Numerical Simulation of the Interaction between the Dryline and ...

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ROBERT B. WILHELMSON. Department of ... LOUIS J. WICKER AND CONRAD L. ZIEGLER .... FIG. 1. Skew T–logp diagram derived from initial model fields at.
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Numerical Simulation of the Interaction between the Dryline and Horizontal Convective Rolls STEVEN E. PECKHAM Cooperative Institute for Research in Environmental Sciences, University of Colorado, and NOAA/Forecast Systems Laboratory, Boulder, Colorado

ROBERT B. WILHELMSON Department of Atmospheric Sciences, University of Illinois at Urbana–Champaign, Urbana, Illinois

LOUIS J. WICKER

AND

CONRAD L. ZIEGLER

National Severe Storms Laboratory, Norman, Oklahoma (Manuscript received 16 September 2003, in final form 19 December 2003) ABSTRACT The results of high-resolution simulations of an idealized dryline environment are discussed. The use of a single high-resolution domain, combined with accurate advection numerics and minimized numerical filtering, allows the explicit resolution of large horizontal convective roll (HCR) circulations and their daytime evolution. The horizontal convective rolls are oriented in the direction of the lower planetary boundary layer (PBL) wind shear. By midafternoon a north–south-oriented dryline develops near the center of the simulation domain with the PBL circulations from both sides intersecting the dryline at multiple locations. West of the dryline, the HCR bands evolve into open convective cell (OCC) structures having stronger and deeper vertical circulations compared to the OCCs and HCRs to the east. The OCCs and HCRs east of the dryline impact the dryline and convective cloud location by modulating the low-level moisture and upslope easterly flow. The interaction between OCC and HCR circulations and the dryline appears primarily responsible for creating a considerable amount of along-line variation in the dryline characteristics. Many shallow convective clouds develop along and west of the dryline over the OCC and HCR updrafts as well as OCC–dryline and HCR–dryline intersection points. The shallow convective clouds evolve into deep convective clouds where OCCs and HCRs to the east intersect the dryline near the same location. When the cumulus clouds move to the east of the dryline and remain over an OCC/HCR updraft, the persistent low-level lifting permits the convective updraft to overcome the cap east of the dryline and directly lift near-surface moisture to its level of free convection.

1. Introduction The dryline has long been known to be an important feature associated with the development of convection in the Great Plains. Rhea (1966) reported that convective storms formed within 200 n mi of the dryline approximately 70% of the time the dryline developed. There was also an indication of a tendency for convective storms to organize into mesoscale convective systems close to the dryline (e.g., Bluestein and Jain 1985). These systems account for a large fraction of the total rainfall during the spring and summer over most of the Great Plains (Fritsch et al. 1986). However, despite the acknowledged importance of the dryline, the processes Corresponding author address: Steven E. Peckham, NOAA/FSL, R/FSI, DSRC, 325 Broadway, Boulder, CO 80305. E-mail: [email protected]

q 2004 American Meteorological Society

that control dryline evolution and its relationship to convective storm formation are not fully understood (Hane et al. 1993). This study examines some of the processes that control the dryline evolution and how they lead to an environmental structure that might favor convective storm formation. Over the past decade, finescale observational studies have revealed that atmospheric boundaries (e.g., sea breezes, convergence zones, drylines) can possess a large amount of along-line variation (Wilson et al. 1992; Wakimoto and Atkins 1994; Atkins et al. 1995, 1998). These studies focused on the relationship between these along-line variations and boundary layer horizontal convective rolls (HCRs) (Etling and Brown 1993; Brown 1980). The HCRs were observed to be nearly periodic in nature and intersect the atmospheric boundaries at nearly right angles. At the intersection locations, undulations developed along the boundary with locally

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enhanced vertical velocities observed at HCR updraft intersection locations. Deep convective clouds developed near or within the region of enhanced ascent. In addition, HCRs were observed on the warmer, more convectively unstable side of the boundary, but on the cooler, more convectively stable side few, if any, HCRs were observed. Atkins et al. (1995) proposed that, for sea breezes, if any HCRs existed on the cool side of the boundary, they were most likely parallel to the boundary. However, dryline observations have shown that convective rolls can exist to the east, or on the cool side, of the boundary and that these rolls can be oriented across the boundary (Atkins et al. 1998). Observation studies have revealed that HCRs often evolve from longitudinal rolls into an open convective cell structure (Agee et al. 1973) during the afternoon over a variety of environmental conditions (Sykes and Henn 1989; Weckwerth et al. 1997, 1999; Kristovich et al. 1999; Cooper et al. 2000). Open convective cells (OCCs) are characterized by a region of weak descending motion that is surrounded by a narrow perimeter of ascending motion and, given sufficient moisture, clouds. The transition from HCRs to OCCs develops as the vertical wind shear within the convective boundary layer (CBL) weakens and the convective instability, or buoyancy of surface air parcels, increases. The observed OCC aspect ratio often ranges between 10 and 15, or about 2 to 3 times larger than the observed HCR aspect ratio. With the increase in computational resources over the past decade, researchers using numerical models are beginning to investigate the interaction between convective rolls and atmospheric boundaries. For example, Dailey and Fovell (1999) demonstrated that high-resolution (500 m) simulations of the sea breeze could reproduce many of the phenomena (e.g., sea-breeze boundary, HCRs) observed during the Convection and Precipitation/Electrification (CaPE) field program. These simulations also demonstrate how the interaction between HCRs and boundaries plays an important role in the formation and evolution of convective clouds along the sea breeze. Numerical studies of drylines interacting with boundary layer HCRs have been recently reported in the literature (e.g., Ziegler et al. 1997; Peckham 1999). These simulations employ a nested domain with horizontal grid spacing that is marginally capable of resolving HCRs (i.e., 1 km), but produce many of the observed phenomena (e.g., a dryline, HCRs, and deep moist convection). Interestingly, these simulations produce HCRs with aspect ratios much larger than those reported in boundary layer observational studies as well as the aspect ratios explained by classic linear theory. One possible explanation for the large aspect ratio is the nonlinear interaction between individual HCRs and/or gravity waves in the free troposphere. Another possible cause of the large aspect ratios comes from the use of too small of a domain in nested grid numerical simu-

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lations. There will be inadequate space in the domain and time for the generation of HCRs at scales resolvable by the nested grid. With the advances in computational power, it is now possible to conduct single-grid high-resolution simulations of the dryline environment to investigate daytime morphology of HCRs near a developing dryline, the HCR–dryline interactions, and subsequent convective cloud formation. Expanding upon previous results from coarse-horizontal-resolution simulations (Peckham and Wicker 2000; Peckham 1999), this investigation examines the formation of HCRs within the daytime dryline environment using single-grid, high-horizontal-resolution (1 km) numerical simulations. This configuration not only permits the HCRs to fully develop across the entire simulation domain, but it also allows HCRs to interact with the dryline. The focus of this study involves the • development of HCRs, OCCs, and the dryline in the numerical model • development of east–west undulations along the dryline and their possible generation by HCR–dryline and OCC–dryline interactions • development and motion of clouds near the dryline and their relationship to HCRs and OCCs. 2. Methodology The Collaborative Model for Multiscale Atmospheric Simulation (COMMAS) is used in this study (Peckham 1999; Wicker and Wilhelmson 1995). COMMAS includes a generalized terrain-following coordinate transformation (Gal-Chen and Sommerville 1975) as well as parameterizations for surface radiation (Benjamin 1983) and land surface processes (Deardorff 1972, 1978). A 1.5-order subgrid-scale turbulence parameterization (Deardorff 1980) is used with the vertical mixing length within the unstable boundary layer following Sun and Chang (1986). A series of 12-h simulations, starting at 1200 UTC (0700 LDT), are performed on a 600 km 3 60 km numerical grid having 1-km horizontal resolution and centered at 34.58N latitude and 1008W longitude. The domain extends vertically to 18 km with a vertical mesh interval smoothly stretching from 50 m at the lowest grid point to approximately 500 m at the domain top. Wave-radiation open lateral boundary conditions (Klemp and Wilhelmson 1978) are applied on the large time step (4 s) at the eastern and western domain edges. A numerical sponge is employed above 14 km MSL to eliminate wave reflection from the rigid domain top. Periodic boundary conditions are specified along the northern and southern boundaries, and the downward shortwave radiation flux is assumed to vary only in the x direction. The use of periodic lateral boundary conditions along the northern and southern boundaries limits the appli-

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TABLE 1. Elevation, virtual potential temperature ( u y ), and relative humidity (RH) values used to initialize the numerical model. Values used along the western boundary Elev (MSL) uy RH 1300 2000 5000 5500 6000 12 000 17 000

302.00 0.30 311.50 0.20 316.20 0.20 318.50 0.20 320.00 0.15 337.00 0.10 390.00 0.10 Values used along the eastern boundary Elev (MSL) uy RH 300 1100 1400 2000 3300 4500 5000 5500 6000 12 000 17 000

300.00 303.00 305.83 311.50 313.00 316.20 317.10 318.50 320.00 337.00 390.00

0.98 0.99 0.40 0.20 0.20 0.20 0.20 0.20 0.15 0.10 0.10

cability of the numerical simulations. In an environment with southerly flow, air parcels in the convective boundary layer are warmed as they are transported northward. As boundary layer air parcels reach the northern boundary, the southern boundary air temperature is reset to the temperature and moisture obtained along the northern boundary. This means that the meridional temperature gradient is not being preserved in the simulations—an assumption that is only valid for approximately one-half of the diurnal cycle. Rather than attempt a simulation with detailed terrain information, the terrain elevation is confined to changes in the x direction and is expressed analytically using a hyperbolic tangent function (Peckham and Wicker 2000). The use of an analytical expression permits the close approximation of the actual terrain without introducing the high-frequency variations associated with realistic topography. Representative profiles of virtual potential temperature and relative humidity at the lateral boundaries are constructed from the 1200 UTC soundings observed at Amarillo, Texas, and Norman, Oklahoma, on 15, 16, 26, and 30 May 1991 (Table 1). These dates coincide with days when convective clouds and storms developed along the dryline during the 1991 Cooperative Oklahoma Profiler Studies (Hane et al. 1993) field program. This numerical study uses profiles similar to those used by Peckham and Wicker (2000), but with additional data at 5.5 and 6 km MSL. The vertical profiles of temperature and humidity preserve the nocturnal inversion below 5 km MSL, the low-level moisture depth, and the elevated mixed layer. Above 5 km MSL the vertical temperature and humidity profiles were assumed to be horizontally homogeneous. The resulting profiles

FIG. 1. Skew T–logp diagram derived from initial model fields at the eastern (gray lines) and western (black lines) boundaries and y 5 30 km.

of temperature and absolute humidity are similar to typically observed atmospheric conditions in dryline environments (Fig. 1). The initial zonal wind has a vertical wind shear of 0.004 s 21 below 4.5 km MSL with the zonal flow equal to 24.8 m s 21 at 0 km MSL, resulting in no zonal wind (u 5 0) at 1200 m MSL. Above 4.5 km MSL the vertical wind shear is incrementally relaxed to zero at 6 km MSL (Fig. 2). The initial y wind component is assumed to be uniformly 6 m s 21 and in geostrophic balance. A lowlevel jet is initialized as an elliptical region of supergeostrophic southerly flow centered 500 m AGL at x 5 425 km with a horizontal radius of 200 km, a vertical radius of 500 m, and a maximum ageostrophic flow of 5 m s 21 .

FIG. 2. Vertical profile of the initial zonal wind component.

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The numerical model is initialized using a three-dimensional technique based upon the interpolation scheme of Ziegler et al. (1995) and Peckham and Wicker (2000). Vertical profiles of potential temperature, zonal momentum, and relative humidity are specified at the eastern and western boundary and are interpolated in the east–west direction with a hyperbolic tangent function using a boundary layer depth of 2 km AGL, a weighting function slope of 5 3 10 23 km 21 , and an inflection point at the domain center. The horizontal wind is assumed to be in geostrophic balance, but north– south variations in virtual potential temperature and momentum are not permitted. Hence, the initial thermodynamic and momentum fields vary in the east–west direction but are uniform in the north–south direction. Reasons for not restricting the initial fields to preserve thermal wind balance are that the simulation domain is subsynoptic scale; therefore, synoptic-scale balances (i.e., thermal wind balance) do not hold. Also, specifying lateral boundary conditions in which thermal wind balance is maintained over space and time (e.g., Richardson and Droegemeier 1996; Richardson 1999) would require restricting the boundary layer evolution and dryline morphology at the boundaries. In addition, any inaccuracy in the parameterization of the boundary layer evolution at the lateral boundaries would result in the development of spurious circulations. These circulations would subsequently be advected into the domain and contaminate the solution. Finally, the use of wave-radiation open boundary conditions along the boundaries restricts boundary layer convective roll orientation to be orthogonal to the boundaries. The use of periodic boundary conditions along the north and south boundaries eliminates the spurious roll orientation along these boundaries at the cost of excluding thermal wind balance. The large distance separating the east and west boundaries from the developing dryline along with the predominantly southerly low-level flow ensures that any possible contamination from the specified boundary conditions has little influence on the simulation results near the domain center. The initial fields (Figs. 3 and 4) depict a low-level inversion below 2 km MSL that increases in height above the surface toward the east while also providing a cap for the low-level moisture. Above the low-level inversion, a conditionally unstable layer extends to the tropopause at 12 km MSL with the lapse rate above the inversion and below 5 km MSL being slightly larger than dry adiabatic. These temperature, moisture, and wind profiles are typical of the dryline environment over the western Great Plains (Schaefer 1973). In addition, a broad, weak convergent flow collocated with a horizontal moisture gradient provides an environment favorable for dryline formation. The initial volumetric soil moisture (hereafter soil moisture) and fractional vegetation coverage (hereafter vegetation) are also specified analytically using the same hyperbolic tangent function as the topography (Fig. 5).

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The easternmost section of the domain is moist and predominantly covered in vegetation, whereas the western domain is initially relatively dry with sparse vegetation. At x 5 600 km (eastern boundary) the soil is assumed to be very moist and heavily vegetated and so the surface soil moisture is initially set at 0.3, the deep soil moisture at 0.4, and the vegetation at 0.7. At x 5 0 km (western boundary) the surface soil is assumed to be dry and sparsely vegetated while the deep soil is moist. Here the surface soil moisture is initially set at 0.075, the deep soil moisture at 0.3, and the vegetation at 0.3. The region of strongest horizontal change in soil moisture and vegetation is located 35 km west of the greatest slope in terrain at x 5 265 km. Further, random spatial fluctuations are applied to the initial soil moisture and vegetation fields to ensure that the HCR and OCC development is physically based and not a product of the computational algorithm (Fig. 5). The surface roughness, vegetation albedo, and soil type are assumed to be uniform and unvarying across the entire domain and are set at 0.1, 0.21, and sandy loam, respectively. 3. Experimental results a. Boundary layer development and morphology Relatively weak (vertical velocities less than 0.25 m s 21 ) counterrotating circulation bands of ascending/descending motion develop in the CBL by 3.5 h (1030 LDT) into the simulation (Fig. 6). These bands are long (10 to 100 km) and narrow (3 to 5 km wide) with convergent low-level flow and a thermally direct circulation. The circulations originate at two distinct regions, near the eastern domain boundary (not shown in Fig. 6) and along the terrain with the steepest slope at the domain’s center. These locations correspond to the locations of greatest surface heating during the first 3.5 h of the simulation—the easternmost region of the domain where the sun rises earliest and where the sloping terrain has the greatest direct shortwave insolation. The updraft bands are oriented in the direction of the lowlevel shear and are oriented at large angles to the local mean boundary layer shear. This result is similar to observed horizontal convective rolls that develop during the morning hours in an unstably stratified boundary layer and are oriented in the direction of the mean boundary layer flow and low-level shear (Weckwerth et al. 1999; Cooper et al. 2000). Therefore, the text refers to these thermally direct circulation bands in the convective boundary layer as horizontal convective rolls (HCRs) despite the inadequate horizontal resolution to completely resolve these features (Dailey and Fovell 1999). The HCR aspect ratio (ratio of distance between consecutive roll updrafts to boundary layer height) ranges from approximately 5 near the eastern domain boundary and just east of the domain center (x 5 330 km) to roughly 10 to the west of the domain center (x 5 270

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FIG. 3. (a) East–west cross section of water vapor mixing ratio (g kg 21 ) along y 5 30 km at t 5 0. Contour interval is 1 g kg 21 . The dark shading represents the topography used in the simulation. (b) Same as (a), but for potential temperature. Contour interval is 1 K. (c) Same as (a), but for the zonal wind component. Contour interval is 1 m s 21 . (d) Same as (a), but for the meridional wind component. Contour interval is 1 m s 21 .

km). The large variation in aspect ratio is due to the nearly horizontal CBL top at about 1.5 km MSL and the sloping terrain combining to make a relatively shallow CBL (roughly 500 m high) in the western domain half and a deeper CBL (about 900 m high) in the eastern domain. The aspect ratios of the rolls produced in the simulation are larger than the aspect ratios of 2 to 4 typically observed as well as predicted by theory (Brown 1980; Etling and Brown 1993; Asai 1970, 1972), possibly due in part to the model’s inability to fully resolve boundary layer circulations at the scale of the horizontal grid resolution. By 5 h into the simulation (noon LDT), the CBL height increases (figure shown later) as the surface warms, and long, narrow HCRs develop across the entire simulation domain (Fig. 7a). The CBL top increases to an elevation of roughly 700 m AGL near x 5 200 km to 1.3 km AGL near x 5 400 km. The HCR updraft

width of 3–5 km and separation distance of roughly 5– 6 km is unchanged from previous hours, but because of the higher CBL top, their aspect ratios decrease to a range between 4 and 7. Though slightly larger than the aspect ratios predicted by theory, the simulated HCR aspect ratios are approaching the theoretical range. The HCR updrafts intensify by 6 h (1300 LDT) for x , 250 km, while east of x 5 300 km the number of HCRs are reduced and an open cell structure is beginning to develop (Fig. 7b). Ascent within the convective rolls west of x 5 250 km has increased to be greater than 2 m s 21 (at elevations of roughly 1 km AGL) along sections of the HCR updraft bands. The CBL tops increase to elevations of 2.0 km AGL for x , 250 km and 1.5 km AGL for x . 250 km (figure shown later). The HCR aspect ratios now range from 2 to 5, in agreement with observations and theory. After 7 h of simulation time (1400 LDT), the bands

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FIG. 4. North–south cross section of the potential temperature along x 5 300 km at t 5 0. Contour interval is 1 K, and the dark shading represents the topography used in the simulation.

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FIG. 6. Horizontal slice at 12 600 s (3 h, 30 min) along vertical grid level 5 (about 300 m AGL) showing regions of ascending motion (black shading). Vectors show the north–south average directions of the wind shear (top) from the surface to the top of the boundary layer (z i ), (middle) from the surface to 0.2 z i , and (bottom) the mean boundary layer wind.

of convective rolls in the boundary layer have transformed to an OCC structure over most of the domain (Fig. 7c). Two parameters have been proposed by previous studies to distinguish between environments that favor horizontal rolls and open cell convection (Sykes and Henn 1989; Weckwerth et al. 1999). The first is

FIG. 5. (a) East–west plot of the initial soil moisture for the surface layer (black) and deep soil layer (gray). (b) East–west plot of the fractional vegetation coverage (vegetation) used for the control simulation.

FIG. 7. (a) Horizontal slice showing regions of ascending motion along vertical grid number 10 (about 600 m AGL) at 18 000 s (5 h) into the simulation. The light gray shading depicts regions of vertical motion between 0 and 2 m s 21 , and the dark gray shading those regions greater than 2 m s 21 . (b) Same as (a), but at 21 600 s (6 h). (c) Same as (a), but at 25 200 s (7 h).

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2z i /L, where 2z i is the CBL height and L is the Monin– Obukhov length. Although each of the previous studies found a study-specific range in 2z i /L that supported rolls, all of the studies reveal that bands of convective rolls are typically observed when 2z i /L , 10, or within slightly unstable convective boundary layers. As values of 2z i /L become greater than 10, indicating that convective instability is dominating wind shear, convective rolls become less organized and evolve to an open cell structure. The second parameter is u*/w*, where u* is the friction velocity and w* is the convective velocity (Sykes and Henn 1989). The friction velocity is defined as u* 5 [C d u] 2 , where C d is the drag coefficient and u is the mean horizontal wind. The convective velocity is defined as w* 5

1

gz i u9u 9 r cp u y

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rolls and waves aloft may become aligned, providing conditions that may support the interaction between HCRs and gravity waves in the overlying stable layer. In the numerical simulation of the dryline the vertical wind shear vector at the top of the CBL is oriented westsouthwest to east-northeast. Gravity waves in the overlying stable troposphere are oriented perpendicular to the shear vector, or north-northwest to south-southeast. The HCRs between x 5 250 km and x 5 300 km are oriented from the southwest to the northeast or skewed approximately 608 to the gravity wave fronts. These orientations most closely match the unidirectional shear case. Therefore, constructive interference between gravity waves and the convective rolls does not appear to be a factor in transformation of HCRs to OCCs.

1/3

,

where g is the gravitational constant, u9u9 is the vertical heat flux, r is the air density, and u y is the virtual potential temperature. Ratio values of u*/w* larger than 0.35 favor the development of HCRs, and as u*/w* decreases (increased convective instability), the HCR bands become less organized and evolve toward an open cell structure. Examination of the simulation results between 5 and 7 h (noon and 1400 LDT) reveals a trend in the simulated conditions from an environment supporting HCRs toward an environment favoring open cell convection. At 5 h, the range in boundary layer stability ratios are 5.9 to 3.2 for 2z i /L and 0.59 to 0.43 for u*/ w*. Over the next 2 h, 2z i /L increases to between 10.8 to 5.8, and u*/w* decreases to a range of 0.49 and 0.40. The ratios obtained from the control simulation do not match previous studies, but there is agreement between the simulation and the predicted trends. Higher values of 2z i /L and lower values of u*/w* are observed to correspond with unorganized convection. Conversely, organized bands of rolls tended to dominate at lower values of 2z i /L and higher values of u*/w*. Several numerical and observational studies have examined the interaction between HCRs and gravity waves in the overlying stable troposphere (LeMone 1990; Balaji and Clark 1988; Lane and Clark 2002). Gravity waves in a stable, vertically sheared troposphere above a convective boundary layer are essentially two-dimensional and oriented perpendicular to the shear vector above the mixed layer. As previously noted, convective rolls are often oriented in the direction of the CBL mean flow or low-level shear vector. Therefore, in the situation in which the environment contains unidirectional shear, boundary layer rolls would be oriented perpendicular to gravity waves in the stable troposphere aloft. Such configuration would prohibit interaction between the tropospheric gravity waves and the HCRs. In the situation in which the environment contains directional shear, such as an Ekman boundary layer, the convective

b. Boundary layer structure Figure 8 shows that at 5 h into the simulation, the HCR updrafts are narrow, shallow, and relatively weak with maximum vertical velocities less than 2 m s 21 . Most of the vertical motions east of x 5 270 km are greater than 0.5 m s 21 because of a higher capping inversion and greater shortwave insolation. Vertical transport of moisture, sensible heat, and momentum prior to this time has been dominated by the subgrid-turbulence parameterization (figure not shown). In the stable troposphere, bands of vertically propagating gravity waves are generated as the HCR circulation updrafts interact with the westerly flow and the overlying stable layer. By 6 h into the simulation, the HCR and OCC updrafts and downdrafts remain narrow, but their increased height and vertical motion permit greater vertical transport of heat, moisture, and momentum (Fig. 9). Lowlevel moisture is enhanced by about 1 to 2 g kg 21 below the HCR and OCC updrafts in the convergent flow. This result is consistent with previous findings (Weckwerth et al. 1996; Ziegler et al. 1995; Shaw et al. 1997). Conversely, low-level mixing ratio values decrease below the HCR/OCC downdrafts as dry air is transported down to the surface by the HCR/OCC downdrafts. In addition, HCR/OCC downdrafts transport higher momentum air toward the surface. This result is counter to the findings of Kristovich (1993), but the simulation does not have the maximum horizontal wind speed near the surface. As previously noted, the interaction between the HCR/ OCC circulations and the overlying free atmosphere generates gravity waves, which gives the plot of vertical velocity a vertical banded appearance. The top of the CBL increases from east to west with the CBL top located around 2 km MSL at x 5 300 km and 3 km MSL at x 5 200 km. A sloping transition region between the two boundary layers occurs between roughly x 5 240 km and x 5 260 km, or near the western edge of the initial soil moisture gradient at x 5 250 km. The HCR and OCC circulations have generated a well-mixed boundary layer by 7 h into the simulation (Fig. 10). Water vapor and virtual potential temperature

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FIG. 8. (a) Cross section of mixing ratio (g kg 21 ) and ascending motion (shading) along y 5 30 km at 18 000 s (5 h) into simulation. Contour interval is 1 g kg 21 . The light gray shading depicts ascending motion between 0 and 2 m s 21 . (b) Same as (a), but for virtual potential temperature. Contour interval is every 1 K. (c) Same as (a), but for the zonal momentum. The contour interval is every 1 m s 21 . (d) Same as (a), but for the meridional momentum. The contour interval is every 1 m s 21 .

tends to be enhanced along and at the base of the HCR/ OCC updrafts because of the convergent, local boundary layer winds. An east–west horizontal gradient in mixing ratio and virtual potential temperature has developed at this time as the boundary layer west of x 5 240 km has grown to a height of roughly 4 km MSL compared to 2 km MSL east of x 5 260 km. Further, differential boundary layer growth and hydrostatic pressure falls result in a lower surface pressure west of x 5 250 km and developing a thermally direct solenoid circulation. The low-level southwesterly flow east of x 5 250 km is accelerated toward the west, turning the CBL winds southeasterly and up the sloping terrain. Meanwhile, downward transport of strong westerly momentum to the west of x 5 250 km has produced west-southwesterly flow, which enhances the convergent low-level flow and the frontogenetic development of a dryline.

c. Dryline development Figure 11a shows the low-level moisture and vertical motions near the center of the domain at 7 h 15 min. OCCs exist west of x 5 250 km, and HCR bands, aligned southwest to northeast, exist east of x 5 250 km. Several locations of east–west horizontal moisture gradients exist at this time; two notable ones are near x 5 240 km and another along x 5 252 km. The horizontal moisture gradient near x 5 240 km is located along the eastern edge of the western convective boundary layer, where the downward transport of dry air has reduced the surface-level moisture. The moisture gradient near x 5 252 km, which will later become the dryline, is located along a north–south zone of ascending motion making the boundary between the western OCCs and the eastern HCRs.

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FIG. 9. (a) Cross section of mixing ratio (g kg 21 ) and ascending motion (shading) along y 5 30 km at 21 600 s (6 h) into simulation. Contour interval is 2 g kg 21 . The light and dark gray shading depicts ascending motion between 0 and 2 m s 21 and greater than 2 m s 21 , respectively. (b) Same as (a), but for virtual potential temperature. Contour interval is every 1 K. (c) Same as (a), but for the zonal momentum. The contour interval is every 2 m s 21 . (d) Same as (a), but for the meridional momentum. The contour interval is every 2 m s 21 .

Fifteen minutes later the structure of the convective rolls remains the same with OCCs west of x 5 252 km and bands of convective rolls to the east of x 5 252 km (Fig. 11b). However, over the 15 min the rolls on either side of x 5 252 km have been transported inward and northward toward x 5 252 km in the convergent, southerly flow. The western CBL horizontal moisture gradient previously located along x 5 240 km has been advected eastward to roughly x 5 247 km while the horizontal moisture gradient developing in the convergent flow along x 5 252 km has intensified. Sections of the latter horizontal moisture gradient (e.g., between y 5 0 km to y 5 15 km along x 5 252 km) exceed 0.5 g kg 21 km 21 , indicating that a segment of the dryline has developed. However, the strong horizontal moisture gradient is not a continuous north–south boundary along x 5 252 km. A strong moisture gradient is located be-

tween y 5 25 km and y 5 35 km at x 5 252 km. This strong gradient weakens north of y 5 35 km (y . 35 km), and the location of strongest horizontal moisture gradient jumps westward to x 5 247 km, producing a double-dryline structure (Hane et al. 1993; Ziegler et al. 1997). The developing dryline segment along x 5 252 km is also associated with a north–south zone of ascending motion. The magnitude of the ascending motion in the updraft band is similar to the HCR updrafts near the boundary. However, the dryline updraft has stronger and deeper vertical motions locally at the HCR–dryline and OCC–dryline intersection locations. The enhancement of the dryline updraft at HCR and OCC intersections is consistent with previous observational and numerical modeling studies (Atkins et al. 1995; Wilson et al. 1992; Dailey and Fovell 1999; Ziegler et al. 1997; Hane et al.

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FIG. 10. (a) Cross section of mixing ratio (g kg 21 ) and ascending motion (shading) along y 5 30 km at 25 200 s (7 h) into simulation. Contour interval is 2 g kg 21. The light and dark gray shading depicts ascending motion between 0 and 2 m s 21 and greater than 2 m s 21 , respectively. (b) Same as (a), but for virtual potential temperature. Contour interval is every 1 K. (c) Same as (a), but for the zonal momentum. The contour interval is every 2 m s 21 . (d) Same as (a), but for the meridional momentum. The contour interval is every 2 m s 21.

1997; Peckham 1999) that found that the intersection of HCR updrafts and environmental boundary circulations (e.g., sea breezes and drylines) leads to stronger ascending motions at the intersection locations. Over the next half hour, the two dryline segments combine into a single continuous north–south-oriented dryline in the convergent flow. For the rest of the afternoon the dryline does not uniformly advance eastward of x 5 252 km (Figs. 11c and 11d), but instead, there are eastward advancement and westward retreats produced by along-line undulations. The along-line undulations are located approximately every 10 to 15 km along the dryline and move northward along the dryline at roughly 3 m s 21 . The along-line propagation speed of the undulations is the same as the mean southerly flow along the dryline boundary (figure not shown). In addition, the undulations appear to be coincident with

the intersection locations of OCCs and HCRs. The dryline tends to bow westward at the intersection of western OCC updrafts bands and eastward at the intersection of western OCC downdraft bands. The HCRs east of the dryline appear to impact the development of along-line undulations by locally enhancing the low-level moisture and easterly flow (see Fig. 10c). Previous studies have provided evidence that gravity waves in the free troposphere can be associated with the development of along-line undulations (e.g., McCarthy and Koch 1982; Koch and McCarthy 1982). In these previous studies, free-tropospheric gravity waves are produced in an environment having weak stability and strong wind shear. In the current study, the gravity waves are instead generated by OCCs and HCRs intruding upon the stable free troposphere (i.e., Lane and Clark 2002; Hauf and Clark 1989; Balaji and Clark

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FIG. 11. (a) Surface mixing ratio (g kg 21 ) and ascending motion at vertical grid level 10 (about 600 m AGL) (shading) at 26 100 s (7 h, 15 min) into simulation. Ascent less than (greater than) 2 m s 21 is shaded light (dark) gray. (b) Same as (a), but at 27 000 s (7 h, 30 min). (c) Same as (a), but at 27 900 s (7 h, 45 min). (d) Same as (a), but at 28 800 s (8 h).

1988). The gravity waves in the free troposphere are oriented north-northwest to south-southeast, or perpendicular to the west-southwesterly wind shear vector (roughly 2608). The wavelength and period of the gravity wave are estimated by gridpoint analysis and Fourier decomposition of the vertical wind to be roughly 6 km and 1500 s (25 min), respectively. This gives a groundrelative phase speed of approximately 4 m s 21 and a southerly phase speed of 0.7 m s 21 . Hence, if the alongline undulations were generated by free-troposphere gravity waves they would move to the north at roughly 0.7 m s 21 , or much slower than what is produced in the simulation. Therefore, the along-line undulations are most likely being transported northward by the low-

level flow and are not directly associated with freetroposphere waves. The premise that a westward undulation is located at a western HCR–OCC updraft intersection does not appear to match the late-afternoon dryline structure. This is due to the ‘‘memory’’ of previous roll–dryline intersections that produce the eastward or westward undulations. For example, at x 5 252 km and y 5 5 km (Fig. 11d), an eastward undulation exists at a western OCC downdraft–dryline intersection, where 15 min prior the eastward undulation was located at a western OCC downdraft (Fig. 11c). Over the 15-min period, this OCC is advected northeastward, but the cell structure of the boundary layer convection results in the apparent south-

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ward propagation of the OCC updraft. The resulting structure is the intersection of an OCC updraft at an eastward dryline undulation [i.e., the OCC updraft at x 5 248 km and y 5 10 km (Fig. 11c) is located at x 5 254 km and y 5 14 km (Fig. 11d)]. The development of along-boundary undulations is in close agreement with Atkins et al. (1998), who hypothesized that the interaction between the dryline and HCRs west of the dryline creates considerable alongline variation. Further, the along-line undulations are similar to the undulations observed along the sea breeze and are reproduced in numerical simulations (Atkins et al. 1995; Dailey and Fovell 1999). The latter study identified two mechanisms that appear to explain the formation of the along-line undulations in the sea breeze as well as the dryline. The first is the interaction between the HCR circulations and the boundary updraft. Along the descending branch of the HCRs the strong westerly flow transports the dryline boundary farther eastward. Conversely, along the ascending branch of the HCRs, the dryline is transported westward by the stronger easterly winds within the eastern CBL. The along-line differences in wind speed combine to create the east–west undulations. The second is the convergence of moisture under the ascending branches of the HCRs. The alternating regions of increased (decreased) moisture within the CBL produce the cross-boundary shift in the sea breeze. Similar mechanisms appear to be responsible for the generation of east–west undulations along the numerically simulated dryline. However, unlike the numerical simulations of sea breezes, the dryline has HCRs located on the eastern side of the boundary that appear to modulate the along-line undulations via enhanced low-level easterly flow and low-level moisture convergence (Figs. 11c and 11d). In addition, the numerically simulated sea breeze had uniformly spaced along-line undulations because of the regular, longitudinal roll structure of the simulated HCRs. In the dryline simulation, the random cell structure of the western OCCs causes the distances separating the along-line dryline undulations to range from 10 to 20 km, resulting in an irregular undulation pattern. d. Cloud development Shallow convective cloud streets develop above the HCRs starting around 4 h into the simulation. The clouds are initially located well east of the developing dryline (figure not shown). After 6 h into the simulation, these shallow convective clouds still exist well east of the domain center. In addition, around this time, as the western CBL height increases to over 4 km MSL, a few convective clouds develop in the western domain half at the top of the western CBL. These forced convective clouds are located over the HCR updrafts and tend to be relatively shallow with bases around 4.25 to 4.75 km MSL and cloud tops penetrating the stable troposphere

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above the CBL to about 5.25 km MSL. The location of the convective cloud development and the need to examine the convective available potential energy (CAPE) and convective instability (CIN; Colby 1984) along the HCR updrafts is consistent with Weckwerth (2000). Around 7.25 h into the simulation, numerous forced convective clouds have developed along and within 10 km west of the developing dryline (Fig. 12a). The convective cloud bases west of the developing dryline are 3.5 to 3.75 km MSL, which is at the top of the western convective boundary layer (Fig. 13a). Surface-based CAPE values (approximately 500 J kg 21 ) are, given the dry, high-shear environment, marginal for deep convective cloud development (Fig. 14a). East of x 5 250 km bands of enhanced CAPE are collocated with the HCR updrafts, where the convergent low-level flow has pooled the low-level moisture under the HCR updrafts. This result is consistent with previous observations of clouds in a convective boundary layer (Weckwerth 2000). The forced convection continues to evolve as individual cloud elements develop and dissipate while propagating toward the east-northeast over the next 15 min. By 7.5 h, a line of convective clouds has developed over the dryline location near x 5 252 km (Fig. 12b). Tracking the evolution and development of the convective clouds reveals that most of the shallow clouds have developed over the past 10 min and the clouds in the previous figure have dissipated (figure not shown). As before, the clouds west of the dryline (x , 250 km) are forced and their location corresponds to the western OCC updrafts. Along the dryline, the convection tends to be located at the OCC–dryline and/or HCR–dryline intersection points. These locations along the dryline correspond to locally enhanced and deeper updrafts. Backward trajectories are computed for a 125 km 3 volume containing the convective cloud at 7 h, 30 min for a 30-min period using 1-min model wind updates. Analysis of the trajectories reveals that the moisture supplying the forced convective clouds originates west of the developing dryline (Figs. 12a and 12b). Surface-based CAPE values along the dryline where clouds exist range from 1000 to 1500 J kg 21 and elsewhere along the dryline CAPE ranges from 500 to 1000 J kg 21 . Thus, the OCCs and the HCRs increase locally the low-level moisture available for cloud development along the dryline at the intersection points of OCC–HCR updraft bands (Fig. 14b). By 7.75 h, the numerous shallow convective clouds along the dryline have evolved into fewer, but larger individual convective elements (Fig. 12c). Most of the convective clouds along the dryline remain shallow and are forced by the merged OCC–HCR and dryline circulations. However, the convective cloud near x 5 253 km, y 5 38 km (Fig. 12c) has become active (Stull 1985) and is developing into a deep convective cloud. Tracking this convective cloud reveals that it existed 15 min prior to the west-southwest of x 5 248 km, y 5

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FIG. 12. (a) Outline of cloud locations (heavy solid lines) and ascending motion at vertical grid level 10 (about 600 m AGL) (shading) at 26 100 s (7 h, 15 min) into simulation. Cloud locations are where the total cloud water mixing ratio in the grid column is greater than zero. Ascending motion greater than 2 m s 21 is shaded dark gray and ascent between 2 m s 21 and zero is shaded light gray. (b) Same as (a), but at 27 000 s (7 h, 30 min). The arrow near the grid center shows the horizontal trajectory of air producing the cloud. (c) Same as (a), but at 27 900 s (7 h, 45 min). (d) Same as (a), but at 28 800 s (8 h). The arrow shows the horizontal trajectory of air producing the cloud.

34 km (Fig. 12b). Despite being active, the cloud bases for the convection remain between 3.5 and 3.75 km MSL, indicating that moisture east of the dryline has not yet been lifted into the active convection (Fig. 13c). Surface-based CAPE values along the dryline where clouds exist continue to range from 1000 to 1500 J kg 21 (Fig. 14c). However, CAPE values at the location of the example showing active convection are in excess of 1500 J kg 21 , suggesting that the enhanced moisture along an HCR east of the dryline is beginning to be drawn into the cloud. By 8 h, the once numerous convective clouds that existed along the dryline have been reduced to only a

few large convective elements (Fig. 12d). Of these isolated clouds, only the convection located near x 5 255 km, y 5 41 km is active and is rapidly developing into a deep convective cloud. Examination of the cloud tracks beyond 8 h reveals that the remaining convective elements along the dryline dissipate while the CBL circulations generate additional convective clouds. The developing deep convective cloud is located east of the surface dryline location, and while it moves to the eastnortheast, it remains located over a HCR updraft where the surface-based CAPE is enhanced (Fig. 14d). The convective clouds that dissipate move northeastward into regions of the eastern CBL that contains descending

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FIG. 13. (a) North–south cross section along x 5 246 km depicting areas of cloud water mixing ratio greater than zero (heavy solid lines), virtual potential temperature, and velocity vectors at 26 100 s (7 h, 15 min) into the simulation. Virtual potential temperature contours are every 1 K. A velocity vector scale (m s 21 ) is provided in the lower-left-hand corner. (b) Same as (a), but the cross section is made at x 5 248 km and at 27 000 s (7 h, 30 min) into the simulation. The air parcel trajectory is shown by the heavy solid arrow. (c) Same as (a), but for cross section at x 5 253 km and at 27 900 s (7 h, 45 min) into the simulation. (d) Same as (b), but for cross section at x 5 255 km and at 28 800 s (8 h) into the simulation.

branches of HCR circulations and lower surface-based CAPE. A vertical slice along x 5 255 reveals that the storm cloud base has lowered by approximately 500 m over the past 15 min to around 3.25 km MSL, which is near the top of the convective boundary layer east of the dryline. In addition, trajectory analysis shows that the strong vertical motion below cloud base is transporting low-level moisture from the east of the dryline up into the deep convective cloud (Figs. 12d and 13d). Therefore, convective clouds originating near the dryline need to remain over a region of low-level ascent (i.e., the ascending branch of an HCR, or dryline circulation updraft) for the convection to become active and develop into deep convection. Otherwise, the descending motion within the CBL will cause a reduction in the low-level moisture and stabilize the environmen-

tal profile causing the convection to dissipate. This is consistent with previous observational and numerical studies (Wilson and Megenhardt 1997; Ziegler and Rasmussen 1998; Hane et al. 1997). 4. Discussion and conclusions The horizontal grid spacing and the north–south domain size used in this study is similar to that used for the inner grid of Ziegler et al. (1997). That is, this study uses a horizontal resolution that is better than most, but not all, previous dryline simulations. However, the combination of periodic lateral boundary conditions along the northern and southern domain boundaries and the extensive domain size appears to allow the OCC- and HCR-scale turbulence to fully develop. So, given the

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FIG. 14. (a) Outline of cloud locations (heavy solid lines) and CAPE (J kg 21 ) (shading) at 26 100 s (7 h, 15 min) into the simulation. CAPE is shaded according to the bottom legend. (b) Same as (a), but at 27 000 s (7 h, 30 min). The arrow points to the active cumulus cloud discussed in the text. (c) Same as (b), but at 27 900 s (7 h, 45 min). (d) Same as (b), but at 28 800 s (8 h).

realistic development of OCCs and HCRs and the subsequent dryline development, what does the simulation(s) reveal that we do not already know from previous studies? The simulation affirms the hypothesis of Atkins et al. (1998) that along-line variations can be products of HCR–dryline interactions. Analogous to the sea breeze (Atkins et al. 1995; Dailey and Fovell 1999), the dryline undulations in the simulations appear to be primarily controlled by the stronger, deeper western OCC circulations. However, HCRs to the east of the dryline can also influence the development of along-line undulations by locally enhancing the low-level easterly flow. In addition, the updrafts were stronger and deeper at the OCC–dryline intersections than elsewhere along the

dryline boundary. Also similar to cloud development at the sea-breeze boundary, active convective clouds develop at the OCC–dryline intersection points. The simulation also shows that the eastern rolls have a significant impact upon the development of cumulus convection by modulating the low-level easterly flow and moisture east of the dryline. Enhanced CAPE values are located along the eastern HCR updrafts where the low-level moisture is pooled. This result is applicable to situations in which there is initially insufficient moisture (CAPE) at the dryline to overcome the environmental convective stability (CIN). The eastern HCRs produce locally enhanced low-level moisture (CAPE), decrease the parcel stability (CIN), and enhance updrafts along the dryline. Deep moist convection develops

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where eastern HCRs and western OCCs intersect the dryline circulation, resulting in a juxtaposition of enhanced low-level moisture and a deep strong updraft. Convection originating at the dryline would need to remain over the ascending branch of an HCR as it moves to the east not only due to the persistent lifting, but also to ensure sufficient low-level moisture for the convection to become active and develop into deep convection. Similar to Ziegler et al. (1997), the low-level air parcels are lifted to their LCL and LFC by the boundary layer circulations. There appears to be little impact by the stable tropospheric motions (e.g., gravity waves) on the formation and evolution of moist convection. The numerical simulation reveals that convective clouds develop over a region of persistent low-level ascending motion (i.e., an OCC or HCR updraft, or their intersection with the dryline). Ziegler et al. (1997) demonstrated that convection develops along a mesoscale updraft band and subsequently moves along the same mesoscale updraft band, maintaining low-level ascent and supporting further convective development. The dryline simulation produces deep moist convection originally west of the dryline over an HCR updraft that becomes active as the convection crosses over the dryline. The convection then continues to grow and intensify into a deep convective cloud east of the dryline as it taps lowlevel moisture from the eastern CBL. Although the convective cloud development process is different from Ziegler et al. (1997), the persistent lifting of low-level moisture under the forced convection that eventually results in deep convective cloud development is similar. The need for deep persistent lift at the dryline to produce deep convection is discussed by Ziegler and Rasmussen (1998), who noted that moist ‘‘boundary layer air parcels must be lifted to their LCL and LFC before leaving the mesoscale updraft to form deep convection.’’ The manner in which deep convective cloud development takes place in the dryline simulation is similar to what is discussed by Ziegler and Hane (1993). They noted that the eastward advection of moisture that was vertically transported by the dryline circulation might aid the development of convective clouds. Ziegler and Hane (1993) state, ‘‘The primary significance of the elevated moist layer is that on other days with convection, clouds and storms could feed from this moisture source before the stage of development where withincloud pressure forces are sufficiently strong to lift air directly from the surface layer and eliminate the capping inversion.’’ In the numerical simulation, forced convective clouds develop over OCCs west of and also along the dryline at OCC–dryline intersections. The convection is fueled by the moisture being lifted along and to the west of the dryline. After passing to the east of the dryline, the dryline circulation updraft continues to provide moisture to fuel the convection until the convective updraft is able to lift the low-level moisture directly from the surface to cloud base.

FIG. A1. The vertical profile of the initial zonal wind component used in the control (dashed line) and the increased zonal flow profile (solid line) used in the sensitivity experiment.

Acknowledgments. This work was part of the first author’s postdoctoral fellowship at the Department of Atmospheric Sciences at the University of Illinois at Urbana–Champaign and was supported by NSF Grants ATM-99866672 and -9633228. Computing support was provided by the National Center for Supercomputing Applications. The authors thank Dr. Steven Koch, Mr. Brent Shaw, and Ms. Nita Fullerton for the technical and editorial comments that improved the manuscript. In addition, the authors thank the three anonymous reviewers that provided numerous constructive comments and greatly improved the manuscript. APPENDIX Sensitivity Tests A series of sensitivity tests are conducted in which a single simulation parameter used to initialize the numerical model is changed to determine its impact on the simulation results. With the exception of a single changed parameter, the configuration and initialization of the sensitivity test simulations are identical to the control simulation. The changes discussed here include a doubling of the north–south domain size to 120 from 60 km, an increase and decrease in the zonal wind magnitude by 2 m s 21 (Fig. A1), and an increase and a decrease in the east–west soil moisture gradient (Fig. A2). Numerical simulations by Peckham and Wicker (2000) found that an increase in the initial zonal wind speed produced large changes in the intensity and height of the dryline circulation. Simulations with a strong westerly flow increase downslope flow and produce a stronger capping inversion and a broader and weaker mesoscale updraft along the dryline boundary, conditions that are believed to be unfavorable to the development of deep moist convection.

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FIG. A2. East–west plot of the initial soil moisture for the surface layer (black) and deep soil layer (gray) for the strong soil moisture gradient (solid lines) and weak soil moisture gradient (dashed lines) sensitivity simulations.

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Each of the sensitivity test simulations exhibit HCR development during the morning hours that is similar to the control (not shown). The HCRs are characterized as long narrow bands of counterrotating thermally direct circulations that are oriented in the direction of the mean low-level flow. Low-level moisture is enhanced under the HCR updrafts in the convergent flow (Figs. A3a, A4a, A5a, and A6a). By 7 h into the simulations, the HCRs evolve to an OCC structure, except between x 5 250 km and x 5 300 km, where the HCRs’ structure remains as longitudinal bands (Figs. A3b, A4b, A5b, and A6b). Also, the OCCs west of the dryline in the increased zonal wind case show more of a longitudinal band appearance compared to the control and other sensitivity simulations. Examination of the sensitivity simulation results between 5 and 7 h (noon and 1400 LDT) also reveals a trend in the simulated conditions from an environment supporting HCRs toward an environment favoring open cell convection (Table A1). During the early afternoon the val-

FIG. A3. Plot of simulation data from the expanded north–south domain simulation. (a) Surface mixing ratio (g kg 21 ), cloud location outline (bold lines), and ascending motion at vertical grid level 10 (about 600 m AGL) (shading) at 25 200 s (7 h) into the simulation. Ascent less than (greater than) 2 m s 21 is shaded light (dark) gray. (b) Same as (a), but at 28 800 s (8 h).

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FIG. A4. Plot of simulation data from the enhanced zonal wind simulation. (a) Surface mixing ratio (g kg 21 ), cloud locations (bold lines), and ascending motion at vertical grid level 10 (about 600 m AGL) (shading) at 25 200 s (7 h) into the simulation. Ascent less than (greater than) 2 m s 21 is shaded light (dark) gray. (b) Same as (a), but at 28 800 s (8 h).

ues for u*/w* decrease from a range between 0.65 and 0.40 to a range of 0.54 to 0.39. Conversely, the range in 2zi/L values increases during the afternoon. At 5 h into the simulation, 2zi/L is between 6.3 and 2.2, and by 7 h it ranges between 10.3 and 4.7. Again, while the ratios obtained from the sensitivity simulations do not match previous studies, there is agreement between the simulation results and the predicted trends. Higher 2zi/L and lower u*/w* ratios are observed to correspond with unorganized convection, but organized bands of rolls tended

to dominate at lower 2zi/L and higher u*/w* values. Here again, the increased zonal wind sensitivity case differs slightly from the other sensitivity cases. This sensitivity case has less variation in the boundary layer parameters because of an enhanced boundary layer wind shear (since u 5 0 at the surface, the increased wind speed aloft produces a stronger low-level shear). Hence, the 2zi/L ratio does not increase (and likewise the u*/w* ratio does not decrease) as much for the increased zonal flow simulation as they do for the control simulation.

FIG. A5. Plot of simulation data from the strong soil moisture gradient simulation. (a) Surface mixing ratio (g kg 21 ), cloud locations (bold lines), and ascending motion at vertical grid level 10 (about 600 m AGL) (shading) at 25 200 s (7 h) into the simulation. Ascent less than (greater than) 2 m s 21 is shaded light (dark) gray. (b) Same as (a), but at 28 800 s (8 h).

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FIG. A6. Plot of simulation data from the weak soil moisture gradient simulation. (a) Surface mixing ratio (g kg 21 ), cloud locations (bold lines), and ascending motion at vertical grid level 10 (about 600 m AGL) (shading) at 25 200 s (7 h) into the simulation. Ascent less than (greater than) 2 m s 21 is shaded light (dark) gray. (b) Same as (a), but at 28 800 s (8 h).

With the stronger soil moisture gradient, fewer HCRs exist east of the dryline than in the control simulation. This difference is possibly due to the stronger dryline circulation and the smaller Bowen ratio, leading to a lower capping inversion height. With the weak soil moisture gradient, random HCR structure is present on both sides of boundary. Dryline development appears delayed because of the weaker east–west horizontal pressure gradient—related to hydrostatic pressure falls and a weaker horizontal moisture gradient. It appears that the soil moisture strongly impacts the CBL development and HCR aspect ratio. A dryline develops in each of the sensitivity tests by 8 h. Each dryline exhibits east–west undulations along the boundary (Figs. A3b, A4b, A5b, and A6b). Although the magnitude of the cross-dryline horizontal moisture gradient varies slightly for each test case, it is the largest

in the strong soil moisture test case and smallest in the weak soil moisture test. Along-line east–west undulations in the dryline boundary are manifested at OCC– dryline intersection locations and appear to be produced primarily by the stronger, deeper western OCC circulations. In addition, the updrafts are stronger and deeper at the OCC–dryline intersections than elsewhere along the dryline boundary. The test simulations reveal that convective clouds tend to develop over a region of persistent low-level ascending motion (e.g., an OCC updraft or OCC–dryline intersection location). To the west of the dryline the convective cloud bases are relatively high (about 4 km MSL) and the CBL is well mixed, indicating that the low-level moisture source for the clouds is from lower levels west of the dryline. Deep convective clouds that develop along and east of the dryline do so at locations

TABLE A1. Range of values for friction velocity (u*) to convective velocity (w*), and boundary layer height (zi) to Monin–Obukhov length (L), for each sensitivity case simulation at 5 h (18 000 s), 6 h (21 600 s), and 7 h (25 200 s) into the simulation. Range of values for friction velocity (u*) to convective velocity (w*) ratio Time (s) Case

18 000

21 600

25 200

Increased N–S domain size 0.63–0.41 0.55–0.45 Increased zonal wind 0.65–0.48 0.63–0.41 Strong soil moisture gradient 0.61–0.40 0.53–0.40 Weak soil moisture gradient 0.63–0.43 0.56–0.39 Range of values for boundary layer height (z i ) to Monin–Obukhov length (L) ratio Case Time (s)

0.48–0.40 0.54–0.53 0.49–0.40 0.49–0.39

Increased N–S domain size Increased zonal wind Strong soil moisture gradient Weak soil moisture gradient

10.3–5.6 6.0–4.7 10.2–5.7 9.6–5.9

5.8–2.8 4.1–2.2 6.3–2.9 5.2–2.8

6.1–4.3 8.7–2.6 8.6–4.4 9.0–4.2

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