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    (a) Composite radar reflectivity factor (dBZ) valid at 1055 UTC 4 Jul 2014 and GOES-13 infrared (10.7 μm) data valid at (b) 0055, (c) 0245, and (d) 0545 UTC 4 Jul 2014. In (a), CI1 is the northern CI event and CI2 is the southern CI event. In (b)–(d), the arrows denote the moisture advected from the west into the CI region.

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    RAP analysis valid at 0600 UTC 4 Jul 2014 for (a) 300-hPa winds (kt; 1 kt = 0.5144 m s−1) and geopotential height (black contours; m), (b) 700-hPa winds (kt) and geopotential height (black contours; m), and (c) 850-hPa winds (kt) and geopotential height (black contours; m). (d) Surface observations valid at 0607 UTC. The white stars denote the CI locations, the zigzag lines in (a) and (b) denote a ridge axis, and the dashed lines in (a) and (b) denote a trough axis. In (c), the red dot denotes the location of the RAP sounding in Fig. 3d.

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    Observed soundings for (a) KDNR at 0000 UTC, (b) KLBF at 1200 UTC, and (c) KDDC at 1200 UTC. (d) RAP model sounding valid at 40.60°N, 100.86°W at 0700 UTC (see red dot on Fig. 2c for the location). In (d), the MAUL is denoted by the black arrow.

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    RAP analysis valid at 0700 UTC 4 Jul 2014 for (a) 700-hPa temperature advection (K day−1), (b) 500-hPa vorticity advection minus the 850-hPa vorticity advection (used to compute the 700-hPa differential vorticity advection), and (c) convergence used to analyze areas conducive to QG ascent.

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    RAP cross section valid at 0900 UTC 4 Jul 2014 through (a) 41.35°N (near CI1) and (b) 37.83°N (near CI2) of potential temperature (horizontal contours), specific humidity (g kg−1; color filled), and winds. The black dashed line denotes the initiation longitude, the red circle denotes the area of elevated convergence (identified based on the downstream decrease in the horizontal wind speed), the red dashed line in (b) denotes the short-wave trough, and the location of each cross section is in the inset in each panel. (c) The 700-hPa RAP analysis of horizontal convergence () valid at 0900 UTC 4 Jul 2014. The black contours identify areas of horizontal convergence greater than .

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    Frederick (KFDR) WSR-88D 0.5° elevation radar reflectivity factor (dBZ) valid at 1018 UTC 11 Jun 2008.

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    As in Fig. 2, but for 11 Jun 2008, and the dashed outlines in (b) denote areas of horizontal convergence (greater than ).

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    GOES-12 infrared (10.7 μm) channel valid at (a) 0631 and (b) 0931 UTC 11 Jun 2008. In (a), the arrow denotes the origin of the orographic wave cloud, and the dashed line denotes the SSW–NNE oriented cloud line. In (b), the arrow denotes the orographic wave cloud, the dashed shape outlines the area where RH increased rapidly (based on IR data), and the solid ellipse denotes the CI location.

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    As in Fig. 4, but for 0900 UTC 11 Jun 2008.

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    Observed soundings at (a) KAMA at 0000 UTC, (b) KLMN (Lamont) at 1200 UTC, and (c) KOUN (Norman, OK) at 1200 UTC. (d) Map indicating the locations of (e)–(i) the RUC model soundings at 0900 UTC. The color of the box surrounding the soundings corresponds to the color of the dot in (d).

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    RUC cross section valid 11 Jun 2008 through 35.21°N [see (a) inset] of potential temperature (solid contours), specific humidity (g kg−1; color filled), and horizontal wind (kt) at (a) 0300, (b) 0600, and (c) 0900 UTC. (d) As in Fig. 5c, but for 600 hPa at 0600 UTC 11 Jun 2008. The initiation location is denoted by the vertical black dashed line, the red arrow denotes the eastward propagating moisture, and the red circles in (b) and (c) denote the area of elevated horizontal convergence.

  • View in gallery

    Amarillo (KAMA) WSR-88D radar reflectivity factor (dBZ) valid at (a) 0954 UTC at 0.5° elevation and (b) 0910 UTC at 3.5° elevation on 15 Jul 2012. The red dashed ellipse in (b) denotes the area of locally enhanced radar reflectivity factor, possibly denoting the ECL.

  • View in gallery

    RAP analysis valid at 0600 UTC 15 Jul 2012 for (a) 300-hPa winds (kt) and geopotential height (black contours; m), (b) 700-hPa winds (kt) and geopotential height (black contours; m), and (c) 850-hPa winds (kt) and geopotential height (black contours; m). (d) Surface observations valid at 0607 UTC. The white stars denote the CI locations.

  • View in gallery

    As in Fig. 4, but for 0700 UTC 15 Jul 2012.

  • View in gallery

    Observed soundings at KAMA valid at (a) 0000 and (b) 1200 UTC. (c) Map indicating the locations of (d)–(i) the RAP model soundings at 0600 UTC. The color of the box surrounding the soundings corresponds to the color of the dot in (c).

  • View in gallery

    (a) As in Fig. 5a, but for 0600 UTC 15 Jul 2012. The red line denotes the region of elevated horizontal convergence (based on the downstream decrease in the horizontal wind speed and the downstream reversal in the wind direction), the cross-sectional location is denoted in the inset, and the black dashed line denotes the CI longitude. (b) As in Fig. 5c, but for 0900 UTC 15 Jul 2012.

  • View in gallery

    GFS analysis of potential temperature on the dynamic tropopause [2 potential vorticity unit (PVU) surface] valid at (a) 1200 UTC 9 Jul, (b) 1200 UTC 11 Jul, (c) 1200 UTC 13 Jul, and (d) 1200 UTC 15 Jul 2012. The black arrows in (b)–(d) denote the PV streamer that broke off during the anticyclonic wave-breaking event. Note that 1 PVU = 10−6 K kg−1 m2 s−1.

  • View in gallery

    Hastings (KUEX) WSR-88D 0.5° elevation radar reflectivity factor (dBZ) valid at 1001 UTC 9 Jul 2005.

  • View in gallery

    RUC analysis valid at 0500 UTC 9 Jul 2005 for (a) 300-hPa winds (kt) and geopotential height (black contours; m), (b) 700-hPa winds (kt) and geopotential height (black contours; m; the ridge axis is denoted by the zigzag line), and (c) 850-hPa winds (kt) and geopotential height (black contours; m). (d) Surface observations valid at 0343 UTC.

  • View in gallery

    (a) North Platte (KLNX) WSR-88D 0.5° radial velocity valid at 0657 UTC 9 Jul 2005, (b) Hastings (KUEX) WSR-88D 2.4° radial velocity valid at 0804 UTC 9 Jul 2005, (c) Goodland (KGLD) WSR-88D 0.5° radial velocity valid at 0632 UTC 9 Jul 2005, and (d) the radial velocity data in (a)–(c) projected onto a map of Kansas and Nebraska. The white arrows in (a)–(c) denote the location of the gravity waves, the black dashed lines in (c) denote the locations of the elevated boundary (straight dashed line) and the outflow boundary (curved dashed line), and the black arrow in (a) and (c) denotes the direction of propagation of the gravity waves as determined by radar-observed motions. The thick dashed white lines in (d) represent the orientation and hypothetical extension of the gravity waves, the thin red-and-white dashed line denotes the orientation of the waves in KUEX, and the white pentagon denotes the CI location.

  • View in gallery

    The 1-min ASOS data valid on 9 Jul 2005 for (a) Goodland, KS (KGLD), and (b) Grand Island, NE (KGRI). The blue curve is the wind speed (m s−1), the brown line is the wind direction (°), the wind barbs are as convection (half barb is 5 kt, and full barb is 10 kt), the red line is the temperature (°C), the green line is the dewpoint (°C), and the black line is the station pressure (hPa). Vertical dashed lines denote the crest of the gravity wave. The inset in (a) depicts the graphical representation of the gravity wave propagation direction similar to that in Gossard and Munk (1954). The data have been smoothed using a seven-point running average.

  • View in gallery

    (a) Observed sounding and (b) calculated Scorer parameter at KLBF at 1200 UTC 9 Jul 2005. (c) Map indicating the locations of (d)–(i) the RUC model soundings at 0600 UTC. The color of the box surrounding the soundings corresponds to the color of the dot in (c).

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Initiation Mechanisms of Nocturnal Convection without Nearby Surface Boundaries over the Central and Southern Great Plains during the Warm Season

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  • 1 School of Meteorology, University of Oklahoma, Norman, Oklahoma
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Abstract

The number of case studies in the literature of nocturnal convection has increased during the past decade, especially those that utilize high-spatiotemporal-resolution datasets from field experiments such as the International H2O Project (IHOP_2002) and Plains Elevated Convection at Night (PECAN). However, there are few case studies of events for convection initiation without a nearby surface boundary. These events account for approximately 25% of all nocturnal convection initiation (CI) events. Unique characteristics of these events include a peak initiation time later at night, a preferred initiation location in northern Kansas and southern Nebraska, and a preferred north–south orientation to linear convective systems. In this study, four case studies of convection that is initiated without a nearby surface boundary are detailed to reveal a number of possible initiation mechanisms, including quasigeostrophic-aided ascent, elevated ascent associated with convergent layers (of unknown causes), the low-level jet, and gravity waves. The case studies chosen illustrate the wide variety of synoptic-scale conditions under which these events can occur.

© 2018 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: Dylan W. Reif, dylanreif@ou.edu

Abstract

The number of case studies in the literature of nocturnal convection has increased during the past decade, especially those that utilize high-spatiotemporal-resolution datasets from field experiments such as the International H2O Project (IHOP_2002) and Plains Elevated Convection at Night (PECAN). However, there are few case studies of events for convection initiation without a nearby surface boundary. These events account for approximately 25% of all nocturnal convection initiation (CI) events. Unique characteristics of these events include a peak initiation time later at night, a preferred initiation location in northern Kansas and southern Nebraska, and a preferred north–south orientation to linear convective systems. In this study, four case studies of convection that is initiated without a nearby surface boundary are detailed to reveal a number of possible initiation mechanisms, including quasigeostrophic-aided ascent, elevated ascent associated with convergent layers (of unknown causes), the low-level jet, and gravity waves. The case studies chosen illustrate the wide variety of synoptic-scale conditions under which these events can occur.

© 2018 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: Dylan W. Reif, dylanreif@ou.edu

1. Introduction

The nocturnal maximum of precipitation during the warm season over the Great Plains has been well documented (e.g., Kincer 1916; Wallace 1975; Easterling and Robinson 1985; Riley et al. 1987; Surcel et al. 2010). Much of that precipitation has been attributed to the eastward propagation of mesoscale convective systems (MCSs) originating in the lee of the Rockies (Fritsch et al. 1986; Carbone et al. 2002; Tuttle and Davis 2006; Carbone and Tuttle 2008). The rest of the precipitation total is from events that initiate in the Great Plains during the night. Reif and Bluestein (2017, hereafter RB17) documented nocturnal convection initiation (CI) events over the Great Plains during the warm seasons of 1996–2015 and identified three location-based CI modes: events that initiate on a surface boundary [at-boundary (AB) CI mode], events that initiate on the cold side of a surface boundary [cold-side (CS) CI mode], and events that initiate with no well-defined surface boundary nearby [no-boundary (NB) CI mode]. The latter mode (NB CI mode) has been referred to in the literature also as “nonfrontal” (Pitchford and London 1962) or as “pristine initiation” (Wilson et al. 2018).

Much less attention has been paid to events in the NB CI mode than to those in the AB CI mode or the CS CI mode even though NB CI mode events account for ~25% of all nocturnal CI events (RB17). These NB CI mode events have unique characteristics such as north–south-oriented linear systems, a relatively late initiation time [0900 UTC or 0400 local daylight time (LDT)], and a preferred initiation location in southern Nebraska and northern Kansas, suggesting that there may be a unique mechanism that could lead to CI. Forecasting the initiation of deep convection remains a challenge (Romine et al. 2016; Keclik et al. 2017), especially NB CI mode events because they occur over a wide range of synoptic environments (as will be discussed in a later section), large-scale circulations may be poorly handled by numerical models (e.g., Davis et al. 2003), and because features above the stable boundary layer, which are not as well observed as surface features, become more important in initiating nocturnal convection (Jankov and Gallus 2004; Trier et al. 2006; Surcel et al. 2010; Squitieri and Gallus 2016). This study describes various features that may be important for the initiation of NB CI mode events, describes mechanisms that may initiate these nocturnal storms, examines case studies to illustrate the relative importance of each of these initiation mechanisms, and highlights the differences in the synoptic environments among the events.

A wide variety of possible initiation features and mechanisms will be suggested in section 2. The purpose of this paper is to identify which of the previously discussed initiation mechanisms may have occurred in some NB CI mode events and to illustrate the different synoptic-scale environments in which these events occur. Cases of NB CI mode events were selected for which the synoptic patterns were all different. The authors had no prior knowledge of the initiation mechanism of the case prior to case selection. The most likely initiation mechanism in each case is determined by a process of elimination: we eliminate the mechanisms that cannot affect the initiation so that the mechanism(s) that remains (remain) is (are) the most likely one(s). This methodology is similar to the ingredients-based forecasting method illustrated by Doswell (1987) and the methodology of Koch et al. (2005) used to identify model-predicted features necessary to form microscale eddies, but rather than focusing on a quantitative evaluation of the ingredients, we focus on a qualitative evaluation of the features present. It is extremely difficult to identify with certainty the exact initiation mechanism, so this study will identify the most likely mechanism based on the evidence presented. An overview of initiation features and mechanisms likely important to NB CI mode events is presented in section 2; the nature of the data sources is noted in section 3; four cases of NB CI mode events are discussed in section 4 to illustrate the different synoptic-scale environments and the different initiation mechanisms associated with these events. A summary of our results and conclusions are given in section 5.

2. Overview of initiation feature and mechanisms

We will focus on initiation-related features and mechanisms that have been mentioned in the literature such as the nocturnal low-level jet (LLJ), quasigeostrophic (QG)-aided ascent, gravity waves, elevated ascent associated with elevated convergent layers, cyclonic potential vorticity (PV) anomalies, isentropic ascent, and the mountain–plains solenoidal (MPS) circulation,1 though the reader should note that this is not an exhaustive list and other features or mechanisms may initiate convection. Many of these may contain aspects of each. For example, midtropospheric perturbations (Wang et al. 2011a,b), cyclonic PV anomalies, and short-wave troughs (all of which may be equivalent to each other) may be explained, at least in part, by quasigeostrophic theory. It is important to note that we are narrowing down the initiation mechanisms to those that are important to the initiation of convection with no nearby surface boundary. Thus, while we acknowledge that features such as cold fronts and drylines can play a role in the initiation of convection at night, we will not consider these features here.

a. Nocturnal low-level jet

The LLJ is a wind maximum within the lowest kilometer above the ground (Bonner 1968; Shapiro et al. 2016). Bonner (1968) classified LLJs based on the relative strength of the low-level wind maximum (Table 1). The LLJ has long been attributed to the initiation of convection under weakly synoptically forced environments (Pitchford and London 1962). Convergence at the terminus of the jet (where the wind speed decreases downstream) may result in CI (e.g., Tuttle and Davis 2006), and temperature and moisture advection over an east–west-oriented surface front destabilizes the environment on the cold side of the front and could result in CI (e.g., Maddox et al. 1979; Trier and Parsons 1993; Moore et al. 2003). However, the latter is not applicable in our cases since there is no nearby surface boundary but would be applicable to events in the CS CI mode.

Table 1.

The extended Bonner (1968) LLJ criterion.

Table 1.

Another explanation for LLJ-related initiation is that associated with convergence on the anticyclonic-shear side of the jet. Bonner (1966) showed that there is ascent on the eastern side of the jet during the early morning hours (~1100 UTC). Recent studies have attempted to explain the ascent on the anticyclonic-shear side of the LLJ. Pu and Dickinson (2014) suggested that as the anticyclonic vorticity on the anticyclonic-shear side of the jet decreased (i.e., cyclonic vorticity increased) as a result of the decreasing jet strength (i.e., after ~0800–0900 UTC), then according to the continuity equation [their Eq. (3)], ascent is possible. More recently, Gebauer et al. (2018) showed that there is moisture advection and convergence atop a strongly veering LLJ. This convergence may result in CI in quasi-north–south-oriented systems later at night when the LLJ is strongest, which is consistent with a subset of NB CI mode events in RB17 for which the orientation of linear systems tends to be along a north–south line.

b. Quasigeostrophic ascent

The physical interpretation of the traditional form of the quasigeostrophic omega equation is that warm advection, vorticity advection becoming more cyclonic with height, and diabatic heating are associated with ascent. Bluestein (1985) suggested that the lifting mechanism of an event similar to an NB CI mode event was a shallow layer of warm advection collocated with a relatively moist layer. Trier et al. (2017) suggested that mesoscale ascent collocated with warm advection can aid in the creation of a moist absolutely unstable layer (MAUL; Bryan and Fritsch 2000) but did not suggest that QG forcing alone could result in an elevated MAUL. The development of this MAUL could result in CI as minimal ascent is required to lift a parcel to its LFC.

The QG approximation is applicable to environments associated with NB CI mode events because the ascent generally occurs above the stable boundary layer and the ageostrophic wind is relatively weak in the absence of strong QG forcing. However, although one must use caution when analyzing the environment near the highly ageostrophic LLJ, a qualitative understanding of vertical motions may still be obtained (Keyser et al. 1992).

Initiation mechanisms noted in the literature including cyclonic PV anomalies, midtropospheric perturbations, and isentropic ascent can be described using a QG framework. The following subsections will outline each of these mechanisms and how QG theory may explain them.

1) Cyclonic PV anomalies

Late-afternoon convection in the lee of the Rockies generates PV anomalies via diabatic processes (e.g., Hoskins et al. 1985; Li and Smith 2010) such as latent heat release from ongoing convection or solar insolation on locally high terrain. Once the diabatic process is removed (e.g., the convection decays and/or the sun goes down), the PV anomalies are advected eastward with the mean flow. Raymond and Jiang (1990) showed that ascent of ~1 cm s−1 occurs on the downshear side of a cyclonic PV anomaly and subsidence occurs on the upshear side of the anomaly. The magnitude of this ascent (typically in the midtroposphere), if persistent, may be enough to lift an elevated parcel to its LFC. If these anomalies are generated during the late afternoon/early evening (~0000 UTC), then by 0900 UTC (the time when initiation of NB CI mode events is most frequent), a parcel would have ascended ~330 m in the absence of any other effects. If one assumed an eastward motion of ~15 m s−1 (Carbone et al. 2002; Carbone and Tuttle 2008; Li and Smith 2010), then the PV anomaly would have moved ~500 km in 9 h. If that PV anomaly originated in the lee of the Rockies, then after 9 h, the anomaly would be in central Kansas/Nebraska. It is also likely that the source air for CI is not affected by the PV anomaly for the entire duration of its eastward drift. For example, the source air for a storm in Nebraska could originate near Oklahoma, while the PV anomaly drifts eastward from the Rockies. There may be some period of time where the air is affected by the anomaly, but it is likely not the entire time.

The phase of the eastward-drifting PV anomalies corresponds well to that of the propagation of nocturnal precipitation (Li and Smith 2010), but it could be that latent heat release from the preexisting convective systems are responsible for the PV signal such that the phase between propagating convection and propagating PV anomalies should be similar. The studies linking these two phases together (e.g., Li and Smith 2010; Carbone et al. 2002; Carbone and Tuttle 2008) do not link the phase of the PV drift to CI in the plains. While the timing and location of the eastward drift of the PV anomaly appears to be located near the climatological maximum in nocturnal CI, it is difficult to say what effect that PV anomaly has on ascent needed for CI. If eastward-drifting PV anomalies are common, why does CI not occur every night? Other features or mechanisms must be playing a role in CI.

2) Isentropic ascent

Under adiabatic conditions, parcels are confined to the isentropic layer in which they reside (Raymond and Jiang 1990). Ascent along an isentropic surface has been used to explain some aspects of elevated convection. Corfidi et al. (2008, p. 1283) state that altocumulus castellanus, a type of cloud associated with elevated convection, “is a visual manifestation of the release of conditional instability as the result of large-scale isentropic ascent of shallow moist layers in the elevated mixed layer.” On an isentropic surface, an isobar represents an isotherm, so pressure advection is equivalent to temperature advection. However, simply stating that a parcel must follow an isentrope under adiabatic conditions could result in the release of condition instability does not say anything about the physical processes responsible for the ascent. These physical processes could be quasigeostrophic-aided ascent through warm advection or cyclonic differential vorticity advection, a frontogenesis-related vertical circulation, or vertical circulations associated with the low-level jet.

c. Gravity waves and bores

Gravity waves, including Kelvin–Helmholtz (KH) waves, convectively induced, deep-tropospheric gravity waves, atmospheric bores, and solitons are common at night (e.g., Fovell et al. 2006; Koch et al. 2008). These waves are associated with couplets of ascent and descent and can result in ascent from 500 m to over 1 km between 1 and 5 h (assuming a vertical velocity of 1–5 m s−1). Fovell et al. (2006) showed that CI can occur downstream of an MCS as a result of diabatically generated high-frequency gravity waves, and Wilson et al. (2018) suggested that gravity waves contributed to the initiation of an elevated isolated cell during the Plains Elevated Convection at Night (PECAN; Geerts et al. 2017) field project.

Atmospheric bores (Koch et al. 1991; Knupp 2006) can also be associated with alternating bands of ascent and subsidence and can result in the semipermanent lift of a layer on the order of ~1 km, making it possible for parcels to achieve their LFC. Environments supportive of bores are those that are either partially blocked or completely blocked (Rottman and Simpson 1989, their Fig. 2), and these conditions are common at night over the Great Plains during the warm season (Haghi et al. 2017). To maintain low-level gravity waves, the energy from the waves cannot propagate vertically, so the energy must be trapped in the lowest levels. Wave trapping occurs when the Scorer parameter [l; Scorer 1949; where , N is the Brunt–Väisälä frequency, and U is the wave-relative wind speed] decreases with height (Crook 1988), which can occur in regions of strong low-level curvature in the vertical profile of the component of the wind normal to the density current (aided by the nocturnal LLJ; Koch et al. 1991, 2008) and the presence of a temperature inversion. Bores have been shown to initiate convection (Wilson and Roberts 2006; Watson and Lane 2016).

d. Other features and mechanisms

1) Elevated convergent layers

An elevated convergent layer (ECL; Banacos and Schultz 2005) is commonly associated with NB CI mode events and may be orographically influenced, the result of an inhomogeneous, strongly veering nocturnal LLJ, or caused by elevated quasigeostrophic forcing. Nearly two-thirds of the elevated initiation events during IHOP_2002 (Weckwerth et al. 2004) were associated with elevated convergent layers (Wilson and Roberts 2006), and some of these ECLs have been associated with a line of storms behind an MCS (known as a bow-and-arrow MCS; Keene and Schumacher 2013). Peters and Schumacher (2015, 2016) discussed how echo training on the back side of a convective system can be associated with an ECL, but that ECL was associated with a preexisting MCS. Few studies have explained the origin of an ECL not associated with a nearby convective system. The feature that produced the ECL is the initiation mechanism, not the ECL itself.

2) MPS circulation

Differential heating between the Rockies and the plains can induce a solenoidal circulation known as the MPS circulation (Tripoli and Cotton 1989; Wolyn and McKee 1994). During the day, a convergence zone develops in the lee of the Rockies, which results in ascent during the day in the lee of the Rockies and subsidence over the plains (potentially limiting daytime convection initiation over the plains). During the night, this circulation reverses sign, and there is broad ascent over the plains (Carbone and Tuttle 2008). Wolyn and McKee (1994) showed that the descending branch of the MPS occurs ~100 km from the base of the mountains. Vertical velocities associated with the MPS have been shown to range between 1 and 5 cm s−1 (Carbone and Tuttle 2008). However, vertical motions from the MPS can be heavily modified by meridionally varying mesoscale features such as ongoing convection, fronts, and the low-level jet, and those features can help focus the precipitation into narrow latitude belts (e.g., Trier et al. 2010).

3. Data

Four NB CI mode events (Table 2) were selected from the RB17 climatology. These cases illustrate the variety of storm and environmental characteristics for which these events can occur. These events can produce severe weather, though not frequently, so only one case study will be of an event that was associated with severe weather reports (severe reports were obtained from the NOAA Storm Event database, available at https://www.ncdc.noaa.gov/stormevents/).

Table 2.

List of events selected for case studies and their respective CI mode and initiation location.

Table 2.

To analyze these cases, data from National Weather Service (NWS) soundings and Weather Surveillance Radar-1998 Doppler (WSR-88D) radars (Crum and Alberty 1993), as well as the Rapid Update Cycle (RUC; Benjamin et al. 2004) and the Rapid Refresh (RAP; Benjamin et al. 2016) operational models are used. The 1-min Automated Surface Observing System2 (ASOS; National Weather Service 1998) wind, temperature, and pressure data are used to identify features associated with the passage of gravity waves. Finally, satellite and infrared data from the Geostationary Operational Environmental Satellite-East (GOES-E; Menzel and Purdom 1994) are used to identify cloud features (gridded satellite data are available at https://www.bou.class.noaa.gov/saa/products/welcome).

4. Case studies

a. 4 July 2014

At 0700 UTC 4 July 2014, convection initiated in south-central Nebraska, forming an arc that extended into the Texas Panhandle by 1030 UTC (Fig. 1). The event in Nebraska (CI1) initiated ~2.5 h before the arc that extended from Kansas to Texas (CI2). The convective system propagated slowly eastward, dissipated by 1600 UTC, and was not associated with any severe weather reports.

Fig. 1.
Fig. 1.

(a) Composite radar reflectivity factor (dBZ) valid at 1055 UTC 4 Jul 2014 and GOES-13 infrared (10.7 μm) data valid at (b) 0055, (c) 0245, and (d) 0545 UTC 4 Jul 2014. In (a), CI1 is the northern CI event and CI2 is the southern CI event. In (b)–(d), the arrows denote the moisture advected from the west into the CI region.

Citation: Monthly Weather Review 146, 9; 10.1175/MWR-D-18-0040.1

At 300 hPa, a ridge in geopotential height was situated over the Rocky Mountains, a trough was located north of the Great Lakes (outside of the figure boundary on Fig. 2a), and there was weak north-northwesterly flow over the Great Plains at 0600 UTC (Fig. 2a). A weak short wave appears from the western northeast panhandle to northeastern New Mexico, but this trough appears to be related to ongoing convection since it is quasi stationary and short-lived. At 700 hPa, there was a short-wave trough upstream of the initiation location (see dashed line in Fig. 2b). A criterion-3 LLJ (Table 1) formed after sunset and spanned a region from the Texas Panhandle to south-central South Dakota (Fig. 2c); CI occurred east of the jet maximum. The surface winds were from the southeast over much of the region, the temperatures were in the mid- to upper 60s °F (~20°C), and the dewpoint temperatures were in the mid- to upper 50s °F (~13°C; Fig. 2d). Since there were no strong [>10 K (100 km)−1] temperature or moisture gradients or wind shift near the initiation location, we conclude that there were no nearby surface boundaries.

Fig. 2.
Fig. 2.

RAP analysis valid at 0600 UTC 4 Jul 2014 for (a) 300-hPa winds (kt; 1 kt = 0.5144 m s−1) and geopotential height (black contours; m), (b) 700-hPa winds (kt) and geopotential height (black contours; m), and (c) 850-hPa winds (kt) and geopotential height (black contours; m). (d) Surface observations valid at 0607 UTC. The white stars denote the CI locations, the zigzag lines in (a) and (b) denote a ridge axis, and the dashed lines in (a) and (b) denote a trough axis. In (c), the red dot denotes the location of the RAP sounding in Fig. 3d.

Citation: Monthly Weather Review 146, 9; 10.1175/MWR-D-18-0040.1

During the late afternoon of 3 July 2014, there were convective storms over the Rockies. This convective activity dissipated by 0430 UTC 4 July (Fig. 1d). The 0000 UTC Denver, Colorado (KDEN), sounding (Fig. 3a) represented the convective environment of the late afternoon when there was convection over the Rockies (Fig. 1b). A well-mixed layer extended from the surface to ~600 hPa, and the environment was relatively moist from 600 to ~370 hPa. This moisture was advected eastward (as seen as clouds on infrared imagery; Figs. 1b–d) and reached the initiation location by 0500 UTC. An examination of a time series of RAP cross sections of specific humidity supports this idea as an area of higher specific humidity between 700 and 550 hPa moves from the Rockies to the plains between sunset and the time of CI (not shown). The 1200 UTC North Platte, Nebraska (KLBF), sounding (Fig. 3b) had a low-level moist layer from 875 to 650 hPa, and farther south, the Dodge City, Kansas (KDDC), sounding (Fig. 3c) had a layer of relatively high moisture that extended from 900 to 600 hPa (though there is a somewhat dry layer between 900 and 750 hPa). The atmosphere was stable below 800 hPa but was conditionally unstable above that (based on the lapse rates in Figs. 3b and 3c). Based on the stability and moisture profiles, it is hypothesized that air parcels ascending to initiate convection were rooted above 800 hPa.

Fig. 3.
Fig. 3.

Observed soundings for (a) KDNR at 0000 UTC, (b) KLBF at 1200 UTC, and (c) KDDC at 1200 UTC. (d) RAP model sounding valid at 40.60°N, 100.86°W at 0700 UTC (see red dot on Fig. 2c for the location). In (d), the MAUL is denoted by the black arrow.

Citation: Monthly Weather Review 146, 9; 10.1175/MWR-D-18-0040.1

There was strong warm advection at 700 hPa and anticyclonic differential vorticity advection (DVA; based on the vorticity advection at 500 hPa minus the vorticity advection at 850 hPa) over much of southwestern Nebraska and northwestern Kansas (Figs. 4a,b). There is evidence of a shallow short-wave trough near 600 hPa (see black dashed line in Fig. 2b and red dashed line in Fig. 5b), but DVA calculated using the 600-hPa vorticity advection instead of 500-hPa vorticity advection was similar to that using 500 hPa (not shown). The strong warm advection in the low to midtroposphere could have resulted in the creation of a MAUL, and RAP soundings taken at 0700 UTC inside an approximately 1.5° latitude by 1.5° longitude grid near the first initiation location (near North Platte) support this idea (an example of one is shown in Fig. 3d). The total quasigeostrophic forcing can be shown using Q-vector convergence (Bluestein 1993). There are two notable areas of Q-vector convergence: western Nebraska and western and southwestern Kansas (Fig. 4c). These locations correspond well with the CI locations.

Fig. 4.
Fig. 4.

RAP analysis valid at 0700 UTC 4 Jul 2014 for (a) 700-hPa temperature advection (K day−1), (b) 500-hPa vorticity advection minus the 850-hPa vorticity advection (used to compute the 700-hPa differential vorticity advection), and (c) convergence used to analyze areas conducive to QG ascent.

Citation: Monthly Weather Review 146, 9; 10.1175/MWR-D-18-0040.1

Fig. 5.
Fig. 5.

RAP cross section valid at 0900 UTC 4 Jul 2014 through (a) 41.35°N (near CI1) and (b) 37.83°N (near CI2) of potential temperature (horizontal contours), specific humidity (g kg−1; color filled), and winds. The black dashed line denotes the initiation longitude, the red circle denotes the area of elevated convergence (identified based on the downstream decrease in the horizontal wind speed), the red dashed line in (b) denotes the short-wave trough, and the location of each cross section is in the inset in each panel. (c) The 700-hPa RAP analysis of horizontal convergence () valid at 0900 UTC 4 Jul 2014. The black contours identify areas of horizontal convergence greater than .

Citation: Monthly Weather Review 146, 9; 10.1175/MWR-D-18-0040.1

Based on the east–west-oriented RAP cross sections (Fig. 5), there may be two different initiation mechanisms. The most likely initiation mechanism for the northern event (CI1) is QG-aided ascent via warm advection (though the LLJ may play a role as well; see convergence in Fig. 5a), while the most likely initiation mechanism for the southern event (CI2; the arc from Kansas to the Texas Panhandle) is LLJ-related ascent. The reasoning for the second event’s initiation mechanism is based on convergence at ~750 hPa (Figs. 5b,c), weaker QG forcing compared to that of the northern event (Fig. 4c), and its location with respect to the LLJ [on the anticyclonic-shear side, a location favorable for the mechanism noted in Gebauer et al. (2018)].

In summary, convection initiated along a southwest–northeast-oriented arc that extended from south-central Nebraska to the Texas Panhandle. It appears that the initiation of the first event (in Nebraska) was aided by quasigeostrophic forcing associated with warm advection in a relatively moist environment (the moisture was likely remnants from prior convection over northeastern Wyoming; see white arrows in Figs. 1b–d), while the initiation of the second event (the arc of convection extending from central Kansas to the Texas Panhandle) was related to elevated convergence that possibly was associated with the LLJ collocated with an area of elevated moisture (that moisture was likely remnants from prior convection over the Rockies; see black arrows in Figs. 1b–d).

b. 11 June 2008

At 0850 UTC 11 June 2008, a north-northeast–south-southwest-oriented line of convection was initiated over southwestern Oklahoma (Fig. 6). It propagated slowly eastward, dissipated by 1630 UTC in eastern Oklahoma (not shown), and was not associated with any severe reports. No low-level features such as parallel bands of enhanced reflectivity forced by gravity waves, which may have aided initiation, were evident in the Frederick, Oklahoma (KFDR), or Twin Lakes (KTLX) radar data.

Fig. 6.
Fig. 6.

Frederick (KFDR) WSR-88D 0.5° elevation radar reflectivity factor (dBZ) valid at 1018 UTC 11 Jun 2008.

Citation: Monthly Weather Review 146, 9; 10.1175/MWR-D-18-0040.1

At 300 (and 500) hPa, there was a trough over the western United States, a ridge was over the Midwest, and there was weakly diffluent, westerly flow over the Great Plains (Fig. 7a). At 700 hPa, the flow was from the southwest over the Great Plains, and the initiation location (indicated by a star) was south of a jet streak (Fig. 7b). There was a strong southwesterly jet at 850 hPa, and CI occurred on the anticyclonic-shear side (eastern side) of the jet (Fig. 7c). On the morning of 10 June, there was a warm front in northern Texas (not shown). As the day progressed, the warm front propagated northward and dissipated by the early evening hours of 10 June. From that time and until the time of CI, there was no surface boundary evident over Oklahoma or southern Kansas (Fig. 7d).

Fig. 7.
Fig. 7.

As in Fig. 2, but for 11 Jun 2008, and the dashed outlines in (b) denote areas of horizontal convergence (greater than ).

Citation: Monthly Weather Review 146, 9; 10.1175/MWR-D-18-0040.1

Prior to CI, moisture that originated from south-central Colorado was advected from the west, appearing as low- to midlevel clouds in the early morning infrared imagery (Fig. 8). The origin of this moisture is uncertain, but it may have been partly the result of moisture remnants from an orographically induced gravity wave. Coincident with this moisture from Colorado, there was a sudden moistening of the environment prior to CI and also appears as clouds in the IR imagery (see dashed outline in Fig. 8b). This sudden moistening likely indicates broad ascent, which could be the result of quasigeostrophic forcing. There was warm advection and weak differential vorticity advection in the area, and the overall QG forcing indicates ascent over the Texas Panhandle into western Oklahoma (Fig. 9). This higher moisture content in the middle troposphere helps destabilize the environment (through a reduction in elevated CIN) such that another feature(s) could initiate convection.

Fig. 8.
Fig. 8.

GOES-12 infrared (10.7 μm) channel valid at (a) 0631 and (b) 0931 UTC 11 Jun 2008. In (a), the arrow denotes the origin of the orographic wave cloud, and the dashed line denotes the SSW–NNE oriented cloud line. In (b), the arrow denotes the orographic wave cloud, the dashed shape outlines the area where RH increased rapidly (based on IR data), and the solid ellipse denotes the CI location.

Citation: Monthly Weather Review 146, 9; 10.1175/MWR-D-18-0040.1

Fig. 9.
Fig. 9.

As in Fig. 4, but for 0900 UTC 11 Jun 2008.

Citation: Monthly Weather Review 146, 9; 10.1175/MWR-D-18-0040.1

An additional cloud feature appears as a south-southwest–north-northeast-oriented line on the IR imagery (see straight dashed line on Fig. 8a). This moisture may have been the result of the advection of moisture from the top of a well-mixed layer seen in the 0000 UTC KAMA sounding (Fig. 10a). The well-mixed layer extended from the surface to ~550 hPa. This relatively narrow band of elevated moisture was advected eastward toward Oklahoma by 0700 UTC [see area of higher relative humidity near 510 hPa in the 1200 UTC Lamont, Oklahoma (KLMN), sounding; Fig. 10b]. This already enhanced cloud feature, collocated with the broad region of ascent, could provide a focus for CI.

Fig. 10.
Fig. 10.

Observed soundings at (a) KAMA at 0000 UTC, (b) KLMN (Lamont) at 1200 UTC, and (c) KOUN (Norman, OK) at 1200 UTC. (d) Map indicating the locations of (e)–(i) the RUC model soundings at 0900 UTC. The color of the box surrounding the soundings corresponds to the color of the dot in (d).

Citation: Monthly Weather Review 146, 9; 10.1175/MWR-D-18-0040.1

There was ~1200 J kg−1 of most unstable CAPE (MUCAPE) in the 1200 UTC KLMN sounding (Fig. 10b) and ~2900 J kg−1 of MUCAPE in the 1200 UTC KOUN sounding (Fig. 10c), but there was −436 and −141 J kg−1 of most unstable CIN (MUCIN) in each sounding, respectively. If ascent of is needed to force a parcel at the initiating level up through the cap, the vertical velocities would have had to have been >29 and 17 m s−1 for each respective sounding, which is unreasonable in an environment with no strong forcing. Therefore, parcels that were lifted must have originated above the most unstable level (~900 hPa for each sounding). Based on the static stability and the moisture, the parcels likely ascended from between 800 and 750 hPa, and the MUCAPE and MUCIN for a parcel lifted between 800 and 750 hPa was ~857 and −249 J kg−1, respectively, for a parcel lifted from 782 hPa based on the 1200 UTC KOUN sounding. RUC model soundings taken at five grid points near the CI location (see map in Fig. 10d) suggest that the model did capture the narrow band of elevated moisture, but the surrounding environment was not moist (Figs. 10e–i).

In the RUC cross section through central Oklahoma at 0600 UTC (Fig. 11b), there was a relatively moist stable layer that extended from the surface to ~700 hPa, and the atmosphere was conditionally unstable above that level (based on the nearly vertically oriented isentropes from 600 and 550 hPa between the initiation location and near 102°W). There was an area of horizontal convergence between 675 and 550 hPa (Fig. 11), well above the low-level stable layer and in the conditionally unstable air. This convergence may be partly the result of ascent associated with a trough upstream of the CI location. Higher wind speeds around the base of the trough converged into relatively weak winds (Fig. 7b), resulting in a southwest–northeast-oriented band of elevated convergence (denoted by the black dashed outline in Fig. 7b). An examination of a time series of specific humidity from RUC cross sections show that an area of higher moisture near 800–750 hPa was advected eastward from the Texas Panhandle into west-central Oklahoma by 0900 UTC (see light blue arrows in Fig. 11).

Fig. 11.
Fig. 11.

RUC cross section valid 11 Jun 2008 through 35.21°N [see (a) inset] of potential temperature (solid contours), specific humidity (g kg−1; color filled), and horizontal wind (kt) at (a) 0300, (b) 0600, and (c) 0900 UTC. (d) As in Fig. 5c, but for 600 hPa at 0600 UTC 11 Jun 2008. The initiation location is denoted by the vertical black dashed line, the red arrow denotes the eastward propagating moisture, and the red circles in (b) and (c) denote the area of elevated horizontal convergence.

Citation: Monthly Weather Review 146, 9; 10.1175/MWR-D-18-0040.1

There was a criterion-3 LLJ (Table 1), and the maximum wind speeds extended from the Texas Panhandle to southeastern Nebraska. However, this LLJ appears to have been synoptically influenced by a low pressure area near 850 hPa in southwestern Nebraska (Fig. 7c). Near the CI location, the jet did not veer strongly with time (not shown), and there was little convergence related to the LLJ (Figs. 7c, 11c). However, as previously noted, there was convergence above the jet, but it does not appear to be related to the veering of the jet.

If the cloud feature in south-central Colorado was in fact an orographically influenced cloud, then air must have ascended up and over the Rocky Mountains. The Froude number is expressed as
e1
where U is the wind speed normal to the obstacle, N is the Brunt–Väisälä frequency, and H is the height of the obstacle; it can be used to determine whether or not flow can ascend up an over an obstacle or must be deflected around the sides. When F > 1, flow can ascend up and over an obstacle. The height at which F = 1 was calculated using RUC data. The average terrain height of the obstacle compared to that of the terrain on the windward side of the obstacle ranges from 1700 to 2000 m. Above 700 hPa, this obstacle height of ~2 km satisfies the condition of F > 1, so flow up and over the obstacle can be expected, possibly resulting in the cloud feature seen in Fig. 8. If it were produced by a vertically propagating gravity wave (Durran 1990), then the Scorer parameter should decrease with height to allow for trapping. The Scorer parameter decreased with height above 700 hPa and leveled off above 650 hPa (not shown).

The initiation mechanism for this case was most likely an elevated convergent layer that was collocated with an area of enhanced midtropospheric moisture. The increased moisture was likely due to broad mesoscale ascent prior to initiation and was perhaps affected by an orographic mountain wave upstream of the CI location. There was a south-southwest–north-northeast-oriented line of enhanced cloudiness that could have been the result of enhanced moisture at the top of an elevated mixed layer that was then advected eastward into Oklahoma throughout the night. This area of enhanced moisture, combined with broad ascent to help saturate the surrounding environment, and possible elevated convergence could have acted to initiate convection in Oklahoma.

c. 15 July 2012

Convection initiated at 0930 UTC 15 July 2012 in the northern Texas Panhandle as a northeast–southwest-oriented convective system (Fig. 12a). The convective system propagated to the southwest and decayed but did not completely dissipate after sunrise. This system was not associated with any severe weather reports. There was an area of enhanced radar reflectivity factor ~30 km southwest of the Amarillo, Texas, WSR-88D in the 3.5° elevation scan (~1.9 km AGL or ~720 hPa; Fig. 12b) at 0910 UTC, which may indicate an area of elevated convergence (based on the radar fine line and absence of surface convergence), even though the reflectivity was enhanced only slightly. The concentric ring of higher dBZ surrounding KAMA likely represents Bragg scattering from the top of the residual layer (Heinselman et al. 2009) or was related to clouds (seen on IR satellite imagery, not shown).

Fig. 12.
Fig. 12.

Amarillo (KAMA) WSR-88D radar reflectivity factor (dBZ) valid at (a) 0954 UTC at 0.5° elevation and (b) 0910 UTC at 3.5° elevation on 15 Jul 2012. The red dashed ellipse in (b) denotes the area of locally enhanced radar reflectivity factor, possibly denoting the ECL.

Citation: Monthly Weather Review 146, 9; 10.1175/MWR-D-18-0040.1

The upper-level synoptic pattern was characterized by a high-amplitude, positively tilted trough located just east of the initiation location at 300, 500, and 700 hPa (Figs. 13a,b). There was a criterion-2 LLJ (Table 1) during the night, but the maximum wind speeds were in northern Kansas and southern Nebraska (reflected in the 850-hPa winds in Fig. 13c). The initiation location was located at the entrance region of the LLJ (a region of horizontal divergence), so based on this weak LLJ forcing, the LLJ did not appear to have played any role in CI in the northern Texas Panhandle on this day as the jet was too far north. The surface winds were from the south-southeast at 10–15 kt (5–8 m s−1) over much of the region, and there was no strong moisture or temperature gradient near the initiation location (Fig. 13d).

Fig. 13.
Fig. 13.

RAP analysis valid at 0600 UTC 15 Jul 2012 for (a) 300-hPa winds (kt) and geopotential height (black contours; m), (b) 700-hPa winds (kt) and geopotential height (black contours; m), and (c) 850-hPa winds (kt) and geopotential height (black contours; m). (d) Surface observations valid at 0607 UTC. The white stars denote the CI locations.

Citation: Monthly Weather Review 146, 9; 10.1175/MWR-D-18-0040.1

There was warm advection at 700 hPa, but the strongest warm advection was farther north in Nebraska and western Kansas, where there was no CI (Fig. 14a). There was weak cyclonic DVA between 500 and 850 hPa (Fig. 14b). Even though there was anticyclonic vorticity advection in the lower to midtroposphere, there was also anticyclonic vorticity advection aloft because of the horizontal wind shear, resulting in negligible differential vorticity advection. Based on the weak QG forcing functions and divergence of the Q vectors (Fig. 14c), it is unlikely that QG ascent played any significant role in CI.

Fig. 14.
Fig. 14.

As in Fig. 4, but for 0700 UTC 15 Jul 2012.

Citation: Monthly Weather Review 146, 9; 10.1175/MWR-D-18-0040.1

Based on the 0000 and 1200 UTC KAMA soundings, as well as RAP soundings valid during the night, the LCLs of the most unstable parcels were located between 700 and 600 hPa (Fig. 15). An examination of proximity RAP soundings indicate that the area of elevated moisture was relatively narrow and near 700 hPa (Figs. 15e–i). There was an elevated convergent layer over the Texas Panhandle (Fig. 16). This elevated convergent layer was associated with the abutment of westerly flow in the midtroposphere, which was likely associated with a short-wave trough over southeastern Colorado, impinging upon northeasterly flow associated with a high-amplitude, positively tilted trough east of the initiation location (Fig. 13b). There was an area of elevated moisture near 700 hPa between 103° and 101°W longitude as well as in the 1200 UTC KAMA sounding (Fig. 15b).

Fig. 15.
Fig. 15.

Observed soundings at KAMA valid at (a) 0000 and (b) 1200 UTC. (c) Map indicating the locations of (d)–(i) the RAP model soundings at 0600 UTC. The color of the box surrounding the soundings corresponds to the color of the dot in (c).

Citation: Monthly Weather Review 146, 9; 10.1175/MWR-D-18-0040.1

Fig. 16.
Fig. 16.

(a) As in Fig. 5a, but for 0600 UTC 15 Jul 2012. The red line denotes the region of elevated horizontal convergence (based on the downstream decrease in the horizontal wind speed and the downstream reversal in the wind direction), the cross-sectional location is denoted in the inset, and the black dashed line denotes the CI longitude. (b) As in Fig. 5c, but for 0900 UTC 15 Jul 2012.

Citation: Monthly Weather Review 146, 9; 10.1175/MWR-D-18-0040.1

An anticyclonic wave-breaking event (Thorncroft et al. 1993; Bosart et al. 2017) occurred between 11 and 13 July. This event resulted in a cyclonic PV streamer that extended from the Great Lakes to the Texas Panhandle during the time of CI (see arrows in Figs. 17b–d). Cyclonic PV anomalies are associated with a cyclonic flow perturbation, so it may be possible that this PV anomaly altered the flow throughout the troposphere and may have aided CI through the formation of the ECL. The cyclonic PV anomaly was centered at 500 hPa above the initiation location, and its horizontal width was approximately 250 km. Assuming a Brunt–Väisälä frequency range of 0.0125 to 0.0175 s−1, the Rossby penetration depth is approximately 1.2 to 1.7 km. At 500 hPa, this would allow perturbations from the anomaly to be felt from ~700 to ~350 hPa. This anomaly retrograded to the west in the presence of northwesterly wind shear of ~20 m s−1 (not shown), which could result in ascent on the southwestern edge of the PV anomaly.

Fig. 17.
Fig. 17.

GFS analysis of potential temperature on the dynamic tropopause [2 potential vorticity unit (PVU) surface] valid at (a) 1200 UTC 9 Jul, (b) 1200 UTC 11 Jul, (c) 1200 UTC 13 Jul, and (d) 1200 UTC 15 Jul 2012. The black arrows in (b)–(d) denote the PV streamer that broke off during the anticyclonic wave-breaking event. Note that 1 PVU = 10−6 K kg−1 m2 s−1.

Citation: Monthly Weather Review 146, 9; 10.1175/MWR-D-18-0040.1

The most likely initiation mechanism for this case was elevated ascent associated with an elevated convergent layer in the midtroposphere (~700 hPa) that was collocated with a narrow and shallow layer of moisture. This ECL was likely the result of the abutment of northeasterly winds from the positively tilted trough east of the initiation location to the westerly winds from a separate trough over southeast Colorado. The QG forcing (temperature advection and DVA) was weak near the initiation location, the LLJ was not a factor in CI (since it was well to the north), and there was no evidence of any gravity waves that could have aided in CI.

d. 9 July 2005

An areal system (a convective system that is neither linear nor single cell; RB17) that was initiated in southwestern Nebraska at 0703 UTC grew upscale into a small MCS by 1000 UTC (Fig. 18) and dissipated by 1900 UTC. This system produced severe weather [1 in. (2.54 cm) hail at 0958 UTC ~120 km southwest of Grand Island, Nebraska] and was associated with a criterion-3 LLJ (reflected at 850 hPa in Fig. 19c) with the maximum wind speed in north-central Kansas. Prior to this initiation, a CI event occurred at ~0400 UTC nearly 130 km to the west (not shown). This cell dissipated by 0530 UTC and was not significant (maximum radar reflectivity factor of 45 dBZ only briefly).

Fig. 18.
Fig. 18.

Hastings (KUEX) WSR-88D 0.5° elevation radar reflectivity factor (dBZ) valid at 1001 UTC 9 Jul 2005.

Citation: Monthly Weather Review 146, 9; 10.1175/MWR-D-18-0040.1

Fig. 19.
Fig. 19.

RUC analysis valid at 0500 UTC 9 Jul 2005 for (a) 300-hPa winds (kt) and geopotential height (black contours; m), (b) 700-hPa winds (kt) and geopotential height (black contours; m; the ridge axis is denoted by the zigzag line), and (c) 850-hPa winds (kt) and geopotential height (black contours; m). (d) Surface observations valid at 0343 UTC.

Citation: Monthly Weather Review 146, 9; 10.1175/MWR-D-18-0040.1

There was a ridge at 300 hPa west of the initiation location, but the winds were weak (<10 m s−1) everywhere over the plains (Figs. 19a,b). The ridge axis at 700 hPa was over the initiation location (Fig. 19b). At 850 hPa, there was a wind maximum southwest of the initiation location (a location shown to have convergence because of the downstream decrease in wind speed). There was a weak surface boundary in northwest Nebraska, but it was >250 km to the west-northwest of the initiation location (see dashed line and star in Fig. 19d).

The Hastings, Nebraska, National Weather Service Forecast Office (GID) noted in the 2030 UTC area forecast discussion on 8 July 2005 that “thunderstorms fired last night across south-central Nebraska as the low-level jet increased with weak warm air advection/theta-e advection at H85. Appears a similar scenario is in place again tonight. The low-level jet again increases this evening with the nose of it pointing into central Nebraska during the overnight hours. Weak warm air advection at H85 and increasing isentropic lift on the 315-K surface may provide enough support to try and generate a few late-night storms. The cap may be a little bit more of a player tonight.” An analysis of the 315-K surface (not shown) indicates possible isentropic ascent near the initiation location between 800 and 750 hPa at 0600 UTC, but the most likely reasoning for this ascent is warm advection by the low-level jet.

Wavelike disturbances appeared in WSR-88Ds data as alternating bands of positive and negative radial velocities (e.g., Koch et al. 2008). Based on the radial velocity data from the Goodland, Kansas (KGLD), radar, there may have been low-level gravity waves3 near the initiation location (see white arrows in Fig. 20c). There are parallel radar fine lines in the North Platte (KLNX) WSR-88D data (Fig. 20a) and are likely a separate set of gravity waves as their orientation is different; however, it is hard to discern the alternating bands of positive and negative radial velocities owing to their orientation with respect to the radar, which was approximately along the radar beam. Finally, a separate set of parallel bands appears in the Hastings (KUEX) Doppler radar velocity data (see white arrows in Fig. 20b) at 0804 UTC. The orientation of these waves is similar to that of the waves seen in KGLD, so it is possible that these waves originated from the same feature. The movement of the radar echoes from KGLD indicate that the speed of the possible gravity waves was ~5 m s−1 from ~320° (from the north-northwest) while the motion of the radar echoes from KLNX indicate that those waves were propagating at ~15 m s−1 from 209° (from the south-southwest). The wavelength of all sets of waves (based on radar data) is ~10 km.

Fig. 20.
Fig. 20.

(a) North Platte (KLNX) WSR-88D 0.5° radial velocity valid at 0657 UTC 9 Jul 2005, (b) Hastings (KUEX) WSR-88D 2.4° radial velocity valid at 0804 UTC 9 Jul 2005, (c) Goodland (KGLD) WSR-88D 0.5° radial velocity valid at 0632 UTC 9 Jul 2005, and (d) the radial velocity data in (a)–(c) projected onto a map of Kansas and Nebraska. The white arrows in (a)–(c) denote the location of the gravity waves, the black dashed lines in (c) denote the locations of the elevated boundary (straight dashed line) and the outflow boundary (curved dashed line), and the black arrow in (a) and (c) denotes the direction of propagation of the gravity waves as determined by radar-observed motions. The thick dashed white lines in (d) represent the orientation and hypothetical extension of the gravity waves, the thin red-and-white dashed line denotes the orientation of the waves in KUEX, and the white pentagon denotes the CI location.

Citation: Monthly Weather Review 146, 9; 10.1175/MWR-D-18-0040.1

Since the speed of the gravity wave near KGLD was 5 m s−1 and the wavelength was about 10 km, the period between waves was ~33 min. There are fluctuations of pressure, temperature, wind speed, and wind direction in the KGLD ASOS station pressure with a period of ~45 min (Fig. 21), slightly longer than that estimated by radar. It is possible that the wave motion was not perpendicular to their orientation, which could account for the difference. Using surface pressure and wind information, one can use the impedance formula (Gossard and Munk 1954) to obtain information about gravity wave propagation direction. Using the wind speed and directions at the pressure perturbation maximum and minimum, a wave propagation direction of 330° was calculated (see inset in Fig. 21a), which corresponds well with that derived from radar data. The Grand Island ASOS data (Fig. 21b) do not suggest the presence of gravity waves near 0800 UTC even though the radar reflectivity factor (not shown) and radial velocity from KUEX show parallel bands suggestive of gravity waves. The banded feature in KUEX does not appear at elevations below 2.4° (~2 km AGL). It is possible that the amplitude of the pressure oscillation is too weak to be resolved by the ASOS pressure sensor; it is also possible that the banded feature is not a gravity wave but instead another feature. However, since the bands are collocated with the CI location, they likely played a role in CI.

Fig. 21.
Fig. 21.

The 1-min ASOS data valid on 9 Jul 2005 for (a) Goodland, KS (KGLD), and (b) Grand Island, NE (KGRI). The blue curve is the wind speed (m s−1), the brown line is the wind direction (°), the wind barbs are as convection (half barb is 5 kt, and full barb is 10 kt), the red line is the temperature (°C), the green line is the dewpoint (°C), and the black line is the station pressure (hPa). Vertical dashed lines denote the crest of the gravity wave. The inset in (a) depicts the graphical representation of the gravity wave propagation direction similar to that in Gossard and Munk (1954). The data have been smoothed using a seven-point running average.

Citation: Monthly Weather Review 146, 9; 10.1175/MWR-D-18-0040.1

The origin of the gravity wave near KGLD is unknown, but it may be related to earlier convection near the initiation location (unlikely since this CI was at 0400 UTC, was weak, and too far west) or an elevated boundary (as the orientation of the gravity waves is parallel to that of the elevated boundary; see straight dashed black line in Fig. 20c). The boundary evident as a fine line in the KGLD radar data is not discernible from the surface data, so the boundary is assumed to be elevated. There is also evidence of an outflow boundary (see curved dashed black line in Fig. 20c). This outflow boundary may have a role in triggering the convection or the gravity wave(s), but the convection was initiated ahead of this boundary, and the gravity wave’s orientation was not parallel to that boundary.

The KLBF sounding had an inversion below 800 hPa (Fig. 22a), a condition supportive of wave reflection (Klemp and Lilly 1975). The Scorer parameter calculated from the 1200 UTC KLBF sounding is noisy owing to the weak gravity wave–relative wind speeds, but there are three notable regions where the Scorer parameter decreases with height: at 950, 900, and 800 hPa (Fig. 22b). However, while this 1200 UTC sounding is supportive of gravity waves, it is unlikely to support convection. An examination of RUC soundings around the CI location indicated a shallow layer of moisture centered at 800 hPa (Figs. 15d–i). Parcels lifted from this level had anywhere from 500 to 2500 J kg−1 of MUCAPE but would also have −137 to −455 J kg−1 of MUCIN. This would require ascent between 16 and 30 m s−1 to overcome the MUCIN. This is unlikely to occur in these weakly forced environments. In the actual environment, there could have been ascent helping to saturate and cool the environment at that level to weaken the cap. It is important to note that the RUC did not initiate convection during this night, so it is likely that RUC model soundings do not represent the actual environment.

Fig. 22.
Fig. 22.

(a) Observed sounding and (b) calculated Scorer parameter at KLBF at 1200 UTC 9 Jul 2005. (c) Map indicating the locations of (d)–(i) the RUC model soundings at 0600 UTC. The color of the box surrounding the soundings corresponds to the color of the dot in (c).

Citation: Monthly Weather Review 146, 9; 10.1175/MWR-D-18-0040.1

It is likely that the ascent needed to cause CI was above 800 hPa (based on the stable boundary layer), and rising motion associated with the banded feature seen in the KUEX radar data was a likely candidate mechanism to force parcels to their LFC, but temperature and moisture advection by the LLJ may have aided the initiation as the ascent needed to force a parcel to its LFC was ~16–30 m s−1 based on MUCIN near the CI location (Figs. 15d–i). One hypothesis for the initiation mechanism is that the intersection of the two gravity waves (Fig. 20c) could provide enough ascent to force a parcel to its LFC. At the time of initiation, the intersection of the gravity waves (assuming that they were able to extend that far) would have been near the initiation location (see thick dashed white lines in Fig. 20d). More observations are needed to see if this could be the case (especially a direct measurement of vertical velocity). The thin red-and-white dashed line in Fig. 20d represents the parallel bands from the KUEX Doppler velocity data at 0804 UTC. If the gravity waves evident in KGLD velocity continued to propagate until 0804 UTC, then their location would be near that represented by the thin red-and-white dashed line. This case represents an especially challenging case to forecast based on the relative minimum conditions supportive of CI. RUC forecast soundings do not seem to support an obvious CI event (and did not initiate convection) and the LLJ was weak compared to that in other events.

5. Summary and conclusions

Four case studies of NB CI mode events (RB17) illustrated a wide variety of synoptic-scale environments and possible initiation mechanisms and features that can result in CI. Examples of the different synoptic-scale environments include an upper-level trough upstream of the CI region, an upper-level ridge axis centered over the CI location, and a high-amplitude upper-level trough just downstream from the CI location. Examples of the different initiation mechanisms include that related to the LLJ, QG-aided ascent, elevated ascent associated with elevated convergent layers (associated with an unknown dynamic mechanism), and gravity waves. These different environments and different initiation mechanisms could pose forecasting problems because there does not appear to be an NB CI mode environment common to all cases. However, features are common to the events examined in this study include a midtropospheric moisture maximum and midtropospheric warm advection, though the magnitudes vary from case to case. While these features may not lead directly to CI, they may prime the environment so that another feature (such as the LLJ, elevated convergent layers, or gravity waves) may initiate convection. Conditional instability is present in all of these cases (and during many other nights without CI), suggesting that conditional instability is a necessary, but insufficient, requirement for nocturnal CI.

The 4 July 2014 (case 1) case represents perhaps the most predictable of the four events presented in part because north–south lines of convection are not uncommon during the night. However, the other three events pose more significant forecasting challenges. The 9 July 2005 (case 4) event was perhaps the most challenging as the only mechanism noted by forecasters was warm advection and convergence at the nose of the LLJ, but other features such as gravity waves possibly triggered by an elevated boundary may have aided CI on that day. Data from recent field experiments that had higher spatiotemporal observations than the current observational network for these NB CI mode events such as BAMEX, IHOP_2002, and PECAN can help fill in some of the deficiencies in the current observational network and help identify initiation mechanisms that the current observational network cannot observe. During PECAN, mobile Doppler radar and lidar data can be used to construct low-level VAD profiles to observe the spatiotemporal evolution of the low-level wind profiles. Soundings were launched every three hours, providing high-spatiotemporal-resolution observation of how the upper-level thermodynamic and wind fields evolve throughout the night. Boundary layer profiling instruments such as microwave radiometers, atmospheric emitted radiance interferometer (AERI) systems, and water vapor differential absorption lidars (DIALs) can be used to better observe gravity waves and determine environment characteristics near the waves. Idealized numerical experiments need to be conducted to identify with more certainty the nature of the dynamical forcing possible for nocturnal NB CI.

Acknowledgments

We extend our thanks to Drs. Dave Parsons, Steven Cavallo, and Alan Shapiro from the University of Oklahoma for their assistance and advice; to Kevin Haghi, Zach Wienhoff, and Chris Riedel for useful discussions and programming assistance; and to Stan Trier, Rit Carbone, and one anonymous reviewer whose reviews helped improve the quality of this manuscript significantly. We also thank Dave Ahijevych at NCAR for maintaining the radar composite archive (available at http://www2.mmm.ucar.edu/imagearchive/). This project was funded by NSF Grants AGS-1237404 and AGS-1262048.

REFERENCES

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1

Other features and mechanisms that may be present but likely do not contribute to initiation of NB CI mode events include conditional symmetric instability (CSI; Schultz and Schumacher 1999) and atmospheric diurnal and semidiurnal tides (Dai et al. 1999).

3

Note that low-level gravity waves were not considered to be a surface boundary in the RB17 climatology.

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