T.O. 6.4.1
A) Describe and assess the stability of atmospheric soundings: local parcel stability;
latent instability; potential instability; differential advection causing changes in
stability
Local Parcel instability:

Most commonly assessed for deep convection

Air parcel is lifted from the surface (or from a level aloft, important for elevated convection); we
assume: 1) its pressure adjusts instantly to that of the environmental air; 2) no heat exchange
between the parcel and ambient air (adiabatic or isentropic). If air is unsaturated, its temperature
decreases with height at the dry adiabatic lapse rate; If air is saturated, its temperature decreases
with wet adiabatic lapse rate (smaller than that of dry air, due to the latent heat release)

At a level, Tparcel > Tenvironment  unstable; Tparcel < Tenvironment  stable; easily seen on a
tephigram. Usually we compare the environmental lapse rate with dry adiabatic rate and wet
adiabatic rate:

 < w  absolute stable (flat or slow Tenv decrease, so parcel always cooler)

 > d  absolute unstable (steep or fast Tenv decrease, so parcel always warmer)

w <  < d  conditional unstable

 = w (or  = d )  neutral equilibrium for saturated (unsaturated) air
Latent instability

For an air parcel, sufficient lifting is required to achieve convection, i.e., it can acquire a
temperature warmer than that of the environment at some point above the layer in which it
originates, it is said to possess latent instability (LI)

LI is realized by the release of latent heat when the available moisture condenses during the
parcel's ascent.
Tc = convective temperature

1)
2)
Two ways to diagnoses:
From the latent instability analysis perspective:
Draw the lowest value wet adiabat (LVWA) tangent to the environmental temperature curve. If
this line intersects the environmental wet bulb temperature curve at some lower level, all parcels
in that layer with w values to the right of the LVWA have latent instability.
From the latent instability layer perspective:
1


If pseudo-adiabats through some point on the environmental wet bulb temperature curve crosses
the environmental temperature curve at higher level, then the point is in an unstable layer and
has latent instability.
LI occurs most frequently in environments with steep lapse rates aloft and high dew points in
lower levels; favorable condition for thunderstorm development
To release LI, a lifting is needed: frontal, topographic, synoptic lift (PVA) etc
Potential instability

A layer of air, if given sufficient ascent, the layer will acquire an unstable lapse rate, then the
layer is said to have potential instability (PI). Important for Acc and elevated thunderstorms

Diagnosis: on the tephigram, compare the lapse rate of Tw curve of the environment to the moist
adiabats. If the environment Tw decreases with height faster than the moist adiabats, then this
layer possesses PI

To realize PI, lift needed by upslope of terrain, synoptic lift etc:
1) For dry air masses, ascent  dry adiabatic cooling of whole layer  vertical stretching of
the layer (p between the top and bottom keeps a constant) steeper lapse rate  less
stable layer
2)
For saturated air (or air that becomes saturated through ascent), differences in humidity of
the layer affect its stability after lift (moister air cools slower), so that layers that are moist at
the bottom will become more unstable than uniform or "top moist" layers
Differential advection causing changes in stability

Warm advection increases with height (more warming at high level)  stability increase

Cold advection increases with height (more cooling at high level)  stability decrease

During night, Radiative cooling of the cloud top  unstable

During daytime, surface heating  unstable

Diagnostics: (use hodograph)
1) warm advection  veering wind with height; cold advection  backing winds with height
2
2)
By comparing the thermal advective rates between two different levels in hodograph, one
can derive the changes in lapse rate and hence, vertical stability.
B) Describe and interpret the vertical wind shear properties of atmospheric
soundings using the hodograph: Vertical wind shear; Storm motion and storm
relative winds; Hodograph curvature
Vertical wind shear

Vertical wind shear is the change of the horizontal wind between two discrete levels, is a vector
related to the horizontal T gradient (thermal wind relationship)

In hodograph, wind vectors at different heights are plotted in a polar coordinate; wind shear is a
vector joining the tips of two wind vectors (pointing to the wind vector tip of a higher level)

Total wind shear over a particular depth is calculated by adding up the entire length of the
segments of the hodograph

Hodograph shape refers to the degree of both speed and directional wind shear and how the
hodograph curves with height (Hodograph shape is very important for TS structure and evolution)

Importance of vertical wind shear:
1) weak wind shear: no updraft tilting, single-cell short-lived storms, strength mainly depends
on CAPE, little new-cell regeneration along gust fronts
2) moderate low-level shear (below 3km), 20kt for new cell generation in downshear side,
expect slightly longer-lived cells/clusters with cell regeneration.
3) strong low-level shear: 30kt shear below 700mb and weak shear above, expect persistent
multi-cell storms with intensity modulated by CAPE.
4) deep wind shear: ~60kt shear below 500mb, ~40kt below 700mb, expect supercells with
enough CAPE
Storm motion

During the early stages, storm generally moves along in the direction and nearly the speed of the
mean wind of SFC - 6km.
1) For straight hodograph, the mean wind use the mid-point of the hodograph below 6km.
2) For a curved hodograph, visually move the origin of the hodograph to the point representing
the SFC wind with the x' axis passing through the 6km wind. Then average all new x' and y'
components of the wind to estimate the mean wind
Storm relative wind

Storm relative wind is "the winds as seen by the storm", i.e., in the storm's frame of reference its
motion is zero. The storm relative winds are simulated by moving the hodograph origin to the tip
of storm motion vector and replotting the wind vectors in the storm's frame of reference.

The wind shear in the storm-relative coordinate will determine how the structure of the storm will
evolve. For a curved hodograph with storm motion off the wind shear vector the storm relative
winds and wind shear are increased.
Hodograph curvature
3


Hodograph curvature refers to the degree of both speed and directional wind shear and how the
resulting hodograph may curve with height. Whether the hodograph is relatively straight or has
clockwise or counterclockwise curvature has implications for thunderstorm evolution.
Assessment of hodograph curvature allows the forecaster to anticipate the motion and type of
storms that may be expected to develop.
C) Explain how convection is dependent on factors such as insolation, moisture,
winds, stability and local geography;
Insolation:
1) provides energy to lift parcels and trigger convection
2) During the day, the insolation heats the ground and evaporates water into the air. The air
near the ground becomes warmer and moister than the air above it  positively buoyant and
statically unstable  upward vertical acceleration of an air parcels rise from the surface
layer
3) Lift air parcel to LCL; if the surface air temperature reaches to convective temperature (Tc)
 set off free convection even without mechanical lift (in such case, convective
condensation level (CCL) = LFC = LCL)

LFC – level of free convection

LCL – lifted condensation level

CCL – convective condensation level
Moisture:
1) moist air is more buoyant than dry air as water vapor is lighter than air, leading to stronger
updraft in the convection
2) An increased amount of water vapor in the lowest layer of the atmosphere results in
increased potential instability
3) For an air parcel, more moisture means  more latent instability
4) Larger moisture increases CAPE and thus the strength of the convection
5) Saturated air parcel cools more slowly than dry air due to latent heat release that supplies
energy for continuous convection
Winds:


horizontal wind convergence produces upward motion, initiate convection
Vertical wind shear:
1) Low-level wind shear important for cold pool interactions and new cell generation
(multicells). In lowest 3km or so can lead to new cell development on downshear side of
cold pool (>=20kts of shear). In some cases low-level shear can lead to organized
multicell storms (>=30kt of shear)
2) Deep shear important for tilting updraft prolonging storm longevity. For supercells, deep
shear necessary for generation of mid-level rotation and vertical perturbation pressure
gradients. In lowest 6 km important for tilting storm updraft prolonging lifetime. In some
cases interactions with updraft can lead to supercell storms (>=30kt of shear)
Stability:

need instability to release latent energy and to characterize updraft strength and modulates
storm intensity

Positive buoyancy (hence, convection) is a result of instability and moisture

Potential instability is a necessary condition for development of thunderstorms. Ordinary
conditional instability, based on the lapse rate of temperature with height, is insufficient. If the
atmosphere becomes unstable in term of the lapse rate, convection develops on a scale too
small to produce a cumulonimbus.
4

If an inversion caps a humid layer near the surface, energy of instability is not released until
higher values of potential instability are reached. Then the convection develops more
vigorously and may form thunderstorms.
Local geography: Provide triggers for convection

Upslope of terrain forces air parcel to rise and lift air parcel to its LFC

Terrain induced mountain-valley breeze, sea-land breeze, barrier jet, sea-breeze fronts and
jet change wind shear and advect moisture

regions of excess surface heating

Provide moisture for convection: Lakes, oceans are moisture sources of convections
D) State and distinguish between the mechanisms by which latent instability and
potential instability are realized
LI



The physical importance of a layer of air possessing LI is that although a layer is locally stable
(the layer may be quite thick as well), parcels of air within it can, if given sufficient lift, become
positively buoyant with respect to their new local environment. LI is realized by the release of
latent heat when the available moisture condenses during the parcel's ascent.
Latent instability is favoured by steep lapse rates aloft and high moisture contents in the low
levels.
Mechanisms to realize LI include:
1) In summer, cloud top cooling, cold pools in mid levels, and Surface heating. For convection
to occur through a deep layer in the atmosphere (e.g., severe convection), the layer of latent
instability should be based near the surface and be of significant depth.
2) In winter, cold air over open water, cloud top cooling, and mid-level steep lapse rates above
cloud layers and a trend in destabilization.
PI:





A layer is said to have PI if, given sufficient lift, the layer will acquire an unstable lapse rate
LI deals with parcel ascent through local lift, while PI tells the potential for decreasing stability
due to large-scale ascent of a layer or of an entire air mass
The lifting mechanisms are required for realizing PI include forced ascent over rising terrain,
PVA ahead of an upper trof, or divergence associated with high level jets.
LI refers to a parcel’s buoyancy relative to the lapse rate of the environment above it. PI looks at
the stability of a particular layer after air mass lift, regardless of the sounding above that layer. LI
requires cold air aloft, PI does not.
However that if there is no potential instability in any part of a sounding there can be no latent
instability. Parcels undergoing finite vertical motions may realize LI but a small scale local lifting
mechanism is required to supply the escape energy. Air mass layers undergoing general lift may
realize PI but sufficient air mass lift is required.
E) Explain latent and potential instability are important to convection



LI and PI indicate that there are available instability exist for convection development, once the
vertical motion of a parcel or a layer is forced upon it, either by local or large-scale synoptic lift,
the instability will be realized for convection.
PI is most commonly found in layers aloft and is often correlated with the presence of ACC or
elevated Cb (this can sometimes include the persistence of nocturnal thunderstorms) assuming
sufficient moisture is present to achieve moist convection.
LI associated with steep lapse rates aloft and high dew points in lower levels favor large positive
buoyant energies and strong updraft, favorable conditions for thunderstorm development. Once
SFC heating allows the convective temperature to be reached, the LI will be realized and
thunderstorm development will depend primarily on the degree of local parcel instability above
LFC.
5


PI is a necessary condition for development of thunderstorms. Ordinary conditional instability,
based on the lapse rate of temperature with height, is insufficient. If the atmosphere becomes
unstable in term of the lapse rate, convection develops on a scale too small to produce a
cumulonimbus.
If an inversion caps a humid layer near the surface, the situation is different. Energy of instability
is not released until higher values of potential instability are reached. Then the convection
develops more vigorously and may form thunderstorms.
6
T.O. 6.4.2
A) Describe the buoyancy processes related to: Updrafts (including Bubble
convection, CAPE, Estimate of updraft strength, CIN, Entrainment); Downdrafts
(including Evaporative cooling, Precipitation loading, Downdraft CAPE, Estimate
of downdraft strength)
Updrafts

Bubble convection:
1) Bubble theory views the active convective elements in a cloud as discrete "bubbles". As a
bubble rises, turbulence mixes environmental air into its interior.
2) The earliest saturated bubbles (thermals) can only penetrate a limited distance upward
because the mixing of environmental air into the bubble reduces the bubble's buoyancy.
With sufficient mixing, the bubble loses its identity and is totally mixed into the
environment. This is known as parcel detrainment.
3) The erosion of these first bubbles can result in a warming and moistening of the local
environment through which they ascend. Clouds may form and dissipate as bubbles rise and
detrain into the environment.
4) Later bubbles, produced by continued heating of the ground, following the same ascent path
will encounter this "enriched" environment and will not erode until encountering drier and
cooler air at higher heights.
5) Bubble convection views the growing cumulus cloud as the result of a train of bubbles rising
through the same channel; each successive bubble rises to greater heights before eroding.
6) Whether or not cumulus clouds will persist and grow to greater heights depends on the net
effect of the rate of enrichment of the environment by the rising bubbles and the rate of
drying by the environment.
7) There are drag forces which work against bubble ascent. These include:
 Loading by Cloud Water, the downward force due to the weight of suspended cloud liquid
water decreases the buoyancy of the cloud bubble.
 Turbulent Drag is the retarding influence of frictional (fluid viscosity) and turbulent
forces. Momentum exchange during entrainment of environmental air which is at rest
represents a major drag force.

CAPE (Convective Available Potential Energy)
1) On a tephi, it refers to the positive area between a lifted parcel T curve and the
environmental T curve
2) Integrated from the LFC to the Equilibrium level (EL)
3) Gives estimate of potential updraft strength
4)
The strongest updraft occurs at the level where the parcel is its warmest with respect to the
environmental air, so a "fat" CAPE indicates stronger vertical accelerations than the skinny"
CAPE, thus produce more hazardous weather

Estimate of updraft strength
1) Assume all CAPE are converted to kinetic energy, then the maximum updraft velocity for a
2)

parcel is Wmax  2CAPE
This is the upper limit to the updraft velocity arising from buoyancy, because the negative
buoyancy effects by entrainment and water loading. Are not considered
CIN (convection inhibition)
1) To release latent instability a parcel must reach its LFC, CIN measures the amount of energy
required to reach level free convection (LFC), it prevents convective initiation.
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2)

On a tephigram, CIN is the “negative area” between lifted parcel T curve and the
environmental T curve integrated between LCL and LFC. Small value of CIN means less lift
required to achieve LFC. CIN values typically increase when the static stability above the
unstable layer increases.
Entrainment
Entrainment refers to mixing of environmental air into the parcel, and can occur both laterally
and at top of updraft. A strong updraft will minimize the amount of entrainment. Entrainment
depends on:
1) Thermal and moisture profile of environment
2) Size of the convective cloud (small clouds  more entrainment) as entrainment occurs
mainly at the periphery of the updraft (large cloud protects updraft core)

Impacts of environment on parcel
1) Environmental air is cooler and drier than a rising saturated parcel
2) Cooler air will lower the sensible heat of the rising parcel ( +ve buoyancy)
3) A “drier” parcel is less buoyant than a moist one ( +ve buoyancy)
4) Drier air leads to cooling as cloud water evaporates in a saturated parcel ( +ve
buoyancy)

Results of entrainment
1) Parcel cools at a greater rate than that predicted by parcel theory
2) As parcel “dries” it becomes more negatively buoyant
3) Net result is an updraft with less buoyant energy (CAPE) and increased negative
buoyancy effects impeding upward accelerations
4) Storms tops are lower than estimated and updrafts are weaker
5) Important for descending parcels (downdraft)

** Summary of updraft
1) when wind shear is weak, buoyancy processes are the dominant control on convective
updrafts. As positive buoyancy increases, upward vertical acceleration increases.
2) buoyant energy arises from temperature differences between a lifted parcel and its
environment. Buoyant energy is maximized in environments characterized by steep midlevel temperature lapse rates (conditional instability) and sufficient moisture within or
below the buoyant layer.
8
3)
4)
5)
6)
positive buoyancy (upward) through a vertical column can be estimated quantitatively
by integrating the parcel excess temperatures between the LFC and the equilibrium
level. This is called CAPE or buoyant energy. Other measures of updraft strength
include various indices such as the Lifted Index and Showalter Index.
unstable parcels must be lifted to their Level of Free Convection (LFC) to release
buoyant energy and produce an updraft.
Convective Inhibition (CIN) is a measure of the amount of lift (energy) required to reach
the LFC. As the CIN increases, more lifting energy is required.
Solar heating is a frequent source of energy for lifting parcels to their LFC. However,
there are many other sources of lift that must be considered before rendering a
convective forecast.
Downdrafts

Why important?
1) Downdraft can cause damaging surface winds
2) Associated with microbursts
3) Significant impact to aviation community
4) Responsible for “cold pool” that is important for new convective cell development
and longevity of multicell storms

Downdraft is enhanced by
1) evaporation cooling
2) precipitation loading
3) momentum transfer

Evaporative cooling: Evaporation of cloud water in an unsaturated layer cools the air, and
produces negative buoyancy to air parcels, and enhances downdraft speeds; enhanced by:
1) availability and entrainment of dry environmental air (increase the evaporation rate)
2) steep environmental lapse rate air (acts to maintain negative buoyancy as the parcel
descends)
3) large liquid water content (provides more precipitation for evaporation)
4) small drop size (Smaller droplets tend to evaporate more quickly than larger drops)

Precipitation loading
1) Downdraft strength is increased by the drag effect of liquid water. Liquid water loading
adds weight to a column, increasing the downward force of gravity. This force
eventually dominates the updraft and initiates the downdraft, usually in the mid-levels of
the storm. As the amount of liquid water in the column increases, water loading
increases, resulting in potentially stronger downdrafts.
2) In general, a stronger updraft will have stronger downdrafts due to the effect of
precipitation loading.
3) Operationally, air-masses with high liquid water content and high buoyant energy are
particularly susceptible to strong downdrafts, particularly when dry air is found at midlevels.

Momentum transfer
Downward transfer of stronger winds aloft towards the surface can also enhance outflow
winds. Storms developing in stronger environmental winds may produce stronger outflows
as a result of downward horizontal momentum transfer.

Downdraft CAPE (dCAPE)
9
1)
2)
3)

Downdraft CAPE (dCAPE) is a theoretical estimate of the max. increase in kinetic energy
available to a downdraft parcel due to negative buoyancy accelerations
On Tephi, it is proportional to the area bounded by a downdraft pseudoadiabat and the
environmental temperature curve below the level at which the downdraft originates. The
level to use as the origin of the downdraft usually is the lowest valued moist adiabat
intersecting the environmental wet bulb curve.
Larger dCAPE implies stronger downdrafts (and outflow winds), and maximized for cold
dry air aloft and warm moist air near the surface.
Estimate of downdraft strength
1) An upper limit to the maximum velocity of the downdraft parcel when it reaches the SFC
2)
3)
can be found using Wmax  2dCAPE
Operationally, look for soundings exhibiting dry air at mid-levels and a warm, moist
boundary layer
The strength and depth of the cold pool at the SFC will be related to the temperature
difference between the downdraft parcel and the ambient air at the surface. High dCAPE
values imply larger negative buoyancy and deeper cold pools.
** Steep lapse rate air contributes to both stronger updrafts and stronger downdrafts (why?)

For updraft, steep environmental lapse rate  fast environment T decrease  larger (Tparcel
– Tenv)  larger positive buoyancy  stronger updraft

For downdraft, buoyancy term doesn’t help explain; but note
1)
stronger updraft caused by steep lapse rate  precipitation  precip. loading
2)
steep lapse rate  cool & drier mid-level air  enhanced evaporation cooling
Both contribute to stronger downdrafts.
10
B) How cold pool and low level shear interactions control storm evolution and
structure





Importance
1) Cold pool and low-level shear interactions responsible for longer lived storms where new
updrafts develop on a preferred flank of the storm
2) Dominant process controlling long-lived mesoscale convective systems (multicell storms)
that can produce widespread severe weather (wind and hail damage)
Cold pool
1) downdraft within storm reaches SFC; cold air spreads out in all directions
2) leading edge acts like cold front providing lift (gust front)
3) horizontal vorticity roll enhances lift over denser cold air
4) If parcels can reach LFC new cells develop
Interaction between cold pool and low-level vertical shear
1) environmental horizontal vorticity is created by shear
2) cold pool vorticity interacts with environmental vorticity
3) upshear side suppressed vertical motion and cell advected over cold pool
4) lift favoured on downshear side of cold pool and cell advected in direction of gust front
5) strength of shear impacts new cell development
Factors for New Cell Development
 Strength of the wind shear vs. that of the new updraft will influence whether cell can
survive
 Speed of gust front and new cell motion a critical factor: Gust front too fast means cell
end up over cold pool and is short-lived; assessing low-level shear will alert you to the
possibility
1) Weak low level wind shear

wind speed above cold pool significantly less than speed of outflow boundary

outflow boundary undercuts updraft of new cell

no new cell development downshear
2) Moderate low level wind shear

wind speed above cold pool and outflow boundary speeds the same

continuous updraft under new cell with no net shear

new cell development downshear

new multicell storms sustain themselves
3) strong low level wind shear

wind speed above cold pool faster than speed of outflow boundary

Enhanced lift on downshear side of pool but cell is “shredded” by wind shear

no new cell development downshear
Operationally

values of 0-3km wind shear of 20kt may be associated with new cell development on the
downshear side of the outflow boundary.

Values of 0-3km shear of 40kt can be associated with long-lived multicells or if combined
with large CAPE can result in bow-echoes, squall lines or derechoes
C) Describe how buoyancy and deep shear interactions control storm evolution and
structure including: Tilting; Hodograph curvature and storm evolution
2)
Tilting
1) In a deep shear environment (0-6km or 500mb), relative high pressure created on upshear
side of cloud; relative low-pressure created on downshear side of cloud
2) Horizontal pressure gradient causes tilting of the updraft (and cloud) tilted downshear
11
3)
4)
5)
6)
7)
3)
The degree of tilting will depend on the vertical velocity of the parcel and the strength of the
wind shear.
The tilting of updraft downshear allows the precipitation particles not falling right over the
updraft and allows the updraft to thrive
perturbation pressure gradient important for supercells
Strong shear will initially act to inhibit storm development. If updraft persists and is strong
enough stronger shear will generally lead to more organized storms. The balance between
the strength of an updraft and the environmental wind shear is critical in determining if a
convective cell will persist and how organized a resulting thunderstorm will be (i.e., singlecell, multicell, supercell).

This may be quantified via the Bulk Richardson Number: BRN = CAPE/WS.

BRN ~ 10-50 supercell; 50-350 multicell; > 350 airmass; too small, no storm
Tilting of the thunderstorm updraft is the first step in the development of a persistent
convective system. Further organization depends on the hodograph shape (strength,
direction and depth of the wind shear).
Hodograph curvature and storm evolution
Vertically wind shear  horizontal vorticity tubes (direction, perpendicular to the wind shear
vector; strength, proportional to the strength of WS)
 Horizontal vorticity is tilted into the vertical in the vicinity of the updraft; a vertical vorticity
couplet forms at mid-levels (~3-8km) of the incipient storm, with low pressure at the center of
rotation (regardless of cyclonic or anti-cyclonic, PGF = centrifugal force, cyclostrophic)
 Stretching of the vertical vorticity in the updraft increases vorticity to a maximum at midlevels and vertical perturbation pressure gradient develops below each of the rotational
centers
 Dynamically induced vertical pressure gradients beneath the rotation centers induce updrafts
on the left and right flanks of the storm. Further storm evolution depends on hodograph shape

Straight Hodograph (unidirectional shear)



Wind shear is unidirectional, the high pressure region on the upshear side extends through the
depth of the column with no vertical pressure gradient (vertically stacked).
This region of high pressure acts to separate the vertically stacked vorticity couplet.
Enhanced by the dynamically induced updrafts on both flanks of the storm the cell separates
into two distinct mirror image storms that begin to move away from one another. The rightmoving (left-moving) storm exhibits cyclonic (anti-cyclonic) rotation. This process is
sometimes aided by the presence of a precipitation-induced downdraft depending on the depth
of the vertical wind shear and liquid water content of the storm.
12

This idealized case for a straight hodograph is a pair of "mirror-image" storms which will
tend to move away from one another as they propagate to the right and left of the mean wind,
with continued development on the updraft flank of each storm. In reality it is much more
common that there is a preferred flank for new development with one storm favoured to
persist and strengthen
Curved Hodograph
Physics:





Wind shear is not unidirectional, high pressure that developed on the upshear side of the
storm rotates as the WS vector, so not vertically stacked
An upward directed perturbation pressure gradient may develop on a preferred (right) flank
of the storm. A rotating updraft (mid-level mesocyclone) is enhanced by both upward
directed pressure gradients (due to mid-level rotation and veering of the horizontal pressure
gradients with the wind shear vector); The left flank of the storm develops a downward
directed pressure gradient force counteracting the updraft induced by the anti-cyclonic
member of the mid-level vorticity couplet
The right flank is now strongly associated with the storm updraft and and the left flank is
associated with the downdraft. The storm begins to deviate to the right of the mean wind
The "right-motion" storm is further enhanced by
1) Contributions of streamwise vorticity to the updraft (resulting in an environment more
conducive to tornadogenesis)
2) Right deviation allows the storm relative inflow into the storm to avoid the main
precipitation area and maintain its buoyant properties
The supercell storm may last for a number of hours. If the cold pool becomes too strong the
cold air will spread out ahead of the updraft and the original supercell will dissipate. Other
factors may lead to regeneration of subsequent rotating updrafts.
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** Description for the whole processes:
1)
2)
3)
4)
Initially there is an updraft present on both flanks of the storm as in the straight hodograph case.
The updrafts are initially enhanced by the low pressure associated with the centers of the midlevel vorticity couplet.
Due to interactions between the developing storm and the environmental wind shear, an upwards
(downwards) directed perturbation pressure gradient develops on the right (left) flank of the
storm. The right flank is now strongly associated with the storm updraft and the left flank is
associated with the downdraft. The storm begins to deviate to the right of the mean wind. The
splitting of the storm as in a straight hodograph environment never really occurs as the left flank
quickly becomes dominated by the storm downdraft.
The combination of dynamic pressure effects due to the mid-level rotation and interaction of the
storm with the environmental wind shear act to strengthen the updraft and associated vertical
velocities beyond those attributed to buoyancy considerations alone. The rotation of the updraft
is enhanced through the stretching of vertical vorticity as the mid-level mesocyclone develops
The storm continues to propagate to the right of the mean winds as new development continues
on the right flank. The "right-motion" enhances the contributions of streamwise vorticity to the
updraft resulting in an environment more conducive to tornadogenesis. Right deviation also
allows the storm relative inflow into the storm to avoid the main precipitation area and maintain
its buoyant properties. The result is an intense, well-organized, and persistent storm capable of
producing large hail, devastating winds, heavy rain, and tornadoes.
** Storm relative helicity

Crosswise Vorticity
Storm relative inflow is perpendicular to the environmental wind shear induced horizontal
vorticity tubes

The main updraft is at the leading edge of the storm, no vorticity contribution to storm (as
horizontal vorticity is tilted into the vertical vorticity couplets which are located on the flanks
of the storm)


Streamwise Vorticity

The storm-relative inflow and the vorticity vector (arising from wind shear generated
horizontal vorticity) are parallel

"spiralling" inflow leading directly to updraft rotation when the inflow (and vorticity) is tilted
into the vertical
14


Storm relative helicity (SRH)

SRH is a measure of the degree of streamwise vorticity present within the inflow
environment of the storm, thus an indicator of the potential for updraft rotation

If crosswise vorticity dominates the storm relative inflow, SRH will be low and there is little
potential for updraft rotation; Conversely, if streamwise vorticity dominates then the
potential for updraft rotation is high

Usually the storm inflow layer is approximated by the 0-3 km layer and SRH is calculated
through this depth

SRH may be represented visually on a hodograph as negative twice the area bounded by the
hodograph and the SFC and 3km storm relative wind vectors
Development of low-level mesocyclone

The tilting and stretching of horizontal vorticity generated locally by the storm itself seems to
play a significant role in low-level mesocyclogenesis (** note tilting and stretching of
environmental horizontal vorticity result in mid-level mesocyclone)

In general, horizontal vorticity will be generated where there are horizontal gradients in
15
buoyancy and resulting vertical motions in cold pool. Within the near-supercell environment
such conditions exist behind the leading edge of the cold pool associated with the forward
flank downdraft (FFD) and rear flank downdraft (RFD)

Note that the vorticity vectors within the cold pool of the FFD generally point towards the
updraft; and the magnitude of these vectors is greater than the northwest-pointing vorticity
vectors (in "warm sector" of wave) associated with environmental wind shear. This
streamwise vorticity increases probability of updraft rotation, and the dynamically induced
upward pressure gradient force (located below the mid-level vorticity center) is able to lift
this air into the storm updraft. This mechanism of streamwise vorticity is thought to play a
major role in the development of the low-level mesocyclone.

Note that the vorticity vectors within the cold pool of the FFD generally point towards the
updraft; and the magnitude of these vectors is greater than the northwest-pointing vorticity
vectors (in "warm sector" of wave) associated with environmental wind shear. This
streamwise vorticity increases

In some cases, low-level rotation may become strong enough to generate a downwarddirected pressure gradient force below the mid-level mesocyclone. This small downdraft
enhancement can locally accelerate the RFD, strengthening the outflow winds wrapping
around the back of the storm. An occlusion of the updraft may occur with possible updraft
regeneration at the point of occlusion. This process is strongly linked to tornadogenesis.

Other mechanisms for tornadogenesis: when storm forms form near SFC lows or SFC
vorticity trofs, or as a storm propagates along a thermal boundary, dryline, or outflow
boundaries from other storms
** About CAPE and Shear
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
Initial stage
Updraft strength is maximized for increasing CAPE while updraft strength decreases for increasing
wind shear. In general, wind shear has a detrimental effect on initial cell growth.

Secondary stage
Updraft strength is maximized for increasing CAPE and intermediate values of wind shear. These
results highlight the enhanced ability of the cold pool to trigger new cells as wind shear increases

Supercell stage
The supercell updraft is maximized for stronger vertical wind shear but again weakens if the shear
is too strong. Interestingly, if the buoyancy is too strong, relative to the wind shear, the updraft
weakens due to the cold pool becoming too strong and moving out too quickly away from the
updraft.
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T.O.6.4.3
A) Describe the life cycle of a single cell thunderstorm, its characteristics, associated
weather, and typical pre-storm environment
Life cycle:

Initial Stage:
 updraft dominated from the surface to cloud top with speed of 10-20 m/s or more
 horizontal convergence maximized near the surface
 entrainment ongoing at cloud boundaries
 cell temperature exceeds the environmental temperature at each level
 positive temperature anomalies in the updraft increasing with time
 greatest concentration of hydrometeors is found at or above the freezing level
 no precipitation reaches the surface

Early mature stage:
updraft continues to feed moisture
precipitation (super-cooled water drops and ice-crystals) forms in upper levels
 anvil forms
 overshooting top and gravity wave may be present
Late mature stage:
 downdrafts initiated at mid-levels and increase in space and with time
 cold pool begins to develop, surface wind gusts develop
 maximum turbulence occurs in regions of greatest vertical motion,
 cloud grows physically, cloud tops (ice crystals present) reach temperatures < -40°C
 cold anomalies occur in the downdrafts, with a maximum in low levels
 warm anomalies still occur in the updrafts
 low-level precipitation is liquid (solid precipitation is possible)
 horizontal boundary of the surface precipitation roughly marks the downdraft boundary
Dissipating stage:
 cell is “collapsing”
 downdrafts spread through the entire cell
 downdraft speeds are less than in the mature stage
 surface wind divergence rapidly decreases




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cell T falls below that of the environment at each level (but eventually equal to Tenv)
ice crystal clouds (cirrus) blow off the cell tops
 shower activity is generally light and decreases
Characteristics:

Typically short-lived (< 1 hr); Small spatial scales (on the order 1km)
Associated weather:
 Server weather rare but not unheard of
 Rain, hail, damaging winds and small tornadoes can occur with strong type storms (with
large CAPE and weak shear)
Typical Pre-storm environment:

In a weakly sheared environment, buoyancy is the dominant physical process driving
thunderstorm evolution

Without shear the storms are vertically oriented

Precipitation has nowhere to go but to fall back through the updraft

Downdraft forms at mid levels due precipitation loading and forms cold pool at the
surface cutting off the storm inflow


B) Describe the life cycle of a multicell thunderstorm, its characteristics, associated
weather, and typical pre-storm environment; differentiate it from a non-severe
thunderstorm;
Multicell thunderstorm:

characterized by a high degree of organization with continued new cell growth on a preferred
flank of the cold pool (usually downshear with respect to 0-3km shear)

Organized Multicells result from interactions between cold pool and environmental shear in
the lowest few km (or orography, fronts, convergence lines or other boundaries)
Isolated multicell storm with mean wind vector in direction of shear vector
Life cycle and structure:

Mean wind (average wind of 0-6km) and low-level (0-3km) shear vector in same direction (cells
move in same direction as storm system)

The growth of new cells occurs on the downshear side of the cold pool. The circulation at the
leading edge of the cold pool is balanced by the environmental horizontal vorticity produced by
low-level wind shear. This results in a persistent “vertical jet” along the leading edge of the cold
pool lifting unstable parcels to their LFC. Cold pool needs to be deep enough and near speed of
winds at cloud base, LFC needs to be near depth of cold pool

The upper portion of precipitation core is positioned slightly downshear, called an echo
“overhang.” This overhang is caused by the differential advection of precipitation by the vertical
wind shear and by the strong lifting force near the leading edge of the cold pool.
19





The area below the overhang is called a “weak echo region” (WER). Often, an overshooting top
is observed above the WER, further implying the existence of a strong updraft in that location
and of an organized storm
Updrafts are created at the leading edge of the spreading cold pool (gust front). Parcels are lifted
and move rear-ward relative to the system with the anvil developing aloft towards the upshear
side, well behind the leading edge of the cold pool.
The downdraft develops at mid-levels within the main precipitation core due to precipitation
loading and evaporation. Downdrafts can be very strong and persistent, fed by a constant supply
of moisture in the neighboring updraft.
The updraft and downdraft remain separated but interact constructively, and can maintain a
strong circulation for hours.
If the gust front has spread out substantially, and the layer of cool air spreads out and becomes
shallower, much of the environmental air is not lifted to its LFC, new cells are no longer likely to
form, the whole multicell system dissipates
Characteristics:

Continual regeneration of cells near the leading edge of the cold pool leads to a lifetime of many
hours

Much larger than single cell storms (on the order of 100-400km)
Associated weather:

Severe weather not uncommon including damaging winds, widespread large hail, heavy rain
(possible flooding if slow moving)
Isolated multicell storm with mean wind direction different than shear vector

More common wind shear is in different direction than mean wind vector

New cells develop on the downshear side (0-3km) of the cold pool but move in a different
direction than the system

Later a new cell will develop in the location of the previous one
Typical pre-storm environment:

The primary process responsible for organizing and sustaining multicell systems is the
interaction of low-level vertical wind shear with the circulation produced by storm-generated
cold pools.
20


When low-level wind shear is sufficient, along with presence of positive buoyancy (CAPE),
regular cell growth will be favored along the down-shear side of the cold pool.
The interaction of the gust front with other features such as orography, frontal boundaries,
convergence lines, or other outflows, also can foster the regeneration of new cells on preferred
flanks.
Differentiate it from a non-severe thunderstorm: (bold + italic denotes multicell TS)

Life time: less than 1 hour; several hours

scale: ~1km; 100-400km

pre-storm environment: weak wind shear and buoyancy is the dominant physical process; lowlevel vertical wind shear and presence of positive buoyancy, interaction between wind shear
and cold pool is dominant physics)

weather: Hail, damaging winds and small tornadoes can occur with strong pulse type storms;
Severe weather occurs as downbursts, moderate size hail, flash floods, and weak tornadoes.
** The dynamics of a convective storm depend upon the environmental CAPE, and shear. When the environmental
vertical shear is relatively weak, only the CAPE and density-current behavior are dynamically significant, and ordinary
cells occur. When the low-level vertical shear is larger and the CAPE is not too small, a succession of ordinary cells, a
multicell storm, can form. If the CAPE is high and the vertical shear is strong, then vertical perturbation pressuregradient forces owing to rotation become dynamically important, and supercell forms.
*The ‘bulk’ Richardson number R, the ratio of CAPE to a quantity proportional to the square of the mean vertical shear
integrated over height (S^2) as an indicator of storm type: R=CAPE/S^2 Where S^2=1/2(u6000m-u500m)^2.
** The behavior of convective storms is highly sensitive to the environmental CAPE and vertical shear, forecasters
must be careful to account for changes in environmental CAPE and vertical shear. On the mesoscale, the shear can
change owing to the passage of shortwaves of jet streaks or fronts and the response of the boundary layer winds to
developing surface cyclones or to mountain-valley or sea-breeze circulations. On the storm scale, shear can change
owing to the environmental response to nearby convection. CAPE can change owing to changes in cloud cover and
differential advection of temperature and moisture.
C) Briefly describe a squall line, a mesoscale convective system (MCS), and a
mesoscale convective complex (MCC)
Squall line:
Thunderstorms are often observed in organized lines, called squall lines
Its structure is irregular, with activity dominated by a few large intense storms and numerous
small, relatively weak ones. Unique phenomena such as bow echoes, line echo wave patterns
(LEWPs), supercells etc. can be found embedded in a squall line.

usually coincident with the line of greatest instability and often organized by large-scale forcing
such as frontal convergence.

Motion of squall line is a result of the advection of cells within the line by the mean wind and the
propagation of the line itself due to the triggering of new cells along its length. For squall lines
larger than about 200 km in length, motion will be perpendicular to its original orientation. For
smaller squall lines, motion is usually in the direction of the low-level shear vector with the
squall line oriented perpendicular to its motion. Local changes in CAPE, shear, or LFC height
can affect line motion as well.
Factors controlling squall line structure and evolution:

The component of the low-level shear vector perpendicular to the orientation of the squall line
itself (called line-normal shear). As the line normal shear increases, organization, longevity and
intensity of system downshear increases.

The size and depth of the cold pool affects system evolution. Strong pools require more shear to
sustain squall lines (to produce deep lifting at the leading edge of the cold pool)

The height of the LFC can influence system strength and longevity. High LFC environments
require more lifting than lower LFC enviromnets. Hence, for a given shear and cold pool
strength, a low-LFC environment is more likely to sustain stronger and longer-lived system.


21
Life cycle and structure:
With weak to moderate shear
Horizontal structure:

Initially, squall lines appear as an elongated band of convective cells, each producing their own
cold pool. Intense convection near leading edge of pool. Convergent along leading edge and
divergent under heaviest precipitation.

During the mature stage, the strongest precipitation echoes remain near the leading edge of the
cold pool and form a continuous line with strong convection along the leading edge. Stratiform
precipitation behind.

By the weakening stage, the convection along the outflow weakens, and recedes from the leading
edge. Cold pool spreads out ahead cutting off inflow. Stratiform precipitation may last for hours.
New line may be triggered out ahead.
Vertical structure: (with weak to moderate shear)

Early phase (1-3): cells mostly vertical

Mature phase (4-5): New cells form along cold pool and advect rearward, updraft tilted upshear
(rearward) feeding stratiform area. Rear inflow jet develops at mid-level, diverging at the
surface.

Dissipating Stage (6-7): The leading cells weaken and the updrafts becomes shallower and tilt
further rearward. Rear-inflow jet surfaces and cold pool spreads out ahead of updrafts
With moderate to strong shear
Horizontal structure:

The evolution time between the stages increases

The leading edge echo gradient is stronger with more intense cells taking on a bowed structure.

The precipitation line is narrower than in the weaker shear case with the intense echo gradient
remaining on the leading edge of the cold pool, implying a perfect balance between the strength
of the shear and cold pool.
22
Large swaths of significant wind damage are likely during the middle phase near the bulging
segments of echoes (bow echoes).
Vertical structure:

Persistent overhangs on downshear side

Updrafts remain strong and vertical and near leading edge of cold pool

As rear inflow jet eventually descends to SFC get sustained strong winds and gust front moves
out ahead

Mesoscale convective system (MCS)






MCS is a much larger convective ‘system’ produced by the interactions of individual cells
(single cells, multicells, supercells) which can appear as a disorganized mass of convective cells
or a highly organized convective line, such as a squall line.
MCSs can evolve from an isolated cell, a group of cells, or initiate rapidly along a linear feature
such as a cold front or surface trough.
MSCs most often develop during the day of evening and continue well into the night, feeding off
of elevated instability. MCSs often last for hours or over night.
All types of sever weather are possible with MCSs, especially damaging winds.
Most common and highly organized model of MCSs: squall lines, bow echoes, and mesoscale
convective complexes (MCCs).
Severity and organization of MCSs increases with increasing values of CAPE and low-level
wind shear.
Mesoscale convective complex (MCC)

MCC is a vast mesoscale convection system formed in such a manner that all the cells interact
and share the same anvil. MCC is often severe weather producers.
23



Presence of strong divergence of horizontal flow at the level of cirrus anvil and the top of the
outflow from the storm (like a tropical cyclone). MCC may rotate in a similar fashion to a
tropical storm.
Physical characteristics of a mesoscale convective complex based on analysis of enhanced
infrared satellite imagery
MCCs exhibit a nocturnal life cycle and are strongly linked with the nocturnal low level jet and
have been found to interact with and modify the larger scale environment.
Definition:
SIZE:
A: Cold cloud shield with continuously low IR temperatures (less than -32°C); area
greater than 100 000 km2.
B: Interior cloud region with temperatures less than -52°C must have an area greater
than 50,000 km2.
INITIATE:
Size definitions A and B are first satisfied.
DURATION:
Size definitions A and B must be met for a period greater than 6 hours.
MAXIMUM:
Contiguous cold cloud shield (IR temperatures less than -32°C) reaches its maximum
size.
SHAPE:
Eccentricity (minor axis/major axis) is greater than 0.7 at time of maximum extent.
Evolution:

Genesis: form late in the afternoon as several convective cells. The synoptic situation usually has
these cells in a region of vertical ascent in association with a short wave trough and pronounced
low level warm air advection.

Development: occurs in the evening hours and is associated with the greatest occurrence of
severe weather. The individual convective cells are at their maximum intensity with the cold
cloud tops grown to satisfy the MCC condition. It is during this time period that the nocturnal
low level jet reaches maximum intensity, enhancing both warm air advection and influx of moist
unstable air.

Mature: the mid to high level cloud shield reaches its maximum extent. As the MCC modifies
the environment, vertical wind shear begins to decrease. Convective cells continue to develop,
however their intensity diminishes reducing severe weather occurrences except for locally heavy
rainfalls.

Dissipation: The dissipation stage begins when convective cells no longer form, thus leaving the
precipitation as strictly stratiform and diminishing with time.




Typically, the first thunderstorms develop during the afternoon with convective complex organization appearing
in the evening.
Under favourable conditions, MCC's persist and reach maximum size around local midnight, that is, about the
time that the nocturnal low-level jet stream both reaches maximum strength and is located nearest the surface.
Thus, the nocturnal jet may act primarily in a large scale sense, providing the MCC its maximum influx of moist
static energy.
Moreover, mature complexes are often associated with upper-level mesoscale regions of pronounced cooling,
high pressure, and strong anticyclonic outflow.
Weather:
Form of local heavy rain, Flash flooding. MCC’s tend to move slowly giving the risk of high
rainfall amounts over a smaller area than with more rapidly moving thunderstorms.
Motion:
The motion of an MCC deviates from the environmental mean wind in the direction of the newly
forming cells. This is generally on the side with the moist low level warm inflow.
24
Bow Echoes
• 20-120km long
• Long-lived symmetric bow-shaped segment
• High CAPE environments (>2500J/kg)with strong low-level shear (0-3km > 30kts)
• Can be isolated or embedded in squall lines
• Long swaths of damaging straight-line winds
• Move in direction of low-level shear vector with speed proportional to strength and speed of
cold pool
Line echo wave pattern (LEWP)
• Long-lived portions of a squall line composed of bow-echoes and supercells
• Are a significant threat to life and property (damaging wind, heavy rain, hail and tornadoes)
Derechos

Large scale damaging wind events due to long-lived bow echoes and LEWPs

Usually associated with tornadoes
Criteria
1) numerous reports of wind damage or wind gusts > 26 m/s along a major axis length of at
least 400 km.
2) the reports must show a pattern of progression.
3) within the area, there must be at least three reports of F1 damage or gusts of at least 33
m/s separated by at least 40 nm.
4) no more than 3 hours between successive events.
Two types
1)
Progressive pattern: This is the most common type of derecho (75% of derechos). It is a
single bow-shaped system oriented perpendicular to the mean wind direction with the
bulge oriented in the direction of the mean flow. This bulging is where the damaging
wind events occur.
2)
Serial Pattern: This is a squall line associated with a strong, often rapidly moving, low
pressure system with a series of smaller bow echoes or LEWPs moving along the line.
The squall line is typically located in the warm sector of an extra-tropical cyclone, ahead
but parallel to the cold front.
25
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