Download Clouds - Dennis Mansfield

Clouds
Why is the Sky Blue?
If you were to travel 20 miles or so above the Earth's surface, the sky
would appear black. What happens during light's descent to Earth that
makes the sky take on a wonderful azure hue?
"White" sunlight passes through our
atmosphere, and molecules in the air,
primarily nitrogen, are just the right size to
scatter light from the blue end of the
visible spectrum. The other colors travel to
the ground with little interference.The
blue light is scattered from molecule to
molecule in the sky, until the light seems
to be coming from every direction.
And Clouds are White Because...?
...the water droplets that make up
clouds are much larger than the
molecules that scatter blue light.
The clouds scatter and reflect all
the visible colors of light that
strike them. Hence, we have white
clouds.
But if the cloud is thick enough, light does not penetrate completely
through the cloud, resulting in dark, heavy-looking cloud bottoms.
Why do clouds form?
Clouds are nothing more than water vapor that condenses and accretes
into a visible form. The usual mechanism is for moisture-laden air near the
Earth's surface to be raised higher into the atmosphere either by an
encroaching air mass or the heat of the sun. As the air is lifted, the
pressure drops and the air is subsequently cooled. The combination of the
two causes water vapor to condense.
CHARACTERISTICS OF AIR PARCELS AND AIR MASSES; CLOUDS
Note: A helpful summary of air parcel behavior can be found at this University of
Georgia web site.
We have now examined the roles of temperature and pressure in determining
many of the aspects of the Earth's weather and climates. We will next explore
how such factors as T and P variations, together with changes in water content of
the atmosphere, movements of the oceanic currents, and the rotation of the
Earth itself, go about producing daily weather at our home locations and beyond.
Later we shall see that this also determine aspects of the varieties of climates
found over the globe. One fact will become obvious: Although the atmosphere's
chemical composition remains nearly constant both spatially (at a given altitude)
and over time, the air itself within the troposphere is subject to small to large
variability in the thermodynamic factors stated above. This provides the driving
forces that lead to the general characteristics of weather and climate.
Local, relatively small volumes of air within the boundary layer whose behavior
we choose to investigate are termed air parcels. Much larger volumes of air (at
least 1600 km [1000 miles] in horizontal dimensions) that can move considerable
distances within the troposphere are called air masses. These are characterized
by these features: 1) they are characteristic of large regions in which properties
of air have small horizontal variation (thus, the surface temperatures are similar
at different places in the region), and 2) boundaries separating air masses are
called fronts.
Any given air parcel or air mass shows relative uniformity (homogeneity) in its
defining characteristics (general temperature range, pressures, moisture content,
etc.). Over very large areas, there are different air masses (properties of one fall
within specific ranges that are in contrast to other masses elsewhere in the
atmosphere) that are usually in motion following definite paths through the
troposphere with respect to other masses existing at the time which are located
adjacent to each specific air mass. These relative movements establish the
weather systems acting at any particular time.
Now, in the real world of the atmosphere air in the boundary layer near the
surface will be heated by solar insolation combined with a rise in ground or water
surface temperatures. This heating causes parcels to expand and rise into the
atmosphere (thermal convection; see below). As this rising air moves upwards it
will lose heat to its surroundings and cool. The drop in temperature that results
will follow some pattern called the Lapse Rate. The Normal Lapse Rate in the
troposphere is 6.5 °C.; this is also called the Environmental Lapse Rate (ELR).
The cooling itself is adiabatic (no more heat enters or leaves). As long as the
parcel is warmer than its surrounding air, it continues to rise. But if, as it
continues to cool, it enters air that is less cool (i.e., warmer than it has become)
and may start to sink or subside. In the examples that follow, it is assumed that
an air parcel does not physically mix with its surroundings (no air exchanged
across boundary between parcel and surroundings). In fact, winds, turbulence,
and other factors do disturb the boundary region as some mixing occurs but not
enough to compromise the parcel's thermodynamic behavior.
In the general case exhibited in this next diagram, the black line shows the
typical temperature curve in the troposphere and part of the stratosphere. For a
parcel having the same temperature as that for the normal ELR (T e = 10°C), at
that altitude (height), the parcel will be stable and stationary. If the temperature
Tp of the parcel is, say,- 10°C, the air will sink; if greater than Te, say, 30°C, the
air will rise.
The next three diagrams further explore this concept:
The above diagram simply shows the expansion of the parcel as it moves
upwards from the surface (1000 mb pressure) to an altitude where the pressure
of the surrounding air is at 500 mb.
In the next pair of diagrams, the top shows how the parcel expands and cools as
it reaches an altitude where its internal pressure (700 mb) equals that of the
surrounding air pressure, at which height it would cease to rise further. In the
bottom situation, the parcel has moved into air whose pressure is < 700 mb (not
properly labeled as such in this diagram) so that it is now denser and begins to
sink, warming as it falls.
The lifting or rising of an air mass can be accomplished in any of 4 ways:
1. Thermal Convection: from surface or lower atmosphere heating.
2. Dynamic Convergence: winds coming together force air to concentrate,
squeezing it and causing it to move upwards.
3. Frontal Collision: two different air masses, one colder than the other, meet,
causing the cold air to wedge under the warm air and driving the latter upward.
4. Orographic Uplift: Air moving laterally meets mountains or other features that
form topographic barriers, causing the air to rise to get past.
For now, we will concentrate on thermal convection; the other three causes of
uplift will be described later. Look first at this diagram that illustrates convective
uplift.
These two charts introduce the concepts of Air Stability and Instability.
Stable air experiences relatively less changes in temperature with height, and
consequently less tendency to move rapidly up or down to new positions.
Unstable air resulting from considerable surface heating moves the air up rapidly
with notable cooling (larger temperature changes) enroute.
If the air has a low moisture content for its temperature, i.e., is unsaturated (does
not contain all the moisture it can hold at the prevailing temperature[s]), its
temperature decrease as it rises (moves upwards) is determined by its Dry
Adiabatic Lapse Rate (DALR). A typical value is -10°C per kilometer upwards into
the atmosphere. If, instead, the air is saturated, the Moist Adiabatic Lapse Rate
(MLR) applies. A typical value may be - 6°C per km of rise. The general
relationship between DALR and MLR is shown here and comments about the
two parameters appear beneath the plot:
We can improve our understanding of lapse rates by looking at examples with
actual numbers. In the left half of the diagram below T d (for an unsaturated air
parcel) denotes the ELR temperature, at 0°C. The air parcel is initially at T = 20°;
upon rising under dry adiabatic conditions the air will cool at 1 km to T = 10°C but
at that altitude the ELR temperature has changed to 2°C, so the air will keep
rising. In the right half, the starting condition for the air has an ELR T d of 20°C
and the air parcel itself has Tp = 20°C. As the air parcel is forced up, it cools but
at a lower rate because condensation releases latent heat. When it reaches 1
km, T = Td and the moisture laden air ceases further rising.
The Actual Lapse Rate, shown in this next diagram, varies in rate of change at
different altitudes. It is similar to the Environmental Lapse Rate (ELR) but shows
actual variations in rates at different altitudes as various factors come into play.
Relative stability can also be expressed by comparing the Dry Lapse Rate to the
Actual Lapse Rates. As the two Actual Rates move from the DLR towards the
curves shown, the air mass becomes either increasing stable (blue curve) or
unstable (red curve).
Note that the Dry Adiabatic Lapse Rate tends to increase at a constant rate
compared with the Moist Lapse Rate which varies with altitude.
In the next few paragraphs as we will expand on what happens to air parcels
upon lifting in terms of their degree of stability.
Lets expand upon these ideas with a related set of diagrams.
A Stable Atmosphere is one that strongly resists change. It occurs whenever the
Dry Adiabatic Lapse Rate is greater (and thus cools more with height) than the
Environmental Lapse Rate. An air parcel under this condition that is forced
upwards cools rapidly (quickly becoming colder than its surroundings) and acts
as though it has negative buoyancy, i.e., it overcomes the lifting force and tends
to sink at some stage to restore equilibrium. Stable air is clear (blue skies) and
devoid of stormy conditions.
In a Neutral Atmosphere, the Dry and Environmental (Actual) Lapse Rates are
the same and the temperatures of both the parcel and its surrounding air are
identical at some neutral layer (which can have a notable vertical thickness,
within which the parcel and surrounding air are in thermal equilibrium [same
temperatures]). A parcel will rise and cool until it reaches this layer where the
temperature balance occurs. This situation usually occurs through some external
lifting force or condition such as convergence or orographic rise.
An Unstable Atmosphere is marked by the DALR being less than the ELR. A
rising parcel remains warmer than its surrounding air and has positive bouyancy
(tends to keep rising). It needs little external force conditions to commence rising.
It will continue to rise until, as it cools, it attains thermal equilibrium with its
surroundings.
An unstable air parcel will behave differently when the amount of moisture it
contains is sufficient to bring the Moist Lapse Rate into play. Consider this figure.
(The term LCL stands for Lifting Condensation Level, an altitude determined by
some pressure at which any moisture in the parcel experiences temperaturecontrolled saturation and begins to condense).
In the above diagram, the moist air parcel is initially 7°C at the surface. It is
initially unsaturated and cools according to the dry adiabatic lapse rate. At the
LCL it is saturated and cools according to the moist adiabatic lapse rate. The
diagram shows that the parcel is warmer than the surrounding air both above and
below the LCL; thus, the entire layer illustrated by this diagram is categorized as
unstable. As we have pointed out, in an unstable layer the initial lifting force is
only needed to get to the parcel going upwards. Immediately after the lifting
begins, the parcel is buoyant and convection will continue on its own. Unstable
environments are common in the afternoons during the summer and are often
responsible for producing late afternoon thunderstorms. A typical set of
conditions is depicted in this diagram.
This next pair reiterates some of the above ideas but puts typical lapse rate
values into the plots. The straight line nature of the plots is often alluded to as
Absolute Stability and Absolute Instability. The plots appear again in the two
diagrams below the pair in which the conditions are displayed pictorially. In the
diagrams, Γ or Gamma refers to the adiabatic lapse rates that are further defined
by the subscripts e = environmental, d = dry, and m = moist.
The next two diagrams pictorialize the conditions associated with Absolute
Stability and Absolute Instability:
The conditions for Absolute Instability occur usually in the lower atmosphere on
warm days. More commonly, the conditions favor Conditional Instability. The
rising air has a lapse rate between the local DALR and MLR. The air initially rises
as stable. Upon reaching the LCL, this air becomes relatively warmer than
surrounding air and will continue to rise without any outside forces (self buoyant).
This situation is especially associated with orographic effects. These diagrams
show this instability variant.
Another condition that affects air parcels in Inversion. This occurs when
conditions reverse the lapse rate so that instead of continually falling
temperatures a layer is reached in which the temperatures actually rise. This can
happen naturally when a warm air mass rides over a cold (as described later) or
when special conditions associated with pollution are made. Diagrammatically,
inversion appears thusly.
Los Angeles is notorious for its smog layers. The combination of hot air coming
from the desert, off the mountains and into the Los Angeles Basin, that is then
underridden by cool air off the ocean, produces this common condition;
Let's apply some of the above ideas to a real case atmosphere. In such an
atmosphere different layers have different stability responses. First look at this
table defining the three prime modes of stability
Γe < Γm < Γd= Absolutely stable
Γm < Γe < Γd = Conditional Instability
Γm < Γd < Γe = Absolutely unstable
------------------------------The symbol Γ (Gamma) refers to adiabatic lapse rate; d = dry; m = moist; e
= environmental
The real case is show in terms of straight line lapse rates that shift their slopes at
different altitudes. When the appropriate calculations are done (not shown here),
the layers involved show these stability modes: 1) absolute instability; 2)
conditional instability; 3) absolute stability; 5) conditional instability; 5) absolute
stability; 6) absolute stability.
We are now ready to shift attention to Air Masses. In one sense, an air mass can
be thought of as a huge air parcel. Once formed it tends to move laterally through
the troposphere (often confined mainly to thick boundary layers) where it meets
and interacts with other air masses it may encounter. An individual air mass
(remember, at least 1600 km wide to qualify as one) will show more internal
variability than the small air parcel but is characterized by sets of T, P, and
moisture conditions that are fairly homogeneous in toto but with discrete changes
of these conditions with altitude and fluctuations from place to place within the
mass.
The air masses can be grouped into five major types based on combinations of
two sets of parameters: 1) cold, cool, warm; and 2) dry and moist. The figure and
table below show these combinations and provide indications of their major
characteristics.
They are further classified in terms of broad geographic places of origin, set up in
these two groupings: 1) Arctic; Polar; Tropic (from high to low latitudes rotational pole to equator); and 2) Maritime and Continental (oceans and large
land masses).
The general starting point regions and directions of movement of the major air
masses onto the North American continent are depicted in this geographic
diagram:
The next map shows air mass distribution at a global scale. While the plotted
positions are typical, their locations are generalized, so that any of the five types
can begin in different regions (but controlled by the latitudinal and land/water
locations described above) and more than the number shown may exist at any
one time.
Within an air mass there will usually be clouds present (can be few in number in
a cold, dry air mass or can be predominant [almost everywhere] in a moist, warm
air mass). Clouds represent condensed forms of water (mainly coarse to fine
droplets; but also very fine mist; ice particles, hail pellets, snowflakes) that may
have finite boundaries as seen from a point locale or may seem to extend
beyond the horizon.
There are four main ways in which clouds form through condensation
mechanisms involving uplift of moist air, as shown in this diagram:
Cloud classifications go back to ancient times but the first modern one was
presented by the Englishman Luke Howard in 1802. Variants of this has since
emerged but the basic factors in classifying clouds remain their heights (several
different ranges), their modes of origin, their appearance, and their relation to
modes of precipitation. Nomenclature is taken from Latin words. All this is
summarized in this chart.
This next diagram (split to appear in top/bottom format) is a pictorial
representations of the different cloud types listed in the table.
From Lutgens and Tarbuck, 1992; adapted from a Ward's Natural Science
Establishment illustration.
Many sites on the Internet show photos of actual clouds. The one I recommend
was prepared by the The Cloudman who happens to be Dr. John Day. (Surf this
site for other information on clouds and other pictures.)
A cloud is a visible mass of condensed droplets, frozen crystals suspended in the
atmosphere above the surface of the Earth or another planetary body, such as a moon.
(Clouds can also occur as masses of material in interstellar space, where they are called
interstellar clouds and nebulae.) The branch of meteorology in which clouds are studied
is nephology.
On Earth the condensing substance is typically water vapor, which forms small droplets
or ice crystals, typically 0.01 mm in diameter. When surrounded by billions of other
droplets or crystals they become visible as clouds. Dense deep clouds exhibit a high
reflectance (70% to 95%) throughout the visible range of wavelengths: they thus appear
white, at least from the top. Cloud droplets tend to scatter light efficiently, so that the
intensity of the solar radiation decreases with depth into the cloud, hence the gray or even
sometimes dark appearance of the clouds at their base. Thin clouds may appear to have
acquired the color of their environment or background, and clouds illuminated by nonwhite light, such as during sunrise or sunset, may be colored accordingly. In the nearinfrared range, clouds would appear darker because the water that constitutes the cloud
droplets strongly absorbs solar radiation at those wavelengths.
Clouds can cast shadows
Clouds and cloud bow above Pacific
Contents
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1 Cloud formation and properties
o 1.1 "Hot ice" and "ice memory" in cloud formation
2 Cloud classification
o 2.1 High clouds (Family A)
o 2.2 Middle clouds (Family B)
o 2.3 Low clouds (Family C)
o 2.4 Vertical clouds (Family D)
o 2.5 Other clouds
o 2.6 Cloud fields
3 Colors
4 Global dimming
5 Global brightening
6 Clouds on other planets
7 See also
8 References
9 External links
[edit] Cloud formation and properties
Global scheme of cloud optical thickness
1. The air is cooled below its saturation point. This happens when the air comes into
contact with a cold surface or a surface that is cooling by radiation or the air is cooled by
adiabatic expansion (rising). This can happen:
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along warm and cold fronts (frontal lift)
where air flows up the side of a mountain and cools as it rises higher into the
atmosphere (orographic lift)
by the convection caused by the warming of a surface by insolation (diurnal
heating)
when warm air blows over a colder surface such as a cool body of water.
2. Clouds can be formed when two air masses below saturation point mix. Examples are:
our breath on a cold day, aircraft contrails and Arctic sea smoke.
3. The air stays the same temperature but absorbs more water vapor into it until it reaches
saturation point.
The water in a typical cloud can have a mass of up to several million tonnes. The volume
of a cloud is correspondingly high and the net density of the relatively warm air holding
the droplets is low enough that air currents below and within the cloud are capable of
keeping it suspended. Conditions inside a cloud are not static: water droplets are
constantly forming and re-evaporating. A typical cloud droplet has a radius on the order
of 1 x 10-5 m and a terminal velocity of about 1-3 cm/s. This gives these droplets plenty
of time to re-evaporate as they fall into the warmer air beneath the cloud.
Cumulonimbus cloud
Most water droplets are formed when water vapor condenses around a condensation
nucleus, a tiny particle of smoke, dust, ash or salt. In supersaturated conditions, water
droplets may act as condensation nuclei.
The growth of water droplets around these nuclei in supersaturated conditions is given by
the Mason equation.
Water droplets large enough to fall to the ground are produced in two ways. The most
important means is through the Bergeron Process, theorized by Tor Bergeron, in which
supercooled water droplets and ice crystals in a cloud interact to produce the rapid growth
of ice crystals; these crystals precipitate from the cloud and melt as they fall. This process
typically takes place in clouds with tops cooler than -15 °C. The second most important
process is the collision and wake capture process, occurring in clouds with warmer tops,
in which the collision of rising and falling water droplets produces larger and larger
droplets, which are eventually heavy enough to overcome air currents in the cloud and
the updraft beneath it and fall as rain. As a droplet falls through the smaller droplets
which surround it, it produces a "wake" which draws some of the smaller droplets into
collisions, perpetuating the process. This method of raindrop production is the primary
mechanism in low stratiform clouds and small cumulus clouds in trade winds and tropical
regions and produces raindrops of several millimeters diameter.
This wave cloud pattern formed off of the Île Amsterdam in the far southern Indian
Ocean
The actual form of cloud created depends on the strength of the uplift and on air stability.
In unstable conditions convection dominates, creating vertically developed clouds. Stable
air produces horizontally homogeneous clouds. Frontal uplift creates various cloud forms
depending on the composition of the front (ana-type or kata-type warm or cold front).
Orographic uplift also creates variable cloud forms depending on air stability, although
cap cloud and wave clouds are specific to orographic clouds.
[edit] "Hot ice" and "ice memory" in cloud formation
In addition to being the colloquial term sometimes used to describe dry ice, "hot ice" is
the name given to a surprising phenomenon in which water can be turned into ice at room
temperature by supplying an electric field of the order of one million volts per meter.[1]
The effect of such electric fields has been suggested as an explanation of cloud
formation. This theory is highly controversial and is not widely accepted as the
mechanism of cloud formation. The first time cloud ice forms around a clay particle, it
requires a temperature of -10 °C, but subsequent freezing around the same clay particle
requires a temperature of just -5 °C, suggesting some kind of "ice memory".[2]
[edit] Cloud classification
Main article: List of cloud types
Cloud classification by altitude of occurrence
Clouds are divided into two general categories: layered and convective. These are named
stratus clouds (or stratiform, the Latin stratus means "layer") and cumulus clouds (or
cumuliform; cumulus means "piled up"). These two cloud types are divided into four
more groups that distinguish the cloud's altitude. Clouds are classified by the cloud base
height, not the cloud top. This system was proposed by Luke Howard in 1802 in a
presentation to the Askesian Society.
[edit] High clouds (Family A)
These generally form above 20,000 feet (6,000 m), in the cold region of the troposphere.
In Polar regions, they may form as low as 16,500 ft (5,030 m); they are denoted by the
prefix cirro- or cirrus. At this altitude, water frequently freezes so clouds are composed
of ice crystals. The clouds tend to be wispy and are often transparent.
Clouds in Family A include:
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Cirrus (CI)
Cirrus uncinus
Cirrus Kelvin-Helmholtz Colombia
Cirrostratus (Cs)
Cirrocumulus (Cc)
Pileus
Contrail, a long thin cloud which develops as the result of the passage of an
aircraft at high altitudes.
[edit] Middle clouds (Family B)
Altocumulus mackerel sky
These develop between 6,500 and 20,000 feet (between 2,000 and 5,000 m) and are
denoted by the prefix alto-. They are made of water droplets and are frequently
supercooled.
Clouds in Family B include:
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Altostratus (As)
Altostratus undulatus
Altocumulus (Ac)
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Altocumulus undulatus
Altocumulus mackerel sky
Altocumulus castellanus
Altocumulus lenticularis
[edit] Low clouds (Family C)
Low clouds
These are found up to 6,500 feet (2,000 m) and include the stratus (dense and grey).
When stratus clouds contact the ground, they are called fog.
Clouds in Family C include:
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Stratus (St)
Nimbostratus (Ns)
Cumulus humilis (Cu)
Cumulus mediocris (Cu)
Stratocumulus (Sc)
[edit] Vertical clouds (Family D)
Cumulonimbus clouds showing strong updrafts
These clouds can have strong up-currents, rise far above their bases and form at many
heights.
Clouds in Family D include:
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Cumulonimbus (associated with heavy precipitation and thunderstorms) (Cb)
Cumulonimbus incus
Cumulonimbus calvus
Cumulonimbus with mammatus
Cumulus congestus
Pyrocumulus
Mammatus cloud formations
[edit] Other clouds
A few clouds can be found above the troposphere; these include noctilucent and polar
stratospheric clouds (or nacreous clouds), which occur in the mesosphere and
stratosphere respectively.
[edit] Cloud fields
A cloud field is simply a group of clouds but sometimes cloud fields can take on certain
shapes that have their own characteristics and are specially classified. Stratocumulus
clouds can often be found in the following forms:
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Open cell, which resembles a honeycomb, with clouds around the edges and clear,
open space in the middle.
Closed cell, which is cloudy in the center and clear on the edges, similar to a filled
honeycomb.
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Actinoform, which resembles a leaf or a spoked wheel.
[edit] Colors
A cloud in a gradient blue sky.
An example of various cloud colors
Colourful cloud formation
Iridescent clouds
Iridescent clouds
Rain bearing clouds
Rain clouds over the North Sea taken from the coast of Herne Bay, Kent
The colour of a cloud tells much about what is going on inside the cloud. Clouds form
when relatively warm air containing water vapor is lighter than its surrounding air and
this causes it to rise. As it rises it cools and the vapor condenses out of the air as microdroplets. These tiny particles of water are relatively densely packed and sunlight cannot
penetrate far into the cloud before it is reflected out, giving a cloud its characteristic
white color. As a cloud matures, the droplets may combine to produce larger droplets,
which may combine to form droplets large enough to fall as rain. In this process of
accumulation, the space between droplets becomes larger and larger, permitting light to
penetrate much farther into the cloud. If the cloud is sufficiently large and the droplets
within are spaced far enough apart, it may be that a percentage of the light which enters
the cloud is not reflected back out before it is absorbed (Think of how much farther one
can see in a heavy rain as opposed to how far one can see in a heavy fog). This process of
reflection/absorption is what leads to the range of cloud color from white through grey
through black. For the same reason, the undersides of large clouds and heavy overcasts
appear various degrees of grey; little light is being reflected or transmitted back to the
observer.
Other colours occur naturally in clouds. Bluish-grey is the result of light scattering within
the cloud. In the visible spectrum, blue and green are at the short end of light's visible
wavelengths, while red and yellow are at the long end. The short rays are more easily
scattered by water droplets, and the long rays are more likely to be absorbed. The bluish
color is evidence that such scattering is being produced by rain-sized droplets in the
cloud.
A greenish tinge to a cloud is produced when sunlight is scattered by ice. A
cumulonimbus cloud which shows green is a pretty sure sign of imminent heavy rain,
hail, strong winds and possible tornadoes.
Yellowish clouds are rare but may occur in the late spring through early fall months
during forest fire season. The yellow color is due to the presence of smoke.
Red, orange and pink clouds occur almost entirely at sunrise/sunset and are the result of
the scattering of sunlight by the atmosphere. The clouds are not that color; they are
reflecting the long (and unscattered) rays of sunlight which are predominant at those
hours. The effect is much the same as if one were to shine a red spotlight on a white
sheet. In combination with large, mature thunderheads this can produce blood-red clouds.
The evening before the Edmonton, Alberta tornado in 1987, Edmontonians observed such
clouds — deep black on their dark side and intense red on their sunward side. In this case
the adage "red sky at night, sailor's delight" was wrong.
[edit] Global dimming
The recently recognized phenomenon of global dimming is thought to be caused by
changes to the reflectivity of clouds due to the increased presence of aerosols and other
particulates in the atmosphere.
[edit] Global brightening
New research From Dimming to Brightening: Decadal Changes in Solar Radiation at
Earth's Surface by Martin Wild et al. (Science 6 May 2005; 308: 847-850) indicates
global brightening trend.
Global brightening is caused by decreased amounts of particulate matter in the
atmosphere. With less particulate matter there is less surface area for condensation to
occur. Since there's less condensation in the atmosphere and increased evaporation
caused by increasing amounts of sunlight striking the water's surface there is more
moisture, causing fewer but thicker clouds.
[edit] Clouds on other planets
Main article: Extraterrestrial_atmospheres
Within our solar system, any planet or moon with an atmosphere also has clouds. Venus'
clouds are composed entirely of sulfuric acid droplets. Mars has high, thin clouds of
water ice. Both Jupiter and Saturn have an outer cloud deck composed of ammonia
clouds, an intermediate deck of ammonium hydrosulfide clouds and an inner deck of
water clouds. Uranus and Neptune have atmospheres dominated by methane clouds.
Saturn's moon Titan has clouds which are believed to be composed largely of droplets of
liquid methane. The Cassini-Huygens Saturn mission has uncovered evidence of a fluid
cycle on Titan, including lakes near the poles and fluvial channels on the surface of the
moon.
[edit] See also
In mountainous areas one often finds the peaks above the clouds as here for the Pico
Ruivo seen from Pico do Arieiro, Portugal.
Weather Portal
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CLAW hypothesis
Cloud albedo
Cloud Appreciation Society
Cloud base
Cloud condensation nuclei
Cloud feedback
Cloud forcing
Cloud seeding
Cloud types
Cloudscape photography
Coalescence
Extraterrestrial skies
Flight ceiling
Fog
Fractus cloud
Mammatus
Mist
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Monsoon
Mushroom cloud
Orographic lift
Precipitation
Thunderstorm
Tornado
Tropical cyclone
Weather lore
Mammatus Clouds
sagging pouch-like structures
Mammatus are pouch-like cloud structures and a rare example of clouds in sinking air.
Sometimes very ominous in appearance,
mammatus clouds are harmless and do not mean
that a tornado is about to form; a commonly held
misconception. In fact, mammatus are usually
seen after the worst of a thunderstorm has
passed.
Photograph by: Manikin
As updrafts carry precipitation enriched air to the cloud top, upward momentum is lost
and the air begins to spread out horizontally, becoming a part of the anvil cloud. Because
of its high concentration of precipitation particles (ice crystals and water droplets), the
saturated air is heavier than the surrounding air and sinks back towards the earth.
The temperature of the subsiding air increases as it descends. However, since heat energy
is required to melt and evaporate the precipitation particles contained within the sinking
air, the warming produced by the sinking motion is quickly used up in the evaporation of
precipitation particles. If more energy is required for evaporation than is generated by the
subsidence, the sinking air will be cooler than its surroundings and will continue to sink
downward.
The subsiding air eventually appears below the
cloud base as rounded pouch-like structures
called mammatus clouds.
Photograph by: NOAA
Mammatus are long lived if the sinking air contains large drops and snow crystals since
larger particles require greater amounts of energy for evaporation to occur. Over time, the
cloud droplets do eventually evaporate and the mammatus dissolve.
Mammatus typically develop on the underside of
a thunderstorm's anvil and can be a remarkable
sight, especially when sunlight is reflected off of
them.
Photograph by: NOAA
CLOUD FORMATION
Clouds form as air rises and cools to its dew point temperature. At this point, the air is
saturated and water vapor begins to condense into the liquid water droplets that make up
a cloud. This occurs at the base of the cloud, which can also be referred to as the
condensation level.
How do we get rising air in the first place? Normally, the atmosphere is in a state of
balance (called hydrostatic balance) where the downward force of gravity and the
upward pressure gradient force of air molecules cancels out - meaning that there is
normally no net force causing air to rise or sink. However, certain processes can
overcome this balance and lead to rising air, and thus cloud formation. These processes of
cloud formation (Figure 5.8 in Ahrens) are:
1. Convection (also called "free convection" or "buoyant lifting"): This occurs when
heated air becomes less dense than the surrounding air and begins to rise. If the
overlying atmosphere is stable, then a rising air parcel will soon become cooler
than its surroundings and cease to rise. The resulting cloud that forms will not
have a large vertical extent, but instead will be somewhat flat (e.g., stratocumulus,
cumulus humilis). If, on the other hand, the atmosphere is unstable or
conditionally unstable, convection may result in a large, towering cloud forming
(i.e. cumulus congestus or cumulonimbus). Daytime heating of the earth's surface
drives the convective activity we see in Tucson.
2. Topographic (also called "orographic lifting" or "forced lifting"): This refers to
rising air motions that result from air flow over a large obstacle such as a
mountain range. Horizontal winds are forced to rise up and over the obstacle (see
Figure 5.12 in Ahrens). This rising motion can be enough to cause cloud
formation and possibly precipitation. On the back (or "leeward") side of the
obstacle, air sinks. Recall that sinking air becomes compressed and heats up,
which prevents cloud formation. This leads to the formation of a rain shadow on
the leeward side of a mountain range.
3. Convergence: Within a low pressure system (also called a cyclone), air flows
counterclockwise (in the Northern Hemisphere) and in toward the center. This
convergence of air at the surface results in rising air motion. For this reason, midlatitude low pressure systems are commonly associated with cloudy and rainy
conditions. Within a high pressure system (or anticyclone), air flow clockwise and
outward from the center. The diverging surface air at the center of an anticyclone
is balanced by downward motion, which prevents cloud formation. Thus high
pressure systems are associated with clear conditions.
4. Frontal lifting: Recall that fronts are boundaries between two different air
masses. A cold front marks the boundary between warm air and advancing cold
air. The colder air, being more dense or "heavier" than the warm air, acts as a
wedge that pushes the warm air up and out of the way. This leads to rising motion
and cloud formation ahead of a cold front. In the case of a warm front, advancing
warmer air encounters a cold air mass and, being "lighter" or less dense, runs up
along the boundary of the warm front. This leads to rising motion and cloud
formation that occurs well ahead of the location of the surface frontal boundary
(see Figure 5.8d).
It should be emphasized that the stability of the atmosphere plays an important role in
determining the type of cloud that forms as a result of any of these processes. Rising
motion in an unstable or conditionally unstable environment will favor vertical
development, and increase the liklihood of precipitation.
3/22/00
PRECIPITAION
We define precipitation as any liquid or frozen water that falls from a cloud and reaches
the ground we it can be measured. Water vapor can be converted into precipitation
through the processes of condensation, freezing, and deposition (i.e., going directly
from vapor to solid form).
When water vapor condenses or freezes, it prefers to condense or freeze onto something it needs a surface (preferable some liquid water or ice already present). Away from the
surface of the earth (in the "free atmosphere"), the only available surfaces are tiny solid
particles suspended in the air (dust, smoke particles, sea salt, etc.) These are called
condensation nuclei or ice nuclei, depending on the process. They are generally quite
small; typical sizes are 2 micrometers (or 0.002 millimeters) in diameter.
The formation of precipitation begins with the process of nucleation, which is the
deposition, freezing, or condensation of water vapor in the free air onto condensation
nuclei. Let's say you manage to get some water vapor to condense onto a condensation
nuclei. What will determine whether this initial water droplet will continue to grow or
not? The answer is that the relative humidity of the air around the droplet will determine
its growth. If we assume the nuclei is "normal" (i.e. neither attracts nor repels water
molecules), then



IF RH < 100 % ---> evaporation exceeds condensation ------> droplet will shrink
IF RH = 100 % ---> evaporation equals condensation ------> droplet will remain
same size
IF RH > 100 % ---> condensation exceeds evaporation ------> droplet will grow
We see that under "normal" circumstances, drop growth by condensation or depostion
will only occur when the environment is supersaturated - when RH > 100 %. This is
fairly rare. How then will a cloud ever form?
The answer is that some condensation nuclei are hygroscpic,i.e., they attract water.
Because they attract water molecules, hygroscopic nuclei allow drop growth to occur
when RH is equal to or less than 100 %. Thus these type of nuclei are very important for
cloud formation. Efforts to enhance precipitation through "cloud seeding" add more
particles to the atmosphere to increase the likelihood of nucleation.
It is common to observe tiny droplets of liquid water in clouds even when the
temperature is below freezing. When liquid water is present at below freezing
temperatures, it is called supercooled water. Remember that water molecules want to
have surfaces to condense (or in this case, to freeze) onto - call them ice nuclei in this
case. Since there are very few ice nuclei in a typical cloud, you end up with supercooled
water droplets hanging around looking for something to freeze on. If it gets cold enough
(at or below -40 C), these supercooled water droplets will freeze even in the absence of
ice nuclei. This is called spontaneous nucleation.
The handout distributed today shows that typical condensation nuclei are 0.1 to 1
micrometer in diameter. Typical cloud droplets are observed to be 1-30 micrometers in
diameter. The smallest precipitation particles are about 0.2 millimeters in diameter (or
about 200 micrometers), and they can be up to 5 millimeters (5000 micrometers) in size.
Condensation and depostion alone can grow typical cloud droplets in a period of 20-120
minutes, but you would have to wait almost forever for these processes to produce a
precipitation-sized particle that is heavy enough to fall. Luckily, there are other ways to
produce precipitation-sized particles that work much faster. The growth of precipitation
particles beyond the cloud droplet stage can be accomplished by the combined effects of
collision and coalescence, and also by the ice crystal process (aka the Bergeron process).
Precipitation Formation by Collision and
Coalescence
Collision and coalescence invlolve interaction of liquid water droplets with other liquid
water droplets - it is most important in clouds with temperatures above freezing - and so
together they are referred to as a "warm rain process". The Bergeron process is most
important in clouds with temperatures below freezing. It involves interactions between
ice particles, supercooled water, and water vapor - hence the name "three-phase
process").
Collision is self-explanatory. Liquid cloud droplets carried by air motions within a cloud
can collide. Obviously, the most effective type of collisions will involve one big droplet
moving through a group of smaller droplets. If the droplets are all moving at different
speeds, that will also increase the likelihood of collisions.
Coalescence refers to the fact that water is "sticky". If two water droplets come into
contact (say by collision) then they may stick together and make one larger droplet. Not
all droplets that collide will stick together - they could bounce off each other. Therefore,
collision and coalescence are not an entirely efficient process. Furthermore, these
colliding droplets will tend to limit the size that a precipitation particle can reach - no
larger than about 5 millimeters in diameter.
3/24/00
PRECIPITATION (cont'd)
The Ice Crystal Process of Precipitation
Formation
Much of the precipitation (frozen or liquid) reaching the earth's surface began high up in
cumulus clouds where the temperature is below freezing. In this environment, the
Bergeron process is much more important for the production of precipitation particles.
The physical principle that drives the Bergeron process is the fact that the saturation
vapor pressure above supercooled liquid water is greater than the saturation vapor
pressure over ice. If you have an ice crystal suspended in a cold cloud which is
surrounded by supercooled water droplets, the difference in saturation vapor pressure
between ice and liquid means that there are more vapor molecules next to the supercooled
droplet and less vapor molecules next to a neighboring ice crystal. This difference drives
a process known as diffusion that moves vapor molecules away from the droplet and onto
the ice crystal, where they deposit themselves. To replace these diffusing molecules and
maintain saturation, water evaporates from the supercooled droplet.
Under these conditions the ice crystal grows at the expense of the supercooled droplets
(see Figure 5.19 in Ahrens). This process is very effective within a cold (i.e. below
freezing) environment due to the fact that there are many more supercooled droplets in a
typical cloud than there are ice crystals. This means there are lots of supercooled droplets
to "feed" each crystal.
Under the Bergeron process, precipitation particles can grow very large very fast. Once
they grow large enough to fall into regions where the temperature is above freezing,
collision and coalescence can also contribute to further growth. For precipitation to
actually reach the surface, however, there has to be sufficient moisture in the cloud for
droplet growth through condensation & deposition, and through collision and
coalescence. There also has to be the right number of ice crystals versus supercooled
water droplets for the Bergeron process to take place. And finally, the air between the
cloud base and the surface has to be moist enough so that the precipitation particles do
not evaporate before reaching the surface. Precipitation particles that fall from a cloud
and evaporate before reaching the surface are called virga.