts are hatched. (a) 1634 CST, 0.5 h before tornado, w > 15 m s−1 shaded; (b) 1647 CST, 0.25 h before tornado, w > 20 m s−1 shaded; and (c) 1722 CST, just after tornado, w > 11 m s−1 shaded.

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2562:TAOSOM>2.0.CO;2

Fig. 4.
Fig. 4.

The maximum vertical velocity vs height within the updraft of the Arcadia storm for various times.

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2562:TAOSOM>2.0.CO;2

Fig. 5.
Fig. 5.

Cimarron radar reflectivity factor (dBZ) and horizontal storm-relative winds at 0.5 km AGL; x and y as in Fig. 3. Gust front positions are marked by a dashed line: (a) 1630 CST; (b) 1643 CST; (c) 1651 CST; (d) 1704 CST, TVS location denoted by a dot and damage path indicated by a heavy line; (e) 1710 CST, TVS location denoted by a dot and damage path indicated by a heavy line; and (f) 1717 CST.

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2562:TAOSOM>2.0.CO;2

Fig. 5.
Fig. 6.
Fig. 6.

Reflectivity factor scan by the Cimarron radar at 0.7° at 1658 CST. Range is given in kilometers.

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2562:TAOSOM>2.0.CO;2

Fig. 7.
Fig. 7.

Time series of the height of the Arcadia storm top, as indicated by the 40-dBZ contour in raw Cimarron radar scans.

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2562:TAOSOM>2.0.CO;2

Fig. 8.
Fig. 8.

The maximum vertical vorticity (10−2 s−1) vs height within the mesocyclone of the Arcadia storm for various times.

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2562:TAOSOM>2.0.CO;2

Fig. 9.
Fig. 9.

Horizontal storm-relative winds and vertical velocity (m s−1) at 1.0 km AGL at 1704 CST; x and y as in Fig. 3. Contours for negative values are dashed. The TVS location is marked with a dot.

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2562:TAOSOM>2.0.CO;2

Fig. 10.
Fig. 10.

Horizontal storm-relative winds and vertical vorticity (10−2 s−1) at 0.5 km AGL for the high-resolution (500-m grid spacing) analyses; x and y as in Fig. 3. Contours for negative values are dashed. The locations of the TVS (closed dot) and tornado damage path (thin solid line segment) are marked: (a) 1704 CST and (b) 1710 CST.

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2562:TAOSOM>2.0.CO;2

Fig. 11.
Fig. 11.

Low-level horizontal storm-relative winds and Cimarron radar reflectivity factor (dBZ) at 1638 CST at 0.5 km AGL; x and y as in Fig. 3. (a) Dual-Doppler analysis winds at 0.5 km AGL at 1638 CST. (b) Time-to-space conversion of winds measured by the instrumented tower at 444 m AGL over a 30-min period centered at 1638 CST. (c) Vector difference between the winds in (a) and those in (b).

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2562:TAOSOM>2.0.CO;2

Fig. 12.
Fig. 12.

Vertical shear of the low-level horizontal winds and Cimarron radar reflectivity factor (dBZ) at 1638 CST at 0.5 km AGL; x and y as in Fig. 3. (a) Vertical shear in the 500–1000-m AGL layer in the dual-Doppler analysis winds at 1638 CST. (b) Time-to-space conversion of vertical shear between 7 and 444 m AGL measured by the instrumented tower over a 30-min period centered at 1638 CST.

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2562:TAOSOM>2.0.CO;2

Fig. 13.
Fig. 13.

Vertical velocities measured by the instrumented tower at 444 m AGL (heavy line) and synthesized in the dual-Doppler analysis at 0.5 km AGL at 1638 CST (contours, m s−1); x and y as in Fig. 3. Contours for negative values are dashed. The time-to-space conversion of tower data covers the period within 15 min of the center time (1638 CST). The heavy straight line indicates the direction of storm motion. The displacement of the curve above (updraft) or below (downdraft) this line indicates the vertical velocity. The maximum and minimum vertical velocities are 6.1 m s−1 and −6.8 m s−1, respectively.

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2562:TAOSOM>2.0.CO;2

Fig. 14.
Fig. 14.

Vertical cross sections constructed from a time-to-space conversion of instrumented tower data between 1635 and 1645 CST. Equivalent distance for storm motion of 11.7 m s−1 is indicated. (a) Equivalent potential temperature (K) and storm-relative winds. (b) Virtual potential temperature (K) and storm-relative winds. (c) Storm-relative winds and contours of the component of wind normal to the cross section (m s−1). Solid (dashed) contours indicate winds into (out of) the page.

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2562:TAOSOM>2.0.CO;2

Fig. 14.
Fig. 15.
Fig. 15.

Vertical profile of equivalent potential temperature (K) in the sounding taken from Edmond, Oklahoma, at 1430 CST (cf. Fig. 2).

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2562:TAOSOM>2.0.CO;2

Fig. 16.
Fig. 16.

Demonstration of how horizontal vorticity associated with vertical shear can be concentrated into a vortex at low levels. The analysis is carried out in terms of a material circuit in the fluid; the process involves tilting of the circuit into the horizontal by a downdraft, followed by convergence of the circuit as it is passed into an updraft. The effects of deformation acting on the material curve are neglected for ease of interpretation. The environmental wind profile is indicated at the left.

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2562:TAOSOM>2.0.CO;2

Fig. 17.
Fig. 17.

Material curves traced backward (using storm-relative trajectories) from a ring with a 2.5-km diameter centered on the mesocyclone at 0.5 km AGL; x and y as in Fig. 3. Velocities are computed by linear interpolations (in time) between the dual-Doppler volume at the time given in the figure and the volume at a previous time. Heights (km AGL) of points along the curve are shown. Contours of vertical vorticity (10−2 s−1) are also shown: (a) followed back 400 s from 1638 CST, and (b) followed back 400 s from 1704 CST.

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2562:TAOSOM>2.0.CO;2

Fig. 18.
Fig. 18.

Hodograph for low-level winds measured by the instrumented tower at 1530 (about 50 km ahead of the updraft) and 1634 CST (2 km ahead of the updraft). Each circle represents 5 m s−1 of wind speed.

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2562:TAOSOM>2.0.CO;2

Fig. 19.
Fig. 19.

Schematic of features at 1.0 km AGL. Downdrafts (w ≤ −1 m s−1) are shaded lightly, updrafts (w ≥ 3 m s−1) are shaded heavily, and mesocyclones (ζ ≥ 0.6 × 10−2 s−1) are denoted by solid curves. The TVS location at 1704 CST is marked.

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2562:TAOSOM>2.0.CO;2

Fig. 20.
Fig. 20.

Perturbation pressure (mb) retrieved from dual-Doppler analyses at 1704 and 1710 CST; x and y as in Fig. 3. Contours for negative values are dashed: (a) vertical cross section along x = 12 km [see (b)], and (b) horizontal cross section at 2.5 km AGL.

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2562:TAOSOM>2.0.CO;2

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