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  • View in gallery

    GOES visible satellite imagery at (a) 1501, (b) 1701, and (c) 1901 UTC 27 March 1994 showing the three mesoconvective systems and subcomponent convective lines. Surface frontal analyses are superimposed on the imagery.

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    Surface winds (whole barb = 5 m s−1, half barb = 2.5 m s−1), temperatures (°C), subjectively analyzed isobars (2-mb intervals), and composite radar reflectivity mosaics for 1500, 1700, and 1900 UTC. The mesolows that were directly associated with three MCSs (see text) are also depicted.

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    Barnes objective analyses of rawinsonde data at 1200 UTC depicting upper-level jet and quasigeostrophic dynamics: (a) 200-hPa winds (isotachs > 50 m s−1 at 10 m s−1 intervals) and geopotential height contours (m), (b) as in (a) except at 300 mb, (c) Q-vector divergence (10−15 m kg−1 s−1, solid contours) and convergence (dashed) at 700 mb, and (d) cross section depicting ageostrophic circulations transverse to upper-level polar (J1) and subtropical (J2) jet streaks and isentropes (solid, 3-K intervals) from North Platte, NE, to Cape Kennedy, FL [see (a) for cross-section location].

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    MASS model 18-h forecasts valid at 1800 UTC of (a) winds at 200 hPa (as in Fig. 3, except isotachs > 45 m s−1 shown), (b) winds at 300 hPa (isotachs > 45 m s−1), (c) geostrophic wind (isotachs > 75 m s−1) and geopotential height contours (m) at 300 hPa, and (d) winds and geopotential height (m) at 850 hPa. Isotachs > 65 m s−1 are shaded in (a) and (b). Also depicted are model forecast locations of surface frontal systems, for ease of reference, and cross sections A–A′ and B–B′ shown in later figures. Note distance scale.

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    Eta and MASS Model forecasts of quasigeostrophic forcing valid for 1800 UTC: (a) 18-h Eta Model forecast Q-vector field at the 500-hPa level, and (b) smoothed 18-h MASS model forecast Q-vector field for the 500–300-mb layer. Q-vector convergence regions are shaded [10−15 m kg−1 s−1 in (a) and 10−14 m kg−1 s−1 in (b)]. The 500-mb geopotential height contours (m) are also depicted in (b).

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    Forecast upper-level jet transverse circulations for (a) 1400 UTC from Omaha, NE, to Tallahassee, FL (Fig. 4b), and (b) 1800 UTC. Vertical cross sections derived from MASS model forecast fields show ageostrophic circulation (vectors), isotachs ≥ 40 m s−1, and isentropes (solid, 2-K intervals). Note horizontal vector scale (top right) and horizontal distance scale (bottom). Forecast surface frontal position is depicted by arrow along abscissa. Large “J” depicts polar jet, small “J” denotes mesoscale jetlet.

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    Objectively analyzed surface frontogenesis fields [shaded at 6°C (100 km)−1 (3 h)−1 intervals], isobars, and frontal analyses at (a) 1200, (b) 1500, (c) 1800, and (d) 2100 UTC.

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    As in Fig. 7 except fields are obtained from MASS model forecasts at the same times. Mesolows are given labels to correspond as closely as possible to those in the observations.

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    (a) Surface frontogenesis function field using the geostrophic wind for comparison to Fig. 8c [shaded at 3°C (100 km)−1 (3 h)−1 intervals], (b) as in (a) except using the ageostrophic wind, and (c) quasigeostrophic frontogenesis analysis using the 900–800-mb layer-averaged Q vectors and Q-vector convergence (shaded > 20 × 10−15 m kg−1 s−1) and potential temperature at 850 mb (solid lines, 3-K intervals). All fields are derived from the MASS model forecasts valid for 1800 UTC.

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    MASS model forecasts of frontal features in a plane normal to the surface cold front, extending from southeastern Missouri to southwestern Georgia (Fig. 4b). (a) and (b) Front-relative circulations, regions of upward motion (shaded at 5 μb s−1 intervals), and isentropes of potential temperature valid for 1400 and 1800 UTC, respectively. (c) and (d) Isentropes of equivalent potential temperature (2-K intervals, dotted lines), absolute geostrophic momentum (10 m s−1 intervals, solid lines), and relative humidity > 90% (shaded). The regions wherein the necessary conditions for conditional symmetric instability are met (see text) are shown by the thick curves. The positions of the predicted surface cold front are denoted by the small arrows at the bottom of the top panels.

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    Conditional symmetric instability (CSI) in the NGM model 18-h forecast for 1800 UTC in a vertical cross section from Omaha, NE, to Tallahassee, FL (along line A–A′ in Fig. 4b). Shown are isentropes of equivalent potential temperature (2-K intervals, thin lines), absolute geostrophic momentum (10 m s−1 intervals, thicker lines), and relative humidity > 90% (shaded). The region wherein the necessary conditions for CSI are met is shown by the very thick curve. The position of the predicted surface cold front is at the arrow appearing at the bottom.

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    Radar-observed precipitation patterns in the local outbreak region (see Fig. 2 for isoecho intensity levels), and their relationships to surface frontal systems at (a) 1600 UTC and (b) 1800 UTC. Compare prefrontal convective band locations at 1600 UTC to the CSI and inertial instability regions seen in 1800 UTC MASS display (Fig. 10d), and to the fine-grid MASS model precipitation forecasts (Fig. 13d).

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    Coarse- (24 km) and fine- (12 km) mesh MASS model forecast fields of precipitation rate (shaded, mm h−1), surface winds (m s−1), and surface equivalent potential temperature (θe) fields (2-K intervals, dotted lines) for (a) 1400 UTC (coarse grid), (b) 1400 UTC (fine grid), (c) 1800 UTC (coarse grid), and (d) 1800 UTC (fine grid). Mesoanalyses are based on the θe and wind fields. Outflow boundaries are denoted by smaller pips.

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    Rawinsonde soundings at 1800 UTC from (a) Centreville, AL, and (b) Jackson, MS, plotted on skew T diagrams and (insets) wind hodographs (m s−1). Surface-based parcel lifted adiabatically is shown by the dashed line, and positive area (CAPE) is represented by shaded regions. Wind barbs on right ordinate use these conventions: flag = 25 m s−1, whole barb = 5 m s−1, half barb = 2.5 m s−1.

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    Evolving relationships between gravity wave troughs A, B, . . . , in mean sea level pressure field (light contour lines) and two mesoscale jetlets evident at 500 mb (revealed by isotachs > 45 or 50 m s−1, thick contour lines) at (a) 1200, (b) 1500, (c) 1800, and (d) 2100 UTC from MASS model forecasts. Vertical motion fields at 600 hPa (upward motions > −20 μb s−1 shaded), surface frontal systems and mesolows, surface wind vectors (same notation as in Fig. 13), and location of 1800 UTC cross section in Fig. 18 are also shown.

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    Spatial bandpass analysis of MASS model forecast mean sea level pressure and surface wind fields (Fig. 15), resulting in (a) perturbation pressure (p′) field valid at 1500 UTC, (b) perturbation wind (u*′) field valid at 1500 UTC, (c) p′ field valid at 1800 UTC, and (d) u*′ field valid at 1800 UTC. Solid (dotted) lines are positive (negative) values at 0.5-mb intervals (for perturbation pressure) and 0.5 m s−1 intervals (for perturbation winds). Gravity wave trough locations copied from Fig. 15 are shown by thick line segments, and surface frontal analyses are also shown.

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    Hourly isochrones of gravity wave troughs A, B, C, D, E, F, G, and H (F12 = forecast hour 12, etc.) obtained from coarse-mesh MASS model forecasts (see Fig. 15). Note horizontal scale in wave A panel.

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    Vertical cross section normal to gravity wave fronts from 1800 UTC coarse-mesh MASS model forecast (Fig. 15). Shown are isentropes of potential temperature (2-K intervals, solid lines), upward vertical motions (shaded at intervals of 4 μb day−1, see gray bar at bottom), and critical level for mean gravity wave speed of 38 m s−1 (dark curve). Locations of surface wave troughs (Fig. 16c) are shown in the cross-sectional plane (G, F, C, E).

  • View in gallery

    As in Fig. 15 except based on fine-mesh MASS model forecast and not depicting the position of the jetlets.

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    Bandpass filter applied to digitized surface barograph data. Response function R(f) has a value of 1.0 at a given frequency when the waves of that frequency are passed without loss of amplitude. The “cutoff frequencies” [at which R(f) = 0.5] are highlighted in gray.

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    Time series of digitized barograph data (mb, light lines) and bandpass-filtered pressure perturbations (mb, dark lines) for Muscle Shoals (MSL) and Huntsville (HSV), AL, covering the period 0500 UTC 27 March–1115 UTC 28 March 1994. Similar phase-shifted pressure fluctuations are evident in the pressure traces.

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    Objectively analyzed pressure perturbation fields (contour interval Δp = 0.2 mb, with p′ < 0 dotted) for (a) 1200, (b) 1500, (c) 1800, and (d) 2100 UTC 27 March 1994.

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Mesoscale Dynamics in the Palm Sunday Tornado Outbreak

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  • 1 Department of Marine, Earth, and Atmospheric Sciences, North Carolina State University, Raleigh, North Carolina
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Abstract

Radar and satellite imagery suggest that strong mesoscale forcing occurred in the Palm Sunday tornado outbreak on 27 March 1994. Parallel lines of severe thunderstorms within each of three mesoscale convective systems developed just ahead of a cold front in Mississippi and Alabama on this date. Analyses of routine meteorological observations, barograph data, and forecasts from the Eta and NGM models and a mesoscale research model (MASS) are used to examine the relative roles of large-scale dynamics and mesoscale processes in triggering and organizing the mesoscale convection.

Quasigeostrophic forcing was absent in the outbreak region. Likewise, a thermally direct circulation system transverse to the upper-level jet that was present to the northwest of the outbreak region was decoupled from the strong low-level ascent occurring in northern Alabama and Mississippi at the time of the outbreak. Strong ageostrophic frontogenesis in the presence of conditional symmetric instability (CSI) was the chief cause for the intense low-level ascent along and behind the front, consistent with the line of severe storms that developed explosively along the front and an observed postfrontal precipitation band. However, the strongest supercells developed in segmented lines 100–200 km ahead of and parallel to the frontal boundary in an atmosphere that the MASS model indicates was inertially unstable due to a mesoscale midlevel jetlet. Analysis suggests that these storms developed in a manner consistent with the predictions of asymmetric inertial instability theory in the presence of convective instability.

Several mesolows were observed to have traveled along the frontal boundary and to have played a key role in focusing the frontogenesis. Similar frontal mesolows were simulated by the MASS model. Strong low-level ascent in the presence of conditional instability helped to deepen the mesolows, but they were strongly modulated by a train of gravity waves propagating on the cold side of the front. A combination of ducting and wave-CISK (conditional instability of the second kind) processes maintained the waves, which remained coupled to the jetlets as they propagated from intense convection in northeastern Texas. A time-to-space conversion objective analysis of bandpass-filtered barograph data reveals that similar waves emanated from this same region.

The lifting patterns produced by the complex interactions between the gravity waves, CSI, asymmetric inertial instability, and frontogenesis satisfactorily explains the development, configuration, spacing, and relative movement of the severe mesoconvective systems on Palm Sunday. All of these mesoscale phenomena were coupled to or strongly influenced by the jetlets, which were produced by strong convection at an earlier time within the region of quasigeostrophic forcing far removed from the tornado outbreak.

Corresponding author address: Dr. Steven Koch, Department of Marine, Earth, and Atmospheric Sciences, North Carolina State University, Campus Box 8208, Raleigh, NC 27695-8208.

Email: Steve_Koch@ncsu.edu

Abstract

Radar and satellite imagery suggest that strong mesoscale forcing occurred in the Palm Sunday tornado outbreak on 27 March 1994. Parallel lines of severe thunderstorms within each of three mesoscale convective systems developed just ahead of a cold front in Mississippi and Alabama on this date. Analyses of routine meteorological observations, barograph data, and forecasts from the Eta and NGM models and a mesoscale research model (MASS) are used to examine the relative roles of large-scale dynamics and mesoscale processes in triggering and organizing the mesoscale convection.

Quasigeostrophic forcing was absent in the outbreak region. Likewise, a thermally direct circulation system transverse to the upper-level jet that was present to the northwest of the outbreak region was decoupled from the strong low-level ascent occurring in northern Alabama and Mississippi at the time of the outbreak. Strong ageostrophic frontogenesis in the presence of conditional symmetric instability (CSI) was the chief cause for the intense low-level ascent along and behind the front, consistent with the line of severe storms that developed explosively along the front and an observed postfrontal precipitation band. However, the strongest supercells developed in segmented lines 100–200 km ahead of and parallel to the frontal boundary in an atmosphere that the MASS model indicates was inertially unstable due to a mesoscale midlevel jetlet. Analysis suggests that these storms developed in a manner consistent with the predictions of asymmetric inertial instability theory in the presence of convective instability.

Several mesolows were observed to have traveled along the frontal boundary and to have played a key role in focusing the frontogenesis. Similar frontal mesolows were simulated by the MASS model. Strong low-level ascent in the presence of conditional instability helped to deepen the mesolows, but they were strongly modulated by a train of gravity waves propagating on the cold side of the front. A combination of ducting and wave-CISK (conditional instability of the second kind) processes maintained the waves, which remained coupled to the jetlets as they propagated from intense convection in northeastern Texas. A time-to-space conversion objective analysis of bandpass-filtered barograph data reveals that similar waves emanated from this same region.

The lifting patterns produced by the complex interactions between the gravity waves, CSI, asymmetric inertial instability, and frontogenesis satisfactorily explains the development, configuration, spacing, and relative movement of the severe mesoconvective systems on Palm Sunday. All of these mesoscale phenomena were coupled to or strongly influenced by the jetlets, which were produced by strong convection at an earlier time within the region of quasigeostrophic forcing far removed from the tornado outbreak.

Corresponding author address: Dr. Steven Koch, Department of Marine, Earth, and Atmospheric Sciences, North Carolina State University, Campus Box 8208, Raleigh, NC 27695-8208.

Email: Steve_Koch@ncsu.edu

1. Introduction

This paper presents a mesoscale analysis of the dynamical environment of the Palm Sunday tornado outbreak, which occurred on 27 March 1994 across the southeastern United States. Approximately 30 tornadoes were spawned by 14 supercell storms that tracked over a narrow region from north-central Alabama to western North Carolina along and just ahead of a slowly moving cold front. These storms exhibited a high degree of mesoscale organization, in that they were organized into front-parallel lines within three large mesoscale convective systems (MCSs). Rapid convective destabilization and extreme values of helicity alerted forecasters to the impending outbreak. Convective available potential energy (CAPE) values across the region increased from <1300 J kg−1 at 1200 UTC to nearly 3000 J kg−1 only 6 h later. Storm-relative helicity is proportional to the strength of the low-level inflow and the vorticity aligned with this flow (the “streamwise vorticity”). Storm-relative helicity values H > 150 m2 s2 have been suggested as a minimum threshold for supercell storms, but H > 450 m2 s2 was calculated from the 1200 UTC sounding at Centreville, Alabama. Such a value is generally associated with violent tornadoes (Davies and Johns 1993).

Despite these obviously important factors, synoptic-scale features associated with “classical” tornado outbreaks (Miller 1955) were clearly absent. There was no well-defined short wave, nor even a surface low pressure center. In fact, the storms developed on the anticyclonic periphery of the polar jet, which was well to the north. The absence of a classical synoptic-scale pattern for a major tornado outbreak and the suggestion of mesoscale organization provide the motivation for this study. In particular, we examine the role of mesoscale processes in altering and organizing the environment in which these storms developed. Mesoscale precipitation bands displaying separations of 25–250 km are not accounted for in the Norwegian cyclone model; rather, they are generally attributable to one or more of the following mesoscale processes in the absence of topographic or land surface forcing: frontal circulations, frontogenetical forcing, gravity waves, and conditional symmetric instability (CSI).

This investigation complements other mesoscale studies in this special issue on the Palm Sunday outbreak. Kaplan et al. (1998) and Hamilton et al. (1998) discuss the role of a simulated mesoscale jet streak in the outbreak, which we show to be related to the presence of frontal mesolows and gravity waves. Langmaid and Riordan (1998) analyze the processes that contributed to frontogenesis along a shallow frontal boundary just ahead of the synoptic cold front in northern Alabama. We examine the interactions between frontogenesis and CSI in the vicinity of these frontal boundaries.

Section 2 presents an overview of the mesoscale organization of convection during the Palm Sunday event. Section 3 examines the quasigeostrophic and jet streak dynamics. Sections 4, 5, and 6 discuss the roles of frontogenetical forcing, CSI/inertial instability, and the interactions between convection and gravity waves, respectively, in the outbreak. Conventional observations, digitally processed barograph data, the Eta Model and Nested Grid Model (NGM) of the National Centers for Environmental Prediction (formerly the National Meteorological Center) and a research mesoscale model [the Mesoscale Atmospheric Simulation System (MASS)] are used in those diagnostic analyses. An earlier version of MASS is documented in Kaplan et al. (1982). Recent updates to the model can be found in Manobianco et al. (1994). The appendix describes the model configuration used here.

2. Convective systems in the Palm Sunday event

Analysis of the Geostationary Operational Environmental Satellite (GOES) imagery and composite radar reflectivity mosaics reveals that three MCSs formed between eastern Mississippi and eastern Louisiana from 1430 to 1730 UTC1 and rapidly intensified as they propagated across northern Alabama and Georgia, before finally dissipating in North Carolina. GOES visible imagery with superimposed surface frontal analyses at 1500, 1700, and 1900 UTC are shown in Fig. 1, and radar mosaics are presented at these same times in Fig. 2. Tornadic thunderstorms composed mesoscale convective system MCS A, which developed explosively from two parallel lines of convection ahead of the surface cold front (SCF). The closest line to the SCF is apparent in both the satellite and radar displays at 1500 UTC along the Mississippi–Alabama border. This line was less than 100 km from the front. The second line within MCS A was most apparent at 1600 UTC (as shown later), although the storm in extreme eastern Alabama in the 1700 UTC radar display (the Cherokee tornadic supercell) is contained within this line. MCS A was by far the most prolific tornado producer of the three systems. MCS B appeared shortly after 1700 UTC in nearly the same original location as MCS A, though it developed right along the front in central Mississippi. While its component storm cells were more numerous than those in MCS A, MCS B was smaller and produced only hail and damaging winds. MCS C was a squall line that developed explosively shortly after 1730 UTC in close proximity to the SCF along the Mississippi–Louisiana state border, and produced numerous reports of damaging wind and large hail. Thus, MCS A was the only system that was not generated precisely along the surface front. The three MCSs are most easily distinguishable in the visible satellite imagery at 1900 UTC (Fig. 1c).

Light precipitation occurred throughout the period over Tennessee and Kentucky behind the SCF and also northward of a shallow stationary frontal boundary (SFB) draped across extreme northern Alabama, Georgia, and South Carolina (Fig. 2). Langmaid and Riordan (1998) show how evaporational cooling from this stratiform precipitation and other processes created and maintained the SFB. The analyses also reveal that three distinct mesolows traveled along the SCF and toward the SFB. Notice that the three MCSs were each associated with one of these mesolow pressure systems (LA, LB, LC). The strongest temperature gradients, vorticity, and moisture convergence were along both the shallow thermal boundary and the synoptic-scale surface front (Langmaid and Riordan 1998).

It is important to observe that the prefrontal tornadic storms were arranged in lines with a similar orientation (240°–60°), which was parallel to the SCF. These supercell storms were long lived even by supercell standards, with one of them persisting for 8 h. The longevity of these storms and the fact that they remained fixed to the lines is highly indicative of the importance of mesoscale forcing. It is also interesting that those storms that formed within lines right along the SCF (MCSs B and C) were generally shorter lived and produced fewer tornadoes. These observations suggest that an understanding of the mesoscale dynamics is essential to explain the behavior of these storms.

3. Upper-level dynamics

An examination of the possible importance of jet streak dynamics and quasigeostrophic forcing at upper levels for the outbreak is presented here. The dynamical importance of upper- and lower-tropospheric jets in severe thunderstorm outbreaks has been elaborated upon by many investigators (e.g., Miller 1955; Beebe and Bates 1955; Palmén and Newton 1969; Weisman and Klemp 1982). Jets are important for their ability to destabilize the atmosphere through vertically differentiable advection processes, and to provide the vertical wind shear necessary for strong, long-lived storms.

Parcels passing through a straight jet experience a clockwise rotation in the ageostrophic velocity. This provides for the familiar four-cell pattern of upper-level divergence and vertical motion symmetric with respect to the jet core—that is, rising air in the right entrance and left exit regions of the jet and descending air in the left entrance and right exit regions. These ageostrophic wind patterns are associated with a thermally indirect transverse circulation within the exit region and a thermally direct circulation within the entrance region of a jet in quasigeostrophic balance (Uccellini and Johnson 1979). The presence of significant curvature and/or thermal advection alters this symmetric vertical motion pattern (Keyser and Shapiro 1986), but the present case is adequately described by the simple straight jet model. In fact, essentially unidirectional southwesterly flow existed throughout the troposphere above 850 mb ahead of the cold front in the outbreak region, whereas in the classical tornado outbreak a large crossing angle between the upper- and lower-level jet streaks typically occurs (Miller 1955; Weisman and Klemp 1982). Severe weather can still develop in such unidirectional flows when relatively strong upper-level divergence associated with the jet streak occurs in the presence of low-level moisture and instability (e.g., McNulty 1978; Koch and Dorian 1988). We now examine whether upper-level divergence was present over the outbreak region stretching from Louisiana to northern Georgia.

An objective analysis (Koch et al. 1983) of the 1200 UTC 27 March rawinsonde data portrays the right entrance region of the polar jet (PJ) (at 300 mb) and the left exit region of the subtropical jet (STJ) (at 200 mb) to be juxtaposed over northwestern Arkansas (Figs. 3a,b). A vertical cross section transverse to the jet axes (Fig. 3d) shows that a thermally direct ageostrophic circulation system existed at this time below the level of the jet, with a circulation system of the opposite sense above this level. This dual circulation is consistent with the conceptual model of PJ–STJ coupling by Uccellini and Kocin (1987), which has been adapted by Kaplan et al. (1998) to the Palm Sunday event. Yet the question remains—is the strong low-level ascent in northern Alabama (at the “×” in the cross section) actually forced by the jet circulation system and/or quasigeostrophic (QG) dynamics?

We evaluated QG forcing of vertical motion using the Q-vector method (Hoskins et al. 1978; Hoskins and Pedder 1980), assuming that convergence (divergence) of the Q vectors implies QG ascent (descent). This approximation works best in the middle troposphere. Quasigeostrophic forcing (Q-vector convergence) at 1200 UTC (Fig. 3c) is found in close proximity to the location of the PJ–STJ juxtapositioning (over northeastern Texas), whereas the strongest ascent is over northern Alabama and is most pronounced at low levels. This suggests that much of the ascent over the outbreak region may actually have been forced by non-QG processes acting in the lower troposphere. This hypothesis is explored further immediately below and in section 4.

Quasigeostrophic forcing of vertical motion at the time of the outbreak (1800 UTC) was evaluated from an analysis of the Q-vector fields and the transverse circulations predicted by the models. The 18-h forecast upper-level wind and 500–300-mb layer-averaged Q-vector fields from the MASS Model are depicted in Figs. 4 and 5b, respectively; the Eta Model 500-mb Q-vector fields are shown in Fig. 5a. The MASS model geopotential and temperature fields were smoothed for this purpose, following the recommendations of Barnes et al. (1996) that this is a necessary prerequisite for examining quasigeostrophic dynamics in high-resolution mesoscale models containing energetic gravity wave and/or convective phenomena (both of these were problematic in the present case). It is apparent that QG forcing occurs only well away from the outbreak region, just as it had 6 h earlier. Comparison of these Q convergence regions with the forecast location of the upper-level jet right entrance region (over southern Missouri) reveals that the strong QG ascent is indeed occurring in close proximity to the right entrance region. Also note that there is no longer any clear separation between the positions of the PJ and the STJ—that is, the jet streak“coupling” emphasized by Kaplan et al. (1998) to have existed at 1200 UTC was inoperative by the time of the outbreak at 1800 UTC.

Vertical cross sections transverse to the jet entrance region are depicted in Fig. 6 at 1400 and 1800 UTC from the coarse grid MASS model forecasts. While a pronounced thermally direct circulation system is apparent above 600 mb, and is attributable to the upper-level jet dynamics, a separate thermally direct cell appears at low levels in association with the cold front. The separation of the upper-level jet ascent region from the frontal updraft increases from 160 km at 1400 UTC to nearly 400 km by 1800 UTC. Three factors explain this increased separation: the surface front (and its attendant updraft jet) propagated southeastward at a speed of 6.9 m s−1, the upper-level jet retrogressed in the cross-sectional plane at a speed of −4.6 m s−1, and the jet ascent region shifted closer to the jet core with time at a speed of −10.2 m s−1. Thus, we conclude that the frontal circulation cell, which is clearly the more pertinent circulation with regard to the tornado outbreak, was not driven by upper-level quasigeostrophic forcing associated with the jet streak.

In summary, there is no evidence from any of the analyses indicating either that synoptic-scale jet streak interaction processes or quasigeostrophic forcing was the cause for the strong ascent seen in the lower troposphere over northern Alabama. As discussed below, the evidence indicates that this ascent was forced by a nonquasigeostrophic frontogenetical circulation in the presence of conditional symmetric instability at low levels. Of course, air that was forced rapidly upward at low levels would continue to ascend in a slantwise fashion over the frontal surface to very high levels as the parcels came under the influence of the upper-level jet circulation system. The entire coupled jet-frontal circulation system can be described by the anafront model, which is a front with rearward-sloping slantwise ascent (Keyser and Shapiro 1986).

One additional feature in the 1800 UTC cross section (Fig. 6b) demands explanation. A third ascent region, a rather deep circulation system, is evident at about 100 km ahead of the surface cold front. Kaplan et al. (1998) discuss how a geostrophic adjustment of the jet streak to convective latent heating far upstream of the outbreak region forced the development of a mesoscale jet streak (hereafter “jetlet”). That study and the idealized model study by Hamilton et al. (1998) suggest that a thermally direct circulation favorable for lifting of the unstable air appeared on the right flank of the jetlet. The ascent region ahead of the SCF in Fig. 6b is related to the presence of this jetlet, which has by this time propagated to the outbreak region in northern Alabama (Fig. 4b). Processes responsible for the generation of the strong ascent on the right flank of this jetlet are discussed by Hamilton et al. (1998). However, we present evidence below for the additional importance of inertial instability in forcing this ascent, which may have been of greater importance in the outbreak than the frontal circulation itself. Having eliminated upper-level QG forcing as a possibility in the tornado outbreak, we now turn our attention to examining the roles played by frontogenesis and other mesoscale processes.

4. Frontogenetical forcing

Frontogenesis is defined as an increase in the horizontal gradient of potential temperature following the motion of an air parcel through a frontal zone. Semigeostrophic theory provides a dynamical basis for understanding how frontogenesis is forced by geostrophic shearing and stretching deformation (Hoskins and Bretherton 1972). However, this theory contains major shortcomings, including its inability to account for the existence of a mesoscale frontal circulation and an intense updraft jet often observed in the lower troposphere at the leading edge of cold fronts (Blumen 1980). Idealized models of frontogenesis have shown that an intense cross-frontal circulation containing a frontal updraft jet can be produced when various physical processes not contained in this theory are included in the model; of particular mention are condensational heating (e.g., Benard et al. 1992) and a cross-frontal gradient of sensible heating (e.g., Koch et al. 1995).

According to Langmaid and Riordan (1998), condensational heating was not a major contributor to frontogenesis in the Palm Sunday case prior to the development of the severe storms in the outbreak region. On the other hand, they discuss the frontogenetical importance of differential sensible heating resulting from the distribution of cloud cover across the SFB in northern Alabama and Georgia during the morning (Fig. 1). According to the Sawyer–Eliassen equation, a strong vertical circulation should develop in response to such intense frontogenesis (Keyser and Shapiro 1986). The importance of such diabatic frontogenesis to prefrontal squall line formation in other cases has been deduced from observations (Koch 1984; Dorian et al. 1988) and shown with three-dimensional mesoscale models (Koch et al. 1997). However, Langmaid and Riordan (1998) conclude that frontogenetically induced ascent was insufficient to overcome the modest convective inhibition present over the outbreak region on Palm Sunday. Furthermore, the supercell storms developed entirely in the warm air well ahead of the SCF and SFB frontal boundaries (Fig. 2). Thus, it would appear that frontogenesis did not play a direct causative role in triggering the prefrontal tornadic storms, though it could have contributed to the rapid development of these storms and/or to the destabilization of the mesoscale environment.

Of particular interest to the present study are the strong interrelationships between mesoscale regions of enhanced frontogenesis along the SCF and SFB, the frontal mesolows, and the mesoconvective sytems. The Petterssen frontogenesis function was computed from objectively analyzed surface data. This function is expressed mathematically as
i1520-0493-126-8-2031-e1
where D is the total deformation in a rotated coordinate system defined with respect to the axis of dilatation, β is the angle measured between this axis and the isentropes, δ is the horizontal divergence, and |pθ| is the magnitude of the horizontal potential temperature gradient. This function ignores the tilting term and cross-frontal gradients of diabatic heating, but nevertheless has been found to be a valuable diagnostic indicator of surface frontogenesis in other case studies (e.g., Koch 1984; Sanders and Bosart 1985; Dorian et al. 1988). The results show that locally enhanced frontogenesis occurs in close association with each of the three mesolows that propagated over the frontal system stretching from northern Louisiana to northern Alabama (Fig. 7). Tornadic supercells constituting MCS A formed in close association with mesolow LA, and other severe storms within MCS B and C later developed in close association with the other two mesolows (Figs. 1–2). Frontogenesis was especially strong in association with mesolow C.

The MASS model forecasts (Fig. 8) display frontal mesolows and associated frontogenesis centers similar to those observed. For example, the observations at 1500 UTC show that mesolow LA was on the Mississippi–Alabama border at the junction of the SCF and SFB, MCS A was developing just to the south of this mesolow, and mesolow LC was in central Louisiana without any attendant convection. By comparison, the coarse grid MASS model forecast is just spinning up mesolow LA, and shows a feature closely corresponding to mesolow LC. Strong frontogenesis occurs in both the model and the observations in association with this mesolow at this time, yet frontogenesis also exists farther northeastward along the frontal boundary into eastern Mississippi, which is where mesolow LB develops immediately thereafter. Excellent agreement exists between all three of the forecast and observed mesolows at 1800 UTC. Mesolow LA is arguably the most important of the three mesolows, since it is the one directly connected to the development of the largely tornadic MCS A. Convection was triggered near mesolow LB in eastern Mississippi just before this time (Fig. 2), and frontogenesis is seen in both the forecast and surface data in direct association with this feature. Finally, by 2100 UTC all three mesoconvective systems have appeared, and once again, we see good agreement between the forecast and observed frontogenesis and mesolows, though the forecast frontogenesis associated with mesolow LC is underdone in the model, while the observations only hint at the existence of LB.

Additional analysis shows that the frontogenesis was almost entirely nonquasigeostrophic. Comparison between the geostrophic (Fig. 9a) and ageostrophic (Fig. 9b) surface frontogenesis fields [i.e., using the appropriate wind in (1)] shows a broken line of strong ageostrophic frontogenesis along the cold front, which when added to the small contribution by the geostrophic component, reproduces exactly the total frontogenesis field (Fig. 8c). Furthermore, there does not appear to be a spatially coherent component of the Q vector along the temperature gradient (directed toward the warmer air) at low levels (Fig. 9c). Since quasigeostrophic frontogenesis can occur only when Q·pθ ≠ 0 (Bluestein 1993), we conclude that there is no evidence for quasigeostrophic forcing in the lower troposphere (nor at the jet stream level, as discussed earlier) in the vicinity of the Palm Sunday tornado outbreak.

5. Conditional symmetric instability and inertial instability

Since frontogenetical forcing cannot explain the development of multiple convective bands, and furthermore, the convective lines in MCS A were prefrontal, we need to consider another mesoscale mechanism that can explain the observed multiple-band organization of deep convection on Palm Sunday. An obvious candidate is CSI, which is manifested as slanted roll circulations with their axes oriented along the thermal wind vector. Observational studies have suggested that CSI (or“slantwise convection”) may be the cause of some precipitation bands with typical widths of ∼30–200 km in convectively stable environments (Bennetts and Sharp 1982; Bennetts and Ryder 1984; Seltzer et al. 1985; Sanders and Bosart 1985; Wolfsberg et al. 1986; and Blakely 1988; Shields et al. 1991; Snook 1992; Moore and Lambert 1993).

Symmetric instability results from an imbalance between the Coriolis, pressure gradient, and buoyancy forces in an atmosphere that is both inertially and buoyantly stable (Emanuel 1979). Conditional symmetric instability arises when a saturated atmosphere is made symmetrically unstable due to the release of latent heat (Bennetts and Hoskins 1979). One method for isolating the presence of CSI is to find regions where the moist Ertel’s potential vorticity is negative. Alternatively, the parcel method can be used for evaluating CSI (Emanuel 1983a,b), and since it is preferred most often by operational forecasters (Snook 1992), we use it here. Given a cross section chosen orthogonal to the thermal wind direction, the necessary conditions for CSI are that a parcel is lifted sufficiently to become saturated, and that the slope of the moist isentropic (θe) surface is steeper than the slope of the geostrophic absolute momentum (Mg) surface, where
Mgυgfx,
υg is the geostrophic wind component normal to the cross section, and colder air is in the −x direction.
Very little attention has been directed to the role of CSI in the development of bands of thunderstorms in convectively unstable atmospheres, the notable exceptions being the case study by Jascourt et al. (1988) and the idealized numerical study of Seman (1994). This is probably due to the fact that it is generally thought that, while CSI and convective instability can coexist, upright convection should dominate because of its faster growth rate (Bennetts and Sharp 1982). Thus, for example, Moore and Lambert (1993) emphasize that care must be exercised to eliminate regions that are convectively unstable from possible CSI consideration. On the other hand, theoretical studies by Emanuel (1985), Thorpe and Emanuel (1985), and Xu (1986, 1989, 1992) suggest that as the moist potential vorticity decreases to zero, the moist ascent in frontal circulations becomes narrow and intense. The moist ascent tends to occur along the direction of least stability to slantwise displacement, which is between the slope of the Mg surfaces and the slope of the θw surfaces (where θw is the wet-bulb potential temperature). Thus, as the convective stability decreases, the circulations tend to asymptotically approach that of upright convection, but be organized by the symmetric instability into bands. Xu (1992) showed that multiple rainbands can be generated by viscous CSI when frontogenetical forcing occurs in the presence of negative moist geostrophic potential vorticity (qw) in a moderately deep saturated layer, where
i1520-0493-126-8-2031-e3
This equation shows that there are several ways that negative qw can arise, but it can occur only in the presence of convective instability (∂θw/∂z < 0) if the absolute geostrophic vorticity η = ςg + f > 0—that is, the atmosphere is inertially stable. Xu (1989) shows that CSI bands will not persist unless they are assisted by frontogenetic forcing. The Palm Sunday event was, of course, characterized by strong frontogenetic forcing.

We conducted CSI analyses upon the MASS, NGM, and Eta models, as well as the rawinsonde data. Vertical cross sections shown in Figs. 10c and 10d depict Mg surfaces (solid lines), equivalent potential temperature θe (dotted lines), and relative humidity RH > 90% (shaded) obtained from the MASS coarse grid forecasts. The cross sections are chosen perpendicular to the mean geostrophic shear vector in the layer between 700 and 800 mb, which is the layer in which the necessary conditions for CSI were most prevalent.

A small region of CSI is seen in the forecast valid for 1400 UTC about 100 km behind the surface cold front, though if the small wiggles in the θe surfaces are ignored, then a condition of near neutrality to slantwise convection appears throughout most of the cold frontal zone. Slantwise neutrality or slight CSI is often observed to coexist with frontogenetical forcing in cases where frontal rainbands are observed (Emanuel 1983b;Sanders and Bosart 1985; Kuo and Reed 1988). However, a distinct difference from those studies is that here a deep layer of convective instability is present ahead of the front and above the shallow CSI region.

The front-relative vector field (Fig. 10a) indicates strong flow directed toward the front at low levels. This flow is related to the low-level jet (Fig. 4d). Strong lifting (24 μb s−1 ∼ 35 cm s−1) occurs at low levels over the front, and general lifting is also evident along the entire sloping frontal surface. Strong upward motion occurring so close to the surface can rapidly spin up low-level vorticity due to stretching effects and could also assist the rapid intensification of the forecast mesolows. This explanation is similar to that drawn by Kuo and Reed (1988) in a numerical study of an explosively developing cyclone, though the spatial and temporal scales are much smaller here.

The frontal updraft jet is better organized and much deeper at 1800 UTC (Fig. 10b). The postfrontal slantwise ascent is also much better revealed at this time. This narrow sloping sheet of rapidly ascending air strongly resembles features seen in idealized models of frontogenesis in the presence of small stability to slantwise convection (Thorpe and Emanuel 1985; Emanuel et al. 1987), though in the present case we find some slantwise instability. These analyses suggest that the region of CSI seen at 1400 UTC was lifted by the ageostrophic frontal circulation to nearly 550 mb by 1800 UTC.2 A postfrontal region of CSI is similarly found in the Eta and NGM model forecasts (e.g., see Fig. 11). It is shown below (section 5a) that in the fine-grid version of the MASS model, lifting of this CSI region in the saturated atmosphere along the frontal surface allowed the instability to be released in the form of a simulated wide cold-frontal rainband about 70 km behind the surface front (Figs. 13c,d), a process that is quite similar to the conceptual picture presented by Knight and Hobbs (1988) in a two-dimensional model of frontogenesis in the presence of CSI.

There is another interesting feature that appears only in the 1800 UTC cross section (Fig. 10d)—a region of inertial instability in a deep layer above the surface front. Inertial instability occurs when the absolute geostrophic vorticity is anticyclonic (Holton 1992), a condition that is equivalent to requiring that in the cross section
i1520-0493-126-8-2031-e4
This condition is clearly met throughout a deep column within ±100 km of the surface front. Notice that this region is the same one occupied by the narrow frontal updraft (Fig. 10b) and that this plume is convectively unstable (Fig. 10d).

We now provide evidence that the strong updraft at 1800 UTC was forced by the inertial instability arising in association with a midlevel jetlet simulated by MASS (though not forecast by the operational models). Evidence for the actual existence of this jetlet is provided by wind profiler data (Kaplan et al. 1998). The forecast jetlet at this time was in northwestern Alabama (weakly evident at the 300-mb level in Fig. 4b). The strong upper-level ascent seen in Fig. 10b is contained within the upward branch of that jetlet’s ageostrophic vertical circulation (Fig. 6b). The jetlet is near the 450-mb level just behind the surface cold front, and is the cause for the extremely large horizontal gradients of Mg (cf. the jetlet location in Fig. 6b to the Mg gradients in Fig. 10d). Inertial instability is present on the right flank of the jetlet, which is also where the strong ascent is located. Neither this jetlet nor the strong inertial instability appeared in a MASS model simulation in which condensational heating was withheld, because the jetlet was forced by a geostrophic adjustment to earlier convection over northeastern Texas (Kaplan et al. 1998) and the inertial instability arose due to the existence of the jetlet.

It is of interest to consider the way in which such a state of inertial instability could arise. This jetlet appeared in response to a geostrophic adjustment process, whereby an imbalance between the pressure gradient and Coriolis forces due to a subgeostrophic momentum anomaly aloft accelerated the flow toward the cool side of the front (Hamilton et al. 1998; Kaplan et al. 1998). The basic process can be further understood with the nonlinear, two-dimensional model of CSI in a convectively unstable atmosphere described by Seman (1994), which showed that momentum transport within the convection tilts the M surfaces into the vertical, resulting in extrema of absolute vorticity (∂M/∂x) along the southern and northern flanks of the convection. He also proved that the addition of large CAPE to the Lagrangian parcel model of CSI (Emanuel 1983a) results in nearly vertical updrafts (instead of the ∼45° tilt characteristic of classical CSI) that are slanted over the cool air (in contrast to ordinary upright convection). This character describes well the structures seen in the MASS model simulation in the region above the frontal zone.

This strong updraft at 1800 UTC (Fig. 10b) is both inertially and convectively unstable, and generates a prefrontal band of convective precipitation in the fine mesh model in eastern Mississippi and western Alabama (Fig. 13d), similar to that observed (Fig. 12a). This naturally raises the question of whether the inertial instability generated convection in the upward branch of the jetlet’s transverse circulation, or whether the convection generated the inertial instability. Admittedly strong nonlinear feedbacks occurred between the jetlet, the inertial instability, and the convection. Nevertheless, a series of cross sections similar to those seen in Fig. 10, but following the jetlet in a semi-Lagrangian sense, revealed that a deep layer of inertial instability and strong ascent at and ahead of the front was associated with the jetlet for several hours prior to the development of convection (as early as 1400 UTC, when the jetlet was forming).

a. Comparisons between rainbands in the MASS model and the radar imagery

The hypothesized role of CSI and inertial instability in the production of the convective bands is further examined by comparing the MASS model fields and the radar imagery to the predictions of theory. The comparisons are summarized in Table 1. Before elaborating upon these comparisons, it is necessary to classify the observed and forecast rainbands. The radar imagery reveals two prefrontal convective bands composed of supercell storms that appear by 1600 UTC (Fig. 12a), a narrow cold frontal rainband from western Alabama to central Mississippi by 1800 UTC (Fig. 12b), and a wide cold-frontal rainband ∼ 100–200 km behind this front stretching from central Mississippi to central Tennessee after 1900 UTC (Fig. 2c). The coarse- and fine-grid MASS model precipitation fields (Figs. 13c,d) both show a similar wide cold-frontal rainband at a distance of 70–100 km behind the front, particularly so in the fine-grid forecast. A narrow cold frontal rainband is also contained in the fine-grid run at 1800 UTC in central Mississippi. Also of considerable interest is a prefrontal convective band in the fine-grid run about 40 km ahead of the cold front in eastern Mississippi and western Alabama at this time, as well as some isolated strong convection in central Alabama, both of which bear similarities to the observed mesoscale convection [though the two convective bands at 1600 UTC (Fig. 12a) are actually at distances of 70 and 140 km ahead of the SCF].

b. The role of conditional symmetric instability in the generation of postfrontal rainbands

The role of CSI in the production of frontal rainbands was further examined by comparing the MASS model fields and the radar imagery to the predictions of viscous CSI theory (Xu 1986). This theory dictates that the slope of the CSI circulation system is dependent upon values of the viscosity and the moist Brunt–Väisälä frequency:
i1520-0493-126-8-2031-e5
Slantwise ascent shows a preference for a slope that is between those of the Mg and θw surfaces. Theory predicts that the slope of the circulations is nearly vertical, that is, essentially that of upright convection, when N2w < 0 (Xu 1989; Persson and Warner 1993). The 1800 UTC special soundings at Centreville, Alabama, and Jackson, Mississippi, displayed CAPE values of 2817 and 2951 J kg−1, respectively, evidencing the strong potential instability, and with no capping inversion (Fig. 14). Similarly, the 1800 UTC MASS forecast at and ahead of the cold front exhibits considerable convective instability (Fig. 10d). Hence, if CSI occurred in the presence of strong frontogenesis at the surface cold front, it might create essentially upright convection in the form of bands of thunderstorms parallel to the thermal wind within the region of CSI.

Unsaturated symmetric instability theory (Bennetts and Hoskins 1979; Emanuel 1979) predicts that the band spacing should equal the ratio of the depth of the conditionally symmetrically unstable region to the slope of the isentropes. However, the most unstable CSI mode has an infinitely narrow updraft when moisture effects are included (Xu 1986). Grid resolution also affects the ability of a numerical model to properly represent CSI modes. Persson and Warner (1993) show that model grid lengths Δx < 30 km are needed to explicitly resolve CSI circulations, and that the most unstable CSI modes are substantially weaker and slower to form in a model that uses Δx > 15 km. This suggests that the coarse-grid MASS model would be able to resolve only the weaker CSI circulations, but that the 12-km fine grid should capably represent even the most unstable CSI modes. The narrowest updraft attained in the mesoscale model simulations of Persson and Warner (1993) was about 70 km, even at grid lengths as small as 6 km. The MASS model produced a wide cold-frontal rainband at a distance of 70 km behind the surface front, in good agreement both with their results, as well as the 80 km predicted from unsaturated symmetric instability theory.3

There is no known relationship for the growth rate of CSI modes. Knight and Hobbs (1988) suggest an analogy to the known growth rate for dry symmetric instability modes as
i1520-0493-126-8-2031-e6
which can be rewritten in terms of the Mg gradients and the slope of the θ surfaces as
i1520-0493-126-8-2031-e7
Application of (7) to the 1400 UTC MASS forecast fields in the CSI region of Fig. 10c, assuming a local θe slope of 0.011 and a latitude of 34°N, gives a predicted growth rate of σ2 = 3 × 10−8 s−2, which corresponds to an e-folding time of τ = 1/σ = 1.6 h. This growth rate is comparable to the 2-h development time for both the strong postfrontal ascent seen in the MASS model (Figs. 13c,d) and the wide cold-frontal rainbands seen in the radar imagery (Fig. 2). In summary, CSI theory predicts a circulation slope, rainband spacing, and band timescale in good agreement with the radar observations and the MASS model fields of the postfrontal bands.

c. The role of convective–inertial instability in the generation of prefrontal convective bands

Two-dimensional roll circulations with axes parallel to the shear vector may occur in the presence of inertial instability. Although similar in that sense to symmetric instability (Emanuel 1979), inertial instability is rarely seen in atmospheric data. While an atmosphere that is both inertially and convectively stable can still be unstable to slantwise displacements (i.e., it can be conditionally symmetrically unstable), an atmosphere that is both inertially and convectively unstable is certainly“ripe” for rapid development of convection into bands parallel to the thermal wind. We tentatively hypothesize that an asymmetric form of inertial instability in the presence of convective instability serves as an adequate explanation for the fact that the long-lasting supercell storms were all arranged in short lines parallel to the low-level shear vector (Fig. 12a).

Stevens and Ciesielski (1986) show that inertially unstable circulations can only occur within very small depths in the presence of dissipative forces, but that asymmetric inertial instabilities—that is, modes that display a structure along the direction of the flow—are preferred over symmetric modes on vertical scales of a few kilometers. This mechanism is particularly attractive because of the finite lengths exhibited by the prefrontal convective lines. This character is easily seen in the 1600 UTC radar imagery (Fig. 12a), in which neither of the two bands has an along-band length larger than 200 km.

The role of asymmetric inertial instability (AII) in the production of bands of severe convection ahead of the cold front was examined using the theory of Stevens and Ciesielski (1986). The methodology follows that used in Ciesielski et al. (1989). We compared the predicted slope of the circulation, the band spacing, the growth rate, and the along-flow length scale to the MASS model fields and to the radar imagery. The predicted circulation would be essentially vertical, since the Mg surfaces are nearly vertical in the region of inertial instability (Fig. 10d). Concerning the growth rate of AII, which is simply 50% larger than the first term in (7), we made two estimates at 1800 UTC—one valid for the AII region at upper levels in the vicinity of the jetlet, and another at 700 mb in the AII region just above the surface front. Calculations indicate that e-folding times of 0.9 h and 1.5 h result in these respective regions. Comparison to the observations (Table 1) indicates good agreement. We note that the conditions for AII were also present 1 h earlier, though the growth rates were smaller. Theory predicts that the band spacing and the along-jet length scales are both related to the width of the AII region, which in this case (Fig. 10d) is about 50 km, compared to the radar-observed band spacing of 70 km (Fig. 12a), MASS model-predicted band spacing of 40 km (Fig. 13d), and the observed and MASS model convective line lengths of 100–200 km. Thus, the predicted band spacing is consistent with the radar and mesoscale model fields, though the band lengths are only in fair agreement.

6. Gravity wave interactions with the jetlets, mesolows, and convection

Given the clear importance played by the surface mesolows in the tornado outbreak, a search was conducted for possible gravity wave activity. Mesoscale gravity waves characteristically display amplitudes of 1–15 hPa, wavelengths of 50–500 km, and periods of 1–4 h (Uccellini and Koch 1987). Most frequently, bands of showery precipitation accompany such gravity waves, but at other times, severe winter weather, damaging winds, or intense thunderstorms can occur. Detailed case studies describing interactions between mesoscale gravity waves and convective storms include those by Uccellini (1975), Stobie et al. (1983), Ferretti et al. (1988), Koch et al. (1988), Lin and Goff (1988), Bosart and Seimon (1988), and Rottman et al. (1992). Mesoscale gravity waves are nearly impossible to detect with standard synoptic hourly observations (Koch and Golus 1988), and they are not routinely predicted by operational forecast models, though mesoscale models have begun to be used to study their dynamics (Powers and Reed 1993; Kaplan et al. 1998). A methodology for the prediction and real-time mesoanalysis of gravity waves using digital mesoscale model data (the MesoEta Model), Automated Surface Observing System 5-min observations, and WSR-88D and wind profiler data has recently been presented by Koch and O’Handley (1997, hereafter referred to as KOH). Our analysis of gravity waves on Palm Sunday utilizes National Weather Service barograph data and the MASS mesoscale model, following the procedures discussed by KOH.

A review of the gravity wave literature suggests that the environment characterizing most mesoscale gravity wave events is one in which an upper-level jet streak separates from the geostrophic wind maximum and approaches an axis of inflection in the height field north of a low-level warm front or stationary front (Uccellini and Koch 1987; KOH). This conceptual model implies that geostrophic adjustment may be an important generation mechanism for such waves, since under such circumstances, transverse ageostrophic motions about the jet would be in a sense opposed to that predicted from quasi- and semigeostrophic theory (Koch and Dorian 1988). The occurrence of waves north of a frontal boundary suggests that wave ducting helps to maintain their coherence. The geostrophic wind maximum in the MASS forecast valid for 1800 UTC (Fig. 4c) is indeed located a large distance upwind of the jet streak, which at that time was over the Great Lakes. Although the axis of inflection is more vaguely defined in the present case than is typically seen in other gravity wave events, nonetheless the jet streak is approaching a large-scale ridge axis located over the east coast. According to the KOH conceptual model, the subsynoptic region that simultaneously defines the entrance region of the jet and the exit region of the geostrophic wind maximum is most conducive to being unbalanced. In this case, such a region is uniquely defined over northeastern Texas and western Arkansas; as we will see below, this is precisely the region from which both the observed and MASS forecast gravity waves emanate. Furthermore, the forecast and observed waves were confined to the cold side of the frontal zone, just as in this conceptual model.

a. Analysis of gravity waves in the mesoscale model

A sequence of three-hourly mean sea level pressure fields from the MASS model coarse-grid forecasts is shown in Fig. 15. A train of propagating pressure disturbances is apparent from northeastern Texas, across northern Louisiana, southern Arkansas, northern Mississippi and Alabama, and throughout Tennessee. The disturbances have a mean wavelength of 160 ± 47 km and a propagation speed of 38 ± 3.4 m s−1, and therefore, a mean wave period of 1.2 h.

These disturbances obey the polarization phase relations describing a gravity wave (Gossard and Hooke 1975). First, the wind perturbations in the northeast-southwest plane of wave propagation (u*′) are in phase with the pressure perturbations (p′): for a wave that is propagating in a northeasterly direction, the troughs (p′ < 0) should be collocated with the strongest northeasterly wind perturbations (u*′ < 0) and the ridges (p′ > 0) should be collocated with the strongest southwesterly wind perturbations (u*′ > 0). This relationship can be thought of as a positive pressure-wind covariance resulting from the wave impedance relation. Close examination of the wind vector and pressure fields (Fig. 15) suggests such a relationship, particularly for waves C, D, E, and F. However, to better facilitate this examination, a spatial bandpass filter (Maddox 1980) was developed and applied to the mean sea level pressure and u* wind fields to obtain quantitative analysis of the mesoscale perturbation fields (Fig. 16). The bandpass-filtered fields show several interesting mesoscale features, including (i) a convergence zone along the Gulf Coast due to differential friction effects, (ii) a mesoscale shear zone parallel to and just behind the frontal system, (iii) a frontal trough (p′ < 0) that is enhanced near the mesolows, and of greatest relevance here, (iv) the wavelike patterns in the p′ and u*′ fields. The filter nicely isolates the wave corridor on the cool side of the front and the strong positive covariance—that is, > 0.

The second polarization phase relation for a gravity wave is between the horizontal u*′ motions and the wave-induced vertical motions, namely a 90° phase difference as a consequence of mass continuity. The strongest upward motions depicted with shading in Fig. 15 are indeed generally found just ahead of each of the wave crests, again consistent with a gravity wave interpretation of the field. In summary, the propagating mesoscale disturbances that modulate the frontal mesolows in the MASS model forecasts are undoubtedly mesoscale gravity waves. The waves generally maintained coherence for several hours as they propagated in a northeasterly direction from their common point of origin (the “wave source region”) centered over northeastern Texas (Fig. 17).

The wave source region is not only where geostrophic adjustment is most likely to have been active, as explained above, but it is also where quasigeostrophic forcing is greatest (Fig. 5). In addition, this region is where intense convection was simulated to have first occurred. This collocation of many different processes may appear to complicate the determination of the actual cause of gravity wave generation in the model. However, since the gravity waves were absent in a run of MASS in which latent heating was not allowed to occur (not shown), this indicates that the waves were either convectively forced or sustained, at least in the model. Further inspection reveals that the waves are clearly confined to the cool side of the surface cold front, which is where strong lower-tropospheric static stability is present, as required for a wave duct.

A vertical cross section chosen to lie along the main wave corridor at 1800 UTC (Fig. 15c) is presented in Fig. 18. Shown are the magnitude of upward velocity (shaded), isentropes (solid lines), and the “critical level” (heavy curve), which is the level at which the wind speed in the direction of wave propagation equals the wave phase speed. The upward motions generally lead the isentropic wave ridges below 500 mb (the waves are propagating from left to right) by approximately one-quarter of a wavelength, which is consistent with a gravity wave interpretation. The locations of the gravity wave troughs in the mean sea level pressure field (Fig. 16c) are also depicted in the cross section. The troughs lie directly underneath isentropic depressions in the low- to mid-troposphere, with the strongest upward motions immediately following the troughs. A fairly close correspondence exists between the critical level and the layer in which the upward motions are most concentrated. Also notice the pronounced upstream (leftward) tilt of the vertical motions above the critical level, and that very little tilt is evident at lower altitudes;this behavior indicates that an energy source for the waves occurred at the critical level (e.g., Gossard and Hooke 1975; Koch et al. 1993). These structural and phase relationships are all consistent with a wave-CISK behavior, according to which (e.g., Raymond 1984) organized convection is forced by convergence associated with a gravity wave, while latent heat release within the convection provides a source of wave energy. Wave-CISK requires that the strongest convective vertical motions exist at the wave critical level, so that the gravity wave and convective system move in tandem together, thereby allowing for optimum transfer of energy to the waves from the convection.

Closer inspection of the isentropic surfaces indicates that the wave motions were primarily confined to a layer of strong static stability below ∼650 mb (although the vertical motions extended above that altitude, deep convection may have affected this behavior). This observation and the fact that the waves exhibit very little tilt below the critical level together suggest that the waves were actually ducted wave-CISK modes, similar to the gravity waves examined in a mesoscale model simulation by Powers and Reed (1993). In other words, energy needed to maintain gravity waves in the presence of vertical leakage of wave energy was apparently provided by convection near the critical level; at the same time, ducting primarily within the lower troposphere reduced the vertical leakage of wave energy. The criteria specified by Lindzen and Tung (1976) for an efficient wave duct were all present in the model in the cool air north of the front. Generally speaking, these conditions are a sufficiently deep layer of large static stability in the lower troposphere, and a critical level within a higher, conditionally unstable layer. The intrinsic (wave-relative) ducted phase speed is given as
i1520-0493-126-8-2031-e8
where D1 is the duct depth, N1 is the Brunt–Väisälä frequency within the stable duct layer, ΘTB) is the potential temperature at the top (bottom) of the stable duct layer, and Θ is the average value within the layer. Generally speaking, computations of the ducted phase speeds support the hypothesis that a disturbance is a ducted gravity wave when these speeds are reasonably close to those observed. The proper choice of the duct layer is not altogether clear, since Fig. 18 shows it to vary horizontally and to not be very well demarcated in terms of an abrupt change of the static stability. A reasonable choice would be the layer below 650 mb (D1 = 3.7 km), which gives C*d = 33 m s−1. When added to the mean wind in the duct layer (∼8 m s−1), this provides for a ground-relative ducted phase speed of 41 m s−1, which closely matches the model wave phase speed (38 m s−1).

These findings lead us to conclude that deep convection in northeastern Texas triggered a train of large-amplitude gravity waves in the MASS model, and that these waves were maintained by both wave-CISK and ducting processes in their northeastward movement along and to the rear of the cold front. It is interesting to note the close correspondence between the wave packets and the two jetlets produced by MASS (Fig. 15). The findings from use of an idealized model with prescribed diabatic heating by Hamilton et al. (1998) suggest that the midlevel jetlet formed in conjunction with a thermally forced gravity wave, both of which propagated downstream, and that a vertical circulation developed on the right flank of the jetlet due to the wave. Gravity waves generated from a heat source have been shown by Lin and Smith (1986) and Lin and Li (1988) to accelerate the flow at the level of maximum heating from the flux convergence of momentum associated with the waves. The MASS simulation suggests that wave-CISK forcing of the jetlet occurred as geostrophic adjustment of the wind to the latent heating was maintained for a considerable period of time. Thus, the frontal mesolows can be visualized as being the end result of a complex interaction between upstream convection, gravity waves, and the jetlets.

Finally, it is worth noting that the existence of the waves is not dependent upon model grid resolution, since in the fine-grid forecasts (Fig. 19) we see a packet of gravity waves very similar to those appearing in the coarser-grid model (Fig. 15). Not only do the locations, appearances, and pressure-wind covariances agree well with those found for the coarse-grid model, but calculations of the phase speed (33.4 m s −1), wavelength (140 km), and period (1.2 h) are also all quite similar.

b. Observational analysis of gravity waves

Barograph data were manually digitized at 15-min intervals, and then subjected to a bandpass filter (Fig. 20) to remove synoptic-scale waves, the diurnal pressure oscillation, and waves with periods shorter than 1 h that are unresolvable with the synoptic barograph spacing of ∼100 km. Examples of such processed barographs are shown from Muscle Shoals (MSL) and Huntsville, Alabama (HSV), in Fig. 21. Notice that the waves with pressure perturbations |p′| ≥ 0.5 mb are seen in both traces, with features in the MSL trace typically leading HSV by about 1 h (suggesting propagating waves).

A problem that immediately appears when attempting to map out the pressure perturbation fields associated with mesoscale gravity waves is that the 100-km average barograph site spacing prohibits an objective analysis of wavelengths as small as 200 km. However, KOH show that it is possible to objectively analyze such wavelengths by using the time series observations to“fill in” data between the sparsely distributed stations. Upon identifying a steadily propagating disturbance, determining its phase velocity C, and using this parameter to “advect” off-time observations, one can effectively create such observations using “time-to-space conversion” (TSC). A TSC version of the Barnes objective analysis scheme (Koch et al. 1983) was used to perform the analyses that follow, with each observation being given a weight that decreased not only with distance from the grid point being analyzed, but also as the time difference from the analysis time.

The selection of an appropriate value for the length of the advection period (about ±45 min), which affects the response characteristics of the objective analysis, is explained by KOH. They also show that wave ducting theory can be of considerable value in predicting the appropriate value of C, in that the vector mean wind in the conditionally unstable layer can be used as a surrogate to obtain a sufficiently accurate estimate of C. Calculations of this mean wind from several soundings in the region extending from Georgia to Missouri produced estimates in the range from C = 231°, 19.6 m s−1 to C = 244°, 42.0 m s−1, with the higher speed values being representative of conditions closer to the polar jet. We also performed a cross-correlation analysis to check on these estimates. The time lag that maximized the cross-correlation function between successive pairs of barograph sites over the outbreak region was computed, and then these were inserted into a least squares program to obtain the phase velocity from the many individual estimates; the method is similar to that described by Koch and Golus (1988). This analysis confirmed that a phase velocity of C = 240°, 22 m s−1 was optimum over the outbreak region from Arkansas into northern Alabama for use in the TSC scheme.

Objectively analyzed maps of p′ were constructed every 15 min from 0500 UTC 27 March to 1115 UTC 28 March. Coherently propagating mesolow and mesohigh disturbances were found from 1200 to 2300 UTC 27 March throughout the outbreak region of northeastern Texas, Arkansas, Louisiana, Mississippi, and northern Alabama and Georgia, as well as into Tennessee and Kentucky (Fig. 22). This period agrees with the time interval during which the frontal mesolows were seen, but the gravity waves obviously affected a much wider area on the cold side of the front. Although the hourly pressure observations are not adequate to reveal the wave activity, there is some hint that significant pressure waves are present deep into the cold air (Fig. 7). The 15-min digitized barograph data are required to adequately capture these mesoscale disturbances. A close comparison between the locations of propagating wave crests and troughs (Fig. 22) and pressure disturbances seen in the surface analyses (Fig. 7) does show fairly good correspondence, despite this sampling limitation.

Comparison of these pressure perturbation maps to the mesoscale model forecast fields (Figs. 15, 16, and 19) reveals that the model has captured well the observed wave corridor region, the wave genesis region, and the wavelengths of ∼150 km, though deterministic forecasts of the exact locations of individual wave troughs and crests is not good. The major difference between the model and the mesoanalysis is that the propagation speed of the observed p′ waves is only ∼22 m s−1, compared to the mesoscale model wave speeds of 38 m s−1. This difference results in a large dissimilarity in wave periods—the observed period was ∼3 h (Fig. 20), whereas the forecast period computes to be ∼1.2 h. Most of the energy in the time series is not contained in such high frequencies, even before the application of the bandpass filter. Such short periods (wavelengths) could not be resolved by the time-to-space conversion objective analysis scheme anyway. We can only speculate as to why MASS produced such high-frequency, rapidly propagating waves. One possibility is that the actual wave source was geostrophic adjustment instead of convection (as seen in the model), since convection would tend to produce higher frequency waves.

Whatever the cause may have been, it is still interesting to see the ability of a mesoscale model to reproduce many of the observed wave characteristics, and to provide understanding of the key role of the associated unbalanced jetlets in the generation of the waves. The MASS model suggests that ultimately, the gravity waves, jetlets, inertial instability, frontal mesolows, and tornadic storms all owed their existence to deep convection over eastern Texas that occurred hours before the supercell storms developed in Alabama! Unfortunately, the observations are inadequate to fully verify all of these processes, though the data at hand do lend support for the hypothesis.

7. Conclusions

The convective precipitation that occurred on 27 March 1994 in the southeastern United States was highly organized on the mesoscale: 1) most important were two prefrontal bands of supercell storms in northern Alabama and Georgia ahead of a slow-moving cold front, 2) a narrow band of severe thunderstorms erupted right along this front from Mississippi to Alabama, and 3) a wide and diffuse rainband was present 100–200 km behind the cold front. The observed banding of convection parallel to the thermal wind shear in this region may have helped to focus ambient vorticity into shear zones. Analyses based upon operational data and models, digitally processed barograph data, detailed radar and satellite imagery, and a research mesoscale model (MASS) suggest that a complex interaction between mesoscale processes was responsible for this spectrum of mesoscale precipitation systems.

The synoptic-scale setting for this tornado outbreak did not exhibit any of the classical ingredients of cyclogenesis, positive vorticity advection associated with a short wave/jet streak, quasigeostrophic forcing, or the upward branch of a jet streak vertical circulation located over the outbreak region. While a pronounced thermally direct transverse circulation was present in the polar jet entrance region, it was separated from the strong low-level vertical circulation associated with the cold front and was about 400 km from the outbreak region.

The MASS model reveals that the low-level frontal updraft jet became established in the presence of strong mesoscale forcing. Frontal mesolows and strong ageostrophic frontogenesis in the presence of large potential instability (CAPE > 2800 J kg−1) created this strong low-level circulation, which probably forced the explosive development of the cold frontal squall line. Rapid lifting occurred as strong flow (associated with a low-level jet) was directed toward the front in a front-relative sense and encountered a steep frontal head. The rapid intensification of the frontal mesolows was fostered by intense stretching within the frontal updraft in the presence of weak thermal stability. The occurrence of the frontogenesis in an environment containing CSI serves to explain the rapid rate of frontogenesis, the narrowness of the severe frontal rainband, and also the wide postfrontal rainband. The spacing (∼100 km) between this weak rainband and the intense frontal band, the rainband growth rates, and the slope of the vertical circulation all agreed well with the predictions from CSI theory.

The dual lines of supercells in the warm air 70 and 140 km ahead of the front in Alabama could not be explained on the basis of either CSI or frontogenetical processes. Rather, these convective lines were shown to have occurred in a kind of dynamical environment that has not been discussed before in tornado outbreaks, namely, one in which both potential instability and asymmetric inertial instability (AII) were present. A thermally direct circulation that was associated with the AII was, in turn, forced by an unbalanced mesoscale jetlet. The theory of convective-AII (Stevens and Ciesielski 1986) explained the near verticality of this circulation, the observed growth rates, and the supercell line spacing and lengths to within a factor of 2.

Bandpass filter analysis of the barograph data and the mesoscale model results indicate that gravity waves displaying wavelengths of ∼150 km were generated upstream of the tornado outbreak region within an area of strong quasigeostrophic ascent and deep convection, though unbalanced dynamics also may have played a role in the generation of the waves, given the similarity of the synoptic flow pattern to the conceptual model of Uccellini and Koch (1987). These gravity waves strongly modulated the frontal mesolows as they propagated coherently northeastward within the cool, stable air north of the cold frontal boundary. The waves were sustained by a supply of energy from wave-CISK, and energy leakage was further prevented by wave ducting in the stable lower troposphere.

These fascinating interactions between gravity waves, frontal mesolows, frontogenesis, conditional symmetric instability, asymmetric inertial instability, and potential instability were strongly influenced or even dependent upon the existence of mesoscale jetlets centered at 450 mb. The jetlets [described by Kaplan et al. (1998) and Hamilton et al. (1998)] were forced in the MASS model by earlier convection in eastern Texas. Inertial instability arose due to the momentum anomaly associated with the strongest jetlet. Geostrophic adjustment to the subgeostrophy created by the upstream convection initiated a train of propagating gravity waves, which organized the frontal mesolows. The horizontal mesoscale circulations associated with the mesolows enhanced moisture convergence and vorticity, resulting in locally enhanced stretching effects, frontogenesis, and low-level vertical circulations. The enhanced frontogenesis led to stronger mesoscale regions of baroclinicity and, according to the thermal wind relation, stronger vertical gradients of absolute momentum just above the surface front. Frontogenesis was further enhanced as the distorted Mg surfaces and steepened equivalent potential temperature surfaces, produced by the strong horizontal transport of high θe by the frontal circulation, came into a configuration necessary for conditional symmetric instability to occur. The ultimate source of all these mesoscale process interactions was the jetlet, which itself was forced by convection far upstream of the tornado outbreak region. Further insight into these scale-interactive processes requires improved understanding and observations of the interactions between convection, gravity waves, and the geostrophic adjustment process associated with unbalanced jetlets.

Acknowledgments

Thanks are extended to Dr. Kevin Knupp at the University of Alabama for contributing the digital radar and satellite imagery, and to Dr. Ken Waight of MESO, Inc., for assistance with running the MASS model. Helpful comments on the manuscript were provided by Ed Delgado at the NWS-Raleigh. Model runs were performed at the North Carolina Supercomputing Center and on the FOAMV cluster, funded by IBM, at North Carolina State University (NCSU). Additional assistance in producing some of the graphics was provided by Gail Andrews, Michael Black, and Brian Potter at NCSU. This research was funded in part by NSF Grant ATM-9319345, from NOAA through the Southeast Consortium on Severe Thunderstorms and Tornadoes under Grant NA27RP029201, and from Air Force Grant F49620-95-1-00226. The views expressed herein are those of the authors and do not necessarily reflect the views of NOAA.

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APPENDIX

Mass Model Configuration

The MASS model is a hydrostatic, terrain-following coordinate modeling system. The configuration used in the present study consists of a Blackadar high-resolution planetary boundary layer scheme, a modified Kuo cumulus parameterization scheme that includes convective-scale downdrafts, and a prognostic explicit microphysics scheme for water and ice particles. MASS was initialized with a first guess provided by a Global Optimum Interpolation (GOI) gridded dataset, reanalyzed with upper-air and surface NMC data using a three-dimensional OI scheme at 0000 UTC 27 March 1994, and further modified using manually digitized radar data to provide a better representation of the moisture field. A 24-h simulation was performed on the “coarse mesh” (24-km grid), which then provided one-way interactive boundary conditions for a “fine mesh” (12-km grid) simulation that was initialized at 0400 UTC and integrated for 20 h of real time. The grid domains are shown in Kaplan et al. (1998).

Fig. 1.
Fig. 1.

GOES visible satellite imagery at (a) 1501, (b) 1701, and (c) 1901 UTC 27 March 1994 showing the three mesoconvective systems and subcomponent convective lines. Surface frontal analyses are superimposed on the imagery.

Citation: Monthly Weather Review 126, 8; 10.1175/1520-0493(1998)126<2031:MDITPS>2.0.CO;2

Fig. 2.
Fig. 2.

Surface winds (whole barb = 5 m s−1, half barb = 2.5 m s−1), temperatures (°C), subjectively analyzed isobars (2-mb intervals), and composite radar reflectivity mosaics for 1500, 1700, and 1900 UTC. The mesolows that were directly associated with three MCSs (see text) are also depicted.

Citation: Monthly Weather Review 126, 8; 10.1175/1520-0493(1998)126<2031:MDITPS>2.0.CO;2

Fig. 3.
Fig. 3.

Barnes objective analyses of rawinsonde data at 1200 UTC depicting upper-level jet and quasigeostrophic dynamics: (a) 200-hPa winds (isotachs > 50 m s−1 at 10 m s−1 intervals) and geopotential height contours (m), (b) as in (a) except at 300 mb, (c) Q-vector divergence (10−15 m kg−1 s−1, solid contours) and convergence (dashed) at 700 mb, and (d) cross section depicting ageostrophic circulations transverse to upper-level polar (J1) and subtropical (J2) jet streaks and isentropes (solid, 3-K intervals) from North Platte, NE, to Cape Kennedy, FL [see (a) for cross-section location].

Citation: Monthly Weather Review 126, 8; 10.1175/1520-0493(1998)126<2031:MDITPS>2.0.CO;2

Fig. 4.
Fig. 4.

MASS model 18-h forecasts valid at 1800 UTC of (a) winds at 200 hPa (as in Fig. 3, except isotachs > 45 m s−1 shown), (b) winds at 300 hPa (isotachs > 45 m s−1), (c) geostrophic wind (isotachs > 75 m s−1) and geopotential height contours (m) at 300 hPa, and (d) winds and geopotential height (m) at 850 hPa. Isotachs > 65 m s−1 are shaded in (a) and (b). Also depicted are model forecast locations of surface frontal systems, for ease of reference, and cross sections A–A′ and B–B′ shown in later figures. Note distance scale.

Citation: Monthly Weather Review 126, 8; 10.1175/1520-0493(1998)126<2031:MDITPS>2.0.CO;2

Fig. 5.
Fig. 5.

Eta and MASS Model forecasts of quasigeostrophic forcing valid for 1800 UTC: (a) 18-h Eta Model forecast Q-vector field at the 500-hPa level, and (b) smoothed 18-h MASS model forecast Q-vector field for the 500–300-mb layer. Q-vector convergence regions are shaded [10−15 m kg−1 s−1 in (a) and 10−14 m kg−1 s−1 in (b)]. The 500-mb geopotential height contours (m) are also depicted in (b).

Citation: Monthly Weather Review 126, 8; 10.1175/1520-0493(1998)126<2031:MDITPS>2.0.CO;2

Fig. 6.
Fig. 6.

Forecast upper-level jet transverse circulations for (a) 1400 UTC from Omaha, NE, to Tallahassee, FL (Fig. 4b), and (b) 1800 UTC. Vertical cross sections derived from MASS model forecast fields show ageostrophic circulation (vectors), isotachs ≥ 40 m s−1, and isentropes (solid, 2-K intervals). Note horizontal vector scale (top right) and horizontal distance scale (bottom). Forecast surface frontal position is depicted by arrow along abscissa. Large “J” depicts polar jet, small “J” denotes mesoscale jetlet.

Citation: Monthly Weather Review 126, 8; 10.1175/1520-0493(1998)126<2031:MDITPS>2.0.CO;2

Fig. 7.
Fig. 7.

Objectively analyzed surface frontogenesis fields [shaded at 6°C (100 km)−1 (3 h)−1 intervals], isobars, and frontal analyses at (a) 1200, (b) 1500, (c) 1800, and (d) 2100 UTC.

Citation: Monthly Weather Review 126, 8; 10.1175/1520-0493(1998)126<2031:MDITPS>2.0.CO;2

Fig. 8.
Fig. 8.

As in Fig. 7 except fields are obtained from MASS model forecasts at the same times. Mesolows are given labels to correspond as closely as possible to those in the observations.

Citation: Monthly Weather Review 126, 8; 10.1175/1520-0493(1998)126<2031:MDITPS>2.0.CO;2

Fig. 9.
Fig. 9.

(a) Surface frontogenesis function field using the geostrophic wind for comparison to Fig. 8c [shaded at 3°C (100 km)−1 (3 h)−1 intervals], (b) as in (a) except using the ageostrophic wind, and (c) quasigeostrophic frontogenesis analysis using the 900–800-mb layer-averaged Q vectors and Q-vector convergence (shaded > 20 × 10−15 m kg−1 s−1) and potential temperature at 850 mb (solid lines, 3-K intervals). All fields are derived from the MASS model forecasts valid for 1800 UTC.

Citation: Monthly Weather Review 126, 8; 10.1175/1520-0493(1998)126<2031:MDITPS>2.0.CO;2

Fig. 10.
Fig. 10.

MASS model forecasts of frontal features in a plane normal to the surface cold front, extending from southeastern Missouri to southwestern Georgia (Fig. 4b). (a) and (b) Front-relative circulations, regions of upward motion (shaded at 5 μb s−1 intervals), and isentropes of potential temperature valid for 1400 and 1800 UTC, respectively. (c) and (d) Isentropes of equivalent potential temperature (2-K intervals, dotted lines), absolute geostrophic momentum (10 m s−1 intervals, solid lines), and relative humidity > 90% (shaded). The regions wherein the necessary conditions for conditional symmetric instability are met (see text) are shown by the thick curves. The positions of the predicted surface cold front are denoted by the small arrows at the bottom of the top panels.

Citation: Monthly Weather Review 126, 8; 10.1175/1520-0493(1998)126<2031:MDITPS>2.0.CO;2

Fig. 11.
Fig. 11.

Conditional symmetric instability (CSI) in the NGM model 18-h forecast for 1800 UTC in a vertical cross section from Omaha, NE, to Tallahassee, FL (along line A–A′ in Fig. 4b). Shown are isentropes of equivalent potential temperature (2-K intervals, thin lines), absolute geostrophic momentum (10 m s−1 intervals, thicker lines), and relative humidity > 90% (shaded). The region wherein the necessary conditions for CSI are met is shown by the very thick curve. The position of the predicted surface cold front is at the arrow appearing at the bottom.

Citation: Monthly Weather Review 126, 8; 10.1175/1520-0493(1998)126<2031:MDITPS>2.0.CO;2

Fig. 12.
Fig. 12.

Radar-observed precipitation patterns in the local outbreak region (see Fig. 2 for isoecho intensity levels), and their relationships to surface frontal systems at (a) 1600 UTC and (b) 1800 UTC. Compare prefrontal convective band locations at 1600 UTC to the CSI and inertial instability regions seen in 1800 UTC MASS display (Fig. 10d), and to the fine-grid MASS model precipitation forecasts (Fig. 13d).

Citation: Monthly Weather Review 126, 8; 10.1175/1520-0493(1998)126<2031:MDITPS>2.0.CO;2

Fig. 13.
Fig. 13.

Coarse- (24 km) and fine- (12 km) mesh MASS model forecast fields of precipitation rate (shaded, mm h−1), surface winds (m s−1), and surface equivalent potential temperature (θe) fields (2-K intervals, dotted lines) for (a) 1400 UTC (coarse grid), (b) 1400 UTC (fine grid), (c) 1800 UTC (coarse grid), and (d) 1800 UTC (fine grid). Mesoanalyses are based on the θe and wind fields. Outflow boundaries are denoted by smaller pips.

Citation: Monthly Weather Review 126, 8; 10.1175/1520-0493(1998)126<2031:MDITPS>2.0.CO;2

Fig. 14.
Fig. 14.

Rawinsonde soundings at 1800 UTC from (a) Centreville, AL, and (b) Jackson, MS, plotted on skew T diagrams and (insets) wind hodographs (m s−1). Surface-based parcel lifted adiabatically is shown by the dashed line, and positive area (CAPE) is represented by shaded regions. Wind barbs on right ordinate use these conventions: flag = 25 m s−1, whole barb = 5 m s−1, half barb = 2.5 m s−1.

Citation: Monthly Weather Review 126, 8; 10.1175/1520-0493(1998)126<2031:MDITPS>2.0.CO;2

Fig. 15.
Fig. 15.

Evolving relationships between gravity wave troughs A, B, . . . , in mean sea level pressure field (light contour lines) and two mesoscale jetlets evident at 500 mb (revealed by isotachs > 45 or 50 m s−1, thick contour lines) at (a) 1200, (b) 1500, (c) 1800, and (d) 2100 UTC from MASS model forecasts. Vertical motion fields at 600 hPa (upward motions > −20 μb s−1 shaded), surface frontal systems and mesolows, surface wind vectors (same notation as in Fig. 13), and location of 1800 UTC cross section in Fig. 18 are also shown.

Citation: Monthly Weather Review 126, 8; 10.1175/1520-0493(1998)126<2031:MDITPS>2.0.CO;2

Fig. 16.
Fig. 16.

Spatial bandpass analysis of MASS model forecast mean sea level pressure and surface wind fields (Fig. 15), resulting in (a) perturbation pressure (p′) field valid at 1500 UTC, (b) perturbation wind (u*′) field valid at 1500 UTC, (c) p′ field valid at 1800 UTC, and (d) u*′ field valid at 1800 UTC. Solid (dotted) lines are positive (negative) values at 0.5-mb intervals (for perturbation pressure) and 0.5 m s−1 intervals (for perturbation winds). Gravity wave trough locations copied from Fig. 15 are shown by thick line segments, and surface frontal analyses are also shown.

Citation: Monthly Weather Review 126, 8; 10.1175/1520-0493(1998)126<2031:MDITPS>2.0.CO;2

Fig. 17.
Fig. 17.

Hourly isochrones of gravity wave troughs A, B, C, D, E, F, G, and H (F12 = forecast hour 12, etc.) obtained from coarse-mesh MASS model forecasts (see Fig. 15). Note horizontal scale in wave A panel.

Citation: Monthly Weather Review 126, 8; 10.1175/1520-0493(1998)126<2031:MDITPS>2.0.CO;2

Fig. 18.
Fig. 18.

Vertical cross section normal to gravity wave fronts from 1800 UTC coarse-mesh MASS model forecast (Fig. 15). Shown are isentropes of potential temperature (2-K intervals, solid lines), upward vertical motions (shaded at intervals of 4 μb day−1, see gray bar at bottom), and critical level for mean gravity wave speed of 38 m s−1 (dark curve). Locations of surface wave troughs (Fig. 16c) are shown in the cross-sectional plane (G, F, C, E).

Citation: Monthly Weather Review 126, 8; 10.1175/1520-0493(1998)126<2031:MDITPS>2.0.CO;2

Fig. 19.
Fig. 19.

As in Fig. 15 except based on fine-mesh MASS model forecast and not depicting the position of the jetlets.

Citation: Monthly Weather Review 126, 8; 10.1175/1520-0493(1998)126<2031:MDITPS>2.0.CO;2

Fig. 20.
Fig. 20.

Bandpass filter applied to digitized surface barograph data. Response function R(f) has a value of 1.0 at a given frequency when the waves of that frequency are passed without loss of amplitude. The “cutoff frequencies” [at which R(f) = 0.5] are highlighted in gray.

Citation: Monthly Weather Review 126, 8; 10.1175/1520-0493(1998)126<2031:MDITPS>2.0.CO;2

Fig. 21.
Fig. 21.

Time series of digitized barograph data (mb, light lines) and bandpass-filtered pressure perturbations (mb, dark lines) for Muscle Shoals (MSL) and Huntsville (HSV), AL, covering the period 0500 UTC 27 March–1115 UTC 28 March 1994. Similar phase-shifted pressure fluctuations are evident in the pressure traces.

Citation: Monthly Weather Review 126, 8; 10.1175/1520-0493(1998)126<2031:MDITPS>2.0.CO;2

Fig. 22.
Fig. 22.

Objectively analyzed pressure perturbation fields (contour interval Δp = 0.2 mb, with p′ < 0 dotted) for (a) 1200, (b) 1500, (c) 1800, and (d) 2100 UTC 27 March 1994.

Citation: Monthly Weather Review 126, 8; 10.1175/1520-0493(1998)126<2031:MDITPS>2.0.CO;2

Table 1.

Characteristics of mesoscale circulations parallel to the shear vector predicted by conditional symmetric instability (CSI) theory, predicted by the theory for asymmetric inertial instability (AII), as seen in the radar observations, and appearing in the MASS model forecasts.

Table 1.
1

All times henceforth will refer to 27 March 1994, unless otherwise noted.

2

The upward motion in the upper troposphere far behind the surface frontal location (in southeastern Missouri) is primarily related to the transverse circulation in the jet streak entrance region described earlier (Fig. 4b).

3

We assume for this calculation that the 900–700 mb (H = 2.0-km deep) layer contains CSI, and determine a frontal slope of α = 4.5 km/180 km = 0.025 based on the θe surfaces defining the front in Fig. 10d, whereupon the predicted wavelength is λ = H/α = 2.0 km × (180/4.5) = 80 km.

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