1. Introduction
Marsham et al. (2011) documented the evolution of a nocturnal mesoscale convective system (MCS) into a daytime squall line over the United States Great Plains on 13 June 2002 during the International H2O Project (IHOP_2002; Weckwerth et al. 2004). The current paper continues this study by using a convection-permitting numerical model to examine physical processes that may have influenced the postsunrise phase of the MCS evolution.
An interesting aspect of the diurnal cycle of Great Plains warm-season convection is its out-of-phase relationship with daytime heating. Carbone and Tuttle (2008) illustrate a maximum frequency of precipitation echoes over the 36°–42°N central Plains corridor that moves eastward from 100°W at 0600 UTC (local midnight) to 95°W by 1200 UTC (near or shortly after sunrise) and a stationary minimum precipitation frequency over the entire region at 1800 UTC (local noon).
It has been long recognized that the Great Plains nocturnal precipitation maximum is influenced by eastward-moving MCSs (e.g., Maddox 1983). Hane et al. (2008) specifically examined the postsunrise evolution of 145 nocturnal MCSs in the southern part of this region and found that about 60% decreased in intensity or dissipated from 1300 to 1700 UTC [local standard time (LST) is 6 h earlier], whereas 12% dissipated from 0900 to 1300 UTC. These results are consistent with the Carbone and Tuttle (2008) 10-yr warm-season precipitation echo climatology. However, while most MCSs dissipate or significantly weaken sometime during the morning, the remaining 28% that either maintain their intensity or strengthen (Hane et al. 2008) represent a challenge for southern Great Plains (SGP) precipitation forecasting.
Among the factors influencing postsunrise MCS evolution over the SGP is the change in mesoscale forcing associated with the cessation of the nocturnal low-level jet (LLJ), including its associated warm advection. Diurnal LLJ changes also alter the vertical wind shear, which itself has been shown to have a strong impact on the strength and organization of deep convection (e.g., Hane 1973; Rotunno et al. 1988; Coniglio et al. 2006).
The postsunrise dissipation of the LLJ over several hours is influenced by thermodynamic stability changes in the developing planetary boundary layer (PBL), which may also directly impact MCSs by altering the vertical distribution of moist static energy. These changes in thermodynamic stability associated with surface heating can also intensify MCS-induced penetrative downdrafts and surface cold pools, which could influence convective reorganization, leading to increased longevity of nocturnal MCSs. On the other hand, Crook and Moncrieff (1988) used idealized two-dimensional simulations to demonstrate that cold pools produced by microphysical processes within deep convection are not necessary to sustain organized long-lived convective systems when sufficient environmental convergence exists. This can result in so-called elevated convection (e.g., Corfidi et al. 2008), where deep moist convection is sustained primarily by air parcels originating from well above the surface. In this situation, such air parcels may be closer to their level of free convection (LFC) than near-surface air parcels (Crook and Moncrieff 1988), which are normally more susceptible to deep convection in environments with significant surface heating that lack organized mesoscale convergence.
The 13 June IHOP case, which we examine using fully three-dimensional simulations, is an excellent example of a nocturnal MCS in an environment with mesoscale convergence that underwent a substantial reorganization in convective structure after sunrise and persisted into the early afternoon (Fig. 1). In the current paper we seek to answer the following questions: how do dynamical features of the mesoscale environment influence the reorganization of the MCS? How is the MCS reorganization impacted by the surface heating after sunrise? To what degree do internal MCS processes, including the development of a downdraft-induced cold pool, influence the overall MCS evolution and its ability to persist through the morning in this region that is climatologically less favorable for daytime convection?
Regional surface analysis at (a) 0900, (b) 1200, (c) 1500, and (d) 1800 UTC 13 Jun 2002 including NEXRAD radar mosaics of maximum reflectivity in a vertical column and objectively analyzed potential temperature (1-K contour intervals) and plots of surface winds (full barb = 5 m s−1, half barb = 2.5 m s−1), potential temperature (K), and water vapor mixing ratio (g kg−1). The barbed symbols depict mesoscale boundaries following the proposed mesoanalysis conventions of Young and Fritsch (1989).
Citation: Journal of the Atmospheric Sciences 68, 12; 10.1175/JAS-D-11-0112.1
In section 2, we summarize the observed MCS structure and environmental characteristics during the postsunrise reorganization of deep convection. A more comprehensive analysis of the MCS life cycle, including its initiation, appears in Marsham et al. (2011). Section 3 describes our numerical model and design of sensitivity studies. In section 4 we use radar, surface mesonet observations, and output from a control simulation to establish the realism of this simulation and to illustrate salient aspects of the postsunrise MCS evolution. Section 5 examines the impact of mesoscale forcing (including that associated with the LLJ and synoptic cold front) on the environment in advance of the deep convection, and uses trajectory analysis to elucidate the evolution of source levels for deep convection in the control run. Model sensitivity studies are used in section 6 to examine likely impacts of various physical processes on MCS evolution.
2. Overview of the 13 June 2002 MCS from surface, sounding, and radar observations
The 13 June IHOP MCS evolved from a multiple banded nocturnal system (Fig. 1b) to an MCS resembling the leading line–trailing stratiform archetype (Fig. 1d) frequently described by others (e.g., Smull and Houze 1985; Johnson and Hamilton 1988; Houze et al. 1990). Also presented in Fig. 1 is the corresponding evolution in surface potential temperature θ, which was constructed using a two-pass Barnes objective analysis of conventional surface airways observations (SAOs) and mesonet observations with the General Meteorological Package (GEMPAK) utility (Koch et al. 1983). The parameters selected for this analysis are chosen to emphasize large mesoscale (L > 100 km) features. Thus, sharp boundaries, such as convectively induced surface outflows and the leading edge of synoptic fronts, are not well resolved. These boundaries are subjectively analyzed and indicated in Fig. 1 using the conventions of Young and Fritsch (1989).
The MCS is influenced by three distinct air masses during its postsunrise transition (Figs. 1b–d). These include (i) a postfrontal air mass over Kansas, (ii) an air mass over western and central Oklahoma that is recovering from stabilization in the wake of a dissipated earlier MCS (cf. Figs. 1a,b) whose initiation is described by Weckwerth et al. (2008), and (iii) a prefrontal air mass that influences the southern end of the MCS by late morning over northern Texas and southeastern Oklahoma (Fig. 1d). The vertical structure of these air masses is respectively characterized by the 1200 UTC Topeka, Kansas (TOP; Fig. 2a), Norman, Oklahoma (OUN; Fig. 2b), and Dallas–Ft. Worth, Texas (FWD; Fig. 2c) radiosondes, whose locations are shown in Fig. 1b.
Skew T–logp plots at 1200 UTC 13 Jun 2002 of National Weather Service rawinsonde data from (a) Topeka, Kansas, (b) Norman, Oklahoma, and (c) Dallas–Ft. Worth, Texas. The locations of these stations relative to the mesoscale boundaries discussed in the text are indicated in Fig. 1b and relevant thermodynamic and kinematic features of these soundings discussed in the text are indicated on the individual panels. The black curved lines indicate examples of paths for near-surface and elevated air parcels predicted by parcel theory. The panels in this figure are modifications to images obtained from the University of Wyoming (from http://weather.uwyo.edu/upperair/sounding.html).
Citation: Journal of the Atmospheric Sciences 68, 12; 10.1175/JAS-D-11-0112.1
A shallow well-mixed PBL already exists by 1200 UTC at FWD (Fig. 2c) south of the dissipating outflow of the previous MCS (Fig. 1b). Within this deepening PBL the air is conditionally unstable with convective available potential energy (CAPE) of about 1300 J kg−1. However, both a weak stable layer immediately above the PBL and a strong and dry midlevel inversion based near 700 hPa (Fig. 2c) inhibit deep convection. The simulations to be reported later indicate substantial subsequent surface heating of this air, resulting in significant increases of CAPE favorable to the MCS longevity later in the diurnal cycle.
The air mass at OUN (Fig. 2b), which most directly influences the MCS during the early stages of its reorganization over Oklahoma (Figs. 1b,c), is characterized by a 20–30-hPa-deep layer of strong stability near the surface that is likely a manifestation of both cooling from cloud microphysical processes associated with the previous MCS and nocturnal surface radiative cooling. Above the surface-based stable layer is a roughly 250-hPa-deep moist and conditionally unstable layer containing CAPE, which includes the remnant nocturnal LLJ near 900 hPa. Deep convection is supported by this elevated conditionally unstable layer with maximum CAPE of about 2000 J kg−1 and significant vertical wind shear (Fig. 2b).
Throughout its life cycle, a significant portion of the MCS remains north of the synoptic cold front (Fig. 1). The 1200 UTC TOP sounding (Fig. 2a) that characterizes this environment shows a roughly 150-hPa deep postfrontal layer. However, this postfrontal layer is only slightly stable and quite moist with CAPE of about 1200 J kg−1 that supports moderately strong deep convection.
3. Numerical model and experiment design
a. Numerical model
Our simulations use version 3.0.1 of the Advanced Research Weather Research and Forecasting (ARW-WRF) model (Skamarock and Klemp 2008). The Thompson et al. (2008) bulk microphysical parameterization, which predicts the mass of cloud water, cloud ice, rain, snow, and graupel, is used in all simulations. Other physical parameterizations include the Rapid Radiative Transfer Model (RRTM) longwave (Mlawer et al. 1997) and Dudhia (1989) shortwave radiation schemes. The PBL parameterization (Janjic 1990, 1994) predicts turbulent kinetic energy (TKE) and governs vertical mixing between model layers. Subgrid horizontal mixing is determined using a Smagorinsky-type first-order closure.
The ARW-WRF is coupled to the Noah land surface model (Ek et al. 2003). The land surface model (LSM) contains a single vegetation canopy layer and predicts volumetric soil moisture and soil temperature in four soil layers. The depths of the soil layers are sequentially 0.1, 0.3, 0.6, and 1.0 m and the root zone is contained in the upper 1 m (top three layers).
b. Model domain and initialization
Our simulations use a single horizontal domain of 800 × 750 grid points (Fig. 3a) with 3-km horizontal grid spacing. The chosen domain size and grid spacing reflect compromises aimed both at capturing relevant large-scale forcings for deep convection and the ability to explicitly simulate the deep convection for a major portion of the MCS life cycle. While not ideal for completely resolving some individual convective cells (Bryan et al. 2003), this horizontal grid spacing resolves well the mesoscale aspects of the deep convection (e.g., Weisman et al. 1997). It is also expected to enable a reasonable assessment of what the relative strength of the convection is at different times and how it is affected by various physical processes (section 3c).
(a) Model domain used in all simulations and smoothed terrain with (b) the initial surface conditions for the CTRL and ADIA simulations started at 0900 UTC 13 Jun 2002 over the part of the model domain indicated by the dashed rectangle in (a). The large and small barbed features in (b) indicate the position of the synoptic cold front and mesoscale outflow boundary from the previous MCS, respectively, discussed in the text.
Citation: Journal of the Atmospheric Sciences 68, 12; 10.1175/JAS-D-11-0112.1
The vertical grid contains 42 levels. Because of the potential importance of adequately resolving surface-based stable layers both within the convection and in its environment (cf. Fig. 2b), the model contains 13 of the vertical levels within its lowest km, with 7 of these levels within the lowest 300 m AGL. Vertical grid spacing is about 1 km near the z ~ 20 km model top. A 5-km-deep absorbing layer (Klemp et al. 2008) mitigates reflection of vertically propagating gravity waves off the model top.
The initial conditions for the atmospheric model are obtained from the National Centers for Environmental Prediction (NCEP) Environmental Data Assimilation System (EDAS) analyses, which have a horizontal grid spacing of about 40 km. Lateral boundary conditions are supplied by the EDAS analyses with 3-h frequency. Motivated by the desire to capture the major convective reorganization of the MCS, which began near sunrise (~1100 UTC), preliminary simulations were initialized using the 0000, 0600, and 0900 UTC 13 June EDAS analyses. Unfortunately, none of these simulations was able to accurately capture the initiation and evolution of the MCS over Oklahoma. Consequently, we use the 1200 UTC 13 June EDAS analyses but begin our control simulation (CTRL) 3 h earlier at 0900 UTC. We speculate that our increased success using the later analysis is partly related to information from the 1200 UTC radiosonde launches (not available in earlier analyses), resulting in a more accurate representation of the environment in the wake of the earlier nocturnal MCS (Fig. 1a), which likely influences our MCS of interest.
The 3-h earlier model start ensures that the initiation of the simulated MCS is well underway by the time significant postsunrise environmental changes related to a generic diurnal cycle begin. We justify our approach by noting that (i) diurnal changes are at their approximate minimum in the hours just before and around sunrise (i.e., 0900–1200 UTC or 0300–0600 LST) and (ii) the positions of relevant surface features in the model initialization, including the synoptic front, the old MCS outflow, and the rain-cooled region in its wake (Fig. 3b) are reasonably similar to their counterparts in the 0900 UTC surface analysis (Fig. 1a). Moreover, we note that our study emphasizes physical processes influencing deep convection in simulations based on an observed case rather than specific quantification of a model’s ability to accurately forecast the case. Still, it is recognized that there are likely to be some artifacts related to the 3-h offset between the model analysis and model start time (section 4).
c. Sensitivity studies
Both environmental forcing and internal dynamics related to precipitation processes are anticipated to play a role in the postsunrise reorganization of the MCS. Changes in environmental forcing may be associated with both (i) mesoscale dynamics associated with the cold front and LLJ and (ii) diurnal variations in surface heating.
An adiabatic simulation, in which all microphysical temperature tendencies are withheld, enables us to examine the effects of mesoscale forcing (including environmental vertical motions) on the convective environment without the complicating effects of latent heat release and resulting buoyant convection.1 This simulation, termed ADIA (Table 1), is otherwise identical to CTRL, including the 0900 UTC initialization.
Listing of simulations discussed in the paper.
To examine the effect of the postsunrise changes in surface heating on the evolution of the MCS and its environment we withhold surface heat and moisture fluxes in the sensitivity simulation NOFX (Table 1). This is accomplished by both excluding the LSM (section 3a) and setting default fluxes in the WRF surface layer parameterization to zero. NOFX is initiated around sunrise at 1100 UTC using t = 2 h model output from CRTL.
One of the hypothesized reasons why typical squall lines can have a structure different from nocturnal MCSs is the greater propensity for development of convection-organizing gravity currents (cold pools) in squall lines (e.g., Fritsch and Forbes 2001, p. 341). The gravity currents are themselves a cumulative effect of cold, negatively buoyant downdrafts penetrating to the surface. In environments with conditionally unstable lapse rates, including the central United States, the most important contributors to cold penetrative downdrafts are latent cooling processes including evaporation, melting, and sublimation. The simulation NOLC (Table 1) tests this organizing effect on the MCS evolution by withholding microphysical temperature tendencies, but, unlike for ADIA, only those resulting from latent cooling processes. Similar to NOFX, NOLC is initiated at 1100 UTC using t = 2 h model output from CRTL. However, unlike for NOFX the delay in the initiation of NOLC is to ensure that a realistic nocturnal MCS forms before eliminating the effects of latent cooling rather than to test effects of a more diurnally dependent process (i.e., environmental surface warming).
4. Overview of the control simulation and comparison with observations
Figure 4 presents the evolution of both the observed MCS and the MCS simulated in CTRL. There is an approximate 1-h lag in the position of the observed MCS from that of the simulated MCS at both near sunrise and during the mid-to-late morning, which is likely influenced by the model being initialized with an analysis that is later than its start time (section 3b). Therefore, model reflectivity is shown 30–60 min before observed reflectivity in Fig. 4, so that the model and observations are at more similar stages of evolution than in simultaneous plots.
The evolution of maximum column reflectivity during the MCS transition (a),(b) revealed by NEXRAD radar mosaics and (c),(d) in simulation CTRL. The locations OK45 and OK15 are Oklahoma Surface Mesonet stations whose time series and those of the closest model grid points are presented in Fig. 5.
Citation: Journal of the Atmospheric Sciences 68, 12; 10.1175/JAS-D-11-0112.1
Although the details differ, around sunrise both the observations (Fig. 4a) and CTRL (Fig. 4c) have a complex reflectivity pattern that includes both well-organized northwest–southeast (NW–SE)-oriented rainbands and less-organized developing northeast–southwest (NE–SW)-oriented rainbands. The NE–SW-oriented rainbands grow upscale, become more organized, and eventually compose the southern part of the MCS as it moves into central Oklahoma by midmorning in both the observations (Fig. 4b) and CTRL (Fig. 4d). One difference is that in the intervening period between sunrise and late morning the observed squall line in south–central Oklahoma (Fig. 4b) evolves from a merger of the easternmost NE–SW-oriented line of developing cells near the western Oklahoma–Texas border and a more established in NE–SW line in the Texas Panhandle (Fig. 4a). In contrast, the simulated squall line (Fig. 4d) grows directly from the developing NE–SW-oriented rainband in western Oklahoma (Fig. 4c).
The approximate 1-h lag in the position of the observed MCS from the simulated MCS noted in Fig. 4 is also evident in a comparison of selected variables from Oklahoma Surface Mesonet (Brock et al. 1995) time series and their counterparts from CTRL (Fig. 5). The mesonet stations OK45 in western Oklahoma (Figs. 5a,c,e,g) and OK15 in central Oklahoma (Figs. 5b,d,f,h) respectively sample the passage of the developing NE–SW rainband closely following sunrise (Figs. 4a,c) and the maturing squall line later that morning (Figs. 4b,d).
Time series of surface weather from Oklahoma Surface Mesonet stations (solid) and the closest model grid point in simulation CTRL (dotted). Locations of the mesonet stations relative to the observed and simulated radar reflectivity are shown in Fig. 4.
Citation: Journal of the Atmospheric Sciences 68, 12; 10.1175/JAS-D-11-0112.1
Despite differences in timing, CTRL captures the MCS-influenced surface conditions during its reorganization period at both of these mesonet locations. For example, both the mesonet stations and CTRL indicate heavy convective rainfall at the leading edge of the MCS for the developing (Fig. 5a) and maturing squall line (Fig. 5b) stages. CTRL also accurately simulates the small (1°–2°C) and gradual surface temperature changes that accompany the convective rainfall in the developing stage (Fig. 5c) and the contrasting larger (5°–6°C) and more abrupt cooling and increases in wind speed that occur with the convection farther east and later in the morning (Figs. 5d,f). The 3-hPa mesohigh pressure perturbation during this later stage (Fig. 5h), which commences with the arrival of the rain (Fig. 5b), surface cooling (Fig. 5d) and wind increases (Fig. 5f), is also well simulated.
The 5-hPa observed surface pressure perturbation in the earlier stages of the transition toward an MCS with a NE–SW-oriented convective region is even stronger than that during the later squall line (cf. Figs. 5g,h) despite much weaker surface cooling (Figs. 5c,d). The mesohigh during these earlier stages was simulated by CTRL but is only about half as strong as observed (Fig. 5g). Despite the underestimation in the early mesohigh strength, subsequent pressure decreases of about 3 hPa that occur during the cessation of lighter rainfall near the back edge of the later squall line (Figs. 4b,d) are well simulated (Fig. 5g). The 1530–1600 and 1630–1700 UTC pressure minima in CTRL and the observations, respectively, reflect the wake low commonly observed in squall-like leading line–trailing stratiform MCSs (e.g., Johnson and Hamilton 1988).
A well-simulated aspect important to MCS organization is the dichotomy in surface changes during different stages. The abrupt wind shift and pressure rises that coincide with the substantial temperature fall with the MCS leading edge passage in the late morning characterize a gravity current outflow. In contrast, the pressure rises and wind speed changes substantially preceding (by about 30–60 min) a much less significant temperature fall earlier in the morning is consistent with a gravity wave outflow influenced by cooling above the surface (e.g., Haertel et al. 2001).
5. Evolution of the control simulation MCS and relationship to its environment
The comparisons in the previous section have established broad similarities in the reorganization stage of the actual MCS and that simulated in CTRL. This motivates us to examine in greater detail our simulation results to elucidate possible key factors influencing this transition of the evolving remnant nocturnal MCS into a daytime squall line.
a. Mesoscale forcing influencing the onset of the NE–SW-oriented convection
A strong regional baroclinic zone extending from north central Oklahoma through the northern Texas Panhandle at the beginning of the CTRL and ADIA simulations is associated with confluence of the southwesterly LLJ and postfrontal northeasterlies (Fig. 6a). During the next 2 h an elongated zone of enhanced moisture at 2 km MSL becomes established in ADIA over a region of warm advection in western and central Oklahoma within the southern part of the broader confluence zone (Fig. 6b). This zone of enhanced moisture at 1100 UTC is located within a time-averaged region of upward motion (ω < 0) from the previous 2 h (Fig. 7a). A NW–SE vertical cross section averaged for 60 km across its orientation indicates 1–2-km-deep upward displacements of the 1100 UTC water vapor mixing ratio qυ contours within the zone of preceding upward motion (Fig. 8), which confirms the major role of the mesoscale ascent on the simulated deeper moisture.
Potential temperature (dashed lines, 1-K contour intervals) and ground-relative horizontal winds at 2 km MSL for the adiabatic simulation at (a) 0900 (t = 0 h) and (b) 1100 UTC 13 Jun (t = 2 h). The thick gray contours in each panel enclose regions with water vapor mixing ratio exceeding 13 g kg−1. The thick arrows are approximate streamlines indicating the deformation zone discussed in the text. The transect NW–SE in (b) indicates the location of the center of the vertical cross section in Fig. 8.
Citation: Journal of the Atmospheric Sciences 68, 12; 10.1175/JAS-D-11-0112.1
(a) 800-hPa potential temperature (dashed as in Fig. 6), ground-relative horizontal winds, and ω time-averaged from 0900 to 1100 UTC for simulation ADIA. The thick solid contour denotes regions where 1100 UTC water vapor mixing ratio exceeds 13 g kg−1. (b) As in (a), but ω is derived from quasigeostrophic theory. In (a) the transect NW–SE indicates the center of the vertical cross section presented in Fig. 8 and the letters A, B, and C indicate the locations of model soundings whose evolution is shown in Fig. 11.
Citation: Journal of the Atmospheric Sciences 68, 12; 10.1175/JAS-D-11-0112.1
Time-averaged (0900–1100 UTC 13 Jun 2002) ω and cross section parallel horizontal winds (thin contours, negative values dashed) and 1100 UTC water vapor mixing ratio (thick contours). This vertical cross section from simulation ADIA is taken along the transect NW–SE in Fig. 7a and is horizontally averaged for 60 km in the direction normal to its orientation.
Citation: Journal of the Atmospheric Sciences 68, 12; 10.1175/JAS-D-11-0112.1
The two bands of mesoscale ascent (Fig. 7a) within larger-scale deformation (Fig. 6), one embedded within the baroclinic zone of the synoptic cold front over Kansas and the other to its south within the LLJ warm advection zone in western Oklahoma, suggest the importance of quasigeostrophic processes. To verify this, quasigeostrophic ω was calculated by solving its Q-vector form [Holton 1992, Eq. (6.35)] subject to a terrain-influenced lower boundary condition. The southernmost band of total ascent associated with the deeper moisture in ADIA (Fig. 7a) has a very similar location and pattern of quasigeostrophic ascent (Fig. 7b). However, the magnitude of the quasigeostrophic ascent for this location is only one-third to one-half that of the total ascent (cf. Figs. 7a,b), which is not surprising given the highly ageostrophic nature of the nocturnal LLJ.
The southern band of mesoscale ascent and associated deepening of the moisture diminishes as the new diurnal cycle progresses and the LLJ decays. However, the onset of the NE–SW convective rainband in CTRL is strongly influenced by these events near sunrise as evinced by the similarity in its nearby pattern of antecedent 2-km MSL moisture (Fig. 9) with that seen in ADIA (Fig. 7a). While the spatial distribution of enhanced lower-tropospheric moisture focuses where new convective development occurs within the MCS, its evolving orientation is likely influenced by the vertical shear associated with the strong LLJ (Fig. 10a).
The 2-km MSL potential temperature (dashed, 1-K contours), ground-relative horizontal winds, and 13 g kg−1 water vapor mixing ratio contour at 1100 UTC 13 Jun 2002 as in Fig. 6a, but for simulation CTRL. The overlaid simulated radar reflectivity is 2 h later (1300 UTC) from simulation CTRL. The letters A, B, and C indicate the locations of model soundings from CTRL whose evolution is shown in Fig. 11.
Citation: Journal of the Atmospheric Sciences 68, 12; 10.1175/JAS-D-11-0112.1
The 0.5-km AGL virtual potential temperature (solid contours), ground-relative 0.5-km AGL horizontal winds (blue barbs), and 0.5–2.5 km AGL vertical wind shear (red barbs) with simulated radar reflectivity (gray shading) for simulation CTRL at (a) 1300 (t = 4 h) and (b) 1500 UTC (t = 6 h) 13 Jun 2002. The horizontal winds and vertical shear within and to the rear of the MCS are eliminated to emphasize conditions in the MCS inflow environment.
Citation: Journal of the Atmospheric Sciences 68, 12; 10.1175/JAS-D-11-0112.1
Previous studies have emphasized the importance of vertical shear to squall line organization in the presence of both gravity currents (e.g., Thorpe et al. 1982; Rotunno et al. 1988) and gravity waves (e.g., Schmidt and Cotton 1990), with enhanced development occurring in the downshear direction. Figure 10a illustrates moderate vertical shear of ΔU ~ 10 m s−1 (red barbs) from 0.5 to 2.5 km AGL oriented locally normal to the developing NE–SW rainband at 1300 UTC. The LLJ (blue barbs), which is still moderately strong (10–15 m s−1) at 1300 UTC, constitutes the base of this vertical shear layer. The direction of the vertical shear varies as one moves southeastward ahead of the NE–SW convective rainband of CTRL (Fig. 10a). This shear variation is also evident in simulation ADIA (not shown), suggesting that it is partly a manifestation of the mesoscale deformation zone into which the lower-tropospheric flow intrudes (Fig. 6b). The line-normal shear ahead of the MCS (Fig. 10a) weakens as the LLJ decays, only after the MCS transition to a NE–SW-oriented squall line on its southern end has become well established (Fig. 10b). We examine the evolving vertical structure of the MCS and its relationship to the vertical shear and other environmental and internal factors in greater depth in section 6.
Further inspection of Fig. 9 indicates that well-organized convection develops at some locations along the deep moisture zone (point A) but not at others (point B). The convection is also delayed in a region that had less moisture at 2 km MSL farther east (point C) despite experiencing similar mesoscale ascent from 0900 to 1100 UTC (cf. Fig. 7a). The 0900–1130 UTC evolution of lower-tropospheric thermodynamic soundings at these locations (Fig. 11) helps clarify reasons accounting for these regional differences in convection initiation.
Evolution of model soundings from simulation CTRL at location (a) A, (b) B, and (c) C depicted in Fig. 9. In (a) the thick lines indicate the approximate path of vertical displacement at location A (Fig. 9) for the time-averaged vertical motion pattern in Fig. 7a.
Citation: Journal of the Atmospheric Sciences 68, 12; 10.1175/JAS-D-11-0112.1
At 0900 UTC the moist conditionally unstable layer in CTRL above the surface at point A (Fig. 11a) is similar to the 1200 UTC observed conditions at OUN (Fig. 2b) located about 150 km to the east. The subsequent further cooling and moistening resulting in the near saturation of this layer in CTRL is consistent with the time-averaged adiabatic ascent, with 2-h time-averaged 800-hPa ω ~ −6 μb s−1 (Fig. 7a) accounting for vertical displacements of about 40 hPa (Fig. 11a). The negligible convective inhibition (CIN) by 1130 UTC in the conditionally unstable layer (Fig. 11a) resulting from mesoscale ascent allows deep convection originating from elevated air parcels to proceed with little or no localized lifting (e.g., Crook and Moncrieff 1988). By 1230 UTC an elevated moist absolutely unstable layer (MAUL; Bryan and Fritsch 2000) was located at point A (not shown).
Although 800-hPa mesoscale ascent is important to convection initiation at point A, not all regions within the ascent band (Fig. 7a) experience such rapid or significant convection initiation (Fig. 9). Point C located to the northeast of point A in Fig. 9 was still influenced at 0900 UTC by strong lower-tropospheric drying (Fig. 11c) that occurred in the wake of the earlier nocturnal MCS (cf. Fig. 1a). In this sounding from CTRL, the characteristic “onion” shape of combined temperature and dewpoint profiles are a manifestation of unsaturated mesoscale downdrafts (Zipser 1977). Here, the lower- tropospheric cooling and moistening that occurs between 0900 and 1130 UTC (Fig. 11c) is consistent with the time-averaged vertical motion (Fig. 7a). However, the dry antecedent conditions prevent point C from experiencing local convection initiation and the simulated MCS arrives here only after initiating and intensifying well upstream. Only short-lived and poorly organized convection develops in CTRL over the Texas Panhandle (Fig. 4c) near point B (Fig. 9). This is related to the near-surface cooling associated with the cold front passage, which locally reduced the equivalent potential temperature beneath 800 hPa (Fig. 11b).
b. Evolution of the thermodynamic environment and source level for convective updrafts
The foregoing analysis has established the importance of mesoscale ascent in the reorganization of the southern part of the MCS into a primary NE–SW-oriented rainband in CTRL and strongly suggested that the onset of this transition involved elevated convection. We now present analyses that quantify how the properties of environmental air parcels entering the convection change after sunrise.
Over our region of interest in Oklahoma, the convective available potential energy of the most unstable air parcels in a vertical column (MUCAPE) within the MCS inflow at 1200 UTC (0600 LST) is maximized over a 100–200-km-wide zone in immediate advance of the MCS (Fig. 12a). This zone of maximum MUCAPE results from air parcels situated well above the surface (Fig. 12b) and is related to horizontal maxima of water vapor (Fig. 12a).
(left) The 1-km AGL MCS-relative horizontal winds and water vapor mixing ratio (thin contours) and MUCAPE for simulation CTRL at (a) 1200 (t = 3 h), (c) 1400 (t = 5 h), and (e) 1600 UTC 13 Jun (t = 7 h). (right) The 1-km AGL ground-relative horizontal winds, potential temperature (K, thin contours) and height (km AGL) of the most unstable parcel at (b) 1200 (t = 3 h), (d) 1400 (t = 5 h), and (e) 1600 UTC 13 Jun (t = 7 h). In each panel the thick contour encloses simulated radar reflectivity exceeding 30 dBZ.
Citation: Journal of the Atmospheric Sciences 68, 12; 10.1175/JAS-D-11-0112.1
Ground-relative winds and θ in Fig. 12b indicate that the remnant NW–SE-oriented rainband from the southern part of the MCS (see also Fig. 4c) is located within the warm advection zone near the northern terminus of the LLJ, as found in the observations (Marsham et al. 2011). Two-hour back trajectories are calculated from 1200 UTC updraft cores at 8 km MSL with maximum vertical velocity wmax > 8 m s−1 (Fig. 13) using 3-min model output and a time step of 30 s. These trajectories comprise elevated source parcels from 1.15 to 1.45 km MSL (~0.7–1.1 km AGL) near the top of the LLJ. Consistent with the mesoscale vertical motion (section 5a), the air along these trajectories gradually rises about 200 m in 1.5 h before entering convective updrafts (Fig. 14a). During this ascent these trajectories become nearly saturated (Fig. 14b) and experience substantial reductions of CIN (Fig. 14c).
Convective 8-km MSL updraft cores from simulation CTRL at 1200 UTC 13 Jun (t = 3 h). The horizontal locations of the NW–SE convective band back trajectories analyzed in Fig. 14 are overlaid on the plot.
Citation: Journal of the Atmospheric Sciences 68, 12; 10.1175/JAS-D-11-0112.1
(a) Trajectory height (km MSL), (b) RH (%), and (c) CIN (J kg−1) along individual back trajectories (color curves) and their average (thick black curve) emanating from 8-km MSL NW–SE band updraft cores with w ≥ 8 m s−1 at 1200 UTC 13 Jun 2002 (t = 3 h) in simulation CTRL. The horizontal paths of these trajectories are shown in Fig. 13.
Citation: Journal of the Atmospheric Sciences 68, 12; 10.1175/JAS-D-11-0112.1
This NW–SE rainband dissipates as the nocturnal LLJ rapidly weakens (Fig. 12d), leaving the NE–SW-oriented convection as the dominant feature in the southern part of the MCS (cf. Fig. 4). The dissipation of the NW–SE rainband may be affected by both the cessation of strong warm advection as the LLJ decays and the weakening of horizontal convergence near its terminus. The NW–SE rainband dissipation is also likely influenced by the general eastward component of MCS motion as a whole toward a more stable region. After a brief decrease, the MUCAPE does eventually increase along the southern end of the MCS (Fig. 12e) as the most unstable parcels in its advance are located near the surface by about 5 h after sunrise (Fig. 12f).
Back trajectories are also calculated from 8-km MSL updraft cores for the NE–SW rainband depicted in Fig. 13, which are stronger than those in the more transient NW–SE rainband. Figure 15 indicates that similar to the NW–SE rainband, the NE–SW rainband has updraft cores composed of elevated source parcels, often originating from even greater heights (Figs. 15a,b). Only by late morning as PBL depths have increased to an average of 1.1 km AGL do the maximum updraft parcels originate from primarily within the PBL (Fig. 15d).
Air parcel origin height for backward trajectories emanating from 8-km MSL updraft cores with w ≥ 8 m s−1 in the NE–SW-oriented rainband from CTRL (ordinate) compared with PBL heights calculated from the model PBL scheme (abscissa). The thin diagonal line indicates where these quantities have equal height values.
Citation: Journal of the Atmospheric Sciences 68, 12; 10.1175/JAS-D-11-0112.1
Corfidi et al. (2008) suggest that transitions between purely elevated and surface-based convection may not always be distinct or easy to predict. Although our simulated maximum updraft cores ingest primarily elevated air until late morning, we find some contribution to MCS convective updrafts from near-surface air begins shortly after sunrise and steadily increases throughout the morning. This is illustrated with forward trajectories released from 200 m AGL in immediate advance of the NE–SW rainband (Fig. 16). At 1200 UTC the majority of these trajectories rise only slightly before descending (Fig. 16b). However, several trajectories reach their LFC and a few participate in troposphere-deep updrafts. By 1400 UTC there is a more continuous spectrum of trajectory depths (Fig. 16d), supporting Marsham et al.’s (2011) conclusion (based on surface and sounding observations) that surface-based convection likely becomes an increasing contributor as the MCS reorganizes into a squall line.
(left) Vertical motion with updrafts (blue contours, with 0.5 m s−1 intervals, starting with w ≥ 0.5 m s−1) and downdrafts (red dashed contours, with 0.5 m s−1 intervals, starting with w ≤ −0.5 m s−1) and ground-relative horizontal winds at model vertical level 6 (~200 m AGL) of simulation CTRL at (a) 1200 (t = 3 h) and (c) 1400 (t = 5 h) UTC 13 Jun 2002. The numbers are identification numbers for forward trajectories released from these locations at model vertical level 6 (~200 m AGL). (right) The heights for these forward trajectories during the following 1.5 h.
Citation: Journal of the Atmospheric Sciences 68, 12; 10.1175/JAS-D-11-0112.1
6. Sensitivity of MCS structure and evolution to selected physical processes
The overall morphology of the simulated reflectivity pattern from 1300 UTC (a few hours after sunrise) to 1600 UTC (late morning), including its horizontal scale and orientation, is not strongly affected by the warming and moistening of the PBL or by latent cooling within the MCS (Fig. 17). This suggests a primary role of mesoscale forcing discussed in the previous section in organizing the convection.
Surface potential temperature, ground-relative horizontal winds, and simulated maximum radar reflectivity in a vertical column for simulations (a),(b) CTRL, (c),(d) NOFX, and (e),(f) NOLC at (left) 1300 and (right) 1600 UTC 13 Jun 2002. The rectangle through the NE–SW-oriented band of large reflectivity in each panel indicates the orientation and location of horizontally averaged (width dimension of rectangle) vertical cross sections displayed in Figs. 19 and 20. Note that the positions of these cross sections are identical among the different simulations but vary between 1300 and 1600 UTC.
Citation: Journal of the Atmospheric Sciences 68, 12; 10.1175/JAS-D-11-0112.1
Some differences in the strength of convective updrafts and downdrafts (Fig. 18) and MCS vertical structure (Figs. 19 and 20) among simulations CTRL, NOFX, and NOLC are nevertheless apparent. In Fig. 18 horizontal model grid points surrounding the MCS are counted where maximum updrafts and downdrafts meet thresholds for both strong (wmax > 10 m s−1, wmin < −2.5 m s−1) and weak (wmax > 1 m s−1, wmin < −1 m s−1) convection. Downdrafts of convective intensity may result from a variety of factors including latent cooling processes, precipitation loading, and gravity waves. Since in the current comparisons among simulations we are most interested in effects of the latent cooling, which are most prevalent at and beneath the melting level, at each horizontal grid point we consider only vertical grid points below 4.5 km MSL for finding wmin. In finding wmax we consider vertical grid points up to 12 km MSL, which is the approximate depth of the troposphere. The gridpoint frequencies in Fig. 18 are affected by both the strength and the coverage of convection, but these factors can be discriminated somewhat by noting differing temporal trends for the different thresholds. For example, in a situation where the number of wmax > 10 m s−1 horizontal grid points is not changing but the number of wmax > 1 m s−1 horizontal grid points is increasing significantly, one could conclude that the coverage of the convection is increasing but its average strength is weakening.
Time series of the model horizontal grid point frequencies of convective updrafts and downdrafts whose strength exceeds the indicated thresholds within a rectangular region surrounding the simulated MCS in CTRL (solid), NOFX (dotted–dashed), and NOLC (dotted).
Citation: Journal of the Atmospheric Sciences 68, 12; 10.1175/JAS-D-11-0112.1
1300 UTC 13 Jun 2002 vertical cross sections for simulations (a),(b) CTRL and (c),(d) NOLC along the transect NW–SE displayed in Figs. 17a and 17e, respectively. These vertical cross sections are averaged for 60 km across the axis of their NW–SE orientation. (a),(c) Vertical velocity with annotated maximum convective updrafts and downdrafts (cm s−1), and virtual potential temperature (2-K contour intervals). (b),(d) Equivalent potential temperature and the component of MCS-relative horizontal winds in the direction parallel to the cross sections with solid contours indicating northwesterly or rear-to-front system-relative flow, and dashed contours indicating southeasterly or front-to-rear system-relative flow.
Citation: Journal of the Atmospheric Sciences 68, 12; 10.1175/JAS-D-11-0112.1
As in Fig. 19, but at 1600 UTC 13 Jun 2002 for simulations (a),(b) CTRL, (c),(d) NOFX, and (e),(f) NOLC along the transect NW–SE in Figs. 17b, 17d, and 17f, respectively. Note that NW–SE is in a different location for these vertical cross sections than in Fig. 19. The rectangles in (a) and (c) indicate the locations where averages are used to compute the environmental vertical shear ΔUmax, cold-pool buoyancies, theoretical gravity current propagation speed, and environmental MUCAPE for CTRL and NOFX, which are displayed in Fig. 22.
Citation: Journal of the Atmospheric Sciences 68, 12; 10.1175/JAS-D-11-0112.1
CTRL and NOFX always have both more convective downdrafts and more strong downdrafts than does NOLC (Figs. 18b,c), which is expected from the experimental design of eliminating latent cooling in NOLC. The more interesting difference prior to 1500 UTC is the existence of more strong updrafts in CTRL and NOFX than in NOLC (Figs. 18a,c).
A line averaged vertical cross section at 1300 UTC (Fig. 19a) taken through the developing NE-SW rainband in CTRL (Fig. 17a) lacks a prominent surface cold pool. Instead, the approximately 90° phase shift between the w and θυ extrema (Fig. 19a) and the flow through the system at low levels (Fig. 19b) are consistent with a gravity wave. Here, the thermal structure resembles that in earlier simulations (Schmidt and Cotton 1990; Schumacher and Johnson 2008; Schumacher 2009) where gravity waves become both amplified and quasi-stationary relative to their convective heat source in environments with vertical shear. The thermal structure also has some similarities to the simulations of Parker (2008) and French and Parker (2010) late in their cycle of nocturnal surface cooling. Vertical cross sections at this stage from NOFX are nearly indistinguishable from those of CTRL and are not presented in Fig. 19. Similar to CTRL, the same average vertical cross section from NOLC (Fig. 19c) lacks a significant surface cold pool in the convective region. However, the thermal perturbations in the gravity wave–like signature above the surface are smaller than in CTRL, which is consistent with weaker cooling (e.g., Haertel et al. 2001; Marsham et al. 2010) in this simulation that excludes temperature tendencies due to melting and evaporation.
The nocturnal LLJ is still evident during this period shortly after sunrise (cf. Fig. 10a) and its contribution to at least moderate values of low-level shear is seen immediately southeast of the deep convection in the plane of the CTRL (Fig. 19b) and NOLC (Fig. 19d) cross sections where ΔU ~ 10 m s−1 through a 2.5-km depth above the LLJ maximum. This vertical shear layer is collocated with a deep layer of high-θe air whose vertical displacement resulted from adiabatic mesoscale ascent (Figs. 7a and 8). The coincidence of these environmental features in the vicinity of the gravity wave–like perturbation is favorable for strong updrafts in CTRL and NOFX at this transition to a single leading convective rainband along the southern end of the MCS (Figs. 17a,b). Note also that the significant line-normal vertical shear exists through the entire depth of the troposphere at this stage (Figs. 19b,d), although some of this deeper shear could be locally enhanced by the MCS upper-level outflow. Schmidt and Cotton (1990) explain how deep vertical shear can enhance lifting associated with gravity waves and Coniglio et al. (2007) show that it is an important predictor for MCS persistence in the more general population of quasi-linear MCSs.
Unlike for NOLC, whose strong updraft coverage increases through 1500 UTC and remains relatively constant during the next 3 h, the number of strong updrafts in CTRL and NOFX continuously declines from 1500 to 1700 UTC (Fig. 18a). This weakening of the convective updrafts in CTRL and NOFX coincides with a change from a gravity wave–like θυ structure (Fig. 19a) to a surface-based cold pool acting as a gravity current (Figs. 20a,c). Concurrently the environmental LLJ weakens, leading to a substantial reduction in lower-tropospheric vertical shear (Figs. 20b,d) from that 3 h earlier (Fig. 19b). Horizontal vorticity imbalances (e.g., Rotunno et al. 1988) that arise primarily from the weakening vertical shear and, to a lesser degree, the strengthening cold pool contribute to the weaker convective updrafts and to intensifying mesoscale regions of ascending front-to-rear and descending rear-to-front flow (Figs. 20a–d). During this updraft and horizontal flow evolution, the simulated leading-edge convective precipitation becomes weaker and lighter precipitation broadens rearward (Fig. 21), as often seen in observed leading-line trailing stratiform squall lines (e.g., Houze et al. 1989, their Fig. 1).
Vertical cross sections of model-derived radar reflectivity (dBZ) and system-relative winds (solid contours with 1 m s−1 interval starting at 1 m s−1, negative values with −1 m s−1 interval starting at −1 m s−1) in CTRL averaged for 60 km in the direction normal to cross section orientation for a portion of the identically averaged cross sections at (a) 1300 (Figs. 19a,b) and (b) 1600 UTC 13 Jun 2002 (Figs. 20a,b), whose horizontal locations are shown in Figs. 17a and 17b, respectively. Note that the aspect ratios of these plots differ from those of the slightly longer vertical cross sections in Figs. 19a,b and 20a,b, which relate to (a) and (b), respectively.
Citation: Journal of the Atmospheric Sciences 68, 12; 10.1175/JAS-D-11-0112.1
The horizontal vorticity imbalance is reflected in the large values of C/ΔU > 1 in both CTRL and NOFX after the approximately 1400 UTC onset of a strong surface cold pool (Fig. 22a), which are characteristic of upshear-tilted squall lines with trailing stratiform precipitation. Here, ΔU is the maximum environmental vertical shear through the depth of the cold pool and C is the theoretical speed of the cold pool obtained by vertically integrating the buoyancy perturbation, 
(a) Time series of environmental CAPE for the most unstable parcel (thin curves) and C/ΔU (thick curves), where C is the theoretical gravity current speed and ΔU is the maximum vertical shear in a layer starting at 200 m AGL and extending through a nominal cold-pool depth of 3.5 km AGL for CTRL and NOFX. The computations are based on averages taken from the rectangles in Figs. 20a and 20c, as in Trier et al. (2006). (b) As in (a), but for the cold-pool buoyancy perturbation at 1400 and 1600 UTC.
Citation: Journal of the Atmospheric Sciences 68, 12; 10.1175/JAS-D-11-0112.1
The postsunrise (1300 UTC)-to-midmorning (1600 UTC) evolution of convective structure in CTRL is independent of near-surface warming in the MCS environment. This is because while surface warming in advance of the CTRL MCS results in much stronger surface buoyancy perturbations in CTRL than in NOFX, these differences never span much of the overall cold pool depth (Fig. 22b). Similarly, the environmental vertical shear differences between CTRL and NOFX are limited to the lowest few hundred meters (Figs. 20b,d). These factors explain the similar C/ΔU values (Fig. 22a) and related average upshear (rearward) vertical tilts of the convective updraft zones (Figs. 23a,b) in CTRL and NOFX, which are consistent with the development of similar ascending (descending) mesoscale front-to-rear (rear-to-front) flows that extend well beyond the convective zone (cf. Figs. 20a,c with Figs. 20b,d). Schumacher (2009) similarly noted development of significant upshear tilts to the convective region of a heavy-rain-producing MCS as organizing mechanism evolved from a diabatically produced gravity wave to a surface cold pool.
Line-averaged vertical velocity contours (thick gray; 1, 2, and 3 m s−1 values) and virtual potential temperature contours (thin black; 1.5-K contour interval) at 1600 UTC 13 Jun 2002 for the leading convective region of (a) CTRL in Figs. 20a,b and (b) NOFX in Figs. 20c,d. The thick dashed black lines highlight the convective updraft zone vertical tilts, which are more highly slanted than depicted due to the 1:2 aspect ratio of the plot.
Citation: Journal of the Atmospheric Sciences 68, 12; 10.1175/JAS-D-11-0112.1
Unlike in NOFX, the number of strong updrafts (Fig. 18a) and downdrafts (Fig. 18b) eventually increases in CTRL, which is likely related to increases in MUCAPE (Fig. 22a), as its convection becomes more surface based. In contrast, there is no late morning increase of MUCAPE in NOFX (Fig. 22a) since surface fluxes are withheld. The increase in the strength of convection in CTRL after 1630 UTC (Figs. 18a,b) is consistent with the observed stronger radar reflectivity (cf. Figs. 1c,d) but is somewhat less pronounced and long lasting. The observed MCS, which exhibits strong intensity at 1800 UTC (Fig. 1d), decays during the early afternoon (not shown).
In contrast to CTRL and NOFX, the NOLC MCS neither develops a strong cold pool nor experiences significant downward transport of rear-to-front momentum (e.g., Mahoney et al. 2009). Therefore the NOLC MCS advances at 9 m s−1 (Fig. 20e) compared to 15 m s−1 for CRTL (Fig. 20a). The lack of upshear MCS tilt (Figs. 20e,f) and the unrealistically slow advancement of NOLC are consistent with the notion that similar errors in some simulations that use cumulus schemes may arise through improper representation of downdrafts (Davis et al. 2003).
The much weaker cold pool in NOLC does not promote upshear tilt and the weakening of updrafts seen in CTRL and NOFX. Furthermore, the slower system speed in NOLC combined with the lack of stabilization in the MCS wake that arises from withholding latent cooling also provides the convection in NOLC (Fig. 20f) with longer access to the deep high-θe air (which resulted from earlier mesoscale lifting in the environment) than in CTRL (Fig. 20b) and NOFX (Fig. 20d). Together, these factors explain the greater number of strong updrafts in NOLC relative to these two other simulations by mid- to late morning (Fig. 18a).
7. Summary and discussion
This study uses a convection-permitting model to successfully simulate and then examine the postsunrise reorganization of an observed nocturnal MCS that occurred during IHOP_2002 on 13 June (Marsham et al. 2011). As such, these simulations offer a relatively rare view of the details by which MCS internal processes and the evolution of its environment can conspire to influence the life cycle of precipitation systems in a region where significant changes are commonly observed following sunrise (e.g., Hane et al. 2008).
Our modeling results illustrate the important role of mesoscale preconditioning of the environment. In the hours immediately before sunrise, a mesoscale band of upward motion coincides with warm advection by the nocturnal LLJ at the southern end of the confluent baroclinic zone associated with a slowly moving synoptic cold front. This vertical motion contributes to a 2-km-deep layer of enhanced water vapor immediately ahead of the evolving MCS.
As in the observations, both NW–SE-oriented convective rainbands and NE–SW-oriented rainbands coexist around sunrise as the MCS begins to reorganize. Trajectories through one of the more transient simulated NW–SE-oriented rainbands reveal that conditionally unstable air originating from near the top of the LLJ is gradually lifted to near saturation over a roughly 100-km zone before participating in deep convection near the northern terminus of the LLJ. Meanwhile the developing simulated NE–SW-oriented convective rainband that reforms the leading edge of the MCS becomes situated within the enhanced water vapor mixing ratio zone associated with the earlier mesoscale upward motion and is initially perpendicular to the local lower-tropospheric vertical shear.
During its early stages after sunrise the simulated NE–SW-oriented convective rainband lacks a significant surface cold pool. Trajectory analysis indicated a large majority of low-level air parcels rise and fall less than about 1 km as they flow through a gravity wave–like thermal perturbation and that the most intense deep updraft cores are fed exclusively by air parcels originating within the moist conditionally unstable layer located from 1 to 2.5 km AGL. Similar inflow branches were found by Schumacher (2009) in his analysis of trajectories encountering a lower-tropospheric gravity wave in idealized simulations of a heavy-rain-producing MCS.
By midmorning (1500–1600 UTC) a cold gravity current surface outflow develops due to penetrative convective downdrafts behind the NE–SW-oriented convective rainband at the leading edge of the simulated MCS. Consistent with observations (Marsham et al. 2011), ascending (descending) front-to-rear (rear-to-front) mesoscale flows had developed by this time. This evolution is consistent with gravity current–driven squall lines that tilt upshear because of horizontal vorticity imbalances between the negative buoyancy within the gravity current and the environmental low-level shear (Rotunno et al. 1988), which in this case had weakened dramatically because of the cessation of the nocturnal LLJ. The simulated MCS strength as measured by the number of strong convective updrafts (w > 10 m s−1) declines during this phase. However, it temporarily increases once again by late morning (1700 UTC) as the surface heating results in CAPE increases and the dominance of surface-based deep convection. Late morning intensification of the MCS is even more evident in the observations.
Sensitivity studies that separately withheld (i) latent cooling processes responsible for strong penetrative convective downdrafts within the MCS and (ii) surface fluxes responsible for daytime PBL warming and moistening in its environment were analyzed. The similar overall orientation and horizontal scale of the MCS in these experiments point to the primary roles of the mesoscale environmental vertical motion and the vertical shear in facilitating its reorganization.
Despite a broadly similar reflectivity pattern, the MCS vertical structure and horizontal motion were, however, strongly influenced by latent cooling processes. When these cooling processes were withheld, maximum vertical motions (both downdrafts and updrafts) within the MCS were initially weaker. Moreover, the MCS in this experiment never develops the observed surface gravity current outflow and a realistic upshear tilt, and it has a horizontal speed that is much slower than observed. Interestingly, the number of strong convective updrafts eventually exceeds that in the control run. This result is broadly consistent the idealized two- and three-dimensional modeling of Crook and Moncrieff (1988) and Schumacher (2009), respectively, which indicated that a significant surface cold pool was not necessary for the maintenance of organized deep convection in conditionally unstable environments with large-scale convergence.
Surface warming has little impact on the MCS through midmorning (1600 UTC), which we attribute to two main factors. First, the antecedent layer of strong stability that is removed by surface heating in the control run was relatively shallow. As a result, differences in both the environmental wind profiles and in the MCS-generated cold pool from the control and the simulation that withheld surface fluxes were restricted to shallow layers near the surface and therefore had little impact on the overall horizontal vorticity imbalance that influenced MCS vertical structure and evolution. Second, although midmorning convective updrafts in the control run comprise a greater mixture of elevated and surface-based air parcels than earlier, the maximum updraft cores are fed by elevated source parcels long after the cold pool develops and thereby ingest air with moist static energy similar to that in the simulation that withholds surface fluxes. Only by late morning, when the convection in the control run becomes primarily surface based, are differences in vertical motion strength between these two simulations significant.
The life cycle of both the observed and simulated MCS differed somewhat from that most typically observed over the southern Great Plains (e.g., Hane et al. 2008; Carbone and Tuttle 2008). Most notably, the MCS in our case study survives into the afternoon, although a significant weakening period occurs in the control simulation. Reasons for the difference in MCS longevity from climatology are a matter of speculation but may be related to particularly strong mesoscale forcing near sunrise in the current case and wake effects associated with an earlier MCS that may have delayed development of new nocturnal convection over the region. Detailed studies of additional cases that could complement existing statistically based studies (e.g., Gale et al. 2002; Coniglio et al. 2007; Hane et al. 2008) are desirable for this postsunrise transitional period when significant MCS evolution often occurs.
Acknowledgments
Sherrie Fredrick (NCAR) is acknowledged for her assistance in running the simulations. The authors thank George Bryan (NCAR) for his helpful internal review of the manuscript. We are also grateful to Russ Schumacher (Colorado State University), Matt Parker (North Carolina State University), and an anonymous referee for their constructive reviews. This research was performed as part of NCAR’s Short Term Explicit Prediction (STEP) Program, which is supported by National Science Foundation funds for the U.S. Weather Research Program (USWRP).
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This simulation is not strictly adiabatic since it retains subgrid mixing and radiative processes.