A Case Study of Nocturnal Rain Showers over the Windward Coastal Region of the Island of Hawaii

2674 MONTHLY WEATHER REVIEW A Case Study of Nocturnal Rain Showers over the Windward Coastal Region of the Island of Hawaii JUN LI ANDYI-LENG CHEN Department of Meteorology, School of Ocean and Earth Science and Technology, University of Hawaii, Honolulu, Hawaii (Manuscript received 7 April 1998, in final form 24 November 1998) ABSTRACT Nocturnal rain showers over the windward side of the island of Hawaii were investigated from the late afternoon of 2 August to the early morning of 3 August 1990 during the Hawaiian Rainband Project (HaRP). Three types of rainbands produce rainfall peaks over the lowland/coastal region during the evening, after midnight, and along the coast in the early morning. In the early evening, an inland rainband develops over the lower slopes as a result of orographic lifting and low-level forcing along the drainage front. As the drainage front progresses toward the coast, new rain cells continue to develop along the drainage front. These cells move westward with the trade winds aloft and dissipate. After the drainage front moves to the Hilo coast, new cell generation along the drainage front ceases. It appears that in the absence of orographic lifting, the shallow ( 0.2 km) katabatic flow offshore is not deep enough to lift the low-level air to saturation. Thus, the rainband weakens and dissipates. It lasts about 6 h and produces the heaviest rainfall over the lowland/coastal region during the analysis period. During the evening transition, rain evaporative cooling aloft deepens the cool pool behind the drainage front over land. After the drainage front moves off the coast around 2300 HST, cold air continues to move from lower slopes and lowlands toward the coast. The offshore flow gradually deepens. By 0207 HST, the depth of the katabatic flow over the coastal region reaches 0.48 km, which is above the level of free convection. A rainband develops along the drainage front offshore. As the rainband moves westward over the deepened katabatic flow, the katabatic flow is replaced by easterly winds. The rainband weakens and dissipates over the lower slopes. During the early morning, two groups of trade wind rain showers move into the coastal region. They are enhanced offshore in the convergent zone between the offshore flow and the trade winds. When the first group of rain showers moves over the katabatic flow, the outflow associated with the rain showers deepens the offshore flow ( 0.6 km) east of the rain showers. The second group of rain showers is more intense than the first group as the rain showers move into the convergent zone and interact with the deep offshore flow. Nevertheless, they weaken rapidly over the deep offshore flow and produce little rainfall along the coast. 1. Introduction Rainfall observations in Hawaii date from the 1840s. Over the years, rainfall distributions over the Hawaiian Islands were studied by numerous investigators (Leopold et al. 1951; Schroeder et al. 1977; Takahashi 1977a; Meisner 1979; Giambelluca et al. 1986, and others). Using rainfall data updated through 1983, Giambelluca et al. (1986) compiled the rainfall atlas for the six largest islands of Hawaii. In general, rainfall maxima correspond to regions of persistent orographic lifting of moisture-laden northeast trade winds up to the windward slopes. Areas of low rainfall are found in leeward areas and atop the highest mountains. High mountain dry ar- Corresponding author address: Dr. Yi-Leng Chen, Department of Meteorology School of Ocean and Earth Science and Technology, University of Hawaii, 2525 Correa Road, Honolulu, HI 96822. E-mail: dave@soest.hawaii.edu eas occur because the moist air is prevented from reaching this elevation by the presence of the trade wind inversion. For mountains with tops well below the inversion, the rainfall maximum occurs near the summit. In addition to orographic lifting, rainfall occurrences over Hawaiian islands are strongly modulated by the diurnal heating cycle (Leopold 1949; Eber 1957; Lavoie 1967; Mendonca 1969; Schroeder et al. 1977; Takahashi 1977a; Garrett 1980; Schroeder 1981, and others). Thermally driven diurnal circulations (e.g., land sea breeze and mountain valley winds) may contribute to rainfall by reinforcing trade orographic lifting, generating areas of low-level convergence with interaction with the prevailing trade winds, or producing orographic lifting in areas not exposed to trade winds (Giambelluca et al. 1986). The island weather is also affected by dynamic blocking as the trade wind flow encounters the island obstacle. The splitting of the trade wind flow by the island of Hawaii has also long been recognized in the literature (e.g., Leopold 1949; Fellbaum 1984; Nick- 1999 American Meteorological Society

NOVEMBER 1999 LI AND CHEN 2675 erson 1979, and others). Leopold noted that the trade wind inversion acts as a lid forcing low-level flow to split around the island. Comprehensive studies of dynamic blocking on the rainfall production on the windward side of the island of Hawaii were carried out by recent modeling and observational studies (Smolarkiewicz et al. 1988; Rasmussen et al. 1989; Carbone et al. 1998). The latent heat transfer from condensation and evaporation of precipitation also feeds back to the island airflow (Chen and Wang 1995; Wang and Chen 1995; 1998; Frye and Chen 1997; Carbone et al. 1998). Prior to the Hawaiian Rainband Project (HaRP), conducted during 11 July to 24 August 1990, most observational studies emphasized the surface thermal forcing as the primary cause for the development of katabatic winds and nocturnal rainfall along the windward coast (Leopold 1949; Lavoie 1967; Schroeder et al. 1977; Takahashi 1977a; Garrett 1980). Cloud bands frequently observed offshore in the early morning were believed to form over the convergent zone between the katabatic/ offshore flow and the incoming trade winds. Once formed, they drift inland resulting in the nocturnal rainfall maximum in the windward coastal region. On the basis of theoretical considerations and numerical modeling, Smolarkiewicz et al. (1988) and Rasmussen et al. (1989) gave a different interpretation from the earlier studies on the rainband formation offshore. For a low Froude number flow regime, the nighttime surface offshore flow is driven by dynamic blocking in their model, they also simulated the generation of offshore cloud bands in the low-level convergent zone along the separation line between the dynamically driven return flow and the incoming trade winds in their model. Rasmussen and Smolarkiewicz (1993) suggest that the nocturnal cooling only slightly modifies the strength of the dynamically driven flow but not the location of the convergent zone. Based on HaRP observations, Chen and Nash (1994) suggest that on the windward side, the evolution of diurnal rainfall patterns is related to a complex interaction among orographic lifting, thermal forcing, and dynamic blocking. Chen and Feng (1995) divided the precipitation over the windward section of the island of Hawaii into three diurnal regimes: the daytime [1100 1900 Hawaiian standard time (HST)] orographic rainfall on the windward slopes, nocturnal (1900 0300 HST) rain showers on the windward lowlands, and coastal precipitation in the predawn and early morning hours (0300 1100 HST). Over the lowland/coastal region, the heaviest 4-h rainfall during the diurnal cycle is observed during 1900 2300 HST. Chen and Nash (1994) suggest that the nocturnal rain showers there are caused by the convergence between the katabatic flow and the incoming trade winds and enhanced by orographic lifting aloft. The mechanism can be referred to as Leopold s (1949) inland model (Feng and Chen 1998). After the onset of the katabatic flow, the nocturnal rainfall on the windward side moves seaward as the katabatic flow progresses toward the coast (Chen and Nash 1994; Chen and Feng 1995). In the early morning, most of the early morning rainbands offshore originate upstream from preexisting trade wind rain showers (Austin et al. 1996; Ochs et al. 1996; Wang and Chen 1998) and are enhanced in the offshore convergent zone where the decelerating trade wind flow encounters the cold offshore flow (Wang and Chen 1998). They weaken as they move westward over the katabatic flow (Wang and Chen 1998). These early morning rain showers along the coast produce much less rainfall on the windward side than nocturnal rain showers in the evening (Chen and Feng 1995). Chen and Nash (1994) documented the diurnal surface circulation and rainfall patterns using HaRP data. For most areas, the surface airflow is dominated by daytime anabatic flow and nighttime katabatic flow. They suggest that in regions where mean winds are weak because of island blocking, the thermally driven diurnal winds become significant. Chen and Wang (1994) found that on the windward side where the mean surface winds are weak because of upstream flow deceleration, the onset of the anabatic (katabatic) flow there in the early morning (late afternoon) is governed by the thermal contrast between the slope surface and the upstream environment at the same altitudes. Clouds and rainfall can modify the surface thermal fields and result in changes in the intensity of the diurnal circulations and the timing of the wind shift from katabatic (anabatic) to anabatic (katabatic) flow at the surface in the early morning (late afternoon) (Chen and Wang 1995). Carbone at al. (1995) suggest that the flow reversal on the windward slopes in the late afternoon is mainly driven by evaporative cooling of orographic precipitation. The composite analyses of the surface airflow and precipitation patterns using HaRP surface data (e.g., Chen and Nash 1994; Chen and Wang 1994, 1995; Chen and Feng 1995) improve our understanding of the general characteristics of surface airflow and trade wind weather over the island of Hawaii. Nevertheless, there are differences among cases. Furthermore, detailed evolution of the katabatic flow and individual rainfall events are smoothed out in these analyses. With large day-today variations in daily trade wind rainfall during HaRP (Chen and Feng 1995; Chen and Wang 1995), case studies using data from all observing platforms will provide further insight on the interactions between trade wind rain showers and the island-induced airflow. The heaviest nocturnal rain showers for the period of 2300 0700 HST during HaRP are recorded on the windward side during the night of 7 8 August (Wang and Chen 1995). For this case, despite the presence of radiative cooling and evaporative cooling, the nocturnal inversion disappears after midnight with increasing surface temperature and moisture content as the potentially warmer air with higher moist static energy aloft mixes with the colder and drier katabatic flow at the surface (Wang and Chen 1995). With evaporative cooling aloft and vertical

2676 MONTHLY WEATHER REVIEW mixing associated with rains, this case is characterized by a deep ( 500 m) stable layer with weak westerly flow ( 3 ms 1 ) in the early morning (Wang and Chen 1995). This is also an extremely strong trade wind ( 11 ms 1 ) case (Frye and Chen 1997) characterized by an elevated Froude number (Fr UN 1 H 1 ) of 0.42 with strong dynamic blocking (Grubišić et al. 1996), where U, H, and N are the upstream wind speed, the averaged mountain height, and the Brunt Väsälä frequency, respectively. Nevertheless, the adiabatic source of dynamically driven return flow on the order of 7 ms 1 predicted by Smolarkiewicz et al. (1988) is not observed (Wang and Chen 1995, their Table 3). In contrast, Frye and Chen (1997) show that for this case the evolution of near-surface katabatic flow is strongly modulated by trade wind rain showers that drift inland, resulting in an oscillation of the drainage front from 5 8 km offshore to 10 km inland. Tethersonde observations also show large day-to-day variations in the strength of katabatic flow (3 9 m s 1 ) related to nocturnal rain shower activities (Wang and Chen 1995). For relatively dry nights, the katabatic flow is stronger and shallower with a pronounced nocturnal inversion in the early morning as compared to wet nights. Feng and Chen (1998) report that for 10 August, a weak (4 5 m s 1 ) trade wind case with little rains after midnight, the katabatic flow reaches 7 m s 1 with a pronounced nocturnal inversion in the early morning. Wang and Chen (1998) analyzed early morning rainbands for the 22 August 1990 case to study the interactions between early morning trade wind rainbands and the katabatic flow. Feng and Chen (1998) performed a case study on the evolution of the katabatic flow throughout the night on the windward slopes and the nearby ocean for a relatively dry night. The objective of this study is to perform a comprehensive analysis of the evolution of nocturnal rain showers and their interactions with the island-induced airflow throughout the entire night under normal trade wind conditions. From the late afternoon 2 August to the early morning 3 August 1990, several rainbands and the katabatic (offshore) flow over land (nearby ocean) were well monitored by Doppler radars and surface mesoscale network. In particular, in addition to monitoring nocturnal rain showers, the range height indicator (RHI) scans were frequently performed by Doppler radars over land and the nearby ocean to document the evolution of katabatic flow throughout the entire night. In section 2, the data used in this study are described. An overview of the largescale and mesoscale conditions is given in section 3. In section 4, we present the formation and evolution of an inland rainband during the evening hours. The influences of katabatic/offshore flow on the inland rainband and the feedback of the rain showers associated with the rainband on the low-level airflow, particularly the depth and strength of the katabatic/offshore flow, are analyzed. In section 5, we document the initiation and evolution of an offshore rainband after midnight and its FIG. 1. Part of the windward side of the island of Hawaii with selected PAM sites (circle with station number), Doppler radars ( ), and the tethersonde site (*) at UHAES. Terrain contours every 500 m. interactions with the island-induced airflow. In section 6, two early morning trade wind rainbands during 0600 0900 HST are described and compared with the inland rainband during the evening and with the offshore rainband after midnight. A summary is given in the last section. 2. Data analysis The data used in this study are primarily from the Portable Automatic Mesonet (PAM) stations and Doppler radars during 2 3 August 1990. Other data sources include upstream soundings obtained by the National Center of Atmospheric Research (NCAR) Electra, tethersonde observations at the University of Hawaii Agriculture Experiment Station (UHAES, Fig. 1), rawinsondes launched by the National Weather Service (NWS) at Hilo airport, satellite images from Geostationary Operational Environmental Satellite and National Oceanic and Atmospheric Administration satellites, and synoptic charts analyzed by NWS. During HaRP, pressure, temperature, wet-bulb temperature, wind direction and speed, and rainfall were measured at 1-min intervals by PAM stations. Stations 15, 16, and 17 (Fig. 1) also recorded radiation variables, including net all-wave radiation (the difference between the incoming and outgoing radiation including both longwave and shortwave radiation), longwave radiation, global solar incoming, diffuse solar radiation, and reflected solar radiation. The HaRP Data Catalog (HaRP 1991) describes details on resolution and accuracy of the sensors used in the experiment. A careful manual inspection was made to eliminate obvious errors

NOVEMBER 1999 LI AND CHEN 2677 and data spikes in the original measurements. In addition, 15-min-average surface data (Chen and Nash 1994) were used to present horizontal distribution of surface fields. The upstream sounding was obtained from the NCAR Electra when the aircraft ascended from 150 m, 20.012 N, 153.960 E, to 4860 m, 20.065 N, 153.785 E, during 0520 0541 HST 3 August. The vertical resolution of the original data was about 15.7 m. To analyze the thermal contrast between the windward slopes and the upstream environment at the same altitudes, the sounding data were interpolated into 1-m intervals. From the lowest flight altitude ( 150 m) to sea level, the sounding data were extrapolated by assuming that the lowest levels were well mixed with constant potential temperature and mixing ratio. Because the upstream sounding represented the large-scale conditions, we assume that the large-scale environment was unchanged from the afternoon of 2 August to the next morning. Based on this assumption, virtual temperature differences (dt v ) for the analysis period, determined by subtracting the upstream virtual temperature at the same altitude from the PAM virtual temperature, were calculated. The procedures for calculating dt v are identical to those used by Chen and Wang (1994). During the late afternoon and the early evening of 2 August, tethersonde observations of winds, pressure, temperature, and wet-bulb temperature were made at UHAES (Fig. 1). During the evening transition period, the tethersonde was launched up to 160 m at 1800, 1830, 1900, 1955, and 2030 HST, respectively. The vertical resolution of the original tethersonde data was about 3.7 m. A careful manual check was used to remove spikes in the data. The accuracy and resolution for the tethersonde observations are discussed by Wang and Chen (1995). In this study, PAM data, virtual temperature differences, and tethersonde data were used to analyze the evolution of the low-level winds and thermodynamic fields over the lower slope/coastal region on the windward side. Two C-band Doppler radars were located at Paradise Park (CP3) and Hilo airport (CP4), respectively (Fig. 1). The CP4 radar was operated continuously from 1400 HST 2 August to 1000 HST 3 August. Most of the radar data analyzed in this study were obtained from surveillance scans and RHI scans by the CP4 radar. A typical set of surveillance volume scan data for the CP4 radar is obtained along three elevation angles (0.5, 2.5, and 4.5 ). For the 4.5 elevation, the radar beam is about 0.8 km (1.6 km) above sea level at a distance of 10 (20) km away from the radar site. To the west of the radar site (Fig. 1), the radar beam (such as along the 270 azimuth) is about 0.3 0.7 km above the sloped surface in the range of 10 20 km from the radar. The 4.5 surveillance scans provide uncontaminated views of radar echoes over the lowland/slope areas for studying the evolution of convective activities from the early afternoon to the early morning. Furthermore, RHI scans at FIG. 2. The large-scale pressure (P-1000 hpa) pattern at 0200 HST 3 Aug 1990. azimuths directed eastward over the ocean and westward over land were used to investigate vertical structure of the katabatic flow and rainbands. 3. Weather conditions From 2 to 3 August, a midlatitude cold front moves toward the mid-pacific Ocean from the northwest. At 0200 HST 3 August, the subtropical high is located at 40 N, 170 W, northwest of the Hawaiian Islands (Fig. 2). The upstream sounding on the early morning of 3 August shows trade winds on the order of 7 9 m s 1 with inversion at 1.8 km (Fig. 3a), which is slightly lower than the typical inversion height ( 2.2 km) during HaRP. Except for a relatively low inversion height, the trade wind conditions during the period are rather typical. The Hilo sounding, launched at 2000 HST 2 August, also indicates similar trade wind conditions over the coastal region (Fig. 3b). For an air parcel at the 1000-hPa level, the Hilo sounding shows that both the lifting condensation level (LCL) and the level of free convection (LFC) are at 967 hpa ( 0.45 km) above the sea level (Fig. 3b) with the convective available potential energy of 45Jkg 1. From the early afternoon of 2 August to the morning of 3 August, there are three major rainfall peaks over the lowland/coastal areas around 2100 HST, 0200 HST, and 0700 HST. The spatial distributions of the 4-h rainfall around these peak rain periods are presented in Fig. 4. For each period, the rainfall amount is slightly below the HaRP median value for the corresponding same period. This may be related to the large-scale undisturbed trade wind conditions with a relatively low inversion height (Chen and Feng 1995) for this particular case.

2678 MONTHLY WEATHER REVIEW FIG. 3. Skew T logp diagram and the wind profile for (a) the aircraft sounding obtained from NCAR Electra, approximately 120 km upstream of the island during 0520 0541 HST 3 Aug, and (b) Hilo sounding observed at 2000 HST 2 Aug. Winds with one pennant, full barb, and half barb represent 5, 1, and 0.5 m s 1, respectively. During the study period, 4-h rainfall has a maximum during 1900 2300 HST (Fig. 4), in phase with the mean diurnal cycle of rainfall on the windward side during HaRP found by Chen and Feng (1995). 4. Inland rainband in the evening a. Katabatic flow onset and rainband initiation During the daytime, the upslope/trade wind flow prevails on the slopes of the island and brings in moisture from low elevations to the slopes (Chen and Nash 1994). PAM station 15 (405 m) and station 16 (1128 m) on the windward slope (Fig. 1) receive daily maximum net radiation in the early afternoon of 2 August (Figs. 5a and 6a). The dt v s between the surface air on the slope and the upstream values at the same altitudes for these two PAM stations reach the daily maximum at 1315 HST. The spatial distribution of the virtual temperature differences at this time shows relatively lower positive values along the 500-m terrain contour over the lower slopes than other regions (Fig. 7a). In the early afternoon, no radar echoes with reflectivities greater than 0 dbz are observed over the slopes (not shown). Similar to the 10 August case analyzed by Feng and Chen (1998), the minimum in spatial distribution of dt v on the lower slopes in the early afternoon is a result of reduced solar heating by orographic clouds. The visible satellite image for 1439 HST shows that the areas of the minimum dt v (Fig. 7b) are well correlated with the regions covered by orographic clouds (Fig. 8). The areas with a large dt v ( 4 K) on the windward side (Fig. 7b) are semiarid regions above the trade wind inversion and along the southeast coast where skies are clear (Fig. 8). For this particular case, the anabatic/trade wind flow turned to a katabatic flow first on the lower slope near station 15 (Figs. 9a,b), consistent with the composite analyses presented by Chen and Nash (1994). The onset of katabatic flow occurs there at 1720 HST with a negative dt v (Figs. 5b,c), in agreement with the composite analysis presented by Chen and Wang (1994). From 1400 to 1800 HST, the surface air temperature at station 15 decreased with small fluctuations (Fig. 5c). The temporal variations of the surface air temperature are well correlated with those of the net radiation (Fig. 5a). These results suggest that, similar to the 10 August case analyzed by Feng and Chen (1998), the decrease

NOVEMBER 1999 LI AND CHEN 2679 FIG. 4. The 4-h rainfall (mm) distribution on the windward area: (a) 1900 2300 HST, (b) 2300 0300 HST, and (c) 0300 0700 HST. Terrain contours every 250 m. in the surface air temperature in the late afternoon is primarily related to the diurnal heating cycle. The change from positive to negative virtual temperature difference first occurs near station 15 (Fig. 9b) with a small negative net radiation value (Fig. 5a). The net radiation increases slightly at 1740 HST and becomes negative again after 1800 HST. During HaRP the net radiation frequently becomes negative more than 1 h before sunset ( 1900 HST) because of the longwave radiative heat loss from the surface. Furthermore, the sun drops behind the mountains before sunset. Rainfall data are missing at station 15. Just before the onset of the katabatic flow at station 15, small weak echoes occur within the valley near and west of station 15. Nevertheless, it is unlikely that the katabatic flow onset at station 15 is a result of orographic rain showers. Note that at the onset of the katabatic flow at station 15 ( 1720 HST), there is no abrupt temperature drop or an increase in the moisture content (Fig. 5c) prior to or accompanying the wind shift. A decrease in both the surface air temperature and dewpoint occurs immediate after the wind shift. The absence of moistening accompanying the cooling after the wind shift suggests that at the initial stage the cold pool behind the drainage front is not primarily caused by rain evaporative cooling as suggested by Carbone et al. (1995) for this particular case. The earliest katabatic flow onset may occur on the slope west of station 15 rather than at station 15. The decrease in both the surface air temperature and dewpoint after the wind shift may be partially caused by the advection from the higher slope surface west of station 15 as the colder and drier katabatic flow moves downslope. At 1730 HST, 10 min after the onset of katabatic flow at station 15, isolated and circular echoes are observed over the lower slopes and the lowland/coastal region (Fig. 9b). Some of the scattered echoes form over the lower slopes ( 0.25 km elevation). The others originate from trade wind cumuli over lowland/coastal region and are pushed toward the slopes by trade winds. In the next hour, these isolated echoes develop rapidly and align along the drainage front (Fig. 9c). By 1841 HST, a band of rain showers forms along the surface drainage front over the lower slopes (Fig. 9d). In this study, we refer to this convective line as an inland rainband. Note that during the initial stage of the inland rainband, the radar echoes are located at the leading edge of the cold pool where dt v equals zero (Figs. 9c,d) rather than over the cold pool. This result provides further support that for this case the cold pool behind the drainage front during the onset of katabatic flow is not primarily caused by rain evaporative cooling.

2680 MONTHLY WEATHER REVIEW FIG. 5. Time series from 1200 to 2300 HST on 2 Aug 1990 at station 15: (a) Net radiation (W m 2 ); (b) wind direction (, heavy line) and speed (m s 1, light line); and (c) rainfall amount (mm/ min, thin vertical bars with scale on the left), temperature ( C, heavy line), dewpoint ( C, light line), and the upstream temperature at the same altitude as the station. Note that the rainfall data for this station were not available during the analysis period. Station 16, which is west of station 15 (Fig. 1), recorded the first orographic rain showers before 1655 HST (Fig. 6c). With daytime anabatic winds and decreasing temperature in the afternoon hours, station 16 becomes saturated just before 1600 HST. In addition to general decreasing trend in the surface air temperature, there are slight abrupt drops in surface air temperature and dewpoint at 1655 HST, 1750 HST, and 1840 HST, corresponding to rain shower activities at these times, respectively (Fig. 6c). Nevertheless, the katabatic flow onset at station 16 does not occur until 1915 HST (Fig. 6b). Furthermore, with saturated conditions, it is unlikely that rain evaporative cooling (Carbone et al. 1995) plays a significant role in the production of negative virtual temperature difference at station 16 for this case. Note that similar to station 15, the temporal variations of the surface air temperature in the afternoon hours are closely related to those of the net radiation. Both the negative virtual temperature difference and the onset of katabatic flow at station 16 occur almost 2 h later than those at station 15. These differences between these two stations are related to higher positive virtual temperature difference at station 16 than station 15 in the early afternoon (Fig. 7) and near saturated conditions with extensive orographic cloud cover at station 16 in the late afternoon with a slow temperature decrease after saturation (Fig. 6c). It is apparent that the cooling in the afternoon hours at station 16 is mainly controlled by the diurnal heating cycle and modified by orographic clouds and scattered rain showers in the late afternoon. Chen and Wang (1995) found that for the rain cases, the extensive orographic cloud cover on the upper slope delays the katabatic flow onset there with a slower decrease in the surface air temperature than the dry cases whereas in the Hilo area, the katabatic flow onset occurs at an earlier time for the rain cases because of the cold pool production by rain evaporation. Based on numerical simulations, Smolarkiewicz et al. (1988) suggest that the island blocking of the trade wind flow is principally responsible for formation of cloud bands in the quasi-stationary low-level convergent zone offshore between the dynamically driven return flow and the incoming trade winds. HaRP observational studies show that the onset of westerly flow on the windward

NOVEMBER 1999 LI AND CHEN 2681 FIG. 6. Same as Fig. 5 but for station 16. side of the island of Hawaii in the late afternoon and early evening is thermally driven rather than dynamically driven (Chen and Nash 1994; Chen and Wang 1994; Carbone et al. 1995). Tethersonde observations at UHAES (Fig. 1) show that within the first hour after the onset of the katabatic flow, its depth is less than 0.15 km (Fig. 10). Analysis of tethersonde data for other HaRP cases also indicates that the katabatic flow is very shallow ( 0.2 km) over the lowlands immediately after the onset (Wang and Chen 1995). Based on the Hilo sounding, an air parcel at the 0.25-km elevation would become saturated only if it is lifted 0.2 km or more. Apparently, immediately after the katabatic flow onset, the drainage front alone is not deep enough to lift the moist air to saturation. Note that the terrain slope is about 0.5 km/10 km over the lowlands and the lower slopes. A surface air parcel will be lifted by the orography above LFC ( 0.45 km) if it travels 10 km westward from the coastal region with the trade wind/upslope flow. The inland rainband forms with additional lifting along the drainage front. Thus, the formation of the inland rainband is a result of the combination of orographic lifting and the low-level forcing along the drainage front over the lower slopes and lowlands as suggested by Chen and Nash (1994). b. Rainband evolution and interaction with the low-level airflow During the initial stage, the inland rainband is located at the leading edge of the drainage front over the lower slopes (Figs. 9d and 11a). The rainband consists of several rain cells that have maximum reflectivities of 10 20 dbz. The width of the rainband (the reflectivity 0 dbz) is about 5 km. From 1941 to 2144 HST, both the intensity and the width of the rainband increase significantly over the lowland/coastal region (Figs. 11b,c). During this period, most of the radar echoes associated with the rainband are west of the leading edge of the drainage front (Figs. 11b,c and 12a,b). Individual rain cells develop along the drainage front, move westward, and weaken. They dissipate on the upper slopes as they move away from the low-level forcing zone along the drainage front. Consistent with the life cycle of rain cells and their westward movement, the horizontal distribution of reflectivities exhibits higher reflectivities on the eastern (forward) edge of the rainband than those on the western side (Figs. 11b,c). With the downward progression of the drainage flow, the rainband as a whole moves eastward to the coast, in agreement with the composite analysis presented by Chen and Nash (1994) and

2682 MONTHLY WEATHER REVIEW FIG. 8. The Advanced Very High Resolution Radiometer satellite visible image over the island at 1439 HST 2 Aug 1990. FIG. 7. Surface winds and differences of the virtual temperature ( C) between PAM stations and the upstream at the same altitudes for (a) 1315 HST and (b) 1430 HST. Winds with one pennant, full barb, and half barb represent 5, 1, and 0.5 m s 1, respectively. Terrain contours every 250 m. the 10 August case analyzed by Feng and Chen (1998). At 2248 HST, the rainband has a maximum reflectivity of 35 dbz and a width of 15 km (Fig. 11d). A CP4 RHI scan along 271 azimuth (looking westward toward the mountain) shows that the echo tops (0 dbz) of the rainband are at 2 km (Fig. 13a). During the evening, the rain showers significantly influence the depth of the katabatic flow over the lowland/coastal region as found by Wang and Chen (1995) and Carbone et al. (1998). Figure 13b shows the vertical cross sections of radial velocities and reflectivities, obtained from a CP4 RHI scan along 90.9 azimuth (looking eastward toward the ocean) at 2252 HST. At this time, the rainband is still over land (Fig. 11d). The leading edge of the katabatic flow has moved off the Hilo coast (about 5 km east of the radar site) (Fig. 13b). Over the coastal region, the katabatic flow ahead (east) of the rainband is shallow ( 0.2 km) and weak ( 2 m s 1 ) (Fig. 13b). The shallow drainage front at the coast is also observed at 2309 HST for the 10 August case (Fig. 17a in Feng and Chen 1998). In contrast, the katabatic flow in the convective region over the lowlands has a depth of 0.4 km with a westerly wind component ( 4 ms 1 ) in the lowest layer (Fig. 13a). Wang and Chen (1995) analyzed the near-surface winds and thermal profiles on the windward area during HaRP. They found that for rainy cases during HaRP, the nocturnal inversion and the nocturnal wind maximum are weaker with a deeper stable katabatic flow than for dry cases. The deepening of the westerly katabatic flow over the convective region is a result of rain evaporative cooling in the trade wind layer aloft (Wang and Chen 1995). Carbone et al. (1998) also show that the rain-cooled westerly flow resembles a density current with a developing cell on the nose. After the katabatic flow moves over the coastal region and off the coast, generation of new cells ceases. Without orographic lifting, the leading edge of the drainage front is not deep enough (Fig. 13b) to lift the low-level air to saturation and to generate convective cells. The inland rainband moves westward and weakens (Fig. 11e). It finally dissipates over the slopes (not shown). The inland rainband lasts about 6 h and produces the heaviest rainfall over the windward lowland/coastal region during the analysis period (Fig. 4a), in phase with

NOVEMBER 1999 LI AND CHEN 2683 FIG. 9. Surface winds and CP4 radar reflectivities (dbz) for the 4.5 elevation angle and differences of the virtual temperature (dt v,in C) between PAM stations and the upstream at the same altitudes for (a) 1626 HST, (b) 1732 HST, (c) 1826 HST, and (d) 1841 HST. Winds with one pennant, full barb, and half barb represent 5, 1, and 0.5 m s 1, respectively. The reflectivity contours start from 10 dbz with 10-dBZ intervals. The shaded scales are shown on the upper right-hand corner. The dt v contours every 1 C. The terrain contours every 250 m. the diurnal variations of rainfall presented by Chen and Feng (1995). 5. Offshore rainband after midnight a. Initiation and evolution of the rainband offshore At 2254 HST, the katabatic flow is deeper over the lowlands ( 0.4 km, Fig. 13a) than at the coast ( 0.2 km, Fig. 13b). During 2200 0300 HST, the surface air is potentially colder over the lowlands than along the coast (Fig. 14), allowing the surface cold air in windward lowlands to move toward the coast (Feng and Chen 1998). With cold air advection from the deep katabatic flow region over the lowlands to the coast, the katabatic flow at the coast gradually deepens. By 0207 HST, the depth of the katabatic/offshore flow there is greater than 0.45 km with a maximum westerly wind component of 5 ms 1 (Fig. 13c). Note that the top of the katabatic/offshore flow is near the LFC ( 0.45 km, Fig. 3). It is apparent that at this time the leading edge of the drainage front offshore is deep enough to initiate free convection. A plan view of the surface airflow shows that at 0130 HST, the katabatic flow is over most of the coastal areas except the eastern corner of the island (Fig. 15a). The surface cold pool over the Hilo region has a thermal contrast (dt v )of 3 K (Fig. 12d). The radar observations at this time indicate that several rain cells develop along the coast (Fig. 15a). These cells rapidly merge into a convective line offshore along the coast (Fig. 15b). At 0207 HST, the offshore rainband has a maximum reflectivity of 38 dbz with echo tops (0 dbz) at 2.2 km (Fig. 13c). High radar reflectivities ( 30 dbz) are located over the low-level convergent zone between the deep katabatic flow at the coast and the incoming trade winds (Fig. 13c). We refer to this rainband as the offshore rainband after midnight. Once formed, the rainband moves westward with the trade wind aloft (Takahashi 1981; Schroeder et al. 1977). It reaches the maximum intensity ( 40 dbz) over the coastal region before 0300 HST (Fig. 15d). After 0300

2684 MONTHLY WEATHER REVIEW observed arc-shaped cloud line (Takahashi 1977b) but insufficient to trigger free convection and rainbands. For the 3 August case, our study shows that deepening of the katabatic flow offshore due to the advection of radiatively and rain cooled air from the lower slopes played an important role in triggering free convection and the offshore rainband. The offshore flow is not dynamically driven as suggested by Smolarkiewicz et al. (1988). FIG. 10. Time series of low-level winds and potential temperatures obtained from the tethersonde observations at UHAES (Fig. 1) during 1830 2030 HST 2 Aug. Winds with one pennant, full barb, and half barb represent 5, 1, and 0.5 m s 1, respectively. Potential temperature contours every 0.5 K. HST, the rainband weakens and finally dissipates over the slopes (Figs. 15e,f). It lasts about 3 h and brings in moderate rainfall over the lowland/coastal region after midnight (Fig. 3). Feng and Chen (1998) analyzed a relatively dry case (10 August 1990). They observe an inland rainband during the evening transition, but no offshore rainbands during 0100 0400 HST. For their case, the Hilo sounding indicates that both the LFC and the LCL for an air parcel at the 1000-hPa level are above the 0.7-km level. The radar observations show that at 0219 HST 10 August, the leading edge of the katabatic flow extends several kilometers offshore, similar to our case. However, the depth of the offshore flow is only about 0.3 km (Fig. 17b in Feng and Chen 1998), shallower than our case. It is apparent that for the 10 August case, the drainage flow offshore is not deep enough to lift the incoming trade wind air to reach the condensation level. It appears that in the absence of orographic lifting, a deep katabatic flow is needed to lift the moist air to saturation and initiate convection along the leading edge of the drainage front offshore. Based on laboratory results (Hunt and Snyder 1980) and modeling studies (e.g., Smolarkiewicz et al. 1988), vertical displacements at the base of an obstacle due to dynamic blocking are on the order of B Fr h, where Fr is the Froude number and h is the height of the obstacle. When related to Hawaiian cases, B is 0.5 0.6 km. (Smolarkiewicz et al. 1988). Comparing B with the LCL and the LFC, Austin et al. (1996) found that lifting at the flow separation line predicted by the model is frequently sufficient to lift the air to the LCL and produce the commonly b. Interaction between the offshore rainband and the katabatic flow As the rainband moves westward into the deep ( 0.45 km) katabatic flow at the coast, both the depth and the intensity of the katabatic flow decreases significantly. At 0252 HST, the rainband is located around the CP4 radar site (Figs. 13d,e). The katabatic flow over the coastal region becomes shallow ( 0.25 km in depth) and weak (less than 2 m s 1, Fig. 13d) as compared to the katabatic flow at the initial stage of the rainband (Fig. 13c). Within 4 km west of the radar site, the easterly winds are observed in the lowest 0.2 km (Fig. 13e). By 0311 HST, the easterly flow dominates with a maximum radial velocity greater than 3 m s 1 in the lowest 0.4-km layer over the coastal region (not shown). The replacement of the katabatic flow by easterly winds during the passage of the offshore rainband is also observed at surface PAM stations over the lowland/coastal region (Fig. 14). Frye and Chen (1997) analyzed rain effects on the katabatic flow for an extremely strong trade wind case (7 8 August) during HaRP and showed that the westward-moving bands of rain showers brought back easterly trade winds when precipitation fell on coastal stations that were under deep katabatic flow. Because the top of the deep katabatic flow over the coastal region is near the cloud base, or the LCL ( 0.45 km, Fig. 3b), before the arrival of the offshore rainband (Fig. 13c), it is unlikely that evaporative cooling beneath the cloud base would result in further deepening of the katabatic/ offshore flow. Instead, katabatic winds in the lowest layer are replaced by easterly winds during rainy periods. The destruction of katabatic winds over the coastal region may be related to the vertical transport of easterly momentum aloft into the lowest katabatic flow layer in convective areas as suggested by Frye and Chen (1997). 6. Early morning trade wind rain showers On the early morning of 3 August, similar to the cases analyzed by Ochs et al. (1996) and Wang and Chen (1998), the trade wind rain showers originate far upstream ( 50 km) of the island and move toward the windward coast. During 0500 0900 HST, two groups of trade wind rain showers move into the coastal region. The first group brings in moderate rainfall along the

NOVEMBER 1999 LI AND CHEN 2685 FIG. 11. Surface winds and CP4 radar reflectivities (dbz) for the 4.5 elevation angle for (a) 1941HST, (b) 2051 HST, (c) 2144 HST, (d) 2248 HST, and (e) 2345 HST. Winds with one pennant, full barb, and half barb represent 5, 1, and 0.5 m s 1, respectively. The reflectivity contours start from 10 dbz with 10-dBZ intervals. The shaded scales are shown on the upper right-hand corner. The terrain contours every 250 m. coast. The second group produces much rainfall offshore during 0700 0900 HST, but little precipitation over the lowland/coastal region. In this section, we examine the evolution of these early morning trade wind rain showers and their interaction with the mesoscale airflow over the coastal region. At 0400 HST, the first group of trade wind rain showers, oriented in the northwest southeast direction, are located 30 km northeast of the Hilo coast (Fig. 16a). Before 0500 HST, shallow ( 0.2 km) katabatic (offshore) flow is observed from both the RHI (Fig. 17a) and the surveillance scans at the lowest elevation angle ( 0.5 ). At 0558 HST, the first group of trade wind rain showers moves into the low-level convergent zone along the surface drainage front, about 18 km northeast of the CP4 radar site. As a result, both the area and the intensity of the radar echoes increase (Fig. 16b). The enhancement of convection associated with trade wind rainbands over the low-level convergent zone offshore in the early morning is in agreement with previous HaRP

2686 MONTHLY WEATHER REVIEW FIG. 12. Same as in Fig. 7 but for (a) 2051 HST, (b) 2144 HST, (c) 2248 HST, and (d) 0130 HST. studies (Frye and Chen 1997; Wang and Chen 1998). The trade wind rain showers are also intercepted by a preexisting northeast southwest-oriented convective line over the lowland/coastal region southeast of Hilo (Fig. 16b). The first group of trade wind rain showers weakens as it moves over the katabatic/offshore flow in the Hilo bay area (Fig. 17b). During this period, a sequence of surveillance scans at the lowest elevation angle ( 0.5 ) show an area of easterly flow ( 2 ms 1 ) associated with the rain showers within the katabatic (offshore) flow region (not shown). Consistent with the horizontal distribution of radial velocities observed on the surveillance scans, an RHI scan along the 90 azimuth at 0628 HST also indicates that easterly winds replace the katabatic/offshore flow on the western side of the rain showers in the lowest 0.5 km, whereas on the eastern side of the rain showers the westerly flow extends to about 20 km east of the radar site (Fig. 17b). Such a flow pattern is very pronounced when the rain showers weaken. At 0714 HST, the rain showers arrive at the coast with radar reflectivities less than 25 dbz (Fig. 17c). In the lowest levels, both the westerly flow on the eastern side and the easterly flow to the west of the rain showers deepen ( 0.6 km) (Fig. 17c), indicating a deep low-level outflow associated with weakening trade wind rain showers as a part of the storm circulation. Above the easterly flow on the western side of the rain shower, a westerly wind component is observed between 0.8 and 1.2 km above the sea level (Fig. 17c). At 0813 HST, the deep ( 0.6 km) offshore flow extends as far as 22 km east of the radar site (Fig. 17d). The second group of trade wind rain showers in the early morning moves toward the coast from the northeast. At 0557 HST, they are located at 40 km northeast of the coast (Fig. 16b). About an hour later, the convection associated with the second group of rain showers is organized into a well-defined rainband (Fig. 16c). At 0741 HST, the rainband reaches its maximum intensity with a width of 20 km and a maximum reflectivity of 45 dbz (Fig. 16d) as it moves into the convergent zone where the deep ( 0.6 km) cold offshore flow converges with the incoming trade winds. At 1813 HST, the echo tops of the rainband are at 2.5 km (Fig. 17d), about 0.7 km above the inversion height ( 1.8 km, Fig. 3a). These results suggest that during the early morning, as found by Wang and Chen (1988) for the August 22 case, deepening of the westerly flow offshore by the first group of trade wind rain showers has a significant influence on the intensification of

NOVEMBER 1999 LI AND CHEN 2687 FIG. 13. RHI scans obtained from the CP4 radar along the 2708 2718 azimuth [(a) and (e), looking westward toward the mountains] or along the 898 918azimuth [(b) (d, looking eastward toward the ocean] for (a) 2254 HST and (b) 2252 HST 2 Aug, and (c) 0207 HST, (d) 0252 HST, and (e) 0256 HST 3 Aug. The left panels show radar reflectivities at 10-dBZ intervals with seven color scales. The right panels show Doppler radial velocities every 1 m s21 with 12 color scales. The positive (negative) radial velocities in (b) (d) represent westerly (easterly) flow. (a) and (e) The positive (negative) radial velocities represent easterly (westerly) flow. For all the velocity panels the warm colors (yellow orange red) and the cold colors (blue purple violet) represent westerly and easterly flow, respectively. The horizontal and vertical spacings (between the 1 symbols) are 1 km and 0.5 km, respectively. the subsequent incoming trade wind rain showers (or rainband). The trade wind rainband weakens as it moves westward over the deep offshore flow (Fig. 18). When it arrives at the coast around 0856 HST, the rainband only has a width less than 10 km and a maximum radar reflectivity less than 20 dbz (Fig. 18). During the weakening period, the easterly winds are observed in the lowest layer on the western (forward) portion of the rainband with weak (#2 m s21 ) westerly flow above (Fig. 17e). The rainband moves into the coastal region during the morning transition and produces little rainfall along the coast. 7. Summary and concluding remarks Nocturnal rain showers on the windward side of the island of Hawaii were investigated from the late afternoon of 2 August to the early morning of 3 August during HaRP. Three types of rainbands are observed. These rainbands exhibit different features in terms of origin (location and timing), evolution, intensity, and their effects on the island-induced airflow. They produce rainfall peaks over the lowland/coast region during the evening, after midnight, and in the early morning, respectively. The heaviest rainfall occurs in the early evening hours.

2688 MONTHLY WEATHER REVIEW FIG. 14. Time series of surface winds and virtual potential temperatures (K) at PAM stations 9, 13, 44, 15, and 16. Winds with one pennant, full barb, and half barb represent 5, 1, and 0.5 ms 1, respectively. The surface air temperatures on the windward slopes reach their daily maxima around noon and decrease in the afternoon hours. The variations of the surface air temperatures are well correlated with the net all-wave radiation at the surface. From the early afternoon to the onset of katabatic flow, the smallest virtual temperature difference between the surface air on the slope surface and its upstream environment occurs on the lower slopes because orographic clouds there reduce the net radiation at the surface. The change from positive to negative virtual temperature difference first occurs near station 15 (elevation 405 m). As a result, the katabatic flow is first observed there during the evening transition period. Immediately after the onset of the katabatic flow at station 15, an inland rainband develops along the drainage front over the lower slopes as a result of orographic lifting and the low-level forcing along the drainage front. During the evening transition, individual cells associated with the inland rainband develop along the surface drainage front. They move westward away from the low-level forcing zone and dissipate. With the downward progression of the shallow ( 0.2 km) katabatic flow toward the coast, new cells are generated east of the existing cells. As a result, the inland rainband as a whole moves eastward down the coast. After the drainage front moves over the coastal region and off the coast, generation of new cells along the drainage front ceases. Without orographic lifting over the coastal region and offshore, the shallow drainage front is not deep enough to lift the low-level air to the level of free convection ( 0.45 km). The rainband lasts about 6 h and produces the heaviest rain showers over the lowland/coastal region during the analysis period. During the evening, most of the rain showers associate with the inland rainband occur over the lowland/ coastal region west of the surface drainage front. The rain evaporative cooling aloft deepens the cool pool behind the drainage front over land with a deeper westerly katabatic flow. During 2200 0300 HST, the surface air is potentially colder over the lowlands than along the coast, allowing the cold air in windward lowlands

NOVEMBER 1999 LI AND CHEN 2689 FIG. 15. Same as in Fig. 11 but for (a) 0130 HST, (b) 0145 HST, (c) 0200 HST, (d) 0300 HST, (e) 0400 HST, and (f) 0503 HST. to move down to the coast. With the cold-air advection from the deep katabatic flow over the lower slopes to the coast, the katabatic flow at the coast gradually deepens. By 0207 HST, the depth of the katabatic/offshore flow at the coast increases to 0.45 km, which is above the LFC. A rainband develops along the drainage front offshore at this time. Once formed, the offshore rainband moves westward toward the island with the trade winds aloft. In the lowest layer, the westerly katabatic flow is replaced by easterly winds as the offshore rainband moves inland. The rainband weakens and dissipates over the lower slopes. It lasts 3 h and produces moderate rainfall over the lowland/coastal region. During the early morning, two groups of trade wind rain showers move into the coastal region from the northeast. As the first group of trade wind rain showers moves into the low-level convergent zone between the cold offshore flow and the trade winds, the trade wind rain showers are enhanced. After they continue to move westward into the offshore flow region near the coast, they weaken with pronounced outflow in the lowest levels. On the east side of the rain showers, the westerly

2690 MONTHLY WEATHER REVIEW FIG. 16. Same as in Fig. 11 but for radar reflectivities for the 2.5 elevation angle: (a) 0400 HST, (b) 0557 HST, (c) 0657 HST, and (d) 0741 HST. flow deepens and extends farther eastward over the ocean. Interacting with a deeper offshore flow, the second group of trade wind rain showers is organized into a well-defined rainband and is more intense as it moves into the convergent zone. The rainband weakens as it moves westward over the deep westerly flow. During the weakening period, the westerly flow in the lowest layer is replaced by easterly flow, whereas on the eastern side, the offshore flow extends farther over the ocean. The nocturnal rain showers in the evening and after midnight during the night of 2 3 August are caused mainly by an inland rainband and an offshore rainband, respectively. Results from this study show that in addition to thermal forcing (nocturnal cooling and rain evaporation), the presence of the mountainous terrain along the windward coast is an important factor for the production of nocturnal rain showers. In the late afternoon, the onset of katabatic flow first occurs on the lower slopes. The development of the inland rainband is caused by the orographic lifting superimposed by the low-level forcing along the drainage front. After midnight, because of the advection of the radiatively and rain-cooled air from the slopes toward the coast, the leading edge of the katabatic flow offshore becomes deep enough to lift the moist incoming trade wind flow to the level of free convection. As a result, the offshore rainband develops. Except for differences in size and shape, the island of Taiwan is under a similar low Froude number (Fr 0.3) flow regime under the prevailing southwesterly monsoon flow during the early summer rainy season (Li and Chen 1998) with a pronounced diurnal cycle in airflow (Chen and Li 1995; Yeh and Chen 1998). Nevertheless, the pronounced nocturnal rainfall along the windward coast found over the island of Hawaii is not evident over the island of Taiwan (Yeh and Chen 1998). The absence of pronounced noc-

NOVEMBER 1999 LI AND CHEN 2691 FIG. 17. RHI scans obtained from the CP4 radar along the 898 918 azimuth (looking eastward toward the ocean) for (a) 0452 HST, (b) 0628 HST, (c) 0714 HST, (d) 0813 HST, and (e) 0858 HST 3 Aug 1990. The left panels show radar reflectivities at 10-dBZ intervals with seven color scales. The right panels show Doppler radial velocities every 1 m s21 with 12 color scales. The positive (negative) radial velocities represent westerly (easterly) flow. The horizontal and vertical spacings (between the 1 symbols) are 2 km and 1 km, respectively. turnal rain showers along the windward coast of Taiwan may be related to the presence of the southwestern plain there. It appears that without steep terrains along the windward coast, the shallow katabatic flow alone may not be deep enough to initiate frequent nocturnal rain showers there. Acknowledgments. We would like to thank those who participated in HaRP, especially the Electra, radar, and tethersonde crews whose dedication made this work possible. This work is supported by the National Science Foundation under Grant ATM-9629886. Acknowledgment is made to the National Center for Atmospheric FIG. 18. Same as in Fig. 11 but for (a) 0815 HST and (b) 0856 HST.