Download Precipitation I: Processes and measurement

Precipitation I: Processes and
measurement
For precipitation to occur, you must have water vapor
Saturation vapor pressure over water
The amount of water vapor that the
atmosphere can carry (the saturation
vapor pressure) is a function of
temperature; the saturation vapor
pressure does not increase linearly
with temperature but rather exponentially
through the Clausius-Clapeyron
equation
Having water vapor in the atmosphere is a necessary but insufficient condition for
precipitation. There must also be a mechanism to promote uplift, cooling and
condensation.
Precipitation requirements and processes
•Condensation occurs when the temperature is ≤ the dewpoint Td
•Without dust in the air to acts as cloud condensation nuclei (CCN),
the air can become supersaturated. You need dust greater than
10-4 mm to form cloud droplets
•Typical CCN = 0.2 μm
•Typical cloud droplets are 20-100 μm
•Rain drops are around 2000 μm - 2 mm
• Downward velocity must exceed uplift velocity for precipitation to fall.
Uplift velocities in convective storms can be large
Relative sizes of CCN, cloud droplets and
rain droplets
http://www.flame.org/~cdoswell/wxmod/wxmod.html
Precipitation requirements and processes (cont).
•Uplift and cooling to the dew point
•Condensation
•Growth of droplets (two basic processes)
1) Bergeron-Findeisen theory (top figure): If ice
and supercooled water droplets exist together in a
cloud, the ice crystals grow at the expense of water
droplets, the reason being that the saturation vapor
pressure over ice is lower than over water. Solid
precipitation that falls may (or may not) melt at a
lower, warmer atmospheric level to become rain.
2) Coalescence (bottom figure): Occurs in warm
(liquid) clouds. Falling droplets have terminal
velocities directly related to their diameter, such
that the larger falling drops overtake and absorb
smaller drops, the smaller drops can also be swept
into the wake of larger drops and be absorbed by
them.
Key processes that provide uplift, cooling and
condensation are convection, frontal uplift and
forced ascent over orographic barriers.
http://schepping.punt.nl/?a=2007-10
http://schepping.punt.nl/?a=2007-10
Bergeron-Findeisen theory (cont.)
Arapahoe Basin – May 2005
http://www.geminibv.nl/tidbits/dauwpunt-druken-temperatuur?set_language=en
http://www.gi.alaska.edu/alison/ALISON_Science_Snow.html
The process of vapor diffusion is most efficient at around -12 to -15oC,
corresponding to the temperature of largest difference between water and
ice saturation vapor pressures. Different combinations of temperature and
supersaturation determine the type of snowflake that forms.
Precipitation requirements and processes (cont).
Import of water vapor is important to maintain precipitation
Consider a thunderstorm
•For a column within the thunderstorm of unit area (m2),
the volume = 10,000 m3
•Typical average water mass is 0.5 g m-3
= 5000 g water in column
= 5000 cm3/10,000 cm2
= 0.5 cm precipitation, which is not much
Thus to maintain precipitation there must be an influx (entrainment)
of vapor-rich air to replace the water that falls from precipitation.
Convective Precipitation
http://www.ux1.eiu.edu/~cfjps/1400/stability.html
http://www.grc.k12.nf.ca/climatecanada/precipfactors.htm
Convective precipitation tends to be
localized. The typical convective
storm occurs when local surface
heating makes air parcels warmer
(less dense) than their environment,
such that they rise and condense.
Keys to convective storm development are low atmospheric stability and ample atmospheric
water vapor. The stronger the decrease in temperature with height (the more negative the
environmental lapse rate), and the more water vapor that is available, the more favorable the
conditions for convective storms.
Lapse rates and stability
Environmental lapse rate: The vertical rate of change
of temperature with height that we would measure
with a thermometer
• Is usually negative (temperature decreases with
height), the reason being that the atmosphere is
heated primarily from the surface and rising air
cools. An increase in temperature with height is
called an inversion.
• Typical values are 6.5oC km-1
Dry adiabiatic lapse rate (DALR) : The rate of cooling
of a rising parcel (a given mass) of air due to
expansion and doing work against the
environment.
• The DALR is 9.8oC km-1
Saturated (wet) adiabatic lapse rate: The lapse rate
after condensation occurs.
• Condensation releases latent heat, such that the
saturated adiabatic lapse rate is les than the
DALR.
• Value varies; a typical value is 5oC km-1
A rising air parcel cools and expands. If
the temperature of the parcel is greater
than (less than) the temperature of the
surrounding environment, it is less (more)
dense than the environmental air and will
continue to rise (it will fall). If condensation
occurs, then the latent heat release makes
the parcel warmer (hence less dense) than
it would otherwise be, making it more likely
that it will stay warmer than the surrounding
environment and continue to rise.
Condensation may result in precipitation.
Key: If the environmental lapse rate
exceeds the DALR, the situation is unstable
and the parcel will rise.
Tropical Cyclones
NOAA composite satellite image of Hurricane Katrina
Dingman 2002 Figure 4-4
Tropical cyclones feed on evaporation and latent heat release. They form between
5-20o latitude in each hemisphere where the sea surface temperature is at least
27oC. The grow through a process called CISK (Conditional Instability of the Second
Kind) through which rising air, condensation and latent heat release results in a drop in
surface pressure, causing inflow (converging winds) and more evaporation, further latent
heat release, and a further drop in surface pressure. Rainfall totals can be impressive.
http://earthobservatory.nasa.gov/IOTD/view.php?id=7079
Frontal Precipitation
Frontal precipitation is associated with
travelling extratropical cyclones (low
pressure systems). They tend to be
strongest in winter. Precipitation is
associated with uplift, cooling and
condensation in warm fronts, cold fronts
and occluded fronts. Extratropical
cyclones should not be confused with
tropical cyclones (hurricanes, typhoons).
Dingman 2002, Figure 4-2
Surface weather analysis from the Canadian Meteorological
Center for Nov. 11, 2010 showing locations of cyclones and their
associated fronts, as well as anticyclones
Patterns of extratropical cyclone activity
Extratratropical cyclone frequency (left) and tracks (right) based on a detection and tracking
algorithm applied to data from the NCEP/NCAR reanalysis for the winter of 1989/1990.
(http://data.giss.nasa.gov/stormtracks/).
Sea level pressure pattern for October 24, 1997, showing
a memorable upslope storm that affected Boulder CO
With a cyclone centered over the Oklahoma panhandle, winds over the eastern plans of
Colorado have a component from east to west, moving “upslope”, promoting cooling,
condensation and precipitation. The system draws in water vapor from the Gulf of Mexico.
Orographic precipitation and chinooks
Because of the latent heat release during ascent,
the air at the top of the mountain barrier is warmer
that it would have been without the latent heat
release. As it descends on the leeward side, it
warms at the adiabiatic lapse rate, and arrives at
a higher temperature than it had before the ascent
process. These warm, leeside conditions are
often associated with strong gusty winds called
chinooks.
Orographic precipitation results from
ascent of air over a mountain barrier,
resulting in adiabiatic cooling,
condensation and precipitation on the
upwind side of the barrier. The leeward
side experiences a rain shadow. This
is in large part why the west slope of the
Front Range of Colorado receives
more precipitation than the plains to the
east. However, the plains can receive
considerable precipitation from “upslope”
storms when a low pressure system lies to
the south. These are most common
in winter.
Some notable temperature changes
linked to chinooks:
Loma, MT: -56oF to 49oF in 24 hours
Spearfish SD: -4oF to 45oF in 2 minutes!
Precipitation quality
Precipitation is not pure, but rather
contains a number of ions. The
straight lines indicate concentration
ranges for continental rain, while the
wavy lines indicate concentrations for
marine rain. Ranges found in high
pollution areas are indicated by
dashed lines.
Dingman 2002 Figure 4-53
Precipitation gauges
http://www.hubbardbrook.org/w6_tour/raingauge-stop/precipitation.htm
http://www.weatherworks.com.au/?p=3082
http://www.novalynx.com/260-2501.html
There are many types of
gauges, which vary in their
catch efficiency. Shielded
gauges, like the one above,
tend to perform better than
unshielded gauges. The liquid
water equivalent of snowfall
is especially hard to measure.
Wyoming snow gauge
Wyoming snow gauges are designed to provide accurate measurements
of snowfall water equivalent. The one pictured above is located on the
drainage divide of Imnaviat Creek, near Toolik Lake, on the North
Slope of Alaska.
Typical SNOTEL sites (summer)
SNOTEL (SNOwpack TELemetry) is an automated system to measure snowpack water
equivalent. There are over 600 SNOTEL sites across the western U.S. Snow pillows
measure the weight of the overlying snowpack. SNOTEL sites complement snowcourses,
where snow water equivalent is measured manually.
Snotel Sites
From Serreze et al., 2001
Areal or gridded estimates of precipitation
A problem often faced in hydrology is using point
measurements of precipitation to come up with a
regional average (such as for a watershed) or gridded
field of precipitation. This generally requires
interpolation. While there are many types of
interpolation, ranging from a simple average of all
stations within a given region to inverse distance
weighting to optimal interpolation, they pretty much all
boil down to the following:
Pinterp= Σ (Pi .wi)/wi
Where Pinterp is the regional or grid point value we wish
to interpolate to, Pi are measurements at each
precipitation gauge (as in the watershed shown at
right) and wi are the interpolation weights. Simply
phrased, the interpolated value is the sum of the station
values times the weights divided by the sum of the
weights. The key is how one determines the weights.
Dingman Figure 4-20
A practical note: If there is a
dense station network, all
interpolation techniques work well.
If the network is sparse, none of
them work well.
Station Density: Contiguous U.S
The station density from the U.S. cooperative network looks rather
dense, but at the scale of medium of small watersheds is can be
quite sparse. The country with the densest station network is Israel,
followed by the UK.
The Arctic: A very sparse network
3
The figure at left shows
monitoring stations north
of 40oN with at least ten
years of record for the
period 1960-1989. Note
the very sparse network
in high northern latitudes.
The situation over the
Antarctic continent is
pathetic. Coverage is
also very sparse over
large areas of the world’s
ocean
From Serreze and Barry, 2005
Getting the weights: Examples
One can develop a function based on the
distance decay of the correlation in precipitation.
(see the left hand figure). While attractive, the
problem is that the correlation length scale of
precipitation tends to be short, especially in
mountainous terrain. One can also make use of
relationships between precipitation and elevation
(hypsometric method). A problem is that in
mountainous terrain, relationships between
precipitation and elevation can be complex.
Dingman 2002 Figure 4-27
Dingman 2002 Figure 4-28
An alternative: The aerological approach
∂W/∂t = ET - P - ∇•Q
∂W/∂t = time change in
precipitable water
(column water vapor)
ET = Evapotranspiration
∇•Q
∂W/∂t
P = Precipitation
∇•Q = Vertically-integrated vapor
flux divergence
Rearrange:
P- ET = - ∇•Q - ∂W/∂t
P
ET
Key: While P and ET may be hard to measure over large areas, we can get
the net precipitation (P-ET) using atmospheric winds and humidities
Annual P-ET from the aerological approach
Aerological estimates of mean annual precipitation minus evapotranspiration (P-ET)
based on NCEP/NCAR data for the period 1970-1999 (mm) for the region north of
60oN. Contours are at every 100 mm up to 500 m (negative values dashed) and at
every 200 mm for amounts of 600 mm and higher [from Serreze and Barry, 2005].
Annual P-ET is positive everywhere except locally in the Norwegian Sea and
southern Barents Sea where evaporation rates are high from autumn through spring
(because of cold, dry air blowing over a fairly warm open ocean).