Multiscale Analysis of the 7 December 1998 Great Salt Lake...

1296 VOLUME 129M O N T H L Y W E A T H E R R E V I E W

q 2001 American Meteorological Society

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

W. JAMES 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 final form 16 October 2000)

ABSTRACT

The large-scale and mesoscale structure of the Great Salt Lake–effect snowstorm of 7 December 1998 isexamined using radar analyses, high-density surface observations, conventional meteorological data, and asimulation by the Pennsylvania State University–National Center for Atmospheric Research fifth generationMesoscale Model (MM5). Environmental conditions during the event were characterized by a lake–700-hPatemperature difference of up to 22.58C, a lake–land temperature difference as large as 108C, and conditionallyunstable low-level lapse rates. The primary snowband of the event formed along a land-breeze front near thewest shoreline of the Great Salt Lake. The snowband then migrated eastward and merged with a weaker snowbandas the land-breeze front moved eastward, offshore flow developed from the eastern shoreline, and low-levelconvergence developed near the midlake axis. Snowfall accumulations reached 36 cm and were heaviest in anarrow, 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 inducingthermally driven circulations and banded precipitation structures similar to those observed in lake-effect regionsof the eastern United States and Canada.

1. Introduction

The prediction of lake-effect snowstorms that developover and downwind of the Great Salt Lake (GSL) is oneof the major forecast challenges facing meteorologistsin northern Utah. Occurring several times each year,Great Salt Lake–effect (GSLE) snowstorms last from afew hours to more than a day, frequently produce snow-falls of 20–50 cm, and have contributed to the staterecord 129-cm lowland storm-total snowfall that wasobserved near Salt Lake City (SLC) from 24 to 28 Feb-ruary 1998 (Carpenter 1993; Slemmer 1998; Steenburghet al. 2000). Despite significant improvement in obser-vational technologies and numerical forecast systems,GSLE snowstorms remain difficult to predict with leadtimes of more than a few hours.

Previous studies have identified the climatologicalcharacteristics, large-scale conditions, and mesoscaleprecipitation structures associated with GSLE snow-storms. Based on lake-effect events identified by visualobservations and spotter reports, Carpenter (1993)found that GSLE snowstorms were associated with post-

Corresponding author address: Dr. W. James Steenburgh, De-partment of Meteorology, University of Utah, 135 South 1460 EastRoom 819, Salt Lake City, UT 84112-0110.E-mail: [email protected]

cold-frontal northwesterly flow at 700 hPa, a lake–700-hPa temperature difference of at least 178C (which ap-proximately represents a dry adiabatic lapse rate), andan absence of stable layers or inversions near or below700 hPa.1 Steenburgh et al. (2000) used observationsfrom a recently installed National Weather ServiceWeather Surveillance Radar-1988 Doppler (WSR-88D)to identify GSLE events between September 1994 andMay 1998. During this period, 16 well-defined GSLEevents were observed, with the synoptic, mesoscale, andconvective characteristics of these events examined us-ing National Centers for Environmental Prediction(NCEP) Rapid Update Cycle version 2 analyses (RUC2;Benjamin et al. 1991, 1994), SLC radiosonde obser-vations, and local WSR-88D radar observations. In ad-dition to supporting the findings of Carpenter (1993),Steenburgh et al. (2000) also found that GSLE eventstend to occur during periods of positive lake–land tem-perature differences, usually exceeding 68C, and aremost active during the overnight and early morninghours. It was hypothesized that the positive lake–landtemperature difference results in the development of

1 Due to the elevation of the GSL (;1280 m above mean sea level),surface and 700-hPa observations are used instead of surface and850-hPa observations as is commonly done in studies of lake-effectsnowstorms over the Great Lakes (e.g., Niziol et al. 1995).

JUNE 2001 1297S T E E N B U R G H A N D O N T O N

FIG. 1. Geographic features of northern Utah. Surface elevation inmeters shaded according to scale at bottom left. Station locationsdiscussed in text are Salt Lake City (SLC), Tooele (TOO), Hat Island(HAT), Gunnison Island (GNI), Great Salt Lake Desert (S17), andthe Salt Lake City NEXRAD radar site (KMTX). Railroad causewayidentified by a dashed line.

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 periodof missing lake-surface temperature data from HAT.

land breezes and low-level convergence that focus thedevelopment of convection over the GSL. The greaterfrequency of lake-effect precipitation during the over-night and early morning hours may be related to thediurnal modulation of the lake–land temperature dif-ference and associated land-breeze convergence, similarto that suggested by Passarelli and Braham (1981) overLake Michigan.

GSLE snowstorms share many similarities with lake-effect snowstorms over the Great Lakes region of the

United States (Carpenter 1993; Steenburgh et al. 2000).Wiggin (1950) described the general characteristics oflake-effect snowstorms in the Great Lakes region, in-cluding their potential for large accumulations and sig-nificant variations in snowfall over short spatial scales.Additionally, Wiggin (1950) noted that such stormswere favored in polar continental air masses during pe-riods of large lake–air temperature differences, near-adiabatic lapse rates, and long overwater fetches. Peaceand Sykes (1966) studied a lake-effect snowband usinga mesoscale surface observing network over the easternend of Lake Ontario. It was found that a narrow con-vergence line accompanied the snowband and it washypothesized that surface sensible heating caused theformation of the snowband, with winds aloft controllingthe location and movement of the band. Subsequentstudies over the Great Lakes have identified a varietyof lake-effect precipitation structures including (i) broadarea coverage, which may include multiple wind-par-allel bands or open cells (Kelly 1982, 1984); (ii) shore-line bands that form roughly parallel to the lee shoredue to the convergence of a land breeze with the large-scale wind field (Ballentine 1982; Braham 1983; Hjelm-felt and Braham 1983; Hjelmfelt 1990); (iii) midlakebands that form when the large-scale flow is parallel tothe long axis of a lake and a lake–land temperaturecontrast exists (Peace and Sykes 1966; Passarelli andBraham 1981; Braham 1983; Hjelmfelt 1990; Niziol etal. 1995); and (iv) mesoscale vortices that form in apolar air mass under conditions of a weak surface pres-sure gradient and large lake–air temperature differential(Forbes and Merritt 1984).

Precipitation during GSLE events is most frequentlycharacterized by the irregular development of radar ech-oes over and downstream of the GSL (Steenburgh et al.2000). The most commonly observed organized precip-itation structures are solitary wind-parallel bands resem-bling midlake bands found over the Great Lakes, and


FIG. 3. Regional RUC2 analyses and observed SLC upper-air sounding at 1200 UTC 6 Dec 1998. (a) Sea levelpressure (every 2 hPa) and 10-m winds (full and half barbs denote 5 and 2.5 m s21, respectively). (b) 700-hPa temperature(every 28C), wind [as in (a)], and relative humidity (%, shaded following scale at upper right). Geopotential heighttrough axis denoted by dashed line. (c) 500-hPa geopotential height (every 60 m) and absolute vorticity (31025 s21,shaded following scale at upper right). Geopotential height trough axis denoted by dashed line. (d) SLC skew T–logpdiagram with temperature and dewpoint (8C) denoted by heavy solid lines. Short-dashed line represents surface parcelascent. Filled circle represents lake temperature. Wind as in (a).

broad-area coverage precipitation shields that form nearthe lee shoreline. In addition, GSLE precipitation some-times occurs in concert with orographic precipitation,or within a broader-scale precipitation shield associatedwith synoptic-scale lifting. Significant enhancement ofGSLE events can occur when lake-induced precipitationfeatures, such as solitary wind-parallel bands, extendover the downstream orography.

Several studies have used numerical models to ex-amine lake-effect snowstorm dynamics (e.g., Lavoie1972; Ballentine 1982; Hjelmfelt and Braham 1983;Hjelmfelt 1990). Using a three-layer primitive equation

model, Lavoie (1972) found that frictional convergencedue to land–water roughness contrasts, and surface sen-sible heating due to lake–air temperature differences,produce upward vertical motion and elevated inversionheights near the lee shoreline of Lake Erie. The lake–air temperature difference was found to be dominant.Hjelmfelt (1990, 1992) examined the importance of low-level instability, lake–land temperature difference, sen-sible and latent heat fluxes, topography, capping inver-sions, and upstream moisture in producing lake-effectsnowstorms over Lake Michigan. He found that bothshoreline-parallel and midlake snowbands were favored


FIG. 4. Same as Fig. 3 except for 0000 UTC 7 Dec 1998.

by strong lake–land temperature differences, weak sta-bility, and the absence of capping inversions at lowelevations. Moderate cross-lake flow enhanced land-breeze-induced convergence, thus strengthening shore-line-parallel bands. Midlake snowbands, however, werefavored by strong wind flow parallel to the long axis ofthe lake. Weaker wind flows combined with strong lake–land temperature differences tended to produce meso-scale vortices instead of midlake bands. Upstream mois-ture was also found to be important in enhancing lake-effect precipitation and land-breeze strength due to la-tent heat release from condensation. Ballentine et al.(1998) described a successful simulation of a Lake On-tario snowband using the Pennsylvania State Univer-sity–National Center for Atmospheric Research fifthgeneration Mesoscale Model (MM5). The simulationreproduced the observed precipitation distribution, al-

though changes in snowband location in response to theevolving synoptic-scale flow had timing errors of a fewhours.

The purpose of this paper, and the companion articleby Onton and Steenburgh (2001), is to describe the evo-lution and physical processes responsible for a GSLEsnowstorm that occurred on 7 December 1998. Snowfallaccumulations of up to 36 cm were produced by theevent, which featured a wind-parallel snowband thatdeveloped near the western shoreline of the GSL andbecame aligned along the midlake axis as it moved east-ward and merged with a weaker snowband. Specificquestions that will be addressed in the two papers in-clude the following.

R What are the underlying mesoscale dynamics respon-sible for the development of GSLE snowbands? Are



solitary wind-parallel bands over the GSL producedprimarily by thermally driven land-breeze conver-gence?

R How important are topographic effects such as oro-graphic uplift and low-level flow blocking and chan-neling? To what degree are GSLE events triggered orenhanced by such local orographic effects?

R How do sensible and latent heat fluxes influence thedevelopment and intensity of lake-effect precipita-tion? Does the hypersaline composition of the GSLsignificantly affect latent heat flux (compared to fresh-water) and snowband evolution or intensity?

R Does frictional convergence due to land–water rough-ness contrasts influence the development of GSLEsnowbands?

R Can present-day mesoscale models accurately simu-late the mesoscale circulations and precipitation pat-

terns observed during GSLE snowstorms? Does the‘‘fixed’’ surface forcing of the lake and surroundingtopography extend predictability, or do small errorsin surface characteristics and the upstream flow char-acteristics limit forecast skill?

The mesoscale structure and evolution of the 7 De-cember 1998 GSLE snowstorm is examined in the pre-sent paper using conventional meteorological data, high-density surface observations provided by MesoWest, acollection of cooperating mesonets in the western Unit-ed States, and a numerical simulation by the nonhy-drostatic MM5. Section 2 describes the regional orog-raphy and unique characteristics of GSL hydrology,composition, and air–lake interactions. Section 3 pre-sents a detailed observational analysis of the 7 Decem-ber 1998 event using RUC2 analyses, radar observa-



tions, and MesoWest surface observations. Then, section4 uses a mesoscale model simulation to further examinethe mesoscale structure and evolution of the event. Asummary and discussion of major results follow in sec-tion 5. Further diagnosis of the dynamics and predict-ability of the 7 December 1998 event is presented inOnton and Steenburgh (2001).

2. The Great Salt Lake and surroundingtopography

There are several unique aspects of the land surfaceproperties and orography of northern Utah that influencethe development of lake-effect precipitation (Fig. 1).These include the region’s intense and complex verticalrelief, and the varying hydrologic structure, thermalcharacteristics, and hypersaline composition of the GSL.

The GSL is the largest body of water in the United Stateswest of the Great Lakes. It currently occupies an areaof ;4400 km2, is about 120 km long and 45 km wide,and has an average (maximum) depth of only 4.8 (10)m. Due to the lack of a drainage outlet, the lake’s sizefluctuates due to interseasonal and interannual variationsin precipitation and evaporation, and has ranged from2500 to 6200 km2 in area and from 1278 to 1284 m insurface elevation since the mid-1850s (Arnow 1980;Wold et al. 1996).

Due to the GSL’s shallow depth, climatological lake-surface temperatures exhibit little lag relative to cli-matological mean air temperatures at Salt Lake City(Steenburgh et al. 2000; see their Fig. 2). The averagelake-surface temperature exhibits a maximum (mini-mum) near 1 August (1 February), similar to the timingof the maximum (minimum) mean air temperature at


FIG. 7. Lowest-elevation (0.58) base-reflectivity analysis from the KMTX WSR-88D radar and MesoWest surfaceobservations at (a) 0400, (b) 0515, (c) 0630, (d) 0815, (e) 1030, (f ) 1315, (g) 1445, and (h) 1900 UTC 7 Dec 1998.Radar reflectivity shaded according to scale at upper left. Station observations include wind barbs (full and half barbsdenote 5 and 2.5 m s21, respectively), temperature (8C; upper left), and three-digit identifier for selected stations (lowerright). 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 throughsummer, the mean lake temperature is similar to themean air temperature at SLC, but during the fall throughearly winter, the mean lake-surface temperature exceedsthe mean air temperature by 28–38C.

Carpenter (1993) suggested that lake-surface tempera-tures may correlate with the preceding week’s mean airtemperature. In the past, estimates of lake-surface tem-perature using this method were necessary for operationalforecasting due to the lack of real-time observations. How-ever, starting in late summer 1998, lake-surface temper-atures have been observed at a MesoWest site installed atHat Island (HAT; see Fig. 1 for location). A comparisonbetween the mean daily lake-surface temperature and mean

daily air temperatures at SLC and HAT for the initial five-month observation period is presented in Fig. 2. This figureclearly illustrates the seasonal decline in both lake-surfaceand mean air temperature. Note, however, that from Sep-tember to December, lake-surface temperatures were gen-erally 28–38C greater than the mean air temperature, whilein January, lake-surface temperatures were similar to themean air temperature. This is in rough agreement with thetwice-monthly observations presented by Steenburgh et al.(2000), although they showed higher lake-surface tem-peratures persisting into mid-January. Also evident in Fig.2 are more rapid lake-surface and mean air temperaturechanges associated with transient synoptic weather sys-tems. Specifically, lake-surface temperature changes of as


FIG. 7. (Continued )

much as 3.38C (58C) in 24 h (48 h) were observed fol-lowing the intrusion of cold air masses into the region inSeptember and October.

The GSL is a terminal lake (i.e., it has no outlet) andcan be up to eight times as saline as ocean water. Currently,the lake is divided by an earthen railroad causeway thatlimits mixing between the northern and southern sections(Sturm 1980; Butts 1980; Newby 1980), named GunnisonBay and Gilbert Bay, respectively. Gunnison Bay has onlylimited freshwater inflow and generally features salinitynear saturation (27%). Salinity in Gilbert Bay, which hasseveral freshwater inlets, has ranged from 6% to 15% andduring December 1998 was near 9%. Due to the highsalinity, the lake never freezes over except near freshwaterinlets. Because the lake never freezes over and can warmrapidly, lake-effect snow is possible from early fall throughlate spring (Steenburgh et al. 2000). The salinity also actsto reduce saturation vapor pressure and latent heat fluxes

compared to those found under similar conditions overfreshwater (Steenburgh et al. 2000; see their Fig. 3). Giventhe current salinity of Gunnison and Gilbert Bays, the ratioof saturation vapor pressure over saline water to saturationvapor pressure over freshwater is approximately 0.70 and0.94, respectively. Due to this reduction in saturation vaporpressure, upward moisture fluxes calculated using a bulkaerodynamic formula [Krishnamurti and Bounoua 1996,their Eq. (8.2)] would be eliminated or negative over Gun-nison (Gilbert) Bay if the difference between the lake-surface temperature and near-surface dewpoint tempera-ture was 58C (0.98C) or smaller.2 The implications of sa-

2 This example was calculated using a lake temperature of 58C, themost common temperature observed during GSLE events (Steenburghet al. 2000). Over the range of lake temperatures observed duringGSLE events, this result varies by 13% or less.


linity on moisture fluxes and precipitation will be exam-ined further in the companion article by Onton andSteenburgh (2001).

Several steeply sloped mountain ranges extending toover 3000 m are located south and east of the GSL (Fig.1). To the east and southeast are the Wasatch Mountains,which are oriented roughly meridionally and rise abrupt-ly to elevations of 2500–3500 m. South of the GSL,the Oquirrh Mountains rise directly from the south shoreto heights of 2500–3250 m, while to the southwest, theStansbury Mountains reach similar altitudes. Lowlandregions between these mountain ranges, including theSalt Lake and Tooele Valleys, are approximately 25 kmwide and feature broadly sloped relief that may alsoproduce orographic precipitation enhancement. For ex-ample, the city of Tooele (TOO), 17 km from the GSLshoreline, is located 215 m above lake level, whilebroadly sloped benches on the western, southern, andeastern sides of the Salt Lake Valley are 150–400 mabove lake level. This lowland relief is comparable tothat found east of Lake Ontario and northern Lake Mich-igan where significant orographic enhancement of lake-effect precipitation occurs (Muller 1966; Hjelmfelt1992; Niziol et al. 1995), while the adjacent mountainranges described above are substantially higher. Otherimportant orographic features include the Great SaltLake Desert, a lowland area west of the lake, and theRaft River Mountains northwest of the lake. Thus, flowfrom the northwest, which is associated with lake-effectstorms (Carpenter 1993; Steenburgh et al. 2000), musttraverse substantial topography before moving over theGSL.

3. Observational analysis of the 7 December 1998snowband

a. Large-scale analysis

To examine the large-scale evolution of the 7 De-cember 1998 snowband event, regional-scale analysesfrom the RUC2 and upper-air soundings from SLC arepresented in Figs. 3–6. At 1200 UTC 6 December 1998,roughly 12 h prior to the onset of lake-effect precipi-tation, a large-scale upper-level trough was located overthe western United States, with the 500-hPa trough axisextending equatorward from eastern Washington intosouthern California (Fig. 3c). The 700-hPa trough axiswas just west of the Utah–Nevada border, with a regionof significant moisture [i.e., relative humidity (RH) .70%] collocated with and upstream of this feature (Fig.3b). There was a weak contrast in temperature acrossthe trough with 700-hPa temperatures over southernUtah near 2128C, compared to 2168C over westernWashington and Oregon. A sea level pressure low centerwas located northwest of Las Vegas beneath a regionof 500-hPa cyclonic absolute vorticity advection (Figs.3a,c). Weak sea level pressure troughing extended north-eastward from the low center into northern Utah. The

observed sounding at SLC showed veering winds withheight from the surface to 700 hPa implying low-levelwarm advection ahead of the trough (Fig. 3d). Condi-tions were not favorable for lake-effect precipitationwith southerly to southwesterly flow, a series of stablelayers, and 58–208C dewpoint depressions evident at lowlevels.

Twelve hours later at 0000 UTC 7 December, shortlyafter the onset of lake-effect precipitation, the 500-hPa(700-hPa) trough axis had moved over (downstream of )SLC (Figs. 4b,c). Although the signature of this troughwas relatively weak at the surface, low-level winds grad-ually became northwesterly to northerly (Fig. 4a), andlow-level cold advection developed over northern Utahas inferred from backing winds in the SLC soundingnear and below 650 hPa (Fig. 4d). In fact, the lowest700-hPa temperatures were now located just upstreamof northern Utah (Fig. 4b). Visible satellite imageryshowed the passage of a band of clouds across the GSLwith the 700-hPa trough between 1400 and 1900 UTC,but no precipitation was reported over northern Utah(not shown). The large-scale pattern described above issimilar to that found at the onset time of lake-effectevents by Steenburgh et al. (2000).

Other characteristics of the environment were alsofavorable for the development of lake-effect precipita-tion. With lake-surface (HAT) and 700-hPa tempera-tures (SLC) of 58C and 215.98C, respectively, the lake–700-hPa temperature difference of 20.98C (12.4 K km21)exceeded the 168C threshold required for GSLE pre-cipitation identified by Steenburgh et al. (2000). Al-though an upper-level sounding upstream of the GSLwas not available, the SLC sounding that was takendownstream of the GSL showed small dewpoint de-pressions throughout most of the troposphere (Fig. 4d).Low-level lapse rates were near moist adiabatic and,although the observed surface parcel at SLC had noconvective available potential energy, a surface parceldefined using air temperature and dewpoint observationsfrom HAT exhibited a limited amount of positive buoy-ancy (not shown). Finally, the lake–land temperaturedifference, calculated using the SLC air temperature andHAT lake-surface temperature, was 88C, near the meanvalue for GSLE events (Steenburgh et al. 2000). Suchconditions favor localized surface heating, boundarylayer destabilization, and the development of land-breeze circulations and low-level convergence over theGSL.

By 1200 UTC 7 December 1998, the 500-hPa troughwas located well downstream of Utah and an upper-level ridge was building over the western United States(Fig. 5c). At this time, lake-effect precipitation was oc-curring in a solitary wind-parallel band extending fromthe GSL into the Tooele Valley. Over northern Utah,moist (RH . 808%) north to northwesterly flow wasevident at 700 hPa with the lowest temperatures at thislevel located just south of the GSL (Fig. 5b). Sea levelhigh pressure was found over eastern Nevada with light


surface winds over northern Utah (Fig. 5a). The sound-ing was moist (dewpoint depressions ,58C) and con-ditionally unstable below 650 hPa, with a strong in-version near 500 hPa (Fig. 5d). In addition, the lake–700-hPa temperature difference was 22.58C (13.0 Kkm21) and the lake–land temperature difference was108C.

By 0000 UTC 8 December 1998, lake-effect precip-itation had ended. At this time, the 500-hPa ridge axiswas moving over northern Utah and the sea level highpressure system was centered over eastern Utah (Figs.6a,c). At 700 hPa, temperatures had climbed to 2128C(Fig. 6b), presumably from large-scale subsidence be-neath the building upper-level ridge and, as can be in-ferred from veering winds with height at SLC (Fig. 6d),warm advection in the lower and middle troposphere.The SLC sounding also shows that the inversion basethat was previously located near 500 hPa had loweredto 700 hPa (cf. Figs. 5d and 6d). In addition, the lake–700-hPa temperature difference was 18.58C (10.5 Kkm21), and the lake–land temperature difference wasunder 58C. These values were near or below the minimaobserved during lake-effect events by Steenburgh et al.(2000). Correspondingly, only shallow, nonprecipitatingcumulus were observed over the region.

b. Mesoscale structure

Observations from the Salt Lake City WSR-88D(KMTX) radar and MesoWest surface networks, presentedin Figs. 7–9, illustrate the mesoscale structure of the GSLEevent. Between 2200 UTC 6 December and 0400 UTC 7December, after the passage of the surface and upper-leveltroughs, disorganized convective cells forming primarilyover the lake and moving downstream to the southeastwere observed in radar analyses (not shown). By 0400UTC 7 December, the last of these cells were drifting intothe Tooele Valley and the first long-lived snowband (snow-band A) began to form near the western shoreline of theGSL (Fig. 7a). This snowband was roughly parallel to thewind flow on Promontory Point (PRP), a mountaintopobserving site approximately 800 m above lake level thatroughly represents a steering-layer wind for lake-effectconvection.3 Weak low-level confluence into the northernend of the snowband, as observed during similar bandedevents over Lakes Michigan and Ontario (e.g., Peace andSykes 1966; Passarelli and Braham 1981; Braham 1983),was suggested by a shift in surface winds from northerlyto northwesterly as the snowband passed over GunnisonIsland (GNI; Fig 8a). Elsewhere, surface winds were gen-erally light and northwesterly to northeasterly. Surface

3 Following Steenburgh et al. (2000) and the experience of localmeteorologists, the steering layer for lake-effect convection is gen-erally considered to be 800–600 hPa. PRP is located at roughly 780hPa.

temperatures ranged from 228 to 2108C, with the highesttemperatures found over and near the GSL.

Over the next 75 min, snowband A intensified and at0515 UTC was located near the western shoreline (Fig.7b). Meanwhile, a second snowband (snowband B) de-veloped over the southernmost arm of the GSL and north-east portion of the Tooele Valley. The wind flow in thenorthern Tooele Valley was confluent toward the northernhalf of this band. Weak confluence into the northern por-tion of snowband A is also suggested by the northwestsurface wind at GNI and north-northeast surface wind atHAT.

By 0630 UTC snowbands A and B were beginningto merge into a solitary snowband (Fig. 7c). SnowbandA had just passed over HAT where winds shifted fromnortherly to northwesterly, suggesting low-level conflu-ence along the northern portion of the snowband axis(Fig. 8b). Significant changes in temperature or dew-point were, however, not observed (cf. Figs. 7b,c; dew-point not plotted). Fifteen minutes later snowband Amoved westward back across HAT resulting in a windshift back to northerly (Fig. 8b; radar analysis notshown). Farther downstream, surface winds beneathsnowband A appeared to be divergent over the westernTooele Valley, perhaps due to convective outflow (Fig.7c). In the eastern Tooele Valley, surface winds re-mained confluent toward the axis of snowband B.

The radar reflectivity analysis for 0815 UTC showsthe solitary snowband that developed from the mergerof snowbands A and B at one of its most organizedstages (Fig. 7d). At this time the snowband extendedfrom just west of HAT southeastward over the TooeleValley and was nearly parallel to the flow at PRP. Re-flectivity values of 20–30 dBZ composed much of thesnowband and likely represent moderate to heavy snow.Isolated reflectivity values approaching 40 dBZ wereobserved within the band over the GSL, Tooele Valley,and western slopes of the Oquirrh Mountains. At thistime, confluent flow that was previously observed overthe Tooele Valley beneath snowband B was weakeningas winds were becoming northerly or northwesterly.

Over the next 135 min the snowband became moremeridionally oriented and by 1030 UTC extended fromnear the center of the GSL southward into the TooeleValley (Fig. 7e). On the mesoscale, surface wind ob-servations continued to suggest that the northern portionof the snowband was associated with low-level conflu-ence. Surface winds at HAT veered from northerly towesterly with snowband passage, as occurred between0600 and 0700 UTC, although the wind shift appearedto follow the passage of the reflectivity band by 15–30min (Fig. 8b). In addition, overlake convergence wassuggested by the westerly wind at HAT and north-north-easterly wind at Antelope Island (Fig. 7e; see Fig. 1 forlocations). This mesoscale wind pattern may have beenrelated to the development of land-breeze circulationsdue to localized heating over the lake surface. Temper-atures over and near the GSL were generally higher than


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 s21, respectively. Arrows mark ap-proximate time that snowband A crossed observation site.

those at surrounding locations, but a lack of wind ob-servations prevented diagnosis of wind flows along thewestern and eastern shorelines. Farther downstream,over the northern Tooele Valley, winds were stronglydiffluent (Fig. 7e), and at the observing site near theGSL shoreline where the Oquirrh Mountains rise abrupt-ly, surface winds had shifted from northeasterly to west-erly (cf. Figs. 7d,e). Although the cause of the diffluentwind pattern over the northern Tooele Valley at this timewas not clear, it is possible that it was produced byconvective outflow associated with precipitation anddiabatic cooling beneath the downstream portion of thesnowband. Compared to the relatively steady confluentflow beneath the upstream portion of the snowband overthe GSL, surface winds throughout the event were morevariable and occasionally diffluent near the downstreamportion of the snowband over the northern Tooele Val-ley.

By 1315 UTC the snowband extended southeastwardfrom the GSL over the western Salt Lake Valley (Fig.7f). With clearing skies, temperatures dropped rapidlyto 2108C or lower in the central and western TooeleValley, substantially lower than temperatures over theGSL. As a result, thermally driven downvalley and off-shore winds developed in this area. Overall, the regionalwind pattern suggests the presence of low-level con-vergence over the GSL and near the axis of the snow-band.

During the next 90 min the snowband gradually de-teriorated into a broad area of precipitation with em-bedded convective cores that was drifting northeastwardby 1445 UTC (Fig. 7g). At this time, surface windsappeared convergent over the GSL, but the snowbandstructure and intensity were beginning to decay for tworeasons. First, the near-steering-layer wind at PRP wasweakening and beginning to veer to westerly (Fig. 8c),a direction with a much shorter overwater fetch. Second,warm advection and subsidence were producing rapidstabilization at midlevels, limiting the depth of surface-based convection (Figs. 5 and 6). By 1900 UTC thenear-steering-layer winds at PRP were west-southwest-erly, the lake-effect precipitation area had drifted east-ward, and new cells were no longer forming (Fig. 7h).

c. Radar composite and snowfall distribution

To summarize the distribution and intensity of snow-fall during this event, a composite radar image was gen-erated from the 155 lowest-elevation (0.58) radar scanstaken from 0000 to 1455 UTC, which encompasses theperiod when lake-effect precipitation was falling overthe Tooele Valley (Fig. 9). This involved computing thepercentage of time that reflectivity values exceeded 10dBZ at each point within each radar scan (hereafter the10-dBZ frequency of occurrence or 10-dBZ FO). Thismethod was originally developed by Slemmer (1998)


FIG. 9. Frequency of occurrence (%) of lowest-elevation angle(0.58) base reflectivity values greater than or equal to 10 dBZ (10-dBZ FO) from 0000 to 1455 UTC 7 Dec 1998 and observed snowfalltotals (cm) from 0000 UTC 7 Dec to 0000 UTC 8 Dec. 10-dBZ FOshaded according to scale at upper left. Topographic contours shownevery 500 m in solid lines (see Fig. 1 for elevations). Lake outlineshown with dashed line.

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

and was used by Steenburgh et al. (2000) to describethe GSLE precipitation distribution as a function of var-ious wind and thermodynamic variables. The compositereflectivity analysis shows that during the event a bandof frequent returns stretched from near HAT to the west-ern slopes of the Oquirrh Mountains (see Fig. 1 forlocations), with a secondary 10-dBZ FO maximum inthe western Salt Lake Valley where the snowband wasresident for a shorter period of time. The highest 10-dBZ FO region (60%–80%) extended in a narrow bandfrom near the southernmost tip of the GSL to TOO.Snowfall totals of 25, 30, and 36 cm (18.8-mm liquidequivalent for the latter) were observed at reporting sitesin this region. Outside this band of heavy snowfall, ac-cumulations were much lower, as indicated by snowfallaccumulations of 5 and 8 cm to the south and west, andreports of trace amounts in the eastern Salt Lake Valley.

4. Model simulation

a. Mesoscale model description

Simulations by the MM5 were used to further ex-amine the evolution of the 7 December 1998 snowbandevent. The MM5 is a nonhydrostatic finite-differenceatmospheric model employing a terrain-following sigmavertical coordinate (Grell et al. 1995). Simulations fea-tured four one-way nested domains with grid spacingsof 54, 18, 6, and 2 km, respectively (Fig. 10). Thirty-six variably spaced full-sigma levels were used in thevertical with resolution varying from approximately 10

hPa in the boundary layer to 30 hPa in the upper tro-posphere.4 Precipitation processes were parameterizedin all four domains using a mixed-phase microphysicalparameterization that included predictive equations forcloud ice, cloud water, rain, and snow and allowed forsupercooled water below 08C and unmelted snow above08C (Grell et al. 1995). The Kain–Fritsch cumulus pa-rameterization (Kain and Fritsch 1993) was used in the54-, 18-, and 6-km domains to represent subgrid-scaleconvective precipitation. Boundary layer processes wereparameterized using the so-called Blackadar scheme thataccounts for the vertical mixing of horizontal wind, tem-perature, mixing ratio, cloud water, and cloud ice in theboundary layer (Blackadar 1976, 1979; Zhang and An-thes 1982). One modification was made to the boundarylayer parameterization to account for the impact of lakesalinity on saturation vapor pressure and surface mois-ture fluxes. North of the railroad causeway (Fig. 1), thesaturation vapor pressure of lake water was set to 70%of that observed for freshwater. This reduction wasbased on recent salinity observations in the northernarm of the GSL (27%) and the saturation vapor pressuremeasurements obtained for lake water by Dickson et al.(1965) and presented in Steenburgh et al. (2000; seetheir Fig. 3). South of the railroad causeway, the satu-ration vapor pressure was set to 94% of that observedfor freshwater based on the observed salinity (9%) andestimates of vapor pressure reduction obtained by Steen-burgh et al. (2000) using Raoult’s law. Other modelparameterizations included a long- and shortwave at-mospheric radiation scheme that accounts for interac-tions with the atmosphere, clouds, precipitation, and sur-

4 Specifically, the full-sigma levels were located at s 5 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, 0.66,0.63, 0.60, 0.57, 0.54, 0.51, 0.48, 0.45, 0.42, 0.39, 0.36, 0.33, 0.30,0.27, 0.24, 0.21, 0.18, 0.15, 0.12, 0.09, 0.06, 0.03, 0.0, with the modeltop at 100 hPa.


FIG. 11. Topography used in the 2-km domain. Elevation (m) shad-ed following scale at bottom.

face (Dudhia 1989), and the Klemp and Durran (1983)radiative upper boundary condition.

Observed terrain data, bilinearly interpolated onto theMM5 grid and filtered with a two-pass smoother/de-smoother, provided the model terrain. For the 6- and2-km domains, a 30-s resolution dataset was used, while10- and 5-min resolution data was used for the 54- and18-km domains, respectively. All land use informationwas derived from a 10-min resolution dataset, thoughthe land use and elevation near the GSL was correctedto match the lake shoreline. The topography for the2-km domain represents most of the major terrain fea-tures of northern Utah, although mountain crest levelsand slopes are somewhat lower and less steep than ob-served (cf. Figs. 1 and 11).

Analyses for initialization, data assimilation, and lat-eral boundary conditions were generated at 12-h inter-vals from 1200 UTC 6 December to 0000 UTC 8 De-cember 1998 in the following manner. First, operationalsurface and upper-level analyses from the NCEP Etamodel (Black 1994; Rogers et al. 1995, 1996), whichwere available at 80-km horizontal and 50-hPa verticalresolutions, were interpolated onto each domain’s hor-izontal grid. This provided a first guess for a modifiedCressman-style analysis (Benjamin and Seaman 1985)that incorporated rawinsonde and surface data. After theremoval of superadiabatic lapse rates below 500 hPa,the analysis was interpolated to sigma coordinates andthe integrated mean divergence was removed to avoidthe production of spurious gravity waves. Sea surfacetemperatures were generated from operational NCEP

analyses that were available on a 18 lat 3 18 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 usedto constrain large-scale error growth in the 54- and 18-km domains. Following Stauffer and Seaman (1990),this involved using Newtonian nudging to relax themodel simulation to the gridded analyses that were gen-erated using the methods described above. Linear in-terpolation 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, whileFDDA was used for the 18-km domain for only the first12 h.

Initial analyses for the 6- and 2-km domains weregenerated by interpolation of analyses from their parentgrids since the density of available observations was notsufficient to adequately resolve features on scales con-sistent with their grid resolutions. Four-dimensional dataassimilation was not used on these domains, althoughdegradation of forecast skill from large-scale errorgrowth should be reduced because of the superior lateralboundary conditions provided by the use of FDDA onthe outer domains (Vukicevic and Paegle 1989). The6-km domain was initialized at the same time as the 54-and 18-km domains (1200 UTC 6 December), while the2-km domain was initialized 12 h later at 0000 UTC 7December. Because of computational resource limita-tions, the 2-km domain was run after the integration ofthe coarser-resolution domains was complete, withboundary conditions provided by hourly output filesfrom the 6-km domain.

b. Simulated large-scale evolution

Analyses from the 18-km domain are presented inFigs. 12–14 to examine the large-scale evolution ofthe model simulation. At 0000 UTC 7 December, thesimulated 500-hPa trough axis extended from Arizonato northern Idaho (Fig. 12c) and the lowest 700-hPatemperatures were located upstream of the GSL (Fig.12b). Over northern Utah, northwesterly flow wasfound at 700 hPa and the surface, with the relativehumidity at the former level exceeding 70% (Figs.12a,b).5 The most notable differences between thesimulation and the RUC2 analyses were the lack of awell-defined 700-hPa trough extending northwardthrough eastern Utah and the placement of the 500-hPa trough axis approximately 50–100 km too farwest in the vicinity of the GSL (cf. Figs. 4b,c and

5 To facilitate comparison with RUC2 surface-wind analyses, whichare for 10 m above ground level (AGL), and MesoWest surface-windobservations, which are generally taken at 10 m AGL, MM5 surfacewinds presented in this paper are 10-m winds that were diagnosedfrom the lowest half-sigma-level wind (;40 m AGL) by assuminga logarithmic wind profile.


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 s21, respectively), and 12-haccumulated precipitation (mm, shaded according to scale at upper right). (b) 700-hPa temperature (every 28C), wind[as in (a)], and relative humidity (%, shaded following scale at upper right). Geopotential height trough axis denotedby dashed line. (c) 500-hPa geopotential height (every 60 m) and absolute vorticity (31025 s21, shaded following scaleat upper right). Geopotential height trough axis denoted by dashed line. (d) SLC skew T–logp diagram with temperatureand dewpoint (8C) denoted by heavy solid lines. Lowest level plotted corresponds to lowest half-sigma level (;830hPa). 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 showednorthwesterly winds extending from the surface to500 hPa, where the winds abruptly backed to south-westerly (Fig. 12d). A conditionally unstable lapserate was found below ;750 hPa. The simulatedsounding agreed well with the observed, althoughsome minor differences were evident (cf. Figs. 4d and12d). In particular, the observed layer of backingwinds near 700 hPa was not found in the model sound-

ing and the simulated surface temperature appearedto be too low (27.08C) compared to the observed(23.38C). The latter was mainly a reflection of theelevation of the model terrain, which for the 18-kmdomain was 422 m (45 hPa) higher than the actualelevation. At a given pressure level, the simulatedtemperature closely resembled the observed temper-ature. The sounding derived from the higher-resolu-tion 2-km domain, in which the terrain was only 34



m (4 hPa) higher than the actual terrain, had a near-surface temperature of 24.38C (not shown).

At 1200 UTC 7 December, the simulated 500-hPatrough axis was located downstream of Utah and anupper-level ridge was building over the western UnitedStates (Fig. 13c). At 700 hPa, the lowest temperatureswere located near northern Utah where northwesterlyflow was found (Fig. 13b). In this region, the simulatedrelative humidity was slightly lower than analyzed bythe RUC2 (Fig. 5b). At the surface, sea level pressurein the higher-resolution MM5 showed more mesoscalestructure than the RUC2 (cf. Figs. 5a and 13a), but therewere no substantial differences in the placement of syn-optic-scale features, including the position of the sealevel pressure high that was centered over the GreatBasin. Comparison of the simulated and observed

soundings (cf. Figs. 5d and 13d) revealed a model warmbias between 500 and 700 hPa and cold bias near thesurface, resulting in a more stable low-level lapse ratethan observed. It should be noted, however, that mod-ification of the low-level temperature and dewpoint dueto heat and moisture fluxes from the GSL was likelyunderrepresented in the 18-km domain since at this gridspacing only 12 grid points represent the GSL.

By 0000 UTC 8 December, the simulated 500-hParidge axis extended from southern California north-eastward to central Montana and was just upstream ofnorthern Utah (Fig. 14c). This position was well fore-cast, although the simulated ridge was slightly less am-plified than analyzed by the RUC2 (cf. Figs. 6c and14c). At 700 hPa, the simulated flow remained north-westerly over northern Utah with temperatures rising to



near 2148C over SLC in response to low-level warmadvection and middle-tropospheric subsidence (Fig.14b). The MM5 correctly centered the sea level pressurehigh over Utah and also produced more mesoscale struc-ture than analyzed by the coarser-resolution RUC2 (cf.Figs. 6a and 14a). The simulated SLC sounding showedveering winds with height at low levels, implying warmadvection, and an isothermal layer between 700 and 600hPa (Fig. 14d). A weaker stable layer was located be-tween a shallow surface-based mixed layer and the baseof the isothermal layer. These features captured the gen-eral character of the SLC sounding, although the staticstability of the simulated isothermal layer was muchweaker than the observed 58C inversion (cf. Figs. 6dand 14d). Low-level temperatures were also ;38C lowerthan observed.

c. Simulated mesoscale structure

To illustrate the simulated mesoscale structure of the7 December 1998 event, Fig. 15 presents analyses fromthe 2-km domain, including the model-diagnosed 10-mwind, lowest half-sigma-level (;40 m AGL) tempera-ture, and vertically integrated precipitation (VIP). Forpurposes of model validation, station plots of wind andtemperature from several MesoWest observing sites areoverlaid on the model analysis. The VIP is the totalmass of parameterized rain and snow in a model columnand is used to illustrate the instantaneous position ofthe snowband at each analysis time. The modeled VIPcan be qualitatively compared to radar reflectivity anal-yses presented in Fig. 7, although it should be notedthat the former is a column-integrated quantity while


FIG. 15. Analyses from the 2-km domain valid at (a) 0400, (b) 0530, (c) 0630, (d) 0830, (e) 1500, and (f ) 2100UTC 7 Dec 1998. Lowest half-sigma-level temperature (every 28C), VIP (kg m22, shaded following scale at upper left),and 10-m wind (full and half barbs denote 5 and 2.5 m s21, respectively). Station plots of observed wind (full and half


←

barbs denote 5 and 2.5 m s21, respectively) and temperature (upper left, 8C) overlaid. Snowband(s) denoted by large capital letter(s). Heavydashed line represents axis of snowband convergence zone.

the latter represents an observation primarily from theradar sample volume and is affected by particle size andshape, as well as other factors such as beam attenuation,refraction, and spreading.

At 0400 UTC 7 December, a region of low-level con-fluence was oriented along the western shoreline of theGSL (Fig. 15a). Model diagnostics at this and subse-quent times showed this region of confluence was con-vergent and will hereafter be referred to as a conver-gence zone. Simulated VIP was located near the south-ern portion of this convergence zone and extendeddownstream along the eastern slopes of the StansburyMountains. Comparison with the corresponding WSR-88D radar reflectivity and mesonet analysis shows thatthis feature represented snowband A, which in the sim-ulation appeared to be forming correctly near the west-ern GSL shore, but with the VIP region located southof the radar reflectivity band (cf. Figs. 7a and 15a). Thisdiscrepancy could be due to model error, although, asnoted above, VIP and radar reflectivity are not entirelyconsistent. Precipitation was also indicated in the radaranalysis over the Tooele Valley. Three weak VIP fea-tures were found in this region. Simulated low-leveltemperatures over the GSL were above 248C, approx-imately 28C greater than over surrounding regions of asimilar elevation (Fig. 15a). This model low-level tem-perature analysis agreed well with observed tempera-tures at most sites, with differences generally less than28C (cf. Figs. 7a and 15a). The most notable differencewas over the Great Salt Lake Desert where the observed(simulated) temperature was 238C (268 to 288C).Wind directions and magnitudes near the convergencezone and over other regions were also in good agree-ment.

Over the next 90 min, simulated snowband A re-mained quasi-stationary and at 0530 UTC the VIP bandextended from about the midpoint of the western GSLshoreline to the southeastern slopes of the StansburyMountains (Fig. 15b). Meanwhile, the second snowband(snowband B) began to organize over the eastern TooeleValley and western slopes of the Oquirrh Mountains.The simulated position of both snowbands was excel-lent, although they did not extend as far poleward asthe corresponding radar reflectivity band (cf. Figs. 7band 15b6). Wind and temperature observations at thistime indicate that the model was in general agreementwith observations, although simulated temperatureswere still too low near the west boundary.

6 Since model output was not available at 0515 UTC, there is a15-min difference between these two figures.

At 0630 UTC (Fig. 15c), snowband A was becomingless organized and diminishing in precipitation intensity,although the convergence zone along the western shore-line was in nearly the same position and possessed asimilar magnitude as at 0530 UTC. The VIP analysisdid not show a continuous band of precipitation; how-ever, the cloud mass associated with snowband B ex-tended poleward toward Antelope Island in a well-or-ganized band (not shown). The initial eastward move-ment of the simulated shoreline convergence zone andmerger of snowbands A and B appeared to be slowerthan observed (cf. Figs. 7c and 15c). Temperatures inmost locations, including the Tooele and Salt Lake Val-leys, were in good agreement, although the simulatedtemperatures in the Great Salt Lake Desert had droppedto well below observed. The modeled wind field verifiedwell against most land-based stations. Wind directionsat GNI and HAT winds were off by roughly 608 due tothe model placing the convergence zone too close to thewestern shoreline.

The simulated precipitation field at 0830 UTC (Fig.15d) was significantly different from observed (Fig. 7d;0815 UTC). At this time, observed snowbands A andB had merged into a solitary snowband that extendedfrom HAT to the Oquirrh Mountains. The simulatedsnowbands, however, were in one of their least orga-nized stages and were just beginning to merge (Fig.15d). Nevertheless, the simulated convergence zone wasstill evident, had moved offshore, and appeared to bewell positioned based on the observation from HAT.

Reintensification and merger of simulated snowbandsA and B occurred over the next few hours in a mannerthat was similar to observed but delayed. This is illus-trated by the evolution of the VIP between 1000 and1300 UTC (Fig. 16), which can be compared with theradar analyses presented in Fig. 7. This sequence illus-trates some of the difficulties of mesoscale quantitativeprecipitation forecasting with existing modeling sys-tems. Although surface winds and temperatures weregenerally well simulated, and the model simulation wasreasonably accurate earlier in the period, errors relatedto the timing of the merger and propagation of the bandswere still apparent.

At 1500 UTC the simulated convergence zone andsnowband were aligned along the major axis of the GSL(Fig. 15e). The overall flow pattern resembled that as-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 lakeshorelines converging near the lake axis. The largestsimulated lake–land temperature differences were foundat this time with a narrow tongue of warm air (.248C)


FIG. 16. VIP (kg m22) 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, ashallow nocturnal inversion had formed and was stron-gest over the Great Salt Lake Desert where near-surfacetemperatures were 2108C or lower (Figs. 15e and 17).Temperatures were similar to observed except at theGreat Salt Lake Desert observing point (S17; see Fig.1 for location) where the simulated temperature wasseveral degrees too low. The persistent model cold biasat this location may be related to errors in the specifi-cation of land surface properties. The Great Salt LakeDesert land surface is composed primarily of salt flats,which at this time of year can be wet enough to forma salt slurry. Such a salt slurry would likely have athermal inertia closer to water (0.06 cal cm22 K21 s21/2)than the desert land surface that was specified in the

model simulation (0.02 cal cm22 K21 s21/2).7 Since theareal coverage of the salt slurry is poorly known, it couldnot be accurately specified in the simulation. During theremainder of the simulation, the area of precipitationdrifted northeastward and weakened as low-level windsbecame southerly to southwesterly and conditions sta-bilized (Fig. 15f).

The total precipitation (liquid water equivalent) pro-duced by the model simulation from 0000 to 1500 UTC7 December 1998 is presented in Fig. 18. In comparison

7 The thermal inertia, x, is defined as x 5 (lCs)1/2, where l is thethermal conductivity of the land surface layer and Cs is the heatcapacity per unit volume (Grell et al. 1995).


FIG. 17. Skew T–logp diagram from the 2-km domain at S17 (seeFig. 1 for location) at 1500 UTC 7 Dec 1998. Temperature and dew-point in heavy solid lines. Heavy dashed line represents surface parcelascent. FIG. 18. Simulated total precipitation (mm) from the 2-km domain

from 0000 to 1500 UTC 7 Dec 1998. Precipitation shaded accordingto scale at upper right. Topographic contours shown every 500 m insolid lines (see Fig. 1 for elevations). Lake outline shown with dashedline.with Fig. 9, the model precipitation band stretching from

just east of Stansbury Island into the eastern TooeleValley was very close to the observed position. Maxi-mum simulated precipitation in this band was 19.3 mm,comparable to, but slightly higher than, the observedmaximum of 18.8 mm at TOO. The simulation alsocaptured the distribution of the precipitation just east ofthe Oquirrh Mountains. Based on radar analyses, themodel appears to have overpredicted precipitation overthe western Tooele Valley and Stansbury Mountains,although no surface snowfall measurements were avail-able for direct validation. Precipitation in this regionoccurred earlier, was shifted farther west, and extendedfarther downstream in the model simulation than wasobserved (cf. Figs. 7b and 15b).

5. Discussion and conclusions

The observational and model-derived analysis de-scribed above illustrates the importance of thermallydriven circulations in producing the 7 December 1998GSLE snowstorm. Specifically, the primary snowbandof the event (snowband A) first formed along a land-breeze front near the western shoreline and eventuallyaligned along the midlake axis as the land-breeze frontpushed eastward, flow along the eastern shoreline be-came increasingly offshore, and convergence developedalong the midlake axis. Thus, despite the relatively smallsize of the GSL and presence of intense vertical relief,the underlying mesoscale dynamics responsible for thisevent appear to be analogous to shoreline and midlakesnowband events over the Great Lakes (e.g., Peace andSykes 1966; Passarelli and Braham 1981; Ballentine

1982; Braham 1983; Hjelmfelt and Braham 1983;Hjelmfelt 1990; Niziol et al. 1995).

The large-scale environment for the event was similarto that identified as favorable for the development ofGSLE precipitation in previous climatological studies(e.g., Carpenter 1993; Steenburgh et al. 2000). Prior tothe onset of lake-effect snow, an upper-level trough axispassed from west to east across the GSL, causing windsbelow 500 hPa to veer from southwesterly to north-westerly, low-level lapse rates to destabilize, and higherrelative humidity air to move into northern Utah. En-vironmental conditions during the event were charac-terized by a lake–700-hPa temperature difference of upto 22.58C, a lake–land temperature difference as largeas 108C, and conditionally unstable low-level lapserates.

Lake-effect precipitation began ;2200 UTC 6 De-cember when unorganized convective cells formed overthe lake and moved downstream to the south and east.At 0400 UTC 7 December, an organized snowband be-gan to form near the western shoreline of the GSL. Thisband was aligned parallel to the steering-layer wind andwas associated with an abrupt wind shift and line ofconfluence produced by a land-breeze front. This ki-nematic structure was analogous to that found duringsimilar events over the Great Lakes (e.g., Peace andSykes 1966; Passarelli and Braham 1981; Braham1983). As the event progressed, a second region of pre-cipitation formed over the southern GSL and easternTooele Valley, and by 0815 UTC merged with the orig-inal snowband to form a solitary midlake snowband.


The snowband was aligned along the surface confluencezone, which was now located near the midlake axis, andwas generally oriented parallel to the steering-layerflow.

By 1445 UTC, the snowband had deteriorated intoan area of precipitation with embedded convectivecores, drifting northeastward over the GSL. Althoughsurface winds appeared to be convergent over the GSL,steering-layer winds were veering to westerly and tem-peratures were increasing aloft as an upper-level ridgedeveloped over the region. Significant lowering of theequilibrium level for convection occurred during thisperiod as the base of a strong inversion that was locatednear 500 hPa at 1200 UTC 7 December lowered to 700hPa by 0000 UTC 8 December. As a result, environ-mental conditions were becoming less favorable forGSLE snowfall due to the shorter overwater fetch andreduced depth of convection (Carpenter 1993; Steen-burgh et al. 2000). By 1900 UTC, precipitation cellswere no longer forming over the GSL.

The heaviest storm-total snowfall was found in a 10-km wide band that extended from the south shore ofthe GSL to the city of Tooele. The maximum observedstorm-total snowfall of 36 cm (18.8-mm liquid equiv-alent) occurred in the city of Tooele. Only trace amountsof snow were reported 30 km from the accumulationband.

The nonhydrostatic model simulation, which featuredan inner nest with 2-km grid spacing and employed four-dimensional data assimilation on the 54-km domain forthe entire simulation, closely matched the large-scaleevolution of the event, with only small timing or place-ment errors of synoptic systems. The model run alsoproduced snowbands that were similar in structure toradar reflectivity patterns observed by the KMTX WSR-88D, although errors in timing of up to 5 h were ob-served and the simulated snowbands appeared to belocated farther downstream than the observed reflectiv-ity bands. Surface winds and temperatures were alsowell simulated compared to surface observations, withthe exception of stronger than observed nocturnal sur-face cooling over the Great Salt Lake Desert that mayhave resulted from errors in the specification of surfaceproperties in that region. The simulated storm-total pre-cipitation agreed well with radar reflectivity compositesand snowfall observations. The maximum simulatedprecipitation was 19.3 mm, slightly greater than the ob-served 18.8 mm, and was found in approximately thesame location as observed. Although the accuracy ofthe quantitative precipitation forecast at 2-km grid spac-ing may raise optimism regarding potential predictiveskill of future high-resolution forecast models, the useof data assimilation to constrain large-scale error growthon the coarser-resolution model grids represents a sig-nificant advantage that would not be available in a real-time environment. Even with this advantage, timing er-rors were observed in the movement and merger of the

snowbands that would affect forecast skill on mesoscaletemporal scales.

The companion paper by Onton and Steenburgh(2001) further describes the processes responsible forthis GSLE snowstorm using model diagnostics and sen-sitivity studies. The predictability of this event in a real-time environment is also examined with a series of sim-ulations incorporating varying environmental condi-tions.

Acknowledgments. This research was supported byNational Science Foundation Grant ATM-9634191 andNOAA Grants NA67WA0465 and NA77WA0572 to theNOAA Cooperative Institute for Regional Prediction atthe University of Utah. Surface observations were pro-vided by MesoWest, a collection of cooperating me-sonets in the western United States. MesoWest data werecollected and processed by John Horel, Mike Splitt, andBryan White of the University of Utah, and Larry Dunnand David Zaff of the National Weather Service. Ad-ditional observational data were provided by the DataSupport Section of the Scientific Computing Divisionof NCAR, which is supported by the National ScienceFoundation. Use of the MM5 was made possible by theMesoscale and Microscale Meteorology Division ofNCAR. Computer time for the model simulation wasprovided by the University of Utah Center for HighPerformance Computing. Special thanks to Justin Cox,Larry Dunn, John Horel, Steve Krueger, Jan Paegle,Tom Potter, Andy Siffert, and David Schultz for theircontributions, advice, and scientific support. We grate-fully acknowledge the efforts of two anonymous re-viewers, whose constructive evaluations greatly im-proved the manuscript.

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