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

Transcription

1 1318 MONTHLY WEATHER REVIEW Diagnostic and Sensitivity Studies of the 7 December 1998 Great 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 using a series of mesoscale model simulations. Localized surface sensible and latent heating are shown to destabilize the boundary layer over the Great Salt Lake (GSL) and to produce mesoscale pressure troughing, land-breeze circulations, and low-level convergence that lead to the development of the primary band of convective clouds and 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, convergence, 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; and R 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 how small errors in the specification of these quantities can impact quantitative precipitation forecasts, potentially limiting the utility of high-resolution mesoscale model guidance. Results are compared to those from studies of lake-effect precipitation over the Great Lakes, and the implications for operational forecasting and numerical weather prediction are discussed. 1. Introduction In the first paper of this series, Steenburgh and Onton (2001) described the structure and evolution of a Great Salt Lake effect (GSLE) snowstorm using observational data and a mesoscale simulation by the nonhydrostatic Pennsylvania State University National Center for Atmospheric Research fifth generation Mesoscale Model (MM5). The event, which occurred on 7 December 1998, featured a wind-parallel snowband that developed along the west shoreline of the Great Salt Lake (GSL), migrated eastward, and eventually merged with a weaker snowband as it became aligned along the midlake axis. Snowfall accumulations reached 36 cm and were heaviest in a narrow, km-wide band extending downstream from the GSL. It was shown that the snowband along the western shoreline formed over a lowlevel convergence zone associated with a land-breeze Corresponding author address: Dr. Daryl J. Onton, Department of Meteorology, University of Utah, 135 South 1460 East Room 819, Salt Lake City, UT djonton@met.utah.edu front. The snowband aligned along the midlake axis as the land-breeze front moved eastward and offshore flow from the east shoreline intensified. The kinematic structure of the event thus appeared to be analogous to events associated with thermally driven land-breeze convergence over the Great Lakes of the eastern United States, including midlake bands produced over Lakes Michigan and Ontario (e.g., Peace and Sykes 1966; Passarelli and Braham 1981; Braham 1983; Hjelmfelt 1990; Niziol et al. 1995). Although the kinematic structure of the 7 December 1998 event was similar to that of midlake bands over the Great Lakes, the event also may have been influenced by unique aspects of the geography of northern Utah, including the presence of intense vertical relief and the hypersaline content of the GSL. As a result, the purpose of the present paper is twofold: (i) to further investigate the physical processes responsible for the 7 December 1998 GSLE snowstorm, and (ii) to examine issues related to the predictability of these events by present-day numerical models. This is accomplished with a detailed analysis of output from the 2-km horizontal resolution domain of the simulation presented by 2001 American Meteorological Society

2 JUNE 2001 ONTON AND STEENBURGH 1319 Steenburgh and Onton (2001) and a series of model sensitivity studies designed to illustrate the relative importance of selected physical and dynamical processes such as surface sensible and latent heat flux contrasts, topographic blocking and channeling, and frictional convergence due to shoreline surface roughness contrasts. Simulations in which lake-surface temperature and upstream moisture are modified are used to illustrate how small errors in the specification of these quantities can impact quantitative precipitation forecasts and potentially limit the utility of high-resolution mesoscale model guidance. The model diagnostic analysis is described in the next section. Then, section 3 describes the sensitivity studies and is followed by discussion and conclusions in section 4. The reader is referred to Steenburgh and Onton (2001) for a description of the geography and topography of northern Utah, characteristics of the GSL, evolution of the 7 December 1998 event, and configuration of the mesoscale simulation. 2. Model diagnostic and trajectory analysis of the 7 December 1998 snowband a. Formative stage At 0300 UTC, the simulated snowband was located along the west shoreline of the GSL where there was low-level convergence between northerly flow over the GSL and northwesterly flow over the west shoreline (Fig. 1a). Detailed analysis of the model simulation is presented for this time because it best describes the structure of the snowband during its early stages of development, although with a timing error of about 2 h (cf. Figs. 1a,b). Model surface analyses showed a tongue of high potential temperature air and a pressure trough located just off of the western shoreline of the GSL (Figs. 2a,b). 1 Immediately west of these features, convergence was maximized along the strong gradients in potential temperature and pressure that were associated with a developing land-breeze front. A similar but weaker feature was located offshore of the eastern shoreline. Figure 3 shows a cross section of circulation vectors (consisting of the horizontal and vertical velocity components in the plane of the cross section), virtual potential temperature, and cloud and precipitation mixing ratio taken perpendicular to the snowband near its farthest upwind extent (line AB in Fig. 4). Below 775 hpa the virtual potential temperature was nearly constant with height over the GSL, indicating near-neutral static stability. Lower virtual potential temperature air was found over land to the west and east of the GSL. Approximately 6 km from the western shoreline, low-level winds converged beneath a 5 Pa s 1 (50 cm s 1 ) updraft that fed a shallow cloud band that produced only neg- 1 Altimeter setting was used to reduce pressure to sea level to avoid aliasing surface temperature gradients into pressure gradients. FIG. 1. (a) Lowest half-sigma-level ( 40 m AGL) temperature (every 2 C), vertically integrated precipitation (kg m 2, shaded according to scale at upper right), and 10-m wind (full and half barbs denote 5 and 2.5 m s 1, respectively) from the 2-km domain at 0300 UTC 7 Dec Thick line denotes lake shoreline. (b) Lowestelevation angle (0.5 ) base-reflectivity analysis from the Salt Lake City WSR-88D (KMTX) (shaded according to scale at upper right) and surface observations at 0515 UTC 7 Dec Station plots denote wind (full and half barbs denote 5 and 2.5 m s 1, respectively) and temperature ( C, upper left). Dashed line denotes lake outline. Topographic contours (solid) every 500 m. ligible amounts of precipitation (precipitation mixing ratios 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 the cloud band was deeper (Fig. 5). Precipitation fell in a narrow shaft that was approximately 10 km in width. The surface pressure pattern evident in Fig. 2b, which featured a midlake ridge separated by two near-shoreline troughs, appeared to be caused by the shoreline geometry. With northerly flow over most of the lake at this time (see Fig. 1a), high virtual potential temperature air and a pressure trough were located over the northern bays of the GSL [Gunnison and Bear River; see Fig. 1 of Steenburgh and Onton (2001) for location]. To the lee of Promontory Point, a peninsula that extends southward into the GSL, lower virtual potential temperature air and higher pressure were found since flow downstream of this feature experienced a shorter overwater fetch and less heating and moistening (Fig. 2). Similarly,

3 1320 MONTHLY WEATHER REVIEW FIG. 2. Simulation from the 2-km domain at 0300 UTC 7 Dec (a) Lowest half-sigma-level ( 40 m AGL) potential temperature (every 0.2 C). (b) Altimeter setting (every hpa) and lowest halfsigma-level convergence ( 10 3 s 1, shaded according to scale at upper right). Dashed line denotes lake shoreline. Phillips (1972) found that overlake isotherms tended to parallel the upwind coastline on Lake Ontario. The strongest pressure gradient was located along the western shoreline and resulted in a substantial wind shift due to the offshore acceleration of the low-level flow. The resulting convergence provided the necessary mechanism to lift low-level air and produce convection along the land-breeze front. Interestingly, the convergence zone and precipitation band were coincident with the pressure gradient, rather than the pressure trough. This might be expected, however, since the airflow over the lake would tend to keep the wind speed and direction nearly constant until significant forcing such as the pressure gradient near the west shoreline could produce some deflection. In a simulation of a shoreline snowband over Lake Michigan, Hjelmfelt and Braham (1983) also obtained a wind field that had the maximum convergence between the pressure minimum and the shoreline. Three-dimensional trajectories beginning at 1800 UTC 6 December and terminating on the lowest halfsigma level [approximately 40 m above ground level (AGL)] at 0300 UTC 7 December further elucidate the processes associated with snowband development (Fig. 6). 2 Trajectories terminating in a line that was roughly perpendicular to the snowband demonstrate that the highest low-level temperatures were associated with trajectories that experienced the greatest overwater fetch (Fig. 6a). Low-level convergence into the snowband was evident with a strong deflection of trajectories ending immediately west of the snowband. This is more clearly illustrated by Fig. 6b, which shows two lines of trajectories terminating to the west and east of the snowband, respectively. Near the western shoreline of the GSL, trajectories were deflected eastward by the strong offshore pressure gradient acceleration. East of the snowband, the flow was primarily meridional. Boundary layer modification by the GSL is illustrated by a series of soundings taken along trajectory 9 of Fig. 6a (Fig. 7). At point A (0000 UTC), just upstream of the GSL, the sounding featured a shallow surface-based mixed layer with a near-surface temperature of 4.8 C (Fig. 7a). After air following trajectory 9 passed briefly over the GSL and temporarily again over land, the sounding at point B (0100 UTC) showed little change (Fig. 7b). Significant boundary layer modification occurred over the next hour as this air moved over the GSL, and the sounding at point C (0200 UTC) showed that during this period the boundary layer deepened 50 hpa while the near-surface temperature increased to 3.2 C (Fig. 7c). Meanwhile, the mean mixing ratio below 775 hpa, the approximate top of the boundary layer, increased from 2.05 to 2.19 g kg 1. As the air moved to point D, where it was beneath the snowband, the near-surface temperature and mean mixing ratio below 775 hpa increased to 2.7 C and 2.20 g kg 1, respectively, as the boundary layer increased slightly in depth (Fig. 7d). This series of soundings shows that sensible heating by the lake warmed, deepened, and destabilized the boundary layer, while latent heating in- 2 Because of the large amount of storage space required to store model output at high temporal resolution, trajectories were calculated using 30-min model output. This time difference was selected since a comparison of trajectories calculated from 5- and 30-min output showed no significant differences.

4 JUNE 2001 ONTON AND STEENBURGH 1321 FIG. 3. Cross section AB from the 2-km domain at 0300 UTC 7 Dec 1998 (see Fig. 4 for cross section location). Virtual potential temperature (solid, every 2 C), total cloud (water and ice) mixing ratio (dashed every 0.2 g kg 1 ), total precipitation (snow and rain) mixing ratio (g kg 1, 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. creased the mean boundary layer moisture. Collectively, the near-surface temperature and mean mixing ratio below 775 hpa increased 2.1 C and 0.16 g kg 1, respectively, with the latter representing an 8% increase compared to the upstream value. It is also interesting to examine the impact of lake salinity on simulated surface latent heat fluxes during this period. North of the railroad causeway (see Fig. 4 for location), where a 30% reduction in the freshwater saturation vapor pressure was specified in the simulation due to the observed 27% salinity (Steenburgh and Onton 2001), latent heat fluxes reached only 100 W m 2 and were significantly lower than over the lake s southern half, where the salinity was 9% and latent heat fluxes reached over 160 W m 2 (Fig. 8). This is in contrast to the typical distribution of surface latent heat fluxes over water bodies of uniform salinity, in which latent heat fluxes decrease downwind of the shoreline (Chang and Braham 1991). The impact of lake salinity on the intensity of this event will be discussed further in section 3. FIG. 4. Cross section locations (thick solid lines) and 2-km domain topography (m, shaded according to scale at bottom). Railroad causeway marked with thick dashed line. Lake shoreline denoted by solid line. b. Mature stage By 1300 UTC the convergence zone and snowband were located near the midlake axis and extended downstream into the Tooele Valley (Fig. 9). As noted by Steenburgh and Onton (2001), the simulated eastward movement of the snowband was delayed, so that the snowband structure at this time was similar to the ob-

5 1322 MONTHLY WEATHER REVIEW FIG. 5. Same as Fig. 3 except along line CD of Fig. 4. served structure at 0815 UTC (cf. Figs. 9a,b). The simulated low-level potential temperature pattern over the lake was dominated by a tongue of warm air that extended down the midlake axis and was flanked by intense potential temperature gradients near the eastern and western shorelines (Fig. 10a). Intensification of the lake land temperature gradient since 0300 UTC (cf. Figs. 2a and 10a) was due primarily to nocturnal cooling over the surrounding land mass, particularly over the Great Salt Lake Desert. A pressure trough was located along the warm tongue and strong pressure gradients had developed near the western and eastern shorelines (Fig. 10b). As a result, the western shore land breeze and the magnitude of the offshore flow near the eastern shoreline had intensified, resulting in the development of a well-defined convergence zone near the midlake axis (Figs. 9a and 10b). The mesoscale circulations responsible for the snowband are further illustrated by the cross section of virtual potential temperature, circulation vectors, and cloud and precipitation mixing ratio presented in Fig. 11 (line AB in Fig. 4). Compared to 10 h previously (Fig. 3), nearsurface virtual potential temperatures to the west (east) of the GSL decreased by 6 8 C (2 3 C), apparently due to nocturnal cooling, resulting in the development of a more dense and stable boundary layer surrounding the GSL. Meanwhile, virtual potential temperatures over the GSL remained relatively constant. The resulting contrast in boundary layer temperature and surface pressure intensified the offshore flow, particularly from the western shoreline. Where the two opposing flows met near the midlake axis, a narrow 6 Pa s 1 ( 60 cm s 1 ) updraft was found beneath the developing cloud and precipitation band. Interestingly, the convergence zone and updraft were again not located directly over the pressure trough, similar to 10 h previously. The overall three-dimensional wind and thermal structure closely resembled that associated with midlake bands over Lakes Michigan and Ontario (e.g., Peace and Sykes 1966; Passarelli and Braham 1981; Braham and Kelly 1982; Hjelmfelt 1990). A cross section along the center of the snowband at 1300 UTC is presented in Fig. 12 (line EF in Fig. 4). Low-level virtual potential temperature increased along the lake axis, indicating reduced stability over the southern half of the GSL, where updrafts formed and produced progressively larger cloud and precipitation mixing ratios as the flow neared the downwind shoreline. Vertical velocities were largest near the downwind shoreline, where updrafts extended to 650 hpa, rather than over the GSL where surface heating was strongest. This was likely due to the downstream advection of convective updrafts in a manner similar to that described by Lin and Smith (1986). At this time, vertical motion in the immediate vicinity of the Oquirrh Mountains was weak compared to that within the snowband, although in section 3 orographically induced ascent will be shown to enhance precipitation at other times. Three-dimensional trajectories, beginning at 0400 UTC 7 December and terminating on the lowest halfsigma level ( 40 m AGL) at 1300 UTC 7 December (Fig. 13) show stronger lake-induced circulations than at 0300 UTC (see trajectories 7 10 in Fig. 6). Although the snowband was aligned along the major lake axis,

6 JUNE 2001 ONTON AND STEENBURGH 1323 FIG. 6. Nine-hour three-dimensional trajectories from the 2-km domain ending at 0300 UTC 7 Dec on the lowest half-sigma level ( 40 m AGL). Trajectory width denotes pressure altitude according to scale at upper right. Lowest half-sigma-level potential temperature contours every 0.2 C shown only near lake level. Terrain heights above 1500 m shaded in gray. Heavy solid line denotes lake shoreline. (a) Trajctories ending in a line perpendicular to the snowband. Parcel locations at 30-min intervals along trajectory 9 denoted with large dots. Locations at 0000, 0100, 0200, and 0300 UTC denoted by letters a, 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 those terminating around the snowband 10 h earlier (Fig. 6). Boundary layer modification along a trajectory that had a relatively long overwater fetch and terminated near the warm tongue (trajectory 18, Fig. 13b) is illustrated by the soundings in Fig. 14. At 1000 UTC, the air following trajectory 18 was approaching the northern shoreline of the Bear River Bay (point a, Fig. 13b). At this location, a shallow surface inversion, presumably produced by nocturnal cooling, was located near the surface and another shallow stable layer was located near 800 hpa (Fig. 14a). The lowest half-sigma level temperature and dewpoint were 8.9 and 10.4 C, respectively. One hour later (1100 UTC), this air was located over the Bear River Bay and had been over water for 30 min (point b, Fig. 13b). At this point, sensible heating over the lake surface had raised the near-surface temperature to 6.3 C and a shallow surface-based mixed layer had developed (Fig. 14b). Near-surface moisture increased slightly with the dewpoint reaching 9.8 C. The stable layer near 800 hpa was still present and had lowered slightly. By 1200 UTC, the air had reached point c, which was located over the less saline southern region of the GSL (Fig. 13a). The near-surface temperature and dewpoint had risen to 4.7 and 9.5 C, respectively, and the stable layer near 875 hpa had weakened (Fig. 14c). During the next hour, the nearsurface temperature and dewpoint increased to 3.8 and 8.4 C, respectively, all stable layers and inversions below 550 hpa eroded away, and the sounding became conditionally unstable up to 675 hpa (Fig. 14d). As an example of the vertical circulations associated with the snowband at this time, three-dimensional trajectories following air that was ingested into the snowband are presented in Fig. 15. These trajectories began at 1000 UTC 7 December and terminated at 1300 UTC 7 December on the sigma level (761 hpa), which was near the top of the snowband. Air following trajectories 1 3 originated at low levels west of the GSL, converged toward the snowband axis, rose rapidly, and moved downstream. Trajectory 1, which was the outermost trajectory west of the snowband at 1300 UTC, was the innermost relative to the midlake axis at the beginning of the trajectory. As air following trajectory 1 approached the low-level convergence zone, it ascended through the snowband. Trajectories 2 and 3 followed air that originated farther to the south and remained near the surface until it rose rapidly when it was ingested in the narrow updraft that supported the snowband (e.g., Fig. 11). Air following trajectories ending in the eastern portion of the snowband followed a slightly different evolution. Trajectories 4 and 5 originated north of the GSL, 80 hpa above lake level, eventually descended to low levels as they converged toward the snowband, then rose rapidly as they were ingested into the snowband. The air following trajectory 6 was ingested into the snowband near its farthest upwind extent, rose to 757 hpa, remained aloft and moved slightly away from the snowband, descended slightly to 803 hpa, converged again toward the snowband, and rose to 761 hpa near the snowband. The evolution of pressure, temperature, water vapor mixing ratio, and cloud water and ice mixing ratio along trajectory 1 are displayed in Fig. 16. At 1000 UTC, the air following trajectory 1 was located west of the shoreline of the GSL at 849 hpa. It descended to 875 hpa ( 5 hpa above lake level) as it followed downward

7 1324 MONTHLY WEATHER REVIEW FIG. 7. Skew T logp diagrams of simulated temperature ( C), dewpoint ( C), and wind (full and half barbs denote 5 and 2.5 m s 1, respectively) along trajectory 9 of Fig. 6a. Pressure scale in hpa. (a) 0000 UTC (point a). (b) 0100 UTC (point b). (c) 0200 UTC (point c). (d) 0300 UTC (point d, end of trajectory). sloping terrain toward the lake shoreline after 1030 UTC. Potential temperature and water vapor mixing ratio increased 3.4 K and 0.4 g kg 1 from 1100 to 1130 UTC as this air moved over the GSL. As it approached the low-level convergence zone and associated updraft prior to 1230 UTC, it began to rise upward with saturation and cloud water development beginning at around 1220 UTC. At approximately 1240 UTC, the air reached its highest elevation and the cloud water mixing ratio was near its peak value. The air then subsided gradually as it began to exit the snowband, while the remaining cloud condensate dissipated and potential temperature and mixing ratio remained nearly constant. The physical picture obtained from the analysis above is consistent with the findings of studies over the Great Lakes that illustrate the role of thermally driven landbreeze circulations in generating solitary wind-parallel snowbands (e.g., Passarelli and Braham 1981; Hjelmfelt and Braham 1983; Hjelmfelt 1990). In the present case, a pressure trough and low-level convergence were induced by upward surface sensible and latent heat fluxes over the GSL. Initially, the convergence zone developed along the boundary between a developing land breeze from the western shoreline and synoptic-scale northerly flow over the GSL. This provided a mechanism to lift low-level air so that, aided by boundary layer heating, moistening, and destabilization over the lake, a band of clouds and precipitation formed roughly parallel to the northwesterly steering-layer flow. As nocturnal cooling increased the lake land temperature contrast, land breezes from the western and eastern shorelines intensified, increasing the strength of the overlake convergence and producing the most organized, mature stage of snowband evolution. The shoreline geometry was also found to influence the pattern of temperature and pressure over the lake, similar to observations over Lake

8 JUNE 2001 ONTON AND STEENBURGH 1325 FIG. 8. Surface latent heat flux (every 10 W m 2 ) from the 2-km domain at 0300 UTC 7 Dec Dashed line denotes lake shoreline. FIG. 10. Same as Fig. 2 except at 1300 UTC 7 Dec Ontario (Phillips 1972), with the highest near-surface temperatures and lowest pressures associated with the longest overwater trajectories. FIG. 9. (a) Same as Fig. 1a except for 1300 UTC. (b) Same as Fig. 1b except for 0815 UTC. 3. Sensitivity experiments a. Experimental design To examine the relative importance of different physical processes on the evolution of the 7 December 1998 snowband, a series of sensitivity experiments was conducted in which either a particular parameter or process was modified or withheld from the 2-km domain (Table

9 1326 MONTHLY WEATHER REVIEW FIG. 11. Same as Fig. 3 except at 1300 UTC 7 Dec ). Unless otherwise noted, all other characteristics of these simulations, including the lateral boundary conditions provided to the 2-km domain, were the same as those used for the control simulation (CTL; described in Steenburgh and Onton 2001). Although the sensitivity experiments presented below provide insight into the relative importance of the physical processes influencing the development, evolution, and predictability of this event, they do not fully quantify the contribution of a particular process to the storm-total snowfall since ultimately it is the nonlinear, synergistic interaction between processes (e.g., Uccellini 1990; Stein and Alpert 1993; Alpert et al. 1995) that yields a snowband of the observed intensity. FIG. 12. Same as Fig. 3 except at 1300 UTC 7 Dec 1998 and along line EF of Fig. 4.

10 JUNE 2001 ONTON AND STEENBURGH 1327 FIG. 13. Nine-hour three-dimensional trajectories from the 2-km domain ending at 1300 UTC 7 Dec on the lowest half-sigma level ( 40 m AGL). Trajectory width denotes pressure altitude according to scale at upper right. Lowest half-sigma-level potential temperature contours every 0.2 C shown only near lake level. Terrain heights above 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 intervals along trajectory 18 denoted with large dots. Locations at 1000, 1100, 1200, and 1300 UTC denoted by letters a, b, c, and d, respectively. b. Effects of surface fluxes and lake salinity Surface fluxes of heat and moisture have been shown to be important physical mechanisms leading to lakeeffect snow in the Great Lakes region (e.g., Lavoie 1972; Passarelli and Braham 1981; Hjelmfelt 1990) and, as described above, appear critical to the development of the 7 December GSLE snowband. As discussed by Steenburgh et al. (2000), by Steenburgh and Onton (2001), and in section 2, the GSL is a hypersaline body of water that has a lower saturation vapor pressure than freshwater. As a result, moisture fluxes are reduced from what they would be for freshwater. In this section, a series of simulations is described that examine the impact of lake salinity, latent heat flux, and sensible heat flux on the evolution of the 7 December 1998 snowband. Knowledge of the sensitivity of mesoscale simulations to the specification of lake salinity is useful given the historical variability of salinity in the GSL (Steenburgh et al. 2000). Figure 17 presents the total precipitation and mean 10-m winds for UTC 7 December from CTL, a simulation with freshwater latent heat fluxes (FRESH), a simulation in which latent heat fluxes over the GSL were ignored (NOLHFLX), and a simulation in which sensible heat fluxes over the GSL were ignored (NOSHFLX). CTL produced 17% less domain-averaged precipitation than FRESH (Table 1; cf. Figs. 17a,b). This difference was due to enhanced precipitation in FRESH during the stage of the simulation where the snowband was resident near the western shoreline of the GSL, as well as a tendency for the snowband to extend farther upstream during most of the simulation. In addition, the precipitation maximum in the eastern Tooele Valley increased from 19.3 to 23.1 mm (Table 1). Thus, by reducing surface moisture fluxes, lake salinity had a significant impact on the amount of snowfall produced during this event. Figure 18a shows the difference in altimeter setting and 10-m wind between the two simulations (i.e., FRESH CTL) and illustrates that, although the FRESH mean wind field is similar to that of CTL, slightly enhanced convergence is evident near the western shoreline. Although not evident in this figure, marginally stronger pressure troughing was evident in FRESH. It appears that the surface moisture fluxes in FRESH led to additional latent heating by condensation and fusion within the cloud band, promoting more vigorous circulations within the snowband. Similar complementary processes in lake-effect storms have been noted by Lavoie (1972), Hjelmfelt and Braham (1983), and Hjelmfelt (1990). NOLHFLX produced 49% less domain-averaged precipitation than the control run (Table 1; cf. Figs. 17a,c). Precipitation was substantially lower than that produced by CTL, did not extend as far upstream over the GSL, and was confined primarily to areas of orographic ascent. Thus, comparison of CTL and NOLHFLX suggests that, although the GSL is hypersaline and small in size, the limited flux of moisture from its surface and resulting increase in low-level mixing ratio was necessary to fully develop the intense lake-induced snowband in this event. This result also illustrates the highly sensitive nature of snowband development since the relatively small increase in low-level moisture over the GSL resulted in a dramatic increase in storm-total precipitation. NOSHFLX produced 28% less domain-averaged precipitation than CTL (Table 1), and, as observed in NOLHFLX, precipitation in the former fell primarily

11 1328 MONTHLY WEATHER REVIEW FIG. 14. Skew T logp diagrams of simulated temperature ( C), dewpoint ( C), and wind (full and half barbs denote 5 and 2.5 m s 1, respectively) along trajectory 18 of Fig. 13b. Pressure scale in hpa. (a) 1000 UTC (point a). (b) 1100 UTC (point b). (c) 1200 UTC (point c). (d) 1300 UTC (point d, end of trajectory). over the Stansbury Mountains, Oquirrh Mountains, and the gradually sloping southern Tooele Valley (Fig. 17d). The limited development of banded lake-effect precipitation structures in NOSHFLX appeared to be due to weakened pressure troughing and decreased low-level convergence over the GSL (Fig. 18b). Although wind directions over the GSL were similar in the two simulations, wind speeds in NOSHFLX were lower and overlake convergence was weaker. The weak convergence that was evident in NOSHFLX arose from nocturnal cooling over the surrounding landmass, which was included in the simulation to prevent inconsistency with the boundary conditions provided by the 6-km resolution mother domain. Thus, the removal of sensible heat fluxes over the GSL decreased the intensity of the lake land temperature difference, mesoscale pressure troughing, thermally driven circulations, and overlake convergence. As a result, precipitation was produced primarily by orographic processes. The weak overlake convergence and banded precipitation structures that appeared for short time periods in NOSHFLX may have been completely eliminated if nocturnal cooling of the surrounding air mass could have been neglected in the simulation. c. Effects of latent heat release Latent heat release due to condensation and fusion in the lake-effect cloud and precipitation band may contribute to an intensification of convective circulations, overlake convergence, and ultimately, precipitation (Ballentine 1982; Hjelmfelt and Braham 1983; Hjelmfelt 1990). To examine the effects of latent heat release, a simulation was run in which latent heat release due to condensation and fusion was neglected, but other moisture and microphysical processes were retained

12 JUNE 2001 ONTON AND STEENBURGH 1329 FIG. 15. Three-hour three-dimensional trajectories from the 2-km domain ending at 1300 UTC 7 Dec on the sigma level ( 760 hpa). Trajectory width denotes pressure altitude according to scale at upper right. Pressure at initial parcel location denoted on figure. Terrain heights above 1500 m shaded in dark gray. Thin solid line represents lake shoreline. Trajectory 1 parcel positions at 30-min intervals labeled. Snowband position at 1300 UTC 7 Dec marked with heavy solid line. FIG. 16. Pressure (hpa), potential temperature (K), mixing ratio (g kg 1 ), and total cloud (water and ice) mixing ratio (g kg 1 ) along trajectory 1 of Fig. 15. (NOLHR). Compared to CTL, precipitation in NOLHR was considerably lower near all three precipitation maxima, with domain-averaged precipitation reduced by 18% (Table 1; cf. Figs. 17a and 19). Although NOLHR and CTL featured similar 10-m wind and precipitation patterns, lake-effect snowbands were weaker and the strength of the low-level inflow and intensity of the convective updrafts were reduced in the absence of latent heat release, as illustrated by the cross sections in Figs. 20a,b. As a result, the precipitation mixing ratio and updraft height were significantly reduced in NOLHR. Thus, latent heat release further intensifies the horizontal and vertical circulations associated with lake-effect snowbands, resulting in additional precipitation. d. Effects of topography Many studies have noted the effects of orographic uplift 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 GSL is surrounded by mountains several times higher than the terrain downstream of the Great Lakes. In order to examine the effects of local topography on GSLE precipitation processes, a sensitivity experiment (FLAT) was run in which the surface elevation was set to the elevation of the GSL ( m), except near the lateral boundaries where the terrain was constrained to match that of the mother domain. To do this, five grid points near the lateral boundaries of the 2-km domain were left unchanged, and the flat topography in the interior was blended with these grid points to avoid creating steep slopes. The removal of topography had a significant impact on the distribution of precipitation (cf. Figs. 17a and 21). Instead of a narrow band of precipitation with a 19.3-mm maximum near Tooele, as was produced by CTL, FLAT produced a broader precipitation region with an 11.1-mm maximum (Table 1). In addition, the band of precipitation that was found downwind from the western shoreline of the GSL also featured a reduced precipitation maximum in FLAT, although its distribution was comparable to CTL. Thus, one effect of the downstream orography, including the Stansbury Mountains, Oquirrh Mountains, and sloping topography within the Tooele Valley, was to enhance precipitation rates within the snowbands, increasing the maximum stormtotal precipitation. Interestingly, 9 mm of precipitation fell near the northeast portion of Stansbury Island in FLAT, while only 2 4 mm fell in CTL. Closer inspection of the model

13 1330 MONTHLY WEATHER REVIEW 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 Description Domain max. precip. (mm) Domain avg. precip. (mm) Control run CTL Full physics simulation Surface fluxes and salinity FRESH Freshwater GSL Surface fluxes and salinity NOLHFLX No surface latent heat flux Surface fluxes and salinity NOSHFLX No surface sensible heat flux Latent heat release NOLHR No latent heat release in condensation Topography FLAT Flat lake-level topography Shoreline roughness contrasts ROUGH Roughness length 10 cm over entire domain Lake temperature T 2 Lake temperature 280 K Lake temperature T 2 Lake temperature 276 K Upstream moisture RH 10 Relative humidity increased by 10% in initial and boundary conditions Upstream moisture RH 10 Relative humidity decreased by 10% in initial and boundary conditions simulations showed that the contrast in precipitation near Stansbury Island was due to differences in the evolution of the precipitation band in the two simulations rather than orographic uplift. Most of the precipitation in FLAT fell between 1200 and 1500 UTC when the precipitation band pivoted near Stansbury Island, rather than moving eastward as occurred in CTL (cf. Figs. 22a,b). Thus, although the total precipitation during this time period was similar between the two simulations, the heaviest precipitation in CTL fell near the south GSL shoreline, whereas it was located near Stansbury Island in FLAT. This result illustrates both the dramatic local variations in snowfall that can arise from subtle differences in snowband placement, and the difficulty in precise forecasting of precipitation occurring on such small scales. Comparison of CTL and FLAT shows that in this event, orographic uplift and channeling were not directly responsible for snowband generation, although orographic enhancement of precipitation totals was found in some regions, such as near the city of Tooele. The lake-induced circulations caused by localized surface fluxes were sufficient for snowband formation. In the FLAT experiment, precipitation extended farther downstream and covered a broader area. Precipitation maxima were also greatly reduced. Similarly, Hjelmfelt (1992) found that overall precipitation rates did not increase significantly when topography was included in simulations of lake-effect snow from Lake Michigan, although local precipitation rates were significantly increased in regions of strong orographic ascent. e. Effects of shoreline roughness contrasts Several studies have found that shoreline roughness contrasts may enhance lake-effect snowfall. Nicosia et al. (1999) documented a lake-enhanced rainband near the Lake Erie shoreline that may have been enhanced by lake land frictional contrasts. They showed theoretically that frictional convergence may occur near a shoreline due to horizontal speed shear in the low-level wind flow and horizontal directional shear resulting from a smaller angle between isobars and wind direction over water than over land. Lavoie (1972) found, using a numerical model, that the effect of shoreline frictional contrasts alone produced elevated inversion heights and enhanced upward velocities over the lee shore of Lake Erie. In CTL, the planetary boundary layer was represented by a parameterization that determines surface fluxes of heat, moisture, and momentum based on similarity theory (Blackadar 1976, 1979; Zhang and Anthes 1982). Over land, a roughness length (z 0 ) of 10 cm was used, while over water, z 0 was calculated as a function of friction velocity using the Charnock relation (Charnock 1955; Delsol et al. 1971; Powers and Stoelinga 2000). For the wind speeds observed in this case, this yielded a z 0 of 0.5 cm. To examine the influence of convergence associated with lake land frictional contrasts on GSLE snowbands, a sensitivity experiment was conducted in which surface momentum fluxes were calculated based on a 10 cm z 0 over both land and water (ROUGH). ROUGH featured reduced low-level wind speeds over the GSL compared to CTL due to greater surface friction, but only minor changes in wind direction were evident and the structure of the snowband along the western shoreline from 0200 to 0700 UTC was changed little (cf. Figs. 17a and 23). Thus, thermally driven circulations appear to be the primary cause for the development of the low-level convergence zone in this region rather than frictionally induced convergence. Greater contrasts in precipitation were, however, observed as the band moved eastward, with the precipitation maximum in ROUGH (10.1 mm) almost half that of CTL (19.3 mm) (Table 1). This difference was due primarily to the reduction in wind speed and associated decrease in surface heat and moisture fluxes due to the greater roughness length over the GSL. f. Sensitivity to lake-surface temperature Although lake-surface temperature is an important variable in the development of lake-effect snowstorms

14 JUNE 2001 ONTON AND STEENBURGH 1331 FIG. 17. Simulated accumulated precipitation (every 2 mm) and vector-averaged 10-m wind (full and half barbs denote 5 and 2.5 m s 1, respectively) from the 2-km domain from 0000 to 1500 UTC 7 Dec 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 Laird 1998), there are significant uncertainties in the specification of the GSL surface temperature that could limit the skill of real-time predictions of lake-effect snowfall. Presently, lake temperature data are collected in real time at only two sites [HAT and GNI; see Fig. 1 of Steenburgh and Onton (2001) for locations] and data from the National Aeronautics and Space Administration s Advanced Very High Resolution Radiometer (AVHRR) on National Oceanic and Atmospheric Administration polar-orbiting satellites has shown that the lake-surface temperature can vary spatially by as much as 5 C. Regular retrieval of lake-surface temperature using this instrument is not yet possible and is sometimes limited by periods of cloud cover. Lake-surface temperature also varies by as much as 2 C diurnally and can change rapidly following cold-air intrusions, as is common prior to lake-effect events (Steenburgh and Onton 2001). Thus, a potential source of numerical forecast error lies in the specification of lake-surface temperature, which can be particularly problematic if the lakesurface temperature changes during the forecast period. To examine the model sensitivity to the specification of lake-surface temperature, simulations were run in which the GSL temperature was set 2 C higher (280 K; T 2) and 2 C lower (276 K; T 2) than in CTL. The

15 1332 MONTHLY WEATHER REVIEW FIG. 18. Difference fields of vector-averaged 10-m wind (scale at lower left) and mean altimeter setting (every 0.1 hpa, zero contour not shown) from 0000 to 1500 UTC 7 Dec Model terrain (m) shaded according to scale at upper right. Dashed line denotes lake shoreline. (a) FRESH minus CTL. (b) NOSHFLX minus CTL. FIG. 19. Same as Fig. 17 except for NOLHR. 2 C perturbation was chosen based on the typical uncertainty in lake-temperature specification the arises from spatial variability, diurnal fluctuations, and largescale airmass changes. The T 2 (T 2) experiment produced a maximum of roughly 34.7 mm (12.0 mm) of precipitation near Tooele and 32% more (24% less) domain-averaged precipitation compared to CTL (Table 1; cf. Figs. 17a and 24a,b), with no major changes in the precipitation pattern and the general evolution of the snowbands. This result suggests that quantitative precipitation forecasts of these events are sensitive to the specification of lake temperature, with perturbations in the specified lake temperature on the order of the known observation uncertainty producing significant changes in the total precipitation. g. Sensitivity to upstream moisture Given the small size of the GSL and the reduction in saturation vapor pressure due to salinity, it is possible that upstream moisture is an important variable in GSLE events. Steenburgh et al. (2000) found that most GSLE events were associated with high values of 700-hPa relative humidity downstream of the GSL, although no representative soundings were available upstream of the GSL. Hjelmfelt (1990) found that vertical velocities, precipitation, and land-breeze strength increase with higher upstream moisture in simulations of lake-effect snowfall over Lake Michigan. This effect was more pronounced with a higher lake land temperature difference and lower stability. To determine the sensitivity of the model forecast to variations in upstream moisture, sensitivity experiments were run with varied relative humidity (RH). The first increased the relative humidity by 10% (up to a maximum of 100%) in the initial and boundary conditions for the 2-km domain (RH 10), while the second (RH 10) decreased the relative humidity by 10% down to a minimum of 5% (RH 10) Compared to CTL (Fig. 17a), snowbands in RH 10 initiated farther upstream and precipitation was more widespread (Fig. 25a). The RH 10 experiment also produced a 27.0-mm precipitation maximum near Tooele, compared to 22.0 mm in

16 JUNE 2001 ONTON AND STEENBURGH 1333 FIG. 20. Cross section CD from the 2-km domain at 1300 UTC 7 Dec 1998 (see Fig. 4 for cross section location). Virtual potential temperature (solid, every 2 C), total cloud (water and ice) mixing ratio (dashed, every 0.2 g kg 1 ), total precipitation (snow and rain) mixing ratio (g kg 1, shaded according to scale at upper left), and vectors of along-section wind and vertical velocity (scale upper right). Lake shorelines denoted by thick solid lines. (a) CTL. (b) NOLHR. CTL, and 91% more domain-averaged precipitation (Table 1). In contrast, precipitation formed farther downstream in RH 10 (Fig. 25b) and at times snowbands were not present at all, despite the fact that a convergence zone formed in the simulation. The slight reduction in available upstream moisture was thus sufficient to greatly reduce the intensity and duration of lake-induced snowbands. As a result, RH 10 produced only 13.8 mm of precipitation in the maximum near Tooele, and 54% less domain-averaged precipitation (Table 1). These results are consistent with those of Hjelmfelt (1990), who found that lake-effect cases with higher upstream relative humidities require less boundary layer moistening to initiate precipitation and generally produce greater accumulations. In the present case, the upstream relative humidity appears to influence how quickly convective clouds and precipitation form over the GSL. Higher upstream relative humidity results in cloud bands that form farther upstream since less boundary layer modification is required to support moist convection. Furthermore, increased upstream relative humidity results in greater storm-total accumulations. The sensitivity studies also illustrate that errors in the upstream relative humidity forecast are likely to affect quantitative precipitation forecasts of GSLE storms, even at high resolution. In present-day numerical models run at the National Centers for Environmental Pre-

17 1334 MONTHLY WEATHER REVIEW FIG. 21. Same as Fig. 17 except for FLAT. FIG. 23. Same as Fig. 17 except for ROUGH. diction, the average root-mean-squared error of the 12-h (36 h) 700-hPa relative humidity forecast over the western United States ranges from 18% to 25% (22% to 28%) (White 1997; Cook 1998; White et al. 1999), substantially larger than the perturbations inserted into the sensitivity studies. Furthermore, radiosonde observations of relative humidity are known to have significant errors and biases that vary among manufacturers (e.g., Elliott and Gaffen 1991; Garand et al. 1992; Connell and Miller 1995). Thus, the observation and subsequent prediction of moisture in the upstream environment is likely to be a significant source of error in high-resolution, quantitative precipitation forecasts of GSLE snowstorms. 4. Summary and conclusions In the companion paper, Steenburgh and Onton (2001) described the structure and evolution of the 7 December 1998 GSLE snowstorm using observational data and a mesoscale simulation. In the present paper, we have extended our investigation through further di- 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.

18 JUNE 2001 ONTON AND STEENBURGH 1335 FIG. 24. Same as Fig. 17 except for (a) T 2 and (b) T 2. agnosis of the control run and a series of sensitivity studies. Analysis of the formative stage of the event (0300 UTC 7 December 1998) showed that the simulated snowband developed along a land-breeze front near the west shoreline where offshore flow was convergent with northerly flow over the GSL. The development of the land-breeze front and convergence zone was associated with localized heating over the lake surface, which produced pressure troughs near the western and eastern shorelines. The complex dual trough structure was produced by the geometry of the upwind (northern) shoreline, which determined the amount of overwater fetch and boundary layer modification. Specifically, isotherms paralleled the upwind shoreline, with low-level warm anomalies and pressure troughs located downstream of the major bays and separated by an intermediate tongue of colder air that was located to the lee of Promontory Point, a peninsula that extends southward into the GSL. Trajectories ending at 0300 UTC showed that the highest low-level temperatures over the GSL were associated with the longest overwater fetches. A series of soundings along a selected trajectory showed that localized heating by the lake warmed, deepened, moistened, and destabilized the boundary layer. Along FIG. 25. Same as Fig. 17 except for (a) RH 10 and (b) RH 10.