Numerical Simulation of Tornadogenesis in an Outer-Rainband Minisupercell of Typhoon Shanshan on 17 September 2006

Wataru Mashiko Meteorological Research Institute, Tsukuba, Japan

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Hiroshi Niino Ocean Research Institute, The University of Tokyo, Tokyo, Japan

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Teruyuki Kato Meteorological Research Institute, Tsukuba, Japan

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Abstract

On 17 September 2006, three tornadoes occurred along the east coast of Kyusyu Island in western Japan during the passage of an outer rainband in the right-front quadrant of Typhoon Shanshan. To clarify the structure of the tornado-producing storms and the mechanism of tornadogenesis, quadruply nested numerical simulations were performed using a nonhydrostatic model with an innermost horizontal grid spacing of 50 m. Several simulated convective storms in the outermost rainband exhibited characteristics of a minisupercell. One storm had a strong rotating updraft of more than 30 m s−1 and a large vertical vorticity exceeding 0.06 s−1. This storm spawned a tornado when the low-level mesocyclone intensified. The tornado was generated on the rear-flank gust front near the mesocyclone center when a secondary rear-flank downdraft (RFD) surge advanced cyclonically around the low-level mesocyclone and overtook the rear-flank gust front at its left-front edge. Backward trajectories and vorticity budget analysis along the trajectories indicate that the secondary RFD surge played a key role in tornadogenesis by barotropically transporting the large streamwise vorticity associated with the environmental low-level veering shear toward the surface. When the secondary RFD outflow surge boundary reached the rear-flank gust front, the horizontal convergence was enhanced, contributing to the rapid amplification of the vertically tilted streamwise vorticity. The diagnostics of the vertical momentum equation and several sensitivity experiments demonstrated that precipitation loading in the area of a hook-shaped precipitation pattern was crucial to the behavior of the RFD and the subsequent tornadogenesis.

Corresponding author address: Wataru Mashiko, Meteorological Research Institute, 1-1 Nagamine, Tsukuba, Ibaraki 305-0052, Japan. Email: wmashiko@mri-jma.go.jp

Abstract

On 17 September 2006, three tornadoes occurred along the east coast of Kyusyu Island in western Japan during the passage of an outer rainband in the right-front quadrant of Typhoon Shanshan. To clarify the structure of the tornado-producing storms and the mechanism of tornadogenesis, quadruply nested numerical simulations were performed using a nonhydrostatic model with an innermost horizontal grid spacing of 50 m. Several simulated convective storms in the outermost rainband exhibited characteristics of a minisupercell. One storm had a strong rotating updraft of more than 30 m s−1 and a large vertical vorticity exceeding 0.06 s−1. This storm spawned a tornado when the low-level mesocyclone intensified. The tornado was generated on the rear-flank gust front near the mesocyclone center when a secondary rear-flank downdraft (RFD) surge advanced cyclonically around the low-level mesocyclone and overtook the rear-flank gust front at its left-front edge. Backward trajectories and vorticity budget analysis along the trajectories indicate that the secondary RFD surge played a key role in tornadogenesis by barotropically transporting the large streamwise vorticity associated with the environmental low-level veering shear toward the surface. When the secondary RFD outflow surge boundary reached the rear-flank gust front, the horizontal convergence was enhanced, contributing to the rapid amplification of the vertically tilted streamwise vorticity. The diagnostics of the vertical momentum equation and several sensitivity experiments demonstrated that precipitation loading in the area of a hook-shaped precipitation pattern was crucial to the behavior of the RFD and the subsequent tornadogenesis.

Corresponding author address: Wataru Mashiko, Meteorological Research Institute, 1-1 Nagamine, Tsukuba, Ibaraki 305-0052, Japan. Email: wmashiko@mri-jma.go.jp

1. Introduction

Significant progress has been made in our understanding of supercell storms through Doppler radar measurements and three-dimensional numerical simulations. However, our knowledge of the dynamics of tornadogenesis in supercell storms is still limited because of difficulties in collecting detailed observational data with good spatial and temporal resolutions and in numerically simulating a tornado that is more than two orders of magnitude smaller than a supercell storm. The existence of a low-level mesocyclone in the supercell alone is not a sufficient condition for tornadogenesis (e.g., Burgess et al. 1993; Trapp 1999; Wakimoto and Cai 2000; Markowski et al. 2002), although tornadogenesis is closely related to it.

Lemon and Doswell (1979) developed a conceptual model of a supercell after synthesizing available surface, visual, and radar observations. The model includes rear-flank downdraft (RFD), forward-flank downdraft (FFD), and associated rear-flank and forward-flank gust fronts that are the respective outflow boundaries from the RFD and FFD. Doppler radar analyses and storm chaser observations have revealed that tornadoes generally form close to the circulation center of the mesocyclone and near the boundary between updrafts and downdrafts. In many cases, a tornado is located along a gradient in vertical motion, between the RFD and the horseshoe-shaped updraft along the leading edge of the RFD (e.g., Lemon and Doswell 1979; Dowell and Bluestein 2002). The RFD has long been surmised to be critical in the genesis of significant tornadoes within supercell storms in many observational, numerical modeling, and theoretical studies (e.g., Markowski 2002; Davies-Jones 2006; Bluestein 2007). However, the precise role of the RFD in tornadogenesis remains somewhat unclear.

Davies-Jones (1982) suggested that when a downdraft forms adjacent to an updraft, vertical vorticity can be generated through the tilting of horizontal vorticity near the ground. Tornadogenesis must thus await the development of a significant downdraft that either tilts initially the horizontal vorticity or simply transports the vertical vorticity downward, if the surface is free slip and no vertical vorticity exists initially near the surface. Markowski et al. (2008) investigated vortex lines through the low-level mesocyclone regions of six supercell thunderstorms using pseudo-dual-Doppler airborne radar observations during the Verification of the Origins of Rotation in Tornadoes Experiment (VORTEX). They illustrated that the vortex lines are strongly suggestive of baroclinic vorticity generation within the hook echo and associated RFD region of the supercells. These vortex lines are subsequently tilted toward the vertical near the surface. This idea that baroclinic vorticity generation in the RFD is important to the development of near-ground rotation in supercells is consistent with Davies-Jones and Brooks (1993), Adlerman et al. (1999), and Straka et al. (2007). Microbursts, which are caused by strong negative buoyancy and tend to generate vortex rings, might therefore be able to trigger tornadoes (Fujita 1989; Davies-Jones 2006). In the meantime, Walko (1993) performed an idealized numerical experiment using a fixed heat sink and source to produce an RFD southwest of an updraft in a westerly shear flow. A tornado-like vortex formed barotropically beneath the updraft on the left-front edge of the advancing cold pool, indicating that the placement of the RFD was crucial for tornadogenesis. Davies-Jones (2006) demonstrated that hydrometeors around the periphery of an updraft dragged the angular momentum down to the ground. The angular momentum was then advected inward toward the center owing to surface friction–induced inflow. The frictional effects also influence the environmental field around tornadic storms. The boundary layer adjacent to the ground is a region of strong horizontal vorticity that significantly influences the development of low-level mesocyclones in the simulated supercells (Wicker 1996).

Recent in situ field observations reveal that tornadogenesis becomes more likely and tornado intensity and longevity increase as the surface buoyancy and equivalent potential temperature in the RFD increase, even though the tornadoes are close to a wind-shift line (Dowell and Bluestein 2002; Markowski et al. 2002; Shabbott and Markowski 2006). Markowski et al. (2003) conducted idealized numerical simulations using an axisymmetric model to show that the simulated tornadoes were stronger and longer-lived when the imposed downdrafts were relatively warm compared to when the downdrafts were relatively cold.

The surface observations and idealized numerical experiments noted above are inconsistent with most of the three-dimensional numerical studies for supercell tornadogenesis conducted from the 1980s to the mid-1990s (e.g., Klemp and Rotunno 1983; Rotunno and Klemp 1985; Wicker and Wilhelmson 1995; Wicker 1996). The results of these simulations revealed that the large vertical vorticity in the low-level mesocyclone and tornado was derived largely from horizontal vorticity generated baroclinically as low-level air approached the storm along the forward-flank cold outflow boundary. The role of baroclinically generated horizontal vorticity in the forward-flank gust front region of a supercell has been called into question by more recent observational studies.

The earlier works noted above studied tornadogenesis in a classic supercell in the midwestern United States. Supercells associated with tropical cyclones often have much smaller dimensions and are thus referred to as “minisupercells” (e.g., McCaul and Weisman 1996; Suzuki et al. 2000). It is of interest whether or not tornadogenesis in minisupercells proceeds in a similar manner as in a classic supercell.

On 17 September 2006, at least three tornadoes hit the east coast of Kyusyu Island in western Japan during the passage of a rainband associated with Typhoon Shanshan. The F2 tornado that hit Nobeoka City (see Fig. 1 for the geographical location) caused severe damage, including three fatalities. To reproduce the parent tornadic storm and the associated tornado simultaneously, quadruply nested numerical simulations were performed using a nonhydrostatic model with an innermost horizontal grid spacing of 50 m. The main objective in the present study is to clarify the mechanism of tornadogenesis in a simulated typhoon minisupercell.

This paper is organized as follows. Section 2 gives a brief overview of the synoptic and mesoscale fields. Section 3 describes the numerical model used in this study. The environmental fields of the tornadic storms are presented in section 4. Section 5 analyzes the structure of a simulated minisupercell. Section 6 details tornadogenesis in the minisupercell using simulation results from a grid having horizontal spacing of 50 m. Sensitivity experiments are provided in section 7. Section 8 discusses the implications of the results. A summary and conclusions are presented in the final section.

2. Overview of synoptic and mesoscale fields based on observational data

The strong Typhoon Shanshan, with a central pressure of 950 hPa, was moving north-northeastward at a speed of 35 km h−1 over the East China Sea when tornadoes hit Nichinan, Hyuuga, and Nobeoka on the east coast of Kyusyu Island at about 1210, 1330, and 1410 Japan standard time (JST; JST = UTC + 9 h), respectively, on 17 September 2006. A radar image at 1400 JST shows that two adjacent rainbands located about 300 km from the typhoon center extended from the north to the east side of the typhoon (Fig. 2). All three tornadoes were generated during the passage of the outermost rainband, which consisted of a number of isolated convective storms with a horizontal scale of 20–40 km. Several storms lasted for more than 2 h and spawned tornadoes. All of the tornadoes were located in the right-front quadrant of the moving typhoon.

A detailed damage survey (Miyazaki District Meteorological Observatory 2006) of the F2 Nobeoka tornado revealed the following points: 1) the tornado moved north-northwestward at a speed of about 90 km h−1 from 1403 to 1408 JST and 2) the damage swath was 7.5 km in length and 150–300 m in width. A pressure drop of about 1 hPa for a duration of about 7 min and a gradual temperature drop of about 1 K were recorded at the Nobeoka Meteorological Observatory, about 1 km west of the tornado path. The pressure drop is considered to be caused by a mesolow associated with the parent storm. The wind profiler at this observatory captured a strong veering shear, especially below 2 km, about 30 min before tornadogenesis. The winds were easterly at the surface and south-southeasterly at about 35 m s−1 at a height of 2 km. The storm-relative helicity (SREH) values (Davies-Jones et al. 1990) over the lowest 1, 2, and 3 km are 484, 641, and 695 m2 s−2, respectively, using the motion of a low-level mesocyclone center from the simulation results (see section 6a). Unfortunately, a power supply failure and technical problems prevented the acquisition of consistent quality data after this time. Moreover, Nobeoka City was not in the detection range of Doppler radars, so detailed features of the tornadic storm were not captured by the observations.

3. Numerical model

The numerical model used in this study is the Japan Meteorological Agency Nonhydrostatic Model (JMANHM; Saito et al. 2006). It is based on fully compressible equations with a map factor. The present study employed the bulk-type cloud microphysics scheme with six water species: water vapor, cloud water, rain, cloud ice, snow, and graupel (Lin et al. 1983; Murakami 1990). The turbulence closure scheme is based on Deardorff (1980) and predicts the turbulence kinetic energy. The surface fluxes are calculated by the bulk method, where the bulk coefficients are determined from the formulas of Kondo (1975) over the sea and of Louis et al. (1982) over land.

Quadruply nested one-way grids are used to conduct high-resolution model integrations. Hereinafter, the experiments with horizontal grid spacings of 5 km, 1 km, 250 m, and 50 m are referred to as NHM5km, NHM1km, NHM250m, and NHM50m, respectively. The model domains of NHM1km, NHM250m, and NHM50m are depicted in Fig. 1. The model specification is presented in Table 1. The vertical coordinate is terrain following: z* = H(zzs)/(Hzs), where zs is the surface height and H is the model-top height. The model contains 50 levels with variable grid intervals from Δz* = 40 m near the surface to 904 m at the top for NHM5km, NHM1km, and NHM250m, and 90 levels with Δz* = 40 m to 304 m for NHM50m. A split-explicit time-integration scheme [see Saito et al. (2006) for more details] was used in this study, and the large time step of NHM50m is 0.5 s.

The initial and boundary conditions of NHM5km are provided from an operational regional analysis of JMA that adopted a four-dimensional variational data assimilation system (JMA 2007). The initial time of NHM5km is 0900 JST on 17 September. The integration time of the innermost NHM50m is 26 min (1414–1440 JST), focusing on the outermost rainband that passed over Nobeoka. Unlike previous numerical studies on tornadoes, the present simulations include complex real topography and surface friction. However, the surface topography over a region less than 5 km square near Nobeoka was removed in order to simplify the analysis of the vorticity budget along the trajectories (section 6). It has been confirmed that the removal of the topography does not significantly influence the simulation results regarding the generation process of a tornado, which occurred about 500 m offshore in the original experiment.

4. Environmental field around the outermost rainband simulated by NHM1km

NHM1km successfully reproduced two adjacent rainbands in the right-front quadrant of the moving Shanshan except that the typhoon’s movement was about 20 min behind (cf. Figs. 3 and 2).

The simulated wind hodograph at Nobeoka just prior to the passage of the outermost rainband, which is assumed to correspond to the observed rainband consisting of tornado-producing convective storms (Fig. 2), indicates that a strong veering shear existed, especially below 2 km (Fig. 4a). The simulated environmental wind was east-southeasterly at about 14 m s−1 at a height of 20 m and south-southeasterly at about 40 m s−1 at a height of 2 km. The SREHs over the lowest 1, 2, and 3 km are 584, 841, and 910 m2 s−2, respectively, using the motion of a low-level mesocyclone center (see section 6a). When storm motion is defined as the mass-weighted mean wind over 6 km (McCaul 1991), these are 609, 684, and 690 m2 s−2, respectively. These values are comparable to the exceptionally large ones in Hurricane Bonnie (1998) (Molinari and Vollaro 2008), which are quite a bit larger than the mean of those in tornado proximity after hurricane landfall (McCaul 1991) or in a strong tornadic supercell environment in the midwestern United States (Thompson et al. 2003). The strong veering shear having large SREH was distributed along the outermost rainband, especially on its eastern (outer) side (not shown). This feature may be partly due to the secondary circulation of the tropical cyclone, which enhances the inflow in the boundary layer toward the outer side of rainbands (Powell 1990; Willoughby 1995).

Convective available potential energy (CAPE) around this rainband was about 1200 J kg−1, which was lower than the 2000 J kg−1 on the southern side of the typhoon (not shown). Although the former value is significantly larger than the average CAPE in a hurricane-tornado environment [e.g., 253 J kg−1; McCaul (1991)], it is still smaller than that of a typical supercell environment over the Great Plains [e.g., 2620 J kg−1; McCaul and Weisman (1996)]. Figure 4b plots the vertical profiles of potential temperature, equivalent potential temperature, and saturation equivalent potential temperature at Nobeoka at 1420 JST. Unlike a supercell environment in the midwestern United States (e.g., Klemp et al. 1981), the atmosphere is highly humid without significant low equivalent potential temperature layers or steep lapse rates aloft.

5. Structure of a minisupercell simulated by NHM250m

Figure 5 depicts the rainband simulated by NHM250m at 1420 JST. The outermost rainband consisted of a number of isolated convective storms with a horizontal scale of 20–40 km, which agrees with the radar observations (Fig. 2). Several convective storms turned out to have characteristics similar to typical supercells, namely, a hook-shaped horizontal distribution of the hydrometeors’ mixing ratio at the southern tip of a convective storm at a height of 1 km (Fig. 6a). The hydrometeor distribution, analogous to the bounded weak-echo region, was collocated with the mesocyclone center accompanied by a pressure drop of about 4 hPa. A storm-relative wind circulated cyclonically around the mesocyclone center.

A vertical cross section of the hydrometeors, vertical velocity, and vertical vorticity along the line A–B in Fig. 6a is given in Fig. 6c. A strong updraft of more than 30 m s−1 formed a “vault” structure of hydrometeors. Regions of large vertical vorticity overlapped with those of a strong updraft: maximum vertical vorticity exceeding 6 × 10−2 s−1 and updraft greater than 30 m s−1 are found at about 1- and 3-km heights. The hydrometeor top capping the vault region was only about 5 km, which is shallower than that of a classic supercell over the U.S. Great Plains. Thus, the simulated storm is similar to a minisupercell associated with a tropical cyclone (e.g., McCaul and Weisman 1996; Suzuki et al. 2000). In the present convective storm of about 30 km in horizontal dimension accompanying the mesocyclone at its southern edge, however, the convection was deeper on the northern side with the hydrometeor top reaching 10 km or more in height. Another feature to be noted is that, though the gust front near the surface was distinguishable by its horizontal wind shear and large vertical vorticity, the potential temperature difference across it was only about 1 K (Fig. 6b).

This simulated storm passed over Nobeoka around 1430 JST. A detailed analysis of the low-level mesocyclone and tornadogenesis will be made by using NHM50m in the next section.

6. Tornadogenesis in the minisupercell simulated by NHM50m

This section examines the tornadogenesis process in the minisupercell using NHM50m. We are not claiming that our simulation succeeded in reproducing the Nobeoka tornado. Rather, we think that the simulation reproduces a representative storm and associated tornado in the observed typhoon environment. Note that this simulation includes full physics and employs realistic initial and boundary conditions using JMA’s operational regional analysis, unlike previous idealized studies (Grasso and Cotton 1995; Wicker and Wilhelmson 1995; Noda and Niino 2005).

a. Evolution of a low-level mesocyclone and a tornado

Figure 7 shows the time series of minimum sea level pressure (SLPmin), maximum near-surface vertical vorticity at z* = 60 m (VORmax), and the horizontal distance between the location of VORmax and the low-level mesocyclone center. The low-level mesocyclone center is defined as the location of the maximum vertical vorticity averaged over a 1-km square at a height of 1 km. A rapid increase of VORmax and sudden drop of SLPmin occurred just after 1427:00 JST. VORmax reached 1.19 s−1, and the pressure drop was about 14 hPa at the peak. In this study, a tornado is conventionally defined as a vortex having a VORmax exceeding 0.65 s−1 in the NHM50m results. A tornado was generated about 700 m away from the low-level mesocyclone center at 1427:18 JST (Fig. 7).

Figures 8a–c illustrate the time–height cross section of the maximum vertical vorticity, the maximum updraft, and the minimum pressure perturbation calculated within a 2.5-km radius of the low-level mesocyclone center at each height. During the period shown in Fig. 8 (1418:00–1436:00 JST), the minisupercell entered its decaying stage, and the storm height was decreasing with time. Meanwhile, a low-level (below 2-km height) mesocyclone with large vertical vorticity, strong updraft, and significant pressure deficit descended and intensified with time, becoming especially prominent from 1425:00 JST. The tornado was initiated at 1427:18 JST. The updraft associated with the low-level mesocyclone with a vertical vorticity of more than 0.4 s−1 reached 50 m s−1 between 0.5 and 1.2 km in height just prior to tornadogenesis. It is apparent that the evolution of the simulated minisupercell is in many respects similar to that of a classic tornadic supercell (e.g., Klemp and Rotunno 1983; Noda and Niino 2005).

Finally, the distance between the VORmax location and the low-level mesocyclone center increased with time, and the tornado dissipated after landfall at 1433:24 JST, at which time the distance exceeded 1400 m (Fig. 7). The tornado lasted for about 6 min.

b. Tornado location in the minisupercell and its structure

To examine the location of the tornado in the minisupercell, horizontal cross sections of vertical velocity at a height of 150 m and potential temperature at z* = 20 m at 1428:30 JST (72 s after tornadogenesis) are presented in Figs. 9a and 9b, respectively. The tornado was located on the rear-flank gust front near the mesocyclone center, close to the intersection between the rear-flank and forward-flank gust fronts (Fig. 9b). Note that the tornado was also located between the left-front edge of the RFD wrapping around the low-level mesocyclone and the strong updraft near the mesocyclone center (Fig. 9a).

A close-up view of the rectangular area in Fig. 9a is shown in Fig. 9c. A strong downdraft of more than 10 m s−1 existed at the eastern edge of the tornado. This corresponds to an occlusion downdraft, caused by a dynamically induced downward pressure gradient force associated with strong low-level rotation (e.g., Klemp and Rotunno 1983; Wicker and Wilhelmson 1995). It is practically indistinguishable from the RFD because it is located near the front edge of the RFD that is wrapping around the developing tornado.

The structure of the simulated tornado is briefly examined here, although the resolution in this simulation is too coarse to resolve details reliably. The diameter of the vortex near the surface was about 500 m, based on the outermost closed pressure contour line (Fig. 9c). The tornado moved north-northwestward at a fast speed of about 90 km h−1, following the movement of the mesocyclone, which agrees well with the observations. A horizontal wind exceeding 50 m s−1 existed only on the right side of the tornado’s track (Fig. 9c). A vertical cross section of the cloud water along the line A–B in Fig. 9a is depicted in Fig. 9d. A funnel-shaped cloud collocated with a pressure deficit and large vertical vorticity is found. The tornado tilted northwestward with increasing height. Vertical profiles of the azimuthally averaged radial and tangential winds are presented in Fig. 10. A frictionally induced shallow inflow of more than 15 m s−1 is found at the lowest model level (z* = 20 m), which caused strong near-surface convergence. A strong tangential wind of more than 30 m s−1 existed between 200 and 250 m in height.

c. Generation process of the tornado in the minisupercell

The generation process of the simulated tornado is carefully examined here. Figures 11a–d illustrate the horizontal distributions of hydrometeors at a height of 1 km with an interval of 30 s from 1426:00 to 1427:30 JST. (Note that the tornado began at 1427:18 JST.) At this height, the hydrometeors consisted of only rainwater. As the low-level mesocyclone intensified, a hook-shaped distribution of hydrometeors became prominent. The RFD (Figs. 11e–h) was collocated well with the hook (Figs. 11a–d) and advanced around the near-surface circulation center. The main updraft lay ahead and to the left of the RFD. The advancing RFD outflow enhanced the low-level convergence at its leading edge, deforming the updraft region into a horseshoe shape with time.

Close-up views of the square areas surrounded by the dashed lines in Figs. 11e–h are shown in Figs. 11i–l. There existed two bands of strong updraft extending southward from the rotating main updraft near the mesocyclone center (Figs. 11i and 11j). The eastern updraft was associated with the rear-flank gust front, while the western updraft was located ahead of the secondary RFD outflow surge. The secondary RFD outflow surge enhanced the low-level convergence, resulting in the western updraft. Recent observational studies have revealed a complex structure in the form of multiple surges within the RFD region. The leading edge of the secondary surge was referred to as the internal RFD outflow surge boundary (Finley and Lee 2008; Lee et al. 2008) or as the secondary rear-flank gust front (Wurman et al. 2007; Marquis et al. 2008). However, the effects of the surges on tornadogenesis or on existing tornadoes have not been clarified. In this study, the leading edge of the secondary RFD surge advanced cyclonically around the low-level mesocyclone, eventually overtaking the rear-flank gust front at its left-front edge (Figs. 11k and 11l) as in the observation by Marquis et al. (2008). Most notable is that the tornado was generated in a region of strong updraft near the mesocyclone center at the left-front edge of the secondary RFD outflow surge, right after it reached the rear-flank gust front (Fig. 11l).

Judging from the evolution of the RFD, it appears that the secondary RFD outflow surge instigated tornadogenesis. To clarify the cause-and-effect relationship between the secondary RFD surge and tornadogenesis in more detail, three sets of backward-trajectory analyses initiating at 1426:30, 1427:00, and 1427:30 JST (48 s before, 18 s before, and 12 s after tornadogenesis) were conducted. Twenty-one parcels are distributed in the region of maximum vertical vorticity located at meandering or closed pressure contour lines over the rear-flank gust front at a height of 150 m (Figs. 12a–c). Their trajectories are obtained by integrating backward in time for 270 s. The three-dimensional model outputs at 1-s intervals are used for the integration, where the Euler scheme with a time step of 0.5 s is adopted. The velocity components at each point on the trajectory are calculated from linear spatial and temporal interpolations of the three-dimensional model outputs. The horizontal wind velocity below the lowest model level (z* = 20 m) is obtained by assuming a logarithmic profile with a surface roughness of 0.1 m over the land and 0.0002 m over the sea. Note that the trajectory calculations were performed in the storm-relative frame.

The trajectory analysis revealed that all parcels in the target region at 1426:30 JST originated only from the front side (northern part) of the storm (Fig. 12d). They started from the low-level northern region and traveled along nearly straight and horizontal paths. Of the parcels that were located in the target region at 1427:00 JST, at which time a slight pressure drop and amplification of vertical vorticity occurred just prior to tornadogenesis, about half originated from the RFD (Fig. 12e). These parcels descended cyclonically through the RFD from 200–500 m in height to a near-surface level and then ascended sharply into the vortex center. The parcels located in the tornado region at 1727:30 JST also originated from two source regions: the RFD and the front side of the storm (Fig. 12f). Trajectory analysis revealed that the tornado was generated just after the parcels traveling through the RFD entered the target region, implying that the secondary RFD surge played a key role in tornadogenesis.

d. Vorticity budget analysis along the trajectory through the RFD

To clarify the mechanism for tornadogenesis, the source and amplifying process of the vertical vorticity in the tornado must be identified. This section discusses the budget analysis for the vertical and horizontal components of the vorticity along a trajectory traveling through the RFD, which is a key factor in tornadogenesis as noted in section 6c. The equation for vertical vorticity ζ is given by
i1520-0493-137-12-4238-e1
where u, υ, and w are the three-dimensional velocity components; ρ is the density; p is the pressure; and Fx and Fy are the x and y components of turbulent mixing. The terms on the right-hand side (rhs) represent the horizontal convergence of vertical vorticity (commonly referred to as vertical stretching), the tilting of horizontal vorticity into the vertical, the solenoidal term, and the frictional term, in that order. The Coriolis force is neglected. Unlike in previous studies, the Boussinesq approximation is not used in this analysis.

Figure 13a depicts a representative path of a parcel that was located at a height of 150 m near the center of the tornado at 1427:30 JST (just after tornadogenesis) and originated from the RFD. The analyses for the other parcels traveling through the RFD gave qualitatively similar results. The validity of our vorticity budget analysis has been verified by comparing the evolution of the vorticity resulting from integrating the Lagrangian derivative [lhs of Eq. (1)], calculated from the sum of the terms on the rhs of Eq. (1), to the vorticity interpolated from the gridpoint values. These two quantities and their evolutions showed good agreement (not shown).

Figure 13b displays the time sequence of the vertical vorticity, the parcel height, and the terms in Eq. (1) during the final 90 s of the trajectory in Fig. 13a. Note that the solenoidal term is not shown since it is always on the order of 10−6 s−2 or less. The parcel first descended while having a small negative vertical vorticity. The vertical vorticity then increased and changed its sign before the parcel reached its nadir at 1427:09 JST. Throughout most of its descent, both the tilting term and the convergence term had small magnitudes. However, the tilting term increased rapidly right before the parcel reached the lowest part of its trajectory, and the vertical vorticity had a positive value of nearly 0.01 s−1 at its nadir. Once the trajectory turned upward, the convergence term became dominant, resulting in a rapid increase in the vertical vorticity. The frictional effect operated negatively after the vertical vorticity intensified.

The evolution of the horizontal vorticity, which is tilted into the vertical and eventually amplified by the horizontal convergence, is examined here.

To this end, it is convenient to write down the necessary equations in seminatural coordinates, where (s, n, k) represent orthonormal basis vectors with the wind vector V = (VH, 0, w) (e.g., Lilly 1982; Adlerman et al. 1999). The equations for the streamwise (ωs) and crosswise (ωn) horizontal vorticities are
i1520-0493-137-12-4238-e2
i1520-0493-137-12-4238-e3
where ψ = tan−1(υ/u) is the horizontal angle of the horizontal velocity vector that increases counterclockwise from the east. The first term on the rhs of both equations represents an exchange between streamwise and crosswise vorticity (without changing the magnitude of vorticity) due to a change in the direction of the horizontal velocity. The second and third terms on the rhs of each equation represent the rate of change of the streamwise/crosswise vorticity due to the convergence (horizontal stretching) and tilting of vortex tubes, respectively. Solenoidal and frictional effects are represented by the fourth and fifth terms on the rhs, respectively.

Figures 13c and 13d display the time evolution of the horizontal vorticity components and the terms in Eqs. (2) and (3) along the same trajectory for 330 s, up until right after tornadogenesis. Streamwise vorticity was the dominant component of the three-dimensional vorticity and increased steadily until the parcel reached the lowest part of its trajectory, while the crosswise vorticity increased less rapidly. Note that the parcel originally had a remarkably large streamwise vorticity of about 0.04 s−1 at 1422:00 JST and that it increased to 0.21 s−1 near its nadir.

Figure 13c clearly demonstrates that the large increase in the streamwise vorticity resulted mainly from the convergence term throughout its descent, followed by the exchange and tilting terms. The frictional term was always negative and worked a little more effectively at the lower level (not shown). Note that the baroclinic generation term (solenoidal term) was less than the order of 10−4 s−2 and thus was small relative to other dominant terms. The tilting and convergence terms in the streamwise vorticity equation rapidly decreased and became negative just before the parcel reached the lowest trajectory location, resulting in a rapid decrease in the streamwise vorticity. At the same time, the streamwise vorticity was tilted into the vertical.

In the crosswise direction, no single forcing term was dominant, but most of the crosswise horizontal vorticity was generated by the convergence and tilting terms. The exchange term tended to effectively reduce the crosswise vorticity with time. The solenoidal term was quite small (not shown). The frictional term did not contribute significantly to the crosswise vorticity, although it increased around the lowest part of the trajectory.

The analysis revealed that most of the streamwise vorticity that was tilted and stretched into the vertical arose principally from amplification of the initial large streamwise vorticity (about 0.04 s−1) due to the convergence term, followed by the exchange and tilting terms. Thus, our concern became, how did the parcel acquire the streamwise vorticity it possessed at the initial time of the trajectory in Fig. 13? Most of the parcels traveled cyclonically through the RFD around the mesocyclone originated from the northern side of the minisupercell between 200 and 500 m in height. The wind hodograph in Fig. 4 represents the storm environment, which had a large horizontal vorticity vector directed west-southwestward between these levels. It is apparent that the large streamwise vorticity of the parcels originated from the strong low-level vertical wind shear in the storm environment.

The RFD, which wrapped around the mesocyclone cyclonically, is important for tornadogenesis because of its downward barotropic transport of parcels with significant streamwise horizontal vorticity associated with environmental vertical shear. In addition, when the secondary RFD outflow surge hit the rear-flank gust front, it caused locally intensified surface convergence there, which significantly amplified the vertically tilted streamwise vorticity. It is inferred that the frictional effect on the cyclonically descended air aided the strong convergence near the surface.

e. What causes the RFD to wrap around the mesocyclone cyclonically?

Although the fact that the RFD is associated with hook echoes is well known, what causes the RFD is still poorly understood. Lemon and Doswell (1979) hypothesized that the RFD is dynamically formed on the upwind side of the updraft and is then enhanced and maintained by evaporative cooling and precipitation drag. Brandes (1981) suggested that the RFD forms as entrained environmental air is cooled evaporatively. Klemp et al. (1981) attributed the RFD in the so-called Del City storm to negative buoyancy caused by precipitation loading and to evaporative cooling, based on a precipitation trajectory analysis.

To clarify what causes the RFD to wrap around the mesocyclone cyclonically, which is a key subject in tornadogenesis, each term in the vertical momentum equation is calculated diagnostically. The equation for the Lagrangian time rate of change of w is
i1520-0493-137-12-4238-e4
where p′ is the pressure perturbation; ρ′ is the density fluctuation from the basic state of ρ, and g is the gravity acceleration. Here, ρ is a 10-km square horizontal average of ρ. The basic state of the pressure satisfies the hydrostatic balance with ρ. The first term on the rhs is the vertical gradient force of the perturbation pressure, which includes dynamic and buoyant contributions. Previous studies using an anelastic or Boussinesq approximation (e.g., Klemp and Rotunno 1983; Finley et al. 2001) revealed that the dynamic pressure perturbation gradient force near the mesocyclone is dominant in the first term, although the buoyant contribution depends on the definition of the base state in conjunction with the second term in (4) (Doswell and Markowski 2004). In this study the second term in (4) is referred to as the buoyancy forcing with approximte treatment. The Coriolis force and diffusional term are ignored since they are negligibly small compared with the other terms.

Horizontal plots of the terms in Eq. (4) at a height of 250 m at 1426:00 JST are presented in Figs. 14a–c. Figure 14a indicates that the perturbation pressure gradient forcing was strongly positive near the mesocyclone circulation center, which agrees with previous studies (e.g., Wicker and Wilhelmson 1995). In the region surrounding the mesocyclone, the perturbation pressure gradient forcing was slightly positive. However, a strong negative pressure forcing existed in a small area to the west of the mesocyclone. Figure 14b indicates that a region of negative buoyancy spread from the northwest to the south of the mesocyclone. This region is collocated with the RFD (cf. Figs. 14b and 11a). The negative buoyancy in the northwest quadrant of the low-level mesocyclone contributed significantly to the downward acceleration of the vertical velocity (Fig. 14c), which caused the trajectories in the RFD to start descending and to accelerate further downward (Figs. 12e and 12f). However, it is apparent from Fig. 14b that the baroclinic generation of the horizontal vorticity is less than the order of 10−4 s−2 as noted in section 6d.

To identify the contribution of precipitation loading to the downward acceleration, the buoyancy term in Eq. (4) was decomposed. Figures 15a and 15b depict the horizontal distribution of the buoyancy due to precipitation loading and the other contributions, respectively. In this analysis, the precipitation essentially consisted of rain at this level. The distributions of the two contributions to the buoyancy exhibited patterns quite similar to each other (cf. Figs. 15a and 15b), and the negative buoyancy around the mesocyclone was nearly collocated with the RFD (cf. Figs. 15 and 11e) and the hook in the hydrometeor field (cf. Figs. 15 and 11a). However, the negative buoyancy caused by precipitation loading contributed more strongly, especially in the northwest quadrant around the mesocyclone. Thus, it is concluded that precipitation loading, rather than evaporative cooling, is the determining factor in the formation of the RFD that cyclonically wraps around the mesocyclone.

At 1428:30 JST (72 s after tornadogenesis), strong downward nonhydrostatic forcing due to the vertical pressure perturbation gradient occurred rapidly near the center of the tornado at a lower level (not shown). This forcing caused a tornado-scale occlusion downdraft (Fig. 9c). The occlusion downdraft simulated in this study was a response to the near-ground rotation of the tornado and therefore was not an instigator of tornadogenesis.

7. Sensitivity experiments

Sensitivity experiments were performed using NHM50m, run from 1414 to 1440 JST, in order to further examine the effects of precipitation loading and evaporative cooling on the RFD and the subsequent tornadogenesis.

In the experiment NOEVP, evaporation from precipitation particles (rain, snow, and graupel) was not permitted. Figure 16a shows the time series of the surface fields in the NOEVP experiment. The tornado was generated almost at the same time (1427:12 JST) as in the control run. However, the near-surface maximum vertical vorticity (max-VORmax) increased to 1.49 s−1 (cf. 1.19 s−1 for the control run) and the minimum sea level pressure (min-SLPmin) was reduced to 973.2 hPa (cf. 979.9 hPa for the control run).

From the time–height plots in Figs. 17a–c, the evolution patterns of the low-level mesocyclone and tornadogenesis were quite similar to those of the control run (Figs. 8a–c), and the tornado lasted 2 min longer (cf. Figs. 16a and 7). The behavior of the RFD and subsequent tornadogenesis were also nearly the same as in the control experiment. However, a small difference is seen in the near-surface temperature field. In the NOEVP experiment, there was almost no temperature deficit around the storm (cf. Figs. 18 and 9b). Minisupercells associated with a tropical cyclone experience only weak evaporative cooling because of the highly humid environment, so the elimination of evaporative cooling had a small impact on the RFD. However, the weak warming of the RFD contributed effectively to the intensification of the tornado. This result is consistent with previous observational studies (e.g., Dowell and Bluestein 2002; Markowski et al. 2002), suggesting that the warmer surface temperature would be favorable for the genesis, intensification, and longevity of a tornado.

In contrast, the NOLOAD experiment, in which the weight of the precipitation (rain, snow, and graupel) is neglected, differs remarkably from the control run. A much weaker tornado (max-VORmax = 0.68 s−1 and min-SLPmin = 987.9 hPa) was generated about 4 min later than in the control run and quickly dissipated (Fig. 16b). Until about 1425:00 JST, the evolution of the low-level mesocyclone in the NOLOAD experiment was almost the same as in the control run (cf. Figs. 8a–c and 17d–f), except for a slightly faster translation speed (not shown). However, the tornado was not generated prior to 1427:30 JST.

A remarkable difference is also seen in the behavior of the RFD. Figures 19a–d show the vertical velocity at a height of 150 m from 1426:00 to 1427:30 JST with an interval of 30 s in the NOLOAD experiment. Unlike the control run, the RFD did not wrap around the low-level mesocyclone cyclonically and existed only in the southwestern portion of the mesocyclone (cf. Figs. 11e–h and 19a–d). The updraft occupied an area on the northwest side of the mesocyclone that was taken up by the RFD in the control experiment. In addition, the secondary RFD outflow surge was relatively weak in comparison with the control run. The difference in the behavior of the RFD is roughly accounted for by the buoyancy due to the precipitation loading, as seen in Fig. 15a.

Another feature to be noted is that the low-level mesocyclone did not decay in the later period. A low-level mesocyclone with a pressure deficit, large vorticity, and strong updraft was sustained even after 1433:00 JST (Figs. 17d–f). Most of the supercells experience a cyclic evolution wherein the original mesocyclone is completely wrapped by the RFD and decays rapidly (e.g., Burgess et al. 1982; Johnson et al. 1987; Adlerman et al. 1999). In the NOLOAD experiment, the RFD did not wrap around the mesocyclone, and this seemed to encourage the greater longevity of the low-level mesocyclone.

8. Discussion

To this point, this study has focused on the RFD for tornadogenesis. However, approximately half of all tornado trajectories came from the front side (northern part) of the storm, near the surface, as mentioned in section 6c. These parcels exhibited a more variable evolution than those originating from the RFD. They traveled close to the surface, so they are more prone to numerical errors in the vorticity budget analysis. To examine the contributions of parcels originating from the RFD and those from the front side of the tornado vortex separately, the evolution of the circulations is analyzed along closed material curves in the tornado at 1427:30 JST (just after tornadogenesis), as described below. The circulation C(t) is defined as
i1520-0493-137-12-4238-e5
which is equal to the area integral of the vorticity component perpendicular to the area enclosed by a curve. Circulations around closed material curves C1 and C2, as shown in Fig. 20a, are calculated. Curve C1 encloses parcels originating from the RFD, and C2 those originating from the front side of the storm and the surrounding area around the tornado.

Figure 20b presents the material curves at 1425:30 JST (about 108 s prior to tornadogenesis) calculated from the backward trajectories. Time series of the circulations, including the curve CALL that encircles the areas enclosed by C1 and C2, are shown in Fig. 20c. At 1427:30 JST, right after tornadogenesis, the contribution of C1 was about one-third of that of CALL. Therefore, the vorticity transported by the RFD does not account for the entire source of the vorticity in the tornado. It is apparent from Fig. 20b that at first the vortex tubes associated with the RFD were directed roughly horizontally and spread widely. Meanwhile, the circuit of C2 stayed close to the ground over a wide region on the northern side of the mesocyclone and contained vertical vorticity–rich air on the forward-flank gust front and the northern part of the rear-flank gust front. The circulation of C2 was always dominant in that of CALL (Fig. 20c). Note that the time change of each circulation is considered to be caused by the frictional effect and/or calculation error, since the baroclinic effect was relatively small, unlike a typical classic supercell in Rotunno and Klemp (1985).

This result implies that a low-level mesocyclone with strong updraft and the vertical vorticity–rich air along the gust fronts set the stage for tornadogenesis in a supercell (e.g., Noda and Niino 2005). However, these two factors are not sufficient conditions for tornadogenesis. The RFD wrapping around the low-level mesocyclone cyclonically is certainly an additional important factor supplying an extra source of the vorticity needed for tornadogenesis. Moreover, the secondary RFD outflow surge intensifies the horizontal convergence and amplifies the vertically tilted vorticity, as noted in section 6.

9. Summary and conclusions

At least three tornadoes hit the east coast of Kyusyu Island in western Japan during the passage of the outermost rainband in the right-front quadrant of Typhoon Shanshan on 17 September 2006. To clarify the structure of tornado-producing storms and the mechanism of tornadogenesis, numerical simulations were performed using quadruply nested grids with horizontal grid spacings of 5 km, 1 km, 250 m, and 50 m. The present simulations included complex realistic topography and surface friction.

The simulations faithfully reproduced the outermost rainband on the right-front quadrant of the typhoon. The environment around the rainband was characterized by strong low-level veering shear and a modest CAPE of about 1200 J kg−1. The rainband consisted of a number of isolated convective storms with a horizontal scale of 20–40 km. Several storms exhibited a hook-shaped pattern and a vaultlike structure of hydrometeors at their southern edge. One storm had a strong rotating updraft of more than 30 m s−1 with vertical vorticity exceeding 0.06 s−1 at about 1- and 3-km heights. The hydrometeor top capping the vault region was only about 5 km, and the near-surface temperature difference across the gust front was very small (about 1 K). These features are characteristics of previously documented tropical cyclone minisupercells.

The innermost simulation with a horizontal grid spacing of 50 m successfully reproduced a tornado spawned by a minisupercell approaching the coast of Nobeoka. The tornado was generated when the low-level mesocyclone intensified. The diameter of the tornado as determined from the pressure field near the surface was about 500 m, and the vertical vorticity exceeded 1.0 s−1.

The tornado was generated on the rear-flank gust front near the mesocyclone center when the secondary RFD outflow surge advanced cyclonically around the low-level mesocyclone and reached the rear-flank gust front at its left-front edge. A backward-trajectory analysis revealed that about half of the parcels located in the tornado vortex just after tornadogenesis originated from the RFD. A vorticity budget analysis for a representative parcel originating from the RFD showed that, when the parcel descended around the mesocyclone, the large streamwise vorticity associated with low-level environmental vertical shear was amplified mainly by the convergence and partly by the exchange and tilting terms. The baroclinic and frictional effects were relatively small in both the streamwise and crosswise vorticity budgets. Just before the parcel reached the lowest level, tilting of the streamwise vorticity into the vertical occurred and the parcel acquired a positive vertical vorticity. Once the trajectory turned upward, the convergence term in the vertical vorticity equation became large and dominant, resulting in a rapid increase in the vertical vorticity and the formation of a tornado.

It can be concluded that the secondary RFD surge, which advances around the mesocyclone cyclonically, plays a key role in tornadogenesis by barotropically transporting the large streamwise horizontal vorticity associated with low-level vertical shear in the environment. It provides an additional source of the vorticity needed for tornadogenesis. Moreover, the secondary RFD outflow surge enhances the horizontal convergence at its leading edge. The horizontal convergence rapidly amplifies the vertical vorticity tilted from the streamwise vorticity and forms a tornado. The behavior of the RFD is dominated by the negative buoyancy due to precipitation loading. The precipitation loading in the area of the hook is crucial to the formation of the RFD, including the secondary RFD outflow surge, which wraps around the mesocyclone in a minisupercell storm.

Acknowledgments

The authors thank three anonymous reviewers for their helpful comments, which contributed to improving the original manuscript. Thanks are extended to Osamu Suzuki of the Meteorological Research Institute (MRI) for his helpful comments. The authors also thank Syugo Hayashi of MRI for his support in conducting the numerical simulations. A topography dataset of the 50-m mesh provided by the Geographical Survey Institute was used for the high-resolution simulations. The numerical simulations in this study were performed using a NEC SX-6 supercomputer system at MRI. This work was conducted as part of the cooperative program (119, 2009) provided by the Ocean Research Institute, the University of Tokyo. This study was supported in part by the Ministry of Education, Culture, Sports, Science and Technology of Japan under the Special Coordination Funds for Promoting Science and Technology program “Mechanism and Prediction of Tornadoes and Countermeasures.”

REFERENCES

  • Adlerman, E. J., K. K. Droegemeier, and R. Davies-Jones, 1999: A numerical simulation of cyclic mesocyclogenesis. J. Atmos. Sci., 56 , 20452069.

    • Search Google Scholar
    • Export Citation
  • Bluestein, H. B., 2007: Advances in applications of the physics of fluids to severe weather systems. Rep. Prog. Phys., 70 , 12591323.

  • Brandes, E. A., 1981: Finestructure of the Del City–Edmond tornadic mesocirculation. Mon. Wea. Rev., 109 , 635647.

  • Burgess, D. W., V. T. Wood, and R. A. Brown, 1982: Mesocyclone evolution statistics. Preprints, 12th Conf. on Severe Local Storms, San Antonio, TX, Amer. Meteor. Soc., 422–424.

    • Search Google Scholar
    • Export Citation
  • Burgess, D. W., R. J. Donaldson Jr., and P. R. Desrochers, 1993: Tornado detection and warning by radar. The Tornado: Its Structure, Dynamics, Prediction and Hazards, Geophys. Monogr., Vol. 79, Amer. Geophys. Union, 203–221.

    • Search Google Scholar
    • Export Citation
  • Davies-Jones, R. P., 1982: Observational and theoretical aspects of tornadogenesis. Intense Atmospheric Vortices, L. Bengtsson and J. Lighthill, Eds., Topics in Atmospheric and Oceanographic Sciences, Springer-Verlag, 175–189.

    • Search Google Scholar
    • Export Citation
  • Davies-Jones, R. P., 2006: Tornadogenesis in supercell storms—What we know and what we don’t know. Preprints, Symposium on the Challenges of Severe Convective Storms, Atlanta, GA, Amer. Meteor. Soc., 2.2. [Available online at http://ams.confex.com/ams/pdfpapers/104563.pdf].

    • Search Google Scholar
    • Export Citation
  • Davies-Jones, R. P., and H. E. Brooks, 1993: Mesocyclogenesis from a theoretical perspective. The Tornado: Its Structure, Dynamics, Prediction, and Hazards, Geophys. Monogr., Vol. 79, Amer. Geophys. Union, 105–114.

    • Search Google Scholar
    • Export Citation
  • Davies-Jones, R. P., D. Burgess, and M. Foster, 1990: Test of helicity as a tornado forecast parameter. Preprints, 16th Conf. on Severe Local Storms, Kananaskis Park, AB, Canada, Amer. Meteor. Soc., 588–592.

    • Search Google Scholar
    • Export Citation
  • Deardorff, J. W., 1980: Stratocumulus-capped mixed layers derived from a three-dimensional model. Bound.-Layer Meteor., 18 , 495527.

  • Doswell III, C. A., and P. M. Markowski, 2004: Is buoyancy a relative quantity? Mon. Wea. Rev., 132 , 853863.

  • Dowell, D. C., and H. B. Bluestein, 2002: The 8 June 1995 McLean, Texas, storm. Part II: Cyclic tornado formation, maintenance, and dissipation. Mon. Wea. Rev., 130 , 26492670.

    • Search Google Scholar
    • Export Citation
  • Finley, C. A., and B. D. Lee, 2008: Mobile mesonet observations of an intense RFD and multiple RFD gust fronts in the May 23 Quinter, Kansas tornadic supercell during TWISTEX 2008. Preprints, 24th Conf. on Severe Local Storms, Savannah, GA, Amer. Meteor. Soc., P3.18. [Available online at http://ams.confex.com/ams/pdfpapers/142133.pdf].

    • Search Google Scholar
    • Export Citation
  • Finley, C. A., W. R. Cotton, and R. A. Pielke, 2001: Numerical simulation of tornadogenesis in a high-precipitation supercell. Part I: Storm evolution and transition into a bow echo. J. Atmos. Sci., 58 , 15971629.

    • Search Google Scholar
    • Export Citation
  • Fujita, T. T., 1989: The Teton–Yellowstone tornado of 21 July 1987. Mon. Wea. Rev., 117 , 19131940.

  • Grasso, L. D., and W. R. Cotton, 1995: Numerical simulation of a tornado vortex. J. Atmos. Sci., 52 , 11921203.

  • Japan Meteorological Agency, 2007: Outline of the operational numerical weather prediction at the Japan Meteorological Agency. JMA, 29–35. [Available from JMA, 1-3-4 Otemachi, Chiyoda-ku, Tokyo 100-8122, Japan].

    • Search Google Scholar
    • Export Citation
  • Johnson, K. W., P. S. Ray, B. C. Johnson, and R. P. Davies-Jones, 1987: Observations related to the rotational dynamics of the 20 May 1977 tornadic storms. Mon. Wea. Rev., 115 , 24632478.

    • Search Google Scholar
    • Export Citation
  • Klemp, J. B., and R. Rotunno, 1983: A study of the tornadic region within a supercell thunderstorm. J. Atmos. Sci., 40 , 359377.

  • Klemp, J. B., R. B. Wilhelmson, and P. S. Ray, 1981: Observed and numerically simulated structure of a mature supercell thunderstorm. J. Atmos. Sci., 38 , 15581580.

    • Search Google Scholar
    • Export Citation
  • Kondo, J., 1975: Air–sea bulk transfer coefficients in diabatic conditions. Bound.-Layer Meteor., 9 , 91112.

  • Lee, B. D., C. A. Finley, and T. M. Samaras, 2008: Thermodynamic and kinematic analysis near and within the Tipton, KS tornado on May 29 during TWISTEX 2008. Preprints, 24th Conf. on Severe Local Storms, Savannah, GA, Amer. Meteor. Soc., P3.13. [Available online at http://ams.confex.com/ams/pdfpapers/142078.pdf].

    • Search Google Scholar
    • Export Citation
  • Lemon, L. R., and C. A. Doswell III, 1979: Severe thunderstorm evolution and mesocyclone structure as related to tornadogenesis. Mon. Wea. Rev., 107 , 11841197.

    • Search Google Scholar
    • Export Citation
  • Lilly, D. K., 1982: The development and maintenance of rotation in convective storms. Intense Atmospheric Vortices, L. Bengtsson and J. Lighthill, Eds., Topics in Atmospheric and Oceanographic Sciences, Springer-Verlag, 149–160.

    • Search Google Scholar
    • Export Citation
  • Lin, Y. H., R. D. Farley, and H. D. Orville, 1983: Bulk parameterization of the snow field in a cloud model. J. Climate Appl. Meteor., 22 , 10651092.

    • Search Google Scholar
    • Export Citation
  • Louis, J. F., M. Tiedtke, and J. F. Geleyn, 1982: A short history of the operational PBL parameterization at ECMWF. Workshop on Planetary Boundary Layer Parameterization, Reading, United Kingdom, ECMWF, 59–79.

    • Search Google Scholar
    • Export Citation
  • Markowski, P. M., 2002: Hook echoes and rear-flank downdrafts: A review. Mon. Wea. Rev., 130 , 852876.

  • Markowski, P. M., J. M. Straka, and E. N. Rasmussen, 2002: Direct surface thermodynamic observations within the rear-flank downdrafts of nontornadic and tornadic supercells. Mon. Wea. Rev., 130 , 16921721.

    • Search Google Scholar
    • Export Citation
  • Markowski, P. M., J. M. Straka, and E. N. Rasmussen, 2003: Tornadogenesis resulting from the transport of circulation by a downdraft: Idealized numerical simulations. J. Atmos. Sci., 60 , 795823.

    • Search Google Scholar
    • Export Citation
  • Markowski, P. M., E. N. Rasmussen, J. M. Straka, R. Davies-Jones, Y. Richardson, and R. J. Trapp, 2008: Vortex lines within low-level mesocyclones obtained from pseudo-dual-Doppler radar observations. Mon. Wea. Rev., 136 , 35133535.

    • Search Google Scholar
    • Export Citation
  • Marquis, J., Y. Richardson, J. Wurman, and P. Markowski, 2008: Single- and dual-Doppler analysis of a tornadic vortex and surrounding storm-scale flow in the Crowell, Texas, supercell of 30 April 2000. Mon. Wea. Rev., 136 , 50175043.

    • Search Google Scholar
    • Export Citation
  • McCaul Jr., E. W., 1991: Buoyancy and shear characteristics of hurricane-tornado environments. Mon. Wea. Rev., 119 , 19541978.

  • McCaul Jr., E. W., and M. L. Weisman, 1996: Simulation of shallow supercell storms in landfalling hurricane environments. Mon. Wea. Rev., 124 , 408429.

    • Search Google Scholar
    • Export Citation
  • Miyazaki District Meteorological Observatory, 2006: On the gust caused by tornado in Miyazaki prefecture associated with Typhoon Shanshan on 17 September 2006 (in Japanese). Meteorological Survey Report on the Natural Disaster, Miyazaki District Meteorological Observatory, 52 pp.

    • Search Google Scholar
    • Export Citation
  • Molinari, J., and D. Vollaro, 2008: Extreme helicity and intense convective towers in Hurricane Bonnie. Mon. Wea. Rev., 136 , 43554372.

    • Search Google Scholar
    • Export Citation
  • Murakami, M., 1990: Numerical modeling of dynamical and microphysical evolution of an isolated convective cloud—The 19 July 1981 CCOPE cloud. J. Meteor. Soc. Japan, 68 , 107128.

    • Search Google Scholar
    • Export Citation
  • Noda, A., and H. Niino, 2005: Genesis and structure of a major tornado in a numerically-simulated supercell storm: Importance of vertical vorticity in a gust front. Sci. Online Lett. Atmos., 1 , 58.

    • Search Google Scholar
    • Export Citation
  • Powell, M. D., 1990: Boundary layer structure and dynamics in outer hurricane rainbands. Part I: Mesoscale rainfall and kinematic structure. Mon. Wea. Rev., 118 , 891917.

    • Search Google Scholar
    • Export Citation
  • Rotunno, R., and J. Klemp, 1985: On the rotation and propagation of simulated supercell thunderstorms. J. Atmos. Sci., 42 , 271292.

  • Saito, K., and Coauthors, 2006: The operational JMA nonhydrostatic mesoscale model. Mon. Wea. Rev., 134 , 12661298.

  • Shabbott, C. J., and P. M. Markowski, 2006: Surface in situ observations within the ouflow of forward-flank downdrafts of supercell thunderstorms. Mon. Wea. Rev., 134 , 14221441.

    • Search Google Scholar
    • Export Citation
  • Straka, J. M., E. N. Rasmussen, R. P. Davies-Jones, and P. M. Markowski, 2007: An observational and idealized numerical examination of low-level counter-rotating vortices in the rear flank of supercells. Electron. J. Severe Storms Meteor., 2 (8) [Available online at http://www.ejssm.org/ojs/index.php/ejssm/article/viewArticle/32].

    • Search Google Scholar
    • Export Citation
  • Suzuki, O., H. Niino, H. Ohno, and H. Nirasawa, 2000: Tornado-producing mini supercells associated with Typhoon 9019. Mon. Wea. Rev., 128 , 18681882.

    • Search Google Scholar
    • Export Citation
  • Thompson, R. L., R. Edwards, J. A. Hart, K. L. Elmore, and P. Markowski, 2003: Close proximity soundings within supercell environments obtained from the Rapid Update Cycle. Wea. Forecasting, 18 , 12431261.

    • Search Google Scholar
    • Export Citation
  • Trapp, R. J., 1999: Observations of nontornadic low-level mesocyclones and attendant tornadogenesis failure during VORTEX. Mon. Wea. Rev., 127 , 16931705.

    • Search Google Scholar
    • Export Citation
  • Wakimoto, R. M., and H. Cai, 2000: Analysis of a nontornadic storm during VORTEX 95. Mon. Wea. Rev., 128 , 565592.

  • Walko, R. L., 1993: Tornado spin-up beneath a convective cell: required basic structure of the near-field boundary layer winds. The Tornado: Its Structure, Dynamics, Prediction and Hazards, Geophys. Monogr., Vol. 79, Amer. Geophys. Union, 89–95.

    • Search Google Scholar
    • Export Citation
  • Wicker, L. J., 1996: The role of near surface wind shear on low-level mesocyclone generation and tornadoes. Preprints, 18th Conf. on Severe Local Storms, San Francisco, CA, Amer. Meteor. Soc., 115–119.

    • Search Google Scholar
    • Export Citation
  • Wicker, L. J., and R. Wilhelmson, 1995: Simulation and analysis of tornado development and decay within a three-dimensional supercell thunderstorm. J. Atmos. Sci., 52 , 26752703.

    • Search Google Scholar
    • Export Citation
  • Willoughby, H. E., 1995: Mature structure and evolution. Global Perspective on Tropical Cyclones, R. L. Elsberry, Ed., WMO/TD 693, 21–62.

    • Search Google Scholar
    • Export Citation
  • Wurman, J., Y. Richardson, C. Alexander, S. Weygandt, and P. F. Zhang, 2007: Dual-Doppler and single-Doppler analysis of a tornadic storm undergoing mergers and repeated tornadogenesis. Mon. Wea. Rev., 135 , 736758.

    • Search Google Scholar
    • Export Citation

Fig. 1.
Fig. 1.

Topography and geographical locations around Kyusyu Island. The area shown corresponds to the model domain of NHM1km (see section 3). The rectangular areas framed by dashed and solid lines indicate the model domains of NHM250m and NHM50m, respectively (see section 3).

Citation: Monthly Weather Review 137, 12; 10.1175/2009MWR2959.1

Fig. 2.
Fig. 2.

(a) Horizontal distribution of the precipitation intensity observed by JMA operational radars at 1400 JST. The cross (×) denotes the estimated center of Typhoon Shanshan at 1400 JST. The solid line shows its track by JMA’s best-track analysis.

Citation: Monthly Weather Review 137, 12; 10.1175/2009MWR2959.1

Fig. 3.
Fig. 3.

Horizontal distribution of the precipitation intensity from 1415 to 1420 JST simulated by NHM1km. The contour lines of the SLP at 1420 JST are drawn at 4-hPa intervals.

Citation: Monthly Weather Review 137, 12; 10.1175/2009MWR2959.1

Fig. 4.
Fig. 4.

(a) Hodograph of horizontal winds at Nobeoka at 1420 JST simulated by NHM1km. Numerals next to the profile indicate heights, and solid squares are plotted at intervals of 1 km above 1-km height. Storm motion (low-level mesocyclone; see section 6a for more details) is also shown. (b) Vertical profiles of the simulated potential temperature (thick solid line), equivalent potential temperature (thin solid line), and saturation equivalent potential temperature (broken line) at Nobeoka at 1420 JST.

Citation: Monthly Weather Review 137, 12; 10.1175/2009MWR2959.1

Fig. 5.
Fig. 5.

Horizontal distribution of the mixing ratio of hydrometeors (sum of rainwater, snow, and graupel) at a height of 1 km at 1420 JST simulated by NHM250m. Arrows denote ground-relative winds.

Citation: Monthly Weather Review 137, 12; 10.1175/2009MWR2959.1

Fig. 6.
Fig. 6.

(a) Close-up view of the rectangular area in Fig. 5. Arrows indicate storm-relative wind vectors, and dashed contour lines denote isobars at intervals of 1 hPa. (b) Horizontal cross section of potential temperature at z* = 20 m. Contour lines indicate vertical vorticity at intervals of 0.004 s−1. (c) Vertical cross section along the line A–B in (a). Solid contour lines denote vertical velocity at intervals of 10 m s−1. The dashed contour line represents a vertical vorticity of 0.06 s−1.

Citation: Monthly Weather Review 137, 12; 10.1175/2009MWR2959.1

Fig. 7.
Fig. 7.

Time series of SLPmin (solid line), VORmax (dashed line), and the distance between the location of VORmax and the low-level mesocyclone center (boldface line) as simulated by NHM50m. The low-level mesocyclone center was determined as the location of the maximum vertical vorticity averaged over a 1-km square at a height of 1 km. Each value was calculated within a 2.5-km radius of the low-level mesocyclone center. The distance is shown only for the period when the vortex satisfied the tornado criterion (see the text for more details).

Citation: Monthly Weather Review 137, 12; 10.1175/2009MWR2959.1

Fig. 8.
Fig. 8.

Time–height cross section of (a) maximum vertical vorticity, (b) minimum pressure perturbation, and (c) maximum vertical velocity in the simulated minisupercell from 1418 to 1436 JST. Each value was calculated within a 2.5-km radius of the low-level mesocyclone center.

Citation: Monthly Weather Review 137, 12; 10.1175/2009MWR2959.1

Fig. 9.
Fig. 9.

(a) Horizontal cross section of vertical velocity at a height of 150 m at 1428:30 JST (shaded). Contour lines denote pressure at 2-hPa intervals. Arrows denote storm-relative wind vectors. Boldface dashed lines indicate zero contours. (b) Horizontal cross section of potential temperature at z* = 20 m. Thin (thick) contour lines correspond to a vertical vorticity of 0.01 s−1 (0.10 s−1). Arrows denote storm-relative wind vectors. The tornado location is indicated by TR. (c) Close-up of the tornadic region in the rectangular area indicated by dashed lines in (a). Note that the grayscale differs from that in (a). Pressure contour lines are drawn at 1-hPa intervals. The heavy dashed contour line corresponds to a ground-relative wind of 50 m s−1. (d) Vertical cross section of cloud water along the line A–B in (a). The contour interval of the pressure perturbation is 300 Pa, and the heavy dashed contour line corresponds to a vertical vorticity of 0.50 s−1.

Citation: Monthly Weather Review 137, 12; 10.1175/2009MWR2959.1

Fig. 10.
Fig. 10.

Vertical profile of the radial (thin solid line) and tangential (thick solid) winds (m s−1) of the simulated tornado. The winds are azimuthally averaged for 100 m < r < 150 m, where r is the distance from the center of the tornado determined as the location of the maximum vertical vorticity at each level. Divergence (dashed line) calculated for r < 150 m is also shown.

Citation: Monthly Weather Review 137, 12; 10.1175/2009MWR2959.1

Fig. 11.
Fig. 11.

Evolution of (a)–(d) the mixing ratio of hydrometeors (sum of rainwater, snow, and, graupel) at a height of 1 km and (e)–(h) vertical velocity at a height of 150 m from 1426:00 to 1427:30 JST at intervals of 30 s until after tornadogenesis. (i)–(l) Close-ups of the rectangular areas in (e)–(h). Solid contour lines indicate pressure at intervals of 2 hPa, and arrows denote storm-relative wind vectors. The dashed contour lines in (i)–(l) denote a convergence of 0.07 s−1.

Citation: Monthly Weather Review 137, 12; 10.1175/2009MWR2959.1

Fig. 12.
Fig. 12.

Three sets of trajectory analyses in the storm-relative frame. (a)–(c) Initial locations of parcels for each trajectory analysis at 1426:30, 1427:00, and 1427:30 JST. The displayed areas in (a)–(c) correspond to the white rectangles in Figs. 11j–l, respectively. All parcels are initially at a height of 150 m. The vertical vorticity is color shaded. Contour lines indicate pressure at intervals of 1 hPa. (d)–(f) Horizontal projections of three-dimensional backward trajectories of the parcels shown in (a)–(c), respectively. Backward trajectories were calculated for 270 s. Colors on each trajectory indicate the parcel heights. Contour lines denote pressure at intervals of 1 hPa at the initial time of each backward trajectory.

Citation: Monthly Weather Review 137, 12; 10.1175/2009MWR2959.1

Fig. 13.
Fig. 13.

Budget analysis of the vorticity equation along the trajectory traveling through the RFD. (a) Horizontal projection of the 330-s backward trajectory for the targeted parcel, located near the tornado center at 1427:30 JST. The markers along the trajectory represent the heights for each 30 s. (b) Time series of the vertical vorticity, the terms in the vertical vorticity equation, and the parcel height for the last 90-s trajectory. (c) Time series of the streamwise vorticity and the terms in the streamwise vorticity equation. (d) As in (c), but for the crosswise vorticity. The solenoidal terms in the vertical and crosswise vorticity equations and the frictional term in the streamwise vorticity equation are not shown since they are relatively small compared to the other terms.

Citation: Monthly Weather Review 137, 12; 10.1175/2009MWR2959.1

Fig. 14.
Fig. 14.

Horizontal cross section of (a) perturbation pressure gradient forcing, (b) buoyancy forcing, and (c) their sum in the vertical momentum equation at a height of 250 m at 1426:00 JST. Heavy dashed lines indicate zero contours. The pressure contour lines are drawn for each 1 hPa. Arrows denote storm-relative wind vectors.

Citation: Monthly Weather Review 137, 12; 10.1175/2009MWR2959.1

Fig. 15.
Fig. 15.

Horizontal cross section of the buoyancy due to (a) precipitation loading and (b) the other contributions at a height of 150 m at 1426:00 JST. Heavy dashed lines indicate zero contours. The contour lines of pressure are drawn for each 1 hPa. Arrows denote storm-relative wind vectors.

Citation: Monthly Weather Review 137, 12; 10.1175/2009MWR2959.1

Fig. 16.
Fig. 16.

As in Fig. 7, but for (a) the experiment without evaporation from the precipitation (NOEVP) and (b) the experiment without precipitation loading (NOLOAD).

Citation: Monthly Weather Review 137, 12; 10.1175/2009MWR2959.1

Fig. 17.
Fig. 17.

As in Fig. 8, but for (a)–(c) the experiment without evaporation from the precipitation (NOEVP) and (d)–(f) the experiment without precipitation loading (NOLOAD).

Citation: Monthly Weather Review 137, 12; 10.1175/2009MWR2959.1

Fig. 18.
Fig. 18.

As in Fig. 9b, but for the experiment without evaporation from the precipitation (NOEVP).

Citation: Monthly Weather Review 137, 12; 10.1175/2009MWR2959.1

Fig. 19.
Fig. 19.

As in Figs. 11e–h, but for the experiment without precipitation loading (NOLOAD).

Citation: Monthly Weather Review 137, 12; 10.1175/2009MWR2959.1

Fig. 20.
Fig. 20.

(a) Initial positions of two closed material curves (C1 and C2) around the tornado at 1427:30 JST. Each material curve consists of several parcels at a height of 150 m. Here, C1 is composed of parcels that travel through the RFD. Pressure contour lines are shown for each 2 hPa. (b) Projection of C1 and C2 at 1425:30 JST calculated by backward trajectories within a storm-relative frame. Markers on the curves denote parcel heights. The contour lines indicate pressure at 1427:30 JST with an interval of 2 hPa. The dashed rectangle denotes the displayed area in (a). (c) Time series of the circulations along C1, C2, and CALL from 1425:30 to 1427:30 JST. Here, CALL represents the entire circle, including C1 and C2.

Citation: Monthly Weather Review 137, 12; 10.1175/2009MWR2959.1

Table 1.

Design of the models.

Table 1.
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  • Adlerman, E. J., K. K. Droegemeier, and R. Davies-Jones, 1999: A numerical simulation of cyclic mesocyclogenesis. J. Atmos. Sci., 56 , 20452069.

    • Search Google Scholar
    • Export Citation
  • Bluestein, H. B., 2007: Advances in applications of the physics of fluids to severe weather systems. Rep. Prog. Phys., 70 , 12591323.

  • Brandes, E. A., 1981: Finestructure of the Del City–Edmond tornadic mesocirculation. Mon. Wea. Rev., 109 , 635647.

  • Burgess, D. W., V. T. Wood, and R. A. Brown, 1982: Mesocyclone evolution statistics. Preprints, 12th Conf. on Severe Local Storms, San Antonio, TX, Amer. Meteor. Soc., 422–424.

    • Search Google Scholar
    • Export Citation
  • Burgess, D. W., R. J. Donaldson Jr., and P. R. Desrochers, 1993: Tornado detection and warning by radar. The Tornado: Its Structure, Dynamics, Prediction and Hazards, Geophys. Monogr., Vol. 79, Amer. Geophys. Union, 203–221.

    • Search Google Scholar
    • Export Citation
  • Davies-Jones, R. P., 1982: Observational and theoretical aspects of tornadogenesis. Intense Atmospheric Vortices, L. Bengtsson and J. Lighthill, Eds., Topics in Atmospheric and Oceanographic Sciences, Springer-Verlag, 175–189.

    • Search Google Scholar
    • Export Citation
  • Davies-Jones, R. P., 2006: Tornadogenesis in supercell storms—What we know and what we don’t know. Preprints, Symposium on the Challenges of Severe Convective Storms, Atlanta, GA, Amer. Meteor. Soc., 2.2. [Available online at http://ams.confex.com/ams/pdfpapers/104563.pdf].

    • Search Google Scholar
    • Export Citation
  • Davies-Jones, R. P., and H. E. Brooks, 1993: Mesocyclogenesis from a theoretical perspective. The Tornado: Its Structure, Dynamics, Prediction, and Hazards, Geophys. Monogr., Vol. 79, Amer. Geophys. Union, 105–114.

    • Search Google Scholar
    • Export Citation
  • Davies-Jones, R. P., D. Burgess, and M. Foster, 1990: Test of helicity as a tornado forecast parameter. Preprints, 16th Conf. on Severe Local Storms, Kananaskis Park, AB, Canada, Amer. Meteor. Soc., 588–592.

    • Search Google Scholar
    • Export Citation
  • Deardorff, J. W., 1980: Stratocumulus-capped mixed layers derived from a three-dimensional model. Bound.-Layer Meteor., 18 , 495527.

  • Doswell III, C. A., and P. M. Markowski, 2004: Is buoyancy a relative quantity? Mon. Wea. Rev., 132 , 853863.

  • Dowell, D. C., and H. B. Bluestein, 2002: The 8 June 1995 McLean, Texas, storm. Part II: Cyclic tornado formation, maintenance, and dissipation. Mon. Wea. Rev., 130 , 26492670.

    • Search Google Scholar
    • Export Citation
  • Finley, C. A., and B. D. Lee, 2008: Mobile mesonet observations of an intense RFD and multiple RFD gust fronts in the May 23 Quinter, Kansas tornadic supercell during TWISTEX 2008. Preprints, 24th Conf. on Severe Local Storms, Savannah, GA, Amer. Meteor. Soc., P3.18. [Available online at http://ams.confex.com/ams/pdfpapers/142133.pdf].

    • Search Google Scholar
    • Export Citation
  • Finley, C. A., W. R. Cotton, and R. A. Pielke, 2001: Numerical simulation of tornadogenesis in a high-precipitation supercell. Part I: Storm evolution and transition into a bow echo. J. Atmos. Sci., 58 , 15971629.

    • Search Google Scholar
    • Export Citation
  • Fujita, T. T., 1989: The Teton–Yellowstone tornado of 21 July 1987. Mon. Wea. Rev., 117 , 19131940.

  • Grasso, L. D., and W. R. Cotton, 1995: Numerical simulation of a tornado vortex. J. Atmos. Sci., 52 , 11921203.

  • Japan Meteorological Agency, 2007: Outline of the operational numerical weather prediction at the Japan Meteorological Agency. JMA, 29–35. [Available from JMA, 1-3-4 Otemachi, Chiyoda-ku, Tokyo 100-8122, Japan].

    • Search Google Scholar
    • Export Citation
  • Johnson, K. W., P. S. Ray, B. C. Johnson, and R. P. Davies-Jones, 1987: Observations related to the rotational dynamics of the 20 May 1977 tornadic storms. Mon. Wea. Rev., 115 , 24632478.

    • Search Google Scholar
    • Export Citation
  • Klemp, J. B., and R. Rotunno, 1983: A study of the tornadic region within a supercell thunderstorm. J. Atmos. Sci., 40 , 359377.

  • Klemp, J. B., R. B. Wilhelmson, and P. S. Ray, 1981: Observed and numerically simulated structure of a mature supercell thunderstorm. J. Atmos. Sci., 38 , 15581580.

    • Search Google Scholar
    • Export Citation
  • Kondo, J., 1975: Air–sea bulk transfer coefficients in diabatic conditions. Bound.-Layer Meteor., 9 , 91112.

  • Lee, B. D., C. A. Finley, and T. M. Samaras, 2008: Thermodynamic and kinematic analysis near and within the Tipton, KS tornado on May 29 during TWISTEX 2008. Preprints, 24th Conf. on Severe Local Storms, Savannah, GA, Amer. Meteor. Soc., P3.13. [Available online at http://ams.confex.com/ams/pdfpapers/142078.pdf].

    • Search Google Scholar
    • Export Citation
  • Lemon, L. R., and C. A. Doswell III, 1979: Severe thunderstorm evolution and mesocyclone structure as related to tornadogenesis. Mon. Wea. Rev., 107 , 11841197.

    • Search Google Scholar
    • Export Citation
  • Lilly, D. K., 1982: The development and maintenance of rotation in convective storms. Intense Atmospheric Vortices, L. Bengtsson and J. Lighthill, Eds., Topics in Atmospheric and Oceanographic Sciences, Springer-Verlag, 149–160.

    • Search Google Scholar
    • Export Citation
  • Lin, Y. H., R. D. Farley, and H. D. Orville, 1983: Bulk parameterization of the snow field in a cloud model. J. Climate Appl. Meteor., 22 , 10651092.

    • Search Google Scholar
    • Export Citation
  • Louis, J. F., M. Tiedtke, and J. F. Geleyn, 1982: A short history of the operational PBL parameterization at ECMWF. Workshop on Planetary Boundary Layer Parameterization, Reading, United Kingdom, ECMWF, 59–79.

    • Search Google Scholar
    • Export Citation
  • Markowski, P. M., 2002: Hook echoes and rear-flank downdrafts: A review. Mon. Wea. Rev., 130 , 852876.

  • Markowski, P. M., J. M. Straka, and E. N. Rasmussen, 2002: Direct surface thermodynamic observations within the rear-flank downdrafts of nontornadic and tornadic supercells. Mon. Wea. Rev., 130 , 16921721.

    • Search Google Scholar
    • Export Citation
  • Markowski, P. M., J. M. Straka, and E. N. Rasmussen, 2003: Tornadogenesis resulting from the transport of circulation by a downdraft: Idealized numerical simulations. J. Atmos. Sci., 60 , 795823.

    • Search Google Scholar
    • Export Citation
  • Markowski, P. M., E. N. Rasmussen, J. M. Straka, R. Davies-Jones, Y. Richardson, and R. J. Trapp, 2008: Vortex lines within low-level mesocyclones obtained from pseudo-dual-Doppler radar observations. Mon. Wea. Rev., 136 , 35133535.

    • Search Google Scholar
    • Export Citation
  • Marquis, J., Y. Richardson, J. Wurman, and P. Markowski, 2008: Single- and dual-Doppler analysis of a tornadic vortex and surrounding storm-scale flow in the Crowell, Texas, supercell of 30 April 2000. Mon. Wea. Rev., 136 , 50175043.

    • Search Google Scholar
    • Export Citation
  • McCaul Jr., E. W., 1991: Buoyancy and shear characteristics of hurricane-tornado environments. Mon. Wea. Rev., 119 , 19541978.

  • McCaul Jr., E. W., and M. L. Weisman, 1996: Simulation of shallow supercell storms in landfalling hurricane environments. Mon. Wea. Rev., 124 , 408429.

    • Search Google Scholar
    • Export Citation
  • Miyazaki District Meteorological Observatory, 2006: On the gust caused by tornado in Miyazaki prefecture associated with Typhoon Shanshan on 17 September 2006 (in Japanese). Meteorological Survey Report on the Natural Disaster, Miyazaki District Meteorological Observatory, 52 pp.

    • Search Google Scholar
    • Export Citation
  • Molinari, J., and D. Vollaro, 2008: Extreme helicity and intense convective towers in Hurricane Bonnie. Mon. Wea. Rev., 136 , 43554372.

    • Search Google Scholar
    • Export Citation
  • Murakami, M., 1990: Numerical modeling of dynamical and microphysical evolution of an isolated convective cloud—The 19 July 1981 CCOPE cloud. J. Meteor. Soc. Japan, 68 , 107128.

    • Search Google Scholar
    • Export Citation
  • Noda, A., and H. Niino, 2005: Genesis and structure of a major tornado in a numerically-simulated supercell storm: Importance of vertical vorticity in a gust front. Sci. Online Lett. Atmos., 1 , 58.

    • Search Google Scholar
    • Export Citation
  • Powell, M. D., 1990: Boundary layer structure and dynamics in outer hurricane rainbands. Part I: Mesoscale rainfall and kinematic structure. Mon. Wea. Rev., 118 , 891917.

    • Search Google Scholar
    • Export Citation
  • Rotunno, R., and J. Klemp, 1985: On the rotation and propagation of simulated supercell thunderstorms. J. Atmos. Sci., 42 , 271292.

  • Saito, K., and Coauthors, 2006: The operational JMA nonhydrostatic mesoscale model. Mon. Wea. Rev., 134 , 12661298.

  • Shabbott, C. J., and P. M. Markowski, 2006: Surface in situ observations within the ouflow of forward-flank downdrafts of supercell thunderstorms. Mon. Wea. Rev., 134 , 14221441.

    • Search Google Scholar
    • Export Citation
  • Straka, J. M., E. N. Rasmussen, R. P. Davies-Jones, and P. M. Markowski, 2007: An observational and idealized numerical examination of low-level counter-rotating vortices in the rear flank of supercells. Electron. J. Severe Storms Meteor., 2 (8) [Available online at http://www.ejssm.org/ojs/index.php/ejssm/article/viewArticle/32].

    • Search Google Scholar
    • Export Citation
  • Suzuki, O., H. Niino, H. Ohno, and H. Nirasawa, 2000: Tornado-producing mini supercells associated with Typhoon 9019. Mon. Wea. Rev., 128 , 18681882.

    • Search Google Scholar
    • Export Citation
  • Thompson, R. L., R. Edwards, J. A. Hart, K. L. Elmore, and P. Markowski, 2003: Close proximity soundings within supercell environments obtained from the Rapid Update Cycle. Wea. Forecasting, 18 , 12431261.

    • Search Google Scholar
    • Export Citation
  • Trapp, R. J., 1999: Observations of nontornadic low-level mesocyclones and attendant tornadogenesis failure during VORTEX. Mon. Wea. Rev., 127 , 16931705.

    • Search Google Scholar
    • Export Citation
  • Wakimoto, R. M., and H. Cai, 2000: Analysis of a nontornadic storm during VORTEX 95. Mon. Wea. Rev., 128 , 565592.

  • Walko, R. L., 1993: Tornado spin-up beneath a convective cell: required basic structure of the near-field boundary layer winds. The Tornado: Its Structure, Dynamics, Prediction and Hazards, Geophys. Monogr., Vol. 79, Amer. Geophys. Union, 89–95.

    • Search Google Scholar
    • Export Citation
  • Wicker, L. J., 1996: The role of near surface wind shear on low-level mesocyclone generation and tornadoes. Preprints, 18th Conf. on Severe Local Storms, San Francisco, CA, Amer. Meteor. Soc., 115–119.