The Vertical Structure of a Tornado near Happy, Texas, on 5 May 2002: High-Resolution, Mobile, W-band, Doppler Radar Observations

Transcription

1 2325 The Vertical Structure of a Tornado near Happy, Texas, on 5 May 2002: High-Resolution, Mobile, W-band, Doppler Radar Observations HOWARD B. BLUESTEIN * AND CHRISTOPHER C. WEISS School of Meteorology, University of Oklahoma, Norman, Oklahoma ANDREW L. PAZMANY Microwave Remote Sensing Laboratory, Department of Computer and Electrical Engineering, University of Massachusetts Amherst, Amherst, Massachusetts (Manuscript received 18 November 2003, in final form 19 April 2004) ABSTRACT A mobile, W-band Doppler radar scanned, at close range, portions of a tornado near Happy, Texas, on 5 May Simultaneous boresighted video images were also recorded, which facilitated correlating the radar-observed features of the tornado with its visual features. Range height indicators (RHIs) of radar reflectivity and Doppler velocity were collected that detail, with high spatial resolution, aspects of the vertical structure of the tornado near the ground. Most of the RHIs showed a column of a weak-echo hole from about 60 m above the ground up to the top of the domain at m above the ground; the hole was roughly 40% broader about 100 m above the ground as it was above, resulting in a characteristic pear-shaped vertical cross section of reflectivity. In this tornado, the condensation funnel was much narrower than the width of the weak-echo hole; the visible debris cloud near the ground was approximately just as wide as the hole above 150 m. The mean depth of the debris cloud was around 200 m. The vertical structure of the Doppler-velocity field exhibited a narrow band of high wind speeds about m above the ground, consistent with airflow inward toward and cyclonically about the tornado. Possible reasons for the observed structure of the tornado are offered. 1. Introduction The vertical variation of wind near the ground in a tornado is of great interest to theoreticians and structural engineers. Since wind-induced forces are generally proportional to the square of the wind speed (e.g., Liu 1993), even relatively small variations in wind speed could result in significant variations in damage potential. It is known from experiments with laboratory models and numerical simulations of laboratory models (Church et al. 1979; Rotunno 1979) and, more recently, from numerical simulations of tornado-like vortices using a large eddy simulation model (Lewellen and Lewellen 1997; Lewellen et al. 2000) that the depth of the bound- * Additional affiliation: Mesoscale and Microscale Meteorology Division, National Center for Atmospheric Research, Boulder, Colorado. The National Center for Atmospheric Research is sponsored by the National Science Foundation. Corresponding author address: Prof. Howard B. Bluestein, School of Meteorology, University of Oklahoma, 100 E. Boyd, Rm. 1310, Norman, OK hblue@ou.edu ary layer and corner regions (where most of the surface inflow and the strongest rising motion are expected to be found) and the character of the wind field in a tornado depend to a large extent on the swirl ratio. The swirl ratio S rm/2q 0 0 /w, (1) pertains to a laboratory vortex simulator in which Q is the volume rate of flow of an updraft that is forced through a hole of radius r 0, 0 is the azimuthal wind component at the outer edge of the updraft hole, M 0 r 0, so that 2 M is the circulation at the outer edge of the updraft, and w is the mean vertical velocity through the updraft hole (Davies-Jones et al. 2001). The swirl ratio is a measure of the relative amount of ambient vorticity compared to the amount of ambient convergence. In a vortex embedded in a flow having a low swirl ratio, for example, there is one circulation cell in the vertical plane cutting through the center, and the depth of the surface inflow layer is only a small fraction of the core radius (Fig. 1). The depth of the surface inflow layer may be only around m for most tornadoes embedded in a flow characterized by a low swirl ratio, and the height of maximum wind speeds may be above or near the top of many dwellings and 2004 American Meteorological Society

2 2326 MONTHLY WEATHER REVIEW VOLUME 132 FIG. 1. Instantaneous vertical cross section of a simulated tornado embedded in a flow having a low swirl ratio, using a large-eddy simulation model. Normalized absolute azimuthal velocity (grayscale) and speed of the wind vector in the r z plane (arrows are interpolated onto a uniform grid for clarity; maximum length corresponds to a ratio of the wind speed in the r z plane to the core speed of 3.8); V c and r c denote the wind speed and radius of the core; x in Lewellen et al. (2000) is denoted as r in this paper. From Fig. 5 in Lewellen et al. (2000). vegetation. Winds measured near the ground or inferred from damage may not represent the highest winds in the tornado. Measurements of the wind field near the surface in tornadoes with accompanying visual documentation are relatively few. Attempts to determine the wind field near the ground by photographically analyzing debris movies or videos are hampered by the difficulty in acquiring them and the impossibility of seeing inside opaque debris clouds. The heights above ground of the maximum wind speeds in tornadoes based on photogrammetric analysis of debris movies range from 15 to 200 m; most of the estimates of maximum wind are at altitudes under 100 m (Bluestein and Golden 1993). Perhaps the most characteristic feature of a tornado as seen by radar is the weak-echo hole, which appears somewhat like a small-scale analog to the radar eye in a hurricane. A weak-echo hole was first noted in the refereed literature by Fujita (1981) in images from a National Weather Service Weather Surveillance Radar (WSR-57) located about 6 km from a tornado on 3 June 1980, near Grand Island, Nebraska. Wakimoto and Martner (1992), using two fixed-site X-band Doppler radars whose antennas had a half-power beamwidth

3 2327 of 0.8, produced, from volume sector scans at constant elevation angle, vertical cross sections of radar reflectivity and Doppler velocity through a nonsupercell tornado in Colorado; this tornado was probed from a range of just over 20 km, on 2 July 1987 during the Convection Initiation and Downburst Experiment (CINDE) (Wilson et al. 1988). Simultaneous photographs were taken from the site of one of the radars, thus permitting an examination of the structural relationship among the reflectivity, Doppler wind, and visual features associated with the tornado. A weak-echo hole was noted in this tornado also. Wakimoto et al. (1996), using data collected by an airborne Doppler radar [Electra Doppler Radar (EL- DORA)], identified a weak-echo hole coincident with a tornado, at a range of 25 km, near Friona, Texas, on 2 June 1995 and also at a range of approximately 25 km, near Kellerville, Texas, on 8 June during the Verification of the Origins of Rotation in Tornadoes Experiment (VORTEX) (Rasmussen et al. 1994). These weak-echo holes extended up to near the top of the storms. Portable and mobile Doppler radars have been used to probe the wind field of tornadoes near the ground at very close range, often well within 5 10 km (Bluestein et al. 1993; Wurman et al. 1996; Bluestein and Pazmany 2000; Wurman and Gill 2000; Wurman 2002; Bluestein et al. 2003a,b). Simultaneous photographic and boresighted video documentation were available for some of the datasets collected. Measurements made with the radars in many tornadoes indicate that the center of a tornado is almost always characterized by a weak-echo hole. However, the existence of a weak-echo hole does not necessarily mean that there is tornado. Wurman and Gill (2000) showed vertical cross sections of radar reflectivity and Doppler velocity, constructed from volume sector scans at constant elevation angle, through a tornado, at about 3-km range, near Dimmitt, Texas, on 2 June 1995 during VORTEX. They used data from the Doppler on Wheels (DOW), an X- band mobile radar whose antenna at the time had a halfpower beamwidth of 1.2. The weak-echo hole seen in the cross sections did not make it all the way down to the ground, but terminated about m above the ground. Resolving the radar reflectivity and Doppler-velocity fields as close as possible to the ground requires a very narrow radar beam and a relatively weak sidelobe pattern. To achieve even higher spatial resolution than afforded by the existing fixed-site radars and mobile X- band radars, a truck-mounted W-band (3-mm wavelength, 95-GHz frequency) mobile radar has been used to probe tornadoes (and dust devils) in the southern plains (Bluestein and Pazmany 2000; Bluestein et al. 2003a,b; Bluestein et al. 2004). The main advantage of this radar is that its half-power beamwidth is only 0.18, so that very high resolution is achieved for tornadoes at close range. Its disadvantage is that attenuation is extreme in heavy precipitation, so that the useful range is often severely restricted. On 5 May 2002, a crew from the University of Oklahoma (OU) and the University of Massachusetts (UMass), using the W-band radar, collected data in a tornado near Happy, Texas, at ranges of km. The dataset is noteworthy because there were some scans in vertical planes [range height indicators (RHIs) at ranges of km] cutting through the tornado near the ground and below cloud base (the first ever made with the W-band radar), thus affording perhaps the highest resolution possible to date in the vertical of the reflectivity and Doppler wind fields in a tornado near the ground. (When volume sector scans are taken, data have to be interpolated to synthesize a vertical cross section, and some resolution is therefore lost.) Furthermore, the tornado passed over a relatively flat surface with low vegetation, so that it was possible for the radar to scan very close to the ground. For the first time, boresighted simultaneous video documentation was available to correlate precisely radar-observed features with visual features. The alignment of the boresighted video camera was marked by locating the top of a chimney tower 2 km away from the radar and then marking its location on the video monitor. The main purposes of this paper are to describe the vertical variation of radar reflectivity and Doppler velocity in the Happy, Texas, tornado and to correlate the radar data with visual features. 2. Data collection and processing a. Description of the radar The characteristics of the W-band radar, which was designed and built at UMass, are detailed in Bluestein and Pazmany (2000). The pulses, whose lengths were 30 m, were gated every 15 m, so that through this oversampling mode of data collection, data in overlapping pulse volumes every 15 m were collected. [It was possible to have used the shorter pulse length of 15 m as in Bluestein et al. (2003a,b), but since only 250 range gates could be processed for each beam, a 3.75-km window would have had to have been imposed. In real time, as the tornado was approaching, it was not possible to select the correct window to view the entire tornado quickly enough. As will be shown soon, attenuation was significant, so that the choice of a longer pulse length fortuitously increased the sensitivity.] At the range of the tornado when RHIs were collected, the azimuthal resolution varied from 10 to 20 m; the pulse volume (width height depth) when RHIs were collected was therefore approximately 15 m 15 m 30 m. Doppler velocity data were available both from a standard pulse-pair algorithm and from polarization diversity pulse-pair processing (PDPP) (Pazmany et al. 1999). The advantage of the latter is that the maximum unambiguous Doppler velocity is 79 m s 1, which is sufficiently high so that wind data are not usually aliased

4 2328 MONTHLY WEATHER REVIEW VOLUME 132 FIG. 3. Photograph of the Happy, Texas, tornado at approximately 1847 CDT, 5 May 2002, from a vantage point 7.2 km east of Happy, Texas. The view is to the west along farm-to-market road (FM) The W-band radar and the first author are visible near the bottom, center. Photograph copyright by M. Kramar. FIG. 2. Approximate path of the Happy, Texas, tornado (estimated from locations of Doppler-vortex signatures and storm reports) superimposed on a road map; the location of the W-band radar is indicated by the dot. The approximate times (CDT) of the tornado are indicated at locations along its track. (in the standard pulse-pair mode the maximum unambiguous velocity was only 8 ms 1 ). The disadvantages of the PDPP technique are that it is noisier than ordinary pulse-pair processing and does not work at ranges closer than about 1.5 km. PDPP does not work at close range, as a result of the finite polarization isolation of the switch network and antenna. b. Data collection Data were collected beginning when the tornado was mature (Golden and Purcell 1978) and passing through and east of Happy, Texas, located 7.2 km to the westsouthwest of the radar. The tornado inflicted extensive damage ( as it moved through the town at about 1945 central daylight time (CDT; all times given in CDT) and it continued to move east-northeastward, in the general direction of the radar (Fig. 2). Damage reported included several homes destroyed and a roof blown off a church; three people were killed and four were injured. Before it struck Happy, Texas, the tornado appeared to be wider and more intense (T. Marshall 2002, personal communication) than it appeared to be later when it was being probed by the W-band radar (Fig. 3). The spatial limits of scans in azimuth and elevation were set electronically by the radar operator, who used boresighted video images as a guide. Eight low-elevation sector scans at approximately constant elevation angle, just above the ground, were collected first when the tornado was km away. A series of eight RHIs were then collected just to the right of, left of, and through the center of the condensation funnel, when the tornado was km away. Finally, four more low-elevation-angle sector scans were collected as the tornado got as close as 1.1 km away and was dissipating, after it had crossed the road the first two authors were on (Figs. 2 and 3). The tornado disappeared on the right side of the road (north of the road), while its parent circulation at cloud base continued to move to the eastnortheast. Because the visible horizon in the boresighted video was tilted, it was determined that the radar truck was not leveled, so that constant elevation scans were not exactly parallel to the ground and RHIs were not exactly in the vertical plane: The radar platform was tilted approximately 5 to the right, or to the north, with height when the radar was looking approximately to the west. It was fortunate that the RHIs were collected at km range. Beyond about 4 km in range, severe attenuation from heavy precipitation limited the areal coverage around the tornado; within 1.6-km range, it was not possible to use the PDPP technique to compute Doppler velocities and it was not possible to unfold the severely aliased standard pulse-pair velocity data. The most significant problem that occurred while data were collected was that elevation angles (for the RHIs) and relative azimuth angles (for the sector scans) were not recorded as a result of a software malfunction. Consequently, the known (constant) scan rate of the antenna and the simultaneous boresighted video documentation were used to deduce the elevation and relative azimuth angles of each beam. Rays collected when the antenna was not scanning or when it was pointed into the ground were, for the most part, discarded. Owing to small uncertainties in the pointing direction, the data collected during RHIs could be shifted slightly upward or down-

5 2329 FIG. 4. Schematic depicting the sector scan of the W-band radar at 1846:38 CDT, 5 May 2002, through the Happy, Texas, tornado, from the location noted in Figs. 2 and 3. The view along the road depicted is to the west. ward; it was generally obvious when this was the case because the bottom of the tornado appeared to be shifted in the vertical with respect to the bottom seen in other scans. c. Data processing Data were processed at UMass and inspected at OU using SOLO software (Oye et al. 1995), which was developed at the National Center for Atmospheric Research (NCAR). The boresighted video documentation and mediumformat (70 mm) color transparencies taken by the first author were used to estimate the dimensions of the funnel cloud, debris cloud, etc. (cf. Fig. 3) using photogrammetric analysis. The distance to the tornado was estimated from the documentation of the times of each image and the known range to the tornado based on the radar data, along with the corresponding time of each radar scan. 3. Discussion of the data analyses a. Sector scans at approximately constant elevation angle Sector scans cut through the top of the visible debris cloud about m above the ground when it was km away (Fig. 4). At these and longer ranges, the data collected were severely attenuated by precipitation (Fig. 5, top left); the far side of the radar echoes exhibited, for the most part, a characteristic monotonic drop-off in reflectivity with range, and heavy precipitation was experienced soon at the radar truck. The sector-scan data are still useful, however, because they provide some information on the diameter of the tornado s features in the reflectivity field (such as its weak-echo eye) and wind field (such as its core diameter; the core is the area of approximate solid-body rotation, characterized by nearly uniform vertical vorticity and marked approximately by the area lying just inside the dipole in Doppler-velocity extrema of its cyclonic vortex signature) relative to the diameter of its debris cloud and condensation funnel, which will be useful later in interpreting the RHIs. The center of the tornado was marked by a m-wide, elliptically shaped hole (the former width is that of the minor axis, which was normal to the radar beam). Since the width of the hole normal to the radar beam was slightly wider than the width of the visible debris cloud (260 m), which extended up to about 190 m above the ground (Fig. 4), the ring of higher reflectivity around the tornado most likely was not associated with the visible debris cloud itself. It is possible, however, that the width of the radar-echo hole was exaggerated slightly, owing to attenuation. A cyclonic vortex signature (Brown and Wood 1991; red/yellow-coded receding Doppler velocities on the right, green/blue approaching Doppler velocities on the left) was found near the inner edge of the ring of higher reflectivity around the tornado (Fig. 5, top right). As a result of the intense attenuation, PDPP velocity data beyond about 4.6-km range were dominated by noise. Since the ordinary pulse-pair processed data were much less noisy (Pazmany et al. 1999) and exhibited an obviously aliased, but coherent cyclonic vortex signature (Fig. 5, bottom left), they were unfolded using the PDPP velocity data as boundary constraints (Fig. 5, lower right). Another cyclonic vortex signature, which was weaker (yellow-coded receding velocities on the right, green approaching velocities on the left), was seen to the left of the tornado (Fig. 5, lower right; marked by an arrow). In analyses of data for other tornado cases, the coexistence of a tornado with weaker vortices, along the rear-flank gust front, have also been noted (e.g., Bluestein et al. 2003b). Highest approaching Doppler velocities were around 40ms 1 ; highest receding velocities were around 10 ms 1. From the series of sector scans, it was estimated that the component of motion of the tornado along the line of sight of the radar was 13 m s 1, toward the radar, or from west-southwest to east-northeast. Therefore subtracting the radial component of tornado motion yields an estimate of peak azimuthal wind speeds in the tornado vortex itself of 27 and 23 m s 1, respectively. (It was not possible to determine the component of motion of the tornado normal to the line of sight because the azimuth positions in each scan are known only in a relative sense. In a qualitative sense, the tornado motion normal to the line of sight was to the north and less in magnitude than the along-the-line-sight component.) It is possible that even higher Doppler velocities were undetected, owing to attenuation; the width of the vortex signature in Doppler velocity (450 m) was therefore probably an overestimate. However, since azimuthal

6 2330 MONTHLY WEATHER REVIEW VOLUME 132 FIG. 5. Portion of the sector scan at 1846:38 CDT (time given here and henceforth for the beginning of each scan), 5 May Range markings are shown every 500 m. Color-coded scales for reflectivity (dbz) and Doppler velocity (m s 1 ) are shown below each panel; positive (negative) Doppler velocities denote motion away from (toward) the radar; (top left) radar reflectivity; (top right) PDPP Doppler velocity; (lower left) ordinary pulse-pair processed Doppler velocity; (lower right) unfolded ordinary pulse-pair processed Doppler velocity. The arrow in the lower-right panel points to a weak cyclonic vortex signature, to the left of the tornado. The tornado was located to the west-southwest of the radar. wind velocities in tornadoes typically fall off rapidly just outside the core (e.g., Bluestein and Pazmany 2000; Wurman and Gill 2000; Wurman 2002; Bluestein et al. 2003a), the core diameter was probably not too much less than m. Although we cannot be certain, then, in relating the core diameter to the diameter of the debris cloud, we can conclude with much higher certainty that the width of the condensation funnel (about 70 m or less) was much less than the width of the core of the tornado. Such a conclusion is significant, because the scanning procedure used to collect RHIs made use of the location of the visible condensation funnel: Scans taken just to the left and right of the condensation funnel therefore were likely inside the core of the tornado and missed the highest possible azimuthal and Doppler wind speeds. It is noted, however, that at 1848:44 CDT (all times given are for the beginning of the scan), the core of the tornado was sampled inadvertently while the antenna was being repositioned between successive RHI scans. In this scan, the antenna motion was intermittent, thus precluding the reconstruction of a true sector-scan image. However, as the radar beam cut across the tornado near the ground, maximum Doppler velocities of 60 (20) ms 1 in the approaching (receeding) direction, to the left (right) side of the tornado, were noted. b. Vertical cross sections through the tornado The high-resolution RHIs cutting through the tornado are the unique aspect of the Happy, Texas, dataset. Since the tornado condensation was often tilted and the radar platform was always tilted slightly, the RHIs do not make perfectly vertical cross sections. In many instances the RHI tilts half the distance across the condensation funnel from the bottom to the top of the RHI (Fig. 6); most of the RHIs during these scans extended from one of the vertical edges of the condensation funnel to the

7 2331 FIG. 6. Schematics depicting the RHI scans (thick solid lines) of the W-band radar on 5 May 2002, through the Happy, Texas, tornado, from the location noted in Figs. 2 and 3, at the times (a) 1848:29, 1849:53, and 1850:12; (b) 1848:52; (c) 1849:18; and (d) 1848:15, 1849:02, and 1849:35 CDT. In (a), (b), and (c), the scans at low-elevation angle passed through the left edge and to the left of the left edge, right edge, and center of the tornado condensation funnel, respectively; in (d) the scans tilled from an edge of the condensation funnel into the middle. Arrows indicate direction of scan; condensation funnel is outlined. The view is to the west-southwest of the radar. center. From Fig. 4, it can be seen that since the radar platform is tilted from the vertical by 5, a beam at the left edge of the condensation funnel near the ground, which is scanned vertically, could end up, at cloud base, about 45 m to the right of the left edge of the funnel, which is a considerable fraction of its 70-m width at cloud base. In those RHIs that appear to cut vertically through the condensation funnel, it is possible that the tornado is tilted in the plane of the RHI itself. In some RHIs the scan did not extend all the way to the ground. None of the RHIs exhibited very high velocities (Figs. 7 10), most likely because none of the vertical cross sections cut across the core, which was farther from the center of the condensation funnel. Figures 7 (1848:29, 1849:53, and 1850:12 CDT), 8 (1848:52 CDT), and 9 (1849:18 CDT) show examples of scans that at low elevation cut through the left edge, right edge, and center, respectively (cf. Fig. 6), of the tornado condensation funnel. In Fig. 10 the rest of the RHIs, which tilted from an edge of the condensation funnel into the middle (1848:15, 1849:02, and 1849:35 CDT), are displayed. The RHIs of radar reflectivity all show evidence of attenuation; that is, the tornado appears to be embedded within a region in which the reflectivity decreases with range. We now highlight and summarize features seen in the reflectivity and Doppler-velocity cross sections (Figs. 7 10). The center of the tornado was coincident with a column of echo-free hole, that, above about 150-m altitude was approximately 250 m in diameter (Fig. 11). The width of the hole was defined subjectively on the basis of both the reflectivity field and Doppler-velocity field. The width of the hole was determined by mea-

8 2332 MONTHLY WEATHER REVIEW VOLUME 132 FIG. 7. Vertical cross sections of radar reflectivity (dbz) and ground-relative Doppler velocity (m s 1 ) from the W-band Doppler radar (the radar is located to the far left in this figure) at the locations noted in Figs. 2 and 3, on 5 May 2002, for the scans depicted in Fig. 6. The cross section cuts from the east-northeast (left-hand side) to the west-southwest (right-hand side), through the left edge and to the left of the left edge of the tornado condensation funnel. Color-coded scales for reflectivity and Doppler velocity are shown below each panel. In this and subsequent figures depicting vertical cross sections of Doppler velocity, positive (negative) velocities denote approaching (receding) motion, unlike in Fig. 5, where the opposite convention is used. (There were signal processing errors that reversed the sign of the Doppler velocity.) Times are (top left) 1848:29, (top right) 1849:53, and (bottom) 1850:12 CDT. The arrow points to a region of relatively high Doppler velocity away from the radar. suring the diameter of the region in which the reflectivity was speckled, was low compared the surrounding reflectivity, was inside a zone of strong reflectivity gradient, and was too low to compute Doppler velocity (this region displayed speckled velocities). The width of the hole was greatest for the scan that cut through the center of the tornado, as would be expected from geometrical considerations. The hole was pear shaped in overall appearance; it was approximately 40% wider at 100 m above the ground as it was above. Since the depth and width of the visible debris cloud were about 190 and 260 m, respectively, it is seen that the width of the debris cloud was approximately the same as that of the hole within the tornado. In a few of the cross sections [1848:52 (Fig. 8) and 1849:18 CDT (Fig. 9)], the echo hole closed up near the ground, so that there was a very thin layer of significant reflectivity. The cross section taken through the tornado at 1849:18 CDT (Fig. 9) is the cleanest one in that much of the scan was oriented parallel to the edges of, and through the center of, the condensation funnel, so that the Doppler velocities represent the radial wind component, if the vortex were perfectly axisymmetric. There is some tilt of the vortex with height toward the radar. However, since the radar antenna was scanning upward, it is likely that much of the tilt was an artifact of the motion of the tornado; in the time it took the radar to complete the RHI ( 10 s or less), the top of the tornado had moved on the order of 10 m closer. At 1849:02 (Fig. 10), on the other hand, when the radar antenna was scanning downward, the tornado appeared to tilt with height away from the radar. In all the vertical cross sections except the last two (1849:52 and 1850:10 CDT), there was a region of much

9 2333 FIG. 8. As in Fig. 7, but at time 1848:52 CDT, through the right edge of the tornado condensation funnel. Solid black line in the bottom, Doppler-velocity panel is the approximate horizon line. higher Doppler wind velocity, pointing in the direction of the tornado about m above the ground. This apparent jet (marked by arrows in the figures) of air flowing inward toward the tornado was as strong as 35 40ms 1, especially at 1848:52 and 1849:02, when the vertical cross sections cut through the right side of the condensation funnel [(Figs. 6, 8, and 10 (top right)]. At 1849:18, when the radar scanned through the center of the tornado (Figs. 6 and 9), a small region of relatively low values of ground-relative Doppler velocity (white yellow region) was found just above and radially outward from the apparent elevated jet (arrow; blue purple region). This small relative minimum in Doppler velocities might indicate frictional return flow. In Fig. 1, it is seen that above the surface inflow layer there are horizontal rolls at the edge of the simulated vortex. It is also possible that the vortex couplet seen in Fig. 9 might be a result of one such roll circulation. In Fig. 9, there is also a region of enhanced inflow in the lowest few elevation scans about 500 m east of the center of the tornado. It is possible that this might actually represent surface inflow. At 1848:52 and 1849:02 only, when the vertical cross sections cut across the right side of the center of the condensation funnel [cf. Fig. 6 and Figs. 8 and 10 (top right)], the scans showed an elevated jet cutting completely across the weak-echo hole. The Doppler velocity fields depicted in the vertical cross sections are therefore consistent with an elevated jet spiraling into and cyclonically around the tornado. 4. Summary and conclusions The Happy, Texas, tornado was the first tornado in which RHI scans were collected by the W-band radar. The vertical reflectivity structure of the tornado was similar to that found by Wurman and Gill (2000). However, the weak-echo hole in the Happy, Texas, tornado was in general broader just above the ground, at the height of the midlevel of the visible debris cloud. We offer the following two possible explanations for this finding: (a) the broadening is a result of centrifuging of scatterers radially outward near or just within the tornado core, or (b) the broadening is a result of the sec-

10 2334 MONTHLY WEATHER REVIEW VOLUME 132 FIG. 9. As in Fig. 7, but at time 1849:18 CDT, through the center of the tornado condensation funnel. Solid black line in the bottom, Doppler-velocity panel is the approximate horizon line. ondary circulation of the tornado, in which there is radial outflow above the surface friction layer advecting scatterers radially outward. The hole narrows below, probably owing to the inward transport of scatterers in the surface inflow layer. These hypotheses could be tested by numerical simulation experiments in which tornado-like vortices representing a broad range of swirl ratios would have particles of various sizes and size distributions injected into them (e.g., Snow 1984; Dowell 2000; Dowell et al. 2001). Such a study, which is beyond the scope of this paper, could explain quantitatively the observed shape of the echo-free eye s reflectivity profile. To estimate the particle sizes in the tornado, it might be possible to use the dual-frequency approach described by Pazmany et al. (2001). To implement this scheme, it will be necessary to scan the same volumes that the W-band radar scans with an X-band mobile Doppler (e.g., Pazmany et al. 2003). The inflow jet at m above the ground was not anticipated. Instead, it was thought that an inflow jet might be found at much lower altitudes, as in Golden and Purcell (1977) and Lewellen et al. (2000). Several possible explanations for an elevated boundary layer jet are as follows: First, evaporatively cooled air was advected around the tornado near the ground, as the tornado was beginning to dissipate. Air flowing into the tornado as a surface-boundary-layer jet was therefore lifted up over the surface cold pool. At the time the W-band radar data were collected, rain was falling around the tornado, but there were no in situ temperature measurements made to confirm the presence of a cold pool. There is some radar evidence, though, that this may have been the case. At 1848:29 (Fig. 7) and 1849: 02 (Fig. 10) there were arclike reflectivity structures ahead (to the left in the figures) of the tornado at low levels; the corresponding Doppler-velocity images show a relative maximum in rear-to-front (i.e., from the right to the left) flow near the ground, behind the arc, which is suggestive of the wind pattern seen behind some thunderstorm gust fronts (e.g., Wakimoto 1982). Another possibility is that there was a surface-boundary-layer jet, but it was asymmetrically distributed about the tornado (Golden and Purcell 1977); the viewing an-

11 2335 FIG. 10. As in Fig. 7, but at times (top left) 1848:15, (top right) 1849:02, and (bottom) 1849:35 CDT, through both the edge of the condensation funnel and the middle of the condensation funnel, depending on the height above the ground. gle of the radar may have been normal to the jet. Movies and videos of tornadoes (viewed at conferences and on television) sometimes show bands of dust feeding into the tornado near the ground from only one direction. The ground surrounding the tornado was relatively wet from precipitation, so that the wind field associated with a boundary-layer jet might not have been visualized. A third possibility is that the surface inflow layer was so shallow that the motion of scatterers in it was not detected by the radar. Even if the second or third hypotheses are true, then the elevated nature of the jet that was detected still needs to be explained. A final possibility is that the elevated jet was really associated with a wave along the edge of the tornado vortex, which would be manifest as a horizontal roll in a vertical cross section (cf. Fig. 1). More datasets need to be collected that provide detailed information on the vertical structure of the reflectivity and wind fields in tornadoes near the ground. To ensure accurate interpretation of the radar scans, the azimuth angles must be recorded properly along with the boresighted video. In the case of the Happy, Texas, tornado, the motion of the tornado had a significant component along the line of sight of the radar. It would be better for data collection if a tornado moved mainly across the line of sight, so that RHIs could be taken at one azimuth from a fixed location; then, as the tornado translates across the plane scanned in the RHI, the spatial resolution across the tornado would be maximized. If the core of a tornado lies outside the edge of its condensation funnel (if it has one) or outside the edge of its debris cloud, then RHIs through the tornado must extend well to the left and right of the visible condensation funnel to ensure that the core of the tornado is sampled. Efforts to minimize the effects of tilting of the RHI plane should be undertaken by more carefully leveling the radar platform. Knowledge of the diameter of the tornado core being sampled is essential. If attenuation is so severe that it is difficult or impossible to determine the core diameter from sector scans, then it may be necessary to increase the length of the radar pulses to enhance the radar s sensitivity at the expense of the along-the-line-of-sight spatial resolution.

12 2336 MONTHLY WEATHER REVIEW VOLUME 132 FIG. 11. Schematic of the composite radar reflectivity field below cloud base, associated with the Happy, TX, tornado, as viewed in the vertical plane. Acknowledgments. This work was supported by National Science Foundation (NSF) Grants ATM and ATM at OU and ATM at UMass. Pat Waukau, Kris Conrad, and Morris Weisman provided computer-related and graphics-related assistance at NCAR. Mark Laufersweiler provided computer support at OU. During field operations, the second author operated the radar, and Brendan Fennell from UMass drove the radar truck. We are indebted to Lou Wicker and to two anonymous reviewers for their helpful comments and suggestions. REFERENCES Bluestein, H. B., and J. H. Golden, 1993: Tornadoes and tornadic storms. The Tornado: Its Structure, Dynamics, Prediction, and Hazards, Geophys. Monogr., No. 79, Amer. Geophys. Union, , and A. L. Pazmany, 2000: Observations of tornadoes and other convective phenomena with a mobile, 3-mm wavelength, Doppler radar: The spring 1999 field experiment. Bull. Amer. Meteor. Soc., 81, , W. P. Unruh, J. LaDue, H. Stein, and D. Speheger, 1993: Doppler-radar wind spectra of supercell tornadoes. Mon. Wea. Rev., 121, , W.-C. Lee, M. Bell, C. C. Weiss, and A. L. Pazmany, 2003a: Mobile Doppler radar observations of a tornado in a supercell near Bassett, Nebraska, on 5 June Part II: Tornado-vortex structure. Mon. Wea. Rev., 131, , C. C. Weiss, and A. L. Pazmany, 2003b: Mobile Doppler radar observations of a tornado in a supercell near Bassett, Nebraska, on 5 June Part I: Tornadogenesis. Mon. Wea. Rev., 131, , C. C. Weiss, and A. L. Pazmany, 2004: Doppler radar observations of dust devils in Texas. Mon. Wea. Rev., 132, Brown, R. A., and V. T. Wood, 1991: On the interpretation of single- Doppler velocity patterns within severe thunderstorms. Wea. Forecasting, 6, Church, C. R., J. T. Snow, G. L. Baker, and E. M. Agee, 1979: Characteristics of tornado-like vortices as a function of swirl ratio: A laboratory investigation. J. Atmos. Sci., 36, Davies-Jones, R. P., R. J. Trapp, and H. B. Bluestein, 2001: Tornadoes and tornadic storms. Severe Convective Storms, Meteor. Monogr., No. 50, Amer. Meteor. Soc., Dowell, D. C., 2000: A pseudo-dual-doppler analysis of cyclic tornadogenesis. Ph.D. thesis, University of Oklahoma, 227 pp., J. Wurman, and L. J. Wicker, 2001: Centrifuging of scatterers in tornadoes. Preprints, 30th Int. Conf. on Radar Meteorology, Munich, Germany, Amer. Meteor. Soc., Fujita, T. T., 1981: Tornadoes and downbursts in the context of generalized planetary scales. J. Atmos. Sci., 38, Golden, J. H., and D. Purcell, 1978: Life cycle of the Union City, Oklahoma tornado and comparison with waterspouts. Mon. Wea. Rev., 106, 3 11., and, 1977: Photogrammetric velocities for the Great Bend, Kansas, tornado of 30 August 1974: Accelerations and asymmetries. Mon. Wea. Rev., 105, Lewellen, D. C., W. S. Lewellen, and J. Xia, 2000: The influence of a local swirl ratio on tornado intensification near the surface. J. Atmos. Sci., 57, Lewellen, W. S., and D. C. Lewellen, 1997: Large-eddy simulation of a tornado s interaction with the surface. J. Atmos. Sci., 54, Liu, H., 1993: Calculation of wind speeds required to damage or destroy buildings. The Tornado: Its Structure, Dynamics, Prediction, and Hazards, Geophys. Monogr., No. 79, Amer. Geophys. Union, Oye, R., C. K. Mueller, and S. Smith, 1995: Software for radar translation, visualization, editing, and interpolation. Preprints, 27th Conf. on Radar Meteorology, Vail, CO, Amer. Meteor. Soc., Pazmany, A. L., J. C. Galloway, J. B. Mead, I. Popstefanija, R. E. McIntosh, and H. B. Bluestein, 1999: Polarization diversity pulse pair technique for millimeter-wave Doppler radar measurements of severe-storm features. J. Atmos. Oceanic Technol., 16, , J. B. Mead, S. M. Sekelsky, D. J. McLaughlin, and H. Bluestein, 2001: Multi-frequency radar estimation of cloud and precipitation properties using an artificial neural network. Preprints, 30th Int. Conf. on Radar Meteorology, Munich, Germany, Amer. Meteor. Soc., , F. J. Lopez, H. B. Bluestein, and M. Kramar, 2003: Quantitative rain measurements with a mobile, X-band, polarimetric Doppler radar. Preprints, 31st Conf. on Radar Meteorology, Seattle, WA, Amer. Meteor. Soc., Rasmussen, E. N., R. Davies-Jones, C. A. Doswell, F. H. Carr, M. D. Eilts, D. R. MacGorman, and J. M. Straka, 1994: Verification of the Origins of Rotation in Tornadoes Experiment: VORTEX. Bull. Amer. Meteor. Soc., 75, Rotunno, R., 1979: A study in tornado-like vortex dynamics. J. Atmos. Sci., 36, Snow, J. T., 1984: On the formation of particle sheaths in columnar vortices. J. Atmos. Sci., 41, Wakimoto, R. M., 1982: The life cycle of thunderstorm gust fronts as viewed with Doppler radar and rawinsonde data. Mon. Wea. Rev., 110, , and B. E. Martner, 1992: Observations of a Colorado tornado. Part II: Combined photogrammetric and Doppler radar analysis. Mon. Wea. Rev., 120, , W.-C. Lee, H. B. Bluestein, C. H. Liu, and P. H. Hildebrand,

13 : ELDORA observations during VORTEX-95. Bull. Amer. Meteor. Soc., 77, Wilson, J. W., J. A. Moore, G. B. Foote, B. Martner, A. R. Rodi, T. Uttal, and J. M. Wilczak, 1988: Convection initiation and downburst experiment (CINDE). Bull. Amer. Meteor. Soc., 69, Wurman, J., 2002: The multiple-vortex structure of a tornado. Wea. Forecasting, 17, , and S. Gill, 2000: Finescale radar observations of the Dimmitt, Texas (2 June 1995), tornado. Mon. Wea. Rev., 128, , J. M. Straka, and E. N. Rasmussen, 1996: Fine-scale Doppler radar observations of tornadoes. Science, 272,