diagnostic and sensitivity studies of the 7 december 1998 ...u0758091/studies/onton2001.pdf · er...

1318 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

Diagnostic and Sensitivity Studies of the 7 December 1998Great Salt Lake–Effect Snowstorm

DARYL J. ONTON AND W. JAMES STEENBURGH

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 processes responsible for the Great Salt Lake–effect snowstorm of 7 December 1998 are examined usinga series of mesoscale model simulations. Localized surface sensible and latent heating are shown to destabilizethe boundary layer over the Great Salt Lake (GSL) and to produce mesoscale pressure troughing, land-breezecirculations, and low-level convergence that lead to the development of the primary band of convective cloudsand precipitation. Model diagnostics and sensitivity studies further illustrate that

R moisture fluxes from the lake surface were necessary to fully develop the snowband;R the hypersaline composition of the GSL did, however, decrease moisture fluxes compared to a body of

freshwater, resulting in a 17% reduction of snowfall;R latent heat release within the cloud and precipitation band intensified overlake pressure troughing, conver-

gence, and precipitation;R orographic effects were not responsible for snowband generation, but they did affect the distribution and

intensity of precipitation in regions where the snowband interacted with downstream terrain; andR surface roughness contrasts across the GSL shoreline did not play a primary role in forming the snowband.

Simulations in which lake-surface temperature and upstream moisture were modified are used to illustrate howsmall errors in the specification of these quantities can impact quantitative precipitation forecasts, potentiallylimiting the utility of high-resolution mesoscale model guidance. Results are compared to those from studies oflake-effect precipitation over the Great Lakes, and the implications for operational forecasting and numericalweather prediction are discussed.

1. Introduction

In the first paper of this series, Steenburgh and Onton(2001) described the structure and evolution of a GreatSalt Lake effect (GSLE) snowstorm using observationaldata and a mesoscale simulation by the nonhydrostaticPennsylvania State University–National Center for At-mospheric Research fifth generation Mesoscale Model(MM5). The event, which occurred on 7 December1998, featured a wind-parallel snowband that developedalong the west shoreline of the Great Salt Lake (GSL),migrated eastward, and eventually merged with a weak-er snowband as it became aligned along the midlakeaxis. Snowfall accumulations reached 36 cm and wereheaviest in a narrow, 10–20-km-wide band extendingdownstream from the GSL. It was shown that the snow-band along the western shoreline formed over a low-level convergence zone associated with a land-breeze

Corresponding author address: Dr. Daryl J. Onton, Department ofMeteorology, University of Utah, 135 South 1460 East Room 819,Salt Lake City, UT 84112-0110.E-mail: [email protected]

front. The snowband aligned along the midlake axis asthe land-breeze front moved eastward and offshore flowfrom the east shoreline intensified. The kinematic struc-ture of the event thus appeared to be analogous to eventsassociated with thermally driven land-breeze conver-gence over the Great Lakes of the eastern United States,including midlake bands produced over Lakes Michiganand Ontario (e.g., Peace and Sykes 1966; Passarelli andBraham 1981; Braham 1983; Hjelmfelt 1990; Niziol etal. 1995).

Although the kinematic structure of the 7 December1998 event was similar to that of midlake bands overthe Great Lakes, the event also may have been influ-enced by unique aspects of the geography of northernUtah, including the presence of intense vertical reliefand the hypersaline content of the GSL. As a result, thepurpose of the present paper is twofold: (i) to furtherinvestigate the physical processes responsible for the 7December 1998 GSLE snowstorm, and (ii) to examineissues related to the predictability of these events bypresent-day numerical models. This is accomplishedwith a detailed analysis of output from the 2-km hori-zontal resolution domain of the simulation presented by

JUNE 2001 1319O N T O N A N D S T E E N B U R G H

FIG. 1. (a) Lowest half-sigma-level (;40 m AGL) temperature(every 28C), vertically integrated precipitation (kg m22, shaded ac-cording to scale at upper right), and 10-m wind (full and half barbsdenote 5 and 2.5 m s21, respectively) from the 2-km domain at 0300UTC 7 Dec 1998. Thick line denotes lake shoreline. (b) Lowest-elevation angle (0.58) base-reflectivity analysis from the Salt LakeCity WSR-88D (KMTX) (shaded according to scale at upper right)and surface observations at 0515 UTC 7 Dec 1998. Station plotsdenote wind (full and half barbs denote 5 and 2.5 m s21, respectively)and temperature (8C, upper left). Dashed line denotes lake outline.Topographic contours (solid) every 500 m.

Steenburgh and Onton (2001) and a series of modelsensitivity studies designed to illustrate the relative im-portance of selected physical and dynamical processessuch as surface sensible and latent heat flux contrasts,topographic blocking and channeling, and frictionalconvergence due to shoreline surface roughness con-trasts. Simulations in which lake-surface temperatureand upstream moisture are modified are used to illustratehow small errors in the specification of these quantitiescan impact quantitative precipitation forecasts and po-tentially limit the utility of high-resolution mesoscalemodel guidance. The model diagnostic analysis is de-scribed in the next section. Then, section 3 describesthe sensitivity studies and is followed by discussion andconclusions in section 4. The reader is referred to Steen-burgh and Onton (2001) for a description of the ge-ography and topography of northern Utah, character-istics of the GSL, evolution of the 7 December 1998event, and configuration of the mesoscale simulation.

2. Model diagnostic and trajectory analysis of the7 December 1998 snowband

a. Formative stage

At 0300 UTC, the simulated snowband was locatedalong the west shoreline of the GSL where there waslow-level convergence between northerly flow over theGSL and northwesterly flow over the west shoreline(Fig. 1a). Detailed analysis of the model simulation ispresented for this time because it best describes thestructure of the snowband during its early stages of de-velopment, although with a timing error of about 2 h(cf. Figs. 1a,b). Model surface analyses showed a tongueof high potential temperature air and a pressure troughlocated just off of the western shoreline of the GSL(Figs. 2a,b).1 Immediately west of these features, con-vergence was maximized along the strong gradients inpotential temperature and pressure that were associatedwith a developing land-breeze front. A similar but weak-er feature was located offshore of the eastern shoreline.Figure 3 shows a cross section of circulation vectors(consisting of the horizontal and vertical velocity com-ponents in the plane of the cross section), virtual po-tential temperature, and cloud and precipitation mixingratio taken perpendicular to the snowband near its far-thest upwind extent (line AB in Fig. 4). Below 775 hPathe virtual potential temperature was nearly constantwith height over the GSL, indicating near-neutral staticstability. Lower virtual potential temperature air wasfound over land to the west and east of the GSL. Ap-proximately 6 km from the western shoreline, low-levelwinds converged beneath a 5 Pa s21 (50 cm s21) updraftthat fed a shallow cloud band that produced only neg-

1 Altimeter setting was used to reduce pressure to sea level to avoidaliasing surface temperature gradients into pressure gradients.

ligible amounts of precipitation (precipitation mixingratios were below the shading threshold in Fig. 3a).Farther downstream, along cross section CD (see Fig.4 for position), a stronger updraft was evident and thecloud band was deeper (Fig. 5). Precipitation fell in anarrow shaft that was approximately 10 km in width.

The surface pressure pattern evident in Fig. 2b, whichfeatured a midlake ridge separated by two near-shorelinetroughs, appeared to be caused by the shoreline ge-ometry. With northerly flow over most of the lake atthis time (see Fig. 1a), high virtual potential temperatureair and a pressure trough were located over the northernbays of the GSL [Gunnison and Bear River; see Fig. 1of Steenburgh and Onton (2001) for location]. To thelee of Promontory Point, a peninsula that extends south-ward into the GSL, lower virtual potential temperatureair and higher pressure were found since flow down-stream of this feature experienced a shorter overwaterfetch and less heating and moistening (Fig. 2). Similarly,


FIG. 2. Simulation from the 2-km domain at 0300 UTC 7 Dec 1998.(a) Lowest half-sigma-level (;40 m AGL) potential temperature (ev-ery 0.28C). (b) Altimeter setting (every 0.025 hPa) and lowest half-sigma-level convergence (31023 s21, shaded according to scale atupper right). Dashed line denotes lake shoreline.

Phillips (1972) found that overlake isotherms tended toparallel the upwind coastline on Lake Ontario.

The strongest pressure gradient was located along thewestern shoreline and resulted in a substantial wind shiftdue to the offshore acceleration of the low-level flow.The resulting convergence provided the necessarymechanism to lift low-level air and produce convectionalong the land-breeze front. Interestingly, the conver-

gence zone and precipitation band were coincident withthe pressure gradient, rather than the pressure trough.This might be expected, however, since the airflow overthe lake would tend to keep the wind speed and directionnearly constant until significant forcing such as the pres-sure gradient near the west shoreline could producesome deflection. In a simulation of a shoreline snowbandover Lake Michigan, Hjelmfelt and Braham (1983) alsoobtained a wind field that had the maximum conver-gence between the pressure minimum and the shoreline.

Three-dimensional trajectories beginning at 1800UTC 6 December and terminating on the lowest half-sigma level [approximately 40 m above ground level(AGL)] at 0300 UTC 7 December further elucidate theprocesses associated with snowband development (Fig.6).2 Trajectories terminating in a line that was roughlyperpendicular to the snowband demonstrate that thehighest low-level temperatures were associated with tra-jectories that experienced the greatest overwater fetch(Fig. 6a). Low-level convergence into the snowband wasevident with a strong deflection of trajectories endingimmediately west of the snowband. This is more clearlyillustrated by Fig. 6b, which shows two lines of trajec-tories terminating to the west and east of the snowband,respectively. Near the western shoreline of the GSL,trajectories were deflected eastward by the strong off-shore pressure gradient acceleration. East of the snow-band, the flow was primarily meridional.

Boundary layer modification by the GSL is illustratedby a series of soundings taken along trajectory 9 of Fig.6a (Fig. 7). At point A (0000 UTC), just upstream ofthe GSL, the sounding featured a shallow surface-basedmixed layer with a near-surface temperature of 24.88C(Fig. 7a). After air following trajectory 9 passed brieflyover the GSL and temporarily again over land, thesounding at point B (0100 UTC) showed little change(Fig. 7b). Significant boundary layer modification oc-curred over the next hour as this air moved over theGSL, and the sounding at point C (0200 UTC) showedthat during this period the boundary layer deepened ;50hPa while the near-surface temperature increased to23.28C (Fig. 7c). Meanwhile, the mean mixing ratiobelow 775 hPa, the approximate top of the boundarylayer, increased from 2.05 to 2.19 g kg21. As the airmoved to point D, where it was beneath the snowband,the near-surface temperature and mean mixing ratio be-low 775 hPa increased to 22.78C and 2.20 g kg21, re-spectively, as the boundary layer increased slightly indepth (Fig. 7d). This series of soundings shows thatsensible heating by the lake warmed, deepened, anddestabilized the boundary layer, while latent heating in-

2 Because of the large amount of storage space required to storemodel output at high temporal resolution, trajectories were calculatedusing 30-min model output. This time difference was selected sincea comparison of trajectories calculated from 5- and 30-min outputshowed no significant differences.


FIG. 3. Cross section AB from the 2-km domain at 0300 UTC 7 Dec 1998 (see Fig. 4 for crosssection location). Virtual potential temperature (solid, every 28C), total cloud (water and ice)mixing ratio (dashed every 0.2 g kg21), total precipitation (snow and rain) mixing ratio (g kg21,shaded according to scale at upper left), and vectors of along-section wind and vertical velocity(following scale at upper right). Lake shorelines denoted by thick solid lines.

FIG. 4. Cross section locations (thick solid lines) and 2-km domaintopography (m, shaded according to scale at bottom). Railroad cause-way marked with thick dashed line. Lake shoreline denoted by solidline.

creased the mean boundary layer moisture. Collectively,the near-surface temperature and mean mixing ratio be-low 775 hPa increased 2.18C and 0.16 g kg21, respec-tively, with the latter representing an 8% increase com-pared to the upstream value.

It is also interesting to examine the impact of lakesalinity on simulated surface latent heat fluxes during thisperiod. North of the railroad causeway (see Fig. 4 forlocation), where a 30% reduction in the freshwater sat-uration vapor pressure was specified in the simulationdue to the observed 27% salinity (Steenburgh and Onton2001), latent heat fluxes reached only 100 W m22 andwere significantly lower than over the lake’s southernhalf, where the salinity was 9% and latent heat fluxesreached over 160 W m22 (Fig. 8). This is in contrast tothe typical distribution of surface latent heat fluxes overwater bodies of uniform salinity, in which latent heatfluxes decrease downwind of the shoreline (Chang andBraham 1991). The impact of lake salinity on the inten-sity of this event will be discussed further in section 3.

b. Mature stage

By 1300 UTC the convergence zone and snowbandwere located near the midlake axis and extended down-stream into the Tooele Valley (Fig. 9). As noted bySteenburgh and Onton (2001), the simulated eastwardmovement of the snowband was delayed, so that thesnowband structure at this time was similar to the ob-


FIG. 5. Same as Fig. 3 except along line CD of Fig. 4.

served structure at 0815 UTC (cf. Figs. 9a,b). The sim-ulated low-level potential temperature pattern over thelake was dominated by a tongue of warm air that ex-tended down the midlake axis and was flanked by in-tense potential temperature gradients near the easternand western shorelines (Fig. 10a). Intensification of thelake–land temperature gradient since 0300 UTC (cf.Figs. 2a and 10a) was due primarily to nocturnal coolingover the surrounding land mass, particularly over theGreat Salt Lake Desert. A pressure trough was locatedalong the warm tongue and strong pressure gradientshad developed near the western and eastern shorelines(Fig. 10b). As a result, the western shore land breezeand the magnitude of the offshore flow near the easternshoreline had intensified, resulting in the developmentof a well-defined convergence zone near the midlakeaxis (Figs. 9a and 10b).

The mesoscale circulations responsible for the snow-band are further illustrated by the cross section of virtualpotential temperature, circulation vectors, and cloud andprecipitation mixing ratio presented in Fig. 11 (line ABin Fig. 4). Compared to 10 h previously (Fig. 3), near-surface virtual potential temperatures to the west (east)of the GSL decreased by 68–88C (28–38C), apparentlydue to nocturnal cooling, resulting in the developmentof a more dense and stable boundary layer surroundingthe GSL. Meanwhile, virtual potential temperatures overthe GSL remained relatively constant. The resultingcontrast in boundary layer temperature and surface pres-sure intensified the offshore flow, particularly from thewestern shoreline. Where the two opposing flows metnear the midlake axis, a narrow 6 Pa s21 (;60 cm s21)

updraft was found beneath the developing cloud andprecipitation band. Interestingly, the convergence zoneand updraft were again not located directly over thepressure trough, similar to 10 h previously. The overallthree-dimensional wind and thermal structure closelyresembled that associated with midlake bands overLakes Michigan and Ontario (e.g., Peace and Sykes1966; Passarelli and Braham 1981; Braham and Kelly1982; Hjelmfelt 1990).

A cross section along the center of the snowband at1300 UTC is presented in Fig. 12 (line EF in Fig. 4).Low-level virtual potential temperature increased alongthe lake axis, indicating reduced stability over the south-ern half of the GSL, where updrafts formed and pro-duced progressively larger cloud and precipitation mix-ing ratios as the flow neared the downwind shoreline.Vertical velocities were largest near the downwindshoreline, where updrafts extended to ;650 hPa, ratherthan over the GSL where surface heating was strongest.This was likely due to the downstream advection ofconvective updrafts in a manner similar to that describedby Lin and Smith (1986). At this time, vertical motionin the immediate vicinity of the Oquirrh Mountains wasweak compared to that within the snowband, althoughin section 3 orographically induced ascent will be shownto enhance precipitation at other times.

Three-dimensional trajectories, beginning at 0400UTC 7 December and terminating on the lowest half-sigma level (;40 m AGL) at 1300 UTC 7 December(Fig. 13) show stronger lake-induced circulations thanat 0300 UTC (see trajectories 7–10 in Fig. 6). Althoughthe snowband was aligned along the major lake axis,


FIG. 6. Nine-hour three-dimensional trajectories from the 2-kmdomain ending at 0300 UTC 7 Dec on the lowest half-sigma level(;40 m AGL). Trajectory width denotes pressure altitude accordingto scale at upper right. Lowest half-sigma-level potential temperaturecontours every 0.28C shown only near lake level. Terrain heightsabove 1500 m shaded in gray. Heavy solid line denotes lake shoreline.(a) Trajctories ending in a line perpendicular to the snowband. Parcellocations at 30-min intervals along trajectory 9 denoted with largedots. Locations at 0000, 0100, 0200, and 0300 UTC denoted by lettersa, b, c, and d, respectively. (b) Trajectories surrounding snowband.

because of the stronger offshore flow, trajectories in Fig.13b actually had a shorter overwater fetch than thoseterminating around the snowband 10 h earlier (Fig. 6).Boundary layer modification along a trajectory that hada relatively long overwater fetch and terminated nearthe warm tongue (trajectory 18, Fig. 13b) is illustratedby the soundings in Fig. 14. At 1000 UTC, the air fol-lowing trajectory 18 was approaching the northernshoreline of the Bear River Bay (point a, Fig. 13b). Atthis location, a shallow surface inversion, presumablyproduced by nocturnal cooling, was located near the

surface and another shallow stable layer was locatednear 800 hPa (Fig. 14a). The lowest half-sigma leveltemperature and dewpoint were 28.98 and 210.48C,respectively. One hour later (1100 UTC), this air waslocated over the Bear River Bay and had been over waterfor ;30 min (point b, Fig. 13b). At this point, sensibleheating over the lake surface had raised the near-surfacetemperature to 26.38C and a shallow surface-basedmixed layer had developed (Fig. 14b). Near-surfacemoisture increased slightly with the dewpoint reaching29.88C. The stable layer near 800 hPa was still presentand had lowered slightly. By 1200 UTC, the air hadreached point c, which was located over the less salinesouthern region of the GSL (Fig. 13a). The near-surfacetemperature and dewpoint had risen to 24.78 and29.58C, respectively, and the stable layer near 875 hPahad weakened (Fig. 14c). During the next hour, the near-surface temperature and dewpoint increased to 23.88and 28.48C, respectively, all stable layers and inver-sions below ;550 hPa eroded away, and the soundingbecame conditionally unstable up to ;675 hPa (Fig.14d).

As an example of the vertical circulations associatedwith the snowband at this time, three-dimensional tra-jectories following air that was ingested into the snow-band are presented in Fig. 15. These trajectories beganat 1000 UTC 7 December and terminated at 1300 UTC7 December on the 0.855 sigma level (761 hPa), whichwas near the top of the snowband. Air following tra-jectories 1–3 originated at low levels west of the GSL,converged toward the snowband axis, rose rapidly, andmoved downstream. Trajectory 1, which was the out-ermost trajectory west of the snowband at 1300 UTC,was the innermost relative to the midlake axis at thebeginning of the trajectory. As air following trajectory1 approached the low-level convergence zone, it as-cended through the snowband. Trajectories 2 and 3 fol-lowed air that originated farther to the south and re-mained near the surface until it rose rapidly when it wasingested in the narrow updraft that supported the snow-band (e.g., Fig. 11). Air following trajectories endingin the eastern portion of the snowband followed a slight-ly different evolution. Trajectories 4 and 5 originatednorth of the GSL, ;80 hPa above lake level, eventuallydescended to low levels as they converged toward thesnowband, then rose rapidly as they were ingested intothe snowband. The air following trajectory 6 was in-gested into the snowband near its farthest upwind extent,rose to 757 hPa, remained aloft and moved slightly awayfrom the snowband, descended slightly to 803 hPa, con-verged again toward the snowband, and rose to 761 hPanear the snowband.

The evolution of pressure, temperature, water vapormixing ratio, and cloud water and ice mixing ratio alongtrajectory 1 are displayed in Fig. 16. At 1000 UTC, theair following trajectory 1 was located west of the shore-line of the GSL at 849 hPa. It descended to 875 hPa(;5 hPa above lake level) as it followed downward


FIG. 7. Skew T–logp diagrams of simulated temperature (8C), dewpoint (8C), and wind (full and half barbs denote5 and 2.5 m s21, respectively) along trajectory 9 of Fig. 6a. Pressure scale in hPa. (a) 0000 UTC (point a). (b) 0100UTC (point b). (c) 0200 UTC (point c). (d) 0300 UTC (point d, end of trajectory).

sloping terrain toward the lake shoreline after 1030UTC. Potential temperature and water vapor mixing ra-tio increased 3.4 K and 0.4 g kg21 from 1100 to 1130UTC as this air moved over the GSL. As it approachedthe low-level convergence zone and associated updraftprior to 1230 UTC, it began to rise upward with satu-ration and cloud water development beginning at around1220 UTC. At approximately 1240 UTC, the air reachedits highest elevation and the cloud water mixing ratiowas near its peak value. The air then subsided graduallyas it began to exit the snowband, while the remainingcloud condensate dissipated and potential temperatureand mixing ratio remained nearly constant.

The physical picture obtained from the analysis aboveis consistent with the findings of studies over the GreatLakes that illustrate the role of thermally driven land-breeze circulations in generating solitary wind-parallelsnowbands (e.g., Passarelli and Braham 1981; Hjelmfelt

and Braham 1983; Hjelmfelt 1990). In the present case,a pressure trough and low-level convergence were in-duced by upward surface sensible and latent heat fluxesover the GSL. Initially, the convergence zone developedalong the boundary between a developing land breezefrom the western shoreline and synoptic-scale northerlyflow over the GSL. This provided a mechanism to liftlow-level air so that, aided by boundary layer heating,moistening, and destabilization over the lake, a band ofclouds and precipitation formed roughly parallel to thenorthwesterly steering-layer flow. As nocturnal coolingincreased the lake–land temperature contrast, landbreezes from the western and eastern shorelines inten-sified, increasing the strength of the overlake conver-gence and producing the most organized, mature stageof snowband evolution. The shoreline geometry wasalso found to influence the pattern of temperature andpressure over the lake, similar to observations over Lake


FIG. 8. Surface latent heat flux (every 10 W m22) from the 2-kmdomain at 0300 UTC 7 Dec 1998. Dashed line denotes lake shoreline.

FIG. 10. Same as Fig. 2 except at 1300 UTC 7 Dec 1998.

FIG. 9. (a) Same as Fig. 1a except for 1300 UTC. (b) Same as Fig.1b except for 0815 UTC.

Ontario (Phillips 1972), with the highest near-surfacetemperatures and lowest pressures associated with thelongest overwater trajectories.

3. Sensitivity experiments

a. Experimental design

To examine the relative importance of different phys-ical processes on the evolution of the 7 December 1998snowband, a series of sensitivity experiments was con-ducted in which either a particular parameter or processwas modified or withheld from the 2-km domain (Table


FIG. 11. Same as Fig. 3 except at 1300 UTC 7 Dec 1998.

FIG. 12. Same as Fig. 3 except at 1300 UTC 7 Dec 1998 and along line EF of Fig. 4.

1). Unless otherwise noted, all other characteristics ofthese simulations, including the lateral boundary con-ditions provided to the 2-km domain, were the same asthose used for the control simulation (CTL; describedin Steenburgh and Onton 2001). Although the sensitivityexperiments presented below provide insight into therelative importance of the physical processes influencing

the development, evolution, and predictability of thisevent, they do not fully quantify the contribution of aparticular process to the storm-total snowfall since ul-timately it is the nonlinear, synergistic interaction be-tween processes (e.g., Uccellini 1990; Stein and Alpert1993; Alpert et al. 1995) that yields a snowband of theobserved intensity.


FIG. 13. Nine-hour three-dimensional trajectories from the 2-kmdomain ending at 1300 UTC 7 Dec on the lowest half-sigma level(;40 m AGL). Trajectory width denotes pressure altitude accordingto scale at upper right. Lowest half-sigma-level potential temperaturecontours every 0.28C shown only near lake level. Terrain heightsabove 1500 m shaded in gray. Heavy solid line denotes lake shoreline.(a) Trajectories ending in a line perpendicular to the snowband. (b)Trajectories surrounding snowband. Parcel locations at 30-min in-tervals along trajectory 18 denoted with large dots. Locations at 1000,1100, 1200, and 1300 UTC denoted by letters a, b, c, and d, re-spectively.

b. Effects of surface fluxes and lake salinity

Surface fluxes of heat and moisture have been shownto be important physical mechanisms leading to lake-effect snow in the Great Lakes region (e.g., Lavoie1972; Passarelli and Braham 1981; Hjelmfelt 1990) and,as described above, appear critical to the developmentof the 7 December GSLE snowband. As discussed bySteenburgh et al. (2000), by Steenburgh and Onton(2001), and in section 2, the GSL is a hypersaline bodyof water that has a lower saturation vapor pressure than

freshwater. As a result, moisture fluxes are reduced fromwhat they would be for freshwater. In this section, aseries of simulations is described that examine the im-pact of lake salinity, latent heat flux, and sensible heatflux on the evolution of the 7 December 1998 snowband.Knowledge of the sensitivity of mesoscale simulationsto the specification of lake salinity is useful given thehistorical variability of salinity in the GSL (Steenburghet al. 2000).

Figure 17 presents the total precipitation and mean10-m winds for 0000–1500 UTC 7 December fromCTL, a simulation with freshwater latent heat fluxes(FRESH), a simulation in which latent heat fluxes overthe GSL were ignored (NOLHFLX), and a simulationin which sensible heat fluxes over the GSL were ignored(NOSHFLX). CTL produced 17% less domain-averagedprecipitation than FRESH (Table 1; cf. Figs. 17a,b). Thisdifference was due to enhanced precipitation in FRESHduring the stage of the simulation where the snowbandwas resident near the western shoreline of the GSL, aswell as a tendency for the snowband to extend fartherupstream during most of the simulation. In addition, theprecipitation maximum in the eastern Tooele Valley in-creased from 19.3 to 23.1 mm (Table 1). Thus, by re-ducing surface moisture fluxes, lake salinity had a sig-nificant impact on the amount of snowfall produced dur-ing this event. Figure 18a shows the difference in al-timeter setting and 10-m wind between the twosimulations (i.e., FRESH–CTL) and illustrates that, al-though the FRESH mean wind field is similar to that ofCTL, slightly enhanced convergence is evident near thewestern shoreline. Although not evident in this figure,marginally stronger pressure troughing was evident inFRESH. It appears that the surface moisture fluxes inFRESH led to additional latent heating by condensationand fusion within the cloud band, promoting more vig-orous circulations within the snowband. Similar com-plementary processes in lake-effect storms have beennoted by Lavoie (1972), Hjelmfelt and Braham (1983),and Hjelmfelt (1990).

NOLHFLX produced 49% less domain-averaged pre-cipitation than the control run (Table 1; cf. Figs. 17a,c).Precipitation was substantially lower than that producedby CTL, did not extend as far upstream over the GSL,and was confined primarily to areas of orographic as-cent. Thus, comparison of CTL and NOLHFLX sug-gests that, although the GSL is hypersaline and smallin size, the limited flux of moisture from its surface andresulting increase in low-level mixing ratio was nec-essary to fully develop the intense lake-induced snow-band in this event. This result also illustrates the highlysensitive nature of snowband development since the rel-atively small increase in low-level moisture over theGSL resulted in a dramatic increase in storm-total pre-cipitation.

NOSHFLX produced 28% less domain-averaged pre-cipitation than CTL (Table 1), and, as observed inNOLHFLX, precipitation in the former fell primarily


FIG. 14. Skew T–logp diagrams of simulated temperature (8C), dewpoint (8C), and wind (full and half barbs denote5 and 2.5 m s21, respectively) along trajectory 18 of Fig. 13b. Pressure scale in hPa. (a) 1000 UTC (point a). (b) 1100UTC (point b). (c) 1200 UTC (point c). (d) 1300 UTC (point d, end of trajectory).

over the Stansbury Mountains, Oquirrh Mountains, andthe gradually sloping southern Tooele Valley (Fig. 17d).The limited development of banded lake-effect precip-itation structures in NOSHFLX appeared to be due toweakened pressure troughing and decreased low-levelconvergence over the GSL (Fig. 18b). Although winddirections over the GSL were similar in the two sim-ulations, wind speeds in NOSHFLX were lower andoverlake convergence was weaker. The weak conver-gence that was evident in NOSHFLX arose from noc-turnal cooling over the surrounding landmass, whichwas included in the simulation to prevent inconsistencywith the boundary conditions provided by the 6-km res-olution mother domain. Thus, the removal of sensibleheat fluxes over the GSL decreased the intensity of thelake–land temperature difference, mesoscale pressuretroughing, thermally driven circulations, and overlakeconvergence. As a result, precipitation was produced

primarily by orographic processes. The weak overlakeconvergence and banded precipitation structures that ap-peared for short time periods in NOSHFLX may havebeen completely eliminated if nocturnal cooling of thesurrounding air mass could have been neglected in thesimulation.

c. Effects of latent heat release

Latent heat release due to condensation and fusion inthe lake-effect cloud and precipitation band may con-tribute to an intensification of convective circulations,overlake convergence, and ultimately, precipitation(Ballentine 1982; Hjelmfelt and Braham 1983; Hjelm-felt 1990). To examine the effects of latent heat release,a simulation was run in which latent heat release dueto condensation and fusion was neglected, but othermoisture and microphysical processes were retained


FIG. 15. Three-hour three-dimensional trajectories from the 2-kmdomain ending at 1300 UTC 7 Dec on the 0.855 sigma level (;760hPa). Trajectory width denotes pressure altitude according to scaleat upper right. Pressure at initial parcel location denoted on figure.Terrain heights above 1500 m shaded in dark gray. Thin solid linerepresents lake shoreline. Trajectory 1 parcel positions at 30-minintervals labeled. Snowband position at 1300 UTC 7 Dec markedwith heavy solid line.

FIG. 16. Pressure (hPa), potential temperature (K), mixing ratio (gkg21), and total cloud (water and ice) mixing ratio (g kg21) alongtrajectory 1 of Fig. 15.

(NOLHR). Compared to CTL, precipitation in NOLHRwas considerably lower near all three precipitation max-ima, with domain-averaged precipitation reduced by18% (Table 1; cf. Figs. 17a and 19). Although NOLHRand CTL featured similar 10-m wind and precipitationpatterns, lake-effect snowbands were weaker and thestrength of the low-level inflow and intensity of theconvective updrafts were reduced in the absence of la-tent heat release, as illustrated by the cross sections inFigs. 20a,b. As a result, the precipitation mixing ratioand updraft height were significantly reduced inNOLHR. Thus, latent heat release further intensifies thehorizontal and vertical circulations associated withlake-effect snowbands, resulting in additional precipi-tation.

d. Effects of topography

Many studies have noted the effects of orographicuplift on lake-effect snowfall in the Great Lakes region(e.g., Muller 1966; Hjelmfelt 1992; Niziol et al. 1995).As described by Steenburgh and Onton (2001), the GSLis surrounded by mountains several times higher thanthe terrain downstream of the Great Lakes. In order toexamine the effects of local topography on GSLE pre-cipitation processes, a sensitivity experiment (FLAT)

was run in which the surface elevation was set to theelevation of the GSL (1279.5 m), except near the lateralboundaries where the terrain was constrained to matchthat of the mother domain. To do this, five grid pointsnear the lateral boundaries of the 2-km domain wereleft unchanged, and the flat topography in the interiorwas blended with these grid points to avoid creatingsteep slopes.

The removal of topography had a significant impacton the distribution of precipitation (cf. Figs. 17a and21). Instead of a narrow band of precipitation with a19.3-mm maximum near Tooele, as was produced byCTL, FLAT produced a broader precipitation regionwith an 11.1-mm maximum (Table 1). In addition, theband of precipitation that was found downwind fromthe western shoreline of the GSL also featured a reducedprecipitation maximum in FLAT, although its distribu-tion was comparable to CTL. Thus, one effect of thedownstream orography, including the Stansbury Moun-tains, Oquirrh Mountains, and sloping topography with-in the Tooele Valley, was to enhance precipitation rateswithin the snowbands, increasing the maximum storm-total precipitation.

Interestingly, 9 mm of precipitation fell near thenortheast portion of Stansbury Island in FLAT, whileonly 2–4 mm fell in CTL. Closer inspection of the model


TABLE 1. Summary of sensitivity experiments. Precipitation (mm) from 0000 to 1500 UTC 7 Dec 1998 from the 2-km domain.Additional information on simulations available in text.

Experiment category Experiment DescriptionDomain max.precip. (mm)

Domain avg.precip. (mm)

Control run CTL Full physics simulation 19.3 0.59Surface fluxes and salinity FRESH Freshwater GSL 23.1 0.71Surface fluxes and salinity NOLHFLX No surface latent heat flux 9.0 0.30Surface fluxes and salinity NOSHFLX No surface sensible heat flux 7.2 0.42Latent heat release NOLHR No latent heat release in condensation 10.7 0.49Topography FLAT Flat lake-level topography 11.1 0.55Shoreline roughness contrasts ROUGH Roughness length 10 cm over entire domain 10.1 0.53Lake temperature T12 Lake temperature 280 K 34.7 0.78Lake temperature T22 Lake temperature 276 K 12.0 0.45Upstream moisture RH110 Relative humidity increased by 10% in initial

and boundary conditions27.0 1.13

Upstream moisture RH210 Relative humidity decreased by 10% in initialand boundary conditions

13.8 0.27

simulations showed that the contrast in precipitationnear Stansbury Island was due to differences in the evo-lution of the precipitation band in the two simulationsrather than orographic uplift. Most of the precipitationin FLAT fell between 1200 and 1500 UTC when theprecipitation band pivoted near Stansbury Island, ratherthan moving eastward as occurred in CTL (cf. Figs.22a,b). Thus, although the total precipitation during thistime period was similar between the two simulations,the heaviest precipitation in CTL fell near the south GSLshoreline, whereas it was located near Stansbury Islandin FLAT. This result illustrates both the dramatic localvariations in snowfall that can arise from subtle differ-ences in snowband placement, and the difficulty in pre-cise forecasting of precipitation occurring on such smallscales.

Comparison of CTL and FLAT shows that in thisevent, orographic uplift and channeling were not di-rectly responsible for snowband generation, althoughorographic enhancement of precipitation totals wasfound in some regions, such as near the city of Tooele.The lake-induced circulations caused by localized sur-face fluxes were sufficient for snowband formation. Inthe FLAT experiment, precipitation extended fartherdownstream and covered a broader area. Precipitationmaxima were also greatly reduced. Similarly, Hjelmfelt(1992) found that overall precipitation rates did not in-crease significantly when topography was included insimulations of lake-effect snow from Lake Michigan,although local precipitation rates were significantly in-creased in regions of strong orographic ascent.

e. Effects of shoreline roughness contrasts

Several studies have found that shoreline roughnesscontrasts may enhance lake-effect snowfall. Nicosia etal. (1999) documented a lake-enhanced rainband nearthe Lake Erie shoreline that may have been enhancedby lake–land frictional contrasts. They showed theoret-ically that frictional convergence may occur near ashoreline due to horizontal speed shear in the low-level

wind flow and horizontal directional shear resultingfrom a smaller angle between isobars and wind directionover water than over land. Lavoie (1972) found, usinga numerical model, that the effect of shoreline frictionalcontrasts alone produced elevated inversion heights andenhanced upward velocities over the lee shore of LakeErie.

In CTL, the planetary boundary layer was representedby a parameterization that determines surface fluxes ofheat, moisture, and momentum based on similarity the-ory (Blackadar 1976, 1979; Zhang and Anthes 1982).Over land, a roughness length (z0) of 10 cm was used,while over water, z0 was calculated as a function offriction velocity using the Charnock relation (Charnock1955; Delsol et al. 1971; Powers and Stoelinga 2000).For the wind speeds observed in this case, this yieldeda z0 of ;0.5 cm. To examine the influence of conver-gence associated with lake–land frictional contrasts onGSLE snowbands, a sensitivity experiment was con-ducted in which surface momentum fluxes were cal-culated based on a 10 cm z0 over both land and water(ROUGH). ROUGH featured reduced low-level windspeeds over the GSL compared to CTL due to greatersurface friction, but only minor changes in wind direc-tion were evident and the structure of the snowbandalong the western shoreline from 0200 to 0700 UTCwas changed little (cf. Figs. 17a and 23). Thus, ther-mally driven circulations appear to be the primary causefor the development of the low-level convergence zonein this region rather than frictionally induced conver-gence. Greater contrasts in precipitation were, however,observed as the band moved eastward, with the precip-itation maximum in ROUGH (10.1 mm) almost half thatof CTL (19.3 mm) (Table 1). This difference was dueprimarily to the reduction in wind speed and associateddecrease in surface heat and moisture fluxes due to thegreater roughness length over the GSL.

f. Sensitivity to lake-surface temperature

Although lake-surface temperature is an importantvariable in the development of lake-effect snowstorms


FIG. 17. Simulated accumulated precipitation (every 2 mm) and vector-averaged 10-m wind (full and half barbsdenote 5 and 2.5 m s21, respectively) from the 2-km domain from 0000 to 1500 UTC 7 Dec 1998. Model terrain (m)shaded according to scale at upper right. Dashed line denotes lake shoreline. (a) CTL. (b) FRESH. (c) NOLHFLX. (d)NOSHFLX.

(Lavoie 1972; Hjelmfelt 1990; Kristovich and Laird1998), there are significant uncertainties in the speci-fication of the GSL surface temperature that could limitthe skill of real-time predictions of lake-effect snowfall.Presently, lake temperature data are collected in realtime at only two sites [HAT and GNI; see Fig. 1 ofSteenburgh and Onton (2001) for locations] and datafrom the National Aeronautics and Space Administra-tion’s Advanced Very High Resolution Radiometer(AVHRR) on National Oceanic and Atmospheric Ad-ministration polar-orbiting satellites has shown that thelake-surface temperature can vary spatially by as muchas 58C. Regular retrieval of lake-surface temperature

using this instrument is not yet possible and is some-times limited by periods of cloud cover. Lake-surfacetemperature also varies by as much as 28C diurnally andcan change rapidly following cold-air intrusions, as iscommon prior to lake-effect events (Steenburgh and On-ton 2001). Thus, a potential source of numerical forecasterror lies in the specification of lake-surface tempera-ture, which can be particularly problematic if the lake-surface temperature changes during the forecast period.

To examine the model sensitivity to the specificationof lake-surface temperature, simulations were run inwhich the GSL temperature was set 28C higher (280 K;T12) and 28C lower (276 K; T22) than in CTL. The


FIG. 18. Difference fields of vector-averaged 10-m wind (scale at lower left) and mean altimeter setting (every 0.1hPa, zero contour not shown) from 0000 to 1500 UTC 7 Dec 1998. Model terrain (m) shaded according to scale at upperright. Dashed line denotes lake shoreline. (a) FRESH minus CTL. (b) NOSHFLX minus CTL.

FIG. 19. Same as Fig. 17 except for NOLHR.

28C perturbation was chosen based on the typical un-certainty in lake-temperature specification the arisesfrom spatial variability, diurnal fluctuations, and large-scale airmass changes. The T12 (T22) experiment pro-duced a maximum of roughly 34.7 mm (12.0 mm) ofprecipitation near Tooele and 32% more (24% less) do-main-averaged precipitation compared to CTL (Table 1;cf. Figs. 17a and 24a,b), with no major changes in theprecipitation pattern and the general evolution of thesnowbands. This result suggests that quantitative pre-

cipitation forecasts of these events are sensitive to thespecification of lake temperature, with perturbations inthe specified lake temperature on the order of the knownobservation uncertainty producing significant changesin the total precipitation.

g. Sensitivity to upstream moisture

Given the small size of the GSL and the reduction insaturation vapor pressure due to salinity, it is possiblethat upstream moisture is an important variable in GSLEevents. Steenburgh et al. (2000) found that most GSLEevents were associated with high values of 700-hPa rel-ative humidity downstream of the GSL, although norepresentative soundings were available upstream of theGSL. Hjelmfelt (1990) found that vertical velocities,precipitation, and land-breeze strength increase withhigher upstream moisture in simulations of lake-effectsnowfall over Lake Michigan. This effect was more pro-nounced with a higher lake–land temperature differenceand lower stability.

To determine the sensitivity of the model forecast tovariations in upstream moisture, sensitivity experimentswere run with varied relative humidity (RH). The firstincreased the relative humidity by 10% (up to a max-imum of 100%) in the initial and boundary conditionsfor the 2-km domain (RH110), while the second(RH210) decreased the relative humidity by 10% downto a minimum of 5% (RH210) Compared to CTL (Fig.17a), snowbands in RH110 initiated farther upstreamand precipitation was more widespread (Fig. 25a). TheRH110 experiment also produced a 27.0-mm precipi-tation maximum near Tooele, compared to 22.0 mm in


FIG. 20. Cross section CD from the 2-km domain at 1300 UTC 7 Dec 1998 (see Fig. 4 forcross section location). Virtual potential temperature (solid, every 28C), total cloud (water andice) mixing ratio (dashed, every 0.2 g kg21), total precipitation (snow and rain) mixing ratio (gkg21, shaded according to scale at upper left), and vectors of along-section wind and verticalvelocity (scale upper right). Lake shorelines denoted by thick solid lines. (a) CTL. (b) NOLHR.

CTL, and 91% more domain-averaged precipitation (Ta-ble 1).

In contrast, precipitation formed farther downstreamin RH210 (Fig. 25b) and at times snowbands were notpresent at all, despite the fact that a convergence zoneformed in the simulation. The slight reduction in avail-able upstream moisture was thus sufficient to greatlyreduce the intensity and duration of lake-induced snow-bands. As a result, RH210 produced only 13.8 mm ofprecipitation in the maximum near Tooele, and 54% lessdomain-averaged precipitation (Table 1).

These results are consistent with those of Hjelmfelt(1990), who found that lake-effect cases with higherupstream relative humidities require less boundary layer

moistening to initiate precipitation and generally pro-duce greater accumulations. In the present case, the up-stream relative humidity appears to influence howquickly convective clouds and precipitation form overthe GSL. Higher upstream relative humidity results incloud bands that form farther upstream since less bound-ary layer modification is required to support moist con-vection. Furthermore, increased upstream relative hu-midity results in greater storm-total accumulations.

The sensitivity studies also illustrate that errors in theupstream relative humidity forecast are likely to affectquantitative precipitation forecasts of GSLE storms,even at high resolution. In present-day numerical mod-els run at the National Centers for Environmental Pre-


FIG. 21. Same as Fig. 17 except for FLAT. FIG. 23. Same as Fig. 17 except for ROUGH.

FIG. 22. Snowband axis at 1-h intervals and accumulated precipitation (inset; every mm) from 1200 to 1500 UTC.Model terrain (m) shaded according to scale at upper right. Dashed line denotes lake shoreline. (a) CTL. (b) FLAT.

diction, the average root-mean-squared error of the 12-h(36 h) 700-hPa relative humidity forecast over the west-ern United States ranges from 18% to 25% (22% to28%) (White 1997; Cook 1998; White et al. 1999), sub-stantially larger than the perturbations inserted into thesensitivity studies. Furthermore, radiosonde observa-tions of relative humidity are known to have significanterrors and biases that vary among manufacturers (e.g.,Elliott and Gaffen 1991; Garand et al. 1992; Connelland Miller 1995). Thus, the observation and subsequentprediction of moisture in the upstream environment is

likely to be a significant source of error in high-reso-lution, quantitative precipitation forecasts of GSLEsnowstorms.

4. Summary and conclusions

In the companion paper, Steenburgh and Onton(2001) described the structure and evolution of the 7December 1998 GSLE snowstorm using observationaldata and a mesoscale simulation. In the present paper,we have extended our investigation through further di-


FIG. 24. Same as Fig. 17 except for (a) T12 and (b) T22.

FIG. 25. Same as Fig. 17 except for (a) RH110 and (b) RH210.

agnosis of the control run and a series of sensitivitystudies. Analysis of the formative stage of the event(0300 UTC 7 December 1998) showed that the simu-lated snowband developed along a land-breeze frontnear the west shoreline where offshore flow was con-vergent with northerly flow over the GSL. The devel-opment of the land-breeze front and convergence zonewas associated with localized heating over the lake sur-face, which produced pressure troughs near the westernand eastern shorelines. The complex dual trough struc-ture was produced by the geometry of the upwind(northern) shoreline, which determined the amount of

overwater fetch and boundary layer modification. Spe-cifically, isotherms paralleled the upwind shoreline, withlow-level warm anomalies and pressure troughs locateddownstream of the major bays and separated by an in-termediate tongue of colder air that was located to thelee of Promontory Point, a peninsula that extends south-ward into the GSL. Trajectories ending at 0300 UTCshowed that the highest low-level temperatures over theGSL were associated with the longest overwater fetches.A series of soundings along a selected trajectory showedthat localized heating by the lake warmed, deepened,moistened, and destabilized the boundary layer. Along


the trajectory, the near-surface temperature and meanmixing ratio below 775 hPa increased 2.18C and 0.16g kg21, respectively, with the latter representing an 8%increase compared to the upstream value. Althoughboundary layer moisture increased over the GSL, sim-ulated latent heat fluxes over the lake’s northern armwere reduced considerably from those over the southernarm because of the former’s hypersaline composition.

At maturity (1300 UTC), the snowband was alignedalong the midlake axis, coincident with a low-level con-vergence zone and a narrow tongue of locally high near-surface temperatures. Nocturnal cooling over land dur-ing the previous 10 h intensified the lake–land temper-ature gradient, pressure gradient, and offshore flow. Be-cause of the stronger offshore flow, trajectories endingat 1300 UTC experienced a shorter overwater fetchcompared with those ending at 0300 UTC. Nevertheless,soundings along an overlake trajectory showed that sen-sible heating rapidly eroded away a shallow surface in-version and all stable layers up to 550 hPa, increasedthe near-surface temperature 5.18C, and increased thenear-surface dewpoint 2.08C, resulting in significantboundary layer destabilization.

The importance of surface heat and moisture fluxesfrom the GSL was examined by comparing the controlsimulation, CTL, to a series of simulations in whichsurface fluxes of heat and moisture were altered. Theseincluded FRESH, where the saturation vapor pressureover the GSL was set to that of freshwater; NOLHFLX,where latent heat fluxes over the lake were neglected;and NOSHFLX, where sensible heat fluxes were ne-glected. The control simulation produced 17% less do-main-averaged precipitation than FRESH, illustratingthat the reduction in saturation vapor pressure by thehypersaline lake composition reduced the magnitude ofthis event. In the absence of moisture fluxes from theGSL (NOLHFX), precipitation was greatly reduced,suggesting that although the flux of moisture from thelake surface was reduced compared to freshwater, it wasnecessary to fully develop the intense snowband in thesimulation. The removal of sensible heat fluxes over theGSL (NOSHFLX) greatly decreased the intensity of thelake–land temperature difference, mesoscale pressuretroughing, thermally driven circulations, and overlakeconvergence. As a result, precipitation was generatedprimarily by orographic processes rather than by lake-induced convergence. This illustrates the importance ofthermally driven circulations induced by the GSL inorganizing low-level convergence and initiating con-vection in the postfrontal environment. In the absenceof thermally driven overlake convergence, the primarymechanism for precipitation generation is the orograph-ic uplift of air that has been moistened by the GSL.

The effects of latent heat release due to condensationwere examined using a simulation in which the latentheat release was ignored (NOLHR). This simulation il-lustrated that latent heat release further intensifies thehorizontal and vertical circulations associated with the

lake-effect snowband, resulting in additional precipi-tation.

A sensitivity study was also conducted to examinethe influence of the topography of northern Utah on thisevent. A simulation with no topography (FLAT) showedthat lake-effect snowbands occurred in the absence oforographic uplift and channeling, but the precipitationarea was broader with weaker maxima. Comparison ofFLAT and CTL showed that the precipitation rates inthe snowbands were enhanced in regions of orographicuplift, including the broadly sloped Tooele Valley. Thus,boundary layer destabilization and lake-induced low-level convergence were sufficient for snowband gen-eration, but significant precipitation enhancement oc-curred in orographically favored areas. Another sensi-tivity study (ROUGH) was used to show that, althoughsome studies have shown that surface roughness con-trasts may generate low-level convergence that initiatesor enhances lake-effect precipitation (e.g., Nicosia et al.1999), such effects were not significant during the 7December 1998 event.

The results described above show that many aspectsof the structure of the primary snowband in this eventwere similar to midlake and shoreline-parallel snow-bands found on the Great Lakes (e.g., Peace and Sykes1966; Passarelli and Braham 1981; Hjelmfelt and Bra-ham 1983; Hjelmfelt 1990; Niziol et al. 1995), but theevent was also influenced by the unique characteristicsof the GSL and topography of northern Utah. Localizedheating over the relatively warm GSL induced boundarylayer destabilization, mesoscale pressure troughing,land-breeze circulations, and low-level convergence, re-sulting in the development of convective updrafts anda wind-parallel band of clouds and precipitation. Thehypersaline content of the GSL was found to reducemoisture fluxes over the lake, and sensitivity studiesillustrated a significant reduction in storm-total precip-itation due to this effect. Circulations induced by thelocal topography were not found to be essential for pro-ducing the snowfall, but they did enhance (reduce) pre-cipitation in regions of orographic ascent (descent).

Although the simulation presented in this paper cap-tured many aspects of the observed event, several fac-tors would limit the skill of the modeling system whenapplied in an operational environment. The presentstudy used observed analyses for boundary conditionsand employed data assimilation on the coarse-resolutiongrids to limit large-scale error growth. Real-time pre-diction would not have such advantages and errors inthe large-scale forecast would likely limit model skill.For example, sensitivity studies showed that an increase(decrease) of 10% in upstream relative humidity resultedin a doubling (halving) of domain-averaged precipita-tion. Such relative humidity changes are considerablysmaller than the errors observed in present-day numer-ical models, which over the western United States fea-ture 12-h relative humidity root-mean-squared errorsover Salt Lake City of 18%–25% (White et al. 1999).


Another concern for operational prediction is the un-certainty in lake temperature that arises from the limitednumber of observations and rapid changes in lake tem-perature that accompany major airmass changes. To il-lustrate the sensitivity of mesoscale model forecasts toerrors in lake temperature, simulations were performedwith the lake temperature increased and decreased by28C (T12 and T22, respectively). Simulation T12(T22) produced 32% more (24% less) domain-averagedprecipitation, but otherwise the distribution of precipi-tation in the simulations were similar to CTL. This resultindicates that inaccurate lake temperatures can producesignificant differences in the quantitative precipitationforecast amount, but not necessarily in the general char-acter and structure of the precipitation patterns, timing,and location.

The upstream moisture and lake-temperature sensi-tivity studies also illustrate some of the predictabilityissues that arise when applying high-resolution meso-scale simulations in an operational environment. Al-though it is often presumed that the forcing of the GSLand surrounding topography is ‘‘fixed,’’ resulting in in-creased predictability for these events, in reality thereare many uncertainties in the specification or predictionof lake temperature and roughness. The influence of therapidly evolving salt slurry on boundary layer temper-atures over the Great Salt Lake Desert, which may ul-timately effect the strength of the land-breeze circula-tions, is also difficult to determine in real time (Steen-burgh and Onton 2001). It is also possible that the in-tense local forcing of the GSL and surroundingtopography may simply exacerbate large-scale modelerrors, greatly reducing forecast utility. For example, amodel forecast producing an upstream relative humiditythat is 20% higher than observed might produce a majorlake-effect snowstorm when, in reality, conditions weretoo dry for snowband development. Future work shouldexamine such predictability issues and determine if theprobabilistic guidance from mesoscale ensembles, runat a sufficient resolution to resolve lake-effect precipi-tation processes, may be required before significantgains in forecast skill are achieved.

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. Use of the MM5 was madepossible by the Mesoscale and Microscale MeteorologyDivision of NCAR, which is supported by the NationalScience Foundation. Computer time for the model sim-ulations was provided by the University of Utah Centerfor High Performance Computing. Surface observationswere provided by MesoWest, a collection of cooperatingmesonets in the western United States. MesoWest datawere collected and processed by John Horel, MikeSplitt, and Bryan White of the University of Utah, andLarry Dunn and David Zaff of the National Weather

Service. Special thanks to Larry Dunn, John Horel,Steve Krueger, Jan Paegle, Tom Potter, David Schultz,and Mike Splitt for their contributions, advice, and sci-entific support. We gratefully acknowledge the effortsof two anonymous reviewers, whose constructive eval-uations greatly improved the manuscript.

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