Multiscale Analysis of the 7 December 1998 Great Salt Lake–Effect

1296 MONTHLY WEATHER REVIEW VOLUME 129

Multiscale Analysis of the 7 December 1998 Great Salt Lake±Effect Snowstorm

W. J AMES STEENBURGH AND DARYL J. ONTON NOAA Cooperative Institute for Regional Prediction, and Department of Meteorology, University of Utah, Salt Lake City, Utah

(Manuscript received 3 May 2000, in ®nal form 16 October 2000)

ABSTRACT The large-scale and mesoscale structure of the Great Salt Lake±effect snowstorm of 7 December 1998 is examined using radar analyses, high-density surface observations, conventional meteorological data, and a simulation by the Pennsylvania State University±National Center for Atmospheric Research ®fth generation Mesoscale Model (MM5). Environmental conditions during the event were characterized by a lake±700-hPa temperature difference of up to 22.5ЊC, a lake±land temperature difference as large as 10ЊC, and conditionally unstable low-level lapse rates. The primary snowband of the event formed along a land-breeze front near the west shoreline of the Great Salt Lake. The snowband then migrated eastward and merged with a weaker snowband as the land-breeze front moved eastward, offshore ¯ow developed from the eastern shoreline, and low-level convergence developed near the midlake axis. Snowfall accumulations reached 36 cm and were heaviest in a narrow, 10-km-wide band that extended downstream from the southern shore of the Great Salt Lake. Thus, although the Great Salt Lake is relatively small in scale compared to the Great Lakes, it is capable of inducing thermally driven circulations and banded precipitation structures similar to those observed in lake-effect regions of the eastern United States and Canada.

1. Introduction cold-frontal northwesterly ¯ow at 700 hPa, a lake±700- hPa temperature difference of at least 17ЊC (which ap- The prediction of lake-effect snowstorms that develop proximately represents a dry adiabatic lapse rate), and over and downwind of the Great Salt Lake (GSL) is one an absence of stable layers or inversions near or below of the major forecast challenges facing meteorologists 700 hPa.1 Steenburgh et al. (2000) used observations in northern Utah. Occurring several times each year, from a recently installed National Weather Service Great Salt Lake±effect (GSLE) snowstorms last from a Weather Surveillance Radar-1988 Doppler (WSR-88D) few hours to more than a day, frequently produce snow- to identify GSLE events between September 1994 and falls of 20±50 cm, and have contributed to the state May 1998. During this period, 16 well-de®ned GSLE record 129-cm lowland storm-total snowfall that was events were observed, with the synoptic, mesoscale, and observed near Salt Lake City (SLC) from 24 to 28 Feb- convective characteristics of these events examined us- ruary 1998 (Carpenter 1993; Slemmer 1998; Steenburgh ing National Centers for Environmental Prediction et al. 2000). Despite signi®cant improvement in obser- (NCEP) Rapid Update Cycle version 2 analyses (RUC2; vational technologies and numerical forecast systems, Benjamin et al. 1991, 1994), SLC radiosonde obser- GSLE snowstorms remain dif®cult to predict with lead vations, and local WSR-88D radar observations. In ad- times of more than a few hours. dition to supporting the ®ndings of Carpenter (1993), Previous studies have identi®ed the climatological Steenburgh et al. (2000) also found that GSLE events characteristics, large-scale conditions, and mesoscale tend to occur during periods of positive lake±land tem- precipitation structures associated with GSLE snow- perature differences, usually exceeding 6ЊC, and are storms. Based on lake-effect events identi®ed by visual most active during the overnight and early morning observations and spotter reports, Carpenter (1993) hours. It was hypothesized that the positive lake±land found that GSLE snowstorms were associated with post- temperature difference results in the development of

Corresponding author address: Dr. W. James Steenburgh, De- 1 Due to the elevation of the GSL (ϳ1280 m above mean sea level), partment of Meteorology, University of Utah, 135 South 1460 East surface and 700-hPa observations are used instead of surface and Room 819, Salt Lake City, UT 84112-0110. 850-hPa observations as is commonly done in studies of lake-effect E-mail: [email protected] snowstorms over the Great Lakes (e.g., Niziol et al. 1995).

᭧ 2001 American Meteorological Society

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United States (Carpenter 1993; Steenburgh et al. 2000). Wiggin (1950) described the general characteristics of lake-effect snowstorms in the Great Lakes region, including their potential for large accumulations and signi®cant variations in snowfall over short spatial scales. Additionally, Wiggin (1950) noted that such storms were favored in polar continental air masses during periods of large lake±air temperature differences, near- adiabatic lapse rates, and long overwater fetches. Peace and Sykes (1966) studied a lake-effect snowband using a mesoscale surface observing network over the eastern end of Lake Ontario. It was found that a narrow convergence line accompanied the snowband and it was hypothesized that surface sensible heating caused the formation of the snowband, with winds aloft controlling the location and movement of the band. Subsequent studies over the Great Lakes have identi®ed a variety of lake-effect precipitation structures including (i) broad area coverage, which may include multiple wind-parallel bands or open cells (Kelly 1982, 1984); (ii) shoreline bands that form roughly parallel to the lee shore due to the convergence of a land breeze with the large- FIG. 1. Geographic features of northern Utah. Surface elevation in scale wind ®eld (Ballentine 1982; Braham 1983; Hjelm- meters shaded according to scale at bottom left. Station locations discussed in text are Salt Lake City (SLC), Tooele (TOO), Hat Island felt and Braham 1983; Hjelmfelt 1990); (iii) midlake (HAT), Gunnison Island (GNI), Great Salt Lake Desert (S17), and bands that form when the large-scale ¯ow is parallel to the Salt Lake City NEXRAD radar site (KMTX). Railroad causeway the long axis of a lake and a lake±land temperature identi®ed by a dashed line. contrast exists (Peace and Sykes 1966; Passarelli and Braham 1981; Braham 1983; Hjelmfelt 1990; Niziol et land breezes and low-level convergence that focus the al. 1995); and (iv) mesoscale vortices that form in a development of convection over the GSL. The greater polar air mass under conditions of a weak surface pres- frequency of lake-effect precipitation during the over- sure gradient and large lake±air temperature differential night and early morning hours may be related to the (Forbes and Merritt 1984). diurnal modulation of the lake±land temperature dif- Precipitation during GSLE events is most frequently ference and associated land-breeze convergence, similar characterized by the irregular development of radar ech- to that suggested by Passarelli and Braham (1981) over oes over and downstream of the GSL (Steenburgh et al. Lake Michigan. 2000). The most commonly observed organized precip- GSLE snowstorms share many similarities with lake- itation structures are solitary wind-parallel bands resem- effect snowstorms over the Great Lakes region of the bling midlake bands found over the Great Lakes, and

FIG. 2. Daily mean lake-surface temperature at HAT (solid), air temperature at HAT (dashed), and air temperature at SLC (dotted) from 2 Sep 1998 to 31 Jan 1999. Large dots demarcate period of missing lake-surface temperature data from HAT.

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FIG. 3. Regional RUC2 analyses and observed SLC upper-air sounding at 1200 UTC 6 Dec 1998. (a) Sea level pressure (every 2 hPa) and 10-m winds (full and half barbs denote 5 and 2.5 m sϪ1, respectively). (b) 700-hPa temperature (every 2ЊC), wind [as in (a)], and relative humidity (%, shaded following scale at upper right). Geopotential height trough axis denoted by dashed line. (c) 500-hPa geopotential height (every 60 m) and absolute vorticity (ϫ10Ϫ5 sϪ1, shaded following scale at upper right). Geopotential height trough axis denoted by dashed line. (d) SLC skew T±logp diagram with temperature and dewpoint (ЊC) denoted by heavy solid lines. Short-dashed line represents surface parcel ascent. Filled circle represents lake temperature. Wind as in (a). broad-area coverage precipitation shields that form near model, Lavoie (1972) found that frictional convergence the lee shoreline. In addition, GSLE precipitation some- due to land±water roughness contrasts, and surface sen- times occurs in concert with orographic precipitation, sible heating due to lake±air temperature differences, or within a broader-scale precipitation shield associated produce upward vertical motion and elevated inversion with synoptic-scale lifting. Signi®cant enhancement of heights near the lee shoreline of Lake Erie. The lake± GSLE events can occur when lake-induced precipitation air temperature difference was found to be dominant. features, such as solitary wind-parallel bands, extend Hjelmfelt (1990, 1992) examined the importance of low- over the downstream orography. level instability, lake±land temperature difference, sen- Several studies have used numerical models to ex- sible and latent heat ¯uxes, topography, capping inver- amine lake-effect snowstorm dynamics (e.g., Lavoie sions, and upstream moisture in producing lake-effect 1972; Ballentine 1982; Hjelmfelt and Braham 1983; snowstorms over Lake Michigan. He found that both Hjelmfelt 1990). Using a three-layer primitive equation shoreline-parallel and midlake snowbands were favored

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FIG. 4. Same as Fig. 3 except for 0000 UTC 7 Dec 1998. by strong lake±land temperature differences, weak sta- though changes in snowband location in response to the bility, and the absence of capping inversions at low evolving synoptic-scale ¯ow had timing errors of a few elevations. Moderate cross-lake ¯ow enhanced land- hours. breeze-induced convergence, thus strengthening shore- The purpose of this paper, and the companion article line-parallel bands. Midlake snowbands, however, were by Onton and Steenburgh (2001), is to describe the evo- favored by strong wind ¯ow parallel to the long axis of lution and physical processes responsible for a GSLE the lake. Weaker wind ¯ows combined with strong lake± snowstorm that occurred on 7 December 1998. Snowfall land temperature differences tended to produce meso- accumulations of up to 36 cm were produced by the scale vortices instead of midlake bands. Upstream mois- event, which featured a wind-parallel snowband that ture was also found to be important in enhancing lake- developed near the western shoreline of the GSL and effect precipitation and land-breeze strength due to la- became aligned along the midlake axis as it moved east- tent heat release from condensation. Ballentine et al. ward and merged with a weaker snowband. Speci®c (1998) described a successful simulation of a Lake On- questions that will be addressed in the two papers in- tario snowband using the Pennsylvania State Univer- clude the following. sity±National Center for Atmospheric Research ®fth generation Mesoscale Model (MM5). The simulation R What are the underlying mesoscale dynamics respon- reproduced the observed precipitation distribution, al- sible for the development of GSLE snowbands? Are

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FIG. 5. Same as Fig. 3 except for 1200 UTC 7 Dec 1998.

solitary wind-parallel bands over the GSL produced terns observed during GSLE snowstorms? Does the primarily by thermally driven land-breeze conver- ``®xed'' surface forcing of the lake and surrounding gence? topography extend predictability, or do small errors R How important are topographic effects such as oro- in surface characteristics and the upstream ¯ow char- graphic uplift and low-level ¯ow blocking and chan- acteristics limit forecast skill? neling? To what degree are GSLE events triggered or enhanced by such local orographic effects? The mesoscale structure and evolution of the 7 De- R How do sensible and latent heat ¯uxes in¯uence the cember 1998 GSLE snowstorm is examined in the pre- development and intensity of lake-effect precipita- sent paper using conventional meteorological data, high- tion? Does the hypersaline composition of the GSL density surface observations provided by MesoWest, a signi®cantly affect latent heat ¯ux (compared to fresh- collection of cooperating mesonets in the western Unit- water) and snowband evolution or intensity? ed States, and a numerical simulation by the nonhy- R Does frictional convergence due to land±water rough- drostatic MM5. Section 2 describes the regional orog- ness contrasts in¯uence the development of GSLE raphy and unique characteristics of GSL hydrology, snowbands? composition, and air±lake interactions. Section 3 pre- R Can present-day mesoscale models accurately simu- sents a detailed observational analysis of the 7 Decem- late the mesoscale circulations and precipitation pat- ber 1998 event using RUC2 analyses, radar observa-

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FIG. 6. Same as Fig. 3 except for 0000 UTC 8 Dec 1998. tions, and MesoWest surface observations. Then, section The GSL is the largest body of water in the United States 4 uses a mesoscale model simulation to further examine west of the Great Lakes. It currently occupies an area the mesoscale structure and evolution of the event. A of ϳ4400 km2, is about 120 km long and 45 km wide, summary and discussion of major results follow in sec- and has an average (maximum) depth of only 4.8 (10) tion 5. Further diagnosis of the dynamics and predict- m. Due to the lack of a drainage outlet, the lake's size ability of the 7 December 1998 event is presented in ¯uctuates due to interseasonal and interannual variations Onton and Steenburgh (2001). in precipitation and evaporation, and has ranged from 2500 to 6200 km2 in area and from 1278 to 1284 m in surface elevation since the mid-1850s (Arnow 1980; 2. The Great Salt Lake and surrounding Wold et al. 1996). topography Due to the GSL's shallow depth, climatological lake- There are several unique aspects of the land surface surface temperatures exhibit little lag relative to cli- properties and orography of northern Utah that in¯uence matological mean air temperatures at Salt Lake City the development of lake-effect precipitation (Fig. 1). (Steenburgh et al. 2000; see their Fig. 2). The average These include the region's intense and complex vertical lake-surface temperature exhibits a maximum (mini- relief, and the varying hydrologic structure, thermal mum) near 1 August (1 February), similar to the timing characteristics, and hypersaline composition of the GSL. of the maximum (minimum) mean air temperature at

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FIG. 7. Lowest-elevation (0.5Њ) base-re¯ectivity analysis from the KMTX WSR-88D radar and MesoWest surface observations at (a) 0400, (b) 0515, (c) 0630, (d) 0815, (e) 1030, (f) 1315, (g) 1445, and (h) 1900 UTC 7 Dec 1998. Radar re¯ectivity shaded according to scale at upper left. Station observations include wind barbs (full and half barbs denote 5 and 2.5 m sϪ1, respectively), temperature (ЊC; upper left), and three-digit identi®er for selected stations (lower right). Snowbands A and B denoted by heavy dashed lines. Topographic contours shown every 500 m in solid lines (see Fig. 1 for elevations). Lake outline shown with dashed line.

SLC on 24 July (5 January). From late winter through daily air temperatures at SLC and HAT for the initial ®ve- summer, the mean lake temperature is similar to the month observation period is presented in Fig. 2. This ®gure mean air temperature at SLC, but during the fall through clearly illustrates the seasonal decline in both lake-surface early winter, the mean lake-surface temperature exceeds and mean air temperature. Note, however, that from Sep- the mean air temperature by 2Њ±3ЊC. tember to December, lake-surface temperatures were gen- Carpenter (1993) suggested that lake-surface tempera- erally 2Њ±3ЊC greater than the mean air temperature, while tures may correlate with the preceding week's mean air in January, lake-surface temperatures were similar to the temperature. In the past, estimates of lake-surface tem- mean air temperature. This is in rough agreement with the perature using this method were necessary for operational twice-monthly observations presented by Steenburgh et al. forecasting due to the lack of real-time observations. How- (2000), although they showed higher lake-surface tem- ever, starting in late summer 1998, lake-surface temper- peratures persisting into mid-January. Also evident in Fig. atures have been observed at a MesoWest site installed at 2 are more rapid lake-surface and mean air temperature Hat Island (HAT; see Fig. 1 for location). A comparison changes associated with transient synoptic weather sys- between the mean daily lake-surface temperature and mean tems. Speci®cally, lake-surface temperature changes of as

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FIG.7.(Continued) much as 3.3ЊC(5ЊC) in 24 h (48 h) were observed fol- compared to those found under similar conditions over lowing the intrusion of cold air masses into the region in freshwater (Steenburgh et al. 2000; see their Fig. 3). Given September and October. the current salinity of Gunnison and Gilbert Bays, the ratio The GSL is a terminal lake (i.e., it has no outlet) and of saturation vapor pressure over saline water to saturation can be up to eight times as saline as ocean water. Currently, vapor pressure over freshwater is approximately 0.70 and the lake is divided by an earthen railroad causeway that 0.94, respectively. Due to this reduction in saturation vapor limits mixing between the northern and southern sections pressure, upward moisture ¯uxes calculated using a bulk (Sturm 1980; Butts 1980; Newby 1980), named Gunnison aerodynamic formula [Krishnamurti and Bounoua 1996, Bay and Gilbert Bay, respectively. Gunnison Bay has only their Eq. (8.2)] would be eliminated or negative over Gun- limited freshwater in¯ow and generally features salinity nison (Gilbert) Bay if the difference between the lake- near saturation (27%). Salinity in Gilbert Bay, which has surface temperature and near-surface dewpoint tempera- several freshwater inlets, has ranged from 6% to 15% and ture was 5ЊC (0.9ЊC) or smaller.2 The implications of sa- during December 1998 was near 9%. Due to the high salinity, the lake never freezes over except near freshwater inlets. Because the lake never freezes over and can warm 2 This example was calculated using a lake temperature of 5ЊC, the rapidly, lake-effect snow is possible from early fall through most common temperature observed during GSLE events (Steenburgh late spring (Steenburgh et al. 2000). The salinity also acts et al. 2000). Over the range of lake temperatures observed during to reduce saturation vapor pressure and latent heat ¯uxes GSLE events, this result varies by 13% or less.

Unauthenticated | Downloaded 09/28/21 02:24 PM UTC 1304 MONTHLY WEATHER REVIEW VOLUME 129 linity on moisture ¯uxes and precipitation will be exam- observed sounding at SLC showed veering winds with ined further in the companion article by Onton and height from the surface to 700 hPa implying low-level Steenburgh (2001). warm advection ahead of the trough (Fig. 3d). Condi- Several steeply sloped mountain ranges extending to tions were not favorable for lake-effect precipitation over 3000 m are located south and east of the GSL (Fig. with southerly to southwesterly ¯ow, a series of stable 1). To the east and southeast are the Wasatch Mountains, layers, and 5Њ±20ЊC dewpoint depressions evident at low which are oriented roughly meridionally and rise abrupt- levels. ly to elevations of 2500±3500 m. South of the GSL, Twelve hours later at 0000 UTC 7 December, shortly the Oquirrh Mountains rise directly from the south shore after the onset of lake-effect precipitation, the 500-hPa to heights of 2500±3250 m, while to the southwest, the (700-hPa) trough axis had moved over (downstream of) Stansbury Mountains reach similar altitudes. Lowland SLC (Figs. 4b,c). Although the signature of this trough regions between these mountain ranges, including the was relatively weak at the surface, low-level winds grad- Salt Lake and Tooele Valleys, are approximately 25 km ually became northwesterly to northerly (Fig. 4a), and wide and feature broadly sloped relief that may also low-level cold advection developed over northern Utah produce orographic precipitation enhancement. For ex- as inferred from backing winds in the SLC sounding ample, the city of Tooele (TOO), 17 km from the GSL near and below 650 hPa (Fig. 4d). In fact, the lowest shoreline, is located 215 m above lake level, while 700-hPa temperatures were now located just upstream broadly sloped benches on the western, southern, and of northern Utah (Fig. 4b). Visible satellite imagery eastern sides of the Salt Lake Valley are 150±400 m showed the passage of a band of clouds across the GSL above lake level. This lowland relief is comparable to with the 700-hPa trough between 1400 and 1900 UTC, that found east of Lake Ontario and northern Lake Mich- but no precipitation was reported over northern Utah igan where signi®cant orographic enhancement of lake- (not shown). The large-scale pattern described above is effect precipitation occurs (Muller 1966; Hjelmfelt similar to that found at the onset time of lake-effect 1992; Niziol et al. 1995), while the adjacent mountain events by Steenburgh et al. (2000). ranges described above are substantially higher. Other Other characteristics of the environment were also important orographic features include the Great Salt favorable for the development of lake-effect precipita- Lake Desert, a lowland area west of the lake, and the tion. With lake-surface (HAT) and 700-hPa tempera- Raft River Mountains northwest of the lake. Thus, ¯ow tures (SLC) of 5ЊC and Ϫ15.9ЊC, respectively, the lake± from the northwest, which is associated with lake-effect 700-hPa temperature difference of 20.9ЊC (12.4 K kmϪ1) storms (Carpenter 1993; Steenburgh et al. 2000), must exceeded the 16ЊC threshold required for GSLE pre- traverse substantial topography before moving over the cipitation identi®ed by Steenburgh et al. (2000). Al- GSL. though an upper-level sounding upstream of the GSL was not available, the SLC sounding that was taken downstream of the GSL showed small dewpoint de- 3. Observational analysis of the 7 December 1998 pressions throughout most of the troposphere (Fig. 4d). snowband Low-level lapse rates were near moist adiabatic and, a. Large-scale analysis although the observed surface parcel at SLC had no convective available potential energy, a surface parcel To examine the large-scale evolution of the 7 De- de®ned using air temperature and dewpoint observations cember 1998 snowband event, regional-scale analyses from HAT exhibited a limited amount of positive buoy- from the RUC2 and upper-air soundings from SLC are ancy (not shown). Finally, the lake±land temperature presented in Figs. 3±6. At 1200 UTC 6 December 1998, difference, calculated using the SLC air temperature and roughly 12 h prior to the onset of lake-effect precipi- HAT lake-surface temperature, was 8ЊC, near the mean tation, a large-scale upper-level trough was located over value for GSLE events (Steenburgh et al. 2000). Such the western United States, with the 500-hPa trough axis conditions favor localized surface heating, boundary extending equatorward from eastern Washington into layer destabilization, and the development of land- southern California (Fig. 3c). The 700-hPa trough axis breeze circulations and low-level convergence over the was just west of the Utah±Nevada border, with a region GSL. of signi®cant moisture [i.e., relative humidity (RH) Ͼ By 1200 UTC 7 December 1998, the 500-hPa trough 70%] collocated with and upstream of this feature (Fig. was located well downstream of Utah and an upper- 3b). There was a weak contrast in temperature across level ridge was building over the western United States the trough with 700-hPa temperatures over southern (Fig. 5c). At this time, lake-effect precipitation was oc- Utah near Ϫ12ЊC, compared to Ϫ16ЊC over western curring in a solitary wind-parallel band extending from Washington and Oregon. A sea level pressure low center the GSL into the Tooele Valley. Over northern Utah, was located northwest of Las Vegas beneath a region moist (RH Ͼ 80Њ%) north to northwesterly ¯ow was of 500-hPa cyclonic absolute vorticity advection (Figs. evident at 700 hPa with the lowest temperatures at this 3a,c). Weak sea level pressure troughing extended north- level located just south of the GSL (Fig. 5b). Sea level eastward from the low center into northern Utah. The high pressure was found over eastern Nevada with light

Unauthenticated | Downloaded 09/28/21 02:24 PM UTC JUNE 2001 STEENBURGH AND ONTON 1305 surface winds over northern Utah (Fig. 5a). The sound- temperatures ranged from Ϫ2Њ to Ϫ10ЊC, with the highest ing was moist (dewpoint depressions Ͻ5ЊC) and con- temperatures found over and near the GSL. ditionally unstable below 650 hPa, with a strong in- Over the next 75 min, snowband A intensi®ed and at version near 500 hPa (Fig. 5d). In addition, the lake± 0515 UTC was located near the western shoreline (Fig. 700-hPa temperature difference was 22.5ЊC (13.0 K 7b). Meanwhile, a second snowband (snowband B) de- kmϪ1) and the lake±land temperature difference was veloped over the southernmost arm of the GSL and north- 10ЊC. east portion of the Tooele Valley. The wind ¯ow in the By 0000 UTC 8 December 1998, lake-effect precip- northern Tooele Valley was con¯uent toward the northern itation had ended. At this time, the 500-hPa ridge axis half of this band. Weak con¯uence into the northern por- was moving over northern Utah and the sea level high tion of snowband A is also suggested by the northwest pressure system was centered over eastern Utah (Figs. surface wind at GNI and north-northeast surface wind at 6a,c). At 700 hPa, temperatures had climbed to Ϫ12ЊC HAT. (Fig. 6b), presumably from large-scale subsidence be- By 0630 UTC snowbands A and B were beginning neath the building upper-level ridge and, as can be in- to merge into a solitary snowband (Fig. 7c). Snowband ferred from veering winds with height at SLC (Fig. 6d), A had just passed over HAT where winds shifted from warm advection in the lower and middle troposphere. northerly to northwesterly, suggesting low-level con¯u- The SLC sounding also shows that the inversion base ence along the northern portion of the snowband axis that was previously located near 500 hPa had lowered (Fig. 8b). Signi®cant changes in temperature or dew- to 700 hPa (cf. Figs. 5d and 6d). In addition, the lake± point were, however, not observed (cf. Figs. 7b,c; dew- 700-hPa temperature difference was 18.5ЊC (10.5 K point not plotted). Fifteen minutes later snowband A kmϪ1), and the lake±land temperature difference was moved westward back across HAT resulting in a wind under 5ЊC. These values were near or below the minima shift back to northerly (Fig. 8b; radar analysis not observed during lake-effect events by Steenburgh et al. shown). Farther downstream, surface winds beneath (2000). Correspondingly, only shallow, nonprecipitating snowband A appeared to be divergent over the western cumulus were observed over the region. Tooele Valley, perhaps due to convective out¯ow (Fig. 7c). In the eastern Tooele Valley, surface winds remained con¯uent toward the axis of snowband B. b. Mesoscale structure The radar re¯ectivity analysis for 0815 UTC shows Observations from the Salt Lake City WSR-88D the solitary snowband that developed from the merger (KMTX) radar and MesoWest surface networks, presented of snowbands A and B at one of its most organized in Figs. 7±9, illustrate the mesoscale structure of the GSLE stages (Fig. 7d). At this time the snowband extended event. Between 2200 UTC 6 December and 0400 UTC 7 from just west of HAT southeastward over the Tooele December, after the passage of the surface and upper-level Valley and was nearly parallel to the ¯ow at PRP. Re- troughs, disorganized convective cells forming primarily ¯ectivity values of 20±30 dBZ composed much of the over the lake and moving downstream to the southeast snowband and likely represent moderate to heavy snow. were observed in radar analyses (not shown). By 0400 Isolated re¯ectivity values approaching 40 dBZ were UTC 7 December, the last of these cells were drifting into observed within the band over the GSL, Tooele Valley, the Tooele Valley and the ®rst long-lived snowband (snow- and western slopes of the Oquirrh Mountains. At this band A) began to form near the western shoreline of the time, con¯uent ¯ow that was previously observed over GSL (Fig. 7a). This snowband was roughly parallel to the the Tooele Valley beneath snowband B was weakening wind ¯ow on Promontory Point (PRP), a mountaintop as winds were becoming northerly or northwesterly. observing site approximately 800 m above lake level that Over the next 135 min the snowband became more roughly represents a steering-layer wind for lake-effect meridionally oriented and by 1030 UTC extended from convection.3 Weak low-level con¯uence into the northern near the center of the GSL southward into the Tooele end of the snowband, as observed during similar banded Valley (Fig. 7e). On the mesoscale, surface wind ob- events over Lakes Michigan and Ontario (e.g., Peace and servations continued to suggest that the northern portion Sykes 1966; Passarelli and Braham 1981; Braham 1983), of the snowband was associated with low-level con¯u- was suggested by a shift in surface winds from northerly ence. Surface winds at HAT veered from northerly to to northwesterly as the snowband passed over Gunnison westerly with snowband passage, as occurred between Island (GNI; Fig 8a). Elsewhere, surface winds were gen- 0600 and 0700 UTC, although the wind shift appeared erally light and northwesterly to northeasterly. Surface to follow the passage of the re¯ectivity band by 15±30 min (Fig. 8b). In addition, overlake convergence was suggested by the westerly wind at HAT and north-northeasterly wind at Antelope Island (Fig. 7e; see Fig. 1 for locations). This mesoscale wind pattern may have been 3 Following Steenburgh et al. (2000) and the experience of local meteorologists, the steering layer for lake-effect convection is gen- related to the development of land-breeze circulations erally considered to be 800±600 hPa. PRP is located at roughly 780 due to localized heating over the lake surface. Temper- hPa. atures over and near the GSL were generally higher than

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FIG. 8. Time series of wind observations from 0300 to 1600 UTC 7 Dec 1998 at (a) GNI, (b) HAT, and (c) PRP. Full and half barbs denote 5 and 2.5 m sϪ1, respectively. Arrows mark ap- proximate time that snowband A crossed observation site. those at surrounding locations, but a lack of wind ob- During the next 90 min the snowband gradually de- servations prevented diagnosis of wind ¯ows along the teriorated into a broad area of precipitation with em- western and eastern shorelines. Farther downstream, bedded convective cores that was drifting northeastward over the northern Tooele Valley, winds were strongly by 1445 UTC (Fig. 7g). At this time, surface winds dif¯uent (Fig. 7e), and at the observing site near the appeared convergent over the GSL, but the snowband GSL shoreline where the Oquirrh Mountains rise abrupt- structure and intensity were beginning to decay for two ly, surface winds had shifted from northeasterly to west- reasons. First, the near-steering-layer wind at PRP was erly (cf. Figs. 7d,e). Although the cause of the dif¯uent weakening and beginning to veer to westerly (Fig. 8c), wind pattern over the northern Tooele Valley at this time a direction with a much shorter overwater fetch. Second, was not clear, it is possible that it was produced by warm advection and subsidence were producing rapid convective out¯ow associated with precipitation and stabilization at midlevels, limiting the depth of surface- diabatic cooling beneath the downstream portion of the based convection (Figs. 5 and 6). By 1900 UTC the snowband. Compared to the relatively steady con¯uent near-steering-layer winds at PRP were west-southwest- ¯ow beneath the upstream portion of the snowband over erly, the lake-effect precipitation area had drifted east- the GSL, surface winds throughout the event were more ward, and new cells were no longer forming (Fig. 7h). variable and occasionally dif¯uent near the downstream portion of the snowband over the northern Tooele Val- c. Radar composite and snowfall distribution ley. By 1315 UTC the snowband extended southeastward To summarize the distribution and intensity of snow- from the GSL over the western Salt Lake Valley (Fig. fall during this event, a composite radar image was gen- 7f). With clearing skies, temperatures dropped rapidly erated from the 155 lowest-elevation (0.5Њ) radar scans to Ϫ10ЊC or lower in the central and western Tooele taken from 0000 to 1455 UTC, which encompasses the Valley, substantially lower than temperatures over the period when lake-effect precipitation was falling over GSL. As a result, thermally driven downvalley and off- the Tooele Valley (Fig. 9). This involved computing the shore winds developed in this area. Overall, the regional percentage of time that re¯ectivity values exceeded 10 wind pattern suggests the presence of low-level con- dBZ at each point within each radar scan (hereafter the vergence over the GSL and near the axis of the snow- 10-dBZ frequency of occurrence or 10-dBZ FO). This band. method was originally developed by Slemmer (1998)

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FIG. 10. MM5 54-, 18-, 6-, and 2-km domains.

FIG. 9. Frequency of occurrence (%) of lowest-elevation angle hPa in the boundary layer to 30 hPa in the upper tro- (0.5Њ) base re¯ectivity values greater than or equal to 10 dBZ (10- posphere.4 Precipitation processes were parameterized dBZ FO) from 0000 to 1455 UTC 7 Dec 1998 and observed snowfall in all four domains using a mixed-phase microphysical totals (cm) from 0000 UTC 7 Dec to 0000 UTC 8 Dec. 10-dBZ FO parameterization that included predictive equations for shaded according to scale at upper left. Topographic contours shown every 500 m in solid lines (see Fig. 1 for elevations). Lake outline cloud ice, cloud water, rain, and snow and allowed for shown with dashed line. supercooled water below 0ЊC and unmelted snow above 0ЊC (Grell et al. 1995). The Kain±Fritsch cumulus parameterization (Kain and Fritsch 1993) was used in the and was used by Steenburgh et al. (2000) to describe 54-, 18-, and 6-km domains to represent subgrid-scale the GSLE precipitation distribution as a function of var- convective precipitation. Boundary layer processes were ious wind and thermodynamic variables. The composite parameterized using the so-called Blackadar scheme that re¯ectivity analysis shows that during the event a band accounts for the vertical mixing of horizontal wind, tem- of frequent returns stretched from near HAT to the west- perature, mixing ratio, cloud water, and cloud ice in the ern slopes of the Oquirrh Mountains (see Fig. 1 for boundary layer (Blackadar 1976, 1979; Zhang and An- locations), with a secondary 10-dBZ FO maximum in thes 1982). One modi®cation was made to the boundary the western Salt Lake Valley where the snowband was layer parameterization to account for the impact of lake resident for a shorter period of time. The highest 10- salinity on saturation vapor pressure and surface mois- dBZ FO region (60%±80%) extended in a narrow band ture ¯uxes. North of the railroad causeway (Fig. 1), the from near the southernmost tip of the GSL to TOO. saturation vapor pressure of lake water was set to 70% Snowfall totals of 25, 30, and 36 cm (18.8-mm liquid of that observed for freshwater. This reduction was equivalent for the latter) were observed at reporting sites based on recent salinity observations in the northern in this region. Outside this band of heavy snowfall, ac- arm of the GSL (27%) and the saturation vapor pressure cumulations were much lower, as indicated by snowfall measurements obtained for lake water by Dickson et al. accumulations of 5 and 8 cm to the south and west, and (1965) and presented in Steenburgh et al. (2000; see reports of trace amounts in the eastern Salt Lake Valley. their Fig. 3). South of the railroad causeway, the saturation vapor pressure was set to 94% of that observed for freshwater based on the observed salinity (9%) and 4. Model simulation estimates of vapor pressure reduction obtained by Steen- a. Mesoscale model description burgh et al. (2000) using Raoult's law. Other model parameterizations included a long- and shortwave at- Simulations by the MM5 were used to further ex- mospheric radiation scheme that accounts for interac- amine the evolution of the 7 December 1998 snowband tions with the atmosphere, clouds, precipitation, and sur- event. The MM5 is a nonhydrostatic ®nite-difference atmospheric model employing a terrain-following sigma vertical coordinate (Grell et al. 1995). Simulations featured four one-way nested domains with grid spacings 4 Speci®cally, the full-sigma levels were located at ␴ ϭ 1.0, 0.99, 0.98, 0.96, 0.93, 0.90, 0.87, 0.84, 0.81, 0.78, 0.75, 0.72, 0.69, 0.66, of 54, 18, 6, and 2 km, respectively (Fig. 10). Thirty- 0.63, 0.60, 0.57, 0.54, 0.51, 0.48, 0.45, 0.42, 0.39, 0.36, 0.33, 0.30, six variably spaced full-sigma levels were used in the 0.27, 0.24, 0.21, 0.18, 0.15, 0.12, 0.09, 0.06, 0.03, 0.0, with the model vertical with resolution varying from approximately 10 top at 100 hPa.

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analyses that were available on a 1Њ lat ϫ 1Њ long grid. The GSL temperature was set to 278 K, the mean lake- surface temperature at HAT during the event period. Four-dimensional data assimilation (FDDA) was used to constrain large-scale error growth in the 54- and 18- km domains. Following Stauffer and Seaman (1990), this involved using Newtonian nudging to relax the model simulation to the gridded analyses that were generated using the methods described above. Linear interpolation in time was used between the analyses, which were at 12-h intervals. For the 54-km domain, FDDA was used during the entire 36-h simulation, while FDDA was used for the 18-km domain for only the ®rst 12 h. Initial analyses for the 6- and 2-km domains were generated by interpolation of analyses from their parent grids since the density of available observations was not suf®cient to adequately resolve features on scales consistent with their grid resolutions. Four-dimensional data assimilation was not used on these domains, although degradation of forecast skill from large-scale error growth should be reduced because of the superior lateral boundary conditions provided by the use of FDDA on the outer domains (Vukicevic and Paegle 1989). The FIG. 11. Topography used in the 2-km domain. Elevation (m) shad- 6-km domain was initialized at the same time as the 54- ed following scale at bottom. and 18-km domains (1200 UTC 6 December), while the 2-km domain was initialized 12 h later at 0000 UTC 7 face (Dudhia 1989), and the Klemp and Durran (1983) December. Because of computational resource limita- radiative upper boundary condition. tions, the 2-km domain was run after the integration of Observed terrain data, bilinearly interpolated onto the the coarser-resolution domains was complete, with MM5 grid and ®ltered with a two-pass smoother/de- boundary conditions provided by hourly output ®les smoother, provided the model terrain. For the 6- and from the 6-km domain. 2-km domains, a 30-s resolution dataset was used, while 10- and 5-min resolution data was used for the 54- and 18-km domains, respectively. All land use information b. Simulated large-scale evolution was derived from a 10-min resolution dataset, though Analyses from the 18-km domain are presented in the land use and elevation near the GSL was corrected Figs. 12±14 to examine the large-scale evolution of to match the lake shoreline. The topography for the the model simulation. At 0000 UTC 7 December, the 2-km domain represents most of the major terrain fea- simulated 500-hPa trough axis extended from Arizona tures of northern Utah, although mountain crest levels to northern Idaho (Fig. 12c) and the lowest 700-hPa and slopes are somewhat lower and less steep than ob- temperatures were located upstream of the GSL (Fig. served (cf. Figs. 1 and 11). 12b). Over northern Utah, northwesterly ¯ow was Analyses for initialization, data assimilation, and lat- found at 700 hPa and the surface, with the relative eral boundary conditions were generated at 12-h inter- humidity at the former level exceeding 70% (Figs. vals from 1200 UTC 6 December to 0000 UTC 8 De- 12a,b).5 The most notable differences between the cember 1998 in the following manner. First, operational simulation and the RUC2 analyses were the lack of a surface and upper-level analyses from the NCEP Eta well-de®ned 700-hPa trough extending northward model (Black 1994; Rogers et al. 1995, 1996), which through eastern Utah and the placement of the 500- were available at 80-km horizontal and 50-hPa vertical hPa trough axis approximately 50±100 km too far resolutions, were interpolated onto each domain's hor- west in the vicinity of the GSL (cf. Figs. 4b,c and izontal grid. This provided a ®rst guess for a modi®ed Cressman-style analysis (Benjamin and Seaman 1985) that incorporated rawinsonde and surface data. After the removal of superadiabatic lapse rates below 500 hPa, 5 To facilitate comparison with RUC2 surface-wind analyses, which the analysis was interpolated to sigma coordinates and are for 10 m above ground level (AGL), and MesoWest surface-wind observations, which are generally taken at 10 m AGL, MM5 surface the integrated mean divergence was removed to avoid winds presented in this paper are 10-m winds that were diagnosed the production of spurious gravity waves. Sea surface from the lowest half-sigma-level wind (ϳ40 m AGL) by assuming temperatures were generated from operational NCEP a logarithmic wind pro®le.

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FIG. 12. Surface, upper-level, and SLC skew T±logp analyses from the 18-km domain at 0000 UTC 7 Dec 1998. (a) Sea level pressure (every 2 hPa), 10-m winds (full and half barbs denote 5 and 2.5 m sϪ1, respectively), and 12-h accumulated precipitation (mm, shaded according to scale at upper right). (b) 700-hPa temperature (every 2ЊC), wind [as in (a)], and relative humidity (%, shaded following scale at upper right). Geopotential height trough axis denoted by dashed line. (c) 500-hPa geopotential height (every 60 m) and absolute vorticity (ϫ10Ϫ5 sϪ1, shaded following scale at upper right). Geopotential height trough axis denoted by dashed line. (d) SLC skew T±logp diagram with temperature and dewpoint (ЊC) denoted by heavy solid lines. Lowest level plotted corresponds to lowest half-sigma level (ϳ830 hPa). Short-dashed line represents surface parcel ascent. Filled circle represents model lake temperature. Wind as in (a).

12b,c). The model-derived sounding at SLC showed ing and the simulated surface temperature appeared northwesterly winds extending from the surface to to be too low (Ϫ7.0ЊC) compared to the observed 500 hPa, where the winds abruptly backed to south- (Ϫ3.3ЊC). The latter was mainly a re¯ection of the westerly (Fig. 12d). A conditionally unstable lapse elevation of the model terrain, which for the 18-km rate was found below ϳ750 hPa. The simulated domain was 422 m (45 hPa) higher than the actual sounding agreed well with the observed, although elevation. At a given pressure level, the simulated some minor differences were evident (cf. Figs. 4d and temperature closely resembled the observed temper- 12d). In particular, the observed layer of backing ature. The sounding derived from the higher-resolu- winds near 700 hPa was not found in the model sound- tion 2-km domain, in which the terrain was only 34

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FIG. 13. Same as Fig. 12 except for 1200 UTC 7 Dec 1998. m (4 hPa) higher than the actual terrain, had a near- soundings (cf. Figs. 5d and 13d) revealed a model warm surface temperature of Ϫ4.3ЊC (not shown). bias between 500 and 700 hPa and cold bias near the At 1200 UTC 7 December, the simulated 500-hPa surface, resulting in a more stable low-level lapse rate trough axis was located downstream of Utah and an than observed. It should be noted, however, that mod- upper-level ridge was building over the western United i®cation of the low-level temperature and dewpoint due States (Fig. 13c). At 700 hPa, the lowest temperatures to heat and moisture ¯uxes from the GSL was likely were located near northern Utah where northwesterly underrepresented in the 18-km domain since at this grid ¯ow was found (Fig. 13b). In this region, the simulated spacing only 12 grid points represent the GSL. relative humidity was slightly lower than analyzed by By 0000 UTC 8 December, the simulated 500-hPa the RUC2 (Fig. 5b). At the surface, sea level pressure ridge axis extended from southern California north- in the higher-resolution MM5 showed more mesoscale eastward to central Montana and was just upstream of structure than the RUC2 (cf. Figs. 5a and 13a), but there northern Utah (Fig. 14c). This position was well fore- were no substantial differences in the placement of syn- cast, although the simulated ridge was slightly less am- optic-scale features, including the position of the sea pli®ed than analyzed by the RUC2 (cf. Figs. 6c and level pressure high that was centered over the Great 14c). At 700 hPa, the simulated ¯ow remained north- Basin. Comparison of the simulated and observed westerly over northern Utah with temperatures rising to

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FIG. 14. Same as Fig. 12 except for 0000 UTC 8 Dec 1998. near Ϫ14ЊC over SLC in response to low-level warm c. Simulated mesoscale structure advection and middle-tropospheric subsidence (Fig. 14b). The MM5 correctly centered the sea level pressure To illustrate the simulated mesoscale structure of the high over Utah and also produced more mesoscale struc- 7 December 1998 event, Fig. 15 presents analyses from ture than analyzed by the coarser-resolution RUC2 (cf. the 2-km domain, including the model-diagnosed 10-m Figs. 6a and 14a). The simulated SLC sounding showed wind, lowest half-sigma-level (ϳ40 m AGL) tempera- veering winds with height at low levels, implying warm ture, and vertically integrated precipitation (VIP). For advection, and an isothermal layer between 700 and 600 purposes of model validation, station plots of wind and hPa (Fig. 14d). A weaker stable layer was located be- temperature from several MesoWest observing sites are tween a shallow surface-based mixed layer and the base overlaid on the model analysis. The VIP is the total of the isothermal layer. These features captured the gen- mass of parameterized rain and snow in a model column eral character of the SLC sounding, although the static and is used to illustrate the instantaneous position of stability of the simulated isothermal layer was much the snowband at each analysis time. The modeled VIP weaker than the observed 5ЊC inversion (cf. Figs. 6d can be qualitatively compared to radar re¯ectivity anal- and 14d). Low-level temperatures were also ϳ3ЊC lower yses presented in Fig. 7, although it should be noted than observed. that the former is a column-integrated quantity while

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FIG. 15. Analyses from the 2-km domain valid at (a) 0400, (b) 0530, (c) 0630, (d) 0830, (e) 1500, and (f) 2100 UTC 7 Dec 1998. Lowest half-sigma-level temperature (every 2ЊC), VIP (kg mϪ2, shaded following scale at upper left), and 10-m wind (full and half barbs denote 5 and 2.5 m sϪ1, respectively). Station plots of observed wind (full and half

Unauthenticated | Downloaded 09/28/21 02:24 PM UTC JUNE 2001 STEENBURGH AND ONTON 1313 the latter represents an observation primarily from the At 0630 UTC (Fig. 15c), snowband A was becoming radar sample volume and is affected by particle size and less organized and diminishing in precipitation intensity, shape, as well as other factors such as beam attenuation, although the convergence zone along the western shore- refraction, and spreading. line was in nearly the same position and possessed a At 0400 UTC 7 December, a region of low-level con- similar magnitude as at 0530 UTC. The VIP analysis ¯uence was oriented along the western shoreline of the did not show a continuous band of precipitation; how- GSL (Fig. 15a). Model diagnostics at this and subse- ever, the cloud mass associated with snowband B ex- quent times showed this region of con¯uence was con- tended poleward toward Antelope Island in a well-or- vergent and will hereafter be referred to as a conver- ganized band (not shown). The initial eastward move- gence zone. Simulated VIP was located near the south- ment of the simulated shoreline convergence zone and ern portion of this convergence zone and extended merger of snowbands A and B appeared to be slower downstream along the eastern slopes of the Stansbury than observed (cf. Figs. 7c and 15c). Temperatures in Mountains. Comparison with the corresponding WSR- most locations, including the Tooele and Salt Lake Val- 88D radar re¯ectivity and mesonet analysis shows that leys, were in good agreement, although the simulated this feature represented snowband A, which in the sim- temperatures in the Great Salt Lake Desert had dropped ulation appeared to be forming correctly near the west- to well below observed. The modeled wind ®eld veri®ed ern GSL shore, but with the VIP region located south well against most land-based stations. Wind directions of the radar re¯ectivity band (cf. Figs. 7a and 15a). This at GNI and HAT winds were off by roughly 60Њ due to discrepancy could be due to model error, although, as the model placing the convergence zone too close to the noted above, VIP and radar re¯ectivity are not entirely western shoreline. consistent. Precipitation was also indicated in the radar The simulated precipitation ®eld at 0830 UTC (Fig. analysis over the Tooele Valley. Three weak VIP fea- 15d) was signi®cantly different from observed (Fig. 7d; tures were found in this region. Simulated low-level 0815 UTC). At this time, observed snowbands A and temperatures over the GSL were above Ϫ4ЊC, approx- B had merged into a solitary snowband that extended imately 2ЊC greater than over surrounding regions of a from HAT to the Oquirrh Mountains. The simulated similar elevation (Fig. 15a). This model low-level tem- snowbands, however, were in one of their least orga- perature analysis agreed well with observed tempera- nized stages and were just beginning to merge (Fig. tures at most sites, with differences generally less than 15d). Nevertheless, the simulated convergence zone was 2ЊC (cf. Figs. 7a and 15a). The most notable difference still evident, had moved offshore, and appeared to be was over the Great Salt Lake Desert where the observed well positioned based on the observation from HAT. (simulated) temperature was Ϫ3ЊC(Ϫ6Њ to Ϫ8ЊC). Reintensi®cation and merger of simulated snowbands Wind directions and magnitudes near the convergence A and B occurred over the next few hours in a manner zone and over other regions were also in good agree- that was similar to observed but delayed. This is illus- ment. trated by the evolution of the VIP between 1000 and Over the next 90 min, simulated snowband A re- 1300 UTC (Fig. 16), which can be compared with the mained quasi-stationary and at 0530 UTC the VIP band radar analyses presented in Fig. 7. This sequence illus- extended from about the midpoint of the western GSL trates some of the dif®culties of mesoscale quantitative shoreline to the southeastern slopes of the Stansbury precipitation forecasting with existing modeling sys- Mountains (Fig. 15b). Meanwhile, the second snowband tems. Although surface winds and temperatures were (snowband B) began to organize over the eastern Tooele generally well simulated, and the model simulation was Valley and western slopes of the Oquirrh Mountains. reasonably accurate earlier in the period, errors related The simulated position of both snowbands was excel- to the timing of the merger and propagation of the bands lent, although they did not extend as far poleward as the corresponding radar re¯ectivity band (cf. Figs. 7b were still apparent. and 15b6). Wind and temperature observations at this At 1500 UTC the simulated convergence zone and time indicate that the model was in general agreement snowband were aligned along the major axis of the GSL with observations, although simulated temperatures (Fig. 15e). The overall ¯ow pattern resembled that as- were still too low near the west boundary. sociated with midlake bands over Lake Michigan (e.g., Peace and Sykes 1966; Braham and Kelly 1982; Hjelm- felt 1990), with land breezes from the opposing lake shorelines converging near the lake axis. The largest 6 Since model output was not available at 0515 UTC, there is a simulated lake±land temperature differences were found 15-min difference between these two ®gures. at this time with a narrow tongue of warm air (ϾϪ4ЊC)

← barbs denote 5 and 2.5 m sϪ1, respectively) and temperature (upper left, ЊC) overlaid. Snowband(s) denoted by large capital letter(s). Heavy dashed line represents axis of snowband convergence zone.

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FIG. 16. VIP (kg mϪ2) from the 2-km domain at (a) 1000, (b) 1100, (c) 1200, and (d) 1300 UTC 7 Dec 1998. Shaded according to scale at upper right.

oriented along the convergence zone axis. Over land, a model simulation (0.02 cal cmϪ2 KϪ1 sϪ1/2).7 Since the shallow nocturnal inversion had formed and was stron- areal coverage of the salt slurry is poorly known, it could gest over the Great Salt Lake Desert where near-surface not be accurately speci®ed in the simulation. During the temperatures were Ϫ10ЊC or lower (Figs. 15e and 17). remainder of the simulation, the area of precipitation Temperatures were similar to observed except at the drifted northeastward and weakened as low-level winds Great Salt Lake Desert observing point (S17; see Fig. became southerly to southwesterly and conditions sta- 1 for location) where the simulated temperature was bilized (Fig. 15f). several degrees too low. The persistent model cold bias The total precipitation (liquid water equivalent) pro- at this location may be related to errors in the speci®- duced by the model simulation from 0000 to 1500 UTC cation of land surface properties. The Great Salt Lake 7 December 1998 is presented in Fig. 18. In comparison Desert land surface is composed primarily of salt ¯ats, which at this time of year can be wet enough to form a salt slurry. Such a salt slurry would likely have a 7 The thermal inertia, ␹, is de®ned as ␹ ϭ (␭C )1/2, where ␭ is the thermal inertia closer to water (0.06 cal cmϪ2 KϪ1 sϪ1/2) s thermal conductivity of the land surface layer and Cs is the heat than the desert land surface that was speci®ed in the capacity per unit volume (Grell et al. 1995).

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FIG. 17. Skew T±logp diagram from the 2-km domain at S17 (see Fig. 1 for location) at 1500 UTC 7 Dec 1998. Temperature and dewpoint in heavy solid lines. Heavy dashed line represents surface parcel ascent. FIG. 18. Simulated total precipitation (mm) from the 2-km domain from 0000 to 1500 UTC 7 Dec 1998. Precipitation shaded according to scale at upper right. Topographic contours shown every 500 m in solid lines (see Fig. 1 for elevations). Lake outline shown with dashed with Fig. 9, the model precipitation band stretching from line. just east of Stansbury Island into the eastern Tooele Valley was very close to the observed position. Maxi- 1982; Braham 1983; Hjelmfelt and Braham 1983; mum simulated precipitation in this band was 19.3 mm, Hjelmfelt 1990; Niziol et al. 1995). comparable to, but slightly higher than, the observed The large-scale environment for the event was similar maximum of 18.8 mm at TOO. The simulation also to that identi®ed as favorable for the development of captured the distribution of the precipitation just east of GSLE precipitation in previous climatological studies the Oquirrh Mountains. Based on radar analyses, the (e.g., Carpenter 1993; Steenburgh et al. 2000). Prior to model appears to have overpredicted precipitation over the onset of lake-effect snow, an upper-level trough axis the western Tooele Valley and Stansbury Mountains, passed from west to east across the GSL, causing winds although no surface snowfall measurements were avail- below 500 hPa to veer from southwesterly to north- able for direct validation. Precipitation in this region westerly, low-level lapse rates to destabilize, and higher occurred earlier, was shifted farther west, and extended relative humidity air to move into northern Utah. En- farther downstream in the model simulation than was vironmental conditions during the event were charac- observed (cf. Figs. 7b and 15b). terized by a lake±700-hPa temperature difference of up to 22.5ЊC, a lake±land temperature difference as large 5. Discussion and conclusions as 10ЊC, and conditionally unstable low-level lapse rates. The observational and model-derived analysis de- Lake-effect precipitation began ϳ2200 UTC 6 De- scribed above illustrates the importance of thermally cember when unorganized convective cells formed over driven circulations in producing the 7 December 1998 the lake and moved downstream to the south and east. GSLE snowstorm. Speci®cally, the primary snowband At 0400 UTC 7 December, an organized snowband be- of the event (snowband A) ®rst formed along a land- gan to form near the western shoreline of the GSL. This breeze front near the western shoreline and eventually band was aligned parallel to the steering-layer wind and aligned along the midlake axis as the land-breeze front was associated with an abrupt wind shift and line of pushed eastward, ¯ow along the eastern shoreline be- con¯uence produced by a land-breeze front. This ki- came increasingly offshore, and convergence developed nematic structure was analogous to that found during along the midlake axis. Thus, despite the relatively small similar events over the Great Lakes (e.g., Peace and size of the GSL and presence of intense vertical relief, Sykes 1966; Passarelli and Braham 1981; Braham the underlying mesoscale dynamics responsible for this 1983). As the event progressed, a second region of pre- event appear to be analogous to shoreline and midlake cipitation formed over the southern GSL and eastern snowband events over the Great Lakes (e.g., Peace and Tooele Valley, and by 0815 UTC merged with the orig- Sykes 1966; Passarelli and Braham 1981; Ballentine inal snowband to form a solitary midlake snowband.

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The snowband was aligned along the surface con¯uence snowbands that would affect forecast skill on mesoscale zone, which was now located near the midlake axis, and temporal scales. was generally oriented parallel to the steering-layer The companion paper by Onton and Steenburgh ¯ow. (2001) further describes the processes responsible for By 1445 UTC, the snowband had deteriorated into this GSLE snowstorm using model diagnostics and sen- an area of precipitation with embedded convective sitivity studies. The predictability of this event in a real- cores, drifting northeastward over the GSL. Although time environment is also examined with a series of sim- surface winds appeared to be convergent over the GSL, ulations incorporating varying environmental condi- steering-layer winds were veering to westerly and tem- tions. peratures were increasing aloft as an upper-level ridge developed over the region. Signi®cant lowering of the Acknowledgments. This research was supported by equilibrium level for convection occurred during this National Science Foundation Grant ATM-9634191 and period as the base of a strong inversion that was located NOAA Grants NA67WA0465 and NA77WA0572 to the near 500 hPa at 1200 UTC 7 December lowered to 700 NOAA Cooperative Institute for Regional Prediction at hPa by 0000 UTC 8 December. As a result, environ- the University of Utah. Surface observations were pro- mental conditions were becoming less favorable for vided by MesoWest, a collection of cooperating me- GSLE snowfall due to the shorter overwater fetch and sonets in the western United States. MesoWest data were reduced depth of convection (Carpenter 1993; Steen- collected and processed by John Horel, Mike Splitt, and burgh et al. 2000). By 1900 UTC, precipitation cells Bryan White of the University of Utah, and Larry Dunn were no longer forming over the GSL. and David Zaff of the National Weather Service. Ad- The heaviest storm-total snowfall was found in a 10- ditional observational data were provided by the Data km wide band that extended from the south shore of Support Section of the Scienti®c Computing Division the GSL to the city of Tooele. The maximum observed of NCAR, which is supported by the National Science storm-total snowfall of 36 cm (18.8-mm liquid equiv- Foundation. Use of the MM5 was made possible by the alent) occurred in the city of Tooele. Only trace amounts Mesoscale and Microscale Meteorology Division of of snow were reported 30 km from the accumulation NCAR. Computer time for the model simulation was band. provided by the University of Utah Center for High The nonhydrostatic model simulation, which featured Performance Computing. Special thanks to Justin Cox, an inner nest with 2-km grid spacing and employed four- Larry Dunn, John Horel, Steve Krueger, Jan Paegle, dimensional data assimilation on the 54-km domain for Tom Potter, Andy Siffert, and David Schultz for their the entire simulation, closely matched the large-scale contributions, advice, and scienti®c support. We grate- fully acknowledge the efforts of two anonymous re- evolution of the event, with only small timing or place- viewers, whose constructive evaluations greatly im- ment errors of synoptic systems. The model run also proved the manuscript. produced snowbands that were similar in structure to radar re¯ectivity patterns observed by the KMTX WSR- 88D, although errors in timing of up to 5 h were ob- REFERENCES served and the simulated snowbands appeared to be located farther downstream than the observed re¯ectiv- Arnow, T., 1980: Water budget and water-surface ¯uctuations of Great Salt Lake. Utah Geolog. 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