Formation and Maintenance of a Long-Lived Taiwan Rainband during 1 3 March 2003

Transcription

1 APRIL 2017 Y U A N D L I N 1211 Formation and Maintenance of a Long-Lived Taiwan Rainband during 1 3 March 2003 CHENG-KU YU Department of Atmospheric Sciences, National Taiwan University, Taipei, Taiwan CHE-YU LIN Department of Atmospheric Sciences, National Taiwan University, and Department of Atmospheric Sciences, Chinese Culture University, Taipei, Taiwan (Manuscript received 3 October 2016, in final form 26 December 2016) ABSTRACT Taiwan rainbands (TRs), defined here as convective lines, which form off the mountainous eastern coast of Taiwan under weakly synoptically forced weather conditions, are a well-known mesoscale phenomenon, but their formative processes remain the subject of debate. This study uses surface and radar observations within the coastal zone of eastern Taiwan and NCEP reanalysis data to document a long-lived TR with a lifetime of ;36 h during 1 3 March 2003 to advance the current general understanding of mechanisms responsible for the TR s formation and maintenance. Detailed analyses indicate that the rainband was initiated by convergence that was produced as low-level environmental northeasterly/easterly onshore flow encountered topographically blocked northerlies that developed nearshore. The northerly blocked flow was observed to weaken and subsequently dissipate because of changing synoptic pressure patterns that caused prevailing southeasterlies/southerlies at low levels. However, colder nearshore air that resulted from the combined effects of orographic blocking, the evaporation of the TR s precipitation, and radiative cooling over coastal land continued to persist and acted to provide a continuing source of lifting for the subsequent maintenance of moist convection. Temporal variations in the precipitation intensity of the studied TR were also shown to be consistent with the theoretical prediction of the interaction between the cold pool and ambient vertical shear. This study suggests that multiple precipitation mechanisms, which involve interactions of diurnally, topographically, and convectively generated circulations along the mountainous coast, may operate and contribute to the longevity of a TR event under suitable circumstances, such as the rapidly evolving synoptic flow observed in the present case. 1. Introduction A wide variety of convective forcings may occur along a coast with significant topography. According to investigations of moist convection off mountainous coasts in different geographical regions, combinations of topographic effects, diurnally induced circulations and waves, and discontinuities in surface roughness between the land and ocean probably determine the nature of coastal convection (Garrett 1980; Houze et al. 1981; Grossman and Durran 1984; Liberti et al. 2001; Mapes et al. 2003; Colle and Yuter 2007; Hassim et al. 2016; Vincent and Lane 2016). Documenting these processes is particularly challenging because of Corresponding author Professor Cheng-Ku Yu, yuku@ ntu.edu.tw complicated interactions between multiple mesoscale circulations. Taiwan is a mountainous island located at ;23.58Nin the northwestern Pacific Ocean. The mountains along eastern Taiwan are extremely steep, with terrain rising sharply to more than 1500 m MSL within km of the shore. The occurrence of moist convection along the mountainous coast of eastern Taiwan is usually related to coastal land sea breezes and circulations that are thermally and/or dynamically induced by the topography (Johnson and Bresch 1991; Sun and Chern 1993). Under weak synoptic forcings, offshore convection in this region frequently takes the form of an elongated, narrow band parallel to the shoreline with multiple embedded rainfall maxima. Such convective bands [herein referred to as Taiwan rainbands (TRs)] are a year-round, well-known mesoscale phenomenon in Taiwan, which DOI: /JAS-D Ó 2017 American Meteorological Society. For information regarding reuse of this content and general copyright information, consult the AMS Copyright Policy (

2 1212 J O U R N A L O F T H E A T M O S P H E R I C S C I E N C E S VOLUME 74 influence the coastal weather and potentially threaten the safety of civil aviation and military activities over this local area (Yu and Lin 2008). However, our understanding of the formation and development of TRs is still rather incomplete because of our limited knowledge of the characteristics of local circulations over this particular region and the uncertainty associated with the responses of coastal forcings to different synoptic conditions. A comprehensive view of TRs was first described in a statistical study by Yu and Lin (2008). According to their analyses, the mean duration of the TRs was ;3.5 h, while the occurrence of relatively long-lived TRs (.6h) was not uncommon. These TRs were found to occur more frequently close to the coast (;10 30 km offshore), and the formative area of the TRs was observed to extend from nearshore regions to at least 100 km offshore. In particular, this study classified these TRs into two types namely, nearshore and offshore TRs based on whether the offshore distance of the formative location was less than or greater than 40 km. Their detailed composite analyses of low-level radar-derived airflow patterns within and adjacent to the TRs indicated that the formation of the nearshore TRs was closely related to the low-level convergence that was produced as coastal offshore flow encountered synoptic onshore flow. In distinct contrast, the formation of the offshore TRs appeared to be more related to the deceleration of prevailing onshore flow due to orographic blocking. These results suggested that the convective forcings for nearshore and offshore TRs would be fundamentally different. A limited number of case studies have attempted to address the possible mechanisms responsible for the formation of TRs, and these previous investigations further suggested a diversity of formative processes for TRs. Yu and Jou (2005) used surface and radar observations to document a nighttime, nearshore TR under weak synoptic southeasterly flow during the mei-yu season (i.e., late spring and early summer months; Chen 1983). Their results indicated that low-level convergence, produced as a cool land breeze (i.e., offshore flow) developing at night encountered the prevailing onshore flow, was an important convective forcing contributing to the formation of the TR. Alpers et al. (2007, 2010) used synthetic aperture radar (SAR) images and numerical simulations and similarly confirmed the importance of offshore-flow-induced convergence on the initiation of their studied offshore TRs. However, these authors argued that the presence of coastal offshore flow may result from the recirculated airflow that is produced as synoptic-scale onshore flow dynamically interacts with steep coastal orography, similar to the formation of cloud bands off the island of Hawaii (Smolarkiewicz et al. 1988). Wang and Huang (2009) utilized high-resolution numerical simulations to investigate the reported case in Yu and Jou (2005) and suggested that a combination of topographically generated return flow and mountain breezes would strengthen the offshore flow and convergence at night and thus initiate the deep convection of TRs. Yu and Hsieh (2009, hereafter YH09) investigated a wintertime, offshore TR by using various available observations over southeastern Taiwan. In distinct contrast to the findings from Yu and Jou (2005) and Alpers et al. (2007, 2010), the detailed diagnosis of YH09 indicated that the initiation of their studied TR did not appear to be related to coastal offshore flow. Instead, an orographically driven northerly flow due to upstream blocking was evident along the coast, and lowlevel convergence generated as the prevailing easterly onshore flow encountered the coastally blocked flow provided lifting conducive to the triggering of moist convection within the TR. The persistence of this blocked flow and the degree of ambient instability were shown to be closely related to the maintenance and intensity of precipitation associated with the TR. Despite the differences and complexity in the details of formative processes, the above observational and modeling studies of TRs indicated that the low-level convergence zone produced by diurnal, topographic, or their combined effects is a common convective forcing to initiate TRs. Nevertheless, little has been learned about the environmental factors or convective processes that control the subsequent development of moist convection within TRs after their initiation. Improving our understanding of the possible causes of these evolving aspects is essential because these factors are fundamentally related to the convective intensity and longevity of TRs. The primary objective of this study is to use various available observations collected over eastern Taiwan to document a TR that occurred during 1 3 March 2003 to improve our knowledge of the mechanisms responsible for this TR s formation and to investigate the possible factors that influenced the convective development of this event. This case exhibited two particularly important aspects. One was the long-lived nature of the rainband, whose environment was characterized by temporally changing, synoptically prevailing winds, as will be illustrated in the following sections. The other was the relatively good coverage of precipitation echoes from two coastal Doppler radars. Both aspects made this TR event unique, enabling us to explore the potential effects of environmental conditions on modifying the intensity of precipitation over the duration of the TR and to identify the coastal forcings and/or processes that led to its initiation and evolution. 2. Data The primary observations used in this study are shown in Fig. 1, including surface observations from selected

3 APRIL 2017 Y U A N D L I N 1213 FIG. 1. Topographic features of Taiwan and the data sources that were used in this study. The terrain height (m MSL) is indicated by shading. The locations of the Hualien (HL) and Green Island (GI) Doppler radars are denoted by triangles. The locations of the chosen surface stations in the coastal zone of eastern Taiwan [Suao (SA), Hualien (HL), Chengkung (CK), Taitung (TT), Tawu (TW), Green Island (GI), and Yonaguni (YO)] and other surface stations around Taiwan Island are denoted by open circles. The location of a select sounding profile over the coastal water of eastern Taiwan (i.e., upstream of the topography) from the NCEP gridded data is denoted by a square. The large circles indicate the 120 (115)-km observational range of the GI (HL) radar. conventional weather stations over Taiwan, a few island stations over the surrounding oceanic area, and measurements from two coastal Doppler radars over eastern Taiwan. The Weather Wing of the Chinese Air Force operational C-band (5 cm) Doppler radar on Green Island (GI) was located ;40 km off the southeastern coast of Taiwan, and the Central Weather Bureau of Taiwan operational S-band (10 cm) Doppler radar on Hualien (HL) was located on the eastern coast of Taiwan (Fig. 1). Details on the characteristics of the GI and HL radars are summarized in Table 1. The GI (HL) radar was operated with a temporal interval of 10 or 20 (10) min between each volume. Both radars included volumetric distributions of reflectivity and radial velocity within a 120- and 115-km (for the GI and HL radars, respectively) maximum observational range (Fig. 1). These radar observations provide good spatial coverage of the precipitation and wind information over the coastal water of eastern Taiwan, where the studied TR formed. In addition to these Doppler radar and conventional surface measurements, the National Centers for Environmental Prediction National Center for Atmospheric Research (NCEP NCAR) reanalysis gridded data and the vertical profiles of wind and thermodynamics from a selected NCEP grid that was located ;100 km off the eastern coast of Taiwan (Fig. 1) were used to analyze the large-scale environmental conditions associated with this studied event. 3. Case overview The evolution of low-level radar-observed precipitation associated with the studied TR during its lifetime (;36 h) is shown in Fig. 2. The mean offshore distance and radar reflectivity, which were averaged along the length of the TR at different times, are also shown in Fig. 3. The offshore distance here is defined as a band-normal distance from the position of maximum reflectivity within the

4 1214 JOURNAL OF THE ATMOSPHERIC SCIENCES VOLUME 74 TABLE 1. Green Island (GI) and Hualien (HL) radar characteristics. Parameter GI HL Longitude (8E) Latitude (8N) Altitude (m) Wavelength (cm) Peak power (kw) Frequency (GHz) Antenna gain (db) Beamwidth (8) Pulse repetition frequency (Hz) Max unambiguous velocity (m s21) Max range (km) Pulse duration (ms) Min detectable signal (dbm) Range resolution (km) Rotation rate (8 s21) Elevations (8) , 0.9, 1.6, 2.5, 3.5, 4.7, 5.8, 6.9, 7.9, 9.4, 10.4, 12.4, , 1.4, 2.4, 3.4, 4.3, 6.0, 9.9, 14.6, 19.5 studied rainband to the coastline. The mean radar reflectivity, which represents the rainband s precipitation intensity, was obtained by averaging maximum reflectivities at different along-band locations. The TR formed ;20 km offshore at approximately 2300 LST 1 March (Figs. 2b and 3a). During the initiation stage, the area of precipitation was relatively limited, with multiple separate reflectivity elements. The TR was narrow (only ;10 km in width) and oriented approximately parallel to the eastern coast of Taiwan. With time, the precipitation FIG. 2. The 1.68 (1.48)-elevation PPI scans of the radar reflectivity (dbz) from the GI (HL) radar, which show the low-level precipitation features of the studied TR from (a) 2100 LST 1 Mar to (t) 1100 LST 3 Mar At grid points where reflectivity data from both the GI and HL radars were available, the maximum reflectivity value was chosen to mitigate the effects of attenuation. The solid curve in each panel marks the coastline of eastern Taiwan. The locations of four coastal surface stations (HL, CK, TT, and TW) are also indicated in each panel for reference.

5 APRIL 2017 Y U A N D L I N 1215 FIG. 3. Temporal variation in (a) the mean offshore distance (km) and (b) the mean radar reflectivity (dbz) averaged along the studied TR from 2300 LST 1 Mar to 1100 LST 3 Mar along the TR intensified rapidly and reached its maximum intensity in the morning around 0700 LST 2 March (Figs. 2f and 3b). Highly organized features of intense radar reflectivities, with peak values exceeding 45 dbz, were observed during this time. The rainband s precipitation tended to weaken considerably after ;0700 LST, followed by some slight oscillations in its intensity during the afternoon and nighttime (Figs. 2g q and 3b). During this stage, the TR became relatively less organized, but a linear pattern of precipitation with convective elements of moderate reflectivities (30 35 dbz) was still evident. In contrast to its quasi-stationary nature during the earlier development of the rainband, a persistent but slow seaward movement (;1ms 21 )was observed on 3 March (Fig. 3a). Interestingly, the studied TR intensified again in the early morning between 0300 and 0900 LST 3 March (Figs. 2r t and Fig. 3b). Finally, the rainband dissipated farther offshore at ;75 km off the coast around 1100 LST (Fig. 3a). As revealed by the surface analysis at 2000 LST 1 March (;3 h before the initiation of the studied TR), the synoptic conditions accompanying this event were primarily characterized by weak high pressure to the north of Taiwan, which brought relatively weak north-northeasterly flow (;5ms 21 ) to the vicinity of Taiwan (Fig. 4a). This weak high pressure continuously progressed eastward over the next 36 h, resulting in the transition of low-level prevailing winds near Taiwan from northeasterlies, easterlies to southeasterlies (Figs. 4b,c). At 0800 LST 3 March, close to the dissipation of the studied TR, the Taiwan area was not influenced by the large-scale circulation associated with the high pressure and instead was dominated by weak southerly/southwesterly flow that prevailed in the warm sector of a newly developing cold front to the north of Taiwan (Fig. 4d). Important aspects of the evolving synoptic winds over the study region can be further depicted by a time height cross section of wind information from the NCEP reanalysis data (location in Fig. 1) over periods encompassing the lifetime of the studied TR, as shown in Fig. 5. A clockwise change in the wind direction from northeasterly, easterly, southeasterly, to southwesterly flow was clearly evident at low levels over the duration of the rainband, consistent with the surface analyses in Fig. 4. This wind transition was primarily confined to the lowest 1 km MSL, which largely reflects the influence of propagating surface high pressure, as described above. The large-scale winds above this level remained similar with time and were dominated by moderate southwesterlies (5 10 m s 21 ) between 1 and 3 km MSL and strong westerlies aloft. General onshore flow (i.e., a coastnormal wind component toward the shore, highlighted with shading in Fig. 5) occurred in the lowest 1 km MSL, except close to the TR s dissipation. The low-level onshore flow intensified from ;2ms 21 during the TR s initiation to greater than 6 m s 21 around 0800 LST 2 March because of the transition of the large-scale prevailing winds. The intensity of the onshore flow began to decrease persistently after 2000 LST 2 March as synoptic flow became more southerly or southwesterly during these periods. The environmental thermodynamics remained similar over the duration of the studied rainband, so a mean NCEP sounding (location in Fig. 1) was calculated for presentation, as shown in Fig. 6. A moist layer (relative humidity of 86% 90%) characterizing the low-level onshore flow was present at low levels, with relatively drier air above 1.5 km MSL (Fig. 6a). Consistent with the low convective buoyancy that is typically observed during the Taiwan cold season (e.g., YH09), the convective available potential energy (CAPE) for the present case was very small (i.e., near zero). However, convective instability (i.e., a decrease in the equivalent potential temperature with height) was observed below 3.5 km MSL (Fig. 6b), which would favor the development of moist convection in the presence of a suitable lifting mechanism at low levels. In terms of airflow terrain interactions, the Froude number (F r 5 U/NH, where U is the upstream wind speed, N is the dry Brunt Väisälä frequency, and H is the representative mountain height) is one of the most important nondimensional parameters that determines

6 1216 JOURNAL OF THE ATMOSPHERIC SCIENCES VOLUME 74 FIG. 4. Central Weather Bureau s surface analysis at (a) 2000 LST 1 Mar, (b) 0800 LST 2 Mar, (c) 2000 LST 2 Mar, and (d) 0800 LST 3 Mar Full wind barbs correspond to 5 m s21; half barbs correspond to 2.5 m s21. The location of Taiwan is highlighted with shading. whether the incident flow tends to climb over or divert around the considered mountain barrier (Smith 1979; Overland and Bond 1993). The mean height of the topography over eastern Taiwan is close to ;2 km MSL (cf. Fig. 1). For an elongated mountain barrier with either infinite or limited length, the degree of orographic blocking has been shown to be most dynamically related to the intensity of the incident flow normal to the barrier (i.e., the onshore flow) (Pierrehumbert and Wyman 1985). For the present case, the onshore flow averaged within the lowest 1 km (MSL) varied from 3.3 to 5.7 m s21 over the duration of the studied TR. The NCEP sounding off the eastern coast (location in Fig. 1) was used to calculate the Brunt Väisälä frequency. The average static stability within the layer of onshore flow ranged from to s21. These calculated values yield a Froude number of The occurrence of upstream blocking as the environmental flow approached the mountainous coast of eastern Taiwan is expected in this relatively low Froude number flow regime (i.e.,,1). 4. Coastal thermodynamics and winds One of the most striking mesoscale aspects observed in the present case is the occurrence of a coastal pressure ridge as the synoptically north-northeasterly and/or northeasterly flow approached and interacted with the mountains of eastern Taiwan. The development of the pressure ridge and its associated thermodynamics can be best described by the temporal variation in the difference in the measured surface pressure and potential temperature between a coastal surface station at Suao (SA in Fig. 1) and an offshore island station at Yonaguni (YO in Fig. 1),

7 APRIL 2017 Y U A N D L I N 1217 FIG. 5. Time height section of the environmental winds observed from the NCEP s gridded data ;100 km off the eastern coast of Taiwan (location in Fig. 1) from 0800 LST 1 Mar to 1400 LST 3 Mar The onshore flow component, which is defined as the wind component perpendicular to the mean orientation of the coastline of eastern Taiwan (;208 clockwise from the north) is indicated by the color shading. Full wind barbs correspond to 5ms 21 ; half barbs correspond to 2.5 m s 21. The thick arrow indicates the passage of the front. The duration of the studied TR is marked by two vertical dashed lines. as shown in Fig. 7. The surface winds observed at YO are also superposed on Fig. 7 to provide the basic context of the upstream oncoming flow. These two particular surface stations were selected for analysis because they were located at a similar latitude, so the influence of large-scale pressure variations with a mainly south north gradient near Taiwan (cf. Figs. 4a and 4b) on the calculated pressure difference could be appropriately minimized. The coastal pressure ridge associated with a temperature deficit of ;1.5 K began to form around 2100 LST 1 March during prevailing north-northeasterly/northeasterly flow, and the ridge reached its maximum intensity (;1.5-mb pressure difference; 1 mb 5 1 hpa) 5 h later at around 0200 LST 2 March. The cooler nearshore air during the initiation of the pressure ridge, as seen from surface observations, was not associated with any precipitation, so its presence at this early stage was more likely caused by adiabatic cooling of ascent due to upstream blocking (Bell and Bosart 1988). The formative timing of the coastal ridging roughly corresponded to the initiation of the studied TR. The coastal ridge weakened continuously FIG. 6. (a) Skew T logp plot of the mean NCEP sounding averaged over the duration of the studied TR. (b) Corresponding vertical profiles of the potential temperature and equivalent potential temperature. after 0200 LST 2 March and dissipated near and after noon on 2 March, with a minor pressure difference of less than 0.5 mb between the coastal and offshore stations, as the large-scale oncoming flow became more southeasterly or southerly. A stronger temperature deficit (;3 4.5 K) persisted near the coast until the dissipation of the studied TR. The colder characteristics along the coastal region adjacent to the rainband segment of the studied TR were similarly evident, as seen from coastal observations

8 1218 J O U R N A L O F T H E A T M O S P H E R I C S C I E N C E S VOLUME 74 FIG. 7. Temporal variation in the differences in the pressure (mb, solid line) and potential temperature (K, dashed line) between the SA and YO surface stations (locations in Fig. 1) from 0800 LST 1 Mar to 1400 LST 3 Mar The surface winds that were observed well offshore at the YO station are indicated by wind flags, with full wind barbs corresponding to 5 m s 21 and half barbs corresponding to 2.5 m s 21. The shading represents the duration of the studied TR. of the surface stations at HL, Chengkung (CK), Taitung (TT), and Tawu (TW) (cf. Figs. 1 and 2). As shown in Fig. 8, the mean potential temperature from these nearshore stations was consistently lower than what was measured by the offshore station at GI (cf. Fig. 1) over the duration of the studied TR. The mean deficit in the potential temperature was relatively small (21 to 21.5 K) in the early morning on 2 March (Fig. 8c). A slightly larger temperature deficit (22 to 23 K) that corresponded to the occurrence of nearshore rainfall began to appear around 0800 LST and tended to persist for the rest of the TR s lifetime (Figs. 8a,c). Radar observations indicated that the coastal rainfall observed by these surface stations was indeed caused by the arrival of precipitation from the studied TR (cf. Fig. 2f). The evaporative cooling of the nearshore precipitation may represent one of the important contributors to the existence of this cooler nearshore air. Note that weak diurnal variation in potential temperature was evident (Fig. 8b), which also modulated the degree of the calculated temperature deficit. For example, the coldest nearshore air occurred around 0500 LST 3 March (black curve in Fig. 8b), which was consistent with the nighttime radiative cooling over coastal land under a relatively clear sky because the clouds and precipitation of the studied TR had moved far away from the coast during this late stage. At the same time, the temperature of the offshore air FIG. 8. Hourly time series of nearshore and offshore surface observations from 0800 LST 1 Mar to 1400 LST 3 Mar (a) Nearshore average rainfall (mm h 21 ) measured at the HL, CK, TT, and TW surface stations (locations in Fig. 1); (b) nearshore average potential temperature (K, black line) measured at the HL, CK, TT, and TW surface stations and the offshore potential temperature (K, gray line) measured at the GI surface station; (c) difference in the potential temperature between the nearshore and offshore surface stations. The duration of the studied TR is marked by two vertical dashed lines. continued increasing (gray curve in Fig. 8b), presumably as a result of the influence of warmer, more southerly large-scale flow (cf. Figs. 5 and 7). These two effects resulted in the maximum temperature deficit (;25K) observed in the early morning on 3 March (Fig. 8c). As will be elaborated in the following sections, cold nearshore air played an essential role in the maintenance of moist convection for the studied case. Further support of the occurrence of a pressure ridge along the eastern coast of Taiwan is provided by a sea level pressure analysis in the vicinity of Taiwan. Figure 9a shows the mesoscale surface analysis at 0300 LST 2 March, close to the maximum intensity of the coastal ridge, as seen from Fig. 7. The mesoscale pressure ridge was evident adjacent to the mountainous coast of eastern Taiwan. There was a relative low along the western coast of Taiwan (i.e., the leeside trough). Such a ridge trough pressure pattern usually occurs as large-scale northeasterly flow interacts with Taiwan Island under a relatively low Froude number flow regime (e.g., Sun et al. 1991). The surface winds along the eastern coast of Taiwan were generally parallel to the

9 APRIL 2017 Y U A N D L I N 1219 FIG. 10. Time series of hourly surface winds from 0800 LST 1 Mar to 1400 LST 3 Mar 2003 at the (a) HL, (b) CK, (c) TT, (d) TW, and (e) YO surface stations. The offshore (onshore) flow component is indicated by a solid line with positive (negative) values. Full wind barbs correspond to 5 m s 21 ; half barbs correspond to 2.5 m s 21. The duration of the studied TR is marked by shading. Note the wider range for the wind speed in (e). FIG. 9. Mesoscale analysis of the sea level pressure at (a) 0300 and (b) 1500 LST 2 Mar Full wind barbs correspond to 5 m s 21 ;half barbs correspond to 2.5 m s 21. A corresponding low-level PPI scan of the radar reflectivity (dbz) is also superposed on each panel for reference. coast and highly ageostrophic, in contrast to the more geostrophic northeasterlies farther offshore. The surface analysis at 1500 LST 2 March (Fig. 9b) indicated a uniform sea level pressure distribution over the Taiwan area, consistent with Fig. 7, which shows the dissipation of the coastal ridge in the afternoon on 2 March. Time series of hourly surface observations from four selected stations (HL, CK, TT, and TW; Fig. 1) along a coastal segment immediately west of the studied TR and a sequence of the lowest PPI scans (0.48 elevation) of radial velocity from the GI radar revealed important aspects of the coastal winds, as presented in Figs. 10a d and Fig. 11, respectively. The surface winds observed well offshore at YO (cf. Fig. 1) werealsoshowninfig. 10e for comparison. The heights of radar measurements from the lowest PPI scan varied from ;0.3 km MSL near the radar to ;1.1 km MSL at its maximum observational range. The coastal surface winds were mostly from the north (i.e., roughly parallel to the coast), and the northerly flow tended to persist during the influence of the coastal pressure ridge (i.e., in the morning on 2 March). The surface northerly flow was somewhat weak (;1 5 m s 21 ), presumably because of friction near the ground. Radar observations indicated

10 1220 J O U R N A L O F T H E A T M O S P H E R I C S C I E N C E S VOLUME 74 FIG. 11. Sequence of the 0.48-elevation PPI scans of radial velocity (m s 21, color shading) from the GI radar from 2300 LST 1 Mar to 1400 LST 2 Mar 2003, which show the development and dissipation of the nearshore northerly blocked flow. Winds observed from the HL, CK, TT, TW, and GI surface stations are also indicated for reference; full (half) wind barbs correspond to 5 (2.5) m s 21. The solid line represents the coastline of eastern Taiwan. stronger radial velocities (210 to 212 m s 21, blue shading in Fig. 11) approaching the radar (i.e., northerly flow) to the north of the radar immediately off the southeastern coast of Taiwan. Note that enhanced, positive radial velocities (i.e., flowing away from the radar) to the south of the radar similarly indicated the presence of the nearshore blocked northerlies. An obvious transition from positive to negative radial velocities to the north-northwest of the radar was evident particularly in the early morning, which would most likely reflect the low-level convergence associated with the studied rainband during these periods. The PPI scans of the radial velocity from different elevations (not shown) indicated that the layer of the enhanced northerly flow (.10 m s 21 ) was primarily confined to the lowest 1 km MSL, which was roughly comparable to the vertical extent of the low-level prevailing onshore flow (Fig. 5) but well below the average height of the coastal mountains (Fig. 1). An available HL sounding released at 0800 LST 2 March (not shown) indicated that the northerly flow was characterized by a higher static stability (N 5;0.02 s 21 ) and convectively neutral conditions (i.e., near-zero variations of equivalent potential temperature in the vertical). A relatively cooler nature (cf. Fig. 7) and stronger stratification have both been recognized to be common thermodynamical characteristics of orographically blocked flow that develops adjacent to the mountain barriers of different geographical locations (e.g., Parish 1982; Overland and Bond 1993; Doyle 1997; Yu et al. 2001; YH09). The surface northerly flow began to disappear (Figs. 10a d) as the coastal ridge dissipated around noon on 2 March (cf. Fig. 7). Radar observations further showed that the coastal northerly flow, as revealed by negative radial velocities (i.e., approaching the radar), continuously weakened after 0200 LST 2 March (Fig. 11). The radar-observed northerly flow almost disappeared at 1400 LST 2 March, which matches the evolving aspects of the coastal surface winds in Fig. 10. The maximum magnitudes of the approaching radial velocities when plotted as a function of time, as shown in Fig. 12, indicated a clear trend corresponding to the temporal variation in the intensity of coastal pressure ridge. These surface and radar analyses support the crucial role of topographically generated pressure ridges in determining the earlier development of coastal winds.

11 APRIL 2017 Y U A N D L I N 1221 onshore flow, with a maximum intensity of 27ms 21 around LST 2 March, were basically consistent with those of the NCEP profiles in Fig Convective forcings FIG. 12. (a) Temporal variation in the radar-observed minimum radial velocity (m s 21, solid line) from the 0.48-elevation PPI scan of the GI radar and the radar-observed maximum radial velocity (m s 21, dashed line) from the 0.58-elevation PPI scan of the HL radar from 2100 LST 1 Mar to 0300 LST 2 Mar (b) Temporal variation in the difference in the sea level pressure (mb) between the SA and YO surface stations, which is as in Fig. 7, but for a different time window. After the collapse of the coastal pressure ridge and its associated blocked flow, coastal surface winds became nearly calm and/or more westerly/northwesterly with generally offshore flow (i.e., positive coast-normal wind component) (Figs. 10a d). The offshore flow tended to intensify slightly during the nighttime and early morning hours, which suggests the importance of diurnal effects on influencing the characteristics of coastal winds. According to the temperature analysis in Fig. 7, radiationcooled air over coastal land is expected to feed into adjacent water areas through the land-breeze-driven offshore flow, which favors the maintenance of preexisting cold nearshore air overnight. A clear wind transition from offshore flow to onshore flow (i.e., from positive to negative coast-normal wind component in Figs. 10a d) was evident in the morning of 3 March. This wind transition after sunrise was accompanied by a sharp increase in surface temperature near the coast and an opposite coast-normal temperature gradient [i.e., warmer (colder) in the nearshore (offshore) regions] (Figs. 8b,c), consistent with the development of sea breeze circulations. In contrast to the characteristics of the nearshore winds, which were largely influenced by either topographic or diurnal effects, the surface winds observed well offshore at YO were dominated by persistent, synoptically driven onshore flow over the duration of the studied TR (Fig. 10e). The evolving variations in the The combination of surface and radar observations, as discussed in the previous sections, suggested that the studied TR would be located along a region of wind transition between the blocked northerlies near the coast and the prevailing onshore flow off the eastern coast of Taiwan. The timing of the rainband s initiation also roughly corresponded to the generation of the coastal pressure ridge and its associated northerly blocked flow (cf. Figs. 7 and 10). The convergence produced as the upstream onshore flow encountered the nearshore cool, blocked flow may represent an important convective forcing conducive to the formation of the studied TR. Detailed wind information near and within the studied rainband derived from available radar observations was presented in this section to further support this scenario. As described in section 3, the precipitation of the observed TR generally exhibited a two-dimensional nature (cf. Fig. 2). The band-normal circulations could be practically derived from the range height indicator (RHI) vertical cross section perpendicular to the band (Wakimoto 1982). The cross-band horizontal velocities were approximated by the horizontal component of adjusted Doppler radial velocities that had the velocity component due to the projection of the particle terminal velocity onto the radar beam removed. The vertical velocity was obtained by calculating the convergence from the cross-band horizontal velocities and then integrating the anelastic continuity equation from a lower-boundary condition of zero vertical velocity at the surface. The reason why the upward integration was applied herein was that the calculated horizontal motions and convergence were less accurate at higher storm levels where larger elevation angles of radar beams were expected. Such a procedure provided better (less) reliability of the derived vertical velocities at lower (upper) levels within the rainband. Given a number of assumptions required for retrieving airflow structures from a single Doppler radar, the result will be interpreted in a more qualitative than quantitative way. In addition, we could also retrieve the horizontal winds along the band-normal direction by using a so-called two-point method (Lawrence et al. 1975; Strauch and Moninger 1978; Schoenberger 1984). In this method, horizontal winds are considered to be the same at two closely adjacent points along the rainband. The crossband and along-band wind components at the central

12 1222 J O U R N A L O F T H E A T M O S P H E R I C S C I E N C E S VOLUME 74 FIG. 13. (a) Low-level PPI scan (1.68 elevation) of the radar reflectivity (dbz, color shading) from the GI radar at 0610 LST 2 Mar The terrain height (m MSL) is indicated by gray shading. The inset box marks the zoomed-in area in (b). The thick solid line along the 3058 azimuth angle marks the location of the vertical cross section approximately normal to the orientation of the studied TR, which is shown in (c). Winds observed from the TT and GI surface stations are also indicated for reference; full (half) wind barbs correspond to 5 (2.5) m s 21. (b) Low-level PPI scan of the radar reflectivity (dbz, color shading) over the inset box in (a) and the low-level horizontal winds (0.5 km MSL) derived from the two-point method along the cross section in (a). (c) RHI display of the radar reflectivity (dbz, color shading) and derived cross-band winds (vectors; key in the upper left) along the cross section. position of the two points can be expressed as a function of the radial velocities at the two points. A primary benefit of this method is its ability to deduce the wind component parallel to the band (i.e., coast-parallel flow for the present case), which can reveal the nearshore blocked flow and its relationship to the rainband s precipitation. The details of these two retrieval methods were also described in Yu and Jou (2005) and YH09. A finescale view of low-level precipitation associated with the studied TR along a coastal segment near the GI radar (Fig. 1) at 0610 LST 2 March is shown in Fig. 13a, and the low-level winds along the 3058 azimuth angle (i.e., in the cross-band direction) derived from the twopoint method and its corresponding cross-band precipitation and airflow structure from the RHI cross section are shown in Figs. 13b and 13c, respectively. As described in section 3, the rainband was quasi-stationary during its earlier development, with highly organized and its strongest precipitation at around 0600 LST 2 March (cf. Figs. 2 and 3a). The horizontal winds in

13 APRIL 2017 Y U A N D L I N 1223 Fig. 13b reveal that the zone of heaviest precipitation (40 50 dbz) in the studied TR coincided with the boundary between the north-northeasterly blocked flow near the coast and the synoptically prevailing easterly flow offshore. Figure 13c further shows that the lowlevel easterly onshore flow (i.e., inflow) entered the precipitation band from the east and began to be lifted near the offshore (eastern) edge of the blocked flow. The precipitation of the studied rainband was characterized by an obvious convective nature, with the 30-dBZ contour exceeding 3 km MSL, enhanced horizontal gradients of radar reflectivity, and strong upward motions. A significant decrease in the intensity of precipitation with height was evident above ;3 km MSL, which may have been related to a convectively stable environment at these higher levels, as shown in Fig. 6b. Some elevated downward motions were found above ;1.5 km MSL on the eastern side of the updraft core. The downdraft was located in regions of relatively low reflectivity so it was probably not caused by water loading. Instead, its occurrence was more likely driven by negative buoyancy associated with evaporation of precipitation and/or by compensating subsidence accompanying the main updraft region. However, the intensities of both updraft and downdraft tend to reach a maximum near the top of the storm, which seems less realistic owing to larger uncertainties of retrieved vertical velocities at these higher levels as described earlier. The actual top of the storm was not adequately captured by this particular vertical cross section because of limitations in the radar scanning. An examination of many other RHI observations indicated that the echo top of the rainband varied in time and space, mostly ranging from 2.5 to 4.5 km MSL. The observed characteristics of relatively lower storm heights (i.e., below the level of 08C; cf. Fig. 6a) suggest the importance of warm rain processes and were also consistent with the layer of convective instability, which was confined to the lowest 3.5 km MSL (cf. Fig. 6). A sequence of similar radar analyses was also performed after the dissipation of the coastal pressure ridge (cf. Figs. 7 and 9b). The radar-derived winds at 2210 LST (Fig. 14) were chosen to illustrate the representative airflow features of the studied rainband during these later periods. The low-level inflow on the eastern side of the rainband became more southerly (Fig. 14b), as expected from the synoptic observations, which showed prevailing south-southeasterly flow at low levels in the afternoon and nighttime of 2 March (cf. Fig. 5). Weak westerly flow was observed on the western side of the rainband, and no evidence of the northerly blocked flow was present (Figs. 14a,b), both of which were consistent with the characteristics of the nearshore surface winds in Fig. 10. The westerly flow was associated with a lower temperature compared to the airflow on the inflow (i.e., eastern) side of the rainband (cf. Fig. 8c). The southerly inflow tended to be lifted upward as it met with the colder, more westerly flow (Figs. 14b,c). As elaborated in section 4, colder air was persistently present nearshore (i.e., to the west of the studied rainband) over the duration of this event, and a stable low-level forcing of lifting required for the maintenance of moist convection is expected to have occurred in the presence of onshore flow (i.e., easterly or southerly inflow). The colder air (i.e., cold pool) existed adjacent to coastal mountain barrier and the environmental onshore flow prevailed during the studied event, both of which would generally prevent the cold pool from spreading away from the nearshore region (e.g., Xu et al. 2012). The eastern edge of the cold pool, as identified by the offshore distance of the studied rainband, began to propagate seaward on 3 March(Fig. 3a). This offshore propagation is consistent with the weakening synoptic-scale onshore flow or even offshore flow observed during these periods (cf. Fig. 5). The results from Figs. 13 and 14 further support that the lifting produced as low-level prevailing onshore flow encountered nearshore blocked/colder flow was an important convective forcing for the studied event. In this context, the temporal difference in the intensity of the rainband s precipitation (cf. Fig. 3b) may have been mostly influenced by the degree of the low-level lifting, given similar environmental thermodynamics over the duration of the rainband, as described in section 3. One of the most prominent thermodynamic features of the nearshore flow was its cooler nature, so the degree of lifting along its offshore (eastern) edge would not have been simply related to convergence between the blocked/colder flow and prevailing onshore flow; instead, the convective intensity would have been dynamically determined by the relative strength between the cold pools and ambient vertical shear, a widely accepted theory of convective dynamics pioneered by Rotunno et al. (1988). In this theory, the ambient vertical shear in the cross-band direction constrains the intensity and tilting nature of the lifted updrafts at the leading edge of cold air, which can be appropriately interpreted with horizontal vorticity balance dynamics. Following this concept, the horizontal vorticity generated by buoyancy changes across the leading edge of the cold pool may be represented by rffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi g C 5 DuH, (1) u o where Du is the cold-pool temperature deficit relative to ambient conditions, u 0 is the base-state potential

14 1224 J O U R N A L O F T H E A T M O S P H E R I C S C I E N C E S VOLUME 74 FIG. 14. As in Fig. 13, but at 2200 LST 2 Mar temperature, and H is the cold-pool depth. The ambient shear-induced vorticity can be represented by Du, which is the vertical shear over the depth of the cold pool H. An optimal state occurs when C 5Du; in this situation, the convective updrafts at the leading edge of the cold pool are upright and the most intense. If the cold pool induced vorticity is not countered by a comparable ambient shear (i.e., either C.Du or C,Du), convective motions would be less intense with upshear or downshear tilt. For the present case, Du can be estimated by the nearshore temperature deficit, as shown in Fig. 8. The cold-pool depth H may be approximated by the vertical extent of the blocked flow (;1 km), as described in section 4. This magnitude is also comparable to or slightly higher than the subcloud layer [i.e., lifting condensation level (LCL), ;0.5 1 km based on the NCEP soundings], where the evaporative cooling of hydrometeors is most prominent. The NCEP profiles in Fig. 5 were used to calculate the ambient vertical shear Du. These calculations can determine the temporal variations in C, Du, and their difference (i.e., C 2Du) over the duration of the studied rainband, as shown in Fig. 15. The results indicated that the ambient vertical shear was generally suboptimal (i.e., C. Du) (Fig. 15a), consistent with the observed updrafts sloping over the cold air (i.e., upshear tilt) at low levels (cf. Figs. 13c and 14c). An overall opposite relationship between the intensity of the rainband s precipitation and the degree of the optimal state was also evident. For example, the

15 APRIL 2017 Y U A N D L I N 1225 FIG. 15. (a) Temporal variation in the buoyancy-induced horizontal vorticity C (solid line) and the ambient vertical shear Du (dashed line) that were calculated below 1 km MSL from 2300 LST 1 Mar to 1100 LST 3 Mar (b) Corresponding temporal variation in the mean precipitation intensity of the studied TR (dbz, gray line), as in Fig. 3b, and the value of C 2Du (dotted line). rainband intensified rapidly after its initiation and reached its maximum intensity at around 0700 LST 2 March, which roughly corresponded to the minimum difference between C and Du (Fig. 15b). This evolution toward increasingly optimal conditions was mostly accompanied by an obvious increase in ambient vertical shear (dashed curve in Fig. 15a). The observed enhanced vertical shear was caused by the intensification of low-level onshore flow (i.e., cross-band flow) owing to the transition of the largescale prevailing winds from northeasterly to more easterly flow during the early morning of 2 March (cf. Fig. 5), as described in section 3. The rainband s intensity weakened in the late morning of 2 March and throughout the rest of the day, which was characterized by an increasing C 2Du (i.e., more suboptimal). The occurrence of the intensified precipitation immediately before the dissipation of this studied rainband, as described in section 3, was also consistent with the significantly decreasing C 2Du (mostly because of the weakening C; cf. Fig. 15a) evident in the morning of 3 March. The dynamical interaction of the cold nearshore air with the environmental vertical shear can reasonably explain the overall evolving aspects of the rainband s intensity. 6. Scale analysis of orographically blocked flow The previous sections indicated the crucial role of northerly blocked flow in the initiation of the studied rainband, in a similar manner to the formative process of an offshore TR documented in YH09. However, the duration of the northerly blocked flow observed in the studied case was relatively short (;12 h; cf. Figs. 10, 11, and 12), in contrast to the relatively steady, long-lived nature of the coastally blocked flow (.24 h) in YH09.In addition, the seaward extent of the blocked flow, as revealed by the offshore distance of the studied rainband in the morning of 2 March (cf. Fig. 3a), was more limited (;10 15 km) compared to that of YH09 (;55 km). A scale analysis of airflow along the mountainous coast of eastern Taiwan was performed for the studied case to seek a fundamental understanding of these observational characteristics and their physical connection to the large-scale environmental conditions. A nondimensional form of the momentum equations in straits or along a mountainous coast, as explored by Overland (1984), may be expressed by dy R l dt 1 C d (u2 1 y 2 ) 1/2 y 1 u 52 p 0 y 2 p (2) y and du R L dt 1 C d (u2 1 y 2 ) 1/2 u 2 y 52 p 0 x 2 p x, (3) where u and y are the nondimensional velocity components in the cross-coast (x) and along-coast (y) directions, respectively; p 0 is the prescribed nondimensional largescale pressure field; p is the nondimensional perturbation pressure from mass adjustment within the mountainous coastal zone; R l and R L are the along-coast and crosscoast Rossby numbers, respectively; and C d is the scaled drag coefficient. Rossby numbers R l and R L measure the nonlinearity of the momentum equations, and C d indicates the relative importance of surface friction to the acceleration terms. These variables are given by R l 5 V fl, (4) R L 5 l2 L 2R, l and (5) C d 5 L D C, d (6) where V is the velocity scale in the along-coast direction, f is the Coriolis parameter, L (l) is the length scale in the along-coast (cross coast) direction, D is the height of the coastal topography, and C d* is the drag coefficient. In this study, the north south length of Taiwan Island and the average mountain half-width of eastern Taiwan (cf. Fig. 1) are used to represent L and l, respectively, and