Case Studies of Midwestern Thundersnow Events

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CASE STUDIES OF MIDWESTERN THUNDERSNOW EVENTS

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A Thesis Presented to the Faculty of the Graduate School University of Missouri-Columbia

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In Partial Fulﬁllment for the Degree Master of Science

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by CHRISTOPHER E. HALCOMB

Dr. Patrick S. Market, Thesis Supervisor

DECEMBER 2001 COMMITTEE IN CHARGE OF CANDIDACY:

Assistant Professor Patrick S. Market Chairperson and Advisor

Professor William B. Kurtz Assistant Professor Anthony R. Lupo

i Acknowledgements

I will begin by thanking my advisor, Dr. Patrick Market, for allowing me the opportunity to study this fascinating subject and allowing me to largely run with it as I saw fit, but giving guidance when I truly needed it. I would like to thank him most for being a great friend and filling my life with humor and laughter. I would also like to thank Dr. Anthony Lupo, who was always willing to help anytime that I asked for it. I also consider him a great friend and I will always treasure the times the three of us have had together. I would like to thank Rebecca Ebert for assisting in the generation of many of the figures that are included in the thundersnow climatology, and for not killing me when I asked her to do just one more thing for me. I would like to acknowledge Jorge Flores for his expertise in writing the computer program to search for all possible combinations in which snow and thunder can occur simultaneously. It would have taken me forever to write such a program. I am deeply appreciative of the support that I have received from my mother and my grandmother, who are two of the kindest and most sincere persons that I have known. They have both been supportive of whatever endeavor that I chose to pursue and through thick and thin. I would also like to thank Cyri Parks for being such a good and supportive friend the last few months. I really am glad that I was able to get to know her. We are so much alike in so many ways. She will always fill a special place in my heart.

ii Contents

1 Introduction 1 1.1 Statement of Thesis ...... 4

2 Literature Review 5 2.1 Studies on Banded Precipitation ...... 5 2.2 Thundersnow-related Climatology ...... 20 2.3 Summary ...... 22

3 Methodology 24 3.1 Climatology ...... 24 3.2 Case Studies ...... 26 3.2.1 9 December 1999, 11 March 2000, and 19 April 2000 ...... 27 3.2.2 5 December 1999 ...... 30

4 Thundersnow Climatology 31 4.1 Spatial and Temporal Patterns ...... 31 4.2 Characteristics of Thundersnow Observations ...... 35

5 Case Studies 40 5.1 5 December 1999 ...... 40 5.1.1 Introduction ...... 40 5.1.2 Surface Analysis ...... 40 5.1.3 Upper Air Analysis ...... 41 5.1.4 Isentropic Analysis ...... 42 5.1.5 Stability Analysis ...... 46 5.1.6 Quasigeostrophic Forcing ...... 46 5.1.7 Conclusions ...... 47 5.2 9 December 1999 ...... 49 5.2.1 Introduction ...... 49

iii 5.2.2 Surface Analysis ...... 49 5.2.3 Upper Air Analysis ...... 50 5.2.4 Isentropic Analysis ...... 51 5.2.5 Stability and Forcing ...... 54 5.2.6 Conclusions ...... 59 5.3 11 March 2000 ...... 60 5.3.1 Introduction ...... 60 5.3.2 Surface Analysis ...... 60 5.3.3 Upper Air Analysis ...... 60 5.3.4 Isentropic Analysis ...... 61 5.3.5 Stability and Forcing ...... 64 5.3.6 Conclusions ...... 68 5.4 19 April 2000 ...... 70 5.4.1 Introduction ...... 70 5.4.2 Surface Analysis ...... 70 5.4.3 Upper Air Analysis ...... 70 5.4.4 Isentropic Analysis ...... 75 5.4.5 Stability and Forcing ...... 75 5.4.6 Conclusions ...... 78

6 Discussion and Conclusions 81 6.1 Discussion of Case Studies ...... 81 6.2 Conclusions ...... 82

References 85

Vita 88

List of Figures

g 2.1 Relationship of e (dashed, K) and M (solid, m s ) in the evaluation of CSI, CI, and WSS (Reproduced from a ﬁgure originally constructed by Dr. James T. Moore and Sean Nolan at Saint Louis University)...... 10

2.2 Lifted indices for various air parcels for North Omaha, Nebraska (OVN) at 0000 UTC on 21 February 1993 (Reproduced from Moore

et al. 1998)...... 11

2.3 T-log P plot of e (K) at North Omaha, Nebraska (OVN) at 0000 UTC on 21 February 1993. (Reproduced from Moore et al. 1998)...... 12

2.4 Conceptual model depicting the frontogenetical region and zone of EPV reduction in a developing cyclone (Reproduced from Nicosia and Grumm 1999)...... 13

2.5 Proposed positive feedback mechanism between frontogenesis and the reduction of EPV (Reproduced from Nicosia and Grumm 1999). .14

2.6 Observed 24-hour snowfall totals (cm) ending at 0000 UTC on 20 January 1995. (Reproduced from Martin 1998b)...... 15

2.7 Cloud-to-ground lightning strikes in a 24-hour period ending at 0000 UTC 20 January 1995 (Reproduced from Martin 1998b)...... 15

v 2.8 (a) 6-hour forecast of 200-m frontogenesis from the UW-NMS model

valid at 0600 UTC 19 January 1995. Shaded regions denote positive

1 1 frontogenesis every 1 K (100 km) day . (b) As in (a) except from a 12-hour forecast valid at 1200 UTC on the same day. (c) As in (a) except from a 18-hour forecast valid at 1800 UTC. (d) As in (a) except from a 24-hour forecast valid at 0000 UTC 20 January 1995. (e) As in (a) except from a 30-hour forecast valid at 0600 UTC 20

January 1995. (Reproduced from Martin 1998a)...... 16 ~ 2.9 Schematic showing the natural coordinate partioning of Q . Dashed lines depict isentropes on an isobaric surface (Reproduced from

Martin 1999)...... 17

~ Q

2.10 The effect of s on horizontal thermal structure. (a) Isentropes ~

(solid lines) in a ﬁeld of Q, with the maximum region of convergence shaded. Dashed line indicates convergence axis, while r depicts

thermal gradient. (b) Thick black arrow depicts original direction of

r , while thick gray arrow depicts the direction of after being ro-

~ Q

tated by s . (c) Oriention of thermal zone in (a) after being rotated

~ Q

by s (Reproduced from Martin 1999)...... 18

Q 5 10

2.11 (a) Convergence of s contoured and shaded every m

1 1 kg s in the 600-900 hPa layer from an 18-hour forecast of the

UW-NMS model valid at 0600 UTC 23 October 1996. (b) As in (a)

~ Q

except for n (Reproduced from Martin 1999)...... 19

2.12 Deﬁntions of several types of instabilities (Reproduced from Schultz and Schumacher 1999)...... 19

2.13 Mean proximity sounding for 13 thundersnow reports during the period of 1968-1971. The thick, solid line depicts the temperature pro- ﬁle. The thick, dashed line depicts the dewpoint proﬁle (Reproduced from Curran and Pearson 1971)...... 21

2.14 Number of hours of thunder at temperatures below 0 C from 1982- 1990. (Reproduced from Holle et al. 1998)...... 22

3.1 Delineation of regions used in the thundersnow climatology...... 25

vi 4.1 Number of thundersnow events for each region during the period of 1961-1990. Refer to regions depicted in Fig. 3.1...... 32

4.2 Number of thundersnow events for each state during the period of 1961-1990. Note that the sum of the events will not equal 375 (Refer to Section 3.1)...... 33

4.3 As in Fig. 4.1 except events (N=375) by category (Refer to Section 3.1 for the description of categories)...... 34

4.4 As in Fig. 4.2 except with normalized values...... 35

4.5 As in Fig. 4.1 except by month (N=375)...... 36

4.6 As in Fig. 4.1 except by Local Standard Time (LST) (N=375)...... 36

4.7 Polar plot showing the location of thundersnow events (Category 1) relative to the position of the center of the parent low (N=247). Direction is given in the traditional meteorological azimuth (degrees) from the postion of the low to the observing station. Distances are given in km...... 37

4.8 As in Fig. 4.1 except by distance (km) from parent low center of initial report (N=375)...... 38

4.9 Representative station model for the average initial report for all thundersnow events (N=375). Temperature and dewpoint are given in degrees F, the wind speed in knots, and the sea level pressure in hPa. The standard deviation for all parameters is also represented. .38

4.10 As in Fig. 4.1 except by snowfall intensity of initial report (N=375). .. 39

5.1 Surface analysis valid at 0600 UTC on 5 December 1999. Isobars drawn every 2 hPa. Red line depicts cross section line...... 41

5.2 Radar mosaic valid at 0600 UTC on 5 December 1999 (obtained from the National Climatic Data Center)...... 42

5.3 850-hPa Rapid Update Cycle initial ﬁeld analysis valid at 0600 UTC

on 5 December 1999. Isotherms (thick lines) drawn every 4 C, height contours (thin lines) every 2 dkm...... 43

vii 5.4 As in Fig. 5.3 except for 300-hPa level. Height contours (thick lines) drawn every 5 dkm, isotachs (thin lines) every 10 knots (2=20 kts, 12=120 kts, etc...... 43

5.5 As in Fig. 5.3 except for 300-hPa divergence (depicted every 2

5 1 10 s )...... 44

5.6 As in Fig. 5.3 except for 700-hPa equivalent potential temperature (depicted every 2 K)...... 45

5.7 As in Fig. 5.3 except dynamic tropopause pressures (depicted every 50 hPa)...... 45

5.8 Vertical cross section analysis of the 0600 UTC Rapid Update Cycle initial ﬁeld from 5 December 1999. Equivalent potential temperature

(dashed lines) are depicted every 3 K, Mg (solid lines) depicted every

5kgms , and relative humidity (dotted lines) depicted every 5%

% for values > 80 . KIAB is located near 97.25 W longitude...... 47

5.9 As in Fig. 5.3 except lifted indices for the 850-700-hPa layer (de-

picted every 1 C)...... 48

~ Q

5.10 As in Fig. 5.3 except for divergence of s in the 850-400-hPa layer

19 1 2 1

(depicted every 5 10 Pa m s ). Dashed lines depict areas

of convergence, while solid lines depict divergence...... 48

~ Q

5.11 As in Fig. 5.10 except for divergence of n ...... 49

5.12 Surface analysis valid at 0900 UTC on 9 December 1999. Isobars depicted every 2 hPa. Red line depicts cross section line...... 50

5.13 Radar mosaic valid at 0900 UTC on 9 December 1999 (obtained from the National Climatic Data Center)...... 51

5.14 850-hPa Rapid Update Cycle initial ﬁeld analysis valid at 0900 UTC

on 9 December 1999. Isotherms (dashed lines) contoured every 5 C, height contours (solid lines) every 30 m...... 52

5.15 As in Fig. 5.13 except 500-hPa heights and absolute vorticity. Ab-

5 1

solute vorticity (dashed lines) contoured every 5 10 s , height contours (solid lines) every 60 m...... 52

viii 5.16 As in Fig. 5.13 except 300-hPa heights, divergence and isotachs. Isotachs (solid purple) contoured every 20 knots and shaded for

speeds > 70 knots, height contours (solid black) every 120 m, and

5 1

divergence (dashed red) every 2 10 s ...... 53

5.17 As in Fig. 5.13 except dynamic tropopause pressure and 700-300-

hPa potential vorticity. Isobars (solid) depicted every 50 hPa, poten-

7 2 1 2

tial vorticity (dashed) every 5 10 m Ks kg ...... 53

5.18 As in Fig. 5.13 except pressure and storm-relative moisture transport vectors on the 294 K isentropic surface. Isobars (solid) de-

picted every 50 hPa, vectors indicate storm-relative moisture trans-

~ ~ ~

C C port (q(V )). calculated to be from 257 at 13.7 m s ...... 54

5.19 Vertical cross section from Clovis, New Mexico (CVS), to Dallas-Ft. Worth, Texas (DFW), from the 0900 UTC RUC initial ﬁeld valid on 09 December 1999. Solid lines depict equivalent potential temperature

(every 2 K), dashed lines depict absolute geostrophic momentum

(every 10 kg m s ). Values of relative humidity greater than 80%

are shaded green, and greater than 90% dark green. LBB is located near the second tick mark from the left edge...... 55

5.20 As in Fig. 5.18 except Petterssen surface frontogenesis (solid lines,

1 1 1

every 4 10 K 100 km 3h ), and (dashed lines, every 3 b 1

s ...... 56

~ Q

5.21 As in Fig. 5.13 except convergence (positive values) of s in the

14 1 1

850-400-hPa layer. Contours drawn every 4 10 mkg s . ... 57

~ Q

5.22 As in Fig. 5.20 except convergence (positive values) of n in the

14 1 1

850-400-hPa layer. Contours drawn every 1 10 mkg s . ... 58

5.23 As in Fig. 5.18 except three-dimensional equivalent potential vortic-

ity calculated with e . Contoured every 1 PVU with negative values shaded...... 58

5.24 As in Fig. 5.18 except three-dimensional equivalent potential vortic-

ity calculated with es . Contoured every 1 PVU with negative values shaded...... 59

ix 5.25 Surface analysis valid at 1200 UTC on 11 March 2000. Isobars depicted every 2 hPa. Red line depicts cross section line...... 61

5.26 Radar mosaic valid at 1200 UTC on 11 March 2000 (obtained from the National Climatic Data Center)...... 62

5.27 850-hPa analysis of the 1200 UTC Rapid Update Cycle initial ﬁeld valid on 11 March 2000. Heights (solid,black) depicted every 30 m,

temperatures (dashed,red) every 5 C...... 62

5.28 As in Fig. 5.27 except 500-hPa analysis. Heights (solid,black) de-

picted every 60 m, absolute vorticity (dashed, green) every 5 10 1 s ...... 63

5.29 As in Fig. 5.27 except 300-hPa analysis. Heights (solid,black) depicted every 120 m. Isotachs (dashed, purple) depicted every 20

knots with speeds of greater than 70 knots shaded. Divergence

5 1

(dashed, red) depicted every 2 10 s ...... 63

5.30 As in Fig. 5.27 except dynamic tropopause and 700-300-hPa poten-

tial vorticity analyses. Tropopause pressures (solid, black) depicted

7 2

every 50 hPa, potential vorticity (dashed, red) every 5 10 m K

1 2 s kg ...... 64

5.31 As in Fig. 5.27 except 500-hPa equivalent potential temperature (depicted every 2K)...... 65

5.32 As in Fig. 5.27 except pressure (depicted every 50 hPa) and storm-

~ ~

relative moisture transport vectors (q(V )) on the 294 K isentropic

1 ~ surface. C calculated to be from 293 at 9.3 m s ...... 65

5.33 Vertical cross section from Omaha, Nebraska (OMA), to Atlanta, Georgia (ATL), from the 1200 UTC RUC initial ﬁeld valid on 11 March 2000. Solid lines depict equivalent potential temperature (every 2

K), dashed lines depict absolute geostrophic momentum (every 10

kgms ). Values of relative humidity greater than 80% are shaded

green, and greater than 90% dark green. STL is located between the ﬁfth and sixth tick marks from the left edge...... 66

5.34 As in Fig. 5.33 except three-dimensional equivalent potential vortic-

ity calculated with e . Contoured every 1 PVU with negative values shaded...... 67

5.35 As in Fig. 5.33 except three-dimensional equivalent potential vortic-

ity calculated with es . Contoured every 1 PVU with negative values shaded...... 67

5.36 As in Fig. 5.33 except Petterssen surface frontogenesis (solid lines,

1 1 1

every 2 10 K 100 km 3h ), and (dashed lines, every 1 b 1

s ...... 68

~ Q

5.37 As in Fig. 5.27 except convergence (positive values) of s in the

14 1 1

850-400-hPa layer. Contours drawn every 2 10 mkg s . ... 69

~ Q

5.38 As in Fig. 5.37 except convergence (positive values) of n in the

14 1 1

850-400-hPa layer. Contours drawn every 1 10 mkg s . ... 69

5.39 Storm total snowfall amounts (inches) for 19 April 2000. (Obtained from the National Weather Service forecast ofﬁce in Rapid City, South Dakota.) ...... 71

5.40 Surface analysis valid at 1200 UTC on 19 April 2000. Isobars depicted every 2 hPa. Red line depicts cross section line...... 71

5.41 Radar mosaic valid at 1200 UTC on 19 April 2000 (obtained from the National Climatic Data Center)...... 72

5.42 850-hPa and 700-hPa RUC initial ﬁelds valid at 1200 UTC on 19 April 2000. 850-hPa heights (solid, black) depicted every 30 m, 700- hPa heights (dashed, red) depicted every 30 m...... 73

5.43 As in Fig. 5.42 except 500-hPa analysis. Heights (solid, black) de-

picted every 60 m, absolute vorticity (dashed, red) depicted every 5

5 1

10 s ...... 73

5.44 As in Fig. 5.42 except 300-hPa analysis. Heights (solid,black) depicted every 120 m. Isotachs (dashed, purple) depicted every 20

knots with speeds of greater than 70 knots shaded. Divergence

5 1

(dashed, red) depicted every 2 10 s ...... 74

xi 5.45 As in Fig. 5.42 except dynamic tropopause and 700-300-hPa poten-

tial vorticity analyses. Tropopause pressures (solid, black) depicted

7 2

every 50 hPa, potential vorticity (dashed, red) every 5 10 m K

1 2 s kg ...... 74

5.46 As in Fig. 5.42 except pressure (depicted every 50 hPa) and storm-

~ ~

relative moisture transport vectors (q(V )) on the 294 K isentropic

1 ~ surface. C calculated to be from 216 at 5.7 m s ...... 75

5.47 Vertical cross section from Worland, Wyoming (WRL), to Pipestone, Minnesota (PQN), from the 1200 UTC RUC initial ﬁeld valid on 19 April 2000. Solid lines depict equivalent potential temperature (every

2 K). Values of relative humidity greater than 80% are shaded green. RCA is located between the ﬁfth and sixth tick marks from the left edge...... 76

5.48 As in Fig. 5.47 except three-dimensional equivalent potential vortic-

ity calculated with e . Contoured every 1 PVU with negative values shaded...... 77

5.49 As in Fig. 5.47 except three-dimensional equivalent potential vortic-

ity calculated with es . Contoured every 1 PVU with negative values shaded...... 77

5.50 As in Fig. 5.47 except Petterssen surface frontogenesis (solid lines,

1 1 1

every 4 10 K 100 km 3h ), and (dashed lines, every 2 b 1

s )...... 78

~ Q

5.51 As in Fig. 5.42 except convergence of s in the 850-400-hPa layer.

14 1 1

Contours drawn every 2 10 mkg s ...... 79

~ Q

5.52 As in Fig. 5.42 except convergence of n in the 850-400-hPa layer.

14 1 1

Contours drawn every 1 10 mkg s ...... 79

xii Chapter 1

Introduction

The occurrence of convective snow has long been a forecast problem for operational meteorologists. Referred to as thundersnow, such events frequently occur in

mesoscale (often meso- ) bands which lie below the resolution of most numerical forecast models. Even with the evolution of mesoscale modeling systems, such as the 40-km Rapid Update Cycle (RUC) Model, the 29-km Meso Eta Model, and the scaleable Mesoscale Atmospheric Simulation System (MASS) Model, the dynamics of thundersnow events are still not well understood, and only in the past 15 to 20 years have such dynamics been addressed in earnest by meteorological researchers. Most of these studies, including Bennetts and Hoskins (1979), Ben- netts and Sharp (1982), Moore and Blakely (1988), and Martin (1998a, 1998b, and 1999), have dealt largely with mesoscale banding of snowfall and not speciﬁcally on studies of thundersnow events. Indeed, some distinction must be made between the terms ‘convective snow’ and ‘thundersnow’. While the presence of thunder necessitates the presence of convection (and the updraft speeds necessary for adequate electrical charge separation), the mere existence of convective overturning is insufﬁcient for the occurrence of charge separation, subsequent lightning, and thunder. Although these terms may be used interchangeably in this thesis (all four heavy convective snow events exhibited thundersnow), the difference between them is understood and

1 will be observed when necessary. Bennetts and Hoskins (1979), and Bennetts and Sharp (1982) attribute the formation of mesoscale banding to the presence of conditional symmetric instability (CSI), which is often present in saturated regions of strong baroclinicity that are considered to be inertially and gravitationally stable. However, several studies of thundersnow events over the past 20 years (including Nicosia and Grumm 1998 and Moore et al. 1998) reveal that the release of convective instability (CI) is also a common underlying culprit in such events, and may produce snowfall rates above those observed with the release of CSI. Nevertheless, the banded environments in which either CSI or CI may be released to produce thundersnow have many similar characteristics, including the advection of warm, moist air in the lower and middle troposphere, strong vertical wind shear, quasigeostrophic frontogenetical forcing at the mid-levels, and strong baroclinicity near the affected region. Moore and Lambert (1993) illustrate that equivalent potential vorticity (EPV) will be negative in either a CSI or CI environment. Nicosia and Grumm (1999) took this a step further when they observed that bands of heavy snow in several case studies coincided directly with the regions of the most negative EPV. As a result, the analysis of EPV may be helpful to operational forecasters in determining the

locations that are most susceptible to mesoscale banding. Related to this, Bennetts 2

and Sharp (1982) were able to show that when growth rates ( ) are calculated to 2

be at least 0.0 h , there is a 75 percent chance that banding will occur, and that

2 2

banding is almost a certainty when is greater than 0.2 h . It is our hypothesis that banded mesoscale precipitation possessing thundersnow appears to occur most commonly in the presence of two scenarios. First, in the northwest sector of an occluded cyclone, and second, ahead of an advancing warm front. Martin (1998a, 1998b, 1999) writes extensively about the dynamics of occluding cyclones. He describes the process of occlusion as less of a frontal pro-

2 cess, and instead as a dynamic middle to upper tropospheric process in which two differing air masses become separated by a thermal ridge. With Martin’s model, the occlusion may reach the surface in time, although this is not an important factor in the development of the precipitation patterns that actually will be observed.

Bands of precipitation, sometimes heavy, will form in association with the thermal

O W AL

ridge aloft, or trowal (TRough f arm air oft), in the northwest sector of an

~ Q occluding cyclone as a result of ascent via synoptic-scale, s forcing (Martin 1999).

The thermal ridge will be represented as a region of high equivalent potential temperatures ( e ) in plan view, which will separate two regions of colder, drier air. In the presence of the release of CSI or CI by frontogenesis, convective bands will form which may result in locally heavy precipitation rates. Whether CSI or CI is released depends, in part, upon the type of curvature-related shear that is present in the middle to upper troposphere. CSI is more likely to exist in regions of anticyclonic shear, while CI is more likely to exist in a cyclonically sheared environment (Bluestein 1986). The latter is due to the advection of warm, moist air in the lower troposphere and cold, dry air in the middle to upper troposphere. In advance of warm fronts, ascent occurs as air is forced up a poleward-sloping baroclinic zone. Weismuller and Zubrick (1998) were able to show that heavy mesoscale precipitation bands formed coincident with the regions of strongest

mid-tropospheric frontogenesis. Frontogenesis in these cases will likely occur as

~ ~

Q Q s a result of frontal-related n forcing, rather than forcing that is observed in the formation of the trowal. Mote et al. (1997), and Moore et al. (1998) state that isentropic upglide is another major component of ascent in these cases, often coincident with formation of a low-level jet. While thundersnow-speciﬁc studies exist, including Curran and Pearson (1971), studies concerning thundersnow-speciﬁc dynamics are few. This thesis will include an extensive 30-year climatology of thundersnow events for the period of

3 1961-1990. The work related to this thesis will also concentrate on the precise dynamics that were involved in the production of thundersnow in four events during the 1999-2000 winter season. The events occurred on 5 December 1999, 9 December 1999, 11 March 2000, and 19 April 2000. Reports of thundersnow in these events were observed at McConnell Air Force Base (IAB) near Wichita, Kansas, Lubbock (LBB), Texas, Ellsworth Air Force Base (RCA) near Rapid City, South Dakota, and St.Louis (STL), Missouri (including several other stations in the metropolitan area), respectively.

1.1 Statement of Thesis

The phenomenon of thundersnow is largely an understudied topic, and this work will serve as one of the ﬁrst to address the speciﬁc dynamic aspects that produce thundersnow. The goals of this thesis are to establish that:

1. Thundersnow occurs most commonly in two locations, in the northwest- sector of occluding cyclones in association with trowal-related frontogenetical forcing, and ahead of a warm frontal zone in a region of maximum frontogenesis in the middle troposphere.

2. CSI will be more prevalent to the northeast of a surface cyclone due to strong anticyclonic shear, while CI will be more prevalent within the occluded-sector.

4 Chapter 2

Literature Review

As previously stated, there have been few studies that have speciﬁcally focused on thundersnow and its dynamics. However, there have been many studies focusing on mesoscale heavy precipitation bands and the underlying instabilities and major sources of dynamic forcing associated with them. As a result of the

meso- banded structure of typical thundersnow events, these studies provide a framework by which thundersnow research can be conducted. This chapter will summarize several of these studies that outline the basic theories that pertain to banded precipitation structure. The studies will be presented in chronological order.

2.1 Studies on Banded Precipitation

The modern concept of CSI was presented by Bennetts and Hoskins (1979). Sym- metric instability was originally used to describe instabilities that arose from “symmetric meridional perturbations of a circular vortex”, and are oriented along the

thermal wind in regions of large horizontal temperature gradients such as frontal

> 0

zones. In regions that are statically stable (N ), where

g @

= ;

N (2.1)

@z 0

5 symmetric instability exists where surfaces are more upright than absolute vorticity vectors. However, conditional symmetric instability cannot occur without the addition of latent heat, which will be released only in a saturated atmosphere. Therefore, the “condition” of CSI is that gravitational instability can be produced by differential motions only in a saturated environment in which the wet bulb po-

tential vorticity,

) r ; q = f ( w

w (2.2)

(where w is the wet bulb potential temperature and is the three-dimensional ab-

solute vorticity vector) is less than zero in a statically stable environment. Now, if

the slope of the w surfaces is greater than that of the absolute vorticity vector, then CSI is present and air parcel ascent will be in a slantwise direction. The absolute

vorticity vectors are also representative of contours of the absolute geostrophic mo-

M = fx + V g mentum, g . In addition, Bennetts and Hoskins (1979) also suggested a three-stage process for the formation of mesoscale rainbands.

1. Parcels move northward and are lifted by a baroclinic wave while the wet bulb potential vorticity decreases due to moisture gradients along the thermal wind.

2. Parcels are lifted to saturation, creating CSI which is manifested by rolls along the thermal wind, leading to a banded cloud structure.

3. As the rolls grow, conditional gravitational instability is created in the middle troposphere.

Bennetts and Sharp (1982) directly applied the concepts discussed in Bennetts and Hoskins (1979) to the prediction of mesoscale frontal rainbands. Moreover, they were interested in the relationship of CSI and CI present in the middle troposphere to precipitation bands generated by cross-frontal ascent. By calculating a

quantity known as the growth rate parameter,

g r r

2 2 w

= f + ;

(2.3)

@ =@ z

0 w

where is the vertical absolute vorticity vector, f is the Coriolis parameter, and r

2 is the horizontal gradient operator, in regions where the relative humidity is 2 greater than 80 percent, they were able to demonstrate a relationship between and the likelihood for rainbands to develop. Their results, which deﬁnitively pro- vided a link between CSI or CI and banded precipitation, are summarized as fol-

lows:

2 2 >

1. if 0.2 h , then frontal precipitation will almost certainly be banded;

2 2 2

2. if 0.2 h > 0.0 h , then there is a 75 percent chance of banding;

2 2 2

3. if 0.0 h , then is of little signiﬁcance as an indicator of banded

2 2

precipitation, unless is greater than -0.1 h , when the probability of

banding is 60 percent. 2

Also, positive values of are indicative of either convective or inertial instability, and negative values indicate CSI. Sanders (1986) performed a case study on a rapidly intensifying cyclone off the New England coast that occurred on 5-6 December 1981. The system was strongly occluded at the surface and a thermal ridge was wrapped cyclonically to the north of the low center. Radar imagery showed a multiple-banded reﬂectivity structure

in the storm’s northwest quadrant within the thermal ridge. An examination of

M g frontogenesis fields and w - cross sections shows that frontogenetical forcing aided in the release of CSI early in the study period, and then CI later in the period. The change from a CSI to a CI environment occurred with the approach of the mid-level cyclone, and as the flow changed from being supergeostrophic (anticyclonic) to subgeostrophic (cyclonic). Also, as the cyclone intensified, the level

7 of strongest frontogenesis descended with time. Moreover, this study exempliﬁed the importance of instability and frontogenesis to the production of heavy precipitation underneath the thermal ridge in the comma-head of an occluded cyclone. Similar results were obtained in a case study of a Midwestern thundersnow event by Moore and Blakely (1988). The event was centered on St. Louis, Mis-

souri, and occurred on 30-31 January 1982. A narrow band of heavy snowfall (> 10 in. or 25 cm) fell within the northern extent of a broad band of very heavy precipitation to the north of a surface and 850 hPa warm front, and under a region

of 300-hPa divergence and anticyclonic shear. Cross sections of Miller frontogen-

M g esis and e - show that heavy snowfall, accompanied in many areas by thunder

and lightning, occurred within a region of intensifying 500-850 hPa frontogenesis,

1 coinciding with the most negative kinematic omegas (< 4 bs ), that aided in the release of borderline CSI/CI near 825 and 600 hPa. Also, regions of frontoge-

nesis coincided with a direct thermal circulation, and were aided by intensifying ~

convergence of Q, where

~ ~

@ V @ V

g g

^ ^

Q =( i; r j ):

r (2.4)

@x @y

Moore and Lambert (1993) concentrated on the use of EPV to locate regions of

CSI. They deﬁned CSI as occurring when e surfaces are steeper than surfaces of M

g in regions where the relative humidity is greater than 80 percent (Fig. 2.1). CSI

is favored in regions of large vertical wind shear, large anticyclonic shear, and low M

static stability. Large vertical shear will ﬂatten g surfaces because of increased

~ V

g gradients, while differential advections can alter the stability proﬁle by forcing

e surfaces closer to the vertical. Anticyclonic shear will cause to approach zero,

creating a weak, inertially stable zone. EPV is deﬁned as

EP V = g ( r );

e (2.5)

8 where g is the gravitational constant and is the three-dimensional vorticity vector.

Assuming geostrophic ﬂow and ignoring terms with ! and variations with respect

to y , then the equation can be expanded to become

@M @ @M @

g e g e

= g [( ) ( )];

EP V (2.6)

@p @x @x @p

AB CD

6 2 1 2

10 where the ﬁnal units are given as 1 m Ks kg or 1 potential vorticity unit (PVU). In an otherwise stable environment, negative values of EPV correspond to regions of CSI, while positive values correspond to regions of conditional sym-

metric stability. If we take the equation term by term, term A is usually negative

~ ~

V M V

g g because g normally increases with height, and is proportional to . Increased

vertical shear will make term A more negative. Term B is normally positive be- M

cause the deﬁnition of g requires that cross sections be oriented normal to the geostrophic or thermal wind. Therefore, the product of terms A and B will usually

be negative. Term C is usually positive except when there is a sharp decrease in

V x

g in the positive direction. Term D, also known as the convective stability term, is negative under conditions of convective stability and positive under conditions of CI. When term D is near zero or is slightly negative, then CSI is more likely to occur. One note of caution is that EPV will be negative in either a CSI or a CI environment. Because doubling rates for CI are signiﬁcantly greater than those with CSI, CI will also dominate over CSI and would likely produce more intense precipitation rates. Doubling time is the period required for a convective element to double its

depth. Upright convection has a doubling time on the order of minutes, while

slantwise convection has a doubling time on the order of hours. As a result, a e cross section will also have to be constructed to ensure that the environment is convectively stable.

g Figure 2.1: Relationship of e (dashed, K) and M (solid, m s ) in the evaluation of CSI, CI, and WSS (Reproduced from a ﬁgure originally constructed by Dr. James T. Moore and Sean Nolan at Saint Louis University).

McCann (1995) expanded the equation for EPV into its full, three-dimensional form. This eliminated the need to orient cross sections normal to the thermal wind,

a task that can be quite difﬁcult when examining well developed systems with M

strongly curved ﬂow. The three-dimensional form is beneﬁcial because g is no

longer present in the EPV formulation. The three-dimensional EPV is expressed as

@ @v @ @u @v @u @

e g e g g g e

EP V = g [ ( + f ) ]:

k (2.7)

@x @p @y @p @x @y @p

With this equation, CSI/CI can be evaluated by plotting EPV and e on a cross

M g section, thus eliminating the need to compare the slopes of e and surfaces. Moore et al. (1998) examined a thundersnow event that produced a band of snowfall of as much as 16 inches across northern Iowa from 20-22 February 1993. Synoptically, the region was 375 km to the north of a quasistationary frontal boundary stretching across northern Missouri. An examination of Fig. 2.2 shows that the LI’s of several parcel displacements at North Omaha (OVN), Nebraska, decreased

10 Figure 2.2: Lifted indices for various air parcels for North Omaha, Nebraska (OVN) at 0000 UTC on 21 February 1993 (Reproduced from Moore et al. 1998).

dramatically above 804 hPa, indicating that an elevated layer of CI was positioned

above a strongly stable layer. A log-P plot of e at OVN (Fig. 2.3) shows two layers of CI from 700 to 651 hPa and from 578 to 540 hPa embedded within a broad layer

of neutrality. Moore et al. (1998) calculated a most unstable CAPE (MUCAPE) at

OVN of 89 J kg 1 at 612 hPa, which is in line with the ﬁndings of Colman (1990).

1 This helped to produce a maximum vertical velocity of -4 bs on the 300 K surface. Moreover, Moore et al. (1998) conclude that this case demonstrates that CI released by isentropic uplift can produce mesoscale bands of heavy precipitation. Weismuller and Zubrick (1998) performed case studies on two heavy snow events (without thunder) that occurred in the Mid-Atlantic region on 26 Febru-

ary 1993 and 30 January 1995. Both events were characterized by meso- bands of locally heavy snowfall and a banded precipitation structure on WSR-88D radars throughout the region. Synoptically, the major features include the presence of an inverted trough and cold-air damming at the surface, warm air advection (WAA) and a southwesterly ﬂow at 850 hPa, and signiﬁcant vertical wind shear. Analyses

Figure 2.3: T-log P plot of e (K) at North Omaha, Nebraska (OVN) at 0000 UTC on 21 February 1993. (Reproduced from Moore et al. 1998). of both events show that the heavy snow bands were located within the regions of the most negative EPV, in which CSI was released by mid-level frontogenesis. This suggests that these inverted trough-cold air damming scenarios are examples of a classic warm-frontal CSI scenario, because of the tendency for anticyclonic shear and large vertical shear to be present with these events. Similar case studies were examined by Nicosia and Grumm (1999) on three heavy snow events (4-5 February 1995, 14-15 November 1995, and 12-13 January 1996) in the Mid-Atlantic and Northeastern states. Using Meso Eta Model gridded

data, the two-dimensional Petterssen frontogenesis and saturated equivalent potential temperature ( es ) were calculated every three hours for each event. Each cyclone exhibited the consistent vertical structure of an evolved cyclone (i.e. stacked),

and produced meso- bands of heavy snowfall. In each case, the heaviest snow was observed in bands coinciding with regions of intensifying mid-level frontogenesis and the most negative EPV. CSI was observed in each case, although embed-

12 Figure 2.4: Conceptual model depicting the frontogenetical region and zone of EPV reduction in a developing cyclone (Reproduced from Nicosia and Grumm 1999). ded regions of CI developed with time on the warm side of the warm frontogenetical region, producing locally intense snow bursts of as much as 6 inches per hour. The authors speculate that CI in conjunction with CSI may help in the formation of a multi-banded precipitation structure. Figure 2.4 shows a conceptual model of a mature cyclone, in which the authors theorize that EPV will be most negative on the warm side of the mid-level frontogenetical region in the vicinity of the dry tongue jet (DTJ). At this point, where the cold conveyor belt (CCB) becomes juxta-

posed with the DTJ, the differential moisture advection is the greatest. This leads

M g to a steepening of e surfaces while surfaces are unaffected. As a result, EPV is reduced and may be further reduced by frontogenesis to produce mesoscale bands (Fig.2.5) As each cyclone became vertically stacked at the end of its development phase, values of EPV increased as a result of decreased vertical shear and increased inertial stability. In general, this study suggests that bands of snowfall with accumulation rates of greater than an inch per hour are more likely within deep layers of highly negative EPV that are intersected by a region of mid-level frontogenesis.

13 Figure 2.5: Proposed positive feedback mechanism between frontogenesis and the reduction of EPV (Reproduced from Nicosia and Grumm 1999).

Perhaps the most extensive study of a single thundersnow event was presented by Martin (1998a, b). This 19-20 January 1995 event occurred in the Midwest and was characterized by a 1100-km long, 65-km wide band of snowfall totaling at least 12 inches (Fig. 2.6). Martin (1998a) attributes the band, which included re-

ports of lightning and thunder (Fig. 2.7), to the interaction of a bent-back front and a narrow wedge of high e air, known as the trowal, in the occluded sector of the cyclone. The bent-back front acted as a warm baroclinic zone and, coupled with a deformation zone, acted as a source of pronounced frontogenesis. Analysis of 200-m frontogenesis ﬁelds shows the process of frontal fracture as time progressed (Fig. 2.8), in which the warm frontogenetical region to the northwest of the low center moved farther from the cold frontogenetical region. The warm frontogene-

sis region was the site of the strongest upward vertical velocities, and was nearly collocated with the region of heavy snowfall. Cross sections of e and EPV reveal that weak symmetric stability (WSS) was present in the lower to middle troposphere over the snowfall region. While a layer of CI was present at the 5-km level, relative humidities were well below saturation. Therefore, it is very unlikely that

any CI would have been released. ~ Martin (1999) used Q to evaluate quasigeostrophic (QG) forcing within the trowal airstream. Trajectory analysis of several occluded cyclones shows that the

14 Figure 2.6: Observed 24-hour snowfall totals (cm) ending at 0000 UTC on 20 Jan- uary 1995. (Reproduced from Martin 1998b).

Figure 2.7: Cloud-to-ground lightning strikes in a 24-hour period ending at 0000 UTC 20 January 1995 (Reproduced from Martin 1998b).

15 Figure 2.8: (a) 6-hour forecast of 200-m frontogenesis from the UW-NMS model

valid at 0600 UTC 19 January 1995. Shaded regions denote positive frontogenesis

1 1 every 1 K (100 km) day . (b) As in (a) except from a 12-hour forecast valid at 1200 UTC on the same day. (c) As in (a) except from a 18-hour forecast valid at 1800 UTC. (d) As in (a) except from a 24-hour forecast valid at 0000 UTC 20 January 1995. (e) As in (a) except from a 30-hour forecast valid at 0600 UTC 20 January 1995. (Reproduced from Martin 1998a).

trowal airstream originates in the warm sector and is then cyclonically wrapped to

the north of the low center, acting as a source of signiﬁcant ascent in the process ~

and focusing precipitation beneath it. Q, which can be used to accurately represent

~ ~

! Q Q s

the QG equation, can be divided into its component terms ( n and ) (Keyser

~ Q

et al. 1988) in a natural coordinate system (Fig. 2.9). n is proportional to the

magnitude of QG frontogenesis and is deﬁned as

Q r

Q =( )^n:

n (2.8)

jr j

~ ~

Q 2rQ n

Positive values of n convergence (- ) are indicative of forcing for ascent as a

~ Q

result of frontal scale processes. s is representative of the rotation of the thermal

ﬁeld and the contributions of synoptic scale processes to forcing for ascent (Fig.

~ Q

2.10). s is deﬁned as

^ ^

Q (k r ) k r

( ): Q =

s (2.9)

jr j jr j

16 ~ Figure 2.9: Schematic showing the natural coordinate partioning of Q. Dashed

lines depict isentropes on an isobaric surface (Reproduced from Martin 1999).

~ ~

Q 2rQ s Positive values for the convergence of (- s ) are indicative of large-scale (synoptic scale) forcing for ascent. Martin‘s (1999) analysis of three occluded cy-

clones shows that ascent within the trowal in the occluded sector occurs primarily

~ Q

due to the convergence of s (Fig. 2.11), thus indicating synoptic scale QG forcing. Schultz and Schumacher (SS); (1999) present a comprehensive overview on many of the concepts and diagnostic methods that relate to symmetric and upright instability. Included in this overview is a comprehensive review of deﬁni-

tions of several types of instabilities (Fig. 2.12). (Note that SS uses the term MPV,

not EPV.) The major contention of SS is that es must be used in order to accurately

w es

diagnose CSI, not e or , because assumes that the environment is saturated

everywhere. This includes its insertion into the EPV equation to replace e , thus

making the term EPV . Otherwise, SS assert that the use of e actually diagnoses potential symmetric instability (PSI). However, we assume that if PSI occurs in a moist region where the relative humidities are greater than 80 percent (Bennetts and Sharp 1982) , then it in effect becomes a CSI environment.

~ Q

Figure 2.10: The effect of s on horizontal thermal structure. (a) Isentropes (solid ~

lines) in a ﬁeld of Q, with the maximum region of convergence shaded. Dashed

line indicates convergence axis, while r depicts thermal gradient. (b) Thick black

arrow depicts original direction of r , while thick gray arrow depicts the direction

r Q

of after being rotated by s . (c) Oriention of thermal zone in (a) after being

~ Q

rotated by s (Reproduced from Martin 1999).

16 1

Q 5 10

Figure 2.11: (a) Convergence of s contoured and shaded every mkg 1

s in the 600-900 hPa layer from an 18-hour forecast of the UW-NMS model valid

~ Q

at 0600 UTC 23 October 1996. (b) As in (a) except for n (Reproduced from Martin 1999).

Figure 2.12: Deﬁntions of several types of instabilities (Reproduced from Schultz and Schumacher 1999).

19 Schultz (1999) also addressed the production of lightning and thunder in lake- effect snowstorms. His hypothesis is that the generation of lightning is dependent upon the presence of relatively higher temperatures and dewpoints in the lower troposphere than those observed in non-thunder producing systems. This may be because the -10o C isotherm charging region is at sufﬁcient elevation to increase low-level vertical velocities to point of being able to separate charge.

2.2 Thundersnow-related Climatology

Curran and Pearson (1971) found 76 reports of TS throughout the contiguous United States for the general time period of 1968 to 1971. After selecting crtieria that a report of TS had to be within 90 nautical miles of a radiosonde station, and that the report had to be within 3 hours of the time of the sounding, a mean proximity sounding was constructed for the remaining 13 reports of TS. The mean sounding (Fig. 2.13) shows that the atmosphere is very moist thoughout the entire depth of the column. A shallow inversion (81 hPa deep), is located just above the surface. Just above the inversion from 800 to 600 hPa, the sounding is very close to moist adiabatic. Colman (1990) presents the results of a four-year climatology of cold sector, elevated thunderstorms, of which thundersnow is a particular type. Colman (1990) identifies three major factors that facilitate the development of elevated thunderstorms - a cold upper troposphere, rapid heating and moisture advection near the boundary layer, and a strongly baroclinic environment. He observed that elevated thunderstorms are favored to the northeast of a surface cyclone in advance of a shallow frontal zone, in a region of cyclonic flow at 850 hPa, anticyclonic flow at 500 hPa, and strong vertical shear. Not surprisingly, the mean Showalter Index (SI) in such events was significantly lower than the mean Lifted Index (LI), 0.6 and 7.4 respectively,, which is indicative of strong inversion layer. One of the most

20 Figure 2.13: Mean proximity sounding for 13 thundersnow reports during the period of 1968-1971. The thick, solid line depicts the temperature proﬁle. The thick, dashed line depicts the dewpoint proﬁle (Reproduced from Curran and Pearson 1971).

21 Figure 2.14: Number of hours of thunder at temperatures below 0 C from 1982- 1990. (Reproduced from Holle et al. 1998). interesting of Colman’s (1990) ﬁndings is that elevated thunderstorms frequently develop in the presence of very small MUCAPE, where the environment is essentially neutral, but in a region of large horizontal stability gradients. A comprehensive climatology of thundersnow and other types of cold weather thunderstorms was conducted by Holle et al. (1998), and encompassed the period of 1982-1990. An examination of Fig. 2.14 shows four preferred regions for the occurrence of thundersnow - the Basin and Range province in the southwestern U.S., the Central Plains, the Great Lakes region, and coastal regions of the Mid-Atlantic and New England states. This study suggests that snowfall rates may increase when accompanied by thunder and lightning, and demonstratively illustrates that the heavy snowfall is more likely when temperatures fall below 5 C, thus making the forecasting of thundersnow quite important.

2.3 Summary

In summary, the preceding review points to a strong correlation between the occurrence of heavy snow bands and the production of strong ascent by frontogenetical

22 forcing. Thundersnow can be produced when CSI or CI is released by frontogenesis, or in the presence of WSS coupled with intense forcing for upward motion. Thundersnow appears to be favored in two regions - to the north of a warm baroclinic zone in association with an open-wave cyclone and southwesterly 700-850 hPa ﬂow (Holle et al. 1998, Weismuller and Zubrick 1998, Nicosia and Grumm 1999) due to the release of CSI, and within the trowal airstream in the northwest or occluded sector of a mature cyclone (Martin 1998a,b) where no type of instability has been shown to be preferred.

23 Chapter 3

Methodology

3.1 Climatology

Data in the form of hourly surface airways (SA) observations were obtained from a CD-ROM of 223 ﬁrst-order stations in the contiguous United States covering the period of 1961 to 1990. Software was written to search for all possible combinations in which thunder and snow were reported simultaneously. These observations of thundersnow were then separated into three geographical regions (East, Central, and West) that are outlined in Fig. 3.1. Using the Digital Atmosphere meteorological software package from WeatherGraphics Technologies, surface maps were produced for each observation of thundersnow (TS) based on a Barnes analysis scheme. The reports were then segregated into events using the following criteria to deﬁne an event:

1. Reports of TS at a single station cannot be separated by more than six hours, or reports from more than one station in the same geographical region cannot be separated by more than six hours.

2. The report of TS should be within 1100 km of the center of a parent low (if present), or be within the cloud shield and region of cyclonic ﬂow around the parent low.

3. If TS was reported by two stations for the same general time period,

24 Figure 3.1: Delineation of regions used in the thundersnow climatology.

then the distance between them must be no more than 1100 km.

(Note: The distance of 1100 km was chosen arbitrarily and is the midpoint between

the low and high ends of the meso- spatial scale.) Each event was then referenced according to its initial report of TS, and then again segregated by the time of day (LST), month, year, intensity of snowfall, geographical region, temperature, dewpoint, sea level pressure (SLP), wind speed and direction, and meteorological direction and distance (km) from the parent low. The direction and distance readings were obtained using a simple screen utility available with Digital Atmosphere. Events were then placed into one of the following seven categories:

1. Associated with a surface low with an identiﬁable position,

2. Lake-effect snow (Great Lakes or Great Salt Lake),

3. Upslope regime,

25 4. General orographic,

5. Associated with a surface low in the Atlantic or Paciﬁc Ocean of uniden- tiﬁable position,

6. Associated with a surface low in Mexico, Canada, or the Gulf of Mexico of unidentiﬁable position,

7. Miscellaneous or unclassiﬁable

Categories 2,3,and 4 include events that are not within the cloud shield or wind ﬂow of a surface cyclone. Category 3 events include those in which the winds were easterly to southeasterly for most of the stations in the Rocky Mountain region that surround a particular TS report. Category 4 events tended to have a westerly ﬂow associated with them. A state-by-state count of thundersnow events was also conducted. If an event encompassed several states, such as Oklahoma, Arkansas, and Missouri, then it was counted as an event for each of those states. However, the event would still be counted as one in the overall count. The events were also normalized according to

the land area of the average state (61,657.18 mi.2 ) in the contiguous U.S. as follows:

(A)= V ;

n (3.1)

61; 657:18mi:

V A

where a is the actual number events for a given state, is the land area for that V

state, and n is the normalized value for that state.

3.2 Case Studies

Five TS events were examined in depth in order to determine the most signiﬁcant reasons for occurrence of thunder. More speciﬁcally, each case was examined in order to determine if any underlying instabilities were present and, if so, what type. The major forcing mechanisms responsible for enhanced ascent were also

26 identiﬁed using isobaric, isentropic, and QG techniques. Even though the goals of each examination are the same, the methods by which these were accomplished differed as a result of the types of data that were available on an individual basis. As a result, the three methods by which the ﬁve cases were examined will be presented in the subsections that follow.

3.2.1 9 December 1999, 11 March 2000, and 19 April 2000

Analyses of these events were accomplished via the use of GEMPAK (GEneral Meteorological PAcKage) on a Sun Ultra60 UNIX workstation. RUC initial ﬁelds and hourly surface data for the time of the observation of TS were utilized for the purpose of diagnosing the cause of thunder in each case. The RUC was chosen

because of its mesoscale resolution (40 km on a 151 113 grid) and because initial ﬁelds are available on an hourly basis. Also, papers by Schwartz et al. (2000) and Smith et al. (2000) have revealed that the RUC is quite effective at resolving parameters that relate to convective activity. Hourly METAR reports were plotted for the time of TS occurrence, and temperature (every 5 F) and pressure (every 2 hPa) were analyzed using both subjec- tive and objective means. The objective analysis was accomplished using a Barnes

scheme on a 55 55 grid with a 75 km grid spacing. The locations of fronts and pressure centers were analyzed subjectively and drawn by hand. When possible, mandatory upper air data (TTAA and TTBB) were plotted for the time of the TS observation and analyzed both objectively and subjectively. Otherwise, RUC initial ﬁelds were contoured. Using GEMPAK’s program for vertical interpolations (GDVINT), RUC isobaric ﬁelds were converted into isentropic coordinates, and kinematic omegas were calculated, where

1000

= T ( )

; (3.2) p

and

! = ! (p ) (rV )dp:

ref (3.3)

p ref

is the temperature a parcel would possess if it were displaced vertically to 1000

T R

hPa, is the temperature at the initial pressure level (p), d is the dry gas constant,

C p ref

p is the molar heat capacity at constant pressure, and is the pressure at some

p ! p ref reference level. ref is usually assumed to be the surface and ( ) is usually

assumed to be zero. This method follows that proposed by O’Brien (1970).

Plan view analyses of e , which is conserved for moist processes, were per-

formed on the mandatory levels in order to locate any possible trowal structures.

e is expressed as

L r

v s

); = exp ( LC L

e (3.4)

C T

p LC L

L v

where LC L is the potential temperature at the lifting condensation level, is the

r T LC L latent heat of vaporization, s is the saturation mixing ratio, and is the tem-

perature at the LCL.

~ ~

= Q Q n Petterssen surface frontogenesis ( ), s (see Equation 2.8), and (see Equa-

tion 2.9) were calculated for each mandatory level in order to diagnose forcing (A

~ ~

Q Q n

GEMPAK script was written in order to calculate s and ), where

= C jr j(E cos 2 D );

= (3.5)

1:08 10 E D

and C is a unit conversion ( ), is the total deformation, is the diver-

V gence (r ), and is the angle between the axis of diliation and the isentropes.

Upper tropospheric influences were ascertained via the analyses of the 300 hPa divergence fields, while traditional analyses of a were also performed. Tropopause-related influences were assessed by the analyses of dynamic tropopause pressures and the 300-700 hPa potential vorticity in isobaric coordinates, which is

expressed as

@~v @

PV = g [ + f +(k )] ;

(3.6)

@ @p

where is the mean layer vorticity. !

Cross sections were constructed for = and in order to assess the impact of

es g

frontogenetical forcing more directly, and for , e , , relative humidity (RH), M , D

and the three-dimensional EPV (EPV3 ) to assess stability. When possible, cross

es sections were oriented normal to the thermal wind so that the slopes of e / and

Mg could be compared in an assessment of PSI/CSI. However, this is not necessary

D 3D

when EPV3 is being utilized. EPV (see Equation 2.7) was calculated via a GEM- D PAK script that was written to compute EPV3 at 50 hPa increments from the 900 hPa to 200 hPa levels. Regions of PSI (SS) that were located within regions with RH of greater than 80 percent are assumed to CSI, using the guideline established by Bennetts and Sharp (1982). If the region of PSI, or any other instability, was located within a region in which the RH was less than 80 percent then the instability was assumed to be untapped, and not a likely contributer to the production of TS.

When possible, stability indices for a nearby sounding (< 50 km from the TS observation) were examined in order to further assess the possibility of gravitational instability, otherwise RUC analyses of these parameters (such as the Lifted Index, Showalter Index, etc.) were contoured from RUC initial ﬁelds. Isentropic analyses of pressure were plotted on isentropic levels along with storm-relative moisture transport vectors in order to ascertain isentropically-related

vertical motion. A level was chosen via the examination of cross sections, to de-

29 termine where the RH was greater than 80 percent, and where the vertical motion

was the strongest in a column above the location of the TS observation. Moisture

~ V

transport vectors (q ) give an indication of not only the ﬂow of moisture, but of

~ ~

moisture convergence as well. The storm-relative component (V ) (Moore et ~ al. 1998) indicates the ﬂow relative to the motion of the system (C ), which was

considered to be the same as the motion of the absolute vorticity maximum that ~ was associated with the parent cyclone. C was calculated as the storm motion over a period of 6 hours.

3.2.2 5 December 1999

Much of same procedure that was administered in Subsection 3.2.1 was used with this case. However, because GEMPAK compatible RUC ﬁelds were not available,

the PCGRIDDS software package was used instead. The RUC ﬁles were obtained

es D

from the National Weather Service Ofﬁce in Pleasant Hill, Missouri. EPV3 , ,

q~v g and could not be calculated, so a e -RH-M cross section was used to assess CSI/PSI. Surface analyses were generated using GEMPAK.

30 Chapter 4

Thundersnow Climatology

4.1 Spatial and Temporal Patterns

The search for TS during the period of 1961 to 1990 uncovered 563 reports for the available data set. Using the criteria previously discussed to identify a TS event, some 375 were established. While TS events certainly occurred within gaps between observing stations in the data set, the 30-year period of record establishes an accurate climatology for the basic synoptic spatial and temporal characteristics of TS. Of the major regions that were identified in Fig. 3.1, the Central Region showed the greatest preference for the occurrence of TS with 164 events, while the East had the fewest (Fig. 4.1). On a smaller scale, Utah and Nevada showed the greatest preference for TS occurrence (Fig. 4.2), while secondary areas of preference were observed over the North Central Plains and the Northeast. While Category 1 type systems were the most common in all three major geographical regions (Fig. 4.3), there are undoubt- edly terrain influences that effect TS occurence, especially in the West and East Regions. Category 3 and 4 events are not uncommon in the West Region, and illustrate the importance of orographic lift to the production of TS in the West. As a result, it is likely that orography plays a significant role in the production of TS in Category 1 events as well, especially in the Basin and Range of Utah and Nevada where mountians are aligned at what are typically high angles to the prevailing

31 Figure 4.1: Number of thundersnow events for each region during the period of 1961-1990. Refer to regions depicted in Fig. 3.1. westerly ﬂow. The presence of large, relatively warm bodies of water (Great Salt Lake and the Great Lakes) also inﬂuence TS occurence in all regions. This is not only due to pure lake-effect snowstrorms, but lake-enhanced vertical motion in Category 1 events as well. A land area-based normalization of the values expressed in Fig. 4.2 reveals a somewhat different spatial pattern (Fig. 4.4). The primary maximum now shifts to the eastern seaboard from the Mid-Atlantic to the New England region, although the values for smaller states, such as Rhode Island and Delaware, are falsely elevated due to their extremely small land areas. An analysis of surface maps of these events reveals that this occurs mainly as a byproduct of intense coastal cyclones. The spatial patterns over the Central and West Regions remain largely un- changed. However, one notable exception is Texas, where normalization reveals a

32 Figure 4.2: Number of thundersnow events for each state during the period of 1961- 1990. Note that the sum of the events will not equal 375 (Refer to Section 3.1).

33 Figure 4.3: As in Fig. 4.1 except events (N=375) by category (Refer to Section 3.1 for the description of categories). much smaller value due to its very large land area. It should be noted that nearly all TS in the state of Texas occurs in the Panhandle, a very small part of the state as a whole. As a result, the Texas Panhandle should still be considered a preferred region for the occurence of TS. Temporally, TS is most common during the month of March in all regions (Fig. 4.5), when synpotic scale sytems are typically the strongest due to strong temperature gradients and upper tropospheric winds. This preference was the strongest in the Central Region, where Wisconsin experienced more TS events than any state for any month (not shown). A weak, secondary temporal maximum was observed for November and December over the East and West regions. In the East, this is the peak of the lake-effect snow season, when the Great Lakes remain largely un- frozen. In the West, this coincides with the transition into the Paciﬁc ‘wet’ season. The analysis of initiaton times for TS events shows a slight diurnal trend for all regions (Fig. 4.6), in which the majority of events occur between 1200 and 2400 Local

34 Figure 4.4: As in Fig. 4.2 except with normalized values.

Standard Time. There is no clear explanation for this since TS is nearly always a result of elevated convection, which should not be affected by diurnal heating. TS also usually occurs in essentially cloudy region with a stable PBL.

4.2 Characteristics of Thundersnow Observations

After the initial observations of TS events were plotted and analyzed, the data illustrates that TS occurs primarily in the the northeastern or northwestern quad- rants of a parent cyclone, if present (Fig. 4.7). This is consistent with the earlier hypothesis of where TS is likely to occur. The initial TS observations occurred at a mean distance of 471 km from the parent low, and showed by far the greatest fre-

quency at distances of 200 to 600 km from the low (Fig. 4.8). For all initial reports, the mean temperature was 31.5 F, the mean dewpoint was 28.5 F, the mean sea

level pressure was 1008.1 hPa, and mean wind was from 1 at 14.3 knots. These

35 Figure 4.5: As in Fig. 4.1 except by month (N=375).

Figure 4.6: As in Fig. 4.1 except by Local Standard Time (LST) (N=375).

36 Figure 4.7: Polar plot showing the location of thundersnow events (Category 1) relative to the position of the center of the parent low (N=247). Direction is given in the traditional meteorological azimuth (degrees) from the postion of the low to the observing station. Distances are given in km. values are represented in the station plot in Fig. 4.9. An analysis of snowfall intensity (Fig. 4.10) shows that light snowfall rates are the most common, but moderate to heavy snow intensities are certainly not uncommon.

37 Figure 4.8: As in Fig. 4.1 except by distance (km) from parent low center of initial report (N=375).

Figure 4.9: Representative station model for the average initial report for all thundersnow events (N=375). Temperature and dewpoint are given in degrees F, the wind speed in knots, and the sea level pressure in hPa. The standard deviation for all parameters is also represented.

38 Figure 4.10: As in Fig. 4.1 except by snowfall intensity of initial report (N=375).

39 Chapter 5

Case Studies

5.1 5 December 1999

5.1.1 Introduction

This event was highlighted by TS at McConnell Air Force Base, Kansas (IAB), occurred at 0600 UTC on 5 December 1999, and included several reports of distant lightning by stations in south central Kansas. Snowfall of 4 to 10 inches fell within a 40-mile wide band extending from southwestern Oklahoma to western Iowa. The RUC 6-hour forecast valid at 0600 UTC included the possibility of convective precipitation near IAB (not shown).

5.1.2 Surface Analysis

A routine surface analysis for 0600 UTC on 5 December 1999 (Fig. 5.1) indicates a closed 1009-hPa closed cyclone over east central Missouri. A cold front trails from the low across southern Missouri across western Arkansas into eastern Texas. A 1038-hPa anticyclone was centered over southeastern Idaho and was pulling very cold, polar air into much of the Plains and Rocky Mountain Region, and tighten- ing the pressure gradient across parts of Kansas. Rain showers were observed in advance of the cold front, while a broad area of warm-frontal type rainfall was occurring from northeastern Missouri into the Great Lakes Region. A narrow band of snow, including a report of TS at IAB, extended west of the low from southwestern

40 Figure 5.1: Surface analysis valid at 0600 UTC on 5 December 1999. Isobars drawn every 2 hPa. Red line depicts cross section line.

Oklahoma into south central Kansas (Fig.5.2).

5.1.3 Upper Air Analysis

An examination of RUC initial ﬁelds for the mandatory levels reveals that the system still possessed a baroclinic tilt with height, indicating that the cyclone was still in its developmental stage. The 850-hPa analysis (Fig. 5.3) shows a 1430-m closed circulation that is centered near the Kansas-Missouri border. Northeast- erly winds were producing warm air advection (WAA) over south central Kansas, while cold air advection (CAA) was occurring over nearly all of Oklahoma. At 700 hPa (not shown), the low was centered slightly to the west, over southeastern Kansas. Again, weak northeasterly ﬂow was producing WAA near IAB. At 500

hPa (not shown), the low was centered near the Kansas-Oklahoma border, not far

5 1 from IAB. A vorticity maximum (34 10 s ) was positioned over northeastern

41 Figure 5.2: Radar mosaic valid at 0600 UTC on 5 December 1999 (obtained from the National Climatic Data Center).

Oklahoma. At 300 hPa (Fig. 5.4), the low (still closed) was centered over south central Kansas. A 120-knot jet streak was located over eastern sections of Oklahoma and Texas, placing IAB under the left exit region of a curved jet streak, and under

an axis of maximum 300-hPa divergence (Fig. 5.5) in which the highest values (9

5 1 10 s ) were found to be over eastern Kansas. Therefore, upper tropospheric di-

vergence was likely a very signiﬁcant factor in the production of ascent near IAB.

5 1

Concurrently, signiﬁcant convergence (-8 10 s ) was occurring in the lower troposphere (not shown), helping to further enhance upward vertical motion by the process of Dyne’s compensation.

5.1.4 Isentropic Analysis

Analysis of e contours for the 0600 UTC RUC initial ﬁeld shows a well deﬁned trowal structure at the 700-hPa level (Fig. 5.6) that extends from Missouri into

42 Figure 5.3: 850-hPa Rapid Update Cycle initial ﬁeld analysis valid at 0600 UTC on

5 December 1999. Isotherms (thick lines) drawn every 4 C, height contours (thin lines) every 2 dkm.

Figure 5.4: As in Fig. 5.3 except for 300-hPa level. Height contours (thick lines) drawn every 5 dkm, isotachs (thin lines) every 10 knots (2=20 kts, 12=120 kts, etc.

Figure 5.5: As in Fig. 5.3 except for 300-hPa divergence (depicted every 2 10 1 s ). south central Kansas. This feature is poorly defined at 850 hPa and is completely absent at 500 hPa. However, this is not surprising since the dynamic tropopause had descended to near the 500-hPa level (Fig. 5.7). As a result, the 500-hPa analysis will be reflective of the intrusion of cold stratospheric air into the middle troposphere. Also, cyclonic potential vorticity advection will contribute to the formation of the trowal below the tropopause, as kinetic energy is created and the flow becomes more meridional (Hoskins et al. 1985). Also, if Martin’s (1999) assertion that occlusions may occur from the top down is correct, then the trowal will develop at the 700-hPa level before it does at 850 hPa. In any event, the close alignment of the 700-hPa trowal feature and orientation of precipitation echoes (Fig. 5.2) is indicative that the trowal was the major feature in the production of TS at IAB.

44 Figure 5.6: As in Fig. 5.3 except for 700-hPa equivalent potential temperature (depicted every 2 K).

250 400

200 300

200

250

200 200

200

500 200 450 200 300

350 200 400 150

991205 0600UTC Tropopause Pressure (hPa)

Figure 5.7: As in Fig. 5.3 except dynamic tropopause pressures (depicted every 50 hPa).

5.1.5 Stability Analysis

M g A e - -RH cross-section (Fig. 5.8) was constructed perpendicular to the 0600 UTC RUC initial ﬁeld 1000-500-hPa thickness analysis (not shown) from near Dodge

City, Kansas, to near Reelfoot Lake in northwestern Tennessee. While IAB was

> % located within a deep layer of high RH ( 90 ), the e contours are nearly horizontal, indicating that the atmosphere was neither symmetrically nor convectively unstable. However, an analysis of the lifted index for the 850-700-hPa layer (Fig. 5.9 shows values slightly below zero. As a result, some weak elevated gravitational instability was present over IAB, which could have easily been released by the abundant middle to upper tropospheric forcing (see next section) that was occurring. Also, the analysis of the RUC initial ﬁelds shows that the K-index was 18, and the total totals index was near 44, both indicative of a marginally unstable atmosphere.

5.1.6 Quasigeostrophic Forcing

~ ~

Q Q n

The analysis of s (Fig. 5.10) and (Fig. 5.11) at 0600 UTC in the 850-400-hPa

19 1

layer shows nearly equal amounts of convergence of quantity (5-10 10 Pa

2 1

~ Q

m s ) over eastern Kansas. s convergence is indicative of the reorientation of

the thermal ﬁeld and synoptic-scale forcing (Martin 1999), and is important to the

~ Q

formation of the trowal feature. In this case, s convergence likely results from

strong upper tropospheric inﬂuences and potential vorticity conversion related to

~ Q

the tropopause anomaly. n convergence is indicative QG frontogenetical forcing within the trowal feature itself. Therefore, these two major mechanisms of forcing acted in tandem to release the weak upright instability that was present over IAB.

46 Figure 5.8: Vertical cross section analysis of the 0600 UTC Rapid Update Cycle

initial ﬁeld from 5 December 1999. Equivalent potential temperature (dashed lines)

are depicted every 3 K, Mg (solid lines) depicted every 5 kg m s , and relative

> % humidity (dotted lines) depicted every 5% for values 80 . KIAB is located near

97.25 W longitude.

5.1.7 Conclusions

The TS in this event occurred within the trowal of an occluding cyclone. The trowal is a region of signiﬁcant frontogenesis, and is mainly the byproduct of tremendous amounts of synoptic-scale forcing. With these two forcing mechanisms working in tandem, and with large amounts of upper tropospheric divergence, this would easily be sufﬁcient to release small amounts of upright instability and thus create

updraft speeds sufﬁcient to create charge separation. The meso- banding of the snowfall observed with this event is common to this type of synoptic setting, and is largely a byproduct of the orientation of the trowal and frontogenesis ﬁelds.

47 Figure 5.9: As in Fig. 5.3 except lifted indices for the 850-700-hPa layer (depicted

every 1 C).

~ Q

Figure 5.10: As in Fig. 5.3 except for divergence of s in the 850-400-hPa layer (de-

19 1 2 1 picted every 5 10 Pa m s ). Dashed lines depict areas of convergence, while solid lines depict divergence.

~ Q

Figure 5.11: As in Fig. 5.10 except for divergence of n .

5.2 9 December 1999

5.2.1 Introduction

This event featured a 99-minute episode of thunderstorms with sleet, that was mixed with snow at times, at Lubbock, Texas (LBB) that occurred at approximately 0900 UTC on 9 December 1999. Snow and sleet accumulation at LBB totaled 7 inches in a 5-hour period, although much of that fell after the thunderstorms had ended. Surface observations of in-cloud and cloud-to-cloud lightning were common at LBB just before and at the onset of precipitation at 0917 UTC.

5.2.2 Surface Analysis

At 0900 UTC on 9 December 1999, the surface analysis (Fig. 5.12) shows a 1010-hPa closed surface cyclone over north central Texas, not far from the Dallas-Ft. Worth metroplex. A cold front trails from the low across south central Texas and then into Mexico. At this time, snow is limited to the northern Texas Panhandle, while a more broad area of rain and thunderstorms was extending from Oklahoma into

49 Figure 5.12: Surface analysis valid at 0900 UTC on 9 December 1999. Isobars depicted every 2 hPa. Red line depicts cross section line. west central Missouri along a warm front (Fig. 5.13). A sharp temperature contrast is observed across Texas, where temperatures were in the upper sixties ( F) over central and southern portions of the state, but only in the upper twenties and thir- ties in the west (Fig. 5.12). A 1030-hPa anticyclone is centered over northwestern Colorado, pulling very cold Polar air into the Plains and Rocky Mountain regions.

5.2.3 Upper Air Analysis

The RUC initial analysis valid at 0900 UTC shows a 1460-m 850-hPa low over northwestern Texas (Fig. 5.14). As a result, WAA is occurring over most of the Pan- handle region, while CAA is occurring over the remainder of the western Texas. A 3020-m low is indicated at the 700-hPa level (not shown), with its center being located just to the south of LBB and its circulation again producing WAA over the

Panhandle. At 500 hPa, a 5560-m low is centered along the Texas-New Mexico

5 1

border, while a vorticity maximum (38 10 s ) is located just to the east of the

50 Figure 5.13: Radar mosaic valid at 0900 UTC on 9 December 1999 (obtained from the National Climatic Data Center).

500-hPa low center (Fig. 5.15). The 300-hPa initial ﬁeld analysis shows a 90-knot

jet streak emerging from the base of a trough over western Texas (Fig. 5.16), plac-

5 1 ing LBB in a region of weak divergence (2 10 s ). Analysis of the dynamic

tropopause shows a 500-hPa tropopause anomaly (Fig. 5.17) near the location of

7 1 1 3 the 500-hPa low (Fig. 5.15), which is associated with a 25 10 Km Pa s potential vorticity maximum located near the position of the 500-hPa low center.

5.2.4 Isentropic Analysis

An isentropic pressure analysis of the 294 K surface (from the 0900 UTC RUC initial field) shows a well-defined trowal feature that originates over southern Texas and then wraps cyclonically into the Texas Panhandle and northeastern New Mexico (Fig. 5.18). An analysis of storm-relative moisture transport vectors (Fig. 5.18) shows that the flow of moisture is oriented very closely to the alignment of the

51 0 1500

-5

1530 -5 5

-5 1530

-5

1560 1530

1470

0 10 10 1530 1500 15 5

1560

991209/0900V000 850 MB HGHT 991209/0900V000 850 MB TMPC Figure 5.14: 850-hPa Rapid Update Cycle initial ﬁeld analysis valid at 0900 UTC on

9 December 1999. Isotherms (dashed lines) contoured every 5 C, height contours (solid lines) every 30 m.

5 5580

5580 5 5

5 15

10 15 5 10

5580 5 35 20 30 5 255640

5700 5820 5

5760 10 0 991209/0900V000 500 MB HGHT 991209/0900V000 500 MB AVORWND (*10**5)

Figure 5.15: As in Fig. 5.13 except 500-hPa heights and absolute vorticity. Absolute

5 1 vorticity (dashed lines) contoured every 5 10 s , height contours (solid lines) every60m.

52 2 50 -2 110 2 -2

90 6 2 0 -6 -4 70 90 -2 6 2 -6 9120 2 0

-6 -2 -4 4 -6 30-2 50 0 70 4

70 2 9240 0 50 -6 70 -4 -2 9480

0 90 0 70 9360 90 2 9480 6 -6 4 0 -4 0 2 991209/0900V000 300 MB MULSPED1.94 991209/0900V000 300 MB HGHT 991209/0900V000 300 MB DIVWND (*10**5)

Figure 5.16: As in Fig. 5.13 except 300-hPa heights, divergence and isotachs.

Isotachs (solid purple) contoured every 20 knots and shaded for speeds > 70

knots, height contours (solid black) every 120 m, and divergence (dashed red)

5 1 every 2 10 s .

2505 200 150

150

15 250 200 250 250 150

250 200

200 5

0 25 150 20 500 450 400 15 350

200 5 300

991209/0900V000 SFC PRES 991209/0900V000 300 : 700 MB PVORTHTA (*10**7) Figure 5.17: As in Fig. 5.13 except dynamic tropopause pressure and 700-300-hPa

potential vorticity. Isobars (solid) depicted every 50 hPa, potential vorticity (dashed)

7 2 1 2 every 5 10 m Ks kg .

53 600

650 750

850 950 700 900

750

800

950

SR-MTVEC 100. m/s 991209/0900V000 294 K PRES Figure 5.18: As in Fig. 5.13 except pressure and storm-relative moisture transport

vectors on the 294 K isentropic surface. Isobars (solid) depicted every 50 hPa,

~ ~ ~

C C

vectors indicate storm-relative moisture transport (q(V )). calculated to be

1 from 257 at 13.7 m s . trowal, which suggests that the trowal airstream was not only the major supplier of Gulf moisture to the system, but was the principle source of isentropic uplift. Isentropic uplift was the strongest over the Texas Panhandle where the pressure gradient is strongest, and where the trowal airstream was forced upward over the warm baroclinic zone.

5.2.5 Stability and Forcing

A cross section constructed from Clovis, New Mexico (CVS), to Dallas-Ft. Worth International Airport, Texas (DFW), shows a layer of elevated CI (Fig. 5.19) that

spans the entire length of the cross section, and extends from near 750 hPa to 550 % hPa above LBB. This layer of CI begins near the top of a layer of high RH (>80 ) that extends from the surface to 700 hPa. So, there would have been adequate moisture for the release of CI to produce TS.

54 200 354 352 350 348 346 344 90 342 340 338 250 336 334 332 80 330 300 328 326 324 350 322 320 90 400 70 318 316 80 314 450 50 60 312 30 40 500 304

550 20

600 10 310 80 650 306 308 700 0 750 312 800 314 -10 316 318 320 328 304 322 850 292 294 300302 324 326 900 296 -20 298 950 991209/0900V000 RELH CVS 991209/0900V000 MSFC DFW 991209/0900V000 THTE

Figure 5.19: Vertical cross section from Clovis, New Mexico (CVS), to Dallas-Ft. Worth, Texas (DFW), from the 0900 UTC RUC initial ﬁeld valid on 09 December

1999. Solid lines depict equivalent potential temperature (every 2 K), dashed lines 1

depict absolute geostrophic momentum (every 10 kg m s ). Values of relative % humidity greater than 80% are shaded green, and greater than 90 dark green. LBB is located near the second tick mark from the left edge.

55 100 0 -4 0

0 4

150

-1

200 0 -2 -3 -4 -4 250 0 4 -5 4 128 -4 0 300 -1

350 4

400 -8 -6 4 450 32 -7 2824 20 500 16 -10 12 8 550 -9 -8 600 -8 -4 650 700 -5 -9 -8 750 800 -3 -28 -4 850 -24-20 -12-16 -5 900 4 -8 0 -4 0 4 -3 -4 128 950 -2 -2 991209/0900V000 OMEG (*10**3) CVS 991209/0900V000 FRNTTHTAUOBS (*10**1) DFW

Figure 5.20: As in Fig. 5.18 except Petterssen surface frontogenesis (solid lines,

1 1 1 1

every 4 10 K 100 km 3h ), and (dashed lines, every 3 bs .

Figure 5.20 shows that upward vertical motion was the strongest (-10 bs )at the 450-hPa level above LBB. A second axis of strong vertical motion was located near the 725-hPa level. Each of these regions of strong upward motion are associated with corresponding regions of frontogenetical maximums at the same levels (Fig. 5.20). The region of frontogenesis in the lower troposphere is associated with the trowal feature, while the region in the middle troposphere is associated with the intrusion of very cold stratospheric air in the tropopause fold. This suggests that these two regions of frontogenesis acted in tandem to release CI, and to en-

hance upward vertical motion.

~ ~

Q Q n

The computation and analysis of s and convergence for the 850-400-hPa

layer above LBB (Figs. 5.21 and 5.22, respectively) shows values of 6 10 m

1 1 14 1 1

kg s and 3 10 mkg s , respectively. A level-by-level analysis of s

~ ~

Q Q s and n convergence shows that convergence was the strongest at 500 hPa, at the height of the tropopause anomaly, and was positive for all points above the

56 0

0 0

-8

20 -4 16 12 8

4 0

-4 -4

0 0

991209(1200V00) 400:850 Layer Qs Convergence

~ Q

Figure 5.21: As in Fig. 5.13 except convergence (positive values) of s in the

14 1 1

850-400-hPa layer. Contours drawn every 4 10 mkg s .

~ Q

850-hPa level (not shown). Meanwhile, n convergence was strongest at the 850- hPa level (not shown). This suggests that large-scale processes were primarily responsible for ascent in this case, as well as for the production of the frontogenesis ﬁelds that acted as releasing mechanisms for CI. While EPV3 was analyzed (Fig. 5.23), it cannot be used as a diagnostic tool for CSI in this case because the atmosphere is not convectively and gravitationally stable. Although a large area of negative EPV3 is evident, it is only indicative that

some type of instability is present (Moore and Lambert 1993). A comparison of

es cross sections of EPV3 that were calculated using e (Fig. 5.23) and (Fig. 5.24) shows that the two are nearly identical. Therefore, either method would have worked well in this case study.

57 1 0 -1 -1

0 -1 0 1 1 0 -1 0

1 -1 0 2 1 -1 -5

0 2 -4 -1 1

2 -1 1 -3 2 5 4 -1 3 -2 -1 1 1 -2

-2 0 1 -1 -1 -1 0

991209(1200V00) 400:850 Layer Qn Convergence

~ Q

Figure 5.22: As in Fig. 5.20 except convergence (positive values) of n in the

14 1 1

850-400-hPa layer. Contours drawn every 1 10 mkg s .

200 4 3

250

6 5 2 300 4 3 -1

350 1

400

450 0 500

550

600 650

700 -1 750 -3 -2 800 1 0 2 2 850 1 3 0 900 950

CVS 991209/0900V000 EPV3 (*10**6) DFW

Figure 5.23: As in Fig. 5.18 except three-dimensional equivalent potential vorticity

calculated with e . Contoured every 1 PVU with negative values shaded.

58 200 4 3

250

6 2 5 300 4 3 1 350 0

400

450

500

550

600 -1 650 700 750 1 0 800 2 3 850 2 3 0 1 900 950

CVS 991209/0900V000 EPVTHES3 (*10**6) DFW

Figure 5.24: As in Fig. 5.18 except three-dimensional equivalent potential vorticity

calculated with es . Contoured every 1 PVU with negative values shaded.

5.2.6 Conclusions

The convective episode associated with this event occurred as elevated CI was released by frontogenesis. Frontogenesis occurred in the lower troposphere in association with the trowal, and in the middle troposphere in association with a tropopause anomaly and its associated stratospheric air intrusion. The two working in tandem helped to create strong ascent, and ideal conditions for the release of elevated upright instability. There can be little doubt that the presence of the tropopause anomaly was also important in the forcing for upward vertical motion. First, the intrusion of stratospheric air would act to lift the warmer, tropospheric air ahead of it, because the stratospheric air is more dense; this is verified by the frontogenesis fields. Sec- ondly, the low static stability associated with the anomaly would enhance the potential for stronger upward vertical velocities (Hoskins et al. 1985). Thirdly, large amounts of kinetic energy would be added to the system, due to conservation potential vorticity, causing the flow to become

59 more meridional (Hoskins et al. 1985). The latter would be important in the cre- ation of the trowal airstream.

5.3 11 March 2000

5.3.1 Introduction

This event featured numerous reports of TS throughout the greater St. Louis (STL) metropolitan area, and lightning imagery (not shown) indicates that a band of TS extended as far south as the Cape Girardeau, Missouri, area. Snowfall totals (not shown) ranged from 4 to 10 inches (10-25 cm) in a band from central Missouri into the St. Louis area.

5.3.2 Surface Analysis

The surface analysis at 1300 UTC on 11 March 2000 (Fig. 5.25) indicates a mod- est, 1010-hPa closed cyclone over middle Tennessee. A warm front extends to the east of the low while a cold front trails to the south of the low. A broad area of rain and snow with embedded thunderstorms is wrapping around the backside of the low center, while showers and thunderstorms are located in the warm-sector (Fig. 5.26). A broad anticyclone (1030 hPa) was centered over extreme western Kansas, and was helping to create a tight pressure gradient and strong winds in the northwest-sector of the cyclone (Fig. 5.25).

5.3.3 Upper Air Analysis

Analyses of the 1200 UTC RUC initial ﬁelds on 11 March 2000 shows a closed 850-hPa low circulation over southeastern Missouri (Fig. 5.27). At this time TS is beginning at STL and weak WAA is occurring at 850-hPa over STL. A closed 700- hPa low is centered over south central Missouri, while a cold pool is forming over southwestern Missouri (not shown). A 500-hPa short wave trough extends along

60 Figure 5.25: Surface analysis valid at 1200 UTC on 11 March 2000. Isobars depicted every 2 hPa. Red line depicts cross section line.

an axis from southwestern Missouri into Louisiana (Fig. 5.28), and its associated vorticity maximum is positioned over central Arkansas. At 300-hPa, a 100-knot jet

streak was located over southern Arkansas (Fig. 5.29) was helping to produce a

5 1

divergence axis (7 10 s ) over eastern Missouri (Fig. 5.29). This divergence axis is nearly colocated with a band of heavier precipitation (Fig. 5.26). In addition, a slightly elongated tropopause anomaly (450 hPa) and its associated PV anomaly are located over central Arkansas (Fig. 5.30).

5.3.4 Isentropic Analysis

Plan view view analyses of e from the 1200 UTC RUC initial ﬁeld indicates that only a very weak e ridge is present at 850 hPa (not shown). However, a hint of a trowal feature is developing at 700 hPa (not shown) and a CAA-WAA couplet has developed. The trowal feature is signiﬁcantly better organized at 500 hPa (Fig.

61 Figure 5.26: Radar mosaic valid at 1200 UTC on 11 March 2000 (obtained from the National Climatic Data Center).

0 1440 -10

1500

1530 0 0

1560

0 -5 5

1440

10 0 1500 1470

1530 000311/1200V000 850 MB HGHT 000311/1200V000 850 MB TMPC Figure 5.27: 850-hPa analysis of the 1200 UTC Rapid Update Cycle initial ﬁeld valid on 11 March 2000. Heights (solid,black) depicted every 30 m, temperatures

(dashed,red) every 5 C.

62 10

5400

5460

15 10

5 15 5

10 5 0 15 5 10 10 5520 10 35

30 5580 25 15 5 205640 5 000311/1200V000 500 MB HGHT 000311/1200V000 500 MB AVORWND (*10**5)

Figure 5.28: As in Fig. 5.27 except 500-hPa analysis. Heights (solid,black) depicted

5 1 every 60 m, absolute vorticity (dashed, green) every 5 10 s .

0 8880 0

50 -2 0 -4 -2 110 4 70 -4 -2 2

-2 -4

0 4 0 6 0 4

2 4 4

9000 -4 90 6 50 30 -2 4 2 70 50 -2 -2 9120 0 0 0 -2 9240 4 2 90

000311/1200V000 300 MB MULSPED1.94 000311/1200V000 300 MB HGHT 000311/1200V000 300 MB DIVWND (*10**5) Figure 5.29: As in Fig. 5.27 except 300-hPa analysis. Heights (solid,black) depicted

every 120 m. Isotachs (dashed, purple) depicted every 20 knots with speeds of

greater than 70 knots shaded. Divergence (dashed, red) depicted every 2 10 1 s .

63 300

250 300

350 10 5 5

300

200

5 5 350 200 200 25 200 20 10400 450 15 300 250 5

Tropopause Heights (hPa) 000311/1200V000 300 : 700 MB PVORTHTA (*10**7) Figure 5.30: As in Fig. 5.27 except dynamic tropopause and 700-300-hPa potential

vorticity analyses. Tropopause pressures (solid, black) depicted every 50 hPa,

7 2 1 2

potential vorticity (dashed, red) every 5 10 m Ks kg .

5.31) where the e ridge is nearly enclosing the cold air pool. The trowal feature at this level also closely matches the precipitation ﬁeld (Fig. 5.26). An analysis of the 294 K isentropic surface, which was chosen because it is at the level of the strongest frontogenesis (to be discussed), and moisture transport vectors shows that signiﬁcant isentropic upglide is occurring over the STL region (Fig. 5.32), with the vectors intersecting the isobars at nearly perpendicular angles.

5.3.5 Stability and Forcing

A cross section was constructed from Omaha, Nebraska (OMA), to Atlanta, Geor- gia (ATL). Figure 5.33 shows that a deep layer of high RH is located above STL and over locations to the southeast. No instability, either upright or slantwise, is shown above STL, although CSI is present southeast of STL and convective instability is present still further southeastward (Fig. 5.33). The three-dimensional EPV analysis shows that values above St. Louis are less than 1 PVU near the 700-hPa level (Fig. 5.34), which is indicative that WSS is present. As in the 9 December 1999

64 300

302

304 316 318 318

316

316 306 308 310 312 314 316

000311/1200V000 500 MB THTE

Figure 5.31: As in Fig. 5.27 except 500-hPa equivalent potential temperature (depicted every 2K).

600

650

800 650 850

700 900 750 950

000311/1200V000 294 K PRES SR-MTVEC

Figure 5.32: As in Fig. 5.27 except pressure (depicted every 50 hPa) and storm-

~ ~ ~

C C

relative moisture transport vectors (q(V )) on the 294 K isentropic surface.

1 calculated to be from 293 at 9.3 m s .

65 100

150 374376 372370 366368 362364 360358 354356 200 350352 348346 344 342340 338336 334332 250 330 90 90 328 326 324 80 300 80 10 322 30 40 90 100 50 70 20 60 110 320 350 318 80 130 400 316

450 312 314 120 310 80 500 308 550 306 90 600 304 314 650 0 302 700 300 750 298 90 318 800 296 294 316 850 290292 284286288 900 280 282 318 80 278 950 318 000311/1200V000 RELH OMA 000311/1200V000 MSFC ATL 000311/1200V000 THTE Figure 5.33: Vertical cross section from Omaha, Nebraska (OMA), to Atlanta, Geor- gia (ATL), from the 1200 UTC RUC initial ﬁeld valid on 11 March 2000. Solid lines

depict equivalent potential temperature (every 2 K), dashed lines depict absolute 1

geostrophic momentum (every 10 kg m s ). Values of relative humidity greater % than 80% are shaded green, and greater than 90 dark green. STL is located

between the ﬁfth and sixth tick marks from the left edge.

es case, a comparison of the EPV ﬁelds using e (Fig. 5.34) and (Fig. 5.35) are very

similar.

Figure 5.36 shows that frontogenesis is occurring at 700 hPa ( 3 10 K

1 1 1

100 km 3h ) and rather vigorous upward vertical motion (! -8 bs )is occurring above the axis of strongest frontogenesis.

Figures 5.37 and 5.38 show that the layer-average convergence of both QG forc-

14 1 1

ing components were quite weak (2 and 1 10 mkg s , respectively), al-

~ Q

though s forcing was slightly greater. The alternating convergence-divergence patterns in 5.37 are similar to the gravity wave signature discussed by Barnes et al. (1996). Jewett et al. (2001) point out that gravity waves may help in the formation

66 200 7 15 12 7 11 6 14 10 5 6 9 4 13 8 3 7 2 6 250 12 11 10 1 300 89 67 5 4 350 3 2 1 400

450

500

550

600 0 650 700 0 750 1 0 800 1 2 850 1 2 900 950

OMA 000311/1200V000 EPV3 (*10**6) ATL

Figure 5.34: As in Fig. 5.33 except three-dimensional equivalent potential vorticity

calculated with e . Contoured every 1 PVU with negative values shaded.

200 15 12 7 7 11 6 10 5 14 9 4 8 3 13 7 2 6 250 12 1 11 6 10 300 89 7 5 0 4 350 3 2 1 400

450

500 0 550

600 650 700 1 0 750 800 1 1 2 850 2 900 950

OMA 000311/1200V000 EPVTHES3 (*10**6) ATL

Figure 5.35: As in Fig. 5.33 except three-dimensional equivalent potential vorticity

calculated with es . Contoured every 1 PVU with negative values shaded.

67 100 0

-2

150 -4 2 0

0 -4 200 -2 -14 0 14 12 -12 10 8 6 4 -10 2 0 250 -16-18 -1 -10-12-14 0 2 -6-8 -1 0 -2 -2 300 -6 -4 -6 8 6 -4 350 4 2 2 0 6 2 400 0 4 450 -20 -18 -16 500 -12-14 0 2 0 -14 -8 -10 550 -6 0 3 600 1 -13 0 -12 -4 -2 2 650 -11 700 0 -10 4 -90 750 -8 0 800 2 -7 850 -6 -6 2 2 4 -4-2 900 0 0 6-58 -2 10 12 -4 1 950 -3 0 202224 16 -2 000311/1200V000 OMEG (*10**3) OMA 000311/1200V000 FRNTTHTAUOBS (*10**1) ATL

Figure 5.36: As in Fig. 5.33 except Petterssen surface frontogenesis (solid lines,

1 1 1 1

! every 2 10 K 100 km 3h ), and (dashed lines, every 1 bs . of convective cells by lowering static stability and increasing upward vertical velocities. Therefore, gravity waves may have helped in the production of TS in this case, although an attempt to show this, conclusively, would be beyond the scope of this thesis.

5.3.6 Conclusions

The TS that occurred across the STL region on the morning of 11 March 2000 occurred primarily as a result of 700-hPa frontogenesis under conditions of WSS. Frontogenesis occurred at 700 hPa with the development of an early trowal structure, which was better organized at 500 hPa. Again, this is suggestive that occlusions may occur from the top-down (Martin 1999). Additional enhancement for ascent was supplied in the form of upper tropospheric divergence, as the STL region was under the left exit region of a curved jet streak. The precipitation patterns coincide very well with the orientation of the trowal, suggesting that the processes

68 0

0 -2 -4

-2

0 2 0

-12 4 2 0 -2 -8-10 2 16 -6 14 -4 12 -4 8 -4 10 -2

000311/12f00 400:850 Convergence of Qs (10e-14)

~ Q

Figure 5.37: As in Fig. 5.27 except convergence (positive values) of s in the

14 1 1

850-400-hPa layer. Contours drawn every 2 10 mkg s .

0 0

1 0

1 2 -1

-2 -1 5

-1 -2 -2 -3 -4 -5 4 -1

000311/12f00 400:850 Convergence of Qn (10e-14)

~ Q

Figure 5.38: As in Fig. 5.37 except convergence (positive values) of n in the

14 1 1

850-400-hPa layer. Contours drawn every 1 10 mkg s .

69 that are involved in the generation of the trowal airstream are the same processes responsible for creating conditions that are favorable for TS.

5.4 19 April 2000

5.4.1 Introduction

The TS in this event occurred at Ellsworth Air Force Base, South Dakota (RCA), on 19 April 2000 at 1200 UTC. The TS occurred as part of a very intense spring storm that produced blizzard conditions over western South Dakota and severe weather over portions of the northern and central Plains. Snowfall amounts of over 30 inches (75 cm) were observed to the west of Rapid City, South Dakota (Fig. 5.39). 24-hour snowfall (18.7 in.) and precipitation (3.87 in.) records were set at RCA, as well as for several other locations in the region. It should be noted that most of the snow fell in less than a 12-hour period.

5.4.2 Surface Analysis

The surface analysis at 1200 UTC on 19 April 2000 shows a tightly wound occluded cyclone along the South Dakota-Nebraska border (Fig. 5.40). An occluded front extends from the low to the triple point in extreme southwestern Iowa. A warm front extends across southern Iowa, while a cold front is cutting through eastern Kansas. A band of locally heavy precipitation, including thunderstorms, extends from western South Dakota into southern Minnesota (Fig. 5.41). The RCA area is on the edge of an area of heavy precipitation, which is in the form of TS at this time.

5.4.3 Upper Air Analysis

Analyses of 1200 UTC RUC initial ﬁelds on 19 April 2000 indicate that 850 and 700 hPa closed cyclonic circulations (Fig. 5.42) are stacked on top of the surface

70 Figure 5.39: Storm total snowfall amounts (inches) for 19 April 2000. (Obtained from the National Weather Service forecast ofﬁce in Rapid City, South Dakota.)

Figure 5.40: Surface analysis valid at 1200 UTC on 19 April 2000. Isobars depicted every 2 hPa. Red line depicts cross section line.

71 Figure 5.41: Radar mosaic valid at 1200 UTC on 19 April 2000 (obtained from the National Climatic Data Center). low (Fig. 5.40), which is indicative of a mature, occluded system. The corresponding thermal structure also indicates that the system is occluded. The 500-hPa low

is centered over the Nebraska Panhandle (Fig. 5.43) while a strong vorticity maxi-

5 1 mum (40 10 s ) is positioned just west of the barotropic circulations below. At 300 hPa, an elongated jet streak is emerging from a trough over the southwestern U.S. (Fig. 5.44). A divergence axis extends from southwestern Iowa into northwestern South Dakota and then into southeastern Wyoming (Fig. 5.44) at the nose of and in the left exit region of the jet streak. A very elongated dynamic tropopause fold extends from Utah into South Dakota (Fig. 5.45), with the lowest tropopause pressures (600 hPa) located over northwestern Nebraska.

72 3000 2970

3000 1470

1440 1440

1470

3030 3030

2880 1320

1500 1350 2910 3030 2940 1380 2970 3060

3000 1410

000419/1200V000 850 MB HGHT 000419/1200V000 700 MB HGHT

Figure 5.42: 850-hPa and 700-hPa RUC initial ﬁelds valid at 1200 UTC on 19 April 2000. 850-hPa heights (solid, black) depicted every 30 m, 700-hPa heights (dashed, red) depicted every 30 m.

5460 10

5520

10 10 10 5

10 5

20 15 10

40 5640 35 30 5 5640 25

5700 10 5520

25 5760

000419/1200V000 500 MB HGHT 000419/1200V000 500 MB AVORWND (*10**5)

Figure 5.43: As in Fig. 5.42 except 500-hPa analysis. Heights (solid, black) de-

5 1 picted every 60 m, absolute vorticity (dashed, red) depicted every 5 10 s .

73 0 8880 00 50 50

9000 -2 -2 0 -2 9120 -2 0 -2 0 -2 70 2 0 -2 -2 2

30 -2 50 -2 -2 0 6 2 -2 8 50 4 6 4 2 4 30 4 0 0 -2 -2 70 9120 -4 -2 4 -2 9240 0 -6 2 -4 6 0 4 2 2 2 50 9360

-4 2 -2 90 110 9480 -2

0 0 -2 0 0 2 0

000419/1200V000 300 MB MULSPED1.94 000419/1200V000 300 MB HGHT Figure 5.44: As in Fig. 5.42 except 300-hPa analysis. Heights (solid,black) depicted

every 120 m. Isotachs (dashed, purple) depicted every 20 knots with speeds of

greater than 70 knots shaded. Divergence (dashed, red) depicted every 2 10 1 s .

10 300

200 250 5

250 5

10 350

300

60025 200 550 500 5 20

250 500 450

15 150

Tropopause Heights (hPa) 000419/1200V000 300 : 700 MB PVORTHTA (*10**7)

Figure 5.45: As in Fig. 5.42 except dynamic tropopause and 700-300-hPa potential

vorticity analyses. Tropopause pressures (solid, black) depicted every 50 hPa,

7 2 1 2 potential vorticity (dashed, red) every 5 10 m Ks kg .

74 600

750

650

700 850

750 850

850

950 900 850 900

800

850 950 750 750 800 750

900 850 750

900

700

000419/1200V000 294 K PRES S-R Moisture Transport Vectors

Figure 5.46: As in Fig. 5.42 except pressure (depicted every 50 hPa) and storm-

~ ~ ~

C C

relative moisture transport vectors (q(V )) on the 294 K isentropic surface.

1 calculated to be from 216 at 5.7 m s .

5.4.4 Isentropic Analysis

A well deﬁned trowal structure is expressed as a long e -ridge at both 700 and 500 hPa (not shown). The trowal shows up well on the analysis of the 294 K isentropic surface (Fig. 5.46), and extends from Missouri into western South Dakota. Storm-relative moisture transport vectors (Fig. 5.46) show that signiﬁcant isentropic upglide and moisture transport is occurring over western South Dakota. The orientation of the trowal is nearly colocated with the position of the precipitation echoes over South Dakota (Fig. 5.41), again suggesting that the trowal acts as the major producer of precipitation in mature cyclones.

5.4.5 Stability and Forcing

A cross section was constructed from Worland, Wyoming (WRL), to Pipestone, Minnesota (PQN). Figure 5.47 shows that RCA was located under a deep layer of

75 200 352 350 348 346 344 342 340 338 336 334 250 332 330

328 326 300 324

322 350

400

450

320 500 80

550

318 600

650 322 316 324 326 700 90 314 90 750 306 312 304 308310 300302 800 298 296 850 80 900 WRL 000419/1200V000 THTE PQN 000419/1200V000 RELH

Figure 5.47: Vertical cross section from Worland, Wyoming (WRL), to Pipestone, Minnesota (PQN), from the 1200 UTC RUC initial ﬁeld valid on 19 April 2000. Solid lines depict equivalent potential temperature (every 2 K). Values of relative humidity

greater than 80% are shaded green. RCA is located between the ﬁfth and sixth tick

marks from the left edge.

> % high RH ( 80 ) at 1200 UTC. The contours of e indicate that the atmosphere was

convectively stable. Mg was not calculated for this cross section because of the heavily curved nature of the thermal wind ﬁeld. The three-dimensional EPV ﬁeld

(Fig. 5.48) shows that areas of WSS (<2 PVU) were located near the surface (above RCA) and above 700 hPa, and separated by a region of fairly strong symmetric

stability (>3 PVU). As in the other cases, the three-dimensional EPV ﬁelds using

es e (Fig. 5.48) and (Fig. 5.49) are very similar.

The analysis of the frontogenesis and ! ﬁelds (Fig. 5.50) indicates a region of

1 1 1

strong frontogenesis ( 16 10 K 100 km 3h ) at 850 hPa, or just above the

surface at RCA). The strongest upward vertical motion (! 16 bs ) is located

at 650 hPa (Fig. 5.50), directly above the level of strongest frontogenesis.

~ ~

Q Q

analysis (850-400 hPa layer) at this time shows s convergence of approxi-

76 200 6 7 6 7 5 5 4 6

5 4 250

300 2

350

400 1

450

500

550 0 600 1 650 0

700 2 750

800 850 0 5 4 900 WRL 000419/1200V000 EPV3 (*10**6) PQN

Figure 5.48: As in Fig. 5.47 except three-dimensional equivalent potential vorticity

calculated with e . Contoured every 1 PVU with negative values shaded.

200 8 6 7 7 6 5 5 4 6 4 5 250

3 2 300

350 1

400

450

500

550 0

600 1 650

700 2 750 3 800 4 850 0 4 5 900 WRL 000419/1200V000 EPVTHES3 (*10**6) PQN

Figure 5.49: As in Fig. 5.47 except three-dimensional equivalent potential vorticity

calculated with es . Contoured every 1 PVU with negative values shaded.

77 200 -4 -20 0 -16 -2 -12 -8 -4 250 4 16 0 12 8 0

300

350 -2 12

400 -4 -6 0 8 -8 4 450

500 0 4 0 4 550

600 16 -10 -8 650 12 -4 -8 20 700 -18 -6 750 4 -16 8 800 8 12 16 -14 4 -4 850 -2 2024 28 0 24 -12 0 20 900 -2 000419/1200V000 OMEG (*10**3) WRL 000419/1200V000 FRNTTHTAUOBS (*10**1) PQN

Figure 5.50: As in Fig. 5.47 except Petterssen surface frontogenesis (solid lines,

1 1 1 1

every 4 10 K 100 km 3h ), and (dashed lines, every 2 bs ).

14 1 1

Q mately 3 10 mkg s (Fig. 5.51 and n convergence near zero (Fig. 5.52). Therefore, the primary forcing mechanisms were on the synoptic scale, and were likely related to the inﬂuence of the trowal. As in the 11 March 2000 event, the Barnes et al. (1996) gravity wave signature is present, indicative that gravity waves may have played a role in the enhancement of vertical motion.

5.4.6 Conclusions

The analysis generated for this event suggests that the TS that was observed at RCA occurred primarily as a result of lower tropospheric frontogenetical forcing in the presence of in the trowal airstream. Weak symmetric stability above the frontogenetical layer would help enhance vertical motions. Upward vertical motion was also enhanced by strong upper tropospheric divergence in the left exit region of a jet streak, and due to processes associated with a nearby tropopause fold. Because of the strong easterly airstream that existed with this system, and

78 0 0

-2 0 0 0 0 0 0 0 0

-2 0 0 0 0 4 0 0 0 0 -2 -2 0 0

0 0 2 0 0 0 12 10 8 4 64 2 -4 -2 4 2 -2 0 2 4 -4 0 -8 -2 -6

6 4 -2 -4

2 -10 -8 -6 -4 0 -2 6 0 -2 4 0 2 0 0 0

Convergence of Qs (400:850) (10e-14)

~ Q

Figure 5.51: As in Fig. 5.42 except convergence of s in the 850-400-hPa layer.

14 1 1

Contours drawn every 2 10 mkg s .

0 1 0 0 0 0

0 0 0 0

-1 1 3 -3 0 0 -2 0 1

-1 0 0 0 -1

1 0 0 0 1 -1 -1 1 0 -1 0 1 0 1 0 -2 -2 2 -1 1 -1 0 3 -1 2 0 1 2 5 -6 -2 4 -5 1 -1 3 -4 2 0 1 -2-3 0 -1 -1 0 3 2 -4 1 0 1 -5 1 -3 -2 1 0 -1

0 0

Convergence of Qn (400:850) (10e-14)

~ Q

Figure 5.52: As in Fig. 5.42 except convergence of n in the 850-400-hPa layer.

14 1 1

Contours drawn every 1 10 mkg s .

79 because of the characteristics of the terrain in western South Dakota, orographic lift could have also played a signiﬁcant role in the production of enhanced vertical motion. However, the strong synoptic-scale inﬂuences that were associated with this event cannot be ignored.

80 Chapter 6

Discussion and Conclusions

6.1 Discussion of Case Studies

Each of the four TS events occurred to the northwest of a surface cyclone, one of which (19 April 2000) had occluded. While only the 19 April event was associated with an occluded surface cyclone, all the events were associated with cyclones that exhibited a well-deﬁned trowal structure at some level. When radar images and the location of trowal structures are compared, it can be stated, conclusively, that the trowal is the major avenue of ascent and precipitation production in these events. This is further illustrated by the analysis of moisture transport vectors done here (Chapter 5), and by the trajectory analyses performed by Martin (1999). The trowal is a region of frontogenesis, which is likely the single most important forcing mechanism in the production of trowal-related precipitation and TS. In each of the four events, the strongest upward vertical motion was located at or above the level of the strongest frontogenesis. None of the four cases exhibited any CSI, with two instances of upright convection from the release of CI (5 December 1999, 9 December 1999) and the other two likely from slantwise convection in the face of WSS (11 March 2000, 19 April 2000). Therefore, instability associated with TS events is likely less important than the forcing for vertical motion. In each event, a signiﬁcant tropopause/PV anomaly was also nearby. This is

81 important for three reasons. First, tropopause folds will reduce the static stability below them, leading to increased upward vertical motion (Hoskins et al. 1985). Second, large amounts of cyclonic PV advection will create large amounts of absolute vorticity below, thus creating conditions in which absolute vorticity advection

increases with height and enhancing QG forcing. This is well represented in the

~ Q

s convergence ﬁeld, which was the dominant forcing mechanism in each event. Third, as PV tries to conserve itself, the ﬂow will become more meridional and the

cyclone will soon become negatively tilted, and the trowal structure will develop.

~ ~

Q Q n s forcing was equal to or greater than forcing in each of the four cases, reinforcing what was discussed in the preceding paragraph, and concurring with the findings of Martin (1999). The frontal-scale processes that occur in these events occur largely due to the forcing by the large-scale flow, and in response to deformation of the thermal field. In regard to the evaluation of CI/CSI, few differences were observed between the methods espoused by Moore and Lambert (1993) and Schultz and Schumacher (1999). The negative EPV regions that are created by the Schultz and Schumacher (1999) method are slightly greater, but had no bearing on the diagnosis of stability at any of the TS locations. The Moore and Lambert (1993) yielded EPV fields with

stronger gradients. This is because the true moisture ﬁeld is used in the calculation

e of e , and strong gradients will be reﬂected as strong gradients of EPV.

6.2 Conclusions

The discussion in this section will be centered on the statements that were outlined in Section 1.1. The ﬁrst contention was that there would be two regions of preferred TS occurrence, in the trowal airstream (due to frontogenesis) in the occluded sector of a surface cyclone, and to the northeast of a surface cyclone in advance of a warm front. The 30-year climatology shows that this is true, at least geographically with

82 regard to the cyclone. Figure 4.7 shows that the northwest and northeast sectors are by far the most preferred regions for the occurrence TS. The case studies illustrate that the TS in each event occurred in the trowal airstream in the occluded sector of the cyclone in regions of abundant frontogenetical forcing. None of the case studies involved TS in the northeast sector, so the dynamics of these types of events cannot be addressed. The second contention was that CSI will be more prevalent in the northeast sector of a cyclone and that CI will be more prevalent in the occluded sector. Again, the first part of this contention cannot be addressed. However, of the events which were studied, only one involved the release of CI, although another (5 December 1999) likely involved release of upright instability, even though no CI was resolved. It is no surprise that the instability was upright in the 5 and 9 December 1999 cases because the middle to upper tropospheric shear is cyclonic and the open-wave cyclone still possessed a baroclinic tilt (Bluestein 1986). The flow in the 11 March 2000 case was less cyclonic, but not really anticyclonic either. Also, the system was also an open baroclinic wave. A region of CSI is present to the southeast of STL (Fig. 5.33), where the flow is more anticyclonic. Therefore, the shear profile was not favorable for the occurrence of either CSI or CI. In the 19 April 2000 event, the shear is cyclonic throughout the troposphere. However, the system was approaching an equivalent barotropic state from the surface to 700 hPa which may have lessened the vertical temperature profile to the point that upright instability could not occur. So, while the second contention could not be proven, it was not disproven, either. Because the systems that produce TS are so complex, the contention that was made may have been oversimplified. This is because the shear and temperature profiles will both effect the stability. A “stacked system” will possess small amounts of vertical wind shear near the cyclone’s center, thus inhibiting the development of symmetric instability (Bluestein 1986). At the same time, the resulting decrease in

83 lapse rates will lessen the likelihood that any upright instability will occur.

84 References

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Bennetts, D.A., and B.J. Hoskins, 1979: Conditional symmetric instability - A possible explanantion for frontal rainbands. Quart. J. Roy. Meteor. Soc., 105, 945-962.

———-, and J.C. Sharp, 1982: The relevance of conditional symmetric instability to the prediction of mesoscale frontal rainbands. Quart. J. Roy. Meteor. Soc., 108, 595-602.

Bluestein, H., 1986: Fronts and jet streaks: A theoretical perpspective. Mesoscale Meteorology and Forecasting, Peter Ray, Ed., Amer. Meteor. Soc., 173-215.

Colman, B.R., 1990: Thunderstorms above frontal surfaces in environments without positive CAPE. Part I: A climatology. Mon. Wea. Rev., 118, 1103-1121.

Curran, J.T., and A.D. Pearson, 1971: Proximity soundings for thunderstorms with snow. Preprints, 7th Conference on Severe Local Storms, Kansas City, MO, Amer. Meteor. Soc., 298-300.

Holle, R.L., J.V. Cortinas, and C.C. Robbins, 1998: Winter thunderstorms in the United States. Preprints, 16th Conf. on Weather Analysis and Forecasting, Phoenix, AZ, Amer. Meteor.Soc., 118-119.

Hoskins, B.J., M.E. McIntyre, and A.W. Robertson, 1985: On the use and signiﬁ- cance of isentropic potential vorticity maps. Quart. J. Roy. Meteor. Soc., 111, 877-946.

85 Jewett, B.F., M.K. Ramamurthy, and R.M. Rauber, 2001: The role of evaporative processes in gravity wave genesis. Preprints, 9th Conference on Mesoscale Pro- cesses, Ft. Lauderdale, FL, Amer. Meteor.Soc., 485-489.

Martin, J.E., 1998a: The structure and evolution of a continental winter cyclone. Part I: Frontal structure and the occlussion process. Mon. Wea. Rev., 126, 303-328.

————., 1998b: The structure and evolution of a continental winter cyclone. Part II: Frontal forcing of an extreme snow event. Mon. Wea. Rev., 126, 329- 348.

————, 1999: Quasigeostrophic forcing of ascent in the occluded sector of cyclones and the trowal airstream. Mon. Wea. Rev., 127, 70-88.

McCann, D.W., 1995: Three-dimensional computations of equivalent potential vorticity. Wea. Forecasting, 10, 798-802.

Moore, J.T., and P.D. Blakely, 1988: The role of frontogenetical forcing and conditional symmetric stability in the Midwest snowstorm of 30-31 January 1982. Mon. Wea. Rev., 116, 2155-2171.

————, and T.E. Lambert, 1993: The use of equivalent potential vorticity to diagnose regions of conditional symmetric instability. Wea. Forecasting, 8, 301-308.

Moore, J.T., A.C. Czarnetzki, and P.S. Market, 1998: Heavy precipitation associated with elevated thunderstorms formed in a convectively unstable layer aloft. Meteorol. Appl., 5, 373-384.

Mote, T.L., D.W. Gamble, S.J. Underwood, and M.L. Bentley, 1997: Synoptic-scale features common to heavy snowstorms in the southeast United States. Wea. Forecasting, 12, 5-23.

86 Nicosia, D.J., and R.H. Grumm, 1999: Mesoscale band formation in three major northeastern United States snowstorms. Wea. Forecasting, 14, 346-368.

O’Brien, J.J., 1970: Alternative solutions to the classical vertical velocity problem. J. Appl. Meteor., 9, 197-203.

Sanders, F., 1986: Frontogenesis and symmetric stability in a major New England snowstorm. Mon. Wea. Rev., 114, 1847-1862.

Schultz, D.M., and P.N. Schumacher, 1999: The use and misuse of conditional symmetric instability. Mon. Wea. Rev., 127, 2709-2732.

Schultz, D.M., 1999: Lake-effect snowstorms in northern Utah and western New York with and without lightning. Wea. Forecasting, 14, 1023-1031.

Schwartz, B.E., S.J. Weiss, and S.G. Benjamin, 2000: An assessment of the Rapid Update Cycle short-range forecast ﬁelds related to convective development. Preprints, 20th Conference on Severe Local Storms, Orlando, FL, Amer. Meteor. Soc., 443-446.

Smith, T.L., S.G. Benjamin, B.E. Schwartz, and G. Grell, 2000: A past and future look at the Rapid Update Cycle for the 3 May 1999 severe weather outbreak. Preprints, 20th Conference on Severe Local Storms, Orlando, FL, Amer. Meteor. Soc., 21-24.

Weismuller, J.L., and S.M Zubrick, 1998: Evaluation and application of conditional symmetric instability, equivalent potential vorticity, and frontogenetic forcing in an operational forecast environment. Wea. Forecasting, 13, 84-101.

87 Vita

Christopher Eric Halcomb was born on 24 February 1967 at Southeast Missouri Hospital in Cape Girardeau, Missouri. After my parents had divorced, my mom (Tracy) and I moved to Poplar Bluff, Missouri, where she married Richard Gant when I was 7 years old. At the age of nine, we moved to Dexter, Missouri, from where I graduated from high school in 1985. After spending some time at the University of Missouri pursuing study in meteorology and respiratory therapy, I received my Bachelor’s Degree in Geoscience from Southeast Missouri State Uni- versity (SEMO) in 1995. Hoping to study earthquakes, I found myself involved in the the unrewarding ﬁeld of hazardous environmental clean-up in Memphis, Tennesee. I then returned home to Cape Girardeau where I started in the BSN program at SEMO. After a year of straight A’s, I decided that nursing was just not for me. So I decided to go back to my true love, meteorology, and returned to Mizzou to get my M.S. My love for meteorology started with a shear fascination for the very active period that was the 1970’s, but was fully realized after a surprise thunder-blizzard paralyzed much of southeastern Missouri and southern Illinois on 25 February 1979, the day after my 12th birthday. Anywhere from 1 to 4 inches of snow was supposed to fall, but up to 24 inches fell instead, with drifts piled up to 8 feet in depth by winds of over 50 mph. This was the highlight of a week that featured severe thunderstorms, heavy rain, huge temperature swings, and a solar eclipse. After that, I was hooked!