Dynamics of a Catalina Eddy Revealed by Numerical Simulation

VOLUME 128 MONTHLY WEATHER REVIEW AUGUST 2000

CHRISTOPHER DAVIS AND SIMON LOW-NAM National Center for Atmospheric Research,* Boulder, Colorado

CLIFFORD MASS University of Washington, Seattle, Washington

(Manuscript received 27 April 1999, in ®nal form 20 December 1999)

ABSTRACT Through numerical simulations with the Pennsylvania State University±NCAR Mesoscale Model the dynamics of a Catalina Eddy event that formed during the period 26±30 June 1988 off the coast of southern California is examined. A strengthening and veering of the low-level synoptic-scale winds, from climatological northwesterlies to northerlies, results in a more pronounced effect of the coastal orography around the California bight region. In particular, relative vorticity formed by ¯ow over the coastal terrain remains offshore. Prior to the formation of a quasi-steady eddy within the bight, the northerlies are strong enough to advect anomalously high vorticity out of the region. The formation of a mature Catalina eddy relies on a rapid deceleration of the synoptic- scale northerlies on 30 June, such that vorticity, once formed, remains in the bight. The eddy is also strongly modulated by the diurnal cycle. Northwesterly ¯ow around 500 m above mean sea level impinging on the mountains north of the bight is enhanced during the late afternoon, mainly as a response to the land±sea thermal contrast. This strengthened ¯ow overlaps temporally with a minimum in low-level strati®cation due to surface heating. The result is air characterized by a relatively high Froude number, which traverses over the coastal mountains and strongly depresses the marine layer over the bight. The depression in the marine layer results in a warm anomaly and cyclonic circulation. Later at night, the incident northwesterlies weaken and the ¯ow becomes more stable, resulting in ¯ow around, rather than over, the coastal mountains. This regime transition yields a wake with little depression of the marine layer and an absence of vorticity generation on the scale of the bight region. Given strong ambient ¯ow, vorticity generated in the evening is swept southward past the bight the following day, but with weak ambient ¯ow, the eddy persists in the bight during daytime, weakening slowly.

1. Motivation and into the Los Angeles basin, an unmistakable sig- nature in visible satellite imagery (Rosenthal 1968). Pre- Disruptions in the climatological ¯ow along the coast- diction of the onset of overcast conditions and deep- line of southern California associated with the Catalina ening of the marine layer on the California coast rep- eddy have long been known to be important for air resents a signi®cant challenge to local forecasters during quality in the Los Angeles basin, as they are associated the otherwise benign warm season. with pronounced increases in the depth of the marine The climatology of eddy events presented in Mass layer (Wakimoto 1987; Thompson et al. 1997). The Cat- and Albright (1989, hereafter MA) demonstrates the im- alina eddy is a quasistationary mesoscale vortex with a portance of changes on the synoptic scale. Following horizontal scale of roughly 100 km, extending through the passage of a low pressure trough into the Paci®c a layer from the surface to between 1 and 2 km above Northwest, the winds near 850 hPa become stronger and mean sea level (MSL). Eddy events are often accom- veer to a northerly or northeasterly direction as the trail- panied by stratus spreading northward along the coast ing anticyclone moves north of the bight region. Cy- clonic circulation is produced over the bight in the lee of the San Rafael Mountains (Fig. 1), and this circulation * The National Center for Atmospheric Research is sponsored by appears to expand slowly with time. According to MA, the National Science Foundation. lee troughing reverses the alongshore pressure gradient, forcing coastally trapped southerly ¯ow from roughly the U.S.±Mexico border to the Los Angeles basin. Ob- Corresponding author address: Christopher A. Davis, National Center for Atmospheric Research, P.O.Box 3000, Boulder, CO 80307- servational case studies by Bosart (1983), Wakimoto 3000. (1987), and MA, as well as modeling studies by Ueyoshi E-mail: [email protected] and Roads (1993, hereafter UR), Ulrickson et al. (1995),

᭧ 2000 American Meteorological Society 2885

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FIG. 1. Locations of the three domains used in our simulations. Inset ®gure depicts the terrain on domain 3 (400-m contour interval) and the locations of various cross sections shown in the paper. Gray box indicates area of averaging to produce time±height cross section in Fig. 7. Gray dot marks point at which time±height cross sections is produced in Fig. 8. and Thompson et al. (1997), all describe cases that gen- is, ``What generates the lee troughing and cyclonic vor- erally ®t the climatology. ticity in the ®rst place?'' It may appear from the literature that the Catalina Although it is apparent that the eddy results from ¯ow eddy is a well-studied phenomenon, but there are some over/around complex terrain, ¯ows in the region in important issues regarding its dynamics that remain. which the eddy forms differs from the uniform ¯ows Most of the previous studies (e.g., Clark 1994) have that have been studied extensively, and about which emphasized the formation of coastally trapped south- most of our knowledge of mesoscale lee vortices and erlies as a de®ning feature of the eddy. However, these lee troughs is based (cf. Smith 1979; Boyer et al. 1987; southerlies depend in part on lee troughing and asso- Smolarkiewicz and Rotunno 1989). For instance, the ciated generation of mesoscale cyclonic vorticity in the strong baroclinicity implied by the contrast between a bight region. Therefore, perhaps the most basic question warm continent and a cool ocean implies strong low-

Unauthenticated | Downloaded 09/25/21 12:38 PM UTC AUGUST 2000 DAVIS ET AL. 2887 level shear. Second, the presence of a marine inversion passing through the Paci®c Northwest, followed by ridg- layer yields a complicated pro®le of static stability. ing in the lower troposphere. As the anticyclone moves Third, the diurnal cycle amplitude over the continent is over the West Coast region, enhanced northerly ¯ow is maximum in the warm season, when eddies are most noted, particularly at Vandenberg Air Force Base (see frequent. The baroclinicity, shear, and stability vary di- Fig. 25 of MA). Just after this time a mature cyclonic urnally implying that the terrain-induced response will eddy is observed in the bight region of southern Cali- vary similarly. fornia. The formation of a quasistationary eddy seems The Catalina eddy is strongly modulated by the di- to coincide with an overall weakening of the synoptic- urnal cycle. Studies of different cases (Bosart 1983; scale northerly ¯ow as the anticyclone moves to the MA; Clark 1994; Thompson et al. 1997) all show an northeast of the region on 30 June. eddy strengthening in the early evening, maturing over- Mesoscale surface analyses (refer to Fig. 22 in MA) night, then weakening the following day. This has been depict an event that requires about four days to unfold. interpreted as an effect of the sea breeze, which disrupts Each day, on average, the eddy appears slightly stronger the eddy circulation during the day. However, UR noted and more extensive than the previous day. However, from their simulation of the 26±30 June 1988 eddy case there is a diurnal variation in the intensity, with sys- that vorticity through a layer roughly 1±2 km deep cen- tematically stronger circulation at night, at least until tered at 950 mb over the bight was systematically larger 30 June when a closed cyclonic circulation persists dur- at night. The diurnal modulation of vorticity aloft is not ing the day as well. During most of this 4-day period, likely the result of the sea-breeze circulation because the eddy appears as a trough elongated northwest to its depth and its horizontal scale (roughly 100 km) great- southeast. At night, southeasterly ¯ow prevails along ly exceed that of the classical sea-breeze circulation. the coastline and somewhat offshore. The ¯ow near and The mechanism responsible for the variation in vorticity slightly inland from the coast appears related to oro- aloft has yet to be determined. graphic blocking by the coastal mountains south of the Because of the relative sparseness of observations, Los Angeles basin. The southeasterlies farther offshore much of the present study will encompass a diagnosis are more nearly geostrophic. To the west of the trough, of the eddy evolution produced by a numerical simu- strong northwesterly ¯ow (10±15 m sϪ1) prevails until lation. Previous modeling studies of this case with high- 30 June. Typical values of relative vorticity associated ly simpli®ed physics and relatively coarse resolution with the trough and subsequent more circular eddy are models have shown reasonable success in simulating on the order of 2±3 ϫ 10Ϫ4 sϪ1. the general features of the Catalina eddy. Our approach The vertical structure of the eddy is largely unknown, will be one of case study, the case being the eddy of owing to the poor sounding coverage offshore. Based 26±30 June 1988, studied by MA, Ueyoshi and Roads on time±height sections of vorticity from their numerical (1993), and Ulrickson et al. (1995). This choice is eco- simulations, which selectively ®ltered out the larger- nomical from the point of presentation, since much of scale evolution, UR revealed the eddy to be about 2 km the observations have been documented in MA and will deep. Thus, the eddy itself appears to be rather shallow, hence only be brie¯y summarized here. Another reason not extending above the peaks of the San Rafael and for selecting this case is that it appears to closely mirror Santa Ynez Mountains. the climatological behavior of eddy events, thus pos- sibly allowing greater generalization of our results. The primary objective of this study is to document 3. Mesoscale model simulation the spatial structure and temporal evolution of the June 1988 eddy event using numerical simulations. We will a. Model speci®cs focus on the source of cyclonic vorticity that comprises the eddy and how this source depends on the diurnal When attempting to use a mesoscale model as a proxy cycle. We will show how the formation of the eddy is for observations, it is always a matter of contention strongly tied to whether air originating over the Central whether a simulation adequately reproduces the obser- Valley of California can traverse the coastal mountains vations and realistically mimics physical processes in to the south and depress the marine inversion layer over the real atmosphere. Because the data are often sparsely the bight region. This happens only during the later distributed, the answer to this question cannot fully be afternoon and evening when the low-level northerlies known. However, it has been repeatedly demonstrated strengthen and the strati®cation is weak upstream from that mesoscale models excel in problems of ¯ow near the coastal terrain. complex terrain (e.g., Seaman et al. 1995; Colle and Mass 2000), primarily because the forcing is known (and stationary). A simpli®cation in the present case is 2. June 1988 case: Overview the relatively minor role played by moist physical pro- As mentioned above, the case of 26±30 June 1988 is cesses. A bene®t of using the model in this instance is well documented in the literature (MA; UR; Ulrickson that we can attempt to isolate various physical processes, et al. 1995). The case features a synoptic-scale trough such as the effect of the diurnal cycle, which has already

Unauthenticated | Downloaded 09/25/21 12:38 PM UTC 2888 MONTHLY WEATHER REVIEW VOLUME 128 been shown to exert a signi®cant in¯uence on eddy ary layer and the Dudhia radiation and simple ice formation. schemes (Grell et al. 1994). Numerics consist of a leap- There are many challenging aspects to simulating frog time scheme with an Asselin ®lter on an Arakawa ¯ows in this region. Perhaps foremost, in the case of B grid, and a semi-implicit treatment of acoustic modes. the Catalina eddy, is the fact that it is a mesoscale phe- The time step for the innermost domain is 13.3 s, shorter nomenon that evolves on synoptic timescales. Thus, it than the usual 20 s used at this resolution in MM5 is essential to correctly predict the synoptic-scale evo- because of the high vertical resolution near steep terrain. lution over the data-sparse Paci®c Ocean while at the Flow past such steep terrain can induce large vertical same time retain horizontal resolution adequate to re- motions that could otherwise violate the Courant±Frei- solve the important topographic features within the drichs±Lewy criterion for numerical stability. coastal zone. Another serious issue is prediction of the marine boundary layer. Because of poor initial data offshore, the marine layer and marine inversion are often b. Control simulation not well represented in analyses or initial conditions. 1) COMPARISON WITH OBSERVATIONS Thus, the model requires time to develop the correct structure. Prediction of the marine layer also represents We ®rst present a sequence of model surface [actually a severe test of the boundary layer scheme in the model. ␴ ϭ 0.997, or about 25 m above ground level (AGL)] In addition, high vertical resolution is necessary to re- winds and sea level pressure corresponding to some of solve the inversion atop the marine layer. the times analyzed in MA (refer to their Fig. 22). From The model used in this study is the Pennsylvania State Fig. 2, it is apparent that the model produces a quasi- University±National Center for Atmospheric Research stationary trough in the lee of the San Rafael Mountains Mesoscale Model, version 5 (MM5). The model is non- during the night of 27/28 June, which weakens during hydrostatic and integrates the fully compressible prim- the following day (1800 UTC 28 June). In contrast, on itive equations. Details may be found in Grell et al. 30 June, the lee trough has become a well-de®ned vortex (1994). In light of the above considerations, we use three that drifts southward through the bight region. domains, consisting of 60-, 20-, and 6.7-km resolution The pressure perturbation associated with the daytime (Fig. 1). This allows us to capture the evolution on the vortex on 30 June (approximately 1.5 mb) is in reason- synoptic scales as well as the local effects of terrain. able agreement with the observations (about 2 mb). The All domains use 45 vertical levels in the terrain-follow- simulated vortex on 30 June has a high degree of circular ing ␴ coordinate, symmetry and ®lls the bight region, similar to the observed eddy. The low-level circulation matches the pr Ϫ ptop available observations fairly well, though some of the ␴ ϭ , (1) ps Ϫ ptop coastal stations are clearly in¯uenced by local effects not resolved by the model. The model produces a re- where pr is the pressure in the hydrostatic reference versal of the climatological pressure gradient along the state, ps is the reference surface pressure, and ptop ϭ southern California coast each night (so that sea level 100 hPa. The ␴ levels are more closely spaced near the pressure increases southward), with the gradient becom- ground, with about 20 levels in the lowest 2 km. ing nearly orthogonal to the coastline during the day. In the control run, all domains are initialized at 1200 There are some noteworthy discrepancies between UTC 27 June (0400 LST). The 60- and 20-km domains the model and observations. For instance, the pressure are run interactively, and to prevent drift of the larger- trough on 28 June is too weak and slightly too far scale ®elds away from reality, we nudge continuously offshore in the model, especially at night. On 30 June, to the 12-hourly National Meteorological Center (now the vortex appears to move southward too rapidly operating as the National Centers for Environmental through the bight. It also appears that the predicted Prediction) analyses on the 60-km domain using the easterlies to the north of the vortex on the night of 30 technique of Stauffer and Seaman (1990). No nudging June are slightly stronger and more extensive than ob- is performed on the 20-km domain. The result is a set served. The model had dif®culty simulating the coastal of lateral boundary conditions (20-km resolution) that stratus, which was observed to accompany the for- force the 6.7-km resolution grid (run separately). To- mation of the eddy (MA's Fig. 18). This problem has pography on the innermost domain is shown in Fig. 1. been noted in other studies using MM5 (Mass and This topography is obtained from a 30-s global dataset, Steenburgh 2000). The presence of clouds, and the interpolated to the 6.7-km domain and smoothed to re- radiative cooling at cloud top, would no doubt help to move wavelengths of twice the grid spacing. The strengthen the marine inversion. However, our results smoothed terrain does not resolve the Santa Ynez and will suggest that a stronger marine inversion does not, San Rafael Mountains as distinct features; hence only by itself, signi®cantly change the evolution of the sim- the San Rafael Mountains are labeled in Fig. 1. ulated eddy. This is consistent with results from UR, The physical parameterizations of primary impor- who obtained an adequate simulation of the June 1988 tance are the Blackadar scheme for the planetary bound- eddy without accurately simulating the marine inver-

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FIG. 2. Sea level pressure and winds at ␴ ϭ 0.997 (25 m AGL) for (a) 18-h simulation, 0600 UTC 28 Jun; (b) 30-h simulation, 1800 UTC 28 Jun; (c) 66-h simulation, 0600 UTC 30 Jun; and (d) 78-h simulation, 1800 UTC 30 Jun. Winds are depicted with the standard barb convention in which a long barb is 5 m sϪ1. A terrain mask (black shading) has been applied to elevations greater than 400 m MSL. Model winds are plotted every third grid point. Large wind symbols indicate observations closest to the time of the model ®elds. As in all horizontal slices, tick marks on edges of plots indicate gridpoint spacing.

sion. Because the cloud ®eld is mainly a response to 2) EDDY STRUCTURE AND EVOLUTION the eddy rather than a cause, poor performance in this regard is not a deterrent to examining the eddy dy- As noted in previous studies, the eddy is warm core; namics; however it obviously represents an important hence it decays with height above the marine inversion. issue for operational prediction. The vortex remains quite strong over the lowest kilo-

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derneath the warm air aloft between 0000 and 0600 UTC. However, the ¯ow also contains numerous smaller-scale transients during this period. By 1200 UTC it is apparent that potential tempera- tures and potential temperature gradients at 500 m have decreased over the bight, but now there is an even more robust surface circulation in phase with a warm anomaly. Although the eddy is warm core, the potential temperature anomaly in the mature eddy is only about half as strong as it is in the early stages of formation. Between 1800 UTC 29 June and 0000 UTC 30 June a transition occurs from a ¯ow regime dominated by north±south elongated ®laments of warm air (Fig. 4b) and vorticity (Fig. 4a) to more concentrated cyclonic vorticity and a broad depression of the marine layer. Figure 5 captures the transition in a vertical cross section roughly orthogonal to the mean northwesterly ¯ow (D±DЈ in Fig. 1). At 1800 UTC, the ¯ow over the bight FIG. 3. Cross section A±AЈ (Fig. 1) depicting wind component more closely resembles a mountain wake as opposed to normal to the cross section (contour interval 2.5 m sϪ1, solid lines a coherent mesoscale eddy (Fig. 4a). A region of stag- for ¯ow out of page, dashed for ¯ow into page), and potential tem- nation occurs directly south of the Santa Ynez and San perature (contour interval, 2 K; gray) for 1800 UTC 30 Jun. Dashed Rafael Mountains. On either side of the wake, the ¯ow lines indicate winds into the cross section. is northerly (Fig. 4b, Fig. 5a), consistent with elongated ®laments of vorticity emanating from the mountains. meter but is barely detectable at 2 km AGL. Thus, most Depression of the marine inversion is weak and con®ned of the circulation is con®ned below the peaks of the to narrow ®laments elongated along the northwesterly Santa Ynez and San Rafael Mountains. The maximum ¯ow. The stagnation rapidly disappears as pronounced intensity of the vortex, in terms of circulation (area- warming ensues during the next 6 h. Instead of ®laments integrated vorticity), occurs between 0.2 and 0.4 km of vorticity, the vorticity is concentrated into a single AGL, generally within or slightly below the marine in- cyclonic anomaly over the northern bight on a scale of version (Fig. 3; see also Fig. 7). 50±100 km. This can be seen in Figs. 4a and 4c and Enhanced cyclonic circulation forms over the north- by noting the change in horizontal shear in Fig. 5 (es- ern bight around 0000 UTC 30 June (Fig. 4c). This sentially the change in relative vorticity). behavior is qualitatively similar to that evinced on the Another perspective on the ¯ow evolution can be seen previous two afternoons; here we focus on the last cycle from vertical cross sections taken roughly along the of vorticity production over the bight, which leads to mean ¯ow below 850 mb (see B±BЈ in Fig. 1) at dif- the mature eddy. As shown by Fig. 4d, this vorticity ferent times (Fig. 6). It is clear that the topographic increase coincides with the onset of pronounced warm- response projects strongly on at least two spatial scales. ing within the marine inversion. As much as a 10-K On the scale of the coastal orography (perhaps 20±30 warming occurs between 1800 and 0000 UTC (Figs. 4b km), there is an abrupt uplift of the isentropes on the and 4d). This warming, on a scale of about 100 km, is windward side and an abrupt descent of isentropes in collocated with pressure falls near the surface, charac- the lee as part of the mountain wave itself. In the model, teristic of lee troughing on this scale. Cyclonic relative the warm air reached nearly to the surface at Santa vorticity forms over the northern bight region, under- Barbara (location of Santa Barbara shown in Fig. 1). neath the warm anomaly. Highly variable winds and rapid warming at this loca- By 0600 UTC, there is a notable displacement be- tion are a characteristic of eddy events (Bosart 1983). tween the position of the center of near-surface vorticity On the scale of perhaps 100 km, there is a depression and circulation, and the position of maxima within the of isentropes in the lee over the bight (note the position warm anomaly at 500 m. Vertical cross sections (not of the highlighted 304-K isentrope) during the late af- shown) reveal that the north±south elongated warm ternoon. This is the source of the mesoscale lee trough- anomaly at 500 m is only 200±300 m deep; hence it ing. Mesoscale subsidence can be discerned over an area manifests itself as only a weak surface pressure trough. extending about 100 km downstream from the mountain The feature is highly transient, probably part of an in- in Figs. 6b and 6d. On the upwind side, prior to 30 June, ertia±gravity wave emitted as the balance adjustment the isentropes slope upward toward the mountain over takes place. Thus the time scale is too short for me- a distance commensurate with mesoscale blocking by soscale vorticity generation through vortex stretching to the terrain (Pierrehumbert and Wyman 1985; Smolar- occur. If one views the ®elds on the scale of the bight kiewicz and Rotunno 1989). By 1200 UTC 30 June (Fig. [O(100 km)], then the cyclonic circulation remains un- 6e), the domain average northwesterly ¯ow is weak, the

Unauthenticated | Downloaded 09/25/21 12:38 PM UTC AUGUST 2000 DAVIS ET AL. 2891 mountain wave nearly disappears and the variation in of the marine layer inversion, increase of geostrophic temperature between the windward and lee side of the vorticity, and subsequent increase of relative vorticity mountain is dominated by weak synoptic-scale baro- occur each day. clinicity and the asymmetry of the diurnal heating. This heating creates a 1-km-deep mixed layer to the north of 3) UPSTREAM CONDITIONS the coastal mountains during the day and but only a shallow mixed layer to the south of the mountains over In addition to an overall weakening during the 3.5- the water. Trajectories [section 3b(3)] suggest that the day period, Fig. 8 shows that the north-northwesterly depression of isentropes in the lee of the coastal terrain winds over the Central Valley north of the San Rafael is predominately adiabatic. and Santa Ynez Mountains maximize their strength at A time±height cross section (Fig. 7) of ®elds averaged night, forming the well-known nocturnal low-level jet over the boxed area in Fig. 1 clearly depicts the eddy in this region (Seaman and Stauffer 1994). The strength- evolution as a multiday period of diurnally oscillating ening of low-level north-northwesterlies in the evening vorticity, strong during the evening, weak during the is a consequence of differential heating due to the land± day. There are similar, though not identical, oscillations sea contrast and the sloping terrain on the west side of in perturbation pressure (Fig. 7a), perturbation potential the Sierra Nevada range. The combined effect is to pro- temperature (Fig. 7a), and geostrophic vorticity (Fig. duce widespread baroclinicity with warmer air to the 7b). Here, the term perturbation refers to the deviation east. Adjustment of this ¯ow toward thermal wind bal- from the initial state. The vorticity (Fig. 7c) tends to ance produces the low-level jet that maximizes a few maximize within the marine inversion, only slightly be- hours after the peak thermal contrast is reached, the low the maximum warming. It is apparent that the de- temporal lag corresponding roughly to the inertial pe- pression of the marine inversion layer and the associated riod. warming precede the vorticity increase in each cycle of In general, the temporal variation of the strength of the oscillation. The lag averages about 3 h but is clearly low-level winds is in phase with the variation of strat- variable. i®cation, with the result being that the characteristic The perturbation pressure ®eld evinces a lag with Froude number (Fr ϭ U/NH) is often less than 0.5, respect to the warming within the marine inversion. Be- approximately the transition between ¯ow mainly cause the ¯ow is very nearly hydrostatic, this lag must around (small Fr) and ¯ow over (large Fr) a three-di- be due to potential temperature changes above 1.5 km. mensional mountain (e.g., Smolarkiewicz and Rotunno Such changes do not appear to strongly in¯uence the 1989). Here, U is the wind speed, N is the Brunt±VaÈisaÈlaÈ spinup of vorticity below 1 km MSL. The reason is that frequency, and H is the topographic height. For ex- these pressure changes are likely occurring on larger ample, when the nocturnal jet is strongest and the strat- scales. This notion is supported by the structure and i®cation is large, an estimate of U ϭ 10 m sϪ1, N ϭ evolution of the Laplacian of the perturbation pressure, 0.02 sϪ1, and H ϭ 1000 m yields Fr ϭ 0.5. During the normalized to have units of vorticity (i.e., the geo- daytime, although the strati®cation is nearly zero, the strophic vorticity). This ®eld has a much more consistent velocity within the boundary layer is also very small. temporal relationship with the vorticity and potential However, careful inspection of Fig. 8 reveals that for a temperature than does the perturbation pressure. In par- brief time in the late afternoon the strati®cation is weak ticular, the lag between geostrophic vorticity and vor- and the wind gradually strengthens such that the char- ticity at 400 m is consistently 2±4 h. This lag is expected acteristic Froude number becomes of order unity or larg- from the fact that in a ¯uid initially at rest, the increase er. Our hypothesis, to be supported further by sensitivity of convergence is proportional to the geostrophic vor- experiments (Section 4), is that the ability of the air ticity. The convergence then leads to vorticity produc- mass below about 1 km MSL to traverse the coastal tion through stretching. mountains north of the bight is the key factor deter- There are some discrepancies between the trends in mining the depression of the marine inversion over the geostrophic vorticity and vorticity and in the potential bight that is so intimately associated with the diurnal temperature. In particular, the warming appears stronger variation of the Catalina eddy. each day, but the vorticity increase becomes smaller each day. The reason for this behavior remains unclear. 4) AIRFLOW Part of the reason could be due to our choice of averaging domain. There is no guarantee that all of the pos- Aspects of the air¯ow around the developing Catalina itive potential temperature anomaly and cyclonic vor- eddy are elucidated by considering two swarms of tra- ticity are contained within the domain at all times. This jectories terminating near the surface (25 m AGLÐthe may be inferred from Fig. 4, which shows a fairly com- lowest model level) (Fig. 9a) and at 950 mb (Fig. 9b). plicated spatial relationship between warming at 500 m We choose 950 mb as those trajectories represent parcels and pressure changes beneath as the ¯ow attempts to that terminate within the marine inversion layer. The adjust toward a balanced state. Our interpretation of Fig. area chosen for each swarm is a rectangle roughly 200 7 is that, practically speaking, a very similar depression km on a side, centered on the bight, upon which parcels

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FIG. 4. (a) Winds, pressure (mb), and vorticity (10Ϫ5 sϪ1) at 50 m MSL for 1800 UTC 29 Jun; (b) winds and potential temperature (gray scale) at 500 m MSL for 1800 UTC 29 Jun; (c) as in (a) but for 0000 UTC 30 Jun; (d) as in (b) but for 0000 UTC 30 Jun; (e) as in (a) but for 0600 UTC 30 Jun; (f) as in (b) but for 0600 UTC 30 Jun; (g) as are chosen at ®ve-gridpoint intervals (about 33-km spac- that move southeastward within the marine layer off- ing). Owing to the relatively high resolution of these shore, exhibiting almost no change in altitude. The par- simulations, model output at 15-min intervals was used cels on the eastern edge of this family curve cyclonically to calculate back trajectories, originating from a regular as the vortex forms. The second family of trajectories grid of points surrounding the eddy at 1200 UTC 30 originates over the continent at a higher altitude and June. The trajectories were computed by further inter- moves southward, traverses the coastal mountains, and polating the winds linearly from the 15-min data to then descend over the bight region. The third family of 5-min time steps. In 5 min, parcels with a velocity of parcels starts at low levels over the coastal zone and 11msϪ1 travel one grid length. moves generally northwestward, clearly experiencing From the total swarm ending at the surface, several orographic trapping. The ensemble of trajectories end- trajectories representing the major airstreams are se- ing over the bight indicates systematic convergence of lected and displayed in Fig. 9a. Three families of tra- air, consistent with a Lagrangian increase of absolute jectories are suggested. The ®rst encompasses parcels vorticity.

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FIG.4.(Continued) in (a) but for 1200 UTC 30 Jun; and (h) as in (b) but for 1200 UTC 30 Jun. For winds, one long barb is 5 m sϪ1. Contour interval for pressure is 0.5 HPa. Relative vorticity greater than 10Ϫ4 sϪ1 (dark) and Ϫ10Ϫ4 sϪ1 (light) are shaded.

Two families of trajectories ending at 950 mb are seen 8). Air originating below 1 km MSL over the Central in Fig. 9b. In the ®rst, parcels move southward, traverse Valley experiences increased blocking and de¯ection the coastal mountains, and descend over the bight. There with time. As night progresses, air descending into the is a suggestion that among the parcels originating to the marine inversion layer over the bight must originate from north of the mountains, those that undergo the greatest increasingly higher altitudes. The issue of blocking will ascent do so at the beginning of their paths, shortly after be investigated more thoroughly in section 4b. 0000 UTC. Parcels arriving at the coastal mountains north The second family of trajectories originates over the of the bight late at night either originate from higher bight at 0000 UTC and experiences little change in al- altitudes or tend to cross the coastal mountains at lower titude. Many members of this family of trajectories re- altitudes than the parcels traversing the mountains earlier. verse their direction and ¯ow northward at later times This behavior is broadly consistent with the fact that the in response to either the cyclonic vorticity generation Froude number characterizing air impinging on the coast- or blocking by the coastal mountains, similar to the al mountains decreases with time during the night (Fig. behavior found at low levels.

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FIG. 5. Cross sections (D±DЈ in Fig. 1) of normal-component wind and potential temperature for (a) 1800 UTC 29 Jun; (b) 0000 UTC 30 Jun. Contour intervals are 2 m sϪ1 and2K.

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FIG. 6. Vertical cross sections (B±BЈ in Fig. 1) showing wind vec- tors in the plane of the section (see scale at lower right) and potential temperature (contour interval, 2 K), for (a) 1500 UTC 28 Jun; (b) 2300 UTC 28 Jun; (c) 1500 UTC 29 Jun; (d) 2300 UTC 29 Jun; and (e) 1200 UTC 30 Jun. Vector scale for all plots is below (e). Contour interval is 2 K.

Integration of potential temperature along trajectories time boundary layer or above it and, thus, are not ex- suggests primarily adiabatic ¯ow for trajectory families posed to heating from below. As parcels traverse the that do not cross the high terrain. For those trajectories coastal terrain, some diabatic cooling occurs, especially that do, the ¯ow on the windward slope is largely adi- just in the lee where parcels pass through the large- abatic. Parcels typically begin near the top of the day- amplitude mountain wave (Fig. 6). The other locus of

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trapped against the coastal orography. Second, we wish to better con®rm the source of vorticity responsible for the transient intensi®cation of the cyclonic vortex. A full Lagrangian budget of the momentum and vorticity equations would require model output with a temporal frequency that is infeasible (essentially every time step). However, some progress can be made by simply summing the individual terms in the momentum equation over short time intervals during the model integration. The momentum equations in MM5 can be written pЈץ *pץ pЈ ␴ץ *d(p*u) mp ϭ p*f␷ ϪϪ ϩD ␴ uץ xץ *xpץ[] dt ␳ pЈץ *pץ pЈ ␴ץ *d(p*␷) mp ϭϪp*fu ϪϪ ϩD , ␴ ␷ץ yץ *ypץ[] dt ␳ Coriolis pressure gradient dissipation (2)

where f is the Coriolis parameter, D represents the combined effects of surface friction and vertical momentum mixing, ␳ is density, and m(x, y) is a map-scale factor. Our calculation accumulates each of the right-hand-side terms over a 1-h period. This yields a more representative picture of the parcel accelerations than might be

FIG. 7. Time±height sections of (a) perturbation pressure (thin) and obtained from a single time step (there are 240 time perturbation potential temperature (heavy), with perturbation refer- steps per hour). ring to deviations from the initial condition; (b) geostrophic relative Figure 10a illustrates the role of the pressure gradient vorticity; and (c) full relative vorticity (heavy) and potential tem- force as the cyclonic circulation in the bight begins in- perature (thin) averaged over a 25 by 25 (167 km by 167 km) gridpoint square over the bight (see Fig. 1). The contour intervals are (a) are creasing (0100 UTC 30 June). Large imbalances are 0.5 hPa and 1 K with positive temperature anomalies greater than 6 obvious as the pressure gradient force tends to dominate K shaded and dashed lines denoting negative values, (b) 5 ϫ 10Ϫ5 the dissipation (frictional) and Coriolis forces, espe- Ϫ1 Ϫ4 Ϫ1 Ϫ5 Ϫ1 s with values greater than 2 ϫ 10 s , and (c) 10 s and2K cially over the northeastern part of the bight. It is ap- with vorticity greater than 6 ϫ 10Ϫ4 sϪ1 shaded. parent that the divergent part of the wind ®eld is initially forced by the pressure ®eld, speci®cally, the Laplacian nonconservation is within the marine inversion. Both of pressure or geostrophic vorticity. This, in turn, mainly locations represent areas where isentropes collapse onto results from the dramatic, localized warming in and one another and are likely the loci of strong mixing, above the marine inversion layer caused by the ¯ow resulting in nonconservation of potential temperature over the mountains to the north. Over the northern part and potential vorticity (PV). Although we have not fo- of the bight, the result is a spinup of cyclonic vorticity. cussed on PV per se in this paper, PV and vorticity have To the west of the terrain lying east of the bight, the similar distributions within the marine inversion. For a pressure gradient force is unopposed near the coast and more detailed discussion of the vorticity and PV per- southeasterly ¯ow results (Fig. 10c). The pressure gra- spectives of ¯ow past complex terrain, the reader is dient force shown in Fig. 10a is nearly constant between referred to Thorpe et al. (1993), Schar and Smith (1993), 0100 and 0400 UTC, during which time the southeast- Persson (1996), and Rotunno et al. (1999). erly ¯ow accelerates. Friction acts to reduce velocities near the surface and approximately accounts for the 4. Additional diagnostics and sensitivity overestimate in local acceleration one obtains from the simulations pressure gradient force alone. Around 0600 UTC (see Fig. 4e for reference) the pressure gradient force grad- a. Momentum budgets and vorticity diagnostics ually becomes oriented orthogonal to the coast as the The motivation for momentum budgets is twofold. eddy strengthens and moves southward through the First we wish to assess whether the southerlies along bight (Fig. 10b). This effectively ends the alongshore the coast of southern California are, at least initially, acceleration.

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FIG. 8. Time±height sections of (a) wind component toward 160Њ and (b) potential temperature for a model grid point over the Central Valley (see Fig. 1 for location of point). The contour intervals are2msϪ1 for wind and 2 K for potential temperature. Winds stronger than 14 m sϪ1 are shaded. Below (b), shaded bars indicate time periods when Fr Ͼ 0.5 (gray) and Fr Ͼ 0.75 (black). Here Fr is de®ned using the mean velocity and strati®cation below 1 km MSL and a depth scale of 1 km.

FIG. 9. Selected back trajectories beginning at 0100 UTC 30 Jun and ending at 1200 UTC 30 Jun on (a) the lowest sigma level of the model (about 25 m AGL) and (b) the 950-mb pressure surface. Each arrow head corresponds to the position of the particle at an hourly interval. Starting pressures are labeled for all trajectories. For those trajectories that experience signi®cant (greater than 20 mb) rise and then fall, the minimum pressure is plotted. For trajectories that end at the surface over land, the ®nal pressure is plotted. Gray rectangles indicate the area over which uniformly spaced trajectory end points were speci®ed.

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FIG. 10. (a) Pressure gradient force (heavy, black arrows), Coriolis force (thin black arrows), and frictional force (vertical mixing, surface friction, and vertical diffusion; gray arrows) at 50 m MSL for 0100 UTC 30 Jun. Forces represent an hourly accumulation from 0000 UTC 30 Jun to 0100 UTC 30 Jun, (b) as in (a) but for 0600 UTC 30 Jun, and (c) vertical cross section (see Fig. 11a) of normal wind component (positive indicating southeasterly ¯ow) and potential temperature (dashed) at 0400 UTC 30 Jun. Contour intervals are 2 m sϪ1 and2K, respectively, with positive section-normal winds shaded according to grayscale table at bottom. b. Sensitivity experiments The cyclonic circulation (average vorticity within the box in Fig. 1) is about 1.5 times stronger than in the 1) REMOVAL OF DIURNAL CYCLE control simulation below 1 km MSL. The mesoscale To test the importance of the diurnal cycle in the warm pool at 1 km MSL is also displaced northward formation of the eddy we perform two sensitivity sim- relative to its location in the control simulation and its ulations. Each of these removes the diurnal cycle by magnitude is enhanced, consistent with a stronger, holding ground temperature ®xed in time at its initial warm-core vortex. value. This effectively preserves the PBL structure for This result suggests two key points. First, the Catalina all time during the simulation. The two start times are eddy is not an example of vortex shedding behind an 1200 UTC 29 June (denoted experiment ND12) and obstacle in low Fr ¯ow. Critical for the strong lee 0000 UTC 30 June (ND00). troughing and vortex spinup is the existence of moderate At 0000 UTC 30 June, the PBL upstream from the northerly ¯ow over a deep, weakly strati®ed layer below coastal mountains has attained its maximum depth of the mountaintop (i.e., large Fr). Weak strati®cation and about 1 km (Fig. 8). By holding ground temperature strong ¯ow allow the air to traverse the mountains and constant after this time (expt ND00), the land±sea ther- produce strong mesoscale sinking and depression of the mal contrast is maintained throughout the simulation. marine layer inversion in the lee. Implicit in the above Consequently, the coastal northerlies remain strong and description is the importance of enhanced strati®cation the large Fr state remains allowing the ¯ow over the in the lee of the mountains simply because a given mountains to persist. The 12-h prediction, valid at 1200 downward displacement can produce greater warming UTC 30 June, is shown in Figs. 11a and 11b (cf. Figs. if the strati®cation is large. However, it is likely, based 4g and 4h). A well-de®ned eddy is evident, one that is on idealized simulations (e.g., Peng et al. 1995; Ska- considerably farther north than in the control simulation. marock et al. 1999), that lee troughing and a robust

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FIG. 11. (a) Winds, pressure (mb), and vorticity (10Ϫ5 sϪ1) at 50 m MSL for 1200 UTC 30 Jun for experiment ND00, (b) winds and potential temperature (gray scale) at 500 m MSL for 1200 UTC 30 Jun for experiment ND00; (c) and (d) are as in (a) and (b) but for expt ND12. Contouring and shading are as in Fig. 4. transient vortex would form even with a typical tro- tained whereas in the control simulation, the low-level pospheric strati®cation over the bight, as opposed to the strati®cation increases north of the mountains and the large strati®cation supplied by the marine inversion lay- ¯ow eventually becomes blocked below crest level, thus er. halting the generation of mesoscale vorticity through Second, it is apparent that the sea breeze does not depression of the marine layer inversion. At 1200 UTC necessarily weaken the eddy or inhibit its formation. 30 June, simulation ND00 more resembles the state of Because the land±sea contrast is maintained in ND00 the control simulation at 0600 UTC 30 June (cf. Figs. throughout, the forcing of the sea breeze is as well. 11a and 4e), the latter being a time before the marine However, the cyclonic vorticity generation overwhelms inversion begins to assume its daytime structure and this effect, as evidence by a near absence of onshore near the end of the period of high Froude number ¯ow ¯ow at 1200 UTC 30 June in ND00. over the coastal mountains (Fig. 8). We speculate that The reason why the eddy in ND00 appears closer to the greater circulation in ND00 is a result of the greater the mountain than in the control simulation is that the time over which the marine layer is depressed. This may forcing, the ¯ow over the mountains (Fig. 12a), is main- allow a more ef®cient adjustment of the wind and mass

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FIG. 12. Cross sections along C±CЈ (see Fig. 1) at 0600 UTC 30 Jun FIG. 13. Flow-normal cross sections along D±DЈ (as in Fig. 5) for with conventions as in Fig. 3: (a) expt ND00 and (b) expt ND12. (a) expt ND00 and (b) expt ND12, both at 0600 UTC 30 Jun.

®elds than in the control simulation, reducing the loss The dramatic difference between ND12 and ND00 is of amplitude through balance adjustment (i.e., the ra- perhaps clearest when viewed in terms of forward tra- diation of inertia±gravity waves). The greater circulation jectories emanating from over the Central Valley. Figure might also result from a greater integrated warm anom- 14 shows the evolution of four representative air parcels aly that can form in ND00 due to the prolonged topo- from each simulation, initialized at 950 mb at different graphic forcing. During the day on 30 June, the eddy points along the ¯ow upwind from the coastal moun- in simulation ND00 moves southward into the bight tains. The trajectories all begin at 0000 UTC 30 June region, but the background ¯ow becomes so weak that and end 12 h later. Three of the four trajectories from it does not progress farther and weakens in place by ND00 are able to traverse the coastal mountains; the 0000 UTC 1 July. fourth is still ascending on the windward slope at the The simulation with no diurnal cycle after 1200 UTC end of the trajectory integration. In ND00, the trajec- 29 June (expt ND12, Figs. 11c and 11d) does not pro- tories cross the mountains and descend, and this descent duce a well-de®ned eddy over the bight region. This is responsible for depressing the marine layer inversion would be expected from the foregoing arguments be- over the bight (Fig. 13a), creating the mesoscale warm cause the thermal contrast between land and water at pool seen in Fig. 11a. Two of the trajectories from ND12 1200 UTC (0400 LST, just before sunrise) is nearly a are completely blocked and reverse direction to form a minimum for the day and the Froude number of the ¯ow well-de®ned southerly current over the eastern portion over the Central Valley is low. By extending this min- of the Central Valley. This reversal forms the ``Fresno imal thermal contrast inde®nitely in time, the low-level eddy'' (Seaman and Stauffer 1994), a mesoscale cy- ¯ow (below 1 km MSL) impinging on the San Rafael clonic vortex that is often observed during the warm and Santa Ynez Mountains remains blocked (Fig. 12b) season and maximizes its intensity late at night. The as the upstream strati®cation becomes large. There is other two trajectories from ND12 are de¯ected westward even more severe blocking of the low-level ¯ow than and forced over the coastal mountains to points well in the control simulation. This creates a wake in the lee removed from the bight region. and long vorticity ®laments downstream from the ¯anks These results strongly support the notion that the for- of the highest terrain. No coherent mesoscale cyclonic mation of the Catalina eddy is initiated when the bulk vorticity is produced on the scale of the bight region as air mass at low levels over the Central Valley is able the marine inversion is not depressed (Fig. 13b). to traverse the coastal mountains. In the control simu-

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major features of the eddy were predictable if the synoptic-scale ¯ow providing boundary conditions to the model was well predicted. This implied predictability several days in advance. It is therefore not surprising that if we initialize the 6.67-km domain from the 20- km simulation at 1200 UTC 29 June instead of 1200 UTC 27 June, the prediction is only slightly different (not shown). Both simulations represented the evolution of the observed eddy reasonably well. The MM5 model is constantly being updated. Since the beginning of this work, a new version of the model was released (version 2), and even this new version has been modi®ed. We performed an additional sensitivity simulation to test to what extent the observed features were reproducible and whether the eddy was simulated better by the updated model. The interested reader is referred to articles by Dudhia et al. (1996) and Dudhia (1998) for a brief description of the changes made in MM5. For our purposes, the most important changes were the reduction of vertical diffusion and the use of potential temperature as the prognostic variable in the adiabatic portion of the thermodynamic equation (Ska- marock et al. 1997). Although the position and intensity of the eddy in the two simulations are similar (not shown), there are some differences in the details of the vertical pro®les of wind and temperature. The boundary layer pro®le upstream from the coastal mountains stabilizes at night, but the strati®cation between 0.5 and 1.0 km MSL on 30 June remains weaker and the near-surface inversion stronger than in the control simulation. The eddy is slightly stronger and slightly farther north in the newer version of the model. The difference in thermal structure appears to be an effect of larger vertical diffusion in the control simulation. This difference makes the version 2 simulation behave a bit like ND00, in the sense that air is FIG. 14. Selected trajectories from ND00 (black) and ND12 (gray) able to ¯ow over the coastal mountains for a longer originating at 950 mb at 0000 UTC 30 Jun. Filled circles mark initial period in the evening than in the control simulation. The location of parcel pairs. Each pair consists of one parcel from each simulation with the newer version of the model agrees experiment initialized at the same location. Arrowheads are spaced at hourly intervals. Pressure labels denote minimum pressure (for slightly better with observations of the placement of the parcels that rise substantially) and ®nal pressure at 1200 UTC 30 Jun. vortex than does the control simulation. Finally, downstream from the mountains the marine layer inversion is sharper in the simulation with the newer version of lation, there is only a narrow window of time during the model, a result of improved numerics (integration late afternoon and evening when this is possible. It is of potential temperature instead of temperature in the during this time that the mesoscale warm pool over the thermodynamic equation) and reduced vertical diffu- bight develops and cyclonic vorticity generation takes sion. It does not appear that the details of the marine place. layer and marine inversion are critical for an accurate simulation of the eddy. This relative insensitivity to the details of the marine layer can also be inferred from the 2) LATER INITIALIZATION AND UPDATED CODE successful simulations by UR utilizing coarser vertical The use of a 3-day simulation at 6.67-km resolution (and horizontal) resolution than in the present study. was motivated mainly to study the entire eddy at a scale at which the coastal mountains could be adequately re- 5. Discussion of dynamical mechanisms solved. However, an obvious question is whether simulations with different initialization times produce sim- Despite the multiday nature of the Catalina eddy life ilar results. Ueyoshi and Roads (1993) discussed the cycle, the pronounced periodicity of the cyclonic vor- issue of predictability at length and decided that the ticity over the bight suggests that the development of

Unauthenticated | Downloaded 09/25/21 12:38 PM UTC 2902 MONTHLY WEATHER REVIEW VOLUME 128 circulation each evening can be thought of as a ``start the type of vorticity ®laments implied in Fig. 5, ones up'' vortex, a balanced, warm-core eddy that results that are characteristic of a wake. from impulsively initiated ¯ow past a mountain at low- Unfortunately, further theoretical guidance as to the to-moderate Rossby number. The Rossby number is here transient response of low-to-moderate Fr, moderate Ro, de®ned Ro ϭ U/fL, with U a characteristic ¯ow speed, ¯ow past an isolated mountain is generally lacking. f the Coriolis parameter, and L the topographic length Cross sections normal to the ¯ow (e.g., Fig. 12) further scale along the ¯ow. In the present case, we may esti- illustrate the complicated nature of the ¯ow impinging mate the Rossby number to be ϳO(1), based on a to- on the San Rafael and Santa Ynez Mountains. In ad- pographic length scale that is several tens of kilometers. dition to the diurnal variations, there is considerable A start-up vortex results from potentially warm air near vertical and horizontal shear within this ¯ow. Such shear mountaintop being forced to lower elevations, locally is a persistent feature of the larger-scale ¯ow within the depressing the isentropes and producing a warm tem- coastal zone. Little is known about how shear modi®es perature anomaly. The ¯ow adjusts to cyclonic balance, the results from the studies based on uniform ¯ow. and the feature is simply swept downstream by the background current, leaving behind an anticyclone over the mountain. In the present case, the relatively sudden de- 6. Conclusions celeration of synoptic-scale northerlies on 30 June al- Our study has focused on the dynamics of the Catalina lows a mature eddy to remain in the bight region once eddy observed to form during the period 26±30 June formed. 1988. As a surrogate for observations, we have con- Despite the temporal lag between thermal contrast centrated on the evolution of the eddy as captured by (temporally coincident with weak strati®cation) and simulations with the MM5 model. While the model re- ¯ow speed, the bulk strati®cation in the boundary layer produces most of the observable features of the vortex upstream from the coastal mountains is still weak when and its evolution, the dearth of upper-air observations the low-level northerlies increase during the late after- on the mesoscale prevents us from verifying many as- noon. This corresponds to a narrow window of time pects of the simulation directly. Nevertheless, the sig- during which ¯ow from the boundary layer air can as- ni®cant ®ndings are the following. cend over the coastal mountains. At all other times, either the ¯ow is weak or strati®cation large, so that the characteristic Froude number (U/NH) is small (around a. Two scales of ¯ow response 0.5 or less), corresponding to ¯ow around the coastal We documented two scales of response to ¯ow over orography. The low Froude number regime exhibits a or around the San Rafael and Santa Ynez Mountains. wake bounded by vortex ®laments. The high Froude First, there is the mountain wave, on a scale of the number regime exhibits a translating mesoscale vortex mountain half-width (about 20 km). Second there is the in the lee. Figure 4 depicts the downstream migration eddy and its associated mesoscale warm pool centered of the warm anomaly, which de®nes the vortex, and the around 500 m MSL with a horizontal a scale of nearly reestablishment of the wake, all on a timescale of about 100 km. 12 h. The association of ¯ow with Fr ϳ 1 and Ro ϳ 1 with ¯ow over the coastal range, as opposed to around the b. Importance of synoptic-scale evolution mountains, is largely an empirical result that is consistent with theoretical results. In the context of quasigeo- The synoptic-scale evolution can be thought of as the strophic theory (small Ro) air ¯ows over topography trigger for the Catalina eddy multiday episode. The regardless of Fr. Thus, a transient vortex is always the strengthening of the ¯ow allows more air to cross the result of an impulsively started ¯ow interacting with an mountain crest and strongly perturb the leeside marine isolated mountain. We believe that this process, shown inversion, creating more lee vorticity. In the present in Peng et al. (1995), qualitatively describes the case, it appears that the weakening of the synoptic-scale strengthening of vorticity over the bight region during ¯ow is the key for allowing the eddy to persist in the the evening hours. However, Ro being unity or larger bight region. and the ®nite height of the topography preclude quan- titative application of quasigeostrophy, and open the c. Importance of coastally trapped ¯ow possibility of other types of response. In Peng et al. (1995) it is shown that decreasing Fr In previous studies, strong emphasis has been placed for Ro near unity results in the formation of a quasi- on the trapping of ¯ow by the coastal mountains, re- steady wake in the lee of a mountain. We believe that sulting in northward acceleration, as the act de®ning the such a wake represents the late-night and daytime ¯ow formation of a mature eddy (MA; Clark 1994). Mo- in the lower troposphere during the period 26±30 June. mentum budget diagnostics suggest that orographically Maps of vorticity at 500 m and 1 km MSL earlier in trapped southerlies exist early in the intensi®cation of the event during the morning hours (not shown) indicate the eddy on 30 June.

Unauthenticated | Downloaded 09/25/21 12:38 PM UTC AUGUST 2000 DAVIS ET AL. 2903 d. Effect of the diurnal cycle Acknowledgments. This work was supported by Grant NOOO14-97-1-0019 from the Of®ce of Naval Research. Emphasis from previous studies has been on the sea The authors bene®ted from discussions with Drs. Ri- breeze as the main diurnal effect modulating the Cat- chard Rotunno, Joseph Klemp, and William Skamarock alina eddy intensity (Bosart 1983; Wakimoto 1987). We of NCAR. ®nd for this case that the sea breeze may be relatively unimportant because simulation ND00, with constant late-afternoon conditions, produced the strongest eddy REFERENCES of any simulation. The diurnal cycle has also been noted to favor Catalina eddy formation at night over the bight Bosart, L. F., 1983: Analysis of a California Catalina eddy event. region (Ueyoshi and Roads 1993), but the cause of this Mon. Wea. Rev., 111, 1619±1633. Boyer, D. L., P. A. Davies, W. R. Holland, F. Biolley, and H. Honji, has not been explained until now. This paper shows that 1987: Strati®ed rotating ¯ow over and around isolated three- an important effect of the diurnal cycle is to modulate dimensional topography. Philos. Trans. Roy. Soc. London, 322A, the strength and strati®cation of the ¯ow impinging on 213±241. the mountains and, hence, the downstream response. Clark, J. H. E., 1994: The role of Kelvin waves in evolution of the Late at night and during much of the day, the low- Catalina eddy. Mon. Wea. Rev., 122, 838±850. Colle, B., and C. F. Mass, 2000: High-resolution observations and level (below 1 km MSL) ¯ow approaching the San Ra- numerical simulations of easterly gap ¯ow through the Strait of fael Mountains from the north is strongly strati®ed or Juan de Fuca on 9±10 December 1995. Mon. Wea. Rev., 128, slow in speed. In the late afternoon this ¯ow begins to 2398±2422. strengthen at the same time that the strati®cation remains Dudhia, J., 1998: Recent changes in MM5-V2 and plans for V3. Preprints, Eighth PSU/NCAR Mesoscale Model Users' Work- low from the daytime heating creating a narrow tem- shop, Boulder, CO, NCAR, 1±2. [Available from C. A. Davis, poral window of higher Froude number ¯ow approach- NCAR, P.O. Box 3000, Boulder, CO 80307-3000.] ing the terrain, Thus, during a period of a few hours in , D. Hansen, and W. Wang, 1996: What's new in MM5 version the evening, the ¯ow is primarily over the San Rafael 2? Preprints, Sixth PSU/NCAR Mesoscale Model Users' Work- Mountains, instead of around them. This results in stron- shop, Boulder, CO, NCAR, 1±3. [Available from C. A. Davis, NCAR, P.O. Box 3000, Boulder, CO 80307-3000.] ger subsidence in the lee, which depresses the marine Grell, G. A., J. Dudhia, and D. R. Stauffer, 1994: A description of inversion, creating a warm anomaly and cyclonic cir- the ®fth-generation Penn State/NCAR mesoscale model (MM5). culation. NCAR Tech. Note 398, 121 pp. It appears that there is an optimal ¯ow strength for Mass, C. F., and M. D. Albright, 1989: Origin of the Catalina eddy. Mon. Wea. Rev., 117, 2406±2436. generating a quasi-steady disturbance in the bight. If the , and W. J. Steenburgh, 2000: An observational and numerical low-level northerly wind approaching the mountains is study of an orographically trapped wind reversal along the west too weak (as during the morning) the air ¯ows around, coast of the United States. Mon. Wea. Rev., 128, 2363±2396. not over, the terrain. If the ¯ow is too strong (as is the Peng, M. S., S.-W. Li, S. W. Chang, and R. T. Williams, 1995: Flow case prior to 30 June), then vorticity once produced will over mountains: Coriolis force, transient troughs and three di- mensionality. Quart. J. Roy. Meteor. Soc., 121, 593±613. simply be advected away. Another way of stating it is Persson, P. O. G., 1996: Low-level potential vorticity anomalies gen- that there is an optimal Rossby number regime (Ro Շ erated by coastal topography: What is their signi®cance? Pre- 1 and Froude number regime (Fr տ 1) that elicits the prints, Seventh Conf. on Mesoscale Processes, Reading, United best mesoscale response on the timescale of about half Kingdom, Amer. Meteor. Soc., 135±137. Pierrehumbert, R. T., and B. Wyman, 1985: Upstream effects of me- the diurnal period. This regime is brought about only soscale mountains. J. Atmos. Sci., 42, 977±1003. brie¯y by the coincidence of a favorable phase of the Rosenthal, J., 1968: A Catalina eddy. Mon. Wea. Rev., 96, 742±743. diurnal cycle with favorable synoptic-scale conditions. Rotunno, R., V. Grubisic, and P. K. Smolarkiewicz, 1999: Vorticity Based on these ®ndings, it is apparent that more the- and potential vorticity in mountain wakes. J. Atmos. Sci., 56, 2796±2810. oretical work is necessary to examine idealized ¯ows Schar, C., and R. B. Smith, 1993: Shallow-water ¯ow past isolated of varying complexity past three-dimensional mesoscale topography. Part I: Vorticity production and wake formation. J. orography. In particular, the coastal zone in many re- Atmos. Sci., 50, 1373±1400. gions of the world consists of extreme topographic var- Seaman, N. L., and D. R. Stauffer, 1994: The Fresno eddy: Numerical investigation of a mesoscale wind feature of the San Joaquin Val- iations and semi permanent large variations in the tem- ley using the Penn State/NCAR MM5. Preprints, Sixth Conf. on perature and winds. Our understanding of coastal phe- Mesoscale Processes, Portland, OR, Amer. Meteor. Soc., 572±575. nomena could be greatly increased with an ability to , , and A. M. Lario, 1995: A multi-scale four-dimensional more strongly link the results of observational and mod- data assimilation system applied in the San Joachin Valley during eling studies with results from more idealized calcula- SARMAP. Part I: Modeling design and basic performance char- acteristics. J. Appl. Meteor., 34, 1739±1761. tions. This obviously points out the need for obtaining Skamarock, W., C. Davis, D. Dempsey, and H.-M. Hsu, 1997: Tem- higher-resolution observations of the low-level ¯ow. In perature or potential temperature?: Comparisons using different the present case, remotely sensed surface winds (derived thermodynamic variables in MM5. Preprints, Seventh PSU/ from scatterometer data) or tracking of low-cloud mo- NCAR Mesoscale Model Users' Workshop, Boulder, CO, NCAR, 80±82. [Available from C. A. Davis, NCAR, P.O. Box 3000, tions, for example, could potentially resolve structures Boulder, CO 80307-3000.] on a scale of tens of kilometers, essential for verifying , R. Rotunno, and J. B. Klemp, 1999: Models of coastally trapped the complicated structures evinced in our simulations. disturbances. J. Atmos. Sci., 56, 3349±3365.

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Smith, R. B., 1979: Some aspects of quasi-geostrophic ¯ow over Thorpe, A. J., H. Volkert, and D. Heimann, 1993: Potential vor- mountains. J. Atmos. Sci., 36, 2385±2393. ticity of ¯ow along the Alps. J. Atmos. Sci., 50, 1573± Smolarkiewicz, P. K., and R. Rotunno, 1989: Low Froude number 1590. ¯ow past three-dimensional obstacles. Part I: Baroclinically gen- Ueyoshi, K., and J. O. Roads, 1993: Simulation and prediction of the erated lee vortices. J. Atmos. Sci., 46, 1154±1164. Catalina eddy. Mon. Wea. Rev., 121, 2975±3000. Stauffer, D. R., and N. L. Seaman, 1990: Use of four-dimensional Ulrickson, B. L., J. S. Hoffmaster, J. Robinson, and D. Vimont, 1995: data assimilation in a limited-area mesoscale model. Part I: Ex- periments with synoptic-scale data. Mon. Wea. Rev., 118, 1250± A numerical modeling study of the Catalina eddy. Mon. Wea. 1277. Rev., 123, 1364±1373. Thompson, W. T., S. D. Burk, and J. Rosenthal, 1997: An investigation Wakimoto, R., 1987: The Catalina eddy and its effect on pollution of the Catalina eddy. Mon. Wea. Rev., 125, 1135±1146. of Southern California. Mon. Wea. Rev., 115, 837±855.

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