Topographic and Diabatic Influences on Baroclinic Storm

TOPOGRAPHIC AND DIABATIC INFLUENCES ON BAROCLINIC STORM

EVOLUTION OVER THE INTERMOUNTAIN WEST

Jason C. Shafer

A dissertation submitted to the faculty of The University of Utah in partial fulfillment of the requirements for the degree of

Doctor of Philosophy

Department of Meteorology

The University of Utah

The evolution of baroclinic storm systems over the western United States is explored through case study and climatological approaches. The case study examines

IOP3 of the Intermountain Precipitation Experiment, and describes the influence of terrain on the evolution of a winter storm from the Pacific coast to northern Utah using a variety of field-program and traditional observations. Results show that the terrain had a dramatic effect on the storm’s low-level front/trough structure, including frontal distortion and discontinuous movement. Specifically, the low-level frontal structure was stripped or removed as the front traversed the Sierra Nevada. Downstream of the Sierra, a new surface pressure trough developed and moved rapidly over Nevada and into northern

Utah, eventually developing characteristics of a weak cold front. The mesoscale precipitation distribution over northern Utah was also strongly influenced by terrain- driven circulations.

The climatological study examined strong surface cold fronts over the western

United States from 1979–2003. Strong cold fronts were identified using standard hourly surface observations and the North American Regional Reanalysis (NARR). Results show that strong cold fronts are mainly continental and that the number of events increases heading eastward from the Sierra Nevada and Cascade ranges and across the Intermoun- tain region. To develop an understanding of how these strong Intermountain cold fronts develop, a NARR composite of the 25 strongest cold fronts at Salt Lake City, UT was produced. Most strong Intermountain cold fronts develop during the day in a deep, dry convective boundary layer, and intensify rapidly during the 12 hours leading up to frontal passage. Results show this intense frontogenesis is driven by a combination of large-scale and diabatic processes, and that the terrain may be important for determining the location of the incipient frontogenesis.

v TABLE OF CONTENTS

ABSTRACT ...... iv

ACKNOWLEDGMENTS ...... viii

Chapter

1. INTRODUCTION ...... 1

2. TERRAIN INFLUENCES ON SYNOPTIC STORM STRUCTURE AND MESOS- CALE PRECIPITATION DISTRIBUTION DURING IPEX IOP3 ...... 3

Introduction...... 3 Data and Methods ...... 6 Results...... 8 Synoptic evolution and vertical structure ...... 8 Mesoscale storm structure over northern Utah ...... 28 Prebaroclinic surface trough period ...... 28 Postbaroclinic surface trough period ...... 41 Analysis of orographic precipitation...... 46 Concluding Discussion ...... 50 Storm evolution over the western United States...... 50 Precipitation distribution over northern Utah ...... 53

3. INTERMOUNTAIN COLD FRONTS...... 56

Introduction...... 56 Data and Methods ...... 59 Data sources ...... 59 Identification of strong cold fronts ...... 61 Results...... 68 Spatial and temporal distribution of strong cold fronts ...... 68 Wind, cloud, and precipitation accompanying strong cold fronts ...... 75 Characteristics of strong Intermountain cold fronts ...... 82 Frontogenesis analysis ...... 108 Discussion and Conclusions ...... 124 4. SUMMARY AND CONCLUSIONS ...... 130

REFERENCES ...... 133

vii ACKNOWLEDGMENTS

I express much gratitude to my advisor, Jim Steenburgh, for his support and guid- ance throughout the past five years. I thank Jim for providing me the opportunity to curi- ously explore Intermountain weather and its paradoxical ways. I’d also like to thank the other members of my committee: Lance Bosart for his encouragement, John Horel for his

GEMPAK wisdom, Jan Paegle for his sincerity, and David Schultz for his tenacity. Spe- cial appreciation is expressed to my father, Harry “Fred” Shafer who opened my eyes to the magnificent natural world through numerous outdoor adventures, and for paying the cable bill so I could watch endless hours of The Weather Channel. I am grateful to my love, Kate Gold, for her tolerance and for her ability to identify graupel and virga.

The past five years in Utah have been quite fascinating and memorable. There are too many people to thank for making the Utah experience entertaining and enjoyable. I thank all of you with whom I’ve shared a passion for the atmosphere, the game of soccer

(go Team Virga!), the mountains, the great food and herbal tea, the snow, education, and life.

This work could not have been accomplished without the help of many. I thank

David White and Paul Neiman of NOAA’s Environmental Technology Laboratory for providing the California wind profiler data presented in Chapter 2. I thank Greg West for helping me get started with GrADS, and Lee Byerle who helped with GrADS arbitrary cross sections while deployed in Iraq. I thank Will Cheng for providing the data used to generate background topography in Chapter 3, Steve Nesbitt and Sally Bensen for sharing their IDL expertise, and Bryan White for keeping my computer alive. I thank Scott

Stephens at NCDC for providing the surface data used in Chapter 3 in a friendly format, and David Schultz for computing DCAPE in Chapter 3. I thank John M. Brown for pointing out important literature, and Elford Astling for offering a different perspective of cold frontal passages at Salt Lake City.

Funding for the analysis of IPEX IOP3 (Chapter 2) was provided by National

Science Foundation Grants (NSF) ATM-0085318 and ATM-0333525, and a series of grants to the Cooperative Institute for Regional Prediction from the NOAA C-STAR program. Comments from the thesis committee, and three anonymous reviewers improved

Chapter 2. The work in Chapter 3 was supported by the aformentioned NSF grants and a

NSF GK-12 fellowship (ATM-0338340) under the Water the Environment Science and

Teaching (WEST) grant.

ix CHAPTER 1

INTRODUCTION

The complex terrain and dry climate of the western United States have a dramatic effect on midlatitude baroclinic systems. The primary goal of this work is to unearth these effects through observational and climatological studies. This is accomplished by examin- ing the orographic modification of a landalling frontal cyclone (Chapter 2), and by developing a climatology of strong Intermountain cold fronts (Chapter 3). These two approaches are useful for better understanding the spectrum of frontal evolution over the

Intermountain West since the case study describes a weakening frontal system, whereas the climatology examines the development of strong cold fronts.

Findings from the case study show that low-level frontal structures undergo a dramatic transformation from the Pacific coast to the Intermountain region. This is the result of interactions with the Sierra Nevada Mountains that act to destroy low-level frontal structures. Findings from the case study also illustrate the importance of mesoscale terrain circulations on the precipitation distribution over northern Utah. Results from the case study have broad implications for understanding frontal evolution over the western United

States and mesoscale precipitation prediction over the Intermountain region. 2

A few times a year, the proper ingredients combine to produce intense, damaging cold fronts over the Intermountain West. Chapter 3 intends to provide the recipe for how such cold fronts develop. Because strong Intermountain cold fronts develop in a deep, dry convective boundary layer, this study has important implications for the life cycle of cold fronts in other arid and semi-arid regions of the world.

This remainder of this dissertation is organized as follows. Chapter 2 shares the published work of Shafer et al. (2005), which describes the effects of terrain on synoptic storm structure and mesoscale precipitation distribution during IOP3 of the Intermountain

Precipitation Experiment. Chapter 3 describes the strong cold front climatology, with emphasis on frontal development over the Intermountain West. Conclusions are provided in Chapter 4. CHAPTER 2

TERRAIN INFLUENCES ON SYNOPTIC STORM STRUCTURE AND

MESOSCALE PRECIPITATION DURING IPEX IOP3

Introduction A major winter storm affected much of the western United States on 12–13

February 2000. Across the lowlands of north-central California, heavy rains produced flooding, the collapse of a Home Depot roof in Colma, and more than 2 million dollars of damage (National Climatic Data Center (NCDC) 2000, p. 29). Mountain snows of 60–120 cm (2–4 ft) fell over the Sierra Nevada, and prefrontal wind gusts reaching 26 m s-1 uprooted trees across parts of southern California (NCDC 2000, p. 26–27). Over northern

Utah, the storm was responsible for 60–90 cm of mountain snow, and up to 2.1 cm of snow water equivalent (SWE) at lowland locations. Significant precipitation gradients were observed over northern Utah. For example, the mountain site Snowbasin (SNI, 2256 m) recorded 4.9 cm SWE, over twice the 2.1 cm which fell at Ogden (OGD, 1360 m), a lowland site only 15 km away. In addition, significant precipitation contrasts were observed between sites at comparable elevations. For example, the 2.1 cm of SWE at

OGD was nearly four times greater than the 0.6 cm observed 45 km away at Salt Lake

City (SLC, 1280 m). The storm system also featured a complex evolution from landfall along the California coast to the interior Intermountain region, including the development 4 of a complex vertical trough structure over the Sierra Nevada and Great Basin (Shafer

2002; Schultz et al. 2002).

This event occurred during the third intensive observing period (IOP3) of the

Intermountain Precipitation Experiment (IPEX, Schultz et al. 2002), a research program designed to improve the understanding and prediction of precipitation over the

Intermountain region of the western United States. IOP3 was characteristic of many

Intermountain cool-season storms where a landfalling Pacific storm system weakened and underwent a complex evolution across the Intermountain region. As will be discussed further within this chapter, the storm’s evolution and precipitation producing mechanisms were intrinsically related to the complex topography of the Intermountain region, which is characterized by basin-and-range terrain where narrow, steeply sloped mountain ranges are separated by broad lowland valleys and basins (Fig. 2.1a). For example, northern Utah features the meridionally oriented Stansbury, Oquirrh, and Wasatch Mountains, which rise

1500–2000 m above the surrounding lowlands (Fig. 2.1b).

This chapter intends to advance the understanding of cool-season Intermountain storm systems through two objectives. The first is to describe the large-scale storm evolution from the Pacific coast to northern Utah, with emphasis on the surface and vertical structure. The second is to examine the precipitation distribution over northern

Utah and its relationship to the interaction of the kinematic and thermodynamic structure of the storm with local topography. The vertical structure of this system underwent dramatic change from California to Utah, and the precipitation distribution across northern

Utah featured both similarities and departures from climatological precipitation versus altitude relationships. Readers are referred to Cox et al. (2005) for a detailed analysis of 5

a) In te r m o u n t S a i RNO in SAC e . r . R RMD r e a g . SFO i N o . e n v a d N a 3000 2500 2000 1500 500 300km

o r t h G e r r e n

a KMTX t

S . a S NI l W t OGD . L a a . s k NSSL5 a NSSL4 e t . c h .

S O SLC e t n a q t Great Salt Lake Basin n . u r

s Salt Lake a i

b r l Tooele r u Valley h r Valley C LN

y M . W

o a

u o s a N u n t n t c a

t h

i a

n 30km i Traverse n s D.PG s Mountains

Figure 2.1. Topography and major geographic features of (a) the western United States and (b) northern Utah. Elevation (m) based on scale at lower left in (a). Boxed area in (a) identifies location of IPEX target area in (b). 6 the kinematic structure of this event near the Wasatch Mountains, and Colle et al. (2005) for a microphysical analysis and numerical simulation of this event.

Data and Methods

The analyses presented in this chapter were performed using a variety of field program and conventional datasets. Surface observations were provided by the MesoWest cooperative networks (Horel et al. 2002), including data from approximately 2500 stations in the western United States and 250 stations across northern Utah. MesoWest observations were quality controlled by comparing the observed values to an estimate based on multivariate linear regression (Splitt and Horel 1998), and by spatial and temporal consistency checks performed while preparing manual analyses. Precipitation observations were quality controlled by Cheng (2001), and generally featured 0.01 inch resolution (amounts were converted to mm for publication). For a description of precipitation gauge characteristics see Cheng (2001). Special 1-sec Automated Surface

Observing System (ASOS) data was also downloaded from the National Climatic Data

Center (NCDC) and used for the surface time series.1 Since MesoWest observations are irregularly spaced horizontally and vertically, the Advanced Regional Prediction System

Data Assimilation System (ADAS, Lazarus et al. 2002) was used to generate 1-km gridded analyses of surface convergence and winds over northern Utah.

Gridded data used for most upper-air analyses was provided by the National

Centers for Environmental Prediction Rapid Update Cycle (RUC2; Benjamin et al. 2004), an hourly assimilation forecast cycle run at 40-km horizontal grid spacing, but available at

1. Although available at high temporal resolution, the special ASOS temperature observations are only to the nearest degree C. 7

60-km horizontal and 25-hPa vertical grid spacing. Comparison with observations and other gridded data suggested that these RUC2 analyses provided the best overall depiction of the storm evolution. However, ETA Data Assimilation System (EDAS) gridded data was also used because it featured a qualitatively more realistic representation of the potential vorticity banners generated downstream of the Sierra Nevada. MesoWest observations, RUC2, and EDAS gridded data were plotted using GEneral Meteorological

PAcKage (GEMPAK) software.

Upper-air data included radiosonde observations at the conventional times (0000 and 1200 UTC) and some special observations at intermediate intervals of three or six hours collected by National Weather Service (NWS) upper-air sites. Over northern Utah,

3-h upper-air soundings were provided by two National Severe Storm Laboratory mobile labs (NSSL4 and NSSL5), which were located 100 km and 10 km upstream of the

Wasatch Crest, respectively (Fig. 2.1b). In addition, 915 MHz wind profiler data was available from Dugway Proving Grounds (DPG, Fig. 2.1b), about 120 km southwest of

SLC, and at sites in California that were operated by the NOAA Environmental

Technology Laboratory (e.g., RMD, SAC, Fig. 2.1a). Flight-level data from the NOAA P-

3 research aircraft and Velocity Azimuth Display (VAD) winds from several Next-

Generation Radar (NEXRAD) sites were also used, but are not presented.

Western United States Weather Surveillance Radar-1988 Doppler (WSR-88D) radar mosaics were generated from data provided in NEXRAD Information

Dissemination Service (NIDS) format (Baer 1991), which has a spatial resolution of 2 km and an approximate reflectivity resolution of 5 dBZ. Many areas of the western United

States, such as central Utah, are not well sampled by the NEXRAD network owing to 8 beam blockage, poor spatial coverage, and overshooting (e.g., Westrick et. al 1999; Wood et al. 2003). Over northern Utah, WSR-88D radar analyses from Promontory Point

(KMTX, see Fig. 2.1b for location) were based on NIDS-formatted data with a spatial resolution of 1 km and an approximate reflectivity resolution of 5 dBZ.

Manual surface analyses were prepared following the methods described in section

3 of Steenburgh and Blazek (2001), where 1500-m pressure was used rather than sea level pressure since 1500 m is near the mean elevation of Intermountain observing sites.

Surface frontal analysis over the Intermountain region is complicated by numerous issues, including but not limited to the effects of thermally and terrain-driven circulations

(Stewart et al. 2002), diabatic processes (Schultz and Trapp 2003), and limitations of the observing network (Horel et al. 2002). Nonetheless, the traditional definition of a frontal zone as an elongated area of strong horizontal temperature gradient (e.g., Bluestein 1986;

Keyser 1986), with the front as the boundary on the warm side of the zone, was implemented with consideration of the aformentioned difficulties. Owing to the higher mean elevation of the Intermountain region, the remainder of this chapter will refer to vertical levels as follows: low-level from the surface to 700 hPa, mid-level 700–500 hPa, and upper-level above 500 hPa.

Results

Synoptic evolution and vertical structure

At 0000 UTC 12 February 2000, an occluded midlatitude cyclone was located off the west coast of the United States with widespread precipitation across central and northern California (Fig. 2.2), and extensive cloudiness extending across much of the 9

42 42 40 40 38

38 40

42 38 44 45 L 46 35 42 48 25 44 46 15 48

Figure 2.2. Surface and radar analyses at 0000 UTC 12 Feb 2000. NEXRAD radar mosaic (reflectivity scale at left) and manual 1500-m pressure analysis. Station reports include 1500-m pressure (tenths of hPa with leading 8 truncated), and wind [full (half) barbs denote 5 (2.5) m s-1]. 10 western United States (Fig. 2.3c). The associated 500-hPa shortwave trough axis and absolute vorticity maximum were positioned upstream of the surface cyclone (Fig. 2.3a).

The surface occluded front extended equatorward from the low center (Fig. 2.2), and, like many California occlusions (e.g., Elliot 1958), behaved more like a cold front. For example, as the occluded front moved across north-central California, 2–4°C surface temperature falls were observed at San Francisco (SFO, Fig. 2.4a) and Sacramento (SAC,

Fig. 2.4b), and radar imagery revealed that a narrow frontal rainband was coincident with the occluded front and embedded within the large-scale precipitation shield (not shown).

The large-scale baroclinity of the occluded front was very weak at 700 hPa (Fig. 2.3b), but

6 h radiosonde data showed about 2°C of cooling at 700 hPa. Southwesterly cross barrier flow produced a lee trough east of the Sierra Nevada (Fig. 2.2).

The vertical structure of the system was complex. Wind profiler observations near

SFO at Richmond (RMD) and SAC (see Fig. 2.1a for locations) indicated backing winds with height within the low-level frontal zone, which sloped rearward with height (Figs.

2.5a and 2.5b). Above the low-level frontal zone, however, the 0300 UTC (all times for 12

February 2000 unless noted) RUC2 analysis showed that the upper-level trough axis sloped forward with height from 600 hPa to 250 hPa, with the leading edge of a very weak upper-level baroclinic zone downstream over western Nevada (Fig. 2.6). The upper-level trough was accompanied by a significant relative humidity gradient, with high relative humidities ahead and lower relative humidities behind. The strongest and deepest upward vertical motion was located downstream of the trough axis and beneath the upper-level baroclinic zone.

By 0600 UTC, the primary 500-hPa vorticity maximum was positioned off the 11

a) 0000 UTC 12 Feb b) 0000 UTC 12 Feb

c) 0000 UTC 12 Feb

Figure 2.3. Upper-level and satellite analyses at 0000 UTC 12 Feb 2000. (a) RUC2 analysis of 500-hPa geopotential height (every 60 m) and absolute vorticity (x10-5 s-1, shaded following scale at bottom). (b) RUC2 analysis of 700-hPa temperature (every 2°C), wind [full (half) barbs denote 5 (2.5) m s-1], and relative humidity (%, shaded following scale at bottom). (c) Infrared satellite image. 12

) San Francisco (SFO) C ° (

p m e T ) a P h (

e r u s s e r P 00 01 02 03 04 05 06 07 08 09 10 11 12 Hour UTC 12 February 2000

) Sacramento (SAC) C ° (

p m e T ) a P h (

e r u s s e r P 00 01 02 03 04 05 06 07 08 09 10 11 12 Hour UTC 12 February 2000

) Reno (RNO) C ° (

p m e T ) a P h (

e r u s s e r P 00 01 02 03 04 05 06 07 08 09 10 11 12 13 14 15 16 17 Hour UTC 12 February 2000

Figure 2.4. Meteograms for (a) San Francisco (SFO), (b) Sacramento (SAC), and (c) Reno (RNO). Full (half) barbs denote 5 (2.5) m s-1. Arrows identify frontal passage in (a) and (b). See Fig. 2.1a for locations. 13

a) Richmond (RMD) L G A

s r e t e M

Hour UTC 12 February

b) Sacramento (SAC) L G A

s r e t e M

Hour UTC 12 February

Figure 2.5. Wind profiler data for 12 February 2000 from (a) Richmond (RMD) and (b) Sacramento (SAC). Full (half) barb denote 5 (2.5) m s-1. Meridional wind component plotted every 3 m s-1 as solid lines. Dashed lines denote wind shear accompanying frontal zone. See Fig. 2.1a for locations. 14

0300 UTC Cross Section

-3 +3 -6 -9 -12

-3 +3 -3 -6 -3

A S B

Figure 2.6. RUC2 cross section along line AB of Fig. 2.2b at 0300 UTC 12 Feb. Potential temperature (thin lines every 3K), vertical motion (ω, thick lines every 3x10-1 Pa s-1 with negative contours dashed), wind [full (half) barbs denote 5 (2.5) m s-1], and relative humidity (light, dark shading denote >70 and 90%, respectively). Upper-level trough axis denoted by heavy dashed line, leading edge of upper-level baroclinic zone denoted by heavy solid line, and surface front position denoted by arrow. Sierra Nevada indicated by S. 15 south-central California coast (Fig.2.7a), and a well-defined cloud shield was evident in satellite imagery (Fig. 2.7c). A large area of moderate to heavy precipitation was found over central California and the Sierra Nevada, while lighter precipitation extended northeastward from north-central Nevada towards southern Idaho (Fig. 2.8). A secondary

500-hPa vorticity maximum (Fig. 2.7a), associated with an upper-level jet streak (not shown), extended downstream across northeast Nevada. As this vorticity maximum approached northern Utah, mid- and upper-tropospheric relative humidities increased

(e.g., 700 hPa, Fig. 2.7b). At 700 hPa, there was little or no temperature advection over central California (Fig. 2.7b), however, wind profiler observations suggested (Figs. 2.5a and 2.5b) that winds were backing with height below 700 hPa, indicating cold advection below 700 hPa. Nearby radiosonde observations showed 2–3°C cooling within this layer from 0000 to 1200 UTC. The 1500-m low center was located over southwest Oregon (Fig.

2.8). The occluded front, which was now detached from the low center, was approaching the windward slopes of the Sierra Nevada (Fig. 2.8).

From 0600-1200 UTC, the surface front approached the Sierra Nevada and its thermodynamic and kinematic structure became less coherent as it split around the Sierra

Nevada, as observed during other events (e.g., Blazek 2000, his Fig. 4.16). Specifically, the eastward movement of the surface front was retarded by the central Sierra. The southern portion of the occluded front was deflected equatorward while the northern portion continued eastward. The front weakened dramatically by the time it moved into the lee of the Sierra. At Reno (RNO, see Fig. 2.1a for location), a station pressure minimum occurred around 0800 UTC with a 2°C temperature decrease between 0900–

1000 UTC, which was coincident with a trace of light rain from 0936–0946 UTC and 16

a) 0600 UTC 12 Feb b) 0600 UTC 12 Feb

c) 0600 UTC 12 Feb

Figure 2.7. Upper-level and satellite analyses at 0600 UTC 12 Feb 2000. (a) RUC2 analysis of 500-hPa geopotential height (every 60 m) and absolute vorticity (x10-5 s-1, shaded following scale at bottom). (b) RUC2 analysis of 700-hPa temperature (every 2°C), wind [full (half) barbs denote 5 (2.5) m s-1], and relative humidity (%, shaded following scale at bottom). (c) Infrared satellite image. 17

44 42 38 40

40 L 42 38 44 45 40 42 46 35

44 48 25

15 46 48 50 50 48

Figure 2.8. Surface and radar analyses at 0600 UTC 12 Feb 2000. NEXRAD radar mosaic (reflectivity scale at left) and manual 1500-m pressure analysis. Station reports include 1500-m pressure (tenths of hPa with leading 8 truncated), and wind [full (half) barbs denote 5 (2.5) m s-1]. 18 veering winds from southeasterly to southerly (Fig. 2.4c). Winds then decreased and shifted to westerly as station pressures began to rise steadily at 1200 UTC. Thus, there was a well-defined wind shift accompanied with a pressure rise at RNO that featured little or no temperature change. The temperature drop preceeded the wind shift and was coincident with light rain and presumably evaporative cooling. This evolution may be related to RNO being immediately downstream of the Sierra Nevada; perhaps the deepest cold air was blocked by the Sierra or never arrived owing to downslope warming, while a precipitation band associated with mid-level vertical motion was able to traverse the Sierra ahead of the large-scale pressure trough.

By 1200 UTC, the upper-level trough was located downstream of the Sierra

Nevada crest over western Nevada, and the upper-level baroclinic zone had weakened (cf.

Figs. 2.6 and 2.9). The upper-level baroclinic zone, mid-level cold advection, and deep tropospheric vertical motion were moving across Nevada and approaching northern Utah.

Precipitation across most of California had ended, and widespread precipitation was falling across much of the Intermountain region, including the IPEX target area (not shown).

A curious aspect of this event was the rapid movement of a surface trough across

Nevada. Between 1200–1800 UTC, a surface trough moved rapidly (20–25 m s-1) across

Nevada, and generally featured veering surface winds and little or no surface baroclinity.

For example, when the surface trough passed Elko (EKO, see Fig. 2.1a for location) around 1400 UTC, winds were calm, surface temperatures remained steady at 2°C, and light snow was falling (not shown). The rapid movement of the surface trough was better correlated with the speed of the upper-level trough axis rather than low-level advection. In 19

1200 UTC Cross Section

-3 +3 -3 +3 -6

-3

A S B

Figure 2.9. RUC2 cross section along line AB of Fig. 2.2b at 1200 UTC 12 Feb. Potential temperature (thin lines every 3K), vertical motion (ω, thick lines every 3x10-1 Pa s-1 with negative contours dashed), wind [full (half) barbs denote 5 (2.5) m s-1], and relative humidity (light, dark shading denote >70 and 90%, respectively). Upper-level trough axis denoted by heavy dashed line, leading edge of upper-level baroclinic zone denoted by heavy solid line, and surface front position denoted by arrow. Sierra Nevada indicated by S, and arrow denotes surface trough position. 20 fact, the surface trough was nearly coincident with the leading edge of the upper-level trough and the deepest upward vertical motion (Fig. 2.10). Thus, the pressure minimum associated with the surface trough was most likely a reflection of upper-level mass divergence and associated vertical motion immediately ahead of the upper-level trough, as observed in other events over the western United States (e.g., Hess and Wagner 1948;

Schultz and Doswell 2000).

At 1800 UTC, the primary 500-hPa vorticity maximum was positioned over southern Nevada and the 500-hPa trough axis had elongated, becoming negatively tilted

(Fig. 2.11a). Although the strongest 500-hPa vorticity was located over southern Nevada, the primary surface low pressure center was positioned well to the north, near the Idaho–

Oregon border (Fig. 2.12), and was associated with a weaker 500-hPa vorticity center near the Oregon-California border. The surface trough, along with widespread clouds and precipitation, was present ahead of and roughly parallel to the 500-hPa trough axis (Fig.

2.11c). At 700 hPa, higher relative humidity and weak cold advection were generally present over the Intermountain region (Fig. 2.11b).

As the surface trough moved across northern Utah between 2000–2200 UTC, it developed baroclinity. The baroclinity, however, was insufficient to call the trough a front, so the trough will be referred to as the baroclinic trough (Sanders 1999a). Ahead of the baroclinic trough, IPEX scientists were perplexed when mesoscale observations revealed a midlevel trough that arrived about 3 h ahead of the baroclinic trough. The midlevel trough was not accompanied by an abrupt change in temperature or moisture, but rather a small (1–3°C over 2 h) decrease in equivalent potential temperature with respect to ice

θ ( ei); see Figure 3 of Cox et al. (2005) for a mesoscale time-height section. The midlevel 21

1800 UTC Cross Section

-3 -3

-6 -3 +3 -6 -3 -6

A S B

Figure 2.10. RUC2 cross section along line AB of Fig. 2.2b at 1800 UTC 12 Feb. Potential temperature (thin lines every 3K), vertical motion (ω, thick lines every 3x10-1 Pa s-1 with negative contours dashed), wind [full (half) barbs denote 5 (2.5) m s-1], and relative humidity (light, dark shading denote >70 and 90%, respectively). Upper-level trough axis denoted by heavy dashed line, leading edge of upper-level baroclinic zone denoted by heavy solid line, and surface front position denoted by arrow. Sierra Nevada indicated by S, and arrow denotes surface trough position. 22

a) 1800 UTC 12 Feb b) 1800 UTC 12 Feb

c) 1800 UTC 12 Feb

Figure 2.11. Upper-level and satellite analyses at 1800 UTC 12 Feb 2000. (a) RUC2 analysis of 500-hPa geopotential height (every 60 m) and absolute vorticity (x10-5 s-1, shaded following scale at bottom). (b) RUC2 analysis of 700-hPa temperature (every 2°C), wind [full (half) barbs denote 5 (2.5) m s-1], and relative humidity (%, shaded following scale at bottom). (c) Infrared satellite image. 23

38 38

44 42 40 36 40 42 L 38 L 36

38 44 40 45 42 35 46 44

46 25 48 48 15 48 48

Figure 2.12. Surface and radar analyses at 1800 UTC 12 Feb 2000. NEXRAD radar mosaic (reflectivity scale at left) and manual 1500-m pressure analysis. Station reports include 1500-m pressure (tenths of hPa with leading 8 truncated), and wind [full (half) barbs denote 5 (2.5) m s-1]. 24 trough appeared to develop along a ridge of high potential vorticity or a high potential vorticity banner that developed over the southern Sierra Nevada. Schär et al. (2003) and

Grubisic (2004) describe the development of similar banners over the Alps and Dinaric

Alps. As shown in Figure 2.13, low and high potential vorticity banners extended downstream from the Sierra at 0600 UTC (Fig. 2.13a). By 1200 UTC, the southwesterly flow had advected the high potential vorticity banner northeastward into northeast Nevada and northwest Utah (Fig. 2.13b). Inspection of cross-sections through the high potential vorticity banner (not shown) showed that the northern portion of the quasi-stationary high potential vorticity banner became mobile when the upper-level trough approached. Thus, it appeared that the northern portion of the high potential vorticity banner became mobile as the upper-level trough approached and thereafter moved over northern Utah (Fig.

2.13c) as the midlevel trough.

By 0000 UTC 13 Feb, the 500-hPa trough had weakened considerably with the primary vorticity maximum in northeast Arizona (Fig. 2.14a). A broad region of low

1500-m pressure extended from the Idaho-Montana border southeastward to Wyoming with weak ridging developing over Utah (Fig. 2.14c). Weak 700-hPa cold advection (not shown), clouds and precipitation were present over much of northern Utah (Figs. 2.14b and 2.14c), but the precipitation event was close to ending.

Figures 2.15a and 2.15b summarize the surface and vertical structure evolution of this event. An occluded midlatitude cyclone made landfall around 0000 UTC 12 Feb

2000. As the system moved inland, the low pressure area became increasingly broad, separated from the surface occluded front, and moved northeastward while the surface occluded front advanced towards the Sierra Nevada. The surface front sloped rearward 25

a) 0600 UTC 12 Feb b) 1200 UTC 12 Feb

= 10 m s-1 = 10 m s-1

c) 1800 UTC 12 Feb

= 10 m s-1

Figure 2.13. ETA Data Assimilation System (EDAS) 700-hPa potential vorticity (every 1x10-7 m2 s-1 K kg-1) and wind (vector scale at lower right). High (low) potential vorticity banners denoted by thick solid (dashed) line. (a) 0600 UTC 12 Feb 2000. (b) 1200 UTC. (c) 1800 UTC. 26

a) 0000 UTC 12 Feb b) 0000 UTC 12 Feb

c) 0000 UTC 12 Feb

44 42 42 40 38 L 36 L 36

45 44 38 40 35 46 42 25 44 15 46

Figure 2.14. Upper-level, satellite, surface, and radar analyses at 0000 UTC 13 Feb 2000. (a) RUC2 analysis of 500-hPa geopotential height (every 60 m) and absolute vorticity (x10- 5 s-1, shaded following scale at bottom). (b) Infrared satellite image. (c) NEXRAD radar mosaic (reflectivity scale at left) and manual 1500-m pressure analysis. Station reports include 1500-m pressure (tenths of hPa with leading 8 truncated), and wind [as in (b)]. 27

a) Synoptic Evolution 00 – 21 UTC

15 18 21 12 L L L 09 L L L . B L 18 21 .L 15 A 12

00 03 06 09 12

2b0)0 Vertical Structure Evolution 00 – 21 UTC 250 00 03 06 09 12 15 18 21 300

400 U

500 18 M 21 700 15 18 00 03 06 12 850 S . . 1000 . A SFO RNO SLC B

Figure 2.15. Summary of surface and vertical structure evolution of IOP3. (a) Surface features every 3 h (conventional symbols) from 0000–2100 UTC 12 Feb 2000. Labels represent hour UTC. (b) Vertical trough structure (roughly along line AB of Fig. 2b) every 3 h. U denotes upper level trough, S surface trough, and M midlevel trough. 28 with height, but the associated upper-level trough was forward sloping. As the system approached the Sierra Nevada, the surface front weakened and split into northern and southern sections, while the upper-level trough and baroclinic zone continued moving eastward. Essentially, the terrain acted to “strip” the low-level occluded front from the upper-level trough. The upper-level trough then approached the Sierra lee trough, and the two features moved rapidly downstream across northern Nevada. The rapid movement of the surface trough suggested close coupling with the mass divergence accompanying the upper-level trough. Over northern Utah, these features were preceded by a midlevel trough that formed along a high potential vorticity banner that developed over the Sierra Nevada.

Mesoscale storm structure over northern Utah

Prebaroclinic surface trough period

Southerly to southwesterly low- and mid-level flow developed over the IPEX target area as the upper-level and surface troughs approached northern Utah. Between

0900–1800 UTC, there was little change to the large-scale kinematic environment that featured southerly low-level flow with an 8–12 m s-1 wind speed maximum at 800 hPa

(Fig. 2.16), which was just below mid-mountain (hereafter defined as 775 hPa or ~2200 m

AMSL). Above this level, winds veered to southwesterly, resulting in a significant cross- barrier flow component at crest level (~700 hPa). In the lowlands, winds were mainly southerly within the valleys upstream of the Wasatch Mountains (Figs. 2.17a and 2.17b).

This southerly low-level along-barrier flow was oriented normal to the 1500-m isobars (cf.

Figs. 2.12 and 2.17b) and was highly ageostrophic. Southwesterly flow over and upstream of the Great Salt Lake became confluent with the along-barrier flow near the eastern shore 29

SLC Time-height section ) a P h (

e r u s s e r P

Day/Hour (UTC)

Figure 2.16. Salt Lake City (SLC) time-height section based on 3-h soundings (thick barbs) θ with RUC2 winds (thin barbs) overlaid. e (solid lines, every 3 K), relative humidity (light and dark shading denote >70 and 90%, respectively), and wind [full (half) barbs denote 5 and (2.5) m s-l]. Dashed lines denote midlevel and surface troughs. 30

a) 1200 UTC 12 Feb

40 km

b) 1800 UTC 12 Feb

Figure 2.17. Northern Utah manual streamline analysis at (a) 1200 and (b) 1800 UTC 12 Feb 2000. Station data and streamlines for elevations between 1280–2200 m AMSL (roughly at and below mid mountain). Full (half) barbs denote 5 (2.5) m s-1. 31 of the Great Salt Lake. ADAS analyses showed that this region of confluent flow was convergent (Fig. 2.18), and it will be referred to as a windward convergence zone.

Precipitation began to fall at mountain locations by ~0900 UTC as the large-scale upward vertical motion and moisture ahead of the upper-level trough overspread northern

Utah (Fig. 2.19a). Precipitation onset was delayed up to 3 h at many lowland sites (e.g.,

SNX and SLC, Fig. 2.19b). This delay appeared to be the result of low-level sublimation and evaporation as suggested by the 0600 and 1200 UTC soundings from SLC which showed a descending cloud base with the 775–600 hPa temperature profile approaching the moist adiabat (Fig. 2.20). By 1200 UTC, vertical profiles of potential temperature (θ)

θ and equivalent potential temperature with respect to ice ( ei) showed that, with the exception of a shallow layer near 800 hPa, θ increased with height from the surface to

θ about 500 hPa, but ei was nearly constant with height up to 575 hPa, indicating approximately neutral moist stability (Fig. 2.21a). By 1800 UTC, this neutral stability layer extended to about 500 hPa at both near-barrier sounding locations (SLC and NSSL5,

Figs. 2.21a,b).

Precipitation during this period was generally widespread with some weak convective elements; 0.5° base reflectivities were no greater than 35dBZ (e.g. Fig. 2.22a).

Animation of radar reflectivity showed quasi-stationary precipitation features, which suggested that precipitation processes were strongly influenced by the regional orography.

For example, radar imagery at 1900 UTC 12 Feb showed an area of locally high reflectivity upstream of the Wasatch Mountains near OGD, whereas lower reflectivities

(often below 5 dBZ) were found over the Salt Lake Valley (Fig. 2.22b). Although beam blockage and other radar limitations prevented analysis of the KMTX reflectivity structure 32

ADAS SFC Divergence

-0.25

-3.0

-6.0

Figure 2.18. ARPS Data Assimilation System (ADAS) surface analysis at 1800 UTC 12 Feb 2000, surface divergence (light and dark shading denote less than -.25x10-4 s-1 and - 3.0x10-4 s-1, respectively) and wind [full (half) barbs denote 5 (2.5) m s-1]. 33

(a) Mountain Precipitation Rates 5 ) 1

- CLN h 4 SNI m m ( 3 e t a R

2 p i c e

r 1 P

06 09 12 15 18 21 00 03 06 Hour UTC (0600 12 Feb – 0600 13 Feb)

(b) Lowland Precipitation Rates 5 ) 1

- SLC

h 4 OGD m SNX m (

3 e t a

R 2

p i c e

r 1 P

06 09 12 15 18 21 00 03 06 Hour UTC (0600 12 Feb – 0600 13 Feb)

m 40 m ( SNI n o

i 30 t a

t OGD i

p 20

i SNX c e r 10

P SLC

06 09 12 15 18 21 00 03 06 Hour UTC (0600 12 Feb – 0600 13 Feb)

Figure 2.19. Precipitation rates and accumulated precipitation at selected locations. (a) Mountain precipitation rates (mm h-1). (b) Lowland precipitation rates (mm h-1). (c) Accu- mulated precipitation (mm). See Fig. 2.1b for locations. 34

Salt Lake City (SLC)

0600 UTC 1200 UTC ) a P h (

e r u s s e r P

Temp (°C)

Figure 2.20. Skew T–logp diagram [temperature (solid) and dewpoint (dashed)] at SLC for 0600 UTC (black) and 1200 UTC (gray) 12 Feb 2000. 35

a) SLC b) NSSL5 1200 1800 1800 2100 ) ) a a P P h h ( (

e e r r u u s s s s e e r r P P θ θ θ θ ei ei

Temp (K) Temp (K)

Figure 2.21. Vertical profiles of potential (left) and equivalent potential temperature with respect to ice (right) at (a) SLC and (b) NSSL5. See Fig. 2.1b for locations. 36

(a) 1200 UTC 12 Feb (b) 1900 UTC 12 Feb

KMTX KMTX

OGD OGD

KMTX

OGD

Figure 2.22. Base (0.5°) radar reflectivity from Promontory Point (KMTX) at (a) 1200 UTC 12 Feb, (b) 1900 UTC 12 Feb, and (c) 2100 UTC 12 Feb 2000. Reflectivity scale (dBZ) on right. Dashed line in (c) denotes position of baroclinic trough. See Fig. 2.1b for location of Salt Lake Valley. 37 directly over the Wasatch Mountains, a narrow reflectivity maximum was observed directly over the barrier crest by the NOAA WP-3D tail radar (see Cox et al. 2005).

Precipitation gauge observations from five northern Utah locations, including two mountain sites (CLN and SNI) and three lowland sites (OGD, SNX, and SLC, see Fig.

2.1b for locations), further illustrate the orographic influences on the precipitation distribution. Although blocking of the large-scale southwesterly flow by the Wasatch and other ranges resulted in southerly along-barrier flow at low levels (e.g., Fig. 2.14, Cox et al. 2005), sufficient (5–10 m s-1) cross-barrier flow at mid-mountain and crest level (Fig.

2.23, NSSL 4 traces) produced a substantial orographic enhancement. By 2100 UTC, SNI and CLN reported 24 mm and 30 mm Snow Water Equivalent (SWE), respectively (Fig.

2.19c), with fairly constant precipitation rates reaching 4 mm h-1 (Fig. 2.19b). At SNI, precipitation was roughly double that at nearby low-level OGD, while that at CLN was about 30 times the precipitation at SLC.

In the lowlands immediately upstream of the Wasatch Mountains, significant precipitation contrasts were observed between OGD, SNX, and SLC. Only 1 mm of rain fell at SLC (surface temperatures were 4–5°C) through 2100 UTC, whereas Ogden (OGD) and Antelope Island (SNX), which are located 30–40 km north of SLC and upstream of the northern Wasatch Mountains, recorded much more with 16 m and 11 mm, respectively

(Fig. 2.19b). At least two factors contributed to this lowland precipitation contrast. One, subcloud evaporation and rain shadowing over the Salt Lake Valley; two, windward convergence and precipitation enhancement in the vicinity of OGD.

Precipitation at OGD and SNX was coincident with a saturated low-level environment as indicated on the 1800 UTC NSSL5 sounding (Fig. 2.24). In contrast, the 38

Cross Barrier Flow ) a P h (

e r u s s e

r Mid-mountain P

Wind Speed (m s-1)

Figure 2.23. Cross-barrier flow magnitude (m s-1) at NSSL4 and NSSL5 for 1800 and 2100 UTC 12 Feb 2000. See Fig. 2.1b for locations. 39

SLC 1800 UTC NSSL5 1800 UTC ) a P h (

e r u s s e r P

Temp (°C)

Figure 2.24. Skew T–log p diagram [temperature (solid) and dewpoint (dashed)] at SLC and NSSL5 for 1800 UTC 12 Feb 2000. 40

1800 UTC SLC sounding revealed a much a drier low-level environment that likely reduced precipitation amounts at SLC. The drier low-level environment at SLC, as well as across much of the Salt Lake Valley, appeared to be partially related to downslope flow from the southern terminus of the Salt Lake Valley (e.g., Fig. 2.17b), which features the

Traverse Mountains that reach more than 2000 m, and/or subsidence produced by rainshadowing to the lee of the Oquirrh Mountains. The presence of a preexisting dry environment (Fig. 2.20) also played a role.

Heavier precipitation was observed at OGD and SNX (Figs. 2.19b,c), which were not strongly influenced by rainshadowing from upstream ranges. These sites were affected also by the windward convergence zone, which enhanced precipitation over the lowlands immediately upstream of the Wasatch Mountains. As illustrated by Figures 2.17 and 2.18, convergence was observed upstream of the Wasatch Mountains where southwesterly flow merged with southerly along-barrier flow. The most significant precipitation enhancement occurred immediately downstream of the windward convergence zone, as illustrated by the region of high radar reflectivity extending approximately 25 km upstream of the

Wasatch Crest near Ogden at 1900 UTC (Fig. 2.22b). Precipitation rates at OGD reached

2.5 mm SWE h-1, which was at times surprisingly similar to the nearby mountain precipitation rate at SNI (cf. Figs. 2.19a,b). The windward convergence zone and associated precipitation region advanced toward the barrier with the approach of the baroclinic trough, as described in detail by Cox et al. (2005) and Colle et al. (2005).

The storm environment was modified only slightly as the midlevel trough, which was marked by veering winds between 700 and 500 hPa, moved across northern Utah between 1800–2100 UTC (Fig. 2.16). P-3 flight-level data between 700 hPa and 500 hPa 41 revealed that there were no abrupt temperature or moisture gradients present across the

θ midlevel trough, but rather a small (1–3°C over 2 h) decrease of mid- and upper-level ei.

No organized precipitation structures were coincident with the midlevel trough, although the confluent flow over the eastern half of the Great Salt Lake advanced toward the

Wasatch Mountains during and following its passage (see Cox et al. 2005). During this period, lowland precipitation rates increased slightly or remained steady, but at most mountain stations (1900 m ASL and above), precipitation rates remained steady or decreased slightly.

Postbaroclinic surface trough period

At 2100 UTC, the baroclinic trough was collocated with a weak east-west band of precipitation over the southern end of the Great Salt Lake (Fig. 2.22c). The baroclinic trough intensified as it moved across northern Utah, and, at SLC, was accompanied by a

2°C temperature drop, surface pressure minimum, and wind shift around 2130 UTC (Fig.

2.25). The strongest temperature changes with the baroclinic trough over northern Utah were coincident with moderate or heavy precipitation; thus, diabatic processes appeared to produce much of the baroclinity accompanying the baroclinic trough as it moved into northern Utah, similar to that described by Schultz and Trapp (2003).

The wind changes accompanying the baroclinic trough were complex across northern Utah. At DPG and the immediately adjacent lowlands, surface winds veered from southerly to westerly with baroclinic trough passage (cf. Fig. 2.17b, Fig. 2.26).

Across the Salt Lake and Tooele Valleys, however, the surface wind shift was weak, or short-lived (Fig. 2.26), with a few exceptions (e.g., SLC, Fig. 2.25). Inspection of 42

Salt Lake City (SLC) ) C ° (

p m e T ) a P h (

e r u s s e r P

07 08 09 10 11 12 13 14 15 16 17 18 19 20 21 22 23 00 01 02 03 04 Hour UTC 12,13 February 2000

Figure 2.25. Meteogram for Salt Lake City (SLC). Full (half) barbs denote 5 (2.5) m s-1. Arrows denote the passage of the baroclinic trough (dashed) and convective line (solid). See Fig. 2.1a for location. 43

2200 UTC 12 Feb

D.PG

Figure 2.26. Northern Utah manual streamline analysis at 2200 UTC 12 Feb 2000. Station data and streamlines for elevations between 1280–2200 m (roughly at and below mid mountain). Full (half) barbs denote 5 (2.5) m s-1. 44

MesoWest stations around 800 m above valley level, however, revealed that a consistent wind shift was associated with the baroclinic trough just above valley level. This elevated wind shift was continuous, moved east-southeast (confirmed with KMTX velocity data, not shown), and was also coincident with the largest temperature falls, which were 2–4°C between 800–600 hPa, with the most significant drop being 4°C at 650 hPa over 2 h. Thus, the most notable temperature and wind direction changes associated with the baroclinic trough were generally elevated, between 800 and 600 hPa, and not at valley level.

Precipitation behind the baroclinic trough grew more convective as the low-levels became increasingly unstable after baroclinic trough passage. Much of the lowland precipitation was associated with two transient mesoscale precipitation areas, one a weakly organized area associated with the baroclinic trough (Fig. 2.22c), and the second a convective line that passed SLC and produced a 4°C temperature drop about 3 h later (Fig.

2.25). In fact, turbulence accompanying this feature was strong enough to abort the P-3’s first landing attempt. The convective line developed in the airmass behind the baroclinic

θ trough, which featured ei decreasing with height up to 675 hPa (Fig. 2.27a). As the baroclinic trough moved southeast, lowland precipitation rates spiked at different times:

OGD around 2100 UTC, SNX at 2200 UTC and SLC at 0000 UTC 13 February (Fig.

2.19b). Precipitation rates at mountain sites, however, peaked after the lowland locations, reaching 4 mm hr-1 at both SNI and CLN between 0000–0100 UTC 13 February (Fig.

2.19b); these increased mountain precipitation rates were associated with destabilization and increased cross-barrier flow behind the baroclinic trough (Fig. 2.27b), but may be an overestimate due to snow buildup on gauge walls falling into the bottom of the gauge, as discussed in Cox et al. (2005). Precipitation enhancement that previously accompanied the 45

a) NSSL4 2100 0300 ) a P h (

e r u s s e r P θ θ ei

Temp (K)

b) Cross Barrier Flow ) a P h (

e r u s s Mid-mountain e r P

Wind Speed (m s-1)

Figure 2.27. Stability and cross barrier flow. (a) Vertical profiles of potential and equivalent potential temperature with respect to ice at NSSL4 at 2100 and 0300 UTC 13 Feb. (b) NSSL5 cross barrier flow at 2100 and 0300 UTC 13 Feb. 46 windward convergence zone over the lowlands upstream of the northern Wasatch (e.g.,

OGD, SNX) ended shortly after passage of the baroclinic trough.

Across the Salt Lake and Tooele Valleys, much of total storm precipitation fell in the postbaroclinic trough environment under northwesterly flow (e.g., SLC 83%) when low-level moisture and instability were maximized. Precipitation at mountain locations, however, was generally evenly distributed throughout the event with CLN and SNI receiving about 40% of their total SWE in the postbaroclinic trough environment.

Precipitation at both mountain and valley locations ended by 0600 UTC 13 February, as a drier environment and subsidence behind the upper-level trough developed.

Analysis of orographic precipitation

Storm-total SWE ratios featured three main spatial precipitation patterns. One, significant precipitation enhancement was observed in the higher elevations. Second, precipitation suppression was observed both to the lee of the Wasatch and Oquirrhs

Mountains, including the Salt Lake and Tooele Valleys. Third, local precipitation enhancement was observed in the lowlands upstream of the northern Wasatch (Fig. 2.28a).

Precipitation generally increased with altitude in IOP3 (Fig. 2.28b), which resembled the nearly linear climatological distribution (Fig. 2.28c). However, unlike the climatological precipitation distribution, IOP3 featured a somewhat bimodal distribution above and below 2100 m or at lowland and mountain locations. This bimodal precipitation distribution suggested that significant variations from climatological (30-Yr annual average) precipitation versus altitude were present during IOP3. Some of these departures from climatology could be explained by two observed spatial precipitation patterns: one, 47

20 a) 20 20 30 40 10

SNI OGD

SNX

10 SLC

40 20 CLN 20

20 10 10

b) IOP3 Observed Precipitation vs. Elevation Mountain ) CLN m (

o SNI i t a v

e Lowland l OGD E

SLC SNX

Liquid Water (cm)

c) Climatological Precipitation vs. Elevation P = 0.0523*A–30.406 CLN ) m (

n o

i SNI t a v e l E OGD SLC,SNX Liquid Water (cm)

Figure 2.28. IOP3 precipitation characteristics. (a) Total accumulated liquid precipitation during IOP3 (every 10 mm) with observations annotated; selected contour lines omitted for clarity. (b) Observed liquid precipitation (cm) versus elevation over the IPEX target area. (c) Climatological (30-Yr annual average) liquid precipitation (cm) versus altitude over the IPEX target area. 48 precipitation suppression in the lee of the Oquirrh Mountains (e.g., SLC), and two, precipitation enhancement in the lowlands upstream of the northern Wasatch (e.g., SNX,

OGD). For example, SNX and SLC, which have a similar elevation (1280 m), orientation relative to the Wasatch, and climatological annual precipitation (SNX, 45.6 cm, SLC, 42 cm), featured nearly a factor of three SWE difference (SNX, 15 mm, SLC, 6 mm).

Mountain sites (e.g., SNI, 2256 m, 49 mm) over the northern Wasatch also featured greater precipitation for their respective elevation compared to higher elevation stations to the south (e.g., CLN, 2945 m, 51 mm).

Precipitation anomaly maps were constructed to further compare climatological and IOP3 spatial precipitation patterns. Since annually averaged precipitation correlates well with altitude over the IPEX target region (Fig. 2.28c), a linear regression equation was developed to describe the climatological precipitation altitude relationship and is given by P = 0.0523*A–30.406, where P is the precipitation in cm and A is the altitude in meters. Analysis of the residuals or anomalies from this predicted value for climatological precipitation are shown in Figure 2.29. Positive anomalies are relatively wet areas for their respective elevations, and include the northern Wasatch Mountains, lowlands immediately upstream of the Wasatch, eastern two thirds of the Great Salt Lake, and the northern and eastern Tooele and Salt Lake Valleys (Fig. 2.29). Negative anomalies, or areas that are relatively dry for their respective elevations, are present over the Great Salt

Lake desert, southern Oquirrh Mountains, and downstream of the Wasatch crest (Fig.

2.29).

In order to illustrate departures from the climatological precipitation-altitude relationship during IOP3, the slope of the climatological linear fit was fixed to the IOP3 49

0 +10 +10 0 -10 -20

+20

+30

-20

-20 Units: cm 0 -10

-20 +10 -10 -20 -10 0 +10

Figure 2.29. Climatological (30-Yr annual average) precipitation anomaly (cm) relative to linear fit. Selected contour lines omitted for clarity. 50 mean precipitation, with residuals from this line presented in Figure 2.30. IOP3 generally featured similar precipitation anomaly patterns to climatology, including positive (wet) anomalies over and upstream of the northern Wasatch and negative (dry) anomalies east of

Wasatch (Fig. 2.30). However, there were differences between IOP3 and climatological precipitation anomaly patterns over the Salt Lake and Tooele Valleys, where the precipitation amounts were smaller than one would expect climatologically.

Concluding Discussion

Storm evolution over the western United States

Very few studies have examined the evolution of a winter storm from the Pacific

Coast to the interior Intermountain West. During IPEX IOP3, a landfalling occluded front weakened and deformed as it approached the Sierra Nevada, similar to that observed previously near the Sierra Nevada by Hoffman (1995), Reynolds and Kuciasuskas (1988), and Blazek (2000), as well as upstream of the Appalachians (Schumacher et al. 1996).

Ultimately, there was little evidence that the baroclinity associated with the surface based occlusion was able to penetrate into the lowlands to the lee of the Sierra Nevada as the upper-level trough continued to move downstream. Thus, the high topography appeared to

"strip" or remove the storm system of its low-level baroclinic structure as it traversed the

Sierra Nevada. Destruction of the low-level baroclinity may have also been aided by adiabatic warming to the lee of the Sierra Nevada, as suggested by Hobbs et al. (1996) for frontal evolution in the lee of the Rocky Mountains.

To the lee of the Sierra, the upper-level trough interacted with the lee trough, which became mobile and moved rapidly downstream as the upper-level trough axis 51

+ 5 0 - 5 - 10 +10

+ 5

+15

0 +15

+ 5 0 - 10 Units: mm - 5 - 5 - 10 - 10 - 5 0

Figure 2.30. IOP3 observed precipitation anomaly (mm) relative to the slope of a linear fit to climatological precipitation versus altitude given in Fig. 2.25c. Selected contour lines omitted for clarity. 52 approached. The downstream movement of the surface trough appeared to be a reflection of mass-divergence associated with the upper-level trough, consistent with early work by

Hess and Wagner (1948) and more recently by Schultz and Doswell (2000).

The scenario described above suggests that the low-level occluded front and associated surface pressure trough did not move continuously across the Sierra Nevada.

Instead, the occluded front was blocked by the Sierra Nevada while the upper-level trough moved downstream and interacted with the lee trough. This evolution contrasts with that of many idealized simulations of fronts traversing topography (see Blumen 1992 and

Egger and Hoinka 1992 for reviews), which typically assume smooth slopes and modest relief so that the front is able to surmount the mountain barrier. It is, however, roughly consistent with the discontinuous low-level evolution described by Dickinson and Knight

(1999) who found that for high mountains, the low-level front is blocked, while the associated upper-level potential vorticity maximum moves downstream and eventually couples with the leeward low-level cyclonic vorticity maximum.

Another curious aspect of IOP3 is the forward-sloping structure of the upper-level trough. This structure resembles that of damping baroclinic waves (e.g., Spencer et al.

1996) and is consistent with the fact that the cyclone was occluded and weakening as it began to interact with the terrain of the western United States.

As the upper-level trough approached northern Utah, it interacted with a high potential vorticity banner that formed downstream of the southern Sierra Nevada. This interaction produced the complex kinematic structure that perplexed IPEX scientists during IOP3 (Schultz et al. 2002) and occurred as follows. Southwesterly flow ahead of the approaching storm system produced high and low potential vorticity banners 53 downstream of the southern and northern Sierra Nevada, respectively. As the mobile upper-level trough approached the quasi-stationary high potential vorticity banner over northeastern Nevada, the northern portion of the high potential vorticity banner became mobile and thereafter moved downstream. As the system reached northern Utah, the midlevel trough was observed. These banners were qualitatively similar to potential vorticity banners that form downstream of the European Alps (e.g., Schär et al. 2003,

Grubisic 2004). Thus, a complex interaction between the terrain and upper-level trough was responsible for the mid-level trough over northern Utah. This suggests that potential vorticity banners that form over the Sierra Nevada may complicate the kinematic structure of winter storms over Nevada and Utah. In this case, however, the midlevel trough had little effect on precipitation processes.

During IOP3, the upper-level trough was able to move relatively unimpeded across the Sierra Nevada and Great Basin ranges while the low-level occluded front was fully blocked, which resulted in discontinuous low-level storm evolution. These observations could be broadly applied to similar events that feature a landfalling mature midlatitude cylone in California and move across the Intermountain region as the accompanying baroclinic wave weakens. Namely, low-level structures of midlatitude cyclones across the

Intermountain region are strongly dependent upon interactions with the underlying terrain, while upper-level features move more or less unimpeded by the underlying terrain.

Precipitation distribution over northern Utah

In general, there is a strong correlation between climatological precipitation and altitude across the Intermountain West. As such, precipitation-altitude relationships are 54 used frequently by meteorologists and hydrologists to predict event or annual precipitation

(e.g., Hevesi et al. 1992). Over the IPEX target region, climatological precipitation increases linearly with height (e.g., Alter 1919; Daly et al. 1994; Fig. 2.25c) with a linear correlation coefficient of 0.70. However, large departures from this relationship have been observed for total storm precipitation (e.g., Horel and Gibson 1994; Cheng 2001; Shafer et al. 2002) and within individual storms (e.g., Steenburgh 2003). In addition, as illustrated by Fig. 2.26a and Peck and Brown (1962), the northern Wasatch Mountains and their immediate upstream lowlands, the eastern two-thirds of the Great Salt Lake, and the northern and eastern Tooele and Salt Lake Valleys are anomalously wet for their elevations, whereas the Great Salt Lake Desert and areas east of the Wasatch Mountains are anomalously dry. Therefore, in addition to a pronounced precipitation-altitude relationship, there are important mesoscale patterns evident in the regional precipitation climatology.

Precipitation during IOP3 occurred under moist southwesterly flow with near moist-neutral stability. Precipitation was dominated by local terrain processes that enhanced/suppressed precipitation at various locations across northern Utah. These terrain-induced processes included not only direct orographic ascent where more precipitation fell in higher terrain, but also enhanced/suppressed precipitation in nearby lowland areas. Three major precipitation patterns were observed: (1) precipitation enhancement with increased altitude, (2) precipitation suppression in the lee of the

Stansbury and Oquirrh Mountains (i.e., the Tooele and Salt Lake Valleys), and (3) precipitation enhancement produced by a windward convergence zone over the lowlands upstream of the northern Wasatch Mountains. 55

These patterns featured important similarities and differences compared to climatology. Precipitation enhancement in the northern Wasatch Mountains and upstream lowlands was consistent with this region being climatologically wet for its elevation. In contrast, IOP3 was dry relative to climatology over the Salt Lake and Tooele Valleys.

These characteristics are consistent with southwesterly flow events based on synoptic experience.

Analyses of departures from climatological precipitation-altitude relationships should be employed to better understand event, seasonal, and long-term precipitation distributions in regions of complex terrain. The construction of precipitation anomaly maps for varying flow directions and stability may be a useful predictive tool since systematic precipitation anomalies can be more easily identified and may help expose the underlying processes responsible for such anomalies. Such anomaly maps may also be useful for identifying important terrain-dominated mesoscale precipitation patterns and bias-correcting numerical model forecasts. These anomaly maps would likely improve upon techniques that rely primarily on the climatological precipitation-altitude relationship.

This study illustrates the importance of the Intermountain West’s topography on synoptic storm evolution and precipitation. Additional research is needed to better understand the passage of baroclinic systems across the Sierra Nevada, including how the resulting terrain-modified structure affects the development and evolution of precipitation over the downstream Intermountain West. CHAPTER 3

INTERMOUNTAIN COLD FRONTS

Introduction

Although the Sierra Nevada may act to weaken approaching Pacific cold and occluded fronts, intense, rapidly developing cold fronts produce a significant threat to lives and property over northern Utah and the adjoining Intermountain West several times a year. High winds, which may occur in the pre or postfrontal environment, can cause major wildfire runs, roads closures due to poor visibility associated with blowing dust, power outages, and other property damage. One strong cold front that moved across northern Utah on 5 June 1995 produced wind gusts in excess of 40 m s-1 and $15 million of damage (NOAA, 1995, p. 356). The cold front accompanying the 2002 Tax Day Storm produced a temperature fall of 19°C in 2 h in the Salt Lake City metropolitan area with wind gusts of more than 30 m s-1 across northern Utah (Shafer 2004). Approximately 50% of the strong cold fronts examined in the present study produced a wind gust that sur- passed the National Weather Service (NWS) wind advisory level (greater than or equal to

20 m s-1). As the population continues to grow rapidly over the Intermountain region

(Nevada and Utah are the fastest and fourth fastest growing states in the United States), so too will the socioeconomic impact related to strong cold fronts. 57

Observational research of strong surface cold fronts over the western United States has been limited primarily to landfalling Pacific fronts (e.g., Elliott 1958; Hobbs et al.

1980; Hobbs and Persson 1982; Reynolds and Kuciauskas 1988; Braun et al. 1997; Colle et al. 1999; Yu and Smull 2000; Chien et al. 2001) and cold fronts along the eastern slopes of the Rocky Mountains (e.g., Shapiro et al. 1985; Marwitz and Toth 1993; Neiman and

Wakimoto 1999; Darby et al. 1999; Neiman et al. 2001). By comparison, the interior western United States between the Sierra Nevada and Continental Divide (i.e., the Intermoun- tain West) has received little attention (e.g., Long et al. 1990; Schultz and Trapp 2003;

Steenburgh 2003).

There is evidence for significant structural modification of colds fronts and their accompanying baroclinic systems as they move across the western United States. For example, near coastal locations, Pacific cold fronts feature the majority of their associated precipitation along and ahead of the surface cold front (e.g., Reynolds and Kuciauskas

1988), whereas Intermountain cold fronts feature mainly postfrontal precipitation (e.g.,

Long et al. 1990; Schultz and Trapp 2003). Furthermore, observational studies of strong

Pacific coast cold fronts show relatively weak surface temperature contrasts, with temperature drops of 2–4°C over 1–4 h (e.g., Braun et al. 1997; Colle et al. 1999; Chien et al.

2001). In contrast, Intermountain cold fronts can feature much stronger temperature falls

(e.g., 7°C in 1 h, Steenburgh and Blazek 2001; 8°C in 1 h, Schultz and Trapp 2003). The most intense temperature fall examined in the present study occurred during the Tax Day

Storm, which featured a dramatic 16°C drop over 1 h.

The Tax Day Storm also featured intense surface frontogenesis, approaching 10°C

2 h-1 100 km-1 over northern Utah. Such rapid frontogenesis can’t be explained solely by 58 large-scale geostrophic deformation (e.g., Eliassen 1959; Stone 1966). Rather, it is hypothesized that diabatic processes, which are enhanced by the dry convective boundary layer, result in highly ageostrophic responses that contribute to the rapid development of strong Intermountain cold fronts.

Wallace Stegner once said, “Aridity, more than anything else, gives the western landscape its character...” (Stegner 1992, p. 46). This aridity results in a high Bowen ratio or large values of sensible heating compared to latent heating. This produces large diurnal temperature ranges that may act to enhance diabatic frontogenesis. The diabatic processes leading to frontogenesis/lysis will be explored within this paper, including differential sensible heating (e.g., Koch et al. 1997; Segal et al. 2004). Diabatic cooling due to evaporation will also be examined since the aridity results in a higher magnitude of low-level evaporative cooling, which has been shown to affect the development and evolution of cold fronts (e.g., Huang and Emanuel 1991; Katzfey and Ryan 1997; Schultz and Trapp

2003). The aridity of the Intermountain region also results in the existence of a deep daytime convective boundary layer. Previous studies involving cold fronts and their interactions with a dry convective boundary layer are limited mainly to Australia (e.g., Reeder

1986; Smith et al. 1995) and show that the evolution of the boundary layer has a strong influence on surface cold fronts. Most of these studies, however, describe preexisting cold fronts interacting with the boundary layer, and not the genesis of a cold front within a dry convective boundary layer, as frequently occurs over the Intermountain West. Numerical studies by Reeder and Tory (2005) and Tory and Reeder (2005) represent one of the first studies to directly explore the effects of a convective boundary layer on the development of surface cold fronts. They found that the daytime convective boundary layer, in the pres- 59 ence of homogeneous surface heating, has an overall frontolytic effect, while the nighttime boundary layer is highly frontogenetic.

This chapter describes the spatial and temporal distribution (seasonal and diurnal) characteristics of strong cold fronts over the western United States, and then examines the processes responsible for their development over the Intermountain West. Results from this study are relevant to operational forecasting since one of the major challenges facing operational meteorology is predicting the timing, location, and strength of surface cold fronts.

In the next section we describe the data and methods used for the study. This is fol- lowed by an analysis of the general characteristics of strong cold fronts over the western

United States and a detailed examination of the processes contributing to their development over the Intermountain West. The chapter concludes with a discussion of results, including a comparison with front studies from other arid regions.

Data and Methods

Data sources

Hourly surface observations were obtained from the National Climatic Data Cen- ter (NCDC) in the DSI-3505 format. This data set includes both hourly and special observations. The 77 stations used in this climatology were chosen by manually checking the data inventories for all first-order stations from 125°W and 100°W longitude; those with approximately 90% or more data during the 1979-2003 period were used (Fig. 3.1).

There were minor problems with these surface data. Data quality varied from station to station, especially 6 h precipitation data. Also, instrumentation and reporting strat- 60

3000 2000 1000 500

Figure 3.1. Stations used for the climatology with 13 km resolution terrain shaded based on scale (m) at lower left. 61 egies were not constant through the 25-year period (1979–2003). The most significant instrumentation change occurred with the installation of ASOS (Automated Surface

Observation System), which began in 1991 and was completed at most locations by 1997.

Another major change occurred 1 July 1996 with the transition from SA (Surface Obser- vation) to METAR (aviation routine weather report) reporting code. After the METAR transition, the average number of special observations per year increased by a factor of two (Fig 3.2a). As a result, the number of strong cold fronts, as identified in this study, increased with the addition of special nonhourly observations that more effectively resolve rapid temperature and pressure changes (Fig. 3.2b).

The National Centers for Environmental Prediction (NCEP) North American

Regional Reanalysis (NARR, Mesinger et al. 2004) was used to generate a composite of strong cold fronts at Salt Lake City. The NARR is a 32 km/45 layer regional reanalysis based on the ETA model 3D-VAR data assimilation system and provides analyses covering North America from 1979–2004. The NARR data were obtained from the National

Climatic Data Center’s Model Data Access National Oceanic and Atmospheric Adminis- tration Operational Model Archive Distribution System Web Interface (http:// nomads.ncdc.noaa.gov/#narr_datasets), and plotted using the Grid Analysis and Display

System (GrADS). NARR composite calculations were also done with GrADS.

Identification of strong cold fronts

A wide and continuous spectrum of atmospheric phenomena exists and any approach to identify “strong” cold fronts is subjective. Numerous salient weather changes may accompany the passage of a cold front including a temperature fall, a pressure rise, a 62

a) s

n 2500 2500 o i t a v r

e 2000 2000 s b O

l a

i 1500 1500 c e p S

f 1000 1000 o

r e b 500 500 m u N

300 300 s t n e v

E 200 200

f o

r e b m u 100 100 N

Figure 3.2. Special observations and the number of cold front events at Salt Lake City (SLC). (a) Number of special (nonhourly) observations at SLC, 1996/7 are not shown owing to inconsistent minutes when observations were reported. (b) Number of strong western United States cold fronts per year. 63 wind direction and/or speed change, a cloud cover increase or decrease, a visibility change, and precipitation. To identify cold fronts in this study, surface temperature and pressure changes are used since these two parameters change most consistently with the passage of a cold front. Thus, temperatures will decrease behind a cold front, and owing to the colder air, a hydrostatic pressure rise will also occur. The magnitude of the pressure rise depends on the depth of the postfrontal air and the speed at which the cold front is moving, with higher postfrontal pressure rises associated with cold fronts whose frontal layer is deep and/or a cold front that moves rapidly. Wind changes are not used to help identify cold fronts owing to local and regional effects of complex topography on wind direction and speed (e.g., Williams 1972; Sanders and Doswell 1995; Steenburgh and

Blazek 2001). Measurements of cloud cover, visibility, and precipitation are not used because of their inherent variations with cold frontal structure.

Two surface criteria are used to identify potential strong cold fronts: a 7°C temperature fall in 2 h or greater but less than 3 h, and a 3 hPa pressure rise in the same period.

This period includes both hourly and special observations. This time frame is chosen since the objective of this study is to sample surface cold fronts with abrupt changes; a longer period may include weaker events that feature broad baroclinity, and a shorter period may include short-lived convective events or data blips.

The temperature fall of 7°C or greater is selected based in part on Sanders (1999a), who defined a strong surface baroclinic zone (i.e., front) as having a gradient of approximately 7°C 100 km-1. If the isotherms are evenly spaced, such a gradient would produce a temperature change of 7°C in 2.5 h if the front moves at approximately 11 m s-1, which is fairly fast. However, if the leading edge of the cold front is sharp, much of the baroclinity 64 may pass within a shorter period. Moreover, examination of average nighttime cooling showed that if a smaller value than ~7°C is chosen, the results may be biased towards days that feature strong diurnal nighttime cooling on calm, clear nights.

The pressure changes are computed using altimeter pressure and the 3 hPa value was selected for two reasons: experience shows that large pressure rises occur behind strong Intermountain cold fronts (e.g., Koppel et al. 2000; Schultz and Trapp 2003), and the 3 hPa value is much larger than the ambient diurnal and semi-diurnal pressure signal, which features early morning and early evening pressure rises of about 0.5 hPa in 2 h over the western United States (Mass et al. 1991). This large pressure rise favors cold fronts with a deep postfrontal layer and/or those that move rapidly.

It is possible that the events identified using only surface temperature and pressure criteria may not be associated with synoptic cold fronts. For example, these events could be produced by gust fronts associated with convection (e.g., Wakimoto 1982). To more effectively constrain the events to “synoptic” cold fronts, a third criteria is used.

This third criteria uses the NARR 700-hPa temperature analysis closest to the time of the frontal passage (NARR data was available every 3 h) to determine if the surface pressure and temperature changes are associated with a synoptic-scale baroclinic zone.

The 700 hPa level is chosen because it is typically near the top of the afternoon planetary boundary layer over the high-elevation western United States and is therefore less sensi- tive to diurnal temperature changes.

The 700-hPa temperature gradients are determined using a hub-and-spoke method that searches in 8 radials (every 45°) from each station at the time closest to frontal passage. The 700-hPa temperature gradient was determined by subtracting the 700-hPa tem- 65 perature at the station from the value 500 km away along each radial (Fig. 3.3). The magnitude of the maximum temperature gradient from the 8 directions is then determined with the following caveat. The three northward pointing radials require a negative value

(i.e., cold air to the north), and the three southward pointing radials require a positive value (Fig. 3.3). This sign limitation for the three southward pointing radials is used primarily because of regional heating contrasts during the monsoon in the desert southwest, which frequently features a cold air temperature anomaly southward of warm temperatures. If this limitation for the three southward-pointing radials is not used, a large number of isolated convective systems would be included in the desert southwest. The east and west radials are allowed to have either negative or positive signs. If these sign conditions were not satisfied, the radial was removed from the computation of the maximum 700-hPa temperature gradient. The magnitude of the maximum 700-hPa temperature gradient is then used to see if the event met the 6°C 500 km-1 criteria.

Events with a 700-hPa temperature gradient less than 6°C 500 km-1 were removed from the climatology. The 6°C 500 km-1 criteria is near the mode of the background temperature gradient (as computed every 3 h from 1979–2003) (Fig. 3.4, black line) and elim- inates 30% of the events identified with only surface temperature and pressure changes

(Fig. 3.4, grey line). 66

Cold W E

S Warm + + +

m k 00 5 +/- +/-

- - -

Figure 3.3. Search scheme used to determine the magnitude and direction of the 700-hPa temperature gradient with station location given as dot. Sign limitation for each radial (+/ –), determined by subtracting the station temperature from the temperature at the end of each search radial. 67

15 15

All Times 12 12 Potential Fronts )

% 9 9 (

y c n e u

q 6 6 e r F

3 3

0 2 4 6 8 10 12 14 16 18 20 22 Temperature Difference (°C / 500km)

Figure 3.4. Histograms of the maximum 700-hPa temperature gradient magnitude for all NARR analyses (black line) and potential cold fronts identified only with surface temperature and pressure criteria (grey line). Potential cold fronts with less than a 6°C 500 km-1 temperature gradient were removed from the climatology. 68

Results

Spatial and temporal distribution of strong colds fronts

There is significant spatial variation in the number of strong cold fronts (hereafter also referred to as events) across the western United States. Coastal locations have very few events, whereas interior regions feature a much greater number of events (Fig. 3.5).

Strong cold fronts are infrequent at coastal locations primarily because landfalling fronts rarely meet the 7°C temperature fall in 2 h criteria. The highest number of events occurs near the eastern slopes of the Rocky Mountains from north-central Montana equatorward into south-central Wyoming and in southeastern Colorado. The greatest number of events

(331) occurs at Casper, WY (CPR), whereas several coastal locations have no events.

Over the Intermountain region, the number of events increases across and downstream of the Sierra Nevada and Oregon Cascades. For example, the number of events increases from 29 at Reno, NV (RNO) to 117 at Salt Lake City, UT (SLC). This pattern suggests that strong cold fronts develop or strengthen over northern Nevada and Utah. The factors responsible for frontal development in this region are described later in this chapter.

Two stations over the Intermountain region are not consistent with this pattern.

The 76 events at Bishop, CA (BIH) is about twice as high as nearby stations. This was likely due to the large ambient diurnal pressure ranges that average 5 hPa and act to enhance the pressure rises accompanying cold fronts. In fact, the average pressure rise with strong cold fronts at BIH was 5.5 hPa, over one standard deviation higher than the

Intermountain average (4.2 hPa). In contrast to BIH, the 50 events at Wendover, NV 69

10 30 50 100 100

200

100

200

3000 2000 1000 500 10 30 50 100

Figure 3.5. Total number of strong cold fronts (1979–2003) with arbitrary contours. Ter- rain (m) shaded according to scale on the left. 70

(ENV) is far fewer than nearby locations. At ENV, northwesterly postfrontal winds experience strong, localized downslope warming, which decreases the temperature gradient accompanying cold fronts. Furthermore, decreased events at ENV can be attributed to missing observations prior to 1996, when approximately 21% of hourly data was missing, with very few special observations (not shown). In comparison, SLC and Elko (EKO) only missed around 2% of hourly data and included an additional 5% of special observations. If adjusted for these missing data, the number of events at ENV would be closer to

60.

For all the western stations, there is a large seasonal cycle in the frequency of strong cold fronts, with a maximum during late spring (Apr-Jun) and minimum in the late fall and early winter (Nov-Feb, Fig. 3.6). This pattern is more pronounced over the Inter- mountain region, with a peak frequency during June.

Geographically, the month of maximum event frequency generally shifts from spring to summer as one moves poleward (Fig. 3.7). An exception is over eastern Colo- rado, where the April peak occurs earlier compared to stations with similar latitudes.

Coastal California and Arizona feature the strongest cool season frequency (Fig. 3.7). This cool season maximum is likely related to the seasonal migration of the polar jet stream and accompanying baroclinic systems, which peak from November–March over the western

United States (Davis and Walker 1992). The high frequency of late spring and summertime events over the northern half of the western United States is not simply explained by the seasonal migration of the jet stream and accompanying baroclinic systems. For example, the number of cold frontal passages peaks during the month of April at SLC (Astling

1984), but the frequency of strong cold fronts, as identified in this study, peaks in late 71

16 All Events 16 Intermountain

) 12 12 % (

y c n

e 8 8 u q e r F 4 4

Jan Feb Mar Apr May Jun Jul Aug Sep Oct Nov Dec

Figure 3.6. Monthly frequency of strong cold fronts for all stations (solid) and Intermoun- tain stations (dashed). 72

GEG/LWS/PDT FCA/MSO/BZN

ACV/BLI/EUG DIK/ISN/GGW LMT/PDX/RDD BKE/BOI BIL/CPR/SHR

SEA/UIL/YKM EKO/NFL/WMC COS/PUB/TAD BIH/RNO

ELY/ENV/SLC

PHX/TUS/YUM

LBB/CVS ABQ/FMN/GUP BFL/DAG/FAT PRB/SAC/SAN SBA/SNS

Figure 3.7. Geographic variability of the monthly frequency of strong cold fronts. Fre- quency inner (outer) circle denote 10% (20%). Month number and stations used for values labeled on outside. 73

May. These results suggest that although the passage of cold fronts decreases in late spring and early summer, the percentage of cold fronts that are “strong,” as defined in this study, is higher during the warm than cool season.

Strong cold frontal passages occur most frequently in the late afternoon and evening, with 78% of all events occurring between 1400 and 2100 LST (Fig. 3.8a). The peak frequency is 14% at 1800 LST, whereas the minimum frequency is 0.7% at 0800

LST. This diurnal cycle is very similar at all locations in the western United States, and to that found by Wiesmueller (1982) for all cold fronts at Denver, Colorado. Although the number of late afternoon events might be enhanced by diurnal cooling, it is also a reflection of frontogenesis associated with daytime diabatic heating, as shown later in this chapter and by other studies. For example, differential sensible heating, due to contrasts in cloud cover, has been shown to strengthen cold fronts during the daytime (e.g., Koch et al.

1995), and the interaction of preexisting cold fronts with a daytime convective boundary layer has been shown to be highly frontogenetic (e.g., Reeder 1986; Physick 1988). The time of maximum frequency also varied by month, tending to peak later in the day during the warm season and earlier in the day during the cool season (Fig. 3.8b).

The direction from which cold fronts came was estimated assuming that the surface front moves in the direction of the maximum 700-hPa temperature gradient analyzed by the NARR. There are situations where this method fails, such as when multiple baroclinic zones exist, the orientation of the surface front departs significantly from the orientation of the 700-hPa isotherms, and/or when fronts have significant along-front wind shear. The latter two may occur when terrain distorts the position of the surface front (e.g.,

Steenburgh and Blazek 2001). Evidence for terrain distortion is given by events in western 74

a) 16 16

12 12 ) % (

y c n

e 8 8 u q e r F 4 4

2400 0300 0600 0900 1200 1500 1800 2100 2400 Local Standard Time (LST)

2100 2100 )

T 2000 2000 S L (

e 1900 1900 m i T 1800 1800 d r a d

n 1700 1700 a t S

l 1600 1600 a c o

L 1500 1500

Jan Feb Mar Apr May Jun Jul Aug Sep Oct Nov Dec

Figure 3.8. Frequency of strong cold fronts. (a) All events by time of day. (b) Peak hourly frequency by month. 75

Montana and vicinity (Fig. 3.9) that feature the maximum baroclinity to the southeast during the time of frontal passage. Cursory inspection of these cases shows the 700-hPa baroclinity advances rapidly downstream while the surface cold front moves slower (not shown).

Over the Intermountain Region, strong cold fronts tend to come from the west and northwest, whereas on the east side of the Rockies, they tend to come from the north (Fig.

3.9). The infrequent passage of cold fronts from the east or northeast over the Intermoun- tain West suggests that cold fronts are rarely intense after they move across the continental divide. This is further illustrated by a closer inspection near the UT-ID-WY border. Cold fronts tend to come from the north at Lander, WY (LND), and to a lesser extent Rock

Springs, WY (RKS) (Fig. 3.10), whereas they come from the west and northwest at SLC,

Burley Idaho (BYI), and Idaho Falls (IDA).

Wind, cloud, and precipitation accompanying strong cold fronts

For all western stations, approximately 13% of events feature at least one wind gust equal to or greater than the National Weather Service high wind warning criteria (26 m s-1) within +/- 12 h of frontal passage. Wind gusts of at least 13 m s-1 occur with about

80% of the events (Fig. 3.11a). The peak wind gust occurs most frequently (69% of the time) during or within an hour of frontal passage (Fig. 3.11b). A similar frequency and magnitude of wind relative to frontal passage is found for Intermountain cold fronts (not shown).

Precipitation (trace or greater) within a 30 h period1 centered on frontal passage occurs during approximately 65% of Intermountain events. After correcting for precipita- 76 tion data quality (i.e., adjusting for missing data), the number of events with precipitation increases by 5–10% and cursory inspection of 80 Intermountain events suggests that another 5–10% of events featured precipitation during their history or within ~200 km of the station (not shown). Thus, precipitation is associated with ~75–80% of Intermountain events. Total precipitation, however, is fairly low with precipitation less than or equal to

5.1 mm 80% of the time, and 15% of the events feature only a trace of precipitation (Fig.

3.12a).

There is a distinct precipitation pattern relative to frontal passage. Measurable precipitation tends to occur (87% of the time) during the 6 h period of frontal passage and two postfrontal periods (Fig. 3.12b). There was little geographic variability to this distribution

(not shown), suggesting similar precipitation structures for strong cold fronts over much of the western United States, although these results are biased to continental regions due to the limited number of strong coastal fronts. Closer inspection of hourly data within 6 hours of 24 SLC strong cold fronts (as identified in this study) from 1998–2003 confirms that the postfrontal precipitation occurs most frequently (63% of the time) around 3 h after frontal passage (Fig. 3.13a, solid line). The geographic variability of hourly precipitation observations relative to cold fronts, however, remains unquantified.

The postfrontal precipitation maximum suggests that the majority of the mid- and low-level clouds occur within the postfrontal environment. This is supported by hourly observations of average cloud height, which show much lower cloud bases in the postfrontal environment (Fig. 3.13a, dashed line). Observations of average cloud cover, however,

show little contrast across the cold front, with a slight decrease 1–2 h ahead of the cold

1. Reliable precipitation data were available only every 6 h at 0600, 1200, 1800, and 0000 UTC so 5 periods were chosen. 77

GEG/LWS/PDT FCA/MSO DIK/ISN/GGW

BIL/CPR/SHR BKE/BOI

EKO/NFL/WMC

LND ELY/ENV/SLC CYS/BFF

PHX/TUS/YUM N NW NE LBB/CVS ABQ/FMN/GUP W E 2040 60 SW SE S

Figure 3.9. Directional frequency of the maximum 700-hPa temperature gradient. Fre- quency given by inner lines (every 20%) for the eight search directions as indicated at lower left with stations used for values annotated. 78

IDA

LND

BYI

RKS N NW NE

W E SLC 2040 60 SW SE S

Figure 3.10. Directional frequency of the maximum 700-hPa temperature gradient. Fre- quency given by inner lines (every 20%) for the eight search directions as indicated at lower right. Location of continental divide depicted by dashed line. 79

a) 1200 1200

900 900 r e

b Not

m Reported u 600 600 N

300 300

0 5 10 15 20 25 30 35 m s-1 (2.5 unit bins)

b) 24 24

18 18

t n e c

r 12 12 e P 6 6

+12 HR +6 HR 0 HR -6 HR -12 HR Hour Relative to Frontal Passage

Figure 3.11. Wind characteristics associated with all strong cold fronts, as defined in this study. (a) Histogram (2.5 m s-1 bins) of maximum wind gust within +/- 12 hours of frontal passage for all events. (b) Frequency of maximum wind gust relative to frontal passage for all events. 80

2000 2000

1500 1500 r e b m

u 1000 1000 N

500 500

0 Trace 2.5 5.1 7.6 10.1 12.7 15.2 17.8 20.3 22.9 25.4 27.9 30.5 Precipitation (2.54 mm bins)

b) 50 50

40 40

t 30 30 n e c r e 20 20 P

10 10

0 0 +12 HR +6 HR 0 HR -6 HR -12 HR Hour Relative to Frontal Passage

Figure 3.12. Precipitation characteristics associated with all strong cold fronts, as defined in this study. (a) Histogram of 30 h precipitation (2.54 mm bins) associated with strong cold fronts. (b) Precipitation frequency relative to frontal passage as given by 6 h precipitation reports. 81

a) t

h 16 16 g i e H 12 12 d u o l C

/ 8 8 r e b m

u 4 4 N

+6 HR +4 HR +2 HR 0 HR -2 HR -4 HR -6 HR

b) OVC OVC

BKN BKN

SCT SCT

FEW FEW +6 HR +4 HR +2 HR 0 HR -2 HR -4 HR -6 HR

Figure 3.13. Cloud and precipitation properties relative to strong SLC frontal passages (1998–2003). (a) Average cloud ceiling (thousands of feet, dashed line), and number of events (out of 24) with precipitation (solid line). (b) Average cloud cover. 82 front (Fig. 3.13b). Despite little contrast in cloud cover between the pre and postfrontal environment, temperature observations show that prefrontal cloudiness does not dramatically inhibit daytime warming.

This cloud and precipitation configuration suggests that diabatic processes are important for the development of strong Intermountain cold fronts. For example, this cloud structure would be frontogenetic during the day since the prefrontal environment will be more strongly heated than the low-cloud postfrontal environment. Such cloud- cover contrasts across cold fronts have been shown to be highly frontogenetic through differential sensible heating (e.g., Koch et al. 1995; Segal et al. 2004). Furthermore, precipitation frequency relative to cold fronts suggests that low-level diabatic cooling associated with evaporative cooling and moist downdrafts may decrease the temperature of postfrontal air, thereby playing an important role in frontal development.

Characteristics of strong Intermountain cold fronts

As shown in the previous section, the frequency of strong cold fronts increases eastward across the Intermountain West, reaching a maximum at SLC. The remainder of this chapter examines the frontal development over the Intermountain West by focusing on cold fronts at SLC. SLC has the highest number of events (117) over the Intermountain region and, compared with other Intermountain stations, has the highest frequency of

“wet” cold fronts (77%), and the highest mean total precipitation with strong cold fronts.

Our approach involves producing a NARR-based composite of the 25 strongest events (based on the magnitude of temperature fall) at SLC (Table 3.1). These 25 events feature a diurnal frequency that is similar to that of all events, and a seasonal frequency 83

Table 3.1: 25 Strongest Cold Fronts at SLC

Hour Temperature Day Month Year (UTC) Fall (°C)

2100 15 Apr 2002 -19

1900 31 Mar 1997 -16

0000 06 Jun 1995 -16

0100 28 May 1994 -14

0200 17 Mar 1989 -13

2300 21 May 1986 -13

0000 28 May 1982 -13

0200 13 Jun 1992 -13

2300 09 Jul 1983 -13

0300 15 Jul 1983 -13

0200 19 Apr 1987 -12

0200 13 Apr 1986 -12

0300 04 May 1993 -12

0100 01 Jul 1997 -12

2300 24 Sep 1992 -12

0000 18 Sep 1989 -12

1800 09 Sep 1986 -12

0000 24 May 1989 -11

0200 09 May 1991 -11

0100 06 Jul 1994 -11

0600 16 Sep 1996 -11

1400 11 Mar 1990 -10

2200 19 Mar 2000 -10

2100 02 Apr 2003 -10

0200 25 Apr 1984 -10 84 that is similar to that over the Intermountain region with events from March to September excluding August (Table 3.1). The composite features spatial variability that is introduced by averaging events from March to September, and to a lesser extent differing large-scale structures with events. The variability among composite members is discussed where nec- essary. Despite this variability, we believe the composite structure and evolution is sufficient for identifying the large-scale structure and conditions under which strong cold fronts develop.

At 24 h prior to the event at SLC (-24 h), no well-defined composite surface cold front is observed. Rather, a broad baroclinic zone is present to the north and west of SLC

(Fig. 3.14a) with the leading edge of the baroclinity bisecting Nevada from southwest to northeast (Fig. 3.14b). Inspection of the composite members showed that 17 of the 25 members feature a similar baroclinic structure (not shown). Winds at 700 hPa are mainly parallel to the isotherms (Fig. 3.14b) and therefore not favorable for confluent frontogenesis. Over the Intermountain region, surface winds are also southwesterly, with warm surface potential temperatures that exceed 32°C over a large area ahead of the baroclinic zone

(Fig. 3.14a). A broad area of low pressure encompasses much of the interior western

United States, and confluent southwesterly surface flow is present over central Nevada and represents the location of frontal development. Aloft, a mobile long-wave upper-level trough approaches the Pacific coast, with stronger 250-hPa winds on the western side of the trough axis (Fig. 3.14d), a favorable pattern for trough digging (Bluestein 1993, p. 71).

At 500 hPa, a broad area of weak ascent is present ahead of the trough axis and within the western or cold side of the baroclinic zone (Fig. 3.14c). Inspection of hour -24 500-hPa height analyses of all the composite members shows that many events feature an upper- 85

a) HR -24: SFC b) HR -24: 700 hPa

.SLC

36 32 28 24 20 c) HR -24: 500 hPa d) HR -24: 250 hPa

Figure 3.14. NARR composite for the strongest 25 cold fronts at hour -24. (a) Mean sea level pressure (every 2 hPa, contoured), wind [full (half) barbs denote 5 (2.5) m s-1], and surface potential temperature (shaded every 4°C according to scale at lower left). (b) 700- hPa temperature (contoured every 2°C), wind [full (half) barbs denote 5 (2.5) m s-1], and relative humidity (shaded and contoured). (c) 500-hPa geopotential height (every 60 m), and vertical velocity (every 3 cm s-1, negative values contoured, positive values shaded and contoured). (d) 250-hPa wind speed (contoured and shaded every 4 m s-1 beginning with 28 m s-1) and wind vectors. Dashed line in (a) denotes position of confluent flow. 86 level trough axis in a similar location to the composite, but that the amplitude and wave- length vary (Fig. 3.15).

At hour -12, the low level (700 hPa and surface) baroclinic zone and pressure trough intensify over central Nevada (Fig. 3.16a) with the leading edge of the composite baroclinic zone oriented from southwest to northeast. The position of the composite cold front was defined as the leading edge of surface potential temperature gradient, which at this time, was coincident with the composite surface wind shift along the confluent flow.

Inspection of composite members shows that the composite baroclinity pattern resembles

21 cases, with the remaining cases featuring either meridionally or zonally oriented cold fronts (not shown). The intensification of the incipient cold front occurs as cold advection advances towards the quasi-stationary leading edge of the low-level confluent flow (Fig.

3.16b). Aloft, the 250-hPa trough digs southeastward (Fig. 3.16d) and ascent intensifies downstream of the 500-hPa trough axis (Fig. 3.16c). This digging occurs during 23 of the

25 composite members, suggesting that the upper-level trough plays a significant role in frontal development. Since upward motion is strongest within the cold side of the baroclinic zone (cf. Fig. 3.16b and Fig. 3.16c), this favors the development of clouds and precipitation behind (westward) the confluent surface flow or location of incipient frontal development.

Confluent surface flow over Nevada represents a boundary between warm, dry

Intermountain air and colder Pacific air. Examination on individual events with the NARR and observations show that the airmass contrasts across the confluent flow are due to temperature, with little contrast in low-level stability and moisture. Orographic processes associated with the Sierra Nevada may contribute to the location of this confluent flow. 87

Composite

Figure 3.15. Composite (top left) versus individual 500-hPa geopotential height analyses (every 60 m) at hour -24. 88 a) HR -12: SFC b) HR -12: 700 hPa

.SLC

36 32 28 24 20 B c) HR -12: 500 hPa d) HR -12: 250 hPa

Figure 3.16. NARR composite for the strongest 25 cold fronts at hour -12. (a) Mean sea level pressure (every 2 hPa, contoured), wind [full (half) barbs denote 5 (2.5) m s-1], and surface potential temperature (shaded every 4°C according to scale at lower left). (b) 700- hPa temperature (contoured every 2°C), wind [full (half) barbs denote 5 (2.5) m s-1], and relative humidity (shaded and contoured). (c) 500-hPa geopotential height (every 60 m), and vertical velocity (every 3 cm s-1, negative values contoured, positive values shaded and contoured). (d) 250-hPa wind speed (contoured and shaded every 4 m s-1 beginning with 28 m s-1) and wind vectors. Dashed line in (a) denotes position of confluent flow. 89

This may occur in response to regional-scale (~500 km) flow distortion by the Sierra

Nevada, but the extent to which the Sierra affects the location and intensity of frontal development, however, remains for future study. Nonetheless, the significance of this confluent flow is apparent at hour -9, when the confluent flow becomes convergent (Fig. 3.17) and highly frontogenetic.

A cross section perpendicular to the developing cold front at hour -9 shows a deep layer of ascent over the developing cold front, which was positioned around 115°W (Fig.

3.18a). This vertical motion is coupled to the surface since is was immediately above a region of low-level convergence where the horizontal ageostrophic flow is confluent (Fig.

3.18b). The vertical motion is related to the approach of the upper-level trough, but it may also be a response to frontogenesis (Sawyer 1956; Eliassen 1962). Separation of the large- scale and frontal circulation and their relative importance, however, remains for future study.

Over the next 3 h, rapid frontal development continues in eastern Nevada, and by hour -6 an intense surface cold front extends from south-central Nevada northeastward into southeastern Idaho (Fig. 3.19). Examination of surface potential temperature at hour -

6 for individual cases shows that approximately 17 events closely resemble the composite cold front position (Fig. 3.20). The remaining events feature cold fronts that are more meridionally or zonally oriented, with the latter featuring a preexisting surface front over northern Utah in 5 cases. Similar to the previous period, the frontal intensification occurs as cold advection advances towards the quasi-stationary warm edge of the baroclinity

(Fig. 3.19b). Furthermore, composite surface potential temperatures ahead of the front increase around 4°C to 36°C suggesting daytime prefrontal heating also acts to intensify 90

HR -9: 700-hPa Divergence

: 10 m s-1

Figure 3.17. 700-hPa divergence at hour -9. Divergence (negative values, every 3x10-5s- 1).Wind vectors according to scale. 91

: 10 m s-1

Figure 3.18. Composite cross sections at hour -9 along AB in Fig. 3.16b. (a) Potential temperature (every 3°C), vertical velocity (positive values shaded and contoured every 0.3 Pa s-1, negative values every 0.1 Pa s-1 with dashed line), and wind [full (half) barbs denote 5 (2.5) m s-1]. (b) Horizontal ageostrophic wind vectors, divergence (negative values shaded every 2x10-5s-1), and potential temperature (every 3°C). 92

a) HR -6: SFC b) HR -6: 700 hPa

.SLC

36 32 28 24 20

c) HR -6: 500 hPa d) HR -6: 250 hPa

Figure 3.19. NARR composite for the strongest 25 cold fronts at hour -6. (a) Mean sea level pressure (every 2 hPa, contoured), wind [full (half) barbs denote 5 (2.5) m s-1], and surface potential temperature (shaded every 4°C according to scale at lower left). (b) 700- hPa temperature (contoured every 2°C), wind [full (half) barbs denote 5 (2.5) m s-1], and relative humidity (shaded and contoured). (c) 500-hPa geopotential height (every 60 m), and vertical velocity (every 3 cm s-1, negative values contoured, positive values shaded and contoured). (d) 250-hPa wind speed (contoured and shaded every 4 m s-1 beginning with 28 m s-1) and wind vectors. Solid line in (a) denotes position of composite cold front. 93

Composite

Figure 3.20. Composite (top left) versus individual surface potential temperature (every 1°C) at hour -6. 94 the front (Fig. 3.19a). This is supported by inspection of composite members, which show that 23 of the 25 events feature prefrontal warming from hour -12 to hour -6, and observations at SLC show that temperatures warm an average of 10°C.

The interaction of the upper-level trough with the low-levels is important to intensify the surface cold front. This is supported by the faster eastward movement of the upper-level trough (Fig. 3.19d) and accompanying vertical motion (Fig. 3.19c) than the surface cold front, which, at this time, results in the strong upper-level vertical motion near the surface cold front. As the upper-level vertical motion approaches the cold front, it likely enhances the cross-front circulation, which, through mass continuity, increases surface convergence and therefore low-level frontogenesis.

The convective boundary layer on both sides of the front deepens with daytime heating (almost all of the events occur in the late afternoon or evening), eventually exceeding 3 km in depth, with the deepest boundary layer near the surface front (Fig.

3.21). Such boundary layers have been shown to enhance prefrontal vertical motions and dramatically increase frontogenesis along coastal fronts interacting with heated land (e.g,

Reeder 1986; Physick 1988; Hakim 1992). In contrast to the observed intensification of cold fronts by a convective boundary layer, investigations of continental Australian subtropical cold fronts, which feature strong homogeneous heating, have shown the daytime boundary layer is moderately frontolytic (Smith et al. 1995; Reeder and Tory 2005), but strongly frontogenetic when nocturnal cooling begins. The latter, however, usually feature little or no horizontal cross-front daytime heating contrasts. The direct role of the convective boundary layer on the incipient development of the composite cold front in this case is 95

a) HR -12: PBL Depth b) HR -9: PBL Depth

.SLC

: 10 m s-1

c) HR -6: PBL Depth d) HR -3: PBL Depth

Figure 3.21. Composite planetary boundary layer depth (km above NARR ground level) and surface wind vectors (scale at upper left) at: (a) -12 h. (b) -9 h. (c) -6 h. (d) -3 h. Loca- tion of confluent flow given by dashed line and cold front solid line. Wind vectors as in (a). 96 unclear, but the indirect effects of allowing strong heating and high potential for evaporation may be quite important.

The prefrontal environment also features strong composite southerly 700-hPa flow that exceeds 15 m s-1 (Fig. 3.22a) and resembles a low-level jet. These prefrontal souther- lies have a strong ageostrophic component that is directed towards the cold front (Fig.

3.22b). The ageostrophic wind converges at the cold front (Fig. 3.22c), and as will be shown, is responsible for a large part of the frontal intensification. In comparison to the postfrontal environment, the prefrontal environment features much broader ageostrophic flow that extends through a deep layer. This ageostrophic flow is not related to an isallobaric response (not shown), rather, it appears to occur in response to frontogenesis.

The prefrontal environment is quite dry, with surface dewpoint depressions around

18°C 3 h prior to the event (Fig. 3.23), resulting in low-level relative humidities around

30% and an average prefrontal lifting condensation level around 4 km above ground level.

Consequently, none of the 25 composite members featured measurable prefrontal precipitation. This dry prefrontal environment allows for strong sensible heating or cooling that may be frontogenetic/lytic depending on the time of day and the heating/cooling rate in the postfrontal environment.

Owing to the limited moisture, composite prefrontal convective available potential energy (CAPE), which is derived using the most unstable parcel in the lowest 180 hPa, is generally less than 100 J kg-1 (Fig. 3.24). However, inspection of 141 prefrontal observed

0000 UTC soundings from the 25 composite members show a somewhat bimodal CAPE

1. Only 14 soundings were used since these cases featured soundings in the prefrontal environment, and two cases had missing data. 97

a) HR -6:700-hPa wind speed b) HR -6: 700-hPa ageostrophic wind A

: 15 m s-1 B

c) ) a P h (

e r u s s e r P

: 15 m s-1

Figure 3.22. Wind characteristics at hour -6. (a) 700-hPa wind speed (shaded and contoured, every 5 m s-1) and wind vectors. (b) 700-hPa ageostrophic wind vectors and potential temperature (every 2°C). (c) Cross section along AB given in (a), potential temperature (every 3°C) and horizontal ageostrophic wind vectors. Solid line in (b) indicates position of surface cold front. 98

Hour -3: SLC NARR Sounding

Temperature (°C)

Figure 3.23. Composite NARR skew T–logp diagram at -3 h, temperature (solid), dewpoint (dashed), and wind [full (half) barbs denote 5 (2.5) m s-1]. 99

Hour -3: CAPE

Figure 3.24. Convective available potential energy (J kg-1) at -6 h. Solid line indicates position of composite surface cold front. 100 distribution with half of the events having less than 100 J kg-1 and four with CAPE greater than 500 J kg-1 (not shown). The members with higher CAPE values occur during the late spring and early summer, suggesting convection is more likely to be associated with cold fronts during the warm than cool season. Furthermore, reports of thunder associated with strong SLC fronts indicate that most of the ~12% of events associated with thunder during the hour of frontal passage or within 6 h after the event occur in the warm season. The low

CAPE values and limited thunder reports suggest that convection associated with strong

Intermountain cold fronts is generally not too intense or deep, but given the low static stability in the daytime well-mixed boundary layer near the cold front, it is suggested that precipitation along or near the cold fronts is convective. In contrast to relatively low

CAPE, values of downdraft convective available potential energy (DCAPE) (Gilmore and

Wicker 1998) are considerably higher, reaching 800 J kg-1 at SLC at -3 h. This means that a high potential exists for evaporatively cooled downdrafts to reach the surface if evaporation and sublimation of precipitation is sufficient to saturate downdraft air. This evaporatively cooled air can strengthen the postfrontal cold air of Intermountain cold fronts (e.g.,

Schultz and Trapp 2003), and in many cases may play an important role during their development.

After a period of quasi-stationary development, the surface cold front increases in speed and passes SLC at 0 h (Fig. 3.25a). This evolution was similar to all composite members where the cold front undergoes quasi-stationary intensification from hour -12 to

-6 and thereafter becomes mobile. The strong surface cold front extends from southeastern

Nevada into northwestern Wyoming (Fig. 3.25a), a configuration that closely resembles

23 of the 25 composite members, with strong cold advection behind the front (Fig. 3.25b). 101

a) HR 0: SFC b) HR 0: 700 hPa

.SLC

36 32 28 24 20

c) HR 0: 500 hPa d) HR 0: 250 hPa

Figure 3.25. NARR composite for the strongest 25 cold fronts at hour 0. (a) Mean sea level pressure (every 2 hPa, contoured), wind [full (half) barbs denote 5 (2.5) m s-1], and surface potential temperature (shaded every 4°C according to scale at lower left). (b) 700- hPa temperature (contoured every 2°C), wind [full (half) barbs denote 5 (2.5) m s-1], and relative humidity (shaded and contoured). (c) 500-hPa geopotential height (every 60 m), and vertical velocity (every 3 cm s-1, negative values contoured, positive values shaded and contoured). (d) 250-hPa wind speed (contoured and shaded every 4 m s-1 beginning with 28 m s-1) and wind vectors. Solid line in (a) denotes position of composite cold front. 102

The strongest composite upper-level vertical motion is observed as the upper-level trough becomes negatively tilted (Fig. 3.25c and Fig. 3.25d). This vertical motion exceeds 9 cm s-

1 immediately overhead the surface cold front, meaning that the large-scale vertical motion or “synoptic forcing” is in phase with the surface front. This is confirmed with examination of composite members, which shows that 500-hPa vertical motion is strongest near the surface front. Inspection of a cross section perpendicular to the front shows that vertical motion is focused at and ahead of the leading edge of the cold front (Fig.

3.26a). However, the strongest upper-level vertical motion does not appear to be entirely coupled to low-level convergence with the strongest upward motion ahead of the largest surface convergence (cf. Figs. 3.26a,b).

By hour +6, the cold front moves downstream of SLC and extends from extreme southern Nevada northeastward toward the WY-UT-CO border (Fig. 3.27a). The northern portion of the front moves faster than the southern portion; this configuration was similar for 16 of the 25 cases. This rapid movement is likely due to strong cross-front postfrontal flow where strong cold advection (Fig. 3.27b) and high pressure rapidly build from the west (Fig. 3.27a). Aloft, the composite upper-level trough axis is over eastern Nevada

(Fig. 3.27) with the strongest upward motion over northern Utah (Fig. 3.27c), but composite vertical motion weakens considerably.

The most rapid movement of the cold front occurs during the 6 h after frontal passage. This movement is accompanied by NARR composite 6 h sea level pressure rises that exceed 8 hPa (Fig. 3.28a), and in one case reaches 16 hPa. These NARR pressure rises were surprising close to those observed. Observations of frontal movement show that many strong cold fronts move faster than the postfrontal winds within the NARR, and are 103

: 10 m s-1

Figure 3.26.Composite cross sections at hour 0 along AB in Fig. 3.16b. (a) Potential temperature (every 3°C), vertical velocity (positive values shaded and contoured every 0.3 Pa s-1, negative values every 0.1 Pa s-1 with dashed line), and wind [full (half) barbs denote 5 (2.5) m s-1]. (b) Horizontal ageostrophic wind vectors, divergence (negative values shaded every 2x10-5s-1), and potential temperature (every 3°C). 104 a) HR +6: SFC b) HR +6: 700 hPa

.SLC

36 32 28 24 20

c) HR +6: 500 hPa d) HR +6: 250 hPa

Figure 3.27. NARR composite for the strongest 25 cold fronts at hour +6. (a) Mean sea level pressure (every 2 hPa, contoured), wind [full (half) barbs denote 5 (2.5) m s-1], and surface potential temperature (shaded every 4°C according to scale at lower left). (b) 700- hPa temperature (contoured every 2°C), wind [full (half) barbs denote 5 (2.5) m s-1], and relative humidity (shaded and contoured). (c) 500-hPa geopotential height (every 60 m), and vertical velocity (every 3 cm s-1, negative values contoured, positive values shaded and contoured). (d) 250-hPa wind speed (contoured and shaded every 4 m s-1 beginning with 28 m s-1) and wind vectors. Solid line in (a) denotes position of composite cold front. 105

a) b)

: 10 m s-1

Figure 3.28. Six-hour surface potential temperature and pressure changes ending at +6 h. (a) Mean sea level pressure rises (every 2 hPa beginning with 4 hPa). (b) Temperature falls (every 2°C less than -4°C). Wind vectors as in (a). Composite cold front position denoted by solid lines. 106 accompanied by strong northwesterly winds within an hour or two of frontal passage.

These northwesterly winds immediately behind the cold front likely have a large cross- front ageostrophic value that helps to control the forward movement of the cold front (e.g.,

Smith and Reeder 1988), but the extent of this remains for future study.

Concurrent with large pressure rises are large temperature falls that exceed 8°C

(Fig. 3.28b). These NARR temperature falls, however, drastically underestimate observed values, which range from 9–24°C and have a mean of 16.5°C (Fig. 3.29a), showing that cold fronts in the NARR are, on average, about half the strength as observed.

The postfrontal environment features higher surface and low-level (700 hPa and below) relative humidity due mainly to colder temperatures and not increased absolute humidity (Fig. 3.29b). The higher relative humidity results in a postfrontal lifting condensation level (LCL) around 1 km AGL, which is 3 km lower than the prefrontal environment (Fig. 3.29b). The lower postfrontal LCL is consistent with cloud base observations

(Fig. 3.13).

The composite evolution and examination of composite members described above provides a qualitative look at the development of strong cold fronts over the Intermountain

West. Most strong surface cold fronts develop rapidly on the leading edge of a preexisting baroclinic zone over central Nevada. The development occurs along confluent flow in central Nevada, which represents the boundary between warm and dry Intermountain air to the south and east and colder Pacific air to the north and west. The location of the confluent flow, in many cases, appears to be dictated by flow interactions with the Sierra

Nevada. The confluent flow quickly becomes convergent, and the cold front intensifies as cold advection, which is associated with a deepening upper-level trough, moves towards 107

a) mean = 16.5°C 8 8 mean = +3°C

6 6 N r e u b m m b u

4 4 e r N

2 2

-24 -21 -18 -15 -12 -9 -5 -2 1 4 7 10 Temperature Change (°C) Dewpoint Change (°C)

b) Hour: -3 Hour: +6 ) a P h (

e r u s s e r P

Temperature (°C)

Figure 3.29. SLC observed temperature and moisture changes versus NARR composite changes. (a) Histogram of SLC observed 6 h surface temperature and dewpoint changes ending at hour +6. (b) SLC NARR composite skew-t diagram for hour -3 and hour +6, temperature (solid), dewpoint (dashed). 108 the quasi-stationary cold front while highly ageostrophic flow impinges from the southeast. The cold front further intensifies as clouds and precipitation, which are forced by large-scale ascent with the upper-level trough and the frontogenesis, develop behind the cold front. The configuration of clouds and precipitation and observed heating relative to the cold front suggests that differential diabatic processes are important for intensifying the cold front, with diabatic frontogenesis strongest during the daytime. After rapid intensification, the cold front then moves rapidly downstream accompanied by large pressure rises and temperature falls.

Frontogenesis analysis

To further diagnose the processes contributing to the development of strong Inter- mountain cold fronts, we calculate frontogenesis (Petterssen 1936), defined as the

Lagrangian rate of change of the magnitude of the horizontal potential temperature (θ) gradient:

d F = ---- ∇ θ (3.1a) dt p

where

∂ ∂ ∂ ∂ --d-- ω = ∂ + u∂ + v∂ + ∂ (3.1b) dt t xp yp p 109

∂ ∂ ∇ p = i∂ + j∂ (3.1c) xp yp

and the subscript p denotes differentiation along a constant pressure surface. For surface frontogenesis calculations, differentiation is done along the NARR surface layer.

Following Miller (1948), equation 3.1a may be written:

F = FW + FT + FD (3.2)

where

1 ∂θ ∂u∂θ ∂u∂θ ∂θ ∂v∂θ ∂v∂θ ------≈------’ -----≈------’ FW = – « + ◊ + « + ◊ (3.3) ∇pθ ∂x ∂x ∂x ∂y ∂y ∂y ∂x ∂x ∂y ∂y

∂θ ∂ω∂θ ∂ω∂θ -----1------≈------’ FT = – « + ◊ (3.4) ∇pθ ∂p ∂x ∂x ∂y ∂y

1 ≈ ∂θ ∂ dθ ∂θ ∂ dθ’ FD = –------«– ------– ------◊ . (3.5) ∇pθ ∂x∂x dt ∂y∂y dt

FW is the frontogenesis due to horizontal confluence and convergence, FT is the tilting frontogenesis, and FD is the frontogenesis due to horizontal gradients in diabatic heating and cooling. Since the boundary layer was deep, well-mixed and ∂θ/∂p was near zero during the period of rapid frontogenesis, FT at the surface was negligible and is not presented. 110

θ To calculate FD, d /dt was estimated as a residual of the thermodynamic energy equation:

θ θ dQ ∂θ ∂θ d------⋅ ∇ θ ω----- = = ∂ + V p + ∂ (3.6) dt CpT dt t p

where Q is the diabatic heating rate, T is the temperature, and Cp is the specific heat of dry air at constant pressure. The local potential temperature change was evaluated with centered time differencing using values at t +/-∆t where ∆t is the 3 h NARR analysis time step. The temperature change due to horizontal and vertical advection was computed using the average advection from t-∆t to t+∆t. Tests using one-sided differencing and instanta- neous horizontal/vertical advection at time t showed limited sensitivity to the approach used for these estimates. However, estimates of advection may contain errors due to averaging of instananous advection values, since the observed advection may differ from the average values.

Frontogenesis calculations are influenced by the poor representation of low-level temperatures in the NARR. Comparison of the observed June mean diurnal temperature range at Intermountain observing sites with the NARR 0000 UTC – 1200 UTC temperature difference suggests the NARR underestimates the diurnal temperature cycle by roughly a factor of two (Fig. 3.30). Therefore, diurnal diabatic frontogenesis effects, par- ticularly near the surface, may not be fully estimated. Frontogenesis calculations are also influenced by the variability of composite members. Although the large-scale temperature pattern is similar for many composite members (Fig. 3.20), actual values feature much 111

a) b)

Figure 3.30. Average diurnal surface temperature range (°C) for June. (a) Observed based on 1971–2000 average. (b) NARR, derived by subtracting the warmest (0000 UTC) and coldest (1200 UTC) temperatures based on 2000-2002 average. 112 variability with standard deviations of 6–7°C at the surface and 700 hPa at the time of frontal passage (Fig. 3.31). Consequently, the magnitude of composite frontogenesis values, in most cases, is smaller than that observed within the composite members.

Frontogenesis calculations are presented beginning at hour -12, which represents a composite mean time of 0900 MST, and the time that the most rapid frontal development commences. At this time, the composite southwest-northeast oriented surface baroclinic zone is located northwest of SLC (Fig. 3.32a). Total surface frontogenesis exceeds 2°C

-1 -1 100 km 6 h along confluent flow in central Nevada (Fig. 3.32a). An area of positive FW

(negative values not shown for clarity) also exists along the confluent flow and near the

Nevada-Idaho border (Fig. 3.32b). Diabatic frontogenesis is strongest ahead of the confluent flow and has a greater magnitude than Fw (Figs. 3.32b and 3.32c). This diabatic frontogenesis results from a minimum in diabatic heating ahead of the confluent flow (Fig.

3.32d), and is produced by reduced sensible heating associated with clouds and precipitation during many of the events. Although beyond the scope of this study, diabatic frontogenesis is large along the Pacific coast, and is produced by large contrasts in heating between land and ocean (Fig. 3.32d). At 700 hPa, the total frontogenesis is weaker, between 0.5–1°C 100 km-1 6 h-1 (Fig. 3.33a). Much of the total frontogenesis along the confluent flow is due to FW (Fig. 3.33b). FT is small in the vicinity of the incipient cold front (Fig. 3.33c), and diabatic frontogenesis is positive well ahead of and behind the confluent flow (Fig. 3.33d).

At hour -6, which represents a composite mean time of 1500 MST, strong frontogenesis is underway. Total frontogenesis exceeds 2°C 100 km-1 6 h-1 ahead of the surface 113

a) b)

Figure 3.31. Standard deviation of composite members at time of frontal passage. (a) Sur- face potential temperature. (b) 700-hPa potential temperature. 114

a) b)

2 1.5 1 0.5 : 10 m s-1

c) d)

Figure 3.32. NARR surface frontogenesis composites at hour -12. (a) Total frontogenesis (shaded every 0.5°C 100 km-1 6 h-1 according to scale on right), potential temperature (every 3°C). (b) Confluent and convergent frontogenesis according to scale in (a). (c) Diabatic frontogenesis according to scale in (a). (d) Diabatic temperature changes. Wind vectors according to scale in (a). Dashed line in (d) indicates confluent flow. 115

a) b)

2 1.5 1 0.5 : 10 m s-1

c) d)

Figure 3.33. NARR 700-hPa frontogenesis composites at hour -12. (a) Total frontogenesis (shaded every 0.5°C 100 km-1 6 h-1 according to scale on right), potential temperature (every 2°C). (b) Confluent and convergent frontogenesis according to scale in (a). (c) Tilting frontogenesis according to scale in (a). Wind vectors according to scale in (a). (d) Diabatic frontogenesis according to scale in (a). 116

-1 - cold front (Fig. 3.34a). FW is largest along the confluent flow, exceeding 1°C 100 km 6

1 from southwestern Nevada to southern Idaho (Fig. 3.34b). Examination of composite members shows that FW is much larger during most events, indicating that composite FW is underestimated. The strongest composite total frontogenesis occurs ahead of the cold front due to large values of FD, which are much larger than FW (cf. Figs. 3.34b,c). Curi- ously, the position of positive FW is not in phase with positive FD, where positive FD is produced by a diabatic heating minimum ahead of the front (Fig. 3.34d). This diabatic heating minimum is roughly coincident with contrasts in NARR sensible heating, especially near the northern portion of the front (Fig. 3.35a), and the reduced sensible heating is likely the result of NARR precipitation, which occurs well into the prefrontal environment (Fig. 3.35b). However, the observed precipitation almost always occurs along or behind the surface cold front (Fig. 3.35c). In addition, as shown previously, the diurnal cycle in the NARR is smaller than observed. For these reasons, the diabatic heating gradient and frontogenesis are likely underestimated and more in phase with the front than suggested by the NARR.

At 700 hPa, the total frontogenesis was similar to that at the surface (Fig. 3.36a) with FW and FD dominating (Figs. 3.36b and 3.36d), but as at the surface, FD was ahead of the cold front. FT was negligible near the cold front due to the lack of a vertical potential temperature gradient within the deep, well-mixed boundary layer, but featured positive values in central California (Fig. 3.36c). These results show that the composite frontal development along the cold front is driven largely by confluent and convergent flow. As at 117

a) b)

2 1.5 1 0.5 : 10 m s-1

c) d)

Figure 3.34. NARR surface frontogenesis composites at hour -6. (a) Total frontogenesis (shaded every 0.5°C 100 km-1 6 h-1 according to scale on right), potential temperature (every 3°C). (b) Confluent and convergent frontogenesis according to scale in (a). (c) Diabatic frontogenesis according to scale in (a). Wind vectors according to scale in (a). (d) Diabatic temperature changes. Solid line in (d) indicates position of composite cold front. 118

a) b)

c) 25 25

20 3.3 3.3 20 1.3 r e

b 15 15 m u

N 10 10 Trace 5 Trace 5

+ 12 HR + 6 HR 0 HR -6 HR -12 HR

Figure 3.35. NARR heating rates and precipitation versus observed precipitation of composite members. (a) Surface sensible heat flux (W m-2) at -6 h. (b) 6 h accumulated NARR precipitation (mm) ending at -6 h. (c) Frequency of precipitation relative to cold frontal passage at SLC for top 25 events with average precipitation amount (mm) annotated. Solid line in (a) indicates composite surface cold front position. 119

a) b)

2 1.5 1 0.5 : 10 m s-1

c) d)

Figure 3.36. NARR 700-hPa frontogenesis composites at hour -6. (a) Total frontogenesis (shaded every 0.5°C 100 km-1 6 h-1 according to scale on right), potential temperature (every 2°C) and wind vectors. (b) Confluent and convergent frontogenesis according to scale in (a). (c) Tilting frontogenesis according to scale in (a). (d) Diabatic frontogenesis according to scale in (a). Wind vectors according to scale in (a). 120

the surface, positive FD is out of phase with positive FW. Case studies and model simulations may be needed to more carefully diagnose the diabatic frontogenesis in these events.

Comparison of geostrophic and ageostrophic frontogenesis at hour -6 (i.e., FW cal- culated using geostrophic and ageostrophic winds, respectively) shows that only a small portion of the 700-hPa FW can be explained by the geostrophic frontogenesis (c.f. Figs.

3.37a,b). Instead, ageostrophic FW dominates because of broad southeasterly ageostrophic inflow towards the cold front. This ageostrophic flow is not related to an isallobaric response (not shown), rather, it appears to occur in response to frontogenesis. Inspection of other times and cross sections (not shown) indicates that the low-level ageostrophic flow is responsible for much of the surface and low-level frontal development.

The strongest composite total surface frontogenesis occurs as the cold front passes

SLC at hour 0, which represents a mean time of 2100 MST (Fig. 3.38a). A band of positive FW extends northeastward from southern Nevada and spreads into two sections over northern Utah with values reaching 1.5°C 100 km-1 6 h-1 (Fig. 3.38b). A broad area of

-1 -1 positive FD that exceeds 2°C 100 km 6 h exists ahead of southern portion of the cold front, but over northern Utah positive FD is coincident with the cold front and positive FW

(Fig. 3.38c). The FD maximum over northern Utah is the result of differential diabatic cooling maximum immediately behind the front (Fig. 3.38d). This cooling may be attributed to evaporation and sublimation, as supported by the frequency and magnitude of postfrontal precipitation (Fig. 3.35c). At 700 hPa, the total frontogenesis exceeds 2°C 100

-1 -1 km 6 h over a broad area (Fig. 3.39), with FW located behind the strongest total frontogenesis (3.39c) and FD ahead (Fig. 3.39d). The high values of 700-hPa FD are due to 121

a) b)

2 1.5 1 0.5 : 15 m s-1

Figure 3.37. Geostrophic and ageostrophic confluent and convergent frontogenesis at hour -6. (a) 700-hPa geostrophic frontogenesis (shaded, every 0.5°C 100 km-1 6 h-1, according to scale on right), potential temperature (every 2°C). (b) As in (a) except using ageostrophic wind. Solid line in (b) indicates composite cold front position. Wind vectors according to scale in (a). 122

a) b)

2 1.5 1 0.5 : 10 m s-1

c) d)

Figure 3.38. NARR surface frontogenesis composites at hour 0. (a) Total frontogenesis (shaded every 0.5°C 100 km-1 6 h-1 according to scale on right), potential temperature (every 3°C). (b) Confluent and convergent frontogenesis according to scale in (a). (c) Diabatic frontogenesis according to scale in (a). (d) Diabatic temperature changes. Solid line in (d) indicates position of composite cold front. Wind vectors according to scale in (a). 123

a) b)

2 1.5 1 0.5 : 10 m s-1

c) d)

Figure 3.39. NARR 700-hPa frontogenesis composites at hour 0. (a) Total frontogenesis (shaded every 0.5°C 100 km-1 6 h-1 according to scale on right), potential temperature (every 2°C). (b) Confluent and convergent frontogenesis according to scale in (a). (c) Tilting frontogenesis according to scale in (a). (d) Diabatic frontogenesis according to scale in (a). Wind vectors according to scale in (a). 124 strong diabatic postfrontal cooling immediately behind the cold front, which is consistent with cooling associated with evaporation and sublimation of precipitation (not shown).

There is again little contribution from tilting frontogenesis (Fig. 3.39c).

By hour +6, which represents a composite time of 0300 MST, the northern half of the surface cold front is moving rapidly into eastern Wyoming, but the southern portion of the front has slowed over extreme southern Utah (Fig. 3.40a), and the surface wind shift is located ahead of northern portion of the cold front. Total frontogenesis is becoming unor- ganized (Fig. 3.40a) with numerous FW and FD maximum (Figs. 3.40b,c). The disorgani- zation of FD results as nocturnal cooling develops (Fig. 3.40d).

Discussion and Conclusions

This chapter presents a climatology of strong cold fronts over the western United

States and identifies them as having: a surface temperature drop of at least 7°C in less than

3 h, a concurrent 3 hPa or greater altimeter pressure rise, and 700-hPa baroclinity greater than or equal to 6°C 500 km-1. Strong cold fronts defined in this manner are primarily continental, with few events occurring near coastal locations due to weak temperature falls. A significant diurnal and seasonal cycle of strong cold fronts is observed, with most events occurring between the late afternoon to early evening and during the late spring and early summer. The direction from which cold fronts come from varies geographically, and in many instances is related to the regional terrain.

The frequency of strong cold fronts increases downstream (eastward) from the

Sierra Nevada and Cascade ranges, reaching a maximum at Salt Lake City. A composite study shows most strong Intermountain cold fronts develop as follows. An amplifying 125

a) b)

2 1.5 1 0.5 : 10 m s-1

c) d)

Figure 3.40. NARR surface frontogenesis composites at hour +6. (a) Total frontogenesis (shaded every 0.5°C 100 km-1 6 h-1 according to scale on right), potential temperature (every 3°C). (b) Confluent and convergent frontogenesis according to scale in (a). (c) Diabatic frontogenesis according to scale in (a). (d) Diabatic temperature changes. Wind vectors according to scale in (a). Solid line in (d) indicates position of composite cold front. 126 upper-level trough approaches a preexisting southwest-northeast oriented baroclinic zone, a pattern that is frequently associated with large-scale confluent frontogenesis (e.g., Pet- terssen 1956, his Fig. 11.9.3); this pattern also resembles that of a developing midlatitude frontal cyclone along a stationary front (Wallace and Hobbs 1977, p. 262). The incipient cold front develops along convergent flow that extends downstream from the Sierra

Nevada across central Nevada. Strong frontogenesis results as cold advection approaches the quasi-stationary front from the west while broad ageostrophic flow impinges from the southeast. The cold front further intensifies as clouds and precipitation, which are forced by large-scale ascent associated with the upper-level trough and strong frontogenesis, develop at and behind the cold front. In many cases, diabatic processes strengthen the cold front where postfrontal cold air is enhanced by diabatic cooling from evaporation, sublimation and cloud shading, and the prefrontal environment experiences strong surface sensible heating. This cloud and precipitation configuration results in the strongest diabatic frontogenesis at the front during the daytime, which supports the observed peak occur- rence of strong cold fronts during the late afternoon and evening. This description of Inter- mountain cold front development is qualitatively similar to that offered by Hoffman

(1995, his Fig. 114).

The diabatic intensification of cold fronts during the daytime over the western

United States is supported by other studies. Indirect evidence is provided by Wiesmueller

(1982) who found that cold frontal passages at Denver, CO are most common during the evening and early night. Direct evidence is offered by Sanders (1999b), who showed rapid daytime development of a cold front in the southwestern United States. 127

Cloud observations confirm the importance of diabatic processes on frontogenesis.

Diabatic frontogenesis is, in part, maximized along the front during the day as a result of cross-front sensible heating contrasts that arise from differential cloud thickness where the postfrontal environment features thick mid- and low-level cloud and the prefrontal environment features relatively thin high clouds. This configuration is highly frontogenetic during the daytime since postfrontal cloud cover decreases sensible heating, whereas the prefrontal environment remains strongly heated, as described by Koch et al. (1995, 1997),

Gallus and Segal (1999), and Segal et al. (2004). The importance of differential sensible heating on cold frontal strength is also supported by Bosart et al. (1972), Physick (1988), and Garratt (1988), who show that coastal fronts reach peak intensity with daytime heating. In these coastal cases, however, the differential sensible heat flux arises mainly from contrasts in the latent heat of evaporation over the land and sea.

A secondary contributor to diabatic frontogenesis is the cooling of postfrontal air by evaporation and sublimation. Although this effect remains unquantified, the abundance of observed postfrontal precipitation within a previously dry boundary layer suggests that evaporation and sublimation are important, as shown for one Intermountain cold front by

Schultz and Trapp (2003). Furthermore, studies of evaporation and frontal development suggest that postfrontal evaporative cooling results in a strengthening and acceleration of cold fronts (Oliver and Holzworth 1953; Bannon and Mak 1984; Huang and Emanuel

1991; Lagouvardos et al. 1993; Katzfey and Ryan 1997), which is consistent with the rapid acceleration of our composite cold front from hour -6 to 0. A mechanism as to how

evaporation may create or modify cold fronts is offered by Seitter and Muench (1985)

who suggest that evaporation beneath a wide frontal rainband evaporates, and produces a 128 low-level cold pool that spreads as a gravity current. This mechanism is similar to that in mesoscale convective systems where the mesoscale pressure dome can be hydrostatically linked to the magnitude of subcloud evaporative cooling (Fujita 1959). Nonetheless, it remains for future study to fully investigate the role of evaporation and sublimation on the development of Intermountain cold fronts, but evidence suggests its magnitude may be greater than that produced by clouds.

Diabatic frontogenesis may not always be directly important for the development of strong Intermountain cold fronts since about 20% of Intermountain events feature no precipitation during their history and likely little cross-front cloud contrasts. This suggests that, in some cases, strong Intermountain cold fronts develop without the direct assistance of diabatic processes. In these dry cases, however, diabatic processes may have an indirect effect on frontogenesis since diabatic heating results in the development of a convective boundary layer. The convective boundary layer may in turn influence the ageostrophic flow, which accounts for more than half of the observed FW. The origins of this ageostrophic flow remain unclear, but it may be related to low-level frictional convergence (Keyser and Anthes 1982), frontal circulations that arise from frontogenesis, or boundary layer effects of directional wind shear with height (Snyder 1998).

A potential analog to Intermountain cold fronts is continental Australian cold

fronts. Paradoxically, observational studies of subtropical Australian cold fronts have

shown they intensify during the night and weaken during the day (Smith et al. 1995). This

is supported by a numerical study of homogeneous heating within a convective boundary

layer by Reeder and Tory (2005), who find that homogeneous daytime heating is slightly

frontolytic while nighttime cooling is frontogenetic. The discrepancy in the diurnal 129 strength of Intermountain and continental Australian cold fronts may be explained by the presence of stronger upper-level troughs with Intermountain cold fronts, which results in a higher frequency of clouds and precipitation. This results in Intermountain cold fronts having strong cross-front diabatic heating gradients, while Australian cold fronts feature homogenous cross-front sensible heating.

This study does not fully quantify the effect of diabatic processes on Intermountain frontogenesis, and it does not describe or explain the small-scale (<100 km) frontal structure, which in many cases resemble gravity currents. Nonetheless, the following conclusions are drawn from this study.

• The NARR dataset effectively resolves synoptic and mesoscale dynamic structures,

but it poorly resolves Intermountain precipitation and surface temperatures.

• As defined in this study, strong cold fronts over the western United States occur prima-

rily over interior regions.

• Strong cold fronts develop over the Intermountain region downstream of the Sierra

Nevada and Oregon Cascades, and are most frequent during the late afternoon and

early evening.

• The intensification of most strong Intermountain cold fronts is associated with an

amplifying upper-level trough.

• The rapid intensification of many strong Intermountain cold fronts is enhanced by dia-

batic processes, which are produced by differential sensible heating due to cloud shad-

ing and postfrontal diabatic cooling by evaporation and sublimation.

• Low-level ageostrophic flow contributes to the majority of the observed frontogenesis. CHAPTER 4

SUMMARY AND CONCLUSIONS

This dissertation has examined the effects of complex terrain and dry boundary layers on synoptic baroclinic systems over the western United States. This was accomplished by an observational study of IOP3 of the Intermountain Precipitation Experiment, and by developing a climatology of strong Intermountain cold fronts.

The analysis of IOP3, which featured a mobile, weakening frontal cyclone, illustrates the importance of terrain on baroclinic storm evolution and Intermountain precipitation systems. These results show that low-level frontal structures undergo a dramatic transformation from the Pacific coast to the Intermountain region as a result of interactions with Sierra Nevada. This transformation occurs as the Sierra Nevada block and dis- tort low-level fronts, essentially destroying or stripping the low-level baroclinity of transient baroclinic systems. In the lee of the Sierra, a discontinuous surface evolution is observed where a lee pressure trough develops, and then moves rapidly downstream with the strongest upper-level vertical motion.

Results from the IOP3 analysis also show that mesoscale (5–50 km) terrain-modified flows have a significant effect on Intermountain precipitation. The terrain may enhance or curtail precipitation in lowlands distant from the terrain producing them, 131 resulting in large deviations from the climatological quasi-linear precipitation-elevation relationship. These terrain-induced deviations, which vary with flow direction, stability, and humidity, have significant implications for operational weather predication. These results show that the construction of precipitation anomaly maps may help precipitation prediction in complex Intermountain terrain by identifying locations that may be biased towards an increased or decreased precipitation during different storm environments or evolutions.

The climatology shows that the frequency of strong cold fronts increases east of the Sierra Nevada, reaching a maximum at Salt Lake City. Thus, under the right conditions, cold fronts do not die in the Intermountain region, they are reborn. In most cases, the genesis of strong Intermountain cold fronts occurs as large-scale confluent frontogenesis and diabatic processes act in concert. Diabatic processes are highly frontogenetical during the daytime, and arise from differential sensible heating caused by cloud shading and diabatic postfrontal cooling from evaporation and sublimation. These clouds and precipitation are largely forced by an amplifying upper-level trough. Results from this climatology should lead to better short-range forecasting of strong Intermountain cold fronts, and provide a foundation for future studies of strong cold fronts across the western United States.

A synthesis of this dissertation suggests the following relationship over the Inter- mountain West. The large-scale structures of baroclinic systems are determined by synoptic-scale dynamics, but the mesoscale and low-level structures of Intermountain baroclinic systems are controlled by interactions with the terrain and dry boundary layer.

Future work should better quantify the extent of diabatic processes on frontal development, including the role of evaporation and sublimation. Future work should also 132 determine the origins of low-level ageostrophic flow, which is responsible for much of the frontal development. These tasks may be accomplished through a combination of case studies and numerical simulations. REFERENCES

Alter, J. C., 1919: Normal precipitation in Utah. Mon. Wea. Rev., 47, 633-636.

Astling, E. G., 1984: On the relationship between diurnal mesoscale circulations and precipitation in a mountain valley. J. Clim. Appl. Meteor., 23, 1635-1644.

Baer, V. E., 1991: The transition from the present radar dissemination system to the NEXRAD Information Dissemination Service (NIDS). Bull. Amer. Meteor. Soc., 72, 29-33.

Bannon, P. R., and M. Mak, 1984: Diabatic quasi-geostrophic surface frontogenesis. J. Atmos. Sci., 41, 2189-2201.

Benjamin, S. G., and Coauthors, 2004: An hourly assimilation-forecast cycle: The RUC. Mon. Wea. Rev., 132, 495-518.

Blazek, T. R., 2000: Analysis of a Great Basin cyclone and attendant mesoscale features. M.S. thesis, Dept. of Meteorology, University of Utah, 122 pp. [Available from Dept. of Meteorology, University of Utah, 145 South 1460 East, Room 819, Salt Lake City, UT 84112-0110.]

Bluestein, H. B., 1986: Fronts and jet streaks: A theoretical perspective. Mesoscale Meteorology and Forecasting. (P. S. Ray, Ed.), Amer. Meteor. Soc., 173-215.

Bluestein, H. B., 1993: Synoptic-Dynamic Meteorology in Midlatitudes. Volume II. Oxford University Press. New York. 594 p.

Blumen, W., 1992: Propagation of fronts and frontogenesis versus frontolysis over orography. Meteor. Atmos. Phys., 48, 37-50.

Bosart, L. F., C. J. Vaudo, and J. H. Helsdon Jr., 1972: Coastal frontogenesis. J. App. Met., 11, 1236-1258.

Braun, S. A., R. A. Houze Jr., and B. F. Smull, 1997: Airborne dual-doppler observations of an intense frontal system approaching the Pacific Northwest coast. Mon. Wea. Rev., 125, 3131-3156. 134

Chien, F., C. F. Mass, and P. J. Neiman, 2001: An observational and numerical study of an intense landfalling cold front along the northwest coast of the United States during COAST IOP 2. Mon. Wea. Rev., 129, 934-955.

Cheng, L., 2001: Validation of quantitative precipitation forecasts during the Intermountain Precipitation Experiment. M.S. thesis, Dept. of Meteorology, University of Utah, 137 pp. [Available from Dept. of Meteorology, University of Utah, 145 South 1460 East, Rm. 819, Salt Lake City, UT 84112-0110.]

Colle, B. A., C. F. Mass, and B. F. Smull, 1999: An observational and numerical study of a cold front interacting with the Olympic Mountains during COAST IOP5. Mon. Wea. Rev., 127, 1310-1334.

Colle, B. A., J. B Wolfe, W. J. Steenburgh, D. E. Kingsmill, J. A. W. Cox, and J. C. Shafer, 2005: High resolution simulations and microphysical validation of an orographic precipitation event over the Wasatch Mountains during IPEX IOP3. Mon. Wea. Rev., 133, 2947-2971.

Cox, J. A.W., W. J. Steenburgh, D. E. Kingsmill, J. C. Shafer, B. A. Colle, O. Bousquet, B. F. Smull, and H. Cai, 2005: The kinematic structure of a Wasatch Mountain winter storm during IPEX IOP3. Mon. Wea. Rev., 133, 521-542.

Daly, C., R. P. Neilson, and D. L. Phillips, 1994: A statistical-to-topographic model for mapping climatological precipitation over mountainous terrain. J. Appl. Met., 33, 140-158.

Darby, L. S., W. D. Neff, and R. M. Banta, 1999: Multiscale analysis of a meso-b frontal passage in the complex terrain of the Colorado Front Range. Mon. Wea. Rev., 127, 2062-2082.

Davis, R. E., and D. R. Walker, 1992: An upper-air synoptic climatology of the western United States. J. Climate., 5, 1449-1467.

Dickinson, M. J., D. Knight, 1999: Frontal interaction with mesoscale topography. J. Atmos. Sci., 56, 3544-3559.

Egger, J., and K. P. Hoinka, 1992: Fronts and orography. Meteor. Atmos. Phys., 48, 3-36.

Eliassen, A., 1959: On the formation of fronts in the atmosphere. The Atmosphere and the Sea in Motion, B. Bolin, Ed., Rockefeller Institute Press, 277-287.

Eliassen, A., 1962: On the vertical circulation in frontal zones. Geofys. Publ., 24, 147-160.

Elliott, R. D., 1958: California storm characteristics and weather modification. J. Meteor. 15, 486-493. 135

Fujita, T., 1959: Precipitation and cold air production in mesoscale thunderstorm systems. J. Meteor., 16, 454-466.

Gilmore, M. S., and L. J. Wicker, 1998: The influence of midtropospheric dryness on supercell morphology and evolution. Mon. Wea. Rev., 126, 943-958.

Gallus, W. A. Jr., and M. Segal, 1999: Diabatic effects on late-winter cold front evolution: Conceptual and numerical model evaluations. Mon. Wea. Rev., 127, 1518-1537.

Garratt, J. R, 1988: Summertime cold fronts in southeast Australia-behavior and low-level structure of main frontal types. Mon. Wea. Rev., 116, 636-649.

Grubisic , V., 2004: Bora-driven potential vorticity banners over the Adriatic. Quart. J. Roy. Meteor. Soc., 130, 2571-2603.

Hakim, G. J., 1992: The eastern United States side-door cold front of 22 April 1987: A case study of an intense atmospheric density current. Mon. Wea. Rev., 120, 2738- 2762.

Hess, S. L., and H. Wagner, 1948: Atmospheric waves in the northwestern United States. J. Meteor., 5, 1-19.

Hevesi, J. A., A. L. Flint, and J. D. Istok, 1992: Precipitation estimation in mountainous terrain using multivariate geostatistics. Part II: Isohyetal Maps. J. Appl. Meteor., 31, 677-688.

Hobbs, P. V., T. J. Matejka, P. H. Herzegh, J. D. Locatelli, and R. A. Houze Jr., 1980: The mesoscale and microscale structure and organization of clouds and precipitation in midlatitude cyclones. Part I: A case study of a cold front. J. Atmos. Sci., 37, 568- 596.

Hobbs, P. V. and P. O. G. Persson, 1982: The mesoscale and microscale structure and organization of clouds and precipitation in midlatitude cyclones. Part V: The sub- structure of narrow cold-frontal rainbands. J. Atmos. Sci., 39, 280-295.

Hobbs, P.V., J.D. Locatelli, and J. E. Martin, 1996: A new conceptual model for cyclones generated in the lee of the Rocky Mountains. Bull. Amer. Meteor. Soc., 77, 1169- 1178.

Hoffman, E. G., 1995: Evolution and mesoscale structure of fronts in the western United States: A case study. M.S. thesis, Dept. of Earth and Atmospheric Sciences, State University of New York at Albany, 287 pp.

Horel, J. D., and C. V. Gibson, 1994: Analysis and simulation of a winter storm over Utah. Wea. Forecasting, 9, 479-494. 136

Horel, J. D., M. Splitt, L. Dunn, J. Pechmann, B. White, C. Ciliberti, S. Lazarus, J. Slemmer, D. Zaff, and J. Burks, 2002: Mesowest: Cooperative mesonets in the western United States. Bull. Amer. Meteor. Soc., 83, 211-225.

Huang, H., and K. Emanuel, 1991: The effects of evaporation on frontal circulations. J. Atmos. Sci., 48, 619-628.

Katzfey, J. J., and B. F. Ryan, 1997: Modification of the thermodynamic structure of the lower troposphere by the evaporation of precipitation: A GEWEX cloud system study. Mon. Wea. Rev., 125, 1431-1446.

Keyser, D. and R. A. Anthes, 1982: The influence of planetary boundary layer physics on frontal structure in the Hoskins-Bretherton horizontal shear model. J. Atmos. Sci., 39, 1783-1802.

Keyser, D., 1986: Atmospheric fronts: An observational perspectives. Mesoscale Meteorology and Forecasting. (P. S. Ray, Ed.), Amer. Meteor. Soc., 216-258.

Koch, S. E, J. T. McQueen, and V. M. Karyampudi, 1995: A numerical study of the effects of differential cloud cover on cold frontal structure and dynamics. J. Atmos. Sci., 52, 937-964.

Koch, S. E., A. Aksakal, and J. T. McQueen, 1997: The influence of mesoscale humidity and evapotranspiration fields on a model forecast of a cold-frontal squall line. Mon. Wea. Rev., 125, 384-409.

Koppel, L. L, L. F. Bosart, and D. Keyser, 2000: A 25-yr climatology of large-amplitude hourly surface pressure changes over the conterminous United States. Mon. Wea. Rev., 128, 51-68.

Lazarus, S. M., C. M. Ciliberti, J. D. Horel and K. A. Brewster, 2002: Near-real-time applications of a mesoscale analysis system to complex terrain. Wea. Forecasting, 17, 971-1000.

Long, A. B., A. W. Huggins, and B. A. Campistron, 1990: Investigations of a winter mountain storm in Utah. Part I: Synoptic analyses, mesoscale kinematics, and water release rates. J. Atmos. Sci., 47, 1302-1322.

Lagouvardos, K., Y. Lemaitre, and G. Scialom, 1993: Importance of diabatic processes on ageostrophic circulations observed during the FRONTS 87 experiment. Quart. J. Roy. Meteor. Soc., 119, 1321-1345.

Marwitz, J., and J. Toth, 1993: The front range blizzard of 1990. Part I: Synoptic and mesoscale structure. Mon. Wea. Rev., 121, 402-415. 137

Mass, C. F., W. J. Steenburgh, and D. M. Schultz, 1991: Diurnal surface-pressure variations over the continental United States and the influence of sea level reduction. Mon. Wea. Rev., 119, 2814-2830.

Mesinger, F., and Coauthors, 2004: NCEP North American Regional Reanalysis. 15th Symp. on Global Change and Climate Variations, Poster 1.1, Combined Preprints CD-ROM, 84th Annual Meeting, Seattle, WA, Jan. 2004.

Miller, J. E., 1948: On the concept of frontogenesis. J. Meteor., 5, 169-171.

NCDC, 1995: Storm Data. Vol. 37, No. 6, 396 pp.

NCDC, 2000: Storm Data. Vol. 42, No. 2, 127 pp.

Neiman, P. J. and R. M. Wakimoto, 1999: The interaction of a pacific cold front with shallow air masses east of the Rocky Mountains. Mon. Wea. Rev., 127, 2102-2127.

Neiman, P. J., F. M. Ralph, R. L. Weber, T. Uttal, L. B. Nance, and D. H. Levinson, 2001: Observations of nonclassical frontal propagation and frontally forced gravity waves adjacent to steep topography. Mon. Wea. Rev., 129, 2633-2659.

Oliver, V. J., and G. C. Holzworth, 1953: Some effects of the evaporation of widespread precipitation on the production of fronts and on changes in frontal slopes and motions. Mon. Wea. Rev., 81, 141-151.

Peck, E. L., and M. J. Brown, 1962: An approach to the development of isohyetal maps for mountainous areas. J. Geophys. Res., 67, 681-694.

Petterssen, S., 1936: Contribution to the theory of frontogenesis. Geofys. Publ., 11, (6), 1- 27.

Petterssen, S., 1956: Weather Analysis and Forecasting. Vol. 1, Motion and Motion Sys- tems, 2d ed. McGraw-Hill. New York. 428 p.

Physick, W. L., 1988: Mesoscale modeling of a cold front and its interaction with a diur- nally heated land mass. J. Atmos. Sci., 45, 3169-3187.

Reeder, M. J., 1986: The interaction of a surface cold front with a prefrontal thermody- namically well-mixed boundary layer. Aust. Met. Mag., 34, 137-148.

Reeder, M. J., and K. J. Tory, 2005: The effect of the continental boundary layer on the dynamics of fronts in a two-dimensional model of baroclinic instability. Part II. Surface heating and cooling. Quart. J. Roy. Meteor. Soc., 131, 2409-2430.

Reynolds, D. W., and A. P. Kuciauskas, 1988: Remote and in situ observations of Sierra Nevada winter mountains clouds: Relationship between mesoscale structure, 138

precipitation and liquid water. J. Appl. Meteor., 27, 140-156.

Sanders, F., and Doswell C. A. III, 1995: A case for detailed surface analysis. Bull. Amer. Meteor. Soc., 76, 505-521.

Sanders, F., 1999a: A proposed method of surface map analysis. Mon. Wea. Rev., 127, 945-955.

Sanders, F., 1999b: A short-lived cold front in the southwestern United States. Mon. Wea. Rev., 127, 2395-2403.

Sawyer, J. S., 1956: The vertical circulation at meteorological fronts and its relation to frontogenesis. Proc. Roy. Soc. London., A234, 346-362.

Schär, C., M. Sprenger, D. Lüthi, Q. Jiang, R. B. Smith, and R. Benoit, 2003: Structure and dynamics of an Alpine potential-vorticity banner. Quart. J. Roy. Meteor. Soc., 129, 825-855.

Schultz, D. M., and C. A. Doswell III, 2000: Analyzing and forecasting Rocky Mountain lee cyclogenesis often associated with strong winds. Wea. Forecasting, 15, 152- 173.

Schultz, D. M., W. J. Steenburgh, R. J. Trapp, J. Horel, D. E. Kingsmill, L. B. Dunn, W. D. Rust, L. Cheng, A. Bansemer, J. Cox, J. Daugherty, D. P. Jorgensen, and J. Meitín, L. Showell, B. F. Smull, K. Tarp, and M. Trainor, 2002: Understanding Utah winter storms: The Intermountain Precipitation Experiment. Bull. Amer. Meteor. Soc., 83, 190-210.

Schultz, D. M, and R. J. Trapp, 2003: Nonclassical cold-frontal structure caused by dry subcloud air in northern Utah during the Intermountain Precipitation Experiment (IPEX). Mon. Wea. Rev., 131, 2222-2246.

Schumacher, P. N., D. J. Knight, and L. F. Bosart, 1996: Frontal interaction with the Appalachian Mountains. Part I: A climatology. Mon. Wea. Rev., 124, 2453-2468.

Segal, M., E. A. Aligo, and W. A. Gallus Jr., 2004: A conceptual and scaling evaluation of the surface wetness effect on daytime moisture convergence along a surface cold front with differential cloud cover. J. Hydromet., 5, 365-371.

Seitter, K. L, and H. S. Muench, 1985: Observation of a cold front with rope cloud. Mon. Wea. Rev., 113, 840-848.

Shafer, J. C., 2002: Synoptic and mesoscale structure of a Wasatch Mountain winter storm. M. S. thesis, Dept. of Meteorology, University of Utah, 65 pp. [Available from Dept. of Meteorology, University of Utah, 135 South 1460 East, Rm. 819, Salt Lake City, UT 84112-0110]. 139

Shafer, J. C., 2004: A different beast: An example of a major midlatitude cyclogenesis event over the Intermountain region of the western United States. 20th Conference on Weather Analysis and Forecasting, Seattle WA, Jan. 2004.

Shafer, J. C., W. J. Steenburgh, J. A. W. Cox, and J. P. Monteverdi, 2005: Terrain influences on synoptic storm structure and mesoscale precipitation distribution during IPEX IOP3. Mon. Wea. Rev., In press.

Shapiro, M. A., T. Hampel, D. Rotzoli, and F. Mosher, 1985: The frontal hydraulic head: A micro-alpha (~1km) triggering mechanism for mesoconvective weather systems. Mon. Wea. Rev., 113, 1166-1183.

Smith, R., K., and M. J. Reeder, 1988: On the movement and low-level structure of cold fronts. Mon. Wea. Rev., 116, 1927-1944.

Smith, R. K., M. J. Reeder, N. J. Tapper, and D. R. Christie, 1995: Central Australian cold fronts. Mon. Wea. Rev., 123, 16-38.

Snyder, C., 1998: Approximate dynamical equations for fronts modified by the planetary boundary layer. J. Atmos. Sci., 55, 777-787.

Spencer, P. L., Carr, F. H., and Doswell, C. A. 1996: Diagnosis of an amplifying and decaying baroclinic wave using wind profiler data. Mon. Wea. Rev., 124, 209-223.

Splitt, M. E., and J. D. Horel, 1998: Use of multivariate linear regression for meteorological data analysis and quality assessment in complex terrain. Preprints, 10th Symp. on Meteorological Observations and Instrumentation, Phoenix, AZ, Amer. Meteor. Soc., 359-362.

Steenburgh, W. J., and T. R. Blazek, 2001: Topographic distortion of a cold front over the Snake River Plain and central Idaho Mountains. Wea. Forecasting, 16, 301-314.

Steenburgh, W. J., 2003: One hundred inches in one hundred hours: Evolution of a Wasatch Mountain winter storm cycle. Wea. Forecasting, 18, 1018-1036.

Stegner, W., 1992: Where the Bluebird Sings to the Lemonade Springs. Wing Books. New York. 227 p.

Stewart, J. Q., C. D. Whiteman, W. J. Steenburgh, and X. Bian, 2002: A climatological study of thermally driven wind systems of the U.S. Intermountain West. Bull. Amer. Meteor. Soc., 83, 699-708.

Stone, P. H., 1966: Frontogenesis by horizontal wind deformation fields. J. Atmos. Sci., 23, 455-465. 140

Tory, K. J. and M. J. Reeder, 2005: The effect of the continental boundary layer on the dynamics of fronts in a two-dimensional model of baroclinic instability. Part I. An insulated lower surface. Quart. J. Roy. Meteor. Soc., 131, 2389-2408.

Wakimoto, R. M., 1982: The life cycle of thunderstorm gust fronts as viewed with doppler radar and rawinsonde data. Mon. Wea. Rev., 110, 1060-1082.

Wallace, J. M., and P. V. Hobbs, 1977: Atmospheric Science. Academic Press Inc. San Diego. 467 p.

Wiesmueller, J. L., 1982: The effect on diurnal heating on the movement of cold fronts through eastern Colorado. NOAA Technical Memorandum, NWS, CR-66.

Williams, P. Jr., 1972: Western Region synoptic analysis–Problems and methods. NOAA NWS Western Region Tech. Memo. NWSTM WR-71, 71 pp. [Available from NOAA NWS Western Region Headquarters, 125 S. State St., Rm. 1311, Salt Lake City, UT 84138-1102.].

Westrick, K. J., C. F. Mass, and B. A. Colle, 1999: The limitations of the WSR-88D radar network for quantitative precipitation measurement over the coastal western United States. Bull. Amer. Meteor. Soc., 80, 2289-2298.

Wood, V. T., R. A. Brown, and S. V. Vasiloff, 2003: Improved detection using negative elevation angles for mountaintop WSR-88Ds. Part II: Simulations of the three radars covering Utah. Wea. Forecasting, 18, 393-403.

Yu, C. and B. F. Smull, 2000: Airborne doppler observations of a landfalling cold front upstream of steep coastal orography. Mon. Wea. Rev., 128, 1577-1603.