AN ABSTRACT OF THE THESIS OF
Shawn P. Bennett for the degree of Master of Science in
Atmospheric Sciences presented on September 10. 1987
Redacted for privacy
Abstract approved: Steven A. Rutledge
On May 28, 1985 a nocturnal esosca1e convective
ystem (MCS) moved into the observational network of the O.K.
PRE-STORM (Oklahoma-Kansas .reliminary egionalxperiment for STORM-Central) Project. Satellite imagery showed an extensive circular cloud shield. Sensing of meteorological parameters by the NCAR (National Center for Atmospheric
Research) Portable Automated Mesonet (PAM II), as well as observations by both conventional and Doppler radar document this MCS as a fast moving squall line.
This MCS developed along the Nebraska-Wyoming border in the warm moist air associated with a quasi-stationary frontal boundary parallel to the Front Range of the Rocky
Mountains. In addition to the lifting caused by the frontal boundary, the upsiope topography of the Central Great Plains as it rises to meet the Front Range of the Rocky Mountains caused additional instability by orographically lifting the convectively unstable air mass. Initial development of the
MCS may have been forced by an indirect circulationin the left-front exit zone of a 25 m jet streak over the region at 300 mb (Ucellini and Johnson, 1979) . At the surface a lee-side trough was found along the Front Range, while aloft a broad long wave ridge was positioned over theCentral U.S.
Changes created in the near surface boundary layer environment during the passage of the MCS are documented by the PAN II mesonetwork. A mesohigh associated with the convective line and a wake lowor "wake depression" (Fujita,
1955; Johnson and Hamilton, 1987) were observed.
Analysis of single and dual-Doppler radar data reveal the kinematic structure and radar reflectivity of the MCS from near the surface boundary layer to the upper troposphere. Several features are identified, including a cyclonic mid-level circulation - 7 km deep (Fritsch and
Chappell, 1980), a plunging rear inflow jet with relative flow speeds of -10-15 ms-, an echo-free notch in the rear of the stratiform region, and a bow echo in the convective line
(Smull and Houze, 1985; 1987a,b)
The 12-hour-plus duration of the MCS may be attributed to a combination of positive feedback mechanisms which reinforced the squall line with positive vertical momentum. Among these positive feedback mechanisms was the configuration of the surface mesoscale pressure systems which may have increased the low-level horizontal convergence and as a result supplied additional mass and moisture to the convective line. This then may build the convective cells within the line and lead to an intensification of the mesoscale pressure systems and induce more convergence, etc
(Fritsch and Chappell, 1980) A Case Study of the May 28, 1985 MesoscaleConvective System Observed During the O.K. PRE-STORM Project
by
Shawri P. Bennett
A THESIS
submitted to
Oregon State University
in partial fulfillment of the requirements for the degree of
Master of Science
Completed September 10, 1987 Commencement June, 1988 APPROVED:
Redacted for Privacy Professor of Atmospheri Sciences in charge of major
Redacted for Privacy Chairman of Depa(entof AtmosphericSciences
Redacted for Privacy Dean of Graduate7choo1
Date thesis is presented September 10, 1987
Presented by Shawn P. Bennett ACKNOWLEDGEMENTS
I would like to thank my MajorProfessor, Dr. Steven
A. Rutledge, for his valuablerecommendations and guidance during the preparation of this thesis. I also appreciate the comments by members of my thesiscommittee Drs. Steven K. Esbensen, Glenn A. Klein, and Peter R. Fontana
I would like to thank Dean A.Vickers for his helpfulness and expertise in computerprogramming, data aquisition, and data analysis techniques. Special thanks are in order for Steven M. Ginn for his helpin computer programming. This research was supported by the NationalScience
Foundation under grants ATM-8317466 and ATM-8521403. TABLE OF CONTENTS
1. INTRODUCTION 1 1.1 General Background and Earlier Studies of Convective Storms 1 1.2 Recent Studies of Mesoscale Convective Systems 6 1.3 Recent Studies Using Doppler Radar Observations 10 1.4 Research Objectives 13
2. PRE-STORM PROGRAM DESIGN AND OPERATIONS PLAN 15 2.1 Program Background and Goals 15 2.2 Observational Program Design 16
3. SYNOPTIC AND SATELLITE OVERVIEW 22 3.1 Surface Synoptic Situation 22 3.1.1 0600 GMT 28 May 1985 meteorological conditions 22 3.1.2 1200 GMT 28 May 1985 meteorological conditions 23 3.1.3 1800 GMT 28 May 1985 meteorological conditions 25 3.2 Upper-Air Long and Shortwave Pattern 27 3.2.1 0000 GMT 28 May 1985 upper-air pattern 27 3.2.2 1200 GMT 28 May 1985 upper-air pattern 30 3.3 Classification and Definition of MCSs 34 3.4 Satellite Imagery of the 28 May 1985 MCS 37
4. MESOSCALE ANALYSIS 47 4.1 Discussion of the WSR-57 Radar Observations 47 4.2 The PAN II Surface Mesonetwork 54 4.2.1 GEMPAK and PAN II data analysis methods 57 4.2.2 Analysis of PAN II data 58 4.2.3 Surface mesoscale pressure systems 74 4.3 Analysis of Single Station Time Series 80
5. SINGLE DOPPLER ANALYSIS 86 5.1 Introduction 86 5.2 Discussion of Vertical Cross-Section Analysis 86 5.3 Discussion of Velocity-Azimuth-Display (VAD) Analysis 106 5.4 Discussion of Vertically-Pointing (VP) Analysis 115
6. DUAL-DOPPLER ANALYSIS OF THE TRAILING STRATIFORN REGION AND ASSOCIATED MESOSCALE CIRCULATION 121 6.1 Discussion of Reflectivity and Kinematic Fields 122 6.2 Summary of Evidence Regarding a Mesoscale Circulation 130
7. CONCLUSIONS 132 7.1 Conclusionsfrom the Observational Analysis 132 7.2 Suggestionsfor Further Research 135
REFERENCES 137 LIST OF FIGURES
1.1. Tracks of MCSs during May, June,July of 1979
through 1983. Dotted lines indicate position of
early storm development. Circled numbers correspond
to the position of themaximum extent of the -32CC
cloud shield from infrared satellitedata. The 'X'1
represents the location where theconvective system
was terminated as a MCS.
(Adapted from Cunning, 1986). 2
1.2 A schematic section through a squallline thunderstorm illustrating the horizontaland
vertical structures of the squall line. The wake
depression, the internal airflow (thin arrows),
and environmental airflow (broad arrows) are
depicted. (Adapted from Fujita, 1955). 5
1.3. Time cross section through the 29 May 1947 squall line during its passage over the Thunderstorm
Project Network near Wilmington, Ohio. Thin solid
lines are wet-bulb potential temperature (deg K);
heavy lines, boundaries of stable layersassociated with squall front on right, and polar front onthe
left. Arrows indicate schematic stormrelative
average circulation. Stippled region is cloud. (Adapted from Newton and Newton, 1959). 7
1.4. Conceptual model of the 22 May 1976 squallline
system at maturity. Outermost scalloped line is
extent of cloud. Outermost solid line is extent
of radar echo useable for Doppler velocity
measurements. Heavy solid lines are more intense
reflectivity features. Mesohigh and wakelow are
located at "H" and "L". Storm relative air motions
are shown by solid arrows. Broad arrow from asterisk indicates hypothesized ice particle
trajectory. (Adapted from Smull and Houze, 1985). 12
1.5. Revised conceptual model based upon new research
findings. Note airflow changes, especially those associated with the convective tower. (Adapted
from Smull and Houze, 1987b). 12
2.1. Overview of the surface observational mesonetwork
in Kansas and Oklahoma during PRE-STORM. 17 3.1. 0600 GMT 28 May 1985 Synoptic analysis 24
3.2 1200 GMT 28 May 1985 Synoptic analysis 24
3.3 1800 GMT 28 May 1985 Synoptic analysis 26
3.4 0000 GMT 28 May 1985 850 mb level analysis 26
3.5 0000 GMT 28 May 1985 700 mb level analysis 29
3.6 0000 GMT 28 May 1985 500 mb level analysis 29
3.7 0000 GMT 28 May 1985 300 mb level analysis 31
3.8 1200 GMT 28 May 1985 mb 850 level analysis 31
3.9 1200 GMT 28 May 1985 700 mb level analysis 33
3.10 1200 GMT 28 May 1985 500 mb level analysis 33
3.11 1200 GMT 28 May 1985 300 mb level analysis 35
3.12Satellite infrared image showing the enhancement curve temperature ranges that correspond tothe various shades of black, gray, and white. Colder
cloud tops infer stronger more intenseconvection.
(Adapted from Maddox, 1980). 36
3.13 0000 GMT 28 May 1985 infrared satelliteimage 36
3.140300 GMT 28 May 1985 infrared satelliteimage 38
3.15 0600 GMT 28 May 1985 infrared satelliteimage 38
3.16 0900 GMT 28 May 1985 infrared satellite image 40
3.17 1101 GMT 28 May 1985 infrared satellite image 40
3.18 1331 GMT 28 May 1985 visible satellite image 42
3.19 1501 GMT 28 May 1985 infrared satellite image 42
3.20 1800 GMT 28 May 1985 infrared satellite image 44
3.21 1830 GMT 28 May 1985 infrared satellite image 44
3.222300 GMT 28 May 1985 infrared satellite image 45 4.1. Observed reflectivity at 1008 GMT 28 May 1985from Wichita NWS WSR-57 radar. The elevation angle of
this PPI and all that follows is10. Light to
darker gray shading every 10 dBZ beginningwith
20 dBZ. Contour interval 10 dBZ beginning with
10 dBZ. 48
4.2 As in fig. 4.1 except for 1127 GMT. 50
4.3 As in fig. 4.1 except for 1356 GMT. 52
4.4 As in fig. 4.1 except for 1457 GMT. 53
4.5 The Oklahoma-Kansas PRE-STORM PAN II mesonetwork. (Adapted from Johnson and Toth, 1986) . 55
4.6 NCAR Portable Automated Mesonetwork Station. (Adapted from Johnson and Toth, 1986). 56
4.7 Mesoanalysis of PAN II data at 1010 GMT 28 May 1985. Reflectivity contouring as in Fig. 4.1. a) pressure
(nib) b) surface divergence(x105 s1). 59
4.7 continued. Mesoanalysis of PAN II data at 1010 GMT 29 May 1985. C) relative vorticity(x105 s-)
d) ground relative surface windfield (m s-),
vector length proportional to speed,shortest =
1 m s1, longest = 9 ms_i. 60
4.8 Mesoanalysis of PAM II data at 1130 GMT 28 May1985. Reflectivity contouring as in Fig. 4.1. a) pressure
(mb) b) surface divergence(x105 s-). 62
4.8 continued. Mesoanalysis of PAM II data at 1130 GNT
29 May 1985. C) relative vorticity(x105 s) d) ground relative surface wind field (ms1),
vectors as in Fig. 4.7. 63
4.9 Mesoanalysis of PAM II data at 1130 GMT 28 May 1985. Reflectivity contouring as in Fig. 4.1.
a) temperature field (°C)b) mixing ratio field
(g/kg). 65
4.9 continued. Mesoanalysis of PAM II data at 1130 GMT
28 May 1985. C) temperature advection (°C hr-i)
d) moisture advection (g/kg hr-i). 66
4.10Mesoanalysis of PAN II data at 1415 GMT 28 May 1985. Reflectivity contouring as in Fig. 4.1.
a) pressure (mb) b) surfacedivergence (x105 s-). 68
4.10continued. Mesoanalysis of PAN II data at 1415 GMT
29 May 1985. c) relative vorticity(x105 s)
d) ground relative surface windfield (m s), vectors as in Fig. 4.7, except longest =6 m s. 69
4.11Mesoanalysis of PAN II data at 1515 GMT 28 May
1985. Reflectivity contouring as in Fig. 4.1.
a) pressure (mb) b) surface divergence(x105 s-). 72
4.11continued. Mesoanalysis of PAN II data at 1515 GMT
29 May 1985. c) relative vorticity(xl05 s) d) ground relative surface wind field (m s),
vectors as in Fig. 4.7, except longest = 12 ms_i. 73
4.12Composite of surface pressure and reflectivity
fields for 28 May 1985. Reflectivity contouring
as in Fig. 4.1. a) 1356 GMT b) 1422 GMT. 76
4.12continued. Composite of surface pressure and reflectivity fields for 28 May 1985. Reflectivity
contouring as in Fig. 4.1. c) 1457 GMT d) 1528 GMT. 77
4.13Tracks of mesohigh and wake low for28 May 1985 squall line from 1000 GMT to 1430 GMT.Number
above the L (wake low) or H (mesohigh)indicates
time, number below indicates departureof pressure,
in mb, at 518 m from 950 mb 79
4.14Schematic diagram of pressure profiles for a
squall line thunderstorm.
(Adapted from Fujita, 1955). 81
4.15Time series depiction of the passage of the28 May 1985 MCS at PAM II station 20. 81
4.16Time series depiction of the passage of the 28 May 1985 MCS at PAM II station 6. 83
4.17Time series depiction of the passage of the 28 May 1985 MCS at PAM II station 24. 83
5.1. 28 May 1985 1003 GMT Vertical Cross-Section
Analysis at CP-3. a) reflectivity structure;
contour interval 5 dBZ, shaded gray for reflectivity > 30 dBZ. b) storm relative airflow;
contour interval 5 ms1, shaded gray for rear-to-front-flow, not shaded for front-to-rear
flow. 88
5.2 As in Fig. 5.1 except for 1033GMT. 91
5.3 As in Fig. 5.1 except for 1120GMT. 94
5.4 As in Fig. 5.1 except for 1148GMT. 97
5.5 As in Fig. 5.1 except for 1225GMT. 99
5.6 As in Fig. 5.1 except for 1314GMT. 101
5.7 As in Fig. 5.1 except for 1148GMTat CP-4. 103
5.8 As in Fig. 5.1 except for 1314GMTat CP-4. 105
5.9. 28 May 1985 1130GMT VADanalysis from the CP-3 1), radar. Horizontal wind speed (m dashed line; horizontal wind direction (deg), solid line;
reflectivity profile (avg dBZ), solid-dashed line;
divergence profile (x i0 1),dashed line; and vertical velocity (m s), solid line. 108
5.10 As in Fig. 5.9 except for 1158 GMT. 108
5.11As in Fig. 5.9 except for 1235 GMT. 113
5.12. 28 May 1985 Vertically-PointingTime-height
Cross-Section at CP-4 beginning 1158 GMT
a) reflectivity (dBZ); contourinterval 5 dBZ, shaded
gray for reflectivity > 30 dBZ b) Doppler velocity (W) (m s); contour interval 1 m s-, shaded grayfor W>0
m s above 0°C level; 2 ms1 contour interval, shaded
gray for W<-6 m s- below0°C. 116
5.13 As in Fig. 5.12 except for; a) particlefallspeed
(Vp) (m 1)b) vertical velocity (w)(m sH) . 117
6.1. Horizontal dual-Doppler Analysis at 13:14 GMT a,c) 1.9 km and 2.9 km vector wind, respectively;
vector length is proportional to wind speed, 1
tick = 12 m s b,d) 1.9 km and 2.9 km vertical
velocity field (w)(m s); darker shading indicates
w>O, lighter shading indicates w<0, refer to
grayscale for magnitude. 123 126 6.2 As in Fig. 6.1 except for 3.9 kmand 4.9 km.
6.3 As in Fig. 6.1 except for 7.9 km. 129 LIST OF TABLES
2.1 Characteristics of the PRE-STOfU4 Doppler radars. (Adapted from Rutledge et al., 1987) . 19
3.1 Size and duration criteria for classification as
an MCC. (adapted from Maddox, 1980). 35
4.1 Accuracy of sensors on the PAM II station.
(Adapted from Johnson and Toth, 1986) . 56 A CASE STUDY OF THE MAY 28, 1985 S0SCALE CONVECTIVE SYSTEM OBSERVED DURING THE O.K. PRE-STORM PROJECT
1. INTRODUCTION
1.1 General Background and Earlier Studies of Convective
Storms Mesoscale convective systems have an important impact on the people who live between the Rocky Mountains and the
Appalachian Mountains of the United States. The impact of these storms on mankind can be tragic, including flash floods and heavy rains, property destruction caused by high winds and tornadoes, and loss of life. Losses in just one such event can be enormous in terms of lives and property. The mesoscale convective system that occurred during 2-4 May 1978 caused 7 deaths, 103 injuries, wind damage, hail, tornadoes,
4-13 inch rains, and severe flooding as it traversed the southern U.S. from west Texas to the Gulf States (Maddox,
1980). Figure 1.1 is a compilation of the tracks MCSs followed over a 5 year period from 1979 - 1983 (Cunning,
1986) . Some MCSs are particularly disastrous, like the 2-4 May 1978 MCS mentioned here, while others are not so destructive, but make a significant contribution to the growing season precipitation in the breadbasket of North
America (Fritsch et al., 1981).
The mesoscale convective system (MCS)is broadly defined as all precipitation systems that have horizontal May 1-15 May 16.31
- 4_.___
..., .- --.
&- -N
June 1-15'-] June 16-30"\J
July 1-15 July 16-31
Fig. 1.1. Tracks of MCSs during May, June, Julyof 1979 through 1983. Dotted lines indicate position of early storm development. Circled numbers correspond to the position of the maximum extent of the -32CC cloud shield frominfrared satellite data. The ttx" represents the location where the convective system was terminated as a MCS. (Adapted from
Cunning, 1986). 3 scales of 50 to 500 km and include significant convection during some part of their lifetimes. Several types of MCS5 have been identified and these storms were of particular interest during the Qklahoma-ansas reliminary egional experiment for STORM-Central (O.K. PRE-STORM) squall lines with new cell formation occurring on their right southern flank, maturing and propagating along the line, then decaying and becoming a stratiform precipitation region toward the northern portion of the line (Newton and Fankhauser, 1964); fast moving squall lines with trailing stratiform precipitation (Newton, 1950; Fujita, 1955); individual supercell storms (Browning and Ludlam, 1962); and systems comprised of short convective lines embedded within a broad region of stratiform rain.
Research programs concentrating on convective storm systems, such as PRE-STORM, had their beginnings in important earlier scientific investigations of the thunderstorm and squall line. Programs like the Thunderstorm Project in Florida during 1946 and in Ohio during 1947, laid the groundwork for subsequent research efforts. This project provided a first-of-its-kind dataset of meteorological field variables on the mesoscale and synoptic scale of the thunderstorm and the squall line. Byers and Braham (1948) reported their theory of thunderstorm circulations based on their analysis of this first complete set of measurements obtained during the Florida phase of the project. They 4 described the thunderstorm circulation as orderly instead of random and chaotic as was previously thought. Newton (1950) completed a synoptic scale study of a prefrontal squall line using data collected during the 1947 field season of the Thunderstorm Project in Wilmington, Ohio.
Newton's study centered on a 29 May 1947 prefrontal squall line that developed in the warm sector ahead of a strong cold front. The resulting squall line was observed to move east faster than the cold front. Newton noted a cool dry tongue of air signified by a low wet-bulb potential temperature descending from behind and toward the squall line front.
This feature is similar to that which Smull and Houze (1985) have described recently as a rear inflow jet.
Continuing the mesoscale analysis of thunderstorms and squall lines begun in the Thunderstorm Project, Fujita (1955) reported on the mesoscale analysis of a cold front over Japan and two squall lines that developed in June 1953 over the states of Kansas and Oklahoma. The mesoscale analysis by Fujita (1955) revealed some features of squall lines that could not be determined in a synoptic scale study. Features such as the thunderstorm high and the wake depression were identified. Figure 1.2 shows the conceptual model developed by Fujita. Newton and Newton (1959) completed a detailed investigation of the effects of vertical shear on the 5
- J\<:L/- 2OOO - - - - DWO - .DWD DW0' I - -;. -I - - FRONT4LS -\ CONO LEVEL LT WRMAIR z-....-: WAE .-
9 co 200 MILES
Fig. 1.2. A schematic section through a squall line thunderstorm illustrating the horizontal and vertical structures of the squall line. The wake depression, the internal airflow (thin arrows), and environmental airflow
(broad arrows) are depicted. (Adapted from Fujita, 1955) kinematic structure of the 29 May 1947 squall line. Figure
1.3 depicts in cross section form the thermodynamic and
kinematic structure of this storm, as well as itsassociated
surface weather conditions. In their analysis, inflow of dry
air into the rear of the storm is present asindicated by a tongue of low air. This flow descended towards the
convective line. This low air originated in the post-storm environment of the middle troposphere(Newton,
1950). Also studied by Newton and Newton (1959) was the self-propagating nature of squall lines and the ability for
squall line type thunderstorms to continue at night when air
mass type thunderstorms tended to die out. A main result of
this study was that the induced hydrodynamic pressure field,
the deviation of the actual pressure p from the hydrostatic
pressure Ph' could produce vertical accelerations that were
comparable in strength with ordinary buoyancy forces.
1.2 Recent Studies of Mesoscale Convective Systems Advancing technologies have provided new tools with
which to investigate the nature of MCSs. Principal among
these is the Doppler radar. Other advances include a more
sophisticated mesonetwork of meteorological stations,
improved satellite imagery, and faster more efficient
computers that offer the opportunity to store, retrieve, and
process the large datasets collected and created during
mesoscale research projects. 7
400 mb 293- 292 F I __-i-292 2!4 2JiiiN\ E4;290 500 289 i ,..-r289 \ rF '5 288..j\ '1 ___N \ 600 92 -.\J -289 LI 290
::: '::
riO
I 19 133 Time (mm.) from 62 I 20 -7 2 188 squall possaqe 44 / 12 -50 p. P.TP TRW+R+
40 3020110 00 90 80 70 60 50 403020 0 0 Distance )m,les)
Fig. 1.3. Time cross section through the 29 May 1947 squall line during its passage over the Thunderstorm Project Network near Wilmington, Ohio. Thin solid lines are wet-bulb potential temperature (deg K); heavy lines, boundaries of stable layers associated with squall front on right, and polar front on the left. Arrows indicate schematic storm relative average circulation. Stippled region is cloud.
(Adapted fromNewton and Newton, 1959) Sanders and Paine (1975) conducted studies on the modification of the mesoscale environment by convection using a simple convective plume model. Their data came from an Oklahoma MCS observed by radiosondes, conventional radar, and a mesonet of surface recording stationsoperated by the
National Severe Storms Laboratory (NSSL) . They stated that initiation of extensive deep convection was triggered by updrafts created solely by a cold front passing through the region. Sanders and Emanuel (1977) studied the hydrologic and momentum budgets of the same May 1970 Oklahoma MCS. They concluded that long lasting convective storms, whether frontal, air mass, or of the squall line type, contain elements of vigorous cumulus convection, as well as an organized and strong mesoscale circulation. They also suggested that the synoptic scale structure of the atmosphere plays an important role in determining the character of the
MCS. Ogura and Liou (1980) provided a case study of a well-organized midlatitude squall line that passed through a mesonetwork deployed by the National Severe Storms Laboratory
(NSSL) on 22 May 1976. Their study identified some important components of the circulation pattern associated with a squall line and trailing stratiform region: low level inflow of moist air along the leading edge of the convective line, inflow into the rear of the stratiform region of low 9 air at middle levels, and a cool downdraft situated below the stratiform cloud. Relative airflow structure clearly showed an upshear tilt in the updraft. An upper-level, mesoscale divergence maximum was found about 70 km to the rear of a mid-level convergence maximum, leading to the presence of a mesoscale updraft situated above the 0°C level and a mesoscale downdraft below. It was found that horizontal momentum was approximately conserved following the air-parcel motion in the updraft region. Momentum from the low-level front-to-rear flow was carried aloft through the convective region where it was met by a sloping rear inflow producing a convergence maximum at middle-levels within the stratiform region. Leary and Rappaport (1983) studied a mesoscale convective system that passed through the data network in
Texas of the High Plains Cooperative Program (HIPLEX) on 8
June 1980. The radar reflectivity distribution revealed an arc-shaped band of intense convection at the leading edge of an extensive area of lighter precipitation, which was stratiform in character. The leading edge of heavy precipitation propagated at an average speed of 20 m In cross section, a distinct radar bright band was observed in the stratiform region. This broad area of light to moderate stratiform precipitation was observed to persist for many hours after the passage of the convective line. Upper air soundings suggested the presence of a mesoscale downdraft in 10 the lower troposphere and a mesoscale updraft aloft.
Striking similarities to radar cross section features described by Houze (1977) in tropical convective systems were noted. Surface analysis indicated that the MCS was initiated along a synoptic scale baroclinic zone.
1.3 Recent Studies Using Doppler Radar Observations The studies reported in Sec. 1.2 have examined squall lines with trailing stratiform regions based on observations of the surface and troposphere obtained by special mesoscale networks including surface and rawinsonde observations, as well as conventional radar. These studies have described the basic circulation pattern of MCSs, but they lack the time and spatial resolution to reveal important details within the mesoscale circulation. Several recent studies by Roux et al.
(1984), Smull and Houze (1985;1987a,b), Srivastava et al.
(1986), and Rutledge et al.(1987), have emphasized the Doppler radar observations made in squall lines with trailing stratiform regions. The use of Doppler radar has enabled these investigators to identify more detailed features in the thermodynamic and kinematic structure of the MCS. This has resulted in refinements to the earlier conceptual model of the MCS structure and circulation.
Smull and Houze (1985; 1987a) reported on the single and dual-Doppler radar analysis of the 22 May 1976 Oklahoma squall line system, the same system studied by Ogura and Liou 11
(1980). Figure 1.4 shows the conceptual model they developed to describe the storm relative airflow and reflectivity structure. Their model shows the front-to-rear flow into the leading edge of the convection with an acceleration of the flow past the convective towers. A rear inflow jet entered the trailing edge of the stratiform region. A mesoscale downdraft extended between -100 km and -20 km. A thunderstorm high (mesohigh) was shown associated with the convective line, while a wake depression (wake low) was analyzed associated with the trailing edge of the MCS.
Further investigations of the rear inflow current into the 22 May 1976 squall line system and other MCSs by Smull and Houze (1987b) showed that flow into the rear of a storm, termed a rear inflow jet by Smull and Houze (1985), is a fairly common feature, although the magnitude of the rear inflow jet can vary considerably. Mid-latitude squall lines typically exhibited the strongest rear inflow, while rear inflow in tropical systems was weak or totally absent.
Figure 1.5 shows their refined conceptual model of the 22 May
1976 squall line system. Analysis of sounding data revealed the rear inflow jet was characterized by relatively low air. A similar thermodynamic structure was reported by
Newton (1950), Newton and Newton (1959), Sanders and Paine
(1975), and Ogura and Liou (1980)
Rutledge eC al.(1987) describe the single-Doppler observations of the 10-11 June 1985 MCS during the PRE-STORN 12
I
IUi
-140 -120 -100 -80 -60 -40 -20 0 20 40 L H HORIZONTAL DISTANCE FROM LEADING SQUALL-LINE ECHO (1(M)
Fig. 1.4. Conceptual model of the 22 May 1976 squall line system at maturity. Outermost scalloped line is extent of cloud. Outermost solid line is extent of radar echo useable for Doppler velocity measurements. Heavy solid lines are more intense reflectivity features. Mesohigh and wakelow are located at "H" and "L". Storm relative air motions are shown by solid arrows. Broad arrow from asterisk indicates hypothesized ice particle trajectory. (Adapted from Smull and
Houze, 1985)
G West - F
I
LU
-I'SU (cU RJ) - OtJ HORIZONTAL DISTANCE FROM LEADING EDGE OF SOIJAIL- LINE EChO (krn)
Sliotiform Reqion Trannilion Convective Reqion Zone
Fig. 1.5. Revised conceptual model based upon new research findings. Note airflow changes, especially those associated with the convective tower. (Adapted from Smull and Houze,
1987b) 13
project. They noted kinematic structures similar to that described by Smull and Houze (1985; 1987 a,b) . Their analysis confirmed the presence of a mesoscale updraft and downdraft, rear inflow jet, and front-to-rear flow. These features were quantified using various single-Doppler techniques. The sloping rear inflow was found to supply an additional negative vertical motion component to the downdraft, whereas the front-to-rear flow provided a positive vertical component to the updraft.
1.4 Research Objectives The focus of this thesis is the analysis of a mesoscale convective system observed during the O.K.
PRE-STORM project on May 28, 1985. The primary objective of this study is to reveal the mesoscale kinematic and reflectivity structure of this fast moving squall line system and its evolution through analysis of single and dual-Doppler radar data. A secondary objective is the analysis of surface data collected by the NCAR Portable Automated Mesonet (PAM II) and conventional WSR-57 radar data, in order to describe the surface characteristics of this squall line system on the mesoscale. Time series portrayals of meteorological field variables observed during the passage of the MCS at various
PAM II stations provide an insight into the various surface characteristics of this convective weather system. 14
The organization of the thesis is as follows. Chapter
2 will provide a brief description of the PRE-STORNproject.
In Chapter 3, an overview of the synopticobservations and satellite imagery is presented. Chapter 4 contains a discussion of the results of the observations and analysis of the PAM II and WSR-57 radar. In Chapter 5, analysis of single doppler data is presented. In Chapter 6, resultsof a dual-Doppler synthesis are presented. Chapter 7 contains a summary of the results of this research andsuggestions for future research. 15
2. PRE-STORM PROJECT DESIGN AND OPERATIONSPLAN
2.1 Program Background and Goals The O.K. (Qklahoma-ansas) PRE-STORNProject of
May-June 1985 was a cooperativeresearch effort between the
National Oceanic and AtmosphericAdministration (NOAA), the
National Center for Atmospheric Research(NCAR), and a number of university groups including theDepartment of Atmospheric
Sciences at Oregon State University (OSU). It was conceived as a combination of theNOAA/Weather Research Program's airborne project to investigate slow-movingnocturnal MCSs and the National Severe StormsLaboratory's (NSSL) convective storm Spring Program (an annualfield program to examine aspects of convective storms). The long-range goals identified forPRE-STOPN were to improve the forecasting of the development,strength, severe weatherprobability, duration and precipitation amounts of
MCSs. Nocturnal MCS5 that typically form in themiddle to late afternoon and became well organizedduring the evening were the main focus of PRE-STORN. The PRE-STORM Project was designed as a preliminary experiment to addresspotential problem areas for the proposed more extensiveSTORM-Central program, as well as its own long-rangegoals. The two potential problem areas that were recognized forthe
STORM-Centzal program were: implementation of newlydeveloped sensing systems and the limited amount of datafrom previous field programs to use in determining precisescientific objectives (Cunning, 1986)
Hence two major goals were identified for the field phase of the program: achieve a reliable and coordinated observing system for investigating MCSs, and collect data necessary to begin preliminary investigationsof the origin, development, dissipation, and structure of MCS5. Among the new sensing systems that were tested todetermine how best to integrate them into the coordinated observational program were wind profilers and airborne-Doppler radar. This thesis, however, will not examine data from these particular observational tools.
2.2 Observational Program Design The preferred region of MCS development and movement
(cf. Fig. 1.1) predicated the location of the observational ground facilities of the PRE-STORM project in the Central
Great Plains. Figure 2.1 is an overview of the entire surface observational network that included NWS conventional radars, digitized NWS conventional radars, NCAR Doppler radars, NWS standard and supplemental rawinsonde sites, wind profiler sites, and PAM II surface mesonet sites.Another part of the surface observational network included a cloud-to-ground lightning location system that was operated by NSSL (National Severe Storms Laboratory) . This network covered a broad area including the states of Kansas and
Oklahoma. In addition to the surface observational network, 17
t'0 1000 950
/-
/ 400 400 / I (an5a5 WlchIta City I ?ot I. . . .Topeka I .f.
IDodg 2TMonen J .: ::,. : Arnahilo*L?I
Qklanoma /
90° 953 105° S..
Existing NWS RawinSonde Sites -- Boundary of MeasurementNetwork ?
Existing NWS WSR-57 Radars PRE-STORM Rawinsonde Sites
NSSL Doppler Radars PJ PRE-STORM Profiler tocattons Existln NWS WSR-57 Digitized Radars PRE-STORM Surlace Mesonet Sites (RADAP II or DtgitIze) -- Dashedline circle indicates range NCAR Doppler Radar Sites of lightning iocaUon sensors
Fig. 2.1. Overview of the surface observational mesonetwork in Kansas and Oklahoma during PRE-STORM. an airborne observational programwas carried out through the deployment of two NOAA P-3 aircraft, oneof which was equipped with a Doppler radar, and a CessnaKing Air from the
University of Wyoming. This thesis concentrates on the analysis of data collected in the Kansas portion of the PRE-STORMnetwork.
The observational platforms important tothis thesis included the NCAR CP-4 5 cm Doppler radar located atCheney State Park
30 km west of Wichita, the NCAR CP-3 5 cm Dopplerradar located near Nickerson, KS, 60 km to the northwestof CP-4, the digitized NWS WSR-57 10 cm radar located atthe Wichita
NWS (National Weather Service) office, and the PANII surface mesonet sites located throughout the state. The characteristics of the NCAR radars are shown in Table 2.1.
The CP-4 site was the coordinating location for the
Kansas portion of the experiment. At this site real-time surface observations from 42 NCAR PAM II stations were available and a color display with time looping capability provided access to the network of NWS WSR-57 radars, as well as infrared satellite observations. Data from these sources provided a much needed nowcasting component of the program and allowed the CP-4 scientists to view storm development beyond the range of the NCAR radars. Communications were maintained via VHF radio with the CP-3 site and the NOAA P-3 aircraft to coordinate Doppler radar scans and aircraft flight patterns, respectively. Table 2.1. Characteristics of the PRE-STORM Doppler radars.
(Adapted from Rutledge et al., 1987)
Parameter Radar Cp-3 C?-4
Wavelength (cm) 5.45 5.49
Maximum range (km) 135 135
Nyquist velocity (msH) 15.37 15.24
Peak power (kW) 400 400
Pulse width (Lsec) 1.0 1.0
Pulse repetition frequency (Hz) 1111 1111
Minimum detectable signal (dEm) -113 -112
Number of range gates 512 512
Azimuthal resolution (deg) 0.8 0.8
Gate spacing (m) 260/l5Ot 260/150t
Number of samples 64/256* 64/256*
tA gate spacing of 150 m was used for the EVAD and.
vertically-point ing scans.
* Number of samples was set at 256 for the
vertically-pointing mode.
A combination of Doppler radar scans wereused in
PRE-STORM. These consisted of 360°, 150°, and 90° sector
scans to cover large and smalldual-Doppler areas, and
specialized scans that included RHI (Range-HeightIndicator), 20
vertically-pointing, and 3-D conical scans. The operating mode of the radars was in one of two basicscanning sequences which depended on the meteorological situation andthe particular scientific objectives being addressed. These modes were defined as the "standard mode" and a"stratiform mode". A "standard mode" scanning sequence consisted of continuous volume scans with a large-area (360° sector) dual-Doppler scan initiated first. After this scan was completed, a series of smaller area sector scans were conducted (either 90° and 150° sectors) to providefiner detailed dual-Doppler data, in both time and space. This cycle was then repeated by returning to a 360°coordinated scan. When a stratiform region associated with a MCS was over one of the Doppler radars, the"stratiform mode" scanning sequence could be initiated. This mode was used if an anvil microphysical study was planned by researchaircraft and/or project scientists agreed to substitute the
"stratiform mode" for the "standard mode" to obtain time-height cross-sections of particle fallspeed and vertical air motion within the stratiform cloud. The "stratiform mode" consisted of one radar being placed in a vertically pointing position to observe particle fallspeed and reflectivity, while the second radar scanned through a relatively large number of elevation angles, which could be 21
used to infer the vertical air motions by the VAD
(Velocity-Azimuth Display) technique. Upon completion of the
VAD scanning procedure, the second radar scannedthrough small elevation steps at low elevation angles orconducted
RHI scans. 22
3. SYNOPTIC AND SATELLITE OVERVIEW
3.1 Surface Synoptic Situation The synoptic setting associatedwith the development and life of the 28 May 1985 MCS wassimilar to other studies done on midlatitude mesoscaleconvective systems. The ingredients necessary for severe stormdevelopment were found in the Nebraska-Wyoming borderregion: low level southerly flow advecting sufficient heat andmoisture into a leeside trough along the Rocky Mountains, a capof very dry air at mid-levels, and an upper level jet streakthat may have forced destabilizing indirect circulation(Beebe and Bates,
1955; Lewis et al, 1974; Maddox andDoswell, 1982; Maddox,
1983) . With these ingredients and the additionalorographic influence of the Rockies, severe convectiondeveloped along the Nebraska-Wyoming border at about1500L on 27 May 1985.
The May 28 MCS was initiated during theevening along the outflow boundaries of this severeconvection as it dissipated in southern Nebraska. Remnants of anvil cloud material from the previous supercell wereincorporated into the newly developing nocturnal MCS (cf.Figs. 3.13, 3.14)
Discussion of the meteorological conditions thatcontributed to the nocturnal MCS developmentfollow, with a subsequent discussion of the satellite imagery of the storm.
3.1.1 0600 GMT 28 May 1985 meteorological conditions At 0600 GMT (0100 L, cf. Fig. 3.1, alltimes are local 23
Central Daylight Time (CDT) on 28 May 1985, unlessotherwise noted, CDT = GMT 5) a surface low pressure troughextended from a thermal low over the Texas Panhandle acrossthe
Rockies to a low over Utah. Lowest pressures were found in a thermal low centered over the Salt Lake Basin. A dynamic warm core high was centered north ofLake Superior, while further south a baroclinic low pressure center was located over Ohio. The Bermuda High dominated the southeast and Gulf Coast States where its characteristic maritime tropicalair mass prevailed. A cold frontal boundary extended east-west from the low over Ohio to the Oklahoma Panhandle where it became a quasi-stationary front that paralleled the Rockies then extended east-west through southern Wyoming into the Utah thermal low. A north-south dry line extended along the Front
Range of the Rockies into Wyoming from the Texas thermal low.
The dry line separated the arid Superior air mass over the desert southwest from the relatively more humid air to the east. A significant moisture gradient is evident across the frontal boundaries, especially the quasi-stationary/dry line front over Wyoming, Colorado, Oklahoma, and Texas.
3.1.2 1200 GMT 28 May 1985 meteorological conditions At 1200 GMT (0700 L, cf. Fig. 3.2) the squall line had moved to central Kansas from northwestern Kansas where it first entered the PRE-STORM observational network at about 24 iøç3 -, j_ Jt4I/ 1 III u1c
- fZ reir
- ? i
'\ 11fi'T H ft '1 .1 ¶3 Fig. .1. 0600 GMT 28 May 1985 Synopticanalysis
I,,-
Fig. 3.2. 1200 GMT 28 May 1985 Synoptic analysis 25
0900 GMT (0400 L) . The squall line was located north of the quasi-stationary portion of the frontal boundary, while the cold front over Illinois and Missouri movedfarther south as the baroclinic high pressure system near LakeSuperior intensified and sagged south over the Great Lakes. The dry line in Texas retreated slightly west, while the remainder of it to the north remained essentially stationary. Thunderstorms and cumulonimbus clouds were observed throughout central Kansas. Frequent in-cloud, cloud-to-cloud, and cloud-to-ground lightning flashes were observed from 0930 GMT through 1200 GMT and logged by radar scientists at the NCAR CP-3 Doppler radar site near
Nickerson. At 1200 GMT the squall line was just about over the CP-4 Doppler radar site at Cheney Lake.
3.1.3 1800 GMT 28 May 1985 meteorological conditions The 1800 GMT (1300 L, cf. Fig. 3.3) synoptic analysis showed that the squall line had passed through central Kansas into northwestern Arkansas. Interestingly, a mesoscale thunderstorm high was observed following the leading edge of the squall line as it passed over a weather reporting station near the Oklahoma-Arkansas border.
The high pressure system over the Great Lakes had moved south of Lake Superior, so that its synoptic scale circulation had begun to draw the warm moist tropical air mass from the Gulf of Mexico north. The northward push of 26
Fig. ------
ai ut' 1.6 ' 6?
\1c
16
Fig. 3.4. 0000 GMT 28 May 1985 850 mb level analysis 27
this tropical air mass created a new warmfront across southeastern Kansas. The squall line was now on the warm air side, being south of the warm frontalboundary.
3.2 Upper-Air Long and Shortwave Pattern Previous studies on thunderstorm initiationcited
differential horizontal advection of heatand moisture as
prerequisites necessary to destabilize thetroposphere and
produce thunderstorms (Means, 1952; Suggand Foster, 1954;
Whiting, 1957) . Maddox (1983) described the synoptic scale conditions associated with the formation of amidlatitude
mesoscale convective complex that formed near KansasCity, MO
on 13 August 1982. In the analysis of that mesoscale convective system, strong horizontal advection of heatand
moisture occurred in the genesis region (GR) of thestorm.
3.2.1 0000 GMT 28 May 1985 upper air pattern Figure 3.4 shows the National Meteorological Center
(NMC) analysis of the 850 mb surface at 0000 GMT. In this
analysis a high amplitude shortwave ridge extended from the
Gulf Coast across the central United States into Canada. The
axis of a shortwave trough extended from southeastern Idaho
to the Texas Panhandle. Cyclonic flow around the trough and into the ridge drew heat and moisture from the southinto the
GR along the Front Range of the Rockies and as far north as
Lake Winnipeg, Manitoba. Note the small dewpoirit depressions observed through this region. Notice also that the largest thermal gradient at the 850 mb level wasfound in the GR.
The 700 mb NMC analysis at 0000 GMT(cf. Fig. 3.5)
shows a broad longwave ridge overthe central US with a shortwave trough aligned with the lee ofthe Rocky Mountains.
Another shortwave trough extended from alongwave trough
centered over Hudson Bay to Mississippi,between these
troughs a shortwave ridge was found.Airflow over the GR and
in southwestern Kansas was from thewest-southwest over the
Rockies and became northwesterly into theshortwave over the
central U.S. Horizontal differential thermal advection was indicated in the GR by theisotherm (-. 3 °C hr )
packing and wind direction. Pockets of moisture were found north of the GR, while arid air was presentin the GR.
The NMC 500 mb analysis at 0000 GMT (of.Fig. 3.6)
indicated that a broad longwave ridge with a weaklee-side
shortwave trough was over the Rocky Mountains. The lee
trough served to flatten the ridge and inducesouthwesterly
flow over the GR. Another shortwave trough extended south from a longwave trough over Hudson Bay throughIllinois to
Mississippi. This shortwave over the Midwest caused northwesterly flow across Iowa, eastern Kansas,and
Louisiana. A pocket of very dry air was present over the central Plains states, and pockets of moisture werefound to
the west over Colorado and Wyoming. Horizontaldifferential
advection of this cooler moist air caused convergenceof 29 -1- - 1 #0.2 03 QL.oq? Q58
- 0 079 13 it-'i #0 "S j0?1 &95 V 13# 01 11 13 q&%1 12 7' #04 1 121 4' Q
:x;i;oi*w&sta0 -i c-7'Ol'-_- #01 Fig. 3.5. 0000 GMT 28 May 1985 700 mb levelanalysis
.1 _i
125?5 4
-c
,
JhJP401 58L1 I *583#O
Fig. 3.6. 0000 GMT 28 May 1985 500 mb levelanalysis 30
water vapor into the GR, hencecreating additional convective instability (Ninomiya, 1971). Note also that a warm dry tongue of air was intruding intoKansas from west Texas.
At 300 mb, the 0000 GMT NMCanalysis (Cf. Fig. 3.7) indicated a 25 m s- jet in the airflow overnorthwestern
Kansas. This places the GR precisely in theleft exit zone
of the jet, a region favorable forintense storm development.
In this region positive verticalmotion induced by the indirect circulation associated with thejet can provide
additional convective instability to initiate severe
convection (Beebe and Bates, 1955; Uccelliniand Johnson,
1979; Maddox, 1983) . A shallow shortwave trough extended
across Montana to Coloradodownstream from the axis of a longwave ridge over much of the western US. Westerly to
northwesterly airflow with cold thermaladvection was
established across the central Plains.
3.2.2 1200 GMT 28 May 1985 upper air pattern Twelve hours later the NMC 1200 GMT 850 mbanalysis
(Cf. Fig. 3.8) indicates that the highamplitude short wave
ridge had shifted eastward. The short wave trough had shifted eastward as well, with its axis extendingtoward the
Texas coast. The southerly flow upstream of the ridge axis continued to advect heat and moisture into thecentral US
primarily along the eastern borders of Texas, Oklahoma,
Kansas, Nebraska, and North and South Dakota. Note that the 31 4oThi
#03
-9 lx ,' 9561 4n4 -,958 962 X#04 960 J958#02-1 27 * Fig. 37. 0000 GMT 28 May 1985 300 mb levelanalysis
ii'o,,
#Q2 1T'k3 I
q N ft;%9 \iiLjç-;
2ht511 1T 519
Fig. 3.8. 1200 GMT 28 May 1985 850 mb level analysis 32
strongest thermal gradient now appeared overcentral Kansas, precisely in the region where the MCS of interest was located. A tongue of warm arid air intrudedinto central Oklahoma, thus helping to tighten the thermaland moisture gradients. At 1200 GMT, the 700 mb NMC analysis (cf.Fig. 3.9) showed that the shortwave trough over themidwest had moved eastward. A less distinctive lee-side trough was foundalong the Front Range and the shortwave ridge wassituated between the troughs. Airflow predominated from the northwest over the entire length of the path traversed by the MCS, exceptin southwestern Kansas where southwest flow prevailed. Thermal and moisture gradients were weaker at this analysis time.
The 1200 GMT NMC 500 mb analysis (cf. Fig. 3.10) indicated that the shortwave trough over the Midwest had shifted eastward. The axis of the broad longwave ridge had also shifted eastward, causing the lee-side trough to fill.
The dry pocket of air had advected east into southern
Illinois, Missouri, Arkansas, and Louisiana. Ample moisture was now found over Nebraska, Iowa, Kansas, and Oklahoma. Confluent airflow is found over central Kansas southeastward to Arkansas, where southwesterly flow induced by the remnants of the lee-side trough in western Kansas and Oklahoma met the northwesterly flow into the shortwave trough to the east.
By 1200 GMT, the 300 mb NMC analysis (Cf. Fig. 3.11) 33 19'-&o'j .tJ
7 TJicoo 7 #01 t'It080 rx #bi
J_-. ..--.
i i-__1.2. ii. 15I 1h18 Q #0! Fig. 3.9. 1200 GMT 28 May 1985 700 mblevel analysis 7--el 4± 16 -0 #02 2-C1 /12 #oo\ 1
11-576 c
3581
I £l.,L ° / -hS8 ___ ..x -.01 X __! .iC#DJ7 '1585 5S3 1
Fig. 3.10. 1200 GMT 28 May 1985 500 nib levelanalysis 34
indicated the broad longwave ridge overthe western U.S. had moved eastward. Preceding it, the shallow shortwavetrough had also translated eastward towardthe Great Lakes. The prevailing direction of the airflow wasnorthwesterly, and
cold thermal advection was no longerpresent over the central
U.S. The 50 knot jet, analyzed within theGR 12 hours earlier, was no longer present.
3.3 Satellite Classification of MCSs Maddox (1980) used enhanced infrared (IR)satellite
imagery to classify one type of MCS, whichhe referred to as
a mesoscale convectivecomplex (MCC) . His classification scheme relied on the physical characteristicsof the cold
cloud shield, its size, duration, andeccentricity, as
measured using enhanced IR satelliteimagery. The physical characteristics Maddox (1980) used to define a MCC are
attributable to the largest mode of MCS development.His MCC
criteria are useful as a framework to describe andclassify a
spectrum of MCS development in comparison tothe longevity
and size of the MCC. Table 3.1 shows his MCC criteria. Figure 3.12 shows an example of how the enhanced IRsatellite
imagery is used to obtain cloud top temperatureinformation
necessary to infer the strength andintensity of the evolving
MCS and classify the MCS according toMaddox's criteria. 35 -!:9;4, ') \ -03 ---- U.,91I1\/ 44.l_9LI2/ /: rrQ ,m'1109q295l-''\Qi_(
n /L!13..9.95OX-L._.----!-oj9i.lL I -3.952 I,
1Lh1.58 12 -01' rfp956 #Q1 95? X -02 xMh-i - 58 / %S<03 - .-#00-___-J9
Fig. 3.11. 1200 GMT 28 May 1985 300 mblevel analysis
Table 3.1.Size and duration criteria forclassific?tion as an MCC. (Adapted from Maddox, 1980) Physical characteristics
Size: ACloud shield with continuously low IR tempera- ture 32°C must have an area 100 000km2 BInterior cold cloudregion with temperature -52°C must have an area 50 000 km2 Iniiae: Size definitions A and B are first satisfied Duration:Size definitions A and B must be met for a period >6h MaximumContiguouscoldcloudshield(IR temperature exen: 32°C) reaches maximum size Shape: Eccentricity (minor axis/major axis) 0.7 at time of maximum extent Terminaic:Size definitions A and B no longer satisfied MEDIM GRAI..- 1L1GMtRAL
I/
Fig. 3.12. Satellite infrared image showingthe enhancement
curve temperature ranges thatcorrespond to the various
shades ofblack, gray, and white. Colder cloud topsinfer
stronger more intense convection. (Adapted from Maddox,
1980)
Fig. 3.13. 0000 GMT 28 May 1985 infrared satellite image. 37
3.4 Satellite Imagery of the 28 May 1985 MCS Visible and enhanced infraredsatellite imagery captured the organization anddevelopment of the 28 May 1985
MCS. At 0000 GMT (1900 L, 27 May 1985) agroup of orogenic supercell thunderstorms (Wetzel etal., 1982; Fankhauser,
1971) had collected over theNebraska Panhandle, between
North Platte and Scott's Bluff, toform the convective component of a MCS with a distinctlycircular cloud shield
(cf. Fig. 3.13) . The Great Plains drop over 1050feet in
less than 150 miles here, creating anup-slope topography,
and a prime location for orogenicthunderstorms. Enhanced infrared images showed cloud top temperatures< -75°C, an
indication of strong mature convection in progress. An arc
cloud line through northeastern Coloradorevealed the
presence of gust front outflowemanating from the decaying
supercells. It is along these outflow boundariesthat the convection associated with the developing MCS wasinitiated.
At this point the MCS had not grownsufficiently large to
meet the classification as establishedby Maddox (1980) By 0300 GMT (2200 L, 27 May 1985) anearly-rounded
circular cloud shield was apparent (cf.Fig. 3.14) . The
extent of the coldest cloud tops at-75°C had decreased in
area from over 10,000km2 to < 3,600 km2, but the area
covered by the growing contiguous circular coldcloud shield
had increased dramatically to > 100,000km2. It was about Fig. 3.14. 0300 GMT 28 May 1985 infraredsatellite image.
Fig. 3.15. 0600 GMT 28 May 1985 infrared satelliteimage. 39
0300 GMT that the MCS had achievedMaddox's (1980) MCC size criteria for initiation. The newly developed MCS had translated southeastward to theKansas-Nebraska border at an average speed of about10 m s.
At 0600 GMT (0100 L, cf.Fig. 3.15) the extent of the coldest cloud tops had increased from< -75°C to about -62°C, an indication that the remnantconvective towers of the tornadic supercell were dissipating. However, there was no decrease over the previous three hoursin the size of the circular cloud shield as the MCS continued todevelop along the outflow boundaries of the decayingsupercell. These outflow boundaries now penetrated into Kansasclose to the northwest corner of the PRE-STORM mesonetwork. Over the next three hours the MCS grew tremendously insize and reintensified. The 0900 GMT (0400 L, cf. Fig. 3.16) enhancedIR satellite image shows the dramatic increasein the size of the circular cloud shield, doubling itsareal coverage and signalling an increase in the intensity of the MCS. The enhanced IR image indicated that the areacovered by the
-62°C minimum cloud top temperatures had enlarged overthe previous three hours from 15,000km2 to 75,000 km2. The concurrent synoptic analysis now carried asquall line through north central Kansas, which would place thesquall line along the southern border of the cloud shield. At about this time, the Wichita WSR-57 radar and PAM IImesonetwork 40
-
.-
\
-' '4
- Fig. 3.16. 0900 GMT 28 May 1985 infraredsatellite image.
--:
Fig. 3.17. 1101 GMT 28 May 1985 infrared satellite image. 41
stations had begun to observe thesquall line as well. Near
0900 GMT the anvil cloudassociated with the MCS had grown to near its maximum sizewith a contiguous cold cloudshield (IR temperatures <-32°C) in area of>225,000 km2, and an eccentricity of about 0.8. The required duration criteriaof maintaining the size characteristics"A" and "B" (cf. Table
3.1) for a period of 6 hours had been met by thistime,
hence this storm could beclassified as a MCC. Satellite
imagery showed that the physicalcharacteristics of the storm
remained in a quasi-steady state asit moved through Kansas
toward northwestern Arkansasuntil its dissipation around
2300 GMT in northern Misssippi. At 1101 GMT (0601 L, cf. Fig.3.17) the satellite
image showed an increase in size of theanvil cloud to about
300,000 km2. The minimum cloud top temperaturehad decreased
to<-70°C. The MCS, as indicated in satelliteimagery, is
now at its most intense stageconvectively and approaching
maturity. The southeastward motion of the MCS hadincreased
from 10 m s at 0300 GMT to about 20 ms1. The visible satellite image at 1331 GNT(0831 L, cf.
Fig. 3.18) showed an arc cloud lineassociated with a gust
front emanating from the squall linein northcentral
Oklahoma. In the wake of the MCS, a wide swathof relatively
cloud free air was observed. This has been explained in
other cases as due mostly to thestabilizing effect of the
mesoscale downdraft (Ogura and Liou,1980; Leary and Houze, 42
i__. -' - -
Fig. 3.18. 1331 GMT 28 May 1985visible satellite image.
- '*_ t
-a
-g
i -
--- - .--
Fig. 3.19. 1501 GMT 28 May 1985infrared satellite image. 43
downdraft in 1979) . Evidence for the presence of a mesoscale this MCS will be presented in Chapter4. Frontal boundary clouds can be seen extending from NewYork toward the Ohio
Valley through southern Illinois tocentral Missouri.
At 1501 GMT (1001 L, cf.Fig. 3.19) the MCS was positioned in the warm moist unstableair found aloft of the quasi-stationary front. The size of the nearly circular
anvil cloud had decreased somewhat toabout 200,000km2. The
areal coverage of the minimum cloud toptemperatures, which
have increased to -62°C from <-70°C, has remained
quasi-steady at about 85,000km2 from 1101 GMT. An arc cloud
line created by outflow from the squallline continued to be
evident along the southern border of theanvil cloud shield.
At 1801 GMT (1301 L, cf. Fig.3.20) satellite image
showed that most of the NCS had moved outof Kansas and had
entered Arkansas, passing into the warm sectorsouth of a
warm front that formed from theearlier quasi-stationary
front. The 1830 GMT (1330 L, cf. Fig. 3.21)visible satellite
image shows a striking example of the relativelycloud free
wake with clouds forming a horseshoe aroundthe wake behind
the MCS. This notch-like cloud pattern may beindicative of
strong rear inflow of dry air.
The enhanced IR satellite images from2300 GMT (1800
L, cf. Fig. 3.22) indicated that the MCShad begun to
dissipate. The area comprised of the cold cloudshield had - i?
: r
. :- 'I r ..'. &4
AJ
I3I4'
:t 4 I.,
1
t
I
I ' ? 'I.
J I.. ... - 45
- .--.
shrunk and was no longer circular. This destruction of the once continuous upper-level cloudshield may be related to the inflow of dry air into the rear of the MCSin the form of a rear inflow jet (Smull and Houze,1987b) . Note the finger-like areas of warmer cloud within the cloud shield, probably associated with the rear inflow of dry air. At this time the system had moved into the synoptic scale cyclonic circulation of a midlevel shortwave in the region (cf.Figs
3.5, 3.6, 3.9, 3.10), thus relinquishing its uniquely mesoscale characteristics to the larger scale. However, its classification as a MCS was discontinued once the system was enveloped by the synoptic scale circulation. The squall line remained intact into northwestern Mississippi. 46
In all, this mesoscale convective systemtraversed a distance of over 2000 km retaining itsidentity in a quasi- steady state for more than 12 hours beforeits dissipation in northwestern Mississippi. 47
4. MESOSCALE ANALYSIS
4.1 Discussion of the WSR-57 Radar Observations The National Weather Service WSR-57 10-cm radar is a conventional non-coherent radar, designed to obtain echo location, and intensity. Non-coherent radars do not utilize the phase differences present between the outgoing and incoming radar waves as do coherent Doppler radars. Doppler radars take advantage of this phase shift to determine particle speed along the direction of the radar beam. The
Wichita WSR-57 radar was digitized for PRE-STORM through the addition of the NOAA/Hurricane Research Division digitizer, which provided for the recording of reflectivities at 2° azimuth intervals and 1 km range intervals.
Contour plots of radar reflectivity (dBZ) depict the growth and decay of the MCS with its squall line and trailing stratiform region. In Fig. 4.1, the low-level reflectivity pattern obtained from a 1° elevation FF1 (plan position indicator) is shown. At this time the storm is characterized by a leading edge of convection (> 30 dBZ), which proceeded a region of weaker echo associated with the developing stratiform region. It should be remembered that a portion of the increase in reflectivity shown in subsequent radar echoes
(cf. Fig. 4.2-4.4) may be due to the echoes moving into closer radar range where they are sampled by a decreasing radar beamwidth, which means that the power returned per square area of the beam coverage will increase as the same 28 May 1985 WSR.57 1008 Z 200
E
Ia
0 0 t 0 C a C a
-200 200 0 200
Dustasco east ot Wichita (turn)
Fig. 4.1. Observed reflectivity at 1008 GMT 28 May 1985 from
Wichita NWS WSR-57 radar. The elevation angle of this PPI and all that follows is 10. Light to darker gray shading every 10 dBZ beginning with 20 dBZ. Contour interval 10 dBZ beginning with 10 dBZ. 49 amount of power is returned through a smaller area. This creates an apparent increase in reflectivity, which is a result of the radar sampling technique and not the growth or increase in strength of the actual radar echo. Forward of the northern part of the line a leading stratiform region is developing that can be attributed to outflow of hydrometeors from the convective towers into an overhanging echo. This feature of severe convection has been described by Browning and Ludlam (1962) and Newton (1950,1966), and is likely associated with strong divergence from the upper levels of the convective cells. Figure 4.2 shows the low-level reflectivity pattern at
1127 GMT. This observation indicated a breakup of the solid convective line echo. Instead of the continuous convective line echo shown in Figure 4.1, two individual convective regions in a north-south orientation are now present. The breakdown of the solid convective line was observed just after sunrise. Sunrise occurred at 1113 GMT in centraal
Kansas. Mesoscale analysis of PAM II data at 1127 GMT, to be discussed in the next section, showed the presence of strong temperature and pressure gradients. It has already been shown in the satellite imagery that the MCS was approaching or was near a stage of maximum convective intensity (Cf.
Fig.3.17). The increased temperature and pressure gradients may be a result of an intensified mesohigh brought about by a strengthened evporatively cooled downdraft as a result of 50
May 1U3 W.I1-! lIZ! Z 200
E
24 2 CP-4 0 0 R- 57 C C
C, C
-200 0 200 Distance east of Wichita (km)
Fig. 4.2. As in Fig. 4.1 except for 1127 GMT. 51
increased convective instability associatedwith the increase in the lower tropospheric lapse rate inthe pre-storm emnvironment just after sunrise.
The stratiform region had increasedtremendously in areal coverage during this period and asecondary band of enhanced reflectivities developed along anorth-south line
from near coordinates -65 km, 55 km to-65 km, 115 km.
Fritsch and Chappell (1980) describe thisincreased low-level
horizontal convergence and therefore thestrengthened mass,
moisture, and temperature advection as "dynamicenhancement".
There was an unfortunate lapse in WSR-57 coverageof
the MCS during the period between 1127 GMTand 1356 GMT due
to hardware problems. The reflectivity distribution at 1356 GMT from the
WSR-57 radar is shown in Fig. 4.3 (the areaoutlined by the
rectangle encompasses the dual-Doppler analysisregion
discussed in Chap. 6) . A continuous convective line echo was
again present. The strongest reflectivities exceeded 50 dBZ
at 55 km, -105 km. No stratiform region was evident along the developing end of the squall line. The stratiform region
extended northward from a bow-like echo. The stratiform
region width had narrowed somewhat from the 1127 GMT
observation, but range effects at the earlier observation
time may have produced an artificially broader stratiform
region. An overhanging anvil preceding the convectiveline 52
28 May 1985 WSR-57 1356Z 200
E
-C
t 0
-200 200 0 200
Distanco Qast ot Wich5a (km)
Fig. 4.3. As in Fig. 4.1 except for 1356 GMT. 53
28 May 1985 WSR.57 1457Z 200
E
0 C 0
0
-200 200 0 200 Distance east ot Wichtta (km)
Fig. 4.4. As in Fig. 4.1 except for 1457 GMT. 54
is apparent from near the bow echo north.
A decrease in the size of the secondary band at1457
GMT (cf. Fig. 4.4) may indicate its dissipation. The maximum reflectivity of the developing cells has decreased to 49 dBZ from 54 dBZ indicating a possible decrease in the strength of the convection. The bow echo continues at 185 km, -55 km with a notch in the echo extending rearward towards 75 km,
-85 km where reflectivity contours indicate an erosion of the trailing stratiform region cloudiness from evaporation associated with the intrusion of the low Orear inflow jet
(Smull and Houze, 1985; 1987a,b) . The presence of a dry rear inflow will be established in the following chapter.
4.2 The PAN II Surface Mesonetwork The experimental design of O.K. PRE-STORM included the deployment of the NCAR second generation Portable Automated
Mesonetwork (PAN II) . Forty-two of these automated meteorological stations were sited throughout the state of
Kansas and along the northern border of Oklahoma (cf. Fig.
4.5) . This siting created a surface array formed in a rectangular grid of approximately 400 X 250 kilometers with a
50 km individual grid spacing. An array of this size is large enough to capture the broadscale features of mesoscale convective systems (Toth and Johnson, 1987)
Each PAN II station contained a host of meteorological instruments that sensed the standard meteorological variables -a 'Q H- 0 96° tjCDrt . 400 40° oIt)ao'1 CDH PRE STORM c_ ;?c PAMILMESONET PAM ii Mesoanalysis coverage :io pi0 - +o NWS Station- PAM Site HLC MIlK 39° o(I) 0:j- NCAR Radar °RSL SLN 7 Doppler radar o"\Conventional (I)p.) 12 14 15 radar(wsR-57) coverage 2 - rtoH '-0(n 380 °GCK f + °H UT Ii CJjtO I +17 °DC +\18 f I +20 +j21 +22 NU .oQ ø 0H 426 28 + 2 '-0 37° -- °LBL +33 4 40 BV I I-ICDFl °GAG 4o p TUL 0(I) 102° 101° 100° 99 2 i 97 I 96 / rtCD t-1 Longi5N63O0 tudi 0 01 56
Table 4.1Accuracy of sensors on the PAM IIstation. PAIl Parametsrs
Parmmeter Sensors ACcurpcv Resoiut:r,
+5% each for 0.1 nS Wind (U, v components) Two orthogonal 2 propeller anemeaent.rS winds > i mm
Dry-bulb temperatur. WcAR psychrometer +0.25 0.05°C
Wet-bulb temperature NCAR psychrometer 0.01°C
?ressure NCAR barometer mb 0.02 ru
Rain Tipping bucket +l5%, 0.25 nm
1. 2 ubOunda. 2. Stalling threshold is approximmt.ly 0.6 s/s. Speed arid direction computation errors are maximized near the cardinal ir.ctions. 3. Plus radiation errors which may at times exceed 0.5 C depending on radiation lsvelm and siting conditions. 4. Plus the 0.25 mm quantizing interval. These figures do not include wind caused errors. S. Operates above 0°C. (AdaptedfromJohn,ot and Coth, (986).
91*(.91: ive.iI.. iWJ'3wits,. not Sm..in mice:,.5(0 .un.
.0 y.ismits p.r SSflIt' CM' e9VMlISifl1° ,..Ia itriem,Ii I i4 ICSQ it'.
10 PoCMe
__ 9W (02: p..rIit a S 40 nuts if pa eras. - ,tae:,pit t0415 paC.Sl7. 'stowsfrei.sitn SUSIe,554 S.OXO..S.Illt..
PW$2 PUT:(itt-el wewPr 0rp .n nut 5.24 fl.Owt.9,551n.IC V...... 1P1LP1eteant at 1 attn. : wtrm,5.lof - .5.51.,.auttic.11y a's,,rflg .9elm.C.Il.ctsintLfmn MW, line.MS ?WISUSU ,.0 TIp,um SUET MIS tS i.0e.t.ilii.. Z II. Jenli htmlm,.awqlc ,.,...n.
Spint cUT9915 55 *.liG tst0 Poll?:flUent S S LIT_i Suite. Pa let0' Client I_i...Pr me. .1 ..tIe. nit .5.0's a.isiftMt(I's* ..n.iw in .:1nu.d iSQ..Sct.Ii,. .1 I_i 'sy. - SQSm, outs. Prim,nn
Fig. 4.6. NCAR Portable Automated Mesonetwork Station.
(Adapted from Johnson and Toth, 1986) 57 of horizontal wind, dry and wet bulb temperatures, pressure, wind maximum, and accumulated rainfall. From these standard variables, important derived quantities such as equivalent potential temperature, dewpoint temperature, potential temperature, mixing ratio, and relative humidity can be derived. To ensure the quality of the PAM II data all datasets collected by the surface array were corrected to remove significant instrument errors and systematic biases in the data (Johnson and Toth, 1986). Sensor accuracies are shown in Table 4.1. The position of the individual sensors on the
10 m PAM II tower is shown schematically in Figure 4.6.
4.2.1 GEMPAK and PAM II data analysis methods
To analyze the data collected by the PAM II surface mesonet, the National Aeronautical and Space Administration's
GEMPAK meteorological data analysis program was used. GEMPAK contains all of the standard data analysis routines employed on surface and upper air data (desjardin et al., 1986)
GEMPAK provided a Barnes objective analysis scheme to produce a gridded dataset of either scalar or vector quantities. The weighting factor was computed using S.
Koch's formula:
weight = (5.051457*(deltan*2./lr)**2)/deltax**2, (1) 4:] where deltan is the station spacing and deltax is the x grid spacing. Grid spacing for the analysis was set at 20 km by 20 km to take advantage of the PAM II deployment of stations in a grid about 55 km by 55 km. It must be remembered during the course of the analysis that the actual data resolution is
55 km. Boundary values were determined using an extended grid containing data outside the data analysis grid. GEMPAK uses centered finite differencing on theinterior grid with backward and forward differencing along the boundaries to calculate values at each grid point.
To input the PAM II data into GEMPAK for analysis it was necessary to format the PAM II data as a surface data file. This was accomplished by an intermediate computer program. Five field variables were input into GEMPAK: temperature, dewpoint temperature, pressure, and the east-west and north-south components of the horizontal wind.
From these five variables all the necessary derived quantities were easily computed with GEMPAK functions.
The figures in the next section correspond in scale and time to the figures in section 4.2 of the WSR-57 radar reflectivity.
4.2.2 Analysis of PAM II data
Figures 4.7a-d show the GEMPAK analysis of the PAM II data at 1010 GMT. In Figure 4.7a the thunderstorm high
(Fujita, 1955) can be seen (in the upper left corner) 59
28 May 1905 WSR-57 1000 Z 200
1
8 CP-4 ! awn- 57
9
-200 a) -200 0 200
Di,I..,c. .al 01 Wchfl m)
20 May 1905 WSI?-57 1008Z 200
I
851?. 5?
I
1.. -200 0 200
Fig. 4.7. Mesoanalysis of PAN II data at 1010 -IMT 28 May 1985
Reflectivity contouring as in Fig. 4.1. a) pressure (mb) b) surface divergence(x105 1) '0 May 1905 WSR-57 1000 Z 200 /)'t
I
4 '0 5,
c) -200 0 200 Oc..aofWch4.N.,I
Iay 1905 WSIT.57 I000Z 200 'ç, N.. . . !S N N.. N. L -(
/N.. iç N. N.. N.. N.. N..4. .4
4.. N. N.. N. N. 4.4-4-c4.44.4.5 4444.4.4.44-4. 1.44.4-4-44.*44-4-4.4 I l 4*4* 4- -'- 4 WSII- 57 '0 1. * 4. .,j .g A 4 <<4.4444 r-7 77 r<<<<<<<<< 4 S p. p. p. F FF FFF4- 4- 4- p P. 1 F
4. NP-P-NrrrFF N NFFFF
N NP-p-Fr4. N NP' P.. N F FF A #- r rrF N N N A?- r
N N 4. N N N N N A A A P. FF
NNN N N N NN N N N * A A p.. p.. r
N N N N Pc N N. N. N. N N a. a. A p.. p. F
s. . fr. . k- N NN N N 4 -200 ----. 200 d) -200 0 0i,4.n. ... .1 W5,4. ..4 Fig. 4.7. continued. Mesoanalysis of PAM II data at 1010 GMT
28 May 1985. c) relative vorticity(x105 1) d) ground relative surface wind field (m s), vector length proportional to speed, shortest = 1 ms1, longest = 9 m s1 61
following the convective line. A strong pressure gradient is found across the developing end of the line,between the mesohigh and a thermal low in the Texas Panhandle. The pressure gradient force in thisregion is 4.5 x io m the largest in the analysis domain. An area of strong surface convergence at -12 x iO s1 is collocated just ahead of the convective line echo. Surface convergence is found along the entire convective line. Behind the convective line the airflow is divergent with maximum values of 8 x 10 s in the trailing stratiform region. This divergence pattern is consistent with the presence of a mesoscale downdraft situated at low-levels in thestratiform region. The existence of a low-level mesoscale downdraft will be confirmed in a later section. Figure 4.7c shows a 16 x105 area of relative vorticity collocatedwith the bow echo, while an area of -6 x 10 s relative vorticity is found 30 km south of the leading edge of the developing convective cells. Figure 4.7d shows the ground relative surface wind field. North-northwesterly flow with cyclonic curvature into the region associated with the bow in the convective line echo is seen. The maximum speed in the north-northwesterly flow is 9 m s at -185 km, 115 km.
Elsewhere across the analysis domain southeasterly flow dominates. The PAM II analysis at 1130 GMT is shown in Figs.
4.8a-d, 4.9a-d. The pressure field (cf. Fig. 4.8a) shows the 62
28 May 1905 WSR-57 1127Z 200-
9
I ,,/1'
I 2 24 - 'I 10
200 0 200
0,1.... Cit 1.4 W.1lit. 0'.l a) 20 May 1985 WSfl-57 ll27Z 200
I
n-u
200 200-200 0 .. b) 0,1.... nit Cl I1d. 01.11 Fig. 4.8. Mesoanalysis of PAM II data at 1130 GMT 28 May
1985. Reflectivity contouring as in Fig. 4.1. a) pressure (mb) b) surface divergence(x105 s) 63
20 May 1905 WSR-57 11272 200-i
I
I
.1
c)
-200 20 0 D.44... ..a o4 Wh4 (k..)
20 May 1905 WSR-57 1127Z 200
-
: : : : : :::::
- t' 7 I
. 4-. 4- 4 4 . b. )- I.- ¶ .
4 & 44 r r "
- r r ¶ kk' I' '4I V < a v - L < a / c,a. ! 0. PISO. 4 44<4-(4< I//I v-I.. 4444<<4< j/ I,_LA>4 t 444<44<4: ,
It . ', I.- 4 4 4 4 4 4 4 4 1 4 4.4.4' <<414<44<4< 1 I- ?- ¶414<4<4< 4.- ¶ 14<4<4< ,, v- a: I.- r 1<444< a: ¶44<44 -200- ...., d) -200 0 200 .4 Wh4. ) Fig. 4.8. continued. Mesoanalysis of PAN II data at 1130 GMT
28 May 1985. c) relative vorticity(x105 s) d) ground relative surface wind field (in 1),vectors as in Fig. 4.7. 64
easterly propagation of the mesohigh with an elongated area
of higher pressure following the leading edge of the MCS.
This mesohigh is similar to that described by Fujita (1955),
cf. Fig. 1.2. An inspection of the analysis of temperature and moisture gradients, as well as temperature and moisture
advection (cf. Figs 4.9a-d) shows that both the gradients and
the advection of moisture and temperature are strong during
this time period. The temperature advection peaks at -18 °C
hr and moisture advection peaks at -2 g/kg hr. The observed advection of cooler drier air into the rear of the
convective line may account for the observed increase in
pressure, as well as the suppression of convective cells in
this portion of the squall line. An observed horizontal
pressure gradient force of 2.5 x iO ms2 is accelerating
the flow outward from the mesohigh toward the leading edge of
the MCS. The first indications of the wake depression or
mesolow (Fujita, 1955; cf. Fig. 1.2) are found in the
northwest corner of the analysis domain where the pressure
decreases from 953 mb to 950 mb. The mesoscale pressure
systems with particular emphasis on the wake low will be discussed in more detail in the section to follow.
Convergence along the developing end of the convective
line has increased since the previous analysis as shown in
Figure 4.8b. The strongest divergence in the stratiform
region at 20 x i0 s1 is collocated with the 30 dBZ contour
of the secondary band. 65
20 May ?905 WSR-57 1127Z 200
Li
3 2 cp2 0-
200- a) -200 0 .. 01 W00. U"0)
20 May 1905 WSII-57 1127Z 200
ty
I
9
I a
-200. 200 b) -200 0 0011.00. ..a .1 Wda. U"") Fig. 4.9. Mesoanalysis of PAIII data at 1130 GNT 28 May
1985. Reflectivity contouring as in Fig. 4.1. a) temperature field (°C)b) mixing ratio field (g/kg) 28 May1985 WSfl-57 1127Z 200..!
a
,lIl1
200 -, 200 C) -200 0 0.,. . .1 a. ...j
28 May $985 WSR-57 1127Z 200
0 I i1i
-200. -200 0 200
..a ol W.th $..) Fig. 4.9. continued Mesoanalysis of PAM II Qata at 1130 GMT
28 May 1985. C) temperature advection (°C hr-) d) moisture advection (g/kg hr-) 67
Figure 4.8c shows the relative vorticity center of -8 x l0 s1 located along the trailing edge of the
stratiform region. This negative vorticity center is found where anticyclonic outflow is occurring below an area of mesoscale subsidence (downdraft)(Ogura and Liou, 1980; Smull
and Houze, 1985) The ground relative wind field is shown in Fig. 4.8d.
From the mesohigh toward the convective line a maximum speed
of 9 m s is found where the strongest pressure gradient
force occurs. Another pressure gradient accelerated flow is
apparent from the mesohigh toward the wake low. The center
of the outflow emanates from the area of themesoscale
subsidence near the secondary band in the stratiform region.
In the pre-storm environment the airflow is
east-southeasterly bringing higher e air toward the
convective line.
Analysis of PAM II data at 1415 GMT is shown in
Figures 4.lOa-d. The pressure field, Figure 4.lOa, indicates that the mesohigh associated with the dissipating cells in the northern portion of the squall line can no longer be maintained. There is no longer any mesohigh associated with the convective line, except for a small portion of its southern developing end. The highest pressure is now ahead of the line where a synoptic scale dynamic high is situated to the northeast. Something of a minor pressure excess associated with the compensating subsidence is found behind 28 May 1905 WSR57 1356 Z 200. N 11K ':7
3
0. )
a) 0 200 .4 WN. .)
20 May 1985 WSfl-57 1356 Z 200
I jo
I
-200 b) -200 0 200 O..4n.. .4d Wt.a. ..N Fig. 4.10. Mesoanalysis of PAM II data at 1415 GMT 28 May
1985. Reflectivity contouring as in Fig. 4.1. a) pressure (mb) b) surface divergence (x 10 s1) Af-,., 1oJc wcn.c7 115, 7 200
c) -200 200 -200 0 D.%I,,c. .flt of Wci3Oa fo,
28 May 1985 WSfl-57 1356 Z 200
I, !
,_ l '_,
"c ' ' ' c < ' <
! 'S S o <
IS 'S 'S 'S 'S J.
'S 'S S 'S ç L L
a E E--:;Ei
4 6 6- 4- 4- 4- 4
4 4- - 4- 4- -- 4- 4- 4- 4-,JI1I 4- 4- 4-4- * 4-4- 4- 4- 4- 4- 4-
I- V 4- 4- 4- 4- 4- 4- 4- 4-
4- 4. 1 'I
4- 4- . 1- -1 4.- 4- 1.- 4- 4- 4- 4- 4- 4-k 44-k4-4444-4-4-4-k 4. k 42'444-4-4-4-4-4-4.4 L- -.1 d) -20- - 200 -200 0 , ffch. Fig. 4.10. continued. Mesoanalysis of PAM II data at 1415
GMT 28 May 1985. C) relative vorticity (x s) d) ground relative wind field (m s), vectors as in Fig 4.7, except longest = 6 m s. 70 the developing end of the squall line. This area of pressure excess is coincident with the developing end of the squall line that lacks a stratiform region. The wake low can be identified by the 948 mb isobar found trailing the stratiform region to the northwest. The wake low appears to be broadening and losing its mesoscale characteristics as it is encompassed in a developing synoptic scale low pressure trough.
Figure 4.lOb shows that the surface divergence field has weakened considerably. The area of peak convergence has decreased 75% from the 1130 GMT analysis and is now located behind the convective line at 15 km, -5 km instead of coincident with its leading edge. Well behind the line in the northwest corner of the domain a small area of convergence associated with a 20 dBZ area of isolated convection was situated. In the area along the convective line divergence was found with peak values on either side of the bow echo that are 60% less than 1130 GMT analysis. This divergence is likely due to the observed pressure gradient acceleration of the flow - 4 x m away from this area.
An area of negative relative vorticity is found near the trailing edge of the stratiform region. Figure 4.lOc also indicates an area of positive relative vorticity associated with the wake low.
Figure 4.lOd shows the wind field at 1415 GMT. An 71 area of subsidence continues under the northernportion of the stratiform region as indicated by a wind minimum at 1 m s. The maximum wind speed, at 6 m s-, is locatedbehind the bow echo at 65 km, -45 km near the trailing edge of the stratiform region. Another area of maximum winds, at 5 m s1, was found where the observed pressure gradient force -6 x io3 m accelerates the flow into the wake low. Analysis of PAM II data at 1515 GMT is shown in Figs.
4.11a-d. Recall that the secondary band of the stratiform region had dissipated somewhat, while the developing end of this squall line continued to produce new convective cells that propagated along the line and diminished (cf. Fig. 4.4).
The satellite imagery shown in Fig. 3.19 indicates that the
MCS continued to look intense, but had decreased in size.
The pressure field in Figure 4.11a is similar to the 1356 GMT analysis in that the mesohigh continued to be associated with developing end of the convective line now just across the border into northern Oklahoma. The mesohigh continued to lead the convective line. The previous north-south pressure gradient had become east-west with a trough of lower pressure through southern Kansas. The portion of the convective line past the bow echo straddles the pressure gradient with the northern portion of the stratiform region into the low pressure trough.
Divergence from the mesohigh at -35 km, -115 km is shown in Figure 4.11b. Flow around the MCS and into the low 72
28 May 1905 WSR-57 I457Z 200
cP:3 :
I
0
I
a)
-20() IF '0 0 200
fl... ..0 04 0h0. 00004
200
I 0 2I ;' 0
I
-200 200 b) -200 0 Oi.Oon.. ..0404W40. 00.) Fig. 4.11. Mesoanalysis of PAN II data at 1515 GMT 28 May
1985. Reflectivity contouring as in Fig. 4.1. a) pressure (mb) b) surface divergence (x iO 1) 73
ray 1985 WSfl-57 1457Z 200
0 N
C) -200 200 0 200
200
4 4 VV V
44 4 0 j V V I 4 F 6 4 4 V V P 4 4 4 4 4 g * F F 44 44 , * k zCP4 4 4
F F 6 6 4 4 4 4 41 a F 4 4 4 4 4 4 4 CF-4 ), / çWSq.s jo F F6 6 6 ç ç i- F FFF F F F F F F FF ! 55FF F F F F F F FF
6- 5 F FF FF 6- 6- 6- 6- 5 e 4.666 FF551$.
' ' ' 4- ' ' 44 *4.4' 4. 4- 4- .
4-4-4-4-4-44 6-6- 4. 424-4- 2_
-200 d) 200 -200 0 0.lon.. 01 WI,0. 10,,) Fig. 4.11. continued. Mesoanalysis of PAM II data at 1515
GMT 28 May 1985. c) relative vorticity (x 1O s) d) ground relative wind field (m s1), vectors as in Fig.
4.7, except longest = 12 m s-. 74 pressure trough had created an area of -12 x105s surface convergence. Figure 4.11c shows the relative vorticity field. An area of 20 x io- s positive relative vorticity is associated with an area of cyclonically curved flow around the MCS toward the wake low (Fujita, 1955; Newton and Newton,
1959; Jessup, 1972). An area of -2 x 10 s negative relative vorticity is found at the trailing edge of the stratiform region near the rear echo notch. An area of 8 x io5 s1 positive relative vorticity is located just past the southern end of the line. Figure 4.11d depicts the wind flow at 1515 GMT. The accelerated flow around the MCS into a wake low is the most dominant feature. Southeasterly flow is found through most of the domain. Comparison of the wind field with the pressure field provides a strong indication that the surface wind direction and speed have been determined by the mesoscale pressure field associated with the MCS.
4.2.3 Surface mesoscale pressure systems In this section a brief description of the surface mesoscale pressure features and their movement in relation to the observed reflectivity structure of the squall line is made. Johnson and Hamilton (1987) have done this type of analysis for the 10-11 June 1985 MCS (see their Figs. 11-15)
They observed a pronounced wake low during the growing phase 75
to the mature phase of the convective system. They attributed the development of the wake low to subsidence warming, and moreover described the wake low as a surface manifestation of the descending rear inflow jet. The rear
inflow jet is described in detail in the next chapter.
Byers and Braham (1948) and Fujita (1955) have shown thatthe
formation of the mesohigh is due to the change in mass
introduced by a precipitation cooled downdraft. Johnson and Hamilton (1987) observed that the wake low present in the 10-11 June MCS moved with the back edge of the
stratiform precipitation area. They point out that this is where no evaporative cooling is available to reduce the
adiabatic warming in the region of focused mesoscale
subsidence associated with the descending rear inflow jet.
In their case the wake low was observed to split into two
identifiable low pressure regions. This splitting of the
main wake low area was associated with a splitting stratiform region, probably caused by the strong intrusion of a dry area
in the form of the mesoscale rear inflow jet.
The position of the wake low with respect to the
stratiform rain region for the 28 May MCS from 1356 GMT to
1528 GMT is shown in Figs. 4.12a-d. Apparently, the entire
wake low was not observed with the PRE-STORM mesonet so that
its more northerly portions are not seen in the analysis. In
the analysis at 1356 GMT and 1422 GMT the wake low (depicted by the 948 mb isobar) is positioned at the back edge of the 76
£0 may :000 V.a,I.J, IOJO L 200
cp-,
N I 1° it!
I
0
DoI.n...... I WotO. to)
N
079'I 9192 b
>' (,
I I 200 -200 0 b) O..l... 1l I W0I0. O0I Fig. 4.12. Composite of surface pressure and reflectivity fields for 28 May 1985. Reflectivity contouring as in Fig.
4.1. a) 1356 GMTb) 1422 GMT. 77
200
Cr-,
I ,J
0 /
'I. 200 I 200 C) 200 0 0.111111.11104 Wtfl.. .,,I
2U May l9U W5H-57 152U 200
Cp,
I eW_ 10 /V 21 ' / /'__
-200 -200 0 200 d/ 0111.nO. 11ol W..hll /lo
Fig. 4.12. continued. Composite of surface pressure and reflectivity fields for 28 May 1985. Reflectivity contouring as in Fig. 4.1. c) 1457 GMT d) 1528 GMT. ilI
northward extending stratiform rain echo. By 1457 GMT no wake low center is present and a synoptic scale trough
(depicted by the 949 mb isobar) is found extending east-west across the northern edge of the stratiformregion. At 1528
GMT the trough continues and no wake low is present. This may signal that upper-level support in the formof focused subsidence warming is no longer available. The physical arrangement and speed of 28 May squall line is quite different from the 10-11 June case, which was a slower, more linear squall line. In the 10-11 June case the wake low remained intact along a track following the stratiform region for a period of 8 hours (see Johnson and
Hamilton, 1987; Fig. 14). The track of the wake low and mesohigh for the 28 May MCS is shown in Fig. 4.13. In this case the wake low did not split, but formed only to the north in conjunction with the most well developed portions of the stratiform region. The track shows that the wake low persisted for about half the time as in the 10-11 June case, moving southeastward with the storm then towards the northeast as the mesohigh began to move rapidly southeastward after 1230 GMT. The wake low moved more slowly after 1230 GMT and dissipated around 1430 GMT, as previously shown (cf.
Fig. 4.12a-d) . Reasons for this abrupt change in movement of the wake low are not fully understood, but may have been the result of the influence of a mid-level mesolow as it descended toward the surface. 79
wake low
- rnesohigh
Fig. 4.13. Tracks of mesohigh and wake low for 28 May 1985 squall line from 1000 GMT to 1430 GMT. Number above the L
(wake low) or H (mesohigh) indicates time, number below indicates departure of pressure, in mb, at 518 m from 950 mb. 4.3 Analysis of PAM II Station Time Series The passage of the squall line at individual PAM II
stations is illustrated in Figs. 4.15-4.17. These time
series were constructed using the five minute observations made at PAM II station 20,6, and 24 respectively. The spatial relationship of these stations to the squall line in
the different stages of its development can be seen in the
WSR-57 radar depictions, Figs. 4.1-4.4. The pressure traces in these time series represent
three of the five stages of the rapidly moving thunderstorm
high as reported by Fujita (1955) and shown in Figure 4.14.
The pressure surge (Fujita, 1955) or pressure jump (Tepper, 1950) illustrated in Fig. 4.14 marks the beginning of the
pressure excess created by the passage of the thunderstorm
high through the undisturbed atmospheric pressure field. The
positive and negative pressure excess formed by the
thunderstorm high (mesohigh) and wake depression (wake low)
can be readily identified (cf. Fig. 4.14). In general, periods of convective line rainfall and
stratiform rainfall are distinguishable in time series
precipitation traces.As a rule, convective line rainfall is identified by the first rapid increase in the trace, while
the stratiform rainfall is identified by a subsequent gradual
climb in the trace. A feature to note in the temperature
trace in each of the time series is the overall increase in
temperature from sunrise at 1113 GMT. Th4UACERSTORMMON FIv(STAGES OFTIlE RAPIOLY MOVING THUNOERSTORM NIGN
PRESSSURGE1 I- WARE O(PRESSION
4
UNOISTIJRSEO PRESS
4
PRESS EXCESS - TIME -
Fig. 4.14. Schematic diagram of pressureprofiles for a squall line thunderstorm. (Adaptedfrom Fujita, 1955)
q.-.IOII e...... o..IIl.., InI I.,.bI MC. WinS G..a*In.. I I..neerfl 1re I.eti . I
- 80
34 24 808- 80
- 10 2s %S !' i ..i'0 ba 18.08
-
*0 0 - 20
- - IC
10 0 820 C 1? 8 *8 00 01 00 08 *0 *1 12 *3 IX 10 0
TI.,.. MT IN.I Fig. 4.15. Time series depiction of the passageof the 28 May
1985 MCS at PAN II station 20. 82
The time series at PAM II station 20 (cf. Fig. 4.15) illustrates the passage of the new convective elements forming at the southern end of the squall line. Heavy convective rain began at 1035 GMT, five minutes after a peak gust of 21 m s- was recorded. Light stratiform rain began at 1055 GMT and ended twenty minutes later contributing 17% of the 7.62 mm of the accumulated rainfall. Rapid cooling commenced with the gust front causing temperatures to decrease 4.2 °C in sixty-five minutes. The pressure profile typifies stage 2, a squall line with no wake depression (cf. Fig. 4.14), indicating that the squall line is in an immature stage with no well-defined stratiform region present behind this portion of the convective line.
At PAM II station 6(cf. Fig. 4.16) the time series showed the passage of weaker convective cells in the leading edge of the convective line and a well-developed trailing stratiform region. Embedded structure within the stratiform region in the form of a secondary band was also evident. A weak gust of 9 m s1 was recorded at 1100 GMT then moderate convective rain began twenty minutes later. Convective rain amounted to 52% of the accumulated rainfall, but occurred over a period thirty-five minutes longer than at PAM II station 20. Stratiform rain began at 1215 GMT and lasted for two hours. A period of enhanced rainfall created by the secondary band began at 1345 GMT accounted for 32% of the Rainfall 000...n.._.tat ton LaoS
-- Pn..o...t I ..b I Mao W no G.... * I r*.. J l.I.t....l
401 301 lb PRESTeRM 25 May 1996 PAM Slat lan .6 0
3. 24 0001- lb
- 70 1
2S hO.t000 S0
24- taI45 I::
'I-', I Il,- J30 ( A 035 -20
'j to
soI L 035-0 OS CP' 05 OS tO 11 12 13 14 15 *5 *7 II 5 Tia. GMT I Fig. 4.16. Time series depiction of the passageof the 28 May
1985 MCS at PAM II station 6. lainfel *aoo...a..tat on 10.11 -- P... I aS I -----MO0 WInd G....t Ia.I --...p..0t... *.I.I....I
40 10 575 tOO III III. PRESTIRM 29 May teas PAM Stat ton .24 0
34 24 555 50
-70 I-
2S- *S - -50 -' ! - ; .i _E0
122 . .
- -3O
035 20
- -to
10 0 .535.0 06 07 OS OS 10 It tO 13 *4 *1 5 17 15 tO
Tt..a OMT th.1 Fig. 4.17. Time series depiction of the passage of the 28 May
1985 MCS at PAM II station 24. 84 total stratiform rainfall and 15% of the total accumulated
rainfall of 21.08 mm.
A gradual cooling occurred after the gust front passed, while a more sudden warming commenced with the minimum negative pressure excess at 1355 GMT.A substantial pressure drop of 8 mb in twenty minutes and vigorouspeak
gust of 25 m s- marked the beginning of the wake depression.
A 4.9 mb resurgence of pressure in eighty minutes with a
refreshening of winds to 16 m s- marked its passage. The highly non-geostrophic flow around a wake vortex, as depicted
in Figure 4.11d, is one possible explanation for this
feature, especially given the sharpness of the pressure drop.
This pressure profile is similar to Fujita's stage 3, a
squall line that has advanced to the mature stage. The wake
low in this time series is quite similar to that reported in the 10-11 June squall line by Johnson and Hamilton (1987)
In their time series analysis (see their Fig. 11), a distinct wake low was observed with a pressure drop =6 mb in about twc hours, which is not as rapid as at PAM II station 6. This difference in the rate of pressure decrease may be a result of the different speed of movement of the two systems. They also observed a similar refreshening of the winds with a pressure rise of =4 mb.
PAM II station 24 time series (cf. Fig. 4.17) depicts the passage of an intermediate portion of the squall line, where mature cells form the leading edge of the convective line followed by a transition zone and stratiform region with an embedded secondary band. At 1325 GMT, a period of heavy convective rain began, lasting for thirty minutes. This precipitation accounted for 81% of the total recorded precipitation at this station. A peak gust to 17 m s- accompanied the heaviest downpour. The transition zone followed lasting twenty minutes and adding 1.63 mm of light
rain. A light stratiform rain started at 1415 GMT with 3.3 mm of precipitation, 13% of the accumulated total of25.4 mm.
The minor temperature drop of 1.5°C accompanying the convective rain was likely lessened by the onset of the diurnal heating cycle. The pressure trace shows a pressure surge followed by a pressure drop of 4.5 mb. A refreshening of the winds to 15 m s- was observed, occurring with the minimum negative pressure excess. A pressure resurgence comparable to that
recorded at PAM II station 6 was not observed. Apparently, a wake low was not encountered at this station. The pressure profile indicates that the squall line has progressed past the mature stage toward dissipation according to the stage 4 pressure profile (cf. Fig. 4.14) classification of Fujita
(1955) 5. SINGLE-DOPPLER ANALYSIS
5.1 Introduction In this chapter we present extensivesingle-Doppler radar data to examine the internal kinematic structureof this squall line storm. A variety of single-Doppler analyses will be shown, including vertical cross-sections (RHI), vertically-pointing time series, and vertical profiles of stratiform vertical air motions by the VAD(Velocity-Azimuth
Display) technique.
5.2 Discussion of the Vertical Cross-Section Analysis To elucidate the kinematic and reflectivityvertical structure of the 28 May MCS, single-Doppler data are presented in a vertical cross-section format. Each cross-section was constructed using data along an azimuth perpendicular to the convective line. System spanning cross-sectional views of the storm structure from the overhanging anvil of the convective line to the back edge of the trailing stratiform region are shown. It should be noted that in discussing the vertical plane analysis, reference to the horizontal plane analysis of the WSR-57 data and PAM II data (cf. Figs. 4.1-4.4, 4.7-4.10) is helpful in providing a three-dimensional conceptual view of the 28 May MCS.
The vertical cross-sections were constructed using the method described by Rutledge et al.(1987) . The velocity data were first interactively unfolded using the NCAR/RDSS (Research Data Support System) at OStJ. The reflectivity and unfolded velocity data were theninterpolated to a Cartesian coordinate system with a grid interval of1.5 km in the horizontal and 500 m in the vertical. Data points were
shifted in the horizontal to correct forthe effects of storm movement that occur during thetime required to complete a
volume scan. Only data with elevation angles <23° were used, hence the radial velocity closelyapproximates the
horizontal velocity. The storm horizontal relative flow was obtained by subtracting the storm speed of17 m s. The
data were then interpolated to thederived cross-sectional
plane. The vertical cross-sections constructedfrom single-Doppler data collected at the NCAR CP-3 radarfor the
period from 1003 GMT to 1314 GMT are shown inFigs. 5.1-5.6.
All cross-sections are along the 290°-110°radials. The
cross-sections of reflectivity and horizontal relativeflow
at 1003 GMT are illustrated in Figs.5.la-b. Fig. 5.la shows
a strong convective cell nearx=45 km located within the
leading convective line. Maximum echo tops extend to over 13
km MSL. The core of the convective cell is identified by the area encompassed by the 40 dBZ reflectivity contour. At
this time a weak trailing stratiform region was located to
the rear of the convective line with most of the surface
precipitation located in the region for x < 105 km.
Figure 5.lb shows a three layered structure in the
storm relative airflow. At low levels an extensive layer of 28-MA Y-85 CP3 RH! 1003 Z 290 DEG
\ ?4. DZADV
9 0
0' x
¶1 4- . 35 120 05 90 75 60 45 30 15 a)
14. UREL -12
\i__ ->0 >9. )
.9.1
Ii
35 120 05 90 75 60 45 30 IS 3
D,Ine lron Rdr (I(m) b) Fig. 5.1. 28 May 1985 1003 GMT Vertical Cross-Section
Analysis at CP-3. a) reflectivity structure; contour interval 5 dBZ, shaded gray for reflectivity > 30 dBZ. b) storm relative airflow; contour interval 5 m s,shaded gray for rear-to-front-flow, not shadedfor front-to-rear flow. rear-to-front flow over 7 km deep entered theback edge of the stratiform region near 120 km. The depth of the rear-to-front flow was over 7 km along theleft portion of the cross-section and then decreasedgradually as it approached the convective region. This rear inflow is similar to that found in other storms with asimilar overall
structure (Smull and Houze, 1987b). The maximum velocities
observed in this strong rear inflow jet exceeded10 m s1
through the trailing stratiform region, butdecreased to < 5
m s- near the back edge ofthe convective region. This flow increased to > 5 m s1 below the convective cell,apparently
in response to the merger with low-level outflowsfrom
convective scale downdrafts. A similar interaction between the rear inflow jet and the convective cells hasbeen
documented by Rutledge et al.(1987) for the 10-11 June
PRE-STORM case. A front-to-rear flow was situated above the rear
inflow jet. As the front-to-rear flow passed the convective cell the flow apparently accelerated to > 25 m s-.
Acceleration of the horizontal wind in the vicinity of
convective lines has been observed in tropical lines by
LeMone (1983) and in a mid-latitude case by Smull and Houze
(1987a) . This acceleration must be a result of a horizontal
pressure gradient force creating additional momentum,since
at no horizontal levels ahead of the linedid we observe
horizontal winds with speeds as large as 25 ms1. Rearward of the convective cells the front-to-rear flow was characterized by speeds 10 m s" This front-to-rear flow was evidently important to the rearwardadvection of ice particles into the stratiform region from the convective cells (Rutledge and Houze, 1987). The third layer of storm relative flow was found at
upper levels above z=8 km betweenx=20 km and x=40 km. This
upper-level rear-to-front flow regime had maximumvelocities
to 15 m s- that were forced by intense divergence present
near the top of the convective cells. This rear-to-front
flow diverted ice particles forward, hence, building the
overhanging anvil. At 1033 GMT (cf. Fig. 5.2a) the squall line had moved
over the CP-3 radar site. The convective region had grown in
width from 20 km to>30 km. The core of the convective cell
continues to be indicated by an area of reflectivities>45
dBZ. The reflectivities in the stratiform region appear to be lower at this time compared to 1003 GMT, which we believe
is a result of attenuation of the radar beam by heavy
precipitation. The storm-relative horizontal flow at 1033 GMT is
shown in Fig. 5.2b. The rear-to-front flow was again
present, entering the system along the back edge of the
stratiform region up to 7 km in height. The location of
maximum rear inflow speed was found in the region immediately
behind the convective line at this time. This is in contrast 91
28-MA Y-85 CP3 RH! 1033 Z 290 DEG
.12 .4 DZADV - -..o_ /
-
9 - 7.1
jJ(2 ::: Ii /t0\1 T T 'AU 105 90 '5 00 45 1) IS
a)
-12 2.4 UREL -10 09 r-\
I 0 :9
-6.4 64 ___,____
/*5sl65,1 :
SLsA - 1.1 1.9
I I I I I IS fl 1*5 t,n I,,. on 95 Sn AS rn
Dinlonne Iron, Ilodo, (19eJ
b)
Fig. 5.2. As in Fig. 5.1 except for 1033 GMT. 92
to the flow at 1003 GMT (cf. Fig. 5.lb)where the maximum rear inflow speeds were foundwithin the stratiforrn region.
The presence of the strongest rear inflow withinthe convective region may have resulted from the merger of convective scale downdrafts. The front-to-rear flow situated above the rear inflow jet showed an overall speedincrease with a 15 m velocity contour extending into the trailing
stratiform region. The 1120 GMT vertical cross-section (cf. Figs. 5.3a,b) was formed by merging vertical planesalong the 2900 and 1100 azimuths, hence allowing a 180° perspective of the
storm. The cross-section of reflectivity (Fig. 5.3a) intersects a vertically-erect convective cell with tops in
excess of 13 km MSL. A small leading upper-level anvil
proceeds the storm, associated with strong upper-level
outflow from the convective region (cf. Fig. 5.3b). A well
developed trailing stratiform region was present at this
time, with the most intense region of stratiform rain located
between x=-30 and x=30. Considerable development was evident
in the intensity of the stratiform region over the last hour
(compare Figs. 5.2a and 5.3a). Within the stratiform region, between x=30 km and x=10 km, an area of enhanced reflectivity
described as a secondary band (Smull and Houze, 1985;
Rutledge and Houze, 1987) illustrates this development. The
secondary band is characterized by the development of an even
more intense radar bright band indicated in Fig. 5.3a near 93
z=3.4 km. The radar bright band is formed whensolid hydrometeors become liquid upon fallingthrough the 0°C level, thereby increasing theirreflectivity, since liquid hydrometeors reflect more power from theintersecting radar beam than frozen hydrometeors (Battan,1959). Rutledge
(1986), Rutledge and Houze (1987), and Smulland Houze (1985) have all noted that the enhanced secondaryband precipitation of the trailing stratiform region can beattributed to the fallout trajectory of hydrometeors as they areadvected rearward by the storm relative airflow and thegrowth of those hydrometeors within a mesoscale updraftlocated at mid-levels within the stratiform anvil. The mesoscale upward motion could contribute to the enhanced reflectivities observed in the secondary band by providing a mechanismfor the growth of ice particles advected into thestratiform anvil. These ice particles advected from the convective cells grow by the deposition of water vapor madeavailable within the mesoscale updraft and by aggregation. Without such a mesoscale updraft, the amount ofrain reaching the surface from the stratiform region would likely be significantly reduced (Rutledge and Houze, 1987) . The establishment of a mesoscale updraft in this storm will be made in the following section. Extensive echo was found aloft and rearward of the stratiform surface rainfall. Apparently, evaporation into the deep rear inflow was responsible for preventing the arrival of hydrometeors at the surface in this region. 94
ItO CEG 290 DEG 28-MA Y-85 CP3 RH! 1120 Z
DZADV 12.4
10.9 10.9 I
94 9.4 { .7.9 7.9
. 64 6.4
40 4.9
14 - .14
19 1.9
.4 .I0 -129
a)
244 UREL /
94.
79-
64-
I9 34. .34
tO- '9
4. I I I I I I I I I I I I hO *03 90 75 60 43 10 IS 0 -10 .90 -43 -Rfl .76 -00 - IllS -1,fl -123
0i0.n. I,o, R.d (Km)
Fig. 5.3. As in Fig. 5.1 except for 1120 GMT. 95
The kinematic structure (cf. Fig. 5.3b) shows a
four-tiered relative airflow pattern. A deep and persistent
rear inflow jet was again present, entering the storm ashigh
as 9 km above MSL. This rear-to-front flow proceeded on a gradual slope, then plunged sharply (near x=0 km) toward the
convective line and became a comparably shallow flow past the
transition zone into the convective cell domain. At this
point (near x=0 km) the rear-to-front flowwas apparently
forced to the surface by relatively dense evaporatively
cooled downdrafts beneath the convective cells and
consolidated into a cold density current (Charba, 1974)
Charba (1974) showed that this kind of cold density flow from
beneath intense squall line thunderstorms preceded the
convective line with the infusion of higher momentum flow
from mid-levels. A backflow in the density current (Goff,
1976; Charba, 1974) was observed near the surface (for x > 10
km) as a front-to-rear flow, deepening toward the back of the system with maximum velocities exceeding 10 m s. Rutledge
et al.(1987) reported a similar sharp plunging in the rear
inflow jet in conjunction with the location of intense convective cells and also concluded that the sharp tilting nature of the jet may be a result of its encounter with intense convective-scale downdrafts associated with heavy precipitation.
A layer of front-to-rear flow capped the rear inflow
jet. This flow exceeded 25 m s1 just above the head of the density current flow along the leading edge of the convective
line. The remainder of the front-to-rear flow was characterized by speeds on the order of 10 m s-. Intense divergence was found at upper levels, resulting in strong
rear-to-front flow ahead of the convective line.
The cross-sections at 1148 GMT (cf. Figs. 5.4a,b) are
again mergers of vertical planes along the 2900 and 110°
azimuths. The cross-section of reflectivity (cf. Fig. 5.4a)
shows the presence of a vertically-erect convective cell with
a vertical extent in excess of 13 km MSL. This cell appears
to be both deeper and more intense than the cells seen at
earlier times. Satellite imagery at 1101 GMT (cf. Fig. 3.17)
showed that the anvil cloud shield had reached its maximum
size about this time, while the PAM II analysis (Cf. Figs.
4.8, 4.9) indicated a strong negative temperature and moisture advection had occurred at the surface during the
storm passage. Both of these indicators are consistent with the observed cell intensity at this time. A well-defined transition zone was now present (Smull and Houze, 1985), as
indicated by the low-level reflectivity minimum between x=-40 km and x=-60 km. The stratiform precipitation region appears to be rather similar to that seen at 1120 GMT although the width of the most intense stratiform rainfall has decreased, probably associated with large evaporation rates in the intense rear-to-front flow. Weak convection appears near x=100 km. This convection was probably driven by convergence 97
IIOOEG 290 OEG 28-MA Y-85 CP3 RH! 1148 Z
DBZ - 12.4 109 109
9-4 94 - .9 - 79
'.4 64
'-9 49 34 t 34 19 19
120 14Q 0 1l 80 8Q 9U 180 140 120 100 80 60 40 20
a)
LJREL
109
94 '-'9, 9-4 -
-. 79
64
19
I II I IHI.' 100 100 14 D,19n9e I,00 flada, (I,)
Fig. 5.4. As in Fig. 5.1 except for 1148 GMT. into the pressure minimum (wake low) located near the rear of the storm (Cf. Figs. 4.10, 4.14) The cross-section of horizontal relative flow (cf.
Fig. 5.4b) is similar to that described at 1120 GMT (Fig.
5.3b). A weaker rear inflow jet (the area with speeds
exceeding 10 m s is substantially reduced or totally
absent) entered the rear of the storm at mid-to-upper
tropospheric levels. Near x=-20 km the rear inflow plunged
sharply and formed a shallow layer of rear-to-front flow in
the transition and convective region. We suspect that the
sudden dip in the rear-to-front flow resulted from the presence of intense convective-scale downdrafts on the edges
of the vertically-erect cells. A deep front-to-rear flow is
present below the rear inflow layer. This flow is both
deeper and more intense compared to earlier times. This flow
was most likely a result of the development of the wake low pressure area, as illustrated in the PAN II analysis (cf.
Figs. 4.10d, 4.14). The front-to-rear flow above the rear
inflow is similar to that discussed at earlier times.
The 1225 GMT vertical cross-section is depicted in
Figs. 5.5 a,b. The elongated vertical core of the convective line shows that the maximum tops continue to extend to near
14 km MSL. Interestingly, the transition zone has widened considerably since 1148 GMT. This occurs in conjunction with a narrowing zone of stratiform rainfall. Erosion of the
stratiform cloud shield and the exaggerated anvil-like 290 OEG 28-MA Y-85 CP3 RH! 1225 Z II000G
2.4 DZADV j't'IA .12, 0.9 ---- to 9.4 - i '0 7.9 ( I 7.9 64
49I
3.4
1.9
35 120 105 90 75 60 45 30 15 0 -IS -30 45 -60 75 -00 105 -120 13
a)
.12. 4 oPEL
10. PS- (
0 9.4 4- \
- S
:1 9.
I I I IIIII (1111- .410 IC fl _IC - AC _o.1 .70 -InC -120
Dislanc, ho.. Rada. (K..)
Fig. 5.5. As in Fig. 5.1 except for 1225 GMT. 100
feature in the trailing stratiform region are evidence of the effects of continued evaporation by the rear inflow jet.
The horizontal velocity structure is shown in Fig.
5.5b. The flow structure at this time is rather similar to that described at 1148 GMT, but there are noteworthy differences. First, the weak rear inflow in the transition and convective zone regions has broadened over time, consistent with the increase in width of the transition
zone reflectivity minimum. Second, the zone of front-to-rear
flow below the rear inflow layer has intensified, probably in respol-ise to the continued development of the wake low.
The shape and location of the rear-to-front flow (or rear inflow jet) is considerably different than that found in the 10-11 June storm (Rutledge et al., 1987) where the jet was largely situated below the 0°C level as it passed throug: the stratiform region. In our case the rear inflow passed through the stratiform region above the 0°C level, hence causing sublimation of ice particles well above the 0°C level. Such a process is consistent with the presence of a narrower secondary band of stratiform rainfall compared to the 10-11 June storm (which obtained widths of nearly 100 km).
The last in the chronological series of vertical cross-sections from the NCAR CP-3 radar is shown in Figs.
5.6a,b. These cross-sections are taken from the vertical plane along the 110° azimuth only. Fig. 5.6a shows that thE 101
28-MA Y-85 CP3 RHI 1314 Z 110 DEG
124 DBZ -12
too -10
94
79
, 64 B 49
34
1,9
-IS -30 -45 -60 '75 -90 -lOS -120 -13
a)
UREL -12
I0I- -10 - -lv. / 1°
- - -60 -90 .90 -lOS -120 - 13
Rd (K..)
b)
Fig. 5.6. As in Fig. 5.1 except for 1314 GMT. 102
convective line had moved out of range of the radar, making the exact width of the transition zone difficult to determine, however, observations from the WSR-57 radar (cf.
Fig. 4.3) show this feature to be similar in width (near 25 km) to the 1225 GMT cross-section. Also, the narrow
stratiform region has changed little in size since 1225
GMT. Apparently, the storm has reached the quasi-steady period of its lifecyle as was indicated in the satellite
imagery. Erosion in the lower levels of the stratiform cloud shield in conjunction with the intrusion of the rear inflow
jet is again evident.
The rear inflow jet shown in the cross-section of horizontal relative flow (Fig. 5.6b) has undergone some interesting changes from the previous analysis. The rear-to-front flow originated at somewhat lower levels, while the once sharply tilting axis of this flow had reverted to a gradual slope. The change in slope of the mesoscale rear inflow could possibly be associated with intense mesoscale subsidence or a transition in the state of the system as a whole to a weaker slanted system as opposed to one dominated by intense convective activity (Rotunno et al., 1987).
Speeds in the front-to-rear flow overlying the rear inflow jet have changed little from 1225 GMT.
The vertical cross-section data from the NCAR CP-4 radar at 1148 GMT (a 2900_1100 vertical plane, Fig. 5.7 a,b) and 1314 GMT (a 110° vertical plane, Fig. 5.8a,b) reveal a 103
290 DEG 28-MA Y-85 CP4 RH! 1148 Z 110 DEG
1 OZADV I2, 0 II I0I -10. .
1:
120 tOO 60 60 40 20 0 - 20 -40 -60 60 -100 -120 -14
a)
24- UREL 12
0.9- -, -to
SO °\
I I I I I I I 40 tOO 00 80 00 40 20 0 20 40 60 00 100 'iOO 14
Distance Iwo, Rad (En,)
Fig. 5.7. As in Fig. 5.1 except for 1148 GMT at CP-4. 104 different aspect of the 3-D structure of the storm. At 1148
GMT the convective line was directly over the radar site, hence the reflectivityand velocity structure (cf. Fig.
5.7b) near the radar is difficult to interpret resulting from the restriction of elevation angles. This observation shows that, in this particular cross-section, the storm had virtually no trailing stratiform region, much different than that observed at CP-3 (cf. Fig. 5.4)
The horizontal relative flow (Fig. 5.7b) shows the
strong rear inflow jet originating near 10 km MSL, then
sloping toward the convective line and at that point plunging
sharply into the rear of the convective cell. The higher originating height of the transition from front-to-rear to rear-to-front flow, and somewhat stronger speeds compared to the CP-3 cross-section (cf. Fig. 5.4), may be responsible for the narrower stratiform region than that observed at CP-3
from the same analysis time. A density current flow is not observable in this cross-section, again in contrast to the
CP-3 cross-section. This suggests the absence of convective-scale downdrafts in the transition zone, or the presence of weaker downdrafts compared to the CP-3 cross-sections. The other recognizable features of the flow include: the low-level front-to-rear flow, the front-to-rear flow overlying the rear-to-front flow structure, and an upper-level rear-to-front flow. 105
28-MA Y-85 CP4 RHI 1314 Z 110 DEG
2 DZyo : to- E
- 7
, 6 2 0 4 / //
- n,1 73 .. 105 120 135
a)
124 (JFiEL -12.
109 6 94
79
64 2 49 1
9
- i ._.. -.-
I I I I 0 -15 -30 -43 -60 -73 - -Inc -IOn -ISO
Distance front Radar (1(m)
Fig. 5.8. As in Fig. 5.1 except for 1314 GMT at CP-4. 106
The reflectivity structure at 1314 GMT (Fig. 5.8a) maximum tops near 13 km. A new convective element can be seen (near x=-125 km) at the 5 kmheight. No transition zone
is present behind the convective line, in agreementwith the
1148 GMT CP-4 cross-section. This is consistent with our earlier argument that convective-scale downdrafts wereeither
absent or weak, hence permitting the fallout ofhydrometeors
in a zone relatively free of deep, sub-saturateddowndrafts. The system-relative velocity pattern (Fig. 5.8b) over
CP-4 is similar in many respects to the pattern over CP-3
(cf. Fig. 5.6b) at this time. However, the more complete four tiered structure of the flow over CP-4 shows the
rear-to-front flow originating near the 7 km height instead
of 8 km. Location and magnitudes of local speed maxima are similar, while differences exist in the continuation of the
rear-to-front flow into the density current outflow below the
convective cells. Local speed maxima in the front-to-rear flow exceed 25 m s1 at the upper-level boundary of the
trailing stratiform echo, as well as in convective region
where the flow appears to be strongly accelerated by the
convective-scale updrafts.
5.3 Discussion of the Velocity-Azimuth Display (VAD)
Analysis Data collected in 3-D conical scans by the NCAR CP-3
radar were used to provide vertical profiles of horizontal 107 wind speed and direction, divergence, vertical air velocity,
and average reflectivity (dBZ). The calculated kinematic properties were determined by the VAD technique described by
Browning and Wexier (1968) . The so called VAD technique (as
described by Browning and Wexler, 1968) uses data acquired
from a rotated radar beam directed through a series of
elevations to determine radial velocity of precipitation particles versus azimuth. From the boundary of a cylinder of
fixed radius (with the radar located directly below the
origin of a cylinder of 20 km radius in this case) the
velocity component normal to the cylinder can be determined
at a series of heights. From the velocity components on the perimeter of the cylinder an estimate of the horizontal wind
vector and mean divergence at the origin of the cylinder can
be obtained. A top-down integration (vertical velocity (w), was set to zero at the top of the profile) of the anelastic
continuity equation using the observed vertical profile of
divergence provides an estimate of the vertical air velocity.
This technique requires that the fields of wind and
precipitation fall speed be horizontally homogeneous, a
condition most often found in the stratiform anvil region.
The results of the VAD analysis technique performed on
data collected at the CP-3 radar at 1120-1130 GMT, 1148-1158
GMT, and 1225-1235 GMT are presented in Figs. 5.9-5.11. This
single-Doppler data acquired at the CP-3 radar site are in
the broadest region of stratiform rain. The vertical profile 1130 GMT VAD Analysis at CP3 on 28 May 1987 Divergence (xlO-3 us) dashed Wind Speed (mis) dashed 40 1 5 10 15 20 25 30 35 -1 -.6 -.2 .2 .6 0 C, I I 10 10 I. S
7.5 7.5 0Cu Cu 5 0
-C .2 2.5 ICu
0 120 102030360 -1 -.6 -.2 .2 .6 1 0 60 Vertical Velocity (mis) solid Wind Direction (deg) solid
I I I I I I I -10 0 10 20 30 40 50 Reflectivity (avg dBZ) solid-dashed
Fig. 5.9. 28 May 1985 1130 GMT VAD analysis from theC2-3
radar. Horizontal wind speed (m s1), dashed line; horizontal wind direction (deg), solid line;reflectivity profile (avg dBZ), solid-dashed line; divergenceprofile (x i- s1), dashed line; and verticalvelocity (m s1),
solid line.
1158 GMT VAD Analysis at CP3 on 28 May 1987 Divergence (xlO-3 ifs) dashed Wind Speed (mis) dashed
I -.6 -.2 .2 .6 1 0 5 10 152025303540 10 10
E "\"s- E 7.5 7.5
Cu
5 5 0
-c ) 2.5 2.5 . I
I -1 -.6 -.2 .2 .6 0 60 120 180 240 300 360 Vertical Velocity (nv's) solid Wind Direction (deg) solid
I I I I I I I -10 0 10 20 30 40 50 Reflectivity (avg dBZ) solid-dashed
Fig. 5.10. As in Fig. 5.9 except for 1158 GMT. 109
of horizontal wind speed (ground-relative) anddirection, divergence, reflectivity, and vertical velocity obtained during 1120-1130 GMT is shown in Fig. 5.9. This profile was computed from data obtained in the forward half of the stratiform rain area, where the rear inflow jet can be seen to tilt sharply toward the surface into alow-level density current (cf. Fig. 5.3). The most noticeable feature in the profile of horizontal wind was the 27 m s speed maxima from
300° at the 1 km level. This speed maxima indicated the possible position of the peak flow axis in the rear inflow
jet and low-level density current. This lower portion of the rear-to-front flow layer extended from the surface to the 2 km level where the wind backed to 245° throughout the remainder of the profile. The divergence profile was typified by low-level moderate divergence to .33 x iO above which strong convergence with a peak value of -. 6 x io s- was found coincident with the upper limit of the reflectivity profile maximum (or bright band, between 2.5 km and 3.4 km).
Rutledge et al.(1987), Smull and Houze (1985), and Ogura and
Liou (1980) have all reported a mesoscale updraft present above the melting level in the layer coincident with the strong convergence. In the layer above 6 km the divergence profile became moderately divergent again.
A maximum negative vertical velocity of -. 8 ms_i (80
cm s-) shown in the lower portion of the profile at the2 km 110
level provided evidence for a mesoscale downdraft. This mesoscale downdraft may have been driven by the melting and evaporation of hydrometeors at and below the 0°C level, as well as hydrostatic mass divergence below this layer. Values on the order of -.6 ms_i (60 cm s-) were reported at the
same level by Rutledge et al.(1987) for a similarly located profile in the 10-11 June MCS. Single-Doppler data from the 1148-1158 GMT volume were processed using the VAD technique and results of the analysis
are shown in Fig. 5.10. The horizontal wind profile (Fig.
5.10) illustrates both the rear portion of the low-level
density current and the mid-level rear inflow jet shown in
the vertical cross-section at 1148 GMT, both speed profiles
are for ground relative flow (cf. Fig.5.4). The increase in
wind speed to 16 m s at 250 meters corresponds to the back
edge of the low-level density current. A sharp decrease in wind speed can be seen above the speed maxima followed by an
increase to 26 m s- at the 4 km level. This mid-level peak
can be attributed to the rear inflow jet traversing the
stratiform region above the bright band. The bright band is clearly evident between z=2.5 km and z=3.5 km in the
reflectivity profile. The reflectivity profile was characteristic of
stratiform precipitation. The weak vertical gradient above the 0°C level is described by Rutledge and Houze (1987) as
indicative of the slow depositional growth of hydrometeors in 111
the stratiform anvil.
The complexity of the divergence profile shown in Fig.
5.10 was apparently a result of radar scans through a mixture
of flow regimes. Strong divergence with peak values to .88
xio3 51 were computed below 1.2 km. The strong divergence
at low-levels can be attributed to divergent flow, as some of
the flow encompassed by the VAD analysis cylinder was carried
forward into the density current, while the remainder was
carried rearward by the low-level front-to-rear flow (cf.
Fig. 5.4b). Next, the weakly convergent layer between 1.0 km and 1.7 km region was apparently a result of the sharp plunging portion of the rear inflow jet where mass is
diverted forward into the density current, as well as toward
the rear in the low-level front-to-rear flow may be creating
an area of mesoscale low pressure above the dif fluent flow.
This low pressure area may be strong enough to slightly
overbalance hydrostatic mass divergence in the layer and form
weak convergence (cf. Fig. 5.4b) . Thirdly, the moderate
divergent layer between 1.7 km and 3.6 km appears to be
similarly associated with the plunging portion of the rear
inflow jet. An influx of mass had apparently been transported into the layer creating a mesohigh and moderate divergence. Evidence for a possible mesohigh at this level
is most clearly shown in Fig. 5.4b, where the rear-to-front
flow enters the narrow gap of the steepest portion of its tilted axis. Above this gap, momentum from the rear inflow 112
jet is forced aloft creating an irregular extension intothe opposing flow. It may be that within this portion of the
flow a building mesohigh forced moderately divergent horizontal winds. Lastly, the layer above 4 km was moderate to weakly convergent in association with theopposing flows
along the tilted axis of the rear inflow jet.
The profile of vertical motion shown in Fig. 5.10
indicates that this location within the stratiform
precipitation region just rearward of the plunging rear
inflow jet was characterized by mesoscale subsidence. The
peak values of subsidence are found above the reflectivity
maximum in contrast to the vertical motion profile shown in
Figs. 5.9 and 5.11. An additional vertical component of motion may have been provided by the descending rear inflow
jet, consistent with its location above the reflectivity
maximum between the 4 to 7 km level.
Vertical profiles derived from the VAD analysis at
1235 GMT are shown in Fig. 5.11. The notch in the back edge
of the stratiform echo was over the radar at this time (cf.
Fig. 5.5a) and the region of enhanced stratiform precipitation was weakening.
At the 1.0 km height the profiles of horizontal wind
speed and direction illustrate the influence of the low-level
front-to-rear flow (note that below 800 meters the wind speed
and direction profile is northwesterly at 16 m s) . Peak
winds (ground relative) are 19 m s from the north-northwest 113
1235 GMT VAD Analysis at CP3 on 28 May 1987 Divergence (xIO-3 us) dashed Wind Speed(rn/a)dashed 15 20 25 30 35 40 .2 .2 .6 1 0 5 10 10 104
E E 7.5 7.5 m V
5 5 0
2.5 .2' .2' 2.5 I Iw
A
120 180 240 300 360 -.6 -.2 .2 .6 1 0 60 Vertical Velocity(rn/a)solid Wind Direction (deg) solid
I I I I -10 0 10 20 30 40 50 Reflectivity (avg dBZ) solid-dashed
Fig. 5.11. As in Fig. 5.9 except for 1235 GMT.
in this layer. Above this winds decrease and back to westerly then increase to 25 ms- at 5 km in the rear-to-front flow. The rear inflow regime becomes front-to-rear above 7 km where the winds decrease and continue to back to the southwest.
A moderately divergent layer was seen below 2 kin in the divergence profile. This is consistent with the VAD analysis cylinder being farther from the strong diffluence found between x=45 km and x=60 km and sampling through a weaker portion of the mesoscale downdraft. Evidently, the mesoscale downdraft associated with this storm was strongest in the forward half of the stratiform region. A layer of convergence was found above this from 2-7 kin with two peaks 114
of moderate intensity. The lower peak was coincident with the reflectivity maximum and was consistent withthe 1130 GMT analysis (cf. Fig. 5.9) . At 5.25 km, the layer of moderate convergence was just above the localspeed maxima characterizing the rear inflow jet. This convergent layer persists from the 1158 GMT analysis (cf. Fig. 5.10) where this second level of the peak convergence was observed750 meters lower. This change in elevation was related to the increase in the height of the interface between the opposing storm relative flows. The layer above 7 km was weakly convergent and became weakly divergent in the storm outflow above 9 km. The vertical motion profile is characterized by mesoscale subsidence throughout, except above 8 km. In this trailing portion of the stratiform region mesoscale downward motion predominates as in the 1158 GMT VAD analysis, except at high levels where weak upward mesoscale motion was found.
This pattern is similar to the 1158 GMT analysis, except that the peak subsidence was below the melting layer as in the
1130 GMT VAD analysis. At 1235 GMT, the strongest subsidence with a'peak value of -.65 m s was seen below the reflectivity maximum near 2 km, where the greatest particle terminal fallspeeds are also found (cf. Fig. 5.13a)
Terminal falispeeds are discussed in the next section.From
6-7.5 km the weak vertical velocity is indicative of a uniform horizontal wind field with very little divergence. 115
Above 7.5 km the positive vertical motion resulted from the weak convergence found at that level.
5.3 Discussion of Vertically-Pointing Analysis
At 1158 GMT the Itstratiform mode" scanning sequence was initiated when the CP-4 radar was placed into the vertically-pointing mode (see section 2.2 for discussion of
"stratiform mode"). The CP-4 radar collected data in this fashion for a period of 52 mm. The time-height
cross-sections of reflectivity (dBZ), Doppler velocity (W),
derived particle falispeed (V)1 and vertical air velocity
(w) derived from this data are shown in Figs.
5.12a,b-5.13a,b, respectively.
The time-height section of reflectivity (Fig. 5.12a) was constructed from averages of individual radar scans over
a two minute period. This 2 minute data was resolved to a grid of 340 m in the horizontal and 150 m in the vertical.
On the left side of the resulting cross-section the back edge of the convective line can be seen, followed by a rather narrow stratiform rain region concurring with the 1148
GMT analysis at CP-4 (cf. Fig. 5.7). On the right side of the cross-section, between the surface and 3.7 km, there is an echo-free notch that Smull and Houze (1985) have attributed to the evaporation of hydrometeors in the rear inflow jet. Above the echo-free area the reflectivity is as large as 20 dBZ, suggesting that hydrometeors grew by 116 28-MA Y-85
14
12.795
10995
9.195 0it to 7.395
5.595 t5'- I 3.795
1.995
.195 31 41 45 1 11 21 Time (minutes) a)
I I
1 11.2 21.4 31.6 41.8 Distance (Km) 28-MA Y-85
14 ;
12.795
10.995
9.195
z 5.595 I 3.795
1.995
.195 21 31 41 45 1 11 Time (minutes)
b) I I 31.6 41.8 1 11.2 21.4 Distance (Km)
Fig. 5.12. 28 May 1985 Vertically-Pointing Time-height
Cross-Section at CP-4 beginning 1158 GMT a) reflectivity (dBZ); contour interval 5 dBZ, shaded gray for reflectivity > 30 dBZb) Doppler velocity (W)(m s1); contour interval 1 m s-, shaded gray forW>0 m s- above 0°C level; 2 m s- contour interval, shaded gray forW<-6 m s1 below 0°C. 117
28-MA Y-85
14
12.795
10.995
9.195
7.395
Z 5.595
3.795
1.995
.195 11 21 31 41 45 Time (minules)
I I I 11.2 21.4 31.6 418 Distance (Km)
28-MA Y-85
'4 I
12795 __
10.995 )I
Time (minutes) b)
11.2 21.4 31.6 41.8 Distance (Km)
Fig. 5.13. As in Fig. 5.12 except for; a) particle falispeed
(Vp) (msH) b) vertical velocity (w)(m s)
L 118
aggregation as they approached the 0°C level, but then evaporated rather rapidly upon fallout.
The time-height cross-section of Doppler velocity (W)
is shown in Fig. 5.12b. This velocity field represents the
sum of the vertical air velocity and particle falispeed.
Regions of positive velocity (dark shading above 4 km) are where upward air motions were sufficient to overcome particle
fallspeeds. Rutledge et al.(1987) described these upper-level positive velocity maxima observed in the 10-11
June squall line as remnants of convective updrafts carried
rearward by the front-to-rear flow. In their analysis (see their Fig. lOb) these positively buoyant regions had largely dissipated before reaching the stratiform region. In this
case two of these buoyant parcels are observed well into the
stratiform anvil and in one noticeable example (above 11 km
in altitude and 41 km in range) contain vertical motions in excess of 1 in s. These two buoyant parcels may be acting
as generating cells (Matejka et al., 1980) seeding
hydrometeors into the stratiform region below. Their presence this far into the trailing stratiform may be
attributed to the greater intensity in the convection-scale updrafts at the time in which these parcels originated between 30-40 minutes prior to 1158 GMT. This would put their origination time close to a period of maximum intensity
as viewed in the satellite imagery (cf. Fig. 3.17) and PAM II data (cf. Fig. 4.8) . A possible origination point for these 119 buoyant parcels can be seen near the 5.5 km level on the left
side of the cross-section. A vertically elongated speed maxima in the front-to-rear flow, similar to the one shown in
Fig 5.4b in association with the convective line, most likely
exists near the origination level of these parcels. This
speed maxima would add an additional component of vertical motion to the parcels fostering their origination. At this
level parcels containing vertical motion in excess of 2 m s-
are seen extending away from the convective line along the
axis of the front-to-rear flow.
The derived reflectivity weighted fallspeed (Vp)is
shown in Fig. 5.13a (see Rutledge et al., 1987 for a description of its derivation). This calculation was only done for levels above the 0°C level (3.4 km) . Below this
level we set V=W. Two regimes were used for the calculation of V based on assumed particle types. In the region where the reflectivity contours exhibited a fairly large slope, we
assumed the presence of graupel. This graupel regime was
implied for time < 13 minutes. For time >13 minutes we
assumed the presence of aggregates of snow, which have a rather weak fallspeed dependence on reflectivity. Fallspeeds on the order of 2.5 m were derived in the graupel region with fallspeeds -1.5 m s- typical in the snow region.
The vertical air velocity field (w)is shown in Fig.
5.13b. This field is a composition of W for z < 3.4 km and the derived vertical velocity (w)for z > 3.4 km. Vertical 120
velocity (w)is derived from,
w=W+V (1)
the addition of W (Fig. 5.12b) and VP (Fig. 5.13a) . The derived vertical air velocity immediately rearward of the convective line (for time < 13 mins.) is dominated by fairly
intense updrafts associated with the convective line. The magnitudes of these updrafts range from 2-4 m s. These drafts are likely associated with the rearward transport of vertical momentum from the convective region, as alluded to earlier. For times > 13 mm., and for z < 11 km the derived vertical velocity is negative, indicating a broad zone of
subsidence throughout much of the observed stratiform echo.
Hence, the ice particles injected from the convective line
are not in a zone favorable for continued growth. This
absence of a mesoscale updraft could be responsible for the relatively narrow stratiform region through the CP-4
cross-section. Above z=ll km the vertical velocity is upward, with the strongest regions of upward motion concentrated in cellular-like features. These features are
likely weakly buoyant elements of the upper portions of convective cells carried rearward by the front-to-rear flow.
These features are not detected in the reflectivity field since they likely contain small particles, which are not readily detected by the radar. 121
6. DUAL-DOPPLER ANALYSIS OF THE TRAILING STRATIFORN
REGION AND ASSOCIATED MESOSCALE CIRCULATION Synthesis of data retrieved from the observations of mid-latitude and tropical lines of convection by multiple
configurations of Doppler radars (usually two or three) has
been reported in papers by Roux et al.(1984), Heymsfield and
Schotz (1985), Smull and Houze (1987a), and others. Studies
of this kind have led to the discovery of structuraldetails
within these convective phenomena that could not be revealed
by rawinsonde networks offering only rather cruderesolution,
or by a single-Doppler radar. In this chapter we will give a brief discussion of mesoscale circulation features observed within the stratiform region associated with the 28 May MCS
by dual-Doppler radar coverage. Data for the dual-Doppler synthesis was collected by
CP-3 and CP-4 in overlapping 1500 sector scans begun at 1314
GMT. The overlapping sectors covered the rear portion of the
trailing stratiform region. For reference purposes the orientation of the horizontal domain of the dual-Doppler analysis within the WSR-57 analysis at 1356 GMT is shown in
Fig. 4.3. The east-west coordinate axis of the dual-Doppler analysis has been rotated to 115° and a storm speed of 17 m
s was subtracted from the derived winds, revealing the
storm-relative wind field. The dual-Doppler synthesis was
carried out using the NCAR/SPRINT and CEDRIC software
packages, which were run interactively on the the Department 122
of Atmospheric Sciences VAX 11/750 computer. The horizontal wind components were obtained by a iterativetechnique. A top-down integration was performed todetermine the vertical velocity field, using an upper-level boundarycondition of w=0 one half grid interval above the highestmeasured
divergence level.
6.1 Discussion of Reflectivity and Kinematic Fields In the dual-Doppler derived wind andreflectivity
field analysis at the 1.9 km height shown in Fig. 6.la, alow
to mid-level N 1-7 km) mesoscale circulationis evident near
the back edge of the stratiform echo (thecirculation center
is marked by an X) . The center is near a comma shaped feature outlined in the 20 dBZ contour of the reflectivity
field. Smull and Houze (1985) reported a similar comma feature in the radar echo and suggested that it was
indicative of cyclonic rotation about the comma heads (see
their Fig. 6). At the 2.9 km height (Fig. 6.lc), the cyclonic rotation has intensified creating a closed circulation about
the vortex center. Two comma-shaped features are outlined in the 20 and 30 dBZ contours along the back edge of the
stratiform echo. The circulation further intensifies at the 3.9 km height (Fig. 6.2a) then begins to weaken at the 4.9 km
height (Fig. 6.2c) and is weaker still, but nonetheless still
identifiable at the 7.9 km height (Fig. 6.3a) . Comparison of 123
28-MAY-85 13:14 DBZ& WINDZ= 1.9 KM 28-MAY-85 13:14 VERTICAL VELOCITYZI.' 60 60 -y
30 30
0 0
-30 -30 I.
-601 I bU 60 -30 0 30 60 -00 r ::- r a) b)
28-MAY-85 13:14 OBZ& WINDZ=2.9KM 28-MA Y-85 13:14 VERTICAL VELOCITY Z = Z9 KM 60 60 4__.
30 30
0 0
-d)-'-9 ------Il o
-30 -30 / L1
-60-I I- -60 -30 0 30 60 -60 60 -60-II L- :::T--a.----- .
C)
Fig. 6.1. Horizontal dual-Doppler Analysis at 13:14 GMT a,c) 1.9 km and 2.9 km vector wind, respectively; vector length is proportional to wind speed, 1 tick = 12 m s1 b,d) 1.9 km and 2.9 km vertical velocity field (w)(m s-); darker shading indicates w>O, lighter shading indicates w the circulation intensity among theselevels establishes that the rotation is strongest at the 3.9 km height(close to the 0°C level) Fitsch and Chappel.1. (1980) suggest an explanationfor this relative-flow circulation, its approximately6 km depth, and its level of maximum intensity near the 3.9km height. The mid-level circulation shown occurred duringthe mature phase of the squall line development. According to Fritsch and Chappell (1980), the mature and late growthphases of squall line development are the most likely timefor the strongest and most well organized circulation to occur. It is during this period that the strongest accelerations act on the flow at low-to-mid levels, while continuedsubsidence warming aloft has caused height falls, mass convergence, and positive vertical motion in the lower levels of the atmosphere. These conditions are conducive to the formation of a cyclonic circulation such as described herein. To begin the cyci.onic rotation, 2-4 hours is required to allow the Coriolis force to act on the mesoscale flow. In this case the time was sufficient since the MCS lasted for more than 12 hours. A secondary role may have been played in the generation of a cyclonic rotation by a strong horizontal shear created where the undulating opposing front-to-rear and rear-to-front flows meet along their sides and break down into vortices that quickly die out as energy cascades from larger to smaller scales. In either case, the circulation 125 that is observed is most likely to be transient in nature and last for a short time once formed. Interestingly, Ogura and Liou (1980) in their study analyzed the vertical vorticity field (see their Fig. 18) and found that upper-levels, generally above 400 mb, were dominated by anticyclonic vorticity, while in the middle to lower layers cyclonic vorticity prevailed. This finding correlates well with the hypothesis made by Fritsch and Chappell (1980) in that it supports the idea that, at least in the latter phases of the squall line development, upper-level subsidence warming creates an anticyclone, while thickness decreases and height falls due to hydrostatic adjustments in the middle to lower layer can create a cyclonic circulation. The vertical velocity component of the relative flow for each of these levels is shown in Figs. 6.1-6.3 b,d. Darker shading indicates positive vertical motion, lighter shading negative vertical motion. At the lowest level (z=l.9) the horizontal axis of the rear inflow jet is difficult to determine, since this level is mostly below the jet, except at the interface between the opposing flows. This interface is located at the -2 m and 2 m vertical velocity contours in close proximity near the right side of the analysis (Fig. 6.1 b).At the upper left portion of the analysis a -2 m s contour is just becoming visible. This is the bottom of the inflow jet between x=-40 and x=-60 126 28-MA Y-85 13: $4 OBZ & WINO Z = 3.9 KM 28-MA Y-85 13:14 VERTICAL VELOCITY Z = 3.9 KM GO 60 0 I1 30 30 - ] 0 0 I, <) -30 -30 ,O 4- -00-GO -30 0 i 30_60 60 ;. ___60 28-MAY-85 13:14 OBZ& WINOZ=4.9KM 28-MAY-85 13:14 VERTICAL VELOCITYZ= 4.9 KM 60 60 --Q _i1 11 .- -. c 1''. - '\1 30 30 0 0 -30 -30 t - 60 GO -30 0 30 60 -60 60 -2 C) d) Fig. 6.2. As in Fig. 6.1 except for 3.9 km and 4.9 km. 127 km in Fig. 5.8. The tilted vertical axis of the rear inflow jet as it crosses the horizontal planes between the 7.9 km level and 1.9 km level is shown in Fig. 5.8. Moving up to the 2.9 km height (Fig. 6.ld) the forward interface of the jet is again visible in a similar position. The area outlined by the -6 m s- contour shows the negative vertical motion component associated with the jet as it passes through the 2.9 kmhorizontal plane just below the 0°C level near 3.4 km. The mean axis of the rear-to-front flow lies along a line connecting these two areas. Having established this mean axis, compare it with the relative-flow shown in Fig. 6.lc, 6.2a,c. On the 2.9 km level the rear inflow jet is seen entering the storm (near the upper left) just past the center of the vortex and the comma-shaped 30 dBZ reflectivity contour. It continues on a left-to-right course into the region of the second comma outlined by the 20 dBZ contour. From there it continues forward toward the convective line, while some of it is deflected into the cyclonic circulation around the vortex joining the front-to-rear flow seen in the upper right portion of the analysis. The continuation of the 30 dBZ contour from the first comma-shaped feature shows the erosion of the stratiform anvil cloud material by evaporation forming a notch in the stratiform echo. This is an indication in the horizontal plane of the entrance of the rear inflow jet into the stratiform region. 128 At the next two levels (Fig. 6.2 b,d) the same general vertical motion pattern is illustrated. The decrease in negative vertical velocity in the area outlined by -6 m s1 at the 2.9 km height to -2 m s- at the 3.9 km height is most likely a result of the reduced evaporative cooling rate above the 0°C level in the atmosphere surrounding frozen hydrometeors as compared to the higher rate in the air surrounding the liquid hydrometeors below the melting level. This would mean that the rear inflow jet would gain an additional negative vertical velocity component from the locally denser air sinking immediately below the 0°C level. This effect would act to strengthen hydrostatic mass divergence from the layers above, which is likely an important element in the formation of a mesolow, and thus, the formation of a mesoscale circulation. Through the 4.9 km and 7.9 km levels (Fig. 6.2d, 6.3b) each horizontal plane slices through successively reduced portions of the rear-to-front flow at a point farther from the speed maxima in its mean vertical axis (cf. Fig. 5.8b), thereby, reducing its contribution to the negative vertical velocity component. This is evident in the reduced size of the local negative vertical velocity maxima previously mentioned. The growth of the overall area of positive vertical velocities at the 7.9 km height illustrates the increasing contribution of the positive vertical velocity component indicative of the front-to-rear flow (cf. Fig. 129 28-MAY-85 73:74 OBZ& WINDZ= 7.9 KM 28-MA Y-85 13:74 VERTICAL VELOCITY 2 7.9 KM 60 30 30 ¶ ¶ S ¶55 os"..'- 555 --''555 5 0'''''' ccc..s ,S,ssss'.'. .ç *. 0 - 0 , . . J ,,.,5,". 0 - '1 - .t. ssss...C... -30 -30 60. . -30 0 30 60 60 .) OF a) b) Fig. 6.3. As in Fig. 6.1 except for 7.9 km. 6.3a) at this level. From the examination of the vertical velocity field and rear inflow jet together it appears very likely that the additional negative vertical motion component supplied by the rear inflow jet is of prime importance to the creation of focused subsidence warming. Focused subsidence warming is a prerequisite to the formation of a mid-level circulation and the mesoscale surface low pressure feature known as the wake low (Fritsch and Chappell, 1980) 130 6.2 Summary of Evidence Regarding a Mesoscale Circulation Vortex development in the lee of thunderstorms and convective lines has been noted in studies by Jessup (1972), Lemon (1976), Smull and Houze (1985), Johnson (1986), and Johnson and Hamilton (1987). Fujita (1955) has shown that wake vortices are established in the mesoscale low pressure region trailing a line of convection (Cf. Fig. 1.2) . In their modelling study of the blocking effects of a circular cylinder in fluid flow, Newton and Newton (1959) have shown that thunderstorms act to some extent as barriers to the relative flow inducing the vertical transport of low-level horizontal momentum, hence, creating vortices in their wake. The signature of a vortex was resolved at the surface by the cyclonic curvature of the winds near the mesoscale low pressure region or wake low following the convective line (cf. Fig. 4.11). The pressure trace at PAM II station 6 indicated the passage of the intense low pressure close to the vortex center by a rapid drop in pressure of 8 nib in twenty minutes (the pressure minimum occurred at 1355 GMT) and a peak gust of wind to 25 ms. A mesoscale circulation was clearly observed at the back edge of the stratiforni region that was most fully developed in the northern portions of this squall line. The level of maximum circulation intensity was found just above the 0°C level near the 3.9 km level. Johnson (1986) report -ed a similar occurrence of a mid-level mesoscale circulation 131 with a maximum circulation intensity just above the 0°C level in the 23-24 June MCS observed in Kansas during PRE-STORN. Mass divergence initiated in an upper-level mesoscale high formed by adiabatic compression and warming in a region of focused mesoscale subsidence could precede significant mid-level convergence that may result in the spin-up of a mesoscale circulation over the region of the subsidence if the subsidence persists for a period of - 2-4 hours allowing the Coriolis force to act on storm-relative flows - 10-15 m s- and/or if sufficient shear exists in the horizontal storm-relative flow. Further, time-series analysis from PAM II station 6 indicates that the transient mid-level circulation may have extended towards the surface deepening the wake low over a small fraction of the mesoscale region and causing the extremely rapid pressure drop observed. Moreover, the 1515 GMT mesoanalysis of the PAM II data (cf. Fig.4.11) revealed that surface winds had attained a sharply cyclonic curvature creating the strongest surface relative vorticity (20 x 10-5 1) analyzed over the PRE-STORM mesonetwork during the passage of the MCS. 132 7. CONCLUSIONS 7.1 Conclusions from the Observational Evidence On May 28, 1985 a nocturnal mesoscale convective system formed near the Nebraska Panhandle along the outflow boundaries of the decaying remnants of a collection of tornadic supercell thunderstorms. Aided by the synoptic scale kinematic and thermodynamic situation a MCS developed and moved southeastward, so that it had begun to enter the northwest corner of the PRE-STORM observational network by the early pre-dawn hours. This thesis, is a case study of the observational data collected by Doppler radar, conventional radar, automated meteorological stations, geo-stationary earth orbiting weather satellite, and standard synoptic weather stations. The synoptic scale analysis has shown that the prerequisites for initiation of the precursor collection of supercell thunderstorms over the Nebraska-Wyoming border existed in the genesis region (GR). These included advection of warm moist air at low-levels with cold advection aloft. Considerable directional and speed shear was found in the airflow over the GR, with southerly flow at low levels, west-northwesterly flow at upper levels, and a 25 m s- jet south of the GR. The GR was located in the forward left-exit zone of a jet streak, a very favorable position for development of intense convection. Scale interactions have not been a subject of this 133 thesis. However, this MCS did move under an upper-level broad high pressure ridge. Why it should not dissipate quickly in a region of synoptic scale subsidence once away from its GB. may be partially explained by the existence of a lee-side trough at lower-levels. Confluence in the steering currents of the mid-level flow was found along the path of the MCS from central Kansas to Arkansas. The additional positive support of the implied mid-level positive vertical motion along this confluence line may have played a part in maintaining the mesoscale updraft within the stratiform anvil cloud. The confluence pattern appears to be a result of the synoptic scale wave pattern in which southwesterly flow is being induced by the lee side trough and northwesterly flow is being induced by a sharp shortwave trough downstream in the airflow (cf. Figs. 3.5, 3.9) Dual-Doppler and PAM II analysis established the existent of a mid-level mesoscale circulation. The origin of this circulation appeared to be related to the formation of an upper-level anticyclone in the region of subsidence warming. Mid-level convergence above the 0°C level in combination with the Coriolis force and horizontal shear in the flow may explain the existence of a mesoscale circulation with its greatest intensity just above the 0°C level. PAM II time-series and mesoanalysis showed that an intense vortex passed through the PRE-STORM observational network. The 134 surface vortex may have been linked to the descent of the mid-level circulation towards the surface. Horizontal observations by conventional radar and vertical cross-section observations by single-Doppler radar confirmed that new convective cells formed along the most southerly portion of the convective line then propagated up the line becoming progressively weaker. A narrow stratiform precipitation region developed rearward of the more intense southerly component of the convective line, while trailing the northerly portions of the line a more substantial stratiform region with embedded areas of heavier stratiform precipitation occurred. The most intense portion of the trailing stratiform region can be attributed to the fallout of hydrometeors as they are advected rearward by the storm-relative airflow into the stratiform region. A notch in the stratiform echo indicated the entrance of a cool, relatively drier, rear-to-front storm-relative flow into the rear of the trailing stratiform region. The notch is created by intense evaporation of hydrometeors. During a period of maximum MCS intensity (as indicated by satellite imagery) the rear inflow jet was observed to descend rapidly in a plunging manner near the back edge of the convective region. Strong convective-scale subsidence in association with heavy convective precipitation generated during this period of intense convection may be responsible for the plunging nature of the flow. The downward slope of 135 the rear inflow jet may have provided an additional negative vertical velocity component to transfer upper-level momentum toward the surface, as well as to focus subsidence warming along the back edge of the stratiform region. A front-to-rear flow was observed to enter the storm along the leading convective line, accelerate past the convective cells and continue into the trailing stratiform region. The front-to-rear flow enters the MCS from the prestorm environment laden with heat, moisture, and momentum, which is transported into the trailing stratiform region. This supply of heat, moisture, momentum by the front-to-rear flow may contribute to the genesis and/or maintenance of the stratiform region. 7.2 Suggestions for Further Research Further study is needed to shed light on the mechanisms involved in the initiation of the mesoscale circulation observed in this case. A time-series of dual-Doppler analyses using data from the 23-24 June MCS and 28 May MCS observed during PRE-STORN should be completed. If both MCS are studied a complete comparison of their horizontal and vertical reflectivity and kinematic structure could be conducted. Also the synoptic scale environment of each case should be evaluated for differences or similarities. Several mesoscale modelling studies could be applied 136 to study different aspects of these two cases. There are two, which I would suggest are most important. First, a microphysical modelling study using wind profiles from each case. And second, a atmospheric dynamical model to study the response of the kinematic flow to changes in theintensity of the convection (i.e., see if a plunging rear inflow jet can be modeled) and see to what extent the rear inflow jet promotes the occurrence of a mesoscale circulation as it passes through the 0°C level. 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