UNIVERSITY OF READING

Department Of Meteorology

28 January 2004: Frontal

Snowband

Jonathan K P Shonk

A dissertation submitted in partial fulfilment of the requirement

for the degree of MSc in Weather, Climate & Modelling

11 August 2004

Abstract

In the UK, both weather fronts and convective lines are very common occurrences.

Usually, cold fronts are associated with a band of cloud and rain, although some-

times they can develop with stronger circulations embedded in them. This can lead

to the development of deeper convection, generating hail, thunder and lightning.

While convective lines are often present at fronts, they can also occur separately,

typically bringing a narrower band of heavier precipitation. Stronger convective lines

can develop into squalls, also bringing stormy weather. Such cases of strong convec-

tion, however, are generally limited to the summer months, when the atmosphere is

warmer and more susceptible to convection.

On 28 January 2004, a cold front passed southwards across the United Kingdom,

bringing a band of snow. Despite it being the middle of winter, the front also brought

hail, thunder and lightning. These features are indicative of vigorous circulation,

with deep convective updraughts and downdraughts, generating gust fronts. In fact,

on account of the presence of these features, the band was not unlike a summer

squall line.

The causes of the unseasonal characteristics of the snowband are considered. The

merging of the cold front with a separate convective line ahead of it is found to be

the reason for its extreme strength. This merge led to intensification and enhanced

frontogenesis over southern England. The band is also analysed for processes that are

known to generate strong circulations. It is found that regions of moist symmetric

instability are present, along with regions showing evidence of sublimation-driven

circulations, although the accuracy of the analysis methods used can be affected by

limitations imposed by the model.

Comparisons are also made between the snowband and similar winter snowband

events occurring in the north-eastern United States. However, the data available is

limited in detail, and the comparison can only be qualitative. Under these limita-

tions, it is hard to see any significant similarities between the UK and US snowbands,

although moist symmetric instability is seen to be an important factor in both cases.

Contents

1 Frontal Snowbands In Winter 3

2 CSI: Conditional Symmetric Instability 7

2.1 Types of instability . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7

2.2 Release of moist symmetric instability . . . . . . . . . . . . . . . . . 11

2.3 Combinations of moist gravitational and moist symmetric instabilities 14

2.4 Common misunderstandings . . . . . . . . . . . . . . . . . . . . . . . 17

2.5 Indicators of conditional symmetric instability . . . . . . . . . . . . . 18

3 Effects Of Sublimation 21

3.1 Sublimation processes . . . . . . . . . . . . . . . . . . . . . . . . . . . 21

3.2 Sublimation-induced circulation and banding . . . . . . . . . . . . . . 24

4 Data: Sources, Methods And Limitations 28

4.1 Met Office data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 28

4.2 The resolution problem . . . . . . . . . . . . . . . . . . . . . . . . . . 31

5 The Synoptic Situation 34

5.1 Reading weather station data . . . . . . . . . . . . . . . . . . . . . . 34

5.2 Surface analysis charts, satellite images and radar . . . . . . . . . . . 37

5.3 Analysis of Met Office model data . . . . . . . . . . . . . . . . . . . . 41

6 Analysis Of Mesoscale Circulations 53

6.1 Evidence of moist symmetric instability . . . . . . . . . . . . . . . . . 53

6.2 Evidence of sublimation . . . . . . . . . . . . . . . . . . . . . . . . . 57

7 Final Discussion 62

Symbols, abbreviations and acronyms 68

References and Acknowledgements 70

1 Frontal Snowbands In Winter

On the evening of 28 January 2004, Reading was struck by a frontal snowband,

bringing hail, thunder, lightning, and plenty of snow. A cold front, embedded in a

strong northerly flow, passed across the UK. It caused widespread disruption over

much of the south of England, and left a scattering of snow over England, Wales

and northern France. Figure 1 is a visible satellite picture from Dundee Satellite

Receiving Station, showing the extent of the snow cover.

Initial considerations directly after the event likened the band to a tropical squall

line. The initial signature of a typical squall line is a sharp gust front, generated by

strong downdraughts striking the ground and spreading out. The downdraughts are

caused by drag exerted on the air by falling hail. It brings cooler, mid-tropospheric

air down to the surface, giving a drop in air temperature. A band of thunder and

lightning follows this, accompanied by downpours of hail, at the point where the

squall line is at its most intense. Behind the hail is a band of persistent precipitation.

While periods of heavier precipitation generally occur in the winter months, thun-

derstorms with hail are more associated with summer. Air temperature in summer

is higher, hence there is an increased tendency for strong convection to occur. A

frontal snowband with lightning and such intense precipitation is certainly not a

common winter event. Even in comparison to summer rainbands, this winter snow-

band seemed fairly severe.

Similar snowband-type events have been seen to occur over the north-east of

the United States. Nicosia and Grumm (1999) studied three snowbands occurring

over Pennsylvania, New York and New Jersey, all occurring in the winter months.

They obtained model data from the mesoscale version of the Eta Model, with a

grid resolution of 29km, along with data from the radar network (Weather Surveil-

lance Radar-1988 Doppler). The weather systems associated with the snowbands

were analysed for evidence of frontogenesis and conditional symmetric instability

(CSI). Analysis for CSI was performed by considering fields of saturated equiva-

lent geostrophic potential vorticity. Each of the snowbands was associated with

3

Figure 1: Visible satellite picture from 10:54 UTC on 29 January, showing snow

cover over England and northern France. Regions of snow cover are indicated by

white areas of land. Data from Dundee Satellite Receiving Station.

4

an extensive low-level cyclone. In two of the three cases that they investigated,

cyclogenesis occurred during the lifetime of the snowband. The surface cyclones

all deepened significantly over the 24-hour observation periods, and the heaviest

snowfall occurred in a narrow band under the comma-shaped cloud head (Nicosia

and Grumm, 1999). Their point of interest was the banding; that is, the structure

and substructure of the snowbands. They claim that, according to Bennetts and

Hoskins (1979), frontogenesis and CSI release work together to form the banding,

and the results of their investigation support this theory. They found that these two

processes play an important part in the generation of the snowbands in the cases

they studied.

Nicosia and Grumm (1999) also mention the limitations of model data. Out of

the two different models available to them, they chose the one with the smallest

gridpoint spacing. They discuss that the structure of the snowbands may be poorly

resolved, even by this higher-resolution mesoscale model. This issue of resolution is

also brought up separately by Clough et al. (2000), who analysed mesoscale banding

in a deepening low pressure system over the North Atlantic. They found that the

circulations responsible for the generation of banding can be practically omitted

on models with larger gridpoint spacing. Clough et al. (2000) also discuss how

the effects caused by sublimation can give rise to banding structure on a front.

Experimentation into the effects of changing the parametrisations of sublimation

on the outcome of a forecast run shows that, along with frontogenesis and CSI,

sublimation is also a key factor in the generation of mesoscale banding.

There are a vast number of questions that arise out of these articles about the

origin, the generation and the processes involved in the weather event on 28 January.

For the purposes of this investigation, the following five will be considered:

◦ What was the synoptic overview during the event?

◦ Why was the band so intense with lightning?

◦ Did conditional symmetric instability play a role in the snowband?

5

◦ Was the circulation driven by sublimation effects?

◦ Were there any similarities to the snowbands occurring in the north-eastern

US?

The next two sections will contain a review of background literature to give an

overview of some of the important processes. Conditional symmetric instability

and its effects will be discussed in section 2, and the effects of sublimation will be

covered in section 3. In section 4, the sources of data will be considered, along with

a discussion of the issue of limitations regarding resolution. The synoptic situation

will be addressed in section 5, where weather maps for 28 January will be analysed

to provide answers to the first two questions. The roles of CSI and sublimation will

then be assessed in section 6, and the results of this investigation will be summarised

and compared with the US snowbands in section 7.

6

2 CSI: Conditional Symmetric Instability

The concept of stability and instability in the atmosphere is widely understood. The

specialised area of conditional symmetric instability, or CSI, is generally not so well

understood. The concept of CSI was first proposed by Bennetts and Hoskins (1979)

as an explanation for the generation of frontal rainbands. They discuss symmetric

baroclinic instability, the system used historically to describe certain atmospheric

processes, and then derive equations predicting growth by this mechanism, using

modal considerations. However, the conditions under which symmetric baroclinic

instability theory is valid limit its usefulness, and the result of this is that symmetric

baroclinic instability is a poor approximation in frontal zones (Bennetts and Hoskins,

1979). They then introduce the existence of a concept that they refer to as CSI.

Since the publication of this article, the number of references to CSI and related

topics in journal articles and literature has increased nearly every year (Schultz

and Schumacher, 1999). Over time, CSI became a fashionable topic to use when

analysing processes in frontal zones. But, as more and more researchers failed to

fully understand the concept of CSI as stated by Bennetts and Hoskins (1979), it

became frequently misused, and as more articles were written with an incomplete

understanding, the misunderstandings were spread even wider. Today, a number of

journal articles are still written with an incomplete understanding of CSI (Schultz

and Schumacher, 1999).

The correct and incorrect interpretations of CSI are best summarised by Schultz

and Schumacher (1999). In this section, CSI is introduced in a similar way to that of

Schultz and Schumacher (1999), starting off from basic instability concepts. Their

notation is followed, with a few minor amendments to completely eliminate the

potential for any ambiguity.

2.1 Types of instability

Before trying to understand conditional symmetric instability, it is worth taking a

step back and considering instability as a whole and seeing where CSI fits in. There

7

are a number of different acronyms associated with instability, which at first can

appear highly confusing when reading a journal article in which the concept of CSI

has not been clearly introduced. These acronyms will be introduced systematically

throughout this subsection.

The concept of stability and instability is best introduced by considering a parcel

of air, with certain properties. Some forcing mechanism acts to push this parcel

into an environment with different properties. If the environment is locally stable,

the force exerted on the parcel by the environment will act to return it to its initial

position. If the environment is locally unstable, however, this force will act to

accelerate the parcel away from its initial position. The important quantities when

considering instability are, for the case of dry instabilities, geostrophic absolute

momentum Mg and potential temperature θ. The quantity geostrophic absolute

momentum is given by the equation:

Mg = vg + fx , (1)

where x is a horizontal co-ordinate, vg is the component of geostrophic wind perpen-

dicular to the positive direction of the x co-ordinate and f is the Coriolis parameter.

In the case of a frontal surface, x is chosen to be the co-ordinate across the front

and y the co-ordinate along the front, hence vg is the geostrophic wind component

along the front. For an air parcel to be unstable when moving in the positive x-

direction, it must move into a region of lower Mg; so must possess more momentum

than its surrounding environment. This type of instability is called inertial instabil-

ity. The condition for inertial instability, expressed mathematically, is (Schultz and

Schumacher, 1999):

∂Mg

∂x< 0 . (2)

Substituting the expression for Mg from (1) into (2) shows that this condition is

equivalent to ξg + f < 0, that is, absolute vorticity being negative.

Gravitational instability in the atmosphere is related to the temperature of the

parcel with respect to its environment. For dry air, a gravitationally unstable envi-

8

ronment possesses a rate of change of temperature T with height that is less than

−Γd, the dry adiabatic lapse rate. Dry gravitational instability, also known as ab-

solute instability, implies the parcel is susceptible to ascent both dry adiabatically

and moist adiabatically (McIlveen, 2003). A parcel moving in a dry gravitation-

ally unstable environment will ascend vertically upwards. This process is referred

to as dry, vertical convection. In terms of potential temperature distribution, the

condition for dry gravitational instability is (Schultz and Schumacher, 1999):

∂θ

∂z< 0 . (3)

The third category of instability is symmetric instability. Under certain atmo-

spheric conditions, it is possible for a parcel to be stable to all horizontal and ver-

tical displacements, but unstable to displacements in a slantwise direction. This

instability combines the effects of gravitational instability and inertial instability.

For dry air, the condition for symmetric instability can be given by either of the

following equations:

(∂θ

∂z

)

Mg

< 0 ;

(∂Mg

∂x

)

θ

< 0 . (4)

In other words, dry symmetric instability is equivalent to inertial instability along

a surface of constant θ, or dry gravitational instability along a surface of constant

Mg. An alternative way of identifying dry symmetric instability is by considering

the gradients of the Mg surfaces and θ surfaces. In the unstable case, the θ surfaces

slope more steeply than the Mg surfaces. If this is the case, the atmosphere is stable

to both horizontal and vertical displacements. The only direction in which ascent

can occur is up the plane between the slopes of the Mg and θ surfaces, that is,

diagonally upwards. This ascent is referred to as slantwise convection, in contrast

to the vertical convection that occurs during the release of gravitational instability.

However, air in the troposphere always has some water vapour content, so dry in-

stabilities alone cannot be applied in every case. There is no moist equivalent of in-

ertial instability, because changes of momentum induced in the parcel will not cause

9

saturation or condensation. Both gravitational and symmetric instability, however,

do have moist versions. Moist instability has two further subdivisions: conditional

instability and potential instability. Conditional gravitational instability (CGI), also

referred to as just conditional instability, is the case where the environmental lapse

rate along a vertical path lies between the dry adiabatic lapse rate and the saturated

adiabatic lapse rate. This is true when the vertical gradient of saturated equivalent

potential temperature is negative, that is (Schultz and Schumacher, 1999):

∂θES

∂z< 0 . (5)

If CGI is present along a surface of constant Mg, by analogy to the dry case, con-

ditional symmetric instability (CSI) is present. For the instability to be conditional,

the air parcel must be saturated. Release of CGI, therefore, gives rise to moist, ver-

tical convection and release of CSI gives rise to moist, slantwise convection. Both

processes will form clouds. The other type of moist instability is potential instabil-

ity. Potential gravitational instability (PGI) occurs in a layer of air where there is a

negative gradient of equivalent potential temperature with height, that is (Schultz

and Schumacher, 1999):

∂θE

∂z< 0 . (6)

For potential instabilities, wet-bulb potential temperature θw can also be used in

place of θE. Potential symmetric instability (PSI) is present where there is a region

of PGI along a surface of constant Mg. The seven different types of instability are

summarised in table 1, adapted from table 1 in Schultz and Schumacher (1999).

The two categories of moist symmetric instability, PSI and CSI, are often referred

to collectively under the joint acronym MSI, and similarly, the moist gravitational

instabilities, PGI and CGI, are collectively referred to as MGI. These six acronyms

will be used throughout this investigation. The terms symmetric instability and

gravitational instability will be used to generalise for all types of each stability,

whether dry, potential or conditional. The word ‘gravitational’ will be retained in

10

Gravitational Symmetric Inertial

Dry Dry gravitational Dry symmetric Inertial

∂θ/∂z < 0 (∂θ/∂z)Mg< 0 ∂Mg/∂x < 0

Conditional Conditional gravitational Conditional symmetric

∂θES/∂z < 0 (∂θES/∂z)Mg< 0

Potential Potential gravitational Potential symmetric

∂θE/∂z < 0 (∂θE/∂z)Mg< 0

Table 1: Summary of the types of instability, adapted from Schultz and Schumacher

(1999).

the instability descriptions of PGI and CGI, to eliminate any possible chance of

ambiguity. Despite this, is it still worth bearing in mind that some journal articles

use PI and CI for potential gravitational instability and conditional gravitational

instability.

2.2 Release of moist symmetric instability

For any form of deep convection to occur, there are three requirements, according to

Doswell III (1987): instability, moisture and lift. For the instability to be released

and clouds to form, saturation is required in the area, along with ascent to the

level of free convection (LFC). The requirement of local saturation implies that

the analysis of conditional instabilities is best to give an idea of susceptibility to

convection. Potential instabilities are a property of layers as opposed to specific

regions. As convection considerations typically involve the ascent of a parcel of air

as opposed to a layer, their use is more limited, although they can still be used

to identify regions where air parcels may rise to give convection. In the case of

11

symmetric instability, the LFC is sometimes referred to as the level of free slantwise

convection. This is analogous to the level of free (gravitational) convection, but for

a slantwise trajectory along a surface of constant Mg. Processes which can cause

ascent to the LFC include forced ascent over orography, and circulations associated

with frontogenesis (Schultz and Schumacher, 1999).

The growth and release of MSI can be considered using two different idealised

models (Schultz and Schumacher, 1999). The first method is by normal modes,

which involves a small-scale perturbation that is allowed to grow. This is applied

in an unforced system where CSI is present, and uses mathematics to predict how

the perturbation will grow and evolve. The second is the release of MSI under

larger-scale circulations, which cause forced ascent, and is usually determined by

considering case studies or numerical modelling. There are two important inhibitors

to MSI release (Schultz and Schumacher, 1999): turbulence in the typically nar-

row slantwise updraught, and the resistance of the descending air in the associated

downdraught. The normal modes method produces circulations that are far too

weak to overcome these inhibiting factors. Therefore, stronger forcing is important

in a realistic model of MSI release, and such forcing occurs readily in frontal zones.

This leads to the conclusion that MSI mostly occurs in the presence of frontogenesis

(Schultz and Schumacher, 1999). In fact, both frontogenesis and slantwise convec-

tion associated with MSI release can give rise to banded precipitation (Emanuel,

1994), although their contributions are found to be inseparable within the bounds

of the Sawyer-Eliassen equation (Schultz and Schumacher, 1999).

Apart from vertical and slantwise, convection also be sub-divided into two further

types, according to Emanuel (1994): activated and statistical-equilibrium. Activated

convection involves a build-up of convective available potential energy (CAPE), or

slantwise convective available potential energy (SCAPE) on a slantwise path. This

instability is then suddenly released by some mechanism that forces ascent. Acti-

vated vertical convection is commonly identified to occur, for example, when CAPE

builds above a low-level inversion. This CAPE is eventually released by heating at

the surface. The activated slantwise convection is not so easily identified by obser-

12

vation, as this would require accurate analysis of SCAPE, which is a complicated

process. Radiosonde ascents take measurements along an approximately vertical

path, so each sounding can give a local measurement of CAPE. The spacing of

upper-air stations, however, is rarely small enough to give sufficient detail of the

upper atmosphere to resolve an entire surface of constant Mg, so the gradients and

positions of the surfaces would have to be inferred.

Statistical-equilibrium convection, in contrast, occurs when the physical processes

that act to destabilise the atmosphere increase the CAPE at approximately the

same rate as that of CAPE release. This implies that CAPE is replenished at ap-

proximately the same rate at which it is consumed, hence CAPE values remain

approximately constant. The slantwise version of statistical-equilibrium convection

is observed far more commonly than its activated counterpart, particularly at frontal

zones (Schultz and Schumacher, 1999), where constant SCAPE values can be found

to persist for several hours. Statistical-equilibrium slantwise convection can con-

tinue even after neutrality has been reached, implying that frontogenesis continues

to force ascent (Schultz and Schumacher, 1999). However, in the mid-latitudes,

statistical-equilibrium vertical convection is found to be a rare occurrence. This

leads to a conclusion about the differences in timescales of forcing and of resultant

circulations. In activated convection, the timescale of the forcing is much longer

than that of the response, hence there is time for the CAPE to build up before

the instability is released. For gravitational instabilities, this is the case. Typical

timescales of forcing are of order hours, while the response occurs on a scale of

minutes, hence vertical convection tends to be activated. For statistical-equilibrium

convection, the timescales of forcing and response must be approximately the same

for the equilibrium to occur, and with symmetric instabilities, this is the case, with

both timescales of order hours. The scale separation is more difficult (Schultz and

Schumacher, 1999), so slantwise convection tends to be statistical-equilibrium.

13

2.3 Combinations of moist gravitational and moist symmet-

ric instabilities

A further complication to the concept of instability is the possibility for the two

types of instability to co-exist, particularly around weather fronts (Schultz and

Schumacher, 1999). Such combination is often referred to as convective-symmetric

instability. From Emanuel (1994), if a completely stable and baroclinic atmosphere

is destabilised by some mechanism, symmetric instabilities will be observed first,

because the growth rate of gravitational instabilities is slower, but will eventually

be overwhelmed by gravitationally-induced convection.

For the development of rainbands, two mechanisms are proposed by Xu (1986).

The author uses θE instead of θES in his article, so is considering potential instabili-

ties as opposed to conditional instabilities. The first mechanism is upscale develop-

ment, which starts off with small-scale vertical convection. Clouds form and start

to organise into bands as the atmosphere tends towards gravitational stability and

symmetric instability is released. This system is most likely to occur outside frontal

zones, particularly on convective lines. A model of upscale development in squall

lines is presented by Jascourt et al. (1988), a schematic of which is shown in figure

2. It is suggested that gravitational and symmetric instabilities can be released al-

most simultaneously in this model. Vertical convection occurs along the front of the

squall line, reducing the MGI and transporting negative equivalent potential vortic-

ity (see section 2.5) into the trailing region of precipitation. This induces MSI release

and slantwise convection, causing enhancement of the region. Vertical updraughts

are predicted along the main squall line, with slantwise circulations developing be-

hind. The combination of the two types of instability in a storm leads to it being

longer-lived and more rain-bearing (Jascourt et al., 1988). However, this challenges

the well-developed idea that the trailing precipitating region behind squall lines is

entirely stratiform precipitation.

Downscale development involves bands created initially by slantwise circulations

in a symmetrically unstable environment. These circulations build and release latent

14

Figure 2: System of upscale development proposed by Jascourt et al. (1988). Dia-

gram initially by Seman (1991). Blue arrow indicates vertical convection; red arrow

indicates slantwise convection.

heat, eventually destabilising the atmosphere gravitationally. This induces vertical

convection, which generates the banded clouds (Xu, 1986). The process typically

occurs along fronts, because larger amounts of moisture increase the chance of gen-

erating gravitational instabilities. Circulations induced by release of MSI can cause

the buckling of surfaces of θE, which creates MGI. A physical model of this setup

is presented in Neiman’s elevator-escalator concept (Neiman et al., 1993), shown

in figure 3. At a warm front, they show regions of stronger ascent at an angle of

about 45◦ among shallower, slantwise-ascending regions at about 10◦ to the hori-

zontal. It is suggested by Neiman et al. (1993) that the steeper regions of ascent,

the so-called ‘elevator’, are gravitationally induced vertical convection, while the

shallower-sloped ascent (the ‘escalator’) is due to slantwise convection caused by

the release of symmetric instability.

Some articles claim that deep convection cannot possibly be associated with moist,

slantwise convection, although, according to Schultz and Schumacher (1999), this is

not necessarily true. The depth of a region of slantwise convection is dependent on

the size of the unstable layer, so it is possible for deep convection and wide slantwise

circulations to develop if the layer is deep enough. Schultz and Schumacher (1999)

claim that updraughts in slantwise convection approximately follow the saturated

15

Figure 3: Elevator-escalator system of downscale development, proposed by Neiman

et al. (1993). Diagram modified from Schultz and Schumacher (1999). Blue arrow

indicates vertical convection; red arrow indicates slantwise convection.

adiabats, or lines of constant θE, while downdraughts follow dry adiabats (in contrast

to Clough et al. (2000); see section 3.2), and the cooling of the descending air

enhances and drives the circulation. Strong downdraughts can be generated by

sublimating snow (Clough et al., 2000).

Whether or not these slantwise circulations can become strong enough to gen-

erate lightning is debated in literature. The generation of lightning typically re-

quires ascent speeds greater than 5 m s−1, and Schultz and Schumacher (1999) are

of the opinion that, since slantwise circulations rarely produce ascent speeds greater

than 5 m s−1, it seems that lightning cannot be generated by slantwise circulation.

However, they identify three scenarios where lightning and slantwise convection are

observed to co-exist: the trailing precipitation region of a squall line, as mentioned

in the previous subsection; in wintertime convection; and in the eyewall of mature

hurricanes. It is proposed by Williams (1991) that charge separation could still oc-

cur outside regions of MGI, although little is known about the scientific reasons for

these three counter-examples (Schultz and Schumacher, 1999). It is also proposed

that slantwise convection could be a source of positive lightning, as it is observed

to occur in all of the three scenarios mentioned above. Normally, lightning is gen-

erated by a vertical charge separation, with negative charges accumulating at the

16

bottom of a cloud. These induce a positive charge at the ground. Slantwise con-

vection could cause horizontal (or near-horizontal) charge separation, exposing the

positively-charged particles to the ground (Schultz and Schumacher, 1999).

2.4 Common misunderstandings

One of the aims of Schultz and Schumacher (1999) is to highlight all the common

misunderstandings about CSI and its related topics, and to minimise their continued

misuse. Some of their recommendations are summarised:

◦ The difference between potential symmetric instability and conditional symmet-

ric instability is commonly overlooked, according to Schultz and Schumacher

(1999), on account of the requirement of saturation for CSI. A number of authors

refer repeatedly to CSI, and identify it using quantities which give information

about PSI instead. The presence of PSI does not imply presence of CSI and,

when analysing moist convection, CSI is a better indicator. Distributions of θE

and θES are rarely identical. Indeed, conditions can occur where the atmosphere

can be conditionally symmetrically unstable, but potentially symmetrically sta-

ble. Schultz and Schumacher (1999) recommend the use of θES when analysing

conditional instabilities, and θE when analysing potential instabilities.

◦ Another common misunderstanding is the interpretation of all the terms in-

volved. Schultz and Schumacher (1999) claim that the meanings of slantwise

convection and MSI are often confused and considered synonymous. This is

incorrect: slantwise convection is not the same concept as MSI. Slantwise con-

vection is the convection arising from the release of the moist symmetric insta-

bility.

◦ The presence of all three requirements for deep, moist convection, according

to Doswell III (1987), cannot be overlooked. Schultz and Schumacher (1999)

emphasise the fact that, if any of the three requirements (instability, moisture

and lift) are missing, moist slantwise convection will not occur.

17

◦ Sometimes geostrophic absolute momentum Mg is approximated by non-

geostrophic absolute momentum M , given by M = v+fx. This approximation

is made by Clough et al. (2000), among others. It is tantamount to assuming

that the atmosphere is geostrophic, which is inconsistent with the requirements

of symmetric instability theory, particularly at weather fronts. Hence Schultz

and Schumacher (1999) recommend that the geostrophic form Mg should al-

ways be used for a full analysis of MSI, as errors introduced by using M could

be non-negligible. However, it is generally inconvenient to use the geostrophic

form of absolute momentum in practical considerations of the atmosphere. An

Mg field is calculated using pressure gradients, and can become noisy, whereas

M fields tend to be smoother and therefore more convenient to analyse.

◦ It is also worth mentioning that the potential temperature quantities θ, θE and

θES should be replaced with the background distributions without perturba-

tions, commonly referred to as θ, θE and θES. Using the hydrostatic approxima-

tion, in reality, makes little difference to the fields, and eliminates the difficulties

involved in separating the θ fields into background distribution and perturba-

tion (the so-called partitioning problem). This approximation has been used

throughout this section, and will continue to be implemented throughout this

investigation.

2.5 Indicators of conditional symmetric instability

The best-known method of identifying regions of CSI is the consideration of gradients

of Mg and θES surfaces on a cross-section through the front, as mentioned in section

2.1. However, there are limitations involved in using this method, and Schultz and

Schumacher (1999) identify three assumptions on the cross-section which must be

made for the method to be valid:

◦ the evaluation cross-section through the front must be perpendicular to the

thermal wind;

◦ the geostrophic wind must be assumed constant along the front;

18

◦ the ageostrophic wind must be small.

In other words, the reliability of this method is dependent on the orientation of the

cross-section that is chosen. An improved method would have no such dependence.

Potential vorticity is an alternative method of analysing instability, and is used by

Nicosia and Grumm (1999). In the case of dry symmetric instability, the important

quantity to consider is geostrophic potential vorticity, referred to as PVg or Pg.

Geostrophic potential vorticity is given by (Schultz and Schumacher, 1999):

Pg = gζg · ∇θ , (7)

where ζg is the three-dimensional absolute vorticity vector. Nicosia and Grumm

(1999) write the equation in an alternative form:

Pg = g

((∂Mg

∂p

∂θ

∂x

)−(∂Mg

∂x

∂θ

∂p

)). (8)

(Nicosia and Grumm (1999) use pressure p as their vertical co-ordinate.) In regions

where Pg is negative, dry symmetric instability is present, in the absence of inertial

instability and dry gravitational instability. This relationship can be extended to PSI

and CSI by considering equivalent geostrophic potential vorticity PgE and saturated

equivalent geostrophic potential vorticity PgES respectively. These are found by

substituting θE or θES in place of θ in (7) or (8). While this method is not limited

by the approximations of the previous method, it fails to discriminate between areas

of symmetric instability and gravitational instability. A region of negative PgES is

not necessarily indicative of a region of CSI; it could be CGI, inertial instability, or

even a negative region generated by some other process (Schultz and Schumacher,

1999).

Another possible method of analysing susceptibility to CSI is to consider SCAPE.

While CAPE gives the maximum kinetic energy available to a vertical updraught;

SCAPE gives maximum kinetic energy for an updraught along an Mg surface, for

both horizontal and vertical motion. However, as mentioned in the previous section,

SCAPE is not easy to compute and as a result, not widely covered in literature.

19

One of the main recommendations made by Schultz and Schumacher (1999) is to

test for gravitational and inertial instabilities first, before considering any symmetric

instability analysis. Despite this, there is no single quantity which can definitely

identify a region of CSI. It is therefore best to use as many different ways of analysing

CSI as possible and assimilate evidence of the presence of CSI, and then attempt to

identify regions where CSI release is probably occurring using this evidence.

20

3 Effects Of Sublimation

The release of conditional symmetric instability is suggested as a method by which

banding of cloud and rain forms in the atmosphere (Schultz and Schumacher, 1999),

particularly in the case of downscale development of rainbands at frontal zones (Xu,

1986). Clough et al. (2000) propose the process of sublimation as a contributing

factor to banding in weather systems. They consider the latent-heating budget of

a weather system with regards to cloud formation and precipitation processes. The

two most important factors in the heating budget are the release of latent heat on

ascent, and cooling by sublimation. While the first effect causes warming of the

atmosphere, the second causes cooling, which can locally exceed the effects of any

latent-heat-induced warming (Clough et al., 2000). The process of sublimation and

its effects on mesoscale circulations and banding are considered in this section.

3.1 Sublimation processes

Ice plays an important role in weather systems. Satellite observation shows the

presence of ice clouds in both convective lines and fronts associated with weather

systems. The lifecycles of these features are typically moderated by the creation

and dissipation of these ice clouds (Clough et al., 2000). Because of the difference in

shape, the behaviour of ice crystals in the atmosphere is far more difficult to model

than that of raindrops, and modern numerical weather prediction (NWP) models

have complicated parametrisations for the processes involved in ice crystal formation

and dispersion. However, ice crystals can have a wide variety of different shapes

and sizes depending on how they form, making it almost impossible to accurately

parametrise every detail of ice crystal behaviour. Equations used for parametrisation

of some of the properties of ice crystals are discussed by Forbes and Clark (1991).

Sublimation is defined as the transfer of molecules from the solid phase as ice

crystals directly to the gas phase, or the reverse process, where ice crystals form

directly from the gas phase (Clough et al., 2000). Both types of transition are

generalised in the atmosphere by (9), which is used by Clough et al. (2000) to

21

describe the rate of change of mass of an ice crystal:

dm

dt= C F

S

(L/KT )(L/RT − 1) +RT/eX. (9)

There are three terms in this equation: the capacitance term C, the ventilation factor

F , and the effect of the environment. This last term is a function of environmental air

temperature T , saturation vapour pressure with respect to ice e, and supersaturation

with respect to ice of the environmental air S. The other terms can be considered

as constants: L is the latent heat of sublimation, K is the thermal conductivity of

air, R is the universal gas constant and X is the diffusivity of water vapour in air.

The quantity in the environmental term of (9) that has the most significant effect

on sublimation rate is the supersaturation term S.

Supersaturation and subsaturation are related to relative humidity (McIlveen,

2003). For the case of water, saturation occurs at a relative humidity value of

(RH) = 1, or 100%. Supersaturation S is the amount by which relative humidity

exceeds saturation, given by S = (RH)−1, and typical values of supersaturation with

respect to water within clouds are between 0 and 0.01 (0% and 1% supersaturation)

(McIlveen, 2003). Subsaturation s is the deficit in relative humidity from saturation,

that is, s = 1− (RH).

For ice, the situation is different. At 0◦C, the saturation points of ice and water

are identical. But as temperature falls, the saturation point of water increases above

that of ice. At a temperature of −40◦C, a relative humidity with respect to water

of 100% corresponds to a relative humidity with respect to ice of about 147%. So,

while supersaturation with respect to water in the atmosphere rarely exceeds 1%,

supersaturation with respect to ice can, in theory, reach 47% in the correct conditions

(Clough et al., 2000). The saturation curves of water and ice are compared in figure

4.

The difference between maximum values of supersaturation and subsaturation

alone suggests that sublimation could potentially happen much faster than crystal

growth by diffusion (Clough et al., 2000). Clough and Franks (1991) mention that

the cooling effects of sublimation must reach a maximum at temperatures near

22

Figure 4: Comparison of saturation curves with respect to water and ice, adapted

from Clough et al. (2000).

0◦C. These results indicate that sublimation could indeed be a rapid process. A

calculation performed by Clough and Franks (1991) shows this to be true. They

found ice sublimation to occur much faster than evaporation of raindrops. At 60%

saturation with respect to ice, they found that the number of ice crystals in a

population decreased by a factor of e in a time-scale of order minutes.

One physical reason for the rapidity of sublimation is the difference in surface area

between raindrops and ice crystals of the same mass, which are found to differ by a

factor of between three and five (Clough and Franks, 1991). For rainfall, most of the

evaporation occurs at the smaller end of the size scale, while for ice crystals, subli-

mation rates are reasonably uniform throughout the whole size spectrum. Smaller

raindrops tend to fall much more slowly than larger ones, hence any latent heat

effects from evaporation will not occur far beyond the lower edge of the precipitat-

ing cloud. Larger ice crystals and snowflakes, however, will sublimate as they fall,

generating a deeper layer of cooling beneath the bottom of the cloud. Sublimation

also causes more intense cooling because of the longer residence time of ice crystals.

Residence time is related to the reciprocal of terminal velocity and is representative

of the time it takes a falling precipitation particle to fall through an arbitrary layer

of atmosphere. A depth-scale can also be calculated (Clough and Franks, 1991) as

23

the vertical height through which a population of the particular precipitation parti-

cle must fall to sublimate or evaporate such that the total mass of ice or water falls

by a certain percentage. They found this depth-scale to be approximately an order

of magnitude smaller for ice sublimation than for rain evaporation.

3.2 Sublimation-induced circulation and banding

As a saturated parcel of air rises, condensation will occur, giving moist adiabatic

ascent. Assuming the condensation is water, if this same parcel, filled with wa-

ter droplets, were caused to descend, the water content within it would start to

evaporate. If this descent were slow enough, it is possible that enough water could

evaporate from the parcel to make the descent also moist adiabatic. In other words,

this process is thermodynamically reversible, and could occur spontaneously (Clough

et al., 2000). However, evaporation of water is a slow process, and this would only

be relevant in very slow downdraughts.

If the particles in the parcel were ice, however, the sublimation-driven down-

draughts could potentially be much faster, and hence stronger. In the correct condi-

tions, it is feasible that circulations could be set up under this reversible, spontaneous

system (Clough and Franks, 1991). A region of ascent occurs, causing condensation

and formation of ice crystals, which fall into a lower layer and sublimate, cooling

it down and forcing it to descend. As it descends, further sublimation occurs to

keep the downdraught descending moist-adiabatically. Therefore, according to Har-

ris (1977), the process of sublimation in the descending air is a result of its tendency

to maintain saturation, despite the warming effects of the subsidence.

Harris (1977) describes how the process of ice sublimation can cause instability and

induce convection. He considers a precipitating cloud with a gravitationally neutral

layer underneath. As the precipitation starts to sublimate, the layer under the cloud

becomes cooled, with the most intense cooling occurring directly beneath the cloud.

This has an effect on the temperature distribution of the layer, and will generate

a stable region on top of an unstable region. As the cooling continues, the unsta-

ble region will eventually become superadiabatic, and convection through this layer

24

will break out. This is consistent with the description of statistical-equilibrium con-

vection, as described by Emanuel (1994), where potential energy is used to produce

ascent and condensation of ice crystals, with the process of sublimation destabilising

the atmosphere at the same time, increasing the available potential energy.

Sublimation-driven circulation will not readily occur, however, if the ascent path

is vertical. Under vertical convection, any falling ice particles will fall back down

into the ascending region. Any sublimation will therefore cool the ascent region and

act to weaken the updraught. Regions of slantwise convection are more prone to

development of sublimation-driven circulation, where a region of slantwise ascent

precipitates into the region underneath, generating a corresponding region of slant-

wise descent (Clough and Franks, 1991). Clough et al. (2000) claim the descent in

such circulations to be moist adiabatic in their article, while in a similar argument

involving sublimation, Schultz and Schumacher (1999) claim the descent to be dry

adiabatic. If the case for dry adiabatic descent were true, however, there would be

no reversibility and hence no spontaneity, so such a circulation would not develop.

These circulations are the basis of the Clough and Franks mechanism, as suggested

in Clough and Franks (1991). A schematic of the mechanism is given in figure 5.

The slantwise updraught contains air that is slightly supersaturated with respect to

ice, while the corresponding downdraught contains air that is slightly subsaturated

with respect to ice. This sets up a relative humidity gradient across the circulation.

The presence of such a gradient, along with the associated regions of slantwise ascent

and descent, can therefore be used as an indicator of sublimation-driven circulations

(Clough et al., 2000).

The generation of banding in frontal clouds is a further extension of this slantwise

circulation, along with some of the features of sublimation mentioned previously in

this section (Clough et al., 2000). A region of slantwise convection causes release

of latent heat and formation of ice precipitation, which falls into the region below,

where the ascent is weaker. The ice starts to sublimate, cooling the layer and

eventually forcing it to descend. As the layer cools, the descent becomes stronger,

with strongest descent occurring near the melting layer, because warmer air can

25

Figure 5: Schematic of the Clough and Franks mechanism, taken from Clough et al.

(2000).

hold more vapour (McIlveen, 2003). Falling precipitation near the melting layer,

combined with the strength of the downdraught, will effectively cut off the source

of the cloud, causing it to collapse and resulting in a break in the cloud. The strong

descent can then act to initiate further convection, and hence proceed to build a

second cloud band. Typical widths of cloud bands are of order 20km to 100km

(Schultz and Schumacher, 1999).

In their article, Clough et al. (2000) analyse a case study from FASTEX, the Fronts

and Atlantic Storm Track Experiment. This took place during January and February

1997 and, through a series of aircraft passes and dropsonde releases, measured cross-

sections of quantities through features of storms in the North Atlantic (Joly et al.,

1997). At least one of the observing periods featured signs of banding, that is,

the development of a separate cloud head next to a mature cloud head (Forbes

and Clark, 1991). Clough et al. (2000) compare the cross-sections generated from

intensive observation period #16 with those generated with NWP models. They

26

perform three modelling experiments, two of which use a mesoscale model with a

12km horizontal resolution and 45 vertical levels. To analyse the importance of

sublimation in the banding, they perform one integration using the sublimation

parameters set as usual, and for the second, they set it to zero, effectively switching

off all sublimation cooling in the model. The banding was only forecast to any extent

in the first experiment, while the second, with sublimation switched off, showed no

significant signs of banding. However, even the first experiment did not show the

banding to the extent that it was found to exist during the observations. The cause

of this is mainly the limitations imposed by model resolution; a problem discussed

further in the next section. According to Clough et al. (2000), these results show

the importance of sublimation in the generation of mesoscale banding.

27

4 Data: Sources, Methods And Limitations

To analyse the snowband event, data from a variety of sources and in a number of

different forms is used. A large amount of data is available on the internet, from

Met Office surface analysis charts to model data derived from NWP models. Use of

such charts is made in the initial considerations about the synoptic situation, along

with observations, particularly those made in the weather station on the Reading

University campus. At this site, daily observations are made, along with continuous

weather monitoring by a set of automated weather sensors measuring temperature,

humidity, heat fluxes, wind speed and wind direction.

4.1 Met Office data

For the analysis of CSI and sublimation, three-dimensional distributions of various

quantities throughout the whole depth of the atmosphere are required. The data

source for this will be data generated from the mesoscale version of the Met Office

Unified Model. This model provides forecasts for a domain covering the UK, Ireland

and parts of Europe (see figure 6). The domain has 146 north-to-south columns and

182 west-to-east rows of gridpoints, spaced at latitude and longitude steps of 0.11◦,

giving a grid spacing of approximately 12km. To keep the resolution as uniform as

possible over the domain, the north pole of the grid is shifted from the geographical

north pole such that the centre of the domain lies on the grid equator. Vertically,

the domain has 38 levels. The vertical co-ordinates used are sigma co-ordinates in

the lower atmosphere, which take into account surface orography, and pressure co-

ordinates higher in the atmosphere, with an increased concentration of levels in the

troposphere. The mesoscale model is run sequentially after the global model, which

covers the whole planet with gridpoints at a lower resolution, approximately 60km

in the mid-latitudes. It has 432 north-to-south columns, 325 east-to-west rows and

38 vertical levels. The Met Office runs the global model forecasts every twelve hours,

while the mesoscale model is run every six hours, using boundary conditions set by

the data in the global model. This gives data sets every six hours, containing one

28

Figure 6: Domain of the mesoscale version of the Met Office Unified Model

analysis and forecasts for the following five hours. The forecasts are made from the

analyses at 00:00, 06:00, 12:00 and 18:00 UTC. For this investigation, the data sets

will be considered every three hours. So the initial analyses will be used, along with

the corresponding three-hour forecasts. Over timescales as short as three hours, it

is reasonable to assume that any differences between forecast and actual state of the

atmosphere will be negligible.

Horizontal and vertical cross-sections of various quantities across the mesoscale

domain will be generated. Initially, the locations of the significant features of the

frontal snowband will be identified, such as the front itself, the tropopause and

29

the associated jet-stream. As the front lies approximately east-west across the UK

throughout the day, all vertical cross-sections through the front will be taken in

the Y direction of the mesoscale model grid. This gives approximately north-south

cross-sections, along the line shown in figure 6, corresponding to X index 75. In all

cross-sections, north is to the right. The quantities used to identify the features will

be discussed in the next section.

The methods chosen for analysing CSI and sublimation are similar to those de-

scribed by Schultz and Schumacher (1999) and Clough and Franks (1991) respec-

tively. For analysis of sublimation, as mentioned in the previous section, a vertical

cross-section through the front will be taken and analysed for presence of the Clough

and Franks mechanism (see figure 5).

The analysis of CSI, however, is not so straightforward, because of various limi-

tations of the Met Office data. As mentioned by Schultz and Schumacher (1999),

the relationship between gradients of geostrophic absolute momentum Mg and sat-

urated equivalent potential temperature θES in a cross-section through the front can

be used as a CSI analysis method. However, the only moist potential temperature

field available is wet-bulb potential temperature θw, which can only be applied to

identify regions of PSI (Schultz and Schumacher, 1999).

The article written by Schultz and Schumacher (1999) and discussed at length

in section 2 gives an understandable definition of conditional instabilities, but ex-

plains very little about potential instabilities, apart from identifying their incorrect

use and highlighting their possible confusion with conditional instabilities. As men-

tioned, potential instabilities are properties of a layer as opposed to a region, and

the instability is released by the lifting of the entire layer (Rogers, 1976). For the

case of a layer of dry air being lifted, there can be no change in the sign of stability,

that is, a stable layer will remain stable and an unstable layer will remain unstable,

both tending towards neutrality. If the layer is moist, however, it could saturate on

ascent and, depending on the water content variation across the layer, it is possible

that an initially stable or neutral layer could become unstable. If this is the case,

the layer is defined as potentially unstable, or convectively unstable, according to

30

Rogers (1976). If such an instability is present along a vertical path through the

layer, it is potentially gravitationally unstable; if the instability is present along a

slantwise path, it is potentially symmetrically unstable. PSI is therefore potentially

more useful in analysing slantwise convection than Schultz and Schumacher (1999)

imply, in that the lifting of this layer through saturation can still lead to cloud

formation.

A piece of program code is used to calculate CAPE by analysing vertical profiles

of the atmosphere, and SCAPE, by deriving profiles that follow momentum con-

tours. To do this, the program must first calculate the absolute momentum field. It

calculates the non-geostrophic version M , as opposed to the geostrophic momentum

Mg. Schultz and Schumacher (1999) claim that using M in place of Mg could be a

source of error. Despite this, the M field will be used for the reasons given in section

2.4.

The presence of PSI will be analysed by considering the gradients of M and θw

fields in vertical cross-sections through the front. For any regions which appear to

show the presence of PSI, it will be borne in mind that many approximations are

involved in the analysis. Fields of SCAPE will also be considered, to find regions of

the atmosphere which could be susceptible to slantwise convection, and also CAPE,

to find regions susceptible to vertical convection. The changes in these two quantities

over time could also be used to identify any downscale development (Xu, 1986).

4.2 The resolution problem

The analysis of circulations involved in snowbands is limited by one more significant

factor: the resolution of the grid used in the NWP model. The effect of grid spacing

on the prediction of mesoscale banding is considered both by Clough et al. (2000)

and by Lean and Clark (2003). Along with the experiment into the importance of

the sublimation parameters in the NWP forecast models, Clough et al. (2000) also

examined the effects of grid spacing on the ability of the model to resolve banding,

observed in the clouds by satellites. They compared results from the same mesoscale

model as mentioned in the section 3.2, with those from the Limited Area Model,

31

run by the Met Office. This covers western Europe and the North Atlantic with a

resolution of 50km and only 19 vertical levels, although has since been discontinued

as the increasing resolution of the global model superceded it. These two models were

both initialised at the same time, using data from the same observation period in

FASTEX, then the predictions for nine hours later were compared with the observed

data. While both models forecast the surface pressure distributions to a reasonable

accuracy, the replication of smaller, mesoscale features was found to vary greatly.

The sharp downdraught, found in the data obtained from the observation period,

was much weaker in the mesoscale model forecast, and practically non-existent in

the Limited Area Model forecast. In terms of banding, comparison of modelled

infra-red satellite pictures with actual pictures showed that the increased resolution

model managed to forecast the banding structure to a reasonable extent, although it

did appear weaker. The Limited Area Model failed to forecast any banding (Clough

et al., 2000).

Whether the horizontal resolution or the vertical resolution limits the develop-

ment of mesoscale banding and circulations is one question addressed by Lean and

Clark (2003). They used data from the same observation period in the FASTEX

experiment, and ran the forecast using a model with a number of different hori-

zontal resolutions, varying from 60km to 2km. Lean and Clark (2003) found that

increasing the horizontal resolution had the effect of allowing smaller-scale circula-

tions and features to develop, but with negligible effect on the larger-scale features,

in agreement with Clough et al. (2000). They also discovered that features with a

length-scale less than about five grid spaces tend to be attenuated. Investigation of

vertical resolution showed a similar effect. They concluded that, to resolve frontal

structure, a horizontal resolution of 24km and 45 vertical layers is required, and

to fully resolve slantwise frontal circulations requires a 2km grid, with 100 vertical

layers. In other words, both horizontal and vertical resolution have an effect on the

ability of the model to resolve small-scale features.

Bearing in mind the size of cloud bands, typically 20km to 100km according to

Schultz and Schumacher (1999), it is reasonable to expect that some banding will

32

be observed by the Met Office mesoscale model predictions, with grid size 12km.

However, it can be expected that features smaller than about 60km may suffer

attenuation, according to Lean and Clark (2003). The Met Office mesoscale model

has sufficient horizontal resolution to resolve the frontal surface, as this quantity

is within the bounds suggested by Lean and Clark (2003). Its vertical resolution

is outside the bounds, although not by a significant amount. A full recreation of

the detail of any slantwise circulation features cannot be expected. Even with these

limits, however, it is fair to assume that there will be some indication of larger-scale

slantwise circulations in the model data.

33

5 The Synoptic Situation

The passing of the snowband over the south of England caused enthusiastic records

to be made by both amateur and professional meteorologists. An overview of the

event from the Reading University campus is given, and then data is considered to

build up a synoptic overview of the snowband event and provide an answer to the

first two questions in section 1.

5.1 Reading weather station data

The weather during the morning and afternoon of 28 January 2004 was not particu-

larly spectacular. There had been a light shower of snow in the night, but by dawn,

the sky was clear, as was the ground. Sunshine persisted through the morning, with

cloud gathering as the day went on. A southward-moving front was forecast to pass

over the UK during the day. In the 00:00 UTC analysis chart on 28 January (figure

7), this front was situated off the north coast of Scotland, and it was moving south

at a speed of about 10 m s−1 in the cold, northerly flow. Weather forecasts were

predicting this front to bring some potentially heavy snowfall over the UK, reaching

southern England by the evening, with the risk of disruption.

Shortly after 17:00 UTC, the front reached Reading. The first sign of the weather

to come was a strong gust front, with a maximum gust recorded at 14.5 m s−1,

and a corresponding veer in the wind direction through about 90◦. This was then

followed by a drop in temperature, from 5.5◦C to 0.5◦C in about five minutes.

Then Reading experienced a passing line of thunderstorms, with lightning and hail.

Behind the line of hail was a broad band of snow, which fell for about an hour.

The air temperature continued to drop, falling below 0◦C. The snow ceased almost

as suddenly as it began, and behind the bands of cloud, the skies were clear. The

temperature continued to fall, reaching a minimum of −3.0◦C during the night,

making it the coldest night in January 2004, according to the data from the Reading

University weather station.

Figures 8 and 9 show the weather data for the day, 28 January, as measured by

34

Figure 7: Extract from the Met Office surface analysis; 00:00 UTC on 28 January.

35

Figure 8: Temperature and relative humidity data from Reading University weather

station, for 28 January.

Figure 9: Wind data from Reading University weather station, with pressure data,

for 28 January. Wind direction scale runs clockwise from north from the axis to the

top of the graph.

36

the automatic weather sensing system at the Reading University weather station,

operated by the meteorology department. The wind data shows the peak in both

average wind speed and gust speed as the gust front passes, as well as the change

in direction across the front. The pressure data, obtained from a separate weather

sensor in a nearby location, shows a dip in pressure as the front passed. The tem-

perature sensors measured the drop in temperature as mentioned above, also with

a rapid increase in humidity as the moist, cold air behind the front swept over the

sensors. A data plot from the heat flux sensors (not shown) indicates the time of

the nearest lightning strike, which acted to kill the sensors completely.

5.2 Surface analysis charts, satellite images and radar

All data analysed in this section is taken from the period between 00:00 UTC on

28 January and 00:00 UTC on 29 January. To eliminate the necessity to keep

reiterating the date, all times stated in the rest of this investigation will be with

respect to midnight on 28 January, with 00:00 UTC on 29 January being referred

to as 24:00 UTC.

The Met Office surface analysis charts in figures 7, 10 and 11 show the situation

at 00:00, 12:00 and 24:00 UTC respectively. The significant features on the charts

are marked with letters for ease of identification. The charts show a low pressure

system persisting over Scandinavia, which moves slowly around the Norwegian coast

towards northern Denmark over the course of the day. There is a strong ridge of

high pressure in the Atlantic, running from Greenland down towards Spain, which

also persists throughout the day. These features act together to maintain a strong

northerly flow over the UK and Ireland, bringing cold air south from the polar

regions. A number of fronts and convective lines are observed being formed in the

Arctic and carried southwards in this northerly flow. Some of the stronger fronts

are associated with waves in the flow, generating small low-pressure centres called

polar lows (McIlveen, 2003). Values of 500mb-geopotential height during the day

are below 5.28km, an indicator that any precipitation will probably fall as snow.

The surface analysis charts show the presence of five marked features in the

37

Figure 10: Extract from the Met Office surface analysis; 12:00 UTC on 28 January.

38

Figure 11: Extract from the Met Office surface analysis; 00:00 UTC on 29 January.

39

northerly flow. Forecasters had predicted a band of snow passing through early

in the morning of 28 January. The source of this snowfall was cold front D. How-

ever, only a small amount of snow was recorded, falling at about 00:30 UTC, as the

front was weaker than predicted. Convective features, including features B and E,

were developing off the Norwegian coast, and progressing slowly southwards over

the course of the day. The important features for the purposes of this investigation

are cold front A and convective line C. Front A is seen to progress southwards over

the country during the day, ending up over northern France by 24:00 UTC. Feature

C, however, only appears on the 12:00 UTC analysis as a convective line ahead of

the main cold front A. One possible reason for this appearance and disappearance

is that the analysis charts are being compiled by different forecasters, who may or

may not identify feature C. The first question to ask is whether this is the case.

The satellite pictures can provide a solution. Figures 12, 13 and 14 show three

satellite pictures for 28 January. The satellite data is infra-red data from Meteosat.

Analysis of the satellite pictures throughout the whole day show that, indeed, there

is a line of cloud developing off the west coast of Scotland at 00:00 UTC, which is the

origin of feature C. This line progresses southwards ahead of front A through until

12:00 UTC. By this time, it lies across the UK at a latitude of 54◦, the latitude of

Manchester (figure 13). Over the next few hours, however, front A starts to catch up

with feature C. By 15:00 UTC, the masses of cloud associated with the two features

are merging on the satellite images, and by 18:00 UTC (figure 14), there is only one

distinguishable feature. So the reason for the lack of presence of feature C on the

00:00 UTC analysis is that the feature was weaker, whereas its absence on the 24:00

UTC analysis is because it has merged with front A. The exact time and location

of the merging of the two features appears, from the satellite pictures, to be just

before 16:00 UTC at the latitude of Birmingham (about 52.5◦).

The satellite pictures, being infra-red, also indicate depth of cloud, as the grey-

scale colours show the temperature, hence the height of the cloud tops. Early in the

day, the cloud bands associated with features A and C are both fairly shallow and

grey. They remain fairly shallow throughout the day, up to about 15:00 UTC. On

40

following pictures, however, as the cloud bands merge, the clouds increase signifi-

cantly in depth, with the cloud tops becoming almost white in the satellite images.

This is indicative of intensification of front A as it merges with feature C.

An animation of Met Office Nimrod radar images from the day supports this

theory of intensification. Figures 15, 16 and 17 show some of the radar images.

Early in the day, there are scattered showers falling in the region of features B

and E, and two bands of light precipitation associated with A and C. These bands

are seen to progress south at the same pace as the cloud features on the satellite

images, and meet just before 16:00 UTC at a latitude of about 52.5◦. There is

then a severe intensification of the band, with precipitation rates measured by the

radar increasing from about 5 mm h−1 to over 20 mm h−1, as the merged front moves

forwards over southern England. Further gradual intensification is seen to occur as

the band proceeds into the English Channel and reaches northern France. These

observations already provide a reason for the severity of the snowband: the fact the

band was made up of two separate features which merged and intensified.

Sferics data, showing the times of lightning strikes across the UK, also shows

the southward progression of the features during the day. This plot is shown in

figure 18. There are some lightning strikes associated with front A striking across

Scotland and northern England before the features merge, although the number of

strikes increases dramatically over southern England after 16:00 UTC and later over

northern France.

5.3 Analysis of Met Office model data

The first field to be considered is wet-bulb potential temperature θw, which is useful

indicator of frontal surfaces. For analysis of fronts over a horizontal cross-section,

a geopotential height a few kilometres off the ground is best, as interactions with

the surface will affect the distribution. A geopotential height of 800mb is chosen.

These cross-sections reinforce the positions of the fronts as marked on the Met

Office surface analysis charts considered in the previous section. However, they do

not show the positions of the separate convective line features, so the presence and

41

Figure 12: Meteosat infra-red satellite image; 06:00 UTC on 28 January.

42

Figure 13: Meteosat infra-red satellite image; 12:00 UTC on 28 January.

43

Figure 14: Meteosat infra-red satellite image; 18:00 UTC on 28 January.

44

Figure 15: Met Office Nimrod radar image; 06:00 UTC on 28 January.

Figure 16: Met Office Nimrod radar image, 11:45 UTC on 28 January.

45

Figure 17: Met Office Nimrod radar image, 18:15 UTC on 28 January.

Figure 18: Sferics data, showing positions of lightning strikes; 28 January.

46

position of feature C during the day cannot easily be found. The steepness of the

θw gradient can be used as a measure of the intensity of the front, and the change

of gradient can give an idea of frontogenesis. In the period before the merge and

intensification, the contours of θw are seen to bunch up as front A moves southwards

over the country. In other words, slight frontogenesis is occurring during the day.

After the merge, however, the contours bunch up much more rapidly, and continue

to do so through the rest of the day. So the effect of the merging of A with C is to

cause a rapid sharpening, followed by an enhanced rate of frontogenesis.

Other features associated with front A can be identified by considering different

fields. A number of fields were considered in vertical north-south cross-section, with

the section passing throughX index 75, as shown in figure 6. This corresponds with a

cross-section lying slightly to the west of the 0◦ longitude line, passing approximately

through Reading. Quantities considered were: wet-bulb potential temperature, dry

potential vorticity, meridional and zonal wind components, and relative humidity

with respect to ice. A vertical cross-section of wet-bulb potential temperature shows

the frontal surface as in the horizontal cross-section, and the vertical sections support

the observation about the merging features leading to increased frontogenesis.

The cross-sections of dry potential vorticity can be used to give the position of

the tropopause. The tropopause is defined as the layer in the atmosphere where

the dry potential vorticity is 2 PVU (or 2× 10−6 K m2 kg−1 s−1). The tropopause is

seen to fold downwards ahead of front A throughout the day. This fold is seen to

be narrow and shallow at first, but becoming much wider and slightly deeper as the

day goes on. Further down the frontal surface, a larger-scale plot of dry potential

vorticity shows a dipole of vorticity, with a positive anomaly region developing below

an enhanced negative anomaly region. This is indicative of the presence of diabatic

heating, and this dipole is seen to develop only after the merging has taken place.

The cross-section of zonal wind (that is, wind through the section, or along the

front) is also a fair indicator of the position of front A, as there is a shift in along-front

wind across the front. Zonal wind could also be used as an indicator of location of

some of the convective features, although the uncertainty in the form of the features

47

makes this identification difficult. The most important feature to be derived from

the zonal wind section is the position of the jet stream, shown as a region at the top

of the front where along-front wind reaches a maximum. Figures 19, 20 and 21 show

frontal surface with θw contours (red), with the position of the tropopause indicated

by the 2 PVU contour (blue), and contours of along-front wind, at 30 m s−1 and

35 m s−1 indicating the jet stream position (black). Over the course of the day, the

jet stream fluctuates in strength, although has no particular correlation to the time

of merging of the features.

Relative humidity with respect to ice and meridional wind cross-sections can be

used to answer another question: whether front A is an ana-front or a kata-front.

Schematics showing the typical forms and features of ana-fronts and kata-fronts

are shown in figure 22. Both types of front are associated with a dry intrusion

behind the frontal zone, and such an intrusion is apparent in the cross-sections of

relative humidity with respect to ice. Figure 23 shows the relative humidity plot for

18:00 UTC, with the dry intrusion present around Y index 75. The main difference

between the two front types is the flow of air ahead of the front. In the ana-front,

the warm air is seen to rise, giving the possibility of deep convection. In the kata-

front, the warm air moves away from the frontal zone and subsides, giving a dry,

cloudless region with a low-level band of stratocumulus. The distribution of relative

humidity with respect to ice shows that, ahead of the frontal zone, there is no large,

upper-level dry feature which would be observed in the case of a kata-front, but

a vast, moist region in which cloud-formation is occurring. Front A is therefore

an ana-front. The distribution of meridional wind, however, does little to reinforce

this prediction. There are too many other circulations occurring, which make the

identification of features typical of an ana-front difficult.

Having considered all these vertical cross-sections through the front, it is noted

that there is little evidence of the existence of feature C ahead of front A. Considering

the satellite pictures in figures 12, 13 and 14, it is apparent that feature C is probably

being missed by the cross-sections. It develops further to the west and its strongest

region moves south-east, and by the time it has merged with front A it is probably

48

Figure 19: Vertical cross-section through the front, showing contours of θw (red),

dry PV (blue) and through-section wind (black); 06:00 UTC on 28 January.

Figure 20: Vertical cross-section through the front, showing contours of θw (red),

dry PV (blue) and through-section wind (black); 12:00 UTC on 28 January.

Figure 21: Vertical cross-section through the front, showing contours of θw (red),

dry PV (blue) and through-section wind (black); 18:00 UTC on 28 January.

49

Figure 22: Features of anabatic and katabatic fronts, taken from Thompson (1998).

Figure 23: Vertical cross-section through the front, showing relative humidity with

respect to ice; 18:00 UTC on 28 January.

50

still too far west to be spotted in the cross-sections through Reading. Plotting a

moving cross-section of wet-bulb potential temperature, however, shows that feature

C can be observed in the wet-bulb potential temperature plots. These north-south

cross-sections start at X index 50 (approximately over the western coast of Scotland)

at 06:00 UTC, moving east by ten indices (about 120km) every six hours for each

new plot, and are shown in figures 24, 25 and 26. A pair of features are seen, with

front A extending high into the troposphere and a shallower, lower-level feature

corresponding to the position of feature C early in the morning. Feature C starts

out fairly weak at the start of the day, but by 12:00 UTC has already undergone

intensification as the rainfall along its length becomes more organised. By 15:00

UTC, feature C is seen to be directly ahead of front A and by 18:00 UTC has indeed

merged with it, as seen on the satellite and radar images.

51

Figure 24: Vertical cross-section showing both features; contours as in figure 19.

06:00 UTC on 28 January.

Figure 25: Vertical cross-section showing both features; contours as in figure 20.

12:00 UTC on 28 January.

Figure 26: Vertical cross-section showing both features; contours as in figure 21.

18:00 UTC on 28 January.

52

6 Analysis Of Mesoscale Circulations

The methods used in analysing the data for evidence of CSI and sublimation have

been discussed at length in section 4.1. In this section, they are applied to the Met

Office data.

6.1 Evidence of moist symmetric instability

The fields to be considered first in the analysis of CSI presence are the CAPE and

SCAPE fields. These are generated from the initial Met Office data fields using

the piece of software mentioned in section 4.1. The process of calculating total

CAPE for an individual gridpoint involves integrating it along a vertical trajectory

through the atmosphere. The program calculates values for CAPE starting from five

different pressure levels: 1,000mb up to 800mb in 50mb steps. The highest of these

five CAPE values is the value the program returns. A similar process is used to

calculate SCAPE, although the integration process is along an M surface. Because

these surfaces do not extend vertically, the program generating the SCAPE plots

only calculates SCAPE for the central region of the domain. This condition has

been included in the program to prevent any errors being introduced by attempting

to integrate SCAPE off the edge of the domain. SCAPE data is therefore only

available over the UK, Ireland and northern France. It is also important to bear

in mind that the CAPE and SCAPE values are given by the positive CAPE and

SCAPE values above the level of free convection, minus the convective inhibitions.

So regions of negative CAPE and SCAPE may occur, where the energy involved in

the convective inhibition is larger than the CAPE and SCAPE above the level of

free convection.

Figures 27 and 28 show a pair of CAPE and SCAPE plots, both for 12:00 UTC.

Throughout the morning, there was a large band of SCAPE stretching across the

country at the approximate position of front A and feature C. The exact position of

the SCAPE regions shows that most of the SCAPE is associated with the convective

feature C, while front A lies on the northern edge of the region, with only small

53

Figure 27: Distribution of SCAPE; 12:00 UTC on 28 January.

Figure 28: Distribution of CAPE; 12:00 UTC on 28 January.

54

amounts of SCAPE associated with it. There is also a region of SCAPE present in

the large area of cloud over Ireland. Both features, however, have virtually no CAPE

associated with them at all. Scattered regions of CAPE exist off the east coast of

Scotland, associated with convective feature E and surrounding vertical convection.

Comparing the CAPE and SCAPE plots with the surface analysis chart for 12:00

UTC (figure 10), it is notable that feature B has practically no CAPE or SCAPE

associated with it at all. Front D, over central France by 12:00 UTC, however, is

seen to possess both SCAPE and CAPE, although much smaller amounts than the

developing features A and C.

This suggests a strong susceptibility to slantwise convection, brought on by release

of symmetric instability along the length of both front A and feature C, although

no susceptibility to vertical convection because of the lack of CAPE. It also suggests

that there is some disorganised vertical convection occurring over the North Sea

around feature E. This is also evident on both the satellite pictures, where patchy,

tall cloud is inferred in the area, and on the radar, where small regions of precipita-

tion are noted. It is also noted that SCAPE values throughout the morning remain

relatively constant, suggesting the occurrence of statistical-equilibrium convection.

The merging of the features, however, causes a significant increase in SCAPE over

south-west England, with values reaching over 800 J kg−1 by 18:00 UTC, contrasting

with values of order 400 J kg−1 earlier in the day. Figures 29 and 30 show SCAPE

and CAPE at this time. This increase is not so marked elsewhere along the front,

although the merging does also seem to give rise to some CAPE generation. This

suggests that some process in the slantwise convection has initiated the build-up

of CAPE, which could potentially lead to vertical convection if the CAPE can be

released. This is reminiscent of the process of downscale development, as described

in section 2.3. After this, however, both the CAPE and SCAPE values on the

front decrease rapidly, apart from a narrow band between the French coast and the

Channel Islands. So the release of any gravitational instability soon causes the rapid

consumption of both CAPE and SCAPE as the snowband further intensifies.

It must be borne in mind, however, that presence of positive CAPE and SCAPE

55

Figure 29: Distribution of SCAPE; 18:00 UTC on 28 January.

Figure 30: Distribution of CAPE; 18:00 UTC on 28 January.

56

does not necessarily imply the release of instabilities. For the analysis of instability

release, the relationship between M and θw is used to find evidence of PSI. To

reiterate, regions where the contours of θw become steeper than contours of M are

indicative of regions of PSI, within the approximations involved (see section 4.1).

Buckling and overturning of the θw surfaces is also indicative of PGI, if the gradient

of θw with height becomes negative. Vertical cross-sections are taken, as before, on

a north-south direction with respect to the grid, along X index 75, which passes

approximately through Reading.

Figures 31 and 32 show cross-sections at 09:00 UTC and 21:00 UTC, comparing

contours of θw and M (the program refers to the meridional component of momen-

tum as N). These cross-sections are centred on the immediate area surrounding

the front. At 09:00 UTC, a clear buckling of M surfaces is apparent for a patch

between Y indices 100 and 105, and for heights between 700mb and 500mb, with θw

surfaces sloping more steeply. This is suggestive of a region of potential symmetric

instability. There is also a small region where the gradient of M becomes negative in

the horizontal direction, indicating a region of inertial instability. Behind the front,

the atmosphere is seen to possess PGI, denoted by the negative gradient of θw with

height.

The region of PSI decreases in size gradually throughout the day, becoming al-

most symmetrically neutral by 15:00 UTC, and remaining so until the end of the

day, although some buckling of M surfaces could indicate small regions of PSI and

inertial instability developing again later after the merging takes place. At this

time, however, the layer of PGI in the lower atmosphere is seen to develop, creating

a shallow potential instability by 21:00 UTC (figure 32).

6.2 Evidence of sublimation

Analysis of the presence of sublimation-driven circulation will be achieved, as men-

tioned previously, by analysing the distributions for evidence of the Clough and

Franks mechanism (figure 5). Cross-sections of relative humidity with respect to

ice and meridional wind (across-front wind) will be taken through the same cross-

57

Figure 31: Vertical cross-section through the front, showing meridional momentum

component and θw contours; 09:00 UTC on 28 January. Momentum has units m s−1.

Figure 32: Vertical cross-section through the front, showing meridional momentum

component and θw contours; 21:00 UTC on 28 January.

58

section as used in the analysis of PSI. Less approximations need to be made in this

case to identify sublimation, and its identification should be easier.

Consideration of the cross-sections shows that sublimation occurs to some extent

throughout the whole day. Possibly the clearest example of the Clough and Franks

mechanism in progress is evident at 09:00 UTC, shown in figure 33. The wind

speeds on the section are positive for winds from left to right, that is, southerly, and

negative for right to left, or northerly. The mean flow in the section is a northerly

flow, as the values on all the contours are negative. The main region of interest

is to the right of Y index 110, and above the 600mb level. Here there is a band

of supersaturated air along the line of the front with relative humidity values with

respect to ice above 100%. While the wind in this band is still northerly, it is less

northerly than its surroundings. Therefore, with respect to the mean flow, the wind

component here is southerly, implying a region of ascent along the slantwise path.

Below the supersaturated region, there is a region of much lower humidity, and a

more northerly component. In other words, there is slantwise descent present in

this region. It is also worth mentioning the larger patch of supersaturation existing

to the left of the band. While the Clough and Franks mechanism is not apparent

around this region, the centre of this region contains the most southerly part of the

flow. This region is where the ice is forming where slantwise convection is at its

strongest. It is then swept upwards along the frontal surface, precipitating out and

causing the resulting slantwise downdraught by sublimation cooling.

For the rest of the day, presence of the Clough and Franks mechanism is not so ap-

parent, with no evidence of circulations quite as strong as in the 09:00 UTC section.

However, there is still a band of increased relative humidity with respect to ice along

a seemingly slantwise path, so it is conceivable that some sublimation processes are

still going on. It is possible that, if the circulation were to get more intense during

the day when the two features merge, that the updraughts and downdraughts of the

resulting circulation would be narrower, and hence the model might not be able to

fully resolve them.

Comparison of the sublimation and PSI analysis sections at 09:00 UTC shows

59

also that the patch of supersaturation, indicative of large amounts of ice formation

and southerly slantwise ascent, coincide well with the region of PSI identified in the

previous subsection. For ice to actually form in such conditions, it is fair to assume

that there has to be some sort of conditional instability present. It is also worth

comparing the cross-section used to identify sublimation with the corresponding

satellite picture to identify the features. The infra-red satellite picture for 09:00

UTC is shown in figure 34. Regions of supersaturation with respect to ice higher up

in the atmosphere are typically indicative of the presence of ice clouds, detectable

on the satellite picture. The line of the cross-section runs just to the left of the 0◦

longitude line on the satellite picture. The region of PSI is situated at the front of

the cloud band associated with front A, while the higher region associated with the

Clough and Franks mechanism forms the cloud shield directly behind it. The other

mass of supersaturation to the top left of figure (left of Y index 95, above 550mb)

corresponds to the edge of the cloud visible as the end of feature C. The regions

of increased humidity trailing downwards from the PSI region in the middle of the

figure coincides with the precipitation region as identified in the radar picture.

60

Figure 33: Vertical cross-section through the front, showing relative humidity with

respect to ice and across-front wind contours; 09:00 UTC on 28 January.

Figure 34: Meteosat infra-red satellite image; 09:00 UTC on 28 January.

61

7 Final Discussion

This investigation has so far attempted to answer four of the questions set out in

the first section. The frontal snowband that passed over the UK on 28 January

2004 was an unexpectedly severe weather event. It displayed some of the properties

associated with squall lines, such as thunder, lightning and hail, and contained

some strong circulations. The presence of the strong gust front and the large drop

in temperature across it supports this. The distribution of relative humidity and

across-front wind indicate that the cold front at the heart of the snowband was an

ana-front, which was still developing in strength as it progressed southwards during

the morning, with frontogenesis occurring throughout. The front was carried in a

strong northerly flow, bringing cold polar air across the UK and was associated with

a small-scale polar low. In the morning, the front brought snowfall across areas of

the UK, with scattered lightning strikes.

The cause of the intensity of the snowband was the merging of the cold front with

the separate convective line directly in front of it. The two features were seen to

drift southwards at different speeds, catching up with each other just before 16:00

UTC. On all the charts, this merge was seen to lead to a sudden sharpening of

the front, followed by an increased rate of frontogenesis, along with a noticeable

deepening in the bands of cloud and a severe intensification of precipitation rates.

The reason for the presence of lightning on the front, however, is not so certain.

As mentioned by Schultz and Schumacher (1999), slantwise circulations are not

normally strong enough to separate the charges to generate lightning and form

hail, particularly lightning and hail as strong as that which occurred within the

snowband. Unless the slantwise circulations were incredibly strong, it seems more

likely that there was some sort of vertical circulation taking place, caused by the

release of some gravitational instability. For this to be the case, the gravitational

instability will have been released by the slantwise circulations in a form of downscale

development. The only doubt about this being the case is the small amount of CAPE

present in the band after the merging takes place. However, the implications of

62

downscale development, according to Xu (1986) in section 2.3, are that the slantwise

circulations destabilise the atmosphere gravitationally, so that vertical convection

can break out. In other words, the destabilisation could be occurring higher up in

the atmosphere. The CAPE program, however, only calculates CAPE lifted from

the lowest few levels of the atmosphere (section 6.1), so it is possible that this CAPE

could be missed by the program.

The vast quantities of SCAPE present along the snowband throughout the day,

combined with the presence of a patch of potential symmetric instability earlier

in the day, are strongly suggestive of the presence of slantwise convection. The

PSI region is seen to decrease in size throughout the morning and reach potential

neutrality even before the merging takes place. While there is no direct evidence

of the release of CSI, it is fair to assume that there is probably some CSI release

occurring in the band during the day, which persists even after potential neutrality

has been reached. It is also noted that the SCAPE during the morning remains

approximately constant, which suggests that the early slantwise convection is in

statistical equilibrium. Potential gravitational instability is seen to exist in the

lower layers of the atmosphere throughout the day, and is still present after the

merge, although it is difficult to determine whether this shallow instability itself

could give rise to deep vertical convection strong enough to generate lightning.

Sublimation-driven circulation is especially noticeable earlier in the day, with

plenty of ice crystals forming in the region of PSI. This effect is seen to decrease

during the day, although it is possible that the strengthening of the front is narrowing

the ascent and descent paths of the circulation, such that their presence is being

missed by the NWP model, as a result of the limiting factor of resolution. Another

question arising out of sublimation is whether or not there is any banding structure.

It is fair to assume under the model limitations mentioned earlier that substructure

of the bands will probably not appear, as the horizontal grid spacing is 12km. One

of the more obvious questions is about the nature of feature C and front A. Their

spacing is a distance not atypical of cloud bands, such as those investigated in

observation period #16 of FASTEX, and there is a band of clearer sky in between.

63

Consideration of the circulations involved in feature C, however, shows no strong

signs of the Clough and Franks mechanism. It is possible, however, that feature

C was formed as a separate cloud band associated with front A when the system

was strengthening in the North Atlantic, but then went on to develop as a separate

convective line, which eventually merged with front A.

The final question set out in section 1 was the comparison of the snowband with

similar cases in the north-eastern US, as investigated by Nicosia and Grumm (1999).

Because of the small amount of data that they present, a full quantitative comparison

is near impossible, although it is possible to qualitatively compare some of the

effects they describe. The synoptic overviews of all three of the US snowbands are

completely different from the UK snowband, as they are all associated with a larger

scale, deep low pressure system as opposed to a single front in a northerly flow.

Despite this, the qualitative description of the snowbands as given by Nicosia and

Grumm (1999) seem similar to the UK case. They describe a region of persistent

snowfall with a strong snowband near the edge, with much larger volumes of snow

falling from it. None of the cases, however, contain thunder and lightning.

A typical plot presented by Nicosia and Grumm (1999) is a cross-section through

the front, giving information about saturated equivalent potential vorticity, sat-

urated equivalent potential temperature and frontogenesis function. Frontogenesis

function can be defined many ways, and is indicative of the strength of frontogenesis,

and shows whether the front is strengthening (positive frontogenesis) or weakening

(negative frontogenesis). As none of these quantities have been analysed in the in-

vestigation into the UK snowband, the drawing of comparisons is not easy. The

US snowbands are associated with deep regions of negative saturated equivalent

potential vorticity, which Nicosia and Grumm (1999) identify as being indicative

almost completely of CSI, with the release of CGI in places, although only on small

scales. Regions of positive frontogenesis are centred on the patches of CSI. While

frontogenesis is observed in the UK cases, and CSI is postulated on evidence of the

existence of PSI, it is hard to determine whether or not the UK snowband is of a

similar structure.

64

The deep region of CSI is strongly suggestive of slantwise circulations being the

only significant circulation existing in the US snowbands, and the result of this

being the formation of a large amount of snow, but no lightning or hail. The fact

that the UK snowband did contain lightning and hail generation is an indicator

that there is some sort of vertical circulation going on. It is worth bearing in mind,

however, that winter storms are one of the three counter-examples in which slantwise

convection and lightning are seen to co-exist (section 2.3), although little is known

about why this the case (Schultz and Schumacher, 1999). Despite this, considering

all the differences in nature and origin of the bands, it seems most likely at this

stage that the processes occurring within the US bands, associated with deepening

low pressure systems and heavy snow, are different from those occurring within the

UK snowband, which was associated with lightning and a small-scale polar low.

CSI (and probably sublimation), however, undoubtedly played some role in both.

A more in-depth investigation into the US snowbands would undoubtedly provide

an answer one way or the other as to whether the same processes are occurring

within them as within the UK snowband. Analysis of a full set of NWP model

data, as opposed to a set of qualitative discussions of the bands, would allow a more

quantifiable solution to be reached.

Weather forecasts issued in the media during the day on 28 January certainly

all mentioned a high probability of snowfall, but no mention of lightning and hail,

so it seems that even the mesoscale version of the Met Office Unified Model was

underestimating the results of the merging of features A and C. Further investiga-

tions would undoubtedly start by re-running the model with observed data to see

how well the model predicted the intensity of the snowband. Forecasts derived from

different initialisation times before the merging at 16:00 UTC could be compared

to find out how early the model forecast the merging, and whether or not lightning

and hail were predicted at any stage.

The quality with which the model forecast the processes occurring during the

merging is the key factor. The main cause of underestimation of the strength of the

snowband after the merging is most likely to be the issue of resolution. Even before

65

the merging of the two features, the evidence of the sublimation-driven circulations

mentioned in the previous section is seen to disappear, when other evidence suggests

that the process continues to occur, but with narrower ascent and descent regions.

Running some of these forecasts for the snowband on a model with a higher reso-

lution could determine if some of these smaller-scale features are actually present.

Ideally, to get full coverage of as many of the small-scale features as possible, a model

with the grid limits as set out by Lean and Clark (2003) is best used. They suggest

a model with a 2km horizontal grid-spacing and 100 vertical levels is required to

resolve all the features of a frontal zone. However, such a high resolution may be

impractical, in terms of calculation times.

An alternative method of identifying whether the sublimation-driven circulation

is occurring in the front is to recreate the experiment of Clough et al. (2000), and

set the sublimation parametrisations in the model to zero, and see whether or not

the front develops in a similar way. Such a change would also answer the question

mentioned earlier in this section, about whether A and C are two separate cloud

bands separated by sublimation effects. If they were, the removal of the sublimation

parametrisation would cause them to develop as a single mass of cloud. Indeed, a

vast number of further modelling experiments could be carried out into the effects of

sublimation here. There are a large number of different parametrisations involved

with ice crystals (Forbes and Clark, 1991), and the snowband could potentially

be used to perform further tests on the sensitivity and effects of modifying these

parametrisations.

Other potential extensions for this investigation include the comparison of the

UK snowband with similar cases that occurred in the UK, for example, comparison

with a summer rainband to find out if similar processes take place in its generation.

It seems a fair suggestion that sublimation may not play such an important role in

a rainband as opposed to a snowband, and such an investigation could test whether

or not this is the case. This could alternatively be performed using modelling. Two

forecast runs could be performed on the snowband; one with the actual data from

January, and one which has been modified such that the fields of quantities in the

66

atmosphere resemble those typical of a summer rainband, in other words, consider

what would have happened to the snowband if it had occurred in a northerly flow

in the middle of summer.

67

Symbols, abbreviations and acronyms

CAPE Convective available potential energy

CGI Conditional gravitational instability

CSI Conditional symmetric instability

FASTEX Fronts and Atlantic Storm Track Experiment

LFC Level of free convection

MGI Moist gravitational instability

MSI Moist symmetric instability

NWP Numerical weather prediction

PVU Potential vorticity unit

RH Relative humidity

SCAPE Slantwise convective available potential energy

C Capacitance term

e Saturation vapour pressure

F Ventilation factor

f Coriolis parameter

g Acceleration under gravity

K Thermal conductivity of air

L Latent heat of vaporisation

M Non-geostrophic absolute momentum

Mg Geostrophic absolute momentum

m Mass

Pg Geostrophic potential vorticity

PgE Equivalent geostrophic potential vorticity

PgES Saturated equivalent geostrophic potential vorticity

p Pressure

R Universal gas constant

68

S Supersaturation

s Subsaturation

T Temperature

t Time

v Along-front wind velocity

vg Along-front geostrophic wind velocity

X Diffusivity of water vapour in air

x Horizontal, across-front co-ordinate

y Horizontal, along-front co-ordinate

z Vertical co-ordinate

Γd Dry adiabatic lapse rate

ζg Geostrophic relative vorticity

θ Potential temperature

θE Equivalent potential temperature

θES Saturated equivalent potential temperature

θw Wet bulb potential temperature

ξg Geostrophic absolute vorticity

69

REFERENCES

Bennetts, D. A. and Hoskins, B. J. (1979). Conditional symmetric instability –

a possible explanation for frontal rainbands. Quarterly Journal Of The Royal

Meteorological Society, 105, 945 – 962.

Clough, S. A. and Franks, S. A. A. (1991). The evaporation of frontal and other

stratiform precipitation. Quarterly Journal Of The Royal Meteorological Society,

117, 1057 – 1080.

Clough, S. A., Lean, H. W., Roberts, N. M., and Forbes, R. M. (2000). Dynamical

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Acknowledgements

Thanks to everyone who provided any information about the snowband, be it

useful meteorological data or just random excited exclamations, like: ‘Snow! Snow!’

Thanks to Alec for the air pressure data, and to all my housemates, Yvonne, Sheetal

and Rob, for helping me maintain sanity throughout the writing of this dissertation.

Without their help, my brain could have become conditionally symmetrically un-

stable. Worse still, it could have frozen, and sublimation-induced circulations could

have developed within it. Thanks also to the designers of LATEX– they saved me

from the indignity of being asked stupid questions by the paperclip in Microsoft

Word. And finally, big thanks to my supervisor, Dr Sue Gray.

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