On the Interaction Extratro~icalformation of an upper-level short wave, 2) column stretching, 3)...

On the Interaction of Extratro~ical Cyclones with ~opggra~hy

by Dov Richard Bensimon

Department of Atmospheric and Oceanic Sciences McGill University, Montreal

February 1997

A thesis submitted to the Faculty of Graduate Studies and Research in partial fulfilment of the requirements of the degree of Master of Science.

O Dov Richard Bensimon 1997

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ABSTRACT

To better understand the role of mountains in lee cyclogenesis, two such cases

which occurred dunng BASE (Beaufort and Arctic Storms Experiment) are simulated

using the Mesoscale Compressible Community mode1 (MC2). Both cases are shown to

satisfj criteria for lee cyclogenesis, despite some ambiguity in its definition. The

successful simulations reveal that lee cyclogenesis involves several processes: 1)

formation of an upper-level short wave, 2) column stretching, 3) enhanced convergence

and increased relative vorticity resulting from adiabatic warming, 4) latent heat release

and, in one case, increased baroclinicity due to low-level blocking by topography.

The results of sensitivity experjments indicate that removal of topography (latent

heat) produces a stronger (weaker) lee cyclone. Topography significantly influences the

distribution of precipitation with climatological consequences for areas in the lee. It is

found that cyclogenesis can still occur in the absence of mountains in the two cases

studied, although mountains modify the cyclogenetic processes.

RESUME

Pour mieux comprendre le rôle des montagnes dans la cyclogenèse orographique,

deux cas tirés de BASE (Beaufort and Arctic Storms Experiment) sont simulés par le

modèle de mésoéchelle compressible communautaire (MC2). Les cas respectent des

critères de cyclogenèse orographique, comportant une certaine ambigüité dans sa

définition. Les simulations réussies révèlent que la cycfogenèse orographique implique

plusieurs processus: 1) formation d'une onde courte en altitude, 2) étirement de colonnes,

3) convergence dans les bas niveaux et une augmentation du tourbillon relatif par

réchauffement adiabatique, 4) chaleur latente, et, dans l'un des cas, intensification de la

zone barocline par blocage de la topographie.

Des tests de sensibilité indiquent qu'ôter la topographie (chaieur latente) produit

un cyclone plus intense (faible). Elle influence de façon importante la distribution des

précipitations avec des conséquences climatologiques pour les régions du côté sous le vent

des montagnes. Ici, la cyclogenèse a lieu même sans les montagnes, mais elles servent

à modifier le processus cyclogénétique.

ii

Table of Contents

AbstractRésumé Table of Contents List of Figures List of Tables Acknowledgernents Chapter 1 : Introduction

1 .1 Previous studies 1.2 Terrninology 1.3 Objectives of the thesis

Chapter 2: Model description and specifications 2.1 Model description 2.2 Modelling design

Chapter 3: Numerical simulations of the SEP24 case 3.1 Verification of control simulation 3.2 Evolution of SEP24 case 3.3 Sensitivity experiments

3.3.1 Effect of latent heating 3.3.2 Effect of topography 3.3.3 Combined effects of latent heating and topography

Chapter 4: Numerical simulations of the SEP8 case 4.1 Verification of control simulation 4.2 Evolution of SEP8 case 4.3 Sensitivity experiments

4.3.1 Effect of latent heating 4.3.2 Effect of topography 4.3.3 Combined effects of latent heating and topography 4.3.4 Effect of domain size and position of

laterd boundaries Chapter 5: Conclusion and summary References

. . II

iii iv

vii viii

1 1 4 5 9 9

13 16 16 17 23 23 23 26 53 53 55 58 58 59 60

iii

List of Figures

Chapter 2

Topography used in control simulations. Units are in m, with contour interval of 200 m. Computational domain used for the SEP24 (SEPI) case outlined by Box A (B). Shading in both boxes indicates location of sponge zone.

Chapter 3

CMC analysis of sea-level pressure at intervals of 2 hPa for (a) 1200 UTC 24, (b) 0000 UTC 25, (c) 1200 UTC 25 and (d) 0000 UTC 26 September 1994; subjectively analyzed frontal positions are also shown. Dashed line in (a) indicates position of continental divide. Sea-level pressure at intemals of 2 hPa (a) 12-h, (b) 24-h and (c) 36-h control simulations, valid at 0000 UTC 25, 1200 UTC 25 and 0000 UTC 26 September 1994, respectively. Line AA' (BB') in (a) ((c)) shows the location of cross sections used in Figs. 3.12, 3.19 and 3.20 (3.5). Distribution of 12-h accumulated precipitation from control simulation contoured at l,5,lO,ZO and 40 mm ending at (a) 00/25- 12, (b) 12/25-24 and (c) 00126-36. Light (heavy) shading indicates accumulations between 1 (5) and 5 (10) mm as well as over 20 (40) mm. (d) Observed precipitation accumulations for 24-h period ending 0000 UTC 26 September 1994. Only arnounts greater than a trace are indicated. Location of surface stations used in verification of precipitation accumulations. Vertical cross section of vertical motion at intervals of 0.1 Pa/s for 0600 UTC 25 September 1994 taken dong line BB' given in Fig. 3.2~. Solid (dashed) contours are for positive (negative) values, indicating descent (ascent). (a) 925-hPa divergence at intervals of 10 X 1û6 s" and (b) 925-hPa relative vorticity at intervals of 2 X 10" s-' from 24-h control simulation valid at 1200 UTC 25 September 1994. Darkened area indicates area with surface pressure less than 925 hPa. Thick, solid line in (a) indicates zone of divergence. Solid (dashed) contours are for positive (negative) valves. Distribution of temperature at intervals of 2°C and horizontal winds at 850 hPa from (a) 12-h, (b) 24-h and (c) 36-h control simulation valid at 0000 UTC 25, 1200 UTC 25 and 0000 UTC 26 September 1994, respectively. Darkened area indicates area with surface

Page

14

pressure less than 850 hPa. Letter 'M' in (c) denotes position of maximum low-level baroclinicity, as discussed in text. Solid (dashed) contours are for positive (negative) values. Distribution of relative hurnidity at intervals of 10% at 700 hPa from (a) 1 2 4 (b) 24-h and (c) 36-h control simulation valid at 0000 UTC 25, 1200 UTC 25 and 0000 UTC 26 September 1994, respectively. Light (heavy) shading indicates relative humidity in excess of 70% (90%). Satellite image at (a) 1754 UTC 24 (b) 0429 UTC 25 (c) 1732 UTC 25 September 1994. Geopotential height (solid) at intervals of 6 dam at 500 hPa and potential vorticity (PV, dashed) at intervals of 0.5 PVU at 300 hPa from (a) O-h, (b) 12-h, (c) 24-h and (d) 36-h control simulation valid at 1200 UTC 24, 0000 UTC 25, 1200 UTC 25, and 0000 UTC 26 September 1994. Light and heavy shading indicate PV greater than 2 and 4 PVU, respectively. Thick "Lu indicates position of surface cyclone centre. Thick, dashed line in (c), (d) indicates position of 500 hPa trough. The 6 hourly positions of the surface cyclone centre (L) and the 300-hPa PV anomaly (X) from control simulation. Dashed lines connect multiple centres of either feature if relevant. Vertical cross section of potential vorticity (solid) at intervals of 0.5 PVU, potential temperature (dashed) at intervals of 3 K, and dong-plane flow vectors, which is taken dong line AA' given in Fig. 3.2a for (a) 12/24-00, (b) W5-12 , (c) 12/25-24 and (d) 00/26-36. Shading indicates potential vorticity greater than 1.5 PVU. Thick solid line in (a) and (b) indicates location of tropopause fold. Scales of vertical motion (Pa s*') and wind speed (m s'l) are indicated in bottom left-hand corner. Position of the surface cyclone is indicated dong the abscissa. Note that lowest vertical level in cross section is 900 hPa. Time series of central pressure for the SEP24 case from Exps. CTL, NLH, NMT and DNM. As in Fig. 3.2, but for Exp. NMT. As in Fig. 3.1 1, but for Exp. NMT. As in Fig. 3.7, but for Exp. NMT. Time series of relative vorticity (X 10'' s-l) calculated over a circular area of 200 km in diameter about the cyclone centre and over the five lowest Gal-Chen levels for Exps. CTL and NMT. As in Fig. 3.3, but for Exp. NMT. Isentropic cross section at intervals of 5 K taken dong line AA' in Fig. 3.2a from Exp. CTL (solid) and from Exp. NMT (dashed) for (a) OOI25- 1 2, (b) 1 2/25-24, (c) 00/26-36. As in Fig. 3.19, but for Exp. CTL (solid) and Exp. NLH (dashed).

CMC analysis of sea-level pressure at intervals of 2 hPa for (a) 0000 UTC 8, @) 0000 UTC 9, (c) 0000 UTC 10 and (d) 0000 UTC 1 1 September 1994; subjectively analyzed frontal positions are also shown. Dashed line in (a) indicates position of continental divide. Sea-level pressure at intervals of 2 hPa (a) O-h, (b) 24-h, (c) 48-h and (d) 72-h control simulations, valid at 0000 UTC OS, 0000 UTC 09,0000 UTC 10 and 0000 UTC 11 September 1994, respectively. Line AA' (BB') in (a) shows the location of cross section used in Fig. 4.4 (4.15, 4.16). Distribution of 24-h accumulated precipitation from control simulation contoured at 1,5,10,20 and 40 mm ending at (a) 00/09- 24, (b) 00/10-48 and (c) 00/11-72. Light (heavy) shading indicates accumulations between 1 (5) and 5 (10) mm as well as over 20 (40) mm. (d) Observed precipitation accumulations for 24-h period ending 0000 UTC 11 September 1994. Only amounts greater than a trace are indicated. Vertical cross section of vertical motion at intervals of O. 1 Pds for 1200 UTC 9 September 1994 taken dong line AA' given in Fig. 4.2a. Solid (dashed) contours are for positive (negative) values, indicating descent (ascent). (a) 850-hPa relative vorticity at intervals of 2 X 10" s-l and (b) 850-hPa divergence at intervals of 10 X 104 S-l and (c) 850-hPa vertical motion at intervals of O. 1 Pais from 48-h control simulation valid at 0000 UTC 10 September 1994. Darkened area indicates area with surface pressure less than 850 hPa. Solid (dashed) contours are for positive (negative) values. Distribution of temperature at intervals of SOC and horizontal winds at 850 hPa from (a) 36-h and (b) 48-h control simulation valid at 1200 UTC 09 and 0000 UTC 10 September 1994, respectively. Darkened area indicates area with surface pressure less than 850 hPa. Shading indicates temperatures above 24°C. Geopotentiai height (solid) at intervals of 6 dam at 500 hPa and potential vorticity (PV, dashed) at intervals of 0.5 PVU at 300 hPa from (a) O-h, (b) 24-h, (c) 48-h and (d) 72-h control simulation vaIid at 0000 UTC 8,0000 UTC 9,0000 UTC 10, and 0000 UTC 11 September 1994. Light and heavy shading indicate PV greater than 2 and 4 PVU, respectively. Thick "Lu indicates position of surface cyclone centre. Thick, dashed line in (c), (d) indicates position of 500 hPa trough. The 12 hourly positions of the surface cyclone centre (L) and the 300-hPa PV anomaly (X) between 00/09-24 and 0011 1-72 from

control simulation. Time senes of centrai pressure for the SEP8 case from Exps. CTL, NLH, NMT and DNM. Sea-level pressure at intervals of 2 hPa (a) 48-h and @) 72-h from Exp. NLH, valid at 0 UTC 10 and ûûûû UTC 11 September 1994, respectively . As in Fig. 4.10, but for Exp. NMT. Time series of relative vorticity (x IO-' s-') calculated over a circular area of 200 km in diameter about the cyclone centre and over the five lowest Gd-Chen Ievels for Exps. CTL and NMT between 12/09-36 and 0011 1-72. 850-hPa contribution of (a) divergence terrn and (b) tilting term at intervals of 1 X 10'' s" to vorticity tendency for OOOO UTC 10 September 1994 from control simulation. Solid (dashed) contours are for positive (negative) values. As in Fig. 4.3, but for Exp. MMT. Isentropic cross section at intervals of 5 K taken dong line BB' in Fig. 4.2a from Exp. CTL (solid) and from Exp. NMT (dashed) for (a) 00109-24, (b) 00/10-48, (c) 00/11-72. As in Fig. 4.15, but for Exp. CTL (solid) and Exp. NLH (dashed).

List of Tables

Chapter I

1.1 Lee cydogenesis cases during BASE

Chapter 2

2.1 Summary of main features in the MC2 mode1

vii

Although only one name appears on the final copy of this thesis, many more

people than just myself have had a hand in it. Firstly, many thanks to both my

supervisors, Prof. Peter Yau and Prof- Da-Lin Zhang without whose help this project

would not have been possible. Acknowledgernent of their hours of explanations, help and

expert advice as well as their continued cornmittement cannot be adequately recognized

in a few words, but 1 thank them for al1 their hard work and efforts. A special thanks

goes to Prof. Chester Newton, whose interest and contributions to my thesis study were

both a privilege and extrernely pertinent.

Many of my peers also spent countless hours discussing, explaining and helping

with the project. In particular, I would like to thank Marco Carrera, Zonghui Huo,

Cinquiang Li and Karl MacGillivray for al1 they helped me with. 1 also benefitted greatly

from the input of Ning Bao, Fanyou Kong, Gary Lackmann, Yubao Liu, and Badrinath

Nagaraj an.

1 would aiso like to thank Michel Desgagne, Pierre Pellerin, Simon Pellerin and

others from the MC2 help desk who spent much of their time helping to solve my

problems. Thanks also go to Gerard Croteau, Serge Filion, Richard Hogue, Serge Nadon

and Vanh Souvanlasy from CMC.

Vicki Loschiavo and Carol Abbott deserve thanks for al1 the help they gave me

in helpiog to sort out the paperwork which accompanies the thesis.

Finally, 1 would like to thank my parents for their continued support and patience

for many long hours spent both at and away from home, without whom 1 could not have

gotten to this point.

Thanks to everyone!

viii

Chapter 1 - Introduction

1.1 Previous studies

The lee sides of major mountain chahs are favourable areas for cyclogenesis

(Petterssen 1956; Chung et al. 1976; Zishka and Smith 1980; Gyakum et al. 1996).

Events have been observed in the Iee of both zonally and meridionally onented mountains

such as the Alps and the Rockies. Despite considerable work in the past decades, the

need for a unified theory of lee cyclogenesis' still exists (Pierrehumbert 1986) and

forecasting of this phenomenon remains a challenge (Tibaldi et al. 1990). Bluestein

(1993) summarized the problem well when he stated: "1s lee cyclogenesis just ordinary

cyclogenesis modified by orography, or is it indeed speciai?".

The influence of the mountains on the circulation of the atmosphere was

recognized as early as the turn of the 20th century (Shaw 1906, Bjerknes et al. 1911).

In the two decades following World War II, several studies have described the synoptic

structure of a Rocky Mountain lee cyclogenesis event (Hess and Wagner 1948; Newton

1956; McClain 1960; Carlson 196 1; Hage 196 1). More recently, atmospheric models

have been used to simuIate this phenomenon. Certain authors considered idealized

scenarios (Lin and Perkey 1989; Bannon 1992; Hayes et al. 1993; Orlanski and Gross

1994), while others exarnined actual cases (Bates 1990; Steenburgh and Mass 1994; Han

et al. 1995). The regions most studied are the European Alps and the American Rockies.

Some attention has also been paid to the Canadian Rockies and the mountains of East

Asia (Chung et al. 1976; Han et al. 1995). However, relatively little research has been

reported for high latitudes.

A number of theories for lee cyclogenesis have been proposed (Smith 1979,

1984a; Pierrehumbert 1986; Hayes et al. 1987). Smith (1984b) put forth 15 different (but

not necessarily separate) rnechanisms to explain this phenomenon (Bates 1990). They

include the "classical" upper-level or jet strearn explanation whereby upper tropospheric

' also known as orographie cyclogenesis; in the present text, the two terms are used interchangeably. For a more detailed discussion on the terrninology, see the section entitled "Objectives of the Thesis"

divergence produces drops in surface pressure and strong upward vertical motion.

According to this theory, low-level stretching of vortices is caused by both upper-level

divergence and downslope motion in the lee (Newton 1956; Speranza 1975; Chung et al.

1976). Other major mechanisms are the rapid changes in incorning wind speeds which

change the orientation of isentropic surfaces (Radinovic 1965), and the effect of the

mountain drag which produces a torque on the atmosphere leading to vorticity production

and the formation of lee cyclones. Smith (1979) concluded that in the lee of the Rockies

and the Alps, cyclogenesis can occur as a result of different mechanisms.

The theory of baroclinic instability has been extended to include the effect of

mountains (Smith 1984a, 1986; Speranza et al. 1985; Pierrehumbert 1986; Hayes et al.

1987). Speranza et al. (1 985) developed a quasi-geostrophic theory which views lee

cyclogenesis as the result of topographic modification of an unstable baroclinic wave.

However, this theory was only able to descnbe the strong deepening following the initial

formation of the lee cyclone.

The theory of conservation of potential vorticity (PV) has also been applied to this

problem (Holton 1979). Fluid colurnns capped by isentropic surfaces will be compressed

as they pass over a mountain, implying that the relative cyclonic vorticity must decrease

to maintain constant PV. Once the mountain has been traversed, the columns stretch and

relative cyclonic vorticity increases, thus making the cyclone "reappear". This mechanism

ignores changes in the earth's vorticity, thus making it less plausible for orographic

cyclogenesis in the vicinity of zonally oriented mountains, such as the Alps.

For cyclogenesis associated with the Rocky mountains, Bannon (1992) outlined

a typical scenario as follows. A surface cyclone is present upstream of the mountain

range, with support from an upper-level trough. This low gradually loses its identity

when passing over the mountains, as a trough develops in the lee. Sustained adiabatic

warming by compression Ieads to the formation of a lee cyclone which initially

propagates to the southeast but later acquires a northward cornponent in its motion. The

initial southeastward path of the low is the result of a combination of pressure Mls (rises)

occurring to the south (north) of the low centre associated with downslope (upslope)

winds and warm air and cyclonic vorticity advection which is strongest to the east of the

low pressure centre. A northeastward shift of the motion occurs once the influence of the

mountains is no longer felt. Pierrehumbert (1984a, 1986) suggested that the reappexance

of the cyclone in the lee of the (Rocky) mountains cm be explained by the temporary

masking of the surface low by the rnountain anticyclone. Recently, Czmetzki and

Johnson (1996) supported this view by proposing that mountain-induced pressure torques,

resulting from asyrnrnetries in the distribution of isobars around the cyclone centre, cm

account for the masking process.

A number of synoptic and mesoscale processes can interact with mountains to

affect cyclogenesis. For instance, low- level baroclinicity can be enhanced by cold air

damming on the windward side and a relatively warm pool of air in the lee. When a

zona1 mountain ridge is located poleward of a large heat source such as a large body of

water (Orlanski and Gross 1994), the release of available potential energy is enhanced by

a thermally direct circulation in the lee as warrn, moist (cold, dry) air is forced to ascend

(descend) the mountain slope. Low-level blocking of the flow by orography can set up

a dipole pressure field, with higher (lower) pressures on the windward (lee) side of the

mountain. However, it should be pointed out that the formation of lee cyclones has also

been observed away h m areas of organized baroclinicity (Karyampudi et al. 1995).

Upper-level jet streaks can cause rapid pressure falls at the surface in the lee,

particularly when the region of interest is below the left exit region of the jet streak

(Achtor and Horn 1986; Mattocks and Bleck 1986). The secondary circulation set up by

the jet streak induces tropopause folding on the cyclonic shear side of the jet, advecting

high values of PV downward into the troposphere (Karyampudi et al. 1995). Tracers such

as ozone and dew-point depressions have confirmed the descent of stratospheric air into

the troposphere during tropopause folding events (Buzzi et al. 1984). In cases of strong

cyclogenesis, upper-level dynarnical processes associated with jet streaks have been found

to be more important than low-level processes (Zupanski and McGinley 1989).

Horizontal advection of PV streamers in the upper troposphere has also been shown to

be an important triggering mechanism for lee cyclogenesis. The strongest cases occur

when the upper-level trough and the associated reservoir of high-PV air move across the .

mountain barrier (Bleck and Mattocks 1984). This upper-level process, together with

low-level orographic blocking of cold air, are believed to be of major importance in lee

cyclogenesis (Mattocks and Bleck 1986).

The advent of numerical modelling has contributed to the improved understanding

of orographic cyclogenesis by allowing sensitivity tests to be performed to isolate the

influence of a particular variable. By removing the topography in a model, it has been

shown that the mendional displacement of the cyclone is dependant on the height and

width of the topography it crosses (Bannon 1992) and the position of the cyclone

indicates a sharp discontinuity as it passes over the mountain. Also the low-level relative

vorticity field is affected as Bates (1990) found higher values of relative vorticity in a run

where the mountains are rernoved. However, by the end of the simulation, when the lee

cyclone had moved well east of the mountains, the absolute vorticity did not differ

significantly in the mountain and no-mountain runs because in the latter case the cyclone

took a more northerly track. The increase in earth's vorticity compensates for the

decrease in relative vorticity noted in the no-mountains simulation.

In passing, we remark that in simulating the interaction of atmospheric systems

with topography, the grid size and terrain resolution are of great importance (McQueen

et al. 1995). When small-scale topography is not properly resolved, the larger-scale flow

dominates the evolution of the atmosphere. Many mesoscale models employ smoothing

to remove 2Ax topographical modes with the results that the quality of simulations of

low-level blocking is degraded. For a better simulation it may be necessary to enhance

the topography to preserve the maximum heights (Pierrehumbert 1984b). Thus, the

proper representation of topography remains an ongoing challenge to numerical modellers.

1.2 Terminology

Some confusion exists on the terrninology in the literature. "Lee cycIogenesis"

and "orographic cyclogenesis" are typically used as synonyms (Tibaldi et al. 1990,

Orlanski and Gross 1994). However, the word "lee" indicates a direction to the incoming

flow, and conveys the notion of a complete understanding of the process (Orlanski and

Gross 1994). For this reason, the term "orographic cyclogenesis" is preferred. Studies

have also indicated that mountains may simply modi@ the development of the cyclone

as opposed to being its cause (Tafferner and Egger 1990), in which case it is even less

clear as to what exactly constitutes an orographic cyclone. Tibaldi et al. (1990) include

in their definition of "orographic cyclogenesis" cases of cyclonic reintensification. A

further confusion arises because "cyclogenesis" has been defined as the tirne and place

of the first appearance of a closed sea-level isobar which must persist for at least 24 hours

(Petterssen 1956). However, different authors have used different increments of sea-level

pressure (SLP) to examine cyclogenesis events, which adds arnbiguity to this definition

(Newton 1996). Pierrehumbert (1986) indicates that cyclogenesis can also be defined as

"a fa11 in pressure of greater than a certain magnitude". Chung et al. (1976) point out that

conflicting definitions of cyclone "development" and "intensification" exist.

"Developmen t u refers to the intensification of a pre-exis ting cyclone, while ffcyclogenesis"

includes both the initiation and intensification of a cyclone. They suggest using

"cyclogenesis" to refer to the process of initial formation of a cyclone and "development"

to its subsequent intensification.

1.3 Objectives of the thesis

In view of the lack of research in orographical cyclogenesis in high latitudes, we

propose to study two such events that occurred during BASE (Beaufort and Arctic Storms

Experiment), conducted between 1 September and 15 October 1994. This experiment was

designed to address some of the issues in the Mackenzie River GEWEX Study (MAGS)

- the Canadian component of the Global Energy and Water Experiment (GEWEX), whose

goal is to better understand the water and energy budgets affecting the global climate

system.

In order to decide on a particular case study, a table listing al1 cases of lee

cyclogenesis during BASE was assembled (Table 1.1). This table results from a

subjective analysis of the tracking of low pressure centres in the lee as indicated on

BASE surface analyses2 (Hudson and Crawford 1995) and is sorted according to

' As low centres are subjectively analyzed in the BASE analyses, and because they were tracked only in the lee, the values of central pressure differ from those discussed for these cases in later chapters of this thesis.

deepening rate3. The entries list both the starting and ending date and time for each case,

the difference in central pressure (deepening) between the two tirnes, the period of

deepening, the associated deepening rate and, if applicable, the Intense Observation Penod

(IOP) during which the case took place. The final column qualitatively indicates the track

of the lee cyclone relative to a central latitude (chosen as 65"N).

The two case studies to be discussed in this thesis are ranked 2nd and 8th in the

table in terms of the deepening rate. The first case occurred between 1200 UTC 24 and

0000 UTC 26 September 1994 (hereafter, the SEP24 case) and involved a relatively

strong parent cyclone giving rise to a weaker lee cyclone. This cyclone passed directly

over the BASE study area, during IOP 7. It represents a case of lee cyclogenesis (Hudson

and Crawford 1995) with one of the strongest lee pressure falls (21 hPa in 36 hours, as

identified from BASE surface analyses) during the BASE period. Even though its

deepening rate is less than the October 4 case, the cyclone remained mostly over land

areas for the period of interest, and is better captured by the observing network. This

storm is also characterized by the fact that the parent cyclone initially located over the

ocean was sufficiently deep that it never lost its identity in the SLP field while traversing

the mountains. Despite filling while approaching the crest, it undenvent a slight

reintensification in the lee. It thus satisfied the criteria of "orographic cyclogenesis" as

defined by Tibaldi et al. (1990).

The second case was chosen as it represents the longest period of lee cyclogenesis

during BASE. Also, its track is relatively far south, thus differing strongly in latitudinal

coverage relative to the SEP24 case. The storm occurred over the 72-h period beginning

0000 UTC 8 September 1994 (hereafter, the SEP8 case). It formed equatonvard of the

first case with an initially weaker parent low giving rise to a reIatively strong lee cyclone

over the Canadian Prairies, with SLP lower than its parent. The parent cyclone

completely lost its identity as it went over the mountains, and subsequently underwent

rapid reintensification in the lee. It thus satisfied more of the general requirements for

The deepening rate is defined as the difference in central pressure at the beginning of the given period and that at the end of the same period divided by the length of the period. It is thus expressed in units of hPa hr".

orographie cyclogenesis. In both cases, closed isobars do appear in the lee and for the

purposes of clarity and consistency, we would associate closed isobars on the windward

(lee) side of the mountains with the parent (lee) cyclone.

This thesis represents a contribution to a better understanding of high-latitude

precipitation systems. It also addresses some of the issues in the Mackenzie River

GEWEX Study (MAGS). Its main objectives are:

a) to study the interaction of two high-latitude extratropical cyclones with

topography using the Mesoscale Compressible Community (MC2) model; and

b) to study the mesoscale structure of these simulated weather systems with

particular emphasis on mesoscale vortices, tropopause folds, and the distribution

of clouds and precipitation.

The thesis is organized as follows. Chapter 2 describes the features of the model

used for this study. Chapter 3 discusses the evolution of the SEP24 case as well as

various sensitivity tests conducted on this case. Chapter 4 examines the SEP8 case and

is structured the same way as the preceeding chapter. The final chapter compares the two

cases, summarizes the work and provides concluding remarks.

Table 1.1 Lee cyclogenesis cases during BASE

-- -

Start Date Central End Date Central Deepe- Time Rate IOP? Route Pressure Pressure ning (hrs) (hPa/hr) (hP@ ( h m (hW

Oct 04,122

Sep 24,122

Sep 07,122

Oct 02,122

Sep 18,002

Sep 21,122

Sep 01,122

Sep 07,122

Oct 13,002

Sep 02,002

Sep 12,122

Sep 14,002

Sep 15,122

Sep 03,122

Oct 05,002

Sep 26,002

Sep 08,122

Oct 05,002

Sep 18,122

Sep 22,122

Sep 03,002

Sep 12,002

Oct 14,002

Sep 03,122

Sep 14,122

Sep 14,122

Sep 16,002

Sep 04,002

South central

North

North

North

Central

North

North

Far south

Central

Central&South

North

South

North central

Northwest - - -

Note: A "*" in the IOP column indicates that the IOP overlaps with the indicated period, but does not entirely cover it. -

8

Chapter 2 - Model description and specifications

2.1 Model description

The numerical model used is the Mesoscale Compressible Comrnunity Model,

known as MC^ or simply, MC2. It onginates from a nested regional hydrostatic model

and has been continually enhanced over the years by rnany different researchers. We

used version 3.2 of MC2 in this thesis.

The model allows for the treatrnent of compressible fluids, and can be run non-

hydrostatically', as was the case in al1 the experirnents described herein. It is based on

the fully elastic Euler equations, and can be applied to a wide range of scales, ranging

from the microscale to the synoptic scale. The goveming equations expressed in a

conformal projection are (Tanguay et al. 1990):

c ' It is also possible to run the MC2 under the condition of hydrostatic balance.

where

U,V,w are the components of the wind dong the X,Y and z cartesian coordinates;

f is the Coriolis parameter (f=2Rsin+, Q=angular speed of rotation of the earth,

@=latitude (degrees));

K is the pseudo kinetic energy: K=?h(u2+v2); S is the metric projection term: s=m2 where m is the map scale facto?;

R is the gas constant for air;

T is the temperature;

q is defined as q=ln(p/p,) with p=pressure and p, being a constant;

Fi represents sources andor sinks of momentum dong either of the three Cartesian

axes;

g is the acceleration due to gravity;

a = RE,, C, being the heat capacity of air at constant pressure;

L represents sources and sinks affecting temperature;

M represen ts moisture;

E represents sources and sinks of moisture;

C represents liquid water content;

B represents sources and sinks of liquid water ~on t en t .~

Note as well that the total derivative, DDt, is defined as follows (Haltiner and

Williams 1980, Chap. 1):

The model employs finite differencing to discretize the basic equations. A

leapfrog in time but serni-implicit differencing scheme (Robert 1969; Kwizak and Robert

The MC2 is run on a uniform resolution grid with a polar stereographic projection.

The notation used follows that given in the Forntulation of the Mesoscale Compressible Community (MC21 Mode1 by Bergeron, Laprise and Caya (1994).

1971) is used. The slower moving waves (Rossby waves) are treated explicitly and the

faster moving gravity waves implicitly. The strategy considerably lengthens the time step

relative to a fully explicit scheme (Haltiner and Williams 1980).

The model is discretized differently in the horizontal and in the vertical (Bergeron

et al. 1994). In the vertical, momentum and thermodynamic levels are defined. Vertical

derivatives are discretized using a second-order centered finite difference method. In the

horizontal, four types of variables are defined: variables related to the geometric terms,

the two velocity components, and the thermodynamic variables. They are set out on a

honzontally staggered grid, which has been shown to yield greater accuracy in the

calculations (Haltiner and Williams 1980).

The vertical coordinate used for integration is the modified Gd-Chen coordinate.

However, data can be input on either pressure or sigma levels, and the model will

interpolate them to the Gd-Chen coordinate. This coordinate, c, is described in Gd-Chen

and Somerville (1 975) with the form:

w here

H is the top of the model atrnosphere (units of length);

b(X,Y) is the height of the topography.

The advantage of using this vertical coordinate is that the surface 6=0 corresponds

to the topography z=h, and the surface c=H is at a constant height z=H. The MC2 uses

a hybrid version of the above coordinate, stretching or cornpressing the coordinate

vertically. This coordinate also has the advantage that the homogeneous kinernatic

conditions at the surface and at the top of the model atmosphere take on a very simple

form: w(c=û)=O and w(<=H)=O.

A serni-Lagrangian advection scheme is also used. It allows for the use of

relatively long time steps while maintaining numerical stability and high accuracy

(Haltiner and Williams 1980). The advantage of combining the semi-Lagrangian method

with the semi-implicit method was demonstrated by Robert et al. (1985). To avoid the

splitting of solutions arising from the computational mode of the leapfrog scheme, a time

filter developed by Robert (1966) and analyzed by Asselin (1972) is applied.

The MC2 is a lirnited-area model. Its boundary conditions must be nested from

a global or hemispheric model or from an analysis which covers a wider area. The

nesting technique is based on a method proposed by Davies (1976) which was validated

by Robert and Yakirniw (1986) and Yakimiw and Robert (1990). In essence, the

technique sets up a fixed number of points at the edge of the integration domain to fom

the "sponge zone". The width of this zone can either be specified or calculated

automatically in proportion to the total number of grid points in the integration. The

values for variables in the sponge zone is graduaily relaxed to the values at the boundary

provided from the larger-domain analysis or from another model. Without this technique,

the transfer of information in and out of the domain of integration cm be problematic,

leading, for exarnple, to spurious reflection of information back into the model domain,

which increases the propagation of numerical noise (Bergeron et al. 1994).

The function of the model is to integrate the goveming equations fonvard in time

to produce a prognostic analysis of the various fields at later times. To launch the model,

a pilot file must be created which contains both the initial conditions as well as the

boundary conditions for subsequent times in the integration. The MC2 requires certain

two-dimensional fields as input (Desgagné et al. 1995) which remain constant in tirne

during the integration: examples include sea surface temperature, deep soi1 temperatures

and snow cover.

The model employs several physics pararnetrizations schemes described in Mailhot

(1994). The standard RPN (Recherche en Prévision Numérique) physics package used

includes a pararnetrization of gravity wave drag described in McFarlane (1987) and

McFarIane et al. (1987). Infrared radiation is handled by the scheme described in Garand

and Mailhot (1990) which is an improved version of that described in Garand (1983).

Solar radiation is represented by the scheme of Fouquart and Bonne1 (1980). The

production of stratifom precipitation is described by a simple condensation scheme.

Isobaric condensation takes place when the relative humidity exceeds a critical value. A

modified Kuo (1974) scheme is used to represent convective processes, as described in

Mailhot et al. (1989).

To achieve an initial geostrophic balance between the mass and wind fields, the

MC2 can be dynarnically initialized. The procedure involves integrating the model

forward in time a certain number of timesteps, followed by a backward integration back

to the initial starting time, b. The dynamic initialization allows time for the emission of

gravity waves to reach a balanced state of geostrophy.

2.2 Modelling design

For each case study, a control experiment and a number of sensitivity tests were

performed. The control simulation for the SEP24 case was perforrned on a 141 X 13 1

polar stereographic grid true at 60°N, with a horizontal resolution of 50 km (denoted as

Box A in Fig. 2.1). AI1 the results'shown for this case are plotted on a smaller domain

than that used for the model integration, as it is necessary to place the boundaries of

limited-area models sufficiently far from the region of interest to prevent contamination

of the simulation by errors propagating inward r o m the boundaries (Chouinard et al.

1994). The computational domain used in the SEP8 case is indicated by Box B in Fig.

2.1. The number of points in the sponge zone for al1 simulations was fixed at six.

In the vertical, 25 computational levels were used, with the model lid set at 25 km.

The Canadian Meteorological Centre (CMC) analyses provided initiai and boundary

conditions for ail the simulations. Topography was obtained from the CMC

climato~ogical data sets, at a horizontal resolution of 50 km4. Interpolation ont0 the

model grid and the subsequent staggering5 of the topography field during the integration

of the model leads to under-representation of the actual heights of the topography,

although the overall shape of the topography is preserved (Fig. 2.1). The dynamic

initialisation feature of the MC2 was used in simulations for both case studies.

A sensitivity test whereby topography from an initially higher resolution data set was interpolated to the 50 km horizontal resolution grid was compared to this method and the mode1 showed itseIf to be insensitive to such a change.

' The MC2 uses a horizontally staggered grid as described in the previous section.

Table 2.1 Sumrnary of main features in the MC2 mode1

Numerics

3D non-hydrostatic Euler equations in (x,y,c)

Semi-implicit time differencing

3D semi-Lagrangian advection scheme

25 computationai levels

Staggered horizontal and vertical grids

Sponge boundary conditions

Ph ysics

Turbulent vertical diffusion pararnetrization

Surface heat and moisture budgets based on force-restore method

Parametrized gravity wave drag

Kuo-type deep convection scheme

Simple shallow convection and grid scale condensation schemes

Infmed and solar radiation schemes account for effects of H20, CO,, O, and

clouds

Fig. 2.1 Topography used in control simulations. Units are in m, with contour interval of 200m. Computational domain used for the SEP24 (SEP8) case outlined by Box A@). Shading in both boxes indicates location of sponge zone.

Chapter 3 - Numerical simulations of the SEP24 case

This chapter starts with a verification of the control simulation (Section 3.1),

followed by its detailed analysis (Section 3.2), and concludes with a discussion of various

sensitivity tests (Section 3.3).

3.1 Verification of control simulation

The SEP24 case began with a deep cyclone with a centrai pressure of 969 hPa

over the Gulf of Alaska (Fig. 3.la). Over the next 12 hours, the cyclone approached the

Coast of Alaska and filled to 980 hPa (Fig. 3.lb). The formation of the Iee cyclone

became evident in the next twelve hours, with only a residual circulation remaining of the

parent cyclone (Fig. 3.1~). The Iee cyclone deepened by 2 hPa in the next twelve hours

(Fig. 3. ld). The simulated SLP fields (Fig. 3.2) indicate that this process is well captured

by the MC2. Twelve hours into the integration, at 0000 UTC 25 September 1996

(hereafter 00/25- 12), the mode1 properly simulates the filling of the parent cyclone, whose

central pressure increases to 976 hPa (Fig. 3.2a). At 12/25-24, as in the analysis, the

results indicate two low pressure centres (Fig. 3.2b); one at 986 hPa in the lee of the

Rockies, the other at 986 hPa being the remnants of the parent cyclone. Finally, by

00/26-36 (Fig. 3.2c), the simulation places the Iee cyclone at the sarne location as in the

analysis, with a central pressure 4 hPa lower than the analyzed value. Also, the MC2

properIy captures the structure of the zonally-oriented SLP trough extending over the

spine of the mountains. Since this SLP trough is a persistent f e a ~ r e during the

simulation (Fig. 3.2), data from the entire BASE period (1 September to 15 October 1994)

were exarnined to determine how frequently the trough occurs over a longer time span.

A subjective analysis indicated that a trough dominated the SLP field over the Rocky

Mountains of Alaska and the Yukon 56% of the time. A ridge or the absence of a trough

was only noted during 16% of the period. In the remaining time (28%), there is either

no clear signature of a ridge or of a trough or that both a ridge and a trough are present

over the region. This result is consistent with the fact that the Gulf of Alaska is an area

climatologically dominated by cyclones (Zishka and Smith 1980).

It should be noted at this point that the SLP fields shown throughout this study

have been diagnosticdly calculated from the geopotential height, temperature, surface

pressure and specific hurnidity of the model output. To filter the diumal variation in SLP

(Mass et al. 1991), we extrapolate the pressure from a Ievel which is 100 hPa less than

the surface pressure to sea level, using the standard atmospheric lapse rate of 6.5O km".

Additionally, the SLP field shown has been passed through a 1-2-1 smoother-desmoother

(Haltiner and Williams 1980) to remove any 2Ax waves which may be present.

Precipitation accumulations (Figs. 3.3a-c) were verified against observations (Fig.

3.3d), and although the bulk of the precipitation in this case fell in data-sparse areas, the

few measurements available compare favourably with the simulation results. Coastal

maxima were generally correct, with 46 mm measured in the 24 hours ending 0000 UTC

26 September 1994 at Yakutat, Alaska (YAK, see Fig. 3.4 for location of stations used

in verification) which is less than the roughly 60 mm predicted by the model during this

time. The pattern, however, is correct and the simulation properly captures the coastal

maximum in precipitation. A subjective verification of other fields such as geopotential

height and cloud cover was performed and confirmed the successful simulation.

3.2 Evolution of the SEP24 case

Previous studies of lee cyclogenesis indicate that the interaction of the low-level

80w with topography can be important in the formation of a lee cyclone. Newton (1956)

concluded while examining a 1ee cyclogenesis event that "...the centre of low-level

development was located in the region where the combined effects of ascent at 500 mb

and descent at the surface gave the greatest verticaI stretching ...Y To examine the

importance of such a mechanism in the SEP24 case, Fig. 3.5 shows a cross section of

vertical velocity taken through the foothills of the Roches. Strong low-level descent was

just beginning in the lee indicated by the maximum of 10.2 x IO-' Pa s", and this

occurred directly below a region of mid-tropospheric ascent of magnitude -5.3 x 10-' Pa

s-l. The superposition of these opposing vertical motions led to stretching of air colurnns

which in turn generated low-level vorticity. Theory shows that if PV is conserved (in the

absence of friction and diabatic effects), a column of air which is stretched in the lee of

a mountain barrier as the separation between isentropes increases would lead to a spin-up

of low-level relative cyclonic vorticity (Holton 1979).

The spin-up of relative vorticity in the lee has also been shown to be associated

with strong low-level convergence (Newton 1956; McClain 1960), although the

propagation of the parent cyclonic circulation into the lee partly increases in the relative

vorticity in this area. The stretching mechanism previously mentioned implies downward

(upward) motion near the surface (aloft); descent at low levels implies divergence through

the continuity equation. At the sarne time, the downward motion leads to adiabatic

warrning through the thermodynarnic equation. This warming by compression leads to

the formation of a lee trough (Figs. 3.2a'b) and also hydrostatically forces an increase in

thickness. The atmosphere reacts to this thickness increase with rising motion, which

counters the initial warming in a manifestation of Lechatelier's principle (Bluestein

1993). This principle States that the atmosphere acts to return to equilibrium when it is

perturbed from this state. This ascent implies favorable convergence at low levels, as

marked by the 'L' in the lee (Fig. 3.6a) dong the position of a relative vorticity

maximum of 14.1 x IO-' S.' (Fig. 3.6b)'.

The preceeding reasoning describes both divergence and convergence occuring in

the lee, albeit not at the same location. In the imrnediate foothills of the mountains,

descent and divergence are present near the surface (indicated by solid line in Fig. 3.6a),

but if one looks further downstream of the mountains, ascent and convergence are present

(indicated by the 'L' in Fig. 3.6b). The area of convergence and cyclonic relative

vorticity is evident at 12/25-24 at a location where this air mass was advected by the low-

level winds (Figs. 3.7b,c) from its position in the previous six hours (Fig. 3.5). The

position of the convergence zone (Fig. 3.6) corresponds to the Iocation of formation of

the lee cyclone (Fig. 3. lc).

Warm advection in the lee at 00/25-12 in advance of the low (as suggested by Fig.

3.7a) contributed to temperature increases, such that a ridge formed in the temperature

Although the formation of a convergence zone in the lee is explained as related to topography, this is not to Say that it was exclusively formed as a result of the terrain. It rnay also have been due, in part, to the propagation of the parent cyclone into the lee.

field (Fig. 3.7a), along which lee troughing in SLP occurred (Fig. 3. lb). This temperature

ridge subsequently amplified as the circulation around the lee cyclone intensified and

enhanced the advection of higher (lower) temperatures poleward (equatorward) (Fig.

3.7b). Amplification of the thermal ridge was greater than that of the corresponding

trough in the wake of the low due to the blocking of southward advection of colder air

by the mountains of central Yukon. Low-level baroclinicity continually increased

between 12/25-24 and 00126-36 in the presence of this cold-air darnming as winds also

began to act frontogenetically (Fig. 3.7~). Although by 00/26-36 the zone of maximum

low-level baroclinicity was to the west of the cyclone centre (indicated by the 'M' in Fig.

3.7c), the lee cyclone had strengthened due to an interaction with this topographically

induced baroclinic zone between 12/25-24 and 00/26-36. The role of topography in

inducing this baroclinic zone is further discussed in section 3.3.

The lower tropospheric vertical motion field was strongly related to topography,

with strong ascent along the coasts of Alaska and British Columbia (not shown) leading

to large precipitation accumulations (Fig. 3.3a). A close inspection of accumulations

indicates that precipitation arnounts generally increased with elevations of topography on

the windward side (not shown), consistent with the observation by Smith (1979).

Immediately below the mountain peaks, however , amount s decreased with height so that

the maxima occur along the slopes of the mountains, and not at the peaks. Conversely,

a lee rain shadow in the Mackenzie River Basin (MRB) is seen in Figs. 3.3b and 3 . 3 ~ .

A second maximum in precipitation is located well downstream of the rnountains (Fig.

3.33, as the low moved eastward. Some topographically induced precipitation also fell

almg the foothills of the Richardson mountains of northern Yukon as the low centre

moved east of this location (Figs. 3.3b7c) and upslope northeasterly winds foliowed. It

is conjectured that when the Beaufort Sea is free of ice, this northeasterly flow constitutes

an important source of moisture for the MRB. Relative minima in precipitation coincide

closely to the location of a tongue of dry air which wrapped into the main circulation of

the systern (Figs. 3.8a-c), suggesting that this mid-tropospheric "dry slot" may partly

expIain local reductions in precipitation. However, sensitivity tests to be described later

in this chapter indicate that reductions of precipitation in the lee of the mountains are due

more to topographic effects than due to the presence of the "dry slot".

Satellite imagery at 1754 UTC 24 September 1994 reveals a discernable spiral

rainband around the cyclone centre (Fig. 3.9a). Convective cloud elements were also

present off the south Coast of Alaska, giving rîsc to the localized heavy precipitation

accumulations previously mentioned. Figure 3.9b depicts the sharply defined cold front,

with clouds trailing over the mountains, onented dong the SLP trough (Figs. 3.lc,d).

When the lee cyclone moved northeastward, a mixture of convective and stratiform

precipitation fell over the mountains of central Yukon and the MRB, as suggested by Fig.

3.9~. Convection was triggered by the destabilization introduced by the advection of

colder air aloft into the area following the passage of the cyclone (not shown).

In the mid-troposphere, a cut-off low was initially embedded within a synoptic-

scale trough (Fig. 3.1 Oa). This closed circulation moved northeastward and opened up

(Rg. 3.10b) and a short-wave trough subsequently formed and amplified in the lee (Figs.

3.10c,d). The role of topography in inducing this shortwave is exarnined in the next

section. The vertical structure of the cyclone was preserved during this penod, with little

westward tilt with height typical of developing cyclones. In fact, the surface cyclone was

at al1 times located directly below the base of the upper-level (500 hPa) trough. The

synoptic-scale flow regime in which this lee cyclone evolved is typical of that associated

with precipitating systems over the MRB (Lackmann and Gyakum 1996).

Although there was little westward tilt with height between the surface cyclone

and the 500-hPa trough, a more discernible tilt was observed with respect to the 300-hPa

PV anomaly, a reference located higher in the troposphere. Figure 3.11 plots the temporal

evolution of the positions of both the upper-ievel (300 hPa) PV anomaly and the surface

cyclone. Over the first 18 hours, the upper-level anomaly was located directly above the

surface cyclone, while the parent system filled. At 12/25-24, the lee cyclone appeared

further downstream from the upper-level PV anomaly (ULPVA), giving rise to a

westward tilt, at which time a deepening of the surface cyclone began. This apparent

"jump" in the position of the surface cyclone was coincident with the end of the decay

of the parent cyclone and the Iee cyclone formation. A sirnilar jump has been noted in

other studies of lee cyclogenesis, e.g. Bannon (1992). Over the last 12 hours, a westward

tilt was still present, but decreased as the disturbance returned to an equivalent-barotropic

structure. Accordingly, the lee cyclone began fîlling after 00/26-36.

Isentropic cross sections taken along the path of the cyclone show a trough at low

levels directly above the location of the surface low at the initial time (Fig. 3.12a). This

trough was located beneath an isentropic ndge in the upper troposphere, so that static

stability was rninimized in a vertical colurnn directly above the cyclone centre. This

vertical structure was largely a result of the latent heat released throughout the lower

troposphere associated with the precipitation which falls along the coast of Alaska, This

trough was advected northeastward, and as it crossed the mountain, a flattening of the

jsentropes occurred (cf., e.g., the 291 K contour in Figs. 3.12a and 3.12b). Once the

cyclone passed over the mountains, the low-level trough became more evident (lower

right-hand corner of Fig. 3.12d). This process is rerniniscent of the masking effect of the

mountains suggested by Pierrehumbert (1 984a, 1986). Although the low-level signature

of the cyclone may be temporarily Iost, the cyclone may be tracked by following the

upper-level signal. Figure 3.12 also indicates that advection of lower (higher) potential

temperatures occurred throughout most of the troposphere on the windward (lee) side of

the mountains (cf., Figs. 3.12a and 3.12d). As shown by Lackmann and Gyakum (1996),

this warm advection provides quasi-geostrophic forcing for ascent, contributing to

precipitation formation in the lee of the mountains.

The passage of the cyclone into the lee of the mountains can be traced by

following the ULPVA associated with it. A PV maximum at 300 hPa was initially

located directly above the surface low (Fig. 3.12a) and was distinct from the stratospheric

reservoir which was within 20" of latitude of this anomaly (Fig. 3.10a). Stratospheric air,

identified by PV values greater than 1.5 potential vorticity units (PVU), (Bluestein 1993)

was found locally at higher pressures than 400 hPa. The downward excursion of the

tropopause fold led to increases in PV throughout the rnid troposphere (Fig. 3.12b). The

folding of the tropopause was more pronounced on the equatonvard side of the jet strearn

(not shown), due to the secondary circulation created by a slight deceleration of the

upper-level flow impinging on the mountains. This ageostrophic circulation induced

downward (upward) motion on the equatorward (poleward) side of the jet Stream,

enhancing (inhibiting) the downward advection of PV. The ageostrophic circulation

described above is the same which is observed in the right (left) exit region of a jet sveak

where descent (ascent) is induced.

It was originally hypothesized that, as the SEP24 case takes place at very high

latitudes, the relatively low tropopause may result in an enhanced exchange between the

troposphere and stratosphere. Initially, the synoptic-scale trough and the polar vortex

were each associated with their own PV anomaly (Fig. 3.10a), but by the end of the 36

hour period, the PV anomaly associated with the lee trough approached the other two

anomalies to form a broad area of high PV air containing a PV minimum which pinched

off from the synoptic-scale ndge to the easi (Fig. 3.10d). The strong forcing provided by

this large area of upper-level high PV rnay have contributed to keep the surface low

deeper than it would have been othenvise. Although in this case, no explicit merger

occurred between the ULPVA associated with the lee cyclone and that associated with

the polar vortex, such a merger may be observed in other cases of high-latitude lee

cyclogenesis.

As previously mentioned, a stationary region of ascent forced by topography was

present throughout the troposphere on the windward side of the mountains (Figs. 3.12a-d).

When this forcing combined with quasigeostrophic forcing for ascent, one would expect

pressures to fa11 at the surface. However, due to low-Ievel blocking of the flow by the

topography, the surface low filled while approaching the coast. Descent was evident in

the immediate Lee of the Rockies (Fig. 3.3, but further downstream, ascent occurred

(Figs. 3.12b,c).

Theoretical studies, such as that of Smolarkiewicz and Rotunno (1989), showed

that strong blocking of a flow by an isolated obstacle c m lead to the formation of lee

vortices. The degree of blocking is assessed using the Froude number, F?, a non-

dimensional quantity proportional to the ratio of the kinetic to potential energy of a given

flow. Although N and h are fixed in a theoretical setting, this is not the case in the

Fr is often defined as U/Nh, where U is the wind speed, N is the buoyancy frequency, and h is the height of an obstacle.

atmosphere, so it is not possible to assign a criticai value of Fr which corresponds to

blocking. Calculations of Fr at low levels in the SEf24 case indicated that values of Fr

decrease as one approaches the mountains, suggesting that blocking may have been

important. However, if an air parcel ascends from a region of weaker static stability into

a region of stronger stratification, it may not be able to surmount the obstacle, despite

having been associated initially with a high value of Fr. Additionaiiy, changes in wind

direction may deflect a parce1 initially targeted to go over a mountain in a different

direction. Therefore, the effect of blocking will have to be elucidated in the next chapter

by examining the resulu of sensitivity experiments where topography is removed.

3.3 Sensitin'ty experiments

The success of the control nin gives us confidence to conduct sensitivity tests to

isolate the effects of different parameters on the lee cyclogenesis.

3.3.1 E f k t of latent heating

Experiments where both stratiform and convective condensation were removed

were conducted to elucidate the importance of latent heat release. The results of this so-

called "dry nin" (hereafter, Exp. NLH) for the SEP24 case are presented in Fig. 3.13, and

indicate higher centrai pressures of the cyclone than in the control simulation (Exp. Cn).

This difference fluctuated with a mean value of 4 hPa throughout most of the 36 hour

pend. The track of the cyclone was unchanged in the dry mn (not shown), and only

rninor differences existed in the structure of the SLP field.

3.3.2 Effect of tupography

As the main purpose of this thesis is to understand the interaction of extratropical

cyclones with topography, the controi simulation is compared next to a simulation where

topography was set identically to zero everywhere in the domain (Exp. NMT). The

roughness length was assiped one value over the ocean (0.1 cm) and another over land

(15 cm). A Iapse rate based on the average rate in the lowest layes above the mountains

was used to replace the volume which they oceupied.

Recognizing that the mass underneath the topography has been replaced by a

fictitious mass field, the results of this experiment for the SEP24 case are presented in

Fig. 3.13 and reveal the presence of a deeper cyclone than in Exp. CTL at al1 times with

differences in central pressure up to 6.5 hPa at 12/25-24. This confirms the importance

of low-level blocking hinted at earlier, as the removd of topography allowed low-level

mass to circulate more freely around the centre of the low. With mountains present, mass

accumulated at low levels, leading to higher SLP. This notion is c o n h e d by the

presence of a slight ridging in the SLP field dong the Coast in Exp. CTL (Fig. 3.2a) but

not in Exp. NMT (Fig. 3.14). Additionaily, the SLP in Exp. NMT exhibited a more

syrnrnetric structure than that of Exp. CTL (cf., Figs. 3.14a and 3.2a). In particuiar, the

SLP trough which lay across the mountains in Exp. CTL both prior to the passage of the

parent circulation (Fig. 3.2a) and following the formation of the lee cyclone (Rg. 3 .2~)

was not as evident in Exp. NMT. The differences in SLP between Exps. CTL and NMT

(not shown) indicated that at al1 locations where the winds had a component blowing

upslope (downslope) in the former, the pressure was lower (higher) in the latter. Despite

these differences, the central pressure of the cyclone in Exp. NMT still followed the sarne

pattern as Exp. CTL, filling for most of the study period, with a slight re-intensification

in the lee (Fig. 3.13). This lee re-intensification was weaker than in Exp. CTL, indicating

that the mountains were largely, but not completely responsible for the lee development

of the cyclone.

It was mentioned in the previous section that an upper-level short wave which

developed in the lee of the mountains provided support for the re-intensification of the

cyclone. The presence of this sarne feature in Exp. NMT (not shown) indicates that this

trough was not generated by the high topography. The fact that an upper-Ievel trough

developed in both simulations is significant, as this suggests that this feature may be

relatively comrnon during cases of lee cyclogenesis. If lee cyclones formed uniquely as

the result of surface spin-up induced by an upper-level short wave, then one would

conclude that lee cyclogenesis is just ordinary cyclogenesis which happens to occur in the

lee of the mountains. However, the vast number of cyclogenesis cases which occur in

the lee of major mountain ranges indicates that the location of these events is not random

and that the previously described physical processes involving topography generally act

in concert with ordinary cyclogenesis mechanisrns to help spin up surface cyclones in the

lee of major mountain chains.

It was previously shown that lee cyclogenesis could be explained by the

development of a westward tilt of the surface cyclone with respect to the ULPVA.

Applying the sarne reasoning to Exp. NMT, Fig. 3.15 shows that the westward tilt which

developed between the ULPVA and the surface cyclone at 12/25-24 in Exp. CTL was no

longer present in Exp. NMT, with the surface cyclone located directly beneath the

ULPVA at this tiine. It is only between 18/25-30 and 00/26-36 that the greatest

separation between the ULPVA and the surface cyclone in Exp. NMT took place,

precisely the six-hour period during which the lee cyclone deepened by 1 hPa (Fig. 3.13 j.

It was already shown that topography enhanced baroclinicity in the area of lee

cyclogenesis owing to low-be l blocking of the cold air moving southward from the

Beaufort Sea. Figure 3.16 is the analog of Fig. 3.7, but for Exp. NMT, and indicates that

even without mountains, cold air advection occurred over northern Yukon and the extreme

northwest Northwest Territories, but the pronounced packing of the isotherms so evident

over central Yukon in Fig. 3 . 7 ~ was absent in Exp. NMT (Fig. 3.16~). An inspection of

the isotherms indicates that temperatures were slightly higher over the area between Great

Bear Lake and Great Slave Lake in Exp. CTL, likely a result of enhanced warming due

to downslope motion to the east of the Rockies.

Lee cyclogenesis is often identified by an increase in low-level relative vorticity

(Bates 1990). In order to detect such an increase, the relative vorticity averaged over a

circle of 200 km in diarneter around the surface cyclone centre and over the 5 lowest

terrain-following surfaces was calculated for both Exps. CTL and NMT (Fig. 3.17). In

both runs, the vorticity actually increased slightly over the first 12 hours because the

location of the relative vorticity maximum, which was initially displaced laterally from

the cyclone centre, became more aligned with the centre of the cyclone. Over the next

12 hours, the vorticity decreased as the parent cyclone decays over the mountains. The

increase in vorticity over the last 12 hours during Exp. CTL signaled the occurrence of

the lee cyclogenesis. This increase in vorticity was conspicuously absent in Exp. NMT.

In light of the above changes in vorticity, it is possible to view the lee cyclogenesis as

follows: a cyclonic disturbance embedded within a zona1 flow interacts with the mountain

anticyclone, reducing its effective relative vorticity. Upon traversing the mountain range,

the initial cyclonic maximum in relative vorticity re-appears and gives rise to a "new" lee

cyclone.

A vorticity budget was calculated for the SEP24 case in order to quantify the

relative importance of the proceses which led to the increases in vorticity in the lee.

According to the vorticity equation (Bluestein 1992, p.254), two terms contributing to the

generation or destruction of vorticity are the divergence and tilting terms. The

contribution of the divergence term (with a maximum of 14 x 10" s -~) was found to be

roughly 3 times as important as the tilting term (maximum of 5 x 10" s-') in generating

cyclonic vorticity at low levels in the lee (not shown).

3.3.3 Combined effects of latent heating and topography

To better elucidate the relative importance of the mountains and latent heat release,

a third sensitivity test was perfomied wherein both parameters were removed (Exp.

DNM). As in the two previous sensitivity experiments, the cyclone's path was essentially

unaffected. Since the removal of latent heat release (topography) tended to increase

(decrease) central pressures with respect to Exp. CTL, Exp. DNM yielded intermediate

pressures, but which were slightly more influenced by the release of latent heating (Fig.

3.13).

The above results show that if a significant arnount of latent heat is released, the

effect can overwhelm the filling of the cyclone due to low-level blocking, as will be seen

in the next case study. Thus, although precipitation is strongly infiuenced by topography,

the associated release of latent heat can provide a feedback and influence the location,

strength and structure of an extratropical cyclone. This consideration suggests that

precipitation patterns were altered in Exp. NMT. A cornparison of precipitation

accumulations between Exps. NMT and CTL for the SEP24 case reveals a rioticeable

reduction in accumulations right along the coasts of Alaska and British Columbia (cf.,

Figs. 3.3 and 3.18). The pronounced rainshadow evident in Exp. CTL (Fig. 3.3b) was

absent in Exp. NMT (Fig. 3.18b), and once the cyclone moved downstream of the

mountains, precipitation actually increased in Exp. NMT. This may be due to the

advection of more moisture into the lee which was previously blocked at low levels, or

due to a redistribution of the moisture which did not becorne orographically induced

precipitation. The absence of a rainshadow in Exp. NMT indicates that its presence in

the control simulation was due more to topography than to the dry intrusion previously

mentioned. In any event, although the lee of major mountain ranges are known to be

much drier than the windward sides, the drying caused by the mountains can extend to

downstream distances much greater than the width of the mountains.

The relative importance of the forcing mechanisms for the deepening of the lee

cyclone can be seen in Fig. 3.13. Between 12/25-24 and 18/25-30, the two experiments

where topography is included (Exps. CTL and NLH) exhibited deepenings of 2 hPa and

1.5 hPa, respectively, whereas the other two (Exps. NMT and DNM) did not. This

indicates that topographie effects were responsible for the deepening of the lee cyclone

during this period. Similarly, between 18/25-30 and 00126-36, the two experiments where

latent heating was included (Exps. CTL and NMT) exhibited slight deepenings of about

1 P a , whereas the other two (Exps. NLH and DNM) did not. Thus, latent heat release

accounted for the deepening of the lee cyclone in the last 6 hours of the simulation. It

was discussed in the preceeding chapter that the topographically-related deepening

mechanisms becarne important as early as 06/25-18. However, the effect of the

topography on the deepening is not apparent in Fig. 3.13 between 06/25-18 and 12/24-24.

The reason is that during these 6 hours, it is the central pressure of the parent cyclone that

is plotted because the decaying cyclone was still the "prirnary disturbance".

The changes in the vertical structure of an orographically modified extratropical

cyclone c m be understood by comparing cross sections of potential temperature from

Exps. NMT and CTL for the SEP24 case (Fig. 3.19). Differences between the two

experiments were largest directly over and downstream of the mountains, with smoother

isentropes in Exp. NMT, indicating that the mountains were responsible for generating

mesoscale structure throughout the troposphere. Slightly higher static stability was found

over the spine of the mountains at 00/25-12 in Exp. NMT (Fig. 3.19a). At 12/25-24, (Fig.

3.19b) weaker static stability was again recorded in a vertical column directly above the

cyclone centre in Exp. CTL. The biggest difference between the two experiments at this

time was found on the upwind side of the mountains of Alaska, where a more pronounced

trough existed in the lower troposphere in Exp. CTL. This trough, which amplified

slightly over the next 12 hours (Fig. 3.19c), was the upper-level counterpart of a surface

trough extending westward across the mountains of Yukon and AIaska from the low

centre located over Banks Island at this time in Exp. CTL (Fig. 3.ld). This feature was

conspicuously absent in Exp. NMT where the cyclone translated northeastward with a

speed slower than that of the control simulation, leading to a slight difference in the

location of the low-level trough by 00126-36 (Fig. 3.19~). At al1 times, the location of

the surface cyclone for both runs was clearly discemable by following the column of

minimum tropospheric static stability accompanied by a local depression of the tropopause

(Fig. 3.19).

The vertical structure of the cyclone changed more significantly in the absence of

latent heat release. Figure 3.20a indicates that by 00125-12, there was already a greater

amplitude in the ridge of isentropes directly above the mountains. By 12/25-24 (Fig.

3.20b), this ridge was amplified even more, and was quite distinguishable from the

relatively flat isentropes of Exp. CTL. Without the warming provided by latent heat, any

given point in the troposphere was colder in Exp. NLH than in Exp. CTL, particularly in

the lower- to rnid-troposphere. This ndging was strongest in the core of the cyclone,

where the largest amount of latent heat was released. This feature also overwhelmed any

flattening of isentropes noted in Exp. CTL, and only amplified the ridging provided by

the mountains. Despite this, once the disturbance moved into the lee, a low-level trough

in the isentropes becarne more evident (Fig. 3.20~). This trough tilted eastward with

height in Exps. CTL and NLH, an unfavourable condition for further development as

confirmed by the filling of the cyclone after 00126-36.

Fig. 3.1 CMC analysis of sea-level pressure at intervals of 2 hPa for (a) 1200 UTC 24; (b) 0000 UTC 25; (c) 1200 UTC 25; (d) 0000 UTC 26 September 1994; subjectively analysed frontal positions are also shown. Dashed line in (a) indicates position of continental divide.

Fig 3.1 (continued)

Fig. 3.2 Sea-level pressure nt intervals of 2 hPa (a) 12-h, (b) 24-h and (c) 36-h control simulations, valid at 0000 UTC 25, 1200 UTC 25 and 0000 UTC 26 September 1994, respectively. Line AA' (BB') in (a) ((c)) shows the location of cross-sections used in Figs. 3.12,3.19and 3.20 (3.5).

(cl

Fig. 3.3 Distribution of 12-h accumulated precipitation from control simulation contoured at 1,5,10,20 and 40 mm ending at (a) 00125-12; (b) 12-25-24 and (c) 00126-36. Light (heavy) shading indicates accumulations between l(5) and 5 (10) mm as well as over 20 (40) mm. (d) Observed precipitation accumulations for 24-h period ending 0000 UTC 26 September 1994. Only arnounts greater than a trace are indicated.

32

Fig. 3.3 (continued)

Fig. 3.4 Location of surface stations used in verification of precipitation accumuIations.

Fig. 3.5 Vertical cross-section of vertical motion at intervals of 0.1 Pa/s for 0600 UTC 25 September 1994 taken dong line BB' given in Fig. 3 . 2 ~ . Solid (dashed) contours are for positive (negative) values, indicating descent (ascent).

Fig. 3.6 (a) 925-hPa divergence at intervals of 10 x 10'~ s-' and (b) 925-hPa relative vorticity at intervals of 2x 1 0 - ~ s-' from 24- h control simulation valid at 1200 UTC 25 September 1994. Darkened area indicates area with surface pressure less than 925 hPa. Thick, solid line in (a) indicates zone of divergence. Solid (dashed) contours are for positive (negative) values.

(c)

Fig. 3.7 Distribution of temperature at intervals of= and horizontal winds at 850 hPa from (a) 12-h, (b) 24-h and (c) 36-h control simulation valid at OOOO UTC 25,1200 UTC 25 and 0000 UTC 26 September 1994, respectively. Darkened area indicates area with surface pressure l e s ~ than 850 hPa. Letter 'M' in (c) denotes position of maximum low-level baroclinicity, as discussed in text. Solid (dashed) contours are for positive (negative) values.

36

Fig. 3.8 Distribution of relative hurnidity at intervals of 10% at 700 hPa from (a) 12-h, (b) 2 4 4 and (c) 36-h control simulation valid at 0000 UTC 25,1200 UTC 25 and 0000 UTC 26 September 1994, respectively. Light (heavy) shading indicates relative humidity in excess of 70% (90%).

~ i ~ . 3-10 Geopotential height (solid) at intervals of 6 dam at 500 hPa and potential vorticity (PV, dashed) at intervals of 0.5 PVU at 300 hPa from (a) O-h, (b) 12-h, (c) 24-h and (d) 36-h control simulation valid at 1200 UTC 24,0000 UTC 25,1200 UTC 25 and OOOO UTC 26 September 1994. Light and heavy shading indicate PV greater than 2 and 4 PVU, respectively. Thick "L" indicates position of surface cyclone centre. Thick, dashed line in (c), (d) indicates position of 500-hPa trough.


41

Fig. 3.11 The 6 hourly positions of the surface cyclone centre (L) and the 300-hPa PV anomaly (X) from control simulation. Dashed lines connect multiple centres of either feature if relevant.

Vertical cross section of potential vorticity (solid) at intervals of 0.5 PVU, potential temperature (dashed) at intervals of 3 K, and dong-plane flow vectors, which is taken dong line AA' given in Fig. 3.2a for (a) 121 24-00, (b) OO/2S- 12, (c) 12/25-24 and (d) 00/26-3 6. S hading indicates potential vorticity greater than 1.5 PVU. Thick solid line in (a) and @)

indicates location of tropopause fold. Scales of vertical motion (Pa s*') and wind speed (m s-') are indicated in bottom left-hand corner. Position of the surface cyclone is indicated dong the abscissa. Note that the lowest vertical level in cross section is 900 hPa.

43

Fig . 3.1 2 (continued)

Fig. 3.13 Time series of central pressure for the SEP24 case from Exps. CTL, NLH, NMT and DNM.

Fig. 3.14 As in Fig. 3.2, but for Exp. NMT.

Fig. 3.15 As in Fig. 3.1 1, but for Exp. NMT.

Fig. 3.16 As in Fig. 3.7, but for Exp. NMT.

Fig. 3.17 Time series of relative vorticity (x 10-~ s-') calculated over a circular area of 200 km in diarneter about the cyclone centre and over the five 1owest'~aLchen levels for Exps. CTL and NMT.

(cl

Fig. 3.18 As in Fig. 3.3, but for Exp. NMT

1 O0

200

250

300

100

500

(a) GnO

Fig, 3.19 Isentropic cross-sections at intervals of 5 K taken dong line AA' in Fig. 3.2a from Exp. CTL (solid) and from Exp. NMT (dashed) for (a) 00125- 12, (b) 12/25-24, (c) 00/26-36.

I

Fig. 3-20 As in Fig. 3.19, but for Exp. CTL (solid) and Exp. NLH (dashed).

Chapter 4 - Numerical simulations of the SEPS case

This chapter starts with a verification of the control simulation (Section 4.1),

followed by a detailed analysis of the case (Section 4.2), and concludes with a discussion

of various sensitivity tests (Section 4.3).

4.1 Verification of control simulations

The SEP8 case began with a cyclone with a central pressure of 1007 hPa over the

eastern Pacific (Fig. 4.la), which moved towards the Coast while filling. Twenty-four

hours later (Fig. 4.lb), a weaker low (1009 hPa) remained over the ocean and a lee

cyclone with a central pressure of 1006 hPa was taking shape over southem Alberta. The

lee cyclone then moved southeastward while deepening (Fig. 4.lc), reaching a central

pressure of 1000 hPa by 0000 UTC 10 September 1994 (hereafter, 00/10-48) with an

inverted trough extending northward from the low centre over al1 of Saskatchewan. The

lee cyclone had a well-defïned circulation by the end of the 72-h integration (00/11-72),

at which time it attained an analyzed central pressure of 997 hPa and was sufficiently far

from the mountains to no longer be significantly affected by their presence. The

difference in central pressure between the lee cyclone at 0011 1-72 and the parent cyclone

at 00/08-00 yields a total deepening of 10 hPa in 72 hours. The simulation results (Fig.

4.2) indicate that evolution of the central pressure of the parent and lee cyclones were

well captured, although some noticeable differences existed in the structure of the SLP

fie1.d. Following the dynamic initialisation, the filtering and smoothing of the SLP field

(see Chapter 3), pressures in the lee of the Rockies were up to 4 hPa lower than in the

corresponding analysis (cf., Figs. 4. l a and 4.2a). The location and central pressure of the

parent cyclone, however, showed good agreement with the analysis. The coastd approach

and gradua1 decay of the parent cyclone over the next 24 hours were well simulated by

the MC2 (Figs. 4.lb and 4.2b), with difference of only 1 hPa in central pressure of the

parent cyclone at 00109-24. In the lee, the analysis indicates a pressure of 1006 hPa in

southeastern Alberta (Fig. 4. lb) which compares to a pressure of 1005 hPa in the vicinity

of Havre, Montana (HVR, see Fig. 3.4) in the control simulation (Fig. 4.2b). At 00110-

48, Fig. 4 . 2 ~ shows only a trough remaining of the parent cyclone over the eastem Pacific

(as per the analysis), as well as a 24-h drop of 5 hPa in centrai pressure of the lee cyclone

centered over Kindersley, Saskatchewan (YKY) at this tirne. Interestingly, the MC2

developed a second low centre at the intersection of the borders of Saskatchewan,

Montana and North Dakota, which was roughly 200 km from the analyzed low centre

over Glasgow, Montana (GGW) and whose central pressure was within 1 hPa of the

analyzed value of 1000 hPa. The most noticeable difference between the simulation and

the analysis at this time is the rnuch weaker inverted trough extending northward from

the lee cyclone centre in the simulation. Finally, at 00/11-72, the MC2 correctly

developed a stronger pressure gradient around the lee cyclone centre than in the previous

time frarne (cf., Figs. 4 . 2 ~ and 4.2d). The cyclone deepened by 3 hPa in the final 24

hours of the simulation, although the location of the centre of the low over north-central

Saskatchewan differs from the analysis which places it in north-central Manitoba. The

now-occluded front which extended southeastward from the low centre in the simulation

(Fig. 4.2d) joined up with its indicated position in the analysis (Fig. 4. ld) over extreme

southern Manitoba.

Figures 4.3a-c give the 24-h accumulations of precipitation for the course of the

entire simulation. Most of the precipitation falling in the first 24 hours of the simulation

did so over much of British Columbia (Fig. 4.3a). Although the maximum of 24.8 mm

in this period is not verifiable since it occurred in a data-sparse region, the accumulation

pattern produced by the mode1 corresponds closely to observations (Fig. 4.3d); e.g. the

"inverse L" pattern over British Columbia (B.C.), the minimum over southeastern B.C.,

the "tongue" of precipitation reaching into north-central Alberta. Over the next 24 hours

(Fig. 4.3b), 1 1.1 mm accurnulated near Fort McMurray, Alberta (YMM) comparing to

average values 10 mm in the observations (not shown). The maximum of 25 mm over

southwestern B.C. compares favourably to the 29 mm reported near Abbotsford, B.C

(YXX) (not shown). Finally, in the last 24 hours (Fig. 4.3c), accumulations in a band

oriented from southwest to northeast occurred over most of northern Saskatchewan and

east-central Alberta. Reported accumulations indicate a similar distribution, with amounts

of close to 25 mm common throughout most of this band and 26 mm reported at Buffalo

Narrows, Saskatchewan (YVT) which is roughly 15 mm less than the maximum of 40.8

mm simulated by the MC2. The maximum of 28.7 mm near Jasper, Alberta (WJW) is

not verified in the observations, although the persistent maximum in southwestern B.C.

of 36.6 mm is corroborated by a report of close to 50 mm in the area. Also of note is

the isolated "patch" of accumulations greater than 5 mm near Baker Lake (YBK) which

also verifies against observations of 6 and 7 mm in the area. This occurred despite the

difference in location of the low centre previously indicated, due to which the simulated

low was centered further West than indicated in the analysis. Verification of other fields

such as temperature and geopotential heights confirmed that the MC2 satisfactorily

reproduced most of the pertinent features of this case of lee cyclogenesis.

4.2 Evolution of the SEPS case

As discussed in the previous chapter, downslope winds in the lee of the Rockies

can generate a low-level spin-up of relative vorticity. This downward motion occurred

in the foothills of the Rockies, as seen in the cross section taken in Fig. 4.4. Also

apparent in this figure is the vertical stretching of columns which was discussed in the

previous chapter, as attested to by the area of ascent imrnediately above the strong descent

in the foothills of the Rockies. As a consequence of this stretching, a low-level spin-up

of relative vorticity occurred and can be seen 12 hours later as a maximum of magnitude

13.5 x 105 s" in a location where it is advected downstream by the low-level winds (Fig.

4.521). The location of this vorticity maximum corresponds closely to the location of

enhanced low-level convergence at the same time (Fig. 4.5b). Strong ascent occurred as

well (Fig. 4.5c), helping to maintain the low-level convergence and increasing the relative

vorticity so that it reached a value of 1.8 x IO4 s-' at 850 hPa by 00/11-72 (not shown).

Warming owing to downslope motion to the east of the Rockies acted in concert

with warm advection (not shown) to produce a ridge in the thermal field (Fig. 4.6). Note

the close correspondence of this thermal ridge with the location of the SLP trough (cf.,

Figs. 4 . 2 ~ and 4.6b). The development of a low-level jet in the SEP8 case helped to

advect continental tropical air into the warm sector of the developing lee cyclone (Fig.

4.6). The presence of a low-level jet over the Great Plains of the United States is well

documented (Wexler 1961 ; Bonner 1968), and has been shown to be strongest (weakest)

early at night (during the early morning) at about 850 hPa (Bluestein 1993). The jet

observed in the SEPS case underwent a similar diurnd variation due to differential

heating on isobaric surfaces near sloping topography. During the nighttime, the

temperature contrat between the eastem slopes of the Rockies (surface pressure = 850

hPa) and the air in the boundary Iayer to the east was reduced compared to the daytime

(Fig. 4.6a), so that the strength of the jet decreased, and winds relaxed to the direction

and speed dictated by the synoptic scale situation. During the daytime, this terrain heated

up more than the air at the same pressure further to the east (Fig. 4.6b), effectively

lowering heights over the mountains, so that southerly winds blowing parallel to the

mountains were strengthened. The exact role of the low-level jet in the lee cyclogenesis

requires further study and will not be discussed here, but is rather mentioned as a feature

of the lee cyclogenesis.

In comparing Figures 4.lc and 4.2c, one notices a difference in location of the

centre of the lee cyclone between the analysis and the simulation. It is interesting to note

that the location of the low centre given in the simulation corresponds exactly to the

location of both the enhanced low-level convergence and the relative vorticity maximum

just discussed (Figs. 4.5a,b). Furthermore, an inspection of the accumulated precipitation

in the final 24 hours of the simulation (Fig. 4 . 3 ~ ) indicates that the location of the low

centre over the same period was close to the maximum in precipitation. In exarnining the

breakdown of precipitation in terms of convective and stratiform contributions to the total

pattern (not shown), it is found that 93% of the maximum near Buffalo Narrows,

Saskatchewan was convective in nature (38 of 41 mm). Satellite imagery (not shown)

confirms the presence of convective cells near the cyclone centre between 00/10-48 and

0011 1-72. The proper simulation of this convection by the MC2 and the associated latent

heat release hydrostatically lowered pressures at the surface, with the result that the MC2

placed the cyclone centre at the location of the most intense convection. This feature of

the lee cyclogenesis will be further examined later in this chapter.

An inspection of precipitation accumulations (Fig. 4.3) does not generally show

a pronounced rainshadow as observed in the previous case. Accumulations in B.C.,

56

however, do tend to mirror the topography, with slight reductions in the Okanagan Valley,

and maxima dong local mountain ranges. The reason the drying in the lee in this case

is not as pronounced as in the SEP24 case is that the circulations of both the parent and

lee cyclones are weaker than in the latter. The weaker pressure gradient throughout most

of this simulation results in weaker winds and a lesser advection of moisture from the

Pacific.

At mid-levels of the troposphere (500 hPa), a planetary-scale trough was initially

present over the eastern Pacific (Fig. 4.7a) with a ridge over the Prairie provinces. Over

the next 24 hours, the trough and its associated PV maximum moved eastward as heights

fell dong the West Coast (Fig. 4.7b). By 00110-48 (Fig. 4.7c), there began to be some

indications of the development of an upper-level shortwave, as suggested in particular by

the presence of two maxima in the PV field - one of magnitude 7.42 PVU over north

Vancouver Island and another associated with the developing shortwave over northeastem

Washington state of magnitude 4.44 PVU. At 00/11-72 (Fig. 4.7d), the presence of the

shortwave trough was more obvious with the minimum height near Buffalo Narrows,

Saskatchewan (YVT). The associated PV maximum now moved to the Alberta-

Saskatchewan border, and decreased in magnitude (on the sarne isobaric surface) to 3.76

PVU. It was distinct from the 6.02 PVU maximum just east of Vancouver associated

with the synoptic-scale trough which was now closely aligned with the West Coast of

North America.

Upon inspection of the relative positions of the ULPVA and the surface cyclone

(Fig. 4.8), it c m be seen that the distance between the two continually decreased between

00/09-24 and 00/11-72. The period of maximum central pressure deepening, however,

took place between 00/09-24 and 00/10-48 (see Fig. 4.2). It thus appears that there was

not a strong correlation between SLP falls in the cyclone centre and the relative

positioning between it and the ULPVA.

4 3 Sensitivity experiments

Having seen that the control simulation successfully reproduced the saiient features

of the SEP8 case, it is now possible to conduct sensitivity tests in order to isolate the

effect of a given parameter.

43.1 Effect of latent heating

A "dry mn" as descnbed in the previous chapter was also conducted on the SEP8

case. Figure 4.9 shows the time series of central pressure for the dry run (Exp. NLH) as

well as for the control simulation (Exp. CIL). The resulü indicate that over the first 48

hours of both simulations, SLP dropped continually, with almost identical values between

the two simulations1. The most noticeable differences in central pressure occurred

between 00110-48 and 00/11-72, with SLP increasing by roughly 4 hPa in Exps. NLH and

SLP decreasing by 3 hPa in Exp. CTL. It was mentioned in section 4.2 that the

development of convection over parts of Saskatchewan was an important feature of the

control simulation. For cornparison, Fig. 4.10 plots the SLP field for 00/10-48 and for

0011 1-72. Perhaps the most striking difference between Figs. 4.10a and 4 . 2 ~ is the

placement of the lee cyclone centre. The removal of latent heat release in Exp. NLH

resulted in the centre k i n g located over northwestem Nonh Dakota (Fig. 4.10a), close

to the location of a closed isobar in Exp. CIZ (Fig. 4 . 2 ~ ) which indicated the presence

of a weaker second centre of Iow pressure. The presence of the convection was thus

crucial in determining both the strength and the location of the lee cyclone. This notion

is reinforced when considering the same situation 24 hours later (cf., Figs. 4.10b and

4.2d). The lee cyclone in Exp. NLH was some 7 hPa higher in central pressure than in

Exp. CTL, as well as k i n g to the northeast of the previous location. The MC2

effectively placed the centre dong the inverted trough which extended northeastward from

the low centre in Exp. Ci l (Fig. 4.2d). Interestingly, the location of the cold-occluded

' In fact, the central pressures indicated in Exp. NLH are slightly lower than in Exp. CïL. This is a r m l t of the filtering of the SLP field previously d d b e d which tends to produce slightly iower pressures at sea-level as compared to the raw mode1 output SLP field for both Exps. CIZ and NLH.

front meridionally oriented across Manitoba was essentially unchanged between the two

simulations (cf., Figs. 4. lob and 4.2d).

4.3.2 Effect of topography

To elucidate the role which topography played in the SEP8 case, a simulation

where it was removed (Exp. NMT, as described in the previous chapter) was performed

and compared to Exp. CTL. Figure 4.9 depicts the time series of central pressure of the

lee cyclone in both Exps. CTL and NMT. The biggest differences between the two

occurred after 00/10-48, in Exp. NLH, where removal of the topography resulted in

central pressures up to 5 hPa lower in Exp. NMT than in Exp. CTL by 0011 1-72. During

the majority of the simulation, however, the central pressure of Exp. CTL was slightly

lower than in Exp. NMT, albeit by less than 1 hPa. This c m be explained by the fact

that in Exp. CTL winds blowing downslope led to adiabatic warming and the associated

lowering of SLP, this did not occur in Exp. NMT, as discussed in the preceding chapter.

In addition to the increased differences in central pressure between Exps. CTL and NMT,

the SLP plot for Exp. NMT over the last 24 hours of the simulation indicates an

important change in the location of the lee cyclone (Fig. 4.1 1). By 00/10-48 (Fig. 4.1 la)

the lee cyclone was displaced slightly to the northwest of its location in Exp. CTL which

is a result of the fact that winds in this location were blowing upslope and low-level

blocking of mass was taking place. This was not the case in Exp. NMT, so lower SLP

resulted. This pattern continued over the next 24 hours so that by 00/11-72, the low

centre in Exp. NMT was just east of Great Slave Lake, north of its location over northem

Saskatchewan in Exp. CTL (cf., Figs. 4.26 and 4.1 lb).

Increases in the lee of low-level relative vorticity were detected by calculating this

quantity around the cyclone centre, as described in the previous chapter. Performing the

same calculation for the SEP8 case (Fig. 4.12) indicates an increase in relative vorticity

between 00/10-48 and 00111-72 which was greater in Exp. CTL than in Exp. NMT'.

The figure shows the trace of relative vorticity only after 12/09-36 because it is only after this tirne that a well-defined lee cyclone centre can be identified.

However, relative vorticity continually increased in Exp. NMT, which indicates that

cyclogenesis occurred in this case even in the absence of the mountains. Therefore, in

this case, topography served to modify the cyclogenesis, but did not create it.

Analyzing the vorticity budget calculations performed on the SEP8 case indicates

that the lee increase in relative vorticity noted in Exp. CTL was due to both contributions

from the divergence term and the tilting term. At 00/10-48, a broad area of positive

vorticity tendency with maxima of 7.84 x 10" s-2 and 13.2 x 10-~ ss.' in the divergence

term was oriented NW-SE across southern Saskatchewan (Fig. 4.13a) and was located

close to an area of positive vorticity tendency in the tilting term with one maximum of

11 x 10~' s-' (Fig. 4.130. The maximum in the tilting terrn acted in concert with the

divergence term maximum of 7.84 x 10" sJ; the stronger maximum in the divergence

term, however, was CO-located with a minimum in the tilting term, both of which were

at the exact location of the lee cyclone in Exp. CTL (Fig. 4.2~). The location of the two

maxima was coincident with the closed isobar described in Fig. 4 . 2 ~ . Subsequent to this

tirne, as the lee cyclone moved away from the mountains, the contribution of the tilting

term decreased relative to that of the divergenceterm. If the stronger contribution of the

tilting terrn was in sorne way related to the presence of the rnountains, it should have

been greatly diminished in Exp. NMT. The vorticity budget for Exp. NMT (not shown)

did show a decrease in the contribution from the tilting term relative to that of the

divergence term, a reduction of about 23% in magnitude. This lessened contribution may

explain why although vorticity increases in Exp. NMT, it did not attain as large values

as in Exp. CTL.

4.3.3 Combined effects of latent heating and topography

A third sensitivity experimeht was conducted wherein both latent heat release and

topography were removed (Exp. DNM). Figure 4.9 shows the central pressure trace for

this experirnent which resembled the trace of Exp. CTL and remained within 3 hPa of it

at a11 times because of the opposing effects of the removal of both latent heating and

topography .

Removing topography has been shown to alter the distribution of precipitation

accumulations in the previous chapter. Figure 4.14 shows the 24 hourly accumulations

of precipitation for Exp. NMT. A cornparison with Exp. CTL indicates that during the

first 24 hours, the general pattern of accumulations as well as the maximum amounts over

north-central B.C. were comparable in both simulations (cf., Figs. 4.3a and 4.14a).

Between 00/09-24 and 00/10-48, the 25 mm maximum in southwestern B.C. in Exp. CTL

(Fig. 4.3b) was reduced in magnitude to just over 10 mm in Exp. NMT (Fig. 4.14b),

whereas a "lee" maximum of 21.2 mm was observed in the latter. As discussed in the

previous chapter, this is likely due to enhanced advection of rnoisture from the Pacific

which was previously blocked by the high topography and "wrung out" along the

rnountains. The same observation is true of the accumuIations in the last 24 hours (cf.,

Figs. 4 . 3 ~ and 4.14c), with the precipitation maxima in Exp. NMT displaced northward

of their location in Exp. CTL, which is consistent with the noted northward displacement

of the lee cyclone centre (cf., Figs. 4.2d and 4.1 Ib). The maximum in precipitation over

the last 24 hours was 36 mm in Exp. NMT (Fig. 4 .14~) which compared to the maximum

of 41 mm over this same period in Exp. CTL (Fig. 4 . 3 ~ ) . This difference represents a

reduction of about 10% in the magnitude of the maximum. If one accounts for the fact

that the cyclone was located father north in Exp. NMT, and considers the overall pattern

of precipitation, the two distributions have a similar orientation, with amounts generally

greater over a wider area in Exp. NMT than in Exp. CTL (cf., Figs. 4 . 3 ~ and 4.14~). This

again relates to the fact that more moisture was advected into this area from the Pacific

and so was available for greater precipitation production.

Isentropic cross sections taken SW-NE through both the upper-level and surface

lows indicate that the trough associated with the lee cyclone tended to amplify with tirne,

as the surface low continually deepened (Fig. 4.15). A cornparison between Exps. CTL

and NMT indicates that the amplitude of the Iee trough was actually greater in Exp. NMT

(e.g. Fig. 4.15b), leading to a more robust surface cyclone (cf., Figs. 4.2d and 4.1 lb). A

slight westward tilt in the trough is noted with increasing height for both Exps. CTL and

NMT. Because of the difference in location of the low between Exps. CTL and NMT,

the lee trough in Fig. 4 . 1 5 ~ also shows a slight difference in location.

As was found with the previous case, comparing isentropic cross sections between

Exps. CTL and NLH (Fig. 4.16) indicates that higher static stability was present in the

latter, especially at lower levels of the atmosphere (Fig. 4.16b,c). As a result of this

increase in static stability, a low-level ridge formed in the isentropes by 00/11-72,

effectively reducing the strength of the lee cyclone. This is consistent with the prior

observation that the release of latent heat was an important factor in the formation of the

SEP8 lee cyclone.

4.3.4 Effect of domain size and position of the lateral boundaries

It is generally desirable to place the lateral boundaries of limited-area models far

enough upstream from the region of interest so that errors at the boundaries do not

contaminate the forecast. This strategy proved to be successful only for the SEP24

simulation. For the SEP8 case, a "picket fence" domain led to the most realistic

simulation. In the "picket fence" experiment, the western (upstream) lateral boundary was

placed dong the W e s t Coast of North America, so that this boundary was driven by

observations. Perforrning the simulation on a larger domain allowed errors from the data-

sparse Pacific to enter the region of interest, which led to a less accurate representation

of the cyclone. Sensitivity tests, which will not be described in any detail here, were

conducted to determine what the best placement for the "picket fence" would be.

Qualitatively, the results of this experimentation indicated that the farther east (west) the

boundary was placed, the lower (higher) the final central pressure of the lee cyclone

became.

The fact that excluding the data-sparse upstream ocean areas improved only one

of the simulations is related to the initial conditions. The SEP24 case begins with a very

strong parent cyclone, which continues to exert a strong signal for the remainder of the

simulation. In contrast, a much weaker parent cyclone is present in the initial conditions

of the SEP8 case and the lee cyclone develops over a rnuch longer period of time (72

hours compared to 36 hours for the SEP24 case).

Fig. 4.1 CMC analysis of sea-level pressure at intervals of 2 hPa of (a) 0000 UTC 8, (b) 0000 UTC 9, (c) 0000 UTC 10 and (d) 0000 UTC 11 September 1994; subjectively analyzed frontal positions are also shown. Dotted line in (a) indicates position of continental divide.

Fig. 4. l (continued)

Fig. 4.2 Sea-Ievel pressure at intervals of 2 hPa (a) O-h, (b) 24-h, ( c ) 48-h and (d) 72-h control simulations, valid at 0000 UTC 08,0000 UTC 09,0000 UTC 10 and 0000 UTC 11 Sep- tember 1994, respectively. Line AA' (BB') in (a) shows the location of cross-section used in Fig. 4.4 (4.15,4.16).


66

Fige 4.3 Distribution of 24-h accumulated precipitation from control simulation contoured at 1,5,10,20 and 40 mm ending at (a) W09-24; (b) 00110-48 and (c) 0011 1-72. Light (heavy) shading indicates accumulations between l(5) and 5 (1 0) mm as well as over 20 (40) mm. (d) Observed precipitation accumulations for 24-h period ending OOûû UTC 1lSeptember 1994. Only amounts greater than a trace are indicated.

67


L4T 40.5 49.3 50.0 50.5 50.8 1.0% 1SO.n -116.7 - 112.7 -108 6 - 104.4

Fig. 4.4 Vertical cross-section of vertical motion at intervals of 0.1 Pals for 1200 UTC 9 September 1994 taken dong line AA' given in Fig. 4.2a Solid (dashed) contours are for positive (negative) values, indicating descent (ascent).

Fig. 4.5 (a) 850-hPa relative vorticity at intervals of 2 x 105 s-' and (b) 850- hPa divergence at intervals of 10 x 1 o - ~ s- l and (c) 850-hPa vertical motion at intervals of 0.1 Pa/s from 48-h control simulation valid at OOOO UTC 10 September 1994. Darkened area indicates area with surface pressure less than 850 hPa. Solid (dashed) contours are for positive (negative) values.

69

Fig. 4.6 Distribution of temperature at intervals of 2 C and horizontal winds at 850 P a for (a) 1200 UTC 09 and (b) OOOO UTC 10 September 1994 from control simulation. Darkened area indicates area with surface pressure l e s ~ than 850 hPa. Shading indicates temperatures above 24 C.

Fig. 4.7 Geopotential height (solid) at intervals of 6 dam at 500 hPa and potential vorticity (PV, dashed) at intervals of 0.5 P W at 300 hPa from (a) O-h, (b) 24h, (c) 48-h and (d) 72-h control simulation valid at 0000 UTC 8, 0000 UTC 9,0000 UTC lOand 0000 UTC 1 1 September 1994. Light and heavy shading indicate PV greater than 2 and 4 PVU, respectively. Thick "L" indicates position of surface cyclone centre. Thick, dashed line in (c), (d) indicates position of 500-hPa trough.

Fig. 4.7(continued)

Fig. 4.8 The 12 hourly positions of the surface cyclone centre (L) and the 300-hPa PV anomaly (X) from control simulation.

CTL- CONTROL

NLH-NO LATENT HEAT

NMT-NO MOUNTAINS

DNM-DRY AND NO MNT.

0 1 2 24 36 48 60 72 Fig. 4.9 Time series of central pressure for the SEP8 case from Exps.

CTL, NLH, NMT and DNM. j

Fig. 4.1 1 As in Fig. 4.10, but for Exp. NMT.

Fig. 4.12 Time series of relative vorticity (x IO-' 8) calculated over a circular area of 200 km in diameter about the cyclone centre and over the five lowest Gd-Chen levels for Exps. CTL and NMT between 12/09-36 and 00/11-72.

Fig. 4.13 850-hPa contribution of (a) divergence term and (b) tilting 9 2 term at intervals of 1 x 10- s' to vorticity tendency for

0000 UTC 10 September 1994 from control simultion. Solid (dashed) contours are for positive (negative) values.

As in Fig. 4.3, but for EXD. NMT.

Fig. 4-15 Isentropic cross-sections at intervals of 5 K taken dong line BB' in Fig. 4.2a from Exp. CïL (solid) and h m Exp. NMT (dashed) for (a) ûû/09-24, @) 00/10-48, (c) 00/11-72..

79

B' Fig. 4-16 As in Fig. 4.15, but for Exp. C'IL (solid) and Exp. NLH

(dashed).

Chapter 5 - Conclusion and summary

The interaction of extratropical cyclones with topography has been investigated

in this thesis based on results from numerical simulations. Two cases of lee cyclogenesis

were analyzed, one which occurred during 24-26 September 1994 (BASE IOP 7), the

other which occurred during 8-1 1 September 1994. The MC2 mode1 is found to

reproduce well the evolution of both cyclones, as well as their associated precipitations.

In both cases, the initial disturbance fills while approaching the coast due to an

equivalent barotropic structure throughout the troposphere and low-level mass build-up

along the coast. It is shown that lee cyclogenesis involves both upper-level and low-level

pracesses. Low-level stretching of air columns, and the associated increase in relative

vorticity plays an important role in both lee cyclogeneses. In the SEP24 case, blocking

by orography increases baroclinicity during the lee cyclogenesis. At upper levels, a

synoptic-scale trough is quasi-stationary over the eastern Pacific, and both lee

intensifications are coincident with the formation of an upper-level shortwave which

emerges from the parent trough. Although not shown for the SEP8 case, the upper-level

shortwave which forms in Exp. CTL is also present in Exp. NMT. The time series

depicted in Figs. 3.10 and 4.7 indicate that the shortwave "breaks off' from the synoptic

scale trough located farther to the West. Alternatively, the ULPVA which provides

support for the lee cyclone becomes detached from the larger T V reservoir" and

propagates to the east. The presence of this feature in both Exps. CTL and NMT in both

cases indicates that topography is not responsible in inducing the upper-level shortwave.

Our analysis shows the formation of a low-level jet in both cases, aithough it is

weaker in the SEP24 case owing to a weaker zona1 temperature gradient along the eastern

slopes of the Canadian Rockies. This occurred because the SEP8 case takes place in the

summer season (versus the faIl for the SEP24 case), and farther south, so that surface

temperatures in the area of lee cyclogenesis are significantly warmer in this case. Surface

temperatures for the SEP8 case range from 2530°C at 002 in the region of cyclogenesis

as compared to surface temperatures for the SEP24 case of 04°C in the analogous region.

AIoft, i.e. 500 hPa, temperatures in the same regions are in the - 10°C to - 15°C (-15°C to -

20°C) for the SEP8 (SEP24) case. Given the lapse rates which result from such

temperature profiles, it is not surprising that convection develops and is a more important

factor in the SEP8 lee cyclogenesis than for the SEP24 case with al1 else being equal (or

similar) as in upper-level forcing.

Previous studies have indicated that the formation of a lee trough and cyclone can

differ substantially from the classical Norwegian cyclone model (Steenburgh and Mass

1994). The same observation cm be made for the lee development of the two cases

studied here. The cyclones in general resemble the classical model once they are

sufficiently far downstream from the mountains. However, closer to the mountains, the '

similarities become much less. In the SEP24 case, the lee cyclone develops along a

stationary front (Fig. 3.lb) which amplifies in time, although a lee trough forms to the

south of the cyclone centre owing to the strong downslope flow (Fig. 3. lc). This trough

line extends southward and joins up with a smaller Iow-pressure centre with its well-

defined warm and cold fronts (Fig. 3.1~). As the main Iow progresses northeastward, a

secondary cold front forms dong the trough extending westward towards the mountains

(Fig. 3.td). The emergence of a warm front to the southeast of the low centre gives it

a more classical appearance. The departures from the classical mode1 for the SEP8 case

are most pronounced during the formation of the lee cyclone (Figs. 4.1 a-c). Although by

00/10-48 (Fig. 4.lc), this lee cyclone has well defined warm and cold fronts which later

begun to occlude (Fig. 4.ld), the initial formation of the lee cyclone does not occur along

any clearly defined frontal boundary, but represents a response to differential vorticity

advection aloft. Upper-level cyclonic vorticity advection contributes to the spin-up of the

SEP24 lee cyclone.

Dry simulations of both cases revealed that the central pressure of the lee cyclone

was higher without the release of latent heat. Cornparison of the control and dry

simulations indicates that latent heating serves to modify the storm evolution, and is

instrumental in the lee cyclogenesis process, without which the lee intensification of the

cyclone becomes much weaker. Additionally, although the track of the low was not

significantly altered by the removal of latent heat release in the SEP24 case, this factor

was very important in the SEP8 case, where the development of convective precipitation

significantly affected both the central pressure and the location of the lee cyclone.

The importance of topography was elucidated in the "no rnountains" simulations.

The results of this experiment confirmed the importance of low-level blocking of the

flow, without which the cyclone achieved much lower central pressures in both cases.

The precipitation field was found to be strongly influenced by topography. In both cases,

accumulations dong the coastal mountain ranges (at inland locations) were reduced

(increased) when the mountains were removed from the simulations. Mountains alter the

distribution of moisture in the atmosphere over distances far greater than their widths,

which has important consequences for studies of the water budget such as the case for

GEWEX.

The question of whether lee cyclogenesis is a special type of cyclogenesis is

discussed. As has been shown, both ordinary cyclogenetic mechanisms and topographic

effects play a role in the lee cyclones studied. Thus, lee cyclogenesis does indeed have

characteristics different from ordinary cyclogenesis, but it also bears many similarities.

In fact, as Exp. NMT attests to, cyclogenesis occurs in both cases even in the absence of

mountains, dthough mountains serve to alter the process. As mentioned earlier, the fact

that so many cases of cyclogenesis are recorded in the lee of major mountain chains is

not indicative of a random occurrence. These areas are preferred sites for cyclogenesis

because of the general cyclolytic effect that mountains have on extratropical cyclones on

their upwind sides. Whether this cyclolysis is viewed as a temporary masking of a

disturbance or whether it is viewed as the dissolution of a parent cyclone and the genesis

of a lee cyclone, both viewpoints implicitly acknowledge that the effect of the rnountains

is to weaken an approaching cyclonic circulation. Conversely, the mountains have a

generally cyclogenetic effect on their downwind sides, where relative vorticity increased

in the lee.

While the parent low in the SEP24 case is initially deeper than its associated lee

cyclone, the opposite is true of the SEP8 case (cf., Figs. 3.1 and 4.1). However, if we

consider pressure decreases in the lee, it is the SEP24 case which demonstrates stronger

falls; SLP drops 21 hPa in 36 hours in this case, while the drop was only 13 hPa in 72

hours in the SEP8 case'. As pointed out in Chapter 1, there is some ambiguity as to

what exactly constitutes lee cyclogenesis. The "classical" cases of lee cyclogenesis are

situations where a parent cyclone over the ocean (identified by closed contours, e.g. in

the SLP field) completely disappears and is later followed by the appearance of closed

isobars in the lee of the rnountains (Bannon 1992). This describes the SEP8 case more

accurately than the SEP24 case, in which no complete dissipation occurs. However,

because the latter case demonstrates both a lee reintensification and stronger SLP falls in

the lee, it satisfies the definition of Pierrehumbert (1986) and Tibaldi et al. (1990) and is

accepted as a case of lee cyclogenesis (Hudson and Crawford 1995). It is largely the

strength of the initial signal which determines whether the evolution of events

downstream of the rnountains is viewed as lee cyclogenesis or lee reintensification. If one

considers lee cyclogenesis to be the reappearance of a masked parent cyclone, it appears

that both cyclones studied here are the manifestation of the same phenomenon, i.e.

extratropical cyclones being influenced by and interacting with topography to different

degrees. Cases that fit the description of the "classical" scenario presumably have

sufficiently weak signals that they can be completely masked by the mountain

anticyclone, and appear as "new" cyclones in the lee. Such cases are to be distinguished

from situations where a parent cyclone remains present over the ocean and a lee cyclone

develops. In these situations, the mountains act to "break up" an initially strong cyclone,

giving rise to several initially smaller and weaker lee cyclones. An example of the

second scenario is the SEP24 case which begins with a very strong depression. The

"masking" of the mountains leads to an apparent weakening of the parent cyclone which

is robust enough to never lose its SLP signature and it undergoes si reintensification in the

lee. Much of the ambiguity in the classification of lee cyclogenesis arises from the fact

that cyclonic circulations of different strengths interact with topography and are affected

to different degrees by it, giving rise to different perceptions of the phenomenon. The

ambiguity also arises due to different definitions and criteria for lee cyclogenesis.

' Because reduction of pressure to sea-level is aIways based on a fictitious lapse rate, the actual value of pressure obtained is hypothetical. However, as the same lapse rate is used everywhere, qualitative cornparisons may still be made between the two cases.

In conclusion, lee cyclogenesis occurs as a result of both topographic effects and

ordinary cyclogenetic rnechanisms. Low-level blocking of the flow is important in filling

cyclones approaching a major mountain chah. The blocking is difficult to quantify, and

future work should examine parce1 trajectories in order to trace the origins of parcels in

the lee. The continued study of other cases of lee cyclogenesis will serve to generalize

the conclusions obtained in this study.

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