956 MONTHLY WEATHER REVIEW VOLUME 129

Evolution of Downslope Flow under Strong Opposing Trade and Frequent Trade- Rainshowers over the Island of

JEFFREY L. FRYE* AND YI-LENG CHEN Department of , School of Ocean and Science and Technology, University of Hawaii at Manoa, Honolulu, Hawaii

(Manuscript received 6 August 1999, in ®nal form 6 August 2000)

ABSTRACT The evolution of downslope ¯ow on the windward side of the island of Hawaii during 7±8 August 1990 is investigated. This period is characterized by atypical strong (ϳ11 m sϪ1) upstream trade winds and frequent nocturnal rainshowers. In the late afternoon, the onset of downslope ¯ow ®rst occurs on the upper slope, which is frequently dominated by orographic clouds. With mean weak surface winds on the windward slopes because of island blocking, the downslope ¯ow onset occurs after the slope surface becomes negatively buoyant. Along the Hilo coastal areas that are well exposed to decelerating trade-wind ¯ow, the offshore ¯ow onset there occurs much later (ϳ2 h) than the Hawaiian Rainband Project (HaRP) mean as the drainage front moves slowly from the windward lowlands toward the coast against the atypical strong incoming trade-wind ¯ow. The downslope ¯ow onset along the coast occurs after a band of trade-wind rainshowers has moved over the coast. These rainshowers produce evaporatively cooled air in the lowest levels and allow the westerly downslope ¯ow to extend offshore. Nevertheless, throughout the night, the horizontal extent of the offshore ¯ow is limited (Ͻ10 km) as the leading edge of the offshore ¯ow encounters increasing trade-wind ¯ow farther offshore. During 0300±0500 Hawaiian Standard Time (HST), a second group of rainshowers moves over the island. With a small horizontal extent and a weak density current structure, offshore ¯ow is temporarily destroyed by rainshowers as a result of vertical mixing. It retreats to windward lowlands as trade-wind rainshowers move onshore. Signi®cant enhancement and focusing of the scattered trade-wind rainshowers occur over the conver- gence zone inland. The surface rainfall forcing is also enhanced by orographic lifting aloft producing the wettest early morning 0300±0700 HST period observed during HaRP. During the early morning, the onset of onshore ¯ow along the coast occurs approximately 1±2 h earlier than the HaRP wet cases. Immediately before the onset of the onshore ¯ow, the surface air along the coast is still negatively buoyant. It appears that the onset of onshore ¯ow is caused by the retrograde motion of the leading edge of the offshore ¯ow. With decreasing buoyancy de®cit after sunrise, the gravity current retreats westward, allowing the warm, moist trade winds offshore to return. The retrograde motion of the density current is more likely to occur in the early morning after sunrise, if the opposing trade winds are stronger.

1. Introduction The Hawaiian Rainband Project (HaRP) was con- ducted during July±August 1990 over the island of Ha- The largest island in the Hawaiian chain, Hawaii is waii to better understand the interaction between steady dominated by two volcanic mountains exceeding 4 km trade-wind ¯ow and an island obstacle. It is the ®rst in elevation, well above the typical height of the trade- time that comprehensive datasets were collected by ad- wind inversion (ϳ2.2 km). In the summer months, the vanced instruments over a subtropical island. During nearly continuous exposure to easterly trade winds HaRP, high-resolution datasets were collected from 50 makes the island ideal for studying air¯ow and precip- National Center for Atmospheric Research (NCAR) itation patterns due to the effects of orographic lifting, Portable Automated Mesonet (PAM) stations, research dynamic blocking, and thermally driven circulations. aircraft (NCAR Electra), tethersondes, rawinsondes, a National Oceanic and Atmospheric Administration (NOAA) boundary layer wind pro®ler, and a dual-Dopp- * Current af®liation: United States Transportation Command, Scott ler network (NCAR CP-3 and CP-4) (Fig. 1). One of Air Force Base, Illinois. the major goals of HaRP was to document the structure and evolution of rainbands frequently observed along the windward coast during the morning hours. Corresponding author address: Dr. Yi-Leng Chen, Department of Meteorology, SOEST, University of Hawaii at Manoa, Honolulu, HI Leopold (1949) described the trade-wind inversion as 96822. a ``lid'' forcing trade-wind ¯ow to move around the E-mail: [email protected] island. He suggested that the formation of band clouds

᭧ 2001 American Meteorological Society

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and depth of the offshore ¯ow. They maintain, however, that the convergence line is primarily a result of dynamic forcing. Analyses of the HaRP data show the effects of thermal forcing on mesoscale circulations. In contrast to mod- eling studies, HaRP analyses show that the nighttime downslope ¯ow on the windward side is not pure dy- namically driven reversed ¯ow. Chen and Nash (1994) present a detailed analysis of the PAM data and show that in regions of weak mean surface winds caused by island blocking, the thermally driven winds become sig- ni®cant. They describe windward rainfall as a complex interaction of orographic lifting, thermal forcing, and dynamic blocking. Chen and Wang (1994) found that on the windward slopes and lowlands with weak mean surface winds, the onset of upslope (downslope) ¯ow is closely related to the thermal contrast between slope surface and upstream environment at the same altitude. Carbone et al. (1995) are in agreement with previous HaRP analyses that show that the westerly ¯ow on the FIG. 1. Windward Hawaii with PAM sites (solid circles with station windward side is primarily a thermally driven circula- number), the tethersonde site at Kaumana Elementary School (open tion. They compared the westerly ¯ow to a classic grav- circle), Doppler radar locations, and pro®ler location. Contour inter- val for elevation is 1000 m. ity current with a 1% density discontinuity across the ¯ow convergence line. Reisner and Smolarkiewicz (1994) studied low (Ͻ1) Fr ¯ow past a three-dimen- offshore of Hilo is caused by the interaction between sional obstacle with uniform heating at the surface and the land breeze and incoming trade-wind ¯ow. Garrett obtained a simple criterion for the transition from the (1980) conducted an observational study of the ¯ow blocked ¯ow to an unblocked ¯ow regime. Chen and pattern over the eastern slopes of Mauna Loa. He com- Wang (1994) show that variations in surface temperature bined previous work from other researchers (Leopold and dew point are related to orography, surface air¯ow, 1949; Eber 1957; Lavoie 1967; Mendonca 1969) with and distributions of cloudiness and rainfall. The am- his observations from a Hilo to Mauna Loa transect to plitudes of the diurnal surface air temperatures are not form a conceptual model of the island-induced circu- uniform over the island. Dry, barren lava soils on the lations. The daytime regime is triggered by differential upper slopes and on the lee sides of mountain ridges heating rates over land and sea, which result in a com- heat up more quickly than moist, well-vegetated sur- bined anabatic and sea breeze wind. At night, the sit- faces on the windward lowlands. After sunrise (before uation reverses, with a thermally forced westerly kat- sunset) the onset of upslope (downslope) ¯ow occurs abatic wind at peak strength just before sunrise. The on the windward slopes where the virtual temperature offshore ¯ow meets the incoming trade winds well off- ®rst becomes warmer (colder) than the upstream envi- shore, resulting in the development of cloud bands off- ronment at the same altitude. In addition, shore of Hilo. and cloud distributions also affect the air¯ow over the In contrast to early observational efforts, the mod- island. Chen and Wang (1995) found that rainshowers eling studies (Smolarkiewicz et al. 1988; Rasmussen et and clouds can affect the timing of the wind shifts from al. 1989; Smolarkiewicz and Rotunno 1990) describe downslope±offshore (upslope±onshore) to upslope±on- the development of westerly ¯ow as a result of a stably shore (downslope±offshore) ¯ow in the early morning strati®ed ¯ow moving over a mountain barrier for (late afternoon) by modifying the thermodynamic ®elds Froude number (Fr ϭ U/Nh, where U is the upstream near the surface. wind speed, N is the Brunt±VaÈisaÈlaÈ frequency, and h is From the analysis of tethersonde and PAM data dur- the height of the barrier) less than unity. For typical ing HaRP, Wang and Chen (1995) take a detailed look trade-wind values (7.5 m sϪ1) upstream of the island of at near-surface winds and thermal pro®les during the Hawaii, the Froude number is between 0.2 and 0.5. On transition periods between katabatic and anabatic ¯ow. the upwind side of the island, the low-level winds are They suggest that the katabatic ¯ow just before sunrise predicted to reverse direction as a result of island block- is characterized by a 50±150-m nocturnal inversion of ing. A convergence line is created where the reverse about a 1±4 K strength. A nocturnal wind maximum is ¯ow encounters the trade winds, resulting in the for- also observed just below the inversion. Wang and Chen mation of Hilo cloud bands. More recent studies by (1995) also note that during cases the nocturnal Rasmussen and Smolarkiewicz (1993) suggest the im- inversion and nocturnal wind maximum are weaker than portance of nocturnal cooling in modulating the strength for dry cases owing to cloudiness and vertical mixing.

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The depth of the downslope ¯ow, however, is deeper Chen and Nash (1994, p. 51) suggest that the nocturnal than in the dry cases owing to evaporative cooling of rainshowers west of Hilo are caused by the convergence the falling raindrops. During rain events, they found a between the downslope ¯ow and the opposing trade- much earlier downslope ¯ow initiation in the afternoon wind ¯ow and are enhanced by orographic lifting aloft. on the windward lowlands, most likely caused by evap- This view is con®rmed by recent radar studies (Carbone orative cooling of raindrops (Carbone et al. 1995) and et. al. 1998; Feng and Chen 1998; Li and Chen 1999). the out¯ow of cool downdrafts from rainshowers. Feng These studies show that in the early evening, an inland and Chen (1998) present a case study of downslope ¯ow rainband frequently initiates along the leading edge of evolution during the night of 9±10 August 1990. For the downslope ¯ow. In addition, some of these nocturnal this relatively dry case under weak to normal trade-wind rainshowers may originate from nonprecipitating trade- strength, they found that downslope ¯ow was initiated wind cumuli that move onshore and interact with the on the upper slope as the surface air temperature there drainage front (Li and Chen 1999). As the evening con- ®rst became colder than the upstream environment in tinues, the maximum rainfall axis shifts to coastal areas the early evening. The surface air temperature decreased as the downslope ¯ow extends toward the coast (Chen gradually in the afternoon in response to the diurnal and Nash 1994). These results are in agreement with heating cycle. At the coast, the offshore ¯ow occurred Takahashi (1977) and Schroeder et al. (1977). These as the drainage front arrived from the lowlands. Mod- studies show that windward coastal stations exhibit noc- eling studies (e.g., Rasmussen and Smolarkiewicz 1993) turnal rainfall maximum around midnight. Furthermore, suggest that the location of the offshore convergence HaRP radar data also show that some of the early morn- zone is mainly determined by the upstream Fr number. ing rainfall along the coast originates far upstream as Nevertheless, as the cold pool at the coast deepened preexisting trade-wind rainshowers (Austin et al. 1996; through the night, offshore ¯ow extended gradually to Wang and Chen 1998) rather than generated within the 15±17 km offshore before sunrise. offshore . These results are in agree- Chen and Nash (1994) compared the surface pressure ment with an early study using satellite data (Larson and air¯ow between the 12 strongest and the 12 weakest 1978). Enhancement of the rainband occurs as it en- trade-wind days during HaRP based on the composite counters the island-induced circulations (Wang and analyses of the PAM data. Their results show that island Chen 1998; Carbone et al. 1998). blocking is more signi®cant when the trade winds up- The primary objective of this study is to document stream are stronger. In addition, the onset times of down- the interaction between the nighttime downslope±off- slope±offshore (upslope±onshore) ¯ow in the late af- shore ¯ow and trade-wind rainshowers under strong ternoon (early morning) in the areas where the diurnal trade-wind conditions based on a case study. The period winds blow in the direction opposing the incoming selected for this study is from 1500 HST 7 August 1990 trade-wind ¯ow also depend on the strength of the up- to 0900 HST 8 August 1990 within the intensive ob- stream trade winds. Under stronger opposing trade-wind servation period (IOP) of HaRP. Some of the strongest ¯ow, the sea breeze starts later in the morning and ends upstream trade winds of the HaRP experiment are ob- earlier in the late afternoon along the Waikoloa coast served on 7±8 August. We will investigate the evening downstream of the Waimea Saddle, whereas along the (morning) transition between the upslope±onshore windward coast and the Hilo area, the offshore ¯ow (downslope±offshore) and downslope±offshore (up- starts later and ends earlier. Cape Kumukahi, located at slope±onshore) ¯ows under atypical strong trade-wind the eastern tip of the island, shows a mean onshore wind conditions. Our results will be compared with the mod- component nearly all night for the 12 strong trade wind eling studies for elevated Fr days (Smolarkiewicz et al. cases during HaRP except for a short period (Ͻ1h) 1988; Grubisic et al. 1996). In addition, the early morn- before sunrise (Chen and Nash 1994). ing of 8 August has the wettest 0300±0700 HST period Prior to HaRP, investigators characterized windward during HaRP (Fig. 2). It is clear from previous studies precipitation by two regimes: afternoon orographically (Chen and Wang 1994, 1995; Carbone et al. 1995) that enhanced rainfall and nighttime showers from the rain- precipitation modi®es the evolution of downslope ¯ow. bands moving onshore (Leopold 1949; Garrett 1980). We will investigate how the relatively large rainfall Nevertheless, based on the analysis of the PAM data amounts observed on the evening and early morning of during HaRP, Chen and Feng (1995, their section 5) 7±8 August affect the evolution of downslope ¯ow. As separated the diurnal rainfall regimes on the windward will be shown later, almost all of the nocturnal rainfall side into three rainfall regimes: the daytime orographic on 8 August is from offshore trade-wind rainshowers rainfall regime, the nocturnal rainfall regime on the drifting onshore. Although some of these rainshowers windward lowlands, and the early morning coastal rain- have bandlike features, it appears that these are not rain- fall regime. For the period of 1900±2300 Hawaiian bands formed in the offshore convergence line discussed Standard Time (HST), the rainfall on the windward low- in previous studies (e.g., Garrett 1980; Smolarkiewicz lands west of Hilo is heavier than any other 4-h period et al. 1988). The enhancement of these trade-wind rain- during HaRP. The early morning rainfall along the coast showers by the downslope ¯ow under strong trade-wind is much less (Chen and Feng 1995; Carbone et al. 1998). conditions will be examined.

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(1995). Tethersonde data during the night of 7±8 August were presented by Wang and Chen (1995; their Fig. 17). The NCAR Electra provided aircraft ¯ight-level data during HaRP. Two ¯ights were conducted for the period of interest. The ®rst began at 2109 HST on 7 August 1990 and ended at 0144 HST on 8 August 1990, and the second went from 0400 to 0735 HST on 8 August 1990. These ¯ights included an upstream sounding tak- en approximately 140 km from Hilo during the ascent of the aircraft. Since some analyses in this study com- pare the virtual temperatures from the upstream sound- FIG. 2. Total rainfall (mm) from 0300±0700 HST for windward ing with values from PAM stations at the same altitude, PAM stations (4, 5, 6, 7, 8, 9, 10, 11, 12, 13, 16, 22, 23, and 44) a comparison was made between the two instrument during HaRP. platforms while the aircraft was on the ground shortly before takeoff and after landing to check for any biases. 2. Data sources The comparison between the PAM station 9 (located at the airport) data and the aircraft data yielded tempera- A network of 50 PAM stations was deployed on the ture and dewpoint values between these two observing island (Chen and Nash 1994) to record surface param- platforms within a range of 0.5 K. No systematic bias eters at 1-min intervals. The standard PAM station mea- was detected. In addition, several low-level (150±200 sured pressure, temperature, wet-bulb temperature, u m) legs were used to evaluate the evolution of low-level and ␷ wind components, and rainfall. From the standard winds offshore. Aircraft ¯ight altitudes were measured measurements, other derived quantities were calculated. by the Sperry radio altimeter because of contamination Dry static energy is de®ned as of the Stewart±Warner radar altimeter at low altitudes (Hawaiian Rainband Project 1991). There are potential S ϭ C T ϩ gz, (2.1) p inertial navigation system errors in turns and particularly where Cp is the speci®c heat of air at constant pressure, in climbing turns that affect the wind measurements. T is temperature, g is the gravitational acceleration, and During HaRP, two C-band Doppler radars were op- z is the altitude. Moist static energy is de®ned as erated at Paradise Park (CP-3) and Hilo Airport (CP-4), respectively. Most of the data used in this study are h ϭ C T ϩ gz ϩ Lq, (2.2) p surveillance scans of re¯ectivity from CP-4 at beam where L is the latent heat of vaporization and q is the elevations of 0.5Њ for long-range observations and 4.5Њ speci®c humidity. A careful manual inspection of the as echoes neared the coastline. The beam elevation of data was used to eliminate obvious errors. The surface 4.5Њ was chosen to reduce ground signals allowing un- wind ®elds are based on 15-min averages after error contaminated viewing of radar echoes over the wind- checking and elimination of suspect data (Chen and ward slopes. Range-height indicator (RHI) scans at an Nash 1994). Some PAM stations were particularly prob- azimuth directed eastward from both CP-3 and CP-4 lematic during HaRP. The tipping bucket rain gauge at were used to investigate near-surface winds using re- station 15 malfunctioned, rendering all rainfall data for ¯ectivity and radial velocity ®elds. this site unusable. Station 5 suffered from poor wind A NOAA 915-MHZ ultrahigh frequency boundary exposure and no wind analyses incorporate this station. layer wind pro®ler was at Paradise Park (station 8). Stations 15, 16, and 17 include additional radiation mea- Resolution and accuracy of the instrument are contained surements including global solar incoming, downward in Grindinger (1992). The pro®ler usually detected infrared, solar diffuse, solar re¯ected, and net ``all wave- clear-air echoes up to 4 km with a 104-m height reso- length'' radiation. The global solar incoming radiation lution. Sampling intervals were every 12 min for re- was missing at station 16 as a result of an instrument ¯ectivity and every 30 min for wind pro®les. Other data problem; however, the net radiation values remained sources include synoptic charts and satellite imagery unaffected. The accuracy and resolution of the PAM from Geostationary Operational Environmental Satellite instruments are contained in the HaRP Data Catalog (GOES) and NOAA satellites. (Hawaiian Rainband Project 1991). Tethersonde operations were conducted at Kaumana 3. Large-scale conditions Elementary School (Fig. 1) during 7±8 August 1990. The tethersonde instrument was manufactured by At- Figure 3 shows the surface analysis at 1400 HST on mospheric Instrument Research (AIR) and measured 7 August 1990. The surface pattern is dominated by two pressure, temperature, wet-bulb temperature, relative high pressure centers north of the island chain separated humidity, and wind speed and direction at about a by a trough axis between them. Weak troughing is ev- 5±10-m resolution. Accuracy and resolution of the in- ident west of the island along 160ЊW. East of the islands strument package are contained in Wang and Chen the surface pressure gradient appears to be tightening,

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FIG. 3. Surface pressure analysis (P-1000 hPa) at 1400 HST on 7 Aug 1990. probably in¯uenced by the tropical storm found far FIG.4.GOES WEST visible satellite image at 1331 HST on 7 Aug south of the islands. Besides the tropical storm and its 1990. effects on the pressure ®eld, the surface pattern is typical for August (Chen and Feng 1995). The visible GOES satellite image at approximately the same time (Fig. 4) Although there are differences in the moisture and in- shows a patch of closed cell stratocumulus far east of version structure, the wind pro®les are very similar, with the island. Closer to the islands, upper-level cloudiness strong trade winds of 11 m sϪ1 (Fig. 7a). can be seen coming from the south, likely associated In summary, the large-scale conditions for 7±8 Au- with the tropical storm. Upstream of the island chain a gust 1990 are dominated by strong trade winds with a mesoscale disturbed area of trade-wind cloudiness can mesoscale disturbed area within the prevailing trade- be seen about 300 km east of Hawaii. The synoptic- wind ¯ow. The Froude number is elevated ϳ0.42 (Gru- scale analyses provide no explanation for its existence. bisic et al. 1996). Besides the apparent tightening of the A sequence of GOES satellite images (not shown) in- surface pressure gradient, the effect of the tropical storm dicates that this area is moving toward the island with to the south is an increase in upper-level clouds that the trade-wind ¯ow. A NOAA image at 0208 HST (Fig. reach and cover the island by about midnight (not 5) shows a better view of the area as it nears the coast. shown). It will be shown in later sections that this area appears to be the source of the observed precipitation during the 7±8 August period. An upstream sounding taken by the NCAR Electra about 140 km east of Hilo at 2100 HST on 7 August (Fig. 6a) shows that the trade winds are strong with peak values of 11 m sϪ1 that represent some of the strongest trade winds observed in the HaRP project. Average HaRP values are closer to 7 m sϪ1 (Chen and Nash 1994). The height of the trade-wind inversion is about 2.8 km in the upstream sounding. From the sat- ellite image at 2130 HST (Fig. 6b), it appears that this sounding was taken within this disturbed area. At 0400 HST on 8 August another upstream sounding (Fig. 7a) shows a lower inversion height of about 1.7 km and a generally drier moist layer than at 2100 HST (Fig. 6a). This sounding was taken in the clear area behind the disturbance as seen in the GOES satellite image at 0430 HST (Fig. 7b). It appears that the disturbed area and its attendant have raised the inversion locally, as found by Grindinger (1992) in her pro®ler studies. In addition, the clear area behind the disturbance may be a result of subsidence, lowering the inversion height. FIG.5.NOAA-11 IR satellite image at 0208 HST on 8 Aug 1990.

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FIG. 6. (a) Temperature and dewpoint pro®les taken approximately FIG. 7. (a) Same as Fig. 6a but for 0400 HST on 8 Aug; and (b) 140 km upstream by the aircraft during ascent starting at 2100 HST same as Fig. 6b but for 0430 HST. on 7 Aug 1990; and (b) GOES WEST IR satellite image at 2130 HST, open dot represents approximate location of aircraft sounding. Wind (m sϪ1) with one pennant, full barb, and half barb representing 5, 1, and 0.5 m sϪ1, respectively. Hereafter, the same convention will be the island as simulated by Smolarkiewicz et al. (1988). used for all ®gures. At 1500 HST, the virtual temperature at the PAM lo- cations is warmer than the upstream sounding at the same altitude (Fig. 8). A minimum of virtual temper- 4. Initiation of downslope ¯ow ature difference exists around stations 16, 10, and 15. a. Afternoon conditions These stations lie at a location dominated by orographic clouds (Fig. 9). It is apparent that orographic clouds The analysis of downslope ¯ow evolution on 7±8 serve to limit incoming solar radiation and result in August 1990 will ®rst begin with a brief discussion of lower surface temperatures than cloud-free locations, as the preceding conditions on the afternoon of 7 August. found by Feng and Chen (1998). Figure 8 shows conditions at 1500 HST with upslope winds occurring over all of windward Hawaii. Maxi- b. Onset of downslope ¯ow on upper slopes mum wind speeds at this time are 4 m sϪ1 observed at several stations (stations 6, 7, and 11), less than that Figure 10 shows that downslope ¯ow ®rst occurs on measured upstream (ϳ11 m sϪ1). In addition to ¯ow the upper slopes (station 16) and then gradually pro- deceleration, winds along the coast are de¯ected around gresses down to the coast. To investigate the downslope

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FIG. 8. Averaged surface winds at 1500 HST and PAM surface FIG. 10. Isochrone analysis of onset of katabatic ¯ow on 7 Aug. virtual temperature (K) at 1500 HST minus the virtual temperature Time is HST. Circle indicates station with no windshift. (K) of the aircraft sounding at the same altitude approximately 150 km upstream.

¯ow onset on the upper slopes, an analysis of high- resolution (1 min) radiation and thermodynamic data from station 16 (elevation 1128 m) was conducted. During the afternoon hours net radiation values are highly variable (Fig. 11a), this is likely due to the ¯uc- tuation of coverage by orographic clouds. Variation in cloudiness affects both insolation and downward long- wave radiation, thus producing large-amplitude changes in net radiation over very short times. Figure 11a shows some phase agreement between the IR down and net radiation values showing the in¯uence of the orographic cloudiness on the surface radiation balance. It is im-

FIG. 11. Time series (1-min interval) of PAM data at station 16 on 7 Aug from 1500 to 0000 HST. (a) Net radiation (W mϪ2) (solid line) and downward infrared radiation (W mϪ2) (dotted line); (b) temper- ature (ЊC) (solid line) and dewpoint (ЊC) (dotted line); (c) moist static energy (kJ kgϪ1) and rainfall (mm); (d) dry static energy (kJ kgϪ1) (solid line) and Lq (kJ kgϪ1) (dotted line); and (e) zonal component of surface wind (m sϪ1) (solid line) and difference in virtual tem- FIG.9.NOAA-11 visible satellite image at 1335 HST on 7 Aug perature (K) (dotted line) between PAM station and upstream aircraft 1990. sounding at the same elevation.

Unauthenticated | Downloaded 09/23/21 08:22 PM UTC MAY 2001 FRYE AND CHEN 963 portant to note that the surface air temperature (Fig. 11b) shows good correlation with net radiation. Clearly, surface temperature is responsive to net radiation ¯ux at the surface. Net radiation begins to decrease sharply before 1700 HST and becomes negative at 1730 HST. A similar drop also occurs in the surface air temperature at this time. At the same time, IR down values drop suggesting a partial clearing of the orographic clouds. Shortly afterward, a sharp increase in IR down is evi- dent. It appears that the orographic clouds are becoming reestablished over the station. After 1730 HST, the sur- face air temperature begins to climb again in response to an increase in net radiation until approximately 1800 HST. The downslope ¯ow onset occurs at about 1915 HST (Fig. 11e). This is within1hofvirtual temperature gradient reversal between the station and the upstream sounding at the same altitude (Fig. 11b) as found by Chen and Wang (1994). There is no evidence that rain evaporation (Carbone et al. 1995) plays any part in the observed cooling before the onset of the downslope ¯ow. If the cooling were primarily related to evaporative cooling, a drop in dry static energy would have accom- panied an increase in speci®c humidity, producing little change in moist static energy (Carbone et al. 1995). In contrast, at the time of downslope ¯ow onset, a steady decrease in dry static energy and speci®c humidity (Fig. FIG. 12. Same as Fig. 11 but at station 15 (rainfall data from teth- 11d) combine to produce a fall in moist static energy ersonde log at Kaumana Elementary School). (Fig. 11c). Furthermore, over the entire windward side of the island between 1400 and 1900 HST, only 0.25 mm of rainfall occurs at station 10 at an elevation 500 and Carbone et al. (1998), atmospheric density current m lower than station 16 (not shown). There are no radar propagation speed, V, may be estimated by echoes in the vicinity of station 16 in the late afternoon 1/2 V ϳ k(gh⌬␪␷ /␪␷ ) ϩ buO, (4.1) (not shown). Since the surface air temperature is re- sponsive to net radiation ¯ux at the surface, the diurnal where k is a constant (ϳ1), g is gravitational acceler- heating cycle is the most likely cause for the observed ation, h is the depth of the cold pool, u 0 is the component decrease in the surface air temperature and subsequent of environmental ¯ow in the direction of gravity current onset of downslope ¯ow as found by Feng and Chen propagation, and b is 0.7 (Simpson and Britter 1980). (1998) for the 10 August case. Quantitative calculations For a katabatic ¯ow with a nocturnal inversion at 50 m of the surface energy balance are not possible because and a virtual potential temperature de®cit (⌬␪␷ /␪␷ )of the surface sensible and latent heat ¯uxes were not mea- 0.8±1.0% (Wang and Chen 1995, their Fig. 17), the sured at the PAM sites during HaRP. estimated gravity current speed forced by buoyancy def- icit [®rst term in Eq. (4.1)] is only about 2.0±2.5 m sϪ1. Thus, it is unlikely that the downward progression of c. Initiation of downslope ¯ow on lower slope the katabatic ¯ow from station 16 to station 15 is caused by the propagation of density current down the slope. The downslope ¯ow onset at station 15 (elevation 405 Similar to station 16, the net radiation data (Fig. 12a) m) occurs at 1945 HST (Fig. 12e), which is about 30 at station 15 exhibit highly variable values during the min later than at station 16. Throughout the late afternoon early afternoon. The surface air temperature (Fig. 12b) and early evening station 16 is potentially warmer than decreases in the late afternoon with a similar decrease station 15 with a higher (ϳ2kJkgϪ1) dry static energy. in net radiation (Fig. 12a), although net radiation re- The downslope ¯ow moving down the terrain from sta- mains positive. This is due to heat loss in the surface tion 16 (elevation 1128 m) must have left the slope sur- layer by upward sensible and latent heat ¯uxes (Stull face before reaching station 15 (Chen and Nash 1994). 1988). Shortly before 1700 HST an increase in net ra- Furthermore, if the gravity current had moved down the diation, IR down, temperature, and dewpoint is ob- slope from station 16 to station 15, it would have had a served. The increase appears to be related to an increase propagation speed of 7 m sϪ1, which is unrealistically in orographic cloud coverage. Net radiation begins to high for a katabatic ¯ow. Following Benjamin (1968) drop sharply after 1715 HST with temperature following

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similarly. Net radiation ®rst becomes negative at 1745 potential temperature (␪␷ ) occurs behind the leading HST, but only brie¯y as continued variability in IR down edge of the downslope ¯ow (Fig. 13a), which marks the results in another increase in net radiation. Finally, at location of the drainage front or the separation of the 1845 HST net radiation becomes negative for the rest cool, dry katabatic ¯ow from the warm, moist trade- of the evening. At the same time temperature and dew- wind ¯ow. The leading edge of the drainage resembles point (Fig. 12b) continue to decrease. At the nearby a density current (Carbone et al. 1998; Feng and Chen tethersonde site, the superadiabatic surface layer dis- 1998). appears around 1800 HST (Wang and Chen 1995, their Interesting changes in the surface ␪␷ ®eld occur be- Fig. 17a). With continued decrease in the surface air tween 2100 and 2200 HST over the coast and windward temperature, the nocturnal inversion starts to form lowlands. Figure 13b shows that the well-de®ned cold around 1900 HST. Similar to the upper slopes, the down- pool over the lowland area west of Hilo has become slope ¯ow onset at station 15 occurs at 1945 HST (Fig. disorganized. In addition, the drainage front has moved 12e) within 1 h after the virtual temperature at the PAM westward back up the slope. Stations 13, 44, and 4 (Fig. site becomes colder than the upstream environment 1) have undergone a wind shift from katabatic ¯ow back around 1900 HST. Downward trends are observed in to an easterly wind direction (Fig. 13b). There is sig- moist static energy (Fig. 12c), dry static energy (Fig. ni®cant cooling along the coast and warming observed 12d), and speci®c humidity (Fig. 12d), suggesting that at inland stations, including those that have undergone the decrease in the surface air temperature is not caused a shift back to easterly winds (Fig. 14a). Signi®cant by rain evaporation. rainfall accumulations are observed over the coast and After the onset of downslope ¯ow, continued cooling lowland areas between 2100 and 2200 HST (Fig. 14b). is observed at station 15 (Fig. 12b) and 16 (Fig. 11b). The cooling/warming pattern and the retreat in the drain- At station 16, temperature and dewpoint (Fig. 11b) age front between 2100 and 2200 HST appear to be steadily decrease until about 2015 HST when a rapid related to rainfall. The satellite imagery at 2130 HST decrease in both values occurs. At the same time a sig- (Fig. 6b) reveals the existence of a disturbed area of ni®cant decrease is observed in IR down, showing a trade-wind cumuli upstream of Hawaii. From radar data, clearing of the orographic clouds formerly over the sta- it is apparent that most of this rainfall is from trade- tion. Afterwards, dewpoint drops dramatically as the wind showers that drift over the island. Low beam el- westerly component reaches 4 m sϪ1 (Fig. 11c). evation (0.5Њ) surveillance scans (Fig. 15) from Hilo are Chen and Nash (1994) show that island blocking is used to examine the rainshowers at far distances off- more signi®cant for the strong trade-wind cases than for shore. Because of the high pulse repetition frequency the weak trade-wind cases with higher positive pressure used to avoid velocity folding, the maximum unambig- anomalies over the windward lowlands (their Fig. 14). uous range is limited to 75 km. At 1800 HST (not The U-component difference between the strong and shown) only a couple of circular echoes exist far north- weak trade-wind cases, Ustrong Ϫ Uweak, shows a small east of CP-4. By 1900 HST (Fig. 15a) more isolated 1 (Ͻ1msϪ ) positive difference on the windward slopes echoes are observed far northeast. The echoes at 1900 for both the daytime and nighttime ¯ow regimes (their HST are probably not new but have moved into the 75- Fig. 15). On the windward slopes, the upslope ¯ow km radar range. The echoes continue to move with the Ϫ1 during the day is slightly weaker (Ͻ1ms ) and the trade winds until they are just offshore of the windward Ϫ1 downslope ¯ow at night is slightly stronger (Ͻ1ms ) coast by 2100 HST (Fig. 15b). There is no evidence of when the trade winds upstream are stronger. For strong any bandlike structure before 2100 HST. trade-wind cases, the mean winds on the windward With the rainshowers found just offshore, it is now slopes remain weak because of signi®cant island block- necessary to use a higher radar beam elevation (4.5Њ). ing. Thus, under strong trade-wind conditions, the onset At 4.5Њ, the radar beam climbs over almost all of the time of the downslope ¯ow on the windward slopes is island topography except for . At 2115 HST still governed by the thermal contrast between the slope (Fig. 16a), some of the showers have already moved surface and the environment at the same altitude, similar over the northern portions of the coast. Some of the to the HaRP mean (Chen and Wang 1994) and a rela- stronger echoes (30 dBZ) are offshore of Paradise Park tively weak trade-wind case (Feng and Chen 1998). (station 8), in addition to a cell very close to CP-4 at Chen and Nash (1994) found that at station 16, the onset Hilo Airport. The wind data reveal that the drainage and cessation times of the downslope ¯ow for the 12 front has not yet reached the coast at this time. Fifteen strong trade-wind cases occur nearly at the same time minutes later (2130 HST), some of the echoes that were as the 12 weak trade-wind cases based on their com- previously offshore are now moving over the coast (Fig. posite analysis of the PAM data (their Fig. 17d). 16b). An area of 30±40-dBZ re¯ectivity is directly over station 8. At 2145 HST, a solid line of 30-dBZ re¯ec- d. Interactions between trade-wind rainshowers and tivity is observed from the northwest of station 13 ex- the downslope ¯ow tending southeast past station 8 (Fig. 16c). This line is At 2100 HST, similar to the 10 August case (Feng along the leading edge of the drainage front. It appears and Chen 1998), an elongated area of minimum virtual that as the rainshowers encounter the katabatic ¯ow,

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FIG. 14. (a) The change in PAM virtual potential temperature (K) from 2100 to 2200 HST and (b) accumulated rainfall (mm) between 2100 and 2200 HST for 7 Aug 1990.

they are enhanced and focused into a line. The line starts to weaken after 2200 HST as it continues to move inland (Fig. 16d). The solid 30-dBZ line has now broken up into three to four major segments. The observed cooling over the entire coastal area be- tween 2100 and 2200 HST (Fig. 14) is apparently related to the trade-wind rainshowers that move onshore. The time series data for station 8 show a large temperature FIG. 13. Averaged surface winds (m sϪ1) and surface virtual po- drop (Fig. 17a) occurring at about 2130 HST when the tential temperature (K) every 1 K at (a) 2100, (b) 2200, and (c) 2300 trade-wind rainshowers with re¯ectivity Ͼ30 dBZ (Fig. HST. The dashed line in (b) denotes the contour for 298.5 K. 15a) arrive. From 2130 to 2145 HST steady increase in the dewpoint (Fig. 17a) is observed as rainshowers move farther inland. A similar signal is observed in the dry static energy and speci®c humidity values (Fig.

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junction with the destruction of katabatic ¯ow in the windward lowlands (Fig. 13b). At station 15, a jump is noticed in temperature and dewpoint (Fig. 12b) during 2215±2230 HST. A similar ramp-up in dry static energy and speci®c humidity (Fig. 12d) is also observed. At approximately the same time as the temperature and moisture increase, a rapid return of the easterly wind component (Fig. 12e) is evident. It seems that the ver- tical mixing associated with rainshowers is transporting potentially warmer air from aloft down into the near- surface katabatic ¯ow layer (Wang and Chen 1995). The observed increase in moist static energy at station 15 (Fig. 12c) occurs near the time that moderate to heavy rainshowers are reported by the tethersonde crew. The continued inland march of the showers is shown in Fig. 18. Weakening of the line is evident at 2215 HST (not shown) and 2230 HST (Fig. 18a) as the 30- dBZ re¯ectivity areas continue to dissipate, re¯ecting an overall weakening of the line. It appears that, al- though the intensity is decreasing, the size is still large, showing light precipitation over most of the windward side of the island. By 2300 HST, the entire echo system begins to dissipate leaving only one 30-dBZ re¯ectivity area north of station 15 (Fig. 18b). At station 15, the rapid return to easterly winds at 2215 HST is followed by a resurgence in downslope ¯ow at 2250 HST with decreasing temperature and dew- point (Fig. 12). At this time, radar echoes associated with the rainband move farther upslope and are west of station 15 (Fig. 18b). At 2300 HST, downslope winds of about 4 m sϪ1 are evident (Fig. 12) and represent the strongest surface downslope ¯ow observed at station 15 throughout the night. At the nearby tethersonde site, the depth of the downslope just before 2200 HST is about 5 hPa with a weak nocturnal inversion (Wang and Chen 1995, their Fig. 15). Its depth increases to more than FIG. 15. CP-4 re¯ectivity at 0.5Њ elevation angle for 7 Aug 1990 15 hPa at 2300 HST after the arrival of the rainband. at (a) 1900 and (b) 2100 HST. Radar re¯ectivity contours every 10 dBZ starting at 10 dBZ. Contour line thicker for re¯ectivity greater In the meantime, the temperature for the lowest 20 hPa than 30 dBZ. or so has a 2.5 K temperature drop, possibly caused by the rain evaporation aloft (Wang and Chen 1995). The strongest (ϳ5msϪ1) downslope winds aloft observed 17c). A signi®cant temperature drop at the time of rain- throughout the night also occur around 2300 HST and shower occurrence followed by a slight temperature in- weakens rapidly after that (Wang and Chen 1995). crease (Fig. 17) is apparently associated with a cold pool from the rainshowers. At station 8, a sharp easterly wind e. Development of the offshore ¯ow along the coast component of near 6 m sϪ1 is observed (Fig. 17d) at the time of the temperature drop around 2130 HST (Fig. Compared to the HaRP mean (Chen and Nash 1994), 17a) as the trade-wind rainshowers arrive (Fig. 16b). the 7±8 August case features much later onset of down- In the windward lowlands, there is an apparent link slope±offshore ¯ow over the windward lowlands and between the destruction of the katabatic ¯ow layer and coastal areas because of the atypically strong trade strong radar re¯ectivity. Strong radar echoes (Ͼ30 dBZ) winds upstream. Coastal locations (stations 13, 9, and tend to cause a switch from katabatic winds to easterly 8) undergo transition from upslope±onshore to down- winds. At 2145 HST, just after the 30-dBZ line passes slope±offshore ¯ow 1±2 h later than the HaRP mean. over station 13, the winds shift from katabatic ¯ow to Cape Kumukahi (station 7) never experiences an off- an easterly direction (Fig. 16c). A similar situation oc- shore ¯ow, remaining well exposed to trade-wind ¯ow curs at 2200 HST for stations 44 and 4 (Fig. 16d) and the entire night. at 2230 HST for station 15 (Fig. 18a). In the windward For the 10 August case, under normal to weak trade lowlands, warming occurs after rainshowers in con- winds compared to the HaRP mean, the onset of the

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FIG. 16. PAM winds and CP-4 re¯ectivity at 4.5Њ elevation angle for 7 Aug 1990 at (a) 2115, (b) 2130, (c) 2145, and (d) 2200 HST. Radar re¯ectivity contours every 10 dBZ starting at 10 dBZ. Contour line thicker for re¯ectivity greater than 30 dBZ. offshore ¯ow along the coast is caused by the arrival ¯ow would move slower toward the coast if the op- of the drainage front from the windward lowlands (Feng posing environmental ¯ow is stronger, delaying the ar- and Chen 1998). The virtual potential temperature def- rival of the drainage front along the windward coast. At icit (⌬␪␷ /␪␷ ) provides the forcing for the drainage front Cape Kumukahi (station 7) on the eastern tip of the to move toward the coast (Carbone et al. 1998). For this island, the onshore ¯ow has a wind component about 5 case, the virtual potential temperature de®cit is about msϪ1, (Fig. 13), which is comparable or slightly greater 0.8%±1.0% (Fig. 17) and the depth of the cold pool than the estimated gravity current speed. The drainage behind the offshore is ϳ200 m (Fig. 19), the estimated ¯ow never reaches the eastern tip of the island the entire gravity density speed is about 4.0±4.5 m sϪ1. With this night. gravity current speed, the drainage front would move While the drainage front moves slowly toward the slowly toward the coast as it encounters the incoming coast against relatively strong incoming trade winds, a trade-wind ¯ow with an onshore component on the order group of rainshowers moves onshore around 2130 HST of 3±5 m sϪ1 over the coastal waters (Fig. 20). For this (Fig. 16). These rainshowers are enhanced offshore by strong trade-wind (ϳ11 m sϪ1) case, the incoming trade- ¯ow deceleration as a result of island blocking (Carbone wind decelerates signi®cantly as it moves toward the et al. 1998; Wang et al. 2000). They further focus into island in agreement with modeling studies (e.g., Smo- a well-de®ned rainband along the drainage front on the larkiewicz et al. 1988). However, over the coastal wa- windward lowlands (Fig. 16; Frye and Chen 1997). The ters, the prevailing trade-wind speed is still comparable surface forcing along the drainage front is enhanced by to the estimated density current speed forced by the the orographic lifting aloft (Chen and Nash 1994). The virtual potential temperature de®cit. Note that with a narrow swath of over 1 mm of rainfall accumulation same virtual potential temperature de®cit, the drainage during 2100±2300 HST (Fig. 21) over the lowlands pro-

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FIG. 17. Time series (1-min interval) of PAM data at station 8 on 7 Aug from 1900 to 0000 HST. (a) Temperature (ЊC) (solid line) and dewpoint (ЊC) (dotted line); (b) moist static energy (kJ kgϪ1) and rainfall (mm); (c) dry static energy (kJ kgϪ1) (solid line) and Lq (kJ kgϪ1) (dotted line); (d) zonal component of surface wind (m sϪ1) (solid line) and difference in virtual temperature (K) (dotted line) between PAM station and upstream aircraft sounding at the same elevation. vides strong evidence that signi®cant enhancement and focusing of the rainshowers has occurred along the drainage front inland. The cold pool associated with the trade-wind rainshowers has provided the temperature drop along the coast (Fig. 17a). Shortly after the tem- perature drop, westerly offshore ¯ow was observed at station 8 after 2215 HST with decreasing temperature and dewpoint (Fig. 17). By 2300 HST, offshore ¯ow is observed at the coast with the exception of Cape Ku- FIG. 18. Same as Fig. 16 except for (a) 2230 and (b) 2300 HST. mukahi (station 7) (Fig. 13c), completing the seaward progression of downslope/offshore ¯ow over the wind- ward side. Numerical simulations of the 7±8 August case with From the low-level (150±200 m) Electra ¯ight leg diurnal varying thermal forcing (Grubisic et al. 1996) conducted from about 2228 to 2300 HST, offshore winds suggest that on elevated Fr days, the development of can be seen north of Hilo Bay and just offshore from the precipitation forcing zone remains con®ned within Hilo Airport (Fig. 20). Farther to the south near Paradise 10 km offshore. This result compares favorably with Park, however, trade winds are observed offshore. The observations. The model also predicts cloud bands to trade-wind ¯ow offshore decelerates signi®cantly as it form approximately 3 h after sunset. It remains station- moves toward the island (Fig. 20). As the offshore ¯ow ary before midnight and starts to move slowly westward extends farther offshore, it would encounter stronger away from the convergence zone after that. Our results environmental winds. The relatively small horizontal show that the offshore ¯ow along the windward coast extent of offshore ¯ow is due to the abnormally strong develops about 4 h after sunset. However, it is driven trade winds. It is unlikely for the offshore ¯ow to move by evaporative cooling after a group of preexisting into an environment with an opposing wind speed much trade-wind rainshowers has moved onshore producing greater than the density current speed forced by the signi®cant rainfall on the windward lowlands. Chen and buoyancy de®cit. For this case, the late onset of the Wang (1995) found that for the 10 wet afternoons in offshore ¯ow along the coast and a small horizontal HaRP, the mean onset of downslope±offshore ¯ow ®rst extent of the offshore ¯ow are related to atypical strong occurs in the Hilo Bay area around 1700 HST with trade winds offshore. evaporatively cooled cold air at the surface. In contrast,

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FIG. 19. Vertical cross section from RHI scan along 80Њ azimuth from CP-4 at 0258 HST on 8 Aug: (a) re¯ectivity (dBZ) and (b) radial velocity (m sϪ1).

for relatively dry afternoons, the offshore ¯ow along the coast occurs more than 3 h later than the rain cases.

5. Evolution of offshore ¯ow after midnight In this section we investigate the evolution of down- slope±offshore ¯ow from about midnight on 8 August 1990 until before the transition to upslope±onshore ¯ow in the morning. This period offers a unique opportunity to study the effects of strong trade winds and frequent trade-wind showers on the downslope ¯ow. We intend to show that mixing associated with showers under the in¯uence of stronger than normal trade winds disrupt the westerly ¯ow and shift the drainage front from off the coast to inland regions. We also will show that with a small offshore extent of the offshore ¯ow and a weak FIG. 20. Aircraft measured winds between 150 and 200 m during density current structure, trade-wind rainshowers inter- 2228±2300 HST. PAM data same as Fig. 7 but at 2245 HST. act with offshore±downslope ¯ow in the coastal area

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From an analytical approach, Smith (1990) explains the reasons why the stably strati®ed ¯ow cannot rise over high ground. He shows that ¯ow stagnation at the lower boundary is associated with increasing pressure along the streamline caused by adiabatic lifting and positive density anomalies aloft. Recent numerical simulations (e.g., Smolarkiewicz et al. 1988; Rasmussen et al. 1989) initialize their model with an unusually dry trade-wind sounding above the 950-hPa level. Orographic clouds were not simulated in the control run of these studies. In their sensitivity tests, orographic clouds are simulated when the upstream trade-wind speed is doubled (ϳ20 msϪ1). Under this situation, the simulated dynamically driven return ¯ow on the windward side reaches ϳ6m sϪ1. These upstream and reversed ¯ows were not ob- served during HaRP. The composite upstream sounding based on 20 HaRP upstream soundings 110±160 km east of Hilo taken by

FIG. 21. Rainfall accumulation (mm) between 2100 and 2300 HST the NCAR Electra during ascent shows that the trade- on 7 Aug 1990. wind layer is moist and conditionally unstable (Chen and Feng 2001). Initialized by this upstream composite sounding, Chen and Feng (2001) simulate orographic and even inland as they move onshore. Rainshowers are clouds on the windward slopes and ¯ow deceleration also enhanced by orographic lifting aloft producing ap- upstream, splitting air¯ow on the windward side and preciable rainfall over the windward lowlands. lee vortices off the Kona coast in their pure dynamical run using the Pennsylvania State University±National a. Near-midnight conditions Center for Atmospheric Research mesoscale model ver- sion 5 (MM5). On the windward side, the air associated As seen in Fig. 13c, by 2300 HST, the downslope± with the simulated orographic clouds is positively buoy- offshore ¯ow is observed over all of the windward side ant. It rises until it reaches the trade-wind inversion with except at Cape Kumukahi (station 7). The strongest convectively driven in¯ow at the lowest levels (Feng downslope winds are observed around midnight hours. and Chen 2001). With the extra degree of freedom in PAM data from station 15 (not shown) and tethersonde data at the nearby Kaumana School site (Wang and Chen the vertical after condensation, the dynamically driven 1995, their Fig. 18b) shows this phenomena. This is a return ¯ow on the upwind side is not required in their departure from HaRP mean conditions. The best de- solution. The convective heating associated with oro- veloped katabatic ¯ow usually occurs near sunrise when graphic clouds feeds back to the island-induced air¯ow. the surface temperature is the coldest (Wang and Chen With convective heating aloft, positive surface pressure 1995). The differences in the 8 August case are the perturbations on the windward side are lower resulting effects of the excessive early morning rainfall. For this in weaker island blocking than without convective heat- case, the diurnal minimum in temperature and dewpoint ing. occurs near midnight (Wang and Chen 1995, their Fig. Feng and Chen (2001) simulate several observed fea- 19), not near sunrise as in the HaRP mean. In the hours tures associated with the nocturnal ¯ow under a small following midnight, a steady increase is observed in Fr number ¯ow regime including the formation of a temperature and dewpoint because of the vertical mixing nocturnal inversion (Wang and Chen 1995), shifting of due to rainshowers and abundant cloudiness that limits the overall cloudy areas from the windward slopes to outgoing longwave radiation (Wang and Chen 1995). In the coastal areas in the evening (Chen and Nash 1994; summary, the strongest downslope winds are observed Feng and Chen 1998; Carbone et al. 1998), gradual when thermodynamic values are at a minimum. For the offshore extension and deepening of the offshore ¯ow 7±8 August case, the coldest air temperature in the low- throughout the night (Feng and Chen 1998), and gen- est levels occurs around midnight. It appears that the eration of early morning band clouds within the con- strength of the downslope winds is related to the thermal vergence zone where the offshore ¯ow meets the in- forcing at the surface. coming decelerating trade-wind ¯ow (Leopold 1949; The solution in Smolarkiewicz et al. (1988) represents Garrett 1980). Feng and Chen (2001) also note that the the ¯ow response to stably strati®ed ¯uid over an iso- nocturnal cooling was underestimated in recent NCAR lated obstacle for a low Fr ¯ow regime. The low-level modeling studies (e. g., Smolarkiewicz et al. 1988; Ras- dynamically driven return ¯ow on the upwind side of mussen et al. 1989; Rasmussen and Smolarkiewicz the island obstacle is an integral part of this solution. 1993) in agreement with the observational evidence pre-

Unauthenticated | Downloaded 09/23/21 08:22 PM UTC MAY 2001 FRYE AND CHEN 971 sented in this study and recent HaRP analyses cited in the introduction. b. Interactions between trade-wind rainshowers and the offshore±downslope ¯ow For the next few hours after midnight, downslope± offshore ¯ow prevails over the windward coast with the only onshore trade winds found at Cape Kumukahi (Fig. 22a). An RHI scan from CP-4 at the Hilo Airport at 0258 HST shows the offshore extent of the offshore ¯ow is only about 6 km. As discussed earlier, the rel- atively small horizontal extent of the offshore ¯ow is related to atypical strong trade-wind offshore. The radial velocity data (Fig. 20) show a gravity head structure at the leading edge of the offshore ¯ow as found by Car- bone et al. (1995). Wind speeds with the head are near 1msϪ1 or less, while offshore components of 3±5 m sϪ1 exist 3±4 km offshore. With strong opposing trade winds offshore, the density current is rather weak com- pared to other HaRP cases (e.g., Carbone et al. 1995; Feng and Chen 1998). For a Fr of 0.42, Smolarkiewicz et al. (1988) predict a nondiabatic source of offshore ¯ow on the order of 6±7 m sϪ1, which is not clearly evident based on our analyses of the 7±8 August data. In addition to rainshowers during 2100±2300 HST, a group of trade-wind rainshowers moves over the island from offshore during 0300±0500 HST. Both of these periods correlate well with the disturbed area of trade- wind cumuli that drift toward the island, with the largest, most developed area coming ashore around 0300±0400 HST. The radar scan at midnight (not shown) shows isolated rain showers near the edge of the 75-km radar range far upstream. It is clear from the echo pattern at midnight and 0100 HST (not shown) that some scattered rainshowers also exist over the lowlands in addition to the rainshowers upstream. At 0200 HST (Fig. 23a) a large echo area with 30-dBZ cores moves into the radar range. This area then moves westward close to the coast by 0300 HST (Fig. 23b). Like the ®rst rainband, these radar echoes appear to be dissipating and regenerating as they track toward the island. This makes tracking individual cells dif®cult. As the rainshowers near the coast at 0315 HST (Fig. 24a), they still have a scattered organization with several areas of 30-dBZ re¯ectivity. The showers cross the coast at 0330 HST (Fig. 24b) and begin to organize into a line at 0345 HST (Fig. 24c). By 0400 HST a clear line develops extending in a north to south orientation from the northern coast to near station 4. The solid 30-dBZ line has isolated cores of near 40 dBZ. Enhancement and focusing of the pre- viously scattered rainshowers are again occurring over FIG. 22. Same as Fig. 13 but at (a) 0300, (b) 0400, and (c) 0600 the windward lowlands. HST. Changes in the surface air¯ow begin to occur around 0400 HST related to the arrival of the second group of trade-wind rainshowers during 0300±0500 HST. Like the previous rainband, wind reversals from westerly to easterly ¯ow begin to occur at coastal stations as the

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¯ow over the coast is temporarily destroyed by rain- showers through vertical mixing causing strong easterly winds to lower to the surface (Fig. 26). For this case, because of the retreat of the downslope±offshore ¯ow, the most pronounced focusing of trade-wind showers in the early morning occurs inland (Fig. 24d) and they are enhanced by orographic lifting aloft. These results are in contrast to the 22 August 1990 case, which is under normal trade-wind conditions. In that case, the well- developed downslope±offshore ¯ow in the early morn- ing extends more than 15 km offshore. The enhancement of the early morning trade-wind rainshowers occurs off- shore when the rainshowers encounter the leading edge of the offshore ¯ow. The rainband weakens after it moves over the well-developed cold, dry offshore ¯ow before reaching the coast, producing little rainfall along the coast (Wang and Chen 1998). It is interesting to note that the early morning of 8 August has the wettest 0300± 0700 HST period during HaRP. At 0400 HST the rainband remains largely intact (Fig. 24d). It starts to weaken after 0415 HST (not shown). The next few scans, at 0430 HST (Fig. 27a), 0445 HST (not shown), and 0500 HST (Fig. 27b) show the rain- band moving upslope with the line continuing to weaken and break up. During this period, westerly ¯ow returns in the windward lowlands west of Hilo after the passage of the second rainband. The rainfall that accumulates from the second rainband (Fig. 28) is much greater than that observed in the ®rst rainband (Fig. 19). It accounts for 40% of the total rainfall between 1900 and 0700 HST. The maximum rainfall axis has a north±south ori- entation with a maximum accumulation of over 10 mm at station 23. The orientation of the maximum rainfall axis agrees well with that of the rainband after individual cells become well organized along the drainage front in FIG. 23. Same as Fig. 15 for 8 Aug 1990 at (a) 0200 and (b) 0300 the windward lowlands. HST. The surface air temperature at station 8 starts to de- crease after 0415 HST because of nocturnal cooling stronger echoes arrive. At 0400 HST (Fig. 22b) easterly (Fig. 25). At 0500 HST (Fig. 27b), another echo area winds are observed inland for the ®rst time since the offshore is poised to move over station 8 over the next evening transition over the windward lowlands. The sig- hour. This echo area appears to be the source of the ni®cant retreat of the cold, dry offshore ¯ow can also rainfall near 0530 HST at station 8 that provides ad- be seen in the ␪␷ ®eld in Fig. 22b. At 0400 HST, the ditional cooling due to rain evaporation and allows an- minimum in ␪␷ is now found much farther inland with other resurgence of offshore ¯ow to the coast (Fig. 25d). its axis along a line described by stations 13, 44, and At 0600 HST, the offshore ¯ow is occurring over the 4. This shows that there has been a temporary disruption entire Hilo Bay area with a cold pool along the coast in the cold pool and the downslope±offshore ¯ow. Dur- (Fig. 22c). ing 0300±0400 HST stations that undergo a brief change from westerly to easterly winds experience an increase 6. Onset of upslope±onshore ¯ow in ␪␷ (stations 9, 13, 4, 44, 8) (Figs. 22 and 25). At station 8 along the coast, a rapid increase in surface air The early morning of 8 August has the wettest 0300± temperature and dewpoint are accompanied by a shift 0700 HST period during HaRP (Fig. 2). Chen and Wang from offshore ¯ow to easterly ¯ow at about 0330 HST (1995) found that before sunrise, rain cases have a high- immediately after rainshowers (Fig. 25). It appears that er surface temperature and a weaker surface downslope potentially warm, moist air aloft mixes with the cool, ¯ow than dry cases. Tethersonde observations on the dry westerly ¯ow near the surface (Wang and Chen windward lowlands show that for rain cases, the noc- 1995). From the wind pro®ler data at the Paradise Park turnal inversion and the westerly wind maximum just in the vicinity of station 8, it appears that the offshore beneath the nocturnal inversion are weaker with a deeper

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FIG. 24. Same as Fig. 16 for 8 Aug 1990 at (a) 0315, (b) 0330, (c) 0345, and (d) 0400 HST. stable downslope ¯ow layer. These differences are re- areas where the nocturnal rainfall is the largest (Figs. lated to rain evaporation cooling aloft, reduced infrared 19 and 28). Nevertheless, for most stations on the wind- radiation heat loss at the lowest level, and vertical mix- ward lowlands and coastal areas, the turning of winds ing associated with rain (Wang and Chen 1995). For the occurs 1±2 h earlier as compared to the mean onset time 8 August case, the nocturnal inversion and the westerly of wet cases reported by Chen and Wang (1995). The wind maximum near the surface at the Kaumana school early turning of winds in the morning of 8 August is are absent in the early morning (Wang and Chen 1995). apparently related to atypical strong trade winds up- After sunrise, rain cases have a slower increase in the stream. surface air temperature than dry cases. Possible causes Wang and Chen (1995) studied the early morning for these differences are 1) reduced insolation by clouds, transition on the windward lowlands based on tether- 2) evaporative cooling of raindrops, and 3) slower sur- sonde and surface data during HaRP. For relatively dry face heating at the wet surface (Chen and Wang 1995). cases, the onset of upslope ¯ow frequently occurs after For rain cases, the latest turning from downslope±off- the disappearance of the nocturnal inversion with a well- shore to upslope±onshore ¯ow occurs in the Hilo coastal mixed surface layer. For rain cases with a deep (30±40 areas (Chen and Wang 1995, their Fig. 13a). The late hPa) stable downslope ¯ow layer, it takes longer for the onset of onshore ¯ow on the windward lowlands and downslope layer to become well mixed than for dry Hilo coastal areas for rain cases is related to a deeper cases after sunrise. The turning from downslope to up- downslope ¯ow layer and a slower surface air temper- slope ¯ow for rain cases usually starts at the lowest ature increase after sunrise (Chen and Wang 1995; Wang levels after the slope surface becomes warmer than the and Chen 1995). For the 8 August case, the latest turning upstream environment at the same level and progresses of downslope±offshore to upslope±onshore ¯ow in the upward. morning occurs in the windward lowlands and coastal The 8 August case is an exception. For this wet morn-

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FIG. 26. Wind data (m sϪ1) (30-min average) from NOAA boundary layer pro®ler at Paradise Park on 8 Aug 1990 from 0000 to 0500 HST.

FIG. 25. Same as Fig. 17 but for station 8 on 8 Aug from 0000 to 0900 HST. ing case, the turning from downslope to upslope ¯ow at the surface on the windward lowlands commences about 30 min before the virtual temperature on the slope surface becomes warmer than the environment (Wang and Chen 1995). The turning starts from the top of the downslope ¯ow layer and progresses downward, show- ing that easterly trade-wind momentum is being carried downward. A similar situation also occurs for the 15 August case, which is also a wet morning (Fig. 2) with strong trade winds aloft (Wang and Chen 1995). The downward transport of the easterly momentum occurs after sunrise with a decrease in the thermal instability in the lowest levels. Wang and Chen (1995) cited the retrograde motion of the density current as another possible cause for the early onset of upslope ¯ow observed on the windward lowlands. The gravity current would be in retrograde motion if its propagation speed became much slower than the opposing trade winds (Simpson and Britter 1980). With surface heating after sunrise, the thermal contrast across the gravity current would decrease, lead- ing to a decrease in the propagation speed of the density current. If propagation speed of the density current falls far below the speed of the opposing trade-wind ¯ow, it would retreat westward. Based on the isochrone analysis of the onset of the upslope±onshore ¯ow (Fig. 29), this process may not be the cause for the early onset of upslope ¯ow on the windward lowlands for this case. If gravity current had retreated westward of the teth- ersonde site, the latest ¯ow reversal would have oc- curred on the windward slopes west of the windward lowlands. Figure 29 shows that the latest onset of up- FIG. 27. Same as Fig. 16 for 8 Aug 1990 at (a) 0430 and (b) 0500 slope ¯ow occurs on the windward lowlands. HST.

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FIG. 28. Rainfall accumulation (mm) between 0300 and 0500 HST FIG. 29. Isochrone analysis of onset of upslope ¯ows on 8 Aug. on 8 Aug 1990. Time is HST. Circle indicates station with no windshift.

Along the windward coastal areas, the shift from off- downslope ¯ow initiation starts on the upper slopes un- shore to onshore occurs more than 1 h earlier than the der an area dominated by orographic clouds and pro- mean onset for the wet mornings reported by Chen and gresses down the slope. The surface winds on the wind- Wang (1995). The onset of onshore ¯ow along the coast slopes are rather weak because of island blocking. The occurs 1±2 h earlier than the ¯ow reversal on the wind- slope surface becomes negatively buoyant in the late ward lowlands (Fig. 29). Examination of the surface afternoon, leading to the onset of downslope ¯ow on data from coastal stations reveals that the shift from the windward slopes. The decrease in the surface air offshore to onshore ¯ow there does not occur after the temperature in the afternoon is most likely related to surface virtual temperature becomes warmer than the the diurnal heating cycle. The onset of katabatic ¯ow upstream value. Rather, along the coast a shift from on the slopes is similar to the 10 August case, which is offshore to onshore ¯ow is accompanied by a rapid under weak to normal trade-wind conditions (Feng and increase in the surface air temperature and moisture Chen 1998). (Fig. 25). No precipitation has fallen during these chang- Along the coastal areas that are well exposed to the es. The isochrone analysis of the upslope±onshore ¯ow decelerating trade-wind ¯ow, the development of the onset shows that the leading edge of the westerly ¯ow offshore ¯ow occurs 1±2 h later than HaRP mean con- retreats inland after sunrise (Fig. 29), allowing the ditions because of the atypical trade winds offshore. warm, moist trade winds to return (Fig. 25). In other Cape Kumukahi never experiences an offshore ¯ow, words, it appears that along the windward coast, the remaining well exposed to trade-wind ¯ow the entire onset of the onshore ¯ow is caused by the retrograde evening. For the 10 August case, the leading edge of motion of the leading edge of the offshore ¯ow after the katabatic ¯ow over the windward lowlands moves sunrise. With the decrease of the thermal contrast across eastward against the incoming trade-wind ¯ow and ac- the drainage front after sunrise, the gravity current prop- counts for the onset of offshore ¯ow along the coast agation speed would decrease. If this speed becomes (Feng and Chen 1998). For this case, the drainage ¯ow less than the speed of the strong opposing incoming moves slowly from the windward lowlands toward the trade-wind ¯ow, the leading edge of the drainage front coast as it propagates against atypical strong opposing near the coast would be in retrograde motion, allowing trade-wind ¯ow. The onset of offshore ¯ow along the the trade-wind offshore to return. The retrograde motion coast is driven by rain evaporative cooling. Around 2130 of the density current is more likely to occur if the HST a group of preexisting rainshowers moves onshore. opposing trade winds are stronger. These rainshowers are enhanced as they encounter the island-induced circulation and focus into a well-de®ned rainband along the drainage front on the windward low- 7. Summary lands. The low-level forcing along the drainage front is One of the goals of this study is to investigate the enhanced by orographic lifting aloft, producing a narrow evolution of downslope±offshore ¯ow on 7±8 August swath of over 1 mm of rainfall accumulation during 1990 in the presence of strong opposing trade winds 2100±2300 HST over the windward lowlands. The rain- and frequent trade-wind rainshowers. For this case, the showers produce an evaporative cooled cold pool at the

Unauthenticated | Downloaded 09/23/21 08:22 PM UTC 976 MONTHLY WEATHER REVIEW VOLUME 129 surface. Subsequently, the downslope ¯ow extends off- is supported by the National Science Foundation under shore. As the westerly ¯ow extends farther offshore, it Grants ATM-9301227 and ATM-9629886. A portion of encounters stronger trade winds. It is unlikely for the this research was completed while the ®rst author at- density current to propagate into an environment with tended the University of Hawaii through the Air Force an opposing wind speed much greater than the density Institute of Technology±Civilian Institution Program. current speed driven by the buoyancy de®cit. The off- Acknowledgment is made to the National Center for shore extent of the downslope±offshore ¯ow remains Atmospheric Research, which is sponsored by the Na- within 10 km of the coast or even inland for this atypical tional Science Foundation, for part of the computing strong trade-wind case throughout the night. resource used in this research. During 0300±0500 HST, a second group of trade- wind rainshowers moves over the island. Because of a REFERENCES small horizontal extent of the offshore ¯ow and a rather weak density current structure under atypical strong Austin, G. R., R. M. Rauber, H. T. Ochs, and L. J. Miller, 1996: trade winds, the interactions between rainshowers and Tradewind cloud and Hawaiian rainband. Mon. Wea. Rev., 124, offshore ¯ow occur in the coastal areas and even inland 2126±2151. Benjamin, T. B., 1968: Gravity currents and related phenomena. J. in the early morning. 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