The Interaction of a Pacific Cold Front with Shallow Air Masses East of the Rocky Mountains

Paul J. Neiman NOAA/Environmental Technology Laboratory, Boulder, Colorado

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Roger M. Wakimoto Department of Atmospheric Sciences, University of California, Los Angeles, Los Angeles, California

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Abstract

The study presented here describes the interactions that occurred between an advancing Pacific cold front and shallow Gulf of Mexico and Arctic air masses situated east of the Rocky Mountains during the Verification of the Origins of Rotation in Tornadoes Experiment (VORTEX) field campaign on 17–18 April over Oklahoma and adjacent states. These interactions were driven largely by the complex topography of the region. Four air masses of distinctly different origin (i.e., Pacific polar, high-altitude continental, Gulf of Mexico, and Arctic), and the boundaries that separated them (i.e., Pacific cold front, dryline, and Arctic front), were observed within the experimental domain. This event produced more than $1 million worth of damage in the experimental domain due to severe weather. A dense network of ground-based in situ and remote observing systems and two research aircraft equipped with in situ sensors and Doppler radars gathered data that allowed the authors to document the passage of a vigorous midtropospheric shortwave trough and associated Pacific cold front, and the interaction of this front with the preexisting Gulf of Mexico and Arctic air masses. The Pacific front intersected the ground to the west of the Arctic frontal boundary and dryline, and subsequently rode over the top of the Gulf of Mexico and Arctic air masses. This study also presents the detailed observational documentation of a dryline–frontal merger by showing the merging or phasing of updrafts associated with the Pacific front and dryline and the subsequent development of a squall line. The behavior of the Arctic front is also explored in detail. Its anomalous southward penetration into the VORTEX domain due to terrain-induced blocking also played a role in producing severe weather.

Corresponding author address: Paul J. Neiman, NOAA/ETL, Mail Code R/E/ET7, 325 Broadway, Boulder, CO 80303-3328.

Email: pneiman@etl.noaa.gov

Abstract

The study presented here describes the interactions that occurred between an advancing Pacific cold front and shallow Gulf of Mexico and Arctic air masses situated east of the Rocky Mountains during the Verification of the Origins of Rotation in Tornadoes Experiment (VORTEX) field campaign on 17–18 April over Oklahoma and adjacent states. These interactions were driven largely by the complex topography of the region. Four air masses of distinctly different origin (i.e., Pacific polar, high-altitude continental, Gulf of Mexico, and Arctic), and the boundaries that separated them (i.e., Pacific cold front, dryline, and Arctic front), were observed within the experimental domain. This event produced more than $1 million worth of damage in the experimental domain due to severe weather. A dense network of ground-based in situ and remote observing systems and two research aircraft equipped with in situ sensors and Doppler radars gathered data that allowed the authors to document the passage of a vigorous midtropospheric shortwave trough and associated Pacific cold front, and the interaction of this front with the preexisting Gulf of Mexico and Arctic air masses. The Pacific front intersected the ground to the west of the Arctic frontal boundary and dryline, and subsequently rode over the top of the Gulf of Mexico and Arctic air masses. This study also presents the detailed observational documentation of a dryline–frontal merger by showing the merging or phasing of updrafts associated with the Pacific front and dryline and the subsequent development of a squall line. The behavior of the Arctic front is also explored in detail. Its anomalous southward penetration into the VORTEX domain due to terrain-induced blocking also played a role in producing severe weather.

Corresponding author address: Paul J. Neiman, NOAA/ETL, Mail Code R/E/ET7, 325 Broadway, Boulder, CO 80303-3328.

Email: pneiman@etl.noaa.gov

1. Introduction

The juxtaposition of the Rocky Mountains, Gulf of Mexico, and Great Plains leads to a variety of regional weather phenomena over the central United States that cannot be adequately portrayed by the landmark Norwegian frontal-cyclone conceptualization (e.g., Bjerknes 1919; Bjerknes and Solberg 1922). These phenomena, which can significantly impact population centers in the region, include cold-air damming of shallow (≲2 km deep) Arctic air on the east side of the Rocky Mountains (e.g., Mecikalski and Tilley 1992; Colle and Mass 1995), and drylines that form east of the Rockies at the confluence of dry flow descending from the Mexican Plateau and moist but shallow (≲3 km deep) flow originating over the Gulf of Mexico (e.g., Carlson et al. 1983; Schaefer 1986; Parsons et al. 1991). These air masses originating over the Arctic and Gulf typically have a potential density that is substantially greater than transient air masses and frontal zones aloft. Hence, the stable stratification capping these shallow air masses reduces the intensity of vertical turbulent-scale mixing across the cap (e.g., Panofsky and Dutton 1984, 144–148) thus decoupling air masses and frontal zones aloft from the surface (e.g., Neiman et al. 1990, 1998). In principle, the top of the shallow Arctic or Gulf air mass acts as the proxy ground level for those air masses and fronts situated aloft (e.g., Neiman et al. 1998). The concept of fronts riding over the top of shallow stable layers, without exchanging air masses but still producing wind shifts through pressure changes at the surface, quite likely explains some of the difficulty encountered in analyzing surface fronts over the United States (e.g., Mass 1991; Uccellini et al. 1992; Sanders and Doswell 1995; Hobbs et al. 1996; Neiman et al. 1998).

Significant weather-producing features that can be masked at the surface by these shallow air masses are cold fronts whose cold sectors originate in the polar air stream over the maritime Pacific Ocean, that is, Pacific cold fronts. Because Pacific fronts and their trailing maritime air masses undergo sensible heating and adiabatic warming while traversing and descending the Rocky Mountains, they typically become potentially warmer than the trapped shallow air masses east of the Rockies and thus cannot penetrate to the surface to produce a significant thermodynamic signature there [though Steenburgh and Mass (1994) show a case to the contrary]. Despite the fact that Pacific fronts often do not possess a thermodynamic signature within Arctic and Gulf air masses, a weak pressure trough and slight wind-direction shift [not associated with a lee trough, e.g., Carlson (1961); Steenburgh and Mass (1994)] at the surface are often observed, as may be expected from the fact that the surface pressure and corresponding wind fields can respond to airmass changes aloft (e.g., Locatelli et al. 1997; Neiman et al. 1998). In addition, the interaction of Pacific fronts with shallow air masses can be important in generating significant weather over the central United States. For example, vertical circulations associated with Pacific fronts can advect shallow moisture upward (e.g., Shapiro 1982), contributing to severe weather such as flooding rains (e.g., Lott 1954), tornadic thunderstorms with large hail (e.g., Locatelli et al. 1995; Neiman et al. 1998), and heavy snowstorms (e.g., Snook 1993).

Because Pacific cold fronts can produce a variety of severe weather across the central United States, it is important to understand how these fronts interact with preexisting shallow air masses, and to accurately portray the position of the leading edge of these fronts aloft on surface analyses [as in Hobbs et al. (1990) and Neiman et al. (1998)]. These issues will be addressed in this study using data taken during the Verification of the Origins of Rotation in Tornadoes Experiment (VORTEX; Rasmussen et al. 1994) field campaign of 1995 that was carried out in Oklahoma and adjacent states. The primary objective of VORTEX was to study tornadogenesis, tornado dynamics and kinematics, and the modulation of tornadic storms by environmental conditions. Although the experiment was not designed specifically to study transient Pacific frontal zones and their interactions with shallow air masses east of the Rockies, the intensive observing period of 17–18 April 1995 provided just such an opportunity.

The event of 17–18 April produced more than $1 million damage (NOAA 1995) in the experimental domain due to severe thunderstorms, including one that spawned several tornadoes in the prefrontal environment that were the main focus of the VORTEX mobile deployment teams. (We defer the study of these prefrontal tornadic storms to others.) During this event, a dense network of ground-based in situ and remote observing systems and two research aircraft equipped with in situ sensors and Doppler radars gathered data within the core of the experimental domain. These data allowed us to document, with great clarity, the passage of a vigorous midtropospheric shortwave trough and associated Pacific cold front, and the interaction of this front with the preexisting modified Arctic and Gulf of Mexico air masses. The event’s evolution is explored through the integrated analysis of data gathered using the many tools summarized in section 2 and in Rasmussen et al. (1994) and Wakimoto et al. (1996). The synoptic-scale structure and evolution are described first, followed by a mesoscale description of the event using finescale horizontal analyses and satellite imagery, wind profiler time–height section analyses, surface time series analyses, cross-section analyses based primarily on in situ aircraft data, and airborne pseudo-dual-Doppler wind and reflectivity analyses.

The observations and analyses show that the Pacific front intersected the ground to the west of the Arctic frontal boundary and dryline (i.e., Gulf of Mexico airmass boundary), and subsequently rode over the top of these shallow air masses. Deep convective precipitation erupted when the Pacific front merged with the dryline. The initiation of intense convection during the approach and/or intersection of fronts with drylines has been discussed by numerous investigators (e.g., Koch and McCarthy 1982; Ogura et al. 1982; Schaefer 1986; Doswell 1987; Parsons et al. 1991; Hane et al. 1993), but few observational studies have actually provided detailed measurements documenting the process. Based on an analysis of surface mesonet data, Koch and McCarthy (1982) proposed that the approach of a front can promote frontogenetical circulations along the dryline that enhance upward motion, which, subsequently, initiates convection. The summary article by Keyser and Shapiro (1986) shows analytically derived, deep tropospheric vertical circulations associated with fronts that could conceivably interact with drylines to initiate convection. Our paper provides detailed observational documentation of a dryline–frontal merger by showing the merging or phasing of their respective updrafts and the subsequent development of convective precipitation.

2. Observing systems

During its field phase in the spring of 1994 and 1995, the VORTEX program used an extensive array of mobile and anchored observing platforms across its experimental domain in Oklahoma and adjacent states (Rasmussen et al. 1994; Wakimoto et al. 1996). The systems that were paramount to our study are described below. By incorporating the data from these systems into horizontal, cross section, time series, and time–height section analyses, the event of 17–18 April is explored. The time-to-space adjustment technique of Fujita (1963) was applied to the appropriate data, when necessary, by invoking Taylor’s (1938) hypothesis (i.e., steady-state weather systems propagating at a fixed phase velocity). Errors in this analysis technique arise when weather systems evolve rapidly with time or exhibit vertical, horizontal, or temporal variations in phase velocity. These errors were limited in this case by using observations taken within ∼1–2 h surrounding an analysis time. Table 1 lists those analyses whose data were subjected to time-to-space adjustment.

The National Center for Atmospheric Research’s (NCAR) Electra research aircraft measured nearly continuous (1 s) standard navigational and meteorological parameters, and utilized its new ELDORA (Electra Doppler Radar) X-band (∼3.15 cm) tail-mounted radar system, during a ∼6 h flight on 17–18 April (Fig. 1). The ELDORA system represents a significant advance compared with the tail-mounted radar on the National Oceanic and Atmospheric Administration’s (NOAA) WP–3D aircraft, because this new system has 1) greater accuracy and sensitivity to observe clear-air phenomena, 2) greater spatial resolution in the along-track direction, and 3) a larger Nyquist frequency to observe a larger unambiguous velocity range. ELDORA’s performance characteristics are described in Hildebrand et al. (1994, 1996) and Wakimoto et al. (1996). Single-Doppler velocities measured by ELDORA’s fore and aft radars were interpolated onto a Cartesian grid with horizontal and vertical grid spacing of 600 and 400 m, respectively. The lowest grid level was ∼300 m. A Cressman filter (Cressman 1959) was used in the interpolation process with a radius of influence of 600 m and 400 m in the horizontal and vertical, respectively. The data were processed within Custom Editing and Display of Reduced Information in Cartesian space (CEDRIC) (Mohr et al. 1986). The hydrometeor fall speeds based on reflectivity were estimated from the relationship established by Joss and Waldvogel (1970) with a correction for the effects of air density (Foote and du Toit 1969). The gridded data were synthesized into horizontal wind fields using established dual-Doppler techniques (e.g., Kropfli and Miller 1976). This pseudo-dual-Doppler synthesis employed a boundary condition of w = 0 m s−1 at the lower boundary, but because the lowest grid level was 300 m above ground level (AGL), a fractional lower boundary condition (e.g., Wakimoto and Kingsmill 1991) was assumed. Vertical air motions were derived from the radar-synthesized horizontal wind fields using an upward integration of the anelastic continuity equation. According to Doviak et al. (1976), the estimated error of the vertical velocity field is approximately 1–2 m s−1 at a height of 2 km AGL, and less at lower levels. Based on the criteria established by Carbone et al. (1985), motions having wavelengths of ∼3.6 km or greater are resolvable in this case. A two-step Leise filter (Leise 1982) was applied to the wind synthesis, which damps wavelengths up to 3.6 km and eliminates those less than 2.4 km.

Ground-based in situ observing platforms that were used in this study include the high spatial-resolution (∼25 km) Oklahoma mesonetwork (Brock et al. 1995) with 5-min sampling, the operational Automated Surface Observing System (ASOS; 5-min resolution) and Surface Aviation Observation (SAO; hourly resolution) system from the National Weather Service (NWS), and operational NWS rawinsondes and mobile CLASS (Cross-chain Loran Atmospheric Sounding System) soundings (see Fig. 1 for those sites within the experimental domain) with temporal resolutions from 12 h down to ∼90 min. The primary ground- and space-based remote sensing platforms that were utilized are NOAA’s GOES-9 satellite, the WSR-88D Doppler radar network (Klazura and Imy 1993), and NOAA’s 404-MHz National Profiler Network (NPN; Ralph et al. 1995). The NPN profilers documented both large- and small-scale phenomena because of their high temporal- and vertical-resolution (6–60 min and 250 m, respectively) measurement capabilities. Winds contaminated by migrating birds were identified by techniques described in Wilczak et al. (1995), and the data were subsequently edited. The few remaining data points that were obviously in error due to some other interference or sampling problem were subjectively removed. The 6-min wind data (used to construct the hourly averages) were used to examine finer-scale phenomena than the hourly wind data could resolve, although such wind data can be less accurate than that of the hourly winds (Ralph et al. 1995). The single-station techniques and limitations used for diagnosing thermodynamic (i.e., horizontal temperature gradients and advections) and precipitation information from profiler data were established in Neiman and Shapiro (1989), Hermes (1991), Neiman et al. (1992), Ralph et al. (1995), and Neiman et al. (1998). The profiler sites within the experimental domain (see Fig. 1) were equipped with surface instruments that measured the basic meteorological variables at 6-min intervals. These surface observations made it possible to relate rapid changes at the surface to comparable timescale changes occurring in the wind and precipitation fields aloft.

3. Synoptic overview of 17–18 April 1995

The synoptic conditions of 17–18 April are presented at 500 mb (Fig. 2) and the surface (Fig. 3) to establish the large-scale context for the detailed mesoscale analyses and interpretations that are presented in sections 4–6. This event was characterized by the eastward migration of a vigorous shortwave trough of polar origin from the eastern Pacific Ocean to the southern and central plains where ambient low-level moisture and moderate convective instability were observed. Four air masses of distinctly different origin (i.e., Pacific polar, high-altitude continental,1 Gulf of Mexico, and Arctic), and the boundaries that separated them (i.e., Pacific cold front, dryline, and Arctic front), are described below.

At 1200 UTC 17 April 1995, the shortwave at 500 mb (Fig. 2a) possessed a closed cyclonic circulation and a cold core of <−30°C centered over northern Arizona and strong westerly component flow to its south. This shortwave was marked at its leading edge by a distinctive cold-frontal baroclinic zone (hereinafter referred to as the Pacific cold front), which sloped downward to the surface (Fig. 3a) in southwestern Colorado through New Mexico and into western Texas and northern Mexico. Farther east, a north–south-oriented dryline at the surface separated prefrontal dry air that originated over the high plateau of the southern Rockies from a moist airstream that originated over the Gulf of Mexico. An Arctic frontal boundary extended southward along the eastern slope of the Rocky Mountains, curved eastward over the Texas Panhandle and then northeastward across central Oklahoma. The southward extension of the Arctic air mass east of the Rockies was due to terrain-induced blocking (e.g., Mecikalski and Tilley 1992; Colle and Mass 1995). Widespread precipitation was situated on the cold side of the Arctic front, and an extensive region of weak echoes was located east of the dryline over eastern Texas that resembled a pre-drytrough rainband (Martin et al. 1995). Convective cells situated beneath the unstable cold core of the shortwave aloft dotted the southern Rockies.

A cross section of potential temperature (θ) and alongfront wind speed (u) (Fig. 4; along line AA′ in Figs. 2a and 3a) highlights the vertical continuity of the Pacific front at the leading and trailing edges of the deep-tropospheric, cold-core Pacific air mass. Enhanced baroclinicity and vertical wind shear associated with the Pacific cold front at the leading edge of the polar air mass extended upward from the surface near El Paso, Texas (ELP), to the vertically depressed tropopause at ∼400 mb over Arizona and Nevada (i.e., over INW and DRA). On the cold side of the front, westerly component flow descended the eastern slope of the southern Rockies, suggesting that this postfrontal air mass experienced adiabatic warming. Because this air mass was also dry (mixing ratio q < 4 g kg−1), most of the incoming solar radiation was converted to sensible heat. Southwesterly jet-stream winds exceeded 60 m s−1 above the Pacific cold front. On the warm side of the front, the shallow (<150 mb deep) and moist (q > 12–16 g kg−1) Gulf of Mexico air mass was bounded by a capping stable layer coinciding with a sharp decrease in moisture with height. The southerly flow within the Gulf air mass veered with height to southwesterly in the prefrontal dry air aloft. At this time, the western boundary of the Gulf air mass [i.e., west of Midland, Texas (MAF)] was east of the surface intersection of the Pacific front and marked the surface dryline; that is, the Pacific front and dryline were distinctly separate surface features at this time.

The Pacific cold front and dryline progressed eastward across western Texas during the following 6 h, remaining distinctly separate at the surface at 1800 UTC 17 April (Fig. 3b). By this time strong postfrontal, downgradient westerly (downslope) flow (>10–15 m s−1) had developed in much of New Mexico and throughout western Texas, with damaging winds approaching 30 m s−1 in the region (NOAA 1995). The northern segment of the Pacific front migrated over the top of the northward-retreating Arctic air mass over northeastern New Mexico. Extensive precipitation persisted on the cold side of the Arctic front and to the east of the dryline at 1800 UTC.

By 0000 UTC 18 April, the 500-mb shortwave circulation had moved northeastward to southeastern Colorado (Fig. 2b), with the leading Pacific cold frontal baroclinicity and its associated jet streak (55 m s−1 at Amarillo) situated over western Texas. The cold core of the Pacific polar air mass warmed by ∼4°C during the 12-h period ending at 0000 UTC. The Pacific cold front, which was merging with the dryline in the vicinity of a developing surface cyclone over north-central Texas at this time (Fig. 3c), extended northward over the top of the shallow Arctic air mass in Oklahoma and Kansas. Postfrontal downgradient westerly (downslope) flow strengthened over eastern New Mexico and western Texas between 1800 and 0000 UTC, and a nascent line of deep convective precipitation paralleled the Pacific front near its leading edge over Texas, Oklahoma, and Kansas. Widespread precipitation continued falling into the northward-retreating Arctic air mass.

Between 0000 and 1200 UTC 18 April the shortwave trough at 500 mb continued migrating northeastward into Nebraska (Fig. 2c), and its cold core moderated by an additional 4°C. The Pacific frontal baroclinicity at this level weakened considerably over the experimental domain by 1200 UTC.

4. Mesoscale frontal evolution—A plan view

Two-hourly surface analyses of virtual potential temperature (θυ) and equivalent potential temperature (θe) between 1800 UTC 17 April and 0000 UTC 18 April (Figs. 5 and 6) highlight the 6-h evolution of the four primary air masses and the three boundaries that separated them. The cool (θυ ≲ 300 K) and moist (θe ≃ 310–325 K) Arctic airmass over northeastern New Mexico, the Texas Panhandle, and the northwestern two-thirds of Oklahoma at 1800 UTC (Figs. 5a and 6a) retreated northward in the following 6 h (Figs. 5b–d, 6b–d), while eroding on its western and eastern flanks over the Texas Panhandle and central Oklahoma, respectively. Easterly component flow was observed within this shallow air mass, primarily east of the Pacific front aloft. At the periphery of this cold dome, collocated horizontal gradients of θυ and θe were concentrated in a 25 to 50 km wide band that defined the Arctic frontal zone. An extreme θe gradient developed within this zone over the Texas Panhandle by 2200 UTC (Fig. 6c) and persisted through 0000 UTC (Fig. 6d). Though this tight surface gradient strongly resembles that associated with pronounced surface drylines in the region (e.g., Schaefer 1974; Parsons et al. 1991; Ziegler and Hane 1993; Hane et al. 1997, among others), in this case it marked the boundary between the moist Arctic cold dome and the dry Pacific postfrontal westerly flow rather than the confluence of moist and dry prefrontal airstreams that produce classical drylines.

The surface dryline in this case extended southward from the Arctic front over central Texas at 1800 UTC (Figs. 5a and 6a) and propagated slowly eastward during the subsequent 6-h period. The gradients of θυ and θe associated with the dryline were of opposite sign, that is, θυ decreased eastward from the dryline and θe decreased westward. This configuration is representative of most drylines (e.g., Schaefer 1974; Parsons et al. 1991; Ziegler and Hane 1993), though the magnitude of the θe gradient in our case was comparatively weak. Surface winds veered from south-southwesterly on the moist side of the dryline to southwesterly on the dry side.

The Pacific cold front moved rapidly eastward from extreme western Texas at 1800 UTC (Figs. 5a and 6a) to western Oklahoma and central Texas at 0000 UTC (Figs. 5d and 6d). Collocated horizontal gradients of θυ and θe trailed the leading edge of the Pacific front to the south of the Arctic air mass. Because the adiabatically warmed Pacific postfrontal airstream was potentially less dense than the shallow Arctic cold dome (θυ ≃ 300–310 K versus 290–300 K, respectively), the Pacific front rode over the top of the cold dome rather than displacing it. Hence, a thermodynamic signature was not observed at the surface as the Pacific front moved eastward aloft. But a wind shift from easterly to westerly component flow coincided with the frontal passage aloft between 2000 and 0000 UTC (Figs. 5b–d, 6b–d), as may be expected from the fact that surface wind fields can respond to airmass changes aloft (e.g., Locatelli et al. 1997; Neiman et al. 1998). Between 1800 and 2200 UTC (Figs. 5a–c, 6a–c), westerly flow in the Pacific postfrontal airstream strengthened to >15 m s−1 south and west of the retreating Arctic air mass. By 0000 UTC 18 April (Fig. 5d), the Pacific front’s θυ gradient had dissipated behind its leading edge due to prefrontal evaporative cooling associated with a developing squall line.

Severe thunderstorms erupted in the region of highest θe air (∼350 K) to the east of the dryline near the Arctic frontal boundary at 2000 UTC (Fig. 6b), and then moved northeastward across the Red River into southwestern Oklahoma. NOAA’s WP-3D flight crew and VORTEX’s ground-based mobile deployment teams focused their resources on this cluster of storms, which produced its first tornado near Temple in southwestern Oklahoma. The storms generated a pool of cool and comparatively dry outflow to the east of the advancing dryline and Pacific front that expanded in areal coverage over southwestern Oklahoma between 2200 UTC (Figs. 5c and 6c) and 0000 UTC (Figs. 5d and 6d). The Pacific front was beginning to migrate eastward over the top of this shallow rain-cooled air mass by 0000 UTC.

Sequential 2-h satellite imagery (Fig. 7) shows the relationship between cloud distributions and the analyzed air masses and frontal boundaries presented above. At 1800 UTC (Fig. 7a) widespread low-level clouds covered the region north of the Arctic front, though this cloud cover decreased significantly to the west of the Pacific front aloft between 2000 and 0000 UTC (Figs. 7b–d), quite likely in response to deep tropospheric subsidence and drying in the postfrontal environment. Jet-stream cirrus at 1800 UTC paralleled the southern portion of the Pacific front and streaked northeastward over the shallow Arctic air. This band of cirrus, which originated in the right entrance region of the jet over northern Mexico, became well-defined west of the surface front between 2000 and 0000 UTC (Figs. 7b–d). Blowing dust in the strengthening Pacific postfrontal westerly airstream was initially observed at 2000 UTC (Fig. 7b) as a region of lighter-toned clear air to the west and southwest of the juncture of the Pacific and Arctic fronts. By 2200 UTC (Fig. 7c) the postfrontal dust cloud became quite distinct and expanded in areal coverage, including over the southern portion of the Arctic air. The thickening dust cloud exhibited cellular patterning by 0000 UTC (Fig. 7d). Farther east, weak convective cells were observed in the post-dryline air to the east of the advancing Pacific front at 2000 UTC (Fig. 7b). This frontally forced convection organized into a line by 2200 (Fig. 7c), but because the line was located in relatively dry conditions behind the dryline, it did not become severe at this time. The prefrontal convection exploded into a squall line (NOAA 1995) by 0000 UTC (Fig. 7d) as the front impinged upon the moist Gulf of Mexico air to the east of the dryline. The propensity of intense convection to develop during dryline-frontal mergers is summarized in Schaefer (1986). Other notable cloud features were also observed in Fig. 7. A north–south cloud band over eastern Texas at 1800 UTC paralleled the dryline to its west, resembling a pre-drytrough cloud shield (Martin et al. 1995). This cloud feature persisted through 0000 UTC. The incipient convective towers of the Temple, Oklahoma, tornadic thunderstorm complex were first observed at 2000 UTC (Fig. 7b) immediately east of the dryline–Arctic-front intersection. This complex grew rapidly and moved northeastward ahead of the prefrontal convective line in the subsequent 4 h. Because this complex is not the focus of this study, it will not be discussed in detail.

5. A time series perspective of the Pacific cold front’s eastward advance

The continuous traces from instruments at individual sites often extend our knowledge about the detailed structure of transient atmospheric boundaries, especially when these traces are placed in the context of analyses derived from intermittent observations taken at relatively widely spaced locations. Therefore, selected time series traces of surface data and time–height section analyses of wind profiler data will be discussed in the context of the mesoscale analyses presented in section 4. The specific sites discussed below are labeled in Figs. 5a and 6a. A common theme contained in each of these traces is the passage of the Pacific cold front. The passage of additional features relevant to individual sites are also described.

A time–height section of data from the Jayton, Texas, wind profiler (JTN; Fig. 8) documented the frontal and dryline boundaries and their adjacent air masses in a region unaffected by significant precipitation. The leading edge of the Arctic air initially moved southward beyond the site at 0500 UTC 17 April when the dewpoint temperature increased by ∼14°C, reflecting the moist character of this shallow air mass. The frontal shear zone initially ascended to 850 mb by 1000 UTC and then retreated north of JTN at 1554 UTC when weak (<2 m s−1) east-southeasterly flow at the surface was replaced by stronger (∼6 m s−1) southerly flow, and the dewpoint depression increased significantly. The dewpoint temperature remained high (∼20°C) following this transition, because the dryline was still west of JTN. The dryline descended from ∼800 mb at 1100 UTC to the surface at 1712 UTC and was marked by a descending layer of vertical directional wind shear in a warm-advection sense. Companion 6-min resolution signal-power data from the profiler’s vertical beam (not shown) contained a layer of enhanced signal power that clearly defined the position of the descending dryline. This enhanced signal-power layer results from the large vertical gradient of moisture (White et al. 1991) associated with the dryline. The dryline passage at the surface was marked by a sharp decrease in dewpoint temperature and a wind-direction shift from ∼6 m s−1 southerly to ∼8 m s−1 southwesterly, followed by the rapid growth of a dry-convective boundary layer, that is, the temporal ascent of enhanced signal power associated with fluctuations of refractive index (e.g., White et al. 1991; White 1993; Angevine et al. 1994). Approximately 3 h later, the Pacific front advanced eastward across the profiler and was marked by a cessation of the diurnal temperature increase and by a surface pressure through. In addition, a rapid 5 m s−1 increase in wind speed, and veering of the wind from southwesterly to westerly, was observed. The along-front wind analysis aloft shows a 10-h temporal ascent of the Pacific frontal shear zone from the surface to the tropopause and a ∼65 m s−1 jet-stream core situated at 200 mb on the warm side of the front. Jet-stream cirrus was situated at ∼325 mb on the anticyclonic shear side of the jet core to the west of the surface intersection of the Pacific front. Strong postfrontal westerly flow above the surface exceeded 20–25 m s−1. A visibility trace from nearby Lubbock (LBB;not shown) shows rapid obscurement immediately following the Pacific frontal passage at the surface, the result of blowing dust that had become airborne over western Texas and eastern New Mexico by the strong postfrontal westerly flow. Given that the blowing dust was observed only within the Pacific postfrontal environment, it serves as a reliable tracer for the passage of the Pacific front at other sites.

Northeast of JTN, the Pacific front advanced eastward over the top of the shallow Arctic air. Surface time series traces from Childress, Texas (CDS; Fig. 9), show evidence of the frontal passage aloft in the absence of convection, and the wind profiler from Vici, Oklahoma (VCI), provided direct observations (Fig. 10). At CDS (Fig. 9) easterly component flow with small dewpoint depressions (<1°C) were observed within the Arctic air, where rainfall reduced the visibility between 0900 and 1430 UTC 17 April and fog reduced it between 1530 and 2300 UTC. The frontal passage aloft was defined at 2125 UTC by a pronounced decrease in pressure fall ahead of the frontal pressure trough and a wind shift from weak (<2 m s−1) southeasterly to steadier (3–4 m s−1) westerly. During this transition the fog dissipated and the visibility increased markedly, consistent with the satellite imagery (Fig. 7) showing a decrease in cloudiness following the frontal passage aloft. A thermodynamic transition was not evident, though the dewpoint depression increased. The Arctic air exited CDS at 0005 UTC when the westerly flow strengthened significantly from 4 to 14 m s−1, the temperature increased briefly by 2.8°C during the initial downward mixing of warmer isentropes from aloft, and the dewpoint temperature plummeted. The visibility decreased briefly following the retreat of the Arctic air because the Pacific postfrontal dust cloud mixed downward to the surface.

Surface traces of the Pacific frontal passage above CDS were quite similar to those at the Vici profiler (Fig. 10), where easterly flow within the moist Arctic air persisted below the lowest profiler range gate (i.e., <500 m deep) until 2342 UTC 17 April. Wind profiler observations of the Pacific frontal passage aloft exhibited a pronounced wind-direction shift from southerly to westerly that sloped upward from the lowest range gate at 2342 UTC to 3 km above mean sea level (MSL) 1 h later. Vici’s signal-power data observed nascent convective precipitation above the leading edge of the front (see also Fig. 7d), that is, a plume of enhanced signal power (>80 dB) aloft associated with Raleigh scattering from precipitation particles (e.g., Ralph et al. 1995). Nearby soundings launched from Woodward, Oklahoma (WWR; situated 43 km northwest of VCI) at 2249 and 2354 UTC (not shown) documented the Arctic frontal inversion at the level of the lowest profiler range gate and showed a 2°C decrease in temperature in the lowest 2 km (above the shallow Arctic frontal inversion) associated with the Pacific frontal passage. Surface traces that show the departure of the Arctic air at 0354 UTC 18 April were similar to those at CDS, though drying occurred more slowly at VCI.

The Pacific front became more difficult to resolve with only surface traces following its merger with the dryline, because the Gulf air mass was becoming potentially more dense with time (due to evaporative processes associated with the developing prefrontal squall-line precipitation) relative to the Pacific postfrontal air mass aloft (which warmed by 8°C at 500 mb between 1200 UTC 17 and 18 April; Figs. 2a–c), similar to that observed in another case by Neiman et al. (1998). In addition, the developing squall line produced convective-scale signatures at the surface that overwhelmed those of the synoptic scale. Therefore, close inspection of wind profiler data aloft proved essential in differentiating the Pacific front from the prefrontal squall line during this phase of the front’s evolution. Observations from the Purcell, Oklahoma, profiler (PRC; Fig. 11) show the squall line’s gust front passing the surface at 0100 UTC 18 April, after which the pressure increased sharply by 5.3 mb with the approach of the rain-cooled, convective-scale mesoridge. The subsequent wind–temperature transition was phase lagged by 10–15 min as expected (Charba 1974; Goff 1976), followed by convective precipitation that totaled 28.8 mm. A deep wake low (Stumpf et al. 1991; Loehrer and Johnson 1995) was subsequently observed at 0205 UTC. Because the squall line contained highly variable convective-scale air motions, the wind profiler could not obtain reliable wind profiles during this 1-h period. The leading edge of the advancing Pacific front reached PRC by ∼0330 UTC and sloped upward from the lowest range gate to 5 km MSL within 3 h. The along-front wind analysis highlights the enhanced frontal shear zone below 1.2 km and above 2.5 km MSL. In between, strong prefrontal westerly inflow into the rear of the squall line disrupted the vertical continuity of this shear zone. Surface traces at PRC show evidence of the frontal passage aloft during a 15-min window between 0320 and 0335 UTC. This window marked the beginning of an extended pressure increase of ∼7 mb, a wind-direction shift from south-southwesterly to west-southwesterly, and a brief 1-h increase in temperature arising from enhanced frontal-scale mixing of isentropes downward into the rain-cooled air mass. Gradual drying that commenced at ∼0800 UTC coincided with an increase in wind speed and brief warming at the surface, and strengthening westerly winds in the lowest kilometer above ground. This would suggest that drier, potentially warmer postfrontal air aloft mixed downward to the surface and replaced the shallow rain-cooled air. Wind profiler observations of the Pacific front and prefrontal squall line at Lamont (LMN) and Haskell (HKL) in Oklahoma (not shown) were similar to those at PRC. Prior to the squall line and frontal passage, the Arctic air retreated north of PRC at 1755 UTC 17 April.

6. Airborne observations of the Pacific cold front’s interactions with shallow air masses

The NCAR Electra played a crucial role in documenting the finescale interactions between the advancing Pacific cold front and the shallow Gulf and Arctic air masses situated east of the Rockies. These interactions are explored below by combining three-dimensional analyses based on the ELDORA observations with analyses of 1-s resolution in situ flight-level data. Special emphasis is given to the merger of the Pacific front with the dryline and the deep convective precipitation that ensued. Though numerous investigators (e.g., Shapiro 1982; Ziegler and Hane 1993; Hane et al. 1997, among others) have used research aircraft, some equipped with Doppler radar, to interrogate the structure and evolution of drylines and fronts in western Texas and adjacent areas, this is the first case in which the clear-air sensitive ELDORA Doppler radar was deployed.

a. Flight-level cross-section analyses

The horizontal mesoscale frontal evolution between 2100 UTC 17 April and 0000 UTC 18 April is summarized in Fig. 12. The vertical structure and evolution of the salient boundaries and air masses during this period are illustrated in a pair of cross sections prepared along lines BB′ and CC′. These sections were constructed by integrating the in situ flight-level data with rawinsonde and surface observations.

The cross sections of θ and u (Fig. 13a) and θe (Fig. 13b) at 2100 UTC show the Pacific frontal zone extending upward from the surface over LBB and then sloping westward with height above 700 mb where strong vertical wind shear was observed. Below this level, winds veered from ∼15 m s−1 southwesterly ahead of the front to stronger west-southwesterly within the frontal zone. At this time, the front was impinging upon the shallow Arctic air mass (below 875–900 mb) situated west of DYS where u was less than 0 m s−1. The Arctic air was characterized by a bubble of high θe that exceeded ∼340 K and a tight dryline-like θe gradient at its western edge. The western portion of this bubble penetrated upward to 830 mb in a ∼40 km wide plume near the leading edge of the advancing Pacific front. Shapiro (1982) showed a similar example of vertical displacement of θe at the western edge of a Gulf air mass, the result of frontally forced ascent with the approach of a cold front. It is therefore logical to conclude in our case that upward motion caused this θe protrusion, even though no observations were taken at the back edge of the Arctic air aloft prior to this time. Because the Arctic air mass was quite shallow and strongly capped in our case, convection did not develop during the Pacific frontal passage here. Section 6b(1) will present detailed airborne radar analyses of this region. Farther east, the dryline extended upward from the surface near DYS and sloped eastward with height, reaching a maximum altitude of 750 mb before sloping downward. The upward bulge of θe was similar to that shown in other dryline studies (e.g., Parsons et al. 1991;Ziegler and Hane 1993). In these studies, the upward motion was forced by low-level convergence near the leading edge of the surface dryline, and it is logical to conclude that the same was true here. The dryline-forced θe plume preceded the frontally forced plume by ∼150 km, that is, the upward motion associated with the dryline and front were decoupled at 2100 UTC. Opposing horizontal gradients of θ and θe characterized the dryline. Because the aircraft transected the dryline in the vicinity of the moist Arctic air, the horizontal θe gradient across the dryline was not well defined. Winds veered with height from southerly in the moist Gulf air to southwesterly in the drier and potentially warmer air aloft.

The 3-h vertical evolution of these boundaries and air masses was captured in cross sections of θ and u (Fig. 14a) and θe (Fig. 14b) at 0000 UTC 18 April. These sections show the shallow, but moist, Arctic dome extending from near GRAN to west of CDS, and they also show enhanced stable stratification above GRAN associated with the thunderstorm-induced cold pool (shown in Fig. 12). The tight dryline-like θe gradient at the back edge of the Arctic air persisted through 0000 UTC. Aloft, the isentropes of the Pacific frontal zone advanced over the top of this shallow layer, and the front’s thermodynamic and wind distributions were similar to those 3 h earlier. Significantly, the leading edge of the front was merging with the dryline between LTS and GRAN. The θe gradient associated with the dryline strengthened dramatically in 3 h and sloped westward with height together with the frontal isentropes. By this time, the upward-motion induced θe plumes associated with the dryline and Pacific front had phased, as evinced by the solitary pre-dryline θe plume that extended considerably deeper into the troposphere than did the two uncoupled plumes 3 h earlier. During this 3-h period, southerly flow within the Gulf air increased from ∼12 to ∼20 m s−1, and the θe within this airstream increased by ∼6 K. By 0000 UTC, a squall line exploded in the vicinity of the merger and coincided with 1) the phasing of the two upward-motion induced θe plumes, 2) the moistening and deepening of the Gulf air mass, and 3) the strengthening southerly flow within this air mass. The dryline-frontal merger will be described in detail in section 6b(2) using ELDORA airborne radar analyses.

b. Airborne radar analyses

The ELDORA is a new airborne Doppler radar platform with increased sensitivity to measure wind-flow patterns within the clear air as well as within precipitating environments. The clear-air observing capabilities of ELDORA have been documented by Wakimoto et al. (1996). Wind-field and reflectivity analyses based on the ELDORA measurements are presented below. The heights of all radar analyses are above ground level.

1) Flow over a shallow Arctic air mass

Horizontal airborne radar analyses over the southern Texas Panhandle highlight the finescale wind and reflectivity distributions in the vicinity of the western edge of the shallow Arctic mass at ∼2052 UTC 17 April 1995 (Fig. 15). At 0.25 km (Fig. 15a), the Arctic front separated easterly component front-relative flow within the Arctic air from southwesterly flow. A band of enhanced clear-air reflectivity (∼3 dBZ) paralleled the front. Aloft (Fig. 15b), southwesterly flow capped the Arctic air. The vertical structure of the Arctic front is illustrated in three cross sections along the line DD′ in Fig 15. The front-relative horizontal winds (Fig. 16a) highlight the vertical wind shear across the Arctic front, and the ground-relative winds (Fig. 16b) show weak westerly component flow in the shallow Arctic air and strong (∼20 m s−1) southwesterly flow aloft. The front-relative wind vectors in the plane of the cross section (Fig. 16c) show the ambient southwesterly flow deflected over the top of the shallow Arctic air mass. The nose-like leading edge and small spatial scale of the transition between these two air masses are similar in form to that described for density-current flows (e.g., Simpson 1969, 1972), despite the fact that the Arctic air mass here was retreating. Convergence associated with the merging of these airstreams below 1 km exceeded −20 × 10−4 s−1 (not shown) above the surface intersection of the Arctic front, resulting in a concentration of clear-air particulates and a vertical plume of enhanced reflectivity (Wilson et al. 1994). Upward motion surpassed 2 m s−1 in this region but decreased with height to zero at 3 km due to divergent flow aloft. Downward motion was observed downstream of the ascent plume above the Arctic air, which led to the depression in the dBZ isopleths there.

A low-level (∼200 m) aircraft transect shown in Fig. 15 documented the Arctic front with 1-s flight-level data. Time series traces (Fig. 17) from this transect show a rapid 3°C increase in temperature and 15°C decrease in dewpoint temperature at ∼2054:50 UTC as the aircraft flew from the Arctic air into the dry southwesterly flow. The ground-relative wind shifted sharply from weak southerly to strong (15–20 m s−1) west-southwesterly across the front. Upward motion of 1 m s−1 accompanied this transition, consistent with the radar-observed upward-motion plume. The rapid change in meteorological variables is consistent with the density-current-like structure observed in Fig. 16c. Approximately 10 km northeast of the Arctic front and 90 s earlier, the aircraft observed a density-current-like wake region within the Arctic air. This was marked by a temporary (∼30 s) increase in temperature and decrease in dewpoint temperature, as warmer isentropes and drier air mixed downward into the Arctic air. The wind shifted from weak southerly to stronger southwesterly for ∼30 s, as the strong southwesterly momentum aloft mixed downward. A downward-motion spike of 6 m s−1 coincided with the wake transition. The radar cross section in Fig. 16c clearly resolved this wake region, depicting a zone of downward motion and an associated downward deflection of radar reflectivity situated ∼10 km northeast of the surface intersection of the Arctic front. This cross section also illustrates the finescale nature of the density-current-like head, whereas the larger-scale, moist isentropic cross section (Fig. 13b) captured the θe plume resulting from the combined effects of Arctic and Pacific frontal ascent.

In many respects, this Arctic front resembled drylines that separate moist Gulf air from dry southwesterly flow. For example, the Arctic front sloped eastward with increasing height and was potentially unstable (e.g., Carlson et al. 1983; Martin et al. 1995). It exhibited a sharp westward decrease in moisture and a westward increase in temperature (e.g., Schaefer 1974; Ziegler and Hane 1993; Hane et al. 1997). The front contained a density-current-like structure and circulation (e.g., Parsons et al. 1991). And, the dry westerly flow was accelerated upward and over the relatively cool surface-based air mass (e.g., Ziegler and Hane 1993). However, deep convection was not triggered at the Arctic front despite the upward motion that was forced here, partly because the capped Arctic air was considerably more shallow than typical Gulf air masses situated east of classical drylines.

The dry westerly component flow that interacted with the Arctic air mass at 2052 UTC was Pacific postfrontal in origin, that is, the Pacific front was east of the radar analysis domain at that time. Several observations support this conclusion, including 1) the ground-relative flow within the Arctic air was westerly (Fig. 16b), consistent with the horizontal mesoscale analyses (Figs. 5 and 6) that showed westerly component flow within the Arctic air following the Pacific frontal passage aloft, 2) the strength of the southwesterly flow in Fig. 16b is similar to that within the Pacific postfrontal air mass shown in cross section BB′ (Fig. 13) and in the mesoscale analyses (Figs. 5 and 6), and 3) the flight-level dewpoint temperature in the southwesterly flow was less than 0°C on average, similar to that observed at surface sites within the Pacific postfrontal airstream.2 Because the Pacific postfrontal westerly airstream was quite dry, it was unable to reach the level of free convection (LFC) during its forced ascent over the Arctic air. Also, because this airstream was characterized by deep-tropospheric postfrontal subsidence (see Fig. 7), the adjoining region of forced ascent situated within the western periphery of the moist Arctic air mass (see Fig. 16c) could not penetrate vertically through the capping inversion to the LFC. Farther east, however, the Arctic front quite likely did act as a forcing mechanism to trigger the Temple, Oklahoma, tornadic thunderstorm complex in the Pacific prefrontal high-θe environment over extreme southwestern Oklahoma (see Figs. 7c,d). The mesoscale analyses in Figs. 5 and 6 show strengthening southerly flow in the moist Gulf air impinging upon the Arctic front between 1800 and 2200 UTC 17 April. Warm-sector mobile soundings that were launched in southwestern Oklahoma (not shown) just prior to the outbreak of severe convection were negatively buoyant but required only ∼500 m of forced ascent in the lower troposphere to destabilize sufficiently to allow deep convection to form. Assuming that the Arctic frontal forcing in this region was similar to that observed by the radar in the dry air (i.e., w = ∼1–2 m s−1), it would have taken only ∼5–10 min to destabilize the atmosphere to the point where deep convection would have erupted.

2) Dryline-frontal merger

The clear-air capabilities of the ELDORA radar proved essential in capturing the wind distributions and associated kinematics during the merger of the Pacific cold front and the dryline. Radar analyses and companion flight-level data from two flight legs are presented. These flight legs are represented by two bold-dotted lines at the bottom of the gray-shaded rectangles in the isentropic cross section in Fig. 14. The first leg emphasizes the front and dryline as separate but adjacent entities in clear-air conditions, while the second (only 22 min later) highlights the imminent merging of these boundaries, the phasing of their ascent plumes, and the subsequent development of convective precipitation. Though the initiation of intense convection during the approach and/or intersection of fronts with drylines has been discussed by numerous investigators (e.g., Koch and McCarthy 1982; Ogura et al. 1982; Schaefer 1986;Doswell 1987; Parsons et al. 1991; Hane et al. 1993), few studies have provided direct measurements documenting the process.

Radar analyses from the first flight leg at ∼2306 UTC (Figs. 18 and 19) show a 20–30-km gap between the Pacific front and dryline. At 0.30 km (Fig. 18a), the dryline separated moist, front-relative south-southeasterly flow from drier southwesterly flow. Enhanced clear-air reflectivity paralleled the dryline, primarily on its moist side. Significant along-dryline variability was observed, similar to that shown in Hane et al. (1997) and Atkins et al. (1998). At 1.10 km (Fig. 18b), a band of enhanced clear-air reflectivity (>3 dBZ) marked the position of the dryline below. During the aircraft’s low-level (∼275 m) westward transect across the dryline at 2306:30 UTC (Fig. 20), the temperature increased by ∼2°C, the dewpoint temperature decreased by ∼6°C, and the ground-relative wind direction shifted from southerly to southwesterly. Less than 30 s later, the aircraft penetrated the shallow Arctic air mass, at which time cooling and moistening commenced. Farther west, the traces reveal a wind-direction shift from southerly to westerly at ∼2310 UTC that is consistent with earlier analyses showing a shift to westerly component flow within the Arctic air following the Pacific frontal passage aloft. Inspection of the 1.10-km radar analysis (Fig. 18b) above the Arctic air shows a front-relative wind-direction shift from southwesterly to west-southwesterly associated with the Pacific front. Because the 275-m flight-level traces did not show a corresponding thermodynamic signature with the front, it is marked as occurring aloft on the 0.30-km radar analysis (Fig. 18a). A band of enhanced clear-air reflectivity at 0.30 km and 1.10 km paralleled the Pacific front on its cold side. The 1-s flight-level data showed a brief period of upward motion approaching 1 m s−1 near the leading edge of the Pacific front aloft.

The vertical structure of the Pacific front and dryline is presented in three cross sections along the line EE′in Fig. 18. The cross sections3 of front-relative and ground-relative horizontal winds (Figs. 19a and 19b, respectively) highlight the cold-advection vertical wind shear through the advancing Pacific front and the warm-advection vertical shear across the dryline. The ground-relative winds exhibited an increase in westerly component flow behind the Pacific front, in agreement with analyses presented earlier. The front did not penetrate downward through the lowest radar range gate. Front-relative wind vectors in the plane of the section (Fig. 19c) reveal a prominent updraft associated with each boundary and a region of prefrontal subsidence cradled between the ascent plumes. Upward motion greater than 1 m s−1 lay above the Pacific front. A noselike upward bulge in clear-air reflectivity coincided with the back edge of this ascent. The lower portion of the reflectivity bulge originated behind the Pacific front. Twenty kilometers farther east, a plume of ascent approaching 3 m s−1 was forced by a 2.5 km deep layer of convergence (not shown) associated with the dryline. Convergence was maximized below 0.5 km in the region of largest clear-air reflectivity (>9 dBZ). At this time prefrontal subsidence was impinging upon the upward motion at the dryline, that is, the vertical circulations associated with these two boundaries were in opposition or out of phase, consistent with the absence of thunderstorm activity. However, because the Pacific front was rapidly overtaking the dryline, these separate regions of forced ascent were about to merge. The phasing of these ascent plumes is highlighted in the pair of moist isentropic cross sections at 2100 and 0000 UTC (Figs. 13b and 14b, respectively).

Radar analyses from the second flight leg at ∼2328 UTC (Figs. 21 and 22) highlight the dryline-frontal merger. The horizontal analysis at 0.7 km (Fig. 21a) shows a ∼10 km separation between the Pacific front and dryline and also shows convective precipitation developing primarily along and east of the dryline. Because the convective cells were evolving rapidly into a solid squall line during the aircraft’s eastward penetration through this region, it was not possible for the aircraft to return westward for additional documentation of the merger process. The dryline at 0.7 km separated the front-relative southerly and southwesterly airstreams. Flight-level traces across the dryline (Fig. 23) show a ∼3°C increase in temperature, a ∼10°C decrease in dewpoint temperature, a 1.7 m s−1 updraft, and a wind-direction shift from south-southwesterly to southwesterly. The undular character of the dewpoint (and other) traces indicate that the western edge of the Gulf air now contained strong wave perturbations, perhaps arising from the developing convection or from the approaching Pacific front. These perturbations strongly resemble gravity waves that were observed with Doppler lidar on a dryline boundary during the TEXEX field experiment (Parsons et al. 1991). Farther west, the transition of front-relative dual-Doppler winds (Fig. 21a) across the Pacific front was more subtle. Postfrontal winds at 0.7 km possessed a larger westerly component than in the prefrontal environment. The superimposed front-relative flight-level winds at 750 m AGL show a more pronounced transition, shifting from prefrontal southerly to stronger postfrontal southwesterly. The 1-s aircraft traces show a well-defined frontal transition that mirrors surface observations of the frontal passage in western Texas, that is, a decrease in dewpoint temperature and a shift to strong (≥15 m s−1) west-southwesterly ground-relative flow on the cold side of the front. At 1.10 km (Fig. 21b), neither the dryline nor the Pacific front exhibited a sharp wind transition, though a broad band of confluence coincided with the dryline below. The Pacific front was more diffuse than 22 min earlier, quite likely because the circulations that were developing in response to the nascent convective activity dominated the frontal-scale circulations. A band of enhanced clear-air reflectivity trailed the Pacific front at 1.1 km, similar to that at 2306 UTC (Fig. 18b).

Cross sections of front-relative and ground-relative horizontal winds at 2328 UTC (Figs. 22a and 22b; along the line FF′ in Fig. 21) show the Pacific front within 10 km of the dryline. These two features retained their defining structural characteristics between 2306 and 2328 UTC, that is, cold-advection vertical wind shear through the advancing Pacific front, an increase in westerly component flow (especially of the ground-relative winds) and a noselike upward bulge in clear-air reflectivity behind the front, and warm-advection vertical shear across the dryline. Front-relative wind vectors in the plane of the section (Fig. 22c) show a broad region of upward motion extending eastward from the leading edge of the Pacific front into the moist air east of the dryline, subsidence behind the leading edge of the Pacific front, and organized ascent deeper in the cold air that mirrored the flight-level data. Comparison of this analysis with its analog 22 min earlier (Fig. 19c) demonstrates that the broad region of prefrontal upward motion represents the merging of the distinctly separate plumes of ascent that were forced kinematically at the Pacific front and dryline. Remnants of these plumes were still resolved at 2328 UTC (Fig. 22c). One plume (>1 m s−1) was located 5 km east of the front, and the other (1–3 m s−1) was centered on the moist side of the dryline. The superimposed reflectivity analysis shows convective precipitation in the moist air east of the dryline where forced ascent was maximized in the region of highest θe (>350 K). The greatest reflectivity occurring aloft indicates that the convection was at an early stage of development, consistent with the incipient phasing of the vertical circulations associated with the front and dryline. The dryline rendering in Fig. 22 emphasizes the vertical breach of the dryline cap by the growing convection.

7. Summary and discussion

Through the integrated analysis of in situ and remote sensing data taken during the VORTEX field campaign on 17–18 April 1995, this study describes the interactions that occurred between an advancing Pacific cold front and the shallow Gulf of Mexico and Arctic air masses situated east of the Rocky Mountains. These interactions were driven largely by the complex topography of the western United States, that is, the juxtaposition of the Rocky Mountains, Gulf of Mexico, and Great Plains. A summary of these interactions, and their significance, are outlined below.

The large-scale evolution of this event was characterized by the eastward migration of a vigorous shortwave trough of polar origin from the eastern Pacific Ocean to the southern and central plains where ambient low-level moisture and moderate convective instability were observed. Four air masses of distinctly different origin (i.e., Pacific polar, high-altitude continental, Gulf of Mexico, and Arctic), and the boundaries that separated them (i.e., Pacific cold front, dryline, and Arctic front), were observed within the experimental domain.

The shortwave trough was marked at its leading edge by a distinctive cold-frontal baroclinic zone, that is, the Pacific cold front. The front moved eastward across the southern Rockies and into the western portion of the VORTEX domain on 17 April, where its baroclinicity extended downward from the upper troposphere to the surface. East of the Rockies, strong postfrontal downslope westerly flow developed at the surface, with damaging winds approaching 30 m s−1 in the region. The detection of the Pacific front at the surface became problematic over the central and eastern portions of the VORTEX domain, because the preexisting Arctic and Gulf of Mexico air masses over the high plains were potentially more dense than the advancing Pacific postfrontal air mass that experienced adiabatic warming and sensible heating while traversing and descending the southern Rockies. Hence, as the Pacific front migrated eastward beyond the western edge of these shallow air masses, its lower terminus intersected the stable stratification capping the top of these air masses rather than the ground, creating a cold front aloft. As a result, the Pacific front did not possess a thermodynamic frontal signature at the surface, though a pressure trough and wind-direction shift were observed, as may be expected from the fact that the surface-pressure and corresponding wind fields can respond to airmass changes aloft (e.g., Locatelli et al. 1997; Neiman et al. 1998). The concept of fronts riding over the top of shallow stable layers, without exchanging air masses but still producing wind shifts through pressure changes at the surface, quite likely explains some of the difficulty encountered in analyzing surface fronts over the United States (e.g., Mass 1991; Uccellini et al. 1992; Sanders and Doswell 1995; Hobbs et al. 1996; Neiman et al. 1998). The presence of precipitation associated with fronts aloft, such as was observed with the Pacific front above the Arctic and Gulf air masses, further compounds the difficulty of surface analysis.

The NCAR Electra research aircraft represented an indispensable remote sensing and in situ observing platform for documenting the thermodynamic and kinematic evolution during the merger of the advancing Pacific cold front and the dryline. This evolution is summarized in a three-panel conceptualization portrayed in Fig. 24. During the first stage (Fig. 24a), the Pacific front and dryline were distinctly separate. A thermally direct vertical circulation spanning much of the troposphere accompanied the Pacific front. The upward-motion branch, whose lower portion was documented with airborne radar analyses, forced a line of weak prefrontal convection in the dry postdryline environment. The downward-motion branch resulted in a marked reduction in low- and midlevel cloud cover in the postfrontal environment. Theoretical work on ageostrophic frontal circulations by Shapiro (1982) and Keyser and Shapiro (1986), among others, support this conceptualization. Cloud cover aloft portrayed in Fig. 24a represents the lower extension of the jet-stream cirrus that was forced in the right entrance region of the polar jet stream. Farther east, the dryline marked the convergence of the dry pre-cold-frontal southwesterly flow and the moist southerly airstream originating over the Gulf of Mexico. The western extension of the moist layer and capping inversion bulged upward due to forced ascent before intersecting the sloping topography. This updraft was capped by prefrontal subsidence aloft, which was observed by airborne radar and inferred in a conceptualization by Shapiro (1982) for a similar case. In effect, the vertical circulations associated with the front and dryline were out of phase, consistent with the absence of thunderstorm activity. The second stage (Fig. 24b) portrays the merging of the Pacific cold front and dryline. Here, the shallow ascent that was forced at the dryline phased with the deep-tropospheric ascent ahead of the front, thus breaching the dryline cap and resulting in the formation of a squall line. This is fully consistent with the fact that intense convection often develops as fronts approach and/or intersect drylines (see the review by Schaefer 1986) and supports the assertion of Koch and McCarthy (1982) that the approach of a front can promote frontogenetical circulations along the dryline that enhance upward motion and initiate convection. In the final stage (Fig. 24c), the Pacific front rode over the top of the potentially more dense, rain-cooled Gulf air mass, while deep convection persisted near its leading edge. Though upward motion was still forced by convergence at the surface intersection of the dryline, postfrontal subsidence aloft prevented new convection from firing. At this time, the anvil from the mature convection coalesced with the jet stream cirrus.

During the event of 17–18 April, the shallow Arctic front and trailing air mass initially progressed equatorward across the central United States, with a southward acceleration east of the Rockies due to terrain-induced blocking (e.g., Mecikalski and Tilley 1992; Colle and Mass 1995). In many respects, the southward extension of the Arctic front over the Texas Panhandle resembled classical drylines that frequent the region. In addition, this portion of the front possessed a scale-contracted density-current-like head structure with forced ascent at its nose. Westerly component flow that was observed interacting with this finescale boundary was Pacific postfrontal in origin. Because the postfrontal airstream was quite dry and was characterized by deep-tropospheric subsidence, it was unable to reach the level of free convection during its forced ascent over the Arctic air. Farther east over Oklahoma, however, forced ascent of high-θe air over the Arctic front quite likely triggered a tornadic thunderstorm complex in the Pacific prefrontal environment over extreme southwestern Oklahoma. In the absence of complex topography, the Arctic front and its plume of forced ascent would have remained farther north in a lower-θe environment and therefore would have precluded the development of tornadic thunderstorms. During the latter phase of this event, the Arctic air mass retreated northward while eroding on its western and eastern flanks.

This study highlights what previous investigators (e.g., Hobbs et al. 1990; Martin et al. 1995; Hobbs et al. 1996; Neiman et al. 1998, among others) have shown, that is, that surface frontal analyses over the central United States can depart significantly from the classical Norwegian frontal-cyclone model because the air masses in this region frequently become modulated by the complex topography. However, we have expanded upon this previous body of work by documenting these departures at an unprecendentedly fine spatiotemporal resolution. These departures, in particular, complicate surface analyses of Pacific cold fronts as they descend the eastern slopes of the Rocky Mountains. It is becoming increasingly apparent that these Pacific fronts often do not intersect the ground over the Great Plains but remain aloft associated with their parent upper-level shortwave troughs. Although these fronts are decoupled from the surface, it is important to note that they can still be associated with severe weather and heavy precipitation over the central United States as illustrated in the present case, and described previously by Hobbs et al. (1990) and Neiman et al. (1998). Fortunately, the availability of continuous wind profiles (and the newer radio acoustic sounding system temperature profiles) from NOAA’s national network of 404-MHz wind profilers (Ralph et al. 1995) can help document the progression of these fronts. These operational remote sensing tools will aid forecasters and researchers alike as they incorporate these ubiquitous terrain-modulated atmospheric features into their surface analyses to improve their short-term forecasts and observational studies, respectively.

Acknowledgments

The implementation and success of the VORTEX field program was made possible by the dedicated participation of many individuals representing numerous organizations. We are grateful to all of them. Special thanks are given to Mel Shapiro for participating in the NCAR Electra research flight. Huaqing Cai helped edit the ELDORA data. Research results presented in this paper were partially supported by the NSF under Grants ATM9422499 and ATM9801720. Jim Adams provided exceptional drafting services.

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Fig. 1.
Fig. 1.

Base map of part of the VORTEX domain showing the primary profiling and airborne observing platforms used in this study: Rawinsondes (triangles), NOAA’s 404-MHz wind profilers (solid dots), and NCAR’s Electra research flight of 1843 UTC 17 April–0055 UTC 18 April 1995 (flight track, thin solid). The position and time (UTC) of the aircraft is marked at 10-min intervals. The three gray-shaded boxes, labeled 1–3, show the domains of the airborne pseudo-dual-Doppler wind analyses in Figs. 15, 18, and 21, respectively.

Citation: Monthly Weather Review 127, 9; 10.1175/1520-0493(1999)127<2102:TIOAPC>2.0.CO;2

Fig. 2.
Fig. 2.

The 500-mb geopotential height (dam, thin solid) and temperature (°C, bold dashed) analysis at (a) 1200 UTC 17 April 1995, (b) 0000 UTC 18 April, and (c) 1200 UTC 18 April. The 500-mb wind vector flags are 25 m s−1, full barbs are 10 m s−1, and half-barbs are 2.5 m s−1. Wind vectors with solid-dot heads are wind profiler observations. The bold dotted line AA′ in (a) is a projection line for the cross section in Fig. 4. The wind profiler site at Jayton, Texas (JTN) is labeled in (a), because data from this profiler are discussed in detail.

Citation: Monthly Weather Review 127, 9; 10.1175/1520-0493(1999)127<2102:TIOAPC>2.0.CO;2

Fig. 3.
Fig. 3.

Sea level pressure (mb, solid) and radar reflectivity (dBZ, shaded) analysis at (a) 1200 UTC 17 April 1995, (b) 1800 UTC 17 April, and (c) 0000 UTC 18 April. The reflectivities (shading: light, 18.5–40.5 dBZ; medium, 40.5–50.5 dBZ; dark, >50.5 dBZ) are based on the NWS hourly radar summaries. Though data from selected radars were not available for each of the three radar summaries, all relevant precipitation distributions were captured. The surface wind barbs, and projection line AA′, are as in Fig. 2. The Arctic front in each panel is represented by the bold line with closely spaced frontal symbols; all other fronts have widely spaced conventional symbols. Open frontal symbols depict fronts above the surface. The dryline is shown as a bold dashed line.

Citation: Monthly Weather Review 127, 9; 10.1175/1520-0493(1999)127<2102:TIOAPC>2.0.CO;2

Fig. 4.
Fig. 4.

Cross section of potential temperature (K, solid), along-front wind speed directed toward 40° (m s−1, bold dashed), and mixing ratio (g kg−1, dotted), along line AA′ of Figs. 2a and 3a at 1200 UTC 17 April 1995. (The along-front wind speed is directed toward 180° at OAK and DRA.) Thin-dashed lines are frontal boundaries and the tropopause. Rawinsonde sites and surface sites are marked with long and short vertical tick marks, respectively. Selected rawinsonde wind vectors are shown. Wind flags and barbs are as in Fig. 2.

Citation: Monthly Weather Review 127, 9; 10.1175/1520-0493(1999)127<2102:TIOAPC>2.0.CO;2

Fig. 5.
Fig. 5.

Mesoanalysis of virtual potential temperature (K, thin solid) at (a) 1800 UTC 17 April 1995, (b) 2000 UTC 17 April, (c) 2200 UTC 17 April, and (d) 0000 UTC 18 April. Frontal boundaries are as in Fig. 3. The dotted line in (c) and (d) is a thunderstorm outflow boundary. Surface wind flags and barbs are as in Fig. 2. Wind profiler (JTN, VCI, and PRC) and surface (CDS) sites, whose data are discussed in detail, are labeled and marked with bold dots in (a).

Citation: Monthly Weather Review 127, 9; 10.1175/1520-0493(1999)127<2102:TIOAPC>2.0.CO;2

Fig. 6.
Fig. 6.

The same as in Fig. 5, except for equivalent potential temperature (K).

Citation: Monthly Weather Review 127, 9; 10.1175/1520-0493(1999)127<2102:TIOAPC>2.0.CO;2

Fig. 7.
Fig. 7.

Visible satellite imagery at (a) 1800 UTC 17 April 1995, (b) 2000 UTC, (c) 2200 UTC, and (d) 0000 UTC 18 April. The frontal analyses are the same as those shown in Figs. 5 and 6.

Citation: Monthly Weather Review 127, 9; 10.1175/1520-0493(1999)127<2102:TIOAPC>2.0.CO;2

Fig. 8.
Fig. 8.

Time–height section (upper panel) and time series (lower panel) of data from the Jayton, Texas (JTN), wind profiler between 2330 UTC 16 April and 1130 UTC 18 April 1995. Upper panel: hourly averaged wind profiles and along-front wind speed directed toward 40° (m s−1, thin solid). Every other wind profile and range gate are shown. Hourly averaged surface winds measured at the profiler are included. Wind flags and barbs are as in Fig. 2. Bold solid (dashed) lines show frontal boundaries (dryline) based on the hourly averaged wind profiles and 6-min spectral moment data (not shown). The shading below 4 km shows the region of boundary layer growth based on the spectral moment data, and the shading between 8 and 10 km portrays the approximate location of the jet-stream cirrus based on infrared satellite imagery. Lower panel: time series traces of 6-min-resolution surface data measured at the profiler [DIR = wind direction (deg); SPD = wind speed (m s−1); T = temperature (°C); Td = dewpoint temperature (°C); PRS = surface pressure (mb)]. The position of JTN is marked in Figs. 2a, 5a, and 6a.

Citation: Monthly Weather Review 127, 9; 10.1175/1520-0493(1999)127<2102:TIOAPC>2.0.CO;2

Fig. 9.
Fig. 9.

Time series traces of 5-min resolution surface data from Childress, Texas (CDS), between 0800 UTC 17 April and 0800 UTC 18 April 1995. The traces are labeled as in Fig. 8, except VIS = visibility (km). The position of CDS is marked in Figs. 5a and 6a.

Citation: Monthly Weather Review 127, 9; 10.1175/1520-0493(1999)127<2102:TIOAPC>2.0.CO;2

Fig. 10.
Fig. 10.

Time–height section (upper panel) and time series (lower panel) of data from the Vici, Oklahoma (VCI), wind profiler between 2200 UTC 17 April and 0600 UTC 18 April 1995. Upper panel: 6-min resolution wind profiles. The bold line marks the approximate position of the Pacific cold front above the surface. The gray-shaded region is signal power greater than 80 dB measured by the profiler’s vertically pointing beam. Every other wind profile is plotted. Lower panel: as in Fig. 8. The position of VCI is marked in Figs. 5a and 6a.

Citation: Monthly Weather Review 127, 9; 10.1175/1520-0493(1999)127<2102:TIOAPC>2.0.CO;2

Fig. 11.
Fig. 11.

Time–height section (upper panel) and time series (lower panel) of data from the Purcell, Oklahoma (PRC), wind profiler between 1530 UTC 17 April and 1130 UTC 18 April 1995. Upper panel:same as in Fig. 8, except every wind profile and range gate are shown. Lower panel: same as in Fig. 8, except PPT = precipitation (mm). The position of PRC is marked in Figs. 5a and 6a.

Citation: Monthly Weather Review 127, 9; 10.1175/1520-0493(1999)127<2102:TIOAPC>2.0.CO;2

Fig. 12.
Fig. 12.

Frontal analysis at 2100 UTC 17 April 1995 (bold boundaries) and 0000 UTC 18 April (thin boundaries) based on Figs. 5 and 6. The frontal symbology is described in Figs. 3 and 5. For the sake of clarity, the Pacific front aloft is not shown here. The dotted line BB′ (CC′) is a projection line at 2100 UTC (0000 UTC) for the cross section in Fig. 13 (14). Rawinsonde data from the LBB, LTS, NS2, and ADM sites (bold) were included in Figs. 13 and/or 14. All rawinsonde and profiler sites in the domain are labeled and are described in Fig. 1.

Citation: Monthly Weather Review 127, 9; 10.1175/1520-0493(1999)127<2102:TIOAPC>2.0.CO;2

Fig. 13.
Fig. 13.

Meso–cross section of (a) potential temperature (K, solid) and along-front wind speed directed toward 40° (m s−1, bold dashed), and (b) equivalent potential temperature (K), along line BB′ of Fig. 12 at 2100 UTC 17 April 1995. Vertical tick marks and selected rawinsonde wind vectors are as in Fig. 4. The rawinsonde soundings are labeled with the time (UTC) of deployment. NCAR Electra time-to-space adjusted flight tracks between 1920 and 2221 UTC 17 April 1995 are depicted by small dotted lines, with selected flight-level wind vectors plotted. Wind flags and barbs are as in Fig. 2. The shaded rectangle shows the domain of the airborne pseudo-dual-Doppler cross section in Fig. 16. The bold dot near the bottom of this shading shows the approximate position of the aircraft during its section-normal transect shown in Figs. 15 and 17.

Citation: Monthly Weather Review 127, 9; 10.1175/1520-0493(1999)127<2102:TIOAPC>2.0.CO;2

Fig. 14.
Fig. 14.

Meso–cross section of (a) potential temperature (K, solid) and along-front wind speed directed toward 40° (m s−1, bold dashed), and (b) equivalent potential temperature (K), along line CC′ of Fig. 12 at 0000 UTC 18 April 1995. Vertical tick marks, selected wind vectors, and time stamps are as in Fig. 13. NCAR Electra time-to-space adjusted flight tracks between 2128 UTC 17 April and 0010 UTC 18 April are shown. The dual-shaded (dark shaded) rectangle shows the domain of the airborne pseudo-dual-Doppler cross section in Fig. 19 (22). The lower (upper) bold dotted line near the bottom of the shading denotes the flight track shown in Figs. 18 and 20 (21 and 23).

Citation: Monthly Weather Review 127, 9; 10.1175/1520-0493(1999)127<2102:TIOAPC>2.0.CO;2

Fig. 15.
Fig. 15.

NCAR Electra radar reflectivity (dBZ, gray-shaded) and pseudo-dual-Doppler front-relative wind vectors (scaled, upper left) between 2051:09 and 2056:30 UTC 17 April 1995 at (a) 0.25 and (b) 1.05 km AGL. The thin line marks the flight track at ∼200 m AGL, with selected front-relative wind barbs (barbs as in Fig. 2) shown and labeled with the times of observation. Time series of data from this track are presented in Fig. 17. The arctic front is shown. The bold line DD′ is the projection line for Fig. 16. The domain of this analysis is shown as box 1 in Fig. 1.

Citation: Monthly Weather Review 127, 9; 10.1175/1520-0493(1999)127<2102:TIOAPC>2.0.CO;2

Fig. 16.
Fig. 16.

NCAR Electra radar cross section of reflectivity (dBZ, gray shaded) along projection line DD′ in Fig. 15: (a) front-relative horizontal wind barbs, (b) ground-relative horizontal wind barbs, and (c) front-relative wind velocity vectors (scaled, upper right) in the plane of the cross section. The arctic front is shown. Wind barbs are as in Fig. 2.

Citation: Monthly Weather Review 127, 9; 10.1175/1520-0493(1999)127<2102:TIOAPC>2.0.CO;2

Fig. 17.
Fig. 17.

Time series traces of 1-s resolution flight-level data at ∼200 m AGL between 2051 and 2157 UTC 17 April 1995. The traces are labeled as in Fig. 8, except w = vertical velocity (m s−1). Wind barbs are as in Fig. 2. The flight track is shown in Fig. 15.

Citation: Monthly Weather Review 127, 9; 10.1175/1520-0493(1999)127<2102:TIOAPC>2.0.CO;2

Fig. 18.
Fig. 18.

NCAR Electra radar reflectivity (dBZ, gray shaded) and pseudo-dual-Doppler front-relative wind vectors (scaled, lower left) between 2300:15 and 2312:30 UTC 17 April 1995 at (a) 0.30 and (b) 1.10 km AGL. The flight track convention is described in Fig. 15, and the flight level was ∼275 m AGL. Time series of data from this track are presented in Fig. 20. The Pacific front and dryline are shown (as in Fig. 3), though the Arctic front is not. The bold line EE′ is the projection line for Fig. 19. The domain of this analysis is shown as box 2 in Fig. 1.

Citation: Monthly Weather Review 127, 9; 10.1175/1520-0493(1999)127<2102:TIOAPC>2.0.CO;2

Fig. 19.
Fig. 19.

NCAR Electra radar cross section of reflectivity (dBZ, gray shaded) along projection line EE′ in Fig. 18: (a–c) same as in Fig. 16. The Pacific front and dryline are shown (as in Fig. 3), though the arctic front is not. Wind barbs are as in Fig. 2.

Citation: Monthly Weather Review 127, 9; 10.1175/1520-0493(1999)127<2102:TIOAPC>2.0.CO;2

Fig. 20.
Fig. 20.

The same as in Fig. 17, except the flight-level data was at ∼275 m AGL between 2300:30 and 2314:00 UTC 17 April 1995. The flight track is shown in Fig. 18.

Citation: Monthly Weather Review 127, 9; 10.1175/1520-0493(1999)127<2102:TIOAPC>2.0.CO;2

Fig. 21.
Fig. 21.

NCAR Electra radar reflectivity (dBZ, gray shaded) and pseudo-dual-Doppler front-relative wind vectors (scaled, lower left) between 2323:30 and 2333:30 UTC 17 April 1995 at (a) 0.70 and (b) 1.10 km AGL. The flight track convention is described in Fig. 15, and the flight level was ∼750 m AGL. Time series of data from this track are presented in Fig. 23. The Pacific front and dryline are shown (as in Fig. 3). The bold line FF′ is the projection line for Fig. 22. The domain of this analysis is shown as box 3 in Fig. 1.

Citation: Monthly Weather Review 127, 9; 10.1175/1520-0493(1999)127<2102:TIOAPC>2.0.CO;2

Fig. 22.
Fig. 22.

NCAR Electra radar cross section of reflectivity (dBZ, gray shaded) along projection line FF′ in Fig. 21: (a)–(c) same as in Fig. 16. The Pacific front and dryline are shown (as in Fig. 3). Wind barbs are as in Fig. 2.

Citation: Monthly Weather Review 127, 9; 10.1175/1520-0493(1999)127<2102:TIOAPC>2.0.CO;2

Fig. 23.
Fig. 23.

The same as in Fig. 17, except the flight-level data were at ∼750 m AGL between 2323 and 2333 UTC 17 April 1995. The flight track is shown in Fig. 21.

Citation: Monthly Weather Review 127, 9; 10.1175/1520-0493(1999)127<2102:TIOAPC>2.0.CO;2

Fig. 24.
Fig. 24.

Conceptual representation of the merger of the Pacific cold front with the dryline, where Δt is approximately 3 h: (a) distinct separation between the advancing front and dryline, (b) merging of the front with the dryline and phasing of their vertical circulations, and (c) decoupling of the front from the surface by the shallow Gulf of Mexico air mass. The Pacific cold frontal zone and Gulf of Mexico air mass are shaded light and dark, respectively. The dashed lines mark the dryline. The gray-shaded arrows portray the vertical motions associated with the Pacific cold front and dryline. Cloud schematics are also shown.

Citation: Monthly Weather Review 127, 9; 10.1175/1520-0493(1999)127<2102:TIOAPC>2.0.CO;2

Table 1.

Summary of analyses that were subjected to the time-to-space adjustment technique. The time window (to the nearest minute) of relevant data for each analysis, and the phase velocity used to adjust the data in each analysis, are also given.

Table 1.

1

This air mass acquired its thermodynamic characteristics over the elevated terrain of the southern Rocky Mountains.

2

The θe plume analyzed at the western edge of the Arctic air mass in Fig. 13b was based on flight-level data taken approximately 20–45 min before the radar analysis at 2052 UTC. At this earlier time the Pacific front had not yet migrated beyond the western edge of the Arctic air mass, thus demonstrating a limitation to the time-to-space adjustment technique.

3

The along-dryline variability shown in Figs. 18 and 21 suggest that the radar cross sections may not be fully representative of the dryline’s vertical structure at other locations.

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  • Angevine, W. M., A. B. White, and S. K. Avery, 1994: Boundary-layer depth and entrainment zone characterization with a boundary-layer profiler. Bound.-Layer Meteor.,68, 375–385.

  • Atkins, N. T., R. M. Wakimoto, and C. L. Ziegler, 1998: Observations of the finescale structure of a dryline during VORTEX 95. Mon. Wea. Rev.,126, 525–550.

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  • Fig. 1.

    Base map of part of the VORTEX domain showing the primary profiling and airborne observing platforms used in this study: Rawinsondes (triangles), NOAA’s 404-MHz wind profilers (solid dots), and NCAR’s Electra research flight of 1843 UTC 17 April–0055 UTC 18 April 1995 (flight track, thin solid). The position and time (UTC) of the aircraft is marked at 10-min intervals. The three gray-shaded boxes, labeled 1–3, show the domains of the airborne pseudo-dual-Doppler wind analyses in Figs. 15, 18, and 21, respectively.

  • Fig. 2.

    The 500-mb geopotential height (dam, thin solid) and temperature (°C, bold dashed) analysis at (a) 1200 UTC 17 April 1995, (b) 0000 UTC 18 April, and (c) 1200 UTC 18 April. The 500-mb wind vector flags are 25 m s−1, full barbs are 10 m s−1, and half-barbs are 2.5 m s−1. Wind vectors with solid-dot heads are wind profiler observations. The bold dotted line AA′ in (a) is a projection line for the cross section in Fig. 4. The wind profiler site at Jayton, Texas (JTN) is labeled in (a), because data from this profiler are discussed in detail.

  • Fig. 3.

    Sea level pressure (mb, solid) and radar reflectivity (dBZ, shaded) analysis at (a) 1200 UTC 17 April 1995, (b) 1800 UTC 17 April, and (c) 0000 UTC 18 April. The reflectivities (shading: light, 18.5–40.5 dBZ; medium, 40.5–50.5 dBZ; dark, >50.5 dBZ) are based on the NWS hourly radar summaries. Though data from selected radars were not available for each of the three radar summaries, all relevant precipitation distributions were captured. The surface wind barbs, and projection line AA′, are as in Fig. 2. The Arctic front in each panel is represented by the bold line with closely spaced frontal symbols; all other fronts have widely spaced conventional symbols. Open frontal symbols depict fronts above the surface. The dryline is shown as a bold dashed line.

  • Fig. 4.

    Cross section of potential temperature (K, solid), along-front wind speed directed toward 40° (m s−1, bold dashed), and mixing ratio (g kg−1, dotted), along line AA′ of Figs. 2a and 3a at 1200 UTC 17 April 1995. (The along-front wind speed is directed toward 180° at OAK and DRA.) Thin-dashed lines are frontal boundaries and the tropopause. Rawinsonde sites and surface sites are marked with long and short vertical tick marks, respectively. Selected rawinsonde wind vectors are shown. Wind flags and barbs are as in Fig. 2.

  • Fig. 5.

    Mesoanalysis of virtual potential temperature (K, thin solid) at (a) 1800 UTC 17 April 1995, (b) 2000 UTC 17 April, (c) 2200 UTC 17 April, and (d) 0000 UTC 18 April. Frontal boundaries are as in Fig. 3. The dotted line in (c) and (d) is a thunderstorm outflow boundary. Surface wind flags and barbs are as in Fig. 2. Wind profiler (JTN, VCI, and PRC) and surface (CDS) sites, whose data are discussed in detail, are labeled and marked with bold dots in (a).

  • Fig. 6.

    The same as in Fig. 5, except for equivalent potential temperature (K).

  • Fig. 7.

    Visible satellite imagery at (a) 1800 UTC 17 April 1995, (b) 2000 UTC, (c) 2200 UTC, and (d) 0000 UTC 18 April. The frontal analyses are the same as those shown in Figs. 5 and 6.

  • Fig. 8.

    Time–height section (upper panel) and time series (lower panel) of data from the Jayton, Texas (JTN), wind profiler between 2330 UTC 16 April and 1130 UTC 18 April 1995. Upper panel: hourly averaged wind profiles and along-front wind speed directed toward 40° (m s−1, thin solid). Every other wind profile and range gate are shown. Hourly averaged surface winds measured at the profiler are included. Wind flags and barbs are as in Fig. 2. Bold solid (dashed) lines show frontal boundaries (dryline) based on the hourly averaged wind profiles and 6-min spectral moment data (not shown). The shading below 4 km shows the region of boundary layer growth based on the spectral moment data, and the shading between 8 and 10 km portrays the approximate location of the jet-stream cirrus based on infrared satellite imagery. Lower panel: time series traces of 6-min-resolution surface data measured at the profiler [DIR = wind direction (deg); SPD = wind speed (m s−1); T = temperature (°C); Td = dewpoint temperature (°C); PRS = surface pressure (mb)]. The position of JTN is marked in Figs. 2a, 5a, and 6a.

  • Fig. 9.

    Time series traces of 5-min resolution surface data from Childress, Texas (CDS), between 0800 UTC 17 April and 0800 UTC 18 April 1995. The traces are labeled as in Fig. 8, except VIS = visibility (km). The position of CDS is marked in Figs. 5a and 6a.

  • Fig. 10.

    Time–height section (upper panel) and time series (lower panel) of data from the Vici, Oklahoma (VCI), wind profiler between 2200 UTC 17 April and 0600 UTC 18 April 1995. Upper panel: 6-min resolution wind profiles. The bold line marks the approximate position of the Pacific cold front above the surface. The gray-shaded region is signal power greater than 80 dB measured by the profiler’s vertically pointing beam. Every other wind profile is plotted. Lower panel: as in Fig. 8. The position of VCI is marked in Figs. 5a and 6a.

  • Fig. 11.

    Time–height section (upper panel) and time series (lower panel) of data from the Purcell, Oklahoma (PRC), wind profiler between 1530 UTC 17 April and 1130 UTC 18 April 1995. Upper panel:same as in Fig. 8, except every wind profile and range gate are shown. Lower panel: same as in Fig. 8, except PPT = precipitation (mm). The position of PRC is marked in Figs. 5a and 6a.

  • Fig. 12.

    Frontal analysis at 2100 UTC 17 April 1995 (bold boundaries) and 0000 UTC 18 April (thin boundaries) based on Figs. 5 and 6. The frontal symbology is described in Figs. 3 and 5. For the sake of clarity, the Pacific front aloft is not shown here. The dotted line BB′ (CC′) is a projection line at 2100 UTC (0000 UTC) for the cross section in Fig. 13 (14). Rawinsonde data from the LBB, LTS, NS2, and ADM sites (bold) were included in Figs. 13 and/or 14. All rawinsonde and profiler sites in the domain are labeled and are described in Fig. 1.

  • Fig. 13.

    Meso–cross section of (a) potential temperature (K, solid) and along-front wind speed directed toward 40° (m s−1, bold dashed), and (b) equivalent potential temperature (K), along line BB′ of Fig. 12 at 2100 UTC 17 April 1995. Vertical tick marks and selected rawinsonde wind vectors are as in Fig. 4. The rawinsonde soundings are labeled with the time (UTC) of deployment. NCAR Electra time-to-space adjusted flight tracks between 1920 and 2221 UTC 17 April 1995 are depicted by small dotted lines, with selected flight-level wind vectors plotted. Wind flags and barbs are as in Fig. 2. The shaded rectangle shows the domain of the airborne pseudo-dual-Doppler cross section in Fig. 16. The bold dot near the bottom of this shading shows the approximate position of the aircraft during its section-normal transect shown in Figs. 15 and 17.

  • Fig. 14.

    Meso–cross section of (a) potential temperature (K, solid) and along-front wind speed directed toward 40° (m s−1, bold dashed), and (b) equivalent potential temperature (K), along line CC′ of Fig. 12 at 0000 UTC 18 April 1995. Vertical tick marks, selected wind vectors, and time stamps are as in Fig. 13. NCAR Electra time-to-space adjusted flight tracks between 2128 UTC 17 April and 0010 UTC 18 April are shown. The dual-shaded (dark shaded) rectangle shows the domain of the airborne pseudo-dual-Doppler cross section in Fig. 19 (22). The lower (upper) bold dotted line near the bottom of the shading denotes the flight track shown in Figs. 18 and 20 (21 and 23).

  • Fig. 15.

    NCAR Electra radar reflectivity (dBZ, gray-shaded) and pseudo-dual-Doppler front-relative wind vectors (scaled, upper left) between 2051:09 and 2056:30 UTC 17 April 1995 at (a) 0.25 and (b) 1.05 km AGL. The thin line marks the flight track at ∼200 m AGL, with selected front-relative wind barbs (barbs as in Fig. 2) shown and labeled with the times of observation. Time series of data from this track are presented in Fig. 17. The arctic front is shown. The bold line DD′ is the projection line for Fig. 16. The domain of this analysis is shown as box 1 in Fig. 1.

  • Fig. 16.

    NCAR Electra radar cross section of reflectivity (dBZ, gray shaded) along projection line DD′ in Fig. 15: (a) front-relative horizontal wind barbs, (b) ground-relative horizontal wind barbs, and (c) front-relative wind velocity vectors (scaled, upper right) in the plane of the cross section. The arctic front is shown. Wind barbs are as in Fig. 2.

  • Fig. 17.

    Time series traces of 1-s resolution flight-level data at ∼200 m AGL between 2051 and 2157 UTC 17 April 1995. The traces are labeled as in Fig. 8, except w = vertical velocity (m s−1). Wind barbs are as in Fig. 2. The flight track is shown in Fig. 15.

  • Fig. 18.

    NCAR Electra radar reflectivity (dBZ, gray shaded) and pseudo-dual-Doppler front-relative wind vectors (scaled, lower left) between 2300:15 and 2312:30 UTC 17 April 1995 at (a) 0.30 and (b) 1.10 km AGL. The flight track convention is described in Fig. 15, and the flight level was ∼275 m AGL. Time series of data from this track are presented in Fig. 20. The Pacific front and dryline are shown (as in Fig. 3), though the Arctic front is not. The bold line EE′ is the projection line for Fig. 19. The domain of this analysis is shown as box 2 in Fig. 1.

  • Fig. 19.

    NCAR Electra radar cross section of reflectivity (dBZ, gray shaded) along projection line EE′ in Fig. 18: (a–c) same as in Fig. 16. The Pacific front and dryline are shown (as in Fig. 3), though the arctic front is not. Wind barbs are as in Fig. 2.

  • Fig. 20.

    The same as in Fig. 17, except the flight-level data was at ∼275 m AGL between 2300:30 and 2314:00 UTC 17 April 1995. The flight track is shown in Fig. 18.

  • Fig. 21.

    NCAR Electra radar reflectivity (dBZ, gray shaded) and pseudo-dual-Doppler front-relative wind vectors (scaled, lower left) between 2323:30 and 2333:30 UTC 17 April 1995 at (a) 0.70 and (b) 1.10 km AGL. The flight track convention is described in Fig. 15, and the flight level was ∼750 m AGL. Time series of data from this track are presented in Fig. 23. The Pacific front and dryline are shown (as in Fig. 3). The bold line FF′ is the projection line for Fig. 22. The domain of this analysis is shown as box 3 in Fig. 1.

  • Fig. 22.

    NCAR Electra radar cross section of reflectivity (dBZ, gray shaded) along projection line FF′ in Fig. 21: (a)–(c) same as in Fig. 16. The Pacific front and dryline are shown (as in Fig. 3). Wind barbs are as in Fig. 2.

  • Fig. 23.

    The same as in Fig. 17, except the flight-level data were at ∼750 m AGL between 2323 and 2333 UTC 17 April 1995. The flight track is shown in Fig. 21.

  • Fig. 24.

    Conceptual representation of the merger of the Pacific cold front with the dryline, where Δt is approximately 3 h: (a) distinct separation between the advancing front and dryline, (b) merging of the front with the dryline and phasing of their vertical circulations, and (c) decoupling of the front from the surface by the shallow Gulf of Mexico air mass. The Pacific cold frontal zone and Gulf of Mexico air mass are shaded light and dark, respectively. The dashed lines mark the dryline. The gray-shaded arrows portray the vertical motions associated with the Pacific cold front and dryline. Cloud schematics are also shown.

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