Radar and Lightning Observations of Deep Moist Convection across Northern Alabama during DC3: 21 May 2012

Retha Matthee Mecikalski Atmospheric Science Department, University of Alabama in Huntsville, Huntsville, Alabama

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Anthony L. Bain Atmospheric Science Department, University of Alabama in Huntsville, Huntsville, Alabama

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Lawrence D. Carey Atmospheric Science Department, University of Alabama in Huntsville, Huntsville, Alabama

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Abstract

The Deep Convective Clouds and Chemistry (DC3) experiment seeks to understand the kinematic and microphysical controls on the lightning behavior of deep moist convection. This study utilized multiple dual-polarization Doppler radars across northern Alabama to quantify microphysical and kinematic properties and processes that often serve as precursors to lightning, such as the graupel echo volume, graupel mass, and convective updraft volume. The focus here was on one multicellular complex that occurred on 21 May 2012 in northern Alabama during DC3. The graupel echo volume and the graupel mass in the charging region correlated well with the total lightning flash rate (FR), and even better than the updraft volumes and maximum updraft velocities. The flash length scales (LS) and flash areas were generally anticorrelated to the FR, while it was correlated to the nonprecipitation ice volume. More specifically, the presence of smaller flashes was associated with a stronger lower positive charge region caused by larger graupel volumes, stronger updraft volumes, and stronger maximum updraft velocities while larger flashes occurred during lower FRs and were associated with a weakened lower positive charge region in combination with a stronger upper positive charge region, weaker updraft velocities, a smaller graupel volume and mass, and an increase in nonprecipitation ice volume.

Corresponding author address: Retha Matthee Mecikalski, Atmospheric Science Department, University of Alabama in Huntsville, National Space Science and Technology Center, 320 Sparkman Dr., Huntsville, AL 35805–1912. E-mail: retha.mecikalski@nsstc.uah.edu

Abstract

The Deep Convective Clouds and Chemistry (DC3) experiment seeks to understand the kinematic and microphysical controls on the lightning behavior of deep moist convection. This study utilized multiple dual-polarization Doppler radars across northern Alabama to quantify microphysical and kinematic properties and processes that often serve as precursors to lightning, such as the graupel echo volume, graupel mass, and convective updraft volume. The focus here was on one multicellular complex that occurred on 21 May 2012 in northern Alabama during DC3. The graupel echo volume and the graupel mass in the charging region correlated well with the total lightning flash rate (FR), and even better than the updraft volumes and maximum updraft velocities. The flash length scales (LS) and flash areas were generally anticorrelated to the FR, while it was correlated to the nonprecipitation ice volume. More specifically, the presence of smaller flashes was associated with a stronger lower positive charge region caused by larger graupel volumes, stronger updraft volumes, and stronger maximum updraft velocities while larger flashes occurred during lower FRs and were associated with a weakened lower positive charge region in combination with a stronger upper positive charge region, weaker updraft velocities, a smaller graupel volume and mass, and an increase in nonprecipitation ice volume.

Corresponding author address: Retha Matthee Mecikalski, Atmospheric Science Department, University of Alabama in Huntsville, National Space Science and Technology Center, 320 Sparkman Dr., Huntsville, AL 35805–1912. E-mail: retha.mecikalski@nsstc.uah.edu

1. Introduction

The Deep Convective Clouds and Chemistry (DC3) experiment occurred during May and June 2012 and was an interdisciplinary study with the goal to investigate and understand the relationship between the microphysical, kinematic, and electrical properties of deep moist convection (DMC; Barth et al. 2015). The properties of these storms and the associated environment can be examined using a variety of remote sensing platforms. The use of dual-polarization Doppler weather radars results in improved observation of hydrometeors and associated microphysical processes that occur within these storms. The microphysical processes can be linked to the kinematic and ultimately electrical structure of DMC. In addition, three-dimensional kinematic flows and estimates of vertical motion are determined by using multi-Doppler wind synthesis techniques. This is particularly desirable to atmospheric chemists, as one of the main objectives of DC3 is to understand DMC’s role in the creation and transport of natural and anthropogenic atmospheric constituents throughout the depth of the troposphere.

Cloud electrification is also known to play a pivotal role in the production of nitrogen oxides (NOx). The electrical characteristics of DMC during DC3 were documented via very high-frequency (VHF) lightning mapping arrays (LMA) and low-frequency (LF) to very low-frequency (VLF) sensors. The motivation behind the careful documentation of lightning during DC3 was twofold: 1) it is hypothesized that lightning can be a useful parameter for describing the intensity of DMC [especially in the absence of weather radar; Byers and Braham (1948); Carey and Rutledge (1996, 2000); Pickering et al. (1998); Lang and Rutledge (2002); Deierling et al. (2008); Barthe et al. (2010)] and 2) it has widespread use in the atmospheric chemistry modeling community. Numerical modeling of lightning flash type, rate, and length are important because of their hypothesized role in the production of NOx (Price et al. 1997; Pickering et al. 1998; Wang et al. 1998; DeCaria et al. 2000, 2005; Dye et al. 2000; Ott et al. 2007; Barthe and Barth 2008; Koshak et al. 2014).

DC3 had three main objectives: 1) compare and contrast environmental differences between DMC with varying microphysical, kinematic, and electrical properties; 2) document microphysical and kinematic processes that are thought to be relevant to cloud electrification and lightning production; and 3) provide a useful radar observable total lightning characteristic relationship that can be used to parameterize the production of NOx via lightning (LNOx) in numerical cloud models that do not include explicit electrification and lightning processes. For this paper, the main purpose is to study a convective storm complex over northern Alabama that occurred on 21 May 2012 during DC3 and provide a detailed synopsis of the environmental conditions as well as the coevolution of the kinematics, microphysics, and lightning behavior of the storm in order to provide data for developing total flash rate (FR) parameterizations that might be helpful in numerical cloud models. The storm on 21 May 2012 is of particular interest as, presently, there are other studies focusing on this storm, which include analysis of airborne chemical and aerosol measurements and modeling of storm structure, flash rates, and LNOx parameterizations (Barth et al. 2015; Carey et al. 2014a,b).

2. Background

As mentioned, DC3 aims to quantify the relationship between DMC and its role in the production and transport of LNOx and other atmospheric constituents throughout the troposphere. The generation of LNOx is outlined by Barthe et al. (2010) as one of the more important atmospheric chemistry goals. Lamarque et al. (1996) and Pickering et al. (1998) illustrated through the use of global chemical models that lightning is one of the leading natural causes of NOx in the upper troposphere.

There is currently some disagreement in the literature as to whether the lightning type [e.g., cloud-to-ground (CG) or intracloud (IC)] has any effect on the amount of LNOx generated per flash. Some modeling studies suggest that there is roughly an order of magnitude difference in LNOx production between CG and IC flashes (Price et al. 1997; Pickering et al. 1998; Koshak et al. 2014). Price et al. (1997) state that CG flashes produce more LNOx than IC flashes based on the stronger energetics associated with the return strokes of CG flashes. Pickering et al. (1998) found that parameterized CG flashes produced larger instantaneous LNOx values versus IC flashes. In a modeling study that utilizes three-dimensional lightning mapping array observations as input, Koshak et al. (2014) found roughly an order of magnitude greater LNOx produced per CG flash than per IC flash. On the other hand, the observational studies of DeCaria et al. (2000, 2005) and Ott et al. (2007) suggest that CG and IC flashes may produce approximately the same amount of LNOx per flash. In a field study by Dye et al. (2000), it was determined that IC flashes were the more significant contributor to LNOx in an isolated Colorado supercell with a high IC to CG (IC:CG) ratio. Climatologically speaking, Boccippio et al. (2001) noted that a region of anomalously high IC:CG ratios existed across portions of northeastern Colorado, southwestern Nebraska, and northwestern Kansas. Elsewhere, including Alabama, the IC:CG ratio is relatively low. Barthe and Barth (2008) performed model-simulated sensitivity studies, which showed that varying the IC:CG ratio yielded very little difference in the LNOx profile.

In addition to flash type possibly being important in the modeling of LNOx, other scientists argue that areal flash length may also be important (Wang et al. 1998, Barthe and Barth 2008). In fact, Barthe and Barth (2008) stated that when the same amount of NO molecules per meter of flash was calculated, 80% of the LNOx was produced by flashes >30 km in extent even though these flashes only composed 30% of the total flashes in their analysis. Furthermore, Carey et al. (2005), Dye and Willett (2007), Hodapp et al. (2008), Ely et al. (2008), Bruning and MacGorman (2013), and Calhoun et al. (2013) showed in one way or another that flashes developing within the convective core and updraft region of mesoscale convective systems (MCSs) and supercells occurred more frequently, but with smaller areal extents than those that developed and/or propagated farther away from the convective core (i.e., in the stratiform/anvil regions of MCS and supercells), which were less frequent but larger in areal extent. Most of these studies either focused on supercells or MCSs and inadvertently found that larger flashes occurred later in the storms’ life cycle and/or outside of the convective core. Bruning and MacGorman (2013), however, specifically studied the flash rate versus flash size relationship in a supercell and explicitly found that these two variables are anticorrelated.

Moreover, larger hydrometeors that develop in the convective regions of MCSs (and other types of storms) will fall out due to gravitational forces and large fall speeds, while smaller (possibly charged) particles with slower fall speeds will be lofted and spread out laterally at the top of the cloud due to divergence, leading to the formation and growth of the stratiform (or anvil) region (Houze 1997). In studies performed by Dye et al. (2007) and Dye and Willett (2007), it was shown that anvils with the reflectivity factor at horizontal polarization (ZH) > 10 dBZ had electric fields >10 kV m−1, indicating that these regions are indeed charged. Currently, there does not seem to be much other research (if any) regarding parameters that control flash size, such as nonprecipitation ice volumes in anvil/stratiform regions. Thus, calculating and comparing the volume of ZH > 10 dBZ in anvil/stratiform regions of ordinary multicellular convection to flash size could show high correlations, making this fruitful research (Weiss et al. 2012; Calhoun et al. 2013).

Numerous studies have also specifically examined the relationship between kinematic and microphysical properties of DMC and their electrical behavior. Goodman et al. (1988), Carey and Rutledge (1996, hereafter CR96), Jameson et al. (1996), Bringi et al. (1997), Carey and Rutledge (2000, hereafter CR00), Lang and Rutledge (2002), Wiens et al. (2005), Kuhlman et al. (2006), Bruning et al. (2007), Deierling and Petersen (2008), and Deierling et al. (2008) have all reported that kinematic and microphysical radar properties trend well with total lightning FR. The most important finding is that ice processes occurring in the mixed phase region (0° to −40°C) are critical to cloud electrification. Zipser and Lutz (1994) examined midlatitude continental, tropical oceanic, and tropical continental convection and suggested that mean updraft velocities of 6–7 m s−1 or maximum updraft velocities of 10–12 m s−1 were sufficient to facilitate cloud electrification. It was hypothesized that updrafts of these magnitudes were capable of lofting rain into the mixed-phase region, where they would freeze and grow via riming into graupel/small hail such that they could participate in noninductive charging (NIC; Takahashi 1978; Saunders 1994). Moreover, it was noted that the vertical motion associated with these updrafts was a source for the continual replenishment of supercooled cloud liquid water.

CR96 examined multicell storms across the Colorado Front Range with the polarimetric Colorado State University–University of Chicago–Illinois State Water Survey (CSU-CHILL) radar. Consistent with NIC, the time evolution of graupel echo volume and FR appeared to be well correlated (CR96). Further, CR00 suggested that graupel particles were required for cloud electrification via NIC and that warm rain clouds (thus clouds without ice) were devoid of lightning. CR00 calculated the ice mass within the mixed-phase region and determined that total ice mass correlated well with the FR. Wiens et al. (2005) examined the kinematic and microphysical behavior of a tornadic supercell across the high plains of Kansas and Colorado. Similar to CR96, Wiens et al. (2005) concluded that the graupel echo volume and FR were well correlated and that the updraft volume >10 m s−1 correlated well with the FR. Kuhlman et al. (2006) simulated the same tornadic supercell, and reproduced similar correlations. In addition, Kuhlman et al. (2006) showed that the updraft mass flux correlated well with the FR but noted that the maximum updraft velocity was not well correlated with the FR for this storm. Deierling and Petersen (2008) noted that lightning can be used as an accurate measure of the updraft intensity and that similar to Kuhlman et al. (2006), the maximum updraft velocity did not consistently correlate well with the mean FR. Deierling and Petersen (2008) also examined the relationship between updraft volume and mean FR. Sensitivity studies examining the updraft volume >0, 5, 10, and 20 m s−1 above the −5°C level revealed that updraft volumes >5 and 10 m s−1 correlated best with the mean FR. Finally, Deierling et al. (2008) showed that mean FR was well correlated with the flux product of hydrometeors within DMC.

Because of the above findings and because cloud-resolving models use parameterization schemes for LNOx production with various assumptions, it is necessary to do a detailed synopsis of the coevolution of the kinematics, microphysics, and electrical behavior of ordinary, multicellular DMC storms across northern Alabama. This study will examine one convective complex that occurred on 21 May 2012 during DC3 and provide a detailed study of the evolution of FR, maximum updraft velocity, updraft volumes >3 and 5 m s−1, graupel mass and volume, nonprecipitating ice aloft as well as flash area and flash length scale (LS), and compare how these variables influence each other and change over time. These results, together with results from other convective complexes (not shown), will eventually be used to develop empirical radar observable FR relationships in an attempt to provide robust expressions that could be used to assist in the accurate depiction of lightning in numerical cloud models that do not explicitly resolve cloud electrification, in order to improve LNOx modeling (e.g., Carey et al. 2014a,b).

3. Data and methods

a. DC3 experimental design and Alabama domain

The field phase of the DC3 experiment was conducted from 15 May to 30 June 2012 across three regions. The three study regions were located across 1) northeastern Colorado, 2) central Oklahoma, and the western Texas Panhandle, and 3) northern Alabama and southern Tennessee. One of the main motivations for the selection of these sites was the varying convective morphology and thus lightning behaviors (Barth et al. 2015). All regions were equipped with dual-polarization radars to document relevant microphysical and kinematic attributes of DMC. For electrical observations of DMC, VHF LMAs were available for detailed three-dimensional mapping of total lightning activity. Data from the National Lightning Detection Network1 (NLDN) consisting of LF to VLF sensors that are located across the United States and used to identify regions of primarily CG lightning activity (Cummins and Murphy 2009) were also used. Environmental upper-air observations from mobile ballooning facilities aided in characterizing both the preconvective and near-storm environments, as well as offered short-term forecasting support for the positioning of remote sensing platforms. The upper-air observations were quality controlled by specialists at NCAR using the techniques outlined in Loehrer et al. (1996). Aircraft observations were performed; however, results from these observations will not be discussed in this study.

Subjective cell identification and tracking was done by a human expert; the multicellular nature of DMC events across DC3 Alabama was problematic (i.e., cell “drop out,” merging, splitting) for automated cell-tracking algorithms. Thus, a subjective Lagrangian approach was used to track the DMC through use of an analysis box that is advected in time and space while calculations of various microphysical, kinematic, and lightning quantities were performed.

b. Overview of WSR-88D and ARMOR radars and methodology

Two polarimetric weather radars across northern Alabama were used in this study: 1) the WSR-88D located at Hytop, Alabama (KHTX); and 2) the Advanced Radar for Meteorological and Operational Research (ARMOR) located at the Huntsville International Airport (KHSV). KHTX is an S-band (10.71 cm) polarimetric radar and operates in a simultaneous slant 45° transmit and receive of the horizontal and vertical channels (NWS Radar Operations Center 2013). Its beamwidth is ~1° with a maximum range of ~230 km and operates in continuous 360° surveillance at fixed elevation angles. ARMOR is a C-band (5.5 cm) polarimetric radar and is co-owned by the University of Alabama in Huntsville (UAH) and WHNT-TV in Huntsville, Alabama (Petersen et al. 2005). ARMOR operates in a slant 45° simultaneous transmit and receive of the horizontal and vertical channels, with a beamwidth of ~1°. With a keen interest on examining upper-level outflow (e.g., the anvil), ARMOR sector volume scan strategies consisted of relatively high maximum elevation angles (highest ~26.8°). ARMOR and KHTX are both capable of measuring the reflectivity factor at horizontal polarization (ZH), Doppler velocity (Vr), differential reflectivity (Zdr), the copolar correlation coefficient (ρhv), and differential phase (Φdp). The specific differential phase (Kdp) for ARMOR is computed using a method outlined in Bringi and Chandrasekar (2001). Additional specifications of ARMOR are discussed in Petersen et al. (2005). The relatively close proximity of ARMOR and KHTX (~70 km) presents the opportunity for three-dimensional wind retrievals within the areas denoted in Fig. 1.

Fig. 1.
Fig. 1.

Map of the Alabama domain during the DC3 field project. The solid red circle represents the location of the ARMOR radar at KHSV. The solid blue circle represents the location of KHTX at Hytop, Alabama. The green triangles show the location of NALMA sensor locations. Note that the northern Georgia sites are excluded. The black dashed lines represent the ARMOR–KHTX multi-Doppler regions. The purple × indicates the location of the radiosonde launch that occurred at 2037 UTC. A distance indicator is shown in the top right of the image.

Citation: Monthly Weather Review 143, 7; 10.1175/MWR-D-14-00250.1

The quality control procedures for ARMOR and KHTX are exhaustive in an attempt to ensure the highest data quality possible. As a result of ARMOR’s wavelength (5.5 cm), propagation effects such as attenuation and differential attenuation, due to the presence of liquid precipitation (including large rain drops and/or melting hailstones) must be corrected. To address the issue of large rain drops, the “raw” ARMOR data collected during DC3 is corrected for attenuation and differential attenuation utilizing a method outlined in Bringi et al. (2001), allowing for the correction of Z data at both horizontal and vertical polarizations. It should be noted here that the method outlined in Bringi et al. (2001) may not work well for melting hail. Though small melting hail was common in the storm, manual visual inspection of the data suggests that propagation effects from melting hail did not affect large swaths of the radar echo. Hence, we are confident that the attenuation and differential attenuation in this region was accurately corrected. Similar results using the same correction procedures were obtained for storms in past studies using ARMOR data (e.g., Deierling et al. 2008).

Ground clutter and second-trip echoes near the convective cells of interest were removed using NCAR’s SoloII software (Lee et al. 1994). The corrected data were then used in the NCAR particle identification (PID) package (Vivekanandan et al. 1999). This algorithm utilizes fuzzy logic to determine hydrometeor type. A total of 17 (14 meteorological and 3 biological) categories are available for the C-band version of the NCAR PID (Deierling et al. 2008). For this study, the graupel/small hail category was used nearly exclusively. To account for the shorter wavelengths and the fact that ARMOR operates in a simultaneous transmit and receive of the horizontal and vertical channels, the version of the PID used in this study was particularly modified for ARMOR by changing membership functions, as performed by Deierling et al. (2008). Both ARMOR and KHTX radar data were gridded using the NCAR REORDER (Oye et al. 1995) software package. This study uses a Cressman weighting scheme (Cressman 1959) to grid the radar data, with ARMOR in the center of the grid and the radius of influence and grid spacing at 1 × 1 × 1 km3. In addition, the NCAR PID information was also gridded to Cartesian space with 1-km grid spacing in the horizontal and vertical dimensions using a nearest neighbor weighting scheme and 1-km radii of influence.

For this study, graupel volume and graupel mass were computed. Consideration was only given to regions between the −10° and −40°C layer. This so-called charging region (CR), as termed by Latham et al. (2004), is hypothesized to be the region in which active NIC of graupel primarily occurs (Takahashi 1978; Saunders 1994). The number of grid boxes associated with graupel particles identified by the NCAR PID were summed over the height layer corresponding to the CR and then multiplied by the gridbox volume to attain the graupel echo volume. For grid boxes identified as containing graupel by the NCAR PID in the same height layer, an estimate of graupel mass was obtained from a ZM relationship found in CR00, which is based on the Rayleigh scattering approximation for an assumed exponential ice particle size distribution. CR00 assumed a typical slope parameter of 4 × 106 m−4 when examining tropical convection (Petersen 1997); while the density of graupel was assumed to approach that of solid ice (ρi = 917 kg m−3). The value of effective reflectivity (Ze), where the NCAR PID determines graupel/small hail within the CR, is then used to compute ZH (in units of mm6 m−3); which is multiplied by a factor of 10−18 such that the appropriate units are satisfied (as in CR00) to obtain the ice mass at a given point. The differences in dielectric constants between water and ice as noted by Smith (1984) were also accounted for. It should be noted that Z-based observations for particles that are in the Rayleigh scattering regime, are dependent on the diameter of the particle to the sixth power (D6). This may result in the dominance of the radar sample volume by only a few large particles and would likely lead to the NCAR PID diagnosing a region of graupel, even though the most common hydrometeor may be ice crystals and could, therefore, result in an overestimate of the graupel echo volume and mass. Furthermore, the exact values given by the ZM relationship should not be taken literally, but rather their trends are what are important here.

In addition to the above precipitation ice and graupel calculations, and as per Dye et al. (2007) and Dye and Willett’s (2007) findings of stronger electric fields in anvils once ZH > 10 dBZ, the mass and volume of particles with ZH > 10 dBZ at temperatures colder than −40°C (nonprecipitation ice mass and volume) was calculated and compared to the change in flash area and LSs over time. The nonprecipitation ice mass was calculated using the Heymsfield and Palmer (1986) ZM relationship: M = 0.089 76Z0.529 [with mass in g m−3 (converted into kg) and Z in mm6 m−3] (Deierling et al. 2008). The dielectric factor of 5.28 (Smith 1984) was incorporated to adjust for differences in ice mass between Z and Ze at horizontal polarization (e.g., ZH). The volume was calculated using the same method as above when calculating the graupel volume. The reason for this comparison is that if the nonprecipitation ice mass and volume of ZH > 10 dBZ is indeed charged (even for anvils of multicellular storms), then an increase in this ice mass and volume should be correlated to an increase in flash size and, as per Barthe and Barth (2008), these larger flashes should produce more LNOx as compared to shorter flashes that occur during the earlier phases of the storm.

After both ARMOR and KHTX radar data are gridded to a common Cartesian plane, NCAR’s Custom Editing and Display of Reduced Cartesian Space (CEDRIC) tool was used for the multi-Doppler wind synthesis (Miller and Frederick 1998). A variational integration method of the mass continuity equation was invoked due to the expected minimization of divergence errors at the upper boundary condition when determining vertical motion from estimates of the U and V components of the horizontal wind as well as estimates of particle fall speed (Gao et al. 1999). Finally, in order to perform a multi-Doppler wind synthesis, the data needed to meet the following criteria: 1) less than 3 min offset between ARMOR and KHTX data, 2) the entire precipitation echo (through at least 10–15 dBZ at storm “top”) needed to be sampled, and 3) Doppler velocities were dealiased properly. The updraft volume is then defined as the product of the pixel count and the volume of that grid box of pixels associated with a given vertical motion value (e.g., 3 or 5 m s−1) within the CR. It should be noted that despite the improvement in estimation of the vertical motion as reported by Gao et al. (1999), errors are still present. These integration errors along with coarse, 1-km grid resolution, likely result in an underestimate of vertical motion.

c. Overview of NALMA and NLDN data and methodology

The North Alabama LMA (NALMA) is owned and operated by NASA’s Marshall Space Flight Center (NASA MSFC) in Huntsville, Alabama. It consists of 11 VHF sensors across northern Alabama (Fig. 1) and 2 VHF sensors in north-central Georgia; the center of the domain is located at the National Space Science and Technology Center (34.72°N, 86.64°W) in Huntsville, Alabama (Koshak et al. 2004; Goodman et al. 2005). NALMA sensors have sampling windows of ~80 μs and record the peak pulse in radiation. Koshak et al. (2004) found that typical horizontal and vertical spatial errors of NALMA-detected VHF sources were ~50–500 m within 100 km from the network center. At ranges in excess of 300 km from the center of the network, these errors increase by an order of magnitude. All cases analyzed during DC3 were within 100 km of the center of NALMA. The McCaul VHF source clustering algorithm was used for this study as the algorithm was designed specifically for NALMA (McCaul et al. 2005, 2009) and takes distance from the NALMA center into account when clustering sources. While filtering of noise is inherently done during the clustering process, only flashes with ≥10 sources were used, following other researchers (e.g., Wiens et al. 2005; Murphy 2006; Schultz et al. 2009; Gatlin and Goodman 2010).

The NLDN is a national lightning network across the United States that operates at the LF to VLF range and was used as a detector of CG lightning flashes. Cummins et al. (1998) and Biagi et al. (2007) noted that the median spatial accuracy error associated with return strokes was ~200–500 m while Vaisala estimates that the ability to detect a CG lightning flash across the continental United States and surrounding waters is ~95% (Cummins et al. 2006; Cummins and Murphy 2009). While NLDN does attempt to discriminate between IC and CG flashes, Cummins et al. (1998, 2006) and Wiens et al. (2005) recommend that a +15-kA threshold be applied for positive CG (+CG) flashes. This dataset uses a 15-kA absolute magnitude threshold to determine CG flashes in general, including negative polarity ground flashes (Fleenor et al. 2009). Lightning FR calculations were then performed by summing the total number of lightning flashes and then dividing by the elapsed radar volume time.

The flash LS was calculated as the square root of the horizontal convex hull (or polygon) area surrounding the NALMA VHF sources in the horizontal for each flash (Fig. 2), as explained in Bruning and MacGorman (2013). The flash area is then just the convex hull area surrounding the NALMA VHF sources (Bruning and MacGorman 2013). It should be noted that the flash LS should not be confused with the “micro-scale channel length that accounts for the tortuous propagation of lightning’s many branches” (Bruning and MacGorman 2013), but that the LS makes the assumption that the “macro-scale potential wells discharged by each flash are roughly planar and axisymmetric” (Bruning and MacGorman 2013). A rudimentary charge analysis, as done by Rison et al. (1999), Shao and Krehbiel (1996), Rust et al. (2005), Wiens et al. (2005), Weiss et al. (2008), Lund et al. (2009), and others was performed on the storm. This is possible because negative leader breakdown in the positive charge region occurs much more impulsively and thus results in many more LMA source points as compared to positive leader breakdown in the negative charge region. As such, a region with many LMA source points is termed a positive charge region while a region with fewer LMA source points is termed a negative charge region (Rison et al. 1999; Shao and Krehbiel 1996; Rust et al. 2005; Wiens et al. 2005; Weiss et al. 2008; Lund et al. 2009, among others).

Fig. 2.
Fig. 2.

Images showing how the convex hull area is calculated through the convex hull (polygon) method: (a) an example of an IC flash and (b) an example of a CG flash [following Bruning and MacGorman (2013)]. The colored circles are each NALMA source location constituting a flash, color coded as shown according to time for the first source (purple) and the last source (dark red). The black star indicates the NLDN stroke location(s).

Citation: Monthly Weather Review 143, 7; 10.1175/MWR-D-14-00250.1

4. Storm evolution results

a. Meteorological overview

The 21 May 2012 case day featured two rounds of DMC across DC3 Alabama. Early in the morning, areas of convection developed upstream of the Tennessee Valley across portions of the mid-Mississippi River valley. Figure 3 depicts the 0600 UTC surface analyses that revealed a cold front stretching from portions of Missouri, southward into Texas. The feature of most importance was the prefrontal trough east of the front that stretched from western Ohio through western Kentucky and Tennessee. It is likely that low-level convergence in conjunction with the weak large-scale ascent beneath the entrance region of the trough was sufficient for the generation of nocturnal DMC. This DMC moved toward the DC3 Alabama domain during the early morning time period on 21 May 2012 and produced numerous outflow boundaries downstream that served to initiate new convection. At ~1100 UTC, a small complex moved through the center portion of the DC3 Alabama network, effectively stabilizing the region. This convection gradually weakened as it moved across eastern portions of the DC3 Alabama network.

Fig. 3.
Fig. 3.

The subjective surface analysis (focusing on the Southeast) from the Hydrometeorological Prediction Center (HPC) at 0600 UTC 21 May 2012. Of note is the cold front stretching from portions of Missouri, southward into Texas, and the prefrontal trough east of the front that stretched from western Ohio through western Kentucky and Tennessee.

Citation: Monthly Weather Review 143, 7; 10.1175/MWR-D-14-00250.1

The 2037 UTC UAH mobile raob (Fig. 4) from near Capshaw, Alabama (Fig. 1), revealed that the air mass across the Tennessee Valley had continued to destabilize as surface-based convective available potential energy (SBCAPE) values increased to ~785 J kg−1 and surface-based convective inhibition (SBCIN) values approached −1 J kg−1. Deep-layer wind shear (or deep-layer wind magnitude difference) was weak with values of 2 × 10−4 s−1 (1.2 m s−1). The 2037 UTC raob revealed modest low-level (surface–3 km) lapse rates on the order of 7.5°C km−1. A summary of environmental parameters obtained from the 2037 UTC raob near Capshaw is shown in Table 1. Under the influence of west-to-northwest flow, convection generally traveled to the east and southeast. Low-level convective outflow progressed toward the west and southwest. This resulted in the generation of new convective cells along the western flanks of multicellular complexes. Initially two distinct updrafts developed as seen in two relative maxima in ZH (Figs. 5a,c). Their individual morphology will be examined before both elements merge at ~2015 UTC.

Fig. 4.
Fig. 4.

2037 UTC UAH mobile raob taken from Capshaw, Alabama.

Citation: Monthly Weather Review 143, 7; 10.1175/MWR-D-14-00250.1

Table 1.

Summary of convective parameters for the 21 May 2012 case day from the 2037 UTC UAH mobile raob. The raob was taken near Capshaw, Alabama, ~18 km NNW of KHSV.

Table 1.
Fig. 5.
Fig. 5.

The ZH CAPPI at 5-km altitude and vertical cross sections in the XZ plane (east–west cut) at various distances north of ARMOR showing the northernmost and southernmost cells that eventually merge to form cell B2. (a) CAPPI at 2001 UTC; (b) vertical cross section at 2001 UTC, 76 km north of ARMOR; (c) CAPPI at 2004 UTC; (d) vertical cross section at 2004 UTC, 76 km north of ARMOR; (e) CAPPI at 2015 UTC; and (f) vertical cross section at 2015 UTC, 65 km north of ARMOR. For the CAPPI images in (a),(c),(e) the color-filled contours represent ZH in dBZ from 10 (dark blue) to 60 dBZ (dark red) and the thick black solid lines represent the vertical motion field in intervals of 2, 4, 6, 8, 10, and 12 m s−1. Horizontal wind vectors are plotted as black arrows and the domain of interest is shown in the dashed red box. For the cross-sectional images in (b),(d),(f) corresponding to the region in the dashed red box in (a),(c),(e), the color-filled contours represent ZH in dBZ from 10 (dark blue) to 60 dBZ (dark red), the black solid lines represent the vertical motion field in intervals of 2, 4, 6, 8, 10, and 12 m s−1 and the dark gray dashed lines represent the Zdr values greater than 0 dB with increments of 1 dB. The thin horizontal black short-dashed lines represent the 0°, −10°, and −40°C isotherms as calculated from sounding data. For (a)–(f), the black solid circles represent NALMA flash initiation points without NLDN ground locations (IC flashes), the gray circles with black borders represent NALMA flash initiation points with NLDN ground locations (CG flashes), and the gray triangles with black borders represent the NLDN ground locations of the return strokes of the CG flashes.

Citation: Monthly Weather Review 143, 7; 10.1175/MWR-D-14-00250.1

b. Development of two separate updrafts (1945–2001 UTC)

Convective elements that would eventually become the storm focus in this study (called storm B2) developed ~80 km north of the ARMOR radar at ~1945 UTC. Development of a mature updraft was hindered due to the weak 850–700-hPa lapse rate (Table 1). Initially, echoes were confined to below the −10°C (5.5 km) level with very weak vertical motion. However, during the 2001 and 2004 UTC ARMOR radar volumes, two distinct maxima in the vertical motion field were observed (Figs. 5a,c). A vertical cross section through the northernmost updraft at 2001 UTC (~76 km north of ARMOR) indicated that vertical motion as large as 4 m s−1 extended up to 6 km (Fig. 5b, solid lines). This vertical motion associated with the northern updraft was of a sufficient magnitude to loft hydrometeors upward and into the 0°C (3.5 km) to −10°C (5.5 km) layer. These hydrometeors exhibited Zdr values on the order of 4–5 dB (Fig. 5b, dashed lines) and were, therefore, likely raindrops that may have been supercooled or this area contained significant water fractions. This feature of higher Zdr values above the 0°C level is known as a Zdr column and has been shown to comprise a mixture of supercooled drops, water-coated ice crystals, large raindrops, and spherical ice particles such as graupel (Illingworth et al. 1987; Conway and Zrnić 1993; Bringi et al. 1997). In addition, Goodman et al. (1988), Jameson et al. (1996), Bringi et al. (1997), CR00, Bruning et al. (2007), and MacGorman et al. (2008) have all shown that the existence of a Zdr column could be an important feature in the electrification process.

Moderate-to-high ZH of 50–55 dBZ in the 0° to −10°C layer (Fig. 5b) confirmed the presence of large hydrometeors associated with the northernmost updraft. Slightly lowered values of ρhv (~0.93, not shown) also suggested some mixture of hydrometeors toward the top of the Zdr column, such as freezing of lofted raindrops (e.g., Hubbert et al. 1998; Smith et al. 1999; Kumjian et al. 2012). The combination of environmental and polarimetric radar data indicated a mixture of supercooled liquid drops, small hail, freezing drops, and graupel in this region. Above the −10°C level, a gradient in both ZH and Zdr exists. Given the values of ZH and Zdr decreasing to 35–45 dBZ and 1–2 dB (Fig. 5b), as well as ρhv values remaining at ~0.93, the number of frozen particles was increasing with height above the −10°C level [similar to Jameson et al. (1996)]. Specifically, this gradient in Zdr suggested a transition, via freezing, to slightly more precipitation ice (likely graupel and/or small hail). NCAR PID output indicated the presence of graupel aloft in the CR (not shown).

c. Lightning production in the northernmost and southernmost updrafts (2004–2012 UTC)

During 2004 UTC, the northernmost updraft located ~76 km north of ARMOR produced its first flashes (Figs. 5c,d). Two of the three flash initiation sources were located just below 8.5 km along a region characterized by low ZH (25–35 dBZ) and low Zdr (~1 dB, Fig. 5d). NCAR PID suggested that this region contained a mixture of graupel and ice crystals. The first VHF radiation source associated with the third flash was detected at an altitude of ~5.2 km and was located in an area of higher ZH (50–55 dBZ) and Zdr (1–2 dB) indicative of graupel and/or small hail, which likely resulted from freezing of supercooled drops. The southernmost updraft (not shown) exhibited a slower rate of electrification than the northern updraft. Neither NALMA nor NLDN detected any lightning activity until this storm cell merged with the northernmost cell.

d. Rapid lightning increase post updraft merger (2015–2023 UTC)

Over the course of the 2015 UTC ARMOR volume, the northern and southern updrafts merged and this complex will now be referred to as B2. The result of this merger between the two updrafts appeared to have been constructive, as vertical motion and the total lightning increased. At 2015 UTC (Figs. 5e,f), a rapid increase in the graupel mass and graupel echo volume occurred, which will be discussed in more detail in section 5. During this period of intensification, radar cross sections at 2015 (Figs. 5e,f) and 2020 UTC (Figs. 6a,b) from ARMOR showed a robust updraft with peak velocities >10 m s−1 and estimates of strong upward vertical motion extended well into the −10°C layer. A strong Zdr column, with Zdr values >5 dB through the 0°C level, together with the 50-dBZ ZH contour through the −10°C level, suggested that the updraft during 2015 UTC (Figs. 5e,f) was efficient at lofting large particles upward. While strong upward vertical motions extend deeper into the mixed-phase region, the large Zdr values observed at 2015 UTC above the 0°C level are weaker by 2020 UTC (Figs. 6a,b). In fact, Zdr values rapidly decreased to 1–3 dB above the 0°C level between 2015 and 2020 UTC, while ρhv lowered to 0.87 (not shown), suggesting a mixture of hydrometeor types due to the freezing of large raindrops. This freezing process leads to the formation of hail embryos, which quickly grow into graupel and hail through the riming process, thus creating isotropic particles that bias Zdr to lower values, lead to Mie resonance (lowering ρhv further) and increase ZH > 55 dBZ above the 0°C level. However, the positive Zdr of 1–3 dB and lowered ρhv values above the 0°C level suggest that the freezing process is not complete, and thus there is a mixture of both frozen (graupel and hail) and liquid precipitation in the cloud at this time. While the presence of hail is not necessarily a requirement for charging, its presence is consistent with the stronger updrafts and alone may be enough to infer that an increase in the size of graupel particles has occurred.

Fig. 6.
Fig. 6.

As in Fig. 5, but for (a),(b) 2020; (c),(d) 2023; and (e),(f) 2044 UTC.

Citation: Monthly Weather Review 143, 7; 10.1175/MWR-D-14-00250.1

Peak FRs, as inferred by NALMA, of ~2 flashes min−1 (with a total of 11 flashes) occurred between 2015 (Figs. 5e,f) and 2020 UTC (Figs. 6a,b; solid black circles indicating NALMA flash initiation points; solid gray triangles indicating NLDN points—see figure caption for more details). At 2015 UTC the majority of the NALMA initiation points occurred at temperatures <0°C, while located in regions of gradients of the vertical motion, ZH and Zdr in the mature (eastern) side of the storm (Figs. 5e,f). By 2020 UTC (Figs. 6a,b), all but two NALMA flash initiation sources were located at temperatures <−10°C; again on the mature (eastern) side of the storm in regions of gradients of vertical motion and ZH. The remaining two flashes were located below 3.5 km (warmer than 0°C). Peak upward vertical motion associated with the strongest updrafts was around 13.9 m s−1 for 2020 UTC (see section 5) and located above the −10°C level (Fig. 6b).

At 2023 UTC (Figs. 6c,d), B2 reached its peak in terms of the NALMA FR (~5 flashes min−1; a total of 14 flashes between 2023 and 2026 UTC). The majority of these flashes initiated in the CR, again on the mature side of the storm, while three flashes initiated between 0° and −10°C, indicating that flash initiation becomes more bimodal over time. As expected from NIC, there is an observed peak in both the graupel echo volume and graupel mass (discussed in more detail in section 5) almost coincident with the peak in the FR. Radar cross sections during 2023 UTC (Figs. 6c,d) revealed that the mid- and upper-level updraft continued to strengthen with upward motions of ~10 m s−1 well into the CR. Interestingly, the contours of vertical motion continued to take on a wedge shape and slope from the west to east (Figs. 6b,d). This likely resulted when surface outflow outpaced the main convective line. Before the start of the collapse of B2 (as of 2029 UTC), the stronger updrafts become collocated with low-to-moderate ZH and small, positive Zdr values (Figs. 6c,d). It is possible during this time that ice crystals, aggregates, and perhaps small graupel particles were still being lofted into colder regions, while larger graupel and small hail particles fell out due to gravity, leading to a strengthening of the electric field and in increase in FRs during this time.

e. Decay and dissipation stages of B2 (2026–2057 UTC)

Consistent with Rutunno–Klemp–Weisman (RKW) theory (Weisman and Klemp 1984; Rotunno et al. 1988) weak environmental shear likely resulted in an unbalanced vorticity between the cold pool and the ambient environment. As a result, the outflow boundary surged south and westward ahead of the main line of convection. The inability of new parcels to be lofted up and over the gust front to their level of free convection likely resulted in a lack of robust updrafts and as a result, the weaker updrafts failed to generate, loft, and suspend larger hydrometeors. Subsequently, ZH decreased with time over the depth of the weakening complex. The ARMOR radar cross section for 2044 UTC (Figs. 6e,f) shows that some residual vertical motion within the CR layer may have supported additional charging. Furthermore, Zdr values were <1 dB within this region; while ZH was 30–45 dBZ (Fig. 6f). This would indicate that residual graupel and or ice particles (such as snow aggregates or small hail) may remain, as pristine ice particles would contribute to larger Zdr values. While not as impressive as earlier in B2’s life cycle, sufficient graupel and ice particles in conjunction with vertical motion were enough to produce seven flashes between 2044 and 2048 UTC, all of which initiated at temperatures colder than −10°C.

5. Integrated results

The evolution of total (IC + CG) NALMA and NLDN CG lighting relative to the dual-Doppler-derived kinematic properties in the CR are shown in Fig. 7. The maximum updraft reached 6.3 m s−1 at 2001 UTC in the CR before lightning occurred at 2004 UTC in B2 (Fig. 7). As stated earlier, this surge in the updraft was associated with the lofting of supercooled drops (i.e., formation of a Zdr column), which later froze and likely participated in rapid NIC prior to the onset of lightning. A maximum updraft ~6 m s−1 and the lofting of supercooled raindrops to T < −10°C prior to first lightning in warm-based clouds is consistent with Goodman et al. (1988), Zipser and Lutz (1994), and CR00. The maximum updraft increased rapidly in advance of the FR and correlated well with the FR with r = 0.85 (Fig. 7). The updraft volumes in Fig. 7 are also well correlated to NALMA FRs; updraft volume >3 m s−1 had r = 0.86 and updraft volume >5 m s−1 had r = 0.87. The PID graupel volume (Fig. 8) corresponds well to the peaks and valleys in the FR, even better than the updraft volumes and maximum updraft speeds (Fig. 7), which is consistent with CR96, Wiens et al. (2005), and other studies and has r = 0.90. Likewise, the graupel mass in the CR is also well correlated to the FR (r = 0.88; Fig. 8), similar to CR00 and Deierling et al. (2008). The FR in B2 was dominated by ICs as the CG FR was much lower (Figs. 7 and 8). Furthermore, the NLDN CG FR was poorly correlated to the NALMA FR, maximum updraft, the >3 and >5 m s−1 updraft volumes, the graupel echo volume, and graupel mass in the CR in B2 (Figs. 7 and 8). The lack of correlation between CG and FR, even in an ordinary multicell storm, is important as some LNOx studies have attempted to estimate total lightning from observations of NLDN CG lightning and a regional IC:CG ratio.

Fig. 7.
Fig. 7.

Evolution of the NALMA total (IC + CG; black solid line) and NLDN CG (black dashed line) lightning FRs (min−1) vs maximum updraft velocity (m s−1; dark blue solid line) and updraft volume (m3) for volumes exceeding updraft thresholds of 3 (green solid line) and 5 m s−1 (green dashed line) in the CR for B2 on 21 May 2012.

Citation: Monthly Weather Review 143, 7; 10.1175/MWR-D-14-00250.1

Fig. 8.
Fig. 8.

Evolution of the NALMA total (IC + CG; black solid line) and NLDN CG (black dashed line) lightning FRs (min−1) vs PID graupel echo volume (m3; dark blue solid line) and PID graupel mass (kg; green solid line) in the CR for B2 on 21 May 2012.

Citation: Monthly Weather Review 143, 7; 10.1175/MWR-D-14-00250.1

The flash LS and nonprecipitation ice volume (for ZH > 10 dBZ at T < −40°C) is provided in Fig. 9 and correlation values comparing the nonprecipitation ice mass and volumes to various flash area and LS calculations are shown in Table 2. In general, the flash LS lagged the FR and other measures of convective vigor, such as graupel mass and updraft volume (Figs. 7, 8, and 9); while the nonprecipitation ice volume aloft was well correlated with the trend in the median flash LS (r = 0.82) and the median flash area (r = 0.77). The nonprecipitation ice mass (not shown) had lower correlation values (r = 0.74 for flash LS and r = 0.69 for flash area; Table 2). Specifically, the NALMA flash LS increased rapidly after an increase in maximum updraft velocity, updraft volume, and graupel volume (cf. Figs. 7, 8, and 9), which is evident when one refers to Fig. 10 [showing all the NALMA source points, color coded per flash from first flash (purple) to last flash (dark red) for the radar volume time period—no constraints were made on the LMA source points and thus all NALMA source points are shown]. At 2004 (Figs. 10a,b) and 2015 UTC (Figs. 10c,d) the peak updrafts were 4.9–13.0 m s−1; while the average flash LS was 3.9–5.3 km. At 2055 UTC (Figs. 10e,f), when the peak updrafts have decreased to ≤7 m s−1 and thus the maximum updraft velocity and updraft volume weakened while nonprecipitation ice increased, the average flash LS was 10.9 km, double that of flashes at 2004 and 2015 UTC. Interestingly, as the storm started weakening and the FR decreased, the majority of the NALMA source points were located between the 0° and −40°C levels (Fig. 10f) while the NALMA source points were more spread out, and many were located below the 0°C level during the developing and mature phases of the storm (Figs. 10b,d). These features seem to compare well with the rudimentary charge analysis that was done for the storm of interest (Fig. 11). For the most part, the storm exhibited an ordinary tripole structure, but there were some variations in this structure over time. In the developing phase of the storm, there was a stronger lower positive charge region and a weak upper positive charge region, while during the mature and dissipating phases of the storm, the lower positive charge region weakened (and for some period almost seemed nonexistent) while the upper positive charge region strengthened as the flash sizes increased. In related research, Carey et al. (2014b) used the NASA Lightning Nitrogen Oxides Model (LNOM) to calculate lightning segment altitude distributions and LNOx production profiles for storm B2. Details on the LNOM LNOx production parameterization can be found in Koshak et al. (2014). A comparison between our results and Carey et al.’s (2014b) LNOM LNOx calculations showed that when the lower positive charge region was strong during the developing and mature phases, the updraft velocity and graupel volume and mass was high (Figs. 7 and 8, respectively), the FR was high (Figs. 7, 8, and 9), the nonprecipitation ice volume was smaller (Fig. 9), the flash sizes were smaller (Figs. 9 and 10a–d), the NALMA sources were not limited to regions colder than 0°C (Figs. 10b,d), and the majority of the LNOx produced below 6-km altitude (warmer than −10°C) was due to CG flashes (Carey et al. 2014b). Once the lower positive charge region weakened and the upper positive charge region strengthened (Fig. 11), the updraft velocity (Fig. 7) and graupel volume and mass (Fig. 8) decreased, the FR decreased (Figs. 7, 8, and 9), the NALMA sources were mostly limited to region colder than 0°C (Fig. 10f), the nonprecipitation ice volume was larger (Fig. 9), the flash sizes increased (Figs. 9 and 10e,f), and there was very little LNOx produced below 6-km altitude (warmer than −10°C) for both IC and CG flashes, thus the majority of LNOx was produced at temperatures colder than −10°C (Carey et al. 2014b). Therefore, these larger flashes could have a considerable impact on the total amount of LNOx produced as found by Barthe and Barth (2008) and Carey et al. (2014b). The FR and flash size anticorrelation also compare well to the findings of Bruning and MacGorman (2013) and as shown by the increase in nonprecipitation ice volume (Fig. 9) and charge analysis (Fig. 11), the upper positive charge strengthens as the anvil develops and grows, leading to the NALMA flash sources propagating into the anvil region, which may relate to Dye et al. (2007) and Dye and Willett’s (2007) findings that anvil regions consisting of ZH > 10 dBZ have electric fields >10 kV m−1.

Fig. 9.
Fig. 9.

Evolution of the NALMA (black solid line) and NLDN (black dashed line) FRs (min−1) compared to the flash LS (km; dark blue box) and the nonprecipitation (anvil) ice volume (m3; green solid line). The nonprecipitation ice volume is calculated for T < −40°C and ZH > 10 dBZ. The dark blue box represents the interquartile range (IQR) (25%–75%) of the flash extent while the horizontal line that divides the box into sections is the median value. The upper whisker is the (75th percentile + 1.5×IQR), or upper inner fence, and the lower whisker is the (25th percentile − 1.5×IQR), or lower inner fence (Wilks 2006). Any values larger than the upper whisker and smaller than the lower whisker are seen as outliers and are shown as circles. When there is only 1 flash in the time frame, it is represented by a black horizontal line.

Citation: Monthly Weather Review 143, 7; 10.1175/MWR-D-14-00250.1

Table 2.

Correlation coefficient values when comparing various flash size distribution parameters to the nonprecipitation ice volume (m3) and nonprecipitation ice mass (kg).

Table 2.
Fig. 10.
Fig. 10.

The ZH CAPPI at 8-km altitude and vertical cross sections in the YZ plane (north–south cut) at various distances east of ARMOR. (a) CAPPI at 2004 UTC; (b) vertical cross section at 2004 UTC, 2 km east of ARMOR; (c) CAPPI at 2015 UTC; (d) vertical cross section at 2015 UTC, 7 km east of ARMOR; (e) CAPPI at 2055 UTC; and (f) vertical cross section at 2055 UTC, 22 km east of ARMOR. In (a),(c),(e) the dashed red box indicates the domain of the storm. For all panels (a)–(f) ZH is in gray levels [from 10 (black) to 60 dBZ (white)]; while the NALMA sources for all flashes are shown as a color-filled circle according to first flash (purple) and last flash (red) during the radar time interval. The NLDN location points are displayed as filled triangles in (b),(d),(f) and these NLDN points have the same color as the flash source points. For instance, if the NALMA flash source points are seen in blue, then the corresponding NLDN ground location of that flash is also in blue.

Citation: Monthly Weather Review 143, 7; 10.1175/MWR-D-14-00250.1

Fig. 11.
Fig. 11.

Rudimentary charge analysis of NALMA sources for storm B2 from 2004:00 to 2054:00 UTC. NALMA sources are color coded by inferred charge locations with red indicating regions of positive charge and blue indicating regions of negative charge. The green source points were undetermined relative to positive or negative charge. The two-dimensional spatial projections are as follows: (a) altitude vs time projection, (b) altitude vs x projection, (c) x vs y projection, (d) source altitude histogram count, and (e) y vs altitude projection.

Citation: Monthly Weather Review 143, 7; 10.1175/MWR-D-14-00250.1

Another interesting feature is that NALMA source points in the anvil only appear in regions with precipitating echo underneath (Fig. 10f). For instance, there are very few NALMA source points in the anvil region between 80 and 100 km north of the radar (Figs. 10e,f) while below this region, there is very little precipitation (seen by a lack of ZH contours in the same region; Fig. 10f). It is possible that the ice crystals in the anvil between 80 and 100 km north of ARMOR are not sufficiently charged because this region is too far away from the convective core, leading to charge dissipation due to turbulence (Bruning et al. 2010) as the ice crystals are transported via divergence into the anvil region. Hence, after flash initiation in the convective region, there is likely not enough charge for the propagation of the flash leader into this region of the anvil and, therefore, no NALMA VHF sources were recorded in the anvil between 80 and 100 km north of ARMOR (Fig. 11).

Finally, because of the flash LS and the spread in these LSs increasing after an increase in the maximum updraft velocity, updraft volumes, and graupel volume and as the FRs decrease, it was decided to investigate the changes in the total flash LS and area. These are all the flash LSs and areas for a given radar volume period, added together (Fig. 12). Both the total flash LS (km; blue) and the total flash area (km2; green) are compared to the FRs (min−1; black). Interestingly, even though the total flash area and LS are highly correlated to each other (r = 0.93) the total flash LS has a higher correlation to the FR (r = 0.95) than the total flash area (r = 0.85) does. Similarly, Mansell and Ziegler (2013) noted that simulated total charge separation had better agreement with total channel length than with total flashes. The small decrease in r for the total flash area and LS comparison to FR is because once one takes the square root of the area (to get the LS), the results are more smooth, especially if there were only a few large flashes compared to a majority of smaller flashes. Ultimately, for this case, both total flash LS and total flash area are well correlated to the FR, yet these high correlations may not always be true, especially in high FR storms. It should be noted, however, that these correlations may simply be an artifact of the way the correlations are computed and that more flashes may lead to higher total flash LS and areas simply by virtue of numbers involved, and therefore further testing on various storm types are needed.

Fig. 12.
Fig. 12.

Comparison of the NALMA (black solid line) FR (min−1) over time to the total flash LS (km; blue line) and the total flash area (km2; green line) for B2. The total flash LS (area) is calculated by summing all the LSs (areas) per radar volume.

Citation: Monthly Weather Review 143, 7; 10.1175/MWR-D-14-00250.1

These comparisons are important because of the methods used in calculating LNOx: if total flash area and LS are always highly correlated to the FRs (different storms and various regions), then the current LNOx methodologies, primarily based on FRs, would likely suffice. If, however, there are variations in these comparisons, then the flash parameterization schemes used in the calculation of LNOx might have to be adjusted to include flash LS or area. More research is required to better understand the relationship between flash LS, flash areas, and the kinematic and microphysical properties of convection. If progress is made on this subject, it is likely to lead to improvements in LNOx parameterizations in models, which currently assume that every flash is the same (i.e., they do not consider flash LS or area).

6. Summary and conclusions

An overview of the environmental controls on lightning properties of a multicellular storm that occurred on 21 May 2012 in northern Alabama during the DC3 campaign has been presented. The primary goal of this DC3 study was to utilize multi-Doppler and dual-polarization radar-inferred observations and analysis products to investigate the microphysical and kinematic control of the FR, flash LS, and flash area. It was demonstrated that radar kinematic (maximum updraft, convective updraft volumes) and microphysical (graupel echo volume, graupel mass) observables in the CR are correlated to the NALMA-observed FR, which is consistent with other previously published work (CR96; CR00; Lang and Rutledge 2002; Wiens et al. 2005; Kuhlman et al. 2006; Deierling and Petersen 2008; Deierling et al. 2008, and many more).

Observations from Doppler polarimetric radars revealed that the maximum updraft velocity, updraft volumes >3 and 5 m s−1, graupel echo volume, and graupel mass correlated well with the FR. Also, flash LS and flash area were shown to be generally anticorrelated to FR as hypothesized and observed by Bruning and MacGorman (2013) in supercell storms. More specifically, smaller flashes occurred during higher FRs and were associated with a stronger lower positive charge region caused by larger graupel volumes, stronger updraft volumes, stronger maximum updraft velocities, a smaller nonprecipitation ice volume, and an LNOx profile that was dominated by CG flashes below 6-km altitude (warmer than −10°C) as shown by Carey et al. (2014b). On the other hand, larger flashes occurred during lower FRs and were associated with a weakened lower positive charge region in combination with a stronger upper positive charge region, weaker updraft velocities, a smaller graupel volume and mass, and an increase in nonprecipitation ice volume, while the majority of LNOx production occurred at temperatures colder than −10°C (Carey et al. 2014b). For this case, the median flash LS and flash area were well correlated with the nonprecipitation ice volume aloft (r = 0.82 and 0.77, respectively) and thus could be one of the possible parameters controlling flash size and length scale. Finally, the larger flashes that occur during the dissipating phase of a storm could dominate the total amount of LNOx produced during a thunderstorm, as quantified by Barthe and Barth (2008) and hypothesized by Wang et al. (1998). In addition, the total flash LS and flash area were both highly correlated to the FR. Future work will explore the above relationships in a variety of storm types. These results relating microphysical and kinematic control of the FR and flash LS, as well as flash area, are likely to lead to improved LNOx parameterizations in models.

Acknowledgments

We wish to recognize funding from the National Science Foundation’s Physical and Dynamical Meteorology (NDF PDM) Program (AGS-1063573), which has supported the DC3 field experiment and associated research. We also wish to thank the many people who made the collection of DC3 observations possible. Finally, we wish to thank two anonymous reviewers as well as Dr. Ted Mansell for comments that have substantially improved the quality of this research paper.

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