Tornado

Lightning in Tornadic Storms' A Review

DONALD R. MACGORMAN

National Severe Storms Laboratory, NOAA, Norman, Oklahoma 73069

1. INTRODUCTION

There have been many reports of unusual lightning characteristics in tornadic storms. For example, eyewitnesses have reported scorching beneath tornado funnels, a steady or rapidly oscillating glow inside funnels, rapidly recurring small patches of light on the side of the thunderstorm, or unusually high or low flash rates [e.g., Church and Barnhart, 1979; Vaughan and Vonnegut, 1976; Vonnegut and Weyer, 1966]. It is difficult, however, to quantify relationships between lightning and tornadic storms from these eyewit- ness reports.

Prior to 1975 most quantitative measurements of lightning in tornadic storms examined the rates and characteristics of

sferics, the electromagnetic noise radiated by lightning. These measurements were difficult to interpret, because results depended partly on the characteristics of the sferics receiver, and because sferics instrumentation provided at best only the bearing to the lightning flash generating the sferics and often provided no location information at all. Furthermore, many of the sferics studies lacked radar data, and only one had any Doppler radar data. Equally important, there was not an adequate appreciation of the variety of storms that produce tornadoes.

Some of these limitations were overcome when systems

were developed to map the location of lightning flashes. For example, beginning in the late 1970s, various organizations began deploying newly developed systems for automatically mapping where lightning channels strike ground over ranges of a few hundred kilometers [e.g., Krider et al., 1980; Orville et al., 1983;Bent and Lyons, 1984; MacGorman and Taylor, 1989]. Now coverage by these systems is continual across the contiguous United States, with errors in strike locations typically less than 10 km [e.g., Mach et al., 1986; MacGor- man and Rust, 1989].

In this paper we first review what was learned about tornadic storms from sferics studies and then consider what

The Tornado: Its Structure, Dynamics, Prediction, and Hazards. Geophysical Monograph 79 This paper is not subject to U.S. copyright. Published in 1993 by the American Geophysical Union.

has been learned from analysis of data from lightning mapping systems.

2. LIGHTNING TERMINOLOGY

This section explains lightning terminology and concepts that will be used in the rest of the paper. Lightning can be classified into two types: (1) Cloud-to-ground lightning has at least one channel spanning from the cloud to the ground. (2) Intracloud lightning does not have a channel to ground. Because the locations of lightning channels relative to visible cloud boundaries are difficult to detect remotely, scientists who study lightning and storm electrification normally group in-cloud, cloud-to-air, and cloud-to-cloud lightning together in the intracloud lightning classification. Cloud-to-ground lightning can be further classified by the polarity of charge that it effectively lowers to ground: Positive cloud-to-ground lightning lowers positive charge; negative cloud-to-ground lightning lowers negative charge. Most cloud-to-ground flashes are negative.

A cloud-to-ground lightning flash usually begins inside the cloud and is first apparent when a faint channel, called the stepped leader, moves from the cloud to the ground in jumps roughly 50-100 m long. When the stepped leader connects with the ground, a bright pulse moves back up the lightning channel in a process called a return stroke. After a pause of roughly 20-150 ms, another leader can travel back down the already established lightning channel, followed by another return stroke. The combination of a leader and return stroke is called a stroke. All strokes going through essentially the same channel to ground make up a single cloud-to-ground flash. There can be anywhere from one to a few tens of strokes in a flash.

Sferics are generated by changes in the vector electric current, the frequency of the radiation depending on the time scale of the changes. Since currents in a lightning channel increase faster than they decay, a lightning channel will generally radiate higher frequencies during an initial current surge than during its decay. The time required for a current pulse to propagate through a channel segment depends on the length of the segment, so longer channels radiate more energy at lower frequencies, if all else is equal.

!74 REVIEW OF LIGHTNING IN TORNADIC STORMS

The spectrum of sferics from a typical return stroke peaks at roughly 5 kHz. Most strong signals at frequencies below about 30 kHz are from cloud-to-ground return strokes and so occur in discrete bursts, separated by tens of milliseconds. At frequencies near or above 1 MHz, the amplitudes of sferics from return strokes, from other processes of a cloud-to-ground flash, and from intracloud flashes are all comparable. Furthermore, at these higher frequencies, sferics are more numerous and more continual throughout a lightning flash, and relatively few are associated with return strokes.

The vertical component of sferics measured at the stirface of the conducting Earth is generally much larger than the horizontal, even if the two components are comparable at the source. Any horizontal component that exists is reduced by currents that are induced in the ground by the stZrics. 'l"hcrefore a sferics receiver at the ground receives little of the energy radiated by horizontal components of current change. Airborne receivers are less affected by currents induced in the ground and so are better able to detect the horizontal components of sferics.

3. SFERICS STUDIES

Initial studies of sferics from tornadic storms examined

radio bands f¾om 10 to 500 kHz. (Sferics normally are classified by the frequency of the receiver used to detect them.) Dickson and McConahy [1956] found that sferics rates detected by a 10-kHz receiver often increased as storms grew rapidly taller, but rates peaked earlier in storms that had more violent severe weather. Rates peaked about 1.5 hour before tornadoes and decreased during tornadoes to about 40% of peak value. At frequencies of 150 kHz and higher, sferics rates were found to increase until tornadoes were produced, and rates became exceptionally high during periods when tornadoes occurred [jones, 1951, I958; Jones and Hess, 1952; Kohi, 1962; Kohl and Miller, 1963]. Kohl and Miller [1963] observed that sferics rates at 150 kHz usually peaked during severe weather and began decreasing prior to the end of severe weather. As might be expected from the above observations, Jones and Hess [1952] found that the ratio of the numbers of higher-frequency to lower- frequency sferics increased for sferics of large amplitude when storms were more severe; the ratio was 1:20 before and after tornadoes and in nonsevere thunderstorms versus

1:1 during tornadoes. Later studies at frequency bands up to 150 MHz [Silberg,

1965; Taylor, 1973; Stanford et al., 1971; Johnson et al., 1977] found that the increase in sferics rates during severe weather was greatest at frequencies above 1 MHz. This is shown in Figure 1, which presents a summary of Taylor's data concerning the dependence of sferics rates on receiver frequency. Silberg [1965] and Taylor [1973] also found that the energy radiated as sferics (not shown) increases at higher frequencies as storm severity increases; radiated energy peaks at roughly 5 kHz for nonsevere storms [Taylor, 1972], but the energy at higher frequencies gradually increases with

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Fig. 1. Average sferics burst rate as a function of receiver frequency for various storm classifications [from MacGorman et al., 1989]. Data are from four tornadic storms, three nontornadic severe storms, and seven nonsevere thunderstorms. Sferics burst rates were measured at 10 receiver frequencies between 10 kHz and 137 MHz, indicated by dots. There was little change above 3 MHz, so higher frequencies are not shown.

increasing severity, until for tornadic storms, the energy is independent of frequency.

There were exceptions to all of these findings. In several cases, high sferics rates did not occur in tornadic storms [Ward et al., 1965; Stan3'brd et al., 1971; Lind et al., 1972; Taylor, 1973; Greneker et al., 1976; Johnson et al., 1977; MacGorman et al., 1989]. Taylor, who examined the largest number of storms, found that roughly 20% of tornadic storms did not have high sferics rates [MacGorman et al., 1989]. It also was possible for high sferics rates to occur in storms without tornadoes [Ward et al., 1965; StanJbrd et al., 1971; MacGorman et al., 1989]. Taylor found that although only 23% of nontornadic severe storms and 1% of nonsevere storms had sferics rates comparable to those of tornadic storms, 70% of the storms with high sferics rates were not tornadic, because the number of nonsevere storms was much larger than the number of tornadic storms [MacGor- man et al., 1989]. Some tornadic storms without high sferics rates had only weak tornadoes [e.g., Johnson et al., !977], and the combined effect of several simultaneous storms

explains some nonsevere storms with high sferics rates [e.g., MacGorman et al., 1989]. However, these situations did not account for all exceptions to the rule that 3-MHz sferics rates are much higher in tornadic storms than in nonsevere storms.

The most obvious inference from the increasing number of sferics at high frequencies and the decreasing or constant number at low frequencies was that intracloud flash rates usually increase while cloud-to-ground lightning return strokes decrease or remain fairly constant during the tornadic stage of storms. This interpretation was further supported by the theoretical work of Stanford [1971] and

MAcGORMAN 175

airborne observations of Shanmugatn and Pybus [1971]. Both reported that horizontal polarization of sferics increased and vertical polarization decreased, indicating a greater preponderance of horizontal lightning channels, as the sferics rate and storm severity increased.

Other observations also supported this interpretation. Jones [ 1958] reported watching a storm when it was produc- ing high sferics rates about 1 hour before a tornado; no ground flashes could be seen, but relatively dim circular patches, about 0.5 km in diameter, were illuminated in rapid succession on the side of the storm. He suggested that these were an unusual form of intracloud lightning. Electric field records for this same storm indicated that the lightning flash rate was 12 min -1 during the tornado [Gunn, 1956]. Von- negut and Moore [1957] inferred from recordings of the closest electric field sensor to the tornado that there were no

cloud-to-ground flashes near the tornado. Unusually high lightning flash rates were also observed in two other tornadic storm systems by Orville and Vonnegut [1977] and Turrnan and Tettelbach [1980], who used satellite-borne optical de- tectors. Films of 17 tornadoes and one funnel cloud docu-

mented low ground flash rates near tornadoes; Davies-Jones and Golden [1975a] (see also Davies-Jones and Golden [1975b, c], Colgate [1975], and Vonnegut [1975] for a discussion) found cloud-to-ground lightning flashes near only two of the 18 filmed vortices. One of these two was a tornado

with 12 ground flashes; the other was a funnel cloud with two ground flashes.

From studies of sferics, however, it appears that the tornado itself is not likely to be the cause of unusual sferics rates. For example, Jones [1965] noted that sferics rates can be unusually high 60-90 min before tornadoes, although they may increase further when tornadoes occur. Scouten et al. [1972] further analyzed the data of Jones [1958] at 150 kHz and concluded that touchdown of the tornado did not affect

the high sferics rate. Although larger-amplitude sferics clustered more closely about a central core near the azimuth of the tornado, sferics as a whole were spread among azimuths spanning most of the storm. A similar lack of correlation with the time and location of tornado touchdown was

reported for sferics in three bands between 300 kHz and 3 MHz [Taylor, 1973] and for sferics at 1-50 kHz [Brown and Hughes, 1978]. Brown and Hughes [1978] also noted a lack of obvious relationships with the tornadic vortex signature [Brown et al., 1978]. In a study of sferics burst rates at high frequencies (30-300 MHz) in two tornadic storms, Johnson [1980] found that the tornadoes began near the time of peaks in the sferics burst rate but were not coincident with any peaks.

Several studies considered changes in tornadic storms that were correlated with the high sferics rates. For example, several investigators found periodicities of 10-60 min in the sferics data and suggested that these resulted from the cyclical intensification and decay of thunderstorm cells [Kohl and Miller, 1963; Stanford et al., 1971; Shantnugam and Pybus, 197!; Lind et al., 1972; Trost and Nornikos, 1975]. Taylor [1973] reported that sferics at high frequencies

originated along azimuths to low-level reflectivity cores and that sferics rates were larger from cores with higher reflectivity. Furthermore, Brown and Hughes [1978] found that the pattern in the number of VLF sferics (which are produced predominantly by return strokes) from the Union City tornadic storm was similar both in time and space to patterns in reflectivity at low levels of the storm.

Although the sferics studies clearly demonstrated that lightning tended to evolve in characteristic patterns in tornadic storms, there were a number of uncertainties. (1) There were uncertainties in extracting lightning flash type and rates from sferics data. (2) No sferics study examined the evolution of mesocyclones or considered differences in the types of storms that produced tornadoes in order to examine why sferics rates evolved differently. (3) No sferics study examined the three-dimensional vector wind field for tornadic storms.

4. LIGHTNING MAPPING STUDIES

Studies using lightning mapping systems have begun to address some of the shortcomings of the sferics studies. However, only two of the mapping studies thus far have included intracloud lightning; all others have examined only cloud-to-ground lightning. Furthermore, mapping studies have found considerable variability in the relationship of tornadoes to cloud-to-ground flash rates. In two tornadic storms, ground flash rates were small before and during tornadoes [Orville et al., 1982; MacGorman et al., 1989]. In three, ground flash rates peaked during tornadoes [Kane, 1991; MacGorman and Nielsen, 1991; Keighton et al., 1991]. In two, tornadoes occurred after ground flash rates had increased but when ground flash rates were not at their maximum [Kane, 1991; MacGorman et al., 1985].

Results from two of the studies suggest some reasons fbr this variability. In the Binger tornadic storm studied by MacGorman et al. [1989], negative ground flash rates within 10 km of the mesocyclone center were less than 1 min -• until the tornado began dissipating. Negative cloud-to- ground flash rates reached a relative maximum after the last tornado, as its mesocyclone core dissipated, and reached an absolute maximum (about 4 min -•) as the last mesocyclone core in the storm dissipated (see Figure 2). Intracloud flash rates reached an absolute maximum of approximately 14 min -• during the most violent tornado and were well correlated with low-level cyclonic shear when the shear was greater than 1 x 10 -2 s- • during the tornado. (Cyclonic shear is one half vertical vorticity for ideal solid-body rotation. In calculating this shear, tornadic winds were ignored; shear was measured across the diameter of maximum tangential wind speed for the mesocyclone, typically a distance of a few kilometers.) After the last tornado, when there was a large decrease in low-level cyclonic shear but the mesocyclone was still strong at middle levels, intracloud flash rates were only 0-2 min -• .

In the Edmond storm studied by MacGorman and Nielsen [1991] there were no intracloud data, but negative

176 REVIEW OF LIGHTNING IN TORNADIC STORMS

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flash rates within 10 km of the mesocyclone center evolved differently than in the Binger storm. Rates increased to a peak of 11 min- • during tornadic activity and appeared to be correlated with cyclonic shear at the 5 km level (Figure 3). Ground strike locations tended to cluster in the vicinity of a reflectivity core north of the mesocyclone during

MAcGORMAN 177

activity and to be more scattered before and after the tornadic stage of the storm (Figure 4). Keighton et al. [1991] observed a similar increase in scattering of lightning strikes after tornadic activity on another day.

MacGorman and Nielsen [1991] suggested that observed differences between the Binget and Edmond storms might account for the differences in the evolution of their lightning flash rates. The Binget storm was a classic supercell storm, with a prominent weak-echo region, a strong, deep updraft core, and a long-lasting mesocyclone having a family of mesocyclone cores [Burgess et al., 1982]. The Edmond storm also was a supercell storm but was weaker than the Binger storm by almost any measure: the duration of its supercell stage was only 30% that of the Binget storm; its mesocyclone was weaker, shallower, and shorter lived; and features such as the weak-echo vault and tornadic vortex

signature were not as pronounced. In fact, the Edmond storm did not form a tornado until the outflow boundary of another storm began to overtake it from the west. MacGor- man and Nielsen [1991] suggested that the increase in flash rates in both storms when the mesocyclone was strongest probably was a result of the strong updraft that existed then. They also suggested, however, that the deeper and stronger updraft core and weak echo region of the Binget storm probably delayed ground flash activity and caused intracloud activity to dominate.

To understand the hypothesis that MacGorman and Nielsen [1991] offered to explain why ground flash rates behaved differently in these two storms, it is necessary first to understand a little about thunderstorm electrification. One important class of thunderstorm electrification mechanisms relies on microphysical particle interactions to place charge of one sign on graupel and charge of the opposite sign on cloud ice. Such mechanisms are affected (1) by factors that control the sizes and numbers of graupel and ice particles that collide, and (2) by factors that control subsequent transport of the charged particles into different regions of a thunderstorm. Factors affecting the number of relevant collisions include the strength and depth of the updraft and the microphysics of the storm. Subsequent transport is dominated initially by the sedimentation of graupel but may become dominated by the relative velocity of the winds in different regions as the graupel and ice crystals move farther apart. Thunderstorm measurements [e.g., KrehbieI, 1981; Byrnne et al., 1983; Chauzy et aI., !985; Dye et al., I986; Koshak and Krider, 1989] suggest that most of the main negative charge at middle levels of the storm is on graupel and precipitation, while the main positive charge is generally on ice crystals at colder temperatures (the next section will discuss one possible exception).

Several investigators have observed that a strengthening updraft at altitudes colder than -20øC increases flash rates (see discussion by MacGorman et al. [!989]). This probably explains the observed increases in flash rates during the tornadic stage of the Binget and Edmond storms. MacGor- man et al. [1989] also observed increasing areas of large reflectivity at heights between 6 km and 8 km shortly before

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and during periods of high flash rates. They suggested that the increased flash rates probably were caused by increased particle interactions leading to reflectivity growth at 7-9 km in and near the strong updraft. Similar correlations of lightning rates with the horizontal area of large reflectivity at midlevels of storms have been observed, for example, by Larson and Stansbury [1974], Lhermitte and Krehbiel [1979], and Keighton et al. [1991].

One of the primary problems posed by the Binger and Edmond storms is explaining why ground flash rates were so small near and during the time when total and intracloud flash rates were large in the BingeT storm and why this apparently was not true for the Edmond storm. MacGorman et al. [1989] suggested that the strong, deep updraft and the resulting pronounced weak-echo region of the Binger storm enhanced production of intracloud flashes and delayed or suppressed production of ground flashes because of effects on the thunderstorm charge distribution.

For example, one primary effect on the charge distribution is to keep negative charge higher than in most storms. This would increase the energy required for lightning to span the distance to ground and would decrease the electric field at most heights below the negative charge. Negative charge is higher for three reasons. (1) The strong updraft core rapidly lifts all but the largest particles to upper levels of the storm. (2) Large updraft speeds do not allow most hydrometeors to remain long in a given layer. The short residence time prevents hydrometeors from acquiring much charge and prevents positive and negative particles from moving far apart after becoming charged. Therefore there is little net negative or positive charge in a given layer of a strong updraft. (3) Temperatures in stronger updraft cores are generally warmer, and this causes the ice processes thought to be responsible for charging (see discussion of II!ingworth [1985]) to occur somewhat higher.

A second important effect of a strong, deep updraft is that regions of net positive and negative charge would be relatively close together near the top of the storm updraft. This occurs for two reasons. (1) As noted above, sedimentation and differing relative velocities of positively and negatively charged particles would not have time to separate the charges very far in a fast updraft. (2) The vertical shear of horizontal winds in a classic supercell storm causes graupel to fall to the side of the updraft core, so much of the storm-scale separation of charge occurs outside the updraft.

When two charged regions are closer together, less net charge is needed in each region to create electric field magnitudes sufficient to initiate lightning between them. This enhances intracloud flash rates. It also reduces the negative charge available for cloud-to-ground lightning near the updraft, because the threshold for initiating lightning regulates the net negative (and positive) charge that can be separated. This, combined with the unusually large height of negative charge discussed above, makes cloud-to-ground lightning much less likely to occur near a very strong, deep updraft core. Ground flash rates increase later, as the negative charge moves closer to ground and as the positive and

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negative charges move farther apart (thereby increasing the amount of net negative charge that can be separated before causing intracloud lightning).

Since MacGorman et al. [ 1989] first suggested this hypothesis, there has been support from two sources. First, the lowest charge detected by the only published balloon-borne electric field sounding through a mesocyclone updraft was a region of negative charge starting at 9.5 km above mean sea level, where the temperature was approximately -37øC in the storm's environment and -31øC in the updraft [Rust et al., 1990]. This lower boundary is 3-5 km higher, and at environmental temperatures 20ø-30øC colder, than the lower boundary normally reported for the main negative charge in continental thunderstorms. Second, Ziegler and MacGor- man [1990] used a kinematic retrieval that included electrification processes [Ziegler et al., 1991] to model the electric field and charge structure of the Binger storm. In agreement with the balloon-borne measurements, the lower boundary of negative charge inside the updraft core in the retrieved charge distribution was elevated to a higher altitude than normal (Figure 5).

MacGorman and Nielsen [1991] suggested that the charge distribution near the updraft of the Edmond storm was more like that of an ordinary thunderstorm and that cloud-to- ground flashes therefore were able to occur frequently when the mesocyclone was strong. They concluded that negative charge was lower, a larger fraction of oppositely charged particles had time to separate, and positive and

MAcGORMAN 179

charge were farther apart because (1) the updraft core was weaker, shallower, and shorter lived, (2) the resulting weak- echo region was shallower and contained larger reflectivi- ties, and (3) the 55-dBZ reflectivity cores extended to a height of only 5-6 km in the Edmond storm versus over 10 km in the Binger storm. Similarly, Keighton et al. [1991] observed another storm in which cloud-to-ground flash rates did not decrease during the tornadic stage and noted that it also had a weaker mesocyclone than the Binger storm.

5. POSITIVE CLOUD-TO-GROUND

LIGHTNING IN TORNADIC STORMS

Data on the less common flashes that effectively lower positive charge to ground (+CG flashes) have been studied in detail for only one tornadic storm, the Edmond storm investigated by MacGorman and Nielsen [1991]. (Reliable positive cloud-to-ground data were not available for the Binger storm or for other tornadic storms before 1983.) In the Edmond storm, positive cloud-to-ground flashes began to occur a few minutes before the tornado touched down,

and positive ground flash rates increased to their maximum value during the tornado, near the time when negative ground flash rates also peaked (see Figure 3). MacGorman and Nielsen [1991] noted from watching ground flash dis- plays for several tornadic storms in real time that positive ground flashes often occurred during tornadoes. Sometimes a few positive ground flashes were clustered on the southern flank of storms in which negative cloud-to-ground lightning still dominated ground flash activity (as in the Edmond storm), and sometimes most of the ground flashes in the storm were positive ground flashes.

Although many severe or tornadic storms do not have large numbers of positive ground flashes, preliminary evi- dence suggests that when relatively high flash rates and densities of positive cloud-to-ground lightning occur, they are associated with severe weather. Reap and MacGorman [1989] reported in a climatological study that there was a correlation between a high density of positive ground flashes and severe weather. Furthermore, in August 1990 and spring 1991, MacGorman and Burgess [1991] observed several storms that were unusual because (1) most cloud-to-ground flashes were positive ground flashes and (2) ground strike points of positive cloud-to-ground lightning occurred in dense clusters, much like the clusters of negative ground flashes observed in most electrically active storms. (Positive ground flashes are generally infrequent and diffuse.) All of the storms with frequent positive ground flashes produced large hail, and many produced tornadoes. Similar observations were reported by Rust et al. [1985], Curran and Rust [1992], and Branick and Doswell [1992].

These storms with unusual positive lightning activity appear to occur in at least two modes. In one, positive ground flashes dominate cloud-to-ground activity throughout the lifetime of the storm, including all tornadic activity. An example of a storm in this mode is shown in Figure 6, which presents ground flash rates for the storm that produced an F4

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Fig. 6. Time series plots of the number of positive and negative cloud-to-ground flashes in each 15-min interval during a storm that produced a violent tornado near Hutchinson, Kansas, on March 26, 1991. Bars indicate the time of tornadoes, labeled with F scale ratings. The approximate time of large hail reports are indicated by the leading edge of Hs.

tornado near Hutchinson, Kansas, on March 26, 1991. This storm produced severe weather throughout much of its life.

In the second mode the dominant polarity of ground flashes in the storm switches from positive to negative, with the peak negative ground flash rates comparable to or larger than the earlier peak positive ground flash rates. An example of the second mode is given in Figure 7, which shows ground flash rates for the storm that produced an F5 tornado in Plainfield, Illinois, on August 28, 1990. Prior to the F5 tornado most ground flashes in the storm were positive. During this period, large hail and four tornadoes rated up to F2 were produced. Three of these tornadoes occurred during the period of maximum positive ground flash rates.

The dominant polarity of cloud-to-ground lightning switched as the Plainfield tornado began. Most subsequent ground flashes were negative, as shown clearly in the time series plots in Figure 7. Also note that there was a decrease in overall ground flash activity shortly before and during the F5 tornado, much as described earlier for negative ground flashes in some tornadic storrns. Negative ground flash rates increased to their largest values at the end of the time series plot, when the storm was merging with other storms to form a squall line (similar to lightning behavior observed


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Goodman and MacGorman [1986] when storms merged to form a mesoscale convective complex).

Visual observations indicate that the character of the

Plainfield storm changed at approximately the time when the last and most violent tornado began and ground flash polarity switched. Initially, the storm grew rapidly into a supercell, and it remained a supercell for much of its life. During this period the storm had a prominent, deep weak-echo region and produced large hail. At about the time that the Plainfield tornado began, heavy precipitation began to wrap around the mesocyclone, filling in the weak-echo region. Observers of the Plainfield storm noted that the tornado was obscured by heavy precipitation.

Curran and Rust [1992] and Branick and Doswell [1992] have observed that low-precipitation supercell storms often produce high densities of positive ground flashes and that nearby supercell storms with heavy precipitation produce the normal preponderance of negative ground flashes. At least some severe storms (including the earlier stages of the Plainfield storm) that have high numbers and densities of positive ground flashes are not low-precipitation storms, but there are no reports of this positive cloud-to-ground lightning signature in high-precipitation supercell storms. Therefore it appears that the transition of the Plainfield storm to a

high-precipitation supercell was related to the transition in ground flash polarity. A similar transition was observed by Curran and Rust [1992].

MacGorman and Nielsen [1991] suggested two reasons why positive ground flashes might occur in severe storms. First, the structure of supercell storms, in which precipitation at middle to low levels is displaced horizontally from the updraft core, might lead to a tilted bipolar charge distribution, with a lower negative charge occurring in the vicinity of precipitation and positive charge occurring on ice crystals at the top of the updraft core, above the weak echo region. The tilt might be sufficient for the upper positive charge to be effectively exposed to the ground, as for winter storms observed by Brook et al. [1982]. Branick and Doswell [1992] have suggested a variation of this: not only does the horizontal displacement of the updraft and precipitation tilt the dipole, but the lack of low-to-middle level precipitation in low-precipitation storms inhibits formation of a large region of negative charge.

The second mechanism discussed by MacGorman and Nielsen [1991] involved a possible region of significant positive charge beneath the main negative charge of the thunderstorm. Such a region could be formed in different ways. For example, several investigators have suggested that a lower positive charge could be formed in thunderstorms by a noninductive mechanism studied in the laboratory by Takahashi [1978], Jayaratne et al. [1983], and others. In most regions in which this mechanism for microphysical charge separation might occur in thunderstorms, collisions between graupel and ice crystals place negative charge on graupel and positive charge on ice crystals. These collisions result in the normal tendency for the main positive charge to be generally above the main negative charge. At temperatures warmer than some threshold between -10 ø and -20øC, however, this is reversed: positive charge is placed on graupel; negative charge, on ice crystals. If enough collisions occur between 0øC and this reversal temperature, they could create a region of positive charge on graupel beneath the main negative charge of thunderstorms.

6. CONCLUSION

Obviously, the studies reviewed in this paper fall far short of fully characterizing the diversity of lightning evolution in tornadic storms. Most studies examined lightning behavior relative just to the existence of tornadoes. Only a few studies examined lightning evolution relative to the evolution of the reflectivity structure, mesocyclone, or wind field of tornadic storms. To test adequately the hypothesized relationship between updraft characteristics and suppression of ground flashes requires a more systematic study of large data sets of lightning and storm characteristics, possibly supplemented by numerical simulations. Such studies also may be able to determine better the underlying relationship that caused the observed close correlation of low-level cyclonic shear and intracloud flash rates in the Binger storm.

If the hypotheses discussed in this paper are correct,

MAcGORMAN 181

expect that more comprehensive studies of tornadic storms will find the following. Almost all storms that have mesocyclones will have an increase in intracloud and total flash rates as the updraft increases in depth, horizontal extent, and speed at altitudes above the -20øC level, thereby increasing the number of hydrometeor collisions that generate charge. Intracloud flash rates will be enhanced even more in classic ,supercell storms, which tend to have very large and strong updrafts that are quasi-steady, in contrast to updrafts in lesser thunderstorms. In some tornadic storms, cloud-to- ground flashes will begin and peak shortly after intracloud lightning (roughly 10-15 rain later), as in most nonsevere storms. However, in storms that have updrafts large and strong enough to create prominent, deep weak-echo regions (with high lower boundaries of negative charge) we expect cloud-to-ground lightning near the updraft and mesocyclone to be suppressed, so that few ground flashes occur when the updraft is strong and ground flash rates peak when the updraft finally weakens.

There are far fewer data on positive cloud-to-ground lightning in severe storms, so observed relationships and suggested hypotheses should be regarded as more tentative. If a tilted charge distribution causes positive cloud-to-ground lightning, then we would expect to see positive ground flashes in storms in which the top of the main updraft is displaced sufficiently far horizontally from the reflectivity core. If the described mechanism for the lower positive charge causes positive cloud-to-ground (+CG) flashes, then we would expect to see + CG lightning in storms in which the number of collisions between graupel and ice crystals is enhanced at temperatures between freezing and roughly -15øC. Such a situation could occur, for example, if there is significant recirculation of graupel, as has been reported for some hail storms.

It may be difficult to determine which of these or other mechanisms are responsible for positive cloud-to-ground lightning. For example, the large separation reported between the mesocyclone and precipitation for some storms in which positive ground flashes dominate [Branick and Doswell, 1992] could result from either scenario. There probably would be significant tilt in the charge distribution, and a significant number of recirculating graupel particles might pass through the layer between 0 ø and -15øC. Distin- guishing which mechanism dominates will require at least some of the following: simultaneous electric field measurements in upper regions of the updraft and in heavy precipitation, improved Doppler radar data, simulations of severe storms, and three-dimensional lightning mapping systems.

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