Transcript Power Point - Tornado Forecasting Workshop

The Meteorology of Tornado
Forecasting:
Basic Concepts
Rich Thompson
SPC
Shut up!
I wanna play the
dryline!
Who am I?
STPs are
high along the warm
front
• Grew up in Houston, TX, with a life-long
interest in weather
• Meteorology graduate from OU
• Worked at NWSFO Houston Roger
1992-94
• Been with SELS/SPC since July 1994
• Lead Forecaster with SPC since 2000
• Rich
Storm chasing experience since 1985
Why am I here?
• Simple goal – I want to share what I’ve
learned over the course of my career,
before I forget everything!
• Want to help people make better forecasts
for severe thunderstorms:
– Appreciation for what goes into SPC forecasts
– Link meteorological concepts to aspects of course work
and operational meteorology
– Better storm chasing strategies
Workshop Organization
1. Observations, convective processes, and
synoptic processes:
 How to read weather maps and sounding diagrams
 “Q-G Theory”, jet streaks, fronts, cyclones
2. Severe thunderstorm ingredients:
 Moisture and lapse rate sources
3. Severe thunderstorm ingredients:
 Lift and vertical wind shear sources
4. Supercell and tornado conceptual models
 Reasons behind parameter development
Workshop Organization Cont’d
5.
6.
7.
8.
9.
Synoptic tornado patterns
Convective mode
Tornado parameter climatology
Numerical models and statistical approaches
Forecasting exercise
Map Plots of Observed Data
• Station model plots
• Contour analysis
• General rules to follow with contours for
various fields
– T, Td, height, wind
Surface Station Model
Wind Direction and Speed
Upper Air Station Model
Wind direction (from) and
speed (knots)
Temperature (C) at
pressure level
Dew point (C) at
pressure level
Height (m above sea level)
of pressure level
Rules for Contour Analysis
• Surface plots:
– Temperature (T) limits the dew point temperature (Td)
– Wind speed related to spacing of pressure contours
(pressure gradient), but wind usually crosses toward lower
pressure due to friction near the ground
• Upper air plots:
– Same limitations as surface with T and Td
– Wind largely parallel to “geopotential height” contours
– Geopotential height ≈ height above sea level of a pressure
surface
Raw SkewT-log P diagram
Constant
pressure
Constant
Temp
Dry adiabat
“skew”
Constant
Mixing
ratio
Features of note in SkewT log P
• Temperature is skewed about 90o from the dry
adiabats
• Pressure decreases as a logarithm of height (faster at
bottom than top)
• Mixing ratio crosses over temperature lines
– It’s a function of pressure, which is why the same dew point
temperature at higher elevation contributes more to buoyancy
• One thing missing is a plot of saturated parcel ascent
Saturated ascent (clouds) –
parcel cools at a rate of
~6.5 C km-1
(pseudo adiabatic “lapse rate”)
Lapse rates
between saturated
and unsaturated –
“conditionally
unstable”
Dry ascent (no clouds) –
parcel cools at a rate of
~9.8 C km-1
(dry adiabatic “lapse rate”)
Lifted Parcel (chunk of air)
• Begin at lifted parcel level (ground)
• Rise “dry adiabatically” until saturation
– Where dry adiabat crosses mixing ratio
– We call this the “lifting condensation level” or LCL
– First guess at cloud base
Moist adiabat
Illustration of lifted parcel
• From LCL, rise “moist adiabatically” to
“level of free convection” or LFC
– “Free convection” begins where lifted parcel becomes
warmer than environment
– Energy resisting lift below LFC is known as “convective
inhibition” or CIN
• From LFC, continue up to the “equilibrium
level” or EL. Accumulated area (energy) from
LFC to EL is known as CAPE
– “Overshoot” above EL
Common Sounding Terms
• Lapse rate – change in temperature with height
– Dry adiabat ≈ 9.8 C km-1
– Moist adiabat ≈ 6.5 C km-1
• Conditional instability – lapse rate between dry
and moist adiabatic
• LCL – lifting condensation level
• LFC – level of free convection
• EL – equilibrium level
• CAPE – buoyancy (positive area)
• CIN – convective inhibition (negative area)
Positive area
(CAPE)
Negative area
(CIN)
What’s the difference between the
lifted parcels?
• Virtual temperature accounts for moisture
– Warmer than measured temperature
– Makes most difference with tropical moisture
• Virtual temperature correction increases CAPE
and reduces CIN
• Which chunk of air to lift?
– Some sort of averaging is usually more representative
– Surface vs. “mixed layer” or “most unstable”
Lifted parcel
(virtual temp)
Lifted parcel
(non-virtual)
Surface and most
unstable parcel are
the same
Lowest 100 mb
mean parcel
Surface parcel
100 mb mean
parcel
Most unstable
parcel
What are we assuming?
• No mixing with environment (not true)
– “entrainment” usually reduces updraft strength from
expectations based on CAPE alone
• All rain falls out instantly (not true)
– Suspended rain particles reduces updraft strength
– That’s why we say “pseudo” adiabatic for saturated parcel
ascent
• “Parcel Theory” is a first guess at a
complicated process!
Always keep in mind what
we don’t know:
• Uncertainty in observations
– “Good” measurements?
– Do they represent what we’re trying to forecast?
• Unknown details with lifted parcels
– What is right layer to view?
– What assumptions are valid, and which might be terribly
wrong?
• Lots of room for error, but the concepts are
useful!
Simple vertical mixing
• Surface heating drives thermals and mixing,
which take heat and moisture both upward
(from surface) and downward (from aloft)
• Usually see surface dew point drop in
afternoon if not offset by moisture advection
(bringing in greater moisture from somewhere
else)
Surface heating drives
vertical mixing
“super adiabatic”
contact layer – lapse
rate steeper than
dry adiabatic drives
vertical mixing
Impact of ascent and moisture
advection
• See moist layer deepen faster than you would
expect with just surface heating and mixing
• “Deep” moist layer and horizontal moisture
advection both combat vertical mixing driven
by surface heating
– Can see moist layer deepen while dew points increase near
surface
Hodograph and Vertical Wind Shear
• Hodograph is the trace connecting the ends of
wind vectors with height
– Shows change in wind speed/direction with height on a
simple plot
• Vertical wind shear can be calculated from the
hodograph
– Vector wind differences and storm-relative winds
Plot tips of vectors with height
1 km
2 km
SFC
Wind Shear from Hodograph
Shear vector 1 km to 9 km
Shear vector SFC to 2 km
Storm-relative Winds
Horizontal vorticity vectors –
perpendicular to hodograph
Storm motion
Surface storm-relative
wind vector
Storm-relative Winds
Storm-relative helicity – area “swept
out” between storm motion and
hodograph
Sounding diagrams are used for…
• Moisture and temperature profiles
• Estimates of CAPE, CIN, Lifted Index, etc.
– Will storms form?
• Vertical wind shear
– What kind of storms will form?
• Many of your favorite thunderstorm
parameters are based in these diagrams, and
subject to the same errors and concerns!
Synoptic Meteorology
• Need to understand what “weather maps”
show
• Need to understand how different processes
are related
– Fronts and cyclogenesis
– “Q-G Theory” and vertical motion
• First, make sure you understand the data plots
and sources for errors!
Quality of Surface Observations?
Courtesy of Oklahoma Mesonet
Standard surface observations
OK mesonet observations at the same time
Dew point analysis for OK mesonet observations
What do these sites have in common?
68
63 68
73
Quality of Observations Aloft?
Understand the Data and Processes!
• Understanding the processes gives you a
sound way to interpret weather data, and
recognize errors
• If you don’t know what you’re using, how do
you know if you’re using it correctly?
– Must consider data quality
• Focus on observations!
The Meteorology of Tornado
Forecasting:
Synoptic Meteorology
Rich Thompson
SPC
Concepts in Synoptic Meteorology
• Synoptic pattern evolution:
– System movement and intensity
– Contributions to vertical motion and vertical wind shear
• Q-G Theory
• Fronts and Jet Streaks
• Vertical motion is atmosphere’s attempt to
restore “hydrostatic” and “geostrophic”
balance!
Geostrophic Wind
• Pressure gradient and Coriolis forces in
balance
– Winds parallel to height contours aloft
– Wind speeds proportional to contour spacing
• Vertical motion is the result of ageostrophic
flow and imbalances in the atmosphere
• Real atmosphere more closely resembles
gradient wind balance
Gradient Wind Balance
• Curvature in flow changes the balance of
forces, adding centrifugal accelerations
• Flow in a trough is subgeostrophic
– Results in divergence downstream from trough
• Flow in a ridge is supergeostrophic
– Results in convergence downstream from ridge
Gradient Wind < Geostrophic Wind
within a low/trough
Gradient Wind > Geostrophic Wind
within a ridge/high
Fastest flow in ridge
Slowest flow in trough
Example
Mass Continuity
• Divergence increasing with height – same as
convergence decreasing with height
Thickness and Thermal Advection
• Warming in layer results in expansion
– Thickness increases around level of max warming
– Density of column decreases
– Height rises above, height falls below
• Cooling in a layer results in contraction
– Thickness decreases around level of max cooling
– Density of column increases
– Height falls above, height rises below
Atmospheric Thickness
Max warming
Layer expands –
thickness increases
Atmospheric Thickness
Max cooling
Layer shrinks –
thickness decreases
Thermal Wind
• Not a real wind
– Relationship between temperature gradient and the
change in geostrophic winds with height
– Veering with height  warm advection
– Backing with height  cold advection
• Winds increase with height above a strong
temperature gradient
– Jet streams/streaks associated with enhanced belts of
~horizontal temperature gradient
Edge of stronger gradient near ground
Temp gradient slopes NW with height
“Vorticity” and “Advection”
• Vorticity = tendency for “spin”
– Think of paddle wheel example
• Advection = transport of an atmospheric
quantity
– Wind bringing air with different temperature,
moisture, or vorticity to a location
Advection Example
AVA
CVA
Quasi-Geostrophic (Q-G) Theory
• Assume mostly geostrophic flow
• Only consider advection by the geostrophic
wind
• Ageostrophic flow allowed as part of the
response to Q-G forcing
– Very simple version of real atmosphere
Q-G Height Tendency Equation
• Height change = Term 1 + Term 2
• Term 1 = geostrophic vorticity advection by
the geostrophic wind
– Referred to as vorticity advection (CVA or AVA)
– CVA leads to height falls
• Term 2 = Differential geostrophic thickness
advection
– Referred to as differential thermal advection
– Height rises above level of max warm advection
– Same idea as thickness change (hypsometric equation)
Q-G Height Tendency
• Height falls with CVA
• Max in thermal advection
– Warm advection max leads to height falls below
and height rises above
• Surface pressure falls with thickness increase above
– Cold advection max leads to height falls above and
height rises below
• Surface pressure rises with thickness decrease above
Q-G “Omega” Equation
• Vertical motion = Term 1 + Term 2
• Term 1 = differential advection of absolute
geostrophic vorticity by the geostrophic wind
– Referred to a differential vorticity advection
• Term 2 = thermal advection
• Both terms scaled by static stability
– Stronger response with steeper lapse rates
Q-G Vertical Motion
• Differential cyclonic vorticity advection (CVA)
– Rising motion where CVA increases with height
• CVA increasing with height = divergence increasing with height,
which is the same as ascent through mass continuity
– Sinking motion where CVA decreases with height
• Thermal advection
– Rising motion with warm advection
– Sinking motion with cold advection
– Easy to see with “isentropic” charts
Frontogenesis
• Strengthening of temperature gradient:
– Response of atmosphere is to weaken gradient
• Ascent on warm side of front (cooling)
• Descent on cool side of front (warming)
• Fronts are zones where thermal advection is
easily enhanced, and fronts are often
preferred corridors for cyclogenesis
– Cyclones tend to move along existing fronts
Jet Streaks with Straight Flow
• Explain patterns of rising and sinking motion
• Divide jet streak into regions
– Entrance (upstream) and exit (downstream) regions
– Left and right side of each region
• Use two physical explanations
– Q-G vertical motion equation
– Frontogenesis
– Both give same answer, which is good!
Q-G Vorticity Explanation
AVA
CVA
CVA
AVA
Jet Streaks and Q-G Theory
• Cyclonic and anticyclonic vorticity (spin) is
created on sides of jet streak
• From Q-G approach, differential CVA is the
same as differential divergence = ascent
• The simple graphic only shows CVA at one
level – assumes increasing with height below
level of strongest jet
Frontogenesis Explanation
Air entering jet
streak – sinking
on cold side,
rising on warm
side.
Air leaving jet
streak – rising
on cold side,
sinking on warm
side.
Jet Streaks and Frontogenesis
• Jet streaks are coincident with stronger
temperature gradients
• Air moving through a jet streak:
– Encounters strengthening gradient entrance region
– Encounters weakening gradient exit region
• Response to frontogenesis:
– Weaken gradient through sinking in cold air and rising in
warm air (entrance region)
– Opposite is true in exit region
Jet Streaks and Curved Flow
• Curved flow much more common
• Change from 4 regions (quadrants) to 2
regions as result of along-stream ageostrophic
flow (recall gradient wind balance):
– Ascent focused close to jet axis, with ascent extending into
right exit region for cyclonically curved jet
– Ascent tends to be stronger than with a straight jet streak
because more atmospheric adjustment necessary to
restore balance
“Baroclinic” Systems
• Temperature advection and system tilt with
height (upstream):
– Often deepening cyclones or rapid movement
– Differential advection leads to destabilization
• Warm advection corresponds to veering winds
with height (thermal wind):
– Large, clockwise-curved hodographs in warm sector
• Jet streaks and fronts:
– Strong speed shear aloft with fronts and jet streaks
Edge of stronger gradient near ground
“Equivalent Barotropic”
• Temperature contours parallel to height
contours
– Little temperature advection
– Usually steady or weakening cyclone
• Without large-scale temperature advection,
will see little veering of winds with height
– Usually weak low-level shear
– Unidirectional flow in warm sector
– Slow system movement
Tying it all together…
• Can explain why ascent occurs downstream
from troughs, along fronts, and in exit regions
of jets
• Areas of ascent are strongest where gradients
are strongest
• Patterns of ascent contribute to cyclone
strength and motion
• All relevant to setting the stage for severe
thunderstorms!
What’s up next?
Lake Charles
Tallahassee
Brownsville
Cancun
Veracruz
Helpful Links
• http://www2.mmm.ucar.edu/people/tomjr/fil
es/realtime/qgomega-usersguide.pdf
• https://www.meted.ucar.edu/bom/qgoe/
• http://www.crh.noaa.gov/images/lmk/QG_Th
eory_Review.pdf
• http://ww2010.atmos.uiuc.edu/%28Gh%29/g
uides/mtr/fw/grad.rxml