Extensive dust outbreaks following the morning inversion breakup in the Taklimakan Desert

Transcription

1 JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 110,, doi: /2005jd005994, 2005 Extensive dust outbreaks following the morning inversion breakup in the Taklimakan Desert Nobumitsu Tsunematsu, 1,2 Tomonori Sato, 1,3 Fujio Kimura, 4,5 Kenji Kai, 6 Yasunori Kurosaki, 1,2,7 Tomohiro Nagai, 8 Hongfei Zhou, 9 and Masao Mikami 10 Received 17 March 2005; revised 17 July 2005; accepted 18 August 2005; published 9 November [1] Extensive dust outbreaks occurred on the late morning of 26 March 2004 in the Taklimakan Desert. An image by the satellite AQUA shows the dust outbreaks as an enormous dust storm extending as far as 1000 km in the direction of east and west across the desert. A ground-based lidar was successful in observing the dust storm. Values of the backscattering ratio from near the ground surface to an altitude of 6 km above sea level rapidly increased as soon as the dust storm covered the lidar observation site. Numerical experiments including a diffusion experiment revealed that strong westerly winds appeared at the surface level following the morning inversion breakup and then induced the dust outbreaks. The nocturnal inversion layer prevented the strong westerly winds from spreading to the surface level until the breakup. Notable nocturnal inversion is considered to be easily formed in the Taklimakan Desert because of its deep basin topography which can accumulate cold air masses. These results show the stability of the planetary boundary layer to be an extremely important factor in dust outbreaks especially in basin deserts, in addition to the intensity of surface winds. This implies that the occurrence frequency of dust storms in the desert has a diurnal cycle. Citation: Tsunematsu, N., T. Sato, F. Kimura, K. Kai, Y. Kurosaki, T. Nagai, H. Zhou, and M. Mikami (2005), Extensive dust outbreaks following the morning inversion breakup in the Taklimakan Desert, J. Geophys. Res., 110,, doi: /2005jd Introduction [2] Dust particles, which reduce the surface net radiation and play a role as active ice nuclei, can influence the global climate, because of the radiative forcing [Tegen et al., 1996; Miller and Tegen, 1998; Sassen, 2002]. Dust particles emitted from arid and semiarid regions in the northwestern part of China are transported to Korea, Japan, the North Pacific, and North America [Iwasaka et al., 1983; Kai et al., 1988; Gao et al., 1992; Chun et al., 2001; Husar et al., 2001; Uno et al., 2001; Uematsu et al., 2002]. The Gobi 1 Japan Science and Technology Agency, Kawaguchi, Japan. 2 Meteorological Research Institute, Tsukuba, Japan. 3 University of Tsukuba, Tsukuba, Japan. 4 Frontier Research Center for Global Change, JAMSTEC, Yokohama, Japan. 5 Graduate School of Life and Environmental Sciences, University of Tsukuba, Tsukuba, Japan. 6 Graduate School of Environmental Studies, Nagoya University, Nagoya, Japan. 7 Now at Center for Environmental Remote Sensing, Chiba University, Chiba, Japan. 8 Meteorological Satellite and Observation System Research Department, Meteorological Research Institute, Tsukuba, Japan. 9 Xinjiang Institute of Ecology and Geography, Chinese Academy of Sciences, Urumqi, China. 10 Atmospheric Environment and Applied Meteorology Research Department, Meteorological Research Institute, Tsukuba, Japan. Copyright 2005 by the American Geophysical Union /05/2005JD Desert and the Taklimakan Desert, shown in Figure 1, are the main sources of dust particles. Dust outbreaks in these deserts, which frequently occur from March to May [Kurosaki and Mikami, 2003], can have serious environmental consequences for East Asia, including the visibility obstruction and health hazards associated with the deposition of large amounts of dust particles. [3] Dust outbreaks in the Gobi Desert are closely connected with synoptic-scale cyclone activity. Qian et al. [2002] showed on the basis of climatological research that the occurrence of dust storms in the desert is strongly related to high-frequency cyclone activity in spring. Shao et al. [2002] and Shao and Wang [2003] also indicated that almost all dust events in the desert arise from the strong northwesterly wind associated with cyclones. Mitsuta et al. [1995] and Takemi [1999] analyzed a severe dust storm in the desert caused by downdrafts emanating from a longlived squall line near a well-developed cold frontal system. These studies emphasize the influence of cold air surges associated with cyclone activity on the dust outbreaks. The degree of the influence is determined by snow cover on the desert [Kurosaki and Mikami, 2004]. [4] The relationship between dust outbreaks and synopticscale cyclone activity in the Taklimakan Desert is less clear [Takemi and Seino, 2005]. They mentioned that dust weather in the desert is affected by cold-air intrusion recognized as the easterly wind prevailing far behind the synoptic-scale cyclone center. Chen and Chen [1987] discovered that the easterly wind is associated with the strong pressure gradient and 1of12

2 Figure 1. Map and topography around the Taklimakan Desert. The solid dot indicates the lidar observation site. The box represents the range of Figures 2, 5, 6, 7, and 10. Sea and lake areas are shaded. important to the occurrence of dust storms in the desert. Tsunematsu [2005] showed that the front of the easterly wind is in the shape of the gravity current, by the analysis of data from lidar observations. This easterly wind usually originates from north of the Tienshan Mountains, then blows along the eastern rim of the mountains toward the desert. The dust particles uplifted by the easterly wind can be entrained to altitudes above 5000 m and transported for long distances by the westerlies [Sun et al., 2001]. Aoki et al. [2005] classified the intruding courses of cold-air inducing dust storm outbreaks in the desert into three patterns. The first pattern is the above-mentioned one (the easterly wind pattern), second is the northerly wind pattern crossing over the Tienshan Mountains, and third is the westerly wind pattern crossing over the Pamirs. These studies revealed the local wind systems to be an important factor for dust outbreaks in the Taklimakan Desert. However, the dynamical aspects of dust outbreaks in the desert have not been clarified completely. There are some cases where the cause of the dust outbreaks cannot be explained only by the local wind systems. Some mechanism other than the local wind systems should contribute to the outbreaks. [5] The present study performed lidar observations in the Taklimakan Desert in March 2004 to observe dust events. The lidar observations were conducted as part of a joint project between Japan and China, the Aeolian Dust Experiment on Climate Impact (ADEC), which aimed at the evaluation of wind erosion and long-range transport processes of dust particles in Asia [Mikami et al., 2002]. The lidar was successful in observing an enormous dust storm on the late morning of 26 March The dust storm was imaged by the Moderate Resolution Imaging Spectroradiometer (MODIS) of the satellite AQUA, as shown in Figure 2a. (This image was obtained from a web page, MODIS Rapid Response System ; gsfc.nasa.gov/.) In the present study, the mechanism of the dust storm outbreak was investigated by reproducing the meteorological conditions over the Taklimakan Desert, using a regional meteorological model. A diffusion experiment was also conducted in order to reproduce the dust storm. 2. Lidar Observations [6] The lidar observations were performed at the Aksu Water Balance Experimental Station of the Xinjiang Institute of Ecology and Geography, Chinese Academy of Sciences. The station is located in the northern part of the Taklimakan Desert (40.62 N, E, 1028 m above sea level (ASL); the solid dot in Figure 1). The Taklimakan Desert is located in a big basin (the Tarim Basin), which is surrounded by high plateaus (the Tibetan Plateau and the Pamirs) and mountains (the Tienshan Mountains) exceeding an altitude of 4000 m ASL. The total area is approximately 400,000 km 2. The desert consists almost entirely of sand and gobi (gravel; grit), except for oasis areas distributed along the rim of the desert (see Figure 2b). The Local Solar Time (LST) in the desert is 6 hours ahead of the Coordinated Universal Time (UTC). [7] The lidar used in the present study is capable of observing the backscattered light of air molecules and aerosol particles at intervals of 5 min with a spatial resolution of 7.5 m. The specifications of the lidar system are listed in the paper by Kai et al. [2002]. This lidar has 2of12

3 Figure 2. Satellite image of the Taklimakan Desert at 0750 UTC 26 March 2004 by MODIS/AQUA: (a) large image and (b) sketch of the large image. An enormous dust storm is happening in the desert. The star in Figure 2a indicates the lidar observation site. The thick solid line in Figure 2b is the outline of the dust storm. The shading and two dotted lines in Figure 2b represent the oasis areas and the boundary of the Tarim Basin, respectively. The horizontal range corresponds to the box in Figure 1. two telescopes as receivers: diameters of 200 mm and 355 mm, in order to measure the backscattered light from near the ground surface to the lower stratosphere. A flashlamp-pumped Nd:YAG laser is transmitted at a 532 nm wavelength. The backscattered light is detected by photomultiplier tubes and processed by the 12-bit A/D converter installed in the transient recorder. Aerosol particles are quantified as the backscattering ratio (R). The value of R at height z is given as follows [Sakai et al., 1997]: Rz ðþ¼ b mðþþb z a ðþ z ; ð1þ b m ðþ z Figure 3. Time-height cross section of the backscattering ratio (R) at the lidar observation site from 0000 UTC 25 March to 0000 UTC 27 March An altitude of the observation site is 1028 m ASL. 3of12

4 where b m and b a are the molecular and aerosol backscattering coefficients. The normalization of R is carried out on the assumption that R equals 1 at altitudes above 15 km ASL where aerosol backscattering is considered to be almost negligible. [8] Figure 3 shows a time-height cross section of the backscattering ratio (R) in the troposphere for the period from 0000 UTC 25 March to 0000 UTC 27 March On 25 March, the values of R at altitudes below 3 km ASL are 2 3, indicating that the quantity of dust particles is relatively large in the planetary boundary layer. The high values (8) above 7 km ASL from 0000 UTC to 0800 UTC 25 March are caused by ice crystals judging from previous lidar observations of ice clouds [Sassen et al., 2003]. [9] The values of R in the lower troposphere suddenly increase at around 0500 UTC 26 March. Height of the high-values layer increases to more than 4 km ASL after 0700 UTC. The high values are due to a large quantity of dust particles associated with the dust storm, as confirmed by the satellite image captured at 0750 UTC 26 March (Figure 2a). In the satellite image, the white areas outside of the Tarim Basin (outside of the dotted lines in Figure 2b) indicate snow cover and clouds. Within the basin, rivers and lakes (or ponds) can be clearly seen in the southwestern part of the Taklimakan Desert and the oasis areas. Whereas, most of the desert is obscured with dust, and surface features cannot be recognized. This confirms the existence of a dust storm extending as far as 1000 km from west to east across the desert. The outer bounds of the dust storm are outlined in Figure 2b. A plump appearance of the dust storm implies that strong surface winds are blowing a lot of dust particles up from the ground surface into the atmosphere. The dust concentration in the east of the dust storm appears to be higher than that in the west where the surface conditions can be vaguely seen through the dust. Thus the sudden increase in the values of R are confidently correlated with the dust storm outbreak. The top of the dust storm reaches to nearly 6 km ASL at around 0900 UTC (Figure 3). Figure 4. The 300 hpa isobaric weather charts around the Taklimakan Desert on 26 March 2004: (a) 0000 UTC, (b) 0600 UTC, and (c) 1200 UTC. The thick solid lines indicate the geopotential heights (m) every 120 m. The arrows indicate the wind directions and velocities (m/s). The thin solid lines represent contour lines of the topography at intervals of 2000 m. The straight line indicates the axis of the trough. The solid dot in Figure 4a indicates the lidar observation site. Sea and lake areas are shaded. The range corresponds to Figure Meteorological Conditions on 26 March 2004 [10] Figure 4 shows the 6-hourly 300 hpa isobaric weather charts including the geopotential heights and wind vectors for the period from 0000 UTC to 1200 UTC 26 March 2004, obtained from the National Centers for Environmental Prediction/National Center for Atmospheric Research (NCEP/NCAR) reanalysis data [Kalnay et al., 1996]. The wind field on this isobaric surface is hardly influenced by the very high mountains and plateaus around the Taklimakan Desert, and therefore the wind field correctly shows the synoptic-scale airflow around the desert. A deep trough extends from north of 50 N to the Taklimakan Desert at 0000 UTC 26 March (Figure 4a). The axis of the trough reaches the east end of the desert at 0600 UTC (Figure 4b). Behind the trough, strong northwest-to-westerly winds prevail over the desert with velocities of more than 30 m/s. By 1200 UTC, the wind directions over the desert generally turn to the west and the velocities slightly decrease as the trough leaves the desert (Figure 4c). [11] Figure 5 shows the 3-hourly observed surface wind fields at 10 m above ground level (AGL) from 0000 UTC to 0900 UTC 26 March 2004, obtained from SYNOP reports 4of12

5 Figure 5. Observed surface wind fields in the Taklimakan Desert on 26 March 2004: (a) 0000 UTC, (b) 0300 UTC, (c) 0600 UTC, and (d) 0900 UTC. The barbs and pennants with a circle represent the wind direction and velocity (m/s) at each observatory. Each barb and pennant indicates the velocity of 1 m/s and 5 m/s, respectively. The solid dot without the barbs and pennants indicates calm. The solid lines represent contour lines at intervals of 1000 m. The star in Figure 5a indicates the lidar observation site. The range corresponds to the box in Figure 1. Table 1. Features of the Model and the Settings Features Basic equation and coordinate nonhydrostatic equations, terrain-following coordinate Calculation domain center of domain: 40.1 N, 82.7 E; number of horizontal grid points: 50 43; interval of horizontal grid points: 25 km; number of vertical layers: 56; intervals of vertical layers: m (varies with height) a Data for initial and boundary conditions NCEP/NCAR 6-hourly data on March 2004 [Kalnay et al., 1996] Physical processes radiation [Nakajima et al., 2000], cumulus convection [Arakawa and Schubert, 1974], cloud microphysics [Walko et al., 1995], turbulence [Mellor and Yamada, 1982], surface process [Louis, 1979], soil model a Lowest layer: 20 m, highest layer: 900 m. Lowest level for wind vectors: Z* = 10 m. 5of12

6 Figure 6. Calculated wind vectors at Z* = 10 m in the Taklimakan Desert on 26 March 2004: (a) 0000 UTC, (b) 0300 UTC, (c) 0600 UTC, and (d) 0900 UTC. The arrows indicate the wind directions and velocities (m/s). The thick arrows represent the wind velocities larger than 6.5 m/s. Interval of the solid contour lines is 1000 m. The thick dashed line in Figure 6a indicates the section of Figure 8. The star in Figure 6a indicates the lidar observation site. The range corresponds to the box in Figure 1. [World Meteorological Organization, 1974] for 10 meteorological observatories located in the Taklimakan Desert. The wind velocities over the desert are small at 0000 UTC (Figure 5a), but slightly increase by 0300 UTC (Figure 5b). An obvious change in the wind field occurred after this time; the wind fields at 0600 UTC and 0900 UTC show strong westerly winds prevailing in the middle of the desert between 39 N and 41 N with velocities of more than 6 m/s (Figures 5c 5d). The strong westerly winds appear to be associated with the strong northwest-to-westerly winds in the upper level, i.e., the synoptic winds (Figure 4b). [12] Detailed analysis of the wind fields over the Taklimakan Desert was performed by numerical simulations. A numerical model used in the present study is a modification of the Regional Atmospheric Modeling System (RAMS) developed at Colorado State University [Pielke et al., 1992]. The modified RAMS was developed at the Terrestrial Environment Research Center (TERC) of University of Tsukuba, Japan (TERC-RAMS). It has been used in many previous studies [e.g., Yoshikane et al., 2001; Okamura and Kimura, 2003; Sato and Kimura, 2003]. Features of the model and the settings in the present study are listed in Table 1. The model is based on compressive nonhydrostatic equations and a terrain-following coordinate. The radiation scheme of the original RAMS was replaced with that of Nakajima et al. [2000]. The horizontal grid 6of12

7 Figure 7. Same as Figure 6 but for Z* = 524 m. interval is set at 25 km. A center of the calculation domain represented by a box in Figure 1 is N, E. The intervals of vertical layers are 20 m at the lowest layer and gradually expand upward up to the maximum value of 900 m at an expansion rate of 1.1 times. The terrain-following coordinate level of z* = 10 m is the lowest level (surface level) for the wind vectors. The initial time of numerical integration is 0000 UTC 24 March The vegetation type is assumed to be desert for the entire domain except for lake areas. The 6-hourly NCEP/NCAR reanalysis data are used for the initial and boundary conditions. The nudging technique is applied to the outer five grids of the domain. [13] Figure 6 shows the calculated wind vectors at the level of z* = 10 m in the Taklimakan Desert from 0000 UTC to 0900 UTC 26 March 2004 at intervals of 3 hours. The wind velocities are small in almost all parts of the desert at 0000 UTC (Figure 6a), but increase in several places by 0300 UTC (Figure 6b). The wind vectors at 0600 UTC clearly show that strong westerly winds prevail over a wide area of the desert (Figure 6c). The observed wind field at this time also shows the strong westerly winds (Figure 5c). The westerly wind velocities slightly decrease by 0900 UTC (Figure 6d). [14] The calculated wind vectors for the period from 0000 UTC to 0900 UTC 26 March 2004 at the level of z* = 524 m are shown in Figure 7. Strong northwesterly winds surge toward the eastern part of the Taklimakan Desert at 0000 UTC (Figure 7a). The strong northwesterly winds have gone down the Tienshan Mountains. The synoptic wind vectors shown in Figure 4a indicate that these northwesterly winds have originated from far to the northwest of the mountains. The northwesterly winds blow hard over the desert at 0300 UTC, turning to the westerly winds in the eastern part of the desert (Figure 7b). The wind velocities 7of12

8 Figure 8. East-west vertical cross sections of calculated potential temperatures (K) on the thick dashed line in Figure 6a (Y = 500 km) on 26 March 2004: (a) 0000 UTC, (b) 0300 UTC, and (c) 0600 UTC. The solid lines indicate the potential temperatures at intervals of 3 K. The shading represents the topography. gradually decrease after this time, as seen in the wind vectors at 0600 UTC and 0900 UTC (Figures 7c 7d). [15] Note that difference in the wind velocities between those levels over the Taklimakan Desert is large until 0300 UTC (Figures 6a, 6b, 7a, and 7b) but relatively small after this time (Figures 6c, 6d, 7c, and 7d). The strong westerly winds prevailing at the surface level (z* = 10 m) after 0300 UTC (Figures 6c and 6d) have already blown hard at the upper level (z* = 524 m) prior to this time (Figures 7a and 7b). These indicate that a transfer of high momentum from the upper level toward the surface level occurred after 0300 UTC. [16] Figure 8 shows east-west vertical cross sections of the calculated potential temperatures from 0000 UTC to 0600 UTC 26 March 2004 at intervals of 3 hours along the section represented by a thick dashed line in Figure 6a. A large vertical gradient of potential temperatures forms at altitudes below 0.5 km AGL over the Taklimakan Desert at 0000 UTC (0600 LST), representing the existence of the notable nocturnal inversion layer (Figure 8a). The potential temperatures above the inversion layer up to 3 km ASL in the western part of the desert are relatively low in comparison with those in the eastern part, indicating that cold air masses penetrate into the desert from the west. The horizontal wind vectors in Figure 7a also show this air current. Thus the nocturnal inversion was considered to be strengthened by the penetration of cold air masses, coupled with the nocturnal radiational cooling. The potential temperatures at 0300 UTC show the beginning of the morning inversion breakup (Figure 8b), and the vertical gradient of potential temperatures at altitudes below 0.5 km AGL decreases to nearly zero by 0600 UTC (Figure 8c). This inversion breakup was caused by the development of the mixed layer due to the daytime ground surface heating rather than a weakening of the cold air current, as demonstrated in Figure 9, which graphs the average surface sensible heat flux on the thick dashed line in Figure 6a. In Figure 9, the sensible heat flux becomes positive between 0100 UTC and 0200 UTC, and then rapidly increases. The strong ground heating affected the inversion breakup, as described in Arya [1988]. [17] These results reveal that the notable nocturnal inversion layer prevented the strong westerly winds from spreading to the surface level until its breakup. Following the breakup, the strong westerly winds spread to the surface level, resulting in the large wind velocities at the level. 4. A Diffusion Experiment on Dust Particles [18] Diffusion experiments on dust particles using the RAMS output are expected to be effective for reproducing Figure 9. Calculated hourly surface sensible heat flux (W/m 2 ) on the thick dashed line in Figure 6a (Y = 500 km) from 0000 UTC to 0600 UTC 26 March Each value was averaged for all grid points on the line. 8of12

9 Figure 10. Calculated dust distributions in the Taklimakan Desert on 26 March 2004: (a) 0200 UTC, (b) 0400 UTC, (c) 0600 UTC, and (d) 0800 UTC. Each dot represents a dust particle. The star in Figures 10a and 10d indicates the lidar observation site. The thick solid line and dotted line in Figure 10d indicate the sections of Figures 11a and 11b, respectively. Interval of the thin solid contour lines is 1000 m. The range corresponds to the box in Figure 1. and investigating the dust storm outbreak through comparison with the satellite image and lidar observations. The quantity of dust particles emitted from the ground (q) is estimated by the following equation: q ¼ Cu ð u tr Þu 2 ; ð2þ where C is a coefficient, u is the surface wind velocity (m/s), and u tr is the threshold wind velocity for sand soil erosion. This equation, followed Gillette [1978], has been frequently used in numerical studies of dust emission and transport [e.g., Tegen and Fung, 1994; Tegen and Miller, 1998; Uno et al., 2001], although some quantitative studies of dust used different equations [e.g., Marticorena et al., 1997; Shao, 2001]. This simple model is sufficient to reproduce and investigate the dust storm outbreak. The value of u is obtained from the wind velocity at the level of z* = 10 m, which is almost equal to 10 m AGL. Kalma et al. [1988] state that the threshold wind velocity u tr at 10 m AGL is 6.5 m/s. The present study adopts this threshold. Wind velocities larger than 6.5 m/s at the level of z* = 10 m are denoted by the thick arrows in Figure 6. [19] In this diffusion experiment, the dust particles are allowed to be emitted from altitudes below 1300 m ASL in the Tarim Basin excluding the oasis areas. The coefficient C is constant for the entire area. The initial time of dust 9of12

10 Figure 11. Vertical cross sections of the calculated dust distribution at 0800 UTC 26 March 2004: (a) east-west vertical cross section along the thick solid line in Figure 10d (Y = 500 km) and (b) northsouth vertical cross section along the dotted line in Figure 10d (X = 450 km). The shading represents the topography. The large arrow in Figure 11b indicates the location of the lidar observation site. emission is set at 0000 UTC 26 March A random walk model presented by Kimura and Yoshikawa [1988] is used in the present study in order to calculate the movement of a large number of the Lagrangian particles. The Runge- Kutta method [Press et al., 1986] is applied for numerical integration of the three-dimensional advection of particles with a time step of 5 s. The horizontal turbulent diffusion of particles is set at a width of 0.1 km per 1.0 km over the distance of advection on the assumption of the Pasquill stability class C for the atmospheric stability, and the concentration follows the Gaussian distribution. The width of vertical turbulent diffusion is determined by the vertical diffusivity obtained from the RAMS prognosis. [20] Figure 10 shows the calculated dust distributions over the Taklimakan Desert for the period from 0200 UTC to 0800 UTC 26 March 2004 at intervals of 2 hours. The dust particles (dots in Figure 10) are projected onto the ground surface, so that the density of particles is proportional to the concentration integrated from the surface to the top of the atmosphere. The dust emission is inactive in the greater part of the desert until 0200 UTC (Figure 10a). Extensive dust outbreaks occur after 0400 UTC (Figure 10b), and some of the dust particles are distributed in the vicinity of the lidar observation site (the star in Figure 10a) at 0600 UTC (Figure 10c). This is consistent with the lidar observation results, Figure 3, which shows the high values of R after 0500 UTC. A large number of the dust particles cover the greater part of the desert with a belt-like distribution at 0800 UTC (Figure 10d). This belt-like distribution closely resembles the shape of the real dust storm imaged by 10 of 12

11 the satellite at 0750 UTC (Figure 2). Figure 10d also shows the high dust concentrations in the eastern part of the dust storm, as recognized in the satellite image. [21] The extensive dust outbreaks could not occur while the notable nocturnal inversion layer existed over the Taklimakan Desert, because the inversion layer prevented the spread of the strong westerly winds to the surface level (Figures 5a, 6a, 7a, 8a, and 10a). Following the morning inversion breakup, however, the westerly winds spread to the surface level and induced the dust outbreaks over a wide area of the desert (Figures 5c, 6c, 7c, 8c, and 10c), resulting in the formation of the enormous dust storm. Many previous studies show that most of dust storms in Chinese deserts occur in association with strong winds behind a cold front [e.g., Littmann, 1991; Sun et al., 2000, 2001]. They mainly emphasizes the effect of the strong winds on the dust storms. However, the numerical experiments conducted here reveal that destabilization of the planetary boundary layer due to the ground surface heating is also an extremely important factor in the dust storm outbreaks. [22] Nocturnal cold air advection such as the drainage flow intensifies the cooling of the planetary boundary layer over the basin bottom in comparison with the boundary layer over flat terrain [Kondo et al., 1989]. The Tarim Basin is deep enough to accumulate the nocturnal cold air easily, and therefore the notable nocturnal inversion layer should be frequently formed over the basin because of the cooling. The notable inversion layer acts as a barrier to the strong synoptic wind, thereby suppressing dust outbreaks in the Taklimakan Desert. The synoptic wind which intruded into the Tarim Basin across the surrounding high mountains appears to mainly glide over the notable inversion layer without diving deep into the layer. In contrast, the nocturnal inversion layer over the flat terrain such as the Gobi Desert is considered to be directly destroyed from the inside by the strong synoptic wind blowing near the surface. These discussions indicate that the dust storms in the Taklimakan Desert are hard to occur during the nighttime. A diurnal cycle of the occurrence frequency of dust storms in the desert is expected to be shown by future studies. [23] Figure 11 depicts east-west and north-south vertical cross sections of the calculated dust distribution at 0800 UTC 26 March 2004 along the thick solid and dotted lines in Figure 10d. The top of the dust storm is nearly 2 km ASL in the direction of east and west (Figure 11a), corresponding to the top of the mixed layer recognized in Figure 8c. Similarly, the north-south distribution shows that the dust particles over the lidar observation site are uplifted to km ASL (Figure 11b). However, the top of the observed dust storm was approximately 5 km ASL at 0800 UTC (Figure 3). This difference indicates that another factor in addition to those considered in the calculation contributed to the rise of the dust particles. The present study hypothesized about this as follows. The mixed layer developed deeper in practice, because the dust layer (dust storm) aloft absorbed the solar radiation and was heated strongly. The dust particles were uplifted higher by the vertical turbulent diffusion in the deep mixed layer. The interaction between the dust layer and the solar radiation may possibly determine the vertical distribution of dust particles in the lower troposphere. In addition, the lidar observation site was located at the edge of the dust storm (Figure 10d). It may be hard for numerical experiments to reproduce the distribution of dust particles at the edge of a dust storm completely, because of their spatial resolution. 5. Summary [24] A ground-based lidar observed an outbreak of dust storm in the Taklimakan Desert on the late morning of 26 March The values of R from near the ground surface to an altitude of 6 km ASL rapidly increased as soon as the dust storm covered the lidar observation site. An AQUA satellite image showed the dust storm extending as far as 1000 km in the direction of east and west. Meteorological analyses and numerical experiments revealed that the enormous dust storm was caused by strong westerly winds originating from far to the northwest of the Tienshan Mountains. The numerical experiments also demonstrated that the notable nocturnal inversion layer formed over the desert, preventing the strong westerly winds from spreading to the surface level. Following the morning inversion breakup, the strong westerly winds spread to the surface level, and then uplifted dust particles over a wide area of the desert, resulting in the enormous dust storm. This phenomenon shows that the stability of the planetary boundary layer has a considerable effect on dust outbreaks, in addition to the intensity of surface winds. The effect can be expected to be applicable to basin deserts such as the Taklimakan Desert, because this topography is conducive to the accumulation of cold air masses and the formation of notable nocturnal inversion. There is a possibility that the nocturnal inversion layer and its breakup determine a diurnal cycle of the occurrence frequency of dust storms in the desert. [25] Acknowledgments. We are grateful to Makoto Abo (Tokyo Metropolitan University), Takatsugu Matsumura (Japan Science and Technology Agency/Meteorological Research Institute), Takuya Matsumoto (Hamamatsu Photonics K. K.), and Makoto Goto (Nagoya University) for the lidar observations. Isao Aoki (CRC Solutions Corporation) contributed to this study for the numerical experiments. The present study was performed through Special Coordination Funds for Promoting Science and Technology of Ministry of Education, Culture, Sports, Science and Technology of Japan. References Aoki, I., Y. Kurosaki, R. Osada, T. Sato, and F. Kimura (2005), Dust storms generated by mesoscale cold fronts in the Tarim Basin, northwest China, Geophys. Res. Lett., 32, L06807, doi: /2004gl Arakawa, A., and W. H. Schubert (1974), Interaction of a cumulus cloud ensemble with the large-scale environment, part I, J. Atmos. Sci., 31, Arya, S. P. (1988), Introduction to Micrometeorology, 420 pp., Elsevier, New York. Chen, G. T.-J., and H.-J. Chen (1987), Study on large-scale features of duststorm system in east Asia, Pap. Meteorol. Res., 10, Chun, Y., K.-O. Boo, J. Kim, S.-U. Park, and M. Lee (2001), Synopsis, transport, and physical characteristics of Asian dust in Korea, J. Geophys. Res., 106, 18,461 18,469. Gao, Y., R. Arimoto, M. Y. Zhou, J. T. Merrill, and R. A. Duce (1992), Relationships between the dust concentrations over eastern Asia and the remote North Pacific, J. Geophys. Res., 97, Gillette, D. (1978), A wind tunnel simulation of the erosion of soil: Effect of soil texture, sandblasting, wind speed, and soil consolidation on dust production, Atmos. Eviron., 12, Husar, R. B., et al. (2001), Asian dust events of April 1998, J. Geophys. Res., 106, 18,317 18,330. Iwasaka, Y., H. Minoura, and K. Nagaya (1983), The transport and spatial scales of Asian dust-storm clouds: A case study of the dust-storm event of April 1979, Tellus, Ser. B, 35, Kai, K., Y. Okada, O. Uchino, I. Tabata, H. Nakamura, Y. Takasugi, and Y. Nikaidou (1988), Lidar observation and numerical simulation of a 11 of 12

12 Kosa (Asian Dust) over Tsukuba, Japan during the spring of 1986, J. Meteorol. Soc. Jpn., 66, Kai, K., S. Hu, H. Zhou, T. Yasojima, B. Xu, T. Nagai, and M. Abo (2002), Development of a new lidar for measuring the Aeolian Dust originated from the Taklamakan Desert: Test observation at Tsukuba, Japan, J. Arid Land Stud., 11, Kalma, J. D., J. G. Speight, and R. J. Wasson (1988), Potential wind erosion in Australia: A continental perspective, J. Climatol., 8, Kalnay, E., et al. (1996), The NCEP/NCAR 40-year reanalysis project, Bull. Am. Meteorol. Soc., 77, Kimura, F., and T. Yoshikawa (1988), Numerical simulation of global scale dispersion of radioactive pollutants from the accident at the Chernobyl nuclear power plant, J. Meteorol. Soc. Jpn., 66, Kondo, J., T. Kuwagata, and S. Haginoya (1989), Heat budget analysis of nocturnal cooling and daytime heating in a basin, J. Atmos. Sci., 46, Kurosaki, Y., and M. Mikami (2003), Recent frequent dust events and their relation to surface wind in East Asia, Geophys. Res. Lett., 30(14), 1736, doi: /2003gl Kurosaki, Y., and M. Mikami (2004), Effect of snow cover on threshold wind velocity of dust outbreak, Geophys. Res. Lett., 31, L03106, doi: /2003gl Littmann, T. (1991), Dust storm frequency in Asia: Climatic control and variability, Int. J. Climatol., 11, Louis, J.-F. (1979), A parametric model of vertical eddy fluxes in the atmosphere, Boundary Layer Meteorol., 17, Marticorena, B., G. Bergametti, B. Aumont, Y. Callot, C. N Doume, and M. Legrand (1997), Modeling the atmospheric dust cycle: 2. Simulation of Saharan dust sources, J. Geophys. Res., 102, Mellor, G. L., and T. Yamada (1982), Development of a turbulence closure model for geophysical fluid problems, Rev. Geophys., 20, Mikami, M., et al. (2002), The impact of Aeolian Dust on climate: Sino- Japanese cooperative project ADEC, J. Arid Land Stud., 11, Miller, R. L., and I. Tegen (1998), Climate response to soil dust aerosols, J. Clim., 11, Mitsuta, Y., T. Hayashi, T. Takemi, Y. Hu, J. Wang, and M. Chen (1995), Two severe local storms as observed in the arid area of northwest China, J. Meteorol. Soc. Jpn., 73, Nakajima, T., M. Tsukamoto, Y. Tsushima, A. Numaguti, and T. Kimura (2000), Modeling of the radiative process in an atmospheric general circulation model, Appl. Opt., 39, Okamura, O., and F. Kimura (2003), Behavior of GPS-derived precipitable water vapor in the mountain lee after the passage of a cold front, Geophys. Res. Lett., 30(14), 1746, doi: /2003gl Pielke, R. A., et al. (1992), A comprehensive meteorological modeling system RAMS, Meteorol. Atmos. Phys., 49, Press, W. H., B. P. Flannery, S. A. Teukolsky, and W. T. Vetterling (1986), Numerical Recipes, 818 pp., Cambridge Univ. Press, New York. Qian, W., L. Quan, and S. Shi (2002), Variations of the dust storm in China and its climate control, J. Clim., 15, Sakai, T., T. Shibata, and Y. Iwasaka (1997), Relative humidity, backscattering ratio and depolarization ratio as derived from Raman lidar observations, J. Meteorol. Soc. Jpn., 75, Sassen, K. (2002), Indirect climate forcing over the western US from Asian dust storms, Geophys. Res. Lett., 29(10), 1465, doi: / 2001GL Sassen, K., P. J. DeMott, J. M. Prospero, and M. R. Poellot (2003), Saharan dust storms and indirect aerosol effects on clouds: CRYSTAL-FACE results, Geophys. Res. Lett., 30(12), 1633, doi: /2003gl Sato, T., and F. Kimura (2003), A two-dimensional numerical study on diurnal cycle of mountain lee precipitation, J. Atmos. Sci., 60, Shao, Y. (2001), A model for mineral dust emission, J. Geophys. Res., 106, 20,239 20,254. Shao, Y., and J. Wang (2003), A climatology of northeast Asian dust events, Meteorol. Z., 12, Shao, Y., E. Jung, and L. M. Leslie (2002), Numerical prediction of northeast Asian dust storms using an integrated wind erosion modeling system, J. Geophys. Res., 107(D24), 4814, doi: /2001jd Sun, J., T. Liu, and Z. Lei (2000), Sources of heavy dust fall in Beijing, China on April 16, 1998, Geophys. Res. Lett., 27, Sun, J., M. Zhang, and T. S. Liu (2001), Spatial and temporal characteristics of dust storms in China and its surrounding regions, : Relations to source area and climate, J. Geophys. Res., 106, 10,325 10,333. Takemi, T. (1999), Structure and evolution of a severe squall line overthearidregioninnorthwestchina,mon. Weather Rev., 127, Takemi, T., and N. Seino (2005), Duststorms and cyclone tracks over the arid regions in east Asia in spring, J. Geophys. Res., 110, D18S11, doi: /2004jd Tegen, I., and I. Fung (1994), Modeling of mineral dust in the atmosphere: Sources, transport, and optical thickness, J. Geophys. Res., 99, 22,897 22,914. Tegen, I., and R. Miller (1998), A general circulation model study on the interannual variability of soil dust aerosol, J. Geophys. Res., 103, 25,975 25,995. Tegen, I., A. A. Lacis, and I. Fung (1996), The influence on climate forcing of mineral aerosols from disturbed soils, Nature, 380, Tsunematsu, N. (2005), Observed dust storm in the Taklimakan Desert on April 13, 2002, Sci. Online Lett. Atmos., 1, Uematsu, M., A. Yoshikawa, H. Muraki, K. Arao, and I. Uno (2002), Transport of mineral and anthropogenic aerosols during a Kosa event over East Asia, J. Geophys. Res., 107(D7), 4059, doi: / 2001JD Uno, I., H. Amano, S. Emori, K. Kinoshita, I. Matsui, and N. Sugimoto (2001), Trans-Pacific yellow sand transport observed in April 1998: Numerical simulation, J. Geophys. Res., 106, 18,331 18,344. Walko, R. L., W. R. Cotton, M. P. Meyers, and J. Y. Harrington (1995), New RAMS cloud microphysics parameterization, part I: The singlemoment scheme, Atmos. Res., 38, World Meteorological Organization (1974), Manual on Codes, vol.i, International Codes, WMO Publ., 306, Geneva. Yoshikane, T., F. Kimura, and S. Emori (2001), Numerical study on the Baiu Front genesis by heating contrast between land and ocean, J. Meteorol. Soc. Jpn., 79, K. Kai, Graduate School of Environmental Studies, Nagoya University, Nagoya , Japan. F. Kimura, Graduate School of Life and Environmental Sciences, University of Tsukuba, Tsukuba , Japan. Y. Kurosaki, Center for Environmental Remote Sensing, Chiba University, Chiba , Japan. M. Mikami, Atmospheric Environment and Applied Meteorology Research Department, Meteorological Research Institute, Tsukuba , Japan. T. Nagai, Meteorological Satellite and Observation System Research Department, Meteorological Research Institute, Tsukuba , Japan. T. Sato and N. Tsunematsu, Japan Science and Technology Agency, Kawaguchi , Japan. (n.tsunematsu@arion.ocn.ne.jp) H. Zhou, Xinjiang Institute of Ecology and Geography, Chinese Academy of Sciences, Urumqi , China. 12 of 12