MODULE 4.2C
SHORT RANGE FORECASTING OF
CLOUD, PRECIPITATION AND
RESTRICTIONS TO VISIBILITY
Boundary Layer
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Table of Contents
TABLE OF CONTENTS
1
1. BOUNDARY LAYER CLOUDS, PRECIPITATION, AND RESTRICTIONS TO
VISIBILITY: INTRODUCTION
2
2. BROAD SCALE FORECASTING
2
3. FORECASTING QUANTITATIVE DETAIL (FINE TUNING)
3
Diurnal Variations
Dissipation of SC/ST at Night
Dissipation of SC/ST Due to Daytime Heating
Formation of SC/ST At Night
Stratus or Fog?
3
4
4
6
6
Upslope Boundary Layer Clouds
7
Lowering of Boundary Layer Clouds in Precipitation
7
Onshore Clouds and Fog
8
Radiation Fog,
12
Ice Fog, Blowing Snow, Haze
Ice Fog
Blowing Snow
Haze
15
15
16
16
Climatology
16
4. BOUNDARY LAYER PRECIPITATION
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1. BOUNDARY LAYER CLOUDS, PRECIPITATION, AND RESTRICTIONS
TO VISIBILITY: INTRODUCTION
Two kinds of forecast problem can be identified when predicting boundary layer clouds and
precipitation. One problem is that of depicting broad scale or general areas of organized
boundary layer weather. The other problem is one that is of particular concern to aviation - the
specification of detail or quantitative information. Before forecasts can be given on the fine
scale aspects, however, the broad scale boundary layer cloud areas need to be specified.
On most occasions, the broad scale or organized boundary layer cloud and precipitation can be
forecast using one approach only - by assessing and anticipating relevant physical processes.
Straight extrapolative techniques generally do not work well with this cloud type. Synoptic
correlation techniques, on the other hand, have merit only if synoptic related processes such as
advections and divergence are at play.
What processes need to be considered in forecasting boundary layer clouds and precipitation?
The important physical processes required for the continued existence of this cloud are a source
of moisture and a means of cooling the low levels of the atmosphere.
Because of the strong emphasis on processes, the first step to predicting boundary layer cloud is
to identify all physical processes that could change the low level moisture supply or bring about
cooling. If boundary layer cloud already exists, the question of whether or not it will continue
consists of anticipating changes that might occur in the main forcing processes. The other more
difficult problem - the forecasting of boundary layer cloud where none exists - must also be
approached from the point of view of anticipating future processes. In either case, the prediction
of boundary layer cloud can prove a complex one always requiring a careful diagnosis of the
situation.
2. BROAD SCALE FORECASTING
The procedure for coming up with a broad scale prognosis of boundary layer clouds,
precipitation, and restrictions to visibility is much the same as that for diagnosis. The prognosis
requires the use of a current analysis, a short range surface prog, and an idea of future changes in
the low level flow aloft. From these, dominant processes at the valid forecast time are identified
as being favourable for either maintenance or dissipation of boundary layer cloud. These
processes can be depicted either mentally or on a blank chart:
a)
Positions of fronts, troughs, lows and the like that might produce low level convergence or
maintain the moisture supply. Alternatively, ridges and highs are useful to consider in the
dissipation process.
b) Other areas of Convergence or divergence that are suggested by the isobars.
c)
Significant upslope or downslope areas. At the same time, areas where onshore or offshore
flows could develop prove significant.
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d) Turbulent mixing areas, indicated generally by where surface winds are expected to be
greater than 15 knots. If a strong pressure gradient is forecast over a hilly area, the potential
for some stratocumulus due to the turbulent mixing is fairly strong.
e)
Areas with moist land surface conditions (e.g. from significant rainfall over the past 24
hours, or from melting show) give the potential either for stratus or for daytime
stratocumulus/cumulus or greater convection.
f)
Areas of low level cooling can lead to stratus and should be suspected when the low level
flow advects over cooler water or a cool land surface. On the other hand, a flow of cooler
air over warm water can increase the moisture content of the air and develop an area of
stratocumulus or of convective cloud.
g) Synoptic scale processes can produce boundary layer cloud through the evaporation of
precipitation and through lift. Synoptic correlation can help in these cases.
Once the broad scale boundary layer processes have been identified, the forecaster must evaluate
other meteorological parameters such as stability and history. These considerations determine
whether the cloud will exist as SC, ST, or have a greater convective nature.
3. FORECASTING QUANTITATIVE DETAIL (Fine Tuning)
A number of techniques are available for forecasting the quantitative details of boundary layer
clouds and restrictions to visibility. Some of these techniques include tools used in timing the
formation and dissipation of cloud or fog. Other techniques include tools for forecasting cloud
bases and restricted visibilities. Because the factors that determine the existence and the type of
this weather are considerable, many of the techniques tend to be quite situation-specific. As a
result, the discussions that follow will cover a number of cloud scenarios. The first portion of
the discussion will outline tools that can be used in evaluating the significance of diurnal
processes to the forecast problem. The remainder of the discussion on fine-tuning techniques
will cover approaches that can be taken when forecasting boundary layer weather for various
dominant processes such as upslope or onshore flow or radiational cooling.
Diurnal Variations
Whether forecasting boundary layer cloud for a site or on the broad scale for an area, diurnal
variations need to be considered. In some situations, especially when dealing with site-specific
weather, diurnal processes can effectively become the sole determinant of variations in
parameters such as local cloud ceilings and visibilities. When dealing with fog, for instance,
diurnal variations in occurrence are definitely skewed towards a maximum in the late overnightearly morning period and a minimum towards the mid afternoon period. For SC and for some
ST, this variation can take the form of a dissipation at night (minima), on the one hand, or a
formation at night due to radiational cooling (maxima). The form that this diurnal variation
causes in boundary layer cloud depends upon the thermal and moisture structure of the lower
atmosphere.
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The diurnal processes that affect boundary layer cloud will be considered at this point. Later
discussions will take up the case of restricted visibilities.
The local dissipation of ST and SC takes place by two processes: either by heating or by mixing
in of drier air. Both processes can be diurnally influenced. The local formation of ST and SC,
on the other hand, is a process mostly supported diurnally by night time cooling. Although
daytime heating is most instrumental in forming low level clouds during the warmer seasons,
these clouds are of a convective nature; their relevant diurnal processes will be covered later
under convective clouds and precipitation.
Dissipation of SC/ST at Night
The process of night time dissipation of boundary layer clouds is assumed to be one of mixing in
of dry air due to destabilization of the cloud top from radiative cooling. This process is most
likely when:
• the cloud sheet is bounded at its top by a dry type inversion, i.e. a rapid decrease of humidity
with height through the region of temperature increase;
• the surface RH is low;
• the cloud is initially thin.
The final point is particularly significant. When the boundary layer cloud is thick, the effect of
night time cooling is as much apt to give thickening of the cloud as it is to result in its dissipation
due to the mixing in of drier air. Hence, diurnal cooling is most likely to succeed in dissipating
boundary layer cloud if the air above the inversion is dry and if the cloud is already thin. When
these conditions apply, nocturnal cooling effectively acts to accelerate the processes that would
have eventually given the cloud dissipation.
Dissipation of SC/ST Due to Daytime Heating
The process here for dissipating boundary layer cloud obviously is solar heating. The clouds that
normally can be dissipated by heating are the more stable types, having moisture trapped by a
capping inversion.
The amount of heating needed to force the cloud dissipation in any situation can be estimated
from a representative tephigram by assuming that the cloud will clear whenever the surface
temperature reaches a value that would establish a dry adiabatic lapse rate up to the level of the
ST top. If a steep inversion is observed, the amount of heating required may be considerably
more than the solar input can provide, and the cloud may not dissipate (particularly in
wintertime). The amount of incoming solar radiation at any time can be a function of many
variables. The most important, of course, are the time of day, the time of year and latitude.
Other factors that can prove significant include consideration for additional cold or warm air
advection, the presence of overlying cloud, and the thickness of cloud to be dissipated.
Although nomograms exist to determine whether or not daytime heating will be able to dissipate
trapped boundary layer cloud, they will not be referenced here. Most of the objective techniques
available tend to step through a number of nomograms that calculate the heat needed to bum off
the capping inversion and the time that the required energy will become available. The objective
techniques for low cloud are very similar to those used in timing the dissipation of radiational
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fog (one described in the nomograms document). Their main differences, if any, tend to lie with
calculations of the reduced insolation (due to fog or cloud).
A handy alternative to the direct use of nomograms is a qualitative assessment of heating effects
using a tephigram representative of the trapped boundary layer cloud. To achieve the
assessment, the forecaster requires both a forecast maximum temperature and a representative
12Z sounding. Armed with these pieces of information, the effect of daytime heating can be
simulated by drawing on the sounding a dry adiabat from the expected maximum surface
temperature to the capping inversion. If the dry adiabat corresponding to the maximum
temperature 'breaks' the inversion, then solar insulation is assumed sufficient to act, during the
course of the day, to dissipate the trapped boundary layer cloud. Although this procedure
appears easy enough, the total number of steps involved in getting the final answer may be
significant if the forecaster first has to estimate the expected maximum temperature.
Given that the forecaster has determined solar insolation to be sufficient to dissipate cloud, the
following rule of thumb can be used to determine the time required for this event. The rule,
derived from observations made in the United States, suggests that the number of hours of
sunshine needed to dissipate ST equals the thickness of the layer in hundreds of metres.
e.g.
ST base
300m
ST tops
800m
thickness = 500m
clearance time: N=5 hours after sunrise
This rule assumes no effect for latitude or time of year, as well as no cloud above the stratus
layer. A table can be used to compensate for a higher layer of cloud. The table is included
below for your reference. Use the table to find the appropriate factor ‘K’, then multiply the
number of hours found above by the factor ‘K’.
Table of the Factor 'K'
AMOUNT OF CLOUD (IN TENTHS)
CLOUD TYPE
1
2
3
4
5
6
7
8
9
10
CI/CS
1.0
1.0
1.1
1.1
1.1
1.1
1.2
1.2
1.2
1.3
AC
1.1
1.1
1.2
1.2
1.3
1.4
1.5
1.7
1.8
2.0
AS
1.1
1.1
1.2
1.3
1.3
1.4
1.6
1.9
2.2
2.5
SC
1.1
1.1
1.2
1.4
1.5
1.6
1.8
2.1
2.4
2.9
Formation of SC/ST At Night
As well as influencing the dissipation of low cloud, diurnal influences can also lead to the
formation of cloud. Daytime heating, for example, can act to dissipate trapped SC or ST or can
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lead to the formation low level clouds of a convective nature. Nocturnal cooling, on the other
hand, can lead to the formation of low level cloud or can assist in the dissipation processes
described earlier. Normally, nocturnal cooling is more instrumental in the formation of cloud
than it is in its dissipation. Techniques for forecasting the nocturnal formation will be described
in the current section while techniques associated with convective processes will be described
later.
The night time formation of ST and SC is more likely if the air is moist and the surface and low
level RH are high. Other favourable conditions include a stable sounding, clear skies, and wind
speeds of the order of 5-9 knots. If surface winds drop lower, fog or a combination of fog and
stratus can be expected. In particular, radiational ST and SC occur relatively commonly at
valley and coastal sites, where the air may remain moist. This radiative cloud has a tendency to
occur near sunrise after a long night of cooling, and to form when the spread is observed near or
less than 2oC.
The problem of determining whether or not low cloud will form at a site due to nocturnal cooling
can be helped by following trends in the spread between surface temperature and dewpoint and
by accepting history. If ST occurred the previous evening and the spread is decreasing again
during the current evening, a repeat performance might he expected. The timing and the ceiling
reached during the previous night might be used as a first guess of expected conditions, provided
that the forecaster keeps in mind the fact that conditions might have changed from those of the
previous night. The forecaster is warned that this first guess often fails, probably because of
slight changes in favourable conditions at a site over time. When history is not available, the
forecaster must maximize use of the trend, taking into account the time of night and the number
of hours remaining before sunrise, local sources of moisture and pollutants, and local drainage
wind peculiarities. Any information that a representative tephigram might yield on moisture in
the low levels can assist in forecasting a ceiling.
An objective technique is described in the nomograms document that calculates the time
required for the formation of low cloud due to nocturnal cooling - provided, of course, that
conditions are favourable. The technique also calculates a probable lowest ceiling. Known as
the Goldman method, the technique can be used to forecast stratus conditions associated with a
couple of situations. The nomograms can be applied to radiational low cloud as well as to
boundary layer cloud formed by the lowering of ceilings in continuous precipitation.
Stratus or Fog?
At times, the forecaster may have difficulty in deciding whether saturation of the air will produce
fog or stratus. In most cases, the controlling factor in determining whether fog or stratus will
result is wind speed. Unfortunately, no critical wind speed values exist to separate the formation
of fog from that of low stratus. Even if proven values did exist, they probably would vary quite
widely from site to site. Typical values of the geostrophic winds for which night time cooling
could lead to stratus rather than fog formation might be in the order of 15-20 knots at a ‘normal’
site and 10-15 knots at an exposed site. In a deep valley with a cross-wind, the critical wind
might exceed 30 knots. Regardless of location, stronger winds imply higher bases - given that
stratus is expected.
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Upslope Boundary Layer Clouds
Forecasting the existence of upslope boundary layer clouds is much easier than the job of
forecasting their bases and tops. One simple approach to the problem of estimating bases and
tops of this cloud is to simulate on a blank tephigram the effect of the upslope or the adiabatic
lifting of air by the underlying terrain. The raw materials needed for this simulation include a
surface analysis, a blank tephigram, and perhaps some surface observations. Using the surface
analysis, the first step is to determine an upstream surface trajectory that some ‘average’ parcel
might have followed in arriving at the forecast site over the course of a period of time (say, 6
hours). The time chosen can be fairly arbitrary, as further discussion will indicate. Once the
surface temperature and dewpoint of the upstream origin to the trajectory are noted, the ‘source
parcel’ is lifted adiabatically on the blank chart, assuming a parcel lift by an amount equivalent
to the difference in station elevations between the forecast site and the origin. It is assumed at
this point that the parcel trajectory is upslope. The point along the lift and the trajectory where
saturation first occurs will be the location where the upslope stratus will first form, provided that
the trajectory does not pass over progressively drier or more moist ground. A simple rule to
remember in applying this procedure is that, for every degree of initial dew point spread,
approximately 300 feet of adiabatic lift are required to give saturation (i.e. the adiabatic lapse
rate). When representative soundings are available, upstream trajectories are not necessary. The
upslope lift then can be applied to appropriate layers of low level moisture on the upstream
tephigram and the resulting bases and tops of the low cloud that form can be used as forecast
numbers for the upslope cloud.
When the air is dry at upstream points, a significant amount of upslope lift and a corresponding
long trajectory may be required before saturation and cloud can result. The length of trajectory
and the time required to give saturation for the parcel are very much a function of the moisture of
the air at source and of the amount of orographic lift that can act. As a result, the period of time
chosen to represent the lifting process is arbitrary, as long as the time is sufficient for saturation
to result.
Forecasting the dissipation of upslope cloud is usually a matter of forecasting a change in the
low level flow from upslope to downslope. Alternatively, a forecast of a lesser component of
upslope flow may improve the upslope cloud conditions.
Lowering of Boundary Layer Clouds in Precipitation
A fairly common problem for the forecaster is that of timing the lowering of ceilings in a
continuous rain area (e.g. in warm frontal situations). The rate at which the ceiling lowers or
drops below airport operating limits holds paramount concern to aviation.
Several approaches are available for estimating both the timing and the height reached by the
lowered ceiling. Included among these is an objective approach known as the Goldman Method.
This method and its associated nomograms were mentioned earlier as objective tools for use in
forecasting the formation of stratus due to radiational cooling. While only the first nomogram
was required earlier for radiational stratus, all of the nomograms associated with the Goldman
Method are needed in the continuous precipitation situation.
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Other approaches beyond the nomograms are available for predicting the timing and bases of
lowering ceilings. One quick, simple, but nonetheless effective approach is simply to look at
upstream locations for an estimate of the length of time required there for ceilings to lower after
initiation of precipitation. Allowing for differences in station elevation, the height of the
upstream stratus ceiling can provide a rough guess of possible conditions that will develop. Of
course, the forecaster always has to consider any other factors such as upslope or onshore flow
that might have influenced the final ceiling at the upstream sites.
When available, a local radiosonde can help in predicting the final ceiling. The use of local
soundings should be limited to 6 hours or less prior to the precipitation beginning; otherwise, a
sounding closer to the precipitation area should be consulted. The guidelines for inferring cloud
from the sounding include the following:
a) The cloud base will form close to the base of the temperature inversion and will
approximate the height where the dewpoint spread is a minimum;
b) Given multiple temperature inversions and/or levels of maximum RH, the ceiling will form
at the highest layer first.
Onshore Clouds and Fog
An onshore flow can result in stable boundary layer clouds such as coastal stratus and sea fog.
At other times, the same onshore flow under conditions representing an unstable thermal
structure can result in convective clouds and in significant precipitation. The processes that
influence each of these results are somewhat similar. The forecasting techniques for handling
each case, on the other hand, can be quite different. The techniques available for forecasting
stable onshore clouds and restricted visibilities will be discussed at this point. The convective
results of an onshore flow will await document 4.3D.
Coastal stratus forms when warm air follows a trajectory across progressively colder water.
Often, the onshore flow will not produce ST over land during the daytime when the flow is light
because surface heating counteracts any inland penetration. When this occurs, the ST may tend
to form inland during the night time. The ST will form inland at a time when radiational cooling
becomes sufficient to produce saturation of air that has been conditioned to high low level
moisture by the onshore flow.
The timing of the advection or the night time formation of coastal ST can be approached from a
few angles. For many forecasters, a knowledge of their local climatology coupled with an
examination of what happened under similar conditions the previous night or day can give fairly
reliable predictions. Alternatively, when marine stratus is advecting into the area, the forecaster
can deduce the time of arrival from a series of satellite IR images. Important clues on timing can
also be found from surface observations. If a few tenths of ST are reported during the onset of
the suspected event, the forecaster generally can expect the formation of a ceiling within the
following 1/2-2 hours, given that the local processes remain operative.
Heights and thicknesses of a marine ST deck can be determined from a representative tephigram.
In practice, tephigrams representative of onshore ST may be hard to find, especially if the low
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cloud has yet to reach shore. An alternative, in this situation, to having a representative
sounding is to consider the 00Z sounding for the area to be affected and to determine the levels
where the onshore moisture already exists. Often, the ST will form where the onshore flow has
conditioned the atmosphere to higher humidities. When available, PIREPs prove invaluable to
the determination of ST tops. In such cases, the observed heights are normally assumed to
persist into the short term unless the forecaster is aware of processes which could otherwise
change conditions. Climatological statistics also can indicate reasonable site-specific ceilings
associated with marine stratus; statistics that indicate probable ceilings as a function of wind and
season often indicate typical marine stratus conditions.
The problem of forecasting sea (or coastal) fog is very much like that of forecasting coastal
stratus. Both phenomena can be forecast using similar techniques. The main differences in
approach between the two lie with any objective or empirical methods used.
As was the case for stratus, one of the best approaches in predicting sea fog is the use of the
latest weather observations. Observed weather trends have particular value when dealing with
the 0-6 hour period. Any reports from ships (note the water temperature), lighthouses, and
coastal stations prove particularly useful in locating the fog. On some occasions, a series of
satellite images can detect and provide guidance on the time of arrival of the fog. Visibilities
normally are forecast to drop quite low once sea fog establishes itself over an area; variations are
usually site-specific and often can be forecast using climatological statistics.
Most objective techniques available for forecasting sea fog originate from empirical studies of
the factors pertinent to its formation. Both of the objective techniques described in this section
are based on empirical studies and tend to be regionally-specific. In practice, however, the
findings revealed by these studies likely can be transferred to other regions, allowing for the odd
modification.
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Figure 1. Example of advection fog formation due to southwest surface flow (after
Osborne).
One such study of sea fog behaviour by Osborne concerned itself with fog in the Maritimes.
This piece of research investigated two aspects of the problem: the formation of sea fog over
water and the advection of the weather onto a coast. Highlights of Osborne's work that could
prove quite useful include the following:
a) When a south to southwest flow prevails as shown in Figure 1, (with an origin from Cape
Cod, for example), the Maritime fogs that form over the western Atlantic initially may be too
shallow to be accurately advected with the surface geostrophic wind. A more reasonable
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advection was found to be 60% of the surface geostrophic gradient and backed 25° from
geostrophic. Osborne produced a nomogram to determine how long a trajectory was needed to
produce or generate sea fog, taking into consideration the geostrophic wind speed and the
upwind dewpoint of the air. The study suggested that sea fog would form when the upstream
dewpoint exceeded the sea temperature by at least 3°C. For stronger flows, a higher upstream
dewpoint would be required to give fog because the warm upstream air (assuming the same
temperature as earlier) then spends less time cooling over the water. Similarly, a colder water
temperature than that normal to the Maritimes in summer would mean that a shorter fetch would
produce the same amount of fog. (See Figure 2). For most sites on the Pacific coast, for
example, where the water is relatively cool in summer and the fetch can be significant, the
forecaster needs to be less concerned with the formation of the sea fog or stratus, and more
concerned with the advection of already thick fog. Forecasters dealing with the Great Lakes, on
the other hand, would need to consider the sea fog formation process since fetches there are
considerably less.
Figure 2. Diagram to determine distance required to form fog in given geostrophic flows
(225° - 290°), for various mean dew point - sea temperature differences (after Osborne).
b) Often, fog in the Maritimes will form in a weak south to southwest flow. If the wind is
forecast to increase, say to 20 knots, the fog frequently can be expected to dissipate or lift to a
ST deck in the initial stages of formation.
Another study, this time by Gurka et al (1982), concerned itself with coastal fog in the area south
of New England. This study produced a series of qualitative techniques regarding the motion of
a well-defined fog bank once it meets the coastline. Highlights of this study include the
following:
1. Like its stratus counterpart, sea fog rarely moves more than a few miles inland during the day.
Instead, the sea fog spreads out parallel to the coast upon hitting the shore in daytime and
spreads at a rate faster than the initial speed.
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2. Sea fog moving nearly parallel to the shoreline before hitting land is more likely to spread
rapidly along the shore than are those fogs that impact from the perpendicular.
3. Most local differences in the rate of growth along the shoreline can be attributed to variations
in the near shore surface temperatures. The rate of growth is faster if the fog is advected over
progressively colder water.
Radiation Fog,
Fog and stratus are weather elements of special interest to the transportation industry since either
can result in delayed schedules affecting both people and equipment. Good visibility is
especially important to aviation for safe efficient operations. Consequently, forecasters need to
provide accurate predictions on all aspects of the fog problem.
The first step required in forecasting any fog event is deciding whether or not fog will occur
during the period of interest. This non-trivial exercise often can be handled using techniques
such as synoptic correlation, modelling the weather, watching observed trends, and use of
representative soundings. A more difficult problem, that of timing the formation and dissipation
of fog, can be approached through use of objective techniques, chart history, observed trends,
and by diagnosis of processes and trends observed from the satellite imagery.
Most of the forecasting techniques mentioned earlier under the discussion of “diurnal variations
in boundary layer cloud” are as applicable to the fog forecasting problem as to the low cloud
problem. In nearly all cases, there is little point in distinguishing radiational stratus and
radiational fog as representing two different boundary layer processes. Whether stratus forms is,
as mentioned earlier, largely determined by a critical wind speed. On average, the critical wind
value is quite variable and is a function of the terrain and, to some extent, the airmass
characteristics.
Techniques for timing the formation and dissipation of radiation fog parallel those for
forecasting ST. As was the case earlier, observed trends at a site can provide important clues as
to whether or not stratus or fog will form and can indicate a possible time of formation.
Likewise, history sometimes can prove a useful approach to timing formation and dissipation.
To a lesser extent, the tephigram can indicate in a rough way the time required to dissipate
radiational fog by daytime heating; the time required often is proportional to the strength of the
inversion. Since the boundary layer processes associated with the formation and dissipation of
radiational stratus and fog are similar, many of the same approaches apply for either problem.
(See notes on the formation of boundary layer cloud due to radiational cooling and the
dissipation due to daytime heating).
When available, local station studies can prove invaluable to the fog forecasting problem. In
addition, basic climatological data can provide rough first guesses on typical visibilities and
times of greater fog frequency for given sites.
The main differences in the techniques for forecasting stratus and fog occur in the application of
objective techniques. When methods rely on an assessment of boundary processes, they can be
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transferred from one forecast problem to another. But, objective techniques cannot since they
tend to be derived from empirical evidence and the relevant nomograms tend to differ.
One of these objective techniques was developed by George to address the fog/no fog decision.
The technique is described in Document 4.3F. The method makes use of radiosonde data and is
applicable for fog prediction over a large area rather than for a specific site.
Other objective techniques are available for forecasting formation and dissipation times of
radiational fog. These techniques include work by Jacobs, Bartham and Kagawa.
Only the highlights of each technique will be given below with the details reserved for 4.3F (for
your reference only). Although useful, the application of these objective techniques is initially
time consuming. With practice and careful selection of the number of sites chosen for use of the
nomograms, however, each of these objective techniques can be applied quickly and can add
considerable improvement or refinement to the timing of fog. Once comfortable with the
operational setting, the forecaster is well advised to practice and to adapt the nomograms to their
‘problem sites’ for fog forecasting.
An objective procedure, developed by Jacobs, can prove useful for forecasting the time of onset
of fog as well as minimum temperatures. This technique was developed from an empirical
finding; namely, that a liquid water content of 0.5 g/kg corresponds to a fairly dense fog. The
technique is based on the premise that air must be cooled below its dew point to condense the
liquid water content and that the amount of extra cooling required is inversely related to the
dewpoint of the air. The following table, based on empirical results, shows the required amount
of cooling needed for a given initial dewpoint to produce a dense fog.
Cooling Below Dewpoint Required for Fog Formation
Dewpoint range (°C)
Extra Cooling (°C)
-2 to 2
4
3 to 6
3
7 to 9
2.5
10 to 14
2
15 to 21
1
Thus, for a sunset situation with T = 15°C and Td = 5°C, the amount of cooling required would
be (15-5+3) = spread + 3 = l3°C. The remaining method and nomograms given in 4.3F are used
to determine if and when this amount of cooling can be achieved and the saturation reached.
Another technique, known as the Bartham Method, is available to forecast the time of onset of
fog formation. This technique was empirically developed and contains calculations of the
cooling required to give saturation and fog, taking into account cloud cover and wind speed.
Although a bit cumbersome to apply on the first try, the method has definite merit and is used
operationally by the occasional forecaster.
The timing for fog clearance can be forecast objectively using a technique known as the Kagawa
Method. In deriving the technique, Kagawa assumed a certain characteristic vertical profile of
temperature and dewpoint associated with radiation fog and calculated the total solar insulation
required to “burn off” the capping inversion and the fog. The Kagawa Method contains a
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number of nomograms for determining the heating required and the time that the insulation
becomes available.
The dissipation of radiation fog can also be forecast qualitatively and quickly using satellite
imagery. One technique, pioneered by Gurka, used the observation that extensive areas of
radiation fog and stratus tend to dissipate from their outer edges inward. (See Figure 4). Gurka
suggested that a process of inward mixing was at least partly responsible for this erosion. A
temperature gradient along the, fog boundary produced by differential surface heating was
claimed to set up a circulation similar to a sea-breeze (see Figure 3), with the result that the
circulation eroded fog along the edges due to the sinking and mixing in of warmer and drier air.
Figure 3. Depiction of circulation that develops near cloud edges due to differential surface
heating (after Gurka).
When viewing fog banks on satellite imagery, care must be exercised. The geometrical centre of
the fog area may not necessarily be the most persistent area, since the thickness and density of
the fog will in general be non-uniform, thus affecting the rate of erosion. The brightest fog areas
on the imagery will tend to be the thickest (or densest) and should dissipate at the slowest rate.
A valley fog, on the other hand, likely will not dissipate according to this model since local
valley circulations should tend to dominate over the sea-breeze type of effect.
Used as a forecasting tool, the relative times of dissipation of radiation fog at various locations
can be determined using a sequence of geostationary satellite images, paying close attention to
the brightness of the fog (related to fog depth) and the particular location of a given station under
the fog bank.
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Figure 4. Schematic of fog clearance process (plan view) over a period of a few hours.
No mention has been given to this point of methods for forecasting fog thicknesses. In general,
the depths of radiation fog will vary considerably. Simple radiation fog, as suggested by a study
by Kagawa, may be only 10-15 mb (300-550 feet) thick. Given valley drainage effects, though,
the fog can be much deeper (up to 50 mb thick).
Typical visibilities in thick radiation fog tend to range from zero to 1/4 mile, normally within
one hour after the first hint of fog formation. The fog may be fairly thin at first, giving only a
partially obscured sky condition. Eventually, if the vertical mixing is sufficient, the fog may
become thick enough to produce total sky obscuration.
In terms of dissipation, the depth of fog is again important to the sequence of events that follow.
Thin fog, for example, will tend to dissipate with a partially obscured sky condition as the
visibility improves rapidly from say, zero to one mile to 15 miles in successive hours. With the
thicker fogs, or those interior to a fog bank, a slower improvement can be anticipated, with
stratus ceilings observed and then breaking up as the morning progresses.
Finally, the forecaster should bear in mind that while the techniques described give definitive
answers, the problem of fog forecasting can be quite frustrating on a case-by-case basis. A slight
change in wind speed or direction, or an unexpected patch of cloud passing through the region
can ruin a seemingly good forecast. The capable forecaster will apply these techniques, in
combination with good synoptic judgment and knowledge gained of the climatology of each
station in the area of responsibility, to produce consistently good radiation fog forecasts.
Ice Fog, Blowing Snow, Haze
Ice Fog
Apart from some local studies, the material from Module 4.1 on ice fog processes summarizes
much of the available operational guidelines.
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Blowing Snow
In general, visibilities in blowing snow are forecast using empirical results and using knowledge
of the relevant processes. The material from Module 4.1 summarizes most of the available
guidelines.
Haze
The discussion from Module 4.1 covers most of the processes and climatology needed to forecast
visibilities in haze.
Climatology
In practice, climatological statistics can prove quite valuable as a weather forecasting tool.
When interpreted properly, these statistics can provide a certain amount of ‘instant’ experience
to the meteorologist newly assigned to a forecast region and can provide a quick recall to the
more experienced individual familiar with the peculiarities of an area.
Climatological statistics are particularly suited for application to site forecasting and prove
helpful in the composition of aviation terminal forecasts. When formatted in meaningful ways,
climatological data can assist the forecaster in determining details such as the most probable
winds for a site in a given synoptic setup or the most reasonable ceiling and visibility at a
location for a given wind. Because the statistics summarize the result of many weather
occurrences, they generally are able to reflect the persistent peculiarities of a site and to indicate
the circumstances under which the peculiarities occur. In a sense, climatological statistics can
help ‘fine tune’ the forecasts and indicate in a reasonable manner the expected variability of
conditions from one site to another given a common broad scale pattern.
Climatology can prove quite helpful to the inexperienced forecaster dealing with boundary layer
weather. Although the statistics can also be useful in some convective situations, their real
strength comes in describing the persistent mesoscale effects that influence weather at a site.
Local boundary layer influences such as a nearby body of water, a paper mill in the area, a hill
on the horizon, surrounding terrain and topography all play a significant role in determining the
mesoscale differences between stations. The statistics can reflect these peculiarities and can
indicate their relative importance for various flows.
4. BOUNDARY LAYER PRECIPITATION
The problem of forecasting boundary layer precipitation areas is very similar to the problem of
forecasting its cloud. While synoptic scale precipitation frequently can be extrapolated directly
or by association with synoptic features, the extrapolation of most forms of boundary layer
precipitation is almost certain of failure. The only time when either of these techniques can be
assured of success occurs when the boundary layer precipitation and cloud can be clearly
associated with synoptic features such as a cold front or fog.
In summer, boundary layer precipitation mostly takes the form of scattered showers associated
with stratocumulus, or drizzle associated with stratus. In winter, boundary layer cloud can
produce a variety of types and frequencies of precipitation ranging from continuous snow to
intermittent freezing drizzle to scattered flurries. On average, the occurrence of boundary layer
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precipitation is considerably less frequent in summer than in winter, provided of course, that
purely convective precipitation is not counted as boundary layer forced. In general, procedures
for forecasting boundary layer precipitation are the same as for the cloud, except that
precipitation may or may not accompany the cloud. Whether precipitation is produced or not
depends on the cloud thickness and temperature and on the amount of lift and moisture
maintaining the cloud. The thicker and cooler the cloud or the greater the lift and moisture
operative, the greater the chances that boundary layer precipitation will occur. In winter
precipitation in the form of light continuous snow can readily occur over a large area from
stratocumulus when temperatures are low and when the cloud thickness is relatively thin. The
minimal amount of lift (and cloud thickness) required to give wintertime boundary layer
precipitation usually turns out to be less than that required to give summertime precipitation.
Since summertime atmospheric temperatures are warmer, boundary layer showers require greater
cloud thicknesses (and lift) to occur during this season. The indirect key to whether or not
boundary layer cloud will produce precipitation appears to be its cloud temperature.
A number of empirical studies have been documented that relate the probability of boundary
layer precipitation occurrence to the properties to its cloud. Some of these results are
summarized in the temperature and height relationships given below. Only a small fraction of
past studies are shown here.
RELATIONSHIPS BETWEEN CLOUD THICKNESS, TEMPERATURE AND
PRECIPITATION (for your reference only)
THICKNESS
0 - 2000 ft
2000 - 3300 ft
3300 - 7500 ft
several km
PRECIPITATION
None
Drizzle Only
Intermittent Light Snow
Moderate Rain or Snow
TEMPERATURE OF CLOUD BASE
Height of Cloud Base Above
0°C Isotherm
Less than 3300 ft
More than 3300 ft
Cloud Thickness
6600 + ft
Any value
Probability of Precipitation
Reaching Ground
90%
5%
THICKNESS AND PRECIPITATION
Clouds whose thicknesses exceed the height of the cloud base by 3300 feet will probably
precipitate. If the thickness is less than the cloud base height, it probably will not. One needs
TCU thicknesses of 7 to 13 thousand feet for coalescence to be effective.
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CLOUD TOP TEMPERATURES AND PRECIPITATION
Rain and snow are unlikely if the cloud top temperature is warmer than -12°C. Rain and snow
usually start to fall, or the fall will increase markedly when the cloud top temperature falls below
-12°C (for ice crystal production, temperatures of -15°C are required).
SHOWERY PRECIPITATION
Cloud depth as calculated by the parcel method should exceed 150 mb.
Wind shear from the ground to the cloud top should be less than 3 knots per 1000 feet. (For air
mass showers).
Only slight chance of showers if cloud top temperature is warmer than -12°C, showers are likely
if temperature is colder than -12°C, and almost certain if the cloud top temperature is below
-40°C.
Showers may also form along a sea breeze front.
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