Orographically Modified Ice-Phase Precipitation Processes during the Olympic Mountains Experiment (OLYMPEX)

Andrew DeLaFrance aDepartment of Atmospheric Sciences, University of Washington, Seattle, Washington

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Lynn McMurdie aDepartment of Atmospheric Sciences, University of Washington, Seattle, Washington

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Angela Rowe bDepartment of Atmospheric and Oceanic Sciences, University of Wisconsin–Madison, Madison, Wisconsin

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Abstract

Over mountainous terrain, windward enhancement of stratiform precipitation results from a combination of warm-rain and ice-phase processes. In this study, ice-phase precipitation processes are investigated within frontal systems during the Olympic Mountains Experiment (OLYMPEX). An enhanced layer of radar reflectivity (Z H) above the melting level bright band (i.e., a secondary Z H maximum) is observed over both the windward slopes of the Olympic Mountains and the upstream ocean, with a higher frequency of occurrence and higher Z H values over the windward slopes indicating an orographic enhancement of ice-phase precipitation processes. Aircraft-based in situ observations are evaluated for the 1–2 and 3 December 2015 orographically enhanced precipitation events. Above the secondary Z H maximum, the hydrometeors are primarily horizontally oriented dendritic and branched crystals. Within the secondary Z H maximum, there are high concentrations of large (>~2-mm diameter) dendrites, plates, and aggregates thereof, with a significant degree of riming. In both events, aggregation and riming appear to be enhanced within a turbulent layer near sheared flow at the top of a low-level jet impinging on the terrain and forced to rise above the melting level. Based on windward ground sites at low, mid-, and high elevations, secondary Z H maxima periods during all of OLYMPEX are associated with increased rain rates and larger mass-weighted mean drop diameters compared to periods without a secondary Z H maximum. This result suggests that precipitation originating from secondary Z H maxima layers may contribute to enhanced windward precipitation accumulations through the formation of large, dense particles that accelerate fallout.

Significance Statement

Precipitation processes are modified within winter storms passing over the Olympic Mountains, often resulting in increased rain and snow on the windward slopes. This study evaluates precipitation characteristics and inferred growth processes related to radar reflectivity maxima within ice layers of clouds, providing insights into ice-phase contributions to windward precipitation. Ground- and aircraft-based measurements indicate that rapid ice growth may occur when branched and platelike crystals are aggregated within turbulent regions that are prevalent over the windward slopes. Large aggregate particles gain mass by collection of supercooled liquid water, which may increase precipitation fallout. While ground measurements suggest that ice-phase growth contributes to windward accumulations, some uncertainty about the processes that occur between the ice layer and the ground remain.

© 2021 American Meteorological Society. For information regarding reuse of this content and general copyright information, consult the AMS Copyright Policy (www.ametsoc.org/PUBSReuseLicenses).

Corresponding author: Andrew DeLaFrance, adelaf@uw.edu

This article is included in the The Olympic Mountains Experiment (OLYMPEX) Special Collection.

Abstract

Over mountainous terrain, windward enhancement of stratiform precipitation results from a combination of warm-rain and ice-phase processes. In this study, ice-phase precipitation processes are investigated within frontal systems during the Olympic Mountains Experiment (OLYMPEX). An enhanced layer of radar reflectivity (Z H) above the melting level bright band (i.e., a secondary Z H maximum) is observed over both the windward slopes of the Olympic Mountains and the upstream ocean, with a higher frequency of occurrence and higher Z H values over the windward slopes indicating an orographic enhancement of ice-phase precipitation processes. Aircraft-based in situ observations are evaluated for the 1–2 and 3 December 2015 orographically enhanced precipitation events. Above the secondary Z H maximum, the hydrometeors are primarily horizontally oriented dendritic and branched crystals. Within the secondary Z H maximum, there are high concentrations of large (>~2-mm diameter) dendrites, plates, and aggregates thereof, with a significant degree of riming. In both events, aggregation and riming appear to be enhanced within a turbulent layer near sheared flow at the top of a low-level jet impinging on the terrain and forced to rise above the melting level. Based on windward ground sites at low, mid-, and high elevations, secondary Z H maxima periods during all of OLYMPEX are associated with increased rain rates and larger mass-weighted mean drop diameters compared to periods without a secondary Z H maximum. This result suggests that precipitation originating from secondary Z H maxima layers may contribute to enhanced windward precipitation accumulations through the formation of large, dense particles that accelerate fallout.

Significance Statement

Precipitation processes are modified within winter storms passing over the Olympic Mountains, often resulting in increased rain and snow on the windward slopes. This study evaluates precipitation characteristics and inferred growth processes related to radar reflectivity maxima within ice layers of clouds, providing insights into ice-phase contributions to windward precipitation. Ground- and aircraft-based measurements indicate that rapid ice growth may occur when branched and platelike crystals are aggregated within turbulent regions that are prevalent over the windward slopes. Large aggregate particles gain mass by collection of supercooled liquid water, which may increase precipitation fallout. While ground measurements suggest that ice-phase growth contributes to windward accumulations, some uncertainty about the processes that occur between the ice layer and the ground remain.

© 2021 American Meteorological Society. For information regarding reuse of this content and general copyright information, consult the AMS Copyright Policy (www.ametsoc.org/PUBSReuseLicenses).

Corresponding author: Andrew DeLaFrance, adelaf@uw.edu

This article is included in the The Olympic Mountains Experiment (OLYMPEX) Special Collection.

1. Introduction

Wintertime rain and snow accumulation in mountainous regions provides a necessary freshwater resource to neighboring lowland areas (Viviroli et al. 2007). Often, mountains receive enhanced precipitation accumulations on the windward slopes and diminished precipitation on the lee slopes (Smith 1979; Roe 2005; Houze 2012). When wintertime midlatitude cyclones encounter mountains, the windward enhancement of stratiform precipitation results from a complex combination of warm-rain processes below the melting level and ice-phase processes above (White et al. 2003; Kingsmill et al. 2006; Massmann et al. 2017; Zagrodnik et al. 2018, 2019). Ice particle growth occurs by diffusion (deposition), collection of other ice particles (aggregation), or accretion of supercooled liquid water drops (riming). Ice particle properties vary due to growth contributions from each process which has implications for the distribution of precipitation along windward and lee slopes of a mountain. Riming- and aggregation-dominated ice growth permits rapid fallout of precipitation on the windward slopes whereas precipitation is advected downstream in a slower depositional growth regime (Hobbs et al. 1973; Colle and Zeng 2004; Colle et al. 2005). A need for additional observational studies of these processes has been demonstrated by increasingly sophisticated approaches to model the microphysics of ice particle growth (Brdar and Seifert 2018; Grabowski et al. 2019).

Hydrometeor properties and growth processes are often inferred from dual-polarization radar measurements. Differential reflectivity (ZDR) is the logarithmic difference between radar reflectivity factors at horizontal and vertical polarizations (ZH and ZV) with near-zero dB values for spherical or randomly oriented hydrometeors and increasing ZDR for decreasing aspect ratio and preferentially horizontally oriented hydrometeors. Enhanced layers of positive ZDR above the melting level near −15°C temperatures have consistently been linked to depositional growth of dendritic or platelike ice crystals (e.g., Kennedy and Rutledge 2011; Andrić et al. 2013; Bechini et al. 2013; Schrom et al. 2015). Below these enhanced ZDR layers, concurrently decreasing ZDR and increasing ZH is an indication of aggregation as particles increase in size and diversity of shape or orientation while falling toward the ground. Radar-based inferences of enhanced dendritic growth with active aggregation have been associated with heavy snowfall rates at the surface (Kennedy and Rutledge 2011; Kumjian and Lombardo 2017). However, polarimetric signatures aloft during winter storms may not be fully explained by dendritic and aggregational growth processes alone (Andrić et al. 2013), and the effects of secondary ice production and riming need to be considered. Indeed, Schrom et al. (2015) describe uncertainty in polarimetric radar signatures associated with the degree of riming and called for additional in situ measurements of particle types and size distributions coincident with radar observations to provide validation of radar-inferred microphysics.

Over complex terrain, several radar-based studies have considered whether orographically modified flow enhances, or at least alters, microphysical growth processes and the accompanying radar signatures aloft (e.g., Houze and Medina 2005; Kingsmill et al. 2006; Medina et al. 2007; McMurdie et al. 2018). Using vertically pointing S-band radar data in the Northern California coastal mountains, Kingsmill et al. (2006) documented a pronounced increase in the rate of change in ZH with height at about 2.5 km above the melting level that was termed a “shoulder.” The “shoulder” occurred near −15°C temperatures and was considered a radar signature of accelerated hydrometeor growth rates, likely by aggregation. Below the “shoulder,” by definition, the rate of increase in ZH toward the ground decreases and in some cases ZH begins to decrease resulting in a local maximum in ZH. They hypothesized that slower diffusional growth rates at warmer temperatures or a decrease in aggregation efficiency could explain this decrease in ZH below the maximum. Kingsmill et al. (2006) proposed that because the “shoulder” was identified both at a mountainous site and at a downstream location, described as free from orographic influences, it was unlikely an orographically forced response, but rather inherent to the large-scale environment.

In contrast, Medina et al. (2007) argued that a secondary ZH maximum above the melting level is likely an orographically forced response or is at least enhanced by flow encountering topography. From a scanning S-band radar, they identified a secondary ZH maximum at a height of approximately 3.5–5.5 km over the windward slopes of the Oregon Cascade Range (~1–2.5 km above the melting level) but not upstream, suggesting an orographic response in ZH aloft. This finding is consistent with McMurdie et al. (2018) where a greater frequency of larger ZH values occurred at all heights between 2 and 8 km over the windward slopes of the Olympic Mountains in Washington State when compared to upstream. The most significant enhancement of ZH over the windward slopes occurred at a height of 4–6 km, comparable to the secondary ZH maximum identified by Medina et al. (2007) and thus bolstering the suggestion of an orographic response aloft.

The physical processes associated with the secondary ZH maximum in orographically modified flow are not well understood. In orographic studies that identified a secondary ZH maximum (Houze and Medina 2005; Kingsmill et al. 2006; Medina et al. 2007; Zagrodnik et al. 2018; McMurdie et al. 2018), polarimetric radar measurements were either not considered or not available. Additionally, while polarimetric radar observations provide significant insight to ice growth processes, in situ measurements are ultimately necessary to validate or refute any radar-based inferences (e.g., Andrić et al. 2013; Schrom et al. 2015; Griffin et al. 2018).

The 2015/16 Olympic Mountains Experiment (OLYMPEX) obtained radar-, aircraft-, and surface-based observations of midlatitude cyclones encountering the Olympic Mountains of Washington State, providing a rich dataset for investigating the orographic effects on ice-phase precipitation processes (Houze et al. 2017; Petersen et al. 2018). An S-band dual-polarization scanning radar collected observations upstream and over the windward slopes of the Olympic Mountains during two intensive observing periods (IOPs): 12 November–19 December 2015 and 4–15 January 2016. In situ observations were collected within the radar domain by the University of North Dakota (UND) Citation aircraft during the first IOP. The purpose of this study is to use dual-polarization radar and in situ measurements to evaluate the microphysical properties of hydrometeors associated with secondary ZH maxima and connect these observations to windward surface precipitation measurements to better understand the effects of orographically modified ice microphysics on windward precipitation. In this study, OLYMPEX data are used to address the following questions:

  1. Is a secondary ZH maximum a signature of orographically modified ice-phase precipitation processes or is it solely a microphysical signature inherent to midlatitude frontal systems?

  2. What hydrometeor properties and growth processes produce the secondary ZH maximum?

  3. How is the windward distribution and intensity of precipitation at the surface impacted by hydrometeor growth processes aloft?

2. Data

a. OLYMPEX radar data

NASA’s S-band dual-polarization Doppler radar (NPOL) scanned nearly uninterrupted during OLYMPEX at a coastal site with an elevation of 139 m on the Olympic Peninsula (Fig. 1). NPOL’s scanning strategy included a series of range–height indicator (RHI) scans extending to 45° in elevation at 3° azimuthal spacing from 210° to 326° over the ocean and at 2° azimuthal spacing from 30° to 60°, broadly covering the Quinault River valley. The NPOL RHIs extended in range to 150 km with a 125-m gate spacing and 0.94° beamwidth. RHIs over a given location repeated every 20 min for a total of 2113 complete radar scan cycles. This study uses version 2 of the NPOL data (Wolff et al. 2017) including calibrated ZH and bias-corrected ZDR (Wolff et al. 2015; Pippitt et al. 2015). For computational efficiency and simplification of the methods used to evaluate NPOL measurements, these data were interpolated to Cartesian coordinates using the Radx2Grid algorithm included in the Lidar Radar Open Software Environment (Bell et al. 2020). Data were gridded at 500-m horizontal and 250-m vertical grid spacing. NPOL data for this analysis are limited to a 60-km maximum range to reduce beam broadening impacts on polarimetric measurements (e.g., Sánchez-Diezma et al. 2000; Brandes and Ikeda 2004; Ryzhkov 2007; Giangrande et al. 2008) and to a minimum range of 10 km to include the full echo depth; echo tops during OLYMPEX rarely extended much higher than 8 km. In this study, we make comparisons between secondary ZH maximum properties over the upstream ocean and the terrain by using detection methods defined by areal fraction thresholds. For these comparisons to be weighted fairly, the larger ocean-based RHI sector was divided into four 30° sectors of equal size to the terrain-focused sector: southwest (210°–240°), west-southwest (240°–270°), west-northwest (270°–300°), and approximately northwest (296°–326°). Because NPOL RHI scans over the northwest sector terminated at 326°, this sector overlaps with the west-northwest sector by four degrees. The southwest sector directly opposes the northeast (30°–60°), terrain-focused scan. The radar sector construction along cardinal orientations permit evaluation based on orientation upstream of the terrain. This approach was motivated by results from McMurdie et al. (2018) which described a pronounced increase in reflectivity values aloft over the terrain during periods of southwesterly low-level flow.

Fig. 1.
Fig. 1.

Map of the Olympic Peninsula showing the position of the NASA S-band dual-polarization Doppler radar (NPOL; 139 m) and three ground sites (Fishery, 52 m; Prairie Creek, 543 m; Wynoochee, 1018 m) on the windward slopes. Terrain elevation is shown in the gray shading. Colored regions indicate radar sectors of 10–60-km radial range: southwest: gray; south-southwest: blue; north-northwest: red; northwest: violet; northeast: green.

Citation: Journal of the Atmospheric Sciences 78, 11; 10.1175/JAS-D-21-0091.1

b. OLYMPEX ground-based data

During OLYMPEX, the windward slopes of the Olympic Mountains were instrumented with Particle Size and Velocity 2 (Parsivel2; Petersen et al. 2017a) disdrometers spanning from near the coast to an elevation of about 1600 m. Drop size distributions constructed from Parsivel2 data in 2-min segments were used to calculate rain rate and mass-weighted mean drop diameter following Zagrodnik et al. (2018). Although this method applies quality control measures to remove contamination of frozen precipitation, we have also removed any period in which the radar bright band was within 500 m of the ground site elevation. Three Parsivel2 ground sites are used in this study to exhibit the change in orographic enhancement of precipitation with elevation (Fig. 1). The Fishery site is in the lower Quinault River valley at an elevation of 51.8 m, at a distance of 19 km (~61° azimuth) from the NPOL site. Prairie Creek is a midelevation site at 542.5 m and 33.5-km distance (~39° azimuth) from NPOL. The Wynoochee site is in the high terrain at 1018 m and 53.6-km distance (~62° azimuth) from NPOL.

Data used to describe atmospheric conditions were obtained from rawinsondes launched at the NPOL site typically every 6 h but as frequently as three hours during storm events (Rutledge et al. 2018). Three-hourly North American Regional Reanalysis (NARR; Mesinger et al. 2006) data at the 32-km horizontal resolution grid point containing the NPOL location are used to fill in the temporal gaps of the soundings. Following McMurdie et al. (2018) and Zagrodnik et al. (2018), integrated water vapor transport (IVT) was used to describe the synoptic-scale moisture flux, 925-hPa wind as an indicator of the low-level flow pattern, and the moist static stability parameter Nm2 (Durran and Klemp 1982) as a measurement of atmospheric stability in three categories: moist unstable (Nm2<0.25×104s2), moist neutral (0.25×104<Nm2<0.25×104s2), and moist stable (Nm2>0.25×104s2). McMurdie et al. (2018) showed that the 925-hPa wind, Nm2, and IVT values of the NARR dataset suitably characterize the synoptic environment during OLYMPEX.

c. OLYMPEX aircraft data

During the first IOP, the UND Citation flew primarily within the NPOL domain over the windward slopes and was equipped with microphysics probes to collect in situ measurements of cloud particles (Poellot et al. 2017). Measurements of two-dimensional projected ice particle images from two optical array probes (OAPs) were used to construct particle size distributions (PSDs) within the ice layer. We use the vertical channel of the 2D Stereo (2D-S) probe and the vertically oriented High Volume Precipitation Spectrometer version 3 (HVPS-3) probe. The University of Illinois/Oklahoma Optical Array Probe Processing Software (UIOOPS; McFarquhar et al. 2018) was used to process the 2D-S and HVPS-3 probe measurements. Following Chase et al. (2018) the 2D-S probe measurements were used for particles 0.225–1 mm in diameter, and the HVPS-3 probe measurements for particles 1–32.5 mm in diameter. The lower threshold of 0.225 mm in the 2D-S measurements minimizes inclusion of supercooled liquid water drops (Chase et al. 2018). UIOOPS defines particle diameter, D, by the minimum enclosing circle. The two-dimensional projected particle size observed from a fixed-orientation probe is an estimation of the particle’s actual maximum dimension (e.g., Petty and Huang 2011). Neglecting any disturbance of the natural particle orientation by the aircraft, sampling of particles that fall with their long axis in a preferentially horizontal orientation with respect to vertically oriented probes means that D approximates the maximum particle dimension.

The spectral ice water content (IWC) is given as
IWC(D)=N(D)m(D),
where N(D) is the number concentration and m(D) is the particle mass at a given D. Reflectivity estimates derived from spectral IWC at size bins, j, following Hogan et al. (2006) are given as
ZPSD=|Ki|20.93(6πρi)2jm(Dj)2N(Dj)dDj.

This method relies on the assumption of spherical particles and Rayleigh scattering properties. Calculations of spectral IWC require that a mass–dimension relationship be defined. Mass–dimension relationships from Brown and Francis (1995), Baker and Lawson (2006), Heymsfield et al. (2004), and a habit-specific relationship of the UIOOPS algorithm were evaluated by comparing ZH measured from NPOL to ZPSD calculated from in situ observations that were collected inside a grid cell within 60 s of the radar sampling time. Although a mass–dimension relationship must be established, we do not compute a total IWC for evaluation. Rather, we prioritize a mass–dimension relationship that provides the best fit for the measurement of primary interest, radar reflectivity, which is a higher-order moment of the PSD than IWC. For the two case studies in section 4b, the power-law relationship, m = 0.0061D2.05, of Heymsfield et al. (2004) provided the best agreement with the NPOL radar having root-mean-square errors of 5.96 and 8.50 dB and absolute errors of −0.27 and −1.40 dB during the 1–2 and 3 December events, respectively.

Objective ice habit classification techniques in UIOOPS were applied to OAP imagery to better understand the microphysical contributions from unique particle habits. Habit classification includes nine ice habit classes: aggregate, dendrite, hexagonal, graupel, spherical, linear, oriented, irregular, and tiny following the algorithm of Holroyd (1987), which was optimized for the 2D-C OAP with a resolution of 25 μm. Our habit-specific analysis focuses largely on the HVPS-3 OAP imagery. Although the thresholds of the algorithm were not explicitly optimized to work with the coarse resolution of the HVPS-3 OAP (150 μm), manual evaluation of several subsets of habit-classified particles indicated that the UIOOPS algorithm generally classified particles from HVPS-3 images as expected. Nonetheless, we have limited our analysis of habit-classified ice particles using HVPS-3 imagery to a qualitative assessment only.

Two of the Citation’s instruments are used to indicate supercooled liquid water. The King hot wire probe uses the power required to evaporate the liquid drops from a constant-temperature hot wire to derive liquid water content (LWC) estimates (King et al. 1978). A minimum threshold of 0.05 g m−3 is used to ensure measurements are indeed liquid water content and not an artifact of sensor noise. The Rosemount Icing Detector (RICE) was used to indicate the presence of supercooled liquid water. Although it is possible for it to be used to derive LWC estimates, the RICE instrument is most effective in discriminating mixed phase from glaciated clouds based on fixed thresholds (Cober et al. 2001) and will be used to support conclusions based on measurements from the King probe. All in situ measurements from the Citation probes used in this study were collected at 1-Hz sampling frequency.

3. Methods

a. Brightband detection

An automated brightband detection algorithm was developed for the NPOL radar data. The bright band typically contains the 0°C level, thus providing segregation of stratiform and convective precipitation and a vertical partition of ice microphysics aloft from warm-rain processes below. At the melting level, reflectivity increases and copolar correlation coefficient, ρHV, the correlation of the horizontal and vertical radar pulses, decreases from a value of one due to increased diversity of hydrometeor shape, orientation, and phase (e.g., melting). Based on these principles, the brightband detection algorithm uses a grid-by-grid search procedure broadly adapted from methods of Gourley and Calvert (2003) and Giangrande et al. (2008) and modified to suit NPOL observations during OLYMPEX. The search procedure is performed independently for each radar sector shown in Fig. 1 as follows:

  1. At each horizontal grid point the height of the column minimum in ρHV is identified, if one exists between values of 0.91 and 0.97. From all grid points in a radar sector, the most frequently occurring (mode) vertical level of ρHV minima is identified.

  2. Within 500 m below to 1000 m above the ρHV minimum mode level, individual gridpoint bright bands are identified as the height of the column-maximum ZH if it exceeds 25 dBZ (Fig. 2a).

  3. If the number of individual grid points that have a bright band, NBB, is more than 35% of the total grid points in the radar sector, individual gridpoint brightband heights are averaged to produce a mean brightband height for that radar sector. For radar scan cycles having NBB fewer than 35% of the sector’s total grid points, no bright band is identified.

Fig. 2.
Fig. 2.

Vertical profiles of ZH illustrating (a) a schematic diagram of the bright band (BB) and secondary ZH maximum (SM) features in a vertical column at a single grid point and (b) radar sector composite profiles from all NPOL scan cycles having a bright band. In the schematic in (a), a light gray vertical line highlights the 25-dBZ threshold used to define a bright band and a dark gray vertical line highlights the 17-dBZ threshold used for detection of a secondary ZH maximum layer. The lower extent of the secondary ZH maximum layer (SMlwr) is defined by the initial level of increase in ZH with height above the bright band. The upper extent of the secondary ZH maximum layer (SMupr) is the uppermost level having a ZH value greater than the level immediately below SMlwr. Composite profiles in (b) were adjusted to be brightband-relative prior to averaging.

Citation: Journal of the Atmospheric Sciences 78, 11; 10.1175/JAS-D-21-0091.1

The ρHV constraint is used to minimize the influence of high ZH values within embedded convective elements or from ground clutter, ensuring that our brightband height is mostly representative of stratiform melting of ice hydrometeors. An asymmetric vertical search window about the ρHV minimum is used because the minimum in ρHV commonly occurs below the ZH maximum at the bright band (Brandes and Ikeda 2004). There is an exemption that allows the use of the second most commonly occurring ρHV minima mode level when it occurs below 4 km, is at a higher elevation than the first, and has a mean layer ZH exceeding 15 dBZ. This exemption was useful in a small number of radar scan cycles where the melting level is either very deep or close to the surface resulting in low ρHV values from ground clutter. We have removed time periods with a brightband height greater than 500 m above or below a rolling mean of 10 consecutive brightband heights to minimize the inclusion of embedded convective elements. Depending on the radar sector, this resulted in one to nine scan cycles (~20-min periods) removed.

Brightband detection is based on three primary input thresholds: ρHV minimum, ZH maximum, and NBB in a prescribed radar sector. The allowable range of ρHV values is similar to Giangrande et al. (2008; 0.90–0.97) but we found that using a lower threshold of 0.91 instead of 0.90 slightly reduced detection of predominantly convective periods. Similar to Gourley and Calvert (2003), there was little sensitivity to the minimum brightband ZH exceedance value. Most radar scan cycles were also relatively insensitive to NBB as stratiform echo was often widespread, encompassing a significant portion of the radar sector. The periods of largest sensitivity to NBB were often associated with a transition phase such as a frontal passage. Through visual confirmation with individual RHI scans, a threshold of NBB exceeding 35% of the radar sector’s grid points was determined to effectively discriminate between a spatially homogeneous bright band and convective or transient features.

Radar-derived brightband heights were evaluated against the 0°C height from soundings launched at the NPOL site and from the NARR dataset within one hour of the radar scan cycle. During the first IOP, radar-derived brightband heights averaged 138 m below the sounding-derived 0°C isotherm with a root-mean-squared error (RMSE) of 284 and 198 m below the NARR 0°C isotherm with a RMSE of 310 m. This performance is similar to that of the algorithm used by White et al. (2002) which produced radar-derived brightband heights in coastal California, on average, 192 m below the sounding-derived 0°C isotherm.

b. Secondary ZH maximum detection

We consider a secondary ZH maximum to be a layer above the bright band where at the lower extent, ZH increases with height and at the upper extent, ZH returns to decreasing with height (Fig. 2a). Thus, secondary ZH maximum detection uses only ZH as an input, intentionally removing the possibility of prescribing any polarimetric or Doppler radar signature. The algorithm searches for a secondary ZH maximum in any radar scan cycle having a bright band as well as in the scan cycle immediately (~20 min) before and after. The search procedure is performed as follows:

  1. At each horizontal grid point within a radar sector, beginning at 1 km above the bright band and searching upward, the algorithm searches for an increase in ZH with height from the level immediately below. The increase must occur over a minimum of two 250-m vertical levels. The initial level of increase in ZH with height establishes the lower extent of the secondary ZH maximum, (SMlwr). The upper extent of the secondary ZH maximum (SMupr) is the highest vertical level with a ZH value exceeding that at the level immediately below SMlwr. Gridpoint secondary ZH maxima must have at least one vertical level that exceeds 17 dBZ.

  2. Within the sector, the most frequently occurring level of the SMlwr layer is identified. This level is then used to identify a horizontal layer in which a secondary ZH maximum commonly occurs over the sector by disregarding grid points with an SMlwr layer beyond 750 m above or below this level.

  3. If the remaining number of grid points that have a secondary ZH maximum, NSM, exceeds 20% of the radar sector total, a secondary ZH maximum is identified for that radar scan cycle and heights of individual gridpoint SMlwr and SMupr layers are independently averaged to produce a secondary ZH maximum layer.

  4. In the event that an additional ZH maximum is identified above the secondary ZH maxima by the same criteria as the first, a decision is made about where the layer of interest exists. If the two layers are separated by less than 500 m, both layers are used and are considered as one larger layer. If, however, there are two distinct layers, the layer at a height nearest to that of the secondary ZH maximum in the previous radar scan cycle is used.

The brightband existence precondition in the secondary ZH maximum search was appropriate for the OLYMPEX dataset for which 0°C heights were always above ground level at the radar. However, without first being modified, this approach would not be suitable in places such as the Intermountain West where freezing temperatures at the ground are expected. Nonetheless, spatial variability in the height of the bright band is not uncommon, especially in complex terrain. Initializing the upward search at 1 km above the average brightband height restricts detection to the cloud’s ice layer, above any brightband variability.

Similar to the use of a minimum ZH threshold for brightband detection, a gridpoint minimum ZH threshold for a secondary ZH maximum was used to avoid detecting slight increases that can occur near echo tops. This threshold was determined by compositing average ZH vertical profiles from all NPOL scan cycles after first being vertically adjusted to a brightband relative sense by placing the zero-height reference at the bright band. Beginning at echo top and moving down in the composite vertical ZH profiles above the bright band from each radar sector, the rate of increase of ZH with decreasing height aloft has a relative maximum approximately at or below 17 dBZ (Fig. 2b). In the averaged profiles, values greater than 17 dBZ contribute to the increased rate of change in ZH aloft. The threshold of NSM exceeding 20% of the total grid points in a radar sector to define a secondary ZH maximum was arrived at in a similar manner as in the brightband detection. Visual comparison of individual RHI scans provided the basis for determining this threshold and, therefore, it is specific to the OLYMPEX dataset and the S-band NPOL radar. As discussed in two case studies in section 4b, secondary ZH maxima tend to be less widespread than the bright band, thus a threshold below 35% is necessary, but if a value much less than 20% was used, undesirable isolated or convective features were also identified.

4. Results

The OLYMPEX field campaign successfully measured multiple midlatitude cyclones that produced widespread and orographically enhanced precipitation over the Olympic Mountains. Secondary ZH maxima were frequently identified in the stratiform radar echo of those storms. Figures 3a and 3b are an example of the reflectivity field from NPOL RHIs during the 3 December 2015 stratiform rain event in a cross-sectional view from southwest to northeast. At this time, the mean brightband height was about 2 km above NPOL. Above the bright band, a secondary ZH maximum layer occurred over the terrain between about 4 and 4.5 km and is most pronounced beyond about 35 km from NPOL (dashed black ellipse in Fig. 3b). Over the upstream ocean, a sufficiently widespread secondary ZH maximum meeting our detection criteria was not identified at this time. Using all NPOL radar scans, we assess how the secondary ZH maximum manifests over the terrain relative to upstream.

Fig. 3.
Fig. 3.

RHI scans from NPOL for (a),(b) reflectivity, (c),(d) differential reflectivity, (e),(f) radial velocity, and (g),(h) spectrum width. The RHI scans occurred at (a),(c),(e),(g) 1441 UTC 3 Dec 2015 at an azimuthal angle of 219° over the upstream ocean and (b),(d),(f),(h) 1453 UTC 3 Dec 2015 at an azimuthal angle of 38° over the windward terrain of the Olympic Mountains. The black dashed ellipse in (b), (d), (f), and (h) represents the approximate location of the secondary ZH maximum layer. Gray shaded regions depict the terrain of the Olympic Mountains.

Citation: Journal of the Atmospheric Sciences 78, 11; 10.1175/JAS-D-21-0091.1

a. Statistical characteristics of the secondary ZH maximum

Of all NPOL scan cycles, a bright band was identified in 20%–23% of scan cycles over each of the four upstream oceanic radar sectors and in 41% of terrain-focused radar sector scan cycles. The median brightband height over all four ocean sectors was approximately 1.7 km above NPOL (Fig. 4a). Over the terrain, the median height decreased to 1.4 km, a common property of stratiform precipitation on windward slopes (e.g., Marwitz 1983, 1987; Medina et al. 2005; Minder et al. 2011). Brightband heights over the terrain were statistically significantly different from each of the four oceanic sectors based on a two-tailed t test at a 95% significance level.

Fig. 4.
Fig. 4.

NPOL radar sector distributions of (a) brightband height, (b) secondary ZH maximum lower- and upper-extent heights above the bright band, (c) secondary ZH maximum lower- and upper-extent temperatures, and (d) column-averaged ZH values from grid points with a secondary ZH maximum. The box represents the 25th and 75th percentiles and the whiskers are the 5th and 95th percentiles of the distributions. The dark horizontal lines are the median values. Hatching in the color fill indicates that the ocean-focused distribution was not statistically significantly different from the terrain-focused distribution based on a two-tailed t test at a 95% significance level.

Citation: Journal of the Atmospheric Sciences 78, 11; 10.1175/JAS-D-21-0091.1

During periods with a bright band, a secondary ZH maximum was identified in 10%–11% of the radar scan cycles over each of the four oceanic radar sectors and in 25% of the terrain-focused radar sector scan cycles. The higher frequency of occurrence over the terrain was most pronounced at larger radial distances. About 25% of the terrain-focused radar scan cycles had a secondary ZH maximum between 40 and 60 km compared to about 20% in the coastal lowland region between 10 and 30 km. Analysis of NARR data within one hour of radar scans (not shown) indicated that secondary ZH maxima were observed in diverse synoptic conditions. However, they were most frequently observed in southeasterly to southwesterly low-level flow having moist neutral stability and IVT greater than about 250 kg m−1 s−1, conditions common to the warm sector of midlatitude cyclones and to atmospheric river events. The structure of the secondary ZH maximum varies by storm but is approximately a 750-m-deep layer between 1.5 and 3 km above the bright band with temperatures between about −6° and −12°C (Figs. 4b,c), and occasionally colder over the terrain due to apparent lifting of the secondary ZH maximum layer as flow encounters the terrain. Column-averaged ZH within individual gridpoint secondary ZH maximum layers were used as a measure of the layer’s intensity. Relative to the upstream ocean, higher average ZH values over the terrain (Fig. 4d) suggest an orographic enhancement of the secondary ZH maximum layer. McMurdie et al. (2018) presented similar findings of a pronounced increase in frequency of large ZH values aloft over terrain when compared to over the ocean, particularly during warm sector conditions. However, by explicitly identifying the secondary ZH maxima feature, we have demonstrated that it can form upstream. This finding suggests that the physical processes that result in a secondary ZH maximum are not necessarily unique to orographically modified flow but are enhanced by flow over the windward slopes. The following section uses dual-polarization radar and in situ measurements to identify the microphysical processes that produce the relatively high ZH in the secondary ZH maximum layer over the windward slopes.

b. Microphysics case studies

The Citation aircraft was deployed within the range of NPOL a total of 20 times during OLYMPEX. The secondary ZH maximum layer was directly sampled by the Citation on six flights, four of which provided in situ observations within a deep portion of the cloud’s ice layer. Analysis is presented below for two diverse storms that both featured a persistent secondary ZH maximum in a similar location over the windward slopes for the duration of the flight. The 3 December flight occurred during the warm sector portion of the midlatitude cyclone having synoptic conditions most similar to those commonly observed during secondary ZH maxima events. The flight during the preceding cyclone on 1–2 December occurred during the prefrontal portion; conditions that were less frequently associated with secondary ZH maxima events. The synoptically different conditions between the warm-sector and prefrontal periods of these two events make for a unique comparison of in situ measurements and inferred dynamic processes associated with secondary ZH maxima development.

1) Case study: 3 December 2015

The 3 December storm featured significant orographic enhancement resulting in 100–200 mm of rain at high elevations on the windward slopes of the Olympic Mountains (Petersen et al. 2017b). The NPOL sounding at 1516 UTC (Fig. 5) shows southerly flow at 925 hPa, south-southwesterly at 700 hPa, near-neutral moist static stability in the 950–850-hPa layer, and IVT of 784 kg m−1 s−1 (Zagrodnik et al. 2019). Over the terrain, a secondary ZH maximum was continuously identified from 1252 to 1652 UTC. During this time, the Citation flew in a series of vertically stacked legs of constant elevation oriented approximately northwest to southeast over the windward slopes (Fig. 6). In situ observations were organized in three distinct vertical layers as illustrated schematically in Fig. 7a. In radar scan cycles coinciding with in situ observations, the SMlwr and SMupr levels respectively ranged from 1.62 to 2 km and 2.27 to 2.62 km above the bright band. To distinguish between layers, a 250-m separation is used between SMlwr and the layer below and between SMupr and the layer above. The flight resulted in 445, 541, and 528 1-Hz in situ observations from above, within, and below the secondary ZH maximum, respectively, largely within a radial range of 40–60 km from NPOL. Temperatures for the top and bottom of each layer obtained from the 1516 UTC sounding indicate that the 1-km layer below terminates just above the bright band at −2.0°C. The layer above (orange shading, Fig. 7a) nearly resembles a temperature range that is commonly associated with dendritic ice growth (Bailey and Hallett 2009).

Fig. 5.
Fig. 5.

Skew T–logp diagrams of temperature (red) and dewpoint temperature (green) from atmospheric soundings launched at the NPOL site at (left) 1516 UTC 3 Dec and (right) 2317 UTC 1 Dec 2015. Figures obtained from http://olympex.atmos.washington.edu/index.html?x=OLYMPEX_Obs.

Citation: Journal of the Atmospheric Sciences 78, 11; 10.1175/JAS-D-21-0091.1

Fig. 6.
Fig. 6.

Flight track between approximately 1400 and 1700 UTC 3 Dec 2015 with altitude in km indicated by color shading. Grayscale shading is the reflectivity field of the lowest-level (0.5°) plan position indicator (PPI) scan between 1447 and 1452 UTC 3 Dec 2015 from the coastal KLGX WSR-88D at Langley Hill, Washington. Blue diamonds indicate the three ground sites used in this study. Green contours depict the terrain elevation at 250-m intervals beginning at 250 m.

Citation: Journal of the Atmospheric Sciences 78, 11; 10.1175/JAS-D-21-0091.1

Fig. 7.
Fig. 7.

In situ measurements on 3 Dec 2015 within the secondary ZH maximum layer (violet), within a 1-km-deep layer above (orange), and within a 1-km-deep layer below (gray) separated by 250 m. (a) Schematic depicting each layer with temperature ranges obtained from the 1516 UTC 3 Dec 2015 sounding and mean LWCs from King probe measurements exceeding 0.05 g m−3. (center) (b) Histograms of the frequency of occurrence of particles by diameter, (c) flight-averaged PSDs, and (d) cumulative reflectivity by particle size constructed from 2D-S (0.225–1 mm) and HVPS-3 (1–32.5 mm) measurements. (right) Habit specific PSDs for (e) dendritic and (f) irregular particles from 2 to 10 mm corresponding to the blue boxed region in (c).

Citation: Journal of the Atmospheric Sciences 78, 11; 10.1175/JAS-D-21-0091.1

At S band, radar scattering is generally in the Rayleigh regime and attenuation effects are small such that radar reflectivity is proportional to the sixth power of particle’s mass-equivalent diameter and, therefore, dominated by the contributions of large particles. Thus, we evaluate how the distribution of particle sizes changes through the ice layer of the cloud. Figure 7b shows the frequency of time that particles of a given size were observed by the 2D-S and HVPS-3 OAPs. The most striking differences between the three layers are at sizes greater than about 2 mm. Above the secondary ZH maximum, particles greater than 2 mm were rarely observed but within and below the secondary ZH maximum, large particles were frequently observed up to about 10 mm in size. The flight-averaged spectral IWC (Fig. 7c) peaks in the secondary ZH maximum layer relative to the layers above and below between about 2 and 10 mm, an indication of rapid hydrometeor growth within this layer. Below the secondary ZH maximum, spectral IWC decreases in the 2–10-mm size range but increases for small particle sizes below about 1 mm. The cumulative reflectivity, ZPSD, increases rapidly with diameter for particles greater than about 2 mm, underscoring the sixth-power dependence on particle diameter (Fig. 7d).

Of the nine possible UIOOPS habit classes, the relatively high spectral IWC for 2–10-mm-sized particles in the secondary ZH maximum layer was dominated by dendritic and irregular habits (Figs. 7e,f). Figure 8 includes imagery of select particles that are broadly representative of the full set of images collected by the 2.3-μm-resolution Cloud Particle Imager (CPI) above, within, and below the secondary ZH maximum layer. The prevalence of dendritic and irregular habit types from UIOOPS is consistent with observations from the CPI (Figs. 8a,b).

Fig. 8.
Fig. 8.

In situ Cloud Particle Imager example imagery from (a) above, (b) within, and (c) below the secondary ZH maximum layers on 3 Dec 2015.

Citation: Journal of the Atmospheric Sciences 78, 11; 10.1175/JAS-D-21-0091.1

To put the in situ measurements in a context of radar-based inferences of ice microphysics, averaged vertical profiles of ZH and ZDR were constructed. Vertical columns of NPOL data at grid points containing a secondary ZH maximum were averaged over the radar sector and then adjusted to a brightband relative framework. Figure 9 shows the averaged vertical profiles coinciding with the 1432–1552 UTC portion of the Citation flight and the average positions of the secondary ZH maximum layer and the layers above and below, shaded by color according to the schematic in Fig. 7a. There is a local maximum in ZDR between 3 and 3.5 km above the bright band indicating the presence of horizontally oriented ice particles (Fig. 9b). The maximum in ZDR occurs near an onset of rapid increase in ZH with decreasing height (within the layer above the secondary ZH maximum, shaded in orange in Fig. 9a). Although the secondary ZH maximum maintains an approximately steady state during the Citation flight, there is some inherent variability of the secondary ZH maximum in both vertical position and magnitude. As a result, averaging over the radar domain results in a smoothed ZH profile with minimal apparent maximum aloft. However, there is a distinct change in the slope of the ZH profile as a result of the ZH maximum between 2 and 2.5 km above the bright band in each scan cycle (top of secondary ZH maximum layer, shaded in violet in Fig. 9a). In the latest two scan cycles (1612 and 1632 UTC), the ZDR maximum layer aloft apparently repositions itself below the secondary ZH maximum layer to approximately 1–1.5 km above the bright band. This time period coincides with a breakup of the upstream stratiform precipitation. While the cause of the apparent repositioning of the ZDR maximum layer is unclear, the secondary ZH maximum persists through the 1652 UTC radar scan cycle.

Fig. 9.
Fig. 9.

Brightband-relative vertical profiles of (a) reflectivity and (b) differential reflectivity from NPOL terrain-focused radar scan cycles coinciding with the 3 Dec 2015 flight. Shaded regions represent the position of the secondary ZH maximum layer and the 1-km layers above and below, colored according to the schematic in Fig. 7a. Layer heights are an average over all radar scan cycles.

Citation: Journal of the Atmospheric Sciences 78, 11; 10.1175/JAS-D-21-0091.1

Above the secondary ZH maximum near the local maximum in ZDR (i.e., 3–3.5 km above the bright band where temperatures are near −15°C; orange shading, Fig. 9b), in situ observations of nearly pristine or lightly aggregated branched and planar crystals (Fig. 8a) suggest that vapor deposition is the primary ice growth process. Below the local maximum in ZDR, there is a sharp downward increase in ZH and decrease in ZDR (orange shading, Figs. 9a,b) and heavily rimed crystals and aggregates were observed in the secondary ZH maximum layer (Fig. 8b). Supercooled liquid water, measured by the King probe, averaged 0.1 g m−3 in both the secondary ZH maximum layer and the layer above (Fig. 7a). These polarimetric signatures could be explained by aggregation alone (Kennedy and Rutledge 2011); however, riming of platelike crystals has a similar polarimetric signature (Schrom and Kumjian 2016). Indeed, the particle imagery indicates that aggregation and riming processes were occurring concurrently (Fig. 8b).

Depositional ice growth requires ice supersaturation and dendritic growth typically occurs when ice supersaturation is near water saturation (Bailey and Hallett 2009). Maintenance of supercooled liquid water necessary for riming also requires high ice supersaturation. Thus, conditions favorable for dendritic growth also favor riming. Previous studies have suggested that supercooled liquid water transport can be effective within turbulent overturning cells that additionally enhance aggregation efficiency by increasing the probability of particle collision (Houze and Medina 2005; Medina et al. 2007). In the radial velocity field of the 1453 UTC RHI, there is evidence of a low-level jet over the upstream ocean that is forced aloft as it encounters the windward slopes (Figs. 3e,f). The secondary ZH maximum occurs over the windward slopes near a sheared layer at the top of the low-level jet and immediately below a layer of enhanced spectrum width, indicative of turbulent air motion (Fig. 3h). This pattern was consistent across other terrain-focused RHIs during the 3 December event. The forced ascent of the low-level jet facilitates the advection of liquid water above the bright band whereby rimed mass is accumulated most efficiently by larger particles (e.g., Johnson 1987). The turbulent region above the secondary ZH maximum is expected to increase particle collisions, enhancing aggregation and therefore, the rate of rimed mass growth.

2) Case study: 1–2 December 2015

In the passage of a weakening upper-level trough from approximately 2000 UTC 1 to 0600 UTC 2 December, windward slopes of the Olympic Mountains received approximately 30–50 mm of stratiform precipitation (Petersen et al. 2017b). The Citation sampled the prefrontal portion of the storm in a vertical spiral over the windward slopes between 3 and 7 km MSL where temperatures were about −5° to −30°C (Fig. 10). The winds in the 2317 UTC 1 December sounding were overall weaker than on 3 December and included a slightly more easterly component at the surface (Fig. 5). The IVT was also less than for the 3 December event albeit relatively high at 500 kg m−1 s−1 (Zagrodnik et al. 2019). However, the most notable difference between these two events was the low-level stability. The 1–2 December event was moist stable in the 950–850-hPa layer (Nm2=1.06×104s2) whereas the 3 December event was moist neutral (Nm2=0.21×104s2; Zagrodnik et al. 2019).

Fig. 10.
Fig. 10.

As in Fig. 6, but for the flight between approximately 2300 UTC 1 Dec and 0200 UTC 2 Dec 2015. The KLGX PPI scan is for 2307–2312 UTC 1 Dec 2015.

Citation: Journal of the Atmospheric Sciences 78, 11; 10.1175/JAS-D-21-0091.1

During the 1–2 December event, a secondary ZH maximum was rarely identified over the upstream ocean but was consistently identified over the windward slopes between 2132 UTC 1 and 0552 UTC 2 December, with the exception of the 0032 and 0132 UTC scan cycles on 2 December. Secondary ZH maxima were most pronounced beyond 30 km from NPOL (dashed black ellipse in Fig. 11b) between the Prairie Creek and Wynoochee ground sites and where the primary ascending and descending Citation spiral occurred (Fig. 10). During the flight, 112, 98, and 487 1-Hz in situ observations were collected from above, within, and below the secondary ZH maximum, respectively. The SMlwr and SMupr levels coinciding with in situ observations respectively ranged from 1.56 to 1.7 km and from 2.05 to 2.24 km above the bright band. The secondary ZH maximum layer spanned −7.0° to −9.9°C (Fig. 12a) at the time of the 2317 UTC sounding. Supercooled liquid water was observed in the secondary ZH maximum layer with an average value of 0.11 g m−3.

Fig. 11.
Fig. 11.

As in Fig. 3, but for RHI scans at (a),(c),(e),(g) 2301 and (b),(d),(f),(h) 2313 UTC 1 Dec 2015.

Citation: Journal of the Atmospheric Sciences 78, 11; 10.1175/JAS-D-21-0091.1

Fig. 12.
Fig. 12.

As in Fig. 7 but for the 1–2 Dec 2015 flight. Temperature ranges for each layer shown in (a) were obtained from the 2317 UTC 1 Dec 2015 sounding.

Citation: Journal of the Atmospheric Sciences 78, 11; 10.1175/JAS-D-21-0091.1

As on 3 December, large ice particles (>2 mm) were observed much more frequently and with a higher flight-averaged spectral IWC within the secondary ZH maximum layer relative to the layers above and below (Figs. 12b,c). At 1 mm, disagreement between the 2D-S and HVPS-3 probe measurements appears significant. Sizing errors increase as particle size becomes small relative to the OAP resolution due to fewer pixels available to resolve the particle (Korolev et al. 1998). Thus, we expect measurements from HVPS-3 near the 1-mm transition size to have the greatest uncertainties. However, spectral IWC at this size range does not significantly affect the computed radar reflectivity estimates (Fig. 12d). Calculated cumulative ZPSD is similar in all three layers up to approximately 2 mm, but at larger sizes, ZPSD is enhanced in the secondary ZH maximum layer as a result of increases in spectral IWC dominated by dendritic and irregular habit classes (Figs. 12e,f).

Despite synoptic differences between the 3 December and 1–2 December events, average PSDs from each flight suggest that similar increases in spectral IWC at large particle sizes resulted in the increased ZH within the secondary ZH maximum layer. Correspondingly, the average vertical ZH profiles for each event are similar, although on 1–2 December, the secondary ZH maximum occurs about 1.75–2.25 km above the bright band (violet shading, Fig. 13a), slightly lower than on 3 December. A maximum in ZDR occurs near cloud top and decreases steadily to a minimum just above the bright band. Although these profiles lack a distinct signature of a dendritic growth zone in the form of a ZDR maximum, individual rimed and unrimed platelike crystals were observed above the secondary ZH maximum (Fig. 14a) where the average ZDR exceeds 0.4 dB, which is expected for a mixed population of dendritic crystals with varied degrees of riming and aggregation. Immediately above the secondary ZH maximum layer, from about 2.5 to 3 km above the bright band, the rapid downward increase in ZH and continued decrease in ZDR (orange shading, Fig. 13b) are consistent with onset of aggregation and possibly riming. The two events had similar LWC values in the secondary ZH maximum layer (~0.1 g m−3) but the LWC that was observed above the secondary ZH maximum layer on 3 December averaged 0.1 g m−3, whereas in the 1–2 December event, no supercooled liquid water was observed. Because riming efficiency increases with LWC (Jensen and Harrington 2015), it is probable that mass growth by riming above the secondary ZH maximum was less efficient on 1–2 December than during the 3 December event. This inference appears to be supported by in situ particle observations from the CPI probe which generally show a lesser degree of riming for the 1–2 December event, especially above the secondary ZH maximum (Fig. 14a). Within the secondary ZH maximum layer, however, riming was apparent in much of the particle imagery (Fig. 14b).

Fig. 13.
Fig. 13.

As in Fig. 9, but for 1–2 Dec 2015.

Citation: Journal of the Atmospheric Sciences 78, 11; 10.1175/JAS-D-21-0091.1

Fig. 14.
Fig. 14.

As in Fig. 8, but for 1–2 Dec 2015.

Citation: Journal of the Atmospheric Sciences 78, 11; 10.1175/JAS-D-21-0091.1

Previous studies have shown strong updrafts to be an effective indicator of riming processes (Moisseev et al. 2009; Schrom and Kumjian 2016). Based on heavier riming observed in the secondary ZH maximum layer and overall higher LWC aloft on 3 December, it is plausible that the updraft region over the windward slopes was either more intense or more vertically extensive than during the 1–2 December event. In the radial velocity field of the 2313 UTC 1 December RHI over the windward slopes, blocked low-level southeasterly terrain-parallel flow is evident by the negative radial velocity indicating flow toward the radar (Fig. 11f). Ascent can be inferred in the lifting of the low-level jet, but the stable layer acts as an upstream barrier extension effectively reducing the slope of the mountain. The mean flow is also notably weaker than on 3 December; thus, the inferred updraft over the windward slopes is also weaker. However, much like the 3 December event, the secondary ZH maximum region coincides with a sheared layer in the radial velocity field immediately below a layer of enhanced spectrum width over the windward slopes beyond 30 km (Figs. 11f,h). Even with higher low-level static stability than the 3 December event, liquid water was apparently transported aloft by orographically forced ascent of a low-level jet. The in situ observations and inferred microphysical processes associated with the secondary ZH maxima of the 1–2 December event are similar to those of the 3 December event. Above the secondary ZH maximum, branched and planar crystals readily aggregated, likely enhanced by turbulent air motion, and subsequently accumulated rimed mass to produce high spectral IWC at large particle sizes. In the following section, these microphysical properties are evaluated within the broader OLYMPEX dataset to assess general properties of the secondary ZH maxima layer.

c. Comparison of secondary ZH maximum and brightb and-only periods

To distinguish properties unique to secondary ZH maxima, we compare observations from all OLYMPEX secondary ZH maxima periods to brightband-only periods (i.e., radar scan cycles with a bright band but no secondary ZH maximum). In situ observations during brightband-only periods were compared to secondary ZH maximum periods by using temperature layers similar to those of the two case studies. Temperatures of −6° to −12°C represent the layer that a secondary ZH maximum would be likely to form, −12° to −18°C, the layer above and 0° to −6°C, the layer below as illustrated in Fig. 15a. During brightband-only periods from all flights, measurements within each layer were used to calculate average PSDs separately for the upstream ocean and the windward terrain (Figs. 15b,c). Observations collected in the 0° to −6°C layer on the 18 December flight were removed due to sampling at the bright band. Unfortunately, no observations were made within a secondary ZH maximum layer over the upstream ocean.

Fig. 15.
Fig. 15.

Average PSDs derived from 2D-S and HVPS-3 measurements on all flights. (a) Temperature ranges of brightband-only-period observations comparable to above (orange), within (green), and below (gray) the secondary ZH maximum layers. (b) PSDs derived from flights over the upstream ocean. (c) PSDs derived from flights over the terrain, with violet corresponding to the secondary ZH maximum layer. Observations from the 18 Dec 2015 were not used in the 0° to −6°C calculations due to sampling at the bright band.

Citation: Journal of the Atmospheric Sciences 78, 11; 10.1175/JAS-D-21-0091.1

Unlike during the 1–2 and 3 December secondary ZH maximum periods, there is no apparent increase in the average IWC for the −6° to −12°C layer relative to the layers above and below over the ocean or over terrain. Averaging over several flights in this manner is subject to a sampling bias depending on which layers were sampled for any given storm. However, each layer’s average PSD is constructed from a minimum of two separate flights. The average spectral IWC is substantially higher at larger sizes in the secondary ZH maximum periods when compared to any other temperature layer during brightband-only periods (Fig. 15c). Ice-phase growth regimes specified by temperature alone cannot sufficiently describe the observed differences in the vertical profile of ZH between secondary ZH maximum and brightband-only periods.

Riming in the secondary ZH maximum layer for both case studies motivates exploring LWC measurements for all secondary ZH maximum and brightband-only periods. Measurements of LWC exceeding 0.05 g m−3 from the King probe in each brightband-only vertical layer are shown in Fig. 16. The most frequent observations of supercooled liquid water over the ocean are found just above the bright band in the 0° to −6°C layer and rarely in the colder layers. Indeed, the RICE instrument indicated icing conditions in 24% of observations in the 0° to −6°C layer and in zero observations in either layer above. Over the terrain, the RICE instrument indicated icing in 13% of the observations in the brightband-only −6° to −12°C layer but 24% in the secondary ZH maximum layer. In general, supercooled liquid water was found more frequently over the terrain when compared to over the upstream ocean in all layers, but especially during secondary ZH maximum periods. The prevalence of riming due to supercooled liquid water in the secondary ZH maximum layer over the windward slopes has implications for the density, and thus fall velocity, of ice-phase particles (e.g., Jensen and Harrington 2015). We hypothesize that the rate of rimed mass growth for some secondary ZH maxima layer particles can be sufficient to overcome windward updraft velocities before being transported to the lee slopes, enhancing windward precipitation. Fallout of relatively large ice-phase hydrometeors implies a larger mass flux through the melting level, thus an increase in brightband ZH values and possibly surface accumulations.

Fig. 16.
Fig. 16.

Histograms of LWC measurements greater than 0.05 g m−3 from the King probe by layer as in Fig. 15a during all brightband-only periods (top) over the upstream ocean and (bottom) over the terrain. The violet histogram over the terrain corresponds to measurements from the secondary reflectivity ZH maximum layer.

Citation: Journal of the Atmospheric Sciences 78, 11; 10.1175/JAS-D-21-0091.1

Over the terrain-focused radar sector, histograms of brightband ZH values at individual grid points during all secondary ZH maxima and brightband-only periods are compared in Fig. 17. During brightband-only periods, ZH was most frequently between 25 and 30 dBZ but during secondary ZH maximum periods, there is a distinct shift in the peak of the distribution to about 35–40 dBZ with a marginal increase in the mean value. Below the bright band, at the surface, precipitation characteristics were evaluated at three windward sites during secondary ZH maxima and brightband-only periods (Fig. 18). Because of the uncertainties in ice particle trajectories over Pacific Northwest coastal mountains (Hobbs et al. 1973; Colle and Zeng 2004; Colle et al. 2005), we do not attempt to relate the location or magnitude of the secondary ZH maximum layer to the ground sites. Rather, we evaluate how the observed precipitation characteristics within the terrain-focused sector differ when a secondary ZH maximum was present and when only a bright band was present.

Fig. 17.
Fig. 17.

Histograms of ZH at the bright band from individual grid points within the terrain-focused radar sector during all secondary ZH maxima (violet) and brightband-only (green) periods. Dashed vertical lines indicate the mean values of the distribution.

Citation: Journal of the Atmospheric Sciences 78, 11; 10.1175/JAS-D-21-0091.1

Fig. 18.
Fig. 18.

Histograms of rain rate and mass-weighted mean drop diameter from OLYMPEX Parsivel2 disdrometer measurements during all secondary ZH maximum (violet) and brightband-only (green) periods at the three windward ground sites shown in Fig. 1. Dashed vertical lines indicate the mean values of each distribution.

Citation: Journal of the Atmospheric Sciences 78, 11; 10.1175/JAS-D-21-0091.1

During secondary ZH maximum periods, rain rate increased, and mass-weighted mean drop diameter was larger at all three ground sites relative to brightband-only periods with the most significant increases at the midmountain site, Prairie Creek (Fig. 18). The greater increases at Prairie Creek compared to Fishery are likely attributable to the preferential formation of secondary ZH maxima over mid- to high-elevation terrain. At the higher elevation Wynoochee site, the depth of the warm-rain layer is shallower than at Prairie Creek, thus the overall contribution of warm-rain processes that produce smaller drops (Zagrodnik et al. 2018) is already reduced. Therefore, the effect of additional melted ice particles during secondary ZH maximum periods is apparently less pronounced than at lower sites. The surface observations are consistent with our hypothesis of windward fallout of large, heavily rimed particles from the secondary ZH maximum layer. However, this analysis does not preclude contributions from concurrent embedded convective cells or enhanced warm-rain processes. Thus, the extent to which the secondary ZH maximum layer is the cause of increased rain rates and drop diameters on windward slopes is uncertain. This interesting finding warrants further investigation in future numerical modeling studies to determine possible trajectories of secondary ZH maxima layer particles, elucidating the effect of precipitation formed in this layer by connecting it to the surface.

5. Discussion and conclusions

During winter storms over North American western coastal mountains, previous studies have documented a secondary ZH maximum layer above the radar bright band (Houze and Medina 2005; Kingsmill et al. 2006; Medina et al. 2007; McMurdie et al. 2018). These previous studies raised several questions about the role of topography on the development of the secondary ZH maximum layer, the microphysical properties associated with this layer, and its effect on windward precipitation. We have used observations over the upstream ocean and over the windward slopes of the Olympic Mountains during OLYMPEX to investigate the microphysical properties of ice particles within secondary ZH maximum layers and discern the influence of orography. This study has resulted in the following conclusions:

  • Secondary ZH maxima layers are observed upstream but its frequency of occurrence and layer-average ZH are enhanced over the windward slopes, indicating an orographic enhancement of ice-phase precipitation processes.

  • Relatively high ZH values of the secondary ZH maximum layer are a result of high concentrations of large (~2–10-mm diameter) hydrometeors typically formed of aggregated dendritic or irregular branched crystals that are often heavily rimed.

  • Analysis of two case studies shows that when a secondary ZH maxima layer is observed over the windward slopes, it commonly occurs near high velocity shear at the top of a cross-barrier low-level jet and below a turbulent layer indicated by a local maximum in spectrum width.

  • The presence of a secondary ZH maxima layer is associated with increased rain rates and larger mass-weighted mean drop diameters at windward ground sites.

Investigations of the microphysics of ice-phase precipitation processes have often been limited to inferences of those processes based largely on polarimetric radar signatures (e.g., Kennedy and Rutledge 2011; Schrom et al. 2015; Kumjian and Lombardo 2017). Enhanced ZDR near −15°C temperatures often indicates a dendritic growth zone and is consistent with in situ particle imagery for the two case studies presented in section 4b. The regularity in formation of the secondary ZH maximum at temperatures just warmer than the dendritic growth zone (−6° to −12°C) and just below enhanced ZDR indicates that a secondary ZH maximum is not so much a signature of dendrites being present as much as an increase in the size of particles due to aggregation. Similar to Houze and Medina (2005), we speculate that aggregation is enhanced within shear-induced turbulent overturning cells during stably stratified and moist–neutrally stable low-level conditions. As aggregation is enhanced, individual particle sizes are increased, thus favoring more efficient accumulation of rimed mass which increases particle fall velocity, and therefore, the rate of fallout on the windward slopes. Orographic lifting of a low-level jet appears to efficiently maintain supercooled liquid water aloft for riming.

The relative impacts of ice-phase and warm-rain processes in orographically enhanced windward precipitation over the Olympic Mountains is a topic that has received significant discussion in recent studies (e.g., Zagrodnik et al. 2018; Conrick and Mass 2019a,b; Zagrodnik et al. 2021). These studies have demonstrated that warm-rain processes significantly contribute to the observed windward precipitation enhancement. Our findings suggest that ice particles formed in secondary ZH maxima layers may additionally enhance windward precipitation. Larger mean drop diameters observed at all three ground sites during secondary ZH maximum periods are an indication of accelerated ice growth aloft that facilitated the fallout of large particles to increase rain rates on the windward slopes.

Within the range of NPOL over the windward slopes, composites of high-resolution numerical modeling simulations by Zagrodnik et al. (2021) indicated that vertical velocities in the ice layer averaged approximately 0.2–1 m s−1 during prefrontal (1–2 December) and warm-sector (3 December) periods. We hypothesize that ice particles unable to grow fast enough to achieve a mass necessary for windward fallout are advected downstream in the low-level cross-barrier flow. Indeed, such differential advection may explain the slight decrease in ZH below the secondary ZH maximum layer. Alternatively, reduced ZH values below the secondary ZH maximum layer may be a result of decreased particle sizes due to breakup. Fragmented ice crystals (i.e., dendrite arms) were occasionally observed in the particle imagery (not shown), an indication of mechanical ice–ice collision fragmentation. Additionally, needles were observed in the particle imagery (Fig. 8c), suggesting that rime splintering (Hallett and Mossop 1974) was likely occurring. Such riming may have also resulted in particle fragmentation from thermal shock (Korolev and Leisner 2020).

Results from this study emphasize riming as a mass growth process complementary to aggregation in the secondary ZH maximum layer that has implications for particle densities. The rate of rimed mass growth increases with particle size (Johnson 1987) and, indeed, flight-averaged PSDs from both case studies in section 4b showed a size dependency in which significant spectral IWC was present at large particle sizes (>~2-mm diameter) within the secondary ZH maximum layer. These results have likely implications for remote sensing retrievals of precipitation over complex terrain. In general, remote sensing retrievals require assumptions about the snow density and shape of the PSD, properties that undergo distinct modification within the secondary ZH maximum layer. Future work will investigate the suitability of these retrieval assumptions and their effect on precipitation rate estimates over complex terrain using coincident in situ and airborne remote sensing retrievals.

Acknowledgments

The authors thank the Quinault Indian Nation and the staff of the Olympic National Park and U.S. Forest Service for permission to install and operate instrumentation on their land. The authors also thank the entire OLYMPEX team for successful data collection during the campaign. Dr. Joseph Finlon assisted with the UIOOPS processing and interpretation of the microphysics probe data. Dr. Joseph Boomgard-Zagrodnik provided code for processing the disdrometer data and assisted in its interpretation. Beth Tully provided graphics support for Figs. 8 and 14. The authors also thank Profs. Roger Marchand and Dale Durran for their useful discussions and feedback on this work. Comments from three anonymous reviewers significantly improved this manuscript. This research was supported by NSF Grant AGS-1657251.

Data availability statement

All data from OLYMPEX used in this study are available from the NASA Distributed Active Archive Center at https://ghrc.nsstc.nasa.gov/uso/ds_details/collections/gpmolyxC.html (Petersen et al. 2018). NCEP reanalysis data were provided by NOAA/OAR/ESRL PSL, Boulder, Colorado, from their website at https://psl.noaa.gov/data/gridded/data.narr.html.

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Save
  • Andrić, J., M. R. Kumjian, D. S. Zrnić, J. M. Straka, and V. M. Melnikov, 2013: Polarimetric signatures above the melting layer in winter storms: An observational and modeling study. J. Appl. Meteor. Climatol., 52, 682700, https://doi.org/10.1175/JAMC-D-12-028.1.

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  • Baker, B., and R. P. Lawson, 2006: Improvement in determination of ice water content from two-dimensional particle imagery. Part I: Image-to-mass relationships. J. Appl. Meteor. Climatol., 45, 12821290, https://doi.org/10.1175/JAM2398.1.

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  • Bechini, R., L. Baldini, and V. Chandrasekar, 2013: Polarimetric radar observations in the ice region of precipitating clouds at C-band and X-band radar frequencies. J. Appl. Meteor. Climatol., 52, 11471169, https://doi.org/10.1175/JAMC-D-12-055.1.

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  • Bell, M. M., M. Dixon, B. Javornik, W.-C. Lee, B. Melli, J. DeHart, and T.-Y. Cha, 2020: nsf-lrose/lrose-cyclone: lrose-cyclone-20200110. Zenodo, https://doi.org/10.5281/zenodo.3604387.

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  • Brdar, S., and A. Seifert, 2018: McSnow: A Monte-Carlo particle model for riming and aggregation of ice particles in a multidimensional microphysical phase space. J. Adv. Model. Earth Syst., 10, 187206, https://doi.org/10.1002/2017MS001167.