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  • View in gallery

    M-GLASS sounding at 0012 UTC 4 Jun 2000 near Bird City, KS.

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    Swath of (a) KGLD composite reflectivity (dBZ) color contours for the period 2210–0121 UTC with a black contour overlaid indicating the regions with updrafts greater than 10 m s−1, and (b) maximum updraft color contours during the dual-Doppler analysis period with a thick black contour overlaid indicating the regions with vertical vorticity greater than 10−2 s−1. Thin black contours at 20 and 40 dBZ from (a) are overlaid onto (b) for reference. Radar locations are denoted with a plus symbol.

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    Time–height contours of total graupel echo volume (solid black contours) and total hail echo volume (grayscale), and maximum updraft time series (dashed black line; values on right axis) for 3 Jun 2000.

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    Time series of updraft volume greater than 10 m s−1 (dashed black line; values on left axis), total graupel echo volume (solid black line; values on left axis), and the total flash rate from the LMA data (solid gray line; values on right axis) for 3 Jun 2000.

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    The average precipitation rate calculated at 3 km MSL for 3 Jun 2000 (solid black) and 29 Jun 2000 (dashed black) for each sequential radar volume in their respective analysis period. The analysis period for 3 Jun was 2210–0121 UTC (with 3–5-min spacing) and for 29 Jun was 2130–0115 UTC (with 5–7-min spacing).

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    Time–height contours of (a) LMA flash initiation height and (b) total LMA sources (grayscale in logarithmic units) with total flash rate overlaid in black for 3 Jun 2000.

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    Lightning mapping of an inverted IC flash at 2302:41 UTC 3 Jun 2000. (top) Altitude of the LMA sources vs time. (middle and bottom) Three different two-dimensional spatial projections along with an altitude histogram of the number of sources. The sources are color coded by time from blue to red. Plus and minus symbols indicate inferred ambient charge regions. This flash initiated near 10-km altitude and progressed downward into an inferred sloping positive charge region at 4–9-km altitude. The delayed sparse sources above the initiation point indicate the inferred negative charge region at 10–11-km altitude.

  • View in gallery

    As in Fig. 7, but for a normal IC flash at 2302:30 UTC 3 Jun 2000. This flash initiated near 9.5-km altitude and progressed upward into an inferred positive charge region at 10–11-km altitude. The delayed sparse sources below the initiation point indicate the inferred negative charge region at 9–10-km altitude.

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    KGLD horizontal radar reflectivity (Zh) at 2301 UTC 3 Jun 2000 at z = (a) 3 and (b) 7 km; vertical cross sections of (c) FHC using polarimetric data from CHILL and (d) KGLD Zh along the slanted black line indicated in (a) and (b). Updraft velocity contours beginning at 10 m s−1 (with a 10 m s−1 contour interval) are overlaid in black in (b) and (c) and in blue in (d). LMA sources of 4 representative flashes between 2302:22 and 2302:41 UTC, including the 2 shown in Figs. 7, 8, have been overlaid as small plus symbols onto (d), color coded by inferred ambient charge (red = positive, green = negative). The first source of each flash is plotted in (d) as diamonds and triangles, where the diamonds (triangles) indicate downward (upward) initial flash propagation. Storm-relative wind vectors have been overlaid onto (a) and (b). FHC categories are LH, SH, HG, LG, VI, wet snow (WS), dry snow (DS), rain (R), drizzle (Drz), and unclassified (UC). Note: the LMA sources atop the storm do not appear to be within reflectivity echo, but the 0- and 10-dBZ contours around the periphery of the storm are partially missing, either due to scanning/gridding geometry or the editing algorithms deleting the echo in low signal-to-noise regions.

  • View in gallery

    As in Fig. 9, but at 2325 UTC. The vertical cross sections in (c) and (d) are along the slanted line shown in (a) and (b), and the FHC used polarimetric data from S-Pol. LMA sources are from 4 representative flashes from 2326:03 to 2326:33 UTC.

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    As in Fig. 9, but at 2344 UTC. The vertical cross sections in (c) and (d) are at y = 71 km, and the FHCused polarimetric data from S-Pol. LMA sources are from 3 representative flashes from 2344:23 to 2344:40 UTC.

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    As in Fig. 9, but at 0026 UTC. The vertical cross sections in (c) and (d) are at y = 53 km, and the FHC used polarimetric data from S-Pol. LMA sources are from 2 representative flashes from 0026:42 to 0027:08 UTC.

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Radar and Lightning Observations of the 3 June 2000 Electrically Inverted Storm from STEPS

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  • 1 Department of Atmospheric Science, Colorado State University, Fort Collins, Colorado
  • | 2 Los Alamos National Laboratory, Los Alamos, New Mexico
  • | 3 Department of Atmospheric Science, Colorado State University, Fort Collins, Colorado
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Abstract

This study addresses the kinematic, microphysical, and electrical evolution of an isolated convective storm observed on 3 June 2000 during the Severe Thunderstorm Electrification and Precipitation Study field campaign. Doppler-derived vertical velocities, radar reflectivity, hydrometeor classifications from polarimetric radar, and Lightning Mapping Array (LMA) charge structures are examined over a nearly 3-h period. This storm, characterized as a low-precipitation supercell, produced modest amounts of hail, determined by fuzzy-logic hydrometeor classification as mostly small (<2 cm) hail, with one surface report of large (≥2 cm) hail. Doppler-derived updraft speeds peaked between 20 and 25 m s−1, and reflectivity was never greater than 60 dBZ. The most striking feature of this storm was its total lack of cloud-to-ground (CG) lightning. Though this storm was electrically active, with maximum flash rates near 30 per minute, no CG flashes of either polarity were detected. The charge structure inferred from the LMA observations was consistent with an inverted dipole, defined as having a midlevel positive charge region below upper-level negative charge. Inverted charge structures have typically been considered conducive to producing positive CG lightning; however, the 3 June storm appeared to lack the lower negative charge layer below the inverted dipole that is thought to provide the downward electrical bias necessary for positive CG lightning.

Corresponding author address: Sarah A. Tessendorf, NOAA/Earth System Research Laboratory, DSRC R/CSD3, 325 Broadway, Boulder, CO 80305. Email: sarah.tessendorf@noaa.gov

Abstract

This study addresses the kinematic, microphysical, and electrical evolution of an isolated convective storm observed on 3 June 2000 during the Severe Thunderstorm Electrification and Precipitation Study field campaign. Doppler-derived vertical velocities, radar reflectivity, hydrometeor classifications from polarimetric radar, and Lightning Mapping Array (LMA) charge structures are examined over a nearly 3-h period. This storm, characterized as a low-precipitation supercell, produced modest amounts of hail, determined by fuzzy-logic hydrometeor classification as mostly small (<2 cm) hail, with one surface report of large (≥2 cm) hail. Doppler-derived updraft speeds peaked between 20 and 25 m s−1, and reflectivity was never greater than 60 dBZ. The most striking feature of this storm was its total lack of cloud-to-ground (CG) lightning. Though this storm was electrically active, with maximum flash rates near 30 per minute, no CG flashes of either polarity were detected. The charge structure inferred from the LMA observations was consistent with an inverted dipole, defined as having a midlevel positive charge region below upper-level negative charge. Inverted charge structures have typically been considered conducive to producing positive CG lightning; however, the 3 June storm appeared to lack the lower negative charge layer below the inverted dipole that is thought to provide the downward electrical bias necessary for positive CG lightning.

Corresponding author address: Sarah A. Tessendorf, NOAA/Earth System Research Laboratory, DSRC R/CSD3, 325 Broadway, Boulder, CO 80305. Email: sarah.tessendorf@noaa.gov

1. Introduction

Based on cloud-to-ground (CG) lightning climatologies, CG lightning lowering negative charge to ground is far more common than positive CG lightning across the majority of the United States. However, storms dominated (>50%) by CG lightning that lower positive charge to ground are indeed observed and appear to be most frequent in the high plains of the United States (Orville and Huffines 2001; Carey and Rutledge 2003). While several hypotheses have emerged to explain the possible charge structure of positive CG-dominated storms (Brook et al. 1982; Seimon 1993; Carey and Rutledge 1998; Lang and Rutledge 2002), few studies have had three-dimensional lightning observations available to infer the general charge structure of such storms. The Severe Thunderstorm Electrification and Precipitation Study (STEPS; Lang et al. 2004) field campaign took place between 17 May and 20 July 2000 in eastern Colorado and western Kansas. A goal of the STEPS campaign was to identify relationships among microphysics, dynamics, and electrification in severe storms on the high plains, and in particular, to investigate positive CG lightning production. A comprehensive network of observing systems was deployed [see Lang et al. (2004) for a complete listing], including two dual-polarization Doppler radars and a three-dimensional Lightning Mapping Array (LMA; Rison et al. 1999).

Until recently, hypotheses offered to explain positive CG-dominated storms and positive CG lightning in general [e.g., the tilted dipole or inverted dipole outlined in detail in Williams (2001)] do not discuss the role of a lower negative charge layer below the lowest positive charge region. The charge structure typically associated with negative CG-producing storms is often referred to as a “normal” tripole, consisting of a main midlevel negative charge region below an upper-level positive charge layer, with a small lower positive charge layer situated below the negative region (Simpson and Scrase 1937; Krehbiel 1986; Williams 1989; Stolzenburg et al. 1998). Several studies (e.g., Jacobson and Krider 1976; Williams et al. 1989) suggest that in these normal tripole storms, the lower positive charge region is required to produce negative CG lightning. The model simulations of storm electrification by Mansell et al. (2002, 2005) also suggest that lower negative charge regions may be necessary for positive CG flashes, consistent with the observations of Wiens et al. (2005). Hence, lower negative charge may play a role in the production of positive CG flashes similar to the role played by lower positive charge in the production of negative CG flashes.

Several studies have already emerged from the STEPS dataset, primarily concerning the 29 June 2000 supercell that produced predominantly positive CG lightning (MacGorman et al. 2005; Tessendorf et al. 2005; Wiens et al. 2005; Kuhlman et al. 2006). The goal of this study is to document properties of a STEPS storm in which no CG flashes were detected, to be used as comparison to the other cases in which CG flashes were observed. This analysis is similar to that presented in Tessendorf et al. (2005), Wiens et al. (2005), and Tessendorf et al. (2007). By studying a broad spectrum of cases we hope to learn more about why some storms are dominated by positive CG lightning.

2. Data and methods

a. Radar data and processing

The triple-Doppler radar network in STEPS was composed of three S-band (10–11-cm wavelength) radars: the Colorado State University (CSU)–University of Chicago–Illinois State Water Survey (CHILL) polarimetric Doppler radar (hereafter CHILL), the National Center for Atmospheric Research (NCAR) S-Pol polarimetric Doppler radar, and the Goodland, Kansas, National Weather Service Weather Surveillance Radar-1988 Doppler (KGLD). The three radars were arranged in a nearly equilateral triangle with approximately 60-km sides (Tessendorf et al. 2005, their Fig. 1). The 3 June storm developed north of the S-Pol radar, near the northwestern edge of the eastern S-Pol–KGLD dual-Doppler lobe, and then tracked to the southeast. It remained within the eastern S-Pol–KGLD dual-Doppler lobe throughout its lifetime.

Wind field syntheses were completed for 25 volume scans during the period 2301 UTC 3 June–0121 UTC 4 June.1 The radar data were interpolated onto a Cartesian grid with 0.5-km horizontal and vertical resolution using NCAR’s Sorted Position Radar Interpolator (SPRINT; Mohr and Vaughn 1979). After the grid interpolation, the velocity data from S-Pol and KGLD were globally unfolded by means of NCAR’s Custom Editing and Display of Reduced Information in Cartesian Space software (CEDRIC; Mohr et al. 1986). In the wind synthesis, the data from each radar were advected to a common time using a manually calculated storm motion vector, and the vertical velocities were obtained using a variational integration of the continuity equation (O’Brien 1970).

Polarimetric data were available from either CHILL or S-Pol between 2210 UTC 3 June and 0121 UTC 4 June. The polarimetric data were edited, gridded, and used in a fuzzy-logic hydrometeor classification scheme (FHC) adapted from Liu and Chandrasekar (2000) and Straka et al. (2000), as in Tessendorf et al. (2005). The FHC algorithm used a temperature sounding from the 0013 UTC 4 June National Severe Storms Laboratory Electric Field Meter (EFM) balloon flight. As in Tessendorf et al. (2005), hydrometeor echo volumes were calculated for each radar scan time by multiplying the number of grid points that satisfied the FHC category of interest by the volume of a grid box (0.125 km3).

b. Lightning data and processing

The two sources of lightning data used in this study are the National Lightning Detection Network (NLDN; Cummins et al. 1998) and the New Mexico Institute of Mining and Technology (NMIMT) LMA (Rison et al. 1999). The NLDN provides the location, polarity, multiplicity, and peak current of CG flashes. The LMA provides measurements of the time and three-dimensional location of very high-frequency (VHF) radiation sources emitted by lightning discharges. For a given lightning flash, the LMA may locate hundreds to thousands of such sources resulting in detailed maps of the total lightning activity. The LMA may suggest a downward-propagating leader to ground, but leaders to ground are poorly resolved by the LMA, particularly in the case of positive CG flashes. Furthermore, the vertical location accuracy is very poor for sources near the ground (Thomas et al. 2004). Hence, CG return strokes cannot be identified with any confidence using only LMA data; thus, CG detection in this study was based on the NLDN data only.

To determine total [CG plus intracloud (IC)] flash rates, we used a sorting algorithm developed at NMIMT (Thomas et al. 2003) that sorts the LMA sources into discrete flashes using a 10-source minimum threshold for the flash grouping. As shown in Wiens et al. (2005), calculating quantitative flash rates with this sorting algorithm is somewhat ambiguous, especially for storms with intense lightning activity. Thus, we will focus our discussion on the trends and the order of magnitude of the flash rates.

To infer charge structure, we analyzed the LMA data on a flash-by-flash basis using the bidirectional discharge model (Kasemir 1960; Mazur and Ruhnke 1993), as in Wiens et al. (2005). For example, we assumed that flashes initiate in strong electric fields between regions of opposite net charge and propagate bidirectionally into the charge regions on either side. The negative breakdown component of a lightning flash is higher in VHF power and thus is more readily detected by the LMA compared with the positive breakdown component. Assuming that negative (positive) breakdown traverses regions of net positive (negative) charge, we infer the qualitative structure of the charge regions involved in the flash based on the temporal evolution of the flash and on the relative number of LMA sources on either side of the flash initiation location. Most IC flashes reveal distinct vertically separated “layers” of charge, with many more LMA sources in the inferred positive layer (see Figs. 7, 8).

3. Overview

a. Environmental conditions

The environment on 3 June 2000 was characterized by strong south-southwesterly surface winds between 8 and 10 m s−1 (gusts to near 13 m s−1) ahead of a surface boundary (not shown). Weaker, northwesterly flow prevailed behind (west of) the boundary. Surface temperatures ahead of the boundary were near 32°C, and dewpoints were near 12°C. Behind the boundary the temperatures were similar, yet the dewpoints were as low as 0°C. Around 1700 UTC, the surface boundary appeared as a convergence line in the radial velocity field and as a weak thin line echo in radar reflectivity (not shown). The boundary was oriented from southwest to northeast and propagated southeastward. The 0012 UTC Mobile GPS/Loran Atmospheric Sounding System (M-GLASS) sampled the inflow environment (within a few kilometers of the storm; Fig. 1) and indicated marginal CAPE (700 J kg−1) and notable drying above 500 hPa, and the upper-level winds were westerly near 25 m s−1.

Just after 2200 UTC, two small cells, referred to as A and B, were observed in southwestern Nebraska along the northern end of this boundary (Fig. 2a). Around 2240 UTC, IC lightning was first detected in cell A. At 2301 UTC (the first dual-Doppler analysis time available), mesocyclonic strength vertical vorticity (>10−2 s−1; Moller et al. 1994) was observed in cell A (Fig. 2b). By 2330 UTC, cell B merged into the forward left flank of cell A. A visible split in the upper-level radar reflectivity echo was observed at 2331 UTC, and the left-moving cell began to diminish soon thereafter (Fig. 2). By 0030 UTC, the maximum updraft in cell A had declined to near 5 m s−1, and after that the radar reflectivity echo continued to decrease over time until the storm had completely dissipated by 0121 UTC.

b. Kinematics and microphysics

At the time of the first dual-Doppler observations at 2301 UTC, updraft speeds were near 20 m s−1 and by 2350 UTC reached a maximum of ∼25 m s−1 (Fig. 3). During the gap in dual-Doppler scans, the maximum updraft speed(s) measured by the South Dakota School of Mines and Technology (SDSMT) T-28 aircraft (in pass 3) was 12.5 m s−1 (Holm 2005), though this may be an underestimate if the aircraft did not penetrate the strongest updraft core. When dual-Doppler observations were available again at 0026 UTC, peak Doppler-derived updrafts were near 13 m s−1 and soon thereafter declined to 5 m s−1 and steadily decreased beyond that. However, the T-28 aircraft measured updraft speeds as high as 18 m s−1 between 0034 and 0037 UTC (pass 7). This discrepancy could be the result of the scale of the aircraft observations, which was smaller than the resolution of the dual-Doppler analysis or due to uncertainties in the dual-Doppler derived winds (Holm 2005).

The storm exhibited mesocyclonic strength vertical vorticity between 2301 UTC and 0026 UTC (Fig. 2b). At 2301 UTC, this vorticity was rather shallow, confined between 7 and 8 km, but by 2310 UTC it lowered in altitude to as low as 4 km and persisted for at least another 20 min based on our dual-Doppler observations (not shown). Notice that the updraft and cyclonic vorticity were collocated with the main reflectivity core for the duration of the dual-Doppler analysis (Fig. 2). Implications of this observation will be discussed further in section 5. Graupel was first detected by the FHC algorithm near 2235 UTC in the midlevels of the storm (near 8 km) and steadily increased in echo volume (EV) until near 2320 UTC, at which time graupel amounts leveled off until near 2350 UTC (Figs. 3, 4). After this time, the graupel EV attained its maximum value of near 1000 km3 at 0002 UTC, most of which was centered near 7 km MSL (Fig. 3).2

Total hail EV was minimal and confined to 7–9 km between 0000 and 0020 (Fig. 3). According to National Climatic Data Center storm data, there was only one large hail (>2 cm) report associated with this storm. The report was at 0015 UTC, immediately after the FHC-inferred hail aloft. The hail EV contours in Fig. 3 do not explicitly show this hail falling out, perhaps because the temporal resolution of the radar data was too coarse or the hail partially melted on its descent and became classified as graupel near the surface. However, it should be noted that most of the FHC-inferred low-level hail EV (i.e., that shown in the contours between 2320 and 0026 UTC that appear constrained to near 3 km) is suspect because of its horizontal stratification just below the melting level. It is possible that these hydrometeors were water-coated graupel (i.e., in the process of melting) that was misclassified as small hail (SH; see section 4b). One other point to note is that the calculated total hail EV is composed of the FHC small hail and large hail (LH) categories, and for this storm, large hail was scarcely detected.

Because of the limited availability of dual-Doppler scans, the trends in the volume of updraft greater than 10 m s−1 (hereafter, UV10) cannot be fully resolved. However, during the portion of the time series with UV10 data, peaks in UV10 were coincident with or preceded those in graupel and hail EV. Shortly after 2310 UTC, UV10 rose sharply and then peaked at 2316 UTC with a value of 97 km3, just before the first peak in graupel EV (Fig. 4). By 2331 UTC, UV10 began to rise sharply again and then achieved its maximum value of 140 km3 at 2344 UTC. The absolute maximum in graupel EV happened within 15 min of the last resolved UV10 peak, but during a period when UV10 was unavailable, and thus it is unknown if UV10 increased, decreased, or remained steady during this gap.

This storm could be considered a supercell based on the single criterion in Moller et al. (1994) that define a supercell as having a deep mesocyclone that persists on the order of tens of minutes. This storm was most similar to a low-precipitation (LP) supercell, however, because its radar echo was relatively small and it did not exhibit a low-level hook echo, nor did it have strong (>60 dBZ) reflectivity (Fig. 2), indicating that it had less overall precipitation than a “classic” supercell (Bluestein and Parks 1983). Average radar-derived rain rates (using an RKdp relationship; Cifelli et al. 2002) on 3 June were a factor of 2 lower than those in the 29 June 2000 STEPS classic supercell during its intense phase (Fig. 5).

c. Charge structure

The relationship between total lightning and graupel EV trends in this storm clearly reinforces the importance of graupel production to the electrification process. There was no lightning detected by the LMA in this storm prior to the initial radar-inferred presence of graupel around 2240 UTC (Fig. 4), nor was there any strong radar reflectivity (nothing greater than 35 dBZ) prior to this time (not shown). Furthermore, the trend in total lightning flash rate closely followed that of graupel EV (Fig. 4). This temporal relationship was confirmed quantitatively by Wiens (2005), who found a statistical correlation coefficient of 0.81 between the time series of total flash rate and graupel EV. The maximum flash rate in this storm was near 36 per minute and occurred at 0002 UTC when the graupel EV reached its peak.

Throughout the duration of the 3 June storm, the vast majority of lightning flashes occurred near the precipitation core of the storm, initiated near 9–11-km altitude, and extended downward producing relatively large numbers of LMA sources below the flash origin and relatively few LMA sources above (Figs. 6, 7). This situation describes what could be termed an inverted dipole, with a negative charge region near 10–12-km altitude (T < −40°C) and positive charge below (Krehbiel et al. 2000; Williams 2001; Rust and MacGorman 2002). Some of these “inverted” IC flashes remained vertically confined to the upper part of the storm, with the positive charge centered near 10 km (T ∼ −30°C) within strong (>30 dBZ) lofted echo. However, most of the inverted flashes extended to much lower altitudes, with the positive charge sloping downward east of the updraft, apparently following the descent of the precipitation (see section 4). Hence, the lower positive charge may have consisted of multiple layers or simply one deep charge layer. There were no flashes that indicated an intervening negative charge region within the positive charge. As the time–height contours of total LMA sources in Fig. 6b indicate, the bulk of the LMA sources were constrained between 5- and 10-km altitude, which is the same altitude range that we consistently identified as the positive charge region of an inverted dipole in our flash-by-flash analysis (see section 4). The LMA source density contours also closely resemble the graupel EV contours in Fig. 3, further emphasizing the relationship between rimed ice and electrification.

Assuming noninductive charging is responsible for electrification, the implication of this inferred charge structure is that larger ice particles (e.g., graupel) received positive charge after rebounding collisions with smaller ice particles (e.g., ice crystals), and the latter received negative charge. Furthermore, according to laboratory studies that base the sign of the charge transferred on temperature and cloud liquid water content (LWC), effective LWC (a combination of the LWC and collision efficiency), or rime accretion rate (Takahashi 1978; Saunders et al. 1991; Saunders and Peck 1998), this would suggest that LWC or rime accretion rates were large enough in this storm for the graupel to acquire positive charge. The maximum adiabatic LWC (calculated from the M-GLASS sounding in Fig. 1) was 3.4 g kg−1 (at 8.9 km). It is well known that entrainment and mixing effects dilute the LWC from adiabatic values; however, aircraft observations in small cumulus have measured near-adiabatic LWC in the cores of updrafts (Lawson and Blyth 1998). The SDSMT T-28 armored aircraft measured maximum in situ LWC in the range of 2–3 g m−3 at an altitude near 6 km (Holm 2005). Holm (2005) performed an LWC analysis that compared adiabatic LWC calculations from a composite EFM/M-GLASS sounding at the height of the aircraft track with the actual SDSMT T-28 LWC measurements. Using the ratio of measured LWC to adiabatic LWC, Holm then adjusted the adiabatic LWC for the inferred regions of noninductive charging and found that the graupel would acquire negative charge using the Takahashi (1978) results but would acquire positive charge using the Saunders et al. (1991) parameters. This reveals the discrepancies between the laboratory studies of noninductive charging, as well as their extreme sensitivity to LWC, and perhaps lends support to effective LWC being an important noninductive charging parameter rather than just LWC alone. Electrification simulations by Mansell et al. (2005) and Kuhlman et al. (2006) have also shown that different charging schemes (based on the different laboratory results) can yield opposite polarity charging, and the rime accretion rate schemes (Saunders and Peck 1998) are more versatile and capable of producing inverted charge structures. Clearly, improvements in our knowledge of noninductive charging parameters are still needed before any conclusions can be made about how the observed charge structures of this storm developed.

4. Detailed evolution

Based on the UV10, graupel EV, and lightning flash rate trends in Fig. 4, we have identified 3 main phases of the storm’s evolution: a developing phase (2210–2310 UTC), a mature phase (2310–0010 UTC), and a dissipating phase (0010–0120 UTC). These three phases of the storm’s life cycle are similar to the three-stage classification defined by Byers and Braham (1949). We will make reference to these phases as we discuss the detailed observations.

a. Developing phase (2210–2310 UTC)

Near 2210 UTC, the 3 June storm (cell A in Fig. 2) was characterized by a high-based (<10 dBZ below 5 km), shallow (<10 dBZ above 9 km) radar echo structure with maximum reflectivity no greater than 30 dBZ (not shown). The storm was not in a location for optimal dual-Doppler analysis throughout most of the developing phase, so diagnosis of the updraft velocity during this time was not possible. The echo base (∼10 dBZ) lowered to near 1.5 km by 2226 UTC, and reflectivity >30 dBZ (inferred to be graupel by FHC) was first observed at 2233 UTC between 7 and 8 km on the west side of the low-level reflectivity echo (not shown). This perhaps indicated the presence of a new and stronger updraft on the west flank of the storm, beneath this lofted echo. After this time, the storm continually exhibited a larger volume of reflectivity >30 dBZ and FHC-inferred graupel echo (Fig. 4). Soon thereafter, near 2240 UTC, the first lightning flashes were observed by the LMA. A second, weaker cell (B in Fig. 2) was first observed on radar at 2239 UTC to the northeast of cell A.

Dual-Doppler observations of the storm were available by 2301 UTC, near the end of the developing phase. At 2301 UTC, both cells A and B had relatively weak reflectivity, with cell A still containing reflectivity just greater than 30 dBZ and some FHC-inferred graupel, while cell B did not exhibit any reflectivity greater than 30 dBZ (Fig. 9). Cell A had two updraft cores at this time: one near the west flank of the reflectivity core, as strong as 20 m s−1, and a shallow and narrow 10 m s−1 updraft in the center of the echo (Fig. 9). Without dual-Doppler observations prior to this time, we cannot diagnose the evolution of the second, smaller updraft, but perhaps it was an older updraft that was dissipating, while the stronger updraft on the west flank was a newer, developing updraft. Low-level inflow at this time was weak and south-southeasterly, with upper-level flow from the northwest (Figs. 9a,b).

The charge structure inferred from the initial flashes, from 2243 to 2255 UTC, is best characterized as an inverted dipole, involving a positive charge region near 8–9 km and an upper-level negative charge region at 10–11 km (not shown). An example inverted flash at 2302 UTC is shown in Fig. 7. In addition to the flashes in this persistent inverted dipole, there was a roughly 30-min time span (2255–2325 UTC) during which several flashes initiated upward from 10–12 km to an inferred upper positive charge region that lay near the upper radar echo boundary of the storm (Figs. 8 –10). Flashes involving the upper positive charge were generally within the anvil, farther downwind (east) of the core. These upper flashes occurred farther and farther from the core of the storm as time progressed. At 2301 UTC, this upper positive charge region was above the inverted dipole and extended downwind into the anvil, centered at a height of 11 km (Figs. 8 and 9). Also at this time, the LMA sources were only in the vicinity of the second, smaller updraft, suggesting that the charge separation mechanisms in the stronger (perhaps newer) updraft had not yet advanced to the point of generating lightning. The charge regions sloped downward away from the updraft core, and most of the midlevel positive charge was in a region of FHC-inferred graupel, while the negative charge aloft was in FHC-inferred snow and vertically oriented ice crystals (VI; Fig. 9). The extreme upper positive charge region was also in FHC-inferred snow and ice crystals.

b. Mature phase (2310–0010 UTC)

During the mature phase of cell A, maximum updraft velocities were near 20 m s−1 and UV10 reached 100 km3 (Figs. 3, 4). By 2325 UTC, cell B had merged into the northern flank of cell A and three 10 m s−1 updraft cores were resolved at 7 km (as seen in Figs. 10a,b). The low-level inflow at this time was still south-south-easterly with northwesterly upper-level flow. The two larger updraft areas (to the west and south) became one broad updraft region by 2331 UTC (not shown), while the northern updraft core remained separate and even developed its own upper-level reflectivity core by 2331 UTC, creating a split in the storm reflectivity echo (also seen at 2344 UTC in Fig. 11b). Because of the enhanced cyclonic vorticity collocated with the updraft during the mature phase of this storm, there was some cyclonic curvature in the flow around the south side of the southern, larger updraft (Figs. 10a,b). The location of this curvature in the flow relative to the updraft, however, likely carried growing particles in the updraft around to the downwind and northern side of the updraft where they fell out away from the storm inflow, thus preventing them from reentering the updraft for further growth.

The reflectivity and FHC fields showed a modest overhang at 2325 UTC (Figs. 10c,d), with graupel inferred in the upper levels of the updraft. The small hail classified by FHC in Fig. 10c near 3 km was likely a misclassification of melting graupel. Note that in this figure, the hydrometeors directly above the region classified as SH are high-density graupel (HG) and low-density graupel (LG). As the LG particles fall through the melting level at 4.5 km, the meltwater on their surfaces promotes higher radar returned power (i.e., radar reflectivity; similar to a bright band), and due to the increasing radar reflectivity, the LG particle is classified as HG, and then as SH, based on the radar variable thresholds in the FHC algorithm (Tessendorf et al. 2005).

The charge structure in the mature phase of this storm was still characterized as an inverted dipole nearest the precipitation core (Fig. 10d). Most of the inferred positive charge was again in the region of FHC-inferred graupel, while the inferred negative charge was in the upper levels where snow and ice crystals were identified by FHC. The single flash in the eastern anvil in Fig. 10 was the last flash that clearly involved the upper positive charge, and it is much farther downwind from the precipitation core than were previous flashes that involved the upper positive charge. Both the negative and upper positive charge layers of this flash were in regions of FHC-inferred snow and ice crystals in the anvil.

By 2344 UTC, the updraft reached its absolute maximum intensity (based on the available dual-Doppler observations) in the main updraft along the southwest flank of the storm (Fig. 11). The northern flank updraft had a distinct reflectivity core that was diverging from the southern updraft over time and responsible for the northern branch of the V-shaped low-level reflectivity. Persistent south-southeasterly inflow and upper-level northwesterly flow was still evident, as well as some cyclonic curvature around the south side of the main updraft (Figs. 11a,b). The persistent inverted dipole (now without the upper positive charge layer) was still the dominant charge structure and, within the updraft region, FHC-inferred graupel (ice crystals) was (were) observed where the LMA indicated positive (negative) charge.

c. Dissipating phase (0010–0120 UTC)

Though the dual-Doppler observations were not available for the 20-min period prior to 0026 UTC, it is apparent that the storm entered its dissipating phase during this time. The graupel EV and total lightning flash rates rapidly diminished near 0010 UTC, and by 0026 UTC dramatically weaker maximum updraft speeds and UV10 were observed (Figs. 3, 4). At 0026 UTC, two distinct cores were observed in the low-level reflectivity field, each corresponding to the northern and southern updraft cores previously discussed at 2344 UTC (Figs. 12a,b). The upper-level reflectivity in the northern core had greatly diminished by 0026 UTC, while the southern upper-level reflectivity core was still >30 dBZ. The updraft in the southern core was still 10 m s−1 at 0026 UTC, but quickly weakened to near 5 m s−1 six minutes later. The low-level inflow was now more southerly, and the upper-level flow was west-southwesterly (Figs. 12a,b).

The charge structure during this phase was still representative of an inverted dipole, with a very deep main positive charge region situated below negative charge (Fig. 12d). The charge regions were elevated within the updraft core and sloped downward away from the updraft into the precipitation core. As in the previous figures, FHC-inferred graupel was observed in regions with positive charge, and ice crystals were inferred aloft in regions of negative charge (Fig. 12). This inverted dipole structure persisted up until the very last observed flash at 0046 UTC (not shown).

5. Discussion

The 3 June storm had moderate updraft speeds (20–25 m s−1), as well as moderate FHC-inferred graupel and small hail, with limited large hail. Compared with the 29 June STEPS supercell (Tessendorf et al. 2005), the 3 June storm formed in an environment with lower CAPE, had half the maximum updraft speeds, and had nearly an order of magnitude less graupel and hail EV. A few possible reasons for the lack of (large) hail in this storm were weaker updrafts, lower UV10, and the collocation of the updraft and cyclonic vorticity (see Fig. 2b). In the case of the 29 June storm, Tessendorf et al. (2005) showed that cyclonically curved flow on the right flank of the updraft was an important ingredient, in addition to sufficient updraft size and intensity, in the production of large (>2 cm) hail. In that storm, the offset of the strong cyclonic flow from the updraft core allowed embryonic particles, which had likely fallen from the upper-level stagnation zone upwind of the updraft, to reenter the updraft for continued growth. With the cyclonic vorticity collocated with the updraft on 3 June, however, the particles grown from scratch were likely exhausted into the anvil or along the north side of the updraft, certainly not in a position to reenter the southeasterly inflow for continued growth (see Fig. 10). Nonetheless, the early evolution (first 2 h) of the 3 June storm had similar orders of magnitude of UV10, graupel EV, and lightning flash rates to the early evolution of the 29 June storm, prior to the latter’s right turn and dramatic intensification (Tessendorf et al. 2005; Wiens et al. 2005). The maximum updraft and hail echo volume on 29 June prior to its right turn, however, was already much higher than that on 3 June.

Maximum total flash rates in this storm were near 30 flashes per minute, which is below the 60 flashes per minute threshold found by Williams et al. (1999) to distinguish nonsevere from severe storms in Florida. However, this threshold may be elevated because of contamination by single-source flashes (whereas we used a 10-source threshold in flash grouping), so the 3 June flash rates may actually have been closer to those observed in severe storms. This storm did have an isolated severe hail report, which technically classifies it as a severe storm, but nonetheless, its total flash rates were somewhat low compared with other severe storms. Furthermore, no CG flashes of either polarity were detected by NLDN. The lightning activity was somewhat similar to that in the early (nonsevere) development of the 29 June supercell. Prior to the right turn and onset of the severe phase in the 29 June storm, the total flash rates were on the same order of magnitude as those on 3 June, and the 29 June storm produced only two CG flashes during the first 2 h of lightning. Once the 29 June storm became severe, flash rates gained an order of magnitude and frequent positive CG flash activity began (Wiens et al. 2005).

Though the mechanism(s) that generated the inverted dipole charge structure with an upper positive charge layer cannot be identified with certainty, we can provide some plausible assumptions. First, we assume that most of the charge separating collisions took place in the main updraft and that the largest hydrometeors resided and fell out closest to the strongest echo. Given the observations that positive charge was also constrained within the strongest echo, the interpretation is that the electrification processes were granting positive charge to the larger particles. The high cloud base near the main updraft and charging region may explain why the particles undergoing growth by riming were consistently charging positively. Following the arguments of Williams et al. (2005), higher cloud base reduces the warm-rain (collision/coalescence) precipitation growth zone, thereby promoting higher supercooled liquid water contents and elevated riming rates in the mixed-phase region where charge-separating collisions occur. According to laboratory studies of the noninductive (collision charging) mechanism (Takahashi 1978; Saunders et al. 1991), graupel (i.e., the rimer) charges positively under high LWC conditions.

There are a few possible explanations for the extreme upper positive charge: 1) it may have been a screening layer above the inverted dipole, because it was observed along the upper storm boundary, or 2) it may have resulted from noninductive charging processes in a different charging regime (i.e., in regions with different ambient temperature and LWC from the main updraft). For example, collisions where LWC would be lower, in either the periphery of the main updraft or in a weaker, older convective updraft, may have resulted in the riming ice receiving negative charge. The observations are most consistent with the latter explanation. The flashes involving the extreme upper positive charge propagated away from the storm core over time, and the last of these flashes, near 2325 UTC (Fig. 10), shows the upper positive LMA sources within noteworthy radar echo, suggesting that this charge layer was perhaps associated with a group of hydrometeors in the anvil and not just a screening layer.

It has been suggested that a lower positive charge region locally enhances the electric field, providing impetus for negative discharges to ground (Jacobson and Krider 1976; Williams et al. 1989; Mansell et al. 2002). Recent studies suggest a similar role for a lower negative charge region in the production of positive CG flashes in inverted polarity storms (Mansell et al. 2002, 2005; Wiens et al. 2005). Like the 3 June storm of this study, the 29 June supercell also exhibited an inverted charge structure. However, unlike the 3 June storm, LMA observations during the severe phase of the 29 June storm showed a lower negative charge region below the inverted dipole (Wiens et al. 2005). Additionally, Kuhlman et al. (2006) simulated the electrification in the 29 June storm and suggested that the observed and simulated lower negative charge layer was crucial to the production of positive CGs. The LMA data never indicated the presence of a lower negative charge layer below the inverted dipole on 3 June. Therefore it would follow that the lack of a sufficient lower negative charge layer on 3 June may have inhibited the production of positive CG flashes.

The reason for the apparent lack of a lower negative charge layer on 3 June is difficult to pinpoint. In the simulation of the 29 June supercell by Kuhlman et al. (2006), the lower negative charge formed by a combination of negative noninductive charging of graupel outside the updraft core, precipitation fallout and recycling, and inductive charging. Based on additional storm electrification simulations, Mansell et al. (2005) suggested that noninductive charging could account for the lower charge layer without inductive charging processes, if the ice crystal concentrations at lower altitudes (i.e., warmer temperatures) were high enough (>50 L−1 in their simulations), but for all other cases, inductive charging was deemed important. Guided by these modeling results, we speculate that the LMA did not indicate a lower negative charge layer on 3 June because of one (or all) of the following factors:

  1. There was a lower negative charge region, but it was too weak to initiate a discharge. Since the LMA cannot reveal a charge region unless that region is involved in lightning, there was no LMA evidence of the lower negative charge. This explanation is partially supported by the 0013 UTC EFM sounding that indicated a weak lower negative charge layer may have been present (Rust et al. 2005).
  2. Inductive charging processes were inhibited, perhaps because of less liquid precipitation in this LP storm.
  3. The lack of precipitation recycling reduced the quantity of riming ice growing at lower altitudes, which suppressed (noninductive or inductive) precipitation-based charge separation processes needed to generate the low-level charge layer.

Though we presume that inverted charge structures favor positive CG lightning, provided a lower negative charge layer is present, negative CGs have been shown to originate from the upper negative charge in an inverted charge structure (Wiens et al. 2005). We suggest that negative CG flashes in the 3 June storm were unlikely, however, because of the presence of a deep positive charge layer between the upper negative charge and the ground. This deep positive charge layer most likely made negative CG flashes less energetically favorable than IC flashes between the two charge layers (Marshall and Stolzenburg 2002).

We feel it is important to make the distinction between storms studied in the literature that have low-CG but high-IC flash rates, or high IC:CG ratios (MacGorman et al. 1989; MacGorman and Burgess 1994; Lang et al. 2000; Lang and Rutledge 2002), and the present storm that exhibited no CG flashes, because the reasons for each type of behavior could be due to different mechanisms. For example, the proposed “elevated charge” hypothesis (MacGorman et al. 1989), which has been previously used to explain low-CG storms, still seems to be a plausible reason for keeping CG flash rates low, while maintaining or enhancing IC flash rates in kinematically intense storms. In fact, even in this storm, the LMA sources and inferred charge layers were observed at higher altitudes nearest the strongest updraft than in the rest of the storm (see Figs. 9 –12). However, we speculate that the absence of a sufficient lower charge layer, both opposite in polarity to the charge region sending charge to ground and of appropriate strength to enhance the electric field and provide the impetus for the discharge to ground, is perhaps a key reason for the lack of CG flashes in otherwise electrically active storms. This suggestion is based upon the two key observations of the electrically active 3 June storm: no CG flashes were detected, and there was no LMA-inferred lower negative charge region.

6. Conclusions

The objective of this study was to examine relationships among the kinematic, microphysical, and electrical aspects of the 3 June 2000 non-CG-producing supercell. The radar coverage on this day was suitable for studying the storm structure evolution; however, the dual-Doppler coverage was not optimal, and therefore we were unable to estimate vertical velocities over part of the storm’s evolution. Nonetheless, the isolated nature of this storm, in addition to its modest flash rates, provided a unique opportunity to study the evolution of charge structure for an inverted storm using the LMA data.

No CG flashes of either polarity were detected in this storm. It exhibited a persistent inverted dipole charge structure, but the LMA data never indicated the presence of a lower negative charge region below the inverted dipole. Much like the lower positive charge region has been deemed important in producing negative CG flashes, these data support the idea that the lack of the lower negative charge layer, which completes the inverted tripole and may locally enhance the electric field allowing for positive CG discharges, may have been a key factor in preventing this storm from producing positive CG flashes. Certainly, more storms that produce IC but not CG lightning need to be examined to evaluate this claim.

Acknowledgments

We thank L. Jay Miller of NCAR for support with the multiple-Doppler analysis. Thanks to Dr. Larry Carey (Texas A&M) for assistance with the polarimetric data processing and hydrometeor identification algorithm suggestions. The S-Pol and KGLD radar data were obtained from NCAR. We thank Dr. Paul Krehbiel, Dr. William Rison, Dr. Ron Thomas, Dr. Tim Hamlin, and Jeremiah Harlin of New Mexico Tech for the LMA data and software. Reviews by Drs. Timothy Lang, Stephen Nesbitt, and Rob Cifelli were also much appreciated and contributed to the improvement of this manuscript. This research was supported by the National Science Foundation (NSF) Physical Meteorology Program under Grant ATM-0309303. The CSU–CHILL facility is supported by the NSF and CSU.

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Fig. 1.
Fig. 1.

M-GLASS sounding at 0012 UTC 4 Jun 2000 near Bird City, KS.

Citation: Monthly Weather Review 135, 11; 10.1175/2006MWR1953.1

Fig. 2.
Fig. 2.

Swath of (a) KGLD composite reflectivity (dBZ) color contours for the period 2210–0121 UTC with a black contour overlaid indicating the regions with updrafts greater than 10 m s−1, and (b) maximum updraft color contours during the dual-Doppler analysis period with a thick black contour overlaid indicating the regions with vertical vorticity greater than 10−2 s−1. Thin black contours at 20 and 40 dBZ from (a) are overlaid onto (b) for reference. Radar locations are denoted with a plus symbol.

Citation: Monthly Weather Review 135, 11; 10.1175/2006MWR1953.1

Fig. 3.
Fig. 3.

Time–height contours of total graupel echo volume (solid black contours) and total hail echo volume (grayscale), and maximum updraft time series (dashed black line; values on right axis) for 3 Jun 2000.

Citation: Monthly Weather Review 135, 11; 10.1175/2006MWR1953.1

Fig. 4.
Fig. 4.

Time series of updraft volume greater than 10 m s−1 (dashed black line; values on left axis), total graupel echo volume (solid black line; values on left axis), and the total flash rate from the LMA data (solid gray line; values on right axis) for 3 Jun 2000.

Citation: Monthly Weather Review 135, 11; 10.1175/2006MWR1953.1

Fig. 5.
Fig. 5.

The average precipitation rate calculated at 3 km MSL for 3 Jun 2000 (solid black) and 29 Jun 2000 (dashed black) for each sequential radar volume in their respective analysis period. The analysis period for 3 Jun was 2210–0121 UTC (with 3–5-min spacing) and for 29 Jun was 2130–0115 UTC (with 5–7-min spacing).

Citation: Monthly Weather Review 135, 11; 10.1175/2006MWR1953.1

Fig. 6.
Fig. 6.

Time–height contours of (a) LMA flash initiation height and (b) total LMA sources (grayscale in logarithmic units) with total flash rate overlaid in black for 3 Jun 2000.

Citation: Monthly Weather Review 135, 11; 10.1175/2006MWR1953.1

Fig. 7.
Fig. 7.

Lightning mapping of an inverted IC flash at 2302:41 UTC 3 Jun 2000. (top) Altitude of the LMA sources vs time. (middle and bottom) Three different two-dimensional spatial projections along with an altitude histogram of the number of sources. The sources are color coded by time from blue to red. Plus and minus symbols indicate inferred ambient charge regions. This flash initiated near 10-km altitude and progressed downward into an inferred sloping positive charge region at 4–9-km altitude. The delayed sparse sources above the initiation point indicate the inferred negative charge region at 10–11-km altitude.

Citation: Monthly Weather Review 135, 11; 10.1175/2006MWR1953.1

Fig. 8.
Fig. 8.

As in Fig. 7, but for a normal IC flash at 2302:30 UTC 3 Jun 2000. This flash initiated near 9.5-km altitude and progressed upward into an inferred positive charge region at 10–11-km altitude. The delayed sparse sources below the initiation point indicate the inferred negative charge region at 9–10-km altitude.

Citation: Monthly Weather Review 135, 11; 10.1175/2006MWR1953.1

Fig. 9.
Fig. 9.

KGLD horizontal radar reflectivity (Zh) at 2301 UTC 3 Jun 2000 at z = (a) 3 and (b) 7 km; vertical cross sections of (c) FHC using polarimetric data from CHILL and (d) KGLD Zh along the slanted black line indicated in (a) and (b). Updraft velocity contours beginning at 10 m s−1 (with a 10 m s−1 contour interval) are overlaid in black in (b) and (c) and in blue in (d). LMA sources of 4 representative flashes between 2302:22 and 2302:41 UTC, including the 2 shown in Figs. 7, 8, have been overlaid as small plus symbols onto (d), color coded by inferred ambient charge (red = positive, green = negative). The first source of each flash is plotted in (d) as diamonds and triangles, where the diamonds (triangles) indicate downward (upward) initial flash propagation. Storm-relative wind vectors have been overlaid onto (a) and (b). FHC categories are LH, SH, HG, LG, VI, wet snow (WS), dry snow (DS), rain (R), drizzle (Drz), and unclassified (UC). Note: the LMA sources atop the storm do not appear to be within reflectivity echo, but the 0- and 10-dBZ contours around the periphery of the storm are partially missing, either due to scanning/gridding geometry or the editing algorithms deleting the echo in low signal-to-noise regions.

Citation: Monthly Weather Review 135, 11; 10.1175/2006MWR1953.1

Fig. 10.
Fig. 10.

As in Fig. 9, but at 2325 UTC. The vertical cross sections in (c) and (d) are along the slanted line shown in (a) and (b), and the FHC used polarimetric data from S-Pol. LMA sources are from 4 representative flashes from 2326:03 to 2326:33 UTC.

Citation: Monthly Weather Review 135, 11; 10.1175/2006MWR1953.1

Fig. 11.
Fig. 11.

As in Fig. 9, but at 2344 UTC. The vertical cross sections in (c) and (d) are at y = 71 km, and the FHCused polarimetric data from S-Pol. LMA sources are from 3 representative flashes from 2344:23 to 2344:40 UTC.

Citation: Monthly Weather Review 135, 11; 10.1175/2006MWR1953.1

Fig. 12.
Fig. 12.

As in Fig. 9, but at 0026 UTC. The vertical cross sections in (c) and (d) are at y = 53 km, and the FHC used polarimetric data from S-Pol. LMA sources are from 2 representative flashes from 0026:42 to 0027:08 UTC.

Citation: Monthly Weather Review 135, 11; 10.1175/2006MWR1953.1

1

The S-Pol radar went down for 20 min prior to the 0026 UTC volume scan and therefore syntheses could not be performed during that period.

2

All heights hereafter will be in kilometers above mean sea level (MSL).

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