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992 MONTHLY WEATHER REVIEW VOLUME 130

Numerical Study of the Diurnal Cycle along the Central Oregon Coast during Summertime Northerly Flow

S. BIELLI,P.BARBOUR,R.SAMELSON,E.SKYLLINGSTAD, AND J. WILCZAK* College of Oceanic and Atmospheric Sciences, Oregon State University, Corvallis, Oregon

(Manuscript received 23 May 2000, in ®nal form 5 September 2001)

ABSTRACT A triply nested mesoscale atmospheric numerical model is used to study the dynamics of the diurnal cycle of the summertime lower atmosphere along the central Oregon coast. Simulations of four consecutive days in September 1998, during which the winds were strong and northerly, are analyzed. Comparisons with pro®ler observations suggest that the model performed well enough to provide a useful estimate of the diurnal circulation. During the four days of interest, the low-level wind pattern has a broad maximum between Cape Blanco and Cape Mendocino, with a large north±south gradient along the Oregon coast. The low-level jet undergoes diurnal horizontal and vertical displacements, which partially resemble previous observational and modeling results along the California coast. In both the model and the pro®ler data, there is a minimum in northerly wind between 1500 and 1800 UTC (0700 and 1000 local time), and a double maximum in offshore ¯ow above the marine boundary layer, with peaks near 0700 and 1600 UTC. At the jet core height, the advection of alongshore momentum is an important component of the alongshore force balance. After 2100 UTC, this advection is the main term balancing the pressure gradient force. Thus, in contrast to the previous results for the California coast, the diurnal circulation is fundamentally three-dimensional in the coastal zone, for several hundred kilometers alongshore and as far as 100 km offshore. The blocking effect of coastal terrain has a strong in¯uence on the diurnal circulation.

1. Introduction ized by a low-level jet with maximum values of 30 m s Ϫ1 at a few hundred meters altitude. The vertical struc- During spring and summer, the west coast of the Unit- ture is characterized by an inversion near or at the ed States is dominated by the Paci®c high centered ap- altitude of the jet maximum. The wind shear was found proximately 1000 km off the north California coast and to be due to thermal wind generated by the large hor- a thermal low inland over California. This regime pro- izontal temperature gradient between the water and the duces persistent northerly (upwelling favorable) winds land. Some observations in the Oregon area have been along the coast interrupted by periods of weak or south- discussed by Neiburger et al. (1961), Johnson and erly ¯ow. The typical summertime coastal meteorolog- O'Brien (1973), Meitin and Stuart (1977), and Elliot ical conditions include a stable marine atmospheric and O'Brien (1977), and more recently by Dorman and boundary layer with northerly wind and a low-level jet near the top of the boundary layer. Winant (1995). The research described here was car- Most observational and numerical studies of this re- ried out as part of the Oregon National Ocean Part- gime have focused on the California coast (e.g., nership Program (NOPP) project The Prediction of Beardsley et al. 1987; Zemba and Friehe 1987; Winant Wind-Driven Coastal Circulation. This project was un- et al. 1988; Bridger et al. 1993; Banta 1995; Dorman dertaken with the recognition that understanding and et al. 1999; Holt 1996; Burk and Thompson 1996; Burk modeling the coastal ocean requires improved knowl- et al. 1999, KoracÆin and Dorman 1999; Dorman et al. edge of coastal atmospheric processes. The atmospher- 2000). For instance, Zemba and Friehe (1987) showed ic component of the Oregon NOPP project had two that along the California coast, the marine atmospheric speci®c objectives: ®rst, to provide an estimate of air± boundary layer during northerly winds was character- sea ¯uxes for use in the ocean modeling component, and second, to study the dynamics of the coastal atmosphere along the Oregon coast, including particu- * Current af®liation: NOAA/ETL, Boulder, Colorado. larly the diurnal cycle. The present contribution addresses the second of these Corresponding author address: S. Bielli, Dept. of Atmospheric two objectives. We focus primarily on mesoscale nu- Sciences, University of Washington, Box 351640, Seattle, WA 98195. merical model simulations centered on Newport for four E-mail: [email protected] consecutive days in September 1998. Section 2 presents

᭧ 2002 American Meteorological Society

Unauthenticated | Downloaded 09/24/21 06:03 AM UTC APRIL 2002 BIELLI ET AL. 993 brie¯y the model, its initialization, and a statistical comparison of modeled and observed ®elds during summer 1999, including results from a wind pro®ler stationed in Newport as part of the Oregon NOPP project. In section 3 we describe the meteorological situation for the Sep- tember 1998 simulations. The results of the control simulation as well as comparison with wind pro®ler data and a 1-day sensitivity study without terrain are presented in section 4. In section 5, we examine the horizontal momentum budget along the coast, and present a ®nal discussion and a summary in section 6.

2. Model description and initial conditions a. The ARPS model Numerical simulations were performed using the Ad- vanced Regional Prediction System (ARPS) described in detail in Xue et al. (1995). The version of the model used for this study is a three-dimensional, nonhydros- tatic, compressible version, with a terrain-following vertical coordinate. Only the warm microphysical processes FIG. 1. The 12- and 4-km domains with the terrain height and the are taken into account, using the Kessler (1969) param- position of the wind pro®ler near Newport (44.7ЊN, 124.07ЊW) and eterization. In all simulations, a stretched vertical co- buoy 46050 (44.62ЊN, 124.53ЊW). The dotted line is the position of the cross sections through Newport. ordinate was used to maximize resolution in the lowest of the 32 levels of the model. Close to the surface, the resolution was 20 m with the ®rst point above ground conditions were imposed from Eta analyses, and the level at 20 m; it progressively increased to an average model was initialized every day at 1200 UTC with a of 450 m. The top of the model is at 13-km altitude, cold start and was run for 24 h. For the 12- and 4-km and a sponge layer is applied above 9 km to minimize domains, initial and time-dependent lateral boundary the re¯ection of internal gravity waves. Increasing the conditions were obtained, respectively, from the 36- and average vertical grid size to 250 m or increasing the 12-km output in a one-way nesting procedure. A sen- altitude of both the sponge layer or the top of the model sitivity study was conducted for 11 September 1998 to did not result in any major differences during a 24-h study the effect of the coastal orography, which con- simulation. Indeed, the sponge layer is relatively low, sisted of removing the terrain but leaving the surface but the main activity is taking place in the low levels characteristics of land unaltered. for this kind of event. Model physical parameterizations also included a 1.5-order turbulent kinetic energy (TKE), fourth-order advection in both horizontal direc- c. Statistical wind pro®ler and buoy veri®cation tions, and radiation and surface ¯ux schemes. The model Before proceeding to the September 1998 case study, topography was derived from the global terrain database we brie¯y summarize a statistical comparison of mod- with 30-s terrain resolution. The model was triply nested eled and observed variables during summer 1999. As with one-way interaction on 60 ϫ 60 grids with 36-, part of the Oregon NOPP project, forecast-mode (36-h 12-, and 4-km resolution. Each domain was centered forecast with cold start at 0000 UTC each day) simu- near Newport, Oregon (44.7ЊN, 124.0ЊW) (Fig. 1). lations were conducted during June through August 1999 with the ARPS model, on the 36- and 12-km do- b. Initial conditions and case study mains. Statistical comparisons of summer 1999 modeled variables from the 12-km grid with National Data Buoy The 4-day period from 1200 UTC September 11 to Center (NDBC) buoy 46050 (44.62ЊN, 124.53ЊW) and 1200 UTC September 15 1998 was selected for this land-based meteorological (Coastal-Marine Automated study. The ¯ow along the Oregon coast experienced Network, CMAN) observations at Newport (Fig. 1) are strong northerly wind during this period and was char- shown in Table 1. Statistics have been calculated based acteristic of a pattern frequently observed during sum- on hourly data and hours 4 to 27 have been used for mer months. the model. The model simulates pressure and northerly The 36-km domain was initialized using data inter- wind variability well, with high correlations between polated from the National Centers for Environmental modeled and observed ®elds. The model shows a rel- Prediction (NCEP) ``early'' Eta Model output (grid 212, atively good agreement with the buoy data for the mean 40-km resolution). Time-dependent lateral boundary pressure, but underestimates it at Newport, which might

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TABLE 1. Statistics of buoy 46050, model, and Newport CMAN (NWO3) data during summer 1999 (Jun±Aug); N is the number of observations, u is cross-shore wind, ␷ is alongshore wind, and std is standard deviation. The last three columns are the correlation, the bias, and the root-mean-square error between the model and the observations. Statistical properties are computed using all days when there were both observations and model simulations. Model variables from forecast hours 4±27 are used for the comparison. The NDBC buoy wind observations, which were at 5 m above sea level, have been corrected to 10 m assuming neutral stability. Buoy 46050 Model N Mean Std Mean Std Corr Bias Rmse Pressure (Pa) 1920 1017.8 3.2 1017.3 3.2 0.94 Ϫ0.56 1.23 Temperature (ЊC) 1931 13.9 1.5 14.8 2.1 0.66 0.95 1.87 u (m sϪ1) 1912 1.2 1.9 2.0 1.9 0.51 0.81 2.04 v (m sϪ1) 1912 Ϫ2.6 4.2 Ϫ3.2 3.8 0.82 Ϫ0.68 2.52 NW03 Model N Mean Std Mean Std Corr Bias Rmse Pressure (Pa) 2031 1018.4 3.1 1013.9 3.2 0.94 Ϫ4.48 4.62 Temperature (ЊC) 2031 12.7 1.8 14.4 2.9 0.56 1.71 3.02 u (m sϪ1) 2031 0.8 1.5 1.5 1.7 0.61 0.67 1.55 v (m sϪ1) 2031 Ϫ1.2 3.9 Ϫ1.8 2.6 0.75 Ϫ0.60 2.65 be in part due to the presence of unresolved terrain. The formance of the model during southerly ¯ow episodes. standard deviations are comparable in all cases. The The comparison between model and pro®ler winds will largest differences are for alongshore wind and tem- be explored further later [section 4c(3)] for the 4-day perature at Newport. The warm bias in the model tem- period of interest in September 1998, for which the perature ®eld at both locations may be due to ocean similarity of modeled and observed diurnal variations upwelling that is not resolved by the Eta sea surface is greater than in these summer 1999 means. The results temperature analysis (Samelson et al. 2002). in Fig. 2 demonstrate that the diurnal cycle is an im- A 915-MHz Radio Acoustic Sounding System portant component of the summertime lower-atmo- (RASS) pro®ler was in place at Newport during this sphere circulation along the Oregon coast, and motivate period, supported by the Oregon NOPP project. Hourly the more detailed analysis that follows. consensus-averaged winds available from the National Oceanic and Atmospheric Administration/Environmen- 3. Synoptic situation tal Technology Laboratory (NOAA/ETL) (www7.etl. noaa.gov/data/archive/realtime) were used to produce We focus on a case study for four days 11±15 Sep- the wind composites. Details on wind pro®ler process- tember 1998, during which a well-developed diurnal ing and accuracy can be found in Weber et al. (1993). cycle was observed. The synoptic pattern during this Also Ralph et al. (2000) demonstrated the capabilities period was fairly typical for the summertime along the of such radars in West Coast summertime studies. Wind west coast of the United States, with the Paci®c high data were ®rst linearly interpolated to constant height located approximately 1000 km off Cape Mendocino, a and then hourly averaged over the summer. The hourly thermal low over California, and north-northwesterly mean diurnal cycles of the Newport pro®ler and model ¯ow along the coast. The sea level pressure and 500- winds for the summer 1999 are shown in Fig. 2. A hPa analyses at 0000 UTC (1700 local time) for 12 and measurable diurnal cycle is present in both of these 15 September 1998 from NCEP Eta data, interpolated summer 1999 means, which include synoptic distur- to the model grid, are displayed in Fig. 3. At 500 hPa, bances, episodes of southerly ¯ow, and other periods the Paci®c Northwest is initially under the in¯uence of when the diurnal cycle is not well developed. The timing a large-scale ridge modi®ed by the presence of a low and the depth of these mean modeled and observed pressure system centered over the California±Nevada diurnal variations are in rough agreement. However, the border (Fig. 3b). At 500 hPa, to the north of Oregon a model overestimates the strength of the wind with max- strong westerly jet extending into British Columbia and imum (minimum) values of 7.25 (0.50) m sϪ1 for the an area of anticyclonic ¯ow produces northerly wind model and 4.75 (0.75) m sϪ1 for the observations, re- over the coasts of Washington, Oregon, and northern spectively. Also note that the jet maximum observed at California. 100±300-m altitude (vs 250-m altitude in the model) An area of high sea level pressure builds in from the occurs 3 h late, and the relatively strong winds near 500 southwest with lower pressure over inland areas of Ne- m persist several hours longer in the observations than vada, California, and eastern Oregon (Fig. 3a). The up- in the model. Finally, the wind minimum observed at per-level closed circulation feature slowly moves south 800±1100-m altitude and 1700±2100 UTC does not ap- and then begins to move east and exit the study region pear in the model. Some of these differences in the mean on 1200 UTC 12 September. Higher pressure builds in summer 1999 ®elds may be due to relatively poor per- its wake over Oregon and the strong westerly ¯ow to

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the next few days, the upper-level ridge moves slowly to the east in association with a deepening trough over the Paci®c. This results in a slow rotation and strength- ening of the winds over Oregon. By 0000 UTC 15 Sep- tember, the winds at 500 hPa are from the southwest over most of Oregon and the Paci®c Northwest and the surface thermal low has shifted north and slightly to the east (Figs. 3c,d). The surface winds along the Oregon coast were northerly and strong throughout the study period, but decreased slowly from 11 September to 15 September.

4. Model results a. Large-scale structure: 36-km simulation In this section we summarize the horizontal distribution of the wind over the 36-km domain during 11± 15 September 1998. Note that, although the model uses terrain-following sigma coordinates, all the ®gures show results converted to a constant height above sea level. Thus, for horizontal cross sections, regions where the model terrain is greater than the cross-section height appear blank. The mean wind speed at 200 m, the approximate height of the wind speed maximum in the low-level jet, has a broad large-scale maximum between Cape Blanco and Cape Mendocino at about 200 km offshore, with a single peak greater than 18 m sϪ1 (Fig. 4b). This wind speed maximum is located in a region where the terrain is higher and closer to the coastline than farther north or south, and may be in part due to the acceleration induced by Cape Blanco, the westernmost point of the Oregon±California coast. Along the Oregon coast, the wind ®eld has a strong north±south gradient, which reverses at Cape Blanco. The mean surface winds are similar, with a maximum value of about 10 m sϪ1. The broad maximum wind speed at this time period is consistent with Special Sensor Microwave/Imager (SSMI) observations of 10-m wind speed (not shown). The corresponding standard deviation of the 200-m wind speed is greater than 5 m sϪ1 and is maximum along the coast, especially at the position of the maximum wind speed (Fig. 4a). At the surface, the standard deviation of wind speed is only 2 m sϪ1, and has a different pattern along the coastline.

b. Diurnal cycle: 12-km simulation FIG. 2. Hourly mean wind speed over summer 1999 (Jun±Aug) from (a) the ARPS model and (b) the Newport wind pro®ler. Values 1) DYNAMICAL STRUCTURE are in m sϪ1 with contour interval of 0.5 m sϪ1. The scale for the wind vectors is beneath. In this section, we summarize the mesoscale structure and variations of the 12-km model wind ®eld in a west± east line crossing Newport and at the Newport pro®ler location. In the 12-km model, the peak alongshore ve- the north dips southward while the thermal low in cen- locity in the low-level jet reaches 16 m sϪ1 at 200 m tral California develops and strengthens. By 1300 UTC and 124.5ЊW (Fig. 5a). The diurnal variation of the wind 13 September, the upper-level winds over the central speed near Newport is clearly apparent in cross sections Oregon coast have a mostly westerly component. Over of 4-day hourly mean wind speed (Figs. 5b,c): during

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FIG. 3. (a) Sea level pressure (hPa) and horizontal wind vector (m s Ϫ1) from 12 Sep 1998 at 0000 UTC, and (b) same as in (a) but for 15 Sep 1998 at 0000 UTC. (c) The 500-hPa height (m) and wind vector (m s Ϫ1) for 12 Sep 1998 at 0000 UTC, and (d) same as in (c) but for 15 Sep 1998 at 0000 UTC from NCEP Eta analyses interpolated to the ARPS grid. The arrow in (c) shows the approximate position of the wind pro®ler near Newport.

the day, the wind speed intensi®es from 12 to 16 m sϪ1 an artifact of the initialization; it appears also in a sim- as it moves toward the coast, with a maximum value at ulation initialized at 0900 UTC instead of 1200 UTC 0300 UTC. At 200 m, the jet core is located at about (not shown). 100 km off the coast in the early morning and moves The axis of maximum wind speed rises from 200 m to about 50 km offshore during the day, when the wind near the coast up to 400 m at about 400 km offshore. speed reaches its maximum (Fig. 5b). The diurnal var- A small secondary wind maximum appears above the iation extends over the entire domain, but the largest coastal mountains and can be seen over the valley by variations are con®ned within the ®rst 200 km offshore 0300 UTC (Figs. 5a,b). The jet core seems to lower (Fig. 5b). Note also the wind speed minimum between slightly down to 200 m at Newport during the morning 1500 and 1800 UTC (Figs. 5b,c). This minimum is not and to rise slightly during the night (Fig. 5c).

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2) THERMODYNAMIC STRUCTURE Maximum temperatures along the coast occur during the afternoon hours in response to solar heating. The maximum heating occurs near the coast, and spans as far as 50 km offshore, a few hours before the wind speed maximum. The water vapor content reaches a minimum on the west side of the jet core at the same time as the wind shows its maximum. At 0300 UTC, the top of the well-mixed marine boundary layer lies at about 200-m altitude near the coast and deepens offshore to about 400 m at 128ЊW (Fig. 6a). The jet core appears at the top of the boundary layer. The marine inversion is weak near Newport with a vertical temperature gradient of only about 3ЊC 100 mϪ1 and a nearly saturated boundary layer below. There is a tongue of dry air at the jet core position associated with weak downward motion and large-scale subsidence (Fig. 6b).

c. Diurnal cycle: 4-km simulation

1) MODEL RESULTS:VERTICAL AND HORIZONTAL STRUCTURE We now proceed to focus on the mesoscale circulation along the central Oregon coast and particularly the diurnal cycle and cross-shore sea-breeze circulation near Newport, including the comparison of 4-km modeled winds at Newport with the RASS pro®ler observations. In the following, we consider only the 4-day hourly means of modeled and observed variables, but the diurnal cycle on each of the four days is similar to this mean. Meridional cross sections at 200-m altitude and vertical cross sections through Newport of the 4-day hourly mean cross-shore (u, zonal) and alongshore (␷, meridional) 12-km wind ®elds are shown in Fig. 7. At this scale, the onshore sea breeze, which develops as the sun rises and the land warms more quickly than the adjacent water, is clearly recognizable (Fig. 7c). Onset of the sea- breeze ¯ow occurs at approximately 1500 UTC (0800 local time) and shutdown at 0800 UTC (Fig. 7a). The sea-breeze intensity increases near the surface, reaching a maximum of 4 m sϪ1 between 2100 and 0000 UTC (1400±1700 local time), with a maximum depth of the ¯ow of about 200 m over the water. At the surface, onshore ¯ow extends as far as 150 km offshore (Fig. 7c). The hourly mean cross-shore wind is easterly up to 4-km altitude near Newport (not shown) during the FIG. 4. (a) Standard deviation of the wind speed in m sϪ1 with simulation period, except when the sea-breeze circula- contour interval of 0.5 m sϪ1 and (b) 4-day mean (11±15 Sep 1998) wind speed at 200-m altitude over the 36-km domain based on hourly tion is present. At 0300 UTC, the return ¯ow associated output in m sϪ1 with contour interval of2msϪ1. The dotted line with the sea-breeze circulation can be separated into two shows the position of the baseline for the cross-coast cross sections components: above the coastal range at 600±800 m, and along the Newport line. offshore at 400±600 m (Fig. 7c). Diurnal variation of both components of the horizontal wind is large at the jet core altitude. The structure of the northerly wind, with a broad maximum between 0000 and 0300 UTC that extends 100 km offshore, is very similar to that in the 12-km simulation (Figs. 5b and 7b). Note however

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FIG. 6. Vertical cross section at 0300 UTC at Newport of 4-day hourly mean (11±15 Sep 1998) for the 12-km domain: (a) potential temperature ␪ in K with contour interval of 0.5 K and (b) water vapor Ϫ1 Ϫ1 content q␷ in g kg with contour interval of 0.5 g kg .

that the magnitude of the wind speed is slightly weaker, and the onshore sea-breeze ¯ow cannot be resolved in the 12-km simulation due to coarser resolution. The maximum cross-shore gradient of the 200-m northerly wind is large and roughly constant near the coast throughout the simulation. The horizontal structure of the wind is quite regular near the coast within the 4-km

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FIG. 5. The 4-day hourly mean wind speed (11±15 Sep 1998) over the 12-km domain: (a) altitude±longitude plot at 0300 UTC and 44.7ЊN (line through Newport; cf. Fig. 1), (b) time±longitude plot at 200 m altitude and 44.7ЊN, and (c) vertical cross section at Newport. Values are in m sϪ1 with contour interval of 1 m sϪ1.

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FIG. 7. The 4-km domain: 4-day hourly mean (11±15 Sep 1998) cross-shore wind (u)inmsϪ1 with contour interval of 0.5msϪ1. (a) Time±longitude cross-section at 200-m altitude and 44.7ЊN (Newport line), (c) altitude±longitude cross section at 0300 UTC at Newport, and (e) latitude±longitude cross section at 200-m altitude and 0300 UTC; (b), (d), and (f), respectively, same as in (a), (c), and (e) but for the alongshore wind (␷)inmsϪ1 with contour interval of 1 m sϪ1.

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FIG. 8. Time±height plot at Newport of 4-day hourly mean (11±15 Sep 1998) of (a) u in m sϪ1 with contour interval Ϫ1 Ϫ1 Ϫ1 of 0.5 m s , (b) ␷ inms with contour interval of 1 m s , (c) ␪ in K with contour interval of 1 K, and (d) q␷ in gkgϪ1 with contour interval of 0.5 g kg Ϫ1 domain at the height of the jet core, with a steady south- breeze at the surface. The boundary mixed layer lowers ward increase of several meters per second in northerly until 2100 UTC, associated with an increase in northerly and offshore ¯ow (Figs. 7e,f). This gradient is the me- wind and an increase of onshore ¯ow (Fig. 8c). Then soscale expression of the large-scale gradient that is the marine boundary layer deepens as the onshore ¯ow evident in Fig. 4b. As shown below, it is large enough builds up, and the onshore wind increases to a maximum to affect the horizontal momentum balance. between 2100 and 0000 UTC in correlation with temperature and water vapor content maxima. Three hours later, the northerly wind attains its maximum whereas 2) MODEL VERTICAL PROFILES AT NEWPORT the offshore ¯ow aloft (above 400 m) reaches a mini- At Newport, the northerly wind at 200 m and below mum. Finally, the northerly wind decreases and the off- has a minimum between 1500 and 1800 UTC, associated shore ¯ow aloft rises again. with increasing onshore ¯ow (Figs. 8a,b). This mini- Above the marine boundary layer, the offshore ¯ow mum occurs progressively later aloft, in phase with de- shows two maxima: around 600 m between 1500 and creasing offshore ¯ow aloft. The maximum offshore 1800 UTC, and around 400 m between 0600 and 0900 ¯ow aloft is coincident with the appearance of the sea UTC. These two maxima are not speci®c to Newport:

Unauthenticated | Downloaded 09/24/21 06:03 AM UTC APRIL 2002 BIELLI ET AL. 1001 time pro®les of the cross-shore wind farther north or south of Newport near the coast (not shown) display similar features, with two minima of different magni- tudes found at two different altitudes. At 600 m, as the offshore wind maximum starts to decrease, the potential temperature increases. Between 200 and 1000 m, the offshore ¯ow maxima are correlated with relatively warm temperatures and low humidities.

3) COMPARISONS WITH WIND PROFILER DATA The model pro®les at Newport discussed above may be compared with 4-day hourly mean horizontal winds for 11±15 September 1998 from the Newport RASS wind pro®ler (Fig. 9). Surface data from the CMAN station have also been included in the wind pro®ler composites. Observed cross-shore wind is primarily offshore up to at least 1200-m altitude with maxima near 1000 m between 1500 and 1800 UTC and near 700 m between 0600 and 0900 UTC (Fig. 9a). These observed offshore ¯ow maxima occur at higher altitude than those for the modeled winds but have similar timing and roughly the same magnitude (Figs. 8a, 9a). The onset of both the observed and modeled sea breeze occurs around 1600 UTC, but for unknown reasons the observed sea breeze persists only until 0700 UTC, 2 h before the model ¯ow reverses. The observations show deep onshore ¯ow between 2300 and 0100 UTC, reaching to 1100 m. The model shows a deep weakening of offshore ¯ow during this period, but no ¯ow reversal above 200 m. The sea breeze reaches a maximum around 2000 UTC in the observations and between 2100 and 0000 UTC in the model (Figs. 8a and 9a). We show below that terrain has a strong in¯uence on the diurnal cycle and especially the cross-shore ¯ow; some of the differences between the model and these observations may be due in part to the incomplete representation and resolution of terrain effects in the model. It is interesting that the double maximum in offshore ¯ow above the boundary layer is found both in the observations and the model. This double maximum is not associated only with these four particular days, and can be identi®ed in the summer 1999 diurnal mean (e.g., near 1100 m in Fig. 2). This variation is not semidiurnal: the two maxima of offshore wind are separated by 15 h. It may involve inertial effects, as the inertial frequency is 17.45 h at this latitude. However, the two maxima appear FIG. 9. Time series at Newport from a composite of wind pro®ler every day at about the same time, so they are apparently and CMAN surface observations for 11±15 Sep 1998 based on hourly phase-locked to the diurnal forcing. The maximum data of (a) 4-day hourly mean cross-shore wind u inmsϪ1 with 1 around 1500 UTC is associated with the weakest cross- contour interval of 0.5 m sϪ , and (b) 4-day hourly mean alongshore wind ␷ in m sϪ1 with contour interval of 1 m sϪ1. coast potential temperature gradient, whereas the maximum around 0600 UTC is associated with a large tem- poral decrease of the cross-coast potential temperature the dynamics of the cross-shore wind above the marine gradient. Attempts to understand this oscillation in terms boundary layer in sections 5 and 6. of simple linear, forced±damped models were inconclu- The observed alongshore wind reaches a minimum sive, although there were consistent indications that in- of northerly ¯ow around 5 m sϪ1 near 1800 UTC, then ertial effects did contribute. We discuss in more detail increases to 13 m sϪ1 around 0300 UTC at 350-m al-

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titude before decreasing again (Fig. 9b). The model to the coastal zone in the control simulation is controlled tends to overestimate the wind speed and underestimate primarily by the topography, and not, for example, by the altitude of the jet core. The observed hourly mean interactions with the low-level alongshore jet. This con- maximum of northerly wind is 4 m sϪ1 weaker than the ®nement evidently intensi®es the diurnal variations in modeled maximum. The difference in altitude of the the coastal zone. maximum is roughly 150 m, but strong diurnal variation of the northerly wind occurs throughout the lower 1000 5. Dynamics of the diurnal cycle: Momentum m in both the model and observations. The pattern of balances the summer 1999 alongshore wind is very similar to the 4-day mean pattern, with northerly wind maximum and In this section, we analyze the dynamics of the model minimum, respectively, around 0100 and 1500 UTC diurnal cycle. We focus on the momentum balances for (Fig. 2). both the cross-shore and the alongshore circulation. One The 4-day and summer 1999 mean pro®ler obser- objective of this analysis is to assess the extent to which vations are generally similar to a 1-month composite of the diurnal cycle can be understood in terms as an es- the alongshore wind at Piedras Blancas, California, pre- sentially two-dimensional circulation, in which along- sented by Ralph et al. (2000), who ®nd large diurnal shore variations are negligible, as might be anticipated variations in alongshore wind up to 1200 m, with min- from the anisotropic geometry and forcing that typically imum and maximum northerly winds around 1700 and characterizes coastal zones, or whether it is intrinsically 0200 UTC, respectively. The 7.5 m sϪ1 amplitude of three-dimensional. diurnal variation of alongshore wind in the 1-month The terms in the equations of motion are diagnosed mean at Piedras Blancas lies between that of the present directly from the 4-km grid ARPS model using the fol- 4-day and the summer 1999 means. There appears to lowing form: be again some evidence of a double maximum in off- Pץ u 1ץ (١u Ϫϩf␷ ϩ Fx (5.1 ´ shore ¯ow in the Piedras Blancas composite, but it oc- ϭϪV xץ t ␳ץ curs above 1500 m and may have a different origin. The || | | | | | | || higher level of this feature might be due in part to the ͦͦ ͦͦͦ uu uuu higher mountains at Piedras Blancas. ta pcm Pץ ␷ 1ץ (١␷ ϪϪfu ϩ F , (5.2 ´ ϭϪV y yץ t ␳ץ d. No-terrain simulation || | | | | | | || The simulation conducted for 11 September 1998 ͦͦ ͦͦͦ ␷␷ ␷␷␷ with the terrain removed shows a number of signi®cant ta pcm differences from the control that illustrate the effect of where the x axis is taken to be the cross-shore direction terrain on the lower-atmospheric circulation. The ini- (west±east) and y the alongshore direction (north±

tialization procedure consisted of interpolating the south). For X ϭ (u, ␷), the terms are Xt, the local ten- NCEP Eta data onto the ARPS grid, applying a smooth- dency term; Xa, the advection term; Xp, the pressure ing, and adjusting the ®elds so that the anelastic mass gradient term; Xm, the friction (mixing) term; and Xc, continuity equation is satis®ed. Without terrain, the po- the Coriolis term. The horizontal and vertical structure sition and the shape of the 36-km wind speed maximum of the main terms are discussed here as well as diurnal are modi®ed with respect to the control simulation. variations at 200 m, height of the low-level jet. Rather than a single peak centered just south of Cape Blanco, two slightly weaker peaks are apparent in the a. Cross-shore momentum balance no-terrain simulation: one south of Cape Blanco and the second south of Cape Mendocino, both with a maximum The cross-shore circulation is driven by diurnal ¯uc- value of about 15 m sϪ1. The shifting of the wind speed tuations in the cross-shore pressure gradient, which arise maximum to the south is more evident at the surface. from contrasts in diurnal heating over land and sea, as Thus, the position of the maximum of wind speed is in a classical sea-breeze circulation. These gradients in¯uenced by the terrain, but the existence of the wind drive an ageostrophic cross-shore circulation, with re- speed maximum is evidently more related to the Paci®c turn ¯ow aloft. This circulation is similar to that found high and the continental thermal low. along the California coast (e.g., Zemba and Friehe 1987; Without terrain, the diurnal variation still extends Beardsley et al. 1987). over the 4-km domain but is weaker. The minimum in At 200 m, the height of the low-level jet, the cross- northerly wind is still present without terrain, but it shore momentum budget at the coast is nearly geo-

appears a few hours later than in the control simulation. strophic (Fig. 10a). Indeed, the Coriolis term uc, which An especially striking difference is the character of the re¯ects the variation of the alongshore wind component sea breeze. In the no-terrain case, the sea breeze prop- and shows a de®nite diurnal variation, is largely bal- agates onshore more than 100 km, in the classical man- anced by the cross-shore pressure gradient. The pressure ner. This suggests that the con®nement of the sea breeze gradient reaches a maximum in late afternoon and early

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FIG. 10. Time series (4-day hourly mean 11±15 Sep 1998) at 200 m of the different terms of the momentum cross- shore equation in cm sϪ2 based on hourly output at (a) Newport and (b) 0.5Њ of longitude offshore, and for the alongshore equation at (c) Newport and (d) 0.5Њ of longitude offshore. Here (u, ␷)a is the advection term, (u, ␷)c the Coriolis term (u, ␷) the pressure gradient term, (u, ␷)m the mixing term, and (u, ␷)t the local tendency term. uץ uץ evening, when the sea-breeze circulation is present and ua ϭϪu Ϫ ␷ . (5.3) yץ xץ -the northerly wind is maximum. Both up and uc mini mum values occur between 1500 and 1800 UTC, cor- |||| ͦͦ related with the northerly wind minimum and the in- uux ␷uy crease of the onshore ¯ow near the coast. Advection plays a role during late afternoon and early evening, At 200 m, the minimum in ua around 0000 UTC is beginning after 1800 UTC, when the onshore ¯ow in associated with a decrease in ␷uy that results primarily the low levels is established, with two peaks: one around from a decrease of the alongshore gradient of the cross- 1800 UTC when the onshore ¯ow is increasing, and the shore wind (uy). The tendency and the friction terms other around 0300 UTC when the onshore ¯ow is de- are negligible at 200-m altitude. Half a degree offshore, creasing. The cross-shore advection offsets some of the the ¯ow is near geostrophic, ua is small and does not pressure gradient increase so that the alongshore wind vary much throughout the day (Fig. 10b), and up attains is not geostrophically balanced during the see breeze its maximum (negative) value about 6 h earlier than at but basically quasigeostrophic. The cross-shore ¯ow is the coast. a thermally driven sea breeze, which is not in geo- The vertical distributions of these quantities at 0300 strophic balance and tends to turn the total ¯ow onshore. UTC in Newport and 0.5Њ longitude offshore are dis-

As vertical advection is small, we can write the ad- played in Fig. 11. The maximum in up occurs at the jet vection term as core height (200 m), beneath which the ¯ow is slowly

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decelerating. At the coast, ua plays a signi®cant role up to 600-m altitude, with a maximum at 200 m for the control simulation (Fig. 11a). Half a degree offshore, the ¯ow is almost geostrophic (Fig. 11b). The balance at the coast for the no-terrain simulation is very similar to the balance offshore for the control

simulation. But up is smaller in the no-terrain case and ua and ut are negligible. Thus the terrain, by blocking the inland penetration of the sea breeze, enhances the cross-shore pressure gradient at the coast. This enhanced pressure gradient is mainly balanced by the advection term.

b. Alongshore momentum balance It might be anticipated from previous work along the California coast (e.g., Zemba and Friehe 1987; Beards- ley et al. 1987) that the interaction of the alongshore jet and the diurnal cross-shore circulation would lead to a ¯uctuation of the alongshore ¯ow that is essentially two-dimensional in character, with the alongshore ¯ow responding primarily to advection and Coriolis forces associated with the cross-shore ¯ow. Surprisingly, this turns out not to be the case. Instead, there are large diurnal ¯uctuations in the alongshore pressure gradient, which evidently arise from intensi®ed heating over the mountains of southern Oregon and northern California, relative to northern Oregon and southern Washington. This alongshore pressure gradient drives an ageostrophic alongshore response that makes the diurnal circulation along the central Oregon coast fundamentally three-dimensional. We ®rst note that the alongshore pressure gradient is weaker than the cross-shore pressure gradient, but is not small (Figs. 10c,d); hence, a two-dimensional sea breeze interpretation will not work. The primary term balanc-

ing ␷ p at 200-m height during the morning hours (1800± 2100 UTC) is ␷ t. At 1600 UTC, ␷ p and the northerly wind reach a minimum. After the onset of the sea-breeze circulation, the dominant balance is between the pressure gradient and the advection term, which varies like the northerly wind. As the onshore ¯ow strengthens, the alongshore ¯ow would tend to decelerate due to cross-

shore turning but it actually accelerates due to ␷ a and ␷ t. Friction and Coriolis terms are negligible at the coast. Geostrophic balance is never reached. Half a degree

offshore, ␷ c is more comparable to ␷ a, but ␷ a is still the largest term balancing ␷ p during early evening, and the diurnal variation is still present, but weaker (Fig. 10d). Close to the surface in the lower half of the marine FIG. 11. Vertical pro®le of 4-day hourly mean (11±15 Sep 1998) boundary layer, the alongshore pressure gradient is bal- of the different terms of the momentum equation cm sϪ2 for the cross- shore budget: (a) in Newport and (b) 0.5Њ of longitude offshore for anced by advection, acceleration, and friction terms the control. when the jet is maximum (Fig. 12). Above the surface, the advection term near the coast is the primary term balancing the pressure gradient force up to 500-m altitude. Between 200- and 600-m altitude, the ¯ow is quite stationary as the tendency term goes down to zero, but it is unsteady in the marine boundary layer, where

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same way as in Eq. (5.3). The dominant term for the advection is vvy at the coast and offshore. This means that the alongshore variation of the alongshore wind plays a signi®cant role in the circulation; thus, the circulation is fully three-dimensional at least in the marine boundary layer. The advection term is also signi®cant in the no-terrain simulation and, thus, does not depend directly on terrain effects, unlike the advection term in the cross-shore momentum budget. The balance near the coast for the no- terrain simulation is very similar to the picture 0.5Њ offshore in the marine boundary layer, although the friction term for the no-terrain simulation has a relatively larger contribution to the budget at the jet core height (Fig. 12b). Above the surface, the advection is the main term balancing the pressure gradient force up to about 400-m altitude. The terrain increases the level up to which the advection term is dominant in the alongshore balance (Figs. 12a,c). Figure 13 shows a vertical cross section at 0300 UTC, and a zonal cross section versus time at 200 m of the

alongshore advection term (␷ a). The vvy term dominates near the coast in the lowest 400 m of the atmosphere and as far as 100 km offshore just above the marine atmospheric boundary layer (Fig. 13a). This term is important during late afternoon and early evening, when the diurnal variation of the alongshore wind is maximum (Fig. 13b). The positive part of the advection term (uvx) is mainly balanced by the tendency term, whereas the negative part (vvy) is balanced by the pressure gradient force. The large positive area over land is due to the offshore ¯ow above the sea-breeze circulation, while the positive area near the coast is due to the sea breeze itself. The advection pattern is similar between 43.7Њ and 45.7ЊN where the coastline is almost straight, except that the advection maximum occurs slightly earlier north of Newport and slightly later south.

6. Discussion and summary Although the model does not reproduce some aspects of the observations, comparisons with available data suggest that the model simulations provide a useful estimate of the lower-atmospheric circulation along the Oregon coast during the simulation period. From these simulations, we construct the following picture of the structure and diurnal cycle of the low-level atmosphere along the central Oregon coast and in particular near Newport during summertime, when strong upwelling favorable winds are present. The cross-shore circulation FIG. 12. Same as in Fig. 13 but for the alongshore momentum is a thermally driven circulation that arises from con- equation. trasts in diurnal heating over land and sea as in a classical sea-breeze circulation. The large alongshore pres- the tendency term is as large as the advection term. The sure gradient variations arise from stronger heating over friction term displays two maxima, at the height of max- the southern Oregon and northern California mountains, imum wind speed and at the surface. relative to northern Oregon and southern Washington; We decompose the advection term for the alongshore they drive an ageostrophic alongshore response and momentum budget into two terms uvx and vvy in the make the diurnal circulation along the central Oregon

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increasing. The main cross-shore balance is geostrophic. A weak secondary jet is also observed at this time. At the beginning of the second stage (1500±1800 UTC), the offshore ¯ow above the marine boundary layer is maximum and the jet is the weakest (Fig. 14a). As the sun rises, the temperature starts to increase. Dur- ing this period, the cross-shore wind in the boundary layer becomes onshore. At the same time, associated with the rapid increase of the cross-shore wind, the northerly wind reaches a minimum. The cross-shore and alongshore pressure gradients are still small and increase rapidly once the northerly wind is minimum. The cross- shore balance is geostrophic but advection starts to increase during this time due to the pressure gradient increase, and all terms in the alongshore contribute to the balance but are small. Once the sea breeze is established, during the third stage (1800±0300 UTC), the northerly wind increases and moves slightly downward and toward the coastline, and the marine boundary layer shallows slightly (Figs. 14b,c). The alongshore pressure gradient continues to increase and the advection, especially for the alongshore circulation, becomes more important. Both ␷ and vy play a signi®cant role in the increase of the alongshore advection. The circulation is fully three-dimensional up to 400-m altitude, as far as 1Њ of longitude offshore, and for at least several hundred kilometers alongshore. Above the marine boundary layer, a minimum of offshore ¯ow is attained. The sea-breeze wind starts then to decrease and the northerly wind continues to increase until about 0300 UTC, close to sundown near Newport. Finally, the northerly wind decreases and the offshore ¯ow above increases again during the fourth stage of the diurnal cycle, from 0300 to 1200 UTC. When the acceleration of the northerly wind reverses, the offshore ¯ow aloft decreases and reaches a maximum value. This maximum in offshore ¯ow aloft is also associated with a minimum in the alongshore pressure gradient and a maximum in tendency term. During this period, the water vapor content is relatively constant, the temperature decreases slightly, and the marine boundary layer deepens slightly. The diurnal variation of the wind near Newport appears to be large and not only con®ned to the coast. The low-level jet behavior near Newport is similar to the jet described by Burk and Thompson (1996) for the Cali- fornia coast during summertime. The altitude of the

FIG. 13. (a) Vertical (m) cross section at Newport of the 4-day model jet core is lower than the altitude reported in other hourly mean (11±15 Sep 1998) alongshore advection term (␷ a)incm studies along the west coast of the United States (e.g., sϪ2. (b) Horizontal time series at 200 m and 44.7ЊN for the same Neiburger et al. 1961; Elliot and O'Brien 1977; Burk term (Newport line). and Thompson 1996), consistent with the fact that the model tends to underestimate the altitude of the jet with respect to the Newport pro®ler wind. There are many coast fundamentally three-dimensional. We decompose similarities with the Burk and Thompson (1996) study this diurnal variation into four stages. of the low-level jet along the California coast, such as During the ®rst stage, early morning (1200±1500 the east±west slope of the marine boundary layer, the UTC), the low-level jet is relatively high and weak and occurrence of the maximum of low-level jet about 5 h the offshore ¯ow above the marine boundary layer is after the maximum of baroclinity, and the movement of

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the jet during its diurnal variation. However, there are also some differences. For example, the maximum of the low-level jet as well as the maximum in baroclinity occurs about 2 or 3 h earlier than in the Burk and Thompson study, which may be related to the three- dimensional nature of the circulation in the present case or it may also be an artifact of there being less large- scale baroclinity in the late afternoon without a hot inland valley like the San Joaquin valley in California. The thermodynamical structure during this 4-day period is consistent with a case study during August 1973 along the Oregon coast, presented by Elliot and O'Brien (1977), having a shallow marine layer where presumed subsidence created a tongue of drier air approximately 10±15 km offshore that extended almost down to the sea surface. A northerly wind minimum during early morning, such as is found here, was also noticed by Holt (1996) in his simulation over California during May 1990. He associated it with the channeling of the land breeze by terrain. This interpretation cannot work for our simulation because the terrain is different. Also, the minimum appears in our no-terrain simulation, though it is weaker. This shows that this minimum is tied to the diurnal cycle, which is stronger with the terrain that produces a larger variation in baroclinity. The early minimum occurs also when the cross-coast thermal gradient is the weakest. The present picture is different from the two-dimensional picture of the sea-breeze circulation that has pre- viously been invoked to explain observations of diurnal variability along the U.S. west coast (e.g., Beardsley et al. 1987; Banta et al. 1993). The large alongshore advection of alongshore momentum, which seems to be a response to the nonzero alongshore pressure gradient, plays an important role along the entire coast of Oregon, including the region north of Cape Blanco, where there are no major cape or headland features. Indeed, we have evidence, discussed elsewhere, that the interaction of hydraulically supercritical marine-layer ¯ow with coastal orography is signi®cant in the region adjacent to and immediately south of Cape Blanco, along the southern Oregon coast. However, orographic features of similar scale are not present along the central Oregon coast and we have not seen evidence of orographic effects of this type in the present calculations and observations. The importance of horizontal momentum advection in the diurnal cycle dynamics does not arise from effects of

←

FIG. 14. Schematic of the cross-coast structure (a) at 1500 UTC when the jet is the weakest, (b) at 0300 UTC when the jet is the strongest, and (c) alongshore structure at 0300 UTC. The alongshore

pressure gradient (␷ p) and alongshore advection (␷ a) are large in the hatched region. The hatched region is about 2Њ of latitude in the alongshore direction and mainly corresponds to the 4-km domain. The gradient of heating has a larger, several hundred km, alongshore scale, which is resolved in the 36-km domain simulation.

Unauthenticated | Downloaded 09/24/21 06:03 AM UTC 1008 MONTHLY WEATHER REVIEW VOLUME 130 this type. The alongshore pressure gradient and the as- Dorman, C. E., and C. D. Winant, 1995: Buoy observations of the sociated alongshore advection forces appear not to be atmosphere along the west coast of the United States, 1981± 1990. J. Geophys. Res., 100, 16 029±16 044. local perturbations. Indeed, they may arise from en- ÐÐ, D. P. Rogers, W. Nuss, and W. T. Thompson, 1999: Adjustment hanced heating over the northern California±southern of the summer marine boundary layer around Point Sur, Cali- Oregon mountains. It would be interesting to examine fornia. Mon. Wea. Rev., 127, 2143±2159. the diurnal circulation along the coast south of this re- ÐÐ, T. Holt, D. P. Rogers, and K. Edwards, 2000: Large-scale structure of the June-July 1996 marine boundary layer along Cali- gion, where the corresponding alongshore momentum fornia and Oregon. Mon. Wea. Rev., 128, 1632±1652. advection presumably must reverse. On the other hand, Elliott, D. L., and J. J. O'Brien, 1977: Observational studies of the there is evidence that the alongshore pressure gradient marine boundary layer over an upwelling region. Mon. Wea. propagates southward along the coast over the course Rev., 105, 86±98. of the diurnal cycle, as the maximum pressure gradient Holt, T. R., 1996: Mesoscale forcing of a boundary layer jet along the California coast. J. Geophys. Res., 101, 4235±4254. tends to occur earlier in the northern part of the domain Johnson, A., Jr., and J. J. O'Brien, 1973: A study of an Oregon sea than in the southern part, so another cause may also be breeze event. J. Appl. Meteor., 12, 1267±1283. possible. It will be necessary to address these issues, Kessler, E., 1969: On the Distribution and Continuity of Water Sub- among others, in order to achieve a comprehensive un- stance in Atmospheric Circulation. Meteor. Monogr., No. 32, Amer. Meteor. Soc., 84 pp. derstanding of the diurnal cycle of the lower atmosphere KoracÆin, D., and C. Dorman, 1999: Marine atmospheric boundary along the U.S. west coast. layer divergence and clouds along California in June 1996. Pre- prints, Third Conf. on Coastal Atmospheric and Oceanic Pre- Acknowledgments. This research was supported by diction and Processes, New Orleans, LA, Amer. Meteor. Soc., the National Oceanographic Partnership Program and 314±318. Meitin, R. J., and D. W. Stuart, 1977: The structure of the marine the Of®ce of Naval Research Grant N00014-98-1-0787. inversion in northwest Oregon during 26±30 August 1973. Mon. We are grateful to three anonymous reviewers for their Wea. Rev., 105, 748±761. thoughtful and constructive reviews, which resulted in Neiburger, M., D. S. Johnson, and C. Chien, 1961: Studies of the numerous improvements to the manuscript. Structure of the Atmosphere over the Eastern Paci®c Ocean in the Summer. 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McPeek, 1993: al Ocean Dynamics Experiment, Pt. 1, Description of the marine Quality controls for pro®ler measurements of winds and RASS boundary layer and atmospheric conditions over a northern Cal- temperatures. J. Atmos. Oceanic. Technol., 10, 452±464. ifornia upwelling region. J. Geophys. Res., 92, 1467±1488. Winant, C. D., C. E. Dorman, C. A. Friehe, and R. C. Beardsley, Bridger, A. F. C., W. C. Brick, and P. F. Lester, 1993: The structure 1988: The marine layer off Northern California: An example of of the marine inversion layer off the central California coast: supercritical channel ¯ow. J. Atmos. Sci., 45, 3588±3605. Mesoscale conditions. Mon. Wea. Rev., 121, 335±351. Xue, M., K. K. Droegemeier, V. Wong, A. Shapiro, and K. Brewster, Burk, S. D., and W. T. Thompson, 1996: The summertime low-level 1995: ARPS version 4.0 user's guide. Center for Analysis and jet and marine boundary layer structure along the California Prediction of Storms, 380 pp. [Available from CAPS, University coast. Mon. Wea. Rev., 124, 668±686. of Oklahoma, Norman, OK 73072.] ÐÐ, T. Haack, and R. M. Samelson, 1999: Mesoscale simulation of Zemba, J., and C. A. Friehe, 1987: The marine atmospheric boundary supercritical, subcritical, and transcritical ¯ow along coastal to- layer jet in the Coastal Ocean Dynamics Experiment. J. Geophys. pography. J. Atmos. Sci., 56, 2780±2795. Res., 92, 1489±1496.

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