Combination of Mesoscale and Synoptic Mechanisms for Triggering an Isolated
Thunderstorm: Observational Case Study of CSIP IOP 1
CYRIL MORCRETTE,* HUMPHREY LEAN,
KEITH BROWNING,* JOHN NICOL,* NIGEL ROBERTS,

PETER CLARK,
ANDREW RUSSELL,# AND ALAN BLYTH@
*Department of Meteorology, University of Reading, Reading, United Kingdom

JCMM, Met Office, University of Reading, Reading, United Kingdom
#School of Earth, Atmospheric, and Environmental Sciences, University of Manchester, Manchester, United Kingdom
@School of Earth and Environment, University of Leeds, Leeds, United Kingdom
(Manuscript received 27 October 2006, in final form 12 January 2007)
ABSTRACT
An isolated thunderstorm formed in the southern United Kingdom on 15 June 2005 and moved through
the area where a large number of observational instruments were deployed as part of the Convective Storm
Initiation Project. Earlier, a convergence line had formed downstream of Devon in the southwest of the
United Kingdom in a southwesterly airflow, along which a series of light showers had formed. The depth
of these showers was limited by a capping inversion, or lid, at around 2.5 km. The deep thunderstorm
convection developed from one of these showers when the convection broke through the lid and ascended
up to the next inversion, associated with a tropopause fold at around 6 km. A series of clear-air reflectivity
RHIs are used to map the height of the capping inversion and its lifting resulting from the ascent along the
convergence line. The origins of the lid are tracked back to some descent from the midtroposphere along
dry adiabats. The strength of the lid was weaker along a northwest-to-southeast-oriented region located
behind an overrunning upper cold front. The transition from shallow to deep convection occurred where
this region with a weaker lid intersected the region with a raised lid, oriented southwest to northeast,
downstream of Devon. A very high resolution forecast model that is being developed by the Met Office
predicted the isolated thunderstorm successfully. This success depended on the accurate representation of
the following two scales: the synoptic-scale and the surface-forced mesoscale convergence line. The inter-
action between these scales localized the convection sufficiently in space and time for the initiation and
subsequent development to be highly predictable despite the relatively poor representation in the model of
processes at the cloud scale.
1. Introduction
Advance knowledge of the location of precipitation
is the main factor limiting the accuracy of forecasts of
river flow (Collier 1996). Flash flooding, which can
cause severe disruption and damage to property, is of-
ten due to heavy or localized convective precipitation,
for example, the flooding in Boscastle, United King-
dom, on 16 August 2004 (Burt 2005; Golding et al.
2005). Consequently, it is important to be able to fore-
cast the timing and location of convective precipitation.
This can be achieved only if the mechanisms leading to
the initiation of convection at a particular time and
place are accurately represented in the forecast model.
Further improvements to model performance will re-
quire a better understanding of these mechanisms. Up
until now, only a small amount of research has been
undertaken in the United Kingdom to observe the at-
mosphere during the initiation of precipitating convec-
tion (Bennett et al. 2006).
The Convective Storm Initiation Project (CSIP) was
organized to obtain detailed observations of the state of
the atmosphere prior to, and during, the early stages of
the development of precipitating convection. The field
campaign associated with CSIP was carried out in June,
July, and August 2005, following a pilot campaign in
July 2004. Browning et al. (2006) summarize the 18
intensive observation periods (IOPs) and show the lo-
cations of the numerous observational instruments de-
ployed in southern United Kingdom during the cam-
Corresponding author address: Dr. C. J. Morcrette, Met Office,
FitzRoy Road, Exeter EX1 3PB, United Kingdom.
E-mail: cyril.morcrette@metoffice.gov.uk
3728 M O N T H L Y W E A T H E R R E V I E W VOLUME 135
DOI: 10.1175/2007MWR2067.1
© 2007 American Meteorological Society
MWR3475

paign. By studying these IOPs we shall improve our
understanding of the mechanisms that are important in
the initiation of convection in an extratropical maritime
environment, such as the United Kingdom. A compari-
son with simulations from a very high resolution nu-
merical weather prediction (NWP) model will allow the
successes and failures of the model to be assessed. This
assessment, together with the improved understanding,
will then lead to informed suggestions for model im-
provement. It is hoped that once the abilities of the
NWP model to forecast the initiation and development
of precipitating convection have been proven, it will be
possible to use it as part of a new nowcasting system
(Golding 2005).
This paper presents an observational case study of an
isolated thunderstorm that formed over southern
United Kingdom during IOP 1. Key observing systems
used in this study were serial rawindondes, released
from several of the special CSIP sites, and the 1275-
MHz and 3-GHz advanced research radars both of
which are mounted on the 25-m dish at Chilbolton (see
locations plotted later in Fig. 2). The Chilbolton radars
alternately made low-elevation plan-position indicator
(PPI) scans and series of range–height indicator (RHI)
scans. Because of the large dish and high power (560
kW), the 3-GHz radar is very sensitive and was used to
map not only areas of very light precipitation but also
clear-air (and cloudy) features, such as the edges of
thermals and developing cumulus convection and the
stable layer capping them. Observational data from the
different instruments are synthesized in this paper to
produce a detailed picture of the atmosphere during the
development of the convection in IOP 1. It shows how
upper-level and boundary layer mechanisms can com-
bine to increase potential instability and raise a capping
inversion locally to a sufficient degree to allow deep
convection to form in a single preferred location. Al-
though we make use of NWP model data to back up our
interpretation of the observations, no modeling experi-
ments will be described here. Modeling experiments to
investigate the sensitivity of the convective develop-
ment to various factors will be presented in subsequent
papers.
2. General situation
The operational Met Office surface analysis for 1200
UTC 15 June 2005 (CSIP IOP 1), is shown in Fig. 1.
There is a low pressure center located to the northwest
of the United Kingdom, which led to a southwesterly
flow over southern United Kingdom where the various
FIG. 1. Operational analysis for 1200 UTC on 15 Jun 2005.
NOVEMBER 2007 M O R C R E T T E E T A L . 3729

observational instruments were deployed. An occluded
front has been analyzed over eastern England and an
upper-level trough is located over the south of Ireland.
In the southwesterly flow between the occluded front
and ensuing trough, a series of light showers, with rain
rates up to 8– 16 mm h1, formed in a line emanating
from Devon, in the southwest of the United Kingdom.
One of these light showers developed into a cluster of
deep convective showers within range of the advanced
meteorological research radars at Chilbolton (51.14° N,
1.44° W; Goddard et al. 1994). The radar observations
collected on this day have allowed us to gain insight
into the evolution of the atmosphere that accompanied
the transition of the convection from small shallow
showers, with modest rain rates, to a small cluster of
deep convective cells, producing heavy rain and light-
ning.
Figure 2 shows the high-resolution visible image from
Meteorological Satellite (Meteosat)-8 at 1200 UTC
(Schmetz et al. 2002). A line of clouds approximately 8
km wide and around 150 km long, with a well-defined
southeastern edge, has formed downwind of Devon.
Figure 3 shows an analysis of the 10-m wind vectors and
convergence at 1000 UTC. These were produced as
part of the Nimrod system by combining surface obser-
vations and NWP model data (Golding 1998). The line
of clouds highlighted in Fig. 2 has formed in a region
where convergent southwesterly flow has persisted for
several hours. Figure 4 shows the precipitation associ-
ated with the clouds that formed along this line. The
PPI of the radar reflectivity factor (hereafter simply
“ reflectivity” ) for the sector to the west of Chilbolton at
1152 UTC shows that most of the showers were pro-
ducing only light rain with reflectivities of little more
than 30 dBZ. Radar vertical cross sections indicate that
these showers were around 3.5 km deep, which is below
the 0° C level, and hence precipitation growth did not
involve the ice phase. Because the differential reflec-
tivity was close to zero, there probably were a relatively
large number of small drizzle drops in these showers, so
that, using an appropriate Z–R relationship (Collier
1996), this implies rainfall rates of up to 5 mm h1. In
comparison, the cluster of showers to the north of Chil-
bolton, which constitutes the storm of interest, had re-
flectivities up to 52 dBZ, and the radar vertical cross
sections indicate that these showers were over 7 km
deep at this time. A comparison with some thermody-
namic soundings suggests that this cluster of showers
was deep enough for ice microphysics to have played a
part in the growth of precipitation particles. Use of the
Marshall and Palmer (1948) relationship would imply a
rainfall rate of around 60 mm h1 for this storm. Data
from the radar network shows the evolution of this
storm’s maximum rain rate (Fig. 5). The increase in rain
rate around 1100 UTC was closely associated with the
transition from shallow to deep convection. A resur-
FIG. 2. High-resolution visible imagery at 1200 UTC 15 Jun 2005 from Meteosat-8. The sites from which
radiosondes were launched are indicated by white crosses: (from left to right) Camborne, Bath, Larkhill,
Reading, Preston Farm, and Herstmonceux. The black circle, which has a radius of 100 km and is
centered on Chilbolton, indicates the maximum range of the 3-GHz radar at Chilbolton. The radial lines
and digits indicate the azimuths along which RHIs were performed. The boundaries of the county of
Devon are indicated by dotted white lines. The dashed white ellipse highlights the cloud line mentioned
in the text.
3730 M O N T H L Y W E A T H E R R E V I E W VOLUME 135

gence of activity occurred at 1200 and 1240 UTC. The
Met Office Arrival Time Difference (ATD) system
(Lee 1990) detected lightning at 1240 UTC.
There are two themes in this study— one is the for-
mation of the shallow convection and light rain show-
ers, while the second is the transition to deep convec-
tion to give an isolated thunderstorm with higher rain
rates. A good forecast for such a case requires correctly
representing the two different convective regimes as
well as the change from one to the other. This is a
challenge for a NWP model but the very high resolution
model used was successful in this case.
3. Use of model diagnostics to support the
observational analysis
In this study we make use of data from four different
NWP models. Two of these, the European Centre for
Medium-Range Weather Forecasts (ECMWF) model
and the global version of the Met Office Unified Model,
are used for large-scale analysis. Two others are used
for more detailed analysis. One of these is the opera-
tional 12-km version of the Met Office Unified Model
(Cullen 1993). This model is nonhydrostatic (Davies et
al. 2005) and has a convection scheme based on Greg-
ory and Rowntree (1990). Finally, we also use data
from a 1.5-km grid length version of the same model,
which, due to its high resolution, is being run without a
convection scheme (Clark and Lean 2006). The 1.5-km
model is the better model for representing the devel-
opment and life cycle of the convective showers be-
cause, unlike the 12-km model with parameterized con-
vection, actual convective circulations can develop, giv-
ing the model some memory of the dynamics of the
storms.
Figure 6a shows an early stage in the development
of the isolated storm in IOP 1 as seen by the opera-
tional radar network at 1100 UTC. The operational ra-
dars have a lower resolution and a higher minimum
detectable threshold than the Chilbolton research ra-
dar, and so some of the weaker showers are not de-
picted in Fig. 6a. Figures 6b,c show the rain-rate fore-
casts from the 1.5- and 12-km versions of the Met Office
Unified Model with a lead time of 11 h. In all three
figures there is a frontal rainband present over the
east of the country, with some showers behind it to
the west. The better representation of the organized
line of showers is given by the 1.5-km model, al-
though this model tends to produce too many showers
along the convergence line when compared to reality
[as shown for other cases by Clark and Lean (2006)].
The convective precipitation in the 12-km model shows
no organization on the meso– beta scale (Orlanski
1975).
FIG. 3. Analysis of 10-m wind speed (m s1) and convergence (s1) at 1000 UTC 15 Jun 2005. The
area depicted is the same as in Fig. 2.
NOVEMBER 2007 M O R C R E T T E E T A L . 3731
Fig 3 live 4/C

Rain-rate data from the weather radar network were
used to track the small cluster of heavy showers from
when it first appeared on the radar network to when it
left the area covered by the 3-GHz radar at Chilbolton
(Fig. 7a). Figure 7b, shows a similar analysis for the
track of the single cluster of heavy showers that formed
in the 1.5-km model. To ensure a fair comparison, the
rain rate from the 1.5-km model was averaged onto a
5-km grid equivalent to that used by the radar network.
Precipitation from the cluster of showers that went on
to become the storm was first detected by the radar
network at 0915 UTC; the model started producing rain
15 min earlier. The location of first precipitation was
about 15 km in error; this error had components of 8
km along and 12 km across the direction of storm mo-
tion, corresponding, respectively, to the errors in the
timing of synoptic features or the detailed initiation,
and the location of the convergence line. The speed of
motion of the storm in the model (47  1 km h1)
agreed with that from the radar network observations
(48  1 km h1) within the limits of measurement error.
Additionally, the storm trajectories deviated by less
than the diameter of the storm after 4 h. The similarities
in the time and location of first precipitation, and mo-
tion of the observed and modeled storms, give us con-
fidence that the 1.5-km model was performing well on
this day. We shall therefore make use of detailed diag-
nostics from the 1.5-km model to support our interpre-
tation of the observations.
FIG. 4. PPI of reflectivity (dBZ ) at 0.5° elevation, at 1152 UTC
from the 3-GHz Chilbolton radar showing the line of rain show-
ers. The radial lines indicate the orientation of the RHI scans
shown in Figs. 10 and 13. The regions of speckled reflectivity less
than around 0 dBZ are due to returns from insects. This scan has
been filtered for ground clutter by removing reflectivities with a
linear depolarization ratio greater than 10 dB.
FIG. 5. Time series from the radar network of the maximum rain rate for the shower of
interest averaged onto a 2-km grid (thin line) or 5-km grid (bold line).
3732 M O N T H L Y W E A T H E R R E V I E W VOLUME 135
Fig 4 live 4/C

4. The lid and its lifting by low-level convergence
downwind of Devon
Figure 8 shows a thermodynamic profile in the vicin-
ity of the line of convective showers. The tephigram for
the radiosonde launched from Bath at 1104 UTC, about
20 km west of the main shower at this time, is shown in
Fig. 8a. Figure 8b shows the profile from the 1.5-km
model for the same time and place. Features of interest
from the profiles include (i) an inversion at 500 hPa in
the observations and a corresponding stable layer be-
ginning at 440 hPa in the model (actually a tropopause
fold, as discussed later); (ii) a dry stratospheric intru-
sion above 500 hPa; and (iii) a dry intrusion between
750 and 550 hPa, the base of which constituted a lid at
750 hPa (725 hPa in the model). The path of a near-
surface parcel is depicted by a gray line in Fig. 8a. Using
a value of w  14° C for this parcel, the convective
inhibition (CIN) of the lid between 750 and 600 hPa
was 85 J kg1. At this stage the height of the showers
was being capped by this lid, limiting their depth to less
than 3 km.
The humidity and temperature gradients found at the
base of the lid correspond to changes in the atmo-
spheric index of refraction (Fig. 9a). Gradients in re-
fractive index can be detected by radar (e.g., Gossard
and Strauch 1983). Figure 9b shows the square of the
vertical gradient of the refractive index and suggests
that clear-air echoes would have been expected from an
altitude of around 2.4 km (Gossard et al. 1998). The
FIG. 7. The track of the developing thunderstorm is shown by the outline of the precipitating region
(a) as measured by the radar network from 0915 UTC when the first echo appears to 1300 UTC when
it moves out of range of the Chilbolton radar and (b) as forecast by the 1.5-km model. The precipi-
tating region is shaded alternately light and dark gray for clarity. Both storms have a speed of around
48 km h1.
FIG. 6. Rain rate at 1100 UTC 15 Jun 2005: (a) as observed by the weather radar network and as forecast by (b) the 1.5-km and (c)
12-km versions of the Unified Model. The circle in (a) shows the 100-km range of the advanced meteorological radar at Chilbolton. The
arrows in (a) and (b) indicate the thunderstorm of interest.
NOVEMBER 2007 M O R C R E T T E E T A L . 3733
Fig 6 live 4/C

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3734 M O N T H L Y W E A T H E R R E V I E W VOLUME 135

profile of static stability N2  (g/0)/z, shown in Fig.
9c, confirms that the clear-air echo would occur at a
similar height as the stable region forming the lid.
Rayleigh scattering from particulates and Bragg scat-
tering resulting from fluctuations in the refractive index
have different wavelength dependence (Gossard and
Strauch 1983). As a result, when considering a dual-
wavelength radar system calibrated for Rayleigh scat-
tering, there will be a difference in the reflectivities
resulting from Bragg scatter at the two wavelengths.
The two radars at Chilbolton operate at 1275 MHz (L
band) and 3.077 GHz (S band). Consequently, reflec-
tivity resulting from Bragg scatter will be 14.0 dBZ
higher at L than at S band. Figures 10a,b show RHIs at
1115 UTC, along 331° , at S and L band. Both scans
show a thin layer of increased reflectivity between the
10- and 18-km range and at around 2-km altitude
(within the black boxes). The convective showers,
reaching a height of 7 km, will be discussed in section 5.
Figure 11 shows a scatterplot of reflectivity at L and S
band for the RHIs at 1115 UTC. Because of the differ-
ent sampling volumes of the two radars (0.75°  66 m
and 0.28°  300 m, respectively), the data have been
linearly averaged to a coarser resolution (0.75°  300
m) for the comparison. This difference in beamwidth
explains why the layer echo seen beyond the 20-km
range by the S-band radar, is no longer seen by the
L-band radar. In Fig. 11, the data collected within the
boxed regions in Figs. 10a,b, corresponding to the layer
of increased reflectivity, are plotted as circles, while all
other data are plotted using crosses. The solid line in
Fig. 11 shows the one-to-one relationship along which
data resulting from Rayleigh scatter are expected to lie.
The dotted line, 14.0 dBZ lower, indicates the expected
location of data resulting from Bragg scatter.
The presence of crosses below the one-to-one line
suggests that Bragg scatter is contributing to the reflec-
tivity in regions other than the layer echo. The other
sources of Bragg scatter include the 1-km-deep ther-
mals between the 20- and 35-km range, and the edges of
the clouds and showers. As a group, the circles from the
boxed region around the layer echo have a reflectivity
16  3 dBZ higher at L band than at S band. This
suggests that the layer of increased reflectivity is due to
Bragg scatter from the inversion seen in the thermody-
namic sounding, and not Rayleigh scatter from particu-
lates such as insects, cloud droplets, or raindrops, which
would result in the same reflectivity at both wave-
lengths.
An RHI of 3-GHz radar reflectivity, along azimuth
FIG. 9. Diagnostics calculated for the radiosonde launched from Bath at 1100 UTC: (a) refractive index, (b)
square of the vertical gradient of the refractive index (103), and (c) static stability (104).
NOVEMBER 2007 M O R C R E T T E E T A L . 3735

283° at 1037 UTC, is shown in Fig. 12a. The RHI shows
some of the showers that later developed into deep
convection located at a range of between 48 and 60 km.
Figure 12a also shows a layer of clear-air echo resulting
from changes in the atmospheric refractive index at the
height of the lid (strictly the echo is mainly from the
base of the lid, but elsewhere for simplicity we ignore
this distinction). Figure 12b shows a vertical cross sec-
tion from the 1.5-km model at 1000 UTC, taken per-
pendicular to the region of raised lid. The shading in-
dicates the dry static stability, and the white contours
represent the cloud liquid water content. The maximum
in dry static stability was due to the large potential
temperature gradient at the height of the lid and cor-
responds to the peak seen at 2.4 km in Fig. 9c. The
height of the clear-air echo at 1037 UTC at the 70-km
range (approximately over Bath) in Fig. 12a is 2.0 km,
which is lower than the height of 2.4 km expected from
the radiosonde launched at 1104 UTC (Fig. 9). How-
ever, the scan along azimuth 310° at 1117 UTC, close to
where the sonde would have drifted to by the time it got
to the height of the lid, shows some clear-air echo at a
height of 2.4 km, in good agreement with the radio-
sonde observations.
Convectively rising boundary layer air had to over-
come this lid in order to achieve the transition to deep
convection with a level of neutral buoyancy located at
the height of the tropopause fold at 500 hPa. Thermo-
dynamic profiles from radiosondes launched from
Larkhill at 1000 and 1200 UTC and from Bath at 1100
and 1200 UTC (e.g., Fig. 8) suggest that if the lid were
to be lifted dry adiabatically by between 500 and 800 m,
the lid would effectively be removed and parcels from
the surface would be able to convect freely up to the
tropopause fold. A key mechanism for such lifting was
the convergence line downwind of Devon. Figure 12a
shows the lid at 2.2 km at the 30-km range, being lifted
to 3.0 km at the 55-km range, at this particular time and
azimuth, and then falling again to 1.9 km at the 70-km
range. Although the convection seen in this RHI had
FIG. 10. (a) RHI of reflectivity (dBZ ) along 331° at 1115 UTC from the 3-GHz, S-band, Chilbolton
radar showing convection having broken through the lid and reaching a height of 7 km. This scan is
perpendicular to the cloud street shown in Fig. 2 and transects the small patch of heavy showers shown
to the north of the radar in Fig. 4. The box between 10- and 18-km range and 1.9- and 2.5-km altitude
identifies the layer of increased reflectivity, which a dual-wavelength comparison showed was due to
Bragg scatter. This scan has been filtered for ground clutter by removing reflectivities with a linear
depolarization ratio greater than 10 dB. (b) RHI of reflectivity (dBZ ) along 331° at 1115 UTC from
the 1275-MHz, L-band, Chilbolton radar. The box between 10- and 18-km range and 1.9- and 2.5-km
altitude identifies the layer of increased reflectivity used in the dual-wavelength comparison. In this
image, a signal-to-noise threshold of 2.2 dBZ has been used to filter residual noise after spectral
processing.
3736 M O N T H L Y W E A T H E R R E V I E W VOLUME 135
Fig 10 live 4/C

not broken through the lid, the height over which the
lid was being lifted suggests that its breaking was im-
minent. This is confirmed by the adjacent RHI, just to
the north (not shown), where convective plumes had
penetrated through the lid.
Convergence lines form downwind of the coast re-
sulting from any of the following three mechanisms: the
different frictional forces over the land and sea (Hunt
et al. 2004), the different surface temperatures of the
land and sea (Simpson 1997), or lifting resulting from
orography (Bader et al. 1995), or a combination of all
three. In southwesterly wind conditions during the sum-
mer, convergence lines tend to form downwind of
Devon and Cornwall in the southwest of the United
Kingdom (Starr 1997). These convergence lines often
lead to lines of cloud and convective showers, and we
believe this to have been the case here, because a com-
parison between Figs. 2, 3, 4, and 6 shows that the lo-
cations of the showers seen by the network and Chil-
bolton radars correspond to the locations of the lines of
clouds emanating from Devon in the satellite imagery
and the location of the convergence line seen in the
10-m wind analysis. The origins of the convergence line
that formed on this day, and a study of the relative
importance of the different mechanisms that contrib-
uted to its formation are described by H. Lean (2007,
unpublished manuscript).
The RHI in Fig. 13 was obtained approximately per-
pendicularly to the convergence line at the time of Figs.
2 and 4. Because the convergence was organized along
a line, the main contribution to this convergence is
likely to have been from confluence at right angles
thereto. The Doppler data in Fig. 13a show confluence
of around 8 m s1 over a distance of 2.5 km near the sur-
face and diffluence of around 16 m s1 over a distance
of 2.5 km at around 3-km altitude. The low-level con-
vergence downwind of Devon was lifting the lid. Close
to the axis of the line, between the 50- and 60-km range
in the RHI, the convection extends up to the height of
the lid, and the lid itself was raised some more as a
result of the ongoing convection. The convection, which
in this section led to rain rates around 10 mm h1,
corresponds to the area of rain showers approximately
40 km southwest of the main storm cluster shown to
FIG. 11. A dual-wavelength comparison of the reflectivity measured by the 3-GHz (S-band) and 1275-MHz
(L-band) radars at 1115 UTC. The reflectivities collected within the boxes shown in Figs. 10a,b, corresponding to
the thin echo layers, are plotted using circles, all other data are plotted using crosses.
NOVEMBER 2007 M O R C R E T T E E T A L . 3737

the north of Chilbolton in Fig. 4. Figure 13b shows that
to the side of this convergence line there is some
boundary layer convection, which is revealed by the
lumpy layer beneath the lid echo (seen most clearly
between the 28- and 48-km range). Although there are
occasional plumes that do reach the height of the lid
(e.g., at the 26- and 33-km range), most of the boundary
layer convection does not.
FIG. 12. (a) RHI of reflectivity (dBZ ) from the 3-GHz radar at Chilbolton at 1037 UTC along an
azimuth of 283° , as indicated by the white dashed line in Fig. 14a. This scan has been partially filtered
for ground clutter by removing reflectivities with a linear depolarization ratio greater than 10 dB. (b)
Vertical cross section of static stability (s2) showing the lifting of the lid in the region of the convergence
line. The white contours, at 0.2 g kg1 intervals, show liquid water content to highlight the location of
the convection.
3738 M O N T H L Y W E A T H E R R E V I E W VOLUME 135

Having presented three examples of 3-GHz radar re-
flectivity RHIs showing thin clear-air-layer echoes
(Figs. 10a, 12a, and 13b), we shall now make use of a
series of RHI scans showing similar clear-air echoes to
map the height of the lid. During CSIP in 2005, the
scanning sequence for the 25-m dish consisted of 20
RHIs at approximately 18° azimuth apart, followed by
a low-level 360° PPI (the whole sequence taking around
43 min). The RHI scans were performed up to an el-
evation of 15° and the elevation scan rate was a func-
tion of elevation with a value of 0.1° s1 at 15° eleva-
tion, decreasing smoothly to 0.02° s1 below 4° eleva-
tion to enable better detection of finescale vertical
structures at long range. The height of the lid, subjec-
tively defined as the height of the maximum in the
clear-air reflectivity layer, was recorded manually at
5-km-range intervals. Because of the time taken to
complete a sequence of 20 RHI scans, the locations of
the data points have been advected by the system ve-
locity (48 km h1 toward the northeast) to create a
distribution at a standardized time. The objective
analysis technique described by Barnes (1964) was then
used to draw smooth contours through the irregularly
spaced data points. This method of producing a hori-
zontal map of the height of a stable layer by looking at
a series of radar RHI scans showing clear-air reflectiv-
ity has already been described by Morcrette et al.
(2006), but they used data from the 1275-MHz (L band)
radar also mounted on the 25-m antenna at Chilbolton.
Although the longer wavelength at L band makes it
preferable for detecting clear-air signals, the 0.75°
beamwidth makes the detection of weak signals at long
range difficult. In contrast, the 0.28° beamwidth of the
S-band radar, when combined with the reduced scan
rate at low elevation used in 2005, improved the ability
of that radar to resolve shallow echoes at low altitude
and enabled useful clear-air signals to be detected over
a wider area.
Figure 14a shows the resulting map of the height of
the lid at 1100 UTC, derived from 20 RHIs collected
between 1028 and 1103 UTC. The digits indicate the
height of the lid in hundreds of meters, while the shad-
ing is the result of the Barnes analysis. The white re-
gions are too distant from observations to allow a reli-
able interpolated value to be calculated. When carrying
out the Barnes analysis to produce the smooth field
from discrete data in Fig. 14a, each data point was set to
exert an influence over a surrounding area defined as a
Gaussian function with a user-specified width. After
experimenting with full-widths-at-half-maximum
(FWHM) of 10, 20, and 30 km, it was found that using
30 km resulted in fields that were too smooth. Using a
FIG. 13. RHI scans from the 3-GHz radar at Chilbolton, along azimuth 310° at 1200 UTC: (a) Doppler velocity
(m s1, positive away from the radar) and (b) reflectivity (dBZ ). This scan is perpendicular to the cloud street
shown in Fig. 2 and transects the light showers shown in Fig. 4. These scans have been filtered for ground clutter
by removing reflectivities with a linear depolarization ratio greater than 10 dB.
NOVEMBER 2007 M O R C R E T T E E T A L . 3739
Fig 13 live 4/C

FWHM of 10 or 20 km confirmed the presence of an
elongated ridge of higher lid height oriented in a south-
west-to-northeast direction. However, using a FWHM
of 10 km caused the radial distribution of data to be-
come apparent in the final fields, while using a FWHM
of 20 km created gradients across the ridge that were
too smooth when compared to the radar RHIs, showing
the rising and falling of the lid (e.g., Fig. 12a). The
FWHM was consequently set to 10 km across the di-
rection of the convergence line and 20 km along it with
a sinusoidal variation in between.
Figure 14a shows that the height of the lid varied
from less than 2 km to the south of the radar to over 2.6
km to the northwest where the convergence line has
lifted the lid over a band downwind of Devon. The
variation in the height of the lid seen in the analysis in
Fig. 14a corresponds to the rise and fall in the height of
the clear-air echo seen, for example, along azimuth 283°
in Fig. 12a, the time-corrected location of which is in-
dicated by the white line in Fig. 14a. The white circles
in Fig. 14a indicate the locations in one of the RHIs
(along azimuth 293° at 1036 UTC) where narrow
convective plumes are observed to have penetrated the
stable layer. The locations of the convective plumes
have also been advected to where they would be at
1100 UTC. These plumes are seen in a region where
the lid has been lifted and weakened. The lifting of
the capping inversion in a band downstream of
Devon is also seen in the 1.5-km model (Fig. 14b). In
the case of the model, the height of the lid has been
diagnosed by searching for the height of the maximum
in dry static stability occurring between the surface and
500 hPa. The region in Fig. 14a where convective
plumes were penetrating the stable layer is close to the
region in the model where the height of the lid was
greatest.
In summary, it has been shown that the light showers
that formed on this day did so along a convergence line
emanating from Devon. The effect of low-level conver-
gence was to raise the height of the lid capping convec-
tion over a band a few tens of kilometers wide. Showers
with mainly modest rain rates developed in this band.
However, the small area of showers seen in Fig. 14a to
have been more vigorous than the others, continued to
intensify and this is described next.
5. Penetration of convection through the lid
Figure 10 shows an RHI at 1115 UTC along 331°
through the same cluster of showers as depicted 38 min
earlier in Fig. 12a, and it shows that the convection had
reached a depth of over 7 km by this time. Figure 15
shows a vertical cross section of dry static stability from
FIG. 14. Plan view of the height (m) of the lid capping the
boundary layer convection at 1100 UTC: (a) derived from 20 RHI
between 1028 and 1103 UTC, and (b) from the 1.5-km model. In
(a) the observed system velocity has been used to correct the
location of the data, and the broken white line indicates the lo-
cation of the vertical cross section along azimuth 283° shown in
Fig. 12a. The numbers (102 m) indicate the observations. The
white dots indicate the locations of convective elements penetrat-
ing the lid. In (b) the large white area is where the boundary layer
is shallow because of the presence of the front.
3740 M O N T H L Y W E A T H E R R E V I E W VOLUME 135

the 1.5-km model shortly before (Fig. 15a) and after
(Fig. 15b) the convection penetrated through the lid. At
1100 UTC (Fig. 15a), the lid had been raised, but it was
still capping the convection and only liquid water was
present in the cloud. By 1200 UTC (Fig. 15b), the con-
vection had broken through the lid and risen to the
height of the tropopause fold, allowing ice to form in
the cloud. Meanwhile the other, lighter showers up-
stream (to the southwest) were still being capped by the
lid and had not gone through the transition from shallow
to deep convection. The reasons for one cluster of show-
ers deepening and not the others will now be discussed.
FIG. 15. Vertical cross sections across the convergence line and through the model’s main storm
showing static stability (shading, s2) and liquid and ice water contents (white and black contours
respectively at intervals of 0.2 g kg1) at (a) 1100 and (b) 1200 UTC, before and after the convection has
penetrated the lid.
NOVEMBER 2007 M O R C R E T T E E T A L . 3741

6. Influence of the upper-tropospheric PV
anomaly
a. Nature of the PV anomaly
The hypothesis to be tested is that the penetration of
the lid and extension of the convection from 3 to 7 km
occurred within the convergence line, but only where it
was modulated by forcing from an upper-level feature.
This occurred in the form of a trough oriented approxi-
mately northwest to southeast in association with a
traveling upper-level potential vorticity (PV) maxi-
mum. Figure 16 shows water vapor imagery for 1200
UTC. There is a dark zone representing relatively dry
air located over the United Kingdom at this time. Dark
zones in the water vapor imagery are known to be as-
sociated with PV maxima, such as occur due to a tropo-
pause depression (Browning 1997). The storm of inter-
est appears as a small bright patch of increased mois-
ture less than 10 km across embedded within the dark
zone. For comparison, Fig. 17a shows the height of the
PV  2 surface at 1100 UTC from the 12-km version of
the Unified Model.
The height of the PV  2 surface can be taken as
being representative of the dynamical tropopause
(Hoskins et al. 1985). The dark zone associated with
drier air in the water vapor imagery corresponds well
FIG. 16. Meteosat Rapid Scan Water Vapor imagery at 1200 UTC.
3742 M O N T H L Y W E A T H E R R E V I E W VOLUME 135

with the PV maximum in the model. This maximum can
be tracked back over 24 h to a PV anomaly (cutoff low)
that separated from a larger PV filament extending
over the Atlantic south from the Labrador Sea. This
large filament preexisted for an additional day over
northeastern Canada.
Figure 17b is a vertical section oriented as shown in
Fig. 17a. It shows the upper-level PV maximum, which
was responsible for the trough and the main upper-level
forcing, to be associated with a distinct tropopause fold.
Beneath it is a shallow layer of high PV corresponding
to the lid that was tending to suppress the convection.
The PV of this layer was between 1.5 and 2 PV units
and it probably originated within, or close to another,
different tropopause-fold region. This has been exam-
ined in more detail by Russell et al. (2007).
Because the lid is the key feature for controlling the
transition from shallow to deep convection, it is of in-
terest to examine how it was formed. Some back tra-
jectories, calculated from ECMWF model data ar-
chived on a 1.25° grid, are used to clarify the origins of
the dry tropospheric streamer between 750 and 550
hPa, the bottom edge of which was acting as the lid
(Fig. 8). A series of back trajectories initialized from
Larkhill (51.2° N, 1.8° W) at 1200 UTC and starting at
50-hPa intervals between 900 and 200 hPa, were calcu-
lated from the ECMWF winds. Figure 18 shows a se-
lection of these trajectories and illustrates the origins of
the air making up different layers in the thermody-
namic profile shown in Fig. 8. These layers (and the
pressure at 1200 UTC 15 June) are (i) the boundary
layer (900 hPa), (ii) the air just below the lid (800 hPa),
(iii) the air just above the lid (700 hPa), and (iv) the
middle troposphere (550 hPa).
Figure 18 shows that the air making up the boundary
layer ascended gradually and advected from over the
Atlantic Ocean, where it will have remained relatively
moist. The midtropospheric air initially descended
slightly, from 500 to 550 hPa. The air that ended up just
above the lid descended from above 500 to 700 hPa,
while the air just below the lid stayed at around 800
hPa. The potential temperature along these latter two
trajectories was fairly constant over the preceding 3
days, 29  1° C and 22  1° C, respectively, suggesting
dry-adiabatic flow. It is the differential vertical motion
along these isentropes that led to the formation of the
lid at around 750 hPa. The formation of the lid by adia-
batic descent from midlevels explains the relative dry-
ness of the capping inversion seen in the radiosonde
profile. The synoptic-scale ascent of the air on either
side of the lid by around 30 hPa over the most recent
8-h period contributed to a weakening of the lid and
this, together with the mesoscale lifting along the axis of
the boundary layer convergence line (downwind of
Devon), enabled the boundary layer convection even-
tually to break through.
b. Structure of the troposphere accompanying the
PV anomaly
To quantify the synoptic-scale changes in tropo-
spheric structure associated with the passage of the PV
anomaly, we have used data from six radiosondes
FIG. 17. (a) Geopotential height (m) of the PV  2 PVU surface, searching downward, at 1100 UTC from the 12-km model. (b)
Vertical cross section of PV at 1100 UTC, taken from SW to NE along the line marked in (a), showing a fold in PV. The lower feature
of high PV (between 700– 750 hPa) corresponds to the temperature inversion (lid) shown in Fig. 8.
NOVEMBER 2007 M O R C R E T T E E T A L . 3743

launched from Larkhill. The location of Larkhill is in-
dicated in Fig. 2 and the radiosondes were launched at
0540, 0815, 1000, 1200, 1400, and 1600 UTC. Figure 19a
shows a time– pressure cross section of the radiosonde
data. Because the system was moving eastward, the
time axis has been reversed to create a pseudo– west–
east cross section. The solid white line near the surface
shows the height of the lifting condensation level
(LCL). The color shading depicts wet-bulb potential
temperature w below the LCL and saturated wet-bulb
potential temperature s above it. The saturated wet-
bulb potential temperature is defined as the wet-bulb
FIG. 18. Three-day back trajectories showing the origin of the air column that was over
Larkhill at 1200 UTC 15 Jun 2005. The trajectories are calculated from the ECMWF model
(archived on a 1.25° grid). (a) Map of location of back trajectories and (b) evolution of
pressure with time. Line style and symbols distinguishes trajectories in terms of their
pressure at 1200 UTC 15 Jun: 900 hPa (solid line and crosses), 800 hPa (dash– dot and
circles), 700 hPa (dashed and triangles), and 550 hPa (dotted and squares). Symbols are
plotted every 24 h in (a), and representative values of potential temperature along each
trajectory are shown in (b).
3744 M O N T H L Y W E A T H E R R E V I E W VOLUME 135

FIG. 19. (a) Time– pressure cross section of s and w constructed from Larkhill radiosonde data. The solid
white line indicates the lifting condensation level (LCL); w is plotted below the LCL and s above it (° C).
The magenta line is the 10% relative humidity contour. The time axis is reversed to ease comparison with
Fig. 20b. (b) Similar to (a) but showing the value of w or s at all places within the cross section taken away
from the surface value of w (° C). The vertical black lines at the top of each panel indicate the launch times
of the six radiosondes used to construct the figures.
NOVEMBER 2007 M O R C R E T T E E T A L . 3745
Fig 19 live 4/C

potential temperature that the air would have if it were
saturated. When comparing parcel and environmental
potential temperatures to assess buoyancy, it is neces-
sary to compare the parcel’s w to the environment’s s.
The magenta line is the 10% relative humidity contour.
In the upper troposphere Fig. 19a shows high values
of s, advected down from the stratosphere, and con-
stituting the tropopause fold, the base of which is seen
descending as low as 600 hPa between 1400 and 1600
UTC. A positive PV anomaly, resulting from a tropo-
pause depression associated with a tropopause fold,
normally has a cold anomaly located below it (Bishop
and Thorpe 1994). Such a cold anomaly is seen in Fig.
19a from 0800 to 1300 UTC between 400 and 650 hPa.
Between 0930 and 1100 UTC, the surface w (13.9° C at
1000 UTC) is greater than the s of the cold anomaly
(13.6° C at 550 hPa), indicating the potential for deep
convection. This convection could not develop at
Larkhill however, due to the presence of two lids. The
lower lid, centered at 850 hPa, was associated with the
frontal zone. The main lid, as referred to elsewhere in
this paper, was centered at 700 hPa. With its low values
of relative humidity, air in this lid had descended from
the midtroposphere, as discussed in the previous sec-
tion. It becomes the primary barrier to deep convection
after 1000 UTC. The observations shown in Fig. 19a
indicate that the s of this lid exceeded the w every-
where in the boundary layer at all times after 1000
UTC. This is as expected because the deep convection
did not pass directly over Larkhill; nevertheless, the
very presence of the deep convection discussed earlier
proves that the lid was penetrated over a limited area
nearby.
The importance of the lid at 700 hPa after 1000 UTC
is revealed even more clearly in Fig. 19b. The informa-
tion in Fig. 19b corresponds to that in Fig. 19a, but we
have plotted instead the surface value of w minus w or
s aloft to give a more direct indication of parcel buoy-
ancy with respect to a parcel originating from the sur-
face. Positive values correspond to regions of positive
area on a tephigram (and vice versa). Notice the sig-
nificant convective inhibition at 700 hPa after 1000
UTC in contrast with the dwindling convective inhibi-
tion associated with the lid at 850 hPa.
We have made use of data from the 1.5-km model to
investigate the key features of the synoptic situation
leading to the spatial variability of the instability and
the lids. It was found that the structure conformed to
the split-front model of Browning and Monk (1982) in
which an upper cold front (UCF) overruns a tongue of
unoccluded warm-sector air at low levels (i.e., the op-
erational analysis in Fig. 1, depicting an occlusion, is not
strictly correct; but, it is common for forecasters, when
concerned with the synoptic scales, to analyze split
fronts as occlusions for simplicity). Our analysis of a
split front is consistent with the radiosonde data from
Larkhill (Fig. 19). The s of the lid was a maximum
along and behind the axis of the warm tongue, and s in
the 600– 500-hPa layer was a minimum above the warm
tongue, due to the presence of the cold anomaly asso-
ciated with the upper-level positive PV anomaly. The
combination of high surface w and low s at midlevels
led to a maximum in potential instability behind the
UCF while the high s of the lid compared to the w at
the surface is evidence of the capping inversion. Figure
20 summarizes the location of these key features at
1000 UTC just prior to the deepening of the convection.
A small surface warm pool (w  14° C) is located
within a narrow south– north low-level warm tongue
(Fig. 20a). This warm pool was around 40 km wide. It
corresponds to the shallow moist zone (SMZ) of the
split-front model, ahead of a surface cold front (not
explicitly marked in Fig. 20a), but behind the UCF. The
convergence line, which formed near the coast, can be
seen in Fig. 20a to have extended inland into the region
of the surface warm pool.
The vertical cross section in Fig. 20b shows the cooler
air (w 13° C) advecting over the surface warm tongue
(i.e., SMZ). The overrunning cold air was important
because it produced the region of potential instability.
Layers of high static stability, or lids, are indicated in
Fig. 20b by dashed lines, the thickness of which repre-
sents the strength of the lid. One of these lids was as-
sociated with the frontal zone. The other lid marked the
base of the previously descended dry intrusion. Cru-
cially, the lid was relatively weak over the surface warm
pool, because the strongest descent had occurred far-
ther upstream and the air was beginning to ascend
again (see Fig. 18b, showing the reascent of the air near
the lid over the most recent 8-h period). This meant
that a further small elevation of the lid, resulting from
a local forcing mechanism such as the convergence line
discussed earlier, was enough to allow convection to
penetrate it. This lifting was concentrated along a line
oriented southwest to northeast whereas the region of a
weaker lid resulting from synoptic-scale ascent behind
the upper cold front was oriented northwest to south-
east. The thunderstorm developed at the intersection of
these two lines.
7. Summary and conclusions
A series of shallow light rain showers formed in a
region of the deepened boundary layer due to a topo-
graphically forced convergence line forming downwind
3746 M O N T H L Y W E A T H E R R E V I E W VOLUME 135

of Devon in the southwest of the United Kingdom. The
convection from the surface was capped by a lid formed
by a streamer of previously subsided relatively warmer
air. One of these showers developed into a small cluster
of deep convective showers with heavy rain rates and
some thunder. The shower that broke through the lid
did so in an area where the lid was weaker. There was
no other development of deep convection along the
convergence line because farther northeast, the frontal
zone prevented convective development while farther
southwest the lid was stronger. The instability experi-
enced by the deep convection was augmented by the
presence of a cold anomaly at midlevels associated with
an upper-level PV anomaly. The location where the lid
broke occurred where two key features intersected: the
region of a weaker lid resulting from the dynamically
forced ascent in the vicinity of, and just behind, the
upper cold front (oriented approximately northwest to
southeast); and the region of lifting resulting from the
topographically forced convergence line (oriented ap-
proximately southwest to northeast).
This study has shown that the 1.5-km model is ca-
pable of representing the evolution of the mesoscale
structure of the atmosphere associated with the transi-
tion from shallow to deep convection. The model suc-
cessfully forecast the synoptic conditions, including the
presence of an upper-level positive PV anomaly and
associated midlevel cold anomaly. The presence of a
tongue of warm w air near the surface led to instability
and the potential for deep convection. The model cor-
rectly represented the presence of a capping inversion,
or lid, and its modulation by a convergence line. As a
result of the model successfully representing the inter-
action between the convergence line, the lid, and the
cold anomaly, it was able to pinpoint the transition
from shallow convective showers to a deep thunder-
storm.
In this case, the success of the model depended on
accurate representation of two scales: the synoptic
(which was well forecast by the coarser-resolution mod-
els driving the 1.5-km model) and the surface-forced
mesoscale convergence line, which was well resolved by
the model and appears to be quite predictable. The
precise mechanism for the formation of this conver-
gence line will be the subject of a subsequent paper.
The interaction between the two scales localized the
convection sufficiently in space and time for the initia-
tion and subsequent development to be highly predict-
able, in spite of the relatively poor representation of
processes at the cloud scale. This is a good example of
nonequilibrium convection, as discussed by Done et al.
(2006), and its association with a high level of predict-
FIG. 20. Summary of the synoptic conditions at 1000 UTC, just prior to the deepening of the convective shower: (a) Analysis of the
relative locations of the convergence line, surface warm tongue, and upper-level cold front in the 1.5-km model. The location of the
convergence line is shown by the dashed line. The position of the upper-level cold front is based on the w  13° C isentrope at 750 hPa.
The line A– B shows the orientation of the cross section in (b). (b) Vertical cross section along the line A– B shown in (a). In both panels,
the shading indicates wet-bulb potential temperature: w 13° C (white), 13° C w 14° C (light gray), w  14° C (dark gray) and
the lowercase letters “ a” and “ b” highlight the location of a shallow pool of warm air. The strong horizontal temperature gradient at
the leading edge of the overriding low-w air has been analyzed as the upper cold front. The dry region, with relative humidity less than
40%, is shown in (b) by the bold gray line. The tropopause has been drawn in (b) along the PV  2 contour. Two regions of high static
stability at low levels have been identified as lids. The thickness of the dashed lines represents the strength of the lid. One of these is
associated with the warm-sector air lagging behind the upper cold front. The other, at about 700 hPa, corresponds to the main lid as
analyzed in this paper.
NOVEMBER 2007 M O R C R E T T E E T A L . 3747

ability once the surface-forced flows are properly re-
solved. The details of the shower, and the later stages of
development of the shower cluster, were more depen-
dent on small-scale processes and less well simulated.
It follows that to improve the forecast via assimila-
tion of observations would require improvement of the
synoptic scale, via the large-scale data assimilation sys-
tem, and/or improved location of the convergence line
before precipitation forms. The latter points to the po-
tential importance of clear-air Doppler radar data, low-
level 2D moisture data (e.g., Fabry et al. 1997), and the
full use of satellite imagery of the early stages of cloud
development.
Acknowledgments. Thanks are due to the Natural
Environment Research Council and the Met Office Na-
tional Meteorology Programme for funding CSIP. We
wish to thank everyone who participated in the CSIP
observational campaign: Judith Agnew, Dave Bamber,
Lindsay Bennett, Barbara Brooks, Ulrich Corsmeier,
Norbert Kalthoff, Sarah Keeley, Darcy Ladd, James
McGregor, Emily Norton, Felicity Perry, and Markus
Ramatschi. Thanks are due to Brian Golding and three
anonymous reviewers for their comments on earlier
versions of this paper. Thanks are due to the Met Office
for the radar network rain-rate data and for launching
extra radiosondes from Larkhill. Thanks are also due to
Eumetsat for the Meteosat-8 high-resolution visible and
water vapor channel data and to the British Atmo-
spheric Data Centre (BADC) for archiving the data
collected during CSIP. The project exploited new in-
struments available through the Universities’ Facility
for Atmospheric Measurement (UFAM), which is
funded by the Natural Environment Research Council
following an initial award from the HEFCE Joint In-
frastructure Fund. The Chilbolton Observatory, around
which the project is based, is managed by John God-
dard and owned by the Council for the Central Labo-
ratory of the Research Councils.
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