1. Introduction
The Tax Day Storm of 15 April 2002 produced the second lowest sea level pressure observed in Utah since observational records began in 1892 and the strongest cold frontal passage at the Salt Lake City International Airport (KSLC; locations of stations and geographical features are shown in Fig. 1) in 25 years (Shafer and Steenburgh 2008, hereafter SS08; based on 2-h temperature fall). Cyclones such as the Tax Day Storm traverse the Intermountain West several times per year (Petterssen 1956; Zishka and Smith 1980; Whittaker and Horn 1981; Lee 1995; Jeglum 2010) and can be accompanied by strong fronts, high winds (>30 m s−1), power outages, blowing dust, dramatic temperature falls, wild fire runs, and/or heavy snow (SS08). Unfortunately, our understanding of these events is limited (Hill 1993) and, in general, forecast skill is lower over the Intermountain West than other regions of the United States (Junker et al. 1989, 1992; McDonald 1998; Yuan et al. 2007).
Intermountain cyclogenesis is a likely consequence of flow interaction with the Sierra Nevada and Cascade Mountains, which form the western boundary of the region (Fig. 1a). In particular, the southern portion of the Sierra Nevada, known as the high Sierra, forms a formidable quasi-continuous barrier 3000–4000 m in elevation. Downstream, basin-and-range topography dominates the Intermountain West, with hundreds of narrow, steeply sloped mountain ranges separated by broad alluvial basins and valleys. These ranges greatly complicate analysis and forecasting across the region.
Remarkably few studies, however, have examined intermountain (i.e., Nevada) cyclogenesis. The most comprehensive is Lee (1995) who found that 75% of intermountain cyclones form in southwesterly-to-westerly large-scale flow, which hydrostatically induces a surface trough to the lee of the Sierra Nevada. Cyclogenesis occurs as quasigeostrophic (QG) forcing for ascent moves over the lee trough. This evolution is broadly consistent with the view of lee cyclogenesis as a two-stage process, the first being the formation of a lee trough as cross-barrier flow develops in advance of an upper-level cyclonic potential vorticity (PV) anomaly (e.g., Buzzi and Tibaldi 1978; McGinley 1982; Tibaldi et al. 1990; Aebischer and Schär 1998). The second involves the slower, quasi-stationary intensification of the cyclone, as the upper-level cyclonic PV anomaly overtakes the lee trough. Surface development can be inferred through QG forcing for ascent (Hoskins et al. 1978), which stretches low-level vorticity within the lee trough, or as the phase locking of an upper-level cyclonic PV anomaly with a low-level cyclonic PV maximum manifest as a surface warm anomaly in the lee of the barrier (e.g., Hoskins et al. 1985; Bleck and Mattocks 1984; Mattocks and Bleck 1986). Theoretical studies suggest that this two-stage process is a reflection of baroclinic wave interaction with orography (e.g., Smith 1984; Tibaldi et al. 1990; Bannon 1992; Davis 1997; Davis and Stoelinga 1999).
The frontal life cycle of intermountain cyclones also remains largely unexplored. The first coherent model of midlatitude cyclone evolution was developed by the Bergen School early in the twentieth century and features the amplification of a frontal wave into an open-wave cyclone, culminating in the formation of an occluded front as the cold front overtakes the warm front (Bjerknes 1919; Bjerknes and Solberg 1922). Although some cases roughly follow this evolution (e.g., Bergeron 1959; Schultz and Mass 1993; Market and Moore 1998; Martin 1998; Schultz et al. 1998), countless deviations have been documented, many of which are related to orography (e.g., Bergeron 1937; Steinacker 1982; Schultz and Mass 1993; Steenburgh and Mass 1994; Hobbs et al. 1996; O’Handley and Bosart 1996; Locatelli et al. 2002; Chien and Kuo 2006).
In one of the few studies of the frontal life cycle of cyclones over the Intermountain West, Horel and Gibson (1994) describe the role of orographic lift in the formation of a warm-core seclusion accompanying a deep tropospheric cyclonic circulation. Beyond this, knowledge of intermountain cyclone evolution has been limited historically by a lack of observations with sufficient spatial and temporal resolution to identify large-scale airmass and circulation changes generated by the region’s complex terrain (Steenburgh and Blazek 2001). Recently, however, the development of the high-density MesoWest cooperative networks (Horel et al. 2002), execution of field programs like the Intermountain Precipitation Experiment (Schultz et al. 2002), and advancement of high-resolution data assimilation and numerical modeling has enabled the detailed study of several intermountain cold fronts (e.g., Steenburgh and Blazek 2001; Schultz and Trapp 2003; Shafer et al. 2006; SS08; Steenburgh et al. 2009). In particular, SS08 identify a dramatic increase in the frequency of strong cold frontal passages between the Sierra–Cascade ranges and northern Utah. This increase appears to be related to the development of an area of confluence that develops downstream of the Sierra Nevada in large-scale southwesterly flow [hereafter the Great Basin confluence zone (GBCZ)] and initiates and/or enhances frontogenesis. Steenburgh et al. (2009) show that the GBCZ can also contribute to discrete frontal propagation.
This paper uses high-resolution analyses and MesoWest surface observations to describe the life cycle and frontal evolution of the Tax Day Storm. In particular, we identify the role of the GBCZ in frontal development and cyclogenesis, examine the contribution of diabatic processes to frontogenesis, and describe the complex frontal evolution that ultimately occurs over the basin-and-range topography of northern Utah.
2. Data and methods
Our analysis employs manual surface analyses, objective upper-air and surface analyses, and satellite and radar imagery. Manual surface analyses use MesoWest observations (Horel et al. 2002), which were quality controlled following Splitt and Horel (1998), with further subjective checks for spatial and temporal consistency during manual analysis. To limit pressure reduction artifacts, we use 1500-m pressure (calculated following Steenburgh and Blazek 2001) instead of sea level pressure since 1500 m is near the mean elevation of intermountain observing sites. Following Sanders (1999), we use the term front to denote the warm edge of a strong baroclinic zone accompanied by a wind shift and pressure trough, and baroclinic trough for a surface feature with a wind shift and/or pressure trough but relatively weak temperature contrast. Following Cohen and Kreitzberg (1997), we classify the GBCZ as an airstream boundary because it is a wind shift line separating two relatively distinct airstreams.
Many of the challenges associated with surface analysis (e.g., Sanders 1999) are exacerbated by the topography and climatology of the Intermountain West (e.g., Rossby 1934; Hill 1993; Schultz and Doswell 2000; Steenburgh and Blazek 2001). In particular, thermally and dynamically driven winds, valley and basin cold pools, and large diurnal temperature cycles complicate frontal analysis. Relying on an objective temperature or potential temperature gradient criteria for distinguishing between fronts, baroclinic troughs, and nonfrontal troughs can be problematic, but, in general, mobile fronts identified in this study are accompanied by a spatially and temporally coherent temperature change of at least 3°C (2 h)−1 [for comparison, SS08 use 7°C (2–3 h)−1 criteria to identify a strong cold front], whereas quasi-stationary fronts are accompanied by a temperature gradient of at least 8°C (110 km)−1 [following the definition of a strong front given by Sanders (1999)]. We use temperature tendency and gradient criteria to extract the maximum amount of information from the spatially and temporally heterogeneous MesoWest network. For example, in some regions observation density is low or there is a lack of stations at a similar elevation, making estimates of the horizontal temperature gradient near the front difficult (many MesoWest stations do not report pressure, precluding the calculation of potential temperature). In these regions, a station reporting at a high temporal frequency may enable identification of the intensity of mobile fronts.
Objective surface analyses are based on a version of the Advanced Regional Prediction System Data Assimilation System (ADAS; Brewster 1996; Xue et al. 2000, 2001, 2003) that was modified at the University of Utah to produce improved surface analyses over complex terrain (Lazarus et al. 2002; Myrick et al. 2005). The ADAS analyses, generated on a 5-km grid, assimilate MesoWest surface observations, with the second operational Rapid Update Cycle (RUC2; Benjamin et al. 2004), available on a 20-km grid, serving as a downscaled background analysis. The resulting ADAS analyses are interpolated to a 0.41° × 0.35° (∼40 km) latitude–longitude grid to provide a smoother analysis for presentation and diagnostic calculations. Subjective analyses of fronts, troughs, and airstream boundaries are overlaid on these plots for ease of interpretation. Trajectories and their related thermodynamic budgets are calculated from the interpolated ADAS analyses using a modified version of the traj.gs script from the Grid Analysis and Display System (GrADS) Script Library (available online at http://www.iges.org/grads/gadoc/library.html). Although these trajectories are two dimensional and confined to the surface, which limits accuracy compared to three-dimensional trajectories, they are still representative of the general airflow due to the high spatial and temporal resolution of the data, and the relatively short trajectory time (e.g., Haagenson et al. 1990). Comparison of trajectories and thermodynamic budgets computed using the interpolated analyses and the original 5-km analyses shows they are quantitatively and qualitatively similar.
Composite radar–satellite imagery analyses are generated by overlaying Next Generation Weather Radar (NEXRAD) level-II lowest-elevation scan (0.5°) radar reflectivity analyses from all available western U.S. radars on 4-km Geostationary Operational Environmental Satellite (GOES) infrared satellite imagery. It is important to note that radar coverage in many portions of the western United States is deficient or incomplete (Westrick et al. 1999; Maddox et al. 2002). This is especially true over portions of central Nevada and central Utah.
3. Results
a. Synoptic and mesoscale overview
At 0000 UTC (times are for 15 April 2002 unless otherwise stated), west-southwesterly flow in advance of an upper-level cyclonic potential vorticity (PV) anomaly and trough extends over the Pacific Northwest and Idaho (Fig. 2a). At 700 hPa, QG forcing for ascent lies over a broad area of northern California and northeast Nevada (Fig. 2b), just north of two weak 1500-m low centers (Fig. 3a). A third 1500-m low center, a so-called thermal low, lies over the Mojave Desert, as is common during the warm season (Rowson and Colucci 1992; Whiteman 2000, p. 165).
A surface-based cold front is advancing southward through southwest Wyoming and northern Utah (Fig. 3a), bringing temperature falls as large as ∼10°C (2 h)−1 (e.g., DPG03, Fig. 4a; station locations in Fig. 1). Differential surface sensible heating arising from inhomogeneous cloud cover may have contributed to frontal sharpening in this region, as described by Koch et al. (1995, 1997), Gallus and Segal (1999), and Segal et al. (2004). For example, the presence of a deep convective boundary layer at KSLC (32°C surface dewpoint depression, Fig. 5) indicates intense prefrontal surface heating; whereas cloud cover (Fig. 6a) likely limited such heating in the postfrontal environment (no soundings are available in this region). Sublimational and evaporative cooling produced by postfrontal precipitation (Fig. 6a) may also have contributed to frontal sharpening, as often occurs with intermountain cold fronts (e.g., Schultz and Trapp 2003; Steenburgh et al. 2009).
A more complex pattern exists farther to the west where the surface baroclinicity is weaker but somewhat enhanced in two areas (Figs. 2c and 3a). The first is immediately north of a baroclinic trough near the northern Nevada border. A lack of postfrontal clouds and precipitation (Fig. 6a) may explain the less abrupt nature of the temperature transition in this region compared to farther east. The second area of enhanced baroclinicity extends across Nevada between the baroclinic trough and a pronounced confluence zone that extends downstream from the high Sierra into northwest Utah (Figs. 2c and 3a). This GBCZ represents an airstream boundary between southerly flow over southeastern Nevada and western Utah and westerly flow over central Nevada and is embedded within a broad lower-tropospheric warm anomaly that extends from the surface to 700 hPa (Figs. 2b,c). Despite the wind shift and warm anomaly, there is little evidence of a coherent 1500-m surface trough, although observations are relatively sparse in this region (Fig. 3a).
The axes of the upper-level cyclonic PV anomaly and trough move over the Pacific coast by 0600 UTC (Fig. 7a). At 700 hPa, the two regions of QG forcing for ascent are now found over the northern California–Nevada border and southeastern Idaho. During this period, the 700-hPa baroclinic zone moves slowly southeastward over Nevada and strengthens as cooler air moves southward over Oregon, California, and northern Nevada (Fig. 7b). Concurrently, the 1500-m low center over northern California dissipates and pressures fall over northwest Nevada, but remain steady over northern Utah. A 1500-m low remains over the Mojave Desert but shifts northwestward as the Mojave cools and crest-level (700 hPa) cross-barrier flow intensifies, leading to pressure falls in the lee of the high Sierra (Figs. 3b and 7b).
The surface cold front becomes stationary over northern Utah shortly after passing KSLC at ∼0200 UTC (Figs. 3b and 4b), but continues to progress slowly southward over southwest Wyoming (Fig. 3b). Concurrently, postfrontal precipitation weakens and becomes more scattered (not shown). To the west, the baroclinic trough moves slowly southward toward the GBCZ, incorporating the intervening baroclinicity (Figs. 3b and 7c). Surface confluence associated with the GBCZ persists over southern and central Nevada, but weakens (cf. Figs. 2c and 7c), perhaps due to the nocturnal decoupling of surface winds.
The upper-level cyclonic PV anomaly and trough compact and dig through 1200 UTC (Fig. 8a) while strong 700-hPa QG forcing for ascent develops over Nevada (Fig. 8b). As a result, 1500-m pressure falls become increasingly concentrated near and along the GBCZ, which is becoming the locus for cyclogenesis (Fig. 3c). As flow across the Sierra Nevada intensifies, the lee trough deepens, shifts northward, and develops a barrier-parallel orientation [similar to, e.g., Steenburgh and Mass (1994) for the Colorado Rockies].
During this period, the 700-hPa thermal wave amplifies (Fig. 8b), and, in response, the former surface cold front over northern Utah and southwest Wyoming retreats slowly northward as a warm front (Figs. 3c and 8c).1 Above the surface-based nocturnal inversion, veering winds and a pronounced stable layer are consistent with the development of warm advection within the frontal zone over KSLC (Fig. 5). This shallow frontal zone extends upward only to 742 hPa (below the mountain ridges surrounding the Salt Lake Valley). Over northern Nevada, comparison of the 0000 and 1200 UTC KLKN soundings reveals cooling throughout the troposphere, but especially below 600 hPa where temperatures have decreased 6°–11°C (Fig. 5). At the surface, the baroclinic trough continues to progress southward toward the GBCZ (Figs. 3c and 8c). In addition, cold advection accompanying the baroclinic trough intensifies as surface winds veer to northerly and strengthen over northwest Nevada. With strong QG forcing for ascent over the region, radar echoes begin to develop above the relatively dry, well-mixed, low-level environment found over northeast Nevada, and a large cirrus shield moves eastward over Utah (Figs. 5, 6b, and 8b).
The upper-level cyclonic PV anomaly and trough take on a negative tilt as they move eastward through 1800 UTC (Fig. 9a). At 700 hPa, the area of QG forcing for ascent becomes more elongated, matching well the orientation of the 1500-m trough and developing cold front (Figs. 3d and 9b). The 1500-m low center is found along the warm front over the Great Salt Lake basin (Fig. 3d), the lowest, broadest topographic basin along the trough axis. As the GBCZ collects and concentrates baroclinicity from the north, including that accompanying the approaching baroclinic trough, a cold front forms over central Nevada and rapidly strengthens as it moves east (Figs. 3d and 9c). Low-level sublimational and evaporative cooling likely also contribute to frontogenesis and frontal scale collapse, as satellite and radar observations show the development of a solid band of postfrontal clouds and precipitation, in contrast to diurnal surface heating occurring in the cloud-free prefrontal environment (cf. Figs. 6c and 9c). Passage of the now rapidly developing cold front is first observed at KTPH at 1545 UTC (Fig. 4c), with a substantially stronger passage at KELY at 1750 UTC (Figs. 3d, 4d, and 9c), the latter observing an 11°C (2 h)−1 temperature fall. Over Utah and southwest Wyoming the entire frontal boundary progresses slowly northward as a warm front that is not yet structurally continuous with the cold front over Nevada (Fig. 3d).
As the upper-level cyclonic PV anomaly and trough advance eastward (not shown), the 1500-m low center reaches its minimum central pressure of 824 hPa at KSLC at 2100 UTC (Fig. 3e), and the corresponding 982-hPa-reduced sea level pressure becomes the second lowest ever reported in Utah. By this time the entire frontal boundary, including the former warm front, reaches its maximum intensity and progresses southward and eastward as a cold front. The accompanying 17°C (2 h)−1 temperature fall at KSLC (Fig. 4b) is the largest in the 25-yr Salt Lake City cold front climatology developed by SS08. At the University of Utah, 12 km east of KSLC, a meteorology student observes a 7°C (10 s)−1 drop with frontal passage using a handheld instrument. Many lowland stations in northern Utah report peak gusts in excess of 25 m s−1 in the pre- and postfrontal environments, with substantially stronger gusts at some stations (e.g., KU24, 32 m s−1; KU42, 35 m s−1; PEM, 47 m s−1). Satellite imagery derived from the Advanced Very High Resolution Radiometer aboard the National Oceanic and Atmospheric Administration (NOAA)-16 polar-orbiting satellite at 2046 UTC reveals widespread blowing dust across much of the Mojave Desert and southern Nevada (Fig. 10). During this period, KLAS reports 20–25 m s−1 wind gusts and ≤400-m visibility under otherwise clear skies. Blowing dust over Utah, evident between the cloud streets within the dry slot, forces the closure of Interstate 15 and leaves a brown layer of dust on the snowpack of the Wasatch Mountains and other ranges. Therefore, we hypothesize that high winds associated with intermountain cyclones contribute to episodic dust deposition onto the mountain snowpack of the region, decreasing albedo, and as discussed by Painter et al. (2007), leading to an earlier melt out of this important natural water resource.
b. Kinematic frontogenesis diagnostics
The analysis above suggests that the cold front accompanying the Tax Day Storm forms as confluent deformation and convergence along the GBCZ, which is initially nonfrontal, collect and concentrate baroclinity from the northern Great Basin. It is along this incipient frontal zone that the 1500-m cyclone forms, implying that the GBCZ plays a central role in both frontogenesis and cyclogenesis. In this section we further investigate the GBCZ.
In the Tax Day Storm, the GBCZ is a persistent, quasi-stationary airstream boundary that develops by 1800 UTC 14 April (not shown). At 0000 UTC 15 April, contraction associated with the GBCZ extends downstream from the high Sierra into northern Utah (Fig. 11a). This pattern persists through 1200 UTC when surface winds are weaker (Figs. 3c and 8c) and contraction associated with the GBCZ is less coherent, but still visible (Fig. 11b). At this time, weak kinematic frontogenesis is found over west-central Utah where baroclinicity associated with the cold front that sagged southward through northern Utah at 0000 UTC (Fig. 11a) merges with the GBCZ (Fig. 11b). By 1500 UTC, the GBCZ is drawing baroclinicity over northern and central Nevada into an increasingly favorable frontogenetical environment (Fig. 11c). In particular, isentropes in this region are oriented nearly parallel to the axes of dilatation along the GBCZ.
The contraction maximum associated with the GBCZ becomes more meridionally oriented over Nevada between 1500 and 1700 UTC but remains quasi-stationary over northern Utah (cf. Figs. 11c,d). Since the GBCZ appears tied to the high Sierra, its orientation appears to be a response to the backing of the ambient crest-level flow as the wave amplifies. By 1700 UTC, the baroclinic zone is collocated with the contraction maxima, the isentropes are aligned along the axes of dilatation, and the resulting kinematic frontogenesis is maximized along the incipient cold front over Nevada and the existing front over northern Utah. These two kinematic frontogenesis maxima are distinct, which likely leads to the lack of structural continuity between the warm and cold fronts observed at 1800 UTC (Fig. 3d).
Kinematic frontogenesis along the GBCZ intensifies rapidly from 1700 to 1800 UTC (cf. Figs. 11d,e) as maximum baroclinicity doubles from ∼15 K (100 km)−1 to ∼30 K (100 km)−1 (not explicitly shown) and kinematic frontogenesis more than triples. Several interrelated processes likely contribute to this apparent frontal scale collapse. The first is an increase in surface winds and associated kinematic frontogenesis as the stable nocturnal boundary layer erodes and the pressure trough deepens. The second is geostrophic adjustment, which has been shown to enhance the cross-front secondary circulation in developing fronts (Hoskins and Bretherton (1972) and Eliassen (1990). Finally, cross-front diabatic heating and cooling contrasts may also contribute, as discussed in the next section and discussed by Koch et al. (1995, 1997), Gallus and Segal (1999), and Segal et al. (2004).
By 2100 UTC contraction, baroclinicity, and kinematic frontogenesis reach their peak intensity (Fig. 11f). The 7°C (10 s)−1 drop accompanying frontal passage at the University of Utah, combined with a frontal speed of 5.6 m s−1, yields a localized temperature gradient of 12.5 K (100 m)−1, somewhat stronger than the 6 K (100 m)−1 reported for another intermountain cold front by Schultz and Trapp (2003).
Averaging contraction and axes of dilatation from 2100 UTC 14 April–2100 UTC 15 April illustrates the cumulative effect of the quasi-stationary GBCZ (Fig. 12a). Contraction associated with the GBCZ initiates frontogenesis as it collects the baroclinicity from the northern Great Basin. This is well illustrated by a time series of average baroclinicity, contraction rate, and kinematic frontogenesis in the region surrounding the GBCZ (Fig. 12b). Kinematic frontogenesis first increases as the initial cold front pushes into northern Utah at ∼0000 UTC, and contraction reaches a maximum with the stronger surface winds in the late afternoon. Overnight, contraction decreases but the mean baroclinicity slowly increases in response to weak kinematic frontogenesis. From 1500 to 2100 UTC contraction and baroclinicity increase nonlinearly as the GBCZ collects and concentrates baroclinicity, kinematic frontogenesis intensifies, and frontal-scale collapse occurs. Kinematic frontogenesis increases most dramatically, being proportional to the product of baroclinicity and contraction. Differential diabatic processes likely also contribute to the frontal development, as discussed in the next section.
c. Trajectories and diabatic processes
As suggested previously, diabatic processes appear to contribute to frontogenesis directly through cross-front contrasts in sensible heating and postfrontal sublimational and evaporative cooling (see Fig. 6c). They may also contribute indirectly by inducing a thermally forced, cross-front circulation that can nonlinearly enhance frontal-scale collapse (e.g., Koch et al. 1995, 1997; Gallus and Segal 1999; Segal et al. 2004). Although it is not possible to fully quantify the direct and indirect effects of differential diabatic heating due to the complex nonlinear feedbacks involved (e.g., Koch et al. 1995), changes in potential temperature along two-dimensional ADAS surface trajectories provide an estimate of direct diabatic contributions to frontogenesis and the potential for enhanced frontal-scale collapse.
Backward 9-h trajectories that begin at 1200 UTC (just prior to sunrise) and terminate on the warm and cold sides of the frontal zone at 2100 UTC (mid-afternoon, time of maximum frontal strength) are shown in Fig. 13a, with beginning, ending, and net change in potential temperature for the trajectories presented in Table 1. Potential temperature increases along all prefrontal trajectories. Trajectories 4–11 and 13–14, which originate over the lower Colorado River Valley and south end of the high Sierra, respectively (Figs. 8c and 9c), experience the largest increases (3–8 K).
In contrast, there is a net decrease or little change in potential temperature along most of the postfrontal trajectories. Trajectories 15–17 begin over the Snake River Plain and experience limited sensible heating and a period of precipitation as they move into northern Utah (Fig. 13 and Table 1), yielding a net increase in potential temperature of <1 K. In contrast, trajectories 18–23 begin with a higher potential temperature but are consistently beneath the cloud and precipitation shield that develops after 1200 UTC (Fig. 6c). As a result, the potential temperature along these trajectories decreases 3–7 K. Farther west, trajectories 24–28 are located near or west of the postfrontal precipitation band (Fig. 6c). Trajectories 24 and 25 experience little change, whereas trajectories 26–28 increase 1–4 K.
Collectively, these trajectories illustrate that differential diabatic heating–cooling makes a substantial direct contribution to frontogenesis along much of the front. Most trajectories in the prefrontal air mass undergo several kelvin of diabatic warming from 1200 to 2100 UTC, whereas postfrontal trajectories beneath the postfrontal precipitation band experience diabatic cooling or little potential temperature change, particularly those terminating over eastern Nevada and western Utah where they remain under the cloud and precipitation shield for a sustained period of time. The estimated direct differential diabatic contribution to frontogenesis during this 9-h period ranges from ∼1 K (100 km)−1 along the southwest end of the front, to ∼5 K (100 km)−1 (∼40% of cross-front baroclinicity) along the central part, to ∼0 K (100 km)−1 at the northeast end.
d. Front–mountain interactions over northern Utah
We now focus on the smaller-scale front–mountain interactions observed over the basin-and-range topography of northern Utah during this event. The surface cold front moves southward into northern Utah at approximately 0000 UTC (Fig. 3a) and by 0400 UTC extends across the basins and valleys south and west of the Great Salt Lake (Fig. 14a).2 The front briefly passes south of LOFUT in the southern Rush Valley at 0515 UTC (Fig. 4e) before retreating just north of LOFUT at 0545 UTC, where it remains stationary through ∼1200 UTC.
As the thermal wave begins to amplify at ∼1200 UTC, the frontal zone begins a complex northward progression as a warm front. Starting with VENU1 at 1225 UTC (Fig. 4f), the warm front moves continuously northward through the Rush Valley (Figs. 14b,c). In contrast, frontal movement over Utah Valley to the east is not continuous. Stations along the valley floor observe abrupt warming at irregular times over a 2.5-h period as the shallow cold air mass is turbulently eroded by increasing southerly and southwesterly flow aloft (1200 UTC KSLC sounding, Fig. 5). For example, QLN (Fig. 4g, 1451 m) warms at 1125 UTC, whereas KPVU (Fig. 4h, 1371 m), 13 km to the south, does not warm until 1335 UTC.
The warm front moves continuously northward through the Salt Lake and Tooele Valleys, but farther west remains stationary through 1500 UTC near the Onaqui Mountains (Fig. 14c). By 1800 UTC the entire frontal boundary is progressing northward and westward as a warm front that stretches from the Great Salt Lake Desert eastward across the northern ends of the Tooele and Salt Lake Valleys (Fig. 14d). This progression continues until ∼2015 UTC when the 1500-m low center moves through the region, and by 2100 UTC the entire front moves southward and eastward as a cold front (Fig. 3e). Along the south shore of the Great Salt Lake, abrupt warming occurs as the warm front approaches LAK from the south, followed immediately by abrupt cooling as the front reverses direction (Fig. 4i). At KSLC the passages of both a warm (1755 UTC) and cold front (2055 UTC) are apparent (Fig. 4b).
As the cold front moves south and east, the movement of the low center along the frontal boundary and the influence of the local basin-and-range terrain results in the western segment of the cold front moving rapidly eastward while the middle and eastern segments move slowly southward (cf. Figs. 14d,e). As a result, over the Rush Valley and portions of Utah Valley, a cold frontal segment passes first from the west, followed by a second cold frontal segment from the north (cf. Figs. 14e,f).
The meteogram from MS6 in the Rush Valley summarizes the entire event well, including the passage of the two frontal segments (Fig. 4j). The antecedent cold front moves southward through northern Utah, passing MS6 at 0310 UTC, after which it becomes stationary just to the south. It begins to move northward at ∼1200 UTC, passing MS6 as a warm front at 1440 UTC. Over the Rush Valley where MS6 is located, the frontal movement is continuous, but over Utah Valley the movement is discontinuous as the shallow cold air mass is removed by turbulent erosion. The warm front moves northward to near the southern shore of the Great Salt Lake, where, at ∼2015 UTC, it reverses directions and becomes a cold front. Because of the interaction of the frontal boundary and low center with the local basin-and-range topography, MS6 observes the passage of two cold frontal segments, one from the west at 2110 UTC and one from the north at 2210 UTC.
4. Conclusions
This paper utilizes high-density surface observations and high-resolution surface analyses to examine the life cycle of an intermountain cyclone and its attendant fronts. The cyclone, known locally as the “Tax Day Storm” produced the second lowest sea level pressure ever reported in Utah and the strongest cold front passage observed at KSLC in the past 25 yr.
As an upper-level PV anomaly and trough approaches the Pacific coast, an airstream boundary, which we have dubbed the Great Basin confluence zone (GBCZ), forms downstream from the Sierra Nevada. Contemporaneously, a baroclinic trough, which was initially draped along northern Nevada border, moves southward. Ultimately the GBCZ collects and concentrates the baroclinity from the northern Great Basin, including that associated with the approaching baroclinic trough. As the upper-level PV anomaly and trough crest the Sierra–Cascade ranges, QG forcing for ascent becomes focused over Nevada and rapid cyclogenesis and frontogenesis occur along the GBCZ. As the system rapidly amplifies, a cloud and precipitation shield develops over the postfrontal air mass, contributing to diabatic frontogenesis and apparently aiding the dramatic frontal-scale collapse that occurs from 1700 to 1800 UTC over Nevada. Over northern Utah, front–mountain interactions lead to a complex frontal evolution including dramatic frontal distortions, apparent discontinuous movement of a warm front due to the turbulent erosion of a shallow, persistent cold pool within Utah Valley, and the passage of multiple cold frontal segments over the Rush Valley.
The GBCZ is a distinctive and important mesoscale feature in this event. Plots of contraction, baroclinity, and frontogenesis illustrate that confluent deformation and convergence associated with the GBCZ collect and organize baroclinicity in the region over an extended period of time. Surface cyclogenesis occurs along the resulting frontal zone as opposed to within the barrier-parallel lee trough as might be inferred from other lee cyclone studies. The quasi-stationary nature of the GBCZ suggests that it is tied to topography, namely the high Sierra. Although other mountain-induced confluence–convergence zones have been identified elsewhere (e.g., Mass 1981; Doyle 1997; Andretta and Hazen 1998; Antonescu and Burcea 2010), the role the GBCZ plays in the genesis of cyclones and fronts over the Great Basin seems to be unique.
Diabatic processes such as differential sensible heating due to cross-front contrasts in cloud cover, and subcloud diabatic cooling from precipitation also contribute significantly to the frontal development in this case. Potential temperature changes along trajectories indicate that direct diabatic contributions, although varying along the front, explain up to 40% of the cross-front baroclinicity. The dramatic scale collapse of the front from 1700 to 1800 UTC likely results from the combined effects of the GBCZ and geostrophic adjustment, which may be nonlinearly enhanced by thermally induced circulations from diabatic processes (as in Koch et al. 1995). This is one of several intermountain front cases where diabatic processes are implicated or shown to be important to frontogenesis (e.g., Schultz and Trapp 2003; SS08; Steenburgh et al. 2009).
Finally, the frontal analyses over northern Utah illustrate just some of the complications of analysis and forecasting over the Intermountain West, including the effects of mesoscale topography on frontal movement and structure. The Tax Day Storm provides an example of a significant deviation from the classical Norwegian and Shapiro and Keyser (1990) cyclone models, and illustrates the complex interactions of large-scale forcing with the terrain of the Intermountain West at various scales. This research furthers our understanding and prediction of these events, and, in particular, provides the first detailed examination of the GBCZ. Uncertainties regarding the underlying mechanisms responsible for the GBCZ, given its role in the formation of cyclones and strong cold fronts that bring threats to life and property in the most rapidly growing region in the United States, warrant further research.
Acknowledgments
We thank Lance Bosart, David Schultz, and Robert Cohen for their insights and assistance during the development of this research. The comments from two anonymous reviewers contributed to an improved manuscript. We gratefully acknowledge the provision of datasets, software, computer time, and/or services by the Unidata Program Center, National Centers for Environmental Prediction, University of Utah MesoWest program and contributors, University of Utah Center for High Performance Computing, Erik Crosman, Sebastian Hoch, Don Murray, and David Taylor. This research is supported by the National Oceanic and Atmospheric Administration CSTAR Program and National Science Foundation under Grant ATM-0627937. Any opinions, findings, and conclusions or recommendations expressed in this material are those of the authors and do not necessarily reflect the views of the National Science Foundation.
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Topography and geography of (a) the Intermountain West and surrounding region [Great Basin boundary from the United States Geological Survey National Wetlands Research Center (2010) indicated by thick line] and (b) the northern Utah study area [indicated by the white box in (a)]. Terrain height (m) shaded according to scale at lower left in (a). Geographic features and observing stations discussed in text are annotated.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
Topography and geography of (a) the Intermountain West and surrounding region [Great Basin boundary from the United States Geological Survey National Wetlands Research Center (2010) indicated by thick line] and (b) the northern Utah study area [indicated by the white box in (a)]. Terrain height (m) shaded according to scale at lower left in (a). Geographic features and observing stations discussed in text are annotated.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
Topography and geography of (a) the Intermountain West and surrounding region [Great Basin boundary from the United States Geological Survey National Wetlands Research Center (2010) indicated by thick line] and (b) the northern Utah study area [indicated by the white box in (a)]. Terrain height (m) shaded according to scale at lower left in (a). Geographic features and observing stations discussed in text are annotated.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
Analyses at 0000 UTC 15 Apr 2002. (a) RUC2 dynamic tropopause (2 PVU) pressure (hPa, shaded following inset scale), wind vectors (m s−1, following inset scale), and isotachs (contoured every 7.5 m s−1). (b) RUC2 700-hPa temperature (contoured every 2°C), wind vectors (m s−1, following inset scale), and 700-hPa QG forcing for ascent (×10−15 K m−2 s−1, shaded following inset scale). (c) ADAS surface analysis with terrain height (m, shaded following inset scale), potential temperature (contoured every 2 K), and wind vectors (m s−1, following inset scale). Manual frontal analysis is from Fig. 3.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
Analyses at 0000 UTC 15 Apr 2002. (a) RUC2 dynamic tropopause (2 PVU) pressure (hPa, shaded following inset scale), wind vectors (m s−1, following inset scale), and isotachs (contoured every 7.5 m s−1). (b) RUC2 700-hPa temperature (contoured every 2°C), wind vectors (m s−1, following inset scale), and 700-hPa QG forcing for ascent (×10−15 K m−2 s−1, shaded following inset scale). (c) ADAS surface analysis with terrain height (m, shaded following inset scale), potential temperature (contoured every 2 K), and wind vectors (m s−1, following inset scale). Manual frontal analysis is from Fig. 3.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
Analyses at 0000 UTC 15 Apr 2002. (a) RUC2 dynamic tropopause (2 PVU) pressure (hPa, shaded following inset scale), wind vectors (m s−1, following inset scale), and isotachs (contoured every 7.5 m s−1). (b) RUC2 700-hPa temperature (contoured every 2°C), wind vectors (m s−1, following inset scale), and 700-hPa QG forcing for ascent (×10−15 K m−2 s−1, shaded following inset scale). (c) ADAS surface analysis with terrain height (m, shaded following inset scale), potential temperature (contoured every 2 K), and wind vectors (m s−1, following inset scale). Manual frontal analysis is from Fig. 3.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
Manual 1500-m pressure (every 2 hPa with leading 8 omitted) and surface frontal analyses at (a) 0000, (b) 0600, (c) 1200, (d) 1800, and (e) 2100 UTC 15 Apr 2002. Cold, warm, and stationary fronts identified with traditional symbols (stationary fronts white). GBCZ and baroclinic trough denoted with dotted and dashed line, respectively. Surface station models include wind (pennant, full barb, and half barb denote 25, 5, and 2.5 m s−1, respectively) and temperature (°C, at upper right).
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
Manual 1500-m pressure (every 2 hPa with leading 8 omitted) and surface frontal analyses at (a) 0000, (b) 0600, (c) 1200, (d) 1800, and (e) 2100 UTC 15 Apr 2002. Cold, warm, and stationary fronts identified with traditional symbols (stationary fronts white). GBCZ and baroclinic trough denoted with dotted and dashed line, respectively. Surface station models include wind (pennant, full barb, and half barb denote 25, 5, and 2.5 m s−1, respectively) and temperature (°C, at upper right).
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
Manual 1500-m pressure (every 2 hPa with leading 8 omitted) and surface frontal analyses at (a) 0000, (b) 0600, (c) 1200, (d) 1800, and (e) 2100 UTC 15 Apr 2002. Cold, warm, and stationary fronts identified with traditional symbols (stationary fronts white). GBCZ and baroclinic trough denoted with dotted and dashed line, respectively. Surface station models include wind (pennant, full barb, and half barb denote 25, 5, and 2.5 m s−1, respectively) and temperature (°C, at upper right).
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
Meteograms of temperature (solid), altimeter setting (dashed, where available), and wind (full and half barb denote 5 and 2.5 m s−1, respectively) at (a) DPG03, (b) KSLC, (c) KTPH, (d) KELY, (e) LOFUT, (f) VENU1, (g) QLN, (h) KPVU, (i) LAK, and (j) MS6. Cold and warm frontal passages denoted by black and gray vertical lines, respectively. See Fig. 1 for station locations.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
Meteograms of temperature (solid), altimeter setting (dashed, where available), and wind (full and half barb denote 5 and 2.5 m s−1, respectively) at (a) DPG03, (b) KSLC, (c) KTPH, (d) KELY, (e) LOFUT, (f) VENU1, (g) QLN, (h) KPVU, (i) LAK, and (j) MS6. Cold and warm frontal passages denoted by black and gray vertical lines, respectively. See Fig. 1 for station locations.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
Meteograms of temperature (solid), altimeter setting (dashed, where available), and wind (full and half barb denote 5 and 2.5 m s−1, respectively) at (a) DPG03, (b) KSLC, (c) KTPH, (d) KELY, (e) LOFUT, (f) VENU1, (g) QLN, (h) KPVU, (i) LAK, and (j) MS6. Cold and warm frontal passages denoted by black and gray vertical lines, respectively. See Fig. 1 for station locations.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
Skew T–logp diagrams (temperature and dewpoint) for KSLC and KLKN at 0000 UTC 15 Apr (black), 1200 UTC 15 Apr (medium gray), and 0000 UTC 16 Apr 2002 (light gray, KSLC sounding not successfully launched). Vertical wind profiles (pennant, full barb, and half barb denote 25, 5, and 2.5 m s−1, respectively) to the right with day/hour (UTC) as indicated.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
Skew T–logp diagrams (temperature and dewpoint) for KSLC and KLKN at 0000 UTC 15 Apr (black), 1200 UTC 15 Apr (medium gray), and 0000 UTC 16 Apr 2002 (light gray, KSLC sounding not successfully launched). Vertical wind profiles (pennant, full barb, and half barb denote 25, 5, and 2.5 m s−1, respectively) to the right with day/hour (UTC) as indicated.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
Skew T–logp diagrams (temperature and dewpoint) for KSLC and KLKN at 0000 UTC 15 Apr (black), 1200 UTC 15 Apr (medium gray), and 0000 UTC 16 Apr 2002 (light gray, KSLC sounding not successfully launched). Vertical wind profiles (pennant, full barb, and half barb denote 25, 5, and 2.5 m s−1, respectively) to the right with day/hour (UTC) as indicated.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
Infrared satellite and lowest-elevation scan (0.5°) radar reflectivity composite at (a) 0000, (b) 1200, and (c) 1800 UTC 15 Apr 2002. Reflectivity scale (dBZ) in lower right of (a). Radar imagery collected 0–10 min after top of hour except at 1153 UTC at KEYX in (b). Radar imagery missing from KMAX and KHDX in (b) and KRTX and KVBX in (c). Manual frontal analyses are from Fig. 3.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
Infrared satellite and lowest-elevation scan (0.5°) radar reflectivity composite at (a) 0000, (b) 1200, and (c) 1800 UTC 15 Apr 2002. Reflectivity scale (dBZ) in lower right of (a). Radar imagery collected 0–10 min after top of hour except at 1153 UTC at KEYX in (b). Radar imagery missing from KMAX and KHDX in (b) and KRTX and KVBX in (c). Manual frontal analyses are from Fig. 3.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
Infrared satellite and lowest-elevation scan (0.5°) radar reflectivity composite at (a) 0000, (b) 1200, and (c) 1800 UTC 15 Apr 2002. Reflectivity scale (dBZ) in lower right of (a). Radar imagery collected 0–10 min after top of hour except at 1153 UTC at KEYX in (b). Radar imagery missing from KMAX and KHDX in (b) and KRTX and KVBX in (c). Manual frontal analyses are from Fig. 3.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
As in Fig. 2, but at 0600 UTC 15 Apr 2002.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
As in Fig. 2, but at 0600 UTC 15 Apr 2002.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
As in Fig. 2, but at 0600 UTC 15 Apr 2002.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
As in Fig. 2, but at 1200 UTC 15 Apr 2002.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
As in Fig. 2, but at 1200 UTC 15 Apr 2002.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
As in Fig. 2, but at 1200 UTC 15 Apr 2002.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
As in Fig. 2, but at 1800 UTC 15 Apr 2002.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
As in Fig. 2, but at 1800 UTC 15 Apr 2002.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
As in Fig. 2, but at 1800 UTC 15 Apr 2002.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
False-color 0.63-μm reflectance from the Advanced Very High Resolution Radiometer aboard the NOAA-16 polar-orbiting satellite at 2046 UTC 15 Apr 2002. Inset: Channel 4–5 (11–12 μm) brightness temperature difference, with large differences typically associated with volcanic or dust aerosols colored (see Iino et al. 2004 for methodology).
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
False-color 0.63-μm reflectance from the Advanced Very High Resolution Radiometer aboard the NOAA-16 polar-orbiting satellite at 2046 UTC 15 Apr 2002. Inset: Channel 4–5 (11–12 μm) brightness temperature difference, with large differences typically associated with volcanic or dust aerosols colored (see Iino et al. 2004 for methodology).
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
False-color 0.63-μm reflectance from the Advanced Very High Resolution Radiometer aboard the NOAA-16 polar-orbiting satellite at 2046 UTC 15 Apr 2002. Inset: Channel 4–5 (11–12 μm) brightness temperature difference, with large differences typically associated with volcanic or dust aerosols colored (see Iino et al. 2004 for methodology).
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
ADAS contraction (×10−4 s−1, shaded according to inset scale), kinematic frontogenesis [black contours every 3 K (100 km)−1 h−1], local orientation of axes of dilatation scaled by magnitude of contraction (×10−4 s−1, according to inset scale), and potential temperature (gray contours every 2 K) at (a) 0000, (b) 1200, (c) 1500, (d) 1700, (e) 1800, and (f) 2100 UTC 15 Apr 2002.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
ADAS contraction (×10−4 s−1, shaded according to inset scale), kinematic frontogenesis [black contours every 3 K (100 km)−1 h−1], local orientation of axes of dilatation scaled by magnitude of contraction (×10−4 s−1, according to inset scale), and potential temperature (gray contours every 2 K) at (a) 0000, (b) 1200, (c) 1500, (d) 1700, (e) 1800, and (f) 2100 UTC 15 Apr 2002.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
ADAS contraction (×10−4 s−1, shaded according to inset scale), kinematic frontogenesis [black contours every 3 K (100 km)−1 h−1], local orientation of axes of dilatation scaled by magnitude of contraction (×10−4 s−1, according to inset scale), and potential temperature (gray contours every 2 K) at (a) 0000, (b) 1200, (c) 1500, (d) 1700, (e) 1800, and (f) 2100 UTC 15 Apr 2002.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
(a) Temporally averaged contraction (×10−4 s−1, shaded according to inset scale) and local orientation of axes of dilatation scaled by magnitude of contraction (×10−4 s−1, according to inset scale) from 2100 UTC 14 Apr–2100 UTC 15 Apr 2002. (b) Area averaged contraction (×10−4 s−1), baroclinicity [K (100 km)−1], and kinematic frontogenesis [×10−1 K (100 km)−1 h−1] within the quadrilateral denoted in (a).
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
(a) Temporally averaged contraction (×10−4 s−1, shaded according to inset scale) and local orientation of axes of dilatation scaled by magnitude of contraction (×10−4 s−1, according to inset scale) from 2100 UTC 14 Apr–2100 UTC 15 Apr 2002. (b) Area averaged contraction (×10−4 s−1), baroclinicity [K (100 km)−1], and kinematic frontogenesis [×10−1 K (100 km)−1 h−1] within the quadrilateral denoted in (a).
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
(a) Temporally averaged contraction (×10−4 s−1, shaded according to inset scale) and local orientation of axes of dilatation scaled by magnitude of contraction (×10−4 s−1, according to inset scale) from 2100 UTC 14 Apr–2100 UTC 15 Apr 2002. (b) Area averaged contraction (×10−4 s−1), baroclinicity [K (100 km)−1], and kinematic frontogenesis [×10−1 K (100 km)−1 h−1] within the quadrilateral denoted in (a).
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
The 9-h backward trajectories (1200–2100 UTC 15 Apr 2002, number indicates position at 2100 UTC) terminating on the warm and cold side of the baroclinic zone. Potential temperature at 2100 UTC contoured every 2 K.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
The 9-h backward trajectories (1200–2100 UTC 15 Apr 2002, number indicates position at 2100 UTC) terminating on the warm and cold side of the baroclinic zone. Potential temperature at 2100 UTC contoured every 2 K.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
The 9-h backward trajectories (1200–2100 UTC 15 Apr 2002, number indicates position at 2100 UTC) terminating on the warm and cold side of the baroclinic zone. Potential temperature at 2100 UTC contoured every 2 K.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
Manual surface frontal analyses over northern Utah at (a) 0400, (b) 1300, (c) 1500, (d) 1800, (e) 2100, and (f) 2200 UTC 15 Apr 2002. Station models are as in Fig. 3; station names are labeled as they are referenced in text.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
Manual surface frontal analyses over northern Utah at (a) 0400, (b) 1300, (c) 1500, (d) 1800, (e) 2100, and (f) 2200 UTC 15 Apr 2002. Station models are as in Fig. 3; station names are labeled as they are referenced in text.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
Manual surface frontal analyses over northern Utah at (a) 0400, (b) 1300, (c) 1500, (d) 1800, (e) 2100, and (f) 2200 UTC 15 Apr 2002. Station models are as in Fig. 3; station names are labeled as they are referenced in text.
Citation: Monthly Weather Review 138, 7; 10.1175/2010MWR3274.1
The 1200 UTC, 2100 UTC, and net change in potential temperature (K) along trajectories displayed in Fig. 13.
The ADAS analysis (Fig. 8c) places the surface baroclinic zone too far south at this time compared to that inferred from MesoWest observations in our manual analysis.
Given the shallowness of the postfrontal air mass, and for ease of presentation, surface fronts are only analyzed over basin and valley regions.