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  • View in gallery
    Fig. 1.

    Computational domain and model terrain (m). Contour interval is 1000 m. The inner box denotes the subdomain over which the model results are presented.

  • View in gallery
    Fig. 2.

    Simulated geopotential height and wind field in CTRL at the 12-h forecast time (valid at 1200 UTC 31 May). Winds (m s−1) with one pennant, full barb, and half-barb denoting 25, 5, and 2.5 m s−1, respectively: (a) 700 hPa, geopotential heights (solid) every 30 m; (b) 700 hPa, isotach every 2.5 m s−1 starting from 12.5 m s−1; (c) as in (a) except for 250 hPa with geopotential heights every 120 m; and (d) 250 hPa with isotach every 5 m s−1 starting from 30 m s−1.

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    Fig. 3.

    Same as Fig. 2 but for objectively analyzed geopotential height and wind field for 1200 UTC 31 May (after Chen et al. 1994).

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    Fig. 4.

    Same as Fig. 2 but for the 24-h forecast time (valid at 0000 UTC 1 June).

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    Fig. 5.

    Same as Fig. 3 but for 0000 UTC 1 June (after Chen et al. 1994).

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    Fig. 6.

    The simulated relative humidity at the 400-hPa level in CTRL at the 24-h forecast time valid at 0000 UTC 1 June. The interval is 10%.

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    Fig. 7.

    The simulated 24-h accumulated precipitation (mm) in CTRL ending at 0000 UTC 1 June. Isopleths are 1, 10, and 25mm. The interval is 25 mm for values greater than 25 mm.

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    Fig. 8.

    The 700-hPa analysis for CTRL at the 24-h forecast time (valid at 0000 UTC 1 June): (a) 24-h differences in the along-contour component of the simulated winds (isolines) and the cross-contour component of the ageostrophic winds (arrows); (b) 24-h differences in the simulated geostrophic winds (isolines) and the isallohypsic winds (arrows). The contour intervals are 4 m s−1. Winds (m s−1) with full barb and half-barb representing 5 and 2.5 m s−1, respectively. The 250-hPa analysis for CTRL: (c) as in (a) with contour intervals in 6 m s−1; (d) as in (b) with contour intervals in 6 m s−1.

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    Fig. 9.

    Simulated geopotential height and wind field in NOLH at the 24-h forecast time (valid at 0000 UTC 1 June). Wind barbs (m s−1) with one pennant, full barb, and half-barb denoting 25, 5, and 2.5 m s−1, respectively: (a) 700 hPa, geopotential heights (solid) every 30 m; (b) 700 hPa, isotach every 2.5 m s−1 starting from 5 m s−1; (c) as in (a) except for 250 hPa with geopotential heights every 120 m; and (d) 250 hPa, isotach every 5 m s−1 starting from 30 m s−1.

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    Fig. 10.

    As in Fig. 9 except for NOSF.

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    Fig. 11.

    As in Fig. 9 except for CTRL40.

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    Fig. 12.

    As in Fig. 9 except for CTRL40-28.

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    Fig. 13.

    Simulated geopotential height and wind field at the 700-hPa level in NOEVA at the 24-h forecast time (valid at 0000 UTC 1 June). (a) Geopotential heights (solid) every 30 m. Winds (m s−1) with one pennant, full barb, and half-barb denoting 25, 5, and 2.5 m s−1, respectively; (b) isotach every 2.5 m s−1 starting from 12.5 m s−1 at the 700-hPa level; (c) the 850–500-hPa thickness difference between CTRL and NOEVA (CTRL minus NOEVA).

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    Fig. 14.

    Potential vorticity at the 250-hPa level with interval every 0.8 PVU (1 PVU = 10−6 m2 K kg−1 s−1) for (a) initial condition (valid at 0000 UTC 1 June); (b) 24-h forecast in CTRL; and (c) 24-h forecast in NOLH.

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    Fig. 15.

    Vertical cross sections of PV (solid lines, every 0.8 PVU) and potential temperature (dashed lines, every 4 K) along the line AA′ in Fig. 14b for (a) 24-h forecast in CTRL and (b) 24-h forecast in NOLH.

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    Fig. 16.

    Differences between CTRL and NOLH (CTRL minus NOLH) at the 24-h forecast time (valid at 0000 UTC 1 June): (a) Thickness between 500- and 250-hPa levels (every 20 m); (b) geopotential height (every 20 m) and vector winds at the 250-hPa level. 250-hPa analysis (solid line is geopotential heights in 120 m): (c) ageostrophic winds due to inertial advection term in CTRL; (d) as in (c) except for NOLH.

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    Fig. 17.

    Observed area-averaged (over the box in Fig. 4b) vertical profile of apparent heat source (Q1) at 0000 UTC 1June.

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    Fig. 18.

    Analysis at 24-h forecast time (valid at 0000 UTC 1 June). (a) As in Fig. 16a except for thickness between 850- and 500-hPa levels (every 10 m); (b) as in Fig. 16b except for the 700-hPa level. 700-hPa analysis (solid lines are geopotential heights every 30 m): (c) ageostrophic wind due to local change of wind in CTRL; (d) as in (c) except for NOLH.

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    Fig. 19.

    Analysis at the 24-h forecast time (valid at 0000 UTC 1 June): 700-hPa relative vorticity (thin lines every 4 × 10−5 s−1) and 850–500-hPa thickness (solid every 30 m) in (a) CTRL; (b) NOLH; and 300-hPa relative vorticity and 500–250-hPa thickness (every 60 m); (c) CTRL; (d) NOLH.

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    Fig. 20.

    The vertical p-velocity contours (every 2 μb s−1) at the 24-h forecast time (valid at 0000 UTC 1 June) for 700-hPa (a) CTRL; (b) NOLH; and for 250-hPa (c) CTRL; (d) CTRL.

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    Fig. 21.

    Area-averaged (over the box in Fig. 4b) vertical distribution of baroclinic conversion term −ωα at the 24-h forecast time (valid at 0000 UTC 1 June) (ω is p-velocity, α is specific volume).

  • View in gallery
    Fig. 22.

    Vertical cross-sectional differences between CTRL and NOLH (CTRL minus NOLH) along the line AA′ in Fig. 14b at the 24-h forecast time (valid at 0000 UTC 1 June). (a) Secondary circulation. Ageostrophic wind differences (m s−1, horizontal component of arrow) and minus omegas (μb s−1, vertical component of arrow). The maximum ageostrophic wind difference along the cross section is 35 m s−1. The contour interval for minus omega is 2 μb s−1; (b) Wind speed normal to the cross-sectional plane (every 2.5 m s−1).

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A Numerical Study of the Low-Level Jet during TAMEX IOP 5

Yi-Leng ChenDepartment of Meteorology, School of Ocean and Earth Science and Technology, University of Hawaii at Manoa, Honolulu, Hawaii

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Xin An ChenDepartment of Meteorology, School of Ocean and Earth Science and Technology, University of Hawaii at Manoa, Honolulu, Hawaii

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Sue ChenNational Center for Atmospheric Research, Boulder, Colorado

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Ying-Hwa KuoNational Center for Atmospheric Research, Boulder, Colorado

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Abstract

The Pennsylvania State University–National Center for Atmospheric Research Mesoscale Model Version 4 is used to simulate the cyclogenesis and the development of a low-level jet (LLJ) that occurred during the Taiwan Area Mesoscale Experiment intensive observing period 5 over southern China. Evaluation of the model results during a 36-h period indicates that the model successfully reproduces most principal features of this event, including cyclone path, intensification of the LLJ, distribution of precipitation, and the secondary circulation across the jet–front system.

Sensitivity tests show that latent heat release is important for the deepening of the cyclone and the development of the LLJ, whereas the model results are not sensitive to boundary layer physics. The lee trough east of the Tibetan Plateau provides the initial low-level vorticity. The initial deepening of the lee cyclone and the development of the low-level southwesterly flow are caused by the vertical motion associated with the upper-level short-wave trough. The potential vorticity and tropopause folding associated with the upper-level front are present in the model simulations even without latent heating. The low-level southwesterly flow transports warm, moist air from the south as the moisture source for condensation. Latent heating results in an increase in the thickness ahead of the short-wave trough in the upper levels, further deepening of the lee cyclone, and a stronger secondary circulation. The LLJ develops through the Coriolis force acting on the cross-contour ageostrophic winds in response to the increased pressure gradients related to the development of the cyclone and is enhanced by latent heating. The dynamic forcing aloft is also enhanced. Condensation heating and warm advection (evaporative cooling and cold advection) exceed adiabatic cooling (warming) ahead of (behind) the cyclone in the lower troposphere, therefore enhancing the low-level baroclinity. These processes interact nonlinearly leading to the further intensification of the LLJ.

* SOEST Contribution Number 4461.

+ Current affiliation: Mesoscale and Microscale Meteorology Division, National Center for Atmospheric Research, Boulder, Colorado.

Corresponding author address: Dr.Yi-Leng Chen, Department of Meteorology/SOEST, University of Hawaii at Manoa, 2525 Correa Road, HIG 331, Honolulu, HI 96822.

Email: dave@kukui.soest.hawaii.edu

Abstract

The Pennsylvania State University–National Center for Atmospheric Research Mesoscale Model Version 4 is used to simulate the cyclogenesis and the development of a low-level jet (LLJ) that occurred during the Taiwan Area Mesoscale Experiment intensive observing period 5 over southern China. Evaluation of the model results during a 36-h period indicates that the model successfully reproduces most principal features of this event, including cyclone path, intensification of the LLJ, distribution of precipitation, and the secondary circulation across the jet–front system.

Sensitivity tests show that latent heat release is important for the deepening of the cyclone and the development of the LLJ, whereas the model results are not sensitive to boundary layer physics. The lee trough east of the Tibetan Plateau provides the initial low-level vorticity. The initial deepening of the lee cyclone and the development of the low-level southwesterly flow are caused by the vertical motion associated with the upper-level short-wave trough. The potential vorticity and tropopause folding associated with the upper-level front are present in the model simulations even without latent heating. The low-level southwesterly flow transports warm, moist air from the south as the moisture source for condensation. Latent heating results in an increase in the thickness ahead of the short-wave trough in the upper levels, further deepening of the lee cyclone, and a stronger secondary circulation. The LLJ develops through the Coriolis force acting on the cross-contour ageostrophic winds in response to the increased pressure gradients related to the development of the cyclone and is enhanced by latent heating. The dynamic forcing aloft is also enhanced. Condensation heating and warm advection (evaporative cooling and cold advection) exceed adiabatic cooling (warming) ahead of (behind) the cyclone in the lower troposphere, therefore enhancing the low-level baroclinity. These processes interact nonlinearly leading to the further intensification of the LLJ.

* SOEST Contribution Number 4461.

+ Current affiliation: Mesoscale and Microscale Meteorology Division, National Center for Atmospheric Research, Boulder, Colorado.

Corresponding author address: Dr.Yi-Leng Chen, Department of Meteorology/SOEST, University of Hawaii at Manoa, 2525 Correa Road, HIG 331, Honolulu, HI 96822.

Email: dave@kukui.soest.hawaii.edu

1. Introduction

Considerable research has been done to better understand the impact of latent heat release on western Atlantic and Great Plains extratropical cyclones. Western Atlantic cyclones have attracted considerable attention (Uccellini et al. 1984; Uccellini et al. 1987; Hadlock and Kreitzberg 1988; Whitaker et al. 1988; Kuo and Low-Nam 1990; Lapenta and Seaman 1990, 1992; Kuo et al. 1992; Nieman and Shapiro 1993a,b; Reed et al.1993a,b; Reed et al. 1994), in part due to the availability of the comprehensive dataset, especially the dataset collected during the Genesis of Atlantic Lows Experiment and the Experiment on Rapidly Intensifying Cyclones over the Atlantic. Latent heat release changes the basic structure of extratropical cyclone systems by increasing upper-level temperatures (Anthes et al. 1982), increasing (decreasing) upper-level (lower-level) geopotential heights (Aubert 1957; Danard 1966; Chang et al. 1982), and by strengthening the vertical motions (Krishnamurti 1968; Tracton 1973; DiMego and Bosart 1982; Chen et al. 1983; Pagnotti and Bosart 1984; Smith et al. 1984; Pauley and Smith 1988). Additional simulations indicate that surface energy fluxes also play a significant role in the development of cyclones (Kuo et al. 1991a).

Despite the large number of case studies on western Atlantic cyclones, corresponding studies on east Asian cyclones in the lee side of the Tibetan Plateau are rare, especially during the early summer rainy season over southern China and the Taiwan area (from May to June) (Chen et al. 1991). During this period, heavy rainfall (>100 mm per day) is frequently accompanied by the development and intensification of the southwesterly flow in the lower troposphere (or low-level jet) (Chen and Yu 1988). Observational studies (Chen et al. 1994, hereafter referred to as CCZ; Chen and Chen 1995) on the Taiwan Area Mesoscale Experiment (TAMEX) low-level jet (LLJ) cases showed that before the seasonal change (Chen 1993), the development of the LLJs is closely related to the developing lee cyclone east of the Tibetan Plateau. The secondary circulation across the jet–front system is characterized by a thermally direct circulation with warm air rising within the southwesterly monsoon flow and cold air sinking in the postfrontal northerlies, and by a weak thermally indirect circulation to the south. Chen and Chen (1995) revealed that when an upstream short-wave trough moved eastward from west or northwest of the Tibetan Plateau, the westerly flow strengthened over the plateau, and a lee trough formed in the lee side with closed isobars in the lower levels. The upstream trough moved farther eastward and merged with the lee trough. The superposition of these two troughs led to the initial lee cyclogenesis. Chen and Chen (1995) also pointed out that the relative position of a midlatitude trough to the merged lee trough played an important role in the further development of the lee trough. When the midlatitude trough was located to the north or northeast of the merged trough, the northerly wind component could bring the cold air to the rear of the merged trough.

In this research, we use a mesoscale numerical model to examine the physical mechanisms responsible for the further deepening of the cyclone after the merger of the upstream eastward-moving short-wave trough with the orographically induced lee trough and its relationship with the intensification of the low-level southwesterly flow during the TAMEX IOP (intensive observing period) 5 (1–2 June 1987). Four sensitivity tests were performed to assess the impacts of latent heat release, boundary layer physics, horizontal resolution, and evaporation of precipitation on the development of the lee cyclone and the subsequent intensification of the low-level southwesterly flow (orLLJ). These results will be compared to previous studies on the LLJ after the seasonal transition. Section 2 describes the model and experiment design; section 3 assesses the control simulation and verification of the model results against observations; section 4 presents the results from the sensitivity tests; section 5 diagnoses the effects of latent heat release on the cyclogenesis and the intensification of the low-level southwesterly flow; and section 6 summarizes the results.

2. Model description and experiment design

a. Model description

The numerical model used in this study is the Pennsylvania State University–National Center for Atmospheric Research (PSU–NCAR) Mesoscale Model Version 4 (hereafter referred to as MM4), which is a hydrostatic, three-dimensional, primitive equation model with a terrain-following sigma (σ) vertical coordinate (Phillips 1957). A detailed description of the model’s governing equations, grid structure, and numerics can be found in Anthes et al. (1987). The computational domain of the model consists of 55 × 83 grid points with an 80-km grid spacing (Fig. 1). It has 23 vertical σ levels defined at 0.025, 0.075, 0.125, 0.175, 0.225, 0.275, 0.325, 0.375, 0.425, 0.475, 0.525, 0.585, 0.645, 0.685, 0.725, 0.775, 0.82, 0.865, 0.91, 0.945, 0.97, 0.985, and 0.995. For display purposes, only a portion of the computational grid will be used to show the simulation results (Fig. 1). The model topography is obtained by analyzing the NCAR 30′ latitude and longitude terrain data with the Cressman (1959) objective analysis scheme.

The planetary boundary layer (PBL) processes are parameterized with a multilayer scheme developed by Blackadar (1979) and tested by Zhang and Anthes (1982) that includes surface heat, moisture, and momentum fluxes. A surface energy budget calculation is used to predict ground temperature, whereas ocean temperatures are independent of time. Short- and longwave radiation are included in the surface energy budget. These radiative fluxes depend upon the model-predicted cloud cover in a parameterization developed by Benjamin and Carlson (1986). Latent heating on the resolvable scales is calculated using explicit prognostic equations for water vapor, cloud water, and rainwater (Hsie and Anthes 1984). Subgrid-scale latent heating is parameterized using an Arakawa–Schubert (1974) scheme, as modified by Grell et al. (1991) to include the effects of convective-scale downdrafts.

The initial conditions are obtained by objectively analyzing the rawinsonde and surface observations with successive scans (Cressman 1959) using the European Centre for Medium-Range Weather Forecasts/Tropical Ocean Global Atmosphere global analysis as the first guess. The lateral boundary conditions are provided by linearly interpolating the 12-h objective analyses.

b. Experiment design

For the control simulation (CTRL) we used full-model physical parameterizations to provide a high temporal and spatial resolution dataset to diagnose the development of the cyclone and the evolution of the LLJ event on 1–2 June 1987. The experiment is initialized at 0000 UTC 31 May 1987 about 24 h prior to the formation of the well-defined cyclone, which is the stage 3 classified by Chen and Chen (1995) during which the eastward-moving upstream troughmerged with the orographically induced lee trough. After the merger, the lee cyclone developed rapidly. This time is chosen to initialize the model because we focus our investigation on the formation of the lee cyclone and the subsequent development of the LLJ. A 36-h integration is performed. Model results during the first 24-h simulation were analyzed in detail to diagnose the physical processes during the rapid developing stage.

Four additional sensitivity tests were conducted to isolate the effects of latent heat release, surface fluxes, the horizontal resolution, and precipitation evaporative cooling on the development of the lee cyclone. For the first test (NOLH), the latent heating associated with both resolvable grid-scale precipitation and subgrid-scale parameterization is not included in the thermodynamic equation to assess the effects of latent heat release. For the second test (NOSF), the latent heat release from both the subgrid-scale and resolvable grid-scale precipitation routines is included, but the surface fluxes are excluded in the planetary boundary layer tendency equations. A comparison of the results between NOSF and CTRL allows us to study the effects of surface fluxes. For the third test (CTRL40), the simulation was performed with the same physical parameterization as CTRL, except that the horizontal grid spacing was reduced to 40 km. For the fourth test (NOEVA), the simulation was performed without the effects of precipitation evaporative cooling. As in CTRL, the four sensitivity tests are initialized at 0000 UTC 31 May, and 36-h simulations are conducted.

3. Control simulation descriptions and verifications

a. Verification of the simulated weather patterns

A detailed case description of the 1–2 June 1987 cyclone–LLJ event can be found in CCZ. Comparison of the 12-h simulated 700-hPa geopotential height fields (Fig. 2a) with observed fields (Fig. 3a) shows that the model has predicted the formation of a lee cyclone with closed isobars in agreement with observations. The predicted cyclone is slightly deeper than the observed, and the location is slightly displaced southwestward compared to observations. The low-level southwesterly flow is strengthening (Fig. 2b) in agreement with CCZ’s analysis (Fig. 3b); however, the simulated wind speed is slightly weaker than observed. West of the cyclone center, the predicted east–west geopotential height gradients (Fig. 2a) are larger than observed (Fig. 3a). As a result, the predicted northerly flow west of the cyclone with a speed greater than 15 m s−1 has a larger area than analyzed by CCZ. The forecast upper-level short-wave trough is slightly deeper (Fig. 2c) than observed (Fig. 3c). In the upper level, a well-defined jet streak with an elongated region of maximum winds exceeding 45 m s−1 extended southeastward from 40°N, 110°E to southern Korea (Fig. 2d), in agreement with observations (Fig. 3d). A secondary curved upper-level jet (ULJ) was observed at approximately 30°N and ahead of the trough axis (Figs. 2d and 3d). Ahead of the short-wave trough, the predicted southwesterly flow is weaker, whereas 30°N and near the trough axis, the predicted westerly flow is stronger (Fig. 2d) compared to CCZ’s analysis (Fig. 3d).

The 24-h forecast valid at 0000 UTC 1 June (Fig.4) is compared with the analyzed weather patterns presented by CCZ (Fig. 5). The model reproduces most principal features of this event, including cyclone path, intensification of the LLJ, upper-level short-wave trough, and the associated ULJ and distribution of precipitation. Although the intensity of the simulated lee cyclone in the lower troposphere (Fig. 4a) is in agreement with observation (Fig. 5a), the position of the simulated cyclone center is slightly west of the observed position. An important development in the model at this time is a strong southwesterly flow in the lower troposphere southeast of the cyclone (Fig. 4b). The intensity and axis of the strong southwesterly flow (Fig. 4b) agree well with observation (Fig. 5b); however, the simulated maximum (Fig. 4b) is north of the analyzed maximum and more compact in nature (Fig. 5b). The simulated northerly flow northwest of the cyclone was stronger than observation. The forecasted maximum wind speed there almost reaches 30 m s−1 (Fig. 5c) compared to a wind speed maximum of 20 m s−1 analyzed by CCZ (Fig. 5c). The forecast upper-level short-wave trough tilted north-northwest to south-southeast and is farther south with more of a closed circulation (Fig. 4c) than analyzed by CCZ (Fig. 5c). The simulated ridging (Fig. 4d) downstream of the upper-level short-wave trough is slightly stronger than observation (Fig. 5d). The simulated ULJ associated with the short-wave trough (Fig. 4d) split into two weaker maxima with wind speeds of 40 m s−1 compared to analyzed 60 m s−1.

The simulated relative humidity fields at the 400-hPa level (Fig. 6) shows a comma-shaped maximum along the southeastern China coast, consistent with the observed comma-shaped cloud pattern observed by the infrared satellite imagery (Fig. 19 in CCZ). The predicted 24-h precipitation pattern (Fig. 7) also agrees fairly well with observed rainfall (Fig. 8a in CCZ).

b. Development of the simulated LLJ and ULJ

Since the evolution of the cyclone and the strengthening of the low-level southwesterly flow (or LLJ) and the ULJ were well simulated in the model, we diagnose the model results to assess whether or not the model successfully captures the physical processes responsible for the intensification of both the LLJ and the ULJ as diagnosed by CCZ from observational data.

At the 700-hPa level, the cross-contour ageostrophic winds are pointing toward lower pressure in the regions east and southeast as well as west of the simulated cyclone (Fig. 8a). These areas experienced a significant increase in the speed of the simulated along-contour wind component in a 24-h period (Fig. 8a). In the southwestern quadrant of the cyclone, where the cross-contour ageostrophic winds are pointing toward higher pressure, the speed of the simulated along-contour wind component decreases. The strengthening or weakening of the simulated along-contour winds are apparently related to the Coriolis force acting on the cross-contour ageostrophic winds. The pattern of the isallohypsic winds (Fig.8b) is similar to that of the cross-contour ageostrophic winds, except in the jet exit and entrance regions where the cross-contour ageostrophic winds caused by the inertial-advection term due to along-stream changes in wind speed (Uccellini and Johnson 1979) are significant. Strong northeasterly cross-contour ageostrophic winds exist in the exit region of the northerly jet behind the 700-hPa trough (Fig. 8a), compared to the much weaker westerly isallohypsic winds (Fig. 8b). Regions where isallohypsic winds are pointing toward lower (higher) pressure experience an increase (decrease) in the geostrophic winds. The pattern of the 24-h differences in the geostrophic wind speed (Fig. 8b) is similar to that of the 24-h differences in the simulated along-contour wind component (Fig. 8a). These features are essentially similar to the results diagnosed from the observational data in CCZ and suggest that the intensification of the LLJ is closely related to the local change in geopotential gradients as the lee cyclone deepens and moves eastward. It appears that results from our numerical experiment have reproduced the basic physical processes for the intensification of the low-level southwesterly flow (or LLJ) diagnosed by CCZ from observational data.

Similar results are found for the 250-hPa level (Figs. 8c and d), except that the 24-h increase in the speed of the ULJ in the direction parallel to geopotential height contours (Fig. 8c) is only approximately one-half of the geostrophic value (Fig. 8d), whereas at the 700-hPa level, the 24-h increase in the speed of the LLJ (Fig. 8a) in the direction parallel to geopotential height contours is comparable to that of the geostrophic winds (Fig. 8b). These results are also in agreement with CCZ. The subgeostrophic flow in the 250-hPa level is apparently caused by the inertial-advective term due to curvature effect (Shapiro and Kennedy 1981). Notable discrepancy between Figs. 8c and 8d occurs in the entrance region of the ULJ where cross-contour southerly ageostrophic winds are in excess of 10 m s−1 (Fig. 8c) compared to weak westerly isallohypsic winds (Fig. 8d). The cross-contour ageostrophic winds in the entrance region of the ULJ are even stronger in CCZ’s analysis (their Fig. 11e) than in CTRL with a maximum speed close to 20 m s−1. Note that, in the jet entrance region, the along-stream acceleration is greater in CCZ’s analysis (Fig. 5d) than in CTRL (Fig. 4d).

4. Sensitivity tests

a. No latent heating

The effects of latent heating on the development of the LLJ are studied by removing the latent heat release associated with both the subgrid-scale and grid-scale precipitation from the thermodynamic equation in the NOLH case. Figure 9 shows the results of the 24-h forecast at both the 700- and the 250-hPa levels. It is evident that latent heating has a significant impact on the development of the cyclone in the lower troposphere. Without latent heating, the lee cyclone in the lower troposphere is weak (Fig. 9a). The maximum speed of the southwesterly flow ahead of the trough is about 10 m s−1 (Fig. 9b), whichis less than one-half of that in the CTRL case (Fig. 4a). In the upper levels, a comparison of Figs. 4c and 9c shows a difference in the shape of the trough because of the enhanced downstream ridging in the CTRL case. The 10 560-m geopotential height contour is located farther south in the NOLH case than in the CTRL case with a stronger ULJ. The beginnings of a closed circulation are apparent in the 250-hPa level wind vectors within the trough in the CTRL case (Fig. 4c) but not in the NOLH case (Fig. 9c). The splitting of the ULJ into two weaker maxima is evident in the CTRL case but not in the NOLH case (Fig. 4d). Although weaker and displaced to the south as compared with the verifying analysis, the ULJ in NOLH (Fig. 9d) resembles the observed ULJ (Fig. 5d) with a well-defined jet core. In the CTRL case, a second jet is observed near the northern edge of the display domain in the verifying analysis but is essentially absent in the NOLH case. These results suggest a strong dependence of the upper-level flow on latent heating.

b. No surface fluxes

Kuo et al. (1991a) found that surface energy fluxes played a significant role for the development of seven marine cyclones in the western Atlantic. To study the role of surface fluxes on the development of this lee cyclone, we performed experiment NOSF by removing surface fluxes in the planetary boundary layer tendency equations. We found that the removal of the surface fluxes has minor impact on the development of this lee cyclone over land (Fig. 10). The model successfully reproduces the formation of a closed lee cyclone in the lower troposphere (Fig. 10a), with its location and intensity compared favorably with those in CTRL. The intensity and orientation of the low-level southwesterly flow (or LLJ) (Fig. 10b) are well simulated; however, the low-level northerly flow behind the cyclone (Fig. 10b) is slightly weaker than that in CTRL (Fig. 4c). This test demonstrates that the intensification of the low-level southwesterly flow is not significantly affected by the surface heat, moisture, and momentum fluxes. Therefore, the evolution of the surface fluxes during the diurnal heating cycle is not essential for the intensification of the low-level southwesterly flow. NOSF experiment also reproduces the deepening of the upper-level short-wave trough (Fig. 10c) and the intensification of the ULJ (Fig. 10d).

c. Higher horizontal resolution

Kuo et al. (1995) discussed the possible effects of resolution on the simulation of a rapid continental mesoscale cyclogenesis. It was stated that the resolution can affect the prediction in at least two ways. The first is the representation of the cyclone on the model grid. The second is that the performance (or the functioning) of a precipitation scheme can be highly grid-size dependent. To test the impact of the grid resolution on this particular case, we performed experiment CTRL40 with identical physics to CTRL, except that the horizontal resolution was increased to 40 km.

At the 700-hPa level, the 24-h simulation shows that the cyclone was about 30 m deeper in CTRL40 (Fig. 11a) than in CTRL (Fig. 4a), whereas the position of the cyclone center is almost the same in these two simulations. The simulatedsouthwesterly (northerly) flow ahead of (behind) the cyclone center is slightly stronger in CTRL40 (Fig. 11b) than in CTRL (Fig. 4b). At the 250-hPa level, the ridging downstream of the upper-level trough is stronger in CTRL40 (Fig. 11c) than in CTRL (Fig. 4c), and the northeastern portion of the ULJ in CTRL40 (Fig. 11d) is weaker with a speed maximum less than 35 m s−1, compared to 40 m s−1 in CTRL (Fig. 4d). For the observed ULJ core around 30°N, 111°E (Fig. 5d), CTRL40 shows slight improvement (Fig. 11d) compared to CTRL (Fig. 4d). Despite these differences, the overall results for the cyclone development and the intensification of the LLJ in the lower troposphere are similar in these two simulations.

To investigate the effect of reduced vertical resolution on the model prediction, we also performed two additional simulations with 16 and 19 sigma levels with 80-km horizontal resolution (not shown). CTRL showed slightly better results than these two simulations. These results are in agreement with Kuo and Low-Nam (1990), who found that increasing the vertical resolution from 15 to 23 levels in their model with a 80-km grid spacing has little impact on their prediction of nine explosive cyclones over the western Atlantic Ocean.

Pecnick and Keyser (1989) and Persson and Warner (1991) examined the consistency conditions for vertical and horizontal resolution for depicting baroclinic zone evolution and showed that spurious gravity waves were produced when optimal ratios were not used, especially for the case when vertical resolution was inadequate. Houghton et al. (1993) made a series of forecasts from different resolution versions of the National Meteorological Center’s (renamed the National Centers for Environmental Prediction) Nested Grid Model to assess an optimal ratio of model vertical and horizontal resolutions. They found that with 40-km resolution, 16 layers in the vertical may not be sufficient. They also suggested that a model formulation with insufficient vertical resolution would be susceptible to spurious oscillations. To examine the vertical and horizontal consistency with higher horizontal resolution, another experiment was performed with 40-km resolution and 28 (CTRL40-28) sigma levels at 0.025, 0.07, 0.11, 0.15, 0.19, 0.23, 0.27, 0.31, 0.35, 0.39, 0.43, 0.39, 0.43, 0.51, 0.55, 0.59, 0.63, 0.67, 0.71, 0.75, 0.79, 0.83, 0.87, 0.91, 0.945, 0.97, 0.985, and 0.995. At the lower troposphere, the simulated cyclone is deeper (Fig. 12a) than the observed (Fig. 5a). Both the LLJ on the southeastern flank of the cyclone center and the northerly jet west of the cyclone center (Fig. 12b) are also stronger than analyzed by CCZ (Fig. 5b). In the upper level, the simulated upper-level trough (Fig. 12c) is sharper than CTRL (Fig. 4c) and than the observed (Fig. 5c). Compared to CTRL, notable improvement in the prediction of the ULJ with a jet core at 29°N, 111°E is evident (Figs. 12d, 4d, and 5d). It is apparent that improvement in the upper-level simulations cannot be achieved by higher horizontal resolution alone. Adequate verticalresolution in the upper levels is also needed. For this case, vertical resolution enhancement has a seemingly positive impact on the structure of the ULJ.

d. No precipitation evaporative cooling

To demonstrate the effects of precipitation evaporative cooling, the no evaporative cooling simulation (NOEVA) was performed by turning off the precipitation evaporative cooling. At the 700-hPa level, the simulated cyclone in NOEVA (Fig. 13a) is about 10 m weaker than in CTRL (Fig. 4a). The simulated southwesterly wind maximum ahead of the cyclone center (Fig. 13b) is approximately 3.0 m s−1 weaker than in CTRL (Fig. 4b). The simulated northerly wind maximum (Fig. 13b) is weaker than that in CTRL (Fig. 4b) by about 5.0 m s−1. As noted earlier, the northerly wind in CTRL is too strong compared to the analysis. These results suggest that the effects of the evaporative cooling may be overestimated in CTRL. Figure 13c shows the 850–500-hPa thickness difference between CTRL and NOEVA. It is apparent that evaporative cooling occurred northwest of the cyclone center (Fig. 13c) and cools the lower troposphere significantly. Southeast of the cyclone center, the effects of evaporative cooling are negligible. These results suggest that the evaporative cooling increases the low-level thermal contrast. At the 250-hPa level, evaporative cooling has no effect on the simulated fields (not shown) because of the fact that the saturated mixing ratio is very small in the upper levels.

5. Effects of latent heating

It has been demonstrated that the removal of latent heating associated with precipitation from the thermodynamic equation in NOLH produces a dramatic impact on the development of the cyclone in the lower troposphere. In this section we will focus on the comparison between CTRL and NOLH to assess the effects of latent heating on the cyclogenesis and the intensification of the LLJ.

a. Potential vorticity

A comparison of the evolution of potential vorticity (PV) at the 250-hPa level between CTRL and NOLH is made to assess the effects of latent heating on tropopause folding and the upper-level front. Advection of positive PV anomaly in the upper levels can induce upward cyclonic circulation at the surface (Hoskins et al. 1985). The PV was computed from the model output on isobaric surfaces according to the expression (Hoskins et al. 1985; Bosart and Bartlo 1991)
i1520-0493-125-10-2583-e1
where f is the Coriolis parameter, Vh is the horizontal wind velocity, ζp is the relative vorticity on an isobaric surface, and p is the two-dimensional gradient operator on the isobaric surface.

Initially, the upper-level PV maximum associated with the deepening upper-level short-wave trough is located at 40°N, 100°E (Fig. 14a). In CTRL, this PV maximum moves southeastward (Fig. 14b). The PV maximum also moves southeastward in NOLH (Fig. 14c) with its intensity and path comparable to those in CTRL. The PV anomaly concentrates in a smaller region in CTRL than in NOLH with a narrower axis extending northeastward. The PV gradients east and northeast of the PV maximum are larger in CTRL than in NOLH. The advection of positive PV anomaly northeast of the PV maximum in CTRL(Figs. 4c and 14b) is more significant than in NOLH (Figs. 9c and 14c). As a result, the cyclonic circulation in the lower troposphere northeast of the PV maximum is more pronounced in CTRL (Fig. 4a) than in NOLH (Fig. 9a). To investigate the tropopause folding, the distributions of potential vorticity and potential temperature for the cross section along the line AA′ in Fig. 14b are presented. This cross section is close to the PV maximum at the 250-hPa level. Tropopause folding associated with the upper-level frontogenesis is clearly shown by the PV cross sections even without latent heating (Fig. 15). Relatively high PV extends downward to the 500-hPa level in both CTRL (Fig. 15a) and NOLH (Fig. 15b). However, high PV extends farther downward, to below the 500-hPa level, in CTRL than in NOLH. Appreciable potential temperature gradients in the upper-level frontal zone are evident in both CTRL (Fig. 15a) and NOLH (Fig. 15b). Nevertheless, the potential temperature gradients across the upper-level front are slightly larger in CTRL than in NOLH. Ahead of the upper-level front, the air is slightly warmer in CTRL than in NOLH because of latent heat release. Behind the upper-level front, the thermal structure is nearly identical between CTRL and NOLH. It is apparent that upper-level dynamic forcing exists in both CTRL and NOLH, and this dynamic forcing is more significant in CTRL than in NOLH.

b. Thermal and dynamic field differences between CTRL and NOLH

To better understand the significance of latent heating on the cyclogenesis and the intensification of the low-level southwesterly flow (or LLJ), the differences in the thermal and dynamic fields between CTRL and NOLH are investigated in this section.

One of the direct influences of the diabatic heating in the upper levels is the increase in thickness ahead of the upper-level trough in CTRL (Fig. 16a). The difference in the 250-hPa winds between CTRL and NOLH shows anticyclonic winds ahead of the trough with higher geopotential heights in CTRL than in NOLH (Fig. 16b). The most significant differences occur along 110°E north of 33°N where the differences in wind vectors have a maximum magnitude close to 40 m s−1. In this area winds are southerlies in CTRL (Fig. 4c) in agreement with observations (Fig. 5c), whereas in NOLH winds are northwesterlies (Fig. 9c). The CTRL trough is sharper than the NOLH trough and exhibits the beginnings of a closed circulation (Figs. 4c and 9c). As a result, southerlies ahead of the trough line extend farther west in CTRL. These differences are caused by the increase in the geopotential height ahead of the trough in CTRL (Fig. 16b) as a result of convective heating. The differences in wind vectors southeast of the maximum differences in geopotential height are northeasterlies (Fig. 16b). These differences result in slightly northwestward displacement of the simulated ULJ axis in CTRL with slightly weaker wind speeds (Fig. 4c) compared to NOLH (Fig. 9c). Immediately northeast of the maximum differences ingeopotential height, the increase in the thickness ahead of the trough also results in stronger westerlies at 26°N, 120°E in CTRL than in NOLH (Fig. 16b). Consequently, the simulated southwesterlies there are stronger in CTRL than in NOLH (Figs. 4d and 9d). Large north–south geopotential height gradients were observed north of the enhanced downstream ridging with a westerly jet north of 40°N in CTRL (Fig. 4d). These features are caused by higher geopotential height ahead of the trough in CTRL than in NOLH (Fig. 16b) as a result of convective heating.

At the 250-hPa level, the ageostrophic winds computed from the inertial-advection effects display anticyclonic flow around the upper-level trough with maximum values near the trough axis and southwesterlies northeast of the trough axis in both CTRL and NOLH (Figs. 16c and 16d). These patterns are remarkably similar to the ageostrophic winds of curved flow for a constant-velocity jet stream embedded within a stationary wave (Shapiro and Kennedy 1981). This flow pattern will give rise to upper-level divergence ahead of the trough in both CTRL and NOLH. In other words, the rising motion ahead of the trough is associated with the upper-level baroclinic wave. The increase in thickness ahead of the trough in CTRL enhances the diffluence of the height contours and modifies the changes in curvature along the flow (Figs. 16c and 16d). The southwesterly ageostrophic winds northeast of the inflection point ahead of the trough and the anticyclonic ageostrophic winds near the trough axis due to inertial advection term are stronger in CTRL than in NOLH (Figs. 16c and 16d) because of the curvature effect. As a result, the upper-level divergence (and hence the rising motion) would be stronger ahead of the upper-level trough around 34°N, 115°E in CTRL than in NOLH. The upward motion there is driven dynamically and is enhanced by latent heat release. From the distributions of both the upper-level PV maximum and the inertial advection term of the ageostrophic winds ahead of the upper-level trough, it is apparent that the dynamic forcing aloft is present in both CTRL and NOLH. It provides the upper-level support needed for the spinup for the lee cyclone and the subsequent development of the LLJ in the lower troposphere. This is one of the main differences between this study and other LLJ studies over southern China (Chen and Chang 1980; Chou et al. 1990; Hsu and Sun 1994; and others). Based on numerical experiments for the LLJ case during 15–16 May 1987, Hsu and Sun (1994) suggested that latent heating associated with shallow stratiform clouds with tops at 5 km is important for the development of LLJ. For our case, the observed vertical profile of apparent heat source (Q1) (Yanai et al. 1973) averaged over the box in Fig. 4b at 0000 UTC 1 June show a pronounced Q1 peak in the upper troposphere (Fig. 17), located at about 350 hPa. It is apparent that for this case most of the nonconvective rainfall is from stratiform clouds associated with deep convection rather than shallow stratiform clouds.

Figure 18a shows the 850–500-hPa thickness difference between CTRL and NOLH. To the southeast (northwest) of the cyclone center, it is warmer (colder) in CTRL than in NOLH. At the 700-hPa level, the southerly (northerly) wind component ahead of (behind) the cyclone center wasstronger in CTRL (Fig. 4a) than in NOLH (Fig. 9a). It appears that the combined effects of condensation heating and warm advection (evaporative cooling and cold advection) ahead of (behind) the cyclone center in CTRL exceeded adiabatic cooling (warming), resulting in warmer (colder) temperature southeast (northwest) of the cyclone in CTRL than in NOLH, and increased the low-level baroclinity. Nagata and Ogura (1991) also pointed out that evaporative cooling beneath stratiform clouds generated a marked frontogenetical forcing and created a cold pool beneath the sloping frontal surface. The differences in the 700-hPa winds between CTRL and NOLH show cyclonic winds near the cyclone center with lower geopotential in CTRL than in NOLH (Fig. 18b).

At 700 hPa, as found by CCZ from the observational data, the pattern of the simulated ageostrophic wind due to the local change of wind (Fig. 18c) is remarkably similar to that of the simulated cross-contour component of the ageostrophic wind field (Fig. 8a), except that the magnitudes of the cross-contour ageostrophic winds are generally larger. The simulated cross-contour ageostrophic winds are primarily related to local changes in the geopotential gradients. In response to the increase in pressure gradients associated with the development and eastward movement of the cyclone, the southwesterly flow ahead of the cyclone strengthens through the Coriolis force acting on the cross-contour ageostrophic winds (see CCZ). In NOLH, without the development of the cyclone, the ageostrophic wind due to local change of wind is very weak and has a weak component crossing isobars ahead of the 700-hPa trough (Fig. 18d). These results show that, with the dynamic forcing aloft, latent heat release played an important role for the deepening of the cyclone and the subsequent strengthening of the low-level southwesterly flow.

c. Secondary circulation

To show the effects of latent heating in strengthening the secondary circulation across the jet–front system, we compare the implicit vertical motion inferred from the quasigeostrophic theory between CTRL and NOLH and present the differences in horizontal winds and vertical motion along the vertical cross sections between CTRL and NOLH.

Charts of the 850–500-hPa (500–200-hPa) thickness were superposed with the 700-hPa (300-hPa) relative vorticity (Fig. 19) to diagnose the distribution of vertical motion in the lower (upper) troposphere implicitly from the quasigeostrophic theory (Sutcliffe 1947; Trenberth 1978). Whenever there is cyclonic (anticyclonic) advection of vorticity by the thermal wind, we can infer ascent (descent). In CTRL, the positive vorticity advection by the thermal wind is evident ahead of the cyclone center in the lower troposphere (Fig. 19a), indicating rising motion there. In NOLH, the thermal wind advection of relative vorticity is insignificant in the lower troposphere (Fig. 19b). In the upper troposphere, without diabatic heating, the thermal wind advection of relative vorticity is evident ahead of the trough (Fig. 19d), suggesting the presence of dynamic forcing associated with the baroclinic wave. With the diabatic heating, the thermal wind advection of relative vorticity is more significant (Fig. 19c).

The distribution of the vertical motion (p-velocity ω) is computed from the kinematic method by the boundary condition at the surface,
ωρgVsh,
where h is the surface topography, ρ the density of the air, and Vs the horizontal wind velocity at the surface. The O’Brien (1970) linear correction method was used to adjust the divergent wind to ensure that the vertical motion vanished at the 100-hPa level. In CTRL, the horizontal distribution of ω at the 700-hPa level (Fig. 20a) reveals rising (sinking) motion ahead of (behind) the cyclone. In the upper level, rising motion is also observed ahead of the trough axis (Fig. 20c). These results are in agreement with the rising motion inferred from the quasigeostrophic theory. The vertical motion implied by the quasigeostrophic theory (Fig. 19) would also lead to a nearly symmetric couplet of upward and downward motion. Nevertheless, Figure 20c shows a prominent upward motion east of the maximum quasigeostrophic forcing and only weak downward motion west of the trough axis. This suggests that terms ignored in the quasigeostrophic theory, such as latent heating and ageostrophic advections, are also significant. In NOLH, the vertical motion (Figs. 20b and 20d) is much weaker than in CTRL.

Warm (cold) air is observed ahead of (behind) the cyclone (Fig. 19) associated with rising (sinking) motion (Fig. 20) suggesting the conversion from the available potential energy to kinetic energy. It is apparent that diabatic heating enhanced the baroclinic conversion in the lower (upper) troposphere by increasing the 850–500-hPa (500–200-hPa) thickness (Fig. 19) and the upward vertical motion (Fig. 20) above (ahead of) the lower-level cyclone center (upper-level trough). Figure 21 shows that the baroclinic conversion term is significantly enhanced in CTRL over NOLH, especially in the upper levels.

The differences along the vertical cross sections between CTRL and NOLH show that the secondary circulation across the jet–front system is much stronger and concentrates in a narrower region in the frontal zone in CTRL than in NOLH (Fig. 22a). It is apparent that latent heat release feeds back to the large-scale flow and strengthens the secondary circulation considerably throughout the entire troposphere with stronger rising (sinking) motion ahead of (behind) the frontal zone in CTRL than in NOLH. The thermally indirect circulation to the south (Fig. 22 in CCZ) is also stronger in CTRL than in NOLH.

The intensity of southwesterly (northeasterly) flow ahead of (behind) the cyclone center is much stronger in CTRL than in NOLH (Fig. 22b). In the lower troposphere and ahead of the frontal zone, the maximum difference in the wind component normal to the cross-sectional plane (or parallel to the LLJ axis) between CTRL and NOLH occurs at the 850-hPa level (Fig. 22b). This is the level where the strengthening of the lower branch of the secondary circulation is most significant in CTRL compared to NOLH (Fig. 22a). The strengthening of the LLJ in CTRL is caused by the Coriolis force acting on the cross-contour ageostrophic winds in the lower troposphere. As discussed earlier, the cross-contour ageostrophic winds near the jet core in the lower troposphere are related to isallohypsic winds (see CCZ) and are enhanced by latent heating. Theisallohypsic winds in the lower troposphere are related to the deepening of the lee cyclone in response to the dynamic forcing aloft as the upper-level short-wave trough approaches.

In the upper troposphere, southwesterly (northerly) flow ahead of (behind) the upper-level trough is weaker in CTRL than in NOLH (Fig. 22b). It is interesting to note that with latent heating, weakening of the upper-level flow northwest (southeast) of the maximum rising motion is most significant at the level where the strengthening of the upper-level return branch of the thermal direct (indirect) circulation is the strongest (Fig. 22a). At the approximate location of the upper-level jet core, which is approximately at the 250-hPa level and just south of the upper-level front in the cross section, the differences in the wind component normal to the cross-sectional plane between CTRL and NOLH are rather small (Fig. 22b).

In a numerical study of the QE II storm, Kuo et al. (1991b) found that in the course of rapid intensification of the storm, diabatic heating associated with latent heat release was the dominant forcing mechanism for the vertical motion of the simulated storm, whereas the diabatic heating also significantly modified the baroclinic structure of the storm to reinforce its adiabatic vertical motion. They also suggested that extratropical cyclogenesis should be viewed in the context of moist baroclinic instability with nonlinear interactions between the baroclinic dynamics and diabatic processes. The results on the TAMEX IOP 5 cyclone–LLJ event suggest that the development of the lee cyclone and the subsequent intensification of the low-level southwesterly flow over southern China prior to the seasonal transition (Chen 1993) is related to moist baroclinic dynamics, rather than a CISK (conditional instability of the second kind) process as suggested for late-season cases over southern China (Chen and Chang 1980) and Mei-Yu systems over the Yangtze River valley (Chen and Dell’Osso 1984).

6. Summary and conclusions

A numerical investigation has been performed with MM4 to study the 1–2 June 1987 cyclone–LLJ event. A comparison between the CTRL and observations demonstrates that the model has successfully reproduced the major features associated with this case, including cyclone path, intensification of the low-level southwesterly flow, upper-level short-wave trough and the associated ULJ, distribution of precipitation, and the secondary circulation across the jet–front system. Sensitivity tests show that latent heating has a significant impact on the cyclogenesis and the intensification of the southwesterly flow in the lower troposphere. The effects of surface fluxes on the cyclogenesis and the intensification of the LLJ are not important.

The spinup of the lee cyclone in the lower troposphere is caused by the arrival of the upper-level short-wave trough, the low-level cold-air advection behind a midlatitude trough (Chen et al. 1994; Chen and Chen 1995), and the positive feedback from latent heat release. The tropopause folding and dynamically driven rising motion ahead of the short-wave trough are present with and without the effects of latent heating. The lee trough provides the initial low-level vorticity. The initial deepening of the lee cyclone and the development of the low-level southwesterly flow are caused by the vertical motion associated with the upper-level short-wave trough. The southwesterly flow transports warm, moist air from the south as the moisture source for condensation. The LLJ ahead of the cyclone center strengthens through the Coriolis force acting on the cross-contour ageostrophic winds inresponse to the increase in pressure gradients and is enhanced by convective heating. Latent heating results in an increase in the thickness in the lower (upper) troposphere above (ahead of) the low-level cyclone (upper-level trough), an enhanced low-level cyclonic circulation. The dynamic forcing aloft is also enhanced. The combined effect of condensation heating and warm advection (evaporative cooling and cold advection) ahead of (behind) the cyclone center exceeds adiabatic cooling (warming) and increases the low-level thermal contrast. The secondary circulation is strengthened and the baroclinic conversion is also enhanced. These processes interact nonlinearly leading to the further deepening of the cyclone and further intensification of the LLJ.

Acknowledgments

We wish to thank G. Barnes and D. E. Stevens for their valuable discussions and H.-C. Yeh for his assistance. We are also grateful to anonymous reviewers for their comments and suggestions. This work was supported by the National Science Foundation under Grant ATM-9206124 and ATM-9421060. Part of the computing resources was supported by the Scientific Computing Division of the National Center for the Atmospheric Research, which is sponsored by the National Science Foundation.

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Fig. 1.
Fig. 1.

Computational domain and model terrain (m). Contour interval is 1000 m. The inner box denotes the subdomain over which the model results are presented.

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2583:ANSOTL>2.0.CO;2

Fig. 2.
Fig. 2.

Simulated geopotential height and wind field in CTRL at the 12-h forecast time (valid at 1200 UTC 31 May). Winds (m s−1) with one pennant, full barb, and half-barb denoting 25, 5, and 2.5 m s−1, respectively: (a) 700 hPa, geopotential heights (solid) every 30 m; (b) 700 hPa, isotach every 2.5 m s−1 starting from 12.5 m s−1; (c) as in (a) except for 250 hPa with geopotential heights every 120 m; and (d) 250 hPa with isotach every 5 m s−1 starting from 30 m s−1.

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2583:ANSOTL>2.0.CO;2

Fig. 3.
Fig. 3.

Same as Fig. 2 but for objectively analyzed geopotential height and wind field for 1200 UTC 31 May (after Chen et al. 1994).

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2583:ANSOTL>2.0.CO;2

Fig. 4.
Fig. 4.

Same as Fig. 2 but for the 24-h forecast time (valid at 0000 UTC 1 June).

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2583:ANSOTL>2.0.CO;2

Fig. 5.
Fig. 5.

Same as Fig. 3 but for 0000 UTC 1 June (after Chen et al. 1994).

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2583:ANSOTL>2.0.CO;2

Fig. 6.
Fig. 6.

The simulated relative humidity at the 400-hPa level in CTRL at the 24-h forecast time valid at 0000 UTC 1 June. The interval is 10%.

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2583:ANSOTL>2.0.CO;2

Fig. 7.
Fig. 7.

The simulated 24-h accumulated precipitation (mm) in CTRL ending at 0000 UTC 1 June. Isopleths are 1, 10, and 25mm. The interval is 25 mm for values greater than 25 mm.

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2583:ANSOTL>2.0.CO;2

Fig. 8.
Fig. 8.

The 700-hPa analysis for CTRL at the 24-h forecast time (valid at 0000 UTC 1 June): (a) 24-h differences in the along-contour component of the simulated winds (isolines) and the cross-contour component of the ageostrophic winds (arrows); (b) 24-h differences in the simulated geostrophic winds (isolines) and the isallohypsic winds (arrows). The contour intervals are 4 m s−1. Winds (m s−1) with full barb and half-barb representing 5 and 2.5 m s−1, respectively. The 250-hPa analysis for CTRL: (c) as in (a) with contour intervals in 6 m s−1; (d) as in (b) with contour intervals in 6 m s−1.

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2583:ANSOTL>2.0.CO;2

Fig. 9.
Fig. 9.

Simulated geopotential height and wind field in NOLH at the 24-h forecast time (valid at 0000 UTC 1 June). Wind barbs (m s−1) with one pennant, full barb, and half-barb denoting 25, 5, and 2.5 m s−1, respectively: (a) 700 hPa, geopotential heights (solid) every 30 m; (b) 700 hPa, isotach every 2.5 m s−1 starting from 5 m s−1; (c) as in (a) except for 250 hPa with geopotential heights every 120 m; and (d) 250 hPa, isotach every 5 m s−1 starting from 30 m s−1.

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Fig. 10.
Fig. 10.

As in Fig. 9 except for NOSF.

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2583:ANSOTL>2.0.CO;2

Fig. 11.
Fig. 11.

As in Fig. 9 except for CTRL40.

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Fig. 12.
Fig. 12.

As in Fig. 9 except for CTRL40-28.

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2583:ANSOTL>2.0.CO;2

Fig. 13.
Fig. 13.

Simulated geopotential height and wind field at the 700-hPa level in NOEVA at the 24-h forecast time (valid at 0000 UTC 1 June). (a) Geopotential heights (solid) every 30 m. Winds (m s−1) with one pennant, full barb, and half-barb denoting 25, 5, and 2.5 m s−1, respectively; (b) isotach every 2.5 m s−1 starting from 12.5 m s−1 at the 700-hPa level; (c) the 850–500-hPa thickness difference between CTRL and NOEVA (CTRL minus NOEVA).

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2583:ANSOTL>2.0.CO;2

Fig. 14.
Fig. 14.

Potential vorticity at the 250-hPa level with interval every 0.8 PVU (1 PVU = 10−6 m2 K kg−1 s−1) for (a) initial condition (valid at 0000 UTC 1 June); (b) 24-h forecast in CTRL; and (c) 24-h forecast in NOLH.

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2583:ANSOTL>2.0.CO;2

Fig. 15.
Fig. 15.

Vertical cross sections of PV (solid lines, every 0.8 PVU) and potential temperature (dashed lines, every 4 K) along the line AA′ in Fig. 14b for (a) 24-h forecast in CTRL and (b) 24-h forecast in NOLH.

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2583:ANSOTL>2.0.CO;2

Fig. 16.
Fig. 16.

Differences between CTRL and NOLH (CTRL minus NOLH) at the 24-h forecast time (valid at 0000 UTC 1 June): (a) Thickness between 500- and 250-hPa levels (every 20 m); (b) geopotential height (every 20 m) and vector winds at the 250-hPa level. 250-hPa analysis (solid line is geopotential heights in 120 m): (c) ageostrophic winds due to inertial advection term in CTRL; (d) as in (c) except for NOLH.

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2583:ANSOTL>2.0.CO;2

Fig. 17.
Fig. 17.

Observed area-averaged (over the box in Fig. 4b) vertical profile of apparent heat source (Q1) at 0000 UTC 1June.

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2583:ANSOTL>2.0.CO;2

Fig. 18.
Fig. 18.

Analysis at 24-h forecast time (valid at 0000 UTC 1 June). (a) As in Fig. 16a except for thickness between 850- and 500-hPa levels (every 10 m); (b) as in Fig. 16b except for the 700-hPa level. 700-hPa analysis (solid lines are geopotential heights every 30 m): (c) ageostrophic wind due to local change of wind in CTRL; (d) as in (c) except for NOLH.

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2583:ANSOTL>2.0.CO;2

Fig. 19.
Fig. 19.

Analysis at the 24-h forecast time (valid at 0000 UTC 1 June): 700-hPa relative vorticity (thin lines every 4 × 10−5 s−1) and 850–500-hPa thickness (solid every 30 m) in (a) CTRL; (b) NOLH; and 300-hPa relative vorticity and 500–250-hPa thickness (every 60 m); (c) CTRL; (d) NOLH.

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Fig. 20.
Fig. 20.

The vertical p-velocity contours (every 2 μb s−1) at the 24-h forecast time (valid at 0000 UTC 1 June) for 700-hPa (a) CTRL; (b) NOLH; and for 250-hPa (c) CTRL; (d) CTRL.

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2583:ANSOTL>2.0.CO;2

Fig. 21.
Fig. 21.

Area-averaged (over the box in Fig. 4b) vertical distribution of baroclinic conversion term −ωα at the 24-h forecast time (valid at 0000 UTC 1 June) (ω is p-velocity, α is specific volume).

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2583:ANSOTL>2.0.CO;2

Fig. 22.
Fig. 22.

Vertical cross-sectional differences between CTRL and NOLH (CTRL minus NOLH) along the line AA′ in Fig. 14b at the 24-h forecast time (valid at 0000 UTC 1 June). (a) Secondary circulation. Ageostrophic wind differences (m s−1, horizontal component of arrow) and minus omegas (μb s−1, vertical component of arrow). The maximum ageostrophic wind difference along the cross section is 35 m s−1. The contour interval for minus omega is 2 μb s−1; (b) Wind speed normal to the cross-sectional plane (every 2.5 m s−1).

Citation: Monthly Weather Review 125, 10; 10.1175/1520-0493(1997)125<2583:ANSOTL>2.0.CO;2

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