Shear-Parallel Mesoscale Convective Systems in a ... - AMS Journals

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Shear-Parallel Mesoscale Convective Systems in a Moist Low-Inhibition Mei-Yu Front Environment CHANGHAI LIU AND MITCHELL W. MONCRIEFF National Center for Atmospheric Research, Boulder, Colorado (Manuscript received 18 April 2017, in final form 24 September 2017) ABSTRACT Numerical simulations are performed to investigate organized convection observed in the Asian summer monsoon and documented as a category of mesoscale convective systems (MCSs) over the U.S. continent during the warm season. In an idealized low-inhibition and unidirectional shear environment of the mei-yu moisture front, the structure of the simulated organized convection is distinct from that occurring in the classical quasi-two-dimensional, shear-perpendicular, and trailing stratiform (TS) MCS. Consisting of four airflow branches, a three-dimensional, eastward-propagating, downshear-tilted, shear-parallel MCS builds upshear by initiating new convection at its upstream end. The weak cold pool in the low-inhibition environment negligibly affects convection initiation, whereas convectively generated gravity waves are vital. Upstream-propagating gravity waves form a saturated or near-saturated moist tongue, and downstreampropagating waves control the initiation and growth of convection within a preexisting cloud layer. A sensitivity experiment wherein the weak cold pool is removed entirely intensifies the MCS and its interaction with the environment. The horizontal scale, rainfall rate, convective momentum transport, and transverse circulation are about double the respective value in the control simulation. The positive sign of the convective momentum transport contrasts with the negative sign for an eastward-propagating TS MCS. The structure of the simulated convective systems resembles shear-parallel organization in the intertropical convergence zone (ITCZ).

1. Introduction In vertically sheared environments, precipitating deep convection tends to organize into distinct forms of mesoscale convective systems (MCSs) such as the springtime squall lines in Oklahoma (Bluestein and Jain 1985), the three linear morphologies in the central United States (Parker and Johnson 2000), and the trailing line with adjoining stratiform (TL–AS) category (Schumacher and Johnson 2005). Gallus et al. (2008) and Jirak et al. (2003) classified warm-season systems in the central United States into 9 and 17 categories, respectively. A comparable variety of MCS occurs during the Asian summer monsoon (mei-yu) period, such as the 7 and 8 morphologies identified by Zheng et al. (2013) and Wang et al. (2014), respectively. The ubiquitous trailing stratiform (TS) MCS and squall line, characterized by a leading convective line

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Corresponding author: Dr. Changhai Liu, [email protected]

and a trailing stratiform region, has received extensive observational, numerical, and dynamical attention. Crucial to this category of mesoscale organization is the lower-tropospheric wind shear (e.g., Thorpe et al. 1982) and its effects on the evaporation-generated cold pools that outflow as density currents (e.g., Rotunno et al. 1988). Other categories of MCS have received less attention although nonlinear dynamical models represent key properties such as propagation and vertical tilt associated with organized convection in sheared environments (Moncrieff 1981). For MCS categories that feature a forward-directed stratiform region, modeling studies by Parker and Johnson (2004a,b) indicated that a strong mid- to upper-troposphere wind enables the downshear advection of condensate to generate weak leading precipitation. Nevertheless, the low-level shear and the cold pool play critical roles as in the squall-type mesoscale systems. Although the occurrence of convective lines with flow-parallel stratiform precipitation is ascribed to the strong along-line flow and shear that facilitate both back-building and along-line hydrometeor transports (Parker 2007a), the near-surface cold outflow

DOI: 10.1175/JAS-D-17-0121.1 Ó 2017 American Meteorological Society. For information regarding reuse of this content and general copyright information, consult the AMS Copyright Policy (www.ametsoc.org/PUBSReuseLicenses).

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is considered a necessary feature. In particular, the nearsurface cold pool significantly affects heavy-rain-producing TL–AS MCS (Peters and Schumacher 2015). Overall, cold pool–shear interaction is a key feature of a large fraction of MCSs observed around the globe. For MCSs in nearly saturated environments, however, the beneficial effects of cold pools are limited by decreased evaporative cooling and a weak density current. As identified by the simulations of Crook and Moncrieff (1988), the effects of mesoscale forcing, gravity wave, and other mesoscale and larger-scale processes are indispensable in that situation. The Schumacher (2009) simulation of a quasi-stationary heavily raining MCS showed that a near-stationary convectively generated low-level gravity wave is responsible for the flow organization. Studies of quasi-stationary convective systems in North America showed that mesoscale lifting and gravity wave forcing are more influential than the cold pool for parts of the MCS lifetime (e.g., Schumacher and Johnson 2008; Schumacher 2009, 2015). The role of convectively generated gravity waves has been addressed by Fovell et al. (2006) in a squall-line context with attention to ‘‘action at a distance’’ convection initiation mechanisms. Lane and Moncrieff (2015) and Moncrieff and Lane (2015) demonstrated the secondary importance of cold pools for upshear- and downshear-propagating mesoscale systems, respectively, for near-saturated tropical environments. We simulate a shear-parallel MCS in a near-saturated environment with minimal convective inhibition. New cells are initiated upstream of the system that has a weak downstream stratiform region. The precipitation pattern resembles MCSs with a flowparallel stratiform region and a back-building structure. This regime of mesoscale organization is often observed in the Asian monsoon period and classified as either back-building/quasi-stationary or parallel stratiform MCSs (e.g., Zheng et al. 2013; Wang et al. 2014). It features significant along-line wind and vertical shear (Wang et al. 2014) and large values of precipitable water (Zheng et al. 2013). Compared to the various categories of MCS in North America, the shear-parallel MCSs in the Asian monsoon period have been little studied despite their frequent occurrence and association with severe weather and extreme precipitation. We investigate the generation and organization of this regime of convective organization in an idealized environment representing the low–convective inhibition conditions in the mei-yu front. This paper is presented as follows. The next section describes the setup and design of the two numerical experiments. Section 3 describes the results of the numerical experiments including the key roles of the weak

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cold pool and convectively generated gravity waves. Section 4 offers concluding remarks.

2. Numerical model and experiment design An idealized modeling framework is utilized to simulate and understand fundamental processes involved in the development of shear-parallel MCS in the mei-yu front partly as a basis for future real-world cases. The WRF Model (Skamarock et al. 2008) is configured with 2401 3 1801 3 101 grid points and horizontal and vertical grid sizes of 0.5 and 0.2 km, respectively. The lateral boundary conditions are periodic in the x direction and open in the y direction. The west–east periodicity has little to no influence on our systems since they are confined to the central part of a large computational domain. Subgrid mixing is treated with a 1.5-order turbulence kinetic energy closure, and the cloud microphysical processes are parameterized with a singlemoment 6-class scheme (Hong and Lim 2006). Tests showed that the simulated convective systems are insensitive to the choice of microphysics parameterization. An explicit sixth-order numerical diffusion scheme is used to control grid-scale numerical noise (Knievel et al. 2007). The initial conditions are derived from a thermodynamic sounding previously used to simulate a heavy-rainfall-producing rainband observed in the 2004 mei-yu season, which is known as baiu in Japan (Yamasaki 2009). Specifically, the initial temperature field is based on a single profile (Fig. 1a) plus a perturbation centered at 1-km altitude with a horizontal (vertical) radius of 100 (1) km of maximum intensity 1 K, given as u(x, y, z) 5 Du cos2 [(p/2) 3 RAD],

RAD # 1;

(1)

where RAD5f[ ( x 2 xc )/100]2 1 [( y 2 yc )/100 ]2 1 (z 2 zc )2 g1/2 , (x, y) 5 (0, 0) is the domain center, xc 5 2100 km, yc 5 0 km, zc 5 1 km, and Du 5 1 K. The initial specific humidity field is defined by the following analytic function: n h io qy (x, y, z) 5 Qy (z) 0:85 1 0:15 exp 2(y=100)2 ,

(2)

where Qy(z) approximates the vertical sounding profile (Fig. 1a) along the central moisture front (y 5 0). The second term inside the braces, a perturbation symmetric to the x axis, idealizes the quasi-stationary mei-yu and similar moisture fronts in other geographical regions. This water vapor distribution defines a 200-kmwide moisture front, corresponding to a surface virtual temperature gradient of about 0.3 K (100 km)21, which

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FIG. 1. (a) Temperature and dewpoint soundings (Yamasaki 2009). (b) Initial wind profile.

is too weak to generate an influential solenoidal circulation. The moisture gradient is not affected by the environmental wind, which is parallel to the specific humidity contours. An important climatological feature of the mei-yu front is its meridionally asymmetric water vapor distribution with significantly higher humidity to the

south. The moisture pattern in Eq. (2) is a useful idealization of the mei-yu front. Note that Yamasaki (2009) used the same temperature sounding but a more complex moisture field to simulate a west-northwest– east-southeast-oriented (nonshear parallel) MCS during the baiu season.

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Two numerical experiments were conducted, namely, the control experiment (CTRL) and a sensitivity experiment. CTRL uses the above thermodynamic conditions and the idealized wind profile of strong low-level westerly shear as shown in Fig. 1b and constant westerly wind above 10 km resembling Yamasaki (2009). It should be pointed out that the wind profile can be physically interpreted as the along-line wind component of the observed MCS, rather than the conventional east– west wind. The sensitivity experiment quantifies the omission of cold pools by suppressing rainwater evaporation cooling in the cloud microphysics parameterizations in order to prevent their development.

3. Simulation results a. Control experiment Precipitating convection develops after about 2 h of integration in response to the warm-bubble perturbation, grows upscale, and subsequently travels eastward, steered by the strong upper-level westerly wind. Figure 2 shows the spatial distribution of surface precipitation over a small portion of the model domain encompassing most of the convective region during 7–12 h. At 7 h, the convection has evolved into an MCS that in the early stage is quasi stationary but subsequently propagates slowly upstream at about 3 m s21 (left panels in Fig. 2). The precipitation pattern, about 100–200 km long, is centered on the moisture front (y 5 0) and parallel to the westerly wind and features heavy convective precipitation upstream and weak stratiform precipitation downstream. The MCS is maintained by convective cells that are initiated continually at its upstream end. The elongated structure of the surface precipitation pattern at the late stage resembles the back-building and parallel stratiform MCS seen in radar observations (e.g., Parker 2007a,b; Zheng et al. 2013; Wang et al. 2014). The quasi steadiness and the upstream evolution involving new cell generation indicate the upstreambuilding character of the convective system. (Herein, ‘‘upstream building’’ refers to cell regeneration at the upstream end, which can be either back building or forward building with respect to the propagation direction of the storm.) Although weak upper-level ascent plays a part, as shown later, the stratiform region is generated by the downstream advection of hydrometeors detrained from the mature convection by the strong mid- and upper-tropospheric westerly wind consistent with Parker (2007a,b). A time sequence of the instantaneous condensate field in Fig. 3 illustrates the cell regeneration near the

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upstream (leading) edge of the convective band that matures after about 7 h of simulation. About 14 cells, alphabetically denoted A–N, develop mainly ahead of the old cells, although cells are also triggered on the flanks. The westerly wind steers the cells downstream where they mature and merge into the MCS, which evolves via a process of consecutive cell formation ahead of the leading edge resembling the TS MCS (e.g., Fovell and Tan 1998). The cross sections of temporally and spatially averaged physical fields in Fig. 4 show the mesoscale kinematics and thermodynamics. The total condensate (Fig. 4a) is part of a downshear-tilted MCS consisting of shallow, intermediate, and deep convection together with a stratiform region. The cloud-top heights in the deep convective region exceed 12 km. The strengths of the low-level mesoscale downdrafts in the stratiform region vary between 20.1 and 20.5 m s21 (Fig. 4b). A distinction from the classical TS MCS is the total absence of a rear-to-front mesoscale circulation underlying the stratiform region (e.g., Zipser 1977; Houze et al. 1989). Consistent with the initial thermodynamic state, the near-saturated environmental airflow in the lowest 2 km uplifts upon approaching and entering the system and exits rearward. The lower troposphere is less moist beneath the stratiform region (Fig. 4b). The positive temperature deviation in the convective and the upper-level stratiform regions is due to enhanced latent heating (Fig. 4c). Cooling in the stratiform region near the 4-km altitude is due to the melting of graupel and snow, and near-surface cooling is due to rainwater evaporation. Weak adiabatic ascent causes minor low-level cooling ahead of the system. While the temperature perturbations otherwise resemble the TS MCS, major disparities are the weakness and shallowness of the cold pool and the weak low-level cooling beneath the stratiform region due to depleted evaporation in the near-saturated environment. The mesolow behind the leading edge (Fig. 4d), a hydrostatic response to the deep warming (Fig. 4c), replaces the typical mesohigh, which is consistent with the absence of a surface gust front (Fig. 4a).

b. Role of the weak cold pool The above description implies the inapplicability of conventional cold pool–low-level shear theory not only for the development and sustenance of the MCS but also the density current theory for its propagation. This key point is validated by a simulation with rainwater evaporation cooling turned off in order to exclude evaporation-driven cold pools, an approach similar to Peters and Schumacher (2016). The system should

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FIG. 2. Snapshots of surface precipitation rate (mm h21) and wind vectors over a subdomain at 7–12 h for (left) CTRL and (right) simulation without rainwater evaporation cooling. The corresponding integration time is displayed at the top of each panel. Wind vectors are based on u and y, and the reference vector (m s21) is shown in the upper-right corner.

feature stronger convection and heavier precipitation because the absent evaporative cooling cannot partly compensate the latent heating. This does occur because the right panels in Fig. 2 have heavier precipitation

along the central moisture front than CTRL. The propagation mechanisms of the two simulations are similar: downstream advection by the westerly flow in the development stage, followed by a quasi-stationary

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FIG. 3. Snapshots of 2-km condensate (g kg21) and wind vectors between 8 h 5 min and 9 h 10 min for a subdomain in CTRL. Cell generations near the leading edge are alphabetically labeled. The corresponding integration time is displayed at the top of each panel. Wind vectors are based on u and y, and the reference vector (m s21) is shown in the upper-right corner.

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FIG. 4. Cross sections for (a)–(d) CTRL and (e)–(h) simulation without rainwater evaporation cooling. Fields are spatially averaged over 5 km across the moisture front and temporally averaged between 9 h 30 min and 10 h 30 min; (a),(e) condensate (shading; g kg21) and u-wind perturbations (contours; 2 m s21 interval); (b),(f) relative humidity (contours) and w (shading; m s21); (c),(g) potential temperature perturbations (8C); and (d),(h) pressure perturbations (Pa). The relative humidity is calculated with respect to a mixed phase between ice and water between 08 and 2208C and ice below 2208C. Wind vectors (m s21) are based on u and w.

stage, and then slow upstream propagation in the mature stage. In the development stage, the system is narrower than CTRL and the mature stage extends much farther (about 400 km) downwind along with

heavier precipitation. (Note the change of distance scale in the right panels of Fig. 2.) Also, except for the upstream leading region, convective cells are generated in situ along the downstream moisture front, in

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association with an enhanced transverse circulation and moisture convergence from the flanks (discussed later). Weak stratiform precipitation is confined to the lateral region of the convective band instead of trailing the leading convection as in CTRL. Both simulations produce a quasi-stationary, shear-parallel, elongated convective system or rainband centered on the moisture front. The along-line cross sections (right panels in Fig. 4) show that the convective cloud tops extend above 14 km and the updrafts and horizontal wind perturbations are much stronger. Tropospheric warming is intense and deep, but the near-ground cold perturbations are negligible in accordance with the absence of rainwater evaporative cooling. The weak cold perturbations near the 08C level stem from the enhanced melting associated with the increased snow and graupel mass. Hydrostatic balance causes stronger and deeper low-pressure perturbations at the low-level leading edge than CTRL. Similar to CTRL, the mesoscale system grows and persists via the forward successive cell initiations. In summary, the formation, maintenance, and persistence of the shear-parallel MCS are fully independent of near-surface cold pool mechanisms.

c. Key role of convectively generated gravity waves It has been long recognized that cumulonimbus clouds are building blocks of MCSs. The development and sustenance of an MCS depend on the successive formation of discrete new convective cells at the leading edge. In the traditional conceptual model of a leading line–trailing stratiform MCS, transient cells are continually initiated by gust-front uplift due to localized low-level convergence between the propagating density current and the near-surface environmental wind. Once initiated at the leading edge, these cells move rearward relative to the gust front, grow to maximum size as cumulonimbus, and then enter the decaying stage. Hydrometeors feed the stratiform region while the convectively driven horizontal pressure gradient initiates and drives the mesoscale rear inflow (e.g., Lafore and Moncrieff 1989). An instantaneous cross section features a cluster of convective clouds at different stages of their life cycle. In contrast, in our simulations, the weak and shallow surface-based cold pool (if it indeed exists) excludes cell regeneration mechanisms essential to the conventional MCS model and the rear inflow. Instead, two categories of gravity waves are responsible for the continual upstream initiation of convection cells. These aspects are reminiscent of Nicholls et al. (1991), Mapes (1993), and Fovell et al. (2006), albeit with significant mechanistic differences as follows.

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The first category of gravity wave corresponds to an upstream-propagating deep mode, excited around 3 h into the simulation when intense convection flares up. The vertical cross sections along the central moisture front in Fig. 5 show the structure of vertical velocity and potential temperature perturbations relative to the initial state at two different stages of the wave activity. At first glance, there is a broad descent, starting from the surface and extending upward and upstream throughout the troposphere ahead of the leading edge of the system at 4 h (Fig. 5a). This descent attenuates upon moving upstream from the convective area and eventually gives way to an ascending branch 1.5 h later (Fig. 5b). However, careful inspection shows that this complex structure is a manifestation of multiple convolved lowfrequency gravity waves in response to convectively generated heating. The leading fast-moving deep subsidence may reflect the gravest mode with a vertical wavelength roughly twice the depth of convective heating. The slower trailing updraft–downdraft couplet in the vertical (i.e., located approximately at x 5 275 km at 4 h and between 2130 and 2100 km and between 250 and 230 km at 5.5 h) is indicative of the second-gravest mode with a vertical wavelength comparable to the convective storm (Nicholls et al. 1991; Mapes 1993; Liu and Moncrieff 2004). The downshear vertical tilt of the MCS is a direct consequence of the organizing effect of environmental wind shear on moist mesoscale dynamics. Of particular interest is the second vertical mode of gravity wave positioned between 250 and 230 km in Fig. 5b. When this disturbance propagates upstream, the associated lower-tropospheric ascent destabilizes the inflow environment, leading to negative potential temperature anomalies at low levels and a shallow cloud deck (see the thick black outline for cloud boundaries in Fig. 5b) embedded within a near-saturated layer ahead of the system (see the thick green line for a relative humidity of 98% in Fig. 5b). These ascending perturbations propagate forward against the background westerly flow and gradually attenuate. After 8 h, the wave signature is minimal but a destabilized environment of a cold moist tongue and an embedded cloud deck is left behind (Fig. 5b), which is favorable for subsequent upstream convection initiation. Note that the gravity wave– induced uplift is a mere few centimeters per second, so a nearly saturated boundary layer and low convective inhibition in the environment (Fig. 1a) seem essential for an effective gravity wave mechanism for the cloud-deck formation. It should be pointed out that the low-level cold anomalies far ahead of the convective area (i.e., near the left boundary) are the remnants of upliftinduced negative temperature anomaly in response to the initial warm-bubble perturbation.

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FIG. 5. Cross sections of vertical velocity (contours), potential temperature perturbation (shading with contours larger than 0.68C being omitted), and wind vectors at (a) 4 and (b) 5.5 h for CTRL. Contour interval is 1 cm s21, and solid (dashed) lines represent upward (downward) motions with thick lines being 0 cm s21. The thick green and black lines in (b) display the relative humidity contour at 98% and cloud boundary, respectively. Fields are spatially averaged over 10 km across the moisture front (y 5 0). Wind vectors (m s21) are based on u and w.

The gravity wave lifting is further illustrated by the trajectories of air parcels originated from a rectangular area centered at (x, y) 5 (250 km, 0 km) at 250 m and 3.5 h (Fig. 6). These parcels, with an initial LCL of 300–400 m and LFC of 400–500 m, ascend slowly and steadily when being advected toward the system and entering the ascending part of the wave at about 4 h. Gravity wave lifting enables a significant number of parcels to reach above their LCL and LFC in advance of the system, although the gustiness caused by the preceding convection could also play a role. Fovell et al. (2006) reported a similar low-frequency gravity wave destabilization that generated a cool moist tongue in the 2–4-km layer ahead of a simulated nocturnal squall line.

The second category is the slow downstream-moving gravity waves directly associated with the intermittent convection regeneration within the preestablished saturated layer by the preceding upstream-propagating waves ahead of the leading edge. These waves are identifiable after the above-discussed low-frequency waves move away from the storm, and their earlier activity roughly between 5.5 and 7 h is more evident some kilometers away from the central moisture front as demonstrated in the condensate and vertical velocity fields at 6 h (Fig. 7). The instantaneous 1-km vertical velocity field at three different times in Fig. 8 exemplifies their occurrence when the storm has evolved into a well-defined elongated shear-parallel morphology (i.e., after 7 h), showing transient wavelike

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FIG. 6. Trajectories of 200 parcels released from a 5 km 3 10 km rectangular area centered at (x, y) 5 (250 km, 0 km) at the 250-m level and 3.5 h, showing (a) altitude changes in the west–east direction and (b) altitude changes with time.

spatially alternating upward–downward motion patterns confined to the nearby central moisture front. Of note is the locally relatively strong updraft cores along the ascending bands, which manifest newly initiated

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convective cells. The wave signatures have a horizontal wavelength of 3–5 km, and their gravity wave attributes are illustrated by the phase-difference quadrature between transient vertical velocity and potential temperature perturbations (defined as the deviation from the initial state) above cloud tops (Fig. 9a). These waves propagate downstream at a Doppler-shifted phase speed of about 5 m s21, significantly slower than the environmental westerly wind (Fig. 9b). They do have an intrinsic upstream phase speed but are advected downstream by the background flow, which is consistent with the slightly upstream-sloped phase lines. The wave excitation concurs with the occurrence of shallow convective clouds in the boundary layer, with each wave updraft being coupled with a cloud cell. In other words, the convective cells arise as gravity waves. As a result of the phase-locked wave and cloud development, the perturbation potential temperature maxima are collocated with the updrafts because of latent heating as explained by Lin et al. (1998). This in-phase temperature and vertical velocity distributions roughly below 1 km, along with the existence of a critical level (i.e., U 2 c 5 0) around z 5 0.7 km (Fig. 9b), meet the conditions for a growing mode according to Lin et al. critical-level argument. The wave development may be understood within the context of critical-level dynamics proposed by Lin et al. (1998). Being nearly vertically oriented, rapidly decaying with height, and confined to the lowest few kilometers

FIG. 7. Snapshot of (a) 1-km-level condensate (g kg21) and (b) vertical velocity (cm s21) over a subdomain near the leading edge at 6 h for CTRL. Wind vectors (m s21) are based on u and y. Dashed ellipses in (a) highlight gravity wave features.

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FIG. 8. Snapshot of 1-km-level vertical velocity (cm s21) and wind vectors near the leading edge at (a) 8.5, (b) 9.5, and (c) 10 h for CTRL. Wind vectors (m s21) are based on u and y.

upon traveling horizontally for significant distances suggest ducted waves. To confirm this speculation, we calculate the Scorer parameter l2 by the classical formula l2 5 N 2 /(U 2 c)2 2 Uzz /(U 2 c) ,

(3)

where N is the Brunt–Väisälä frequency, U is the background horizontal wind, Uzz is the second derivative of U with height, and c is the phase speed. Figure 10 presents the spatial pattern of the Scorer parameter along the upstream moisture front. The parameter maximizes below 1.5 km, decreases rapidly

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FIG. 9. Cross section of (a) vertical velocity (contours) and potential temperature perturbation (shading) and (b) x-axis wind component (m s21) along the moisture front over a subdomain at 10 h for CTRL. Contour interval in (a) is 1 cm s21, and solid (dashed) lines represent upward (downward) motions, with thick lines being 0 cm s21. The green lines illustrate cloud boundaries. Wind vectors (m s21) are based on u and w.

above, and attains a minimum in the vicinity of the 8.5-km altitude, with little horizontal variations. This vertical structure is favorable for wave trapping or ducting. Above 4 km, l2 is less than 1 3 1026 m22 and thus capable of trapping the gravity waves displayed in Figs. 8 and 9; that is, l2 , k2, where k is the horizontal wavenumber (Lane and Moncrieff 2015). In summary, the upstream convective cell regeneration is attributed to the joint action of two categories of gravity wave. The preceding upstreampropagating wave destabilization promotes a moist layer and an embedded shallow cloud deck, while the ensuing downstream-propagating waves are concurrent with the episodic triggering of boundary layer cumulus clouds within the preexisting saturated layer.

This upstream-building process resembles the Fovell et al. (2006) action-at-a-distance mechanism for convection initiation involving interaction between lowfrequency and ducted high-frequency gravity waves. However, there are interesting disparities between our simulated shear-parallel system and the TS MCS simulated by Fovell et al. (2006). The first disparity concerns the distinct role of gravity waves in storm evolution and persistence. In Fovell et al. (2006), the new cells in the established storm are triggered by cold pool lifting. While gravity waves trigger cells far ahead of the gust front, and occasionally form a new storm, they do not benefit and may harm the parent storm. In contrast, in our simulations, the weak cold pool is incapable of triggering new cells. Instead, the two

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FIG. 10. Cross section over a subdomain ahead of the leading edge at 10 h for CTRL, showing the Scorer parameter (1026 m22) averaged along the y axis 10 km across the moisture front. Wind vectors (m s21) are based on u and w.

categories of gravity wave regulate cell regeneration near the leading edge and are essential for storm sustenance. A second disparity from Fovell et al. (2006) is that the gravity waves originate from the parent

storm, move upstream, and initiate new convective cells upon progressing through the forward environment, while high-frequency waves in our simulation are excited upstream in attendance with boundary layer

FIG. 11. Cross sections of line-normal wind and wind vectors averaged between 9.5 and 10.5 h for (a) CTRL and (b) simulation without rainwater evaporation cooling. Fields are spatially averaged from 50 to 150 (250) km along the x axis for CTRL (simulation without rainwater evaporation cooling). Contour interval is 1 m s21, and thick curves represent 0 m s21. Wind vectors (m s21) are based on y and w.

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cumulus cloud development and have rearward Doppler-shifted phase speeds. Finally, in our simulations, the wave-ducting mechanism is attributed to the strength and vertical curvature of the background flow, while in Fovell et al. (2006), it is associated with the upper-tropospheric forward outflow from the MCS.

4. Concluding remarks As far as we know, this is the first simulation of a poorly understood category of MCS consisting of quasistationary, shear-parallel convective bands in a moist, low–convective inhibition environment in strong unidirectional vertical shear. Observed examples include shear-parallel MCSs along the quasi-stationary mei-yu front in the Asian summer monsoon and over the central United States during the warm season. New convection cells are initiated near the upstream edge (either back building or forward building depending on the direction of propagation) and travel downshear and mature into a longlasting, shear-parallel, downshear-tilted MCS. The negligible contribution of the cold pool identifies an important disparity from the archetypal shear-perpendicular model of MCS (Moncrieff 1992). Our MCS differs substantially from the MCSs in the Fovell et al. (2006) simulations, which had substantial evaporation-driven density currents. In place of the dominant role of cold pool dynamics in the classical shear-perpendicular MCS, our convective cell regeneration mechanism evokes gravity wave dynamics. In particular, two distinct modes of gravity wave regulate continual upstream convective initiation in the lowinhibition environment. The deep, upstream-propagating mode, triggered by the initial strong convection, destabilizes the upstream environment and results in a shallow saturated or near-saturated layer that is favorable for subsequent convection initiation. This is consistent with the findings of Nicholls et al. (1991) and Mapes (1993). The ensuing short-wavelength ducted mode advects slowly downstream and is directly linked to the intermittent initiation of convective cells near the upstream edge of the convective band. The sensitivity experiment with the rainwater evaporation deactivated confirms the minimal role of cold pool dynamics in CTRL. More active convection centers on the moisture front, which leads to a much stronger, more elongated, and quasi-steady shear-parallel MCS. Moreover, as well as the continual upstream cell regeneration in the neighborhood of the leading edge that occurs in the control simulation, intense convection occurs downstream of the leading edge because of the enhanced transverse (line normal) circulation and moisture convergence from both flanks of the system (Fig. 11). Unlike CTRL, the stratiform precipitation develops at the sides

FIG. 12. Vertical profiles of line-parallel momentum transport (m2 s22) per unit mass averaged along the moisture front and between 9.5 and 10.5 h for CTRL (black curve) and simulation without rainwater evaporation cooling (blue curve). The spatial average is from 0 to 125 (200) km along the x axis for CTRL (simulation without rainwater evaporation cooling).

of the active convection rather than trailing the upstream convection. Enhanced convective heating caused by the deactivated evaporative cooling increases the moisture convergence and intensifies both the transverse circulation and the MCS. The along-line vertical transport of uwind (line parallel) momentum (Fig. 12) is about double that in CTRL and significantly deeper. The positive momentum transport (Fig. 12), due to the downshear tilt of the airflow, contrasts with the negative value by an eastward-propagating TS MCS. This underscores the difference between the two categories of organization. The mean-flow acceleration [westward (eastward) in the lower (upper) part of the convective layer], namely, the negative of the vertical gradient of momentum transport, reduces the low-level vertical shear. The trajectory analysis for air parcels originating ahead of the system (Figs. 13a–d) confirms the basic three-dimensional character of the MCS. Figure 13e shows the four branches of system-relative flow, namely, a jump updraft, a jump downdraft, throughflow, and up–downflow. The rear-to-front mesoscale downdraft that characterizes the classical two-dimensional MCS is entirely absent. The unidirectional pattern of airflow shows that the MCS propagates in a wavelike manner. The downshear-tilted jump updrafts transport boundary layer parcels to the 4–12-km layer. A proportion of the near-surface parcels initially ascend

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FIG. 13. (a)–(d) Trajectories of 200 parcels released from an upstream area (the boxed region in Fig. 2) at 6 h and different altitudes for CTRL. The trajectories are integrated for 6 h under the massless and frictionless assumption. (e)–(g) Schematic of the key features of the simulated MCS. (e) Side view of airflow in the MCS consisting of four interacting branches: a jump updraft (red), a jump downdraft (blue), an up–downdraft (purple), and throughflow (yellow). Thickness of the trajectories is proportional to the strength of the airflow. (f) End view of the MCS and the transverse circulation (green). (g) Plan view of the elongated structure involving convection cells initiated at the upshear end of the band and traveling downshear to sustain the MCS.

1–3 km then return downstream to the lower levels in the ‘‘up–down’’ branch. Other parcels below 6 km descend in the jump downdraft to the lower troposphere, and most parcels above 6 km traverse the system in the throughflow branch with small vertical displacement. In conclusion, the three-dimensional spatial structure, wavelike propagation, and downshear tilt of the airflow trajectory patterns in our simulations resemble those in Moncrieff and Lane (2015). The meridional extent of the moisture front is limited by the confluence of the convectively generated transverse circulation, and the ‘‘carrot shape’’ is a consequence of the increased horizontal scale of an MCS during its mature stage (Figs. 13f,g). Finally, the upstream-building and shear-parallel evolution of the MCS and the development of a transverse circulation resemble the Dudhia and Moncrieff (1987) and Khouider and Moncrieff (2015) simulations of shear-parallel rainbands embedded in the ITCZ.

Acknowledgments. We are grateful for support from the NASA Grant NNX13AO39G: diagnostic analysis and cloud-system modeling of organized tropical convection in the YOTC-ECMWF database to develop climate model parameterizations. We thank Dr. John Peters and two anonymous reviewers for their constructive comments. The National Center for Atmospheric Research is sponsored by the National Science Foundation. REFERENCES Bluestein, H. B., and M. H. Jain, 1985: Formation of mesoscale lines of precipitation: Severe squall lines in Oklahoma during the spring. J. Atmos. Sci., 42, 1711–1732, doi:10.1175/ 1520-0469(1985)042,1711:FOMLOP.2.0.CO;2. Crook, N. A., and M. W. Moncrieff, 1988: The effect of large-scale convergence on the generation and maintenance of deep moist convection. J. Atmos. Sci., 45, 3606–3624, doi:10.1175/ 1520-0469(1988)045,3606:TEOLSC.2.0.CO;2. Dudhia, J., and M. W. Moncrieff, 1987: A numerical simulation of quasi-stationary tropical convective bands. Quart. J. Roy. Meteor. Soc., 113, 929–967, doi:10.1002/qj.49711347711.

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