The Structure, Evolution, and Dynamics of a Nocturnal Convective System Simulated Using the WRF-ARW Model

AUGUST 2017 B L A K E E T A L . 3179

BENJAMIN T. BLAKE,DAVID B. PARSONS,KEVIN R. HAGHI, AND STEPHEN G. CASTLEBERRY University of Oklahoma, Norman, Oklahoma

(Manuscript received 20 September 2016, in ﬁnal form 24 April 2017)

ABSTRACT

Previous studies have documented a nocturnal maximum in thunderstorm frequency during the summer across the central United States. Forecast skill for these systems remains relatively low and the explanation for this nocturnal maximum is still an area of active debate. This study utilized the WRF-ARW Model to simulate a nocturnal mesoscale convective system that occurred over the southern Great Plains on 3–4 June 2013. A low-level jet transported a narrow corridor of air above the nocturnal boundary layer with convective instability that exceeded what was observed in the daytime boundary layer. The storm was elevated and associated with bores that assisted in the maintenance of the system. Three-dimensional variations in the system’s structure were found along the cold pool, which were examined using convective system dynamics and wave theory. Shallow lifting occurred on the southern ﬂank of the storm. Conversely, the southeastern ﬂank had deep lifting, with favorable integrated vertical shear over the layer of maximum CAPE. The bore assisted in transporting high-CAPE air toward its LFC, and the additional lifting by the density current allowed for deep convection to occur. The bore was not coupled to the convective system and it slowly pulled away, while the convection remained in phase with the density current. These results provide a possible explanation for how convection is maintained at night in the presence of a low-level jet and a stable boundary layer, and emphasize the importance of the three-dimensionality of these systems.

1. Introduction et al. 2003; Clark et al. 2007), significant challenges still exist even with convection-permitting ensembles (Clark This work is motivated by the need to improve fore- et al. 2009). Our study focuses on a high-resolution, casts of summertime, nocturnal convective systems that convection-allowing simulation of a nocturnal meso- occur over the Great Plains of North America. The scale convective system (MCS) over the southern problem is challenging, as studies (e.g., Heideman and Great Plains. Fritsch 1988; Uccelini et al. 1999; Weckwerth et al. 2004) Fritsch et al. (1986) asserted that MCSs account for have argued that the accuracy of quantitative pre- 30%–70% of the warm season precipitation over the re- cipitation forecasts is consistently lower during the gion between the Rocky Mountains and the Mississippi warm season as a result of the increased role of deep River, with larger contributions during the months of convection. Researchers have demonstrated that there June–August. The rainfall that is generated from these is a pronounced nocturnal maximum in rainfall across MCSs is essential to mitigating drought and enhancing the central United States (e.g., Kincer 1916; Wallace midsummer crop growth, which increases overall agricul- 1975) and that nocturnal convection has been difficult to tural production. However, MCSs also herald the potential represent in numerical weather prediction (NWP) for severe weather including hail, damaging winds, torna- modeling systems, particularly those with parameterized does, and flash floods (Maddox 1980; Jirak et al. 2003). convection (e.g., Surcel et al. 2010). Climate models The nocturnal maximum over the Great Plains can be have also had difficulty in accurately representing these attributed in part to convective systems that originate nocturnal convective systems (Bukovsky and Karoly over the higher plains and the mountains to the west 2009; Pritchard et al. 2011). Although convection- (Cotton et al. 1983; Wetzel et al. 1983), which then be- allowing models have demonstrated some skill (Davis come associated with an eastward-moving envelope of deep convection (e.g., Carbone et al. 2002; Ahijevych Corresponding author: Benjamin Blake, [email protected] et al. 2004; Parker and Ahijevych 2007; Carbone and

DOI: 10.1175/MWR-D-16-0360.1 Ó 2017 American Meteorological Society. For information regarding reuse of this content and general copyright information, consult the AMS Copyright Policy (www.ametsoc.org/PUBSReuseLicenses). Unauthenticated | Downloaded 10/01/21 05:50 PM UTC 3180 MONTHLY WEATHER REVIEW VOLUME 145

Tuttle 2008; Keenan and Carbone 2008). The nocturnal 2006; Parker 2008; Peters and Schumacher 2015a,b, environment over the Great Plains associated with these 2016). However, another possibility is that weak and storms is often characterized by the presence of a stable transient cold pools can generate gravity waves in the boundary layer (SBL) and a low-level jet (LLJ) (e.g., stable air that cause persistent lifting of conditionally Blackadar 1957; Holton 1967; Shapiro et al. 2016). The unstable air (Crook and Moncrieff 1988; Parker 2008; LLJ is often associated with an elevated maximum in Schumacher 2009; Marsham et al. 2010). A bore is a equivalent potential temperature ue and the advection of gravity wave response that can be generated as a warm, moist air into the region (Maddox 1983; Wetzel density current intrudes a low-level stable layer (e.g., Crook et al. 1983; Fernando and Weil 2010). In addition, the 1988; Rottman and Simpson 1989). Surface observations LLJ is frequently present during nocturnal convective of density currents and bores both include hydrostatic events (Bonner 1968; Maddox 1983) and is associated pressure jumps during the passage of their frontal with heavy rainfall (Arritt et al. 1997; Tuttle and boundaries; however, the surface temperature rises or Davis 2006). does not change as a bore passes by, and the pressure The combination of an SBL and the most unstable air change is permanent (Koch et al. 1991). Recent obser- aloft in association with the LLJ results in a greater vational studies (Weckwerth et al. 2004; Knupp 2006; occurrence of elevated convection at night, which Wilson and Roberts 2006; Koch et al. 2008a,b; presents an additional challenge for accurately predict- Tanamachi et al. 2008; Martin and Johnson 2008; ing nocturnal storms (Wilson and Roberts 2006). Here, Hartung et al. 2010; Marsham et al. 2011) have revealed we define convection to be elevated if it occurs in an that atmospheric bores are generated by outflows from environment with the greatest instability located above nocturnal convection and can, in turn, trigger strong Earth’s surface (e.g., Corfidi et al. 2008). A few previous vertical motions, resulting in intense upward displace- authors have used different definitions for elevated ments of air parcels (;0.5–1.5 km) in the lowest ;3km convection, such as Parker (2008), whose definition was and a sustained elevation of the wave duct. Whether the that a convective system should ingest no parcels from mechanism is a bore or a gravity wave, a wave duct can the surface layer. Elevated storms, often taking the form trap the waves by preventing the vertical propagation of large, nocturnal MCSs, can be initiated from air lo- of significant wave energy out of the ducting layer cated above the boundary layer or they can originate (Lindzen and Tung 1976). from surface-based systems that become elevated over Recent highly idealized modeling studies have dem- time (Corfidi et al. 2008; Parker 2008). onstrated the potential role of bores in the evolution of Several previous observational studies have de- squall lines in the presence of an SBL (Parker 2008; scribed a preferential location for the formation of French and Parker 2010). Parker (2008) found that as nocturnal MCSs. Maddox et al. (1979) observed a dis- the stability increases and the convection becomes ele- tinct nocturnal preference of storm initiation on the cool vated, the mechanism responsible for lifting the parcels side of a quasi-stationary surface front that was located flowing into the system evolves from a surface cold pool downstream from the axis of maximum low-level flow. to a feature Parker characterized as a bore with some Colman (1990) developed a climatology of elevated gravity wave characteristics. Eventually, the interaction thunderstorms across the United States and found the between the bore lifting drives the squall line, which typical elevated thunderstorm to occur north of a sur- then propagates at the speed of the bore (French and face warm front in a region with northeasterly surface Parker 2010). In a related mechanism for the mainte- winds. Trier and Parsons (1993) also found the favored nance of convection, Fovell et al. (2006) proposed that position for the development of large, organized, noc- lifting from high- and low-frequency gravity waves turnal convective events over the central United States propagating ahead of the storm promotes new cell de- to be north of a quasi-stationary frontal boundary due to velopment by prepping the advance environment. the LLJ overrunning the front. However, nocturnal The idealized two-dimensional framework for the MCSs can occur well to the south of a frontal zone (e.g., maintenance of nocturnal convection outlined by Koch et al. 2008a). French and Parker (2010) does not address the three- In contrast to theory that guides the upscale growth dimensional variations of these nocturnal storms. The and structure of surface-based convection (e.g., Rotunno goal of our investigation is to advance our knowledge of et al. 1988), a comprehensive theoretical framework of the dynamics, structure, and evolution of nocturnal nocturnal convection is lacking (Trier et al. 2006). convective systems through examination of a simulated Some recent studies have demonstrated that many convective system from the perspectives of theory as nocturnal MCSs produce surface cold pools and re- applied to atmospheric bores (e.g., Scorer 1949; Rottman spond to them as daytime MCSs would (Trier et al. and Simpson 1989; Baines 1995) and frameworks often

Unauthenticated | Downloaded 10/01/21 05:50 PM UTC AUGUST 2017 B L A K E E T A L . 3181 utilized to explain the dynamics of convective systems (e.g., Rotunno et al. 1988; Bryan and Rotunno 2014; Peters and Schumacher 2016). Our approach allows for a more in- depth investigation of convective dynamics and the linkages to wave theory than the previously described idealized studies. Specifically, we investigated a nocturnal, elevated MCS event that occurred over the southern Great Plains during the night of 3–4 June 2013. The MCS initiated in the afternoon as surface- based storms associated with a well-defined surface cold pool. As the boundary layer stabilized at night and the LLJ intensified, the system transitioned from being surface based to elevated and it produced waves, including bores. The MCS was simulated with the Weather Research and Forecasting (WRF) Model, allowing us to test the ability of a numerical model to recreate this observed transition and associated wave structure.

The paper is organized as follows. A description of the FIG. 1. Plots of composite reflectivity (dBZ) from NCAR’s chosen case is given in section 2. Section 3 outlines the Mesoscale and Microscale Meteorology Laboratory (MMM) Im- model configuration used in this study. Section 4 pro- age Archive valid at (a) 1700, (b) 2000, (c) 2330, (d) 0130, (e) 0400, vides an overview of the simulation. Convective system and (f) 0700 LST. The color scale for the radar reflectivity is shown on the rhs. State labels for OK, KS, CO, and TX are displayed in dynamics and wave theory are applied in section 5 to (a). In (b), the red triangle marks the location of the Buffalo explain the results presented in section 4, followed by mesonet site and the green triangle marks the location of the conclusions in section 6. Chickasha mesonet site. In (c), the white arrow points to the bore from the earlier convection, which is located just behind it. The light blue arrow points to the convection that evolves into the 2. Description of the observed case nocturnal, elevated MCS seen in (d)–(f). A detailed synoptic analysis of this 3–4 June 2013 event can be found in Blake (2015). The storm was associated with a short-wave trough that moved through allowed us to document the evolution of surface quan- northern Oklahoma with a low pressure center over tities such as temperature, pressure, rainfall, wind speed, southeastern Colorado and western Kansas. A quasi- and wind direction and to determine whether the MCS stationary, warm frontal boundary was observed across was associated with a density current and/or a bore. central Kansas with a well-defined LLJ developing over Data from the Buffalo, Oklahoma, mesonet site parts of Oklahoma and Texas. The nocturnal MCS oc- (Fig. 2a) show that a well-defined surface cold pool was curred well to the south of this frontal boundary, which present early in the system’s life cycle as convection allows for insight into the mechanisms responsible for strengthened, moved to the east-southeast, and then the initiation and maintenance of nocturnal convection turned more toward the south-southeast with time apart from frontal ascent. (Figs. 1b,c). The convection in this study lasted for approximately Within the NOAA Rapid Refresh (RAP) model, a 22 h, from the early afternoon of 3 June until the late strong gradient in convective available potential energy morning of 4 June. The system became organized (CAPE) was evident in the southerly flow at 1 km AGL, (Fig. 1a) with storms initiating in southwestern Kansas the layer with the maximum CAPE (Fig. 3). The region and the Oklahoma Panhandle around 1700 local stan- of maximum CAPE at 1 km AGL also corresponded to dard time (LST)1 on 3 June near the intersection of the reduced convective inhibition (CIN) (Fig. 3). Between dryline and an outflow boundary from earlier convec- 2300 and 0100 LST, the initial system dissipated as it tion in northwestern Kansas. Data from the Oklahoma moved out of this zone of high CAPE into less favorable Mesonet (Brock et al. 1995; McPherson et al. 2007) conditions in central Oklahoma. Data from the Chickasha, Oklahoma, mesonet site, located in the central part of the state (Fig. 2b), show that from ;0000 1 This study is concerned with the diurnal cycle; hence, we used to 0030 LST a pressure jump was accompanied by a LST, where LST 5 UTC 2 6h. minimal change in surface temperature, indicating the

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FIG. 2. Time series data of surface temperature (8C; red curve), surface pressure (mb; dark yellow curve), ac- 2 cumulated rainfall (mm; shaded green), wind speed (m s 1; cyan curve), and wind direction (8; magenta curve) for (a) the Buffalo mesonet site from 1700 to 0400 LST and (b) the Chickasha mesonet site from 2300 to 0500 LST. The light blue arrows in (a) and (b) mark the passage of a cold pool and the orange arrows in (b) mark the passage of a bore.

2 feature that moved through was likely a bore. The sur- moistening of 6–7 g kg 1 of water vapor at heights be- face observations of the bore correspond to the radar tween 500 m and 2 km AGL associated with the first two ‘‘fine line’’ depicted in Fig. 1c, which passed over wave fronts of the bore (Fig. 4b). The joint presence of Chickasha shortly after 0000 LST. Following the passage moistening and cooling suggests lifting in a stable envi- of the bore, the temperature decreased between 0030 ronment where moisture decreases with height. Vertical and 0130 LST, consistent with the passage of the weak displacements were also calculated using the observed convective cold pool (Fig. 2b). There was no rainfall temperature changes and assuming that any lifted par- recorded at the Chickasha mesonet site because the cels would follow the dry-adiabatic lapse rate (Fig. 4c). earlier convection had dissipated. The maximum vertical displacements with the bore ap- As the initial convective system died out, the bore proached 900 m between 1 and 3 km with a net upward continued to propagate toward the southeast. Further displacement of ;400 m with the passage of the bore. exploration of this bore was undertaken through the Figure 4d depicts the evolution of the potential tem- utilization of retrieved temperature and humidity pro- perature field with height and shows how the colder air files obtained by the MP-3000A microwave radiometer in the SBL was indeed displaced upward by the bore. located on the roof of the National Weather Center in The magnitude of the lifting was consistent with the Norman, Oklahoma. The general lack of cloud cover for previously mentioned studies, although the lifting asso- this event increased the reliability of the microwave ciated with the bore extended well beyond the SBL into retrievals (Castleberry 2014). The retrieved profiles the lower troposphere. Such deep lifting is likely to bring show pronounced cooling through a deep layer extend- parcels closer to their level of free convection (LFC), ing from 1 to 4 km above ground level (AGL) (Fig. 4a). which was ;4 km AGL over this region for parcels lifted 2 The water vapor mixing ratio (g kg 1) field reveals from 1 km AGL, and reduce any CIN present in the

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precipitation behind the leading edge between 0400 and 0600 LST (Figs. 1e,f). The MCS gradually dissipated between 0800 and 1000 LST. Unfortunately, the close proximity of this second bore to the MCS and its precipitation meant the event was associated with signiﬁ- cant water vapor and high liquid water path values, preventing an accurate determination of the proﬁles of temperature and humidity from the microwave radiometer (Castleberry 2014).

3. Design of the simulation The WRF is an NWP model that was developed as a collaborative effort among several institutions (e.g., National Center for Atmospheric Research, National Oceanic and Atmospheric Administration, Naval Re- search Laboratory, Federal Aviation Administration) and various university scientists (Skamarock et al. 2008). Our investigation used version 3.6.1 of the WRF with the advanced research core (WRF-ARW). The model employs terrain-following hydrostatic pressure co-

21 ordinates on an Arakawa C grid. Our simulations were FIG. 3. Horizontal cross section of CAPE (J kg ) at 1 km AGL valid at 0130 LST. The black contours, along with the relative conducted for 22 h, beginning at 1200 LST 3 June and 2 minima and maxima, represent CIN (J kg 1) at 1 km AGL. This ending at 1000 LST 4 June, spanning the three domains plot is a 30-min forecast from the 0100 LST RAP model analysis shown in Fig. 5. Three domains were used in order to data. The state boundaries are indicated in the cross section with resolve convective processes for the feature of interest the latitude and longitude shown along the axes. Wind vectors are across the inner domain while maintaining computa- plotted in the background. tional efficiency. We utilized a two-way nest, in which the outer domain provided boundary values for the in- environment. This lifting and its relationship to CAPE ner grid, and the inner grid could then influence the and CIN will be explored later within the context of our coarse outer domain. The outer domain contained 121 3 simulations. 121 horizontal grid points with 9-km grid spacing, the During the dissipation of the initial system, new con- middle domain contained 301 3 301 grid points with vection started to develop behind the leading system 3-km grid spacing, and the inner domain contained around 2200–2300 LST (Fig. 1c), a phenomenon known 499 3 529 grid points with 1-km grid spacing. The as rearward off-boundary development (Peters and location of the 1-km domain was chosen to capture as Schumacher 2014). This new convection subsequently much of the MCS evolution as possible given computing developed into a nocturnal, elevated MCS, which will be restraints. The 3-km domain was also crucial, as many of the primary focus of our study. Between 0100 and the wave features produced by the convection propa- 0200 LST the system organized itself into two convective gated out of the 1-km domain by 0800 LST. Rayleigh bands (Fig. 1d). The strongest storms occurred when gravity wave damping was utilized aloft, beginning these two convective bands merged between 0200 and around 11 km AGL. We used 100 vertical levels for all 0400 LST. In this region, a second bore formed along the three domains. The vertical grid spacing ranged from leading edge of the system. The data from the Chickasha ;65 m near the surface to ;200 m beyond 1.5 km AGL mesonet site indicate that sharp increases in surface with the first model level at ;32 m AGL. pressure were accompanied by changes in wind di- The lateral boundary conditions for the simulation were rection, an increase in wind speed, and an increase in updated every hour utilizing RAP output obtained from surface temperature immediately followed by a sharp the National Operational Model Archive and Distribution decrease and the onset of rainfall (Fig. 2b). Based on the System (NOMADS) server. The RAP data did not contain expected structure of these systems, the surface data soil information. Thus, we employed the Noah land surface again most likely depict a bore propagating just ahead of model (Ek et al. 2003), taken from the North American the density current. The radar reflectivity illustrates the Land Data Assimilation System, for these simulations. The development of a broad region of trailing stratiform physical parameterizations were chosen based on the

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FIG. 4. Time–height cross sections of (a) retrieved perturbation temperature (K), 2 (b) perturbation water vapor mixing ratio (g kg 1), (c) vertical displacements (km), and (d) total potential temperature (K) derived from the MP-3000A microwave radiometer retrievals [(a)–(c) taken from Castleberry (2014)]. The small symbol at ;0030 LST indicates a mea- surement and subsequent retrieval that may be impacted by an optically thick cloud with a high 2 liquid water path ($500gm 2). Vertical displacements in (c) were estimated using the perturbation temperatures and the dry adiabatic lapse rate. The displacements are calculated relative to the temperature at the initial times [the ﬁrst ;10 min of T(z) data] in this cross section.

results of various sensitivity analyses and are outlined in 4. Overview of the simulation Table 1. The Mellor–Yamada–Nakanishi–Niino (MYNN) model was chosen for the PBL scheme because the The WRF simulation generally performed well when Mellor–Yamada–Janjic´ (MYJ) version did not accurately compared to observations, as can be seen from com- recreate the storm and Yonsei University (YSU) approach parisons of Figs. 1 and 6. In the observations and in the produced a warm surface temperature bias. The reader is simulation, small, unorganized cells form around 1330 LST referred to Blake (2015) for further results of the 3 June with the system organizing into an MCS sensitivity tests. around 1700 LST (Fig. 6a). The observations and the

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TABLE 1. The parameterization schemes utilized by the WRF-ARW.

Atmospheric Parameterization process scheme Reference and notes Longwave RRTM Mlawer et al. (1997) radiation Shortwave New Goddard Chou and Suarez (1999) radiation Cloud Morrison Morrison et al. (2009); microphysics double-moment scheme Land surface Noah Ek et al. (2003) Convection BMJ Janjic´ (1994); 9-km domain only PBL MYNN level 2.5 Nakanishi and Niino (2004) Surface layer MYNN Nakanishi and Niino (2004)

2 2 initially at 1200 J kg 1 with a CIN of 2175 J kg 1 (Fig. 8). As the night progressed, the surface flow sta- 2 bilizes and CAPE decreases to 400 J kg 1, while the 21 FIG. 5. Domain configuration used by the WRF Preprocessing magnitude of CIN increases to 2500 J kg until 2100 LST, System (WPS) for the model runs. when the surface inflow lacks significant CAPE. The sta- bilization of the surface conditions is expected from nocturnal radiational cooling and the formation of model have the convection moving to the east and an SBL. southeast with the initial line dying out between 2300 The conditions aloft at 1 km AGL are quite different and 0100 LST, and a second line of storms forming as the (Fig. 8). The inflow region at this height is located nearly first line is diminishing, with the nocturnal MCS dissi- due south of the MCS, with CAPE and CIN initially pating between 0900 and 1000 LST. One subtle differ- quite close to the surface values at 1700 LST, suggesting ence in storm structure occurs near 0400 LST (Figs. 1e that the 1-km values were within the convective and 6e) when the model produces a relatively broad area boundary layer. Subsequently, CAPE decreases and the 2 of convection with reflectivities greater than 50 dBZ.In magnitude of CIN only increases to 2200 J kg 1 so that the observations, the region with reflectivities of 50 dBZ by 2130 LST, there is no longer any significant CAPE or greater was confined to a narrower west-southwest– within the 1-km inflow. This pronounced decrease in the east-northeast-oriented band at this time. This difference CAPE precedes the dissipation of the first system. After is a known problem with microphysical parameteriza- convective cells develop into the second MCS, CAPE 2 tions (Morrison and Milbrandt 2015). increases to over 2000 J kg 1 and the magnitude of CIN 2 As expected, given the similarities discussed above, decreases to 280 J kg 1 by 0230 LST. For a given the model and the observations both depict similar re- amount of lifting, it would be easier to initiate and sus- gional environments, including a south-to-north corri- tain deep convection within these nocturnal conditions dor of high-CAPE values located over western than those observed in the PBL during the day. By 0630 LST Oklahoma and the Texas Panhandle that is in place from there is no longer any inflow with significant CAPE the start of the simulation until the passage of the noc- or CIN into the MCS at 1 km. turnal MCS (Figs. 3 and 7b). Since the observations The high values of CAPE at 1 km are primarily due to suggest a transition from convection with surface-based large values of the water vapor mixing ratio within the convergence associated with a cold pool to lifting aloft strong wind speeds associated with the LLJ (Fig. 9). The partly associated with a bore, we investigated the evo- time of peak wind speeds, ;0230 LST, corresponds to lution of the CAPE and CIN of the inflow air both at the the time at which CAPE in the 1-km inflow is the surface and at 1 km AGL. We utilized the 3-km domain greatest. As discussed earlier, the LLJ is known to and focused on inflow conditions ahead of the system provide an elevated source of moist, unstable air. Thus, (i.e., Fig. 7) over a time span from 1700 LST 3 June to advection associated with the nocturnal LLJ brought 0630 LST 4 June. The various storm inflow regions were moist, unstable air into the storm, which played a crucial determined subjectively. The inflow region for the sur- role in creating a high-CAPE inflow and likely impacted face flow is located southeast of the MCS with CAPE the longevity of the MCS.

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FIG. 6. Composite reﬂectivity (dBZ) generated from the WRF Model at (a) 1700, (b) 2000, (c) 2330, (d) 0130, (e) 0400, and (f) 0700 LST in the 3-km domain. Compare to the observed reﬂectivities in Fig. 1. The black triangle in (d) represents the location of Chickasha.

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21 FIG. 7. Horizontal cross section of reflectivity (dBZ), wind vectors, and CAPE (J kg ) in the 3-km domain at 1 km AGL valid at (a) 1930 and (b) 0130 LST. Reflectivity contours are drawn from 0 to 70 dBZ with a contour interval of 10 dBZ. The thick black line represents the line used to calculate CAPE and CIN in the 1-km inflow region.

To examine the structure of the storm relative to the 1 km AGL. The potential temperature difference was advection of CAPE with the LLJ, we investigated the calculated within the 3-km domain in a framework potential temperature difference between the cold pool moving with the MCS (Fig. 10). From 1700 LST 3 June and the ambient air ahead of the storm at the surface and to 2030 LST 4 June, the difference is stronger at the

FIG. 8. Three curves that illustrate how CAPE and CIN evolve in three distinct inflow regions with time at the surface and at 1 km AGL. An inflow region is defined as the region containing the air flowing into the storm, and its location changes with height. A 130–150-km line was manually drawn across the various inflow regions every 30 min from 1700 LST 3 Jun to 0630 LST 4 Jun. CAPE and CIN were calculated by averaging their values at nine points along each line, roughly every 15 km. The red curve is for the surface inflow region for the first system, the blue curve is for the 1-km inflow region for the first system, and the cyan curve is for the 1-km inflow region for the second system.

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21 FIG.9.AsinFig. 7b, but for the mixing ratio (g kg ).

FIG. 10. Horizontal cross section of temperature (8C) and black contours of reflectivity (dBZ) valid for 0700 LST in the 3-km do- surface than at 1 km AGL (Fig. 11a). As the night main at 1 km AGL. Two lines were manually drawn every 30 min from 1700 LST 3 Jun to 0930 LST 4 Jun, consisting of an 85-km line progresses, the situation changes. Beginning at 2100 LST, passing through the cold pool and a 170-km line passing through the potential temperature difference is greater at 1 km the ambient air ahead of the cold pool. The cold pool line is located AGL than at the surface. The difference increases to as within 40 km behind the leading edge of the outflow and the am- large a magnitude as 68C between 0300 and 0400 LST bient air line is located within 40 km ahead of it. The temperature during the passage of the nocturnal, elevated MCS. values in Fig. 11 were obtained by averaging the values at five points along the cold pool line and nine points along the ambient Between 0000 and 0030 LST, the surface cold pool from air line. The black line with red endpoints is the cold pool line and the first convective system dies out and the cold pool the black line with blue endpoints is the ambient air line. from the second system does not form until 0100 LST. This gap in the occurrence of the cold pool warranted the exclusion of difference values at 0000 and 0030 LST. pressure, accumulated rainfall, wind speed, and wind At 1 km AGL, the cold pool from the first system also direction evolved with time at the first model level diminishes around 0000 LST, but the cold pool from the at Chickasha. The model output (Fig. 13) indicates a second system has formed by that time. The second 1.6-mb (1 mb 5 1 hPa) pressure jump between 0345 and convective system is the result of elevated convection 0350 LST that is associated with no appreciable change 2 initiation, which explains why it possesses a stronger in temperature, wind speeds of ;15 m s 1, and an abrupt potential temperature difference at 1 km AGL and later shift in wind direction. This feature is followed by a develops a surface-based cold pool. A comparison be- 2.58C increase in temperature from 0350 to 0405 LST, tween the surface potential temperature difference ob- which is then followed by a 3.78C drop in temperature served in the RAP BUFR data and simulated by the and a 1.38C increase in dewpoint from 0405 and 0410 LST WRF (Fig. 11b) suggests extremely close agreement, with the onset of rainfall a few minutes later. Ac- lending credence that the simulated MCSs have char- cording to past observations of bores (Koch et al. 1991), acteristics similar to the observed convection. these data suggest that the feature is most likely a bore To explore the generation of the cold air aloft, we just ahead of the density current and the convection, examined the evolution of the reflectivity and vertical which is also in agreement with the Chickasha surface velocity fields at 1 km AGL. The horizontal cross sec- observations (Fig. 2b). From past observational studies, tions reveal a wave pattern in the vertical motions that the warming with and following the passage of the wave forms during the mature stage of the nocturnal, elevated is likely due to the downward mixing of stable air. MCS (Fig. 12). Data at 5-min intervals from WRF were Vertical cross sections of buoyancy were created in two used to observe how variables such as temperature, directions using the 1-km domain to get a better sense of

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FIG. 11. The evolution of the potential temperature difference. (a) Comparison of the potential temperature difference at the surface and at 1 km AGL. (b) Comparison of the surface potential temperature difference observed in the RAP BUFR data and in the WRF.

how the system structure varied from south to southeast. where u and qy represent the potential temperature and The orientations selected for the vertical cross sections water vapor mixing ratio of the ambient air ahead of the 0 are north–south (N–S) and northwest–southeast (NW– density current, u is the difference between u and u, qc is SE) (Fig. 12a). We deﬁned the buoyancy ﬁeld as follows: the cloud water mixing ratio, qr is the rainwater mixing ratio, and q is the mixing ratio of frozen condensate. u0 i [ 1 : 2 2 2 2 The terms in the buoyancy equation were calculated as a B g 0 61(qy qy) qc qr qi , (1) u function of height relative to an ambient air mass well

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21 FIG. 12. Horizontal cross sections of vertical velocity (cm s ) and reflectivity (dBZ) in the 1-km domain at 1 km AGL at (a) 0430 and (b) 0700 LST. Reflectivity contours are drawn from 0 to 70 dBZ with a contour interval of 10 dBZ. The two black lines with cyan endpoints in (a) represent the paths of the cross sections in Figs. 14–16. ahead of the system. The ambient air terms were calcu- The vertical cross sections of vertical velocity in Fig. 15 lated in the 1-km domain at each model level using an that correspond to the two vertical cross sections in area average over a 10 km 3 10 km box centered in the Fig. 14 provide insight into differences found in the ambient air, which was defined to be at least 10 km ahead buoyancy field. In both vertical cross sections, the lifting of the leading edge of the lifting in the cross section. Note extends well above the nocturnal SBL, which was gen- that the vertical acceleration due to the effective buoy- erally less than 500 m [refer to Fig. 53 in Blake (2015) and ancy also includes the vertical gradient in the buoyancy the potential temperature fields in Figs. 14 and 15]. Along perturbation pressure gradient (e.g., Doswell and the N–S cross section (Fig. 15a), the strongest ascent Markowski 2004). The calculated distribution of buoy- at the leading edge of the system occurs below ;2km, ancy would, along with the dynamic terms, contribute to with additional lifting extending over the lowest 5 km. pressure perturbations that will further impact the orga- This lifting is paired with strong downward motion fol- nization and structure of the system. The perturbation lowed by an alternating ascent and descent pattern above pressure field was not diagnosed in this study. the density current in the lowest ;3 km. The pattern of The relatively deep layers of negatively buoyant air descent and ascent near the leading edge of the system (Fig. 14) ahead of the density current at the surface in extends well into the troposphere. The vertical motion both vertical cross sections are consistent with the net pattern is more complex aloft within the layers of weak upward displacement of a stable air mass, as would be positive buoyancy. In contrast, the convectively active, expected from a borelike disturbance. The boundary NW–SE vertical cross section (Fig. 15b) reveals that the layer air in Fig. 14 is indeed stable, as evidenced by the height of the strongest ascent increases with rearward sharp increase of potential temperature with height, but extent, with the largest vertical motions near the leading the stability varies between the two cross sections with a edge of the positive buoyancy aloft (see Fig. 14b). A more deeper nocturnal stable layer evident along the N–S in-depth analysis of the evolution and dynamics of the vertical cross section. The conclusion of a wave response wave field will be undertaken in the future. We will with the negative buoyancy associated with the lifting of continue to refer to this wave as a bore, since the wave is stable air is supported by the displacement of the isen- associated with net displacements of stable air in the tropes in these figures. Significantly more positive lower troposphere, forms along and moves ahead of the buoyancy is found above 4 km in the NW–SE cross leading edge of the convection, and produces variations section, with the location of the buoyant plume aloft in the surface fields consistent with bores. collocated with the leading edge of the density current, Recent observational (Bryan and Parker 2010) and corresponding to a stronger updraft in this region. modeling (Bryan and Morrison 2012) studies have

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sufficient for the CAPE to be fully realized. For the NW–SE orientation (Figs. 16b,d), the ambient environment shows large CIN below ;1 km associated with the nocturnal PBL topped by large values of elevated CAPE above ;1.5 km. While the NW–SE orientation has more CIN in the SBL, the magnitude of CIN within the layer of maximum CAPE is actually smaller than in the N–S cross section. The lifting by the bore again re- duces the CIN, as parcels are brought closer to their LFC. In this NW–SE cross section, the net displacement of the high-CAPE layer leads to buoyant convection. Previous studies (e.g., Koch et al. 2008a; Coleman and Knupp 2011) have suggested that lifting by bores creates favorable conditions by reducing CIN. However, these studies utilized values of CAPE and CIN either at the surface or within the stable boundary layer, which does not detect the important processes and highly favorable conditions aloft in association with the LLJ. A related study (Peters and Schumacher 2016) showed that flow with an SBL interacting with a cold pool led to an increase in CIN and a decrease in CAPE as the flow en- countered and passed over the outflow boundary. In our study, all of the CAPE is ingested by the system on its southeastern flank, and the CAPE and CIN go to 0(Figs. 16b,d). On the southern flank, CAPE decreases as the flow encounters and passes over the cold pool, while CIN remains about the same (Figs. 16a,c). We will now further explore the observed variation in structure around the cold pool through the application of convective system dynamics and wave theory in an effort to better understand the structure and maintenance of FIG. 13. As in Fig. 2, but for 5-min WRF output for the Chickasha mesonet site from 0330 to 0430 LST. nocturnal convective systems. shown that cold pools with a depth of 3–4 km do occur in 5. Dynamical inferences surface-based storms, so the deep layers of negative a. Convective system dynamics buoyancy in Fig. 14 do not necessarily indicate the system is elevated. Previous investigations (e.g., Colman Rotunno et al. (1988), often referred to as RKW (for 1990) have somewhat arbitrarily selected a few layers Rotunno, Klemp, and Weisman), proposed a theory for aloft to determine if convective instability is present for the optimal conditions for a vertically oriented updraft elevated convection rather than calculating the full along outflow boundaries. The optimal state for con- vertical profile of CAPE and CIN. Such approaches vection, consisting of vertically oriented updrafts, occurs would often underestimate or even miss the presence of when the negative horizontal vorticity produced baro- other elevated unstable layers. In this study, vertical clinically by the cold pool is balanced by the positive cross sections of CAPE and CIN for both orientations horizontal vorticity associated with the ambient vertical were calculated from the model data to demonstrate wind shear, resulting in a steady-state momentum force that the simulated MCS is elevated (Fig. 16). In the N–S balance (Bryan and Rotunno 2014). This balance is in- orientation (Figs. 16a,c), the ambient environment de- deed favorable for convection, as it produces deeper and picts high CIN confined to the lowest ;300 m, with the stronger ascent of warm air at the leading edge of the highest CAPE located in the layer between ;300 m and cold air (Parsons 1992). While the interaction between ;1 km. As the bore passes by, the high-CAPE layer is the vertical shear and the cold pool has a pronounced displaced upward to ;1.5–2 km and the CIN decreases impact on the vertical motions along the cold pool, a as is expected from this ascent. The lifting is not number of factors can influence the intensity and

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22 FIG. 14. Vertical cross sections of the buoyancy field (m s ) in the 1-km domain at 0430 LST for the (a) N–S and (b) NW–SE orientations. The contours represent isentropes, and the ground-relative wind vectors projected onto the path of the cross section are also drawn. The horizontal black lines denote the density current head and the vertical cyan lines mark the leading edge of the convection. The arrows indicate the density current height do and the inversion height ho. The path of each cross section is shown in Fig. 12a. duration of convective systems (e.g., magnitude of the the wind speed difference across the shear layer, there CAPE and CIN, midlevel humidity, distribution of di- will be an upward parcel displacement. Negative c/DU abatic sources within the convective system, deep ver- ratios would produce much smaller parcel displace- tical shear) so that the use of RKW to determine the ments (Moncrieff and Liu 1999). French and Parker intensity or longevity of convection has been questioned (2010) discuss how the shear in the effective inflow layer (e.g., Stensrud et al. 2005; Coniglio et al. 2012). containing a maximum in CAPE and minimum in CIN is The analysis of contrasting wind shear scenarios that more important for elevated systems than the shear in follows is similar to that of Trier et al. (2010) and Peters the SBL. Therefore, we will focus on the integrated and Schumacher (2015a,b). In order for there to be vertical wind shear over the layer of maximum CAPE. A positive horizontal vorticity, dU/dz must be greater than positive integrated dU/dz will be classified as favorable, 0, where U represents the component of the wind speed while a negative integrated dU/dz will be classified as perpendicular to the cold pool in the cross section. unfavorable. This favorable and unfavorable classifica- Within this framework, positive dU/dz means that it is tion is similar to that of Peters and Schumacher (2015a,b), oriented away from the cold air. Figure 17 from Bryan who focused on the orientation of DU with respect to the and Rotunno (2014), illustrates that for any positive outflow boundary. These studies found that DU was c/DU ratio, where c is the cold pool intensity and DU is oriented toward the cold pool on the southwestern flank

FIG. 15. As in Fig. 14, but for vertical velocity.

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21 21 FIG. 16. As in Fig. 14, but for (a) N–S CAPE (J kg ), (b) NW–SE CAPE, (c) N–S CIN (J kg ), and (d) NW–SE CIN. The vertical line denotes the leading edge of the convection.

of the MCS (unfavorable), but DU pointed toward the calculated for the southeastern flank to clarify why the warm air on the southeastern flank (favorable). bore eventually moves out ahead of the convection and Vertical profiles of the wind speed and vertical wind why the convection stays coupled to the elevated density shear in the ambient environment along the N–S and current. The vertical displacements of the isentropes NW–SE orientations are shown in Fig. 17 for 0400 LST. were utilized to obtain an estimate for how far upward Both orientations depict the expected vertical shear the parcels with the highest CAPE are displaced as they profile for an elevated system with negative vertical ascend over the bore. If the upward displacement of wind shear associated with the LLJ beneath positive parcels along the bore is smaller than their DZLFC values vertical wind shear. However, the integrated vertical ahead of the outflow, then the bore is insufficient to wind shear over the layer of maximum CAPE is quite trigger convection, and parcels instead continue to different for each orientation. On the south flank of the travel rearward of the bore until the density current lifts system, most of the CAPE is located between 0 and them to their LFCs. The parcels with the highest CAPE 1.5 km AGL (Fig. 16a), and the integrated dU/dz ahead of the system are located ;2 km AGL for the 2 is 20.217 s 1 over this layer. Conversely, the strongest NW–SE orientation (Fig. 16b), and they are displaced CAPE along the system’s southeast flank is concentrated approximately 500 m by the bore. The LFC ahead of the between 1.5 and 3 km AGL (Fig. 16b), corresponding to outflow in the NW–SE orientation is approximately 21 an integrated vertical wind shear of 0.077 s .TheNW– 3 km, yielding a DZLFC of 1 km for the parcels originat- SE orientation has a positive value for integrated dU/dz, ing at 2 km AGL. The lifting associated with the bore which is quite favorable for strong lifting along the out- occurs primarily below 2 km AGL on the system’s flow/bore, while the N–S orientation has a negative value southeastern flank, and since most of the CAPE is above for integrated dU/dz, which is unfavorable. this layer, the bore does not trigger convection and it To supplement our vertical wind shear analysis, slowly pulls away from the system. Destabilization by

DZLFC, the distance a parcel must be vertically displaced the bore drops the LFC from 3 to 2.75 km. The parcels at to reach its LFC, was computed for both orientations 2.5 km AGL, which now have a DZLFC of 0.25 km, are (Peters and Schumacher 2016). First, DZLFC was then displaced 1 km by the elevated density current, as

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21 FIG. 17. Vertical proﬁles of the horizontal wind speed (m s ) of the ambient air in the 1-km domain at 0400 LST 2 for the (a) N–S and (b) NW–SE orientations. The vertical wind shear (s 1) is color contoured in the background, and the integrated vertical wind shear is shown over the layer of maximum CAPE. indicated by the isentropes. This displacement is greater Koch et al. 1991). The reader is referred to Fig. 16 in Koch than DZLFC, and convection is observed as a result et al. (1991) for a more detailed description of the ﬂow (Fig. 18a). types. Bores will be initiated in the partially or completely

Here DZLFC was also calculated for the southern flank blocked flow regimes. It has been shown that the possible to illustrate why convection does not continue to the time-independent solutions of the shallow water equa- south (Fig. 18b). Ahead of the outflow on the N–S ori- tions are determined by two nondimensional parameters, entation, the LFC is approximately 3 km. Parcels with Do and F (Houghton and Kasahara 1968): the highest CAPE are located ;0.5 km AGL (Fig. 16a), d D 5 o (2) yielding a DZLFC of 2.5 km. The parcels are displaced o ho 1 km by the bore, which is a larger displacement than on (U 2 C ) (U 2 C ) the NW–SE orientation yet still not sufficient to reach F 5 dc 5 sffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffidc , (3) C the LFC, which drops from 3 to 2.75 km after the passage gw Du g vw h u o of the bore. The parcels at 1.5 km, which now have a vw DZLFC of 1.25 km, are then displaced 0.5 km by the density current. Therefore, convection is not favored to where Do is the nondimensional height, defined as the move to the south, which is correct as it primarily con- ratio of the density current height do to the height of the tinues to the southeast. inversion of the ambient air ho. The value of do was set equal to the height above the density current at which u b. Hydraulic theory e became greater than the surface ue of the ambient air. If Hydraulic theory is utilized to show that bores are an there were no clouds present, then the virtual potential expected response from the interaction between the cold temperature would have been sufficient. The level at 2 pool and the ambient environment. We will approximate which du/dz first drops below 0.005 K m 1 is represented 21 our environment as a two-layer, inviscid flow over an by ho. The lapse rate must also be less than 0.005 K m obstruction. Following Koch et al. (1991), we will assume over a layer at least 200 m thick. This criterion elimi- that the density current acts in an analogous manner to an nates noise in the vertical profile of du/dz from quali- obstruction. Specifically, four flow regimes are possible fying as the inversion height. In addition, F is the Froude when a density current encounters a surface-based stable number, which is the ratio of the density current speed layer: supercritical, subcritical, partially blocked, and Cdc to the phase speed of a long, internal gravity wave completely blocked flow (Rottman and Simpson 1989; (Cgw) confined within the stable layer, and Cdc was

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FIG. 18. (a) Conceptual figure illustrating how convection was able to occur on the NW–SE orientation. The jagged black line labeled ‘‘air parcel at 2 km AGL’’ represents an isentrope for the most unstable parcels, which are lifted 0.5 and 1 km by the bore and density current, respectively. The horizontal black line represents the LFC of air parcels ahead of the outflow. The clouds indicate the approximate location of convection. (b) Conceptual figure illustrating why convection did not occur on the N–S orientation. The jagged black line labeled ‘‘air parcel at 0.5 km AGL’’ represents an isentrope for the most unstable parcels, which are lifted 1 and 0.5 km by the bore and density current, respectively. estimated for both orientations using vertical cross sec- wave energy out of the layer (Crook 1988; Koch and tions of potential temperature to track how far the in- Clark 1999). One mechanism by which waves can hori- terface between the warmer, ambient air and the colder, zontally propagate is through internal reflection. The density current air moved over time. The virtual po- equation for the vertical wavenumber m2 is given by tential temperature at the surface for the ambient air is ›2U uvw, and Duvw is the inversion strength, which is the N2 › 2 difference in uy from the surface to the inversion top of m2 5 2 z 2k2 , (4) 2 2 (U 2 C ) the ambient air. (U Cbore) bore Table 2 provides a list of the flow regimes, estimated 2 Cdc, do, and mean values of ho, Duvw, Do, and F for both where N is the square of the Brunt–Väisälä frequency orientations. According to these calculations, a partially and the other terms have been previously defined (Scorer blocked flow regime is occurring in both the N–S and 1949; Crook 1988). The Scorer parameter l2, given by the NW–SE orientations at 0415 LST (Fig. 19), which means first two terms on the right-hand side of the following that a bore should be generated. In addition to do, Do, F, equation, is often utilized to diagnose the probability of a and Cdc, the bore strength bstr, bore depth h1), and wave duct. Given a packet of waves encompassing a phase-corrected bore speed Cbore are also calculated. spectrum of wave modes (each with its own specific Refer to Blake (2015) for values of these quantities. horizontal and vertical wavelength), as l2 decreases with height, some of the wave modes are trapped under an c. Linear wave theory exponentially decaying layer, while the rest of the wave According to linear wave theory, in order to maintain a modes continue to vertically propagate. long-lived wave, some mechanism must exist in order to The Scorer parameter was plotted with height at 0415 LST trap wave energy and prevent the vertical propagation of in order to determine if an appropriate wave duct is

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TABLE 2. The flow regimes, estimated Cdc, do, and mean values of layer is not infinite in the simulations, and, in accordance ho, inversion strength, Do, and F for both orientations. with theory, the waves associated with the bore do not Orientation N–S NW–SE appear to be completely trapped within the wave duct for the NW–SE orientation, as the vertical motions ex- h (m) 979.65 1285.81 o tend up to ;8 km with the wave ducts located at ;0.75– Duvw (K) 8.6 11.28 21 ; Estimated Cdc (m s ) 13.33 15.71 1.25 and 2.25–2.75 km AGL. This finding is also Estimated do (m) 537.94 1977.32 supported by the observations provided by the micro- Do 0.55 1.54 wave radiometer retrievals (Fig. 4), which also suggest F 1.67 1.75 lifting throughout the lower troposphere. We do note, Flow regime (SP, PB, CB, SB) PB PB however, that the retrievals become less dependable with increasing elevation. One noticeable difference exists between the two in place. The following phase-corrected bore speeds orientations. The N–S orientation contains seven layers 2 were utilized: 13.33 m s 1 for the N–S orientation and of positive Scorer below negative Scorer extending up to 2 15.71 m s 1 for the NW–SE orientation (Blake 2015). 8 km, while the NW–SE orientation contains three Both the N–S and NW–SE orientations contain a layer layers of positive Scorer below negative Scorer located of positive Scorer parameters in the lowest kilometer, below 4.5 km. In addition, the layer of negative Scorer at topped by a layer with a negative Scorer parameter ;4.5 km AGL for the NW–SE orientation is very weak. centered around 1 km (Fig. 20). If we estimate the mean Our observations of complex, layered wave motion 2 5 3 25 positive and negative Scorer layers to be l1 0.5 10 particularly in the N–S cross section, which we described 2 52 3 25 and l2 0.12 10 for the N–S orientation and in section 4, appear consistent with multiple positive– 2 5 3 25 2 52 3 25 l1 0.6 10 and l2 0.09 10 for the NW–SE negative Scorer layers. Further analysis based on linear orientation, both with z1, the height of the wave duct, wave theory to explain these multiple layers and how ;800 m, and the square of the horizontal wavenumber they affect vertical motions and wave structures will be 2 ; 3 27 of kx, 0.4 10 , then according to (4) a wave with a conducted in the future. Given that our previous anal- wavelength of 10 km should be trapped, which is roughly ysis revealed the southern flank of the storm is less fa- consistent with the simulations (Figs. 14 and 15). How- vorable for lifting, these results confirm that the ever, Baines (1995) assumed that the second negative southern edge of the MCS is more likely to be charac- Scorer layer is infinite; thus, the vertical wave motion in terized by a lack of convection, but potentially more the infinite layer should approach 0. The negative Scorer pronounced wave activity.

FIG. 19. Plot of the nondimensional height (Do) and the Froude number (F) at 0415 LST for the N–S (red circle) and NW–SE (blue circle) orientations. The Do, F parameter space indicates the ﬂow regime that is occurring: a, supercritical; b, partially blocked; c, completely blocked; and d, subcritical.

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25 22 FIG. 20. Two lines that illustrate how the Scorer parameter (310 m ) changes with height at 0415 LST for the N–S (red line) and NW–SE (blue line) orientations. The vertical black line 2 represents a Scorer parameter of 0 m 2.

6. Summary and conclusions potential temperature difference became significantly stronger aloft than at the surface as the MCS transi- Forecast skill for nocturnal convection remains rela- tioned into its nocturnal, elevated stage. This observa- tively low, and one reason is the difficulty in accurately tion is also consistent with previous studies that claim simulating storm systems associated with the greater the SBL inhibits deep penetrative downdrafts from occurrence of elevated convection at night. This study reaching the surface, preventing the formation of a utilized version 3.6.1 of the WRF-ARW Model to surface cold pool and indicating the lesser importance simulate a nocturnal MCS event that occurred over the of a surface cold pool in maintaining convection at night southern Great Plains on 3–4 June 2013. The system was (e.g., Trier and Parsons 1993; Parker 2008). well simulated, as evidenced by the surface data, radar, The storm’s outflow varied in three dimensions and and microwave radiometer retrievals being in relatively our application of theories of convection and wave good agreement with the simulation. dynamics provided some insight into these variations. The MCS in this study began as a surface-based sys- Shallower ascent was likely along the southern flank of tem driven by the lifting of high-CAPE and low-CIN the MCS, as expected from the vertical shear and boundary layer air. As the night progressed, nocturnal DZLFC analysis. The convectively active southeastern radiational cooling stabilized the PBL, diminishing the flank had a layer of elevated CAPE that interacted CAPE and CIN in the storm’s surface inflow region. with a deep (4–5 km) zone of negative buoyancy to However, the advection of moist, unstable air associated produce lifting. While the bore on the southeastern with the nocturnal LLJ created an inflow region of high flank did not trigger convection by itself, it assisted in CAPE and low CIN at the 1-km level, and conditions transporting high-CAPE air toward its LFC, and the aloft became increasingly favorable for convection. This DZLFC analysis revealed that the additional lifting by finding is in agreement with previous observations of the elevated density current allowed for deep convec- nocturnal convection and the idealized work (e.g., tion to occur. As a result, the bore was not coupled to French and Parker 2010), which found the source of the convective system and it slowly pulled away, while parcels with high CAPE to be 1 km AGL or higher for the convection remained in phase with the elevated elevated convection. From the cold pool analysis, the density current.

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Several differences exist between our results and pre- Follow-on studies to this work include a systematic vious idealized simulations. For example, in the Parker examination of conditions over the Great Plains show- (2008) and French and Parker (2010) studies, the con- ing that partially blocked flow regimes with favorable vection became elevated and associated with a borelike wave trapping and bores are common over the region feature only after nocturnal cooling eliminated CAPE (HPS) and a more general exploration of the charac- within the SBL. In our case, a bore response was gener- teristics of lifting in association with bores and its re- ated in the presence of low-level CAPE. In addition, the lationship to deep convection (Parsons et al. 2017, soundings used in the idealized Parker simulations were manuscript submitted to J. Atmos. Sci.). Also of interest for severe convection, while our case was more consistent is repeating the simulations using other model parame- with the class of less severe nocturnal MCSs that were terizations (i.e., the NSSL two-moment microphysics investigated in the Haghi et al. (2017, hereafter HPS) scheme, which also produced an accurate representation systematic study of IHOP_2002 (Weckwerth et al. 2004). of the event). The recent Plains Elevated Convection At The MCS in this study was associated with a well-defined Night (PECAN) field campaign (Geerts et al. 2017) will maximum in CAPE and low CIN above the SBL, with allow us to apply this analysis to a greater number greater CIN in the SBL compared to the Parker simula- of cases. tions, favoring elevated convection in our case and surface-based convection in the Parker simulations. Acknowledgments. We are grateful that this work was Another key difference is that the Parker simulations supported by NSF Grant AGS-1237404 with NSF Grant had convection staying ‘‘locked on’’ to the singular wave AGS-1229181 supporting the analysis of the microwave response, which was not generated by the cold pool until radiometer data. We also appreciate data from the quite late in the simulations. In our result, the presence Oklahoma Mesonet (available at https://www.mesonet. of partially blocked flow generated a bore response org). The WRF-ARW Model was run using the Boomer much earlier on in the simulation that moved slowly supercomputer at the University of Oklahoma. We greatly away from the density current. This difference suggests appreciate the comments of colleagues, especially the that the Parker simulations do not fall in the partially ever-patient Alan Shapiro, who was a frequent sounding blocked flow regime that is frequently observed in the board for our dynamic interpretation. Finally, we thank nocturnal environment. In unpublished work by the three anonymous reviewers for their helpful suggestions second author and colleagues (Parsons et al. 2017, and constructive feedback during the review process. manuscript submitted to J. Atmos. Sci.), it is shown that the idealized Parker simulations generally fall in the supercritical flow regime, which is represented by region REFERENCES AinFig. 19 and is located outside the distribution of Adams-Selin, R. D., and R. H. Johnson, 2013: Examination of gravity observed flow regimes in the HPS systematic study. waves associated with the 13 March 2003 bow echo. Mon. Wea. Forward propagation speeds for MCSs differ sub- Rev., 141, 3735–3756, doi:10.1175/MWR-D-12-00343.1. stantially, depending on whether they are coupled with a Ahijevych, D. A., C. A. Davis, R. E. Carbone, and J. D. Tuttle, 2004: Initiation of precipitation episodes relative to elevated bore or a density current. The density current speeds in terrain. J. 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