December 1989 K. Tsuboki, Y. Fu jiyoshi and G. Wakahama 985. Doppler Radar Observation of Convergence Band Cloud Formed on

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1 December 1989 K. Tsuboki, Y. Fu jiyoshi and G. Wakahama 985 Doppler Radar Observation of Convergence Band Cloud Formed on the West Coast of Hokkaido Island. II: Cold Frontal Type By Kazuhisa Tsuboki, Yasushi Fujiyoshi and Gorow Wakahama Institute of Low Temperature Science, Hokkaido University, Sapporo 060, Japan (Manuscript received 29 May 1989, in revised form 6 September 1989) Abstract A convergence band cloud moved southwestward through the Ishikari Plain on 20 January, In this study, we mainly used a single Doppler radar to examine the kinematic and radar echo structures of a convergence band cloud of the cold frontal type. The band cloud, approximately 40km in width, shifted southwestward and accompanied an advancing land-breeze front. The front was geometrically and dynamically similar to that of a gravity current found in laboratory experiments. Kelvin-Helmholtz instability waves with wavelength of 1-2km were observed along the interface of the land breeze and the northwesterly monsoon wind. Two small (*6km) vertical circulations similar to those seen in the warm frontal type were embedded in the mesoscale (*15km) vertical circulation. A low-level convergence and a resultant updraft which were stronger than that of the warm frontal type were located ahead of the front. A distinct subsidence, which is not found in the warm frontal type, followed the front. The updraft caused strong convective echoes ahead of the front, while, owing to the subsidence, the echoes behind the front rapidly weakened. The height of maximum updraft coincided with that of the potential instability layer. The updraft and the potential instability may be responsible for the formation of the convergence band cloud. Based on our observations, we constructed a conceptual model of the formation of the convergence band cloud. 1. Introduction During outbreaks of winter monsoon air, a broad single cloud band, known as a convergence band cloud, occasionally forms off shore along the west coast of Hokkaido, Japan. The convergence band cloud is several hundred kilometers in length and several tens of kilometers in width; it is usually maintained for 1-2 days and brings local heavy snowfall at its landing place. When the convergence band cloud develops, the weather is clear on its east side, i. e. over the island of Hokkaido. Based on ground observations and satellite imagery, Okabayashi and Satomi (1971) pointed out that the convergence band cloud is accompanied by a discontinuous line of wind and temperature. Muramatsu et al. (1975) reported that a northeasterly wind with a speed of 4-5m s-1 blows on the east side of a convergence band cloud. Recently Kobayashi of al. (1987) concluded that the formation of a shear line between a monsoon wind and the wind from the island of Hokkaido was coincident with the movement of the southern edge of the convergence band cloud. The convergence band cloud is considered to be formed between the northwesterly monsoon wind and the c 1989, Meteorological Society of Japan wind from the land, i. e. a land breeze. Similar cloud bands which are often referred to as snowbands and a land breeze are frequently observed in the Great Lakes region. Passarelli and Braham (1981) observed shoreline-parallel snowbands and showed that a shallow winter land breeze (only a few hundred meters deep) has an important role in organizing the low-level convergence and in the formation of the snowbands. Ballentine (1982) indicated by numerical simulations that a snowband coincides with a narrow band of upward motion which results from the convergence between a land breeze and an easterly wind over Lake Michigan, and has also shown that latent heat release plays an important role in intensifying the land breeze circulation. Using a mesoscale numerical model, Hjelmfelt and Braham (1983) found that latent heat release plays an important role in strengthening convection, and that the basic mesoscale circulation pattern, including a land breeze, is caused by the lake-land surface temperature difference. Schoenberger (1984) applied a gravity current model to the motion of the land-breeze front and found good agreement between the observed and predicted speeds of advance. Ishihara et al. (1989) examined the characteristics and structures of snowbands in the west-

2 986 Journal of the Meteorological Society of Japan Vol. 67, No. 6 Table 1. Characteristics and parameters of the Doppler radar. ern Hokuriku district of Japan and found that the snowbands were produced in the convergence zone between a land breeze and the northwesterly monsoon wind. The land breeze along the west coast of Hokkaido may be responsible for the formation and maintenance of the convergence band cloud. However, detailed vertical and horizontal structures of a landbreeze front have not yet been determined because both the land breeze and the land-breeze front are very shallow and their time-variation is rapid. Based on an analysis of satellite imagery, Fujiyoshi and Wakahama (1987) classified the appearance process of a convergence band cloud into two types: a retreating or warm frontal type and an advancing or cold frontal type. The former is a type in which the convergence band cloud intensifies and moves toward the land with a small curvature, while the latter type moves away from the land with significant curvature. Fujiyoshi et al. (1988) used a Doppler radar to reveal the radar echo and kinematic structures of the convergence band cloud of the warm frontal type. This type has a simple radar echo and kinematic structures, since the land breeze was at a decaying stage. We used a single Doppler radar to observe a typical case of the convergence band cloud which passed over the radar during the period from 2000JST 20 January to 0200JST 21 January, It was a cold frontal type and was different from the warm frontal type which was studied by Fujiyoshi et al. (1988). The band cloud shifted southwestward and accompanied an advancing land-breeze front. In this paper, we will offer a detailed picture of the radar echo and kinematic structures of the convergence band cloud and of the airflow associated with the land-breeze front. 2. Data and procedures of analysis a. Data and instruments The Doppler radar was designed for the observation of snow clouds. Table 1 summarizes its characteristics and parameters. The location of the radar Fig. 1. Topographical map of the observational area and location of facilities; the solid square indicates the observational point of the Doppler radar, the full circles indicate the AMeDAS data points. was inland *12km from the coastline and the detectable range was 40km. The radar can obtain both the radar reflectivity factor (Z) and the mean Doppler velocity (VR) simultaneously. Its threedimensional scans yield a CAPPI display of the radar reflectivity. We observed the temperature, humidity, pressure, visibility, and wind at the radar site and their one minute averaged values were digitally recorded by a computer. Snowfall intensity was measured by an electrobalance and one minute averaged data were

3 December 1989 K. Tsuboki, Y. Fujiyoshi and G. Wakahama 987 Fig. 2. Surface synoptic analyses for (a) 0900JST on 20 January, 1986, (b) 2100JST on 20 January, 1986, and (c) 0900JST on 21 January, recorded (Konishi et al., 1988). The unit of the visibility is the attenuation rate of light (%) by snow particles. In addition, we also used satellite (GMS-3) imagery, rawinsonde data, and AMeDAS (Automated Meteorological Data Acquisition System) data supplied by the Japan Meteorological Agency. Figure 1 shows the topography of the observational area and the location of the Doppler radar. b. VAD analyses The kinematic properties of a wind field are derived from the observation of its radial velocities as a function of the azimuth at a constant elevation angle (the so-called Velocity Azimuth Display or VAD). Assuming that the horizontal wind field is representable by its first-order Taylor expansion, we can describe the mean Doppler velocity VR(*) fully by the zeroth, and by the first and the second harmonic component of a Fourier series (Browning and Wexler, 1968). VR(*), therefore, may be written as: VR(*)=A1+A2sin*+A3cos* +A4sin2*+A5cos2* (1) where * is the azimuth angle which is measured clockwise from the north. In order to determine the five coefficients of Eq. (1), we adopted a leastsquares-fitting method. The method works correctly even though there are gaps in the distribution of Doppler velocities owing to a lack of precipitation particles. A detailed description of the method is given by Tsuboki and Wakahama (1988). The coefficient of the zeroth harmonic component (A1) gives the horizontal divergence DivVH, A2 and A3 give the horizontal wind speed and wind direction, and A4 and A5 give deformation fields. The vertical component of wind can be estimated from the integration of horizontal divergence of the continuity equation. Under an anelastic approximation, the vertical velocity W(Z) at a height of Z is where W0 is the vertical velocity at a height of Z0, and H is the scale height of the air density taken to be km. We used the VAD analysis to obtain the kinematic structure of the convergence band cloud and the airflow associated with the passage of the land-breeze front. Elevation of the radar antenna used for the VAD analyses was 20*. 3. Synoptic situation and surface observations The surface synoptic weather maps (Fig. 2) indicate that a synoptic scale low passed over the island of Hokkaido and moved eastward. A synoptic scale cold front passed over the Ishikari district at 0900 JST 20 January 1986 (Fig. 2a). At 2100JST (Fig. 2b), a typical Japanese winter pressure pattern had set in : a low was located in the east and a high in the west. A northwesterly monsoon wind began to prevail. On the west coast of Hokkaido, however, the pressure gradient was rather small, and a mesoscale cyclone was generated, with a central surface pressure of 1012mb and a scale of *200km in diameter. The synoptic scale low moved eastward while its central pressure deepened, and was located *600 km to the east of Hokkaido at 0900JST 21 January (Fig. 2c). The pressure gradient on the west coast of Hokkaido was still small. The mesoscale cyclone moved southward and was located over the Ishikari district. Figure 3 shows a time-height cross section of equivalent potential temperature and wind derived from serial rawinsondes launched at Sapporo. After

4 988 Journal of the Meteorological Society of Japan Vol. 67, No. 6 Fig. 3. Time-height cross section of equivalent potential temperature (*e, in K) and wind derived from serial rawinsonde observations made in Sapporo. Winds are plotted according to standard meteorological convention; a half barb indicates 5 knots, a full barb 10 knots, and a flag 50 knots. The shaded areas indicate regions of potential instability. *1200JST on 20 January, a deep cold air advection associated with a northwesterly monsoon wind followed the synoptic followed the synoptic scale cold front. A potential instability layer (which is shaded in the figure) was mainly present from *1km to 2km in altitude during the period only from 0900 * JST to 2400JST on 20 January. The south-easterly wind and a close packing of isentrope below *500 m in altitude during the period 0000JST and 2100 JST 21 January indicate that another shallow cold air mass had penetrated beneath the north-westerly monsoon wind. The cold air mass advected from inland Hokkaido; this is the land breeze. In Wakkanai, similar conditions were observed: cold air advection following the cold front, a mixed layer below *2.5 km in altitude and a deep stable layer above this level, shallow cold advection of an easterly wind under the northwesterly monsoon wind, and a weak potential instability at *1km in altitude at 2100 JST 21 January. Figures 4a and 4b show satellite imagery at 2100 JST on 20 January and 0100JST on 21 January, 1986, respectively. The satellite imagery revealed that the convergence band cloud was located *50 km off shore along the west coast of Hokkaido. The southernmost edge of the convergence band cloud curved and extended to the northeast region of the Ishikari Plain at 2100JST (Fig. 4a). The convergence band cloud then moved southwestward across the Ishikari Plain and was located in the southwest region of the plain at *0100JST on 21 January, 1986 (Fig. 4b). Figure 5 shows the time changes of the surface weather conditions observed at the radar site on Fig. 4. Satellite imagery of the convergence band cloud: (a) 2100JST on 20 January 1986 and (b) 0100 JST on 21 January 1986.

5 December 1989 K. Tsuboki, Y. Fujiyoshi and G. Wakahama 989 Fig. 5. Time changes of the surface weather conditions at the radar site. The passage of the land-breeze front is indicated by the arrow January, The arrow in the figure indicates the passage of the land-breeze front accompanied by a significant wind shift at the surface. The wind direction at the radar site changed clockwise from northwest to southeast and the wind speed decreased. The passage of the convergence band cloud caused a rather strong snowfall at the radar site during the period from 2230 to 2400JST; its maximum peak value was 4mm hr-1 at 2310JST. The minimum peak of visibility and the maximum peak of snowfall rate at 2310JST coincided with the surface wind shift. The air temperature decreased from -6.0* at 2300JST to -7.0* at 0000JST. After 0130JST, the air temperature decreased steadily and reached its local minimum value (-12.4*) at 0400JST. 4. Structure of the land breeze The passage of the convergence band cloud coincided with the appearance of the land breeze in the Ishikari Plain. Doppler radar observation revealed that the land-breeze front passed the radar site at 2315JST on 20 January, In this section we give an account of the kinematic structure of the land breeze and the movement of the advancing Fig. 6. Time series of horizontal wind speed in the vertical cross section whose azimuth angle is 300*; these are derived from the RHV displays: (a) 2305JST, (b) 2315 JST, and (c) 2326JST on 20 January, The shaded areas indicate the land breeae. land-breeze front as revealed by the Doppler radar. a. Kinematic structure of airflow Figure 6 shows vertical cross sections of horizontal wind speed derived from the RHV (Range Height Velocity) scans of the Doppler radar. The azimuth angle was 300*, which is approximately parallel to the direction of the movement of the front at the radar site. If we assume that the radial component of vertical velocity is small enough to ignore when the elevation of the radar is less than 60*, the horizontal wind speed VH in the vertical cross section can be obtained from VH=VR/cos*e, where *e is an elevation angle of the radar. We interpolated the horizontal wind speed above the radar from the wind speeds of both sides. The regions of negative

6 990 Journal of the Meteorological Society of Japan Vol. 67, No. 6 Fig. 7. Time-height cross section of horizontal wind velocities derived from the VAD analysis. Winds are plotted according to standard meteorological convention; a half barb indicates 5 knots and a full barb 10 knots. The thick solid line indicates the depth of the land breeze. speed indicate a land breeze in which the wind blew from the right to the left in the figure. The figure shows that the land breeze advanced under the northwesterly monsoon wind. The advancing speed of the front in the vertical section was approximately 3m s-1, and the direction was northwestward when it passed the radar site. The frontal surface was steep, almost vertical near the front. The vertical features of the land breeze and the vertical wind shear above it varied rapidly with time. In contrast, a land-breeze front of a warm frontal type has a gentle slope of the frontal surface and the time-variation of the shape is small (Fujiyoshi et al., 1988). Like a thunderstorm outflow, a land breeze is also a gravity current. Simpson (1969, 1972, 1982), Simpson and Britter (1979), and Britter and Simpson (1978,1981) have described many of the characteristic properties of gravity currents based on laboratory studies. Charba (1974) has shown a dynamic similarity between gravity currents and thunderstorm outflows. Mueller and Carbone (1987) have pointed out five major features of gravity currents: the head, the nose, the turbulent wake, the body, and the undercurrent. The land breeze had features of a gravity current. The head which is the elevated region of cold air at the land-breeze front, is pronounced in Figs. 6b and 6c. At 2315JST, the height of the head was *500m, which was 150m higher than the region just behind the head. The nose which is the region of cold air overrunning the warm air is notable in Fig. 6b. The height of the body was m. The undulation of the interface of the land breeze and the northwesterly monsoon wind, which is indi- Fig. 8. Time-height cross section of the absolute vertical shear of horizontal wind derived from the VAD analysis. The thick line indicates the depth of the land breeze. cated by a 0m s-1 isotach, suggests the occurrence of Kelvin-Helmhotz instability waves with a wavelength of 1-2km. Figure 7 shows a time-height cross section of the horizontal wind velocities above the radar. At 2316JST, the wind direction started to veer below 300m in altitude and the wind speed to decrease; * this indicated the passage of the land-breeze front. The depth of the land breeze, which was obtained from numbers of RHV scans whose azimuth were 300*, increased with time up to *600m at 0018JST. The wind speed of the land breeze was 2-4m s-1 and its maximum speed was located at the middle level of the land breeze layer. Above *1km in altitude, the wind speeds were almost constant with height. We refer to the layer as a prevailing wind layer. Prior to 2100JST, the wind direction below *4 km in altitude was approximately northwest. When the convergence band cloud passed over the radar site during the period from 2200JST 20 January to 0100JST 21 January, the wind direction in the prevailing wind layer below *1.5km in altitude was north (not shown). In the prevailing wind layer the wind direction backed counterclockwise with height; this coincided with the cold air advection in the lower level. The vertical shear of horizontal wind at the radar site can be obtained from the horizontal wind velocities. The absolute vertical shear is given by where VX and Vy are the X - and Y- components of horizontal wind, and the height difference *Z is taken to be 86m. Figure 8 shows a time-height cross

7 December 1989 K. Tsuboki, Y. Fujiyoshi and G. Wakahama 991 Fig. 9. Time-height cross section of (a) the horizontal divergence obtained from the VAD analysis (the shaded areas indicate the region of divergence), (b) the vertical velocities derived from the intergration of the divergance (the shaded areas indicate the region of downward motion), and (c) the echo intensity derived from RHI scans. The thick line indicates the depth of the land breeze. section of the absolute vertical shear. Prior to 2315 JST, there was no significant shear in the prevailing wind layer. A layer of strong shear was located along the frontal surface of the land breeze after 2315JST. We refer to this layer as a shear layer. The thickness of the shear layer increased with an increase in the depth of the land breeze. The development of shear layer suggests the occurrence of strong mixing: a momentum exchange between the northwesterly monsoon wind and the land breeze. Whereas the land breeze was less than *500m in depth, the shear layer developed up to *1km in altitude behind the front. The shear layer of the land breeze must have a strong influence on the modification of snow clouds. Figures 9a and 9b show respectively the timeheight cross sections of the horizontal divergence and of the vertical velocity. Prior to the passage of the land-breeze front over the radar site, there was strong convergence below *1km in altitude at 2300JST. Consequently, a strong updraft was located ahead of the front; its maximum speed was more than 50cm s-1 and it was located km in altitude. The height of the maximum updraft was three times as high as than the depth of the land breeze. Prior to 2350JST, an updraft was present below *2km in altitude. Downward motion was present up to *2km in altitude after 2350JST. Assuming a steady state of this phenomenon, we can consider that the main upward air motion occurred just ahead of the front while downward air motion followed it. There were some small structures of divergence and of vertical air motion. During the period from 2330 to 2345JST there was devergence below *300 m in altitude corresponding to a weak downdraft of -10cm s-1 from 2330 to 2345JST below *500 * m in altitude; this was followed by a decreasing of the depth of the land breeze layer. There was convergence during the period from 2345 to 2355JST which caused a weak updraft below *500m. The echo intensity (Fig. 9c) was strong just ahead of the land-breeze front, while it weakened rapidly and the echo top defined by a contour of 10dBZ decreased behind front. Figure 10 shows a time-height cross section of wind vectors. The horizontal components are projections of horizontal wind velocities on the azimuth of 300*. The direction is approximately parallel to that of the movement of the front at the radar site. If we assume a steady state of the phenomenon, Fig. 10 depicts the airflow in a vertical cross section. The horizontal distance scale at the base of the figure assumes that the land-breeze front was moving at a constant speed of 3m s-1. There was a mesoscale vertical circulation of which the horizontal scale was 15km: an updraft region *2310JST and a down * draft region at *0010JST. Two small vertical cir-

8 992 Journal of the Meteorological Society of Japan Vol. 67, No. 6 culations whose horizontal scale were *6km were superimposed on the mesoscale vertical circulation; one was present from 2305 to 2340JST and the other from 2340 to 0010JST centered at *300m. Fujiyoshi et al. (1988) also observed small vertical circulations in the warm frontal type. The horizontal scale of these two small vertical circulations is nearly the same as that of the cut-off vortex of an advancing sea breeze (Simpson et al., (1977)). b. Movement of the land-breeze front We used VAD measurements, of which the elevation angle was 1.5*, in order to detect a land-breeze front. A front is defined as a convergence line of radial velocities or an azimuthal shear. Our method of detecting a front is the same as that described by Uyeda and Zrnic (1986). In the case of the landbreeze front observed on 20 January, 1986, three dis- Fig. 10. Time-height cross section of the two dimensional wind vectors: the horizontal and vertical components. The horizontal components are projections of horizontal wind velocities on the azimuth of 300*. The thick line indicates the depth of the land breeze. The distance scale at the base of the figure assumes that the land-breeze front was moving at a speed of 3m s-1. tinct discontinuous lines of Doppler velocity were identified simultaneously on the mountain slope to the north of the Ishikari Plain (Fig. 11). The average gradient of the slope is 0.5* -1*. They were approximately parallel at a distance of *5km from each other and advanced southwestward. We refer to the southernmost one as the leading front, as this was the leading edge of the cold air mass. Until now, a land-breeze front and a sea-breeze front have been identified as a single leading edge of denser air (Simpson, 1977; Nakane and Sasano, 1986). In this case, however, the land breeze had two discontinuous lines of wind behind the leading front. We refer to the former as Discontinuous line A and to the latter as line B. Figure 12 shows the movements of the leading front and of the two discontinuous lines. Prior to 2259JST, the leading front (Fig. 12a) advanced southwestward. At 2311JST, the leading front was distorted by the topography and a current toward the coast was induced. As a result, the leading front passed the radar site from east to west at 2315 JST and the wind direction at the radar site veered clockwise from northwest to southeast over a short period of time. The front of the current toward the sea, which was a part of the leading front, advanced northwestward and the wind direction of the current was southeast. The leading front encountered the mountain range to the southwest of the Ishikari Plain and stagnated there. Discontinuous lines A (Fig. 12b) and B (Fig. 12c) also advanced southwestward. Discontinuous lines A and B disappeared, however, at 2137 and 2151JST, respectively, and we were not able to track them in the Ishikari Plain. Figure 13 shows the displacements of the leading front and of the two discontinuous lines along the XY-line; this line is approximately perpendicular to the leading front in Fig. 12. The ordinate means the horizontal distance measured from the point X in Fig. 12. The advancing speed of the leading front decreased from 7.5m s-1 to 4.0m s-1 at 2134JST after the front reached the plain. The leading front Fig. 11. Time series of three discontinuous lines of wind detected by the Doppler radar at the same time. L.F., A, and B are the leading front, Discontinuous line A and B, respectively.

9 December 1989 K. Tsuboki, Y. Fujiyoshi and G. Wakahama 993 stagnated during the period from 2204 to 2245 JST at *5km northeast of the radar site and then started to advance again at a speed of 4.9m s-1. The front of the seaward current which was a part of the leading front advanced at a speed of 3m s-1 and passed the Doppler radar. Discontinuous lines A and B advanced at a speed of 8.6m s-1 and 8.2 m s-1, respectively; these were *1m s-1 faster than that of the leading front. These two lines did not overtake the leading front, however, but disappeared when they reached the plain. 5. Snowband developed in the convergence band cloud Figure 14 shows a time-height cross section of echo intensity at the radar site. The vertical profiles of the echo intensities above the radar were derived from RHI scans. The arrow in the figure indicates that the land-breeze front passed the radar site at 2315JST. Prior to the passage of the front, the level of echo top defined by 10 dbz increased with time up to *3.5km at its maximum height. There were two peaks of echo top and two intense echo regions (> 24dBZ) at 2215JST and 2300JST. Fujiyoshi et al. (1988) have also reported the presence of two bands with intense radar echo in a convergence band cloud. After the front passed by, the level of echo top and the echo intensity decreased rapidly. The time-series of CAPPI at 1250m in altitude (Fig. 15) shows the horizontal echo structure and the movement of the snowband. The snowband was *40 km in width and extended inland more than 50km from the coast. If we compare the intense echo region in Fig. 15 with the location of the leading front in Fig. 12a, we can see that the band-like intense echo was always located ahead of the land-breeze front. The direction and speed of each cell in the snowband were nearly the same as those of the prevailing wind above 2km in altitude, while the snowband shifted southwestward. The movement of the snowband coincided with the advance of the landbreeze front. When the land-breeze front encountered the mountain to the southwest of the Ishikari Plain and stagnated, the snowband also stagnated and started to weaken. Fig. 12. Movements of the discontinuous lines: (a) the leading front, (b) Discontinuous line A, and (c) Discontinuous line B.

10 994 Journal of the Meteorological Society of Japan Vol. 67, No. 6 Fig. 13. Displacements of the leading front (full circles), Discontinuous line A (full triangles), and Discontinuous line B (full square) along the XY-line in Fig. 12. The horizontal distance is measure from the point X in Fig. 12. Fig. 14. Time-height cross section of the echo intensity (dbz) above the radar, which were derived froth RHI scans. Figure 16 shows a series of vertical cross sections of echo intensity perpendicular to the snowband. The position of the vertical cross sections are indicated in the CAPPI displays (Fig. 15). An arrow at the bottom of the figure indicates the position of the land-breeze front. The front moved from right to left in the figure. Intense convective cells always developed ahead of the front. In Fig. 16b, a convective cell ahead of the front can be seen developing while a cell behind the front is decaying. The level of echo top was high ahead of the front and was low behind the front. The convective activity of the cloud was vigorous ahead of the front, but lost its strength behind the front. The snowband advanced southwestward, while the convective cells which composed the snowband Fig. 15. Time series of the CAPPI displays of the snowband developed along the convergence band cloud at a height of 1250m at (A) 2226JST, (B) 2256JST, and (C) 2330JST on 20 January The radar echoes are contoured at 10dBZ, the lightly shaded areas indicate 19-24dBZ and the black areas more than 24dBZ. moved southeastward. In order to measure the advancing speed of the snowband, we made a timerange cross section (Fig. 17) whose azimuth angle was parallel to the XY-line in Fig. 12, i.e. approximately parallel to the direction of advance of the front and approximately perpendicular to the snowband. During the period from 2100 to 2200JST, the snowband shifted at a constant speed of 7.2m s-1. The speed was approximately equal to the advancing speed of the leading front on the slope (7.5m s-1).

11 December 1989 K. Tsuboki, Y. Fujiyoshi and G. Wakahama 995 Fig. 15. (continued) During the period from 2230 to 2300JST, the speed of the snowband was 5.3m s-1; this was close to that of the leading front on the flat plain (4.0*4.9 ms-1). We can conclude that the movement of the snowband was dominated by the movement of the front. 6. Discussion Sakakibara et al. (1988a) classified the mesoscale snowfall systems observed in western Hokuriku into eight types. They found a type of vortex of convective coluds and two types of band-shaped clouds; one is parallel to the winds and the other is normal to the winds. As we can see in Fig. 4, the convergence band cloud of the present paper had characteristic features both of the band-shaped clouds and of the vortex. It was particularly broad and intense in comparison with other band-shaped clouds (Fig. 4a). Sakakibara et al. (1988b) used a Doppler radar to examine two convective snowbands in western Hokuriku and found that their structure was similar to those of tropical and midlatitude squall lines. They observed gust fronts during the passage of the snowbands and concluded that the gust fronts are important to the maintenance of the snowbands. In contrast, the convergence band cloud had no remarkable gust front. We infer that the landbreeze front was important to maintenance of the snowband of the present study. It caused a horizontal divergence and lifted up the warm air of the northwesterly monsoon wind. Ishihara et al. (1989) observed mesoscale snowbands which developed between the northwesterly monsoon wind and a land breeze in the western coast of the Hokuriku district. They resembled the convergence band cloud in the formation and maintenance mechanisms, but they Fig. 16. Time series of vertical cross sections of echo intensity, which are perpendicular to the snowband at (A) 2226JST, (B) 2256 JST, and (C) 2330JST on 20 January, The positions of the cross section are shown in Fig. 15. were multi-bands which formed over the sea and moved toward the coast. When the snowband moves toward the coast, the land-breeze front also retreats toward the cosat by the pumping of the land breeze air into the snowband. On the other hand, the convergence band cloud was a single band and its movement was dominated by the movement of the landbreeze front. The southwestward movement of the convergence band cloud coincided with the advance-

12 996 Journal of the Meteorological Society of Japan Vol. 67, No. 6 Fig. 18. Schematic of conceptual moaei of formation of the convergence band cloud. Fig. 17. Time-range cross section of the echo intensity, whose azimuth is 25*. ment of the land-breeze front. As Fujiyoshi et al. have reported, a land-breeze front has characteristic properties which are similar to a gravity current in a laboratory tank apparatus. The front which they studied was a warm frontal type. The head of the land breeze was therefore poorly defined and the slope of the frontal surface was not so steep. In contrast, the land-breeze front in this study was a cold frontal type. The head and the nose were clearly identified and the slope of the frontal surface was steep, especially near the front. The updraft associated with the cold frontal type is therefore larger than that of the warm frontal type. The shape of the frontal surface was highly variable with time in comparison with that of warm frontal type. The structural features of this type were similar to that of a thunderstorm outflow. The horizontal wind profile of the land breeze was similar to the velocity profile behind a gravity current head measured in a laboratory model (Simpson and Britter, 1979): a jet-shaped pattern in the land breeze layer, the presence of a shear layer and a constant wind speed in the prevailing wind layer. The laboratory experiments (Simpson, 1969, 1972, 1982; Britter and Simpson, 1978) show that the mixing of dense fluid and ambient fluid is caused by Kelvin-Helmholtz billows generated at the front of the head. Using a two-dimensional numerical model, Droegemeier and Wilhelmson (1987) studied thunderstorm outflow and found that turbulent mixing is associated with breaking Kelvin-Helmholtz billows which form at the shear interface between the two fluids. Simpson et al. (1977) suggested that the mixing is due to Kelvin-Helmholtz instability on the sea breeze fronts. Fujiyoshi et al. (1988) and Tsuboki et al. (1989) have reported that Kelvin- Helmholtz instability waves occure along a landbreeze frontal surface and that the wavelength is *1 km. In the case of a cold frontal type, the RHV display showed the occurrence of Kelvin-Helmholtz instability waves whose wavelength was 1-2km. The densities of the northwesterly monsoon wind and the land breeze at a height of 300m were 1.31 and 1.30kg m-3, respectively. The difference of wind speed was 5-7m s-1 between the land breeze and the northwesterly monsoon wind. The critical wavelength of Kelvin-Helmholtz instability waves is thus 1-2km, which is within the range of the observed scale. Kelvin-Helmholtz instability waves play an important role in the mixing of the two air masses, namely in the exchange of momentum, heat and moisture. The shear layer is generated by the momentum exchange, and a thermodynamical stable layer is formed due to the mixing of heat and moisture. The propagation velocity Vp of a gravity current is given by (Benjamin, 1968) VP=(2gH*/*L)1/2 (4) where H is the depth of gravity current, g is the acceleration of gravity, * is the potential temperature difference between two fluids, and *L is the potential temperature of gravity current. We estimated the potential temperatures of the northwesterly monsoon wind and of the land breeze and the potential temperature difference at the height of 300m from rawinsonde data. Substitution of values H=300m, *=3 K and *L=261 K into Eq. (4) yields Vp =8m s-1. The Doppler radar analysis shows that the leading front of the land breeze moved southwestward at a speed of 7.5m s-1 before 2134JST on 20 January (Fig. 13); this agrees well with the theoretical value given by Eq. (4); it also

13 December 1989 K. Tsuboki, Y. Fujiyoshi and G. Wakahama 997 Fig. 19. Distribution of the surface wind obtained from the AMeDAS data. Winds are plotted according to standard meteorological convention; a half barb indicates 5 knots and a full barb 10 knots. suggests that the hypothesis of gravity current can account for the movement of the land-breeze front. After 2134JST the advancing speed of the front decreased owing to the topography. Based on our observations, we constructed a conceptual model for the formation of the convergence band cloud (Fig. 18). When a land breeze flows out from the land to the sea, an updraft that results from the low-level convergence occurs ahead of the land-breeze front and the potential instability layer just ahead of the front is lifted up. If the layer is humid and the height difference between the layer and its lifting condensation level is small, potential instability is actualized and convection occurs along the front. In fact, the Doppler radar observation and sounding analysis revealed that the updraft coincided with a layer of potential instability and that the height difference between the layer and its lifting condensation level was m. The subsidence, which was revealed by the VAD analysis, followed the land-breeze front. Satellite imagery shows that the weather was clear on the east side of the convergence band cloud; this would be due to the subsidence behind the front. No distinct subsidence has been observed in the warm frontal type. Sasaki and Deguchi (1988) demonstrated by numerical simulation a circulation pattern which is similar to that shown in Fig. 10: a shallow land breeze which flows out from the land to the sea under a northwesterly monsoon wind, a strong updraft which results from a low-level convergence at the land-breeze front, and a subsidence behind the front. They point out that the sea-land surface temperature difference is important for the formation of the low-level convergence. Figure 19 shows the distribution of surface winds measured by the AMeDAS at 0100JST on 21 January, The wind at the west coast was approximately east. At the southernmost edge of the convergence band cloud, convergence of the surface wind was present: northeasterly on the east side and northwesterly on the west side in the Ishikari Plain. In the northernmost part of Hokkaido, the surface winds during the period of our observation were approximately easterly-northeasterly along both coasts; this indicates that the winds in that region were blowing from the Okhotsk sea. The sea surface temperature of the Okhotsk sea is always colder than that of the Japan sea. In fact, the Okhotsk sea was covered with sea ice on 20 January, We suspect, therefore, that the easterlynortheasterly wind was also a shallow cold advection. We infer that the land breeze and the easterlynortheasterly winds are responsible for formation and maintenance of the convergence band cloud. In contrast, the land-breeze front of the warm frontal type retreated, although there was still an air temperature difference between the land breeze and the northwesterly monsoon wind. The reason would be that both the easterly-northeasterly wind and the subsidence behind the front were no longer strong to supply enough cold air to the land breeze layer. 7. Summary Satellite imagery showed that a convergence band cloud passed over the Ishikari Plain on the night January, Doppler radar observation revealed that an intense snowband developed in the convergence band cloud and that an advancing landbreeze front coincided with the snowband. In this paper we have described the kinematic and radar echo structures of the convergence band cloud which was a cold frontal type, and the airflow associated with the advancing land-breeze front. This type of the convergence band cloud was found when the land breeze was developing and had different structures from that of a warm frontal type. The RHV displays showed that the advancing land-breeze front has characteristic features similar to a gravity current found in laboratory experiments: the head, the nose and the turbulent wake. The depth of body was m. The advancing speed of the front was *3m s-1 and the direction was northwestward when it passed the radar site. The shape of the frontal surface varied considerably with time in comparison with that of a warm frontal type. Doppler radar observation shows the occurrence of Kelvin-Helmholtz instability waves along the in-

14 998 Journal of the Meteorological Society of Japan Vol. 67, No. 6 terface between the land breeze and the northwesterly monsoon wind, which have the same wavelength, 1-2km, as those observed in the warm frontal type. Kelvin-Helmholtz instability waves played an important role in the mixing of the northwesterly monsoon wind and the land breeze. Profiles of horizontal wind derived from the VAD analysis were divided into three layers; the land breeze layer, the shear layer and the prevailing wind layer. The shear layer developed along the frontal surface of the land breeze. The updraft which resulted from the low-level convergence was present ahead of the front and subsidence was present behind the front. We observed two small (*6km) vertical circulations which were embedded in the mesoscale (*15km) vertical circulation. Although distinct subsidence did not occur in the warm frontal type, small vertical circulations were also present in the warm frontal type. Rawinsonde data revealed that a stronger updraft than that present in the warm frontal type coincided with the potential instability layer; this might be responsible for the formation of the convergence band cloud. Owing to the subsidence, the weather was clear on the east side of the convergence band cloud. Three discontinuous lines of wind, the leading front, Discontinuous lines A and B, were identified at the same time by the VAD display of a 1.5* elevation angle. The leading front, Discontinuous lines A and B advanced southwestward at a speed of 7.5, 8.6 and 8.2m s-1, respectively. The speed of the leading front decreased to 4.0m s-1 after reaching the plain. Owing to the topography, a seaward current was induced. The front of the current advanced northwestward at a speed of 3m s-1 in the coastal region. At the radar site, the surface wind veered clockwise from northwest to southeast when the land-breeze front passed by. Part of the leading front stagnated against the mountains to the southwest of the Ishikari Plain. We examined the radar echo structure of the snowband which developed in the convergence band cloud. The snowband was *40km in width and reached inland more than 50km from the coast. The snowband shifted southwestward and its displacement coincided with the advance of the land-breeze front. In the cold frontal type as well as in the warm frontal type, strong convective cells developed ahead of the front owing to the low-level convergence. The echo behind the front rapidly weakened owing to the subsidence. The movement of the snowband was dominated by the movement of the land-breeze front. Based on our observations, we constructed a conceptual model of the formation of the convergence band cloud. When a land breeze blows, the lowlevel convergence between the land breeze and the northwesterly monsoon wind occurs and the potential instability layer just ahead of the front is lifted up. Potential instability is then actualized, and the convergence band cloud develops along the front. Because there is subsidence behind the front, the weather behind the front is clear. Acknowledgment The authors would like to thank the staff of the Japanese Meteorological Agency, Sapporo, for supplying us with the rawinsonde data. We also thank Dr. T. Yamada for his support of the AMeDAS data collection. References Ballentine, R.J., 1982: Numerical simulation of landbreeze-induced snowbands along the western shore of Lake Michigan. Mon. Wea. Rev., 110, Benjamin, T.B., 1968: Gravity currents and related phenomena. J. Fluid Mech., 31, Britter, R.E. and J.E. Simpson,1978: Experiment on the dynamics of a gravity current head. J. Fluid Mech., 88, Britter, R.E. and J.E. Simpson,1981: A note on the structure of the head of an intrusive gravity current. J. Fluid Mech., 112, Browning, K.A. and R. Wexler,1968: The determination of kinematic properties of a wind field using Doppler radar. J. Appl. Meteor., 7, Charba, J., 1974: Application of gravity current model to analysis of squall-line gust front. Mon. Wea. Rev., 102, Droegemeier, K.K. and R.B. Wilhelmson, 1987: Numerical simulation of thunderstorm outflow dynamics. part 1: Outflow sensitivity experiments and turbulence dynamics. J. Atmos. Sci., 44, Fujiyoshi, Y. and G. Wakahama, 1987: Classification of appearance process of a convergence band cloud formed on the west coast of Hokkaido Island as revealed by GMS imagery. Low Temperature Science, 46, (in Japanese). Fujiyoshi, Y., K. Tsuboki, H. Konishi and G. Wakahama, 1988: Doppler radar observation of convergence band cloud formed on the west coast of Hokkaido Island (I): warm frontal type. Tenki, 35, (in Japanese). Hjelmfelt, M.R. and R.R. Braham, Jr., 1983: Numerical simulation of the airflow over Lake Michigan for a major lake-effect snow event. Mon. Wea. Rev., 111, Ishihara, M., H. Sakakibara and Z. Yanagisawa 1989: Doppler radar analysis of the structure of mesoscale snow bands developed between the winter monsoon and the land breeze. J. Meteor. Soc. Japan (to be published). Kobayashi, F., K. Kikuchi and T. Motoki,1987: Studies on the convergence band cloud formed in the midwinter seasons on the west coast of Hokkaido Island, Japan (I). Geophys. Bull. Hokkaido Univ., 49, (in Japanese with English summary). Konishi, H., T. Endoh and G. Wakahama, 1988 : A new snow gauge using an electric balance. Seppyo, 50, 3-7 (in Japanese with English abstract).

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