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    (a) Surface elevation map of the IMPROVE-2 study area with the flight track of the NOAA P3 (solid red line) overlain. Locations of the sounding site (UW), profiler (IB), vertically pointing NOAA/ETL S-band (MB), and NCAR S-Pol radar (SPL) are also indicated. Plotted flight track encompasses the period 2300 UTC 13 Dec–0100 UTC 14 Dec 2001. (b) Map of the Pacific Northwest with the 36-, 12-, 4-, and 1.33-km MM5 domains shown.

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    MM5 4-km simulated wind vectors (black arrows), potential temperature (color shading), pressure (blue dotted contours), and topography (gray contours) from the 4-km domain simulation. Analyses are at heights of 1 and 3 km MSL, respectively at (a), (b) 2000 UTC 13 Dec; (c), (d) 2300 UTC 13 Dec; and (e), (f) 0200 UTC 14 Dec 2001. The blue dashed box represents the region where airborne P3 dual-Doppler measurements were acquired. The black line shows the position of the east–west cross section in Fig. 3. The leading edge of the cold front along the cross section is indicated by CF.

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    East–west cross section (position shown in Fig. 2) of MM5 4-km simulated θ (color shading), θe (solid blue lines), and wind barbs (m s−1) for (a) 2000 UTC 13 Dec, (b) 2300 UTC 13 Dec, and (c) 0200 UTC 14 Dec 2001. The solid brown contour indicates the 0°C isotherm. The location of cold front given by gray dashed line with the shaded region in (b) showing a region of enhanced vertical gradients of θ and θe.

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    The U- and V-component winds from the NCAR ISS profiler (located at IB in Fig. 1a) at 30-min intervals (dashed color lines) from 2247 to 0047 UTC 13–14 Dec 2001. The solid black lines represent the average of the five individual profiles.

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    The U-component speed, dry and moist squared Brunt–Väisälä frequency (N2d and N2m), and inverse Froude number (Nh/U), where h is the mean height of the Cascade crest (h = 1.68 km) from the special UW sounding location (UW in Fig. 1a) at 0000 UTC 14 Dec 2001 and the MM5 1.33-km simulation.

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    Composite grid of (a), (c), (e) P3 dual-Doppler wind speed (m s−1, color shading) and wind vectors and (b), (d), (f) corresponding radar reflectivity at elevations of (a), (b) 1.5-, (c), (d) 3-, and (e), (f) 4.5-km MSL. Red line indicates flight track of P3 over the 2-h period from 2300 to 0100 UTC 13–14 Dec 2001. The black box in (e) and (f) outlines region over which data were averaged to create the mean east–west cross sections in Figs. 7 and 11, respectively. Dotted lines depict the north–south mean crest line.

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    Average east–west cross section of airborne dual-Doppler winds for the (a) U and (c) V component (color shading, m s−1). Black dots in represent the position of the P3 north–south legs. Dashed lines delineate the boundaries separating individual Doppler analysis grids included in the composite. Dashed black box in (a) shows the region averaged to create the “upstream” profile used to calculate Fig. 8. Averaged east–west cross section of the MM5 1.33-km simulated (b) U and (d) V component with contours of simulated equivalent potential temperature (θe: black contours) and potential temperature (θ: dashed white contours). Model fields were averaged in time from 2300 to 0100 UTC 13–14 Dec 2001 (forecast hours 10–13) using model output at 15-min intervals. White × marks in (b) are the ending locations of the trajectories shown in Fig. 9.

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    Dual-Doppler (a) U ′ and (b) V ′ in the east–west cross section corresponding to Figs. 7a,c. Black dots represent the position of the P3 north–south legs. Dashed lines delineate the boundaries separating individual Doppler analysis grids included in the composite.

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    (a) Map of backward trajectories terrninating over the windward slopes of the Cascades (at point −122.8°N, 44.5°W) at 2300 UTC 13 Dec 2001 at four different heights (key lower right) as shown in Fig. 7b. (b) Lagrangian traces of pressure (hPa), potential temperature (K), equivalent potential temperature (K), and relative humidity (%) following the parcel tracks shown in (a), plotted against the distance (km) from the endpoint.

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    Meridionally averaged east–west cross sections of MM5-generated equivalent potential temperature (θe) overlaid on (a), (c), (e) U-component flow and (b), (d), (f) vertical velocity for three different 1.33-km MM5 simulations using the (a), (b) MRF-PBL; (c), (d) Eta-PBL; and (e), (f) Yonsei-PBL. Model fields were averaged in time from 2300 to 0100 UTC 13–14 Dec 2001 (forecast hours 10–13) using model output at 15-min intervals.

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    Average east–west cross section of MM5 1.33-km modeled (a) CLW (qw) and (b) snow (qs) overlaid on airborne dual-Doppler derived radar reflectivity (dBZ; color shading, key at right). The modeled CLW and qs shown were averaged in time from 2300 to 0100 UTC 13–14 Dec 2001 (forecast hours 10–13) using model output at 15-min intervals.

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    Average MM5 east–west cross sections of MM5 CLW (g kg−1) overlaid on airborne Doppler reflectivity for three different 1.33-km MM5 simulations: (a) control, (c) no-coastal mountains (Noco), and (e) smoothed terrain (Smooth). Model θe overlaid on model vertical velocity for (b) control, (d) no-coastal mountains, and (f) smoothed terrain. The modeled fields were averaged from 2300 to 0100 UTC 13–14 Dec 2001 (forecast hours 10–13) using model output at 15-min intervals.

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    The U, V, and W components (m s−1) at 2.5 km MSL along leg 2 of the P3 flight track (location shown Fig. 1b) from the MM5 1.33-km model simulation, P3 in situ measurements, and as derived from dual-Doppler airborne radar (key at right). The P3 in situ winds were averaged using a 10-s running mean providing a horizontal resolution of about 1 km. The Doppler winds along the flight track were obtained through bilinear interpolation. The fourth panel shows CLW measurements along leg 2 as measured by the King probe aboard the P3 and as predicted by the MM5 1.33-km simulations. The bottom panel is the underlying terrain along the flight track.

  • View in gallery

    The north–south cross section of airborne dual-Doppler and MM5-predicted (a), (b) U- and (c), (d) V-component flow along leg 2 of the P3 flight track (cf. Fig. 1a). Contours of Doppler-derived vertical velocity are displayed at intervals of 0.2 m s−1 in (a) and (c), while contours of model predicted vertical velocity are displayed in intervals of 0.3 m s−1 in (b) and (d). Negative contour values are dashed.

  • View in gallery

    The north–south cross sections along leg 2 of (a) airborne dual-Doppler derived vertical velocity (m s−1) overlaid on reflectivity and (b) contours of modeled snow, graupel, and rain mixing ratios in g kg−1. Shaded regions in (b) denote areas where model-predicted CLW mixing ratios exceeded 0.2 g kg−1.

  • View in gallery

    (a) Modeled accumulated precipitation from 2200 to 0100 UTC 13–14 Dec 2001 (color shading, key at right) and terrain elevation (black contours at intervals of 250 m) for the 1.33-km resolution MM5 control simulation and (b) for the smoothed-terrain version of the 1.33-km resolution MM5 simulation. Black dashed line in (a) and (b) indicates the position of leg 2 of the P3 flight track and cross section in Fig. 17b. White dashed line represents the orientation of the meridionally averaged east–west cross section displayed in Fig. 17a.

  • View in gallery

    (a) The east–west cross section of meridionally averaged QPF from 2200 to 0100 UTC 13–14 Dec 2001 from the control simulation (solid black line) and smoothed run (dashed gray line) with underlying terrain of the control simulation (black filled area) and smoothed terrain (thick gray dashed line). Orientation of cross section is shown by white dashed line in Fig. 16. (b) Same as in (a) but with QPF and underlying terrain displayed for the south–north-oriented cross section along leg 2 of the P3 track (cf. Fig. 1a). Orientation of the cross section is shown by the black dashed line in Fig. 16. Thick dashed gray line delineates the mean terrain profile for the smoothed-terrain simulation.

  • View in gallery

    Same as in Fig. 17, but over the extended interval 1400 UTC 13 Dec–0800 UTC 14 Dec 2001.

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    Three-dimensional idealized schematic of topography and wind flow over the IMPROVE study area from 2300 to 0100 UTC 13–14 Dec 2001. Blue arrows show strong southerly low θe airflow at low levels along the windward (west facing) slopes of the Cascade range which was subsequently involved in wave generation over multiple small-scale east–west-oriented ridges–valleys within the Cascade foothills. Red arrows show the high θe cross-barrier flow that surmounted the low θe air and exhibited a vertically propagating mountain-wave structure anchored to the mean north–south Cascade crest.

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Multiscale Mountain Waves Influencing a Major Orographic Precipitation Event

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  • 1 Department of Atmospheric Sciences, University of Washington, Seattle, Washington
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Abstract

This study combines high-resolution mesoscale model simulations and comprehensive airborne Doppler radar observations to identify kinematic structures influencing the production and mesoscale distribution of precipitation and microphysical processes during a period of heavy prefrontal orographic rainfall over the Cascade Mountains of Oregon on 13–14 December 2001 during the second phase of the Improvement of Microphysical Parameterization through Observational Verification Experiment (IMPROVE-2) field program. Airborne-based radar detection of precipitation from well upstream of the Cascades to the lee allows a depiction of terrain-induced wave motions in unprecedented detail.

Two distinct scales of mesoscale wave–like air motions are identified: 1) a vertically propagating mountain wave anchored to the Cascade crest associated with strong midlevel zonal (i.e., cross barrier) flow, and 2) smaller-scale (<20-km horizontal wavelength) undulations over the windward foothills triggered by interaction of the low-level along-barrier flow with multiple ridge–valley corrugations oriented perpendicular to the Cascade crest. These undulations modulate cloud liquid water (CLW) and snow mixing ratios in the fifth-generation Pennsylvania State University–National Center for Atmospheric Research (PSU–NCAR) Mesoscale Model (MM5), with modeled structures comparing favorably to radar-documented zones of enhanced reflectivity and CLW measured by the NOAA P3 aircraft.

Errors in the model representation of a low-level shear layer and the vertically propagating mountain waves are analyzed through a variety of sensitivity tests, which indicated that the mountain wave’s amplitude and placement are extremely sensitive to the planetary boundary layer (PBL) parameterization being employed. The effects of 1) using unsmoothed versus smoothed terrain and 2) the removal of upstream coastal terrain on the flow and precipitation over the Cascades are evaluated through a series of sensitivity experiments. Inclusion of unsmoothed terrain resulted in net surface precipitation increases of ∼4%–14% over the windward slopes relative to the smoothed-terrain simulation. Small-scale waves (<20-km horizontal wavelength) over the windward slopes significantly impact the horizontal pattern of precipitation and hence quantitative precipitation forecast (QPF) accuracy.

Corresponding author address: Matthew F. Garvert, Department of Atmospheric Sciences, University of Washington, 408 ATG Building, Box 351640, Seattle, WA 98195-1640. Email: mgarvert@atmos.washington.edu

Abstract

This study combines high-resolution mesoscale model simulations and comprehensive airborne Doppler radar observations to identify kinematic structures influencing the production and mesoscale distribution of precipitation and microphysical processes during a period of heavy prefrontal orographic rainfall over the Cascade Mountains of Oregon on 13–14 December 2001 during the second phase of the Improvement of Microphysical Parameterization through Observational Verification Experiment (IMPROVE-2) field program. Airborne-based radar detection of precipitation from well upstream of the Cascades to the lee allows a depiction of terrain-induced wave motions in unprecedented detail.

Two distinct scales of mesoscale wave–like air motions are identified: 1) a vertically propagating mountain wave anchored to the Cascade crest associated with strong midlevel zonal (i.e., cross barrier) flow, and 2) smaller-scale (<20-km horizontal wavelength) undulations over the windward foothills triggered by interaction of the low-level along-barrier flow with multiple ridge–valley corrugations oriented perpendicular to the Cascade crest. These undulations modulate cloud liquid water (CLW) and snow mixing ratios in the fifth-generation Pennsylvania State University–National Center for Atmospheric Research (PSU–NCAR) Mesoscale Model (MM5), with modeled structures comparing favorably to radar-documented zones of enhanced reflectivity and CLW measured by the NOAA P3 aircraft.

Errors in the model representation of a low-level shear layer and the vertically propagating mountain waves are analyzed through a variety of sensitivity tests, which indicated that the mountain wave’s amplitude and placement are extremely sensitive to the planetary boundary layer (PBL) parameterization being employed. The effects of 1) using unsmoothed versus smoothed terrain and 2) the removal of upstream coastal terrain on the flow and precipitation over the Cascades are evaluated through a series of sensitivity experiments. Inclusion of unsmoothed terrain resulted in net surface precipitation increases of ∼4%–14% over the windward slopes relative to the smoothed-terrain simulation. Small-scale waves (<20-km horizontal wavelength) over the windward slopes significantly impact the horizontal pattern of precipitation and hence quantitative precipitation forecast (QPF) accuracy.

Corresponding author address: Matthew F. Garvert, Department of Atmospheric Sciences, University of Washington, 408 ATG Building, Box 351640, Seattle, WA 98195-1640. Email: mgarvert@atmos.washington.edu

1. Introduction

Experience indicates that mesoscale models, when run at sufficiently high resolution, are often capable of accurately simulating the complex kinematic structures and underlying dynamics associated with moist flow over mountainous terrain (Colle and Mass 1996; Steenburgh and Mass 1996; Doyle et al. 2000; Cairns and Corey 2003). Additionally, increased model resolution (down to 4 km) has been shown to improve the distribution of surface precipitation over major orographic barriers (Colle and Mass 2000; Colle et al. 1999, 2000). Nevertheless, at higher resolutions there are still major problems with quantitative precipitation forecasts (QPFs), with some studies suggesting significant overprediction on the windward slopes, and often underprediction in the lee (Colle et al. 1999, 2000). Colle and Mass (2000) identified the need to make improvements to the model’s microphysical parameterizations to correct these shortcomings.

To investigate these errors in model QPF, several studies have focused on mesoscale models’ bulk microphysical parameterizations (BMPs). Such studies often have utilized two-dimensional models to investigate sensitivities of BMPs, or simply focused on the differences in surface precipitation and associated mixing ratios aloft by contrasting BMPs (Colle 2004; Thompson et al. 2004). Most of these studies, however, have lacked the comprehensive set of observations necessary to isolate errors in model parameterizations of microphysical processes from problems with basic-state kinematic fields.

Previous field studies such as the Coastal Observation and Simulation with Topography Experiment (COAST), the California Land-falling Jets Experiment (CALJET), and the Sierra Project have highlighted the importance of flow kinematics, including upstream blocking and resulting barrier jet formation, in determining the intensity and distribution of orographic precipitation. Marwitz (1987a, b) documented the presence of flow blocking and a pronounced barrier jet accompanying widespread stratiform precipitation over the windward slopes of the Sierra Mountains. Such events were associated with the approach of a baroclinic zone and strong, stable prefrontal flow impinging upon the 2-km-high barrier of the Sierras. Marwitz (1987b) and Rauber (1992) found that upstream convergence due to blocking created areas of high cloud liquid water (CLW) amounts and an associated displacement of maximum precipitation well upstream of the Sierra crest. Evidence of upstream blocking and its impact on precipitation processes has also been documented in the European Alps (Rotunno and Ferretti 2001; Medina and Houze 2003; Bousquet and Smull 2003a, b) and the Wasatch Mountains of Utah (Cox et al. 2005).

Studies over the Mongollon Rim area of Arizona showed that smaller-scale gravity waves (horizontal wavelengths of ∼20 km) generated by topography influenced cloud structure and modified precipitation distributions. Bruintjes et al. (1994) observed enhanced small-scale vertical motion, associated with the gravity waves that extended high into the atmosphere. These wave motions were excited by hills upwind of the main barrier, and produced high CLW independent of the large-scale baroclinic storm system. The enhanced CLW regions were shown to greatly affect both the rate and horizontal distribution of precipitation.

These and other studies indicate the necessity of identifying errors in the simulated kinematic fields in order to isolate possible errors in model BMPs. To that end, the second phase of the Improvement of Microphysical Parameterization through Observational Verification Experiment (IMPROVE-2) provides a unique opportunity to examine kinematic, dynamic, and microphysical processes and associated flow and precipitation structures over topography, and hence to assess the validity of present-day numerical model BMPs. The IMPROVE-2 field experiment, which took place in November–December 2001 over the Oregon Cascades, collected a uniquely comprehensive and detailed set of observations of the thermodynamic, kinematic, and microphysical structure of many precipitation events (Stoelinga et al. 2003). The collection of kinematic and thermodynamic data concomitantly with microphysical measurements provides a unique opportunity to isolate potential problems in the model BMP from model errors in the dynamic representation of such storm systems.

An exceptional means of mapping detailed flow structure within precipitation-laden layers over complex terrain is provided by airborne Doppler radar. IMPROVE-II is one of only two experiments [the other being the Mesoscale Alpine Program (MAP); e.g., Bousquet and Smull 2003a] in which airborne Doppler radar data were systematically collected during the passage of precipitating cloud systems over a major mountain barrier. The 13–14 December 2001 event was particularly well suited to mapping the flow across terrain because (i) systematically designed flight tracks were repeated on both the windward and leeward sides of the Cascade crest (Fig. 1a), providing multiple views of the slowly evolving flow/precipitation structure across the barrier, and (ii) the extension of detectable radar echo into the lee of the Cascades allowed unusually complete illumination of the subsiding branch of the mountain-wave circulation.

The 13–14 December storm system has been extensively examined in a variety of papers (e.g., Garvert et al. 2005a, b; Colle et al. 2005; Woods et al. 2005, Medina et al. 2005). The present paper builds upon this foundation by providing a far more detailed view of the three-dimensional kinematic and reflectivity fields from the Willamette Valley (∼100 km upstream of the Cascade crest) to the lee slopes. This paper utilizes the unique airborne dual-Doppler dataset to qualitatively assess the amplitude and strength of mountain waves on several distinct scales over both the leeside and windward slopes from a fully three-dimensional perspective. Output from the fifth-generation Pennsylvania State University–National Center for Atmospheric Research (PSU–NCAR) Mesoscale Model (MM5) is evaluated alongside the observations, and then employed to examine the evolution of these standing waves. A broader suite of ground-based radar measurements, special soundings, wind profiler measurements, and flight-level data is used to further evaluate the MM5’s performance. Model sensitivity studies assist in analyzing the complex interaction between the upstream flow field and the mountain wave anchored to the Cascade crest. Finally, details of the microphysical structures aloft and surface precipitation patterns are examined as they relate to these diverse mountain-wave structures.

2. Methodology and data sources

a. Airborne dual-Doppler measurements

The capability of airborne dual-Doppler radar to map 3D wind fields over complex terrain is well established (Yu and Smull 2000; Bousquet and Smull 2003a, b). Using the fore/aft scanning technique (Frush et al. 1986; Jorgensen and Smull 1993; Hildebrand 1998), 3D reflectivity and radial velocity data for the 2-h period from 2300 UTC 13 December to 0100 UTC 14 December 2001 were gathered within five overlapping ∼40-km-wide volumes centered along each of five north–south-oriented flight legs (Fig. 1a). Each of these individual legs was flown at a constant altitude and was approximately 130–140 km long, with 40-km interleg spacing in the cross-barrier (east–west) direction. Each leg required approximately 30 min to complete. The first leg was flown over the Willamette Valley at an altitude of 2 km [all heights are mean sea level (MSL)], with the three subsequent legs increasing in altitude to a maximum of 4 km for the leg flown immediately leeward of the Cascade crest. The fifth and final leg was flown at a height of 3.2 km well to the lee of the Cascade crest.

The regular spacing of these five legs provided continuous, high-resolution dual-Doppler coverage over the IMPROVE-2 study area, and is thus extremely well suited for comprehensive evaluation of mesoscale simulations of terrain-modified airflow and precipitation processes. While airborne Doppler measurements are generally regarded as quantitatively useful to a range of ∼40 km, the relatively tight spacing of the P3 legs allowed restriction of analyzed data to a zone within ∼20 km to either side (i.e., east and west) of each track, ensuring a high-fidelity analysis. Prior to interpolation, radial velocity and reflectivity measurements were subjected to automated and manual editing, as described by Bousquet and Smull (2003a), to remove ground clutter contamination, noise, and other artifacts. Radial data were then interpolated to a composite Cartesian grid measuring 240 km × 170 km in the horizontal with 1-km grid spacing in x and y, and 0.25-km resolution in the vertical from immediately above the terrain to a maximum height of 10 km MSL (or echo top, if lower). Multiple radial views from “fore” and “aft” scans were synthesized following Jorgensen and Smull (1993) to yield continuous fields of horizontal airflow (U, V) and reflectivity (dBZ) in regions of detectable echo. Synthesized reflectivity and component airflow fields were subjected to a two-pass Leise filter (Leise 1982) to eliminate poorly resolved small-scale features and to assure smooth, continuous transitions across “seams” separating data from individual flight legs. The absence of notable discontinuities at these locations attests to the success of this approach.

Rudimentary estimates of vertical air motion were performed by downward integration of the anelastic continuity equation. Downward integration serves to minimize error growth (Ray et al. 1980) and these vertical velocity estimates were uniformly bounded and well behaved at the lowest levels. The resulting composite analysis utilized in this study provides a largely uninterrupted view of kinematic and precipitation fields fully spanning the Cascade crest, extending from the Willamette Valley (located >100 km upstream) eastward to the leeside desert plateau of central Oregon. Due to the relatively slow evolution of flow and precipitation structures during this period of prefrontal P3 sampling (Garvert et al. 2005b; Medina et al. 2005), the composite analysis may be viewed as a steady-state depiction of the precipitation and flow pattern during this 2-h period.

b. Mesoscale model setup and description

The MM5 version 3.6 was employed in nonhydrostatic mode to simulate the 13–14 December 2001 system. A 36-km outer domain covering a large area of the eastern Pacific and Pacific Northwest was run for 36 h (Fig. 1b), with a 12-km nest centered over the Pacific Northwest. The model was initialized at 0000 UTC 13 December 2001 by interpolating a modified National Centers for Environmental Prediction (NCEP) Aviation Model (AVN)1 initialization for 0000 UTC 13 December 2001 to the outermost MM5 grid. The 0000 UTC 13 December AVN gridded analysis was improved by incorporating surface and upper-air observations using a Cressman-type analysis scheme (Benjamin and Seaman 1985). Additional analyses were generated every 6 h using similarly modified AVN initialization and forecast grids, and then linearly interpolated in time to provide continuous lateral boundary conditions for the 36-km domain. To ensure the most accurate simulation, four-dimensional data assimilation (FDDA) was employed during the first 24 h of the forecast on both the 12- and 36-km domains. The FDDA scheme (Stauffer and Seaman 1990; Stauffer et al. 1991) applies a Newtonian relaxation technique to nudge the MM5’s wind, temperature, and moisture fields toward modified AVN surface and upper-air initialization grids at 1200 UTC 13 December 2001 and 0000 UTC 14 December 2001.

Thirty-two unevenly spaced full-sigma levels were used in the vertical, with maximum resolution in the boundary layer. The simulation used the updated (version 3.6) explicit moisture scheme of Reisner 2 (Thompson et al. 2004), Grell cumulus parameterization (Grell 1993), and the Medium-Range Forecast Model (MRF) planetary boundary layer scheme (Hong and Pan 1996). The configuration of the model was identical to the original MM5 simulation presented in Garvert et al. (2005a) except that nudging on the outer domains was extended to 24 h, as opposed to 12 h in the original simulation. As a result, the simulated strength of midlevel wind speeds and the timing of the surface front were both improved. Yet despite the differences between these simulations, the deduced QPF biases were similar, suggesting that the magnitude and influence of errors in the model’s bulk microphysical parameterizations overshadowed those rooted in the kinematic fields.

In addition to the outer 36- and 12-km domains, a separate 4-km simulation with a 1.33-km nest centered over the central Oregon Cascades was run for 24-h starting at 1200 UTC 13 December 2001. The 4-km grid was initialized by linearly interpolating forecasts from the 12-km MM5 simulation. The extent of the 4- and 1.33-km domains are also shown in Fig. 1b. The 1.33-km domain was expanded from the original dimensions employed by Garvert et al. (2005a) in order to encompass a larger area to the east (leeward) of the crest, as well as the coastal mountains to the west. These inner domains did not employ nudging or cumulus parameterization.

c. 13–14 December 2001 case overview

Garvert et al. (2005a) present a detailed account of the evolution of the synoptic fields for the 13–14 December 2001 storm event, including validation against the MM5 control simulation. Although this present paper will predominately focus on a 2-h period from 2300 UTC 13 December to 0100 UTC 14 December, a brief overview of the kinematic and thermodynamic evolution of the storm is included here to provide a context for the mesoscale features described in the subsequent sections.

At 2000 UTC 13 December 2001, a large area of high potential temperature (θ) was located offshore of the West Coast at heights of 1 and 3 km (Figs. 2a,b). East of the high-θ region, there was a weak gradient in θ, with values over western Oregon being 3–4 K lower than those offshore. Coincident with this gradient was a transition in wind direction from westerly offshore to more south-southwesterly near the coast. A second stronger gradient in θ was evident entering the domain from the northwest, and was associated with the approaching forward-tilted cold front as described in Garvert et al. (2005a) and Woods et al. (2005).

An 800-km-long east–west cross section (location shown in Fig. 2) was constructed to illustrate the vertical structure and evolution of the large-scale thermodynamic and kinematic fields (Fig. 3). The cross section showed that at 2000 UTC a cold front (position indicated by dashed gray line) was well offshore (Fig. 3a). West of the front, cold air advection was occurring. To the east of the front, a relatively broad and uniform area of warm air advection was present marked by veering of wind direction with height. There was no indication in this cross section of enhanced southerly flow (i.e., a barrier-parallel jet) over the windward slopes as would be expected in a case of significant blocking. A significant mountain wave was evident, with persistent sharp descent and subsequent downstream rebound of the θ contours to the lee of the Cascade crest. Radar measurements at this time showed a large shield of stratiform precipitation over the study area (cf. Fig. 3a of Garvert et al. 2005a).

Three hours later at 2300 UTC, the low-level θ gradient near the coast had intensified and progressed onshore with southwesterly flow persisting over the study area (Fig. 2c). The leading edge of the cold advection had moved eastward but still was about 150-km west of the study area at 1 km and near the coast at 3 km (Figs. 2c,d). The east–west vertical cross section at this time (Fig. 3b) shows that east of the cold front, the slope of the θ contours had increased significantly. Additionally, over the windward slopes of the Cascades, enhanced vertical gradients in θ (gray shading in Fig. 3b) and equivalent potential temperature (θe) were present between 2 and 4 km. Substantial directional shear was present within this zone, with winds veering from southerly at low levels to a more westerly direction above the enhanced thermal (θ and θe) gradient.

Measurements of the upstream flow profile, provided by the NCAR Integrated Sounding System (ISS) profiler located at Irish Bend (IB in Figs. 1a and 3b), support the MM5’s depiction of substantial low-level veering in the winds during the prefrontal period. Figure 4 displays profiles of U- and V-wind components at IB at 15-min intervals from 2247 UTC 13 December to 0047 UTC 14 December 2001, as well as an average profile for the entire 2-h period. Although there was some variance in the U and V components, these profiles remained relatively unchanged over the 2-h period. Consistent with the simulated veering at low-level winds, the U component at IB increased from about 5 m s−1 near the surface to an average of 30 m s−1 at 2.5 km, while the V component reached a maximum speed of ∼23 m s−1 at 1.5 km before weakening aloft. Above 2.5 km both the U and V components were relatively uniform with a modest 4–5 m s −1 increase.

By 0200 UTC the cold front had entered the study area (Figs. 2e,f). The vertical cross section shows that winds throughout the lowest 6 km were from the west behind the front (Fig. 3c). Due to the modification of the θ and θe gradients by mountain waves over the Cascades, the position of the cold front at upper levels is difficult to ascertain.

3. Observed and modeled flow kinematics over the Oregon Cascades

a. Upstream conditions

The basic structure of mountain waves is determined by the size and shape of the underlying terrain and by the vertical profiles of temperature, cross-barrier flow speed, and moisture of the impinging flow. A special sounding (located at UW in Fig. 1a) taken at 0000 UTC 14 December 2001 measured environmental conditions just upstream of the Cascade crest (Fig. 5). Although this sounding was likely somewhat modified by the Cascades and coastal mountains (Medina et al. 2005), it provides the best available source of information on the thermodynamic structure directly upstream of the elevated terrain of the Cascades.

Similar to the IB profiler, the observed U component increased significantly from the surface to a height of 2.9 km, reaching a peak speed of approximately 38 m s−1. Above this strongly sheared layer, U was relatively uniform up to 4.5 km, above which it again began to increase with height. The MM5 1.33-km simulation essentially replicated this structure, but slightly overpredicted the strength of U within the shear layer below 2 km, while underpredicting the magnitude of U above, resulting in an underprediction of the vertical shear magnitude over the windward slopes.

The dry Brunt–Väisälä frequency (N2d) calculated from the sounding indicates the profile was stable throughout the lowest 6 km (Fig. 5). An additional calculation of the moist Brunt–Väisälä frequency (N2m; Durran and Klemp 1982; Einaudi and Lalas 1973) was also performed, since the atmosphere was saturated below 600 hPa (Garvert et al. 2005a). The inclusion of moisture in the moist Brunt–Väisälä calculation accounts for the latent heat of condensation and thus provides a better indication of the stability of the upstream flow in a saturated environment (Durran and Klemp 1982). Throughout most of the profile, N2m remained weakly positive, indicating neutral to stable flow. A layer of increased stability near the top of the U-shear layer was present between 2 and 3 km, where the N2d values peaked near 2 × 10−1 s−2 (Fig. 5). The 1.33-km MM5 simulation also reproduced the observed stratification, including the enhanced stability between 2 and 3 km.

Linear theory can be used to predict the character of mountain waves provided that the barrier height is small compared with the vertical wavelength of the resulting wave (Durran 2003). The “nonlinearity parameter” or inverse Froude number [viz. Nh/U, where N is the dry or moist Brunt–Väisälä frequency, h is the average height of the mountain (assumed here to be 1.68 km), and U is the cross-barrier wind] was calculated for the sounding in Fig. 5. In the surface layer below 0.75 km, Nmh/U was greater than unity while above this layer the nonlinearity parameter remained close to 0.5 for both moist and dry N2 values. Theory predicts that a nonlinearity parameter value of 0.5 is associated with a linear (i.e., “flow over”) mountain-wave regime (Smith 1989).

b. Windward slopes

At a height of 1.5 km MSL, composite airborne Doppler measurements showed southwesterly flow over the Willamette Valley (Fig. 6a). As this flow encountered the elevated terrain of the Cascade foothills, its speed markedly decreased. Observed reflectivity showed relatively heavy precipitation over the entire study area, with pockets of enhanced reflectivity reaching 40 dBZ over the windward slopes and a sharp reduction of values leeward of the crest. (Fig. 6b). Winds at 3 km (i.e., above crest level) were more westerly (Fig. 6c), with a pronounced maximum in wind speeds to the lee of the Cascade crest. The observed veering of winds between 3.0 and 4.5 km (cf. Figs. 6c,e) was consistent with warm advection ahead of the frontal system, as previously seen in Fig. 3b. Reflectivity patterns seen in Figs. 6b,d,f will be more extensively discussed in section 4.

Averaged east–west cross sections of the radar-observed U- and V-component flow provide a general view of kinematic fields across the Cascades (Figs. 7a,c). These sections were constructed by meridionally averaging the gridded 1-km dual-Doppler data over the 140-km north–south region between 44.0° and 45.0°N (boxed region in Fig. 6e). Analogous mean cross sections of simulated flow (Figs. 7b,d) were created using the U and V fields from the MM5 1.33-km simulation. The MM5 cross sections were averaged for the time period of the P3 flight, from 2300 to 0100 UTC 13–14 December 2001 (forecast hours 11–13) using model output at 15-min intervals.

As previously suggested by the UW sounding and IB profiler, airborne dual-Doppler data indicate that the incoming flow was strongly sheared upwind of the crest, with the U component increasing more than 20 m s−1 from the surface to the top of the shear layer at approximately 2-km MSL (Fig. 7a). The depth of this shear layer remained relatively constant as flow ascended the windward slopes. Above the shear layer, a jet of stronger (>36 m s−1) cross-barrier flow centered at 3–4 km progressively increased strength over a 90-km interval upstream of the mean crest, ultimately reaching speeds greater than 40 m s−1 over and immediately leeward of the crest. The U-wind (cross barrier) flow structure documented over the windward Cascade slopes differs appreciably from cases of profoundly blocked flow observed adjacent to taller mountain ranges such as the Sierras (Marwitz 1987a) and Alps (Bousquet and Smull 2003a, b). In those events, radar and in situ measurements depicted the cross-barrier shear layer as being mainly concentrated below the crest level, with the depth and strength of the shear layer decreasing with proximity to the crest (cf. Fig. 3c of Marwitz 1987a; Fig. 8b of Bousquet and Smull 2003b).

In the present case, strong V-component (along barrier) flow was coincident with the aforementioned U-component shear layer and decreased with height above this layer (Fig. 7c). There was a ∼4 m s−1 enhancement in the V flow over the Willamette Valley (∼100 km upstream of the crest), but the absence of a clearly defined V maximum (or “barrier jet”) below crest level further distinguishes this IMPROVE-2 case from aforementioned examples of profoundly blocked flow.

To better illustrate the alteration of the wind components by terrain, “upstream” U and V profiles were created by averaging the dual-Doppler wind fields over a section of the Willamette Valley between −123.30° and −123.06°W (dashed boxed region at left edge in Fig. 7a). Since the elevation of the Willamette Valley is near sea level, the upstream profile’s wind values in MSL are roughly equal to those above the ground level (or AGL). The upstream U (and V) AGL profiles were subtracted from the corresponding U (and V) values AGL over the rest of the cross section in Fig. 7 to create the U ′ and V ′ plots in Fig. 8. Using this methodology, the frictional deceleration of the U flow following the sloping terrain is largely eliminated, allowing other terrain influences such as blocking to be better illuminated.2

Immediately above the foothills, U ′ values were predominantly between −1 and 1 m s−1, indicating that, excluding height-dependent frictional effects, the cross-barrier flow at the lowest radar-sampled levels was relatively unperturbed (Fig. 8a). In contrast, large positive U ′ values were concentrated over the lower lee slopes and extended westward to a height of 6 km over the crest. This pattern can be attributed to a vertically propagating mountain wave that will be discussed in more detail in the next section. For the along-barrier flow, a small area of positive V ′ values of 5–7 m s−1 was centered 100 km upstream of the crest below 2.0 km (Fig. 8b). Otherwise, V ′ values over the windward slope were close to zero, corroborating observations in Fig. 7c that failed to indicate a pronounced barrier jet or other signs of pronounced upstream blocking. The observed flow patterns and computed flow perturbations thus support the calculations of the nonlinearity parameter in Fig. 5, which predicted a flow-over regime in which blocking effects are minimal. Based upon idealized high-resolution 2D simulations, however, Medina et al. (2005) concluded that that some modification of the cross-barrier during 13–14 December 2001 may be attributed to nonlinear interactions between stably stratified flow and the orographic barrier.

MM5-averaged east–west cross sections of U and V are displayed in Figs. 7b,d. The MM5 successfully simulated the broad characteristics of the flow, including the low-level shear in the U component and core of stronger U above. The model also correctly simulated the core of stronger V component at low levels rising over the windward slopes, as well as the observed decrease of V with height above 2–3 km. Over the windward slopes, the strength of U shear was underpredicted, with low-level U values being too strong while those above 2 km were 5–8 m s−1 too weak. Additionally, the height of the modeled U-wind shear layer over the windward slopes was about 0.5 km less than observed. Low-level V-component flow was 4–6 m s−1 stronger than that observed, with the elevated core of maximum V values extending farther east toward the lee than observed. The implications of these errors on the structure and strength of mountain waves will be explored in the section 3c.

Simulated θ and θe contours are also plotted in Figs. 7b,d. Since the environment during the 13–14 December case was near saturation up to 4 km (Garvert et al. 2005a), θe can be treated as a tracer of the steady-state flow, except in areas where diabatic processes such as melting are active. Enhanced vertical gradients of θ (dashed white lines) and θe (solid black lines) were found surmounting the zone of weaker low-level U winds and within the zone of enhanced U aloft (Fig. 7b). This corresponds well with the UW sounding shown in Fig. 5 (position of UW labeled in Figs. 7a,b), which shows the strongest N2 values (and increased stability) between 2 and 3 km coinciding with the top of the shear layer. Stronger V-component flow was centered below the θe gradient and rapidly decreased with height above (Fig. 7d). Decreasing values of θ toward the crest are consistent with reflectivity observations from the P3 radar that showed a tendency for the bright band (and the associated 0°C level) to descend with approach to the Cascade crest, as will be shown in section 4. A MM5 simulation in which diabatic effects of melting were excluded (not shown) indicated, however, that melting had a minimal influence on θ values on the windward slopes.

Trajectories were calculated using the 4-km domain MM5 simulation to determine the origin of the θe and θ gradient present in Figs. 7b,d. Figure 9a shows the paths of four trajectories ending at 2300 UTC 13 December (forecast hour 12). All terminate at the same horizontal location (44.3°N, −122.3°W) but at heights ranging from 1 to 4 km (cf. Fig. 7b). These trajectories were calculated using 600-s time steps from MM5 4-km hourly output. Lagrangian values of pressure, θ, θe, and relative humidity values for the four trajectories are displayed in Fig. 9b.

The parcel ending at 1 km moved northward nearly parallel to the Cascades, at a relatively constant pressure level of 910 hPa from 1200 to 2300 UTC 13 December. This parcel was subsaturated from 1200–1800 UTC 13 December, although relative humidities increased from 75%–90% along its northward track. After 1800 UTC, the parcel’s relative humidity remained close to 100% as it neared the zone of enhanced orographic precipitation over the IMPROVE-2 study area. The values of θe and θ for this parcel showed a 5–6 K increase during its northward trajectory, apparently as a result of terrain-induced mixing. The parcel ending at 2 km followed a similar course, originating to the south-southwest off the coast of California. This parcel remained around 950 hPa between 1200 and 2000 UTC, but then rose to level of 784 hPa as it encountered the higher terrain of the Cascade foothills. The θ and θe values along this trajectory also experienced a 5 K increase during the 4-h interval plotted, again indicative of warming induced by mixing.

The two parcels that ended at 3 and 4 km both originated within a distinctly different air mass and region than the lower trajectories. Both originated to the west-southwest of the study area over the open waters of the Pacific, in a nearly saturated air mass characterized by higher θe. The 3- and 4-km parcels experienced significantly more ascent than the lower parcels, rising over 225 hPa in less than 5 h. The θe of these elevated parcels remained relatively steady, while θ dramatically warmed due to release of latent heat of condensation. These results imply that the θ and θe gradients over the windward slopes in the prefrontal regime were the product of dissimilar airmass origins in combination with sharply veering flow upstream of the Cascade terrain.

c. Leeside region

The extension of radar-detectable precipitation into the lee of the Cascades presents a unique opportunity to examine the strength and amplitude of a standing wave anchored to a mountain barrier. In association with strong cross-barrier flow surmounting the Cascades, dual-Doppler observations showed this high-momentum air plunging downward immediately to the lee of the crest (Fig. 7a). The observed 32 m s−1 contour, which approximately delineates the top of the U-shear layer, dropped from a maximum height of 3.0 km just upstream of the crest to 2.1 km MSL (1.0 km AGL) approximately 10 km downstream (Fig. 7a). The 32 m s−1 contour then rebounded slightly, before leveling out at about 2.2 km. Independent ground-based observations around this time indicated strong winds reaching the surface, with reports of downed trees in the immediate lee of the crest (Medina et al. 2005). The aforementioned positive U ′ perturbations in the lee of the mountain crest (Fig. 8a) are consistent with the presence of a vertically propagating mountain wave (Durran 2003).

The MM5 also depicted a core of strong U momentum ascending the windward slopes and plunging downward in the lee of the Cascades (Fig. 7b). The modeled 32 m s−1 contour line dropped by approximately 1.2 km over a horizontal distance of 10 km, which is comparable in amplitude to the radar-observed displacement. However, the simulated core of strong U momentum penetrated much closer to the ground and spanned a larger horizontal downstream distance than was observed. Such errors in the model depiction of the mountain wave can increase the amount of precipitation spillover reaching the ground, and may contribute to an overprediction of precipitation in the leeside region as discussed in Garvert et al. (2005a).

Previous research has shown that the structure and amplitude of mountain waves are sensitive to the depth of the boundary layer and upstream shear (Smith 1979; Peng and Thompson 2003). As discussed above, the model underpredicted the strength and depth of the shear layer over the windward slopes of the Cascades. Since the depth of the modeled shear layer was less than that observed, this upstream effect might account for subsequent errors in the model depiction of the mountain-wave structure over and leeward of the crest.

To assess the impact of the planetary boundary layer (PBL) representation and associated shear on the mountain wave, a variety of sensitivity tests were conducted using different PBL parameterization schemes available in MM5. As noted in section 2, the control simulation was run with the MRF-PBL, which utilizes the Troen–Mahrt representation of the countergradient term (Hong and Pan 1996). A second run was performed that was identical to the control except that the Eta-PBL was used on the inner 4- and 1.33-km domains. The Eta-PBL, which is used in the operational Eta Model (Janjic 1990, 1994), is a 1.5-order closure scheme in which turbulent kinetic energy (TKE) is predicted through a prognostic energy equation, as adapted from the Mellor–Yamada scheme.

The vertical extent of the low θe and associated depth of the shear layer over the windward slopes was further reduced in the Eta-PBL simulation compared to the control run (Figs. 10a,b), contrasting even more with the observations Fig. 7a. Yet, the cross-barrier shear was stronger than that in the control simulation, better matching observations. In contrast to the dual-Doppler observations, the Eta-PBL simulation brought unrealistically stronger U-momentum values even closer to the surface and over a wider area to the lee of the Cascade crest. Additionally, the wave structures as depicted by vertical velocity in the Eta simulation were significantly different than those in the control, with higher amplitude and downward displaced vertical motions in the lee (Figs. 10b,d).

A third simulation was performed using a modified version of the MRF-PBL (referred to an Yonsei PBL), as outlined in Noh et al. (2003). Inclusion of the Yonsei-PBL produced results similar to the Eta-PBL simulation, reducing the depth of the upstream shear layer and increasing the amplitude of the lee-wave response (Figs. 10e,f). Inclusion of the Yonsei-PBL, however, did not correct the underprediction in the strength of the upstream shear layer as significantly as the Eta simulation.

4. Observed reflectivity patterns and modeled precipitation fields

a. Larger-scale (>20 km) precipitation features

A meridionally averaged east–west cross section of radar reflectivity was constructed from the composite airborne Doppler analysis (Figs. 6b,d,f) and is shown in Fig. 11a. The observed slope of the distinct bright band indicates that the melting layer decreased in height from 2 km over the Willamette Valley to just 1.5 km over the elevated terrain, consistent with the presence of cooler air on the windward side of the Cascades (as previously suggested by MM5-based depictions; cf. Fig. 7b). Radar reflectivities decreased rapidly downstream of the crest, indicative of large evaporation and/or fallout rates. Locally increased reflectivity values were observed over the Willamette Valley (∼100 km upstream of the crest), the windward foothills (centered ∼70 km upstream of the crest), and within a broad area commencing ∼44 km upstream of the crest and extending into the lee. These structures are also clearly evident in the composite airborne Doppler radar analyses at 3 km (Fig. 6d).

To explore the origins of these prominent reflectivity features, model-predicted CLW and snow (qs) fields were averaged using 15-min interval MM5 output from 2300 to 0100 UTC 13–14 December 2001 and overlaid on the observed Doppler reflectivity fields (Figs. 1la,b). By comparing these model-based mixing ratios with the Doppler radar observations, the origins of these precipitation structures can be better ascertained. The time-averaged model fields indicated increases in the amount of CLW at 100 and 70 km upstream of the crest, over virtually the same regions where the reflectivity perturbations were observed (Fig. 11a). The fact that these perturbations in CLW remain clearly evident even after averaging over a 2-h period indicates the stationary and persistent nature of these perturbations, an aspect also evident in ground-based radar data collected during this interval (cf. Figs. 8a,b of Medina et al. 2005).

Maximum values in the simulated snowfield (>0.9 g kg−1) extended from 40 km upstream of the crest to the immediate lee, and were collocated with the observed higher reflectivity values aloft (Fig. 11b). The vertically pointing National Oceanic and Atmospheric Administration (NOAA) Environmental Technology Laboratory (ETL) S-band radar (MB in Fig. 1a), positioned about 20 km upstream of the Cascade crest, also showed a layer of enhanced reflectivity between 4 and 5 km that remained relatively stationary over the radar site for the duration of the prefrontal precipitation event (cf. Fig. 2 of Garvert et al. 2005b). These high snow mixing ratios were then advected over the crest by the strong cross-barrier flow before being shunted earthward in the immediate lee of the Cascades (Fig. 11b) by the strong downward motion (P3 in situ measurements reaching −4 m s−1; Fig. 12 of Garvert et al. 2005a) within the deep, vertically propagating gravity wave. Additionally, as snow reached the 0°C level, enhanced reflectivity associated with melting likely contributed to the observed reflectivity maximum in the immediate lee. In an analysis of the modeled and observed microphysical processes associated with the 13–14 December 2001 case, Colle et al. (2005) noted that in the immediate lee of the Cascades melting snow and graupel were the predominant contributions to rain generation and precipitation at the surface. The simulated large gradient of snow mixing ratios in the lee of the Cascades corresponds quite well to the rapid reduction of observed reflectivity values.

To further trace the origins of these observed reflectivity structures, the depictions of the east–west cross section of simulated CLW were expanded westward by 80 km to include the coastal mountains, as shown in Fig. 12a. Two additional MM5 1.33-km simulations were performed, one with the coastal mountains removed (designated as “Noco” in Figs. 12c,d), and another employing a lower 36-km resolution version of the topography (“smooth” in Figs. 12e,f). These runs were identical to the control simulation except for the differences in the underlying topography.

Figures 12a,b show an enhanced area of CLW and strong upward motion between 180 and 140 km upstream of the crest in association with the incoming flow encountering the coastal mountains. The lack of observations over this area prevents an independent evaluation of the veracity of this particular simulated feature. An area of descent and reduced model CLW was collocated with the zone of lower-reflectivity values observed in the lee of the coastal mountains (i.e., 120–140 km upstream of the Cascade crest). A significant rebound both in the simulated CLW and observed dBZ fields was evident over the relatively flat Willamette Valley, centered ∼100 km upstream of the Cascade crest. As seen in Figs. 12c,d, the strong vertical velocity oscillations apparent between 180 and 120 km (i.e., over the coastal mountain zone) in the control run were absent in the Noco simulation, being replaced by a broad but weaker (0–0.2 m s−1) updraft. Nonetheless, there is a significant response in the model CLW field at low levels below 2 km between 160 and 180 km upstream of the crest, which is attributed to frictional convergence as the flow over the Pacific slowed over the land. Without the coastal mountains, the upward vertical velocity and CLW maximum over the Willamette Valley (100 km upstream of crest) failed to develop. It is also noteworthy that within the smooth run, which contains a rudimentary representation of the coastal mountains, the enhancement of CLW over the Willamette Valley remained evident. These differences suggest a lee-wave response downwind of the coastal mountains effectively enhancing precipitation over the Willamette Valley.

Over the Cascade foothills, the CLW and vertical velocity fields in the Noco simulation are almost identical to those in the control simulation, indicating that the perturbations in the observed reflectivity fields over these areas cannot be attributed to the coastal mountains (Figs. 12c,d). The areas of enhanced model CLW, vertical velocity, and observed reflectivity were all collocated with a significant rise in the underlying topography located approximately 70 km upstream of the Cascade crest. The vertical velocity and CLW fields from the smooth run (Figs. 12e,f) show fewer short-wavelength features than seen in the control simulation (Figs. 12a,b). The smoothed-terrain run thus illustrates the sensitivity of windward mountain waves and microphysical responses to the local steepness and details of the underlying terrain.

b. Smaller-scale (<20 km) precipitation features

Thus, far this paper has focused on the larger-scale (horizontal length scales >20 km) precipitation and vertical velocity perturbations within a meridionally averaged east–west cross section spanning the Cascade crest. Yet as established by Garvert et al. (2005a, b), the strong southerly flow at low levels interacted with the complex terrain of the windward slopes and significantly impacted CLW amounts at smaller scales (<20 km). The top three panels in Fig. 13 show in situ, dual-Doppler-derived, and model-predicted U, V, and W components, respectively, at a height of 2.5-km along leg 2 of the P3’s flight track (as shown in Fig. 1a and by black dot 60 km upstream of the mean crest in Figs. 7a,c). The in situ and Doppler-derived U and V components along leg 2 indicate that the winds varied by 5–7 m s−1 over horizontal distances of 10–15 km in this region. These variations can be directly linked to details of the underlying terrain, with positive V-component perturbations collocated with local peaks and localized stronger U maxima located above valley locations (Garvert et al. 2005a). The MM5 correctly simulated these variations, except for slightly overpredicting their amplitude relative to that observed.

In situ vertical velocities along leg 2 also show evidence of significant terrain-induced variations. As discussed in Garvert et al. (2005a), upward vertical velocities coincided with areas in which the strong V-component flow impinged on higher terrain while negative downward perturbations were preferentially located on the lee (north) side of the ridges and coincident with the stronger V-component flow. Such perturbations have significant impacts on the CLW field, as discussed in Garvert et al. (2005b) and illustrated in Fig. 13. The model has also been shown to accurately represent the magnitude and positions of these vertical velocity oscillations but overpredict the resulting enhancement of CLW (Garvert et al. 2005b).

As described in section 2b, crude dual-Doppler estimates of vertical air motion (w) were derived by downward integration of the anelastic continuity equation subject to a boundary condition of w = 0 immediately at echo top. In the presence of large-amplitude vertically propagating waves, this assumption may be violated and represent a source of error. Nonetheless, comparison of these w estimates by independent observational measures (and simulation results as well) lends increased confidence in their accuracy. In keeping with the inherently coarser spatial resolution of airborne Doppler radar data after required filtering (viz. resolved horizontal scales >8–10 km), the derived vertical velocities were significantly smoother than those directly obtained via flight level measurements, but still showed generally good agreement in terms of the phase and sign of longer-wavelength features of the simulated flow (Fig. 13). In the following discussions comparisons of Doppler-derived vertical motions with simulation results are purely qualitative, relating the approximate location and phase (sign) of vertical motions relative to terrain corrugations.

Doppler-derived vertical velocities are overlaid on the Doppler-observed U- and V-component fields along the north–south cross section of leg 2 in Figs. 14a,c. The strongly sheared U-component flow is clearly depicted below 3 km, with near-surface values of 8 m s−1 increasing to 36–40 m s−1 at a height of 3.5 km. Within this strongly sheared layer, the observed V component reached nearly 30 m s−1. As this strong V flow impinges on the foothill terrain, the strong vertical velocity perturbations seen in the in situ data are also clearly evident in the Doppler-derived w field. Despite the imposed rudimentary boundary condition of w = 0 at echo top, positive w > 0.2 m s−1 are registered at heights of 5 km, indicating the significant upward penetration of these small-scale mountain waves generated over the Cascade foothills.

Due to the presence of strong vertical shear and variable thermodynamic conditions in this case, an accurate calculation of the Scorer parameter (relevant to vertical wave propagation) is difficult to accomplish. Using available sounding and Doppler radar data, values of Nm and V-component flow within the lowest 3 km are estimated to have been near 0.0077 and 22 m s−1, respectively, yielding a Scorer parameter (l = N/V) of 3.5 × 10−4 m−1, and a corresponding critical horizontal wavelength of ∼18 km. Flow undulations generated by the underlying Cascade foothill topography having wavelengths of <18 km are thus predicted to be evanescent and decay rapidly with height. As seen in Fig. 14, the terrain within the north–south cross section along leg 2 exhibits significant small-scale variability, implying that most waves triggered over the Cascade foothills were likely evanescent. Yet there is some evidence, especially near 30–40 and 70–80 km in Fig. 14, that waves induced by the widest ridge–valley pairs may have exceeded this cutoff wavelength. Moreover, the presence of marked meridional flow accelerations and collocated subsidence in the lee of these localized foothill barriers is consistent with theoretical depictions of vertically propagating mountain-wave behavior (e.g., Smith 1979).

The vertical penetration of such waves may also be influenced by shear. Assuming hydrostaticity, Smith (1980) demonstrated that wave components having phase lines perpendicular to the flow penetrate farther aloft (and consequently generate more upslope condensation) than those that are more oblique. The 13–14 December 2001 storm embodied a strongly sheared southerly wind component at low levels, which would tend to favor limited upward penetration of topographically generated waves having east–west axes (such as those sampled along leg 2). At somewhat higher levels, where prevailing winds veered to become more westerly, damping of these waves would be expected to be consistent with patterns seen in Fig. 14. At the same time, this elevated layer would be expected to favor upward penetration of waves triggered at higher elevations and having north–south axes, such as the profound mountain wave anchored to the Cascade crest shown in Fig. 10. Possible impacts of these contrasting behaviors on surface precipitation rates are discussed in the following section.

The MM5 U- and V-component flows along leg 2 (Figs. 14b,d) are in good agreement with observations, showing significant shear both in U and V below 3 km. As mentioned previously, however, the simulation exhibited notable errors in this layer, with the shear layer in U being shallower and weaker than observed and low-level V winds being 4–8 m s−1 too strong below 3 km. Despite these errors, the model’s vertical velocity field compares well with observations, simulating the dramatic oscillations associated with the strong V values interacting with the complex terrain. These simulated w oscillations also show greater vertical penetration than observed by the Doppler-derived values, which is probably closer to reality given the artificial boundary conditions w = 0 at echo top imposed on the dual-Doppler analysis.

To illustrate the impact of the wave-induced motions on precipitation processes, Fig. 15a displays Doppler-derived vertical velocities overlaid on reflectivity along leg 2. Consistently high dBZ values are located near a height of 1.8 km in conjunction with the radar bright band. Enhancement of radar reflectivities repeatedly occurred at distances 5–8 km north (downstream) of strong positive vertical velocities. For example, at 15 km from the southernmost edge of the cross section and near a height of 2.0 km, a strong vertical velocity maximum is evident where the strong V-component flow impinged upon a 1-km-high ridge. At a distance 5–8 km north of this positive vertical velocity perturbation, an enhancement in the dBZ values is observed, implying increased precipitation rates triggered by this ascent.

Model depictions of the precipitation and CLW fields along leg 2 showed the complex microphysical interactions produced by the strong vertical velocity perturbations (Fig. 15b). Pockets of high CLW (mixing ratios >0.4 g kg−1) are present over the individual ridges coincident with simulated regions of ascent of the areas of strong upward vertical velocity. A large snowfield above these CLW perturbations is simulated, with some enhancement in the snow mixing ratios appearing above the individual ridges. Where the snowfield intersected with the high CLW pockets, riming of the snow and graupel formation above the freezing level was diagnosed. The higher fall speeds of the rimed snow and graupel evidently resulted in the enhancement of the observed bright band over and immediately leeward of the crest as these particles fell through the melting layer. Additionally, as precipitation particles were advected northward in the low-level flow, they encountered strong downdrafts in the lee of the ridges and rapidly fell below the melting layer (Fig. 15a). A corresponding increase in surface precipitation rates is illustrated by higher mixing ratios of rain (>0.4 g kg−1) over the crest and in the immediate lee of several of the ridges along leg 2. The precipitation distribution at the surface was evidently very sensitive to the amount of riming, the fall speed of the particles, and the phase and amplitude of these small-scale terrain-induced waves.

c. Surface precipitation patterns

Using an MM5 simulation virtually identical to the one considered here, Colle et al. (2005) calculated a windward precipitation efficiency (PE; total hydrometeor fallout over the windward slopes divided by the total amount water vapor loss in this region) of 50% for the 13–14 December case. This percentage was less than the PE of 80% calculated over the Sierra Mountains (Colle and Zeng 2004a, b), with the difference attributed to the narrower width of the Cascades and strong cross-barrier flow during the 13–14 December event. To assess the influence of orography on cumulative precipitation at smaller spatial scales, as required for forecasts within individual drainage basins, the results in section 4b indicate that one must consider the impact that mountain waves and associated microphysical processes exert on surface precipitation rates. For this purpose, the 3-h cumulative QPF from the 1.33-km MM5 control simulation during the interval 2200–1000 UTC 13–14 December 2001 is presented in Fig. 16a, while a 1.33-km smoothed-terrain simulation’s QPF for the same 3-h period is shown in Fig. 16b. The terrain used for the smoothed 1.33-km model simulation is identical to that used in the 12-km outer domain. Previous studies have shown that in order for the model to simulate correctly the small-scale waves over the windward Cascade slopes, use of a model grid spacing of 1.33 km with high-resolution terrain is required (Garvert et al. 2005a). By smoothing the terrain and holding the grid spacing of both simulations constant at 1.33 km, variations in precipitation amounts can be more directly attributed to the differences in terrain.

The control simulation shows regions of higher precipitation along the coastal mountains at the western edge of the domain, while over the relatively flat Willamette Valley amounts are reduced and more uniform (Fig. 16a). Over the windward slopes of the Cascades, significant precipitation variations are indicated, with MM5-derived totals ranging from >3.0 cm over the local crests to 1–1.5 cm in some of the valleys (Fig. 16a). The precipitation distribution over the windward slopes is consistent with the findings from a linear model presented in Smith et al. (2005) that used a 1-km grid spacing and prescribed environmental values of wind direction, moist stability, wind speed, and moisture. Smith et al. (2005) noted that the small-scale structure of the precipitation fields (see their Fig. 6) with significant variability on scales of 20 km over the windward slopes is similar to that shown in Fig. 16a. Smith et al. concluded that for waves of horizontal scales on the range of 10–30 km, the effect of condensate advection slightly dominates any upstream wave tilt that may occur, with the resulting maximum precipitation being located over the highest terrain, albeit with some spillover similar to that seen in Fig. 16a. The east–west alignment of the MM5 precipitation maxima over the windward slopes is consistent with the linear model’s precipitation distribution when using a prescribed upstream wind sounding from a southerly direction (see Fig. 6c of Smith et al. 2005). The V component of the wind favors the upward penetration of wave components with phase lines perpendicular to the airflow (east–west direction) generating more upslope condensation and precipitation (Smith et al. 2005).

The smoothed-terrain simulation shows far more uniform spatial distribution of precipitation over the windward slopes, with precipitation totals consistently near 1.5–2 cm (Fig. 16b). A slight reduction is evident 10–20 km upwind of the crest in both the control and smoothed simulations, while immediate leeward of the Cascade crest precipitation totals are in excess of 3 cm for the control simulation and 2.5–3 cm for the smoothed simulation. Both model runs show a significant dropoff in precipitation amounts farther to the lee, with a sharper gradient appearing in the control simulation.

To highlight differences in the precipitation distributions for these two runs, a meridionally averaged east–west cross section of QPF was calculated for the smoothed-terrain and control simulations (Fig. 17a). Despite significant local differences in the detailed horizontal distributions of precipitation for these two simulations, integrated domainwide precipitation amounts resulted in total precipitation amounts that were very similar: 294 cm for the control simulation and 289 cm for the smoothed simulation. This indicates that the differences in topography did not significantly alter the total amount of precipitation over the domain as a whole, but instead horizontally redistributed this precipitation.

A significant difference in precipitation totals for the two simulations was evident over the windward slopes between 100 and 20 km upstream of the crest, where the control simulation’s meridionally averaged precipitation totals were 4%–17% higher than those in the smoothed simulation. Integrating over the windward slope region (area delimited in Fig. 17a), a 7% increase was observed in the control’s precipitation amount when compared with the smoothed-terrain simulation.

A south–north cross section of the 3-h QPF along leg 2 (black dashed line in Figs. 16a,b) as derived from both the control and smoothed simulations is displayed in Fig. 17b. The control simulation shows significant local variations in QPF, with higher totals over the ridges and lower QPF within the valleys. These precipitation extremes correspond closely to the variations in mixing ratios and Doppler reflectivities seen in Fig. 15. By contrast the smoothed simulation’s QPF pattern along leg 2 failed to resolve the smaller-scale precipitation perturbations caused by the complex terrain. When integrating the QPF along leg 2, the accumulated precipitation total for the control simulation was 171 cm, as compared to the smoothed-terrain run’s total of 161 cm. This represents a 6% increase in precipitation amounts that, ignoring any large-scale differences in the kinematic fields between the two simulations, may be attributed to the small-scale terrain-induced waves and ensuing microphysical processes.

Another interesting feature of the precipitation distribution is the reduction in totals near the crest and concomitant increase in the immediate lee within the control simulation (Fig. 17a). This modulation is apparently caused by the strong mountain wave and associated acceleration of the cross-barrier flow over the Cascade crest, which acted to rapidly advect precipitation toward the lee slopes. The increased QPF in the lee occurs within the area where the MM5 simulated the strong U flow reaching closer to the surface than observed. However, the QPF minimum near the crest and maximum to the lee is consistent with the Doppler reflectivity patterns described in Fig. 11b. The smoothed run showed a similar pattern in the surface precipitation, with a lessened contrast in precipitation between the crest and the lee (Figs. 16b and 17a). This may be due to the reduced wave amplitude near the crest for the smoothed run (Figs. 12b,f), resulting in reduced spillover of precipitation into the lee. Due to a lack of detailed surface observations with sufficient temporal resolution immediately surrounding the Cascade crest, detailed verification of this precipitation distribution is not possible. Yet an examination of observed precipitation amounts from 1400 to 0800 UTC 13–14 December 2001 (Fig. 16 of Garvert et al. 2005a) does indicate a tendency for reduced precipitation totals near the crest as compared to those over the immediate windward and leeward slopes.

To illustrate the importance of the local terrain details during the entire event, Fig. 18a shows a east–west cross section of meridionally averaged QPF from 1400 to 0800 UTC 13–14 December 2001. This extended time period includes prefrontal, frontal, and postfrontal precipitation regimes, and has previously been compared against observations in Garvert et al. (2005a) and Colle et al. (2005). The integrated precipitation amount over the entire cross section showed a net increase of 8% in the control (987 cm) versus the smoothed simulations (911 cm). This increase in the integrated precipitation amounts differs from that computed for the 3-h prefrontal period, during which integrated precipitation amounts were roughly equal for both the smoothed-terrain and control simulations. Therefore, the impact of the increased terrain resolution on precipitation distributions and amounts is apparently sensitive to the thermodynamic and flow properties of the upstream flow. Specifically relatively shallow precipitating features characteristic of postfrontal regimes would be expected to be more sensitive to details of the underlying flow and terrain.

Over the windward slopes, the model predicted a 12% increase in the control’s simulation-integrated precipitation (585 cm) versus the smoothed terrain (518 cm), while over the crest and lee minimal differences in the integrated precipitation amounts were seen (322 versus 320 cm). Increased precipitation amounts over the windward slopes appear to be tied to the presence of the smaller-scale wave perturbations. For example, along a north–south section (leg 2) the QPF within the control simulation is 17% greater than that in the smoothed-terrain simulation (551 versus 471 cm) over the entire 18-h period, indicating the robustness and persistence of the small-scale perturbations identified during leg 2 (Fig. 18b). Meridionally averaged model precipitation for the period from 1400 to 0800 UTC 13–14 December 2001 also exhibited a precipitation minimum along the crest and maximum to the lee, but with much less amplitude than found during the 3-h period from 2200 to 0100 UTC when cross-barrier flow speeds and moisture transport were near their maximum. (Fig. 18a). The reduced difference between the crest and lee QPF values is consistent with model and observational evidence indicating that the mountain wave was strongest from 2300 to 0100 UTC, resulting in a period of increased spillover of precipitation into the lee found during the 3-h time period of P3 aircraft observations.

5. Summary and conclusions

This paper illustrates the value of dual-Doppler radar data over complex terrain in evaluating the kinematic and precipitation fields produced by a mesoscale model. A detailed three-dimensional analysis of the flow and precipitation fields during a heavy precipitation event over the central Oregon Cascades on 13–14 December 2001 has been presented. Dual-Doppler data were collected systematically during this event, and provided a unique opportunity to compare detailed kinematic measures of terrain-induced perturbations over complex terrain with those evident in a high-resolution mesoscale model simulation. This paper expands on previous 13–14 December 2001 IMPROVE studies (Garvert et al. 2005a, b; Colle et al. 2005; Medina et al. 2005) by providing a fully three-dimensional, uninterrupted view of reflectivity and kinematic fields spanning the complex terrain of the Cascades extending from the Willamette Valley to the lee slopes of eastern Oregon.

A high-resolution PSU–NCAR MM5 model was nested down to 1.33 km to simulate the event. The model accurately captured the general characteristics of the flow over the windward slopes, including the strongly sheared U component at low levels, the core of maximum U atop this shear layer, and a zone of strong V-component flow at low levels. The model-derived low-level flow underestimated both the depth and strength of the U-shear layer. However, the model captured the general mesoscale characteristics of the flow, and thus the simulation is suitable for describing the evolution and underlying dynamic mechanisms responsible for many of the observed flow and associated precipitation structures.

The kinematic structure of the 13–14 December 2001 case, which exhibited a tendency for maximum along-barrier flow at low levels and a strong cross-barrier jet atop a sloping shear layer, was also present in other IMPROVE and MAP cases (Medina et al. 2005). In idealized high-resolution 2D simulations, Medina et al. concluded that this strongly sheared cross-barrier flow pattern may result from a combination of influences including 1) synoptic-scale influences (e.g., geostrophic warm advection and associated veering of the upstream wind profile with height), 2) frictional effects over complex topography, and 3) upstream blocking and nonlinear interactions between stably stratified flow and a steep orographic barrier. Relative contributions of these distinct factors in more realistic (observed) situations are difficult to measure, but may certainly vary from case to case and from one mountain range to another.

Dual-Doppler observations showed that as the strong U-component flow surmounted the barrier, it rapidly descended in the immediate lee of the Cascade crest in a manner consistent with the presence of a vertically propagating mountain wave. These observations represent one of the highest-resolution depictions of mountain-wave behavior obtained to date. To assess the impact of the upstream shear profile and boundary layer on the simulated mountain wave, sensitivity tests were completed using three different PBL parameterization schemes: the MRF-PBL, Eta-PBL, and a modified version of the MRF-PBL referred to as the Yonsei-PBL. It was shown that all of these schemes underpredicted the vertical depth of the cross-barrier shear layer and produced flow speeds that were too strong near the surface. These tests clearly indicate that the amplitude and strength of the mountain wave are extremely sensitive to the upstream shear layer and choice of PBL parameterization. Incorrect representation of the upstream shear layer could be an important source of errors in the simulated strength and amplitude of the modeled mountain wave anchored to the Cascade crest. It is therefore imperative that errors in the model depiction of the upstream shear layer and boundary layer parameterizations be corrected in order to further improve quantitative precipitation forecasts over complex terrain.

An average east–west cross section of observed radar reflectivity was also constructed from the composite airborne Doppler analysis to more comprehensively examine precipitation fields over the Oregon Cascades. A strong Föehn wall-like gradient in the reflectivity field was apparent over the lee slopes, with radar reflectivity decreasing rapidly a short distance downstream from the crest. Over the windward slopes, substantial modulation in the reflectivity was clearly evident, with three distinct areas of higher reflectivity: one located over the Willamette Valley (∼100 km upstream of the crest), a second over the windward foothills (centered ∼70 km upstream of the crest), and a third larger area of enhanced reflectivity commencing ∼40 km upstream of the mean crest and extending into the immediate lee of the Cascade crest.

To explore the origins and persistence of these perturbations, model-predicted microphysical fields were compared to the observed radar reflectivity pattern. Experiments were also identified to assess the influence of the coastal mountains and terrain resolution on the reflectivity and precipitation over the upstream areas. Differences between the simulations with and without the coastal mountains suggest that a lee wave produced by the coastal mountains acted to enhance precipitation over the Willamette Valley. A smoothed-terrain simulation illustrated the sensitivity of the windward mountain waves and microphysical responses to the steepness and other details of the underlying terrain.

Collocated with the plume of higher reflectivity values over the Cascade crest, simulated maximum snow mixing ratios (>0.9 g kg−1) extended from 40 km upstream of the crest to the immediate lee. These high snow mixing ratios were advected east of the barrier by the strong cross-barrier flow before being shunted downward in the immediate lee of the Cascades within a zone of strong subsidence associated with a high-amplitude mountain-induced gravity wave. This process resulted in a reduction of modeled precipitation amounts near the crest and an area of markedly enhanced precipitation amounts in the immediate lee. Although this precipitation distribution could not be clearly verified, an examination of precipitation amounts from 1400 to 0800 UTC 13–14 December 2001 (cf. Fig. 16 of Garvert et al. 2005a) indicates some tendency for reduced precipitation totals near the crest versus over the windward slopes and immediate lee.

Significant vertical velocity perturbations were also present over the Cascade foothills at horizontal scales of ∼15–20 km (Fig. 19). Their development has been traced to the interaction of strong low-level southerly flow with multiple ridge–valley terrain corrugations extending west from the more meridional Cascade crest. Model depictions of the precipitation and CLW fields along a P3 flight track in this zone show that complex microphysical interactions were produced by these locally strong vertical velocity perturbations that enhanced precipitation over the crest and immediate lee of several of the windward ridges. The precipitation distribution at the surface was very sensitive to the amount of riming, the resultant fall speed of the particles, and the phase and amplitude of the small-scale waves. The kinetic energy of these waves was concentrated at horizontal scales far greater than the turbulence-induced transient vertical motions shown by Houze and Medina (2005), which were found to develop within a layer of enhanced vertical shear. Those small-scale oscillations, whose inferred horizontal wavelengths were nearer 5 km, represent yet another source of vertical motions that might potentially contribute to orographic enhancement of precipitation.

An analysis of the model-simulated precipitation, focusing on the period of heavy prefrontal precipitation between 2200 and 0100 UTC, showed the small-scale waves increased the modeled precipitation amount over the windward slopes by 4%–14% over those in a corresponding smoothed-terrain model simulation. Yet during this same 3-h period, the smoothed-terrain resolution did not significantly alter the total amount of precipitation integrated over the entire domain, but rather acted mainly to redistribute rainfall relative to individual ridge–valley features. When a longer 18-h period was considered (which included prefrontal, frontal, and postfrontal regimes), the integrated precipitation amount over the windward slopes for the control simulation showed a net increase of 12% over the smoothed-terrain simulation. Examination of microphysical fiends showed this increase was directly tied to an increase in precipitation over the windward slopes caused by the smaller-scale wave perturbations. Additional work will be required to understand more fully the impact of these and other recently identified smaller-scale wave motions on the orographic precipitation process. Specifically, idealized model simulations that can better isolate and quantify the microphysical processes and resulting surface precipitation patterns will likely be required, and airborne Doppler radar observations or other suitable means of comprehensive remote sensing over even more extended periods will be highly desirable.

Acknowledgments

This research was supported by the National Science Foundation Grant ATM 9979494. The authors thank Amy Haase, Socorro Medina, Stacy Brodzik, Dave Ovens, Brian Colle, Dale Durran, and Jon Locatelli for their valuable contributions to this work. We extend additional thanks to the P3 flight crews who provided the critical observations that made this study possible. The comments of two anonymous reviewers enhanced the quality of the manuscript.

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Fig. 1.
Fig. 1.

(a) Surface elevation map of the IMPROVE-2 study area with the flight track of the NOAA P3 (solid red line) overlain. Locations of the sounding site (UW), profiler (IB), vertically pointing NOAA/ETL S-band (MB), and NCAR S-Pol radar (SPL) are also indicated. Plotted flight track encompasses the period 2300 UTC 13 Dec–0100 UTC 14 Dec 2001. (b) Map of the Pacific Northwest with the 36-, 12-, 4-, and 1.33-km MM5 domains shown.

Citation: Journal of the Atmospheric Sciences 64, 3; 10.1175/JAS3876.1

Fig. 2.
Fig. 2.

MM5 4-km simulated wind vectors (black arrows), potential temperature (color shading), pressure (blue dotted contours), and topography (gray contours) from the 4-km domain simulation. Analyses are at heights of 1 and 3 km MSL, respectively at (a), (b) 2000 UTC 13 Dec; (c), (d) 2300 UTC 13 Dec; and (e), (f) 0200 UTC 14 Dec 2001. The blue dashed box represents the region where airborne P3 dual-Doppler measurements were acquired. The black line shows the position of the east–west cross section in Fig. 3. The leading edge of the cold front along the cross section is indicated by CF.

Citation: Journal of the Atmospheric Sciences 64, 3; 10.1175/JAS3876.1

Fig. 3.
Fig. 3.

East–west cross section (position shown in Fig. 2) of MM5 4-km simulated θ (color shading), θe (solid blue lines), and wind barbs (m s−1) for (a) 2000 UTC 13 Dec, (b) 2300 UTC 13 Dec, and (c) 0200 UTC 14 Dec 2001. The solid brown contour indicates the 0°C isotherm. The location of cold front given by gray dashed line with the shaded region in (b) showing a region of enhanced vertical gradients of θ and θe.

Citation: Journal of the Atmospheric Sciences 64, 3; 10.1175/JAS3876.1

Fig. 4.
Fig. 4.

The U- and V-component winds from the NCAR ISS profiler (located at IB in Fig. 1a) at 30-min intervals (dashed color lines) from 2247 to 0047 UTC 13–14 Dec 2001. The solid black lines represent the average of the five individual profiles.

Citation: Journal of the Atmospheric Sciences 64, 3; 10.1175/JAS3876.1

Fig. 5.
Fig. 5.

The U-component speed, dry and moist squared Brunt–Väisälä frequency (N2d and N2m), and inverse Froude number (Nh/U), where h is the mean height of the Cascade crest (h = 1.68 km) from the special UW sounding location (UW in Fig. 1a) at 0000 UTC 14 Dec 2001 and the MM5 1.33-km simulation.

Citation: Journal of the Atmospheric Sciences 64, 3; 10.1175/JAS3876.1

Fig. 6.
Fig. 6.

Composite grid of (a), (c), (e) P3 dual-Doppler wind speed (m s−1, color shading) and wind vectors and (b), (d), (f) corresponding radar reflectivity at elevations of (a), (b) 1.5-, (c), (d) 3-, and (e), (f) 4.5-km MSL. Red line indicates flight track of P3 over the 2-h period from 2300 to 0100 UTC 13–14 Dec 2001. The black box in (e) and (f) outlines region over which data were averaged to create the mean east–west cross sections in Figs. 7 and 11, respectively. Dotted lines depict the north–south mean crest line.

Citation: Journal of the Atmospheric Sciences 64, 3; 10.1175/JAS3876.1

Fig. 7.
Fig. 7.

Average east–west cross section of airborne dual-Doppler winds for the (a) U and (c) V component (color shading, m s−1). Black dots in represent the position of the P3 north–south legs. Dashed lines delineate the boundaries separating individual Doppler analysis grids included in the composite. Dashed black box in (a) shows the region averaged to create the “upstream” profile used to calculate Fig. 8. Averaged east–west cross section of the MM5 1.33-km simulated (b) U and (d) V component with contours of simulated equivalent potential temperature (θe: black contours) and potential temperature (θ: dashed white contours). Model fields were averaged in time from 2300 to 0100 UTC 13–14 Dec 2001 (forecast hours 10–13) using model output at 15-min intervals. White × marks in (b) are the ending locations of the trajectories shown in Fig. 9.

Citation: Journal of the Atmospheric Sciences 64, 3; 10.1175/JAS3876.1

Fig. 8.
Fig. 8.

Dual-Doppler (a) U ′ and (b) V ′ in the east–west cross section corresponding to Figs. 7a,c. Black dots represent the position of the P3 north–south legs. Dashed lines delineate the boundaries separating individual Doppler analysis grids included in the composite.

Citation: Journal of the Atmospheric Sciences 64, 3; 10.1175/JAS3876.1

Fig. 9.
Fig. 9.

(a) Map of backward trajectories terrninating over the windward slopes of the Cascades (at point −122.8°N, 44.5°W) at 2300 UTC 13 Dec 2001 at four different heights (key lower right) as shown in Fig. 7b. (b) Lagrangian traces of pressure (hPa), potential temperature (K), equivalent potential temperature (K), and relative humidity (%) following the parcel tracks shown in (a), plotted against the distance (km) from the endpoint.

Citation: Journal of the Atmospheric Sciences 64, 3; 10.1175/JAS3876.1

Fig. 10.
Fig. 10.

Meridionally averaged east–west cross sections of MM5-generated equivalent potential temperature (θe) overlaid on (a), (c), (e) U-component flow and (b), (d), (f) vertical velocity for three different 1.33-km MM5 simulations using the (a), (b) MRF-PBL; (c), (d) Eta-PBL; and (e), (f) Yonsei-PBL. Model fields were averaged in time from 2300 to 0100 UTC 13–14 Dec 2001 (forecast hours 10–13) using model output at 15-min intervals.

Citation: Journal of the Atmospheric Sciences 64, 3; 10.1175/JAS3876.1

Fig. 11.
Fig. 11.

Average east–west cross section of MM5 1.33-km modeled (a) CLW (qw) and (b) snow (qs) overlaid on airborne dual-Doppler derived radar reflectivity (dBZ; color shading, key at right). The modeled CLW and qs shown were averaged in time from 2300 to 0100 UTC 13–14 Dec 2001 (forecast hours 10–13) using model output at 15-min intervals.

Citation: Journal of the Atmospheric Sciences 64, 3; 10.1175/JAS3876.1

Fig. 12.
Fig. 12.

Average MM5 east–west cross sections of MM5 CLW (g kg−1) overlaid on airborne Doppler reflectivity for three different 1.33-km MM5 simulations: (a) control, (c) no-coastal mountains (Noco), and (e) smoothed terrain (Smooth). Model θe overlaid on model vertical velocity for (b) control, (d) no-coastal mountains, and (f) smoothed terrain. The modeled fields were averaged from 2300 to 0100 UTC 13–14 Dec 2001 (forecast hours 10–13) using model output at 15-min intervals.

Citation: Journal of the Atmospheric Sciences 64, 3; 10.1175/JAS3876.1

Fig. 13.
Fig. 13.

The U, V, and W components (m s−1) at 2.5 km MSL along leg 2 of the P3 flight track (location shown Fig. 1b) from the MM5 1.33-km model simulation, P3 in situ measurements, and as derived from dual-Doppler airborne radar (key at right). The P3 in situ winds were averaged using a 10-s running mean providing a horizontal resolution of about 1 km. The Doppler winds along the flight track were obtained through bilinear interpolation. The fourth panel shows CLW measurements along leg 2 as measured by the King probe aboard the P3 and as predicted by the MM5 1.33-km simulations. The bottom panel is the underlying terrain along the flight track.

Citation: Journal of the Atmospheric Sciences 64, 3; 10.1175/JAS3876.1

Fig. 14.
Fig. 14.

The north–south cross section of airborne dual-Doppler and MM5-predicted (a), (b) U- and (c), (d) V-component flow along leg 2 of the P3 flight track (cf. Fig. 1a). Contours of Doppler-derived vertical velocity are displayed at intervals of 0.2 m s−1 in (a) and (c), while contours of model predicted vertical velocity are displayed in intervals of 0.3 m s−1 in (b) and (d). Negative contour values are dashed.

Citation: Journal of the Atmospheric Sciences 64, 3; 10.1175/JAS3876.1

Fig. 15.
Fig. 15.

The north–south cross sections along leg 2 of (a) airborne dual-Doppler derived vertical velocity (m s−1) overlaid on reflectivity and (b) contours of modeled snow, graupel, and rain mixing ratios in g kg−1. Shaded regions in (b) denote areas where model-predicted CLW mixing ratios exceeded 0.2 g kg−1.

Citation: Journal of the Atmospheric Sciences 64, 3; 10.1175/JAS3876.1

Fig. 16.
Fig. 16.

(a) Modeled accumulated precipitation from 2200 to 0100 UTC 13–14 Dec 2001 (color shading, key at right) and terrain elevation (black contours at intervals of 250 m) for the 1.33-km resolution MM5 control simulation and (b) for the smoothed-terrain version of the 1.33-km resolution MM5 simulation. Black dashed line in (a) and (b) indicates the position of leg 2 of the P3 flight track and cross section in Fig. 17b. White dashed line represents the orientation of the meridionally averaged east–west cross section displayed in Fig. 17a.

Citation: Journal of the Atmospheric Sciences 64, 3; 10.1175/JAS3876.1

Fig. 17.
Fig. 17.

(a) The east–west cross section of meridionally averaged QPF from 2200 to 0100 UTC 13–14 Dec 2001 from the control simulation (solid black line) and smoothed run (dashed gray line) with underlying terrain of the control simulation (black filled area) and smoothed terrain (thick gray dashed line). Orientation of cross section is shown by white dashed line in Fig. 16. (b) Same as in (a) but with QPF and underlying terrain displayed for the south–north-oriented cross section along leg 2 of the P3 track (cf. Fig. 1a). Orientation of the cross section is shown by the black dashed line in Fig. 16. Thick dashed gray line delineates the mean terrain profile for the smoothed-terrain simulation.

Citation: Journal of the Atmospheric Sciences 64, 3; 10.1175/JAS3876.1

Fig. 18.
Fig. 18.

Same as in Fig. 17, but over the extended interval 1400 UTC 13 Dec–0800 UTC 14 Dec 2001.

Citation: Journal of the Atmospheric Sciences 64, 3; 10.1175/JAS3876.1

Fig. 19.
Fig. 19.

Three-dimensional idealized schematic of topography and wind flow over the IMPROVE study area from 2300 to 0100 UTC 13–14 Dec 2001. Blue arrows show strong southerly low θe airflow at low levels along the windward (west facing) slopes of the Cascade range which was subsequently involved in wave generation over multiple small-scale east–west-oriented ridges–valleys within the Cascade foothills. Red arrows show the high θe cross-barrier flow that surmounted the low θe air and exhibited a vertically propagating mountain-wave structure anchored to the mean north–south Cascade crest.

Citation: Journal of the Atmospheric Sciences 64, 3; 10.1175/JAS3876.1

1

The AVN is currently known as the Global Forecast System (GFS) model.

2

For example, at a windward slope location where the terrain height was 0.75 km, a height of 2.00 km MSL would be equal to 1.25 km AGL (location A in Fig. 7a). The U value at A (∼22 m s−1) would then be subtracted from the “upstream” U value at B (∼18 m s−1) resulting in a U ′ value of 4 m s−1 (C = A − B) as shown in Fig. 8a. Due to lack of data at the lowest levels over the Willamette Valley (as a result of the radar ground-clutter elimination process), the average upstream wind profile (at point A) extends no lower than 0.75 km AGL. As a result, regions below 0.75 km AGL in Figs. 8a,b are left blank.

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