• Brady, R. H., and E. J. Szoke, 1989: Case study of a nonmesocyclone tornado development in northeast Colorado: Similarities to waterspout formation. Mon. Wea. Rev.,117, 843–856.

  • Davies, J. M., C. A. Doswell III, D. F. Burgess, and J. F. Weaver, 1994: Some noteworthy aspects of the Hesston, Kansas, tornado family of 13 March 1990. Bull. Amer. Meteor. Soc.,75, 1007–1017.

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

  • ——, D. Burgess, and M. Foster, 1990: Test of helicity as a tornado forecast parameter. Preprints, 16th Conf. on Severe Local Storms, Kananaskis Park, AB, Canada, Amer. Meteor. Soc., 588–592.

  • Dutton, J. A., 1976: The Ceaseless Wind. An Introduction to the Theory of Atmospheric Motion. McGraw-Hill, 579 pp.

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  • Markowski, P. M., E. N. Rasmussen, and J. M. Straka, 1998a: The occurrence of tornadoes in supercells interacting with boundaries during VORTEX-95. Wea. Forecasting,13, 852–859.

  • ——, J. M. Straka, E. N. Rasmussen, and D. Dowell, 1998b: Observations of low-level baroclinicity generated by anvil shadows. Mon. Wea. Rev.,126, 2942–2958.

  • National Oceanic and Atmospheric Administration, 1995: Storm Data. Vol. 37, No. 5, 221 pp.

  • Purdom, J. F. W., 1976: Some uses of high resolution GOES imagery in the mesoscale forecasting of convection and its behavior. Mon. Wea. Rev.,104, 1474–1483.

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

  • ——, J. M. Straka, R. P. Davies-Jones, C. A. Doswell III, F. H. Carr, M. D. Eilts, and D. R. MacGorman, 1994: The Verifications of the Origins of Rotation in Tornadoes Experiment: VORTEX. Bull. Amer. Meteor. Soc.,75, 997–1006.

  • 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 M. L. Weisman, 1988: A theory for strong, long-lived squall-lines. J. Atmos. Sci.,45, 463–485.

  • Straka, J. M., E. N. Rasmussen, and S. E. Frederickson, 1996: A mobile mesonet for fine-scale meteorological observations. J. Atmos. Oceanic Technol.,13, 921–936.

  • Wakimoto, R. M., and J. W. Wilson, 1989: Nonsupercell tornadoes. Mon. Wea. Rev.,117, 1113–1140.

  • ——, and N. T. Atkins, 1996: Observations on the origins of rotation:The Newcastle tornado during VORTEX-94. Mon. Wea. Rev.,124, 384–407.

  • ——, C.-H. Liu, and H. Cai, 1998: The Garden City, Kansas, storm during VORTEX-95. Part I: Overview of the storm’s life cycle and mesocyclogenesis. Mon. Wea. Rev.,126, 372–392.

  • Wilczak, J. M., T. W. Christian, D. E. Wolfe, R. J. Zamora, and B. Stankov, 1992: Observations of a Colorado tornado. Part I: Mesoscale environment and tornadogenesis. Mon. Wea. Rev.,120, 497–520.

  • Wood, V. T., and R. A. Brown, 1997: Effects of radar sampling on single-Doppler velocity signatures of mesocyclones and tornadoes. Wea. Forecasting,12, 928–938.

  • View in gallery

    Conceptual model of flow near a boundary. From Maddox et al. (1980 their Fig. 2).

  • View in gallery

    Skew T–logP diagram of soundings from Amarillo at 1103 UTC (solid), Texas Tech University (Lubbock) at 2015 UTC (thin broken) and 2348 UTC (thick broken). CAPEs are shown in the legend.

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    Hodographs from Texas Tech University (Lubbock) from 1715 UTC (thick solid), 2015 UTC (thin broken), and 2348 UTC (thick broken). Velocities in m s−1. Heights in km AGL. Values in the legend are SRH (m2 s−2) estimates spanning the range of observed storm motions.

  • View in gallery

    GOES-8 visible satellite image for 1815 UTC 2 Jun 1995.

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    GOES-8 visible satellite image for 2115 UTC 2 Jun 1995.

  • View in gallery

    Surface observations from 1800 UTC 2 Jun 1995. Full wind flags represent 5 m s−1; half flags 2.5 m s−1. Temperatures and dewpoints in °C. Sites in the Oklahoma panhandle are Oklahoma mesonet sites; all others NWS/FAA. Approximate boundary location marked with the bold line. Additional towns mentioned in the text are shown with small dots.

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    As in Fig. 6. but for 2100 UTC. Site marked P1 is an observation from the Probe-1 VORTEX vehicle.

  • View in gallery

    Isochrone analysis of the leading edge of the narrow cloud line associated with the prominent boundary. Position every 30 min, based on 15-min GOES-8 visible image analysis. Unlabeled filled circles are locations of communities named on other maps herein.

  • View in gallery

    Echo centroid tracks for five 1-h periods depicted with the various line types shown in the legend. Thick lines represent boundary positions according to the symbols in the legend, and apply to the start of the 1-h periods. Numbers in circles correspond to tornadoes listed in Table 1.

  • View in gallery

    Graph of low-level mesocyclone differential velocity (difference between inbound and outbound Doppler radial velocity peaks on base scan) vs distance from the prominent boundary (positive toward cooler air). Each dot represents one detection; all data at 10-min intervals are included.

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    Time–height diagram of the maximum reflectivity (dashed lines) and mesocyclone strength (½ the equivalent solid-body vertical vorticity; solid lines) for the Friona, TX storm. The storm crossed the prominent boundary during the period denoted with the heavy bar labeled B. The tornado occurred during the period shown with the heavy arrow. Rotation contoured and labeled in black (×10−4 s−1). Reflectivity contoured and labeled in gray (dBZ). Gray shading denotes time and heights at which rotation exceeded 2 × 10−2 s−1.

  • View in gallery

    As in Fig. 11 except for the Dimmitt, TX, tornadic storm, and with X denoting a reflectivity maximum and N a minimum.

  • View in gallery

    Analysis of mobile mesonet data (subset of observations from vehicle P1) during a transect of the boundary southwest of Amarillo, TX. Heavy line represents the boundary with the orientation established through analysis of GOES-8 data. Black lines are major highways; gray lines are other roads. Wind flags drawn with full barb = 5 m s−1; half barb = 2.5 m s−1. Temperature in degrees C; mixing ratios in 0.1 g kg−1. True north is toward the top of the map.

  • View in gallery

    As in Fig. 13 except for a boundary transect near Bovina, TX.

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    WSR88D reflectivity at 2102 UTC. Boundary and azimuth of the cross section in Fig. 16 is indicated with the heavy arrow. Reflectivity in 3-dB increments (increasingly darker shades) ranging from 0 to 15 dBZ.

  • View in gallery

    Cross section of Lubbock WSR-88D radial velocity at 2102 UTC (contoured; ms−1). Arrows are position of radar beams. Boundary-parallel component of horizontal vorticity s−1 represented by plotted values at the top of the figure, computed from interpolated radial velocity at 2 km AGL minus velocity at 0.5°, divided by the actual height difference. Location of radar fine line represented with heavy black bar between 45- and 49-km range.

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    As in Fig. 16 but for 2202 UTC.

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    Low-level skew T–logp diagrams of soundings taken at various locations with respect to the prominent boundary.

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    The 0–4 km AGL hodographs for the same locations and times as in Fig. 18. Wind speeds in m s−1. Surface points from surface analysis. Heights in km AGL.

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The Association of Significant Tornadoes with a Baroclinic Boundary on 2 June 1995

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  • 1 Cooperative Institute for Mesoscale Meteorological Studies, National Severe Storms Laboratory, and University of Oklahoma, Norman, Oklahoma
  • 2 Center for the Analysis and Prediction of Storms, University of Oklahoma, Norman, Oklahoma
  • 3 School of Meteorology and Center for Analysis and Prediction of Storms, University of Oklahoma, Norman, Oklahoma
  • 4 School of Meteorology and National Severe Storms Laboratory, University of Oklahoma, Norman, Oklahoma
  • 5 Cooperative Institute for Mesoscale Meteorological Studies, National Severe Storms Laboratory, and University of Oklahoma, Norman, Oklahoma
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Abstract

On 2 June 1995, the large-scale environment of eastern New Mexico and western Texas was generally favorable for the occurrence of supercells because of the presence of strong deep shear and storm-relative helicity, as well as sufficient convective available potential energy (CAPE). Indeed, many supercells occurred, but the only storms to produce tornadoes were those supercells that crossed, or developed and persisted on the immediate cool side of a particular outflow boundary generated by earlier convection. Surface conditions, vertical vorticity, and horizontal vorticity near this boundary are documented using conventional and special observations from the VORTEX field program. It is shown that the boundary was locally rich in horizontal vorticity, had somewhat enhanced vertical vorticity, and enhanced CAPE. Theoretical arguments indicate that the observed horizontal vorticity (around 1 × 10−2 s−1), largely parallel to the boundary, can be readily produced with the type of buoyancy contrast observed. It is hypothesized that such local enhancement of horizontal vorticity often is required for the occurrence of significant (e.g., F2 or stronger) tornadoes, even in large-scale environments that appear conducive to tornado occurrence without the aid of local influences.

Corresponding author address: Erik Rasmussen, 3450 Mitchell Lane, Bldg. 3, Rm. 2034, Boulder, CO 80301.

Email: rasmussen@nssl.noaa.gov

Abstract

On 2 June 1995, the large-scale environment of eastern New Mexico and western Texas was generally favorable for the occurrence of supercells because of the presence of strong deep shear and storm-relative helicity, as well as sufficient convective available potential energy (CAPE). Indeed, many supercells occurred, but the only storms to produce tornadoes were those supercells that crossed, or developed and persisted on the immediate cool side of a particular outflow boundary generated by earlier convection. Surface conditions, vertical vorticity, and horizontal vorticity near this boundary are documented using conventional and special observations from the VORTEX field program. It is shown that the boundary was locally rich in horizontal vorticity, had somewhat enhanced vertical vorticity, and enhanced CAPE. Theoretical arguments indicate that the observed horizontal vorticity (around 1 × 10−2 s−1), largely parallel to the boundary, can be readily produced with the type of buoyancy contrast observed. It is hypothesized that such local enhancement of horizontal vorticity often is required for the occurrence of significant (e.g., F2 or stronger) tornadoes, even in large-scale environments that appear conducive to tornado occurrence without the aid of local influences.

Corresponding author address: Erik Rasmussen, 3450 Mitchell Lane, Bldg. 3, Rm. 2034, Boulder, CO 80301.

Email: rasmussen@nssl.noaa.gov

1. Introduction

It recently has been established that the majority of significant (>F1) tornadoes in the Verifications of the Origin of Rotation in Tornadoes Experiment (VORTEX;Rasmussen et al. 1994) domain in 1995 occurred on the cool side of, and in proximity to, boundaries (Markowski et al. 1998a). Numerous other case studies (e.g., Purdom 1976; Maddox et al. 1980; Davies et al. 1994;Wakimoto and Atkins 1996; Wakimoto et al. 1998) have documented the occurrence of tornadic supercell storms in proximity to boundaries. However, the recent findings from VORTEX are making it apparent that boundaries such as outflows, synoptic fronts, various radar-detectable fineline features of unknown origin, and even boundaries established by differences in the radiation balance across the edge of an anvil shadow (Markowski et al. 1998b) may be generally important in tornado production.

Several plausible roles for boundaries in enhancing the likelihood of significant supercell tornadoes are explored in this paper. These are investigated using data from VORTEX collected on 2 June 1995. This case presents a unique dataset for examination of the nature and role of boundaries because many supercells occurred in west Texas and eastern New Mexico on this day, and a prominent outflow boundary was present. It will be shown that the supercells evolved quite differently depending on their location with respect to the boundary. Storms that remained on the warm side of the boundary, despite substantial convective available potential energy (CAPE) and storm-relative helicity (SRH), did not produce tornadoes, and in general did not have significant low-level [1–3 km above ground level (AGL)] mesocyclones. Storms that crossed the boundary from warm to cool in general produced significant tornadoes (e.g., the tornadoes at Dimmitt and Friona, Texas), and evolved very rapidly upon crossing the boundary. Storms that formed on or near the boundary, and then moved farther into the cool air mass, evolved relatively slowly, but did produce low-level mesocyclones as well as generally weaker tornadoes than the boundary-crossing storms.

In section 2, the possible roles of boundaries in enhancing the potential for low-level mesocyclone development are explored. Next, a case study of the boundary and storm evolution on 2 June 1995 is presented in section 3. The evolution of storm-scale rotation in two of the storms that produced violent tornadoes (Dimmitt and Friona storms) is examined in section 4, with emphasis on the events immediately surrounding the times that the storms crossed the boundary. In section 5, data are presented quantifying the relative magnitudes of horizontal and vertical vorticity, and the importance of other local changes near the boundary. Conclusions are presented in section 6.

2. Mesoscale influences of boundaries

a. Vertical vorticity enhancement

Maddox et al. (1980) examined four cases in which tornadoes occurred along thermal boundaries that had been produced by earlier thunderstorm activity. It was concluded that thermal boundaries are “a favored location for the occurrence of severe tornadic storms.” Based upon their observations, a physical model (Fig. 1) of boundary layer wind fields was developed to explain the intensification of thunderstorms as they move across and along thermal boundaries. Point A is within the warm, dry, well-mixed air mass where the winds veer slightly with height and speed increases slowly. Point B is in the hot, moist, and conditionally unstable air mass where the winds veer rapidly through the subcloud layer. Point C is in a region of cool, moist, and stable thunderstorm outflow where the winds veer slowly with height until the transition into the warmer air mass above, where pronounced veering occurs.

For conditions like those shown in Fig. 1, considering the mean subcloud winds, the vertical wind profiles were modified by the presence of the boundary so as to maximize the meso-γ-scale convergence and cyclonic vertical vorticity within a narrow zone between points B and C. In addition, the mean subcloud-layer convergence and vorticity between points A and B are actually smaller in magnitude than they are for the surface winds alone (Maddox et al. 1980). Maddox et al. suggested that this mesoscale intensification effect is most likely to cause minioutbreaks of intense tornadoes along thermal boundaries when conditions do not favor a widespread outbreaks of tornadoes.

b. Horizontal vorticity enhancement

Though not examined by Maddox et al., the production of horizontal vorticity by a thermal boundary cannot be neglected. Horizontal vorticity is readily stretched horizontally by storm-induced accelerations to the flow, and then tilted into, and stretched by the storm updraft. The amount of horizontal vorticity that might be produced near the ground can be estimated as follows. The two-dimensional horizontal vorticity equation for inviscid Boussinesq flow, neglecting the Coriolis effect (e.g., Rotunno et al. 1988), is
i1520-0493-128-1-174-e1
where x is the horizontal distance normal to the baroclinic zone (positive x toward the colder air), and buoyancy (neglecting the water loading term gq1), B, is
i1520-0493-128-1-174-e2
where gravitational acceleration is g and virtual temperature is Tυ with overbar denoting the environmental condition. The horizontal vorticity, η, which points toward +y (i.e., along the baroclinic zone), is
i1520-0493-128-1-174-e3
where u and w are the horizontal wind component normal to the boundary and the vertical wind component, respectively. In this simplified framework, horizontal gradients of buoyancy are the sources of horizontal vorticity.

Some comments about the limitations of this framework are required. The two-dimensional approximation can fail near storms for a number of reasons, the foremost being that η can be greatly increased by along-boundary horizontal stretching owing to storm inflow. Further, this discussion neglects the generation of boundary-normal horizontal vorticity, chiefly because we are unable to quantify it using the 2 June 1995 observational dataset. It should be noted that the Coriolis acceleration has the effect of turning the line-parallel horizontal vorticity, generated through buoyancy gradients, into the line-normal direction. However, the line-parallel component should generally remain dominant owing to the ongoing effects of generation.

Vorticity that is initially mainly horizontal can contribute to quasi-vertical vortices such as supercell mesocyclones through reorientation and stretching. Horizontal vorticity that is streamwise (i.e., vorticity and storm-relative velocity vectors parallel) produces net updraft rotation (Davies-Jones 1984) upon tilting. It is worthwhile to note that the tilting of crosswise horizontal vorticity (horizontal vorticity and velocity vectors perpendicular) also produces vertical vorticity. In the case of purely crosswise horizontal vorticity, integration of vertical vorticity over an entire updraft yields no net rotation; however, a pair of vortices, one cyclonic and the other anticyclonic, will result.

Another important, and related parameter that can be used operationally from a sounding is SRH. Relative to a moving storm,
i1520-0493-128-1-174-e4
where V is the velocity, c is the storm motion vector, and h is typically chosen to be 3 km, approximately the level-of-free convection and top of the inflow layer (Davies-Jones et al. 1990).

The vector of the horizontal vorticity generated in (1) is directed across the buoyancy gradient (along the buoyancy isopleths). Therefore, for a wide range of typical storm motion, the generation of horizontal vorticity due to buoyancy gradients will increase SRH to the degree that the flow is also along the buoyancy (temperature) isopleths. Away from active convection where boundaries have become quasi-stationary, buoyancy gradients typically exist in regions with a base-state motion normal to the gradient vector [i.e., flow parallel to the boundary; e.g., the present study and Maddox et al. (1980)]. Even if the flow was at one time largely perpendicular to the boundary, Coriolis acceleration will lead to flow increasingly parallel to the boundary with time.1 In these instances, the baroclinically generated horizontal vorticity will be largely streamwise.

From (1), the longer the residence time of a parcel in a boundary, the greater the amount of horizontal vorticity generated (Dutton 1976). For a 10-km wide zone with a 5°C temperature gradient, /dt is on the order of 2 × 10−5 s−2. If a parcel, with no vorticity initially, resides in such a gradient for 5 min, it can attain a horizontal vorticity of 6 × 10−3 s−1, which, if increased through stretching by a factor of 1.7 or more subsequent to tilting, would support a mesocyclone (∼1 × 10−2 s−1). However, actual horizontal vorticity in the atmosphere would be smaller depending on the degree of diffusion.

It also is important to note that vorticity can remain with parcels after a boundary loses its identity (its horizontal temperature gradient). The vorticity remains until dissipated by turbulence, which could conceivably be hours after a boundary ceases to exist in the temperature field depending on the mixing timescale (in turn dependent upon the degree of cloudiness and time of day). Moreover, boundaries must be viewed as three-dimensional sloping surfaces. Their intersections with the ground mark the surface boundary. Thus, anywhere there is a sloping region of temperature contrast, horizontal vorticity is generated, and this zone is typically a region that slopes from the surface boundary back over the cooler air mass.

3. The 2 June 1995 VORTEX events

a. Large-scale overview

Rapid evolution, especially of the flow, in the large-scale environment occurred on 2 June 1995. This is illustrated in Fig. 2 using the National Weather Service (NWS) sounding from Amarillo (hereafter all geographical references are in Texas unless otherwise noted; 1107 UTC), and VORTEX Cross-chain Loran Atmospheric Sounding System soundings from Texas Tech University (Lubbock; 1715, 2015, 2348 UTC). Under the assumption of homogeneity in the early morning air mass across west Texas, it appears that the entire depth of the troposphere warmed during the course of the day, with about 3°C midtropospheric warming occurring between 1200 and 0000 UTC. At Lubbock, a gradual deepening of the moist boundary layer occurred during the afternoon. CAPE, computed using the 1000-m mean near-ground parcel and the virtual temperature correction, increased from very small values in the early morning to values in the range of 1500–2000 J kg−1 during the afternoon, mainly due to heating in the boundary layer. The boundary layer mixing ratio at Lubbock was around 12 g kg−1 most of the afternoon (a short-duration drop to 10 g kg−1 occurred around 2315 UTC). At some locations near the prominent outflow boundary, mixing ratios were nearly 14–15 g kg−1, which would yield CAPE values of 2500–3000 J kg−1 (details discussed in section 4).

The evolution of the flow during the day was very pronounced. All upper winds at Amarillo at 1200 UTC were less than 22 m s−1. By 0000 UTC (3 June), winds near 250 mb had increased to 40 m s−1; at 500 mb winds increased from 20 to 42 m s−1. Details of this strengthening at Lubbock are shown in Fig. 3. Tendencies are strong at all levels between 1715 and 2348 UTC. Storm-relative helicity for observed storm motions increased from about 50 m2 s−2 to about 500 m2 s−2. The shear vector between the lowest 500 m and 5 km AGL increased from about 19 m s−1 to about 33 m s−1. All of these changes seemed to be associated with the passage of a midlevel shortwave ridge axis across the region during the morning, followed by the advance of a deep tropospheric jet into the region. Rasmussen and Blanchard (1998) show that when significant (F2 or greater damage rating) tornadoes occur, they are often associated with similar combinations of the observed SRH and CAPE near or above 2000 J kg−1. However, a major point made in that study is that much more often environments with these conditions fail to be associated with significant tornadoes. The present study would tend to indicate that mesoscale enhancements to these large-scale conditions may be important in favoring tornado occurrence.

During the late morning (1815 UTC satellite image in Fig. 4), moderate to strong thunderstorms developed in the area north of Amarillo and moved southeastward. Although weak storm rotation was observed in several of these cells and severe hail (>1.9 cm diameter) occurred, the organization was generally multicellular and dominated by outflow, consistent with the rather weak upper flow still present near midday. These storms produced copious outflow, generating an extensive boundary that advanced southwestward (2115 UTC satellite image in Fig. 5). Surface observations showing the bulk movement of this boundary and the larger-scale evolution are shown in Figs. 6 and 7.

b. Chronology of boundary evolution

The prominent outflow boundary continued to advance west and southward across west Texas during the period 1800–2100 UTC (Fig. 8). Its location was determined by tracking a narrow band of cumulus2 and stratocumulus at 15-min intervals using GOES-8 data. It was corroborated as the “armada” of VORTEX vehicles passed through the boundary just southwest of Amarillo (see section 3c) and by noting the backing of surface winds and falling temperatures in surface observations from Amarillo and Dalhart, and the Oklahoma Mesonet sites in the western Oklahoma Panhandle.

During the period 2100–2230 UTC, some parts of the boundary became obscured by cirrus anvils. However, GOES-8 and KLBB WSR-88D observations clearly show the boundary in the Texas south plains between Lubbock and Plainview (these observations are discussed in more detail in section 5). During this time, the boundary in that region slowed considerably (to a few meters per second). The portion of the boundary in the Texas Panhandle continued advancing westward and was located using surface observations at Dalhart;Clayton, New Mexico; and Oklahoma Mesonet stations in the Oklahoma Panhandle. It was found to pass through Clayton, New Mexico, between 2000 and 2100 UTC. This was consistent with positions extrapolated using the last obtainable satellite-derived motion.

In the Texas south plains (region of west Texas generally surrounding Lubbock) between 2230 and 0120 UTC, the boundary was continuously visible in KLBB WSR-88D data. Animations of reflectivity imagery indicate that the boundary was nearly stationary until shortly before 0100 UTC when it began moving northeastward. The western part of the boundary could only be located with precision at one time (2320 UTC) when mobile mesonet (Straka et al. 1996) data revealed it to be about 1 km east of Friona (details section 3c).

c. Chronology of storm evolution

Thunderstorms began forming in two areas at around 2100 UTC. Several radar echoes first formed on the boundary between Floydada and Dimmitt, while others formed from Clovis, New Mexico, northward through extreme eastern New Mexico (Fig. 9). As the afternoon and evening progressed, cells formed successively farther south through eastern New Mexico, with the mean northeastward motion bringing them into west Texas. No additional significant storms formed on the boundary after the early initiation. Most of the cells that formed near the diffuse dryline in eastern New Mexico had lifetimes of around 1 h, although several persisted for more than 3 h. Two of the storms that formed on the outflow boundary were supercells with lifetimes of over 4 h.

Cell reflectivity maxima and 25-dBZ echo outlines were plotted from KLBB WSR-88D data at 10-min intervals for the period 2100–0200 UTC. Cells were tracked, with motion verified by viewing animations, to produce the tracks shown in Fig. 9. At least four storms that formed near the dryline moved northeastward across the prominent boundary. Several other storms were too far north to be observed with the KLBB radar (Amarillo and Cannon Air Force Base, New Mexico, WSR-88D radars were not operating on this day). However, it is known that two of these supercells that formed near the dryline produced significant tornadoes within about 30 km of the estimated position of the boundary on the cool side (tornadoes 1, 3, and 4 in Table 1). On the warm side of the boundary, motions and speeds were consistently to the northeast at about 8.5–12.5 m s−1 (30–45 km h−1). A few of these supercells deviated slightly to the right, moving toward the east-northeast. By contrast, the storms that crossed the boundary or formed on the boundary and moved into the cooler air had much more deviant motion compared to the average motion on this day. Cell centroids meandered east or northeastward at 2.8–5.6 m s−1 (10–20 km h−1), with several storms becoming stationary for periods of 15–30 min. It is plausible that the difference in cell motions between supercells near the boundary and those to the south can be explained by the fact that low-level rotation was stronger near the surface in storms interacting with the boundary (see section 4), contributing to larger perturbation pressure forces (Rotunno and Klemp 1982).

Consistent with the degree of CAPE and shear present across the entire region, most of the storms that formed near the dryline and persisted for more than an hour became supercells with persistent mesocyclones at middle levels (∼4–8 km AGL). At 2200 UTC, four storms near the New Mexico border had mesocyclones; by 2300 UTC, all five storms that had moved eastward into west Texas had mesocyclones. Similarly, all of the storms that formed on the boundary and remained within about 60 km of it had persistent mesocyclones (a storm that moved fairly quickly northward into the cooler air to the north of Dimmitt at 2200 UTC had a short-lived middle-level mesocyclone only). Counting all of the middle-level (above 3 km AGL) mesocyclones that were in existence at 1-h intervals (commencing 2100 UTC), there are nearly equal numbers on the warm side (15) and the cool side (16) of the boundary.

In contrast to the widespread occurrence of supercells with middle-level (3–7 km AGL) mesocyclones across the region, low-level (1–3 km AGL) mesocyclones were not so evenly distributed geographically. For this analysis, all storms were examined at 10-min intervals using data from the 0.5° elevation base scan of the KLBB WSR-88D. Most of the storms were about 75–125 km from Lubbock, with a base scan height ranging from about 1000 to 2000 m AGL. The occurrence of low-level mesocyclones, quantified in terms of differential velocity (maximum inbound to maximum outbound radial velocity; generally spanning several azimuthal samples) and distance from the boundary, is shown in Fig. 10 (note that the location of the boundary was known to within ±3 km). Although the occurrence of middle-level mesocyclones was nearly evenly divided across the boundary, over two-thirds of the low-level mesocyclones (50/73) occurred on the cool side of the boundary. More importantly, the mesocyclones on the cool side of the boundary were clustered in a narrow region extending about 50 km into the cool air and were considerably stronger,3 in general, than the those on the warm side. Low-level mesocyclones exceeding 35 m s−1 in differential velocity were found exclusively within 50 km of the boundary on the cool side. Low-level mesocyclones were also much more persistent on the cool side of the boundary, with storms crossing the boundary exhibiting very rapid increases in rotation. Of the three storms that crossed the boundary, none had low-level mesocyclones prior to crossing the boundary and two (examined in section 4) developed strong low-level rotation and tornadoes within 30 min of crossing the boundary.

The occurrence of tornadoes on 2 June 1995 (Table 1) has been established through several means. First, Storm Data (NOAA 1995) was examined. Six reports were discarded because they could be refuted through VORTEX observations, were completely inconsistent with Doppler radar velocity and reflectivity measurements, or were clearly the same tornado viewed from two locations. Reports from experienced storm chasers were processed to verify and/or refute the other reports found in Storm Data. Two tornadoes were added based on observations by VORTEX and other storm chasers. Except for two of the tornadoes that were clearly small and weak, all of the tornadoes in this dataset featured fairly large condensation and debris clouds. It should be noted that Storm Data contained 9/10 of the actual tornado occurrences, and that 6/15 of the Storm Data reports were deemed to be false (all F0).

Tornado occurrences (except for numbers 1, 3, and 4) are plotted in circles in Fig. 9. All of the Texas tornadoes occurred within about 60 km of the surface boundary, on the cool side, with the most intense tornadoes (Friona and Dimmitt) occurring within about 15–25 km of the boundary. Because the boundary location in New Mexico is less certain, the locations of tornadoes with respect to the boundary in that region are known with less confidence, but it appears likely that these also occurred near the boundary on the cool side.

4. Evolution of storms that crossed the boundary

In this section, the evolution of two storms that crossed the prominent boundary is examined to gain further insight into the possible role that the boundary played in the intensification of low-level rotation. The Friona tornadic storm developed in the vicinity of Clovis, New Mexico, around 2300 UTC, at the southern end of a short line of supercells. The evolution of reflectivity and rotation is shown in Fig. 11. The time–height analysis of reflectivity (shown in gray) shows a deepening storm with 40-dBZ reflectivity reaching 10 km by t =1000 s (∼2317 UTC), indicative of strong updrafts. The storm crossed the boundary at approximately 2320–2325 UTC (marked B in the figure) near Bovina at which time it turned toward the right and commenced moving eastward. Just prior to this time, low-level maximum reflectivity decreased while reflectivity increased further at middle levels. This could be indicative of strengthening low-level updrafts. By 1750 s (∼2329 UTC) a deep (>8 km) core of reflectivity in excess of 65 dBZ had formed and apparently descended with time. This core was generally located just to the north of a developing mesocyclone.

Rotation, as measured by the differential velocity divided by the horizontal separation between Doppler maxima (½ of the vertical vorticity of the equivalent solid body rotation), showed a gradual increase (black contours) around 6–8 km AGL as the storm moved toward the boundary. At about the same time as the storm crossed the boundary, rotation near 2–4 km AGL increased significantly. This strong mesocyclone developed both upward and downward through the storm, with the tornado forming at about 2342 UTC (2500 s). The causes of the downward development and tornado formation are beyond the scope of this paper. However, it is interesting that intense rotation first developed above the inflow layer but below the preexisting middle-level mesocyclone. This would tend to indicate that the low-level mesocyclone intensified owing to tilting of streamwise horizontal vorticity from below the 2–4 km AGL layer (very strong radial velocity convergence existed below this developing mesocyclone implying strong stretching in the lower levels as well). Had the low-level mesocyclone been due to convergence, which is generally the strongest in the boundary layer, acting on a preexisting near-ground mesoscale vortex (e.g., Wakimoto and Wilson 1989) or preexisting vertical vorticity otherwise associated with the boundary (Maddox et al. 1980), one would expect the intensification to proceed first nearer to the ground and then develop upward with time.

One other observation is pertinent to this discussion. Just prior to the storm crossing the boundary, the KLBB WSR-88D indicated storm inflow (storm-relative flow directed toward the updraft) at 1.1 km AGL of about 30 m s−1. By 2332 UTC (1920 s), inflow on the base scan had increased to at least 42 m s−1. The intense inflow was observed by some vehicles in the VORTEX armada to cause nearly zero visibility in blowing dust and damage to power lines. Conditions were so bad that these vehicles could not proceed toward the storm and had to use an alternate route. The acceleration of the inflow occurred as the low-level mesocyclone was intensifying in the 2–4 km AGL layer. It is plausible that pressure falls in this layer, associated with the strong rotation, induced the large inflow acceleration. In turn, the horizontal stretching associated with the localized inflow would have further increased the storm-relative helicity of the inflow parcels. Details of the evolution and morphology associated with the Friona tornadic supercell will be the topic of a future analysis.

The other supercell that produced a major tornado after crossing the boundary was the storm that affected Dimmitt around 0100 UTC. A time–height analysis of vorticity has been performed for this storm as well (Fig. 12). The evolution of the Dimmitt storm is more complex than that of the Friona supercell because a sharp, very short duration, weakening of rotation commenced near the ground about 15 min after the storm crossed the boundary, followed by an abrupt reintensification leading to tornadogenesis. However, several features of the evolution are similar to the Friona event, including the reflectivity evolution. The rotation first increased at about 2–3 km AGL immediately after the boundary was crossed, with the reflectivity maximum forming aloft. Then, strong rotation spread upward and downward with time concurrent with the descent of the reflectivity maximum. Doppler velocities recorded at the KLBB WSR-88D indicated that an inflow surge also developed immediately after boundary crossing, with inflow winds increasing from ∼30 m s−1 to 38 m s−1.

5. Analysis of the boundary

In the introduction, the finding of Maddox et al. (1980) that boundaries are rich in low-level vertical vorticity, as well as the role of buoyancy gradients in producing horizontal vorticity, were discussed. The relative magnitudes of the horizontal and vertical vorticity components on 2 June 1995 can be assessed from VORTEX and conventional data. First, vertical vorticity is considered. On the meso-β scale, the presence of the boundary seemed to have little effect on vertical vorticity; the surface winds during the afternoon were unusually uniform from 110° to 130° at about 12 m s−1. However, on the meso-γ scale, the winds were perturbed by the boundary (as one would expect owing to the presence of the local solenoidal circulation).

The VORTEX armada passed through the boundary twice, first at around 1949 UTC (Fig. 13) and later at 2320 UTC (Fig. 14). Examining these transects, it is apparent that winds became fairly uniform beyond about 5 km from the boundary itself. Given the fairly straight geometry of the boundary, the flow can be simplified as two uniform flow fields, one on either side of the boundary, with a narrow transition zone between. Thus, to estimate vertical vorticity, the simplification will be introduced that the alongboundary variations in the wind are much smaller than the cross-boundary variations. Rotating the coordinate system [as in Eq. (1)] so that y is parallel to the boundary and x is normal and points toward cooler air, the equation for vertical vorticity under the foregoing assumption becomes
i1520-0493-128-1-174-e5
where υ is the y component of the wind in the rotated coordinate system.

For the early transect, southwest of Amarillo (Fig. 13), the alongboundary wind component on the cool side is about −1 m s−1, given the winds from 050° at 8 m s−1 and the boundary orientation established from satellite imagery. The observation locations have been transformed using time-to-space conversion and the known boundary motion to allow for accurate space-scale estimates. The alongboundary component on the warm side estimated from the observations east of Canyon is about 4.5 m s−1. These winds are observed over a Δx of approximately 1 × 104 m, yielding an estimated vertical vorticity of −5.5 × 10−4 s−1. Thus we find that during the early stages of the progress of this boundary, the vertical vorticity is actually negative—a finding not too surprising in that insufficient time has elapsed to allow for the development of positive vorticity through the effects of Earth’s rotation. Note that the temperature gradient magnitude is approximately 1 K km−1.

The second transect occurred near Bovina around 2320 UTC as the boundary became stationary and was being crossed by the pretornadic Friona supercell. Estimating the alongboundary wind components (Fig. 14) yields 16 m s−1 on the cool side and about 6.5 m s−1 on the warm side. Using 10 km as the horizontal length scale yields vertical vorticity of 9.5 × 10−4 s−1. Thus the vertical vorticity near the boundary, provided that these two datasets are representative, changed sign and increased considerably between 1949 and 2320 UTC. The magnitude of the temperature gradient was as large as 4 K km−1 at this latter time.

Interestingly, the latter dataset also shows that the temperature on the cool side did not increase significantly during the course of the afternoon, despite nearly clear conditions close to the boundary. However, the mixing ratio increased from about 12 g kg−1 to around 15 g kg−1 between the two observations. These changes can probably be attributed to the fact that several storms to the east of Friona rained into the near-ground air mass, and the resulting evaporation sufficiently cooled the air mass to offset heating associated with afternoon insolation north of the boundary. The combination of enhanced humidity and warm temperatures quite close to the boundary produced a maximum in equivalent potential temperature at the boundary; values reached nearly 357 K at the boundary, but were only 348 K 5 km south of the boundary, and 353 K 10 km east of the boundary. This enhancement of equivalent potential temperature along, and to a lesser extent, on the cool side of the boundary, may explain some of the observed storm intensification upon boundary crossing (also observed by Maddox et al. 1980) because of increased CAPE associated with these changes. CAPE computed from a sounding obtained east of Bovina at 2058 UTC, modified using the mobile mesonet surface observations only, is about 2370 J kg−1 south of the boundary, 2830 J kg−1 on the boundary, and 2620 J kg−1 in the cooler air 10 km to the east.

Horizontal vorticity produced by the boundary, detected in reflectivity data (Fig. 15), can be estimated using WSR-88D radar data from KLBB. To do this, an azimuth was selected at which the radar beam was nearly normal to the boundary (roughly 025°). The radial velocity data from the lowest few elevation angles were plotted in range versus height format, and contoured for two different times: 2102 UTC (Fig. 16) and 2202 UTC (Fig. 17). Because of the choice of azimuths, these analyses show the boundary-parallel component of horizontal vorticity; the magnitude of the horizontal vorticity vector must be greater than, or equal to, this single-component magnitude. However, it is likely that most of the horizontal vorticity is in this component. Further, as seen in the mobile mesonet analysis above, the low-level winds are largely parallel to the boundary at this time (Fig. 14) implying that this large horizontal vorticity vector is mainly streamwise. The horizontal vorticity vector well to the south of the boundary can be estimated from special fixed-site CLASS soundings (Fig. 3) from Lubbock, as well as VAD winds from the KLBB radar. These indicated the ambient horizontal vorticity magnitude was about 6 × 103 s−1, so the boundary augmented this value by about 75% to 10.5 × 10−3 s−1. In the mobile mesonet analyses above, it was shown that most of the wind variation associated with the boundary occurred in a band about 5 km across. Likewise, the WSR-88D cross sections show the greatest wind perturbations in a narrow band almost collocated with the enhanced reflectivity identifying the boundary itself, and that the boundary-parallel horizontal vorticity magnitude remains somewhat elevated well into the cool air mass.

Further evidence of the structure of the boundary is obtained through three fortuitously positioned soundings. These were obtained at 2348 UTC about 15 km south of the boundary at Ralls, at 2248 UTC about 15 km north of the boundary at Lockney, and at 2355 UTC about 50 km north of the boundary west of Silverton. The low-level thermodynamic soundings are shown in Fig. 18. The temperature differential across the boundary is around 4°C, and some cooling is evident to heights of around 1500 m. Low-level mixing ratios are much higher just north of the boundary, with values approaching 16 g kg−1, and the air is nearly saturated as low as about 700 m AGL. Farther north, mixing ratios are not as large, but are still higher than those found south of the boundary. The hodographs vary markedly across the boundary as well (Fig. 19). The hodograph with the largest 0–2-km shear and storm-relative helicity is the one obtained 15 km north of the boundary, with the shear magnitude being about 8.5 × 10−3 s−1. These relative spatial changes in sounding-derived horizontal vorticity are consistent with the Doppler-derived horizontal vorticity. The shear magnitude south of the boundary is about 6.3 × 10−3 s−1. Thus, for this segment of boundary (the sounding cross section is within 20 km of the Doppler analysis cross section), it appears that the horizontal vorticity is augmented about 35% on ∼20 km scales by the presence of the boundary.

Comparing the horizontal component of vorticity in the immediate vicinity of the boundary, which is at least 1 × 10−2 s−1, and the vertical, which is about 1 × 10−4 s−1 to 1 × 10−3 s−1, indicates that the horizontal component must be considered as the primary source for the enhanced low-level updraft rotation that occurs as storms pass into the cool side of the boundary. Indeed, simply tilting the quasi-horizontal vortex tubes present along the boundary would result in mesocyclone-strength vertical vorticity with or without amplification through stretching. It must be noted that the mobile mesonet winds are 3 m AGL measurements, and the winds deeper in the boundary layer are considerably stronger. If the wind directions are similar to the 3-m winds at greater heights, vertical vorticity may be several times larger at greater height. However, it is unlikely that they exceed more than about one-third of the horizontal component. Arguments could be made that the horizontal component was even stronger near the tornadic storms (about 100 km northwest of where the radar observations were obtained) due to increased baroclinity because of the outflow contributions of intervening storms. In fact, such is likely the case considering the observed temperature differential near the Dimmitt storm (21°C in the inflow sector versus 30°C south of the boundary at Lubbock).

The data collected by VORTEX on 2 June 1995 do not permit an assessment of the existence and role of low-level mesovortices (⩽10 km scale) that could have been present along the boundary in the low-level intensification of supercell rotation. These vortices have been documented by Brady and Szoke (1989), Wakimoto and Wilson (1989), and Wilczak et al. (1992) in northeastern Colorado, and Wakimoto and Atkins (1996) and Wakimoto et al. (1998) in other VORTEX cases. It can be argued that such mesovortices in general should not contain vertical vorticity as large as the horizontal vorticity generally expected along boundaries according to the discussion in section 2. Further, with the large vertical mass transport observed in the lower levels of tornadic supercells, which persists along with the low-level mesocyclone for perhaps at least several tens of minutes, a rather large mesovortex would be required to sustain the low-level mesocyclone, and the mesovortex would have to move along with the storm after it crossed the boundary. A much more plausible explanation for the observed low-level mesocyclones is the tilting of low-level, intense horizontal (quasi-streamwise) vorticity from the storm inflow.

6. Conclusions

This paper has presented evidence that out of a large number of supercells on 2 June 1995, only those that crossed, or developed on the immediate cool side of a prominent outflow boundary, produced tornadoes. The outflow boundary was shown to be a region of enhanced CAPE, enhanced horizontal (quasi-streamwise) vorticity, and modestly enhanced vertical vorticity.

Maddox et al. (1980) provided comprehensive documentation that, in some cases, intense tornadoes were associated with storms crossing mesoscale boundaries, with tornado longevity related to the amount of time the parent storm persisted in the favorable air mass on the immediate cool side of the boundary. Although the Maddox et al. analysis related tornado occurrence to enhanced convergence and vertical vorticity near the boundaries, this paper has examined another potential source of storm rotation near boundaries: augmented horizontal vorticity. Both studies identified enhanced CAPE near the boundaries that owed to various mechanisms.

In the present study, synoptic conditions were favorable for supercells by any commonly used measure. As shown in section 2, by late afternoon the soundings at Lubbock had CAPE of 1500–2000 J kg−1 and storm-relative helicity in excess of 500 m2 s−2. There is no evidence that these favorable conditions did not exist throughout the warm, humid air mass ahead of the dryline and south of the boundary, in which about 15 supercells occurred. Yet, tornado occurrence was confined to a much smaller region along and on the cool side of the prominent boundary.

Maddox et al. argued that boundaries were effective in those cases in which “macroscale conditions did not favor large family outbreaks of storms.” Based on the present analysis, as well as the summary work of Markowski et al. (1998a), the following hypothesis4 is permitted: in general, significant tornadoes require augmentation of storm-relative helicity (or perhaps the combination of helicity and CAPE or other quantities) beyond what is usually thought to be associated with environments conducive to tornado outbreaks, and this augmentation occurs on the meso-β scale, especially in conjunction with baroclinic boundaries. In other words, it is plausible that the large scale rarely provides sufficient vorticity and CAPE for significant tornadoes. There is no clear evidence in previous case studies of the occurrence of significant tornadoes without boundaries being present, even in major outbreaks. Future research should attempt to determine whether or not significant tornadoes can occur without local augmentation of horizontal vorticity in the near-storm environment to values significantly larger than previously thought required. Although the data in this study strongly indicate that the most important role of the outflow boundary was to augment horizontal vorticity, the roles of locally increased CAPE as well as boundary layer humidity also merit further research.

Acknowledgments

We are grateful for the encouragement and grant support of Dr. Stephen Nelson of the National Science Foundation. Partial support for this research was provided by the National Science Foundation through Grants EAR95-12541, ATM-9120009, and ATM-9617318. Additional support was provided by NOAA through a Presidential Early Career Award for Scientists and Engineers. We wish to acknowledge the contribution of Dr. Charles Doswell III through several helpful conversations and a review of this manuscript. We also acknowledge Dr. Robert Maddox for his patience in reviewing two early versions of this paper. Ms. Janelle Janish processed WSR-88D data for its inclusion herein. VORTEX data were collected through the untiring efforts of over 100 participants, many of them volunteers, without which the project could not have succeeded. Mike Smith is acknowledged for several stimulating discussions regarding the potential importance of boundaries to tornadic supercells. Jim Purdom and John Weaver are acknowledged for useful discussions regarding the role and nature of boundaries in general, as well as the behavior of the 2 June 1995 boundary.

REFERENCES

  • Brady, R. H., and E. J. Szoke, 1989: Case study of a nonmesocyclone tornado development in northeast Colorado: Similarities to waterspout formation. Mon. Wea. Rev.,117, 843–856.

  • Davies, J. M., C. A. Doswell III, D. F. Burgess, and J. F. Weaver, 1994: Some noteworthy aspects of the Hesston, Kansas, tornado family of 13 March 1990. Bull. Amer. Meteor. Soc.,75, 1007–1017.

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

  • ——, D. Burgess, and M. Foster, 1990: Test of helicity as a tornado forecast parameter. Preprints, 16th Conf. on Severe Local Storms, Kananaskis Park, AB, Canada, Amer. Meteor. Soc., 588–592.

  • Dutton, J. A., 1976: The Ceaseless Wind. An Introduction to the Theory of Atmospheric Motion. McGraw-Hill, 579 pp.

  • Maddox, R. A., L. R. Hoxit, and C. F. Chappell, 1980: A study of tornadic thunderstorm interactions with thermal boundaries. Mon. Wea. Rev.,108, 322–336.

  • Markowski, P. M., E. N. Rasmussen, and J. M. Straka, 1998a: The occurrence of tornadoes in supercells interacting with boundaries during VORTEX-95. Wea. Forecasting,13, 852–859.

  • ——, J. M. Straka, E. N. Rasmussen, and D. Dowell, 1998b: Observations of low-level baroclinicity generated by anvil shadows. Mon. Wea. Rev.,126, 2942–2958.

  • National Oceanic and Atmospheric Administration, 1995: Storm Data. Vol. 37, No. 5, 221 pp.

  • Purdom, J. F. W., 1976: Some uses of high resolution GOES imagery in the mesoscale forecasting of convection and its behavior. Mon. Wea. Rev.,104, 1474–1483.

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

  • ——, J. M. Straka, R. P. Davies-Jones, C. A. Doswell III, F. H. Carr, M. D. Eilts, and D. R. MacGorman, 1994: The Verifications of the Origins of Rotation in Tornadoes Experiment: VORTEX. Bull. Amer. Meteor. Soc.,75, 997–1006.

  • 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 M. L. Weisman, 1988: A theory for strong, long-lived squall-lines. J. Atmos. Sci.,45, 463–485.

  • Straka, J. M., E. N. Rasmussen, and S. E. Frederickson, 1996: A mobile mesonet for fine-scale meteorological observations. J. Atmos. Oceanic Technol.,13, 921–936.

  • Wakimoto, R. M., and J. W. Wilson, 1989: Nonsupercell tornadoes. Mon. Wea. Rev.,117, 1113–1140.

  • ——, and N. T. Atkins, 1996: Observations on the origins of rotation:The Newcastle tornado during VORTEX-94. Mon. Wea. Rev.,124, 384–407.

  • ——, C.-H. Liu, and H. Cai, 1998: The Garden City, Kansas, storm during VORTEX-95. Part I: Overview of the storm’s life cycle and mesocyclogenesis. Mon. Wea. Rev.,126, 372–392.

  • Wilczak, J. M., T. W. Christian, D. E. Wolfe, R. J. Zamora, and B. Stankov, 1992: Observations of a Colorado tornado. Part I: Mesoscale environment and tornadogenesis. Mon. Wea. Rev.,120, 497–520.

  • Wood, V. T., and R. A. Brown, 1997: Effects of radar sampling on single-Doppler velocity signatures of mesocyclones and tornadoes. Wea. Forecasting,12, 928–938.

Fig. 1.
Fig. 1.

Conceptual model of flow near a boundary. From Maddox et al. (1980 their Fig. 2).

Citation: Monthly Weather Review 128, 1; 10.1175/1520-0493(2000)128<0174:TAOSTW>2.0.CO;2

Fig. 2.
Fig. 2.

Skew T–logP diagram of soundings from Amarillo at 1103 UTC (solid), Texas Tech University (Lubbock) at 2015 UTC (thin broken) and 2348 UTC (thick broken). CAPEs are shown in the legend.

Citation: Monthly Weather Review 128, 1; 10.1175/1520-0493(2000)128<0174:TAOSTW>2.0.CO;2

Fig. 3.
Fig. 3.

Hodographs from Texas Tech University (Lubbock) from 1715 UTC (thick solid), 2015 UTC (thin broken), and 2348 UTC (thick broken). Velocities in m s−1. Heights in km AGL. Values in the legend are SRH (m2 s−2) estimates spanning the range of observed storm motions.

Citation: Monthly Weather Review 128, 1; 10.1175/1520-0493(2000)128<0174:TAOSTW>2.0.CO;2

Fig. 4.
Fig. 4.

GOES-8 visible satellite image for 1815 UTC 2 Jun 1995.

Citation: Monthly Weather Review 128, 1; 10.1175/1520-0493(2000)128<0174:TAOSTW>2.0.CO;2

Fig. 5.
Fig. 5.

GOES-8 visible satellite image for 2115 UTC 2 Jun 1995.

Citation: Monthly Weather Review 128, 1; 10.1175/1520-0493(2000)128<0174:TAOSTW>2.0.CO;2

Fig. 6.
Fig. 6.

Surface observations from 1800 UTC 2 Jun 1995. Full wind flags represent 5 m s−1; half flags 2.5 m s−1. Temperatures and dewpoints in °C. Sites in the Oklahoma panhandle are Oklahoma mesonet sites; all others NWS/FAA. Approximate boundary location marked with the bold line. Additional towns mentioned in the text are shown with small dots.

Citation: Monthly Weather Review 128, 1; 10.1175/1520-0493(2000)128<0174:TAOSTW>2.0.CO;2

Fig. 7.
Fig. 7.

As in Fig. 6. but for 2100 UTC. Site marked P1 is an observation from the Probe-1 VORTEX vehicle.

Citation: Monthly Weather Review 128, 1; 10.1175/1520-0493(2000)128<0174:TAOSTW>2.0.CO;2

Fig. 8.
Fig. 8.

Isochrone analysis of the leading edge of the narrow cloud line associated with the prominent boundary. Position every 30 min, based on 15-min GOES-8 visible image analysis. Unlabeled filled circles are locations of communities named on other maps herein.

Citation: Monthly Weather Review 128, 1; 10.1175/1520-0493(2000)128<0174:TAOSTW>2.0.CO;2

Fig. 9.
 Fig. 9.

Echo centroid tracks for five 1-h periods depicted with the various line types shown in the legend. Thick lines represent boundary positions according to the symbols in the legend, and apply to the start of the 1-h periods. Numbers in circles correspond to tornadoes listed in Table 1.

Citation: Monthly Weather Review 128, 1; 10.1175/1520-0493(2000)128<0174:TAOSTW>2.0.CO;2

Fig. 10.
Fig. 10.

Graph of low-level mesocyclone differential velocity (difference between inbound and outbound Doppler radial velocity peaks on base scan) vs distance from the prominent boundary (positive toward cooler air). Each dot represents one detection; all data at 10-min intervals are included.

Citation: Monthly Weather Review 128, 1; 10.1175/1520-0493(2000)128<0174:TAOSTW>2.0.CO;2

Fig. 11.
Fig. 11.

Time–height diagram of the maximum reflectivity (dashed lines) and mesocyclone strength (½ the equivalent solid-body vertical vorticity; solid lines) for the Friona, TX storm. The storm crossed the prominent boundary during the period denoted with the heavy bar labeled B. The tornado occurred during the period shown with the heavy arrow. Rotation contoured and labeled in black (×10−4 s−1). Reflectivity contoured and labeled in gray (dBZ). Gray shading denotes time and heights at which rotation exceeded 2 × 10−2 s−1.

Citation: Monthly Weather Review 128, 1; 10.1175/1520-0493(2000)128<0174:TAOSTW>2.0.CO;2

Fig. 12.
Fig. 12.

As in Fig. 11 except for the Dimmitt, TX, tornadic storm, and with X denoting a reflectivity maximum and N a minimum.

Citation: Monthly Weather Review 128, 1; 10.1175/1520-0493(2000)128<0174:TAOSTW>2.0.CO;2

Fig. 13.
Fig. 13.

Analysis of mobile mesonet data (subset of observations from vehicle P1) during a transect of the boundary southwest of Amarillo, TX. Heavy line represents the boundary with the orientation established through analysis of GOES-8 data. Black lines are major highways; gray lines are other roads. Wind flags drawn with full barb = 5 m s−1; half barb = 2.5 m s−1. Temperature in degrees C; mixing ratios in 0.1 g kg−1. True north is toward the top of the map.

Citation: Monthly Weather Review 128, 1; 10.1175/1520-0493(2000)128<0174:TAOSTW>2.0.CO;2

Fig. 14.
Fig. 14.

As in Fig. 13 except for a boundary transect near Bovina, TX.

Citation: Monthly Weather Review 128, 1; 10.1175/1520-0493(2000)128<0174:TAOSTW>2.0.CO;2

Fig. 15.
Fig. 15.

WSR88D reflectivity at 2102 UTC. Boundary and azimuth of the cross section in Fig. 16 is indicated with the heavy arrow. Reflectivity in 3-dB increments (increasingly darker shades) ranging from 0 to 15 dBZ.

Citation: Monthly Weather Review 128, 1; 10.1175/1520-0493(2000)128<0174:TAOSTW>2.0.CO;2

Fig. 16.
Fig. 16.

Cross section of Lubbock WSR-88D radial velocity at 2102 UTC (contoured; ms−1). Arrows are position of radar beams. Boundary-parallel component of horizontal vorticity s−1 represented by plotted values at the top of the figure, computed from interpolated radial velocity at 2 km AGL minus velocity at 0.5°, divided by the actual height difference. Location of radar fine line represented with heavy black bar between 45- and 49-km range.

Citation: Monthly Weather Review 128, 1; 10.1175/1520-0493(2000)128<0174:TAOSTW>2.0.CO;2

Fig. 17.
Fig. 17.

As in Fig. 16 but for 2202 UTC.

Citation: Monthly Weather Review 128, 1; 10.1175/1520-0493(2000)128<0174:TAOSTW>2.0.CO;2

Fig. 18.
Fig. 18.

Low-level skew T–logp diagrams of soundings taken at various locations with respect to the prominent boundary.

Citation: Monthly Weather Review 128, 1; 10.1175/1520-0493(2000)128<0174:TAOSTW>2.0.CO;2

Fig. 19.
Fig. 19.

The 0–4 km AGL hodographs for the same locations and times as in Fig. 18. Wind speeds in m s−1. Surface points from surface analysis. Heights in km AGL.

Citation: Monthly Weather Review 128, 1; 10.1175/1520-0493(2000)128<0174:TAOSTW>2.0.CO;2

Table 1.

The 2 June 1995 tornadoes.

Table 1.
1

This effect is maximum at one-fourth of the inertial period, or 0.5 πf−1, where f is the magnitude of the Coriolis acceleration. This is approximately 5.5 h at the latitude of this case study. Later, the Coriolis acceleration would turn the flow into the line-normal direction, toward the cooler air.

2

Comparing satellite with surface and WSR-88D reflectivity data indicates that the cloud band tracked on satellite was typically to the rear (toward the cool side of the boundary) compared to the surface location.

3

Note that some error is inherent in these estimates because of the problems of sampling vortex flow at great range (Wood and Brown 1997). However, the signal here is fairly obvious and likely not an artifact of sampling problems.

4

This hypothesis is presented to motivate further research, and not as a conclusion derived from the analysis of this paper. An anonymous reviewer of this manuscript provided a list of violent tornadoes that the reviewer thought occurred without the presence of a boundary, and the reviewer may be correct. However, based on the analysis herein, it is arguable that boundaries cannot be detected in many historical datasets because they lack mesonetworks, WSR-88D, and high space–time resolution satellite imagery.

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