Convective clouds
1
Convective Clouds
Introduction
Convective clouds are formed in vertical motions that result from the instability of the
atmosphere. This instability can be caused by:
a.
heating at the bottom of an air layer;
b.
cooling at the top of an air layer;
c.
lifting or saturation of a potentially unstable layer;
d.
a combination of all the above.
Over land the diurnal variation of convective clouds is often clearly visible. On clear mornings,
when the sun is able to heat the earth’s surface the conditions for Cumulus formation are
favourable. If the instability and relative humidity is large enough Cumulus clouds begin to form
early. There is no Cu-formation is the atmosphere is only slightly unstable or stable and the
relative humidity is low. After reaching a maximum in the afternoon the Cu activity decreases
and finally the clouds dissolve in the late afternoon or early evening. Over sea the diurnal
variation of Cu-formation usually is very small. If there is any activity it is near the end of the
night.
The ascending motions of convection decrease or stop when these motions encounter stable
layers, or especially inversions. The form into which Cu-clouds develop mainly depends on the
vertical distance between the condensation level (the level where temperature equals the dew
point) and the cloud base of the stable layer, from the stability and the thickness (vertical
distance from top to bottom) of this layer: if it is very stable the cumulus tops may be forced to
spread horizontally. If the layer is not very thick only some tops will spread horizontally and some
will shoot through the layer and may continue to grow.
Observations have shown that the initiation of convection and its weather is only seldom a local
phenomenon. On satellite photos it is often apparent that convective conditions are occurring
simultaneously over extensive areas and are a result of the vertical structure of an air mass and
the characteristics of the actual dominant large-scale weather systems.
The convection process
In Figure 1 the environmental lapse rate (γ) in the early morning (at the time of the minimum
temperature) is indicated by the thick in the θs,p-diagram; the dew point curve is indicated by the
thick dashed line. If an air parcel with temperature T and dew point Td was lifted from the
surface, then on level A it could reach its lifting condensation level (LCL) . On that level the
temperature of the air parcel is lower than the temperature of the environment so that it will sink
back and condensation will not occur.
Absorption of radiation at the surface will increase the surface temperature (T1, T3 etc.) creating
unstable conditions in the lowest layers of the atmosphere. Therefore air parcels will rise and
subsequently cool dry adiabatically; as a result the environmental lapse rate near the surface will
change (e.g. the curve at T3). The LCL will increase but the temperature of the environment at
this level will be higher than the temperature of the air parcel so that no clouds will form.
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Figure 1. Changes in the environmental lapse rate during the day.
The dew point temperature will change only slightly: usually some water, e.g. dew on the soil or
on vegetation will evaporate into the atmosphere and mix with the lowest air layers. It appears
that as a rule the dew point will slightly increase except when the near-surface air is too dry. A
good estimate of the surface dew point that will be reached can be made by drawing on the
θs,p-diagram the average maximum mixing ratio line (<χ> in Figure 1) of a layer with a thickness
of 50 mb from the surface upwards using equal areas ; the intersection of the line with the
surface pressure isobar will give the dew point.
When T3 is reached an air parcel with temperature T3 and dew point Td3 will lift from the surface.
However, before it reaches LCL3 the air parcel again has reached a temperature by cooling that
is lower than the temperature of the environment, hence LCL3 will not be reached.
Ongoing radiation will increase the temperature even further and T4 will be reached. Air parcels
that rise from the surface will reach level C where the line <χ> is intersected (LCL4); at that
height condensation will occur and consecutive ascent will be along a saturated-adiabat. In the
case of Figure 1 the environmental lapse rate above level C is located to the left of the saturated
adiabat the situation is unstable (in this case is het conditionally unstable). Hence as soon as the
temperature T4 has been reached cumuliform clouds will appear.
The value of the surface temperature where convective clouds start to develop is called the
convection-temperature (Tc), level C is called the convective condensation level (CCL). In
Figure 1 the dry adiabat through Tc and the average maximum mixing ratio line <χ> intersect in
the convective condensation level.
As long as the atmosphere is colder than the rising air parcel, i.e. in Figure 1 as long as the
saturated adiabat γs is located to the right of the environmental lapse rate (γ) the parcel will
continue to rise until the intersection of γs with γ (level D). At this height the vertical motion will
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stop. The area between γs and γ between levels C and D is a measure for the amount of energy
(in J/kg) available for the parcel to rise (increase its potential energy). This area is called CAPE
(Convective Available Potential Energy).
The formation of convective clouds
In the previous text the explanation of rising air parcels changed into an explanation of
convective clouds. With the so-called parcel method the assumption was made that the
processes within the parcel were adiabatic and that the atmosphere was not perturbed by the
motion of the parcel If a whole layer is involved these assumptions may not be true:
a.
in measurements within Cu clouds the saturated adiabatic lapse (γs) rate is never
observed; temperatures and amount of water are lower than expected from the parcel
method; the lapse rate within the cloud is often larger than γs,
b.
rising large air parcels create downdrafts in the surrounding air,
c.
evaporation takes place at cloud boundaries.
Entrainment
The effects mentioned under a are a result that the rising air also mixes with environmental air
while also some air from the parcel is expelled. The characteristics of the air in the parcel will
change continuously. This phenomenon is called "entrainment". The consequences of
entrainment are:
a.
Sometimes the environmental lapse rate suggests strong formation of cumulus or even
showers; in practice only small cumulus (Cu humilis) will form and no showers develop.
b.
The radiosonde data suggest the cloud tops to be at the intersection of γ and γs (D in
Figure 1); in reality cloud tops may be considerably lower.
If the radiosonde data indicate cloud tops at a level higher than 500 mb, then almost always
showers will form if the average relative humidity between the CCL and the 500 mb level is larger
than 75%. If the average relative humidity is less than 30% then showers are very unlikely.
These “rules” only cover a part of the possible cumulus clouds: it is very well possible to have
showers with clouds not even reaching the 500 mb level. Nevertheless these rules illustrate the
influence of entrainment.
Compensating downdrafts
A large amount of air that ascends rapidly creates in the immediate environment compensating
downward motions. These downdrafts reduce the temperature difference between cloud and
environment; hence they also decrease the amount of CAPE (Figure 2). AB is the environmental
lapse rate, AC is the saturated-adiabat. On level P the temperature difference between cloud
and environment is TC-TD. As a results of the downdraft air surrounding the cloud is brought from
level P1 to level P dry-adiabatically, reaching a temperature TE. From the figure it appears that
TC-TE < TC-TD. Therefore the downdraft also lowers the cloud tops.
Evaporation
Especially at the top of the ascending cumulus clouds turbulent mixing with the environment will
occur; there most evaporation will take place. The energy needed for evaporation is supplied by
cooling of the cloud top. Suddenly the dissipating top may vanish completely. However,
evaporating cumulus tops will moisten the dry layer making it easier for consecutive cloud tops to
survive and grow.
From observations it appears that shortly after the top of a cumulus congestus dissipates a
shower may form. In this case cloud droplets of a critical size for active coalescence were first
moved to the cloud top with the rest of the air. The evaporative cooling at the top either creates a
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downdraft or reduces the upward motions enabling these cloud droplets to descend into the
cloud and start the coalescence and precipitation processes.
Figure 2. The effect of downdraft around a rising cloud.
Forecasting convective clouds
Considering the diurnal behaviour of the formation of convective clouds a forecast made at e.g.
08 UTC for the afternoon of the same day must contain a statement of the chance of the
formation of cumuliform clouds. Figure 1 serves as a starting point:
a.
determine the maximum mixing ratio <χ> averaged over the lowest 50 mb using the dew
point curve of the radiosonde data of 00 UTC; the intersection of this line with the
environmental lapse rate will give you the CCL;
b.
from the CCL follow the dry adiabat to the surface pressure isobar and read TC; when
this temperature is reached convective clouds will appear;
c.
decide if the air temperature will exceed TC; and if so at what time this will happen. The
diurnal temperature changes in the same air mass of previous days may be an important
indication;
d.
if TC will be exceeded then the height of the cloud base must be determined. This should
coincide with the CCL because that is where condensation will start; due to a required
supersaturation the cloud base may be somewhat higher. Furthermore, during ascend
some entrainment will occur decreasing the humidity of the air parcel and increasing the
cloud base slightly above the CCL. The cloud base may also be higher than the CCL
because air parcels may not lift directly once TC has been reached: a so-called
superadiabatic lapse rate is the result. The air parcel “sticks” to the surface and only
starts ascending when the wind disconnects it from the earth’s surface. Therefore a wind
speed of a few knots is necessary. If such an air parcel having T > TC leaves the surface,
it will follow a different dry adiabat which intersects <χ> at a higher level than the CCL.
On average the cloud base is approximately 25 mb higher than the CCL. An empirical
formula for the height of the CCL (in m) is hCCL(m) = 125 (T - Td) where T and Td are the
temperature and dew point at the surface, respectively.
e.
Theoretically the cloud top is located in D (Figure 1). In practice, due to the effects
mentioned before, the cloud top is at least 25 mb below level D.
f.
Forecasting the cloud cover (N) of the convective clouds is not easy. In practice the
situation of the previous days, if similar, will give some guidance. An empirical rule, giving
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a rough indication is:
4 U

N =  CCL − 6 
5 6

where UCCL is the relative humidity (in %) at the CCL.
For forecasting precipitation from convective clouds it is important to realize that a cloud can only
produce precipitation if the cloud layer is thick enough and/or the temperature in the upper part
of the cloud is low enough. In the Netherlands cumulus clouds must have a considerable vertical
extent and the 0°C-level must be exceeded in order to produce precipitation. If the cloud top
exceeds the -7°C-level and the relative humidity between the CCL and the cloud top height is
large enough, then surely showers are possible. The higher the cloud top and the larger the area
of instability on the thermodynamic diagram the more intense the precipitation will be.
An important aspect in considering the environmental lapse rate is the fact that there is less
development of convective clouds in and area with anticyclonically curved isobars. Of course
subsiding motions are important, especially when a subsidence inversion is present which may
gradually sink and oppose the vertical development of cumuliform clouds.
After reaching the maximum temperature convection will diminish gradually and will stop once a
surface inversion develops. Remaining clouds will evaporate creating downdrafts because of
evaporative cooling of the air. Cumulus clouds will dissolve and now Sc cumulogenitus may be
observed. Cloud dissipation will occur more rapidly if the air is dry. Especially in the winter
cumuliform clouds, originating over the North Sea, may remain present during the night. With
strong NW-winds these clouds may penetrate a considerable distance inland before dissolving.
Forecasting thunderstorms
The most important conditions for the formation of thunderstorms are:
A large amount of CAPE that can be released by door convection.
Vertical windshear over the entire convective layer.
The intensity of thunderstorms depends on the number of clouds where CAPE is released: many
small clouds or a few large ones.
Windshear has the following effects:
in the absence of shear, updrafts and downdrafts will coincide. Initially the cloud will grow
and the updrafts become stronger; next precipitation will fall from the upper parts of the
cloud. The precipitation will fall trough the region with updrafts, weakening the updrafts
(evaporative cooling) and the cloud dissipates rapidly.
if the wind speed, but not the wind direction, varies with height, then updrafts are directed
at an angle with the vertical and precipitation will fall next to the updraft region. Updrafts
and downdrafts exist next to one another and the cloud will not dissipate as quickly as in
the situation without windshear. Updrafts may be stopped by lateral spreading of cold air
when the downdrafts reach the surface.
if there is shear both in wind direction and wind speed then complex mesoscale systems
may develop. These may live for several hours, more or less independent of surface
heating.
Forecasting thunderstorms
For the forecast of thunderstorms a number of rules and indices have been developed. A brief
overview is presented here.
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- Indication for thunderstorms from the cloud top height:
top below 4000 m:
thunderstorms not likely
top between 4000 and 5000 m: thunderstorms is likely
top over 5000 m:
thunderstorms very likely.
- Boyden Index
The Boyden Index (I) can be calculated from the radiosonde data:
I = (Z-200) - T
Here Z is the thickness of the 1000-700 hPa layer (in dam!) and T is the temperature (in °C) at
700 hPa. Thunderstorms are likely if I >= 94-95.
This index is only valid in or near Western Europe.
- Total Totals Index
This index is commonly used as a severe weather indicator and is defined by:
TT = T850 + Td,850 -2×T500
TT Index
< 45
45 - 55
> 55
Thunderstorm potential
weak
moderate
strong
- S Index
This index is primarily used to indicate thunderstorm potential from April through September. It is
defined by:
S{TT} = TT - (T-Td)700 - K
where TT is the Total Totals Index (see above) and
K is defined as
0 when T850 - T500 ≥ 25
2 when T850 - T500 > 22 and < 25
6 when T850 - T500 ≤ 22
S{TT}
< 40
40 - 46
> 46
Thunderstorm potential
none
possible
likely
- KO index
Defined by:
KO = [(Θe,500 + Θe,700)/2] - [(Θe,850 + Θe,1000)/2]
where Θe,ppp is the potential equivalent temperature at level ppp hPa. NB Θe,ppp can be found in
the “RAOB Data” window as ePT.
This index was developed to estimate thunderstorm potential in Europe. It is more sensitive to
moisture than other more traditional indices and is best used in cooler, moist climates. If the
surface pressure is below 1000 hPa then Θe, surface is used.
KO Index
>6
2-6
<2
Thunderstorm potential
weak
moderate
strong
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- Lifted Index
A stability index used to determine thunderstorm potential. The LI is calculated by lifting an
air parcel adiabatically from a level 50 hPa above the surface to 500 hPa. The difference
between the temperature of the air parcel and the environmental temperature at 500 hPa
represents the LI.
LI
> -3
-3 to -5
< -5
Thunderstorm potential
weak
moderate
strong
- K Index
Defined by:
KI = T850 - T500 + Td,850 - (T700 - Td,700)
KI
< 20
20 - 25
26 - 30
31 - 35
> 35
Thunderstorm potential
none
weak: isolated thunderstorms
moderate: widely scattered thunderstorms
moderate: scattered thunderstorms
strong: numerous thunderstorms
The above indices must always be used in conjunction with large-scale indicators of
thunderstorms. Extra convective energy may depend on:
Position and movement of upper-air troughs or lows. Thunderstorms are usually found
along or just in front of upper-air troughs or lows.
The existence and movement of low-level convergence lines, such as fronts.
Elevated areas may receive extra heating.
Other useful synoptic tools are:
Analysis of the dew point and/or Θw. Tongues with high dew points or Θw may help to
define areas with a high risk of thunderstorms;
Difference in Θw on 500 and 850 hPa
Cyclonic curvature of surface isobars.
Conditions favourable for heavy thunderstorms:
Heavy showers are mesoscale systems that might develop into heavy thunderstorms. It is
important to note that:
some of the heaviest and most extensive thunderstorm activity did not occur before 1218 hours after the temp showed a warm layer over a convective layer. Such a warm layer
may prevent the release of convective energy, thereby building up the amount of energy
which is subsequently released explosively.
In heavy thunderstorms the dew point is often higher than 13°C and may even reach
18°C. At 850 hPa analogous values for Θw are found. Winds are from the SE to SSW
with speeds between 20-30 kts in a narrow tongue.
Advection of dry air at mid-atmospheric levels. With potential instability Θw at 500 hPa
may 2-5°C lower than at 850 hPa. Winds at 500 hPa must have veered 20-40 degrees
relative to the winds at 850 hPa, speeds 35-50 kts.
Further veering above 500 hPa, with 300 hPa winds from SSW-W and speeds of 50-85
kts are good conditions to have the downdrafts in a favourable position for generating
heavy thunderstorms.