Diurnal Winds in the Himalayan Kali Gandaki Valley. Part II: Modeling

Transcription

1 1062 MONTHLY WEATHER REVIEW Diurnal Winds in the Himalayan Kali Gandaki Valley. Part II: Modeling GÜNTHER ZÄNGL, JOSEPH EGGER, AND VOLKMAR WIRTH Meteorologisches Institut, Universität München, Munich, Germany (Manuscript received 27 December 1999, in final form 20 June 2000) ABSTRACT The Penn State NCAR mesoscale model MM5 is used to simulate and better understand the wind observations in the Kali Gandaki Valley reported in the first part of this paper. The Kali Gandaki River originates in Nepal near Tibet, flows southward through the Mustang Basin, crosses the Himalayas in a gorge, and descends to the lowlands of Nepal. Extremely strong diurnal upvalley flow in the gorge and the basin alternates with rather weak drainage flow in the night. As proposed in Part I, the Mustang Basin and the Tibetan Plateau can be considered as an elevated heat source driving the upvalley flow during the day. However, the extreme strength of the diurnal upvalley winds and the order-of-magnitude asymmetry between day and night cannot be explained with a simple plateau circulation theory. The model is successful in simulating almost all aspects of the observations. The simulations strongly suggest that the observed acceleration of the upvalley winds near the entrance to the Mustang Basin is linked to a supercritical-like flow pattern. Gravity waves induced by the ridges protruding into the valley appear to contribute to this flow structure. Humidity is found to be essential for simulating the strength of the observed day night asymmetry because of its impact on the boundary layer structure above the Himalayan foothills, especially due to the evaporation of rain. In addition, advection of relatively stable air from the foreland into the basin is important for the formation of the gravity waves and also explains part of the asymmetry. The Plateau of Tibet appears to have a small but positive impact on the flow speeds in the valley. 1. Introduction Recent observations of the diurnal wind system of the Kali Gandaki Valley with its rather intense upvalley flow and its weak downvalley winds have been presented in Part I of this paper (Egger et al. 2000, referred to as KG1 in what follows). The Kali Gandaki River originates near the town of Lo Manthang close to the border of Nepal. The river flows southward in the relatively wide Mustang Basin where several tributaries join the main river. The Himalayan range is crossed between Marpha and Ghasa. The valley is narrow, winding, and more than 5000 m deep in this section. The river rushes down in a gorge to the south of Ghasa (see KG1 for further details) to reach the Pokhara area at an altitude of 1000 m above mean sea level. Figure 1 gives the main topographic features of the Kali Gandaki Valley and the location of the major villages for the two innermost nesting domains of the model simulations, which will be outlined in section 3. In fall 1998 pibal observations of the winds were made at eight locations all the way from Lete near the Corresponding author address: Günther Zängl, Meteorologisches Institut der Universität München, Theresienstraße 37; München, Germany. guenther@meteo.physik.uni-muenchen.de river s exit from the Himalayas up to Lo Manthang close to the source (KG1). The main findings can be summarized as follows. During the day, relatively weak upvalley winds are observed near Ghasa. There is upvalley acceleration between Lete and Marpha in the narrow part of the valley. There appears to occur further acceleration when the flow enters the wide basin of Mustang, where typical wind speeds are m s 1 and typical depths of the inflow layer are m. The winds weaken in the northern part of the basin, but the upvalley flow extends at least up to Lo Manthang. The upvalley flow regime forms first near the ground late in the morning but full depth of the inflow layer is attained within about 1 h. Late in the night, weak drainage flow sets in with velocities well below 5 m s 1 and a layer depth of less than 1000 m, so that there is a pronounced diurnal asymmetry. The Kali Gandaki Valley wind differs much from other valley wind systems studied so far. Upvalley winds observed in Alpine valleys rarely attain surface wind speeds above 5 m s 1. In the Austrian Inn Valley, for example, maximum upvalley wind speeds are typically between 3 and 4ms 1 in summer (Dreiseitl et al. 1980). Similar speeds are found in the Dischma Valley, a small valley in Switzerland (Hennemuth 1987). In the free atmosphere, upvalley speeds up to 7ms 1 have been observed in the Inn Valley (Pamperin and Stilke 1985) American Meteorological Society

2 MAY 2001 ZÄNGL ET AL FIG. 1. Orography (m) as prescribed in domains (a) d4 and (b) d5 of the mesoscale simulation model and the locations of the villages mentioned in the text. Distances in km; contour interval 500 m. (a) Solid frame, subdomain of d4 shown in Figs. 6 and 15; dashed dotted frame, location of d5; LM, Lo Manthang. (b) Solid frame, subdomain of d5 shown in Fig. 4; dashed dotted frame, smaller subdomain shown in Fig. 8; Ta, Tatopani; G, Ghasa; L, Lete; Tu, Tukuche; M, Marpha; J, Jomsom; K, Kagbeni; C, Chuksang; dashed lines, cross sections TL (Tatopani Lete) and TK (Tukuche Kagbeni). Moreover, the day night asymmetry observed in the Kali Gandaki Valley appears to be unique. An upvalley wind that is at least 5 times stronger than the downvalley wind despite a sunshine duration of close to 12 h has never been reported for any other valley. In Alpine valleys, upvalley and downvalley winds are commonly of similar strength in spring and autumn. However, a strong day night asymmetry with reversed sign has been observed at the mouth of the Inn Valley into the Bavarian foreland (Pamperin and Stilke 1985). At a station just off the mouth, a nocturnal low-level jet has been found reaching peak speeds of 14 m s 1 at a height of 200 m above ground while diurnal upvalley speeds hardly exceed 2 m s 1. Similar phenomena have also been observed in other valleys ending abruptly at a plain (Whiteman 1990). They are likely due to a shooting hydraulic flow that is established at the point where the valley fans out into the plain. It is the goal of this paper to gain a better understanding of this unique and exceptionally strong valley wind in the Kali Gandaki. To achieve this goal, simulations with the Pennsylvania State University National Center for Atmospheric Research (Penn State NCAR) fifth-generation Mesoscale Model (MM5) are performed. First, a reference run with realistic orography and very high resolution is conducted in order to reproduce the Kali Gandaki wind system as realistically as possible. Then, sensitivity experiments are performed in order to further explore the dynamics of the flow. Most of them are carried through with idealized orography to elucidate the importance of particular orographic features and to reduce computational costs. In addition, the impact of moisture, vegetation, and the interaction between radiation and moisture is investigated. The ability of mesoscale models to simulate airflow over highly complex terrain has been demonstrated by several studies in recent years. Most of them focus on downslope windstorms or strong gap winds. For example, Saito (1993, 1994) investigated downslope windstorms over Shikoku Island, Japan, and Clark et al. (1994) examined a windstorm at the Colorado Front Range in the region of Boulder. The flow over and around the Pyrenées and its diurnal cycle was studied by Georgelin et al. (1996). Schmid and Dörnbrack (1999) performed simulations of a gravity wave breaking over the Alps during the South Foehn and studied the predictability of such events, which may be quite hazarduous to aviation. Several case studies of gap flow in the Washington State Cascade Mountains were performed by Colle and Mass (1998a,b). All these simulations were successful in reproducing the essential features of the observations and provided deeper insight into the dynamics of the storms. There has also been a large number of simulations of valley winds, but most of them were restricted to idealized or rather simple realistic orography. Channeling effects in large valleys were studied, for example, by Whiteman and Doran

3 1064 MONTHLY WEATHER REVIEW (1993) and Eckman (1998). Bossert and Poulos (1995) conducted simulations of nocturnal drainage flow and compared idealized orography of increasing complexity with realistic orography. Daytime upslope winds were studied, for example, by Atkinson and Shahub (1994). Finally, idealized simulations of the full diurnal cycle of valley winds were performed by Li and Atkinson (1999), focusing on the transition regimes between upvalley and downvalley winds. A more thorough review of studies related to valley winds is given in Li and Atkinson (1999). In summary, there is good evidence that state-of-the-art mesoscale models are appropriate for simulating the valley wind in the Kali Gandaki. A feature of the MM5 that is particularly important for the simulations in the present study is the interactive nesting. It is likely that the Kali Gandaki Valley wind is influenced not only by the Mustang Basin but also by the heat low over the Tibetan Plateau (e.g., Reiter and Tang 1984). Thus, the integration domain should cover the whole Plateau of Tibet. On the other hand, the central part of the Kali Gandaki Valley is very narrow so that a grid distance of no more than 1 km is required in order to resolve the orography with sufficient accuracy. These requirements cannot be met without a nesting technique. The basic mechanisms which have been proposed to explain valley winds are summarized in section 2. Moreover, a simple estimate for the strength of the Kali Gandaki Valley wind is derived. The setup of the MM5 simulations is given in section 3. The results of the reference run are described in section 4. A discussion follows in section 5 where several sensitivity experiments are presented. The paper closes with a set of conclusions (section 6). 2. Valley wind mechanisms The most frequently cited valley wind mechanism is based on the so-called volume factor or topographic amplification factor (TAF; Steinacker 1984). Given a certain horizontal area A and a certain height h, it simply states that the volume of air over the plain V p Ah is larger than the corresponding volume of air within the valley (V ), where part of the total volume is filled with mountains. Thus, assuming that both areas receive the same amount of solar energy, the air in the valley is heated more strongly. Quantitatively, the heating ratio is equal to V p /V, that is, the topographic amplification factor. During the night, the situation is essentially reversed although one has to keep in mind that longwave cooling may be reduced in narrow valleys due to low sky view factors. It is important to note that this valley wind mechanism does not require a sloping valley axis. Thus, it is appropriate for explaining the valley wind system of the Inn Valley, for example, which has very little slope in its lower part. According to Steinacker (1984), it is necessary to include all tributary valleys into the calculation of the TAF because considering the main valley alone would underestimate the actual volume factor. If the valley axis is sloping, the strength of the valley wind may be enhanced because the solar energy is released at a greater height in the upper part of the valley. This is essentially the same mechanism that drives slope winds. If the lower part of a valley slopes very steeply, as is the case for the Kali Gandaki Valley, the valley can be considered to exit into the free atmosphere to a first approximation, whence a kind of plateau effect becomes active. While the valley air is heated during the day through absorption of solar radiation and convective mixing, the air at corresponding heights over the plain may exhibit very weak diurnal temperature variations because it is well above the ground. Moreover, the pressure in the free atmosphere can be expected to increase during the day due to the heating of the underlying air masses. As a consequence, an intense valley wind circulation is to be expected regardless of the volume factor or the slope of the valley axis. Since the latter mechanism may be of considerable importance for the Kali Gandaki Valley wind, we want to estimate the valley wind strength resulting therefrom. Let us assume that the height difference between the valley bottom and the plain is 2000 m and that a relatively narrow valley tube connects the Mustang Basin with the free atmosphere over the plain (cf. Fig. 1b). One can expect that the free atmosphere 2000 m above ground undergoes little diurnal temperature changes although the daytime convective boundary layer may be deeper. Radiative cooling of the free atmosphere does not exceed a few Kelvin per day except at the top of a cloud layer. On the basis of the numerical experiments reported later on, we assume that the depth of the mixed layer in the Mustang Basin, H M, is 1000 m and that the temperature difference, T, between the basin and the free atmosphere over the plain is 5 K. Then, assuming that T is constant throughout the mixed layer, the hydrostatic equation yields a corresponding pressure difference of T p 0gH 1.6 hpa, (1) M T 0 where the density 0 and the temperature T 0 have been setto1kgm 3 and 300 K, respectively. For a sinusoidal pressure wave, this yields a day night difference of about 3 hpa, which is in good agreement with the observations reported in KG1. If one assumes negligible wind speeds at the southern end of the valley tube, Bernoulli s equation yields a velocity of 2 p 2gHM T 1 18 m s (2) 0 T0 at the entrance into the Mustang Basin. This is close to the maximum surface wind speeds observed at Jomsom

4 MAY 2001 ZÄNGL ET AL Expt TABLE 1. List of simulations discussed in the text. Description Discussed in section REF Reference run (see sections 3a,b) 4, 5 REF-SU As in REF but for summer solstice conditions 5c REF-WI As in REF but for winter solstice conditions 5c REF-D As in REF but for dry physics; temperatures 5a as close to REF as possible S 1 Idealized orography as shown in Fig. 11, 5a large-scale fields as in REF S 1 -V As in S 1 but with different vegetation (see 5a section 5a) S 1 -M As in S 1 but with reduced moisture 5a S 1 -MV As in S 1 but with vegetation as in S 1 -V 5a and reduced moisture as in S 1 -M S 1 -D As in S 1 but with large-scale fields as in 5a REF-D S 2 -D As in S 1 -D but with constant valley 5a width; broad valley cross section everywhere S 3 Orography as in S 2 -D but with dam at the 5a southern exit; moist physics as in REF S 4 -D As in S 1 -D but with no Tibetan Plateau 5b S 5 -D As in S 1 -D but with lower plateau height (3000 m instead of 4500 m) 5b and Kagbeni although frictional effects have been neglected here. A serious deficiency of this simple consideration is that it does not explain the observed day night asymmetry. Although the weakness of the nocturnal winds observed in the Mustang Basin may be explained with the cold air being piled up to the north of the narrow valley tube, Eq. (2) predicts strong downvalley winds in the steep part of the valley, that is, between Lete and Tatopani. It is one of the central goals of this paper to resolve this issue. 3. Setup of the simulations a. Model description and initialization The nonhydrostatic version of the Penn State NCAR MM5, version 2.7 (Grell et al. 1995), is used to carry out the simulations. The model equations are based on a terrain-following sigma coordinate. An explicit moisture scheme (Dudhia 1989) and a radiation scheme, which takes into account moisture and clouds (Grell et al. 1995), are used in the reference run REF and in some of the sensitivity experiments. All experiments discussed below are listed in Table 1. Convection is resolved explicitly in the innermost nested domains (see below), which cover the region of interest. Therefore, no parameterization scheme for convection is used. The planetary boundary layer is parameterized using the Blackadar scheme (Zhang and Anthes 1982). Klemp and Durran s (1983) upper radiative boundary condition is applied in order to prevent the reflection of upward propagating gravity waves. It is important to realize that the radiation scheme of the model does not take into account the slope of the terrain. Given a certain position of the sun the energy received by a grid box is independent of the slope of the terrain. In the same way, longwave radiation is emitted as if the terrain were flat. Some additional sensitivity experiments are carried out in a dry mode with a very simple radiation scheme. The MM5 simulations have been conducted both with realistic orography and with idealized orography. The latter is used in several sensitivity experiments to explore the relative importance of certain orographic features. The simulations with realistic orography have been performed on five interactively nested domains with horizontal resolutions of 64.8 km (domain 1, d1), 21.6 km (d2), 7.2 km (d3), 2.4 km (d4), and 0.8 km (d5), respectively. In the vertical, 39 unevenly spaced full-sigma levels are used, corresponding to 38 halfsigma levels at which most of the variables are calculated. The lowest half-sigma level ( 0.994) is about 45 m above ground and will be referred to as surface level in the following. The vertical distance between the sigma levels is slightly less than 100 m near the ground and increases up to 800 m near the top (100 hpa). The model orography of domains 1 and 2 (3 5) has been interpolated from terrain data of 5 (30 ) resolution using a Cressman analysis scheme. It is filtered by a twopass smoother desmoother. Domain 1 is centered at 35 N, 85 E and covers an area of roughly 4000 km 4000 km (see Fig. 2). Domain 2 covers most of the Tibetan Plateau, and domain 3 extends from 27.5 to 30.5 N and from 82 to 86 E. The orography of domain 4 and domain 5 is displayed in Fig. 1. As can be seen, domain 4 covers all of the Kali Gandaki Valley and extends into Tibet while the northern border of domain 5 is located to the south of Lo Manthang. Some extreme peaks (e.g., Dhaulaghiri and Annapurna) have been capped in order to prevent numerical problems. Land use information has been interpolated from a 10 data grid provided by NCAR. Unfortunately, the actual resolution of these data is only 1 in the region of interest, which is too coarse for the present purpose. Therefore, we decided to modify the land use data in domains 3 5. Land use is specified as a function of height: tropical forest between 500 and 2500 m, grassland between 2500 and 3000 m, and desert above 3000 m. This corresponds to moisture availability values of 50%, 15%, and 2%, respectively. It may be noted that this specification is a good approximation along the Kali Gandaki Valley (see KG1). In particular, the north south asymmetry in vegetation across the Himalayas is implicitly taken into account by this simple scheme thanks to the north south asymmetry in surface altitude. In the idealized simulations, the Tibetan Plateau is replaced by a rectangular block with side lengths of 1100 km 1100 km and a height of 4500 m. The boundaries of the block are smoothed with a cosine function over a distance of 220 km. The Himalayan

5 1066 MONTHLY WEATHER REVIEW FIG. 2. Potential temperature (contour interval 3 K) and horizontal wind (full barb 5ms 1 ; circles, 1.25 m s 1 )att 15hina subdomain of domain 1 in REF: (a) surface, the straight line gives the position of the cross section TB (Tibet); (b), z 8 km, the position of domain 4 is indicated by the rectangle. The shading indicates terrain heights h: light shading, 1000 h 3000 m; medium, 3000 h 5000 m; dark, h 5000 m. massif is represented by a smooth mountain range with a height of 5300 m, and the Himalayan foothills are also accounted for. The idealized Kali Gandaki Valley connects, as in reality, the Himalayan foreland with the Tibetan Plateau (see Fig. 11 below). Unlike the realistic simulations, the idealized ones are performed on four domains only. Their horizontal resolutions are 40.5, 13.5, 4.5, and 1.5 km, respectively. The land use specification is similar to that mentioned above: tropical forest below 2500 m, grassland between 2500 and 3000 m, and desert above 3000 m. The initialization of the flow in the presence of the extremely steep orography is difficult. The standard technique to initialize the MM5 model is to interpolate largescale analysis fields to the model domains. However, the orography used in the analysis schemes is much smoother than that prescribed here. An interpolation from such fields to the extremely deep Kali Gandaki Valley is bound to fail. The resulting strong surface winds would excite intense and unrealistic gravity waves above the steep mountain slopes. As an alternative, one could start from an atmosphere at rest. However, the north south heating difference leads to the establishment of a thermal wind that may affect the flow in the Kali Gandaki Valley in an unrealistic manner. To circumvent these problems, an idealized large-scale flow is specified. The idea is to construct a flow field that exerts as little influence as possible on the valley wind and that remains quasi-stationary. Quasi-stationarity is assumed when the temperature drift obtained in long-term test runs is below 2 K in 96 h everywhere in domain 1. The initial temperature field is specified by imposing a latitude-dependent surface temperature and by prescribing a vertical temperature gradient. This gradient depends on latitude in the upper troposphere in order to simulate a sloping tropopause. In addition, a warm anomaly is specified over the Tibetan Plateau. Geopotential heights are calculated by specifying a constant height of the 700-hPa surface and solving the hydrostatic equation in the vertical. This way we ensure that mean winds are small in the Kali Gandaki area. Winds are assumed geostrophic. The Coriolis parameter at 30 N is chosen as representative for all domains; that is, we use an f-plane model. Tropospheric humidity is set to a quite low value (30% 40%) over most of the domain so as to avoid moist convection in regions where the horizontal resolution is too coarse to resolve it. Higher humidity values (50% 70%) are specified only in the Himalayan foreland where they are needed to produce the observed moist convection. Such quasi-stationary fields have been created for summer solstice, equinox, and winter solstice. Since the field experiment (see KG1) was carried out in September and October, the equinox field is taken for the reference run (REF; see also Table 1 for a list of experiments). The basic temperature data of this field are as follows. At 25 N, a temperature of 28 C is prescribed at 1000 hpa. The vertical temperature gradient is set to 8 K km 1 between 1000 and 700 hpa, to 6Kkm 1 between 700 and 500 hpa, and to 7.5 K km 1 in the rest of the troposphere. The tropospheric meridional temperature gradient is 2 K (1000 km) 1. The warm anomaly over the Tibetan Plateau is set to 9 K at the surface. A vertical temperature gradient of 9 Kkm 1 is assumed over the plateau until ambient temperatures are met. The warm anomaly specified over Tibet is also geostrophi-

6 MAY 2001 ZÄNGL ET AL FIG. 3. Potential temperature (contour interval 2 K) in cross section TB (see Fig. 2a); t 15 h. cally balanced. A weak surface low is prescribed in the summer and autumn fields to keep the fields realistic. In winter, no temperature anomaly and a very weak surface high are specified over Tibet. All simulations start at 1800 UTC. This time corresponds approximately to local midnight as the Kali Gandaki Valley is located near 84 E. In the following, the time UTC 6 h will be used as local time although the actual local time in Nepal is UTC 5¾ h. The simulations are continued until t 42 h except for some sensitivity tests, which end at t 30 h. This relatively short integration period is sufficient. For example, the fields in the afternoon of day 1 are almost identical to those of day 2. Results from day 2 are displayed only for nocturnal and morning conditions, otherwise the results from day 1 are taken. b. Large-scale fields Before presenting the valley wind simulations, we briefly discuss the large-scale fields obtained in REF (Figs. 2, 3). Figure 2a shows the surface wind and the surface potential temperature in the afternoon (15 h). A pronounced warm anomaly is located above the Tibetan Plateau. There, the potential temperature is about 325 K, which is 7 10 K more than that at comparable heights over India. This warm anomaly is similar to that prescribed at the beginning of the initialization and reflects the heat low over Tibet. The related pressure perturbation at a height of 5 km is about 2 hpa in the afternoon (not shown). During the night shallow layers of cold air form both above the mountains and the lowlands (not shown). The Tibetan heat low weakens during the night but does not disappear, because the warm anomaly lasts throughout the night above the shallow cold layer. The flow into the heat low in Fig. 2 is strongest along the southern slopes of the Tibetan Plateau with speeds of 5 ms 1. There is vigorous convective mixing above the plateau during the day where resolved vertical wind speeds can be as large as 0.4ms 1 despite the relatively low resolution (21.6 km in domain 2). Nocturnal downslope winds typically have speeds of 2 3 m s 1. Their relative weakness partly reflects the persistence of the heat low. The flow in the upper troposphere (Fig. 2b; z 8 km) is dominated by the westerly winds induced by the north south gradient of the temperature. Moreover, there is pronounced anticyclonic outflow above the plateau that balances the inflow at low levels. The anticyclone is associated with a jet to the north of the plateau that is, however, too weak when compared to observations. This deficiency is not surprising because the observed jet is part of the general circulation of the atmosphere while that in Fig. 2b owes its existence to local factors only. The isentropes at that height (not shown) are essentially oriented from west to east and show only a small impact of the Tibetan Plateau. The large-scale flow in the Kali Gandaki region, as marked by the frame in Fig. 2b, is weak both at the surface and at z 8 km. The structure of the boundary layer in the afternoon can be inferred from Fig. 3, where isentropes are plotted for a section across the Tibetan Plateau. There is a convectively mixed layer above Tibet with a depth of 3 4 km. This layer is almost adiabatic because there is little moist convection in the model somewhat in contrast to what is observed. The convection above the Indian lowlands is moist with a condensation level at 2.5 km. Above this level, the stratification is stable. The temperature differences between the Plateau and the free atmosphere are 7 10 K. 4. Results of the reference run In this section, the valley wind in the Kali Gandaki Valley and its diurnal cycle are discussed as obtained in REF. Note that the results are taken from the second simulation day until late in the morning (t 11 h) in order to exclude spinup effects. The surface winds at the end of the night (t 6 h) are shown in Fig. 4a. Throughout the valley, the downvalley wind is very weak with speeds hardly exceeding 2 m s 1. The same is true for the whole Mustang Basin (not shown). There are, however, stronger winds at altitudes above 5000 m. These appear to be slope winds. Winds and isentropes along the cross-section line TL (see Fig. 1b) are displayed in Fig. 5a. 1 The drainage 1 The sigma-coordinate system of the MM5 is based on a constant reference state that is defined by temperature and pressure. The actual pressure is the sum of the reference pressure and the perturbation pressure, which is displayed in Figs. 5a and 10a.

7 1068 MONTHLY WEATHER REVIEW FIG. 4. Surface winds (full barb 5ms 1 ; circles, 1.25 m s 1 ) in domain 5 in the reference experiment REF at (a) t 6 h (second simulation day), (b) t 11 h (second day), and (c) t 15 h (first day). Topographic shading starts at 3000 m with steps of 1000 m; darkest shading: h 6000 m.

8 MAY 2001 ZÄNGL ET AL FIG. 6. Surface wind (barbs) and isentropes (contour interval 3 K) at t 15 h in domain 4 in REF. FIG. 5. Cross-section parallel wind (vectors), potential temperature (solid; contour interval 1 K), and perturbation pressure [dashed; contour interval 0.05 hpa; (a) only] in cross section TL in REF: (a) t 6 h (second day); maximum wind speed 3 m s 1. Note that all pressure values are negative; and (b) t 15 h; maximum wind speed 12ms 1 ; shading marks wind speeds 10 m s 1. flow layer is very shallow. A layer with upvalley flow is seen between 500 and 1500 m above ground. Further above, there is again downvalley flow. Wind speeds are less than 5 m s 1 at all heights below 7 km. This result is rather close to the observations, which show similarly weak downvalley flows (Figs. 7 and 9 of KG1). The isentropes are nearly horizontal in Fig. 5 and there is virtually no pressure gradient. During the following hours, winds remain weak in the valley. First indications of a convectively mixed layer are seen at 9 h, but there is no upvalley flow. Wind speeds along the valley axis are below 2 m s 1 up to a height of 5 km. By 10 h, upvalley flow has set in. The velocity maximum is 7 ms 1 near Marpha in the narrowest part of the valley. The depth of the upvalley flow layer is about 1 km in most parts of the valley, and the largest speeds occur near the ground. One hour later (Fig. 4b), there is vigorous upvalley flow with maximum speeds of 15 m s 1. The maximum has shifted slightly toward the northeast. A jet is seen extending from Marpha toward the northeast. Wind speeds are clearly below 10 m s 1 to the south of Tukuche so that substantial accelerations take place to the north of Tukuche. All these features are again in good agreement with observations (see KG1). The rapid extension of the upvalley flow regime toward Tibet as seen in Fig. 4b is presumably not quite so realistic. Strong winds have been observed in Lo Manthang in the afternoon only. The strongest upvalley winds are obtained in the afternoon between 13 and 16 h. The surface wind field at t 15 h is displayed in Fig. 4c. Maximum wind speeds of 20 m s 1 occur now in the Jomsom area. The zone of strong winds extends farther to the north than at 11 h. Wind speeds are 15 m s 1 at the northern boundary of domain 5. Figure 6 reveals that winds with speeds between 10 and 13 m s 1 blow up to the border of Tibet. The wind has also strengthened between Ghasa

9 1070 MONTHLY WEATHER REVIEW and Tukuche where it now reaches some 10 m s 1. Winds are still weak to the south of Ghasa. This wind pattern is again in accordance with observations with maximum winds between Jomsom and Kagbeni and wind speeds above 10 m s 1 up to Lo Manthang. However, the maximum speed simulated around Jomsom is somewhat higher than reported in KG1. It is only on 28 September that velocities above 20 m s 1 were observed in the afternoon. In the Mustang Basin, a significant upvalley increase in potential temperature is evident from Fig. 6. Since this temperature pattern remains fairly stationary during the day, it can be inferred that temperature advection and local heating are in approximate balance. The isentropes in the TL cross section are sloping downward toward the north (Fig. 5b) up to a height of about 5 km, the slope being generally larger in the northern part. The condensation level is between 2.5 and 3 km in the southern parts (see also Fig. 3). The convective heating of the free atmosphere above the Himalayan foothills and the Indian lowlands is not as intense as that above the valley and therefore the temperature increases toward the north with concomitant pressure fall and acceleration. The flow structure in the TK cross section is striking (Fig. 7a). There is a pronounced descent of the isentropes between Marpha and Jomsom within a layer of at least 2000-m depth. Farther north, the isentropes are nearly parallel to the ground. The air is accelerated strongly in the descending part of the flow. The surface wind speed increases from 6 to 20 m s 1 over this distance. Correspondingly, the surface pressure decreases by 2 hpa between Tukuche and Jomsom (Fig. 7b). This is in good agreement with what one would obtain from the Bernoulli equation [see Eq. (2)], indicating that momentum balance is dominated by the pressure gradient and acceleration terms in this section. Farther toward the north, the pressure increases again but remains clearly lower than in Tukuche. Comparing the pressure fields of 6 and 15 h, we find a pressure fall of 3.5 hpa in Jomsom. This value is rather close to the observed daily pressure range of hpa (KG1). The flow pattern in Fig. 7a is strongly reminiscent of hydraulic flow through a horizontal contraction (e.g., Armi and Williams 1993). The pronounced inversion in Fig. 7a allows us to define a reduced gravity g g / 0.25 m s 2, where 8 K is the increase of the potential temperature across the top of the upvalley flow layer of depth h. For the Froude number V(g h) 1/2 to be close to unity, the flow speed V must be 16 m s 1 if h is assumed to be 1000 m. This means that the conditions for supercritical flow to occur may be satisfied near Marpha (Arakawa 1969). Of course, a shallow water model as used in the standard theories is not fully applicable in the Kali Gandaki Valley. Moreover, there are also wavelike features in this flow. Vertical motion at the surface follows the orography with downward motion over the downwind slopes of the ridges FIG. 7. Cross section TK at t 15 h in REF. (a) Cross-section parallel wind (vectors) and potential temperature (contour interval 1 K). Shading: light, wind speeds m s 1 ; medium, m s 1 ; dark, 20 m s 1 ; maximum wind speed: 23 m s 1. (b) Perturbation pressure, contour interval 0.25 hpa, negative values dashed. Shading denotes absolute wind speed as in (a). protruding into the valley (Fig. 8a). This wave pattern can be identified at 750 m above the surface, too (Fig. 8b). However, comparing the positions of the vertical wind extrema reveals that the wave pattern is shifted upwind at 750 m above ground. This phase shift suggests that vertically propagating gravity waves are excited over the ridges. In particular, the downward motion between Marpha and Jomsom shown in Fig. 7a can be traced back to these gravity waves as they are propagating both vertically and horizontally toward the valley

10 MAY 2001 ZÄNGL ET AL axis 2 (Smith 1980; Zängl 1999). Of course, the vertical propagation of the gravity waves is limited by the depth of the upvalley flow layer, which provides a kind of a critical level on its top. Durran and Klemp (1987) have shown that the transition from a gravity wave pattern to hydraulic type flow is likely if a critical level is present in stratified flow. In this case, the wave energy is reflected at the critical level, leading to very strong winds below. However, the overall flow situation as encountered near Marpha is more complicated than what has been envisaged so far in theories of orographic flows. A future investigation will be devoted to this phenomenon. At the moment it is difficult to decide how close the flow in Figs. 7 and 8 is to reality. The observations allow one to allocate the section with the highest speeds to the Marpha Kagbeni section in support of the simulations. However, they do not indicate a reduction of the depth of the strong wind layer as suggested by Fig. 7a. On the other hand, it is almost impossible to derive such a depth from the observations of the rapidly drifting balloons with sufficient accuracy. The only measurements that are possibly suitable to verify the present simulations stem from a motorglider experiment that took place in February 1985 (Neininger and Reinhardt 1986). The data collected during the flight Nepal 11 on 7 February 1985 indicate descending isentropes between Marpha and Jomsom. At a height of 3500 m, an increase in potential temperature by 3 K over a distance of about 10 km was found. Moreover, maximum wind speeds of 15 m s 1 were observed near Jomsom. Both values are very close to those obtained in the winter simulation (REF-WI; see section 5). However, the observations do not reveal whether a gravity wave pattern as shown in Fig. 7a is really present because there were only two flight sections, one at 3200 m and one at 3500 m. Additional observations would be needed to resolve this issue. Later in the day, the valley wind remains strong in REF until 21 h between Marpha and Kagbeni, with maximum speeds of 14 m s 1 near Jomsom. Such strong winds have not been observed so late in the day. Farther north in the Mustang Basin, however, the surface wind weakens soon after sunset and has disappeared by 21 h. This is in agreement with observations. Another feature that is supported by observations is that the upvalley wind weakens earlier near the surface than farther above. This feature is seen most clearly at t 24 h (Fig. 9). The elevated upvalley flow layer appears to be a combined effect of surface cooling, which tries to establish a downvalley flow, and the persistent Tibetan heat low, which drives the upvalley flow above. It may 2 Note that the hump between Marpha and Jomsom visible in the cross section is due to the bending shape of the valley. The valley bottom has an almost constant slope in this region and cannot be made responsible for any wave phenomenon. FIG. 8. Horizontal wind (barbs) and vertical wind velocity (positive values solid; negative dashed) in inner subdomain of D5 (see Fig. 1b) at t 15 h in REF. (a) surface; contour interval 1 m s 1, zero line omitted; (b) ( 750 m above ground); contour interval 0.5ms 1, zero line omitted. be noted that the flow field at 6 h (Fig. 5a) still shows weak upvalley flow between 500 and 1500 m above ground. Moreover, Fig. 9 suggests that the upvalley flow still reaches the ground near Marpha at t 24 h. This is not inconsistent with the observations presented in KG1. They indicate upvalley flow at t 24 h in Marpha, even though these measurements were probably influ-

11 1072 MONTHLY WEATHER REVIEW a. Asymmetry between day and night FIG. 9. As Fig. 7a but for t 24 h. Maximum wind speed 7ms 1. enced by a synoptic-scale disturbance. Thus, the question of how realistic the midnight upvalley wind at Marpha is must be left open. The atmosphere above the lowlands is rapidly stabilized in the evening below the condensation level because rain is falling at that time in REF. Evaporation is accounted for in the model, so that the rain cools the subsaturated boundary layer. It is pointed out that the broad features of the modelled rain distribution are fairly realistic. The rain is largely restricted to the area south of Lete where the climate is, indeed, very humid (see KG1, especially Table 1). In this region, rain was frequently observed during the campaign. 5. Sensitivity experiments and discussion The agreement of observations and the MM5 numerical results is encouraging. Given the accuracy and type of data available, no major discrepancy between observations and simulations can be found. This leaves us with three major problems to be further discussed by aid of the sensitivity experiments 1) Why is the nocturnal downvalley flow so weak? 2) Why is the diurnal upvalley flow so intense? 3) What is the role of the Tibetan Plateau? There is another problem that has not been discussed so far. In winter, it is by no means obvious to what extent the Mustang Basin can be considered as an elevated heat source. Tibet is covered by a low-level anticyclone at that time (Murakami 1981). Thus there is one more question: 4) How does the valley flow vary with the seasons? There are several factors that may contribute to the pronounced day night asymmetry of the Kali Gandaki Valley wind. Certainly, the fact that the Tibetan heat low does not fully disappear during the night plays an important role in this context. As noted above, an elevated weak upvalley flow layer persists throughout the night in the reference simulation. Thus, downvalley flow can form only in a shallow layer near the surface due to local cooling within the valley. It will be shown at the end of this section that the day night asymmetry is reduced significantly in winter where a high instead of a heat low is present over Tibet. Further possible factors are related to moist processes. Latent heating and evaporational cooling have strong influence on the stratification of the atmosphere. Moreover, clouds may reduce nocturnal cooling preventing a drainage layer to be formed. To demonstrate the influence of moisture, a dry run REF-D has been conducted in parallel to REF. Moist processes are switched off completely in this run. The large-scale fields prescribed in REF-D have been chosen as close as possible to REF. The fields are identical to REF above z 5 km, and the initial surface temperature is also the same. It is primarily the structure of the boundary layer that is affected by the different model physics. Due to the absence of moisture, all of the solar energy received by the ground is transferred to the atmosphere as sensible heat. This leads to the generation of a well-mixed boundary layer of 4 km depth over India capped by an inversion. The convective layer in REF is deeper over India than in REF-D but the condensation level is about 2.5 km above the ground and the air is more stably stratified above that level than in the dry run. Correspondingly, there is a strong inversion on top of the convective layer in REF-D but almost none in the moist run. The nighttime flow (6 h, second day) in the TL cross section of REF-D is displayed in Fig. 10a (to be compared to Fig. 5a). The downvalley flow is much stronger than in REF. Vigorous drainage flow with speeds up to 8ms 1 evolves to the south of Ghasa where even a weak hydraulic jump is seen. Compared to REF (Fig. 5a), significant differences in the structure of the isentropes and in the pressure field are found. In REF, the isentropes are horizontal and there is almost no pressure gradient to drive a downvalley flow. In REF-D, however, the isentropes are sloping down toward the south, and a pressure drop of 0.3 hpa over a distance of 5 km is seen around Ghasa. Moreover, the static stability of the free atmosphere is clearly lower in REF-D than in REF. This is consistent with the differences in drainage flow strength since buouyancy forces act against the formation of strong slope winds in a stable environment. Although the nocturnal drainage flow is much more intense in REF-D than in REF, a substantial asymmetry remains between day and night. The upvalley wind in

12 MAY 2001 ZÄNGL ET AL FIG. 11. Idealized orography as used in the S 1 sensitivity experiments, domain 4. The straight line marks cross section I. The inner part of section I confined by the dashes is section II. FIG. 10. (a) As in Fig. 5a but for dry run REF-D; maximum wind speed 8ms 1. Note again that all pressure values are negative. (b) As in Fig. 7a but for dry run REF-D; darkest shading: 25 m s 1 ; maximum wind speed 27 m s 1. the TK section (Fig. 10b) is more intense than in REF (Fig. 7a) with a maximum speed of 27 m s 1 and shows indications of a hydraulic jump. The slope of the isentropes is steeper and the maximum horizontal temperature difference across the supercritical flow is now 9 K as compared to 5KinREF.Correspondingly, the pressure difference between Tukuche and Jomsom amounts to 3 hpa instead of 2 hpa. This increase in strength appears to be related to the existence of the stable layer on top of the convective boundary layer in REF-D mentioned above. To gain further insight into the physical processes underlying the differences between REF and REF-D, we switch to the idealized model orography. Since our idealized orography does not require such a high horizontal resolution as realistic orography, a larger number of experiments can be carried out with the same computational costs. Moreover, idealized orography allows one to assess the importance of orographic details. The orography of the innermost domain (domain 4) of the S 1 runs (see Table 1) is displayed in Fig. 11. The valley is south north oriented and the topography is symmetric with respect to the axis of the valley. There is a basin akin to the Mustang Basin and a narrow section of the valley. First, the situation at night is considered. Five experiments are conducted and compared to each other. Experiment S 1 uses the same initialization and the same physics parameterizations as REF; S 1 -V differs from S 1 in the vegetation specified over the lowlands. Instead of tropical forest, grassland is specified at all heights below 3000 m, corresponding to a moisture availability of 15% instead of 50%. This test allows for assessing the importance of soil moisture, which governs the Bow-

13 1074 MONTHLY WEATHER REVIEW FIG. 12. Isentropes (contour interval 1 K) and cross-section parallel wind (vectors) in section II at t 6 h (second day) in the runs (a) S 1, (b) S 1, -M, (c) S 1 -MV, and (d) S 1 -D. Shading: light, wind speeds 5 10 m s 1 ; medium, 10 m s 1. Maximum wind speeds: (a) 6, (b) 7.5, (c) 11, and (d) 11 m s 1. en ratio, that is, the ratio of sensible to latent heat fluxes transmitted to the atmosphere. Experiment S 1 -M again uses the same physics as S 1, but the enhanced initial moisture values to the south of the Himalayas (see section 3a) are removed. The combination of low humidity and modified vegetation yields S 1 -MV. Finally, S 1 -D uses dry physics as does REF-D. Despite somewhat larger wind speeds, the flow obtained in S 1 (Fig. 12a) is very similar to REF. A shallow drainage flow layer is present in the gently sloping part of the valley as well in the uppermost 500 m of the steeply sloping part. Farther downward, however, positive buoyancy decelerates the downvalley flow near zero. As in REF, the stable stratification over the Himalayan foothills can be traced back to the evaporation of rain that falls from late afternoon till evening. Other stabilizing processes could not be identified in this run because they are overridden by evaporational cooling. In particular, the reduction of surface cooling due to the cloud cover over the lower part of the valley is overcompensated by evaporation. Experiment S 1 -V yields essentially the same flow field as S 1 and is not shown here. Although the amount of rain falling in S 1 -V is somewhat lower than in S 1, the same evaporational cooling is attained since the relative humidity reaches 100% in both cases. In S 1 -M (Fig. 12b), no precipitation falls

14 MAY 2001 ZÄNGL ET AL along the valley axis, but there is still significant precipitation (3 8 mm) along the southern rim of the Himalayas. The weak westerly wind that is present over the Himalayan foothills advects the stabilized air to the lower Kali Gandaki Valley so that the static stability is only slightly weaker than in S 1. Correspondingly, the valley wind progresses somewhat farther down the steep part of the valley than in S 1, but overall the situation has not changed much. More significant differences are found in S 1 -MV (Fig. 12c). Due to the higher Bowen ratio, the daytime convective layer is 1 2 K warmer and about 600 m deeper than in S 1 -M (3.3 instead of 2.7 km). Moreover, the amount of rain falling in the late afternoon is negligible. As a result, the nocturnal static stability of the free atmosphere is significantly lower than in S 1 and S 1 -M, and the downvalley wind is much stronger (11 instead of 7.5 m s 1 in S 1 -M) extending over the whole steep part of the valley. Nevertheless, the static stability of the free atmosphere is far from dry adiabatic, indicating that stabilizing processes other than the evaporation of rain exist. A close inspection of the results suggests that downward advection and vertical stretching of the stable air flowing out of the Kali Gandaki Valley is at least partly responsible for the stabilization. In the neighborhood of the valley, the slope wind circulation along the Himalayan massif acts in the same way. Another possible stabilizing process is differential radiative cooling. Since the water vapor mixing ratio increases toward the ground, radiative cooling of the free atmosphere can be expected to increase toward the ground, thereby stabilizing the atmosphere. Inspection of the cooling rates computed by the radiation scheme shows that longwave cooling, indeed, increases toward the ground. However, the corresponding stabilizing effect cannot be dominating because the cooling rates of the free atmosphere are too low overall not exceeding 4 K day 1 in the absence of clouds. This is confirmed by an experiment in which S 1 -MV was repeated with a very simple radiation scheme, parameterizing longwave cooling as a function of temperature only. Although the vertical gradient of longwave cooling provided by this scheme is much weaker than that obtained from the more realistic scheme, the results were almost identical. The dry run S 1 -D (Fig. 12d) is very similar to S 1 - MV. Stability and maximum velocity hardly differ from S 1 -MV, and the downvalley wind zone extends down to the lowlands. From this set of experiments, we infer that moisture contributes to the weakness of the nocturnal downvalley wind to the degree that convective rain stabilizes the atmosphere over the Himalayan foothills by evaporative cooling. Moreover, the moisture provided by the lush vegetation on the foothills reduces the condensation level and therefore the depth of the daytime dry-adiabatic layer. Thus, the thickness of the layer that has to be stabilized during the night is not as large as in the drier runs. However, the contribution of the latter effect to the nocturnal stratification can hardly be separated from evaporational cooling. Comparing the results obtained with idealized orography with those based on realistic orography (REF and REF-D), we find that realistic orography yields weaker downvalley flow than our idealized orography. It is suggested that the stepwise and bending descent of the valley from Tukuche to the foothills is less susceptible to strong downvalley flow than the steep and straight descent in the idealized orography. To complete the analysis of the day night asymmetry of the Kali Gandaki Valley wind, the reasons for the extreme strength of the daytime upvalley wind have to be explored further. As has already been suggested during the discussion of REF and REF-D, vertically propagating gravity waves generated by the protruding mountain ridges play a crucial role in establishing the high wind speeds obtained around Jomsom. Thus, the question arises what the flow would be like without these mountain ridges. Moreover, the reason for the stable stratification in the Kali Gandaki Valley has to be clarified. Since a deep convectively mixed layer is present over both the Himalayan foothills and the Tibetan Plateau (see Fig. 3), one may wonder why this is not the case in the Kali Gandaki. Obviously, the combination of vertically propagating gravity waves and supercritical flow obtained in REF and REF-D would not be possible in a deep dry-adiabatic boundary layer. Again, idealized model orography is used to further investigate these questions. First, the S 1 runs already discussed above are reconsidered focusing now on the daytime upvalley flow. Note that the width of the valley is constant between the southern exit and the point where the widening to the idealized Mustang Basin begins (Fig. 11). This rules out the generation of gravity waves by protruding ridges as in REF (see Fig. 8). In S 1 -D (Fig. 13a), we obtain again a valley jet with strong acceleration where the valley widens to the basin, but there are only small indications of gravity wave activity. The maximum speed is about l6 m s 1 as compared to 27ms 1 in REF-D. This indicates that orographic details are important for the exceptional strength of the winds around Jomsom although other factors may also contribute to the differences between REF-D and S 1 -D. The overall difference of the topographies is rather large so that one cannot uniquely attribute the changes in wind speed to the specific change of topography in the narrow part of the valley. The advection of cold air up to Tibet can be seen clearly in the boundary layer above the slope. Interestingly, all the different S 1 runs discussed above show almost the same maximum wind speed. Despite large differences in the depth of the dry-adiabatic layer over the Himalayan foothills, the maximum upvalley velocity obtained in S 1 is only 1 m s 1 lower than in S 1 -D (not shown). Although the differences are somewhat larger with realistic orography, it is concluded that