Numerical Study of the Diurnal Cycle along the Central ... - CiteSeerX

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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 final 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 profiler 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 profiler data, there is a minimum in northerly wind between 1500 and 1800 UTC (0700 and 1000 local time), and a double maximum in offshore flow 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 influence on the diurnal circulation.

1. Introduction During spring and summer, the west coast of the United States is dominated by the Pacific high centered approximately 1000 km off the north California coast and a thermal low inland over California. This regime produces persistent northerly (upwelling favorable) winds along the coast interrupted by periods of weak or southerly flow. The typical summertime coastal meteorological conditions include a stable marine atmospheric boundary layer with northerly wind and a low-level jet near the top of the boundary layer. Most observational and numerical studies of this regime have focused on the California coast (e.g., Beardsley et al. 1987; Zemba and Friehe 1987; Winant et al. 1988; Bridger et al. 1993; Banta 1995; Dorman et al. 1999; Holt 1996; Burk and Thompson 1996; Burk et al. 1999, Korac˘in and Dorman 1999; Dorman et al. 2000). For instance, Zemba and Friehe (1987) showed that along the California coast, the marine atmospheric boundary layer during northerly winds was character* Current affiliation: NOAA/ETL, Boulder, Colorado. Corresponding author address: S. Bielli, Dept. of Atmospheric Sciences, University of Washington, Box 351640, Seattle, WA 98195. E-mail: [email protected]

q 2002 American Meteorological Society

ized by a low-level jet with maximum values of 30 m s 21 at a few hundred meters altitude. The vertical structure is characterized by an inversion near or at the altitude of the jet maximum. The wind shear was found to be due to thermal wind generated by the large horizontal temperature gradient between the water and the land. Some observations in the Oregon area have been discussed by Neiburger et al. (1961), Johnson and O’Brien (1973), Meitin and Stuart (1977), and Elliot and O’Brien (1977), and more recently by Dorman and Winant (1995). The research described here was carried out as part of the Oregon National Ocean Partnership Program (NOPP) project The Prediction of Wind-Driven Coastal Circulation. This project was undertaken with the recognition that understanding and modeling the coastal ocean requires improved knowledge of coastal atmospheric processes. The atmospheric component of the Oregon NOPP project had two specific objectives: first, to provide an estimate of air– sea fluxes for use in the ocean modeling component, and second, to study the dynamics of the coastal atmosphere along the Oregon coast, including particularly the diurnal cycle. The present contribution addresses the second of these two objectives. We focus primarily on mesoscale numerical model simulations centered on Newport for four consecutive days in September 1998. Section 2 presents

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briefly the model, its initialization, and a statistical comparison of modeled and observed fields during summer 1999, including results from a wind profiler stationed in Newport as part of the Oregon NOPP project. In section 3 we describe the meteorological situation for the September 1998 simulations. The results of the control simulation as well as comparison with wind profiler 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 final discussion and a summary in section 6. 2. Model description and initial conditions a. The ARPS model Numerical simulations were performed using the Advanced 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, nonhydrostatic, compressible version, with a terrain-following vertical coordinate. Only the warm microphysical processes are taken into account, using the Kessler (1969) parameterization. In all simulations, a stretched vertical coordinate 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 first point above ground level at 20 m; it progressively increased to an average of 450 m. The top of the model is at 13-km altitude, and a sponge layer is applied above 9 km to minimize the reflection of internal gravity waves. Increasing the average vertical grid size to 250 m or increasing the altitude of both the sponge layer or the top of the model did not result in any major differences during a 24-h simulation. Indeed, the sponge layer is relatively low, but the main activity is taking place in the low levels for this kind of event. Model physical parameterizations also included a 1.5-order turbulent kinetic energy (TKE), fourth-order advection in both horizontal directions, and radiation and surface flux schemes. The model topography was derived from the global terrain database with 30-s terrain resolution. The model was triply nested with one-way interaction on 60 3 60 grids with 36-, 12-, and 4-km resolution. Each domain was centered near Newport, Oregon (44.78N, 124.08W) (Fig. 1). b. Initial conditions and case study The 4-day period from 1200 UTC September 11 to 1200 UTC September 15 1998 was selected for this study. The flow along the Oregon coast experienced strong northerly wind during this period and was characteristic of a pattern frequently observed during summer months. The 36-km domain was initialized using data interpolated from the National Centers for Environmental Prediction (NCEP) ‘‘early’’ Eta Model output (grid 212, 40-km resolution). Time-dependent lateral boundary

FIG. 1. The 12- and 4-km domains with the terrain height and the position of the wind profiler near Newport (44.78N, 124.078W) and buoy 46050 (44.628N, 124.538W). The dotted line is the position of the cross sections through Newport.

conditions were imposed from Eta analyses, and the model was initialized every day at 1200 UTC with a cold start and was run for 24 h. For the 12- and 4-km domains, initial and time-dependent lateral boundary conditions were obtained, respectively, from the 36- and 12-km output in a one-way nesting procedure. A sensitivity study was conducted for 11 September 1998 to study the effect of the coastal orography, which consisted of removing the terrain but leaving the surface characteristics of land unaltered. c. Statistical wind profiler and buoy verification Before proceeding to the September 1998 case study, we briefly summarize a statistical comparison of modeled and observed variables during summer 1999. As part of the Oregon NOPP project, forecast-mode (36-h forecast with cold start at 0000 UTC each day) simulations were conducted during June through August 1999 with the ARPS model, on the 36- and 12-km domains. Statistical comparisons of summer 1999 modeled variables from the 12-km grid with National Data Buoy Center (NDBC) buoy 46050 (44.628N, 124.538W) and land-based meteorological (Coastal-Marine Automated Network, CMAN) observations at Newport (Fig. 1) are shown in Table 1. Statistics have been calculated based on hourly data and hours 4 to 27 have been used for the model. The model simulates pressure and northerly wind variability well, with high correlations between modeled and observed fields. The model shows a relatively good agreement with the buoy data for the mean 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, y 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

Pressure (Pa) Temperature (8C) u (m s21 ) v (m s21 )

Model

N

Mean

Std

Mean

Std

Corr

Bias

Rmse

1920 1931 1912 1912

1017.8 13.9 1.2 22.6

3.2 1.5 1.9 4.2

1017.3 14.8 2.0 23.2

3.2 2.1 1.9 3.8

0.94 0.66 0.51 0.82

20.56 0.95 0.81 20.68

1.23 1.87 2.04 2.52

NW03

Pressure (Pa) Temperature (8C) u (m s21 ) v (m s21 )

Model

N

Mean

Std

Mean

Std

Corr

Bias

Rmse

2031 2031 2031 2031

1018.4 12.7 0.8 21.2

3.1 1.8 1.5 3.9

1013.9 14.4 1.5 21.8

3.2 2.9 1.7 2.6

0.94 0.56 0.61 0.75

24.48 1.71 0.67 20.60

4.62 3.02 1.55 2.65

be in part due to the presence of unresolved terrain. The standard deviations are comparable in all cases. The largest differences are for alongshore wind and temperature at Newport. The warm bias in the model temperature field at both locations may be due to ocean upwelling that is not resolved by the Eta sea surface temperature analysis (Samelson et al. 2002). A 915-MHz Radio Acoustic Sounding System (RASS) profiler was in place at Newport during this period, supported by the Oregon NOPP project. Hourly consensus-averaged winds available from the National Oceanic and Atmospheric Administration/Environmental Technology Laboratory (NOAA/ETL) (www7.etl. noaa.gov/data/archive/realtime) were used to produce the wind composites. Details on wind profiler processing and accuracy can be found in Weber et al. (1993). Also Ralph et al. (2000) demonstrated the capabilities of such radars in West Coast summertime studies. Wind data were first linearly interpolated to constant height and then hourly averaged over the summer. The hourly mean diurnal cycles of the Newport profiler and model winds for the summer 1999 are shown in Fig. 2. A measurable diurnal cycle is present in both of these summer 1999 means, which include synoptic disturbances, episodes of southerly flow, and other periods when the diurnal cycle is not well developed. The timing and the depth of these mean modeled and observed diurnal variations are in rough agreement. However, the model overestimates the strength of the wind with maximum (minimum) values of 7.25 (0.50) m s 21 for the model and 4.75 (0.75) m s 21 for the observations, respectively. Also note that the jet maximum observed at 100–300-m altitude (vs 250-m altitude in the model) occurs 3 h late, and the relatively strong winds near 500 m persist several hours longer in the observations than in the model. Finally, the wind minimum observed at 800–1100-m altitude and 1700–2100 UTC does not appear in the model. Some of these differences in the mean summer 1999 fields may be due to relatively poor per-

formance of the model during southerly flow episodes. The comparison between model and profiler winds will be explored further later [section 4c(3)] for the 4-day period of interest in September 1998, for which the similarity of modeled and observed diurnal variations is greater than in these summer 1999 means. The results in Fig. 2 demonstrate that the diurnal cycle is an important component of the summertime lower-atmosphere circulation along the Oregon coast, and motivate the more detailed analysis that follows. 3. Synoptic situation We focus on a case study for four days 11–15 September 1998, during which a well-developed diurnal cycle was observed. The synoptic pattern during this period was fairly typical for the summertime along the west coast of the United States, with the Pacific high located approximately 1000 km off Cape Mendocino, a thermal low over California, and north-northwesterly flow along the coast. The sea level pressure and 500hPa analyses at 0000 UTC (1700 local time) for 12 and 15 September 1998 from NCEP Eta data, interpolated to the model grid, are displayed in Fig. 3. At 500 hPa, the Pacific Northwest is initially under the influence of a large-scale ridge modified by the presence of a low pressure system centered over the California–Nevada border (Fig. 3b). At 500 hPa, to the north of Oregon a strong westerly jet extending into British Columbia and an area of anticyclonic flow produces northerly wind over the coasts of Washington, Oregon, and northern California. An area of high sea level pressure builds in from the southwest with lower pressure over inland areas of Nevada, California, and eastern Oregon (Fig. 3a). The upper-level closed circulation feature slowly moves south and then begins to move east and exit the study region on 1200 UTC 12 September. Higher pressure builds in its wake over Oregon and the strong westerly flow 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 Pacific. This results in a slow rotation and strengthening of the winds over Oregon. By 0000 UTC 15 September, the winds at 500 hPa are from the southwest over most of Oregon and the Pacific 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 figures 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 s 21 (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 field 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 s 21 . 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 s 21 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 s 21 , 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 profiler. Values are in m s 21 with contour interval of 0.5 m s 21 . The scale for the wind vectors is beneath.

the north dips southward while the thermal low in central California develops and strengthens. By 1300 UTC 13 September, the upper-level winds over the central Oregon coast have a mostly westerly component. Over

1) DYNAMICAL

STRUCTURE

In this section, we summarize the mesoscale structure and variations of the 12-km model wind field in a west– east line crossing Newport and at the Newport profiler location. In the 12-km model, the peak alongshore velocity in the low-level jet reaches 16 m s 21 at 200 m and 124.58W (Fig. 5a). The diurnal variation of the wind speed near Newport is clearly apparent in cross sections 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 21 ) 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 21 ) 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 profiler near Newport.

the day, the wind speed intensifies from 12 to 16 m s 21 as it moves toward the coast, with a maximum value at 0300 UTC. At 200 m, the jet core is located at about 100 km off the coast in the early morning and moves to about 50 km offshore during the day, when the wind speed reaches its maximum (Fig. 5b). The diurnal variation extends over the entire domain, but the largest variations are confined within the first 200 km offshore (Fig. 5b). Note also the wind speed minimum between 1500 and 1800 UTC (Figs. 5b,c). This minimum is not

an artifact of the initialization; it appears also in a simulation initialized at 0900 UTC instead of 1200 UTC (not shown). The axis of maximum wind speed rises from 200 m near the coast up to 400 m at about 400 km offshore. A small secondary wind maximum appears above the coastal mountains and can be seen over the valley by 0300 UTC (Figs. 5a,b). The jet core seems to lower slightly down to 200 m at Newport during the morning 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 1288W (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 38C 100 m 21 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: STRUCTURE

FIG. 4. (a) Standard deviation of the wind speed in m s 21 with contour interval of 0.5 m s 21 and (b) 4-day mean (11–15 Sep 1998) wind speed at 200-m altitude over the 36-km domain based on hourly output in m s 21 with contour interval of 2 m s 21 . The dotted line shows the position of the baseline for the cross-coast cross sections along the Newport line.

VERTICAL

AND HORIZONTAL

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 profiler 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 (y , meridional) 12-km wind fields 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 seabreeze flow 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 s 21 between 2100 and 0000 UTC (1400–1700 local time), with a maximum depth of the flow of about 200 m over the water. At the surface, onshore flow 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 simulation period, except when the sea-breeze circulation is present. At 0300 UTC, the return flow associated with the sea-breeze circulation can be separated into two components: above the coastal range at 600–800 m, and 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 u in K with contour interval of 0.5 K and (b) water vapor content q y in g kg 21 with contour interval of 0.5 g kg 21 .

that the magnitude of the wind speed is slightly weaker, and the onshore sea-breeze flow 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 ← 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 44.78N (line through Newport; cf. Fig. 1), (b) time–longitude plot at 200 m altitude and 44.78N, and (c) vertical cross section at Newport. Values are in m s 21 with contour interval of 1 m s 21 .

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FIG. 7. The 4-km domain: 4-day hourly mean (11–15 Sep 1998) cross-shore wind (u) in m s 21 with contour interval of 0.5 m s 21 . (a) Time–longitude cross-section at 200-m altitude and 44.78N (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 (y ) in m s 21 with contour interval of 1 m s 21 .

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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 s 21 with contour interval of 0.5 m s 21 , (b) y in m s 21 with contour interval of 1 m s 21 , (c) u in K with contour interval of 1 K, and (d) q y in g kg 21 with contour interval of 0.5 g kg 21

domain at the height of the jet core, with a steady southward increase of several meters per second in northerly and offshore flow (Figs. 7e,f). This gradient is the mesoscale expression of the large-scale gradient that is evident in Fig. 4b. As shown below, it is large enough to affect the horizontal momentum balance. 2) MODEL

VERTICAL PROFILES AT

NEWPORT

At Newport, the northerly wind at 200 m and below has a minimum between 1500 and 1800 UTC, associated with increasing onshore flow (Figs. 8a,b). This minimum occurs progressively later aloft, in phase with decreasing offshore flow aloft. The maximum offshore flow aloft is coincident with the appearance of the sea

breeze at the surface. The boundary mixed layer lowers until 2100 UTC, associated with an increase in northerly wind and an increase of onshore flow (Fig. 8c). Then the marine boundary layer deepens as the onshore flow builds up, and the onshore wind increases to a maximum between 2100 and 0000 UTC in correlation with temperature and water vapor content maxima. Three hours later, the northerly wind attains its maximum whereas the offshore flow aloft (above 400 m) reaches a minimum. Finally, the northerly wind decreases and the offshore flow aloft rises again. Above the marine boundary layer, the offshore flow shows two maxima: around 600 m between 1500 and 1800 UTC, and around 400 m between 0600 and 0900 UTC. These two maxima are not specific to Newport:

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time profiles of the cross-shore wind farther north or south of Newport near the coast (not shown) display similar features, with two minima of different magnitudes 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 flow maxima are correlated with relatively warm temperatures and low humidities. 3) COMPARISONS

WITH WIND PROFILER DATA

The model profiles at Newport discussed above may be compared with 4-day hourly mean horizontal winds for 11–15 September 1998 from the Newport RASS wind profiler (Fig. 9). Surface data from the CMAN station have also been included in the wind profiler 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 flow 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 flow reverses. The observations show deep onshore flow between 2300 and 0100 UTC, reaching to 1100 m. The model shows a deep weakening of offshore flow during this period, but no flow 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 influence on the diurnal cycle and especially the cross-shore flow; 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 flow 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 identified 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 every day at about the same time, so they are apparently phase-locked to the diurnal forcing. The maximum around 1500 UTC is associated with the weakest crosscoast potential temperature gradient, whereas the maximum around 0600 UTC is associated with a large temporal decrease of the cross-coast potential temperature gradient. Attempts to understand this oscillation in terms of simple linear, forced–damped models were inconclusive, although there were consistent indications that inertial effects did contribute. We discuss in more detail

FIG. 9. Time series at Newport from a composite of wind profiler and CMAN surface observations for 11–15 Sep 1998 based on hourly data of (a) 4-day hourly mean cross-shore wind u in m s 21 with contour interval of 0.5 m s 21 , and (b) 4-day hourly mean alongshore wind y in m s 21 with contour interval of 1 m s 21 .

the dynamics of the cross-shore wind above the marine boundary layer in sections 5 and 6. The observed alongshore wind reaches a minimum of northerly flow around 5 m s 21 near 1800 UTC, then increases to 13 m s 21 around 0300 UTC at 350-m al-

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titude before decreasing again (Fig. 9b). The model tends to overestimate the wind speed and underestimate the altitude of the jet core. The observed hourly mean maximum of northerly wind is 4 m s 21 weaker than the modeled maximum. The difference in altitude of the maximum is roughly 150 m, but strong diurnal variation of the northerly wind occurs throughout the lower 1000 m in both the model and observations. The pattern of the summer 1999 alongshore wind is very similar to the 4-day mean pattern, with northerly wind maximum and minimum, respectively, around 0100 and 1500 UTC (Fig. 2). The 4-day and summer 1999 mean profiler observations are generally similar to a 1-month composite of the alongshore wind at Piedras Blancas, California, presented by Ralph et al. (2000), who find large diurnal variations in alongshore wind up to 1200 m, with minimum and maximum northerly winds around 1700 and 0200 UTC, respectively. The 7.5 m s 21 amplitude of diurnal variation of alongshore wind in the 1-month mean at Piedras Blancas lies between that of the present 4-day and the summer 1999 means. There appears to be again some evidence of a double maximum in offshore flow in the Piedras Blancas composite, but it occurs above 1500 m and may have a different origin. The higher level of this feature might be due in part to the higher mountains at Piedras Blancas.

to the coastal zone in the control simulation is controlled primarily by the topography, and not, for example, by interactions with the low-level alongshore jet. This confinement evidently intensifies the diurnal variations in the coastal zone. 5. Dynamics of the diurnal cycle: Momentum balances In this section, we analyze the dynamics of the model diurnal cycle. We focus on the momentum balances for both the cross-shore and the alongshore circulation. One objective of this analysis is to assess the extent to which the diurnal cycle can be understood in terms as an essentially two-dimensional circulation, in which alongshore variations are negligible, as might be anticipated from the anisotropic geometry and forcing that typically characterizes coastal zones, or whether it is intrinsically three-dimensional. The terms in the equations of motion are diagnosed directly from the 4-km grid ARPS model using the following form:

|

]u 1 ]P 5 2V · =u 2 1 f y 1 Fx ]t r ]x z

|

|

ut

d. No-terrain simulation The simulation conducted for 11 September 1998 with the terrain removed shows a number of significant differences from the control that illustrate the effect of terrain on the lower-atmospheric circulation. The initialization procedure consisted of interpolating the NCEP Eta data onto the ARPS grid, applying a smoothing, and adjusting the fields so that the anelastic mass continuity equation is satisfied. Without terrain, the position and the shape of the 36-km wind speed maximum are modified with respect to the control simulation. Rather than a single peak centered just south of Cape Blanco, two slightly weaker peaks are apparent in the no-terrain simulation: one south of Cape Blanco and the second south of Cape Mendocino, both with a maximum value of about 15 m s 21 . The shifting of the wind speed maximum to the south is more evident at the surface. Thus, the position of the maximum of wind speed is influenced by the terrain, but the existence of the wind speed maximum is evidently more related to the Pacific high and the continental thermal low. Without terrain, the diurnal variation still extends over the 4-km domain but is weaker. The minimum in northerly wind is still present without terrain, but it appears a few hours later than in the control simulation. An especially striking difference is the character of the sea breeze. In the no-terrain case, the sea breeze propagates onshore more than 100 km, in the classical manner. This suggests that the confinement of the sea breeze

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z

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ua

z

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up

z

|

|

uc

z

(5.1) |

um

]y 1 ]P 5 2V · =y 2 2 f u 1 Fy , ]t r ]y z

yt

|

|

z

ya

|

|

z

|

yp

|

z

yc

|

|

z

(5.2)

|

ym

where the x axis is taken to be the cross-shore direction (west–east) and y the alongshore direction (north– south). For X 5 (u, y ), the terms are X t , the local tendency term; X a , the advection term; X p , the pressure gradient term; X m , the friction (mixing) term; and X c , the Coriolis term. The horizontal and vertical structure of the main terms are discussed here as well as diurnal variations at 200 m, height of the low-level jet. a. Cross-shore momentum balance The cross-shore circulation is driven by diurnal fluctuations in the cross-shore pressure gradient, which arise from contrasts in diurnal heating over land and sea, as in a classical sea-breeze circulation. These gradients drive an ageostrophic cross-shore circulation, with return flow aloft. This circulation is similar to that found along the California coast (e.g., Zemba and Friehe 1987; Beardsley et al. 1987). At 200 m, the height of the low-level jet, the crossshore momentum budget at the coast is nearly geostrophic (Fig. 10a). Indeed, the Coriolis term u c , which reflects the variation of the alongshore wind component and shows a definite diurnal variation, is largely balanced by the cross-shore pressure gradient. The pressure 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 crossshore equation in cm s 22 based on hourly output at (a) Newport and (b) 0.58 of longitude offshore, and for the alongshore equation at (c) Newport and (d) 0.58 of longitude offshore. Here (u, y )a is the advection term, (u, y )c the Coriolis term (u, y ) the pressure gradient term, (u, y )m the mixing term, and (u, y )t the local tendency term.

evening, when the sea-breeze circulation is present and the northerly wind is maximum. Both u p and u c minimum values occur between 1500 and 1800 UTC, correlated with the northerly wind minimum and the increase of the onshore flow near the coast. Advection plays a role during late afternoon and early evening, beginning after 1800 UTC, when the onshore flow in the low levels is established, with two peaks: one around 1800 UTC when the onshore flow is increasing, and the other around 0300 UTC when the onshore flow is decreasing. The cross-shore advection offsets some of the pressure gradient increase so that the alongshore wind is not geostrophically balanced during the see breeze but basically quasigeostrophic. The cross-shore flow is a thermally driven sea breeze, which is not in geostrophic balance and tends to turn the total flow onshore. As vertical advection is small, we can write the advection term as

u a 5 2u |

]u ]u 2 y . ]x ]y

z

uux

|

|

z

(5.3)

|

y uy

At 200 m, the minimum in u a around 0000 UTC is associated with a decrease in yuy that results primarily from a decrease of the alongshore gradient of the crossshore wind (uy). The tendency and the friction terms are negligible at 200-m altitude. Half a degree offshore, the flow is near geostrophic, u a is small and does not vary much throughout the day (Fig. 10b), and u p attains its maximum (negative) value about 6 h earlier than at the coast. The vertical distributions of these quantities at 0300 UTC in Newport and 0.58 longitude offshore are displayed in Fig. 11. The maximum in u p occurs at the jet core height (200 m), beneath which the flow is slowly

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decelerating. At the coast, u a plays a significant role up to 600-m altitude, with a maximum at 200 m for the control simulation (Fig. 11a). Half a degree offshore, the flow 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 u p is smaller in the no-terrain case and u a and u t 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

FIG. 11. Vertical profile of 4-day hourly mean (11–15 Sep 1998) of the different terms of the momentum equation cm s 22 for the crossshore budget: (a) in Newport and (b) 0.58 of longitude offshore for the control.

It might be anticipated from previous work along the California coast (e.g., Zemba and Friehe 1987; Beardsley et al. 1987) that the interaction of the alongshore jet and the diurnal cross-shore circulation would lead to a fluctuation of the alongshore flow that is essentially two-dimensional in character, with the alongshore flow responding primarily to advection and Coriolis forces associated with the cross-shore flow. Surprisingly, this turns out not to be the case. Instead, there are large diurnal fluctuations in the alongshore pressure gradient, which evidently arise from intensified 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 first 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 balancing y p at 200-m height during the morning hours (1800– 2100 UTC) is y t . At 1600 UTC, y 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 flow strengthens, the alongshore flow would tend to decelerate due to crossshore turning but it actually accelerates due to y a and y t . Friction and Coriolis terms are negligible at the coast. Geostrophic balance is never reached. Half a degree offshore, y c is more comparable to y a , but y a is still the largest term balancing y 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 boundary layer, the alongshore pressure gradient is balanced by advection, acceleration, and friction terms 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 flow 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 significant role in the circulation; thus, the circulation is fully three-dimensional at least in the marine boundary layer. The advection term is also significant 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 noterrain simulation is very similar to the picture 0.58 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 (y 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 flow 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 43.78 and 45.78N 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

FIG. 12. Same as in Fig. 13 but for the alongshore momentum equation.

the tendency term is as large as the advection term. The friction term displays two maxima, at the height of maximum wind speed and at the surface. We decompose the advection term for the alongshore momentum budget into two terms uvx and vvy in the

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 is a thermally driven circulation that arises from contrasts in diurnal heating over land and sea as in a classical sea-breeze circulation. The large alongshore pressure gradient variations arise from stronger heating over the southern Oregon and northern California mountains, relative to northern Oregon and southern Washington; they drive an ageostrophic alongshore response and make the diurnal circulation along the central Oregon

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FIG. 13. (a) Vertical (m) cross section at Newport of the 4-day hourly mean (11–15 Sep 1998) alongshore advection term (y a ) in cm s 22 . (b) Horizontal time series at 200 m and 44.78N for the same term (Newport line).

coast fundamentally three-dimensional. We decompose this diurnal variation into four stages. During the first stage, early morning (1200–1500 UTC), the low-level jet is relatively high and weak and the offshore flow above the marine boundary layer is

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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 flow above the marine boundary layer is maximum and the jet is the weakest (Fig. 14a). As the sun rises, the temperature starts to increase. During 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 crossshore 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 y and vy play a significant role in the increase of the alongshore advection. The circulation is fully three-dimensional up to 400-m altitude, as far as 18 of longitude offshore, and for at least several hundred kilometers alongshore. Above the marine boundary layer, a minimum of offshore flow 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 flow 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 flow aloft decreases and reaches a maximum value. This maximum in offshore flow 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 confined to the coast. The low-level jet behavior near Newport is similar to the jet described by Burk and Thompson (1996) for the California coast during summertime. The altitude of the model jet core is lower than the altitude reported in other studies along the west coast of the United States (e.g., Neiburger et al. 1961; Elliot and O’Brien 1977; Burk and Thompson 1996), consistent with the fact that the model tends to underestimate the altitude of the jet with respect to the Newport profiler wind. There are many similarities with the Burk and Thompson (1996) study of the low-level jet along the California coast, such as the east–west slope of the marine boundary layer, the occurrence of the maximum of low-level jet about 5 h 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 threedimensional nature of the circulation in the present case or it may also be an artifact of there being less largescale 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 previously 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 flow with coastal orography is significant 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 (y p ) and alongshore advection (y a ) are large in the hatched region. The hatched region is about 28 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.

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this type. The alongshore pressure gradient and the associated alongshore advection forces appear not to be local perturbations. Indeed, they may arise from enhanced heating over the northern California–southern Oregon mountains. It would be interesting to examine the diurnal circulation along the coast south of this region, where the corresponding alongshore momentum advection presumably must reverse. On the other hand, there is evidence that the alongshore pressure gradient propagates southward along the coast over the course of the diurnal cycle, as the maximum pressure gradient tends to occur earlier in the northern part of the domain than in the southern part, so another cause may also be possible. It will be necessary to address these issues, among others, in order to achieve a comprehensive understanding of the diurnal cycle of the lower atmosphere along the U.S. west coast. Acknowledgments. This research was supported by the National Oceanographic Partnership Program and the Office of Naval Research Grant N00014-98-1-0787. We are grateful to three anonymous reviewers for their thoughtful and constructive reviews, which resulted in numerous improvements to the manuscript. REFERENCES Banta, R. M., 1995: Sea breeze shallow and deep on the California coast. Mon. Wea. Rev., 123, 3614–3622. ——, L. D. Olivier, and D. H. Levinson, 1993: Evolution of the Monterey Bay sea-breeze layer as observed by pulsed Doppler lidar. J. Atmos. Sci., 50, 3959–3982. Beardsley, R. C., C. E. Dorman, C. A. Friehe, L. K. Rosenfeld, and C. D. Winant, 1987: Local atmospheric forcing during the Coastal Ocean Dynamics Experiment, Pt. 1, Description of the marine boundary layer and atmospheric conditions over a northern California upwelling region. J. Geophys. Res., 92, 1467–1488. Bridger, A. F. C., W. C. Brick, and P. F. Lester, 1993: The structure of the marine inversion layer off the central California coast: Mesoscale conditions. Mon. Wea. Rev., 121, 335–351. Burk, S. D., and W. T. Thompson, 1996: The summertime low-level jet and marine boundary layer structure along the California coast. Mon. Wea. Rev., 124, 668–686. ——, T. Haack, and R. M. Samelson, 1999: Mesoscale simulation of supercritical, subcritical, and transcritical flow along coastal topography. J. Atmos. Sci., 56, 2780–2795.

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Dorman, C. E., and C. D. Winant, 1995: Buoy observations of the atmosphere along the west coast of the United States, 1981– 1990. J. Geophys. Res., 100, 16 029–16 044. ——, D. P. Rogers, W. Nuss, and W. T. Thompson, 1999: Adjustment of the summer marine boundary layer around Point Sur, California. Mon. Wea. Rev., 127, 2143–2159. ——, T. Holt, D. P. Rogers, and K. Edwards, 2000: Large-scale structure of the June-July 1996 marine boundary layer along California and Oregon. Mon. Wea. Rev., 128, 1632–1652. Elliott, D. L., and J. J. O’Brien, 1977: Observational studies of the marine boundary layer over an upwelling region. Mon. Wea. Rev., 105, 86–98. Holt, T. R., 1996: Mesoscale forcing of a boundary layer jet along the California coast. J. Geophys. Res., 101, 4235–4254. Johnson, A., Jr., and J. J. O’Brien, 1973: A study of an Oregon sea breeze event. J. Appl. Meteor., 12, 1267–1283. Kessler, E., 1969: On the Distribution and Continuity of Water Substance in Atmospheric Circulation. Meteor. Monogr., No. 32, Amer. Meteor. Soc., 84 pp. Korac˘in, D., and C. Dorman, 1999: Marine atmospheric boundary layer divergence and clouds along California in June 1996. Preprints, Third Conf. on Coastal Atmospheric and Oceanic Prediction and Processes, New Orleans, LA, Amer. Meteor. Soc., 314–318. Meitin, R. J., and D. W. Stuart, 1977: The structure of the marine inversion in northwest Oregon during 26–30 August 1973. Mon. Wea. Rev., 105, 748–761. Neiburger, M., D. S. Johnson, and C. Chien, 1961: Studies of the Structure of the Atmosphere over the Eastern Pacific Ocean in the Summer. University of California Publications in Meteorology, Vol. 1, No. 1, 1–94. Ralph, F. M., P. J. Neiman, P. O. G. Persson, J. M. Bane, M. L. Cancillo, J. M. Wilczak, and W. Nuss, 2000: Kelvin waves and internal bores in the marine boundary layer inversion and their relationship to coastally trapped wind reversals. Mon. Wea. Rev., 128, 283–300. Samelson, R., and Coauthors, 2002: Wind stress forcing of the Oregon coastal ocean during the 1999 upwelling season. J. Geophys. Res., in press. Weber, B. L., D. B. Wuertz, D. C. Welsh, and R. McPeek, 1993: Quality controls for profiler measurements of winds and RASS temperatures. J. Atmos. Oceanic. Technol., 10, 452–464. Winant, C. D., C. E. Dorman, C. A. Friehe, and R. C. Beardsley, 1988: The marine layer off Northern California: An example of supercritical channel flow. J. Atmos. Sci., 45, 3588–3605. Xue, M., K. K. Droegemeier, V. Wong, A. Shapiro, and K. Brewster, 1995: ARPS version 4.0 user’s guide. Center for Analysis and Prediction of Storms, 380 pp. [Available from CAPS, University of Oklahoma, Norman, OK 73072.] Zemba, J., and C. A. Friehe, 1987: The marine atmospheric boundary layer jet in the Coastal Ocean Dynamics Experiment. J. Geophys. Res., 92, 1489–1496.