Some Observations of Rotating Updrafts in a Low-Buoyancy, Highly Sheared Environment

Paul M. Markowski School of Meteorology, University of Oklahoma, Norman, Oklahoma

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Jerry M. Straka School of Meteorology, University of Oklahoma, Norman, Oklahoma

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Abstract

The authors document some of the unusual rotating updrafts (one of which produced a tornado) that developed over central Oklahoma on 28 October 1998 in an environment of strong (1.8 × 10−2 s−1) low-level (0–3 km) mean shear. The maximum convective available potential energy (including virtual temperature effects) a “storm” could have realized was approximately 300 J kg−1; however, most of the storms probably realized less than 100 J kg−1. Average (maximum) parcel virtual temperature excesses were estimated to be 0.4–1.2 K (1.8–2.8 K). Echo tops were measured from less than 5 km to 11.2 km above ground level (AGL), although visual observations and radar data suggested echoes that extended above approximately 5–6 km AGL were not associated with significantly buoyant cloud elements. Radar characteristics of many of the storms were similar to supercell storms (e.g., weak echo regions, echo overhang, velocity couplets, hook echoes), as were some of the visual characteristics near cloud base (e.g., wall clouds, rain-free bases, and striated low-level updrafts); however, visual characteristics in middle to upper portions of the storms were not characteristic of typical severe storms, supercells, or previously documented “minisupercells.” Furthermore, the buoyancy realized by the updrafts was estimated to be considerably less than environments associated with the aforementioned minisupercells.

* Additional affiliation: Center for the Analysis and Prediction of Storms, Norman, Oklahoma.

Corresponding author address: Paul Markowski, University of Oklahoma, Sarkeys Energy Center, 100 E. Boyd St., Room 1310, Norman, OK 73019.

Email: marko@rossby.ou.edu

Abstract

The authors document some of the unusual rotating updrafts (one of which produced a tornado) that developed over central Oklahoma on 28 October 1998 in an environment of strong (1.8 × 10−2 s−1) low-level (0–3 km) mean shear. The maximum convective available potential energy (including virtual temperature effects) a “storm” could have realized was approximately 300 J kg−1; however, most of the storms probably realized less than 100 J kg−1. Average (maximum) parcel virtual temperature excesses were estimated to be 0.4–1.2 K (1.8–2.8 K). Echo tops were measured from less than 5 km to 11.2 km above ground level (AGL), although visual observations and radar data suggested echoes that extended above approximately 5–6 km AGL were not associated with significantly buoyant cloud elements. Radar characteristics of many of the storms were similar to supercell storms (e.g., weak echo regions, echo overhang, velocity couplets, hook echoes), as were some of the visual characteristics near cloud base (e.g., wall clouds, rain-free bases, and striated low-level updrafts); however, visual characteristics in middle to upper portions of the storms were not characteristic of typical severe storms, supercells, or previously documented “minisupercells.” Furthermore, the buoyancy realized by the updrafts was estimated to be considerably less than environments associated with the aforementioned minisupercells.

* Additional affiliation: Center for the Analysis and Prediction of Storms, Norman, Oklahoma.

Corresponding author address: Paul Markowski, University of Oklahoma, Sarkeys Energy Center, 100 E. Boyd St., Room 1310, Norman, OK 73019.

Email: marko@rossby.ou.edu

1. Overview

During the evening of 28 October 1998, several shallow (echo tops generally below 6 km) convective clouds developed in central Oklahoma. A photograph from 2340 UTC (540 pm CST; Fig. 1a) shows three of these updrafts from distances of 16, 31, and 50 km, respectively, looking north-northeast. The echo tops of clouds 1, 2, and 3 [as determined from Weather Surveillance Radar-1988 Doppler (WSR-88D)] were 6.2, 11.2, and 6.0 km above ground level (AGL), respectively (hereafter, all heights are AGL); however, the cloud tops were estimated photogrammetrically to be at 6.1, 10.5, and 8.8 km, respectively.1 Both clouds 1 and 2 prompted tornado warnings. An earlier storm2 spawned a brief tornado at 2255 UTC (10 km north of Meridian, Oklahoma;3 the tornado was documented by a local television crew; Fig. 1c). Locally heavy precipitation (>2 cm) was recorded at several Oklahoma Mesonet (Brock et al. 1995) sites, owing to storms that moved repeatedly over some locations. A few intracloud lightning strikes were observed near and after sunset (M. Branick 1998, personal communication); however, no cloud-to-ground strikes were recorded by the National Lightning Detection Network during the event.

The moist convection was initiated in a region of moderately strong surface convergence [0.5–1.0 × 10−4 s−1 as analyzed by the Advanced Regional Prediction System (ARPS) Data Analysis System (ADAS; Brewster et al. 1994)] along a slowly moving dryline in central Oklahoma (Fig. 2). Aloft, a negatively tilted short-wave trough extended from western Nebraska to central Oklahoma at 500 mb at 0000 UTC 29 October. The convection did not produce elongated anvils; visible satellite imagery suggested a line of towering cumulus along the dryline that cast shadows near sunset (Fig. 3).4

We document this event because we believe these storms may represent a weaker variety (in terms of their size and inferred updraft strength) of storms that can be called supercells (Doswell and Burgess 1993). Despite their unimpressive visual appearance at middle and upper levels (with respect to storm depth; Fig. 1a), the storms displayed low-level visual and radar characteristics similar to larger supercells, including wall clouds (Fig. 1b; Fujita 1960), striations (Bluestein 1984; Bluestein and Parks 1983), clear slots (Fig. 1b; Beebe 1959;Moller et al. 1974; Peterson 1976; Burgess et al. 1977;Lemon and Doswell 1979), hook echoes (Stout and Huff 1953; Fujita 1958; Browning 1964; Lemon et al. 1978;Lemon and Doswell 1979; Ray et al. 1981; Forbes 1981), midlevel echo overhang (Marwitz 1972; Chisholm 1973), and weak echo regions (Browning and Donaldson 1963; Browning 1964, 1965; Burgess et al. 1977; Lemon 1977).

Browning (1964) was the first to use the term supercell. Weisman and Klemp (1984) later proposed that storms with a high degree of updraft and vertical vorticity correlation be called supercells. The definition now most widely accepted seems to be that a supercell is a storm that contains a “deep, persistent mesocyclone” (Johns and Doswell 1992; Doswell and Burgess 1993; Doswell 1996).

Relatively recently a variety of low-topped (tops <10 km) supercells, often referred to as “miniature” supercells, have been documented (e.g., McCaul 1991, 1993;Davies 1993; Kennedy et al. 1993; Stalker et al. 1993;Monteverdi and Quadros 1994; Knupp et al. 1998). These supercells are shallower than their taller (tops typically >12 km) counterparts that were documented originally (e.g., Browning and Donaldson 1963; Browning 1964, 1965, 1977), although some of the visual (Davies 1993; Stalker et al. 1993), radar (Kennedy et al. 1993; Stalker et al. 1993), and dynamical characteristics (McCaul 1993; Wicker and Cantrell 1996; McCaul and Weisman 1996) have been shown to be similar.

It will be shown that the storms documented in this paper probably should be called supercells because of the persistent (all lasted longer than 1 h), deep mesocyclones (in terms of fractional depth of the updrafts; all extended over >50% of the updraft depths) detected by Doppler radar. And because the storms were so shallow (most echo tops <6 km), strictly speaking, they were miniature. However, we will show that the environment of these minisupercells was considerably different from the minisupercell environments previously documented. It will be shown that the convective available potential energy (CAPE)5 realized by the storms probably was not greater than 300 J kg−1, with less than 100 J kg−1 likely realized by most of the updrafts—several times less than in all previously documented miniature supercell events.

This case serves as an example of how the radar appearance of a storm can sometimes be vastly different from what one might anticipate based on the visual appearance of a storm. While the exact low-level processes associated with tornadogenesis are not known, they apparently were present in the 28 October storms despite the lack of realized CAPE and shallow depth—the majority of “typical” supercells do not contain the low-level processes necessary for tornadogenesis [i.e., most supercells, possibly greater than 80% (D. Burgess 1997, personal communication), are nontornadic].

2. Visual and radar appearance

The storms lacked elongated anvils, although storms 2 and 3 (Fig. 1a) had flat tops. Careful analysis of WSR-88D data suggested that these tall (∼8–11 km), flat tops were short-lived (<30 min). These were not nascent anvils about to begin expanding downstream, nor did they represent a steady state of the convection. Echo tops were between 5–6 km for most of the lifetimes of the updrafts (as in storm 1 in Fig. 1a), with cloud parcels only occasionally attaining the higher altitudes of storms 2 and 3. The wispy, glaciated appearance of the upper halves (above approximately 5 km) of storms 2 and 3 suggested that the parcels in the upper portions of the clouds had little buoyancy.

Storm 1 and the lower portions of storm 2 (lower portion of storm 3 not visible in Fig. 1a) had better-defined, cumuliform cloud edges, suggestive of parcel buoyancy from approximately 1.5 km (estimated visually and from sounding data) to approximately 5 km (estimated from Figs. 1 and 4). The lowest portions of the updrafts were observed by the authors to be striated and laminar (no photograph available), suggestive of updraft rotation and forced ascent of negatively buoyant air parcels. Furthermore, wall clouds and clear slots were observed at the bases of storms 1 and 2.

Cloud-top rise rates from WSR-88D algorithm time–height output [analysis performed using the WSR-88D Algorithm Testing and Display System (WATADS 1998)] as well as scrutiny of plan position indicators (PPIs) showed maxima of cloud-top (0 dBZ when done manually with PPIs) rise rates typically <4–8 m s−1, with the largest about 16 m s−1. Examination of the WSR-88D PPI data and algorithm output also showed narrow (<10 km), often short-lived (<10 min) echo tops, such that cloud-top divergence calculations could not be made without serious errors. All of the data suggested that from a radar viewpoint, the vertical motions were significantly (50%–75%) less than the maximum that might be expected based on the realizable CAPE (<100 to 300 J kg−1; to be discussed in section 3). Furthermore, algorithm output and hand calculations using WSR-88D PPI data indicated that by the time parcels approached cloud top, little mass flux was present (little echo production by the anvil on radar using the 0 dBZ boundary or visually in photographs).

Radar data also revealed many structures that have been associated with supercell storms. A 2339 UTC radar image, obtained from approximately the time of the photograph in Fig. 1a, showed the presence of mesocyclones (all extending through half of the cloud depths or more) in storms 1, 2, and 3 (Fig. 4b). A small hook echo was associated with storm 1 and a strong reflectivity gradient was observed in storm 2 (Fig. 4a). The strongest horizontal shear was detected in storm 1 (8.6 × 10−3 s−1). Reflectivity cross sections (Figs. 4c and 4d) showed that the top of storm 1 was near 6 km;however, in storm 2, low (<30 dBZ) reflectivity extended to 11.2 km. This taller echo region did not appear to be associated with vigorous convection (Fig. 5). In the next section it will be shown that air parcels probably were not significantly buoyant in the region where this echo was found.

WSR-88D PPIs at a number of elevation angles are presented in Fig. 6 for 2339 UTC. The 0 dBZ echo was chosen as the threshold to include in the plots as a surrogate for visible cloud boundaries. Some studies have found the visible cloud edge to correspond to the reflectivity as large as 10 dBZ in 10-cm radar data, so our choice of the 0 dBZ threshold should be considered conservative (Knight and Miller 1993; Wakimoto and Bringi 1988). If indeed the visible cloud boundaries were reasonably well approximated by the 0 dBZ echo contours, then Fig. 6 reveals that most of the cloud tops were below 6 km AGL. However, if the photogrammetric analysis errors were within tolerable limits, the top of storm 3 (∼8.8 km) extended above the radar-detected echo top (∼6.0 km) [the radar-detected echo tops of storms 1 and 2 agreed well with the photogrammetric analysis (Fig. 5)].

Although additional photographs were not available, many other storms displayed supercellular structures. A sample of radar images in Fig. 7 reveals some of the more prominent radar signatures observed. Radar imagery from 2249 UTC (a few minutes before the tornado near Meridian, Oklahoma) showed an inflow notch on the pretornadic storm (Fig. 7a). Another storm to the southwest was associated with a reflectivity appendage. Velocity data at this time (Fig. 7b) also showed horizontal shears associated with mesocyclones (9.0 × 10−3 s−1 in the storm near Meridian; 2.1 × 10−2 s−1 in the in the storm southwest of the Meridian storm; 8.5 × 10−3 s−1 in the storm northeast of the Meridian storm). Echo tops of the storms included in Figs. 7a and 7b ranged from 4.9 to 6.0 km.

The storm labeled in Fig. 1a as storm 2 (which prompted a tornado warning) is depicted in an earlier base reflectivity image from 2314 UTC (Fig. 7c). A prominent hook echo is visible and a cross section of the same storm revealed a well-defined bounded weak echo region. Also visible in the cross section are low reflectivity echoes (<30 dBZ) that extend to roughly 10 km, similar to those that were observed later in time (a cross section of this storm from 2339 UTC was shown in Fig. 4d).

3. Sounding and hodograph characteristics

The 0000 UTC 29 October, Norman, Oklahoma, National Weather Service rawinsonde was released approximately 50 km ahead of the broken line of rotating storms while the storms were in progress. The sounding probably meets most criteria for being considered a representative “proximity” sounding (e.g., Darkow 1969;Brooks et al. 1994; Rasmussen and Blanchard 1998). Moreover, since upward mass flux in the storms was probably relatively small, it is not likely that the sounding would have been affected significantly by compensating subsidence 50 km away (Weisman et al. 1998).

The surface dewpoint on the sounding (Fig. 8) was modified from 18° to 19°C to reflect the dewpoint of the Oklahoma City surface airways observation at 2300 and 0000 UTC, as well as the dewpoint observed by the Oklahoma Mesonet (Brock et al. 1995) site at Norman6 from 2300–0000 UTC. It is possible that the low dewpoint bias of the rawinsonde (D. Andra and C. Doswell 1998, personal communication) may have affected measurements above the surface as well. This would have resulted in an underestimate of the mean mixing ratio in the lowest 50 mb, which was used in some of the CAPE and convective inhibition (CIN) calculations.

An undiluted, lifted surface parcel had 1510 J kg−1 of CAPE following “pure” parcel theory (parcel process curve A in Fig. 8). CAPE was substantially less (729 J kg−1) if a parcel with the mean thermodynamic characteristics of the lowest 50 mb was lifted according to pure parcel theory (parcel process curve B in Fig. 8).

Two additional parcel process curves are displayed in Fig. 8. Curve C is our best hand-drawn estimate of the parcel process curve associated with the storm having the highest observed echo top (11.2 km; 220 mb), storm 2 in Fig. 1a. Dilution of surface parcels en route to cloud base would have been likely owing to strong vertical shear in the subcloud layer,7 thus we assumed the lifted parcel possessed the mean thermodynamic characteristics of the lowest 50 mb (as assumed for curve B). Parcel process curve C deviates from parcel theory presumably because of entrainment, water loading, and other effects (Emanuel 1994). Since it is essentially impossible to accurately account for all of the individual thermodynamic effects on an ascending parcel, curve C was reasonably approximated by constraining the storm top to be located near the level of the highest observed echo top. The CAPE associated with this process curve, which we believe represents the upper limit of the amount of CAPE realized, was 305 J kg−1. The average parcel virtual temperature (Tυ) excess for this curve [from the level of free convection (LFC) to the equilibrium level (EL)] was 1.2 K. The maximum parcel Tυ excess was 2.8 K and occurred at 7.6 km (375 mb).

Similarly, we estimated a parcel process curve (curve D in Fig. 8) that may have been associated with the majority of the storms, which were observed to have echo tops from 5 to 6 km. The CAPE associated with this curve was 46 J kg−1. The average parcel Tυ excess for curve D (from the LFC to the uppermost EL) was 0.4 K. The maximum parcel Tυ excess was 1.8 K and occurred at 3.2 km (665 mb). Cloud and echo tops observed to range from approximately 6–8 km (475–360 mb; e.g., Fig. 7d) may have contained neutrally buoyant parcels that “coasted” through this layer (in which the lapse rate was nearly 9 K km−1).

Strong mean shear (1.8 × 10−2 s−1; Rasmussen and Wilhelmson 1983) was present in the lowest 3 km (Fig. 9). Storm-relative helicity (SRH; Davies-Jones 1984; Lilly 1986) also was large [195 m2 s−2 0–1 km SRH; 285 m2 s−2 0–2 km SRH; 397 m2 s−2 0–3 km SRH; note the particularly large 0–1 km SRH (Markowski et al. 1998)]. The bulk Richardson number (Weisman and Klemp 1982, 1984) for CAPE ranging from 46 to 305 J kg−1 was 0.4–2.7 (compared to 6.5–13.4 for CAPE ranging from 729 to 1510 J kg−1). The vertical shear present at low levels would have been favorable for development of a strong upward-directed dynamic vertical pressure gradient (Rotunno and Klemp 1982, 1985;Klemp 1987). Forced ascent owing to dynamic effects would have been required for parcels to overcome the CIN (35 J kg−1 for parcel process curves B, C, and D) present on the 0000 UTC sounding (Fig. 8). The visual observations by the authors of laminar, striated clouds at low levels were consistent with a sounding containing CIN and strong shear. In contrast to low levels, only weak, disorganized vertical shear was present from 4 to 8 km (Fig. 9). The appearance of an erect storm 2 (which extended to ∼11 km) in the photograph (Fig. 1a) is a visual manifestation of the lack of midlevel shear. [Although supercells often stand erect in the presence of strong shear owing to large dynamic vertical pressure gradient contributions to vertical velocity, the weak vertical velocities inferred in the 28 October storms (based on echo top growth analyses and the low CAPE) probably would have led to updrafts with significant tilt had strong shear been present.]

4. Final remarks

In the first section we proposed that the storms in central Oklahoma on 28 October 1998 might represent the most marginal form of supercell, although the events of the day showed how significant some of these storms could be—one was sufficiently organized for tornadogenesis. The environment of the storms presented in this paper had little in common with the previously documented environments of miniature supercells. The most notable difference was the lack of CAPE, which was considerably less than in the cases presented by Davies (1993), Stalker et al. (1993), Kennedy et al. (1993), Monteverdi and Quadros (1994),8 Knupp et al. (1998), and in numerical simulation studies by McCaul (1993), Wicker and Cantrell (1996), and McCaul and Weisman (1996). The CAPE present in lake-effect snow and rain events is commonly larger (up to 500–1000 J kg−1 or more; Miner and Fritsch 1997) than on the 0000 UTC Norman sounding. In previous studies of shallow supercells, CAPE was on the order of 1000 J kg−1, and low-level parcel accelerations were comparable in magnitude to the larger versions of supercells (Fig. 10; cf. Fig. 8).9 At the locations of maximum buoyancy, parcels were typically 5 K warmer than their surroundings, in contrast to a maximum (average) parcel Tυ excess of 1.8 K (0.4 K) in the present case for parcel process curve D. Furthermore, CAPE was relatively small and storm tops were relatively low in most of the previously documented minisupercell cases mainly because of a low tropopause, not because of low buoyancy throughout the troposphere. [McCaul (1991) studied hurricane environments that exhibited low buoyancy throughout the troposphere; however, the majority of storm tops observed on 28 October were lower than the storm tops typically observed in the tornado-producing convection associated with landfalling hurricanes.]

The synoptic-scale environment on 28 October 1998 in which the low-buoyancy, rotating updrafts formed also was different from most of the large-scale minisupercell environments previously documented (e.g., Kennedy et al. 1993; Monteverdi and Quandros 1994), in which the low-topped storms formed beneath cold, upper lows, often in the cold sector behind surface cold fronts; however, the storms were observed near the “possible secondary area for small supercells” suggested by Davies (1993).10 The 28 October 1998 storms formed in the warm sector under a high tropopause (near 200 mb) and relatively warm temperatures aloft (e.g., −8°C at 500 mb). One might speculate that if the warm pocket from 650 to 500 mb was absent, storms might have had a better chance of realizing the larger CAPE available to parcels that behaved according to pure parcel theory (the warm pocket did not allow for significant parcel accelerations owing to buoyancy, thus entrainment may have had a more detrimental effect in that layer than if parcel accelerations had been large). If this had been the case, then they might have resembled the taller, more “classic” supercell storms. Another question to be asked is whether something was special about the strength or structure of the convergence zone (e.g., Ziegler and Rasmussen 1998), since storms are rarely observed and cannot be simulated (Weisman and Klemp 1982, 1984, 1986) in such high-shear (∼2 × 10−2 s−1), low-CAPE (<100 J kg−1) environments. Or are these storms more common than believed?

The visual appearance of the 28 October 1998 storms was in at least a few ways different from those appearing in past literature. Previously documented minisupercells looked like shortened versions of the larger supercell, with long anvils and cumuliform clouds having the crisp appearance typically associated with buoyant moist convection. The storms presented herein did not produce elongated anvils (Fig. 3), nor was the convection what an experienced observer would characterize as vigorous (the observations of echo top growth also suggested weak updrafts). Visual similarities with previously documented supercells included the laminar, striated cloud bases and wall clouds.

Finally, this paper has shown that the visual appearance of a storm can be deceiving. The radar characteristics associated with these storms seemed surprising, given their relatively unimpressive visual appearance. This case demonstrates that, in some instances, the radar characteristics of a storm may be seemingly independent of the storm’s mid- and upper-level visual characteristics.

Acknowledgments

This paper was inspired by conversation with Jason Lynn (NSSL) as the events of the day unfolded. We appreciate thought-provoking discussions with Dr. Erik Rasmussen (CIMMS/NSSL) and Dr. Chuck Doswell (NSSL). We also are grateful to Gary England (KWTV, Oklahoma City), Daphne Zaras (NSSL), Greg Stumpf (NSSL), Arthur Witt (NSSL), Mike Eilts (NSSL), Dale Morris (Oklahoma Climate Survey), Dr. Ken Crawford (Oklahoma Climate Survey/OU), and Dr. Ron Holle (NSSL) for providing photographs, satellite, radar, surface, and lightning data. We also thank Rob Carver (OU) and Mark Askelson (OU) for their assistance, as well as the three anonymous reviewers. This work was partially funded by NSF Grants EAR-9512145, ATM-9120009, and ATM-9617318.

REFERENCES

  • Beebe, R. G., 1959: Notes on the Scottsbluff, Nebraska tornado, 27 June 1955. Bull. Amer. Meteor. Soc.,40, 109–116.

  • Bluestein, H. B., 1984: Further examples of low-precipitation severe thunderstorms. Mon. Wea. Rev.,112, 1885–1888.

  • ——, and C. R. Parks, 1983: A synoptic and photographic climatology of low-precipitation severe thunderstorms in the Southern Plains. Mon. Wea. Rev.,111, 2034–2046.

  • Brewster, K., F. Carr, N. Lin, J. M. Straka, and J. Krause, 1994: A local analysis system for initializing real-time convective-scale models. Preprints, 10th Conf. on Numerical Weather Prediction, Portland, OR, Amer. Meteor. Soc., 596–598.

  • Brock, F. V., K. Crawford, R. Elliot, G. Cuperus, S. Stadler, H. Johnson, and M. Eilts, 1995: The Oklahoma Mesonet: A technical overview. J. Atmos. Oceanic Technol.,12, 5–19.

  • Brooks, H. E., C. A. Doswell, and J. Cooper, 1994: On the environments of tornadic and nontornadic mesocyclones. Wea. Forecasting,9, 606–617.

  • Browning, K. A., 1964: Airflow and precipitation trajectories within severe local storms which travel to the right of the winds. J. Atmos. Sci.,21, 634–639.

  • ——, 1965: Some inferences about the updraft within a severe local storm. J. Atmos. Sci.,22, 669–678.

  • ——, 1977: The structure and mechanisms of hailstorms. Hail: A Review of Hail Science and Hail Suppression, Meteor. Monogr., No. 38, Amer. Meteor. Soc., 1–39.

  • ——, and R. J. Donaldson, 1963: Airflow and structure of a tornadic storm. J. Atmos. Sci.,20, 533–545.

  • Burgess, D. W., R. A. Brown, L. R. Lemon, and C. R. Safford, 1977:Evolution of a tornadic thunderstorm. Preprints, 10th Conf. on Severe Local Storms, Omaha, NE, Amer. Meteor. Soc., 84–89.

  • Chisholm, A. J., 1973: Radar case studies and airflow models. Alberta Hailstorms, Meteor. Monogr., No. 36, Amer. Meteor. Soc., 1–36.

  • Darkow, G. L., 1969: An analysis of over sixty tornado proximity soundings. Preprints, Sixth Conf. on Severe Local Storms, Chicago, IL, Amer. Meteor. Soc., 218–221.

  • Davies, J. M., 1993: Small tornadic supercells in the central Plains. Preprints, 17th Conf. on Severe Local Storms, St. Louis, MO, Amer. Meteor. Soc., 305–309.

  • Davies-Jones, R. P., 1984: Streamwise vorticity: The origin of updraft rotation in supercell storms. J. Atmos. Sci.,41, 2991–3006.

  • Doswell, C. A., 1996: What is a supercell? Preprints, 18th Conf. on Severe Local Storms, San Francisco, CA, Amer. Meteor. Soc., 641.

  • ——, and D. W. Burgess, 1993: Tornadoes and tornadic storms: A review of conceptual models. The Tornado: Its Structure, Dynamics, Prediction, and Hazards, Geophys. Monogr., No. 79, Amer. Gepophys. Union, 161–172.

  • ——, and E. N. Rasmussen, 1994: The effect of neglecting the virtual temperature correction on CAPE calculations. Wea. Forecasting,9, 625–629.

  • Emanuel, K. A., 1994: Atmospheric Convection. Oxford University Press, 580 pp.

  • Forbes, G. S., 1981: On the reliability of hook echoes as tornado indicators. Mon. Wea. Rev.,109, 1457–1466.

  • Fujita, T. T., 1958: Mesoanalysis of the Illinois tornadoes of 9 April 1953. J. Meteor.,15, 288–296.

  • ——, 1960: A detailed analysis of the Fargo tornadoes of June 20, 1957. U.S. Weather Bureau Research Paper No. 42, 67 pp. [Available from National Technical Information Service, Springfield, VA 22161.].

  • Johns, R. H., and C. A. Doswell, 1992: Severe local storms forecasting. Wea. Forecasting,7, 588–612.

  • Kennedy, P. C., N. E. Westcott, and R. W. Scott, 1993: Single-Doppler radar observations of a mini-supercell tornadic thunderstorm. Mon. Wea. Rev.,121, 1860–1870.

  • Klemp, J. B., 1987: Dynamics of tornadic thunderstorms. Annu. Rev. Fluid Mech.,19, 369–402.

  • Knight, C. A., and L. J. Miller, 1993: First radar echoes from cumulus clouds. Bull. Amer. Meteor. Soc.,74, 179–188.

  • Knupp, K. J., J. Stalker, and E. W. McCaul Jr., 1998: An observational and numerical study of a mini-supercell storm. Atmos. Res.,49, 35–63.

  • Lemon, L. R., 1977: Severe thunderstorm evolution: Its use in a new technique for radar warnings. Preprints, 10th Conf. on Severe Local Storms, Omaha, NE, Amer. Meteor. Soc., 77–83.

  • ——, and C. A. Doswell, 1979: Severe thunderstorm evolution and mesocyclone structure as related to tornadogenesis. Mon. Wea. Rev.,107, 1184–1197.

  • ——, D. W. Burgess, and R. A. Brown, 1978: Tornadic thunderstorm airflow and morphology derived from single-Doppler radar measurements. Mon. Wea. Rev.,106, 48–61.

  • Lilly, D. K., 1986: The structure, energetics, and propagation of rotating convective storms. Part II: Helicity and storm stabilization. J. Atmos. Sci.,43, 126–140.

  • Markowski, P. M., E. N. Rasmussen, and J. M. Straka, 1998: A preliminary investigation of the importance of helicity “location” in the hodograph. Preprints, 19th Conf. on Severe Local Storms, Minneapolis, MN, Amer. Meteor. Soc., 230–233.

  • Marwitz, J. D., 1972: The structure and motion of severe hailstorms. Part I: Supercell storms. J. Appl. Meteor.,11, 166–179.

  • McCaul, E. W., Jr., 1991: Buoyancy and shear characteristics of hurricane tornado environments. Mon. Wea. Rev.,119, 1954–1978.

  • ——, 1993: Observations and simulations of hurricane-spawned tornadic storms. The Tornado: Its Structure, Dynamics, Prediction, and Hazards, Geophys. Monogr., No. 79, Amer. Geophys. Union, 119–142.

  • ——, and M. L. Weisman, 1996: Simulations of shallow supercell storms in landfalling hurricanes. Mon. Wea. Rev.,124, 408–429.

  • Miner, T. J., and J. M. Fritsch, 1997: Lake-effect rain events. Mon. Wea. Rev.,125, 3231–3248.

  • Moller, A., C. A. Doswell, J. McGinley, S. Tegtmeier, and R. Zipser, 1974: Field observations of the Union City tornado in Oklahoma. Weatherwise,27, 68–77.

  • Monteverdi, J. P., and J. Quadros, 1994: Convective and rotational parameters associated with three tornado episodes in northern and central California. Wea. Forecasting,9, 285–300.

  • Peterson, R. E., 1976: The Sunray tornado. Bull. Amer. Meteor. Soc.,57, 805–807.

  • Rasmussen, E. N., and R. B. Wilhelmson, 1983: Relationships between storm characteristics and 1200 GMT hodographs, low-level shear, and stability. Preprints, 13th Conf. on Severe Local Storms, Boston, MA, Amer. Meteor. Soc., J5–J8.

  • ——, and D. O. Blanchard, 1998: A baseline climatology of sounding-derived supercell and tornado forecast parameters. Wea. Forecasting,13, 1148–1164.

  • Ray, P. S., B. C. Johnson, K. W. Johnson, J. S. Bradberry, J. J. Stephens, K. K. Wagner, R. B. Wilhelmson, and J. B. Klemp, 1981: The morphology of several tornadic storms on 20 May 1977. J. Atmos. Sci.,38, 1643–1663.

  • Rotunno, R., and J. B. Klemp, 1982: The influence of the shear-induced pressure gradient on thunderstorm motion. Mon. Wea. Rev.,110, 136–151.

  • ——, and ——, 1985: On the rotation and propagation of simulated supercell thunderstorms. J. Atmos. Sci.,42, 271–292.

  • Stalker, J. R., K. R. Knupp, and E. W. McCaul Jr., 1993: A numerical and observational study of an atypical “miniature” supercell storm. Preprints, 17th Conf. on Severe Local Storms, St. Louis, MO, Amer. Meteor. Soc., 191–195.

  • Stout, G. E., and F. A. Huff, 1953: Radar records Illinois tornadogenesis. Bull. Amer. Meteor. Soc.,34, 281–284.

  • Wakimoto, R. M., and V. N. Bringi, 1988: Dual-polarization observations of microbursts associated with intense convection: The 20 July storm during the MIST project. Mon. Wea. Rev.,116, 1521–1539.

  • WATADS, 1998: Reference guide for version 10.0. 72 pp. [Available from Storm Scale Applications Division, National Severe Storms Laboratory, 1313 Halley Circle, Norman, OK 73069.].

  • Weisman, M. L., and J. B. Klemp, 1982: The dependence of numerically simulated convective storms on vertical wind shear and buoyancy. Mon. Wea. Rev.,110, 504–520.

  • ——, and ——, 1984: The structure and classification of numerically simulated convective storms in directionally-varying wind shears. Mon. Wea. Rev.,112, 2479–2498.

  • ——, and ——, 1986: Characteristics of isolated convective storms. Mesoscale Meteorology and Forecasting, P. S. Ray, Ed., Amer. Meteor. Soc., 331–358.

  • ——, M. S. Gilmore, and L. J. Wicker, 1998: The impact of convective storms on their local environment: What is an appropriate ambient sounding? Preprints, 19th Conf. on Severe Local Storms, Minneapolis, MN, Amer. Meteor. Soc., 238–241.

  • Wicker, L. J., and L. Cantrell, 1996: The role of vertical buoyancy distributions in miniature supercells. Preprints, 18th Conf. on Severe Local Storms, San Francisco, CA, Amer. Meteor. Soc., 225–229.

  • Ziegler, C. L., and E. N. Rasmussen, 1998: The initiation of moist convection at the dryline: Forecasting issues from a case study perspective. Wea. Forecasting,13, 1106–1131.

 Fig. 1.
Fig. 1.

(a) Photograph from 2340 UTC 28 Oct 1998 (540 pm CST) from Oklahoma City, Oklahoma, looking north-northeast. Clouds labeled as 1, 2, and 3 are also labeled as such in Fig. 4. Clouds 1, 2, and 3 were associated with echo tops of approximately 5.2, 11.2, and 6.0 km AGL, respectively, as determined by WSR-88D data (however, the cloud tops were estimated photogrammetrically to be at 6.1, 10.5, and 8.8 km AGL, respectively). The upper halves of clouds 2 and 3 visually appear to lack significant buoyancy. Local television crews observed rotating wall clouds at the bases of updrafts 1 and 2. Both storms prompted tornado warnings that were issued by the National Weather Service in Norman, Oklahoma. (Photograph by P. Markowski.) (b) Video frame from 2250 UTC 28 Oct 1998 (450 p.m. CST) showing a wall cloud and clear slot near Meridian, Oklahoma. (Courtesy of Gary England, KWTV.) (c) Video frame from 2255 UTC 28 Oct 1998 (455 pm CST) of a tornado near Meridian, Oklahoma. (Courtesy of Gary England, KWTV.)

Citation: Monthly Weather Review 128, 2; 10.1175/1520-0493(2000)128<0449:SOORUI>2.0.CO;2

Fig. 2.
Fig. 2.

(a) Surface and upper-air composite at 0000 UTC 29 Oct 1998. Geopotential height contours (gray) are labeled in dam. Surface isobars (black) are labeled in mb. Fronts are depicted using conventional symbology, and the dryline is depicted as the boundary with unshaded scallops. The lightly shaded region encloses 500-mb wind speeds of 25 m s−1 or greater as analyzed by ADAS. The region enclosed by the dashed rectangle is enlarged in (b). Note that the analysis in (b) is valid at 2345 UTC 28 Oct. Temperatures and dewpoints in (b) are in °F. Each full wind barb is 5 m s−1 (10 kt). Radar echoes >25 dBZ detected on the lowest tilt by the KTLX (Twin Lakes, Oklahoma) WSR-88D are lightly shaded in (b).

Citation: Monthly Weather Review 128, 2; 10.1175/1520-0493(2000)128<0449:SOORUI>2.0.CO;2

Fig. 3.
Fig. 3.

GOES-8 1-km resolution visible satellite image from 2315 UTC 28 Oct 1998.

Citation: Monthly Weather Review 128, 2; 10.1175/1520-0493(2000)128<0449:SOORUI>2.0.CO;2

 Fig. 4.
Fig. 4.

(a) Base reflectivity (1.4° elevation angle) from the KTLX (Twin Lakes, Oklahoma) WSR-88D at 2339 UTC 28 Oct. Storms labeled as 1, 2, and 3 are the same storms appearing in Fig. 1a under the same labels. Tornado warnings were in effect for storms 1 and 2 at this time. The photograph in Fig. 1a was taken from the location marked X. The lines through storms 1 and 2 show the paths along which the cross sections of (c) and (d) were constructed. (b) As in (a) but base velocities are shown. (c) Cross section of storm 1 at 2339 UTC. The abscissa and ordinate are labeled in km. (d) As in (c) but for storm 2. The KTLX radar is to the south [outside of the lower boundary of the region visible in (a) and (b).]

Citation: Monthly Weather Review 128, 2; 10.1175/1520-0493(2000)128<0449:SOORUI>2.0.CO;2

Fig. 5.
Fig. 5.

As in Fig. 1a but with reflectivity (every 5 dBZ contoured) detected by the KTLX WSR-88D superimposed. Note that middle and upper portions of cloud 2 are associated with reflectivity <30 dBZ, with much of the upper portion having reflectivity <20 dBZ. (Photograph by P. Markowski.)

Citation: Monthly Weather Review 128, 2; 10.1175/1520-0493(2000)128<0449:SOORUI>2.0.CO;2

Fig. 6.
Fig. 6.

PPI plots (0.5°, 5.2°, 6.2°, 7.5°, 8.7°, and 12.0° elevation angles) from the KTLX WSR-88D from 2339 to 2343 UTC 28 Oct. Reflectivity 0 dBZ and greater is shaded light gray. On the 0.5° elevation angle PPI, regions of reflectivity exceeding 30 dBZ are shaded with a darker gray. On each range ring, the first number is the range from KTLX and the second number is the beam height at that range. The 0 dBZ echo threshold was chosen as a conservative approximation of the visible cloud boundaries. The 12° elevation angle PPI was the highest PPI on which echo associated with storm 3 was detected. The small square in the lower right corner of the 0.5° elevation angle PPI shows a 4 km × 4 km square, which represents the resolution of infrared satellite imagery; thus, cloud-top temperatures were not able to be reliably determined from satellite data.

Citation: Monthly Weather Review 128, 2; 10.1175/1520-0493(2000)128<0449:SOORUI>2.0.CO;2

 Fig. 7.
Fig. 7.

A sample of some of the supercellular radar (KTLX) structures observed on 28 Oct 1998. Base reflectivity and base velocity (0.5° elevation angle) at 2249 UTC are shown in (a) and (b), respectively, a few minutes prior to the tornado sighting near Meridian. (c) Base reflectivity (0.5° elevation angle) of a storm that prompted a tornado warning (appears as storm 2 in Fig. 1a 26 min later) at 2314 UTC. A cross section through this storm in (d) shows a well-defined bounded weak echo region.

Citation: Monthly Weather Review 128, 2; 10.1175/1520-0493(2000)128<0449:SOORUI>2.0.CO;2

Fig. 8.
Fig. 8.

Skew T-logp diagram depicting the 0000 UTC 29 Oct 1998 Norman, Oklahoma, sounding. Pressures are in mb, mixing ratios are in g kg−1, and temperatures are in Celsius. Each full wind barb is 5 m s−1 (10 kt) and each triangle is 12.5 m s−1 (25 kt). The surface dewpoint was modified (to 19°C) to equal the value of the 0000 UTC Oklahoma City surface airways observation (which also equaled the dewpoint observed at 2300 UTC, as well as the dewpoint observed in Norman by the Oklahoma Mesonet from 2300 to 0000 UTC), due to a low bias in the rawinsonde surface dewpoint (D. Andra and C. Doswell 1998, personal communication). The four parcel process curves are for (a) an undiluted surface parcel (parcel theory), (b) a parcel lifted with the mean potential temperature and mixing ratio of the lowest 50 mb (parcel theory), (c) an estimated parcel process curve associated with the storm having the highest observed echo top, and (d) an estimated parcel process curve associated with the majority of the storms, which were observed to have echo tops from 5 to 6 km.

Citation: Monthly Weather Review 128, 2; 10.1175/1520-0493(2000)128<0449:SOORUI>2.0.CO;2

Fig. 9.
Fig. 9.

Hodograph for the 0000 UTC 29 Oct 1998 Norman, Oklahoma, sounding. Labels along the hodograph are heights AGL in km.

Citation: Monthly Weather Review 128, 2; 10.1175/1520-0493(2000)128<0449:SOORUI>2.0.CO;2

Fig. 10.
Fig. 10.

Skew T-logp diagram for 0000 UTC 12 Mar 1990 constructed using the Dodge City and Topeka, Kansas, soundings (from Davies 1993). The shaded positive area was obtained by lifting an undiluted surface parcel. This minisupercell environment is considerably different than the environment sampled on 28 Oct 1998 in central Oklahoma (cf. Fig. 8). Maximum parcel buoyancy was comparable to the maximum parcel buoyancy characteristic of the environments of more “classic” supercells; however, the relatively low tropopause allowed storm tops to extend only slightly above 400 mb.

Citation: Monthly Weather Review 128, 2; 10.1175/1520-0493(2000)128<0449:SOORUI>2.0.CO;2

2

Hereafter, clouds 1, 2, and 3 are referred to as storms 1, 2, and 3.

3

Meridian, Oklahoma, is approximately 50 km north-northeast of the location of the photograph shown in Fig. 1a.

4

Cloud-top temperatures could not be reliably determined from infrared satellite imagery owing to the 4-km spatial resolution of such data; the updrafts and cloud tops (estimated from WSR-88D cross sections and plan position indicators to be generally less than 5–8 km) would have each occupied only one or two pixels at most.

5

All CAPE calculations include virtual temperature effects (following Doswell and Rasmussen 1994).

6

This observation site is approximately 1 km from the rawinsonde site.

7

Strong vertical shear could have at least two effects, both of which tend to reduce buoyancy: 1) enhancing the mechanical production of turbulence and associated mixing, and 2) lengthening of subcloud layer parcel trajectories (Ziegler and Rasmussen 1998).

8

One of the soundings in Monteverdi and Quadros (Fig. 4 in their paper) showed more similarity to the 0000 UTC Norman sounding than the other minisupercell soundings, in terms of the location of the EL and CAPE.

9

We believe it is unlikely that only a small fraction of the CAPE cited in the previously documented shallow supercell studies was actually realized, since the observed storm tops in those studies were close to what parcel theory would predict. In the present case, observed storm tops were considerably lower than what pure parcel theory would have predicted based on the parcel theory CAPE values of O(1000) J kg−1.

10

This secondary area was allowed only if cold air was present aloft (Davies 1993); on 28 October 1998, the mid- and upper-tropospheric temperatures were anomalously warm (e.g., −8°C at 500 mb).

1

Only one landmark suitable for Photogrammetric analysis was present in the photograph; thus, these cloud-top estimates may contain errors of several hundred meters. Furthermore, cloud widths could not be estimated photogrammetrically.

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  • Beebe, R. G., 1959: Notes on the Scottsbluff, Nebraska tornado, 27 June 1955. Bull. Amer. Meteor. Soc.,40, 109–116.

  • Bluestein, H. B., 1984: Further examples of low-precipitation severe thunderstorms. Mon. Wea. Rev.,112, 1885–1888.

  • ——, and C. R. Parks, 1983: A synoptic and photographic climatology of low-precipitation severe thunderstorms in the Southern Plains. Mon. Wea. Rev.,111, 2034–2046.

  • Brewster, K., F. Carr, N. Lin, J. M. Straka, and J. Krause, 1994: A local analysis system for initializing real-time convective-scale models. Preprints, 10th Conf. on Numerical Weather Prediction, Portland, OR, Amer. Meteor. Soc., 596–598.

  • Brock, F. V., K. Crawford, R. Elliot, G. Cuperus, S. Stadler, H. Johnson, and M. Eilts, 1995: The Oklahoma Mesonet: A technical overview. J. Atmos. Oceanic Technol.,12, 5–19.

  • Brooks, H. E., C. A. Doswell, and J. Cooper, 1994: On the environments of tornadic and nontornadic mesocyclones. Wea. Forecasting,9, 606–617.

  • Browning, K. A., 1964: Airflow and precipitation trajectories within severe local storms which travel to the right of the winds. J. Atmos. Sci.,21, 634–639.

  • ——, 1965: Some inferences about the updraft within a severe local storm. J. Atmos. Sci.,22, 669–678.

  • ——, 1977: The structure and mechanisms of hailstorms. Hail: A Review of Hail Science and Hail Suppression, Meteor. Monogr., No. 38, Amer. Meteor. Soc., 1–39.

  • ——, and R. J. Donaldson, 1963: Airflow and structure of a tornadic storm. J. Atmos. Sci.,20, 533–545.

  • Burgess, D. W., R. A. Brown, L. R. Lemon, and C. R. Safford, 1977:Evolution of a tornadic thunderstorm. Preprints, 10th Conf. on Severe Local Storms, Omaha, NE, Amer. Meteor. Soc., 84–89.

  • Chisholm, A. J., 1973: Radar case studies and airflow models. Alberta Hailstorms, Meteor. Monogr., No. 36, Amer. Meteor. Soc., 1–36.

  • Darkow, G. L., 1969: An analysis of over sixty tornado proximity soundings. Preprints, Sixth Conf. on Severe Local Storms, Chicago, IL, Amer. Meteor. Soc., 218–221.

  • Davies, J. M., 1993: Small tornadic supercells in the central Plains. Preprints, 17th Conf. on Severe Local Storms, St. Louis, MO, Amer. Meteor. Soc., 305–309.

  • Davies-Jones, R. P., 1984: Streamwise vorticity: The origin of updraft rotation in supercell storms. J. Atmos. Sci.,41, 2991–3006.

  • Doswell, C. A., 1996: What is a supercell? Preprints, 18th Conf. on Severe Local Storms, San Francisco, CA, Amer. Meteor. Soc., 641.

  • ——, and D. W. Burgess, 1993: Tornadoes and tornadic storms: A review of conceptual models. The Tornado: Its Structure, Dynamics, Prediction, and Hazards, Geophys. Monogr., No. 79, Amer. Gepophys. Union, 161–172.

  • ——, and E. N. Rasmussen, 1994: The effect of neglecting the virtual temperature correction on CAPE calculations. Wea. Forecasting,9, 625–629.

  • Emanuel, K. A., 1994: Atmospheric Convection. Oxford University Press, 580 pp.

  • Forbes, G. S., 1981: On the reliability of hook echoes as tornado indicators. Mon. Wea. Rev.,109, 1457–1466.

  • Fujita, T. T., 1958: Mesoanalysis of the Illinois tornadoes of 9 April 1953. J. Meteor.,15, 288–296.

  • ——, 1960: A detailed analysis of the Fargo tornadoes of June 20, 1957. U.S. Weather Bureau Research Paper No. 42, 67 pp. [Available from National Technical Information Service, Springfield, VA 22161.].

  • Johns, R. H., and C. A. Doswell, 1992: Severe local storms forecasting. Wea. Forecasting,7, 588–612.

  • Kennedy, P. C., N. E. Westcott, and R. W. Scott, 1993: Single-Doppler radar observations of a mini-supercell tornadic thunderstorm. Mon. Wea. Rev.,121, 1860–1870.

  • Klemp, J. B., 1987: Dynamics of tornadic thunderstorms. Annu. Rev. Fluid Mech.,19, 369–402.

  • Knight, C. A., and L. J. Miller, 1993: First radar echoes from cumulus clouds. Bull. Amer. Meteor. Soc.,74, 179–188.

  • Knupp, K. J., J. Stalker, and E. W. McCaul Jr., 1998: An observational and numerical study of a mini-supercell storm. Atmos. Res.,49, 35–63.

  • Lemon, L. R., 1977: Severe thunderstorm evolution: Its use in a new technique for radar warnings. Preprints, 10th Conf. on Severe Local Storms, Omaha, NE, Amer. Meteor. Soc., 77–83.

  • ——, and C. A. Doswell, 1979: Severe thunderstorm evolution and mesocyclone structure as related to tornadogenesis. Mon. Wea. Rev.,107, 1184–1197.

  • ——, D. W. Burgess, and R. A. Brown, 1978: Tornadic thunderstorm airflow and morphology derived from single-Doppler radar measurements. Mon. Wea. Rev.,106, 48–61.

  • Lilly, D. K., 1986: The structure, energetics, and propagation of rotating convective storms. Part II: Helicity and storm stabilization. J. Atmos. Sci.,43, 126–140.

  • Markowski, P. M., E. N. Rasmussen, and J. M. Straka, 1998: A preliminary investigation of the importance of helicity “location” in the hodograph. Preprints, 19th Conf. on Severe Local Storms, Minneapolis, MN, Amer. Meteor. Soc., 230–233.

  • Marwitz, J. D., 1972: The structure and motion of severe hailstorms. Part I: Supercell storms. J. Appl. Meteor.,11, 166–179.

  • McCaul, E. W., Jr., 1991: Buoyancy and shear characteristics of hurricane tornado environments. Mon. Wea. Rev.,119, 1954–1978.

  • ——, 1993: Observations and simulations of hurricane-spawned tornadic storms. The Tornado: Its Structure, Dynamics, Prediction, and Hazards, Geophys. Monogr., No. 79, Amer. Geophys. Union, 119–142.

  • ——, and M. L. Weisman, 1996: Simulations of shallow supercell storms in landfalling hurricanes. Mon. Wea. Rev.,124, 408–429.

  • Miner, T. J., and J. M. Fritsch, 1997: Lake-effect rain events. Mon. Wea. Rev.,125, 3231–3248.

  • Moller, A., C. A. Doswell, J. McGinley, S. Tegtmeier, and R. Zipser, 1974: Field observations of the Union City tornado in Oklahoma. Weatherwise,27, 68–77.

  • Monteverdi, J. P., and J. Quadros, 1994: Convective and rotational parameters associated with three tornado episodes in northern and central California. Wea. Forecasting,9, 285–300.

  • Peterson, R. E., 1976: The Sunray tornado. Bull. Amer. Meteor. Soc.,57, 805–807.

  • Rasmussen, E. N., and R. B. Wilhelmson, 1983: Relationships between storm characteristics and 1200 GMT hodographs, low-level shear, and stability. Preprints, 13th Conf. on Severe Local Storms, Boston, MA, Amer. Meteor. Soc., J5–J8.

  • ——, and D. O. Blanchard, 1998: A baseline climatology of sounding-derived supercell and tornado forecast parameters. Wea. Forecasting,13, 1148–1164.

  • Ray, P. S., B. C. Johnson, K. W. Johnson, J. S. Bradberry, J. J. Stephens, K. K. Wagner, R. B. Wilhelmson, and J. B. Klemp, 1981: The morphology of several tornadic storms on 20 May 1977. J. Atmos. Sci.,38, 1643–1663.

  • Rotunno, R., and J. B. Klemp, 1982: The influence of the shear-induced pressure gradient on thunderstorm motion. Mon. Wea. Rev.,110, 136–151.

  • ——, and ——, 1985: On the rotation and propagation of simulated supercell thunderstorms. J. Atmos. Sci.,42, 271–292.

  • Stalker, J. R., K. R. Knupp, and E. W. McCaul Jr., 1993: A numerical and observational study of an atypical “miniature” supercell storm. Preprints, 17th Conf. on Severe Local Storms, St. Louis, MO, Amer. Meteor. Soc., 191–195.

  • Stout, G. E., and F. A. Huff, 1953: Radar records Illinois tornadogenesis. Bull. Amer. Meteor. Soc.,34, 281–284.

  • Wakimoto, R. M., and V. N. Bringi, 1988: Dual-polarization observations of microbursts associated with intense convection: The 20 July storm during the MIST project. Mon. Wea. Rev.,116, 1521–1539.

  • WATADS, 1998: Reference guide for version 10.0. 72 pp. [Available from Storm Scale Applications Division, National Severe Storms Laboratory, 1313 Halley Circle, Norman, OK 73069.].

  • Weisman, M. L., and J. B. Klemp, 1982: The dependence of numerically simulated convective storms on vertical wind shear and buoyancy. Mon. Wea. Rev.,110, 504–520.

  • ——, and ——, 1984: The structure and classification of numerically simulated convective storms in directionally-varying wind shears. Mon. Wea. Rev.,112, 2479–2498.

  • ——, and ——, 1986: Characteristics of isolated convective storms. Mesoscale Meteorology and Forecasting, P. S. Ray, Ed., Amer. Meteor. Soc., 331–358.

  • ——, M. S. Gilmore, and L. J. Wicker, 1998: The impact of convective storms on their local environment: What is an appropriate ambient sounding? Preprints, 19th Conf. on Severe Local Storms, Minneapolis, MN, Amer. Meteor. Soc., 238–241.

  • Wicker, L. J., and L. Cantrell, 1996: The role of vertical buoyancy distributions in miniature supercells. Preprints, 18th Conf. on Severe Local Storms, San Francisco, CA, Amer. Meteor. Soc., 225–229.

  • Ziegler, C. L., and E. N. Rasmussen, 1998: The initiation of moist convection at the dryline: Forecasting issues from a case study perspective. Wea. Forecasting,13, 1106–1131.

  • Fig. 1.

    (a) Photograph from 2340 UTC 28 Oct 1998 (540 pm CST) from Oklahoma City, Oklahoma, looking north-northeast. Clouds labeled as 1, 2, and 3 are also labeled as such in Fig. 4. Clouds 1, 2, and 3 were associated with echo tops of approximately 5.2, 11.2, and 6.0 km AGL, respectively, as determined by WSR-88D data (however, the cloud tops were estimated photogrammetrically to be at 6.1, 10.5, and 8.8 km AGL, respectively). The upper halves of clouds 2 and 3 visually appear to lack significant buoyancy. Local television crews observed rotating wall clouds at the bases of updrafts 1 and 2. Both storms prompted tornado warnings that were issued by the National Weather Service in Norman, Oklahoma. (Photograph by P. Markowski.) (b) Video frame from 2250 UTC 28 Oct 1998 (450 p.m. CST) showing a wall cloud and clear slot near Meridian, Oklahoma. (Courtesy of Gary England, KWTV.) (c) Video frame from 2255 UTC 28 Oct 1998 (455 pm CST) of a tornado near Meridian, Oklahoma. (Courtesy of Gary England, KWTV.)

  • Fig. 2.

    (a) Surface and upper-air composite at 0000 UTC 29 Oct 1998. Geopotential height contours (gray) are labeled in dam. Surface isobars (black) are labeled in mb. Fronts are depicted using conventional symbology, and the dryline is depicted as the boundary with unshaded scallops. The lightly shaded region encloses 500-mb wind speeds of 25 m s−1 or greater as analyzed by ADAS. The region enclosed by the dashed rectangle is enlarged in (b). Note that the analysis in (b) is valid at 2345 UTC 28 Oct. Temperatures and dewpoints in (b) are in °F. Each full wind barb is 5 m s−1 (10 kt). Radar echoes >25 dBZ detected on the lowest tilt by the KTLX (Twin Lakes, Oklahoma) WSR-88D are lightly shaded in (b).

  • Fig. 3.

    GOES-8 1-km resolution visible satellite image from 2315 UTC 28 Oct 1998.

  • Fig. 4.

    (a) Base reflectivity (1.4° elevation angle) from the KTLX (Twin Lakes, Oklahoma) WSR-88D at 2339 UTC 28 Oct. Storms labeled as 1, 2, and 3 are the same storms appearing in Fig. 1a under the same labels. Tornado warnings were in effect for storms 1 and 2 at this time. The photograph in Fig. 1a was taken from the location marked X. The lines through storms 1 and 2 show the paths along which the cross sections of (c) and (d) were constructed. (b) As in (a) but base velocities are shown. (c) Cross section of storm 1 at 2339 UTC. The abscissa and ordinate are labeled in km. (d) As in (c) but for storm 2. The KTLX radar is to the south [outside of the lower boundary of the region visible in (a) and (b).]

  • Fig. 5.

    As in Fig. 1a but with reflectivity (every 5 dBZ contoured) detected by the KTLX WSR-88D superimposed. Note that middle and upper portions of cloud 2 are associated with reflectivity <30 dBZ, with much of the upper portion having reflectivity <20 dBZ. (Photograph by P. Markowski.)

  • Fig. 6.

    PPI plots (0.5°, 5.2°, 6.2°, 7.5°, 8.7°, and 12.0° elevation angles) from the KTLX WSR-88D from 2339 to 2343 UTC 28 Oct. Reflectivity 0 dBZ and greater is shaded light gray. On the 0.5° elevation angle PPI, regions of reflectivity exceeding 30 dBZ are shaded with a darker gray. On each range ring, the first number is the range from KTLX and the second number is the beam height at that range. The 0 dBZ echo threshold was chosen as a conservative approximation of the visible cloud boundaries. The 12° elevation angle PPI was the highest PPI on which echo associated with storm 3 was detected. The small square in the lower right corner of the 0.5° elevation angle PPI shows a 4 km × 4 km square, which represents the resolution of infrared satellite imagery; thus, cloud-top temperatures were not able to be reliably determined from satellite data.

  • Fig. 7.

    A sample of some of the supercellular radar (KTLX) structures observed on 28 Oct 1998. Base reflectivity and base velocity (0.5° elevation angle) at 2249 UTC are shown in (a) and (b), respectively, a few minutes prior to the tornado sighting near Meridian. (c) Base reflectivity (0.5° elevation angle) of a storm that prompted a tornado warning (appears as storm 2 in Fig. 1a 26 min later) at 2314 UTC. A cross section through this storm in (d) shows a well-defined bounded weak echo region.

  • Fig. 8.

    Skew T-logp diagram depicting the 0000 UTC 29 Oct 1998 Norman, Oklahoma, sounding. Pressures are in mb, mixing ratios are in g kg−1, and temperatures are in Celsius. Each full wind barb is 5 m s−1 (10 kt) and each triangle is 12.5 m s−1 (25 kt). The surface dewpoint was modified (to 19°C) to equal the value of the 0000 UTC Oklahoma City surface airways observation (which also equaled the dewpoint observed at 2300 UTC, as well as the dewpoint observed in Norman by the Oklahoma Mesonet from 2300 to 0000 UTC), due to a low bias in the rawinsonde surface dewpoint (D. Andra and C. Doswell 1998, personal communication). The four parcel process curves are for (a) an undiluted surface parcel (parcel theory), (b) a parcel lifted with the mean potential temperature and mixing ratio of the lowest 50 mb (parcel theory), (c) an estimated parcel process curve associated with the storm having the highest observed echo top, and (d) an estimated parcel process curve associated with the majority of the storms, which were observed to have echo tops from 5 to 6 km.

  • Fig. 9.

    Hodograph for the 0000 UTC 29 Oct 1998 Norman, Oklahoma, sounding. Labels along the hodograph are heights AGL in km.

  • Fig. 10.

    Skew T-logp diagram for 0000 UTC 12 Mar 1990 constructed using the Dodge City and Topeka, Kansas, soundings (from Davies 1993). The shaded positive area was obtained by lifting an undiluted surface parcel. This minisupercell environment is considerably different than the environment sampled on 28 Oct 1998 in central Oklahoma (cf. Fig. 8). Maximum parcel buoyancy was comparable to the maximum parcel buoyancy characteristic of the environments of more “classic” supercells; however, the relatively low tropopause allowed storm tops to extend only slightly above 400 mb.

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