Mobile Mesonet Observations on 3 May 1999

Paul M. Markowski Department of Meteorology, The Pennsylvania State University, University Park, Pennsylvania

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Abstract

Two long-lived tornadic supercells were sampled by an automobile-borne observing system on 3 May 1999. The “mobile mesonet” observed relatively warm and moist air, weak baroclinity, and small pressure excess at the surface within the rear-flank downdrafts of the storms. Furthermore, the downdraft air parcels, which have been shown to enter the tornado in past observational and modeling studies, were associated with substantial convective available potential energy and small convective inhibition. The detection of only small equivalent potential temperature deficits (1–4 K) within the downdrafts may imply that the downdrafts were driven primarily by nonhydrostatic pressure gradients and/or precipitation drag, rather than by the entrainment of potentially cold environmental air at midlevels.

Corresponding author address: Dr. Paul Markowski, The Pennsylvania State University, 503 Walker Building, University Park, PA 16802. Email: marko@mail.meteo.psu.edu

Abstract

Two long-lived tornadic supercells were sampled by an automobile-borne observing system on 3 May 1999. The “mobile mesonet” observed relatively warm and moist air, weak baroclinity, and small pressure excess at the surface within the rear-flank downdrafts of the storms. Furthermore, the downdraft air parcels, which have been shown to enter the tornado in past observational and modeling studies, were associated with substantial convective available potential energy and small convective inhibition. The detection of only small equivalent potential temperature deficits (1–4 K) within the downdrafts may imply that the downdrafts were driven primarily by nonhydrostatic pressure gradients and/or precipitation drag, rather than by the entrainment of potentially cold environmental air at midlevels.

Corresponding author address: Dr. Paul Markowski, The Pennsylvania State University, 503 Walker Building, University Park, PA 16802. Email: marko@mail.meteo.psu.edu

1. Introduction

Direct measurements of meteorological variables near tornadoes have been relatively scarce, owing to the rarity of the phenomenon and the spatial and temporal frequency of standard observations. A few fortuitous datasets have been analyzed by Tepper and Eggert (1956) and Fujita (1958). Storm intercept field programs, first organized in the 1970s, have obtained additional observations within supercell storms (e.g., Golden and Morgan 1972; Bluestein 1983; Davies-Jones 1986); however, our collection of in situ measurements within supercells remains miniscule.

A fleet of instrumented automobiles was designed by Straka et al. (1996) for use in the Verification of the Origins of Rotation in Tornadoes Experiment (VORTEX; Rasmussen et al. 1994), conducted on the U.S. Great Plains during the springs of 1994 and 1995. Coined the “mobile mesonet,” this platform collected surface data within severe storms with unprecedented spatial (100–1000 m) and temporal (10–60 s) resolution. The mobile mesonet also has been used in subsequent years following VORTEX, with the spring of 1999 being exceptionally fruitful for operations. This paper summarizes a few noteworthy observations made on 3 May 1999, during a significant outbreak of tornadoes in Oklahoma and Kansas.

During the evening of 3 May, the mobile mesonet intercepted a pair of tornadic supercells in southwestern and central Oklahoma (Fig. 1). The first supercell was intercepted approximately 25 km north of Fort Sill, Oklahoma, shortly after 2200 UTC (this storm will be referred to as storm A to be consistent with National Weather Service surveys of the event). A second supercell was intercepted farther west around 0000 UTC (this storm will be referred to as storm B). Detailed accounts of these intercepts appear in section 3, following a description of the analysis techniques in section 2. A few final comments are provided in section 4.

2. Mobile mesonet data

The mobile mesonet senses temperature, relative humidity, pressure, and wind velocity. Time and position are recorded using a global positioning system receiver. In 1999, data were recorded at 2-s intervals. A complete list of instrument specifications (including response times and errors) is available from Straka et al. (1996).

Vehicle velocities were removed from the wind velocity data when vehicle accelerations were small. In the presence of significant vehicle accelerations (>1 m s–2), accurate wind data could not be obtained. Furthermore, field operations relied on the use of radio communication. On occasion, radio frequency interference caused large errors to arise in the meteorological measurements. Data with these gross errors were removed prior to analysis. In addition, biases in the data were removed by way of intercomparisons between vehicles. The intercomparisons were made in relatively quiescent weather conditions while the vehicles were moving as a caravan (e.g., en route to a storm).

The quality-controlled observations used in the analyses were averaged over 12-s intervals. Data were plotted relative to radar echoes using time-to-space conversion, assuming that the features being analyzed did not change significantly over the time interval during which the measurements were made (the “Taylor hypothesis”). Supercells are not steady. If they were, then tornadogenesis could not occur. However, steadiness assumptions, at least for short time intervals, are virtually unavoidable in observational studies (e.g., multiple Doppler radar analysis). For the research herein, features were assumed to be steady for ±3 min, with respect to an analysis time. This is approximately the length of time that it takes for a Weather Surveillance Radar-1998 Doppler (WSR-88D) to complete a volume scan. There is some confidence that such steadiness assumptions were not too severe, for the analyzed fields tended to be free of noise [inappropriate steadiness assumptions lead to the artificial creation of warm and cold (or moist and dry) pockets in the analysis following a time-to-space conversion].

In addition to the raw thermodynamic data recorded by the mobile mesonet, several derived variables were computed. Virtual potential temperature was computed, with the inclusion of liquid water effects, in addition to the contribution from water vapor. The liquid water content was parameterized from the WSR-88D reflectivity at the lowest elevation angle, using the method of Rutledge and Hobbs (1984). Equivalent potential temperature θe was computed by lifting a surface parcel adiabatically to 100 hPa, where the potential temperature of the parcel was assumed to be equal to the θe of the parcel. Pressure p was reduced to the average height of the vehicle observations (407 m) using U.S. Geological Survey Level-2 Digital Elevation Model data. Convective available potential energy (CAPE) and convective inhibition (CIN) were computed for parcels by inserting surface thermodynamic measurements obtained from the mobile mesonet into the 0000 UTC 4 May sounding at Norman, Oklahoma (located approximately 25 km south of Oklahoma City; see Fig. 2).1 The buoyancy integrations in the CAPE calculations were terminated at 500 hPa, owing to the loss of the sounding data at upper levels. Virtual temperature effects were not included in the CAPE and CIN calculations. Whereas (more precisely, its fluctuation) can be considered a measure of parcel buoyancy, CAPE and CIN values can be viewed as measures of the potential buoyancy of a parcel.

The fluctuations of , θe, and p (denoted by a prime, e.g., ) from a base state (denoted by an overbar, e.g., , where = ) were analyzed instead of the absolute quantities themselves. The base state was defined at the storm location by interpolating from smooth subjectively analyzed contours obtained from Oklahoma Mesonet stations (Brock et al. 1995; Fig. 3). All definitions of a “base state” are arbitrary (and potentially problematic); the method used herein is similar to that used by Fujita (1955) and Charba and Sasaki (1971).

The uncertainty of perturbation variables (denoted by δ) depends on the propagation of instrument errors (e.g., temperature, relative humidity, and pressure errors), uncertainties associated with the estimation of liquid water in the case of , and specification of the base state. A detailed analysis of these effects appears in Markowski (2000). On 3 May 1999, it has been estimated that δp′ ≈ 0.8 hPa, δ ≈ 0.5 K, δθe ≈ 2.5 K, δCAPE ≈ 110 J kg–1, and δCIN ≈ 9 J kg–1.

The mobile mesonet comprised three vehicles on 3 May 1999, in contrast to the much larger assembly used for VORTEX. The goal of the field operations of 1999 was to sample the rear-flank downdraft (RFD) region of supercells (Lemon and Doswell 1979), rather than attempt to sample a broad inflow region, as was the case during VORTEX (Rasmussen et al. 1994). Thus, a smaller array of vehicleborne instruments could accomplish the sampling goal. Furthermore, a smaller fleet had some logistical advantages; for example, coordination among mobile mesonet crews and with the nowcaster was considerably easier.

The mobile mesonet RFD sampling strategy was favored for several reasons. In environments initially absent of vertical vorticity (i.e., vortex lines are quasi-horizontal), a downdraft is necessary for intense vertical vorticity to arise at the ground (Davies-Jones 1982; Davies-Jones and Brooks 1993; Walko 1993). Even as early as 1975, Fujita (1975) hypothesized that the angular momentum transport by a downdraft may be critical to tornadogenesis. Burgess et al. (1977) and Barnes (1978) also made similar speculations. Many studies have found that the air parcels that supply the tornado pass through the RFD. For example, observations by Brandes (1978), Lemon and Doswell (1979), Rasmussen et al. (1982), and Jensen et al. (1983) have shown or implied a near total occlusion of the low-level mesocyclone by the RFD prior to tornadogenesis. Furthermore, Brooks et al. (1993), Wicker and Wilhelmson (1995), and Adlerman et al. (1999) have found that the trajectories entering the near-ground circulations in their numerical simulations passed through the hook echo and RFD. Given the prior emphasis on the RFD in the tornadogenesis process and the apparent consensus that RFD air parcels enter the tornado, it is believed that the thermodynamic characteristics of hook echoes and RFDs naturally assume importance. For this reason, the focus of the next section is on the observations obtained largely within the RFDs of the intercepted tornadic supercells.

3. Observations

a. Storm A

The mobile mesonet approached storm A around 2200 UTC. The supercell was associated with two brief, weak tornadoes at 2151 and 2155 UTC, before an intercept could be engaged. At 2220 UTC, the supercell produced its first significant tornado.2 The tornado persisted for 15 min, although it did not do significant damage along its 10-km track (F1 Fujita-scale rating). The RFD associated with the circulation was sampled by the mobile mesonet west-southwest through east-southeast just prior to tornadogenesis, and during the lifetime of the tornado all three vehicles trailed the tornado to the southwest, within the RFD.

At 2219 UTC, 1 min before tornadogenesis, the mobile mesonet detected relatively small deficits within the RFD and inflow of less than 2 K (Fig. 4). The RFD parcel temperatures were 1.0–1.5 K warmer than the inflow parcels. The lowest θe values, located within the hook echo, were approximately 4 K less than the average inflow values (Fig. 5); however, these θe values (∼345 K) were larger than any θe values observed on the 0000 UTC Norman sounding (Fig. 2), because of the sounding being launched east of the axis of maximum θe values (Fig. 3). Therefore, estimates cannot be made of the altitude from which RFD parcels may have descended (nor could estimates be made even if θe could be assumed to be conserved in the absence of entrainment). The relatively small and θe deficits (and temperature excess) of the downdraft air parcels also were associated with small CIN and substantial CAPE (Fig. 6). CIN values as small as 2 J kg–1 were detected in the hook echo, and CAPE values in the RFD generally exceeded 900 J kg–1 in the lowest 500 hPa (it is estimated that the total CAPE associated with the RFD air parcels was ∼3000 J kg–1).

The mobile mesonet also sampled a small pressure excess (<1 hPa) in the RFD at 2219 UTC, except within a few hundred meters of the intensifying circulation center, where a small pressure deficit (about −1 hPa) was detected (Fig. 7). A small pressure deficit also was sampled in the inflow east of the gust front. Streamlines revealed difluence at the surface within the RFD, with the strongest difluence (and divergence) apparently situated in the hook echo (Fig. 4). Anticyclonic vertical vorticity also was present within the RFD to the south and east of the incipient tornado.

The mobile mesonet was unable to sample the RFD within 1 km of the tornado at later times; however, data were collected within the RFD approximately 2–5 km southwest of the tornado until its demise. At 2229 UTC (6 min prior to tornado dissipation), only small deficits were detected again in the RFD; however, larger θe deficits were sampled (and correspondingly smaller, yet significant, CAPE values of ∼200–400 J kg–1 below 500 hPa), with deficits as large as 7–8 K recorded a few kilometers west-southwest of the tornado (Fig. 8). By 2234 UTC (1 min prior to tornado dissipation), the and θe deficits within the RFD appeared to increase slightly, with deficits exceeding 2 K and θe deficits as large as 10 K being detected within 3 km of the tornado to its west and southwest (Fig. 9).

The relatively long-lived tornado dissipated at 2235 UTC, and another tornado was produced by storm A at 2246 UTC. The tornado produced F3 damage and lasted until 2310 UTC, but the tornado could not be closely intercepted for logistical reasons. The road network did not allow sampling within several kilometers of the tornado, and many roads became obstructed by debris. Storm A went on to produce another significant tornado near 2323 UTC, which produced F5 damage in the suburbs of Oklahoma City. This tornado also was not intercepted for logistical reasons (e.g., increasing amounts of traffic near the metropolitan area, debris-filled roadways). Instead, the mobile mesonet abandoned storm A in pursuit of another tornadic supercell to the west (storm B).

b. Storm B

The mobile mesonet arrived at storm B at approximately 0000 UTC (4 May). Storm B produced several brief, weak tornadoes from 2236 to 2324 UTC. Another tornado was reportedly on the ground for 21 min from 2338 to 2359 UTC, but this tornado dissipated before data could be collected. From 0000 to 0100 UTC, numerous additional tornadoes (as many as eight) were reported, most of which were brief and weak (F0–F1 damage ratings), although two tornadoes persisted for longer than 5 min. Data were collected within a few kilometers of the surface circulation centers at several times from 0000 to 0100 UTC.

At 0026 UTC, storm B was not producing a tornado, but brief tornadoes were reported just before and after the analysis time (at 0020 and 0034 UTC). The mobile mesonet obtained data within 0.5–2.0 km of the mesocyclone center at the surface. Within 3 km of the circulation center, the most negatively buoyant air sampled within the hook echo and RFD region was associated with a deficit of less than 1 K (Fig. 10). Furthermore, the baroclinity at the surface within the hook echo was very weak, with maximum |∇h| values of less than 0.5 K km–1 (where ∇h is the horizontal gradient operator). The relatively warm air within the downdraft also was associated with small θe deficits (Fig. 11). Within the hook echo, θe values were nearly the same as in the inflow, and θe values were only 1–2 K less than inflow values in a zone that appeared to wrap around the north side of the circulation center. As was the case for storm A, all of the surface θe values measured by the mobile mesonet in storm B were larger than the maximum θe value observed on the 0000 UTC Norman sounding. It also can be inferred that the CAPE (CIN) within the RFD of storm B at 0026 UTC was large (small).

The mobile mesonet made additional penetrations of the hook echo and RFD near 0043 UTC. A tornado, which formed at 0037 UTC, was observed at the analysis time. The tornado dissipated at approximately 0048 UTC. Again, only small and θe deficits were detected within the RFD and hook echo (Fig. 12), with substantial CAPE and small CIN also being present within the downdraft at the surface (not shown). Baroclinity within the hook echo also was weak (maximum |∇h| of ∼1 K km–1).3

At 0052 UTC, the RFD and hook echo region of storm B were again sampled relatively well, and another tornado was in progress at the time (the tornado formed at 0047 UTC and dissipated at 0100 UTC). As at earlier times, only small and θe deficits (<2 K) were observed, and baroclinity was weak (Fig. 13; maximum |∇h| of ∼1 K km–1 detected north of the tornado). Streamlines diverged within the hook echo, and a couplet of vertical vorticity appeared to straddle the hook echo. Anticyclonic vertical vorticity, estimated to be approximately −5 × 10–3 s–1, was observed at the surface on the side of the hook echo opposite the stronger, cyclonic vertical vorticity associated with the tornado. CIN values were very insignificant (<10 J kg–1 within 5 km of the tornado in the RFD), and RFD parcels also were associated with large CAPE (>800 J kg–1 below 500 hPa; Fig. 14).

Surface pressure gradients were weak at 0052 UTC at distances of more than 500 m from the tornado, similar to observations in storm A (Fig. 15). A small pressure excess (<1 hPa) was detected within the RFD, and a small pressure deficit (also <1 hPa) was measured within the region of cyclonic vertical vorticity associated with the tornado parent circulation. The relatively weak pressure gradients are somewhat consistent with observations made by the mobile mesonet crews of “light winds” within a few kilometers of the tornadoes, both in storms A and B (J. Straka 1999, personal communication).

Storm B produced at least six additional tornadoes, some strong, after dark, which arrived between 0100 and 0130 UTC. Data collection operations by the mobile mesonet were terminated near 0130 UTC. The supercell dissipated in northern Oklahoma around 0400 UTC.

4. Discussion and closing remarks

The observational data presented herein have several limitations. First, road networks do not allow continuous sampling of moving updrafts for periods longer than about 5 min before repositioning of the vehicles requires that they temporarily forfeit data collection in critical regions of the storm; therefore, the time evolution of features is difficult to document. Furthermore, steadiness reluctantly was assumed for durations as long as 6 min (±3 min from the analysis reference times) in constructing the analyses, to maximize the coverage of data (gathered by a finite number of vehicles). Second, thermodynamic fields and their gradients cannot be ascertained above the surface by direct means. At best, only the sign of the gradients can be inferred above the surface, based on assumptions of the lapse rates beneath and at a distance from the storm. Third, time histories of air parcels are important, possibly more than 30 min prior to tornadogenesis. It was not possible to compute trajectories at the surface because of inadequate observation density.

Despite the above impediments, substantial evidence was presented of the following characteristics of the 3 May 1999 tornadic supercells:

  1. RFDs were associated with small and θe deficits,

  2. RFDs contained substantial amounts of CAPE and small amounts of CIN,

  3. hook echoes lacked strong baroclinity at the surface, and

  4. pressure fluctuations and gradients were small at distances greater than approximately 250 m from the circulation centers.

It is perhaps most intriguing that the RFDs were associated with what one might consider to be “small” signatures, both kinematically and thermodynamically.

The relatively small and θe deficits within the RFDs may imply a descent largely forced by nonhydrostatic pressure gradients, rather than the entrainment of low-θe environmental air at midlevels, and subsequent generation of negative buoyancy by evaporative chilling, as in long-standing conceptual models (e.g., Browning and Ludlam 1962; Browning and Donaldson 1963). Visual observations by mobile mesonet volunteers (e.g., J. Straka 1999, personal communication) suggested that the hook echoes were composed of only a thin veil of mainly large raindrops; thus, precipitation loading may only have been of secondary importance in the downdraft forcing. Furthermore, it is worth reiterating that the observations of RFD thermodynamic properties similar to the properties representative of the inflow do not necessarily imply small vertical downward excursions for RFD air parcels. The depth of descent cannot be ascertained from the surface θe measurements alone.

Some evidence also was presented of vertical vorticity couplets straddling the divergence maxima within the RFDs (roughly collocated with the hook echo), as was presented by numerous other studies at low levels (e.g., Ray et al. 1975, 1981; Brandes 1977, 1978, 1981; Fujita and Wakimoto 1982; Wurman et al. 1996). This probably is evidence that RFDs are involved in a downward displacement of vortex lines, perhaps necessarily supplying angular momentum to the tornado, as many others have previously conjectured (section 1).

It is speculated that some of the above intriguing features of the 3 May 1999 tornadic supercells, particularly the relative warmth and moistness of the downdrafts, may be relevant to the problem of tornadogenesis. Does tornadogenesis probability, longevity, and intensity increase as RFD parcels become more buoyant? If buoyant RFDs are propitious for tornadogenesis, are there any large-scale environmental conditions from which we may anticipate warm, moist RFDs? It is generally accepted that only a relatively small percentage of supercells are tornadic; yet, on 3 May 1999, nearly all were tornadic. This fact may suggest that large-scale factors may exist, at least in some cases, from which favorable RFD thermodynamic characteristics can be inferred. Additional cases and implications and, it is hoped, some answers to the above-posed questions will be the subject of a series of companion papers.

Acknowledgments

I am most grateful to all of the volunteers whose countless personal sacrifices made data collection possible. I am indebted to Drs. Jerry Straka and Erik Rasmussen for their support during field operations and for stimulating discussions related to supercells and tornadogenesis during the last several years. Data-quality checks were performed by Mr. Al Pietrycha, and indispensable assistance in obtaining digital elevation model data was provided by Mr. Rob Carver. Dr. Joshua Wurman provided data from the Doppler on Wheels mobile radars, which were used to help to identify the position of the gust front in storm A. Mr. Mark Shafer provided Oklahoma Mesonet data. I also thank Dr. Chuck Doswell and two anonymous reviewers for helping to improve the manuscript. NSF Grant ATM-9617318 partially supported this research.

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Fig. 1.
Fig. 1.

The tracks of storms A and B, with the regions indicated where mobile mesonet data were collected. The radar echoes outline the 30-dBZ base reflectivity values at the 0.5° elevation angle of the Twin Lakes, OK, KTLX WSR-88D; times (UTC) also are included for each radar echo. Tornado paths are gray (storm A) and black (storm B). The surface stations SWO, CSM, OKC, and FSI are Stillwater, Clinton, Oklahoma City, and Fort Sill, respectively

Citation: Weather and Forecasting 17, 3; 10.1175/1520-0434(2002)017<0430:MMOOM>2.0.CO;2

Fig. 2.
Fig. 2.

(left) Skew T–logp diagram from Norman, OK, at 0000 UTC 4 May 1999. The hodograph is shown in the inset, with 10 m s–1 speed rings and numerals along the hodograph indicating heights AGL (km). (right) Vertical profile of θe derived from the Norman sounding.

Citation: Weather and Forecasting 17, 3; 10.1175/1520-0434(2002)017<0430:MMOOM>2.0.CO;2

Fig. 3.
Fig. 3.

Meso-α-scale hourly subjective analyses of (K), θe (K), and pressure [hPa; reduced to the mean elevation of the mobile mesonet observations (407 m MSL)] from 2200 to 0100 UTC 3–4 May. Data were obtained from the Oklahoma Mesonet. Station models display (reading anticlockwise, beginning with the top-left value) temperature (°C), dewpoint (°C), (K), θe (K), and reduced pressure (hPa). Wind barbs are relative to the ground, with each full (half) barb being equal to 5 (2.5) m s–1

Citation: Weather and Forecasting 17, 3; 10.1175/1520-0434(2002)017<0430:MMOOM>2.0.CO;2

Fig. 3.
Fig. 4.
Fig. 4.

Subjective analysis of (K) in storm A at 2219 UTC 3 May 1999 ( = 304.4 K). Contours are dashed where significant uncertainty exists owing to sparseness of observations. The analysis time is 1 min prior to tornadogenesis. Mobile mesonet station models include (reading anticlockwise, beginning with the numeral at the top left) temperature (in degrees Celsius to the nearest 0.1°C with the decimal omitted), dewpoint temperature (in degrees Celsius to the nearest 0.1°C with the decimal omitted), (in kelvins to the nearest 0.1 K with the decimal omitted), and θe (in kelvins to the nearest 1 K). Wind barbs depict storm-relative winds [each full (half) barb equals 5 (2.5) m s–1]. A few streamlines have been drawn in gray. Mobile mesonet observations have been averaged over 12-s intervals, and a steadiness was assumed for a period of ±3 min (with respect to the analysis time) in the time-to-space conversion. Observations obtained more than 1 min before or after the analysis reference time are “flagged” with a vertical bar through the center of the station model. Storm-scale fronts are depicted using conventional frontal symbology, and their placement has been aided by inspection of mobile radar data provided by J. Wurman. The M indicates the position of the mesocyclone center on the lowest radar elevation angle available. Base radar reflectivity data were obtained from the 0.5° elevation angle of the KTLX WSR-88D, with the reflectivity scale (dBZ) included at the bottom. The region analyzed is indicated in the larger-scale inset

Citation: Weather and Forecasting 17, 3; 10.1175/1520-0434(2002)017<0430:MMOOM>2.0.CO;2

Fig. 5.
Fig. 5.

As in Fig. 4 but θe (K) is analyzed (θe = 350.0 K).

Citation: Weather and Forecasting 17, 3; 10.1175/1520-0434(2002)017<0430:MMOOM>2.0.CO;2

Fig. 6.
Fig. 6.

As in Fig. 4 but CIN (J kg–1) is analyzed. CAPE (below 500 hPa) and CIN values appear to the left and right, respectively, in each station model

Citation: Weather and Forecasting 17, 3; 10.1175/1520-0434(2002)017<0430:MMOOM>2.0.CO;2

Fig. 7.
Fig. 7.

As in Fig. 4 but p′ (hPa) is analyzed (p = 953.1 hPa). Reduced pressure values, with the leading 9 omitted, appear in each station model. Values are to the nearest 0.1 hPa, with the decimal omitted

Citation: Weather and Forecasting 17, 3; 10.1175/1520-0434(2002)017<0430:MMOOM>2.0.CO;2

Fig. 8.
Fig. 8.

As in Fig. 4 but at 2229 UTC 3 May 1999. The analysis time is 9 min after tornadogenesis and 6 min prior to tornado dissipation. The T indicates the tornado location

Citation: Weather and Forecasting 17, 3; 10.1175/1520-0434(2002)017<0430:MMOOM>2.0.CO;2

Fig. 9.
Fig. 9.

As in Fig. 4 but at 2234 UTC 3 May 1999. The analysis time is 14 min after tornadogenesis and 1 min prior to tornado dissipation

Citation: Weather and Forecasting 17, 3; 10.1175/1520-0434(2002)017<0430:MMOOM>2.0.CO;2

Fig. 10.
Fig. 10.

As in Fig. 4 but storm B is analyzed at 0026 UTC 4 May 1999 ( = 304.0 K). No tornado was in progress at the analysis time, but tornadoes were observed shortly before and after 0026 UTC (at 0020 and 0034 UTC)

Citation: Weather and Forecasting 17, 3; 10.1175/1520-0434(2002)017<0430:MMOOM>2.0.CO;2

Fig. 11.
Fig. 11.

As in Fig. 10 but θe (K) is analyzed (θe = 349.0 K)

Citation: Weather and Forecasting 17, 3; 10.1175/1520-0434(2002)017<0430:MMOOM>2.0.CO;2

Fig. 12.
Fig. 12.

As in Fig. 10 but at 0043 UTC 4 May 1999. The analysis time is 6 min after tornadogenesis and 5 min prior to tornado dissipation

Citation: Weather and Forecasting 17, 3; 10.1175/1520-0434(2002)017<0430:MMOOM>2.0.CO;2

Fig. 13.
Fig. 13.

As in Fig. 10 but at 0052 UTC 4 May 1999. The analysis time is 5 min after tornadogenesis and 8 min prior to tornado dissipation (this is a different tornado than that which appears in Fig. 12). A few streamlines also have been drawn

Citation: Weather and Forecasting 17, 3; 10.1175/1520-0434(2002)017<0430:MMOOM>2.0.CO;2

Fig. 14.
Fig. 14.

As in Fig. 13 but CIN (J kg–1) is analyzed. Station models are as in Fig. 6

Citation: Weather and Forecasting 17, 3; 10.1175/1520-0434(2002)017<0430:MMOOM>2.0.CO;2

Fig. 15.
Fig. 15.

As in Fig. 13 but p′ (hPa) is analyzed (p = 951.5 hPa). Station models are as in Fig. 7.

Citation: Weather and Forecasting 17, 3; 10.1175/1520-0434(2002)017<0430:MMOOM>2.0.CO;2

1

It always is possible to debate the representativeness of a sounding. The 0000 UTC Norman sounding was obtained during the lifetimes of the tornadic supercells and from within 75–100 km of the storms that were intercepted; however, it will be evident later that the environment sampled by the Norman sounding, at least near the surface, differed from the environment to the west in which the supercells were occurring (Fig. 3).

2

The tornado is referred to as significant because of its longevity.

3

The adjective “weak” is used to describe the baroclinity because a value of |∇h| ∼ 1 K is considered to be small on the storm scale. However, by some standards (e.g., on the synoptic scale), a gradient of this magnitude would be considered to be enormous.

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  • Fig. 1.

    The tracks of storms A and B, with the regions indicated where mobile mesonet data were collected. The radar echoes outline the 30-dBZ base reflectivity values at the 0.5° elevation angle of the Twin Lakes, OK, KTLX WSR-88D; times (UTC) also are included for each radar echo. Tornado paths are gray (storm A) and black (storm B). The surface stations SWO, CSM, OKC, and FSI are Stillwater, Clinton, Oklahoma City, and Fort Sill, respectively

  • Fig. 2.

    (left) Skew T–logp diagram from Norman, OK, at 0000 UTC 4 May 1999. The hodograph is shown in the inset, with 10 m s–1 speed rings and numerals along the hodograph indicating heights AGL (km). (right) Vertical profile of θe derived from the Norman sounding.

  • Fig. 3.

    Meso-α-scale hourly subjective analyses of (K), θe (K), and pressure [hPa; reduced to the mean elevation of the mobile mesonet observations (407 m MSL)] from 2200 to 0100 UTC 3–4 May. Data were obtained from the Oklahoma Mesonet. Station models display (reading anticlockwise, beginning with the top-left value) temperature (°C), dewpoint (°C), (K), θe (K), and reduced pressure (hPa). Wind barbs are relative to the ground, with each full (half) barb being equal to 5 (2.5) m s–1

  • Fig. 3.

    (Continued)

  • Fig. 4.

    Subjective analysis of (K) in storm A at 2219 UTC 3 May 1999 ( = 304.4 K). Contours are dashed where significant uncertainty exists owing to sparseness of observations. The analysis time is 1 min prior to tornadogenesis. Mobile mesonet station models include (reading anticlockwise, beginning with the numeral at the top left) temperature (in degrees Celsius to the nearest 0.1°C with the decimal omitted), dewpoint temperature (in degrees Celsius to the nearest 0.1°C with the decimal omitted), (in kelvins to the nearest 0.1 K with the decimal omitted), and θe (in kelvins to the nearest 1 K). Wind barbs depict storm-relative winds [each full (half) barb equals 5 (2.5) m s–1]. A few streamlines have been drawn in gray. Mobile mesonet observations have been averaged over 12-s intervals, and a steadiness was assumed for a period of ±3 min (with respect to the analysis time) in the time-to-space conversion. Observations obtained more than 1 min before or after the analysis reference time are “flagged” with a vertical bar through the center of the station model. Storm-scale fronts are depicted using conventional frontal symbology, and their placement has been aided by inspection of mobile radar data provided by J. Wurman. The M indicates the position of the mesocyclone center on the lowest radar elevation angle available. Base radar reflectivity data were obtained from the 0.5° elevation angle of the KTLX WSR-88D, with the reflectivity scale (dBZ) included at the bottom. The region analyzed is indicated in the larger-scale inset

  • Fig. 5.

    As in Fig. 4 but θe (K) is analyzed (θe = 350.0 K).

  • Fig. 6.

    As in Fig. 4 but CIN (J kg–1) is analyzed. CAPE (below 500 hPa) and CIN values appear to the left and right, respectively, in each station model

  • Fig. 7.

    As in Fig. 4 but p′ (hPa) is analyzed (p = 953.1 hPa). Reduced pressure values, with the leading 9 omitted, appear in each station model. Values are to the nearest 0.1 hPa, with the decimal omitted

  • Fig. 8.

    As in Fig. 4 but at 2229 UTC 3 May 1999. The analysis time is 9 min after tornadogenesis and 6 min prior to tornado dissipation. The T indicates the tornado location

  • Fig. 9.

    As in Fig. 4 but at 2234 UTC 3 May 1999. The analysis time is 14 min after tornadogenesis and 1 min prior to tornado dissipation

  • Fig. 10.

    As in Fig. 4 but storm B is analyzed at 0026 UTC 4 May 1999 ( = 304.0 K). No tornado was in progress at the analysis time, but tornadoes were observed shortly before and after 0026 UTC (at 0020 and 0034 UTC)

  • Fig. 11.

    As in Fig. 10 but θe (K) is analyzed (θe = 349.0 K)

  • Fig. 12.

    As in Fig. 10 but at 0043 UTC 4 May 1999. The analysis time is 6 min after tornadogenesis and 5 min prior to tornado dissipation

  • Fig. 13.

    As in Fig. 10 but at 0052 UTC 4 May 1999. The analysis time is 5 min after tornadogenesis and 8 min prior to tornado dissipation (this is a different tornado than that which appears in Fig. 12). A few streamlines also have been drawn

  • Fig. 14.

    As in Fig. 13 but CIN (J kg–1) is analyzed. Station models are as in Fig. 6

  • Fig. 15.

    As in Fig. 13 but p′ (hPa) is analyzed (p = 951.5 hPa). Station models are as in Fig. 7.

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