August 1990 H. Kondo 435. A Numerical Experiment on the Interaction between Sea Breeze and

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1 August 1990 H. Kondo 435 A Numerical Experiment on the Interaction between Sea Breeze and Valley Wind to Generate the so-called "Extended Sea Breeze" By Hiroaki Kondo National Research Institute for Pollution and Resources, MITI, 16-3, Onoga* wa, Tsukuba 305, Japan (Manuscript received 16 January 1990, in revised form 6 June 1990) Abstract A two-dimensional numerical experiment with sea, plain, slope and plateau and a three-dimensional numerical experiment with sea, plain, slope, valley and plateau were executed to investigate the interaction between the sea breeze and the thermally and topographically induced flow. Generation of the so-called "extended sea breeze (ESB)" was also investigated. In the two-dimensional experiment with sea, plain, slope and plateau, a weak wind towards inland appeared between the sea breeze front and inland slope over the plain whose length was km. The topography was changed to a valley in the three-dimensional model. The mouth and end of the valley were located 20 km and km from the coastline, respectively. The addition of this valley caused the surface wind in the area between the sea breeze front and the end of the valley to become stronger, and showed the long-range penetration of the ESB. The depth of the wind system in the valley was equal to or greater than the depth of the mixing layer in the valley. The additional experiments in which the width of the valley was changed showed that the wind in the valley was strongest in the narrowest valley. The mechanism was basically consistent with the discussion by Defant (1951). In the case that the distance between the coastline and inland topography was short, the sea breeze was enhanced in the morning; however, the same topography worked to stop the advance of the sea breeze in the late afternoon. The non-linear effect was important to the development of the combined sea breeze and valley wind and the generation of a unified circulation (i. e. the ESB). 1. Introduction The local wind with diurnal variation whose horizontal and vertical scales are much greater than the usual sea breeze is often found over the Kanto Plain in Japan. The direction of the wind, which is very complicated in the morning according to the complicated coastline and orography, is quickly unified into a southern wind in the early afternoon throughout the Kanto Plain (Kawarnura, 1977; Fujibe and Asai, 1979). The thickness of the wind layer reaches km (Tsuruta, 1986; Yoshikado and Kondo, 1989). This large-scale local wind with diurnal variation has been often called the Extended Sea Breeze (ESB) over the Kanto Plain. Kondo (1990; hereafter abbreviated as KO) showed two possible mechanisms of this phenomenon; one occurs without a general wind, the other occurs under the condition of a SW general wind. KO concluded that ESB-like wind, which occurred with a SW general wind, had no relation with "sea" breeze. In the case without a general wind, both the sea breeze and the thermally and topographically induced flow are responsible for 1990, Meteorological Society of Japan * the generation of the ESB. So far, most of the scientists who have studied the ESB over the Kanto Plain have supposed that the ESB was a kind of combined phenomena of the sea breeze and the thermally and topographically induced flow such as up-slope wind (Kikuchi et al., 1981) or that the ESB was the flow blowing into the thermal-low developed over the central mountainous area in Japan (Kurita et al., 1988). The interaction of sea breeze and up-slope wind of the mountain was first investigated by Mahrer and Pielke (1977) in a numerical model. Asai and Mitsumoto (1978) calculated the interaction of sea breeze and up-slope wind over two-dimensional sea, plain, slope and plateau. Asai and Mitsumoto considered the topography around the Inland Sea of the Seto region, where the length of the coastal plain was only 15 km. The results showed that the scale of the sea breeze became a little greater than a pure sea breeze and that the sea breeze was a little enhanced. Ookouchi et al. (1978) obtained similar results. Because these calculations were twodimensional and the introduced topography was too close to the coastline, these models were inadequate to investigate the mechanism of the ESB over the

2 436 Journal of the Meteorological Society of Japan Vol. 68, No. 4 Kanto Plain. The two-dimensional sea breeze shown in KO was able to advance only 50 km before sunset. Observational results in the Kanto Plain showed that the wind further inland had already become strong before sunset. Also the surface wind over the plain described in Mannouji (1982), who investigated the local wind over the plain, slope and plateau, was not as strong as the observation in the south of the Kanto Plain. These differences suggest that the observed ESB over the Kanto Plain is not explained only by the sea breeze phenomena or only by the plain-to-plateau wind. The results of KO have suggested that the three-dimensional valley topography was effective in forming the ESB. In the present paper, the combined phenomena of the sea breeze and the thermally and topographically induced wind are investigated and the mechanism of the generation of the ESB is discussed. 2. Interaction between the sea breeze and the thermally and topographically induced flow Using the same mesoscale model as KO, the results of five numerical experiments were investigated first. Five experiments were executed over the topography of (1) two-dimensional land and sea (RUN 1), (2) two-dimensional plain-slopeplateau (RUN 2), (3) two-dimensional sea-plainslope-plateau (RUN 3), (4) three-dimensional Plain and valley (RUN 4) and (5) three-dimensional seaplain-valley (RUN 5), respectively. The configuration and scales of RUN 1 to RUN 5 are shown in Fig. la-le. In all the RUNS, *x = (*y =) 6.67 km, and 60 grid points were used for the vertical (z) profile. In the horizontal direction, 60 grid points were used for longitudinal (x) direction for both the two- and three-dimensional model, and 30 grid points were used for the transverse (y) direction of the valley for the three-dimensional cases. The height of the plain both in side and outside of the valley was the same as that of the sea surface. The top boundary was set at 5900 m (Fig. 1). The lateral boundaries were all closed and insulated in all the twodimensional cases. In the three-dimensional cases, calculations were only made for one side of the valley, so that *u/*y = 0 was assumed on the lateral boundary through the center of the valley. This topography is similar in scale to that of the Pacific Ocean an* the inland mountainous are along P1, P2 and P3 in the Kanto Plain shown in KO. The Coriolis force was neglected except for RUN 42, which will be discussed in the last part of the present paper. Other surface parameters and initial conditions were set the same as for the three-dimensional model in KO. Figure 2 shows the cross sections of the wind component of the five RUNS across the coastline at16:00. Fig. 1. Topography and its scale used in the present experiment. (a) RUN 1, (b) RUN 2, (c) RUN 3, (d) RUN 4 and (e) RUN 5. D in (e) is the sampling point for the potential temperature profile (see Fig. 11). For the three-dimensional models, the cross section is shown through the center of the valley. Figure 2a, the sea breeze case, shows that the sea breeze front was located approximately 33 km from the coastline at 16:00. The wind further inland was very weak. Figure 2b, in which the sea surface was replaced with the land surface and the slope and plateau were added, shows a strong wind that was similar to the plain-to-plateau wind shown by Mannouji (1982). Figure 2c, in which the sea surface was added to RUN 2, shows that there were two strong wind regions. This pattern may appear to be a superposition of Fig. 2a and 2b. The inland-ward wind was generated over the plain between the sea breeze front and the slope wind. Since the direction of the wind extending from the coast to the plateau was heading inland, the wind that appeared over the plain may be identified as the ESB, if we define the ESB as a local wind the direction of which has a diurnal vari-

3 August 1990 H. Kondo 437 Fig. 2. Distributions of vector wind on the x - z plane in (a) RUN 1, (b) RUN 2 and (c) RUN 3, and on the x - z plane through the center of the valley in (d) RUN 4 and (e) RUN 5 at 16:00.* indicate the coastline. The indicated distance is that from the side boundary. ation similar to sea breeze, and the scale of which is larger than that of the sea breeze. The thickness of the inland directed wind between the sea breeze front and the slope wind was more than 1000 m, although the velocity was much lower than the observations of the ESB in the Kanto Plain. Figure 2d, in which a three-dimensional valley was introduced without the sea surface, shows that the wind over the plain in the valley was stronger than shown in Fig. 2b. The wind between the sea breeze front and the slope flow was further increased (Fig. 2e) when the sea surface was added. The sea breeze front in Fig. 2e was not as clear as that seen in Fig. 2c. The results of Fig. 2 suggest the following: in the formation of the ESB, it is important that the wind in the region between sea breeze front and the inland up-slope or valley wind be further enhanced. From the observational results over the Kanto Plain, this enhancement had already occurred by the early afternoon (Kawamura, 1977; Tsuruta, 1986; Yoshikado and Kondo, 1989). Some comparisons of the velocity of the wind perpendicular to the coastline over the plain or in the valley among the five RUNS are given in Fig. 3. Figure 3a shows the distribution of the ratio,* on the x-z plane at 16:00, where ul is the horizontal wind velocity component perpendicular to the coastline in RUN 1, and u2 is that in RUN 2. u3 is that in RUN 3. The near surface wind near the inland side of the sea breeze front in RUN 3 was intensified by 40 %, in comparison with the simple superposition of the sea breeze (RUN 1) and up-slope wind (RUN 2). Figure 3b is similar to Fig. 3a except that

4 438 Journal of the Meteorological Society of Japan Vol. 68, No. 4 Fig. 2. (Continued) u2 is result of RUN 4 and that u3 is that of RUN 5. The near surface wind near the inland side of the front was intensified by 50%. Figure 3c is similar to Fig. 3a except for the ratio of the wind velocity of RUN 5 (three-dimensional) to that of RUN 3 (twodimensional). On the inland side of the sea breeze front, the wind was intensified by % up to a 1000m elevation. Also, the depth of the wind system had become greater near the slope region. The three-dimensional valley structure enhances the wind in the valley more than the enhancement over the plain in the two-dimensional case. Figure 4 indicates the relative depression of the surface pressure over the plain or in the valley in five RUNS. The depression of RUN 21, in which the case with narrow width of the valley was calculated (see 3.2) is added in this figure for the sake of later discussion. The pressure at the coastline in each RUN was taken as the reference value. In the figure, 2.8x 10-4 for it (Exner function) was assumed to be 1hPa. In the case for no sea breeze front (RUN 2, 4), the total pressure depression was small and the pressure was linear throughout the plain. However, in the case with the sea breeze, the depression was great near the coastline and the horizontal pressure gradient varied across the frontal zone. The total depression was much greater than that in RUN 2 and 4. The pressure difference which was calculated at a distance of 33.3km is listed in Table 1 for five Runs. Each difference was calculated between the coastline and 33.3km inland and between 80km and 113.3km inland, respectively. For RUN 5, in which the frontal structure was not so clear, the difference of the pressure gradient across the frontal zone was smaller than that in RUN 1 and RUN 3. This is consistent with the result obtained by Linden and Simpson (1986) that the non-linear density gradient (here pressure gradient) undergoes frontogenesis and that the large difference of the gradient causes frontogenesis to occur rapidly. 3. The effect of the scale of a valley The results of KO showed that the ESB was seen not only over the Kanto Plain. The ESB was found to occur similarly over the Noubi Plain and Nasunohara under the condition with no general wind. There are common topographical features of these three regions. The north and west sides of the Kanto Plain are surround by high mountains. The Noubi Plain is also surrounded by high mountains except in the south. Nasunohara is a valley with mountains along its west and east sides. These facts suggest that a three-dimensional valley structure is important for the formation of the ESB. The result that the wind velocity over the plain between the sea breeze front and the slope of RUN 3 was much lower than that in the valley of RUN 5 also demonstrates that a three-dimensional effect is important for the formation of the ESB over the Kanto Plain. The effect of the scale of the three-dimensional valley, such as height of the inland plateau, the width and the length of the valley on the strength of the local wind

5 August 1990 H. Kondo 439 Fig. 4. The relative surface pressure on the plain or on the center of the valley in RUN 1, 2, 3, 4, 5 and 21. The indicated distance is that from the coastline. Fig. 3. Distribution of the relative intensity of horizontal wind component perpendicular to the coastline on x - z plane at 16:00. (a) the ratio of RUN 3 to RUN 1 +RUN 2, (b) the ratio of RUN 5 to RUN 1 +RUN 4 and (c) the ratio of RUN 5 to RUN 3.* indicates the coastline and the indicated distance is that from the coastline. Table 1. The pressure difference between the coastline and 33 km inland and between 80 km and 113 km inland for five RUNS (hpa). system which appears over the plain or in the valley is investigated in this section. The strength of the local wind systems seems to be a very important factor in the generation of the ESB. 3.1 Height of the plateau The effects of the difference in height of the plateau was investigated using the same model as that in RUN 5 except for the different plateau heights. The height was set at 500 m and 1000 m (RUN 11, 12), respectively, while the height in RUN 5 was 1400 m. The wind distributions of RUN 11 and 12 at 16:00, similar to the simulation shown in Fig. 2, are shown in Fig. 5. The plain-to-plateau wind pointed out by Mannouji (1982) appeared in both cases. The wind was weaker in RUN 11 in comparison with RUN 5, and 12. The wind velocity in the valley was slightly stronger in RUN 12 than that in RUN 5, while the velocity for RUN 11 was half of that for RUN 5. The depth of the inlandward wind was not horizontally uniform along the valley in all the RUNS. In RUN 5, the wind became thicker towards inland and the maximum was 1700 m which was attained near the end of the valley. The wind layer in RUN 11 became thinner towards inland, contrary to the result of RUN 5. The maximum thickness was attained near the inland side of the sea breeze front, i. e. at 1200 m. The thickness of RUN 12 was nearly uniform but became a little thicker over the end of the valley and the maximum thickness was 1400 m. The two-dimensional result (RUN 3; Fig. 2c) showed that the thickness did not depend on the distance from the coastline. The difference of the dependency of the thickness on the distance from the coastline between the results of RUN 5 and 11 was due to the vertical motion of the wind in the valley. The inland-ward wind in the valley was slightly ascending in RUN 5 and 12, while the wind was descending near the end of the valley in RUN 11. The descending motion may imply a rather weak connection between the sea breeze and the valley wind in RUN 11. Time variations of the potential temperature profiles (Fig. 6) and a time-height cross section of the horizontal wind (Fig. 7) at 80 km inland from the coastline in the center of the valley are shown for

6 440 Journal of the Meteorological Society of Japan Vol. 68, No. 4 Fig. 5. Distributions of vector wind on x - z plane through the center of the valley in (a) RUN 11 and (b) in RUN 12. * and indicate the coastline and the point 80 km inland from the coastline. RUN 5, RUN 11 and 12. From the potential temperature profiles (Fig. 6a, 6b, 6c) the layer between 1000 m and 1500 m was warmer in RUN 11 than in RUN 5 in the afternoon, which was the result of the adiabatic heating through the descending motion there. The mixing layer depth was estimated from the figure as the layer in which the vertical gradient of the potential temperature had a negative value. The broken line shown in Fig. 7 is the mixing layer depth thus estimated. The inland-ward wind started to develop after 10:00 in RUN 11 and 9:00 in RUN 5. The wind became stronger after 12:00 in RUN 5 and after 15:00 in RUN 11. The thickness of the inland-ward wind coincided well with the depth of the mixing layer in RUN 11. However the wind layer was thicker than the mixing layer depth both in RUN 5 and 12. This suggests that the thickness of the wind is equal to the depth of the mixing layer when the height of the plateau is lower than the mixing layer, while the thickness of the wind layer becomes greater than the depth of mixing layer when the height of the plateau is higher than the depth of the mixing layer. However, even for the case with a high plateau, Fig. 5, Fig. 7b and 7c indicate that the inland-ward wind was strong in the mixing layer. 2 Width o f the valley 3. The length of the valley and height of the plateau were fixed as the same as those in RUN 5, and the half width of the valley at the valley mouth was changed to 53.3 km (RUN 21) and km (RUN 22), while that in RUN 5 was 93.3 km. The results show the velocity of inland-ward wind changed little over the plateau (plain-to-plateau wind) and *ear the sea side of the sea breeze front but it changed more in the valley. There was little difference among RUNS 5, 21 and 22 in depth of the wind system in the valley. Figure 8 shows the ratio of horizontal wind component along the center line of the valley perpendicular to the coastline within the valley in RUN 21 and 22 to the component in RUN 5. The wind in RUN 21 was 50-60% stronger than that in RUN 5 in the region between sea breeze front and the end of the valley. On the other hand, the wind velocity there in RUN 22 was 20-30% less than in RUN 5. In RUN 3, the wind velocity was 40% less than that in RUN 5 (Fig. 3c). The opening angle of the valley toward the sea was 57, 90, 119 and 180 degrees for RUN 21, 5, 22 and 3, respectively. It may be concluded that a narrower valley with the same along valley length causes the wind in the valley to be stronger. Figure 9 shows the potential temperature profiles in RUN 5 and 21 at 80 km inland from the coastline on the center line of the valley at 16:00. The potential temperature of the layer between m in RUN 21 was higher than that of RUN 5. The reason of this difference will be discussed in Section

7 August 1990 H. Kondo 441 Fig. 6. Potential temperature profiles and their time variation at 80 km from the coastline (a) in RUN 5, (b) in RUN 11 and (c) in RUN The mixing layer depth in RUN 21 was less than 1 km throughout the day. 3.3 Horizontal scale of the valley The topography was similar to that in RUN 5 except that the length of the valley was reduced to 46.7 km (RUN 31) The angle of the valley was the same as RUN 5, so that the half width of the valley at the mouth of the valley was reduced to 53.3 km. The time-height cross section of the horizontal wind at km and 66.7 km inland from the coastline in RUN 5 and 31, respectively, (these locations were the same distance (6.7 km) from the end of the val- Fig. 7. The time-height cross section of the horizontal wind at 80 km inland from the coastline. (a) RUN 5, (b) RUN 11 and (c) RUN 12.

8 442 Journal of the Meteorological Society of Japan Vol. 68, No. 4 Fig. 8. Distributions of the ratio of the relative intensity of the horizontal wind component perpendicular to the coastline on x - z plane at 16:00. The ratio of RUN 21 to RUN 5 (top) and that of RUN 22 to RUN 5 (bottom). * indicates the coastline and the indicated distance is that from the coastline. ley) indicates that the thickness of the inland-ward wind was 1400m at 13:00 in RUN 31 and it was reduced after 15:00; although, the maximum thickness was 1750 m in RUN 5 and was attained at 15:00. The maximum wind velocity in RUN 31 was 5m/s at 18:00, while the maximum in RUN 5 was 6m/s at 20:00. The maximum velocity appeared earlier and was weaker in the smaller scale valley. Figure 10 shows the horizontal wind component near the surface at 20:00 in RUN 5 and 31. Although in both figures in Fig. 10 there was an inland-ward wind, the front of the plain (sea)-to-plateau flow along the center of the valley advanced 50km more inland in RUN 5 than in RUN 31. That is, the valley with the larger horizontal scale had a larger horizontal scale wind system extending from the coastal area to an area inland in this scale range. However, if the valley and coastline are too far apart, the sea breeze cannot reach the valley. There should be an optimum scale of the coastal plain and the valley to cause the maximum horizontal extent of the wind system. On the coastal plain outside of the valley, where there was a slope near the coastline, the wind became weaker at 20:00 in both RUN 5 and 31. Although the sea breeze became the strongest at 20:00 in the two-dimensional sea breeze run (RUN 1), the slope at 20 km inland from the coastline blocked the sea breeze by this time. The wind at the foot of the side slope in the valley turned its direction towards the center of the valley at 20:00, indicating there was no direct forcing for the wind to go over the slope at this time. These results demonstrate that Fig. 9. Potential temperature profiles of RUN 5 and 21 at 80 km inland from the coastline at 16:00. the topography located near the coastline works to stop the advance of the sea breeze in the late afternoon even though the same topography enhances the sea breeze in the morning till early afternoon, as pointed out by Mahrer and Pielke (1977) and Asai and Mitsumoto (1978). 4. Discussion It is concluded from the results of the preceding section that the surface wind in the area between the sea breeze front and the end of the valley was enhanced by the topography of the valley. The inlandward wind appeared to be unified into one flow extending from the coastline to the plateau, thus suggesting the generation of the ESB. The wind became stronger in the narrower valley. In this section, the mechanism of the valley wind and the process of the formation of the ESB is discussed in detail. The distributions of the relative surface pressure depression along the center of the valley in RUN 21 at 16:00 was added in Fig. 4. Among RUNs 3, 5 and 21, in which a sea breeze appeared, the region of large pressure gradient stayed relatively near the coast both in RUN 3 and RUN 5, while the gradient in the inner part of the valley became rather great in RUN 21. The difference of the pressure gradient between the coastal side and the valley side was smallest in RUN 21. Considering the fact that the surface pressure in the valley showed the largest decrease in RUN 21, the airmass over the valley in RUN 21 should be warmer than that of RUN 3 and 5. Actually, the potential temperature of the layer between the 1000m and the 2000 m elevations was the warmest in RUN 21 at 16:00. The maximum of the potential temperature difference between RUN 5 and RUN 21 at 80 km inland from the coastline was 0.8 K at 16:00 in the layer, while the differencee of the layer lower than 1000m was only 0.2 K (Fig. 9). The heating

9 August 1990 H. Kondo 443 Fig. 10. The vector wind distribution about 50 m above the surface in RUN 5 (left) and (right) at 20:00. rate of the layer during 8:00 to 16:00 on the point was always greater in RUN 21 than that in RUN 5 except for 8:00. Table 2 shows the total heating of the layer between 1000m and 2000m during the same 8 hour period, and the contribution of each fraction through the advection terms and vertical diffusion term. The table shows that the most of the heat was contributed by adiabatic heating through the descending motion and that the heating by this descending motion was twice as great in RUN 21 as that. in RUN 5. The narrower valley produces stronger downward motion over the valley near noon, which warms the atmosphere there. Hence the surface pressure was most reduced in the narrowest valley. This mechanism also explains the reason why the largest depression appeared at the end of the valley. The mechanism for generating the valley wind was found to be basically the same as that considered by Defant (1951). That is, first the up-slope wind develops over the side slope and then the valley-toplateau wind develops across the valley slope. After the above, the counter flow of both winds descends over the valley and heats the air there adiabatically. Because this warming reduces the surface pressure in the valley more than that of outside of the valley, a wind into the valley is generated by the surface pressure gradient. The pressure gradient (see Fig. 4) was important in the generation of the ESB. The results of RUN 1, RUN 3 and RUN 5 were compared. It should be noted that the pressure gradient was largest in the inner half of the valley in RUN 5 among others, although the gradient was large near the sea breeze front in all the RUNS. Then the spatial variation of the pressure curve in RUN 5 was more linear than that in RUN 1 or RUN 3. Hence one unified circulation, that is, the ESB was generated from the coastal area to the plateau. The pressure depression in RUN 5 over the plain was not the simple summation of those in RUN 1 and 4. For instance, the depression near 30 km from the coastline in RUN 5 was smaller than that in RUN 1. This means the pressure depression in RUN 5 was not the result of a linear combination of RUN 1 and RUN 4, even though the topography of RUN 5 was a linear combination of them. Figure 11 shows (a) the potential temperature profiles at the coastline, at 40km inland from the coastline and at 93.3 km from the coastline in RUN 3 and (b) those at the same location along the center of the valley in RUN 5 at 16:00. The profile over the point of inland plateau (point D in Fig. le) is added in Fig. lib. The difference in the potential temperature over the plateau and that over the other locations at elevations above 1400m caused to plain-to-plateau wind (Mannouji, 1982). The difference in the mean potential temperature of the mixing layer between at 40 km and at 93.3km inland from the coastline was greater in RUN 5 than in RUN 3. This is consistent with the greater pressure gradient in RUN 5 than in RUN 3 at the inland valley. The time variations of the heating rate of the air near the surface at 40km from the coastline is shown in Fig. 12 during the period from 8:00 to 16:00. This figure indicates the following. At 40km, the air below 1200 m was heated less in RUN 5 than in RUN 3. This lesser heating was caused by the

10 444 Journal of the Meteorological Society of Japan Vol. 68, No. 4 Table 2. Averaged heating (K) in the layer between 1000 m and 2000 m at 80.0 km inland from the coastline during 8:00 to 16:00. Fig. 12. Time variation of the heating rate of the layer between the surface and 1200 m elevation at 40 km from the coastline in RUN 5 (left) and in RUN 3 (right). A is total heating, B is heating by the advection of w-component, C is that of u-component, D is the heating by the vertical diffusion and E is that by the horizontal diffusion. Fig. 11. Potential temperature profiles at 16:00. A, B and C are at the coastline, at 40 km and at 93.3 km from the coastline, respectively, (a) in RUN 3 and (b) in RUN 5. D in (b) is that over the plateau (point D in Fig. le). larger cooling through the advection of inland-ward flow. The cooling started at 10:00 in RUN 5, which was two hours earlier than RUN 3. This was consistent with lower averaged potential temperature in RUN 5 at 40 km inland. That is, the valley topography caused an early start of strong and cool (heavy) inland-ward flow, which brought along a relatively high surface pressure from the coast. Thus, the surface pressure gradient near the coast became smaller, while that of the inland valley became larger than those in RUN 1 and RUN 3. Then the surface pressure gradient from the coastline to the end of the valley became nearly linear in RUN 5. Finally the effect of Coriolis force was investigated. The region was extended in y-direction, and whole valley topography was included in the model, although vertical resolution was reduced to half that for previous RUNS, because of the limited computational capacity. The topography was similar to that for RUN 5, and the results were compared between f at the equator (RUN 41) and at 35 N (RUN 42). The horizontal wind distribution near the surface at 16:00 is shown in Fig. 13. With Coriolis force, the horizontal wind direction slightly turns toward right. However there was little difference between the results of RUN 41 and Conclusion The interaction between the sea breeze and the thermally and topographically induced flow was investigated, using both a two- and three-dimensional

11 August 1990 H. Kondo 445 Fig. 13. The vector wind distribution about 50 m above the surface in (a) RUN 41 and (b) RUN 42. mesoscale model. Three cases were calculated using a two-dimensional model, in which the effects for land-sea, plain-slope-plateau and sea-plainslope-plateau were examined, respectively. Two cases were calculated using a three-dimensional model in which the effects for plain-valley-plateau and sea-plain-valley-plateau were examined, respectively. The results indicated that the wind in the valley between the sea breeze front and the up-slope wind at the end of the valley was more enhanced than that over the two-dimensional plain, demonstrating the generation of the ESB. The depth of the inland-ward wind over the plain or in the valley was equal to the mixing layer depth for cases where the height of the plateau was lower than the mixing layer height. When the plateau height was greater than the mixing layer depth, the depth of the wind system became greater than the depth of the mixing layer. Even for these cases, the wind in the mixing layer was stronger than that over the mixing layer. In the three-dimensional model, a narrower valley generated a stronger wind in the valley. When slope and plateau were located near the coast, the sea breeze in the morning was enhanced by the topography, but the advance of the sea breeze was blocked by the same topography in the afternoon. A circulation developed from over the side slope in the valley and to over the plateau (valley-to-plateau wind) reduced the surface pressure in the valley due to the adiabatic warming by its descending motion over the valley. Because the pressure difference was greater with the narrower valley, the largest depression developed at the end of the valley. The valley wind started in the early morning and gradually extended to the sea side. There was large pressure gradient in the zone of the sea breeze front. Since the valley wind that developed along the sea breeze partly carried the cool and heavy air from over the sea inland with the development of the sea breeze, the pressure gradient in the sea breeze zone with a three-dimensional valley became smaller than that for a pure sea breeze. Then the difference in the pressure gradient between the sea breeze zone and that of the inner valley became smaller and a zone with a large pressure gradient which extended from the coastal area into the valley appeared, making one unified circulation. The ESB was thus generated through the non-linear combination between the sea breeze and the valley wind. The Coriolis force did not have as much of an effect on the daytime results. Acknowledgement The calculation was executed on FACOM M780 of the Research Information Processing System in the Agency of Industrial Science and Technology. The comments on the manuscript by Dr. H. Niino, of the Meteorological Research Institute and Dr. S. Mitsumoto, of the National Institute for Environmental Studies and Mr. H. Yoshikado, of the National Research Institute for Pollution and Resources are greatly appreciated. The comments by Dr. A. Huber, of the U.S. Environmental Protection Agency, were helpful in improving the manuscript. References Asai, T. and S. Mitsumoto, 1978: Effects of inclined land surface on the land and sea breeze circulation. A numerical experiment. J. Meteor. Soc. Japan, 56, Defant, F., 1951: Local winds. Compendium of Meteorology, Fujibe, F. and T. Asai, 1979: A study of local winds in the Kanto district. Part 1: Structures of wind systems with diurnal variation. Tenki, 26, (in Japanese; correction in Tenki, 28, 202). Kawamura, T., 1977: The climate of the land-sea breezes. Report of the Atmospheric Environmental Experiment in the South Kanto area I. Japan Meteorological Agency (in Japanese).

12 446 Journal of the Meteorological Society of Japan Vol. 68, No. 4 Kikuchi, Y., S. Arakawa, F. Kimura, K. Shirasaki and Y. Nagano, 1981: Numerical Study on the effects of mountains on the land and sea breeze circulation in the Kanto district. J. Meteor. Soc. Japan, 59, Kondo, H. 1990: A numerical experiment of the "extended sea breeze" over the Kanto Plain, J. Meteor. Soc. Japan, 68, Kurita, H., H. Ueda and S. Mitsumoto, 1988: Threedimensional meteorological structure of long-range transport of air pollution under light gradient wind. Tenki, 35, (in Japanese). Linden, P.F. and J.E. Simpson, 1986: Gravity-driven flows in a turbulent fluid. J. Fluid Mech., 172, Mahrer, Y. and R.A. Pielke, 1977: The effects of topography on sea and land breezes in a two-dimensional numerical model. Mon. Wea. Rev., 105, Mannouji, N. 1982: A numerical experiment on the mountain and valley winds. J. Meteor. Soc. Japan, 62, Ookouchi, Y., M. Uryn and R. Sawada, 1978: A numerical study of the effects of mountains on the land and sea breezes. J. Meteor. Soc. Japan, 56, Tsuruta, H., 1986: The process of transportation of the pollutants from coastal area to inland area. Report of "Environmental Science", B280-R11-2, (in Japanese). Yoshikado, H. and H. Kondo, 1989: Inland penetration of the sea breeze in the suburban area of Tokyo. Boundary Layer Meteor., 48,