Regional and Mesoscale Meteorology Branch

2. Improve short term forecasts of convective initiation.
At the end of this session, you should be aware of the most recent theoretical work on boundary characteristics that modulate the initiation of convection. You should also be aware of several techniques that you can perform using AWIPS to apply these theories to short term forecasting of thunderstorms.

2. Limited to initiation directly resulting from surface boundaries.
We exclude mentioning initiation resulting from elevated convection caused by warm air ascending well behind a boundary such as a warm front. Surface-based convection forms from direct parcel ascent from near the surface (within 100’s of meters AGL). However, we consider only the surface-based convection that forms within 50 km or 40 minutes of an identifiable boundary rooted in the Planetary Boundary Layer (PBL).

3. Initiation mechanisms are considerations, not hard fact.
Convective initiation is one of the least observed and least understood phenomena in meteorology. What is presented to you in this session are the currently understood theories and considerations on what boundary characteristics are important in convective initiation. In the next couple years, the theories presented here may change significantly based on new modeling and observational studies. There is a planned field project called TIMEX (Thunderstorm Initiation Mobile Experiment) in 2002 that will sample boundaries with unprecedented detail. We should learn more from this one experiment than we know about convective initiation to date.

The TIMEX web page is at: www.nssl.noaa.gov/timex

1. Assessing CAPE/CIN
A. Be aware of how mesoscale ascent can radically change Mixed Layer CAPE.
B. Understand the concept of the vertical lifting distance to the LFC and its vertical profile. Understand what is a favorable vertical profile for convective initiation.

2. Assessing low-level convergence across a boundary.

3. Measuring depth of ascent zone in a boundary.
A. Depth relative to the LCL and LFC.

4. Estimating potential depth of ascent zone.
A. Boundary-relative steering-layer flow
B. Boundary-orthogonal shear across boundary.

5. Identifying along-boundary variations and interactions.
A. Boundary mergers, collisions.
B. Interactions with free convective cells, rolls.

Mention this is a table of contents on what we’ll be covering.

Take a look at sounding or model based CIN fields in the warm season and you will see vast areas of zero CIN and positive CAPE with no thunderstorms present. Clearly removing CIN is not a sufficient condition to initiate convection.

Considerations:

• Strong mesoscale updrafts can force air through large CIN. However, initiation not likely in zero CIN unless there is some forced updraft.
What is meant by forced updraft is a mesoscale region of ascending air that helps provide a support base for subsequent parcels to reach the LFC. This mesoscale updraft can arise from free atmospheric convective cells such as rolls, rayleigh convection. Mostly, boundaries, topography create the mesoscale ascent.

• Initial cumulus and TCU rarely follow the theoretical surface-based moist adiabat.
This will be shown in frame 7 where we examine where entrainment forces the parcel thermodynamics to deviate from the theoretical curve.

• Virtual temperature corrections have an uncertain impact on CAPE/CIN. The next slide shows an example of this.

This is not surprising since Mixed Layer CAPE uses a layer and moisture needs to be well mixed . What maybe surprising and a bit sobering is that MLCAPE can double or more depending on whether the sounding was launched inside out outside of a boundary. Even Horizontal Convective Rolls (HCRs) can lead to the same differences in CAPE. Evening soundings half a roughly 50/50 chance of being launched inside or outside an HCR updraft zone.

Surface air will be displaced upward at least 1.5 km

Air at higher initial levels also is forced upward but with less total displacement.

Convection is likely if an airparcel at some level is lifted to the LFC corresponding to that parcel.

Imagine air flowing over a solid obstacle such as a mountain range or, in this case, a sharp outflow boundary. Air trajectories would likely follow the paths outlined in this figure. The lowest level air would be forced upward the most. However, as the initial parcel level increases, the forced vertical displacement diminishes. This kind of lifting profile is a rough estimation of reality derived from laboratory observations by Simpson (1969) (QJR. Meteorl Soc., 104, pp 413)and radar observations by Wakimoto (1982)(MWR, 110, pp 1060)

The flow patterns may not match what is shown here but the vertical displacements relative to the LFC can be illustrated here. The LFC changes according to the sounding and the level of the parcel origin. In this diagram, the parcel lifted along the yellow trajectory successfully reaches its LFC to initiate deep convection.

Foote (1984) used the theoretical lifting profile of air forced upward by a gustfront to the vertical profile of lifting distance needed to air parcels at various levels to reach their respective LFCs. The application, which will be shown in the next few frames, shows which vertical moisture profiles are favorable for boundaries to initiate convection.

As an aside….The first frame of the loop shows the ceiling heights at about 6500 ft AGL which, if is the base of cumulus clouds and the boundary layer was well mixed, would yield a dewpoint depression of 30° F. The actual dewpoint depression was 37° F.

• Consider the amount of lifting required to bring an air parcel from its initial height to its LFC called H_lfc-z. For example, a surface air parcel would need to be lifted 2.1 km. Therefore, H_lfc-z = 2.1.

• Then consider a vertical profile H_lfc-z and compare it to a theoretical vertical profile of vertical air displacements by Foote, 1984.

We will lift air parcels up to their LFCs starting at the surface, then at 925mb, 900mb,…, 725mb. Then we can compute a vertical profile of H_lfc-z . In two frames there will be a graph showing the vertical profile of H_lfc-z and a theoretical vertical lifting distance profile from the outflow boundary. But first…

The next slide will be a modification of the KEDW sounding.

• What would happen now if a gust front lifted this air?

• We’ll compare both H_lfc-z profiles to a theoretical lifting profile by Foote, 1984.

• The unmodified sounding would remain dry represented by the solid blue profile.

The ordinate is the parcel starting pressure level. The abscissa is the vertical lifting distance H_lfc-z.

The area between the green dashed lines represent the potential vertical lifting profile from a barrier-like outflow boundary. For example, the air initially residing at 850 mb would be lifted from 1 to 1.3 km.

• Favorable convective soundings occur where:
• LFC heights are low
• Boundary depth is deep relative to the Hlfc-z profile.

A decreasing lifting distance with height Hlfc-z indicates the sounding is getting closer to saturation and that air from multiple levels can initiate deep convection. Low LFC heights makes it more possible for weaker boundaries to initiate deep convection.

• The theoretical lifting profile:

• Over a gust front may not be valid.

• Does not apply to non-density current boundaries.

This is a rather old theoretical treatment of lifting over a density current boundary. Air parcel trajectories may differ greatly and in fact we’ll see that shear and mean flow profiles relative to boundary motion can impact the depth of the ascending zone. Non density current boundaries will likely have a completely different set of air trajectories within the ascending zone.

• LFC distance profiles (Hlfc-z) may differ considerably in the space of a few km.

Consider the large moisture profile differences from within a convective roll updraft to that just outside.

An AWIPS user has a few observational tools to make broad estimations on the depth of a boundary and its ascent zone.

Reflectivity should then not be as much of a function of distance for points closer than A. The only thing that can affect reflectivity is the signal to noise ratio for weakly returning fineline echoes.

• The fineline partially fills the beam between points A and B.

• Beyond point B, there is too little echo for detectability. However, there could still be echo in the beam volume. The echoes might fall below the signal to noise ratio.

• Once the beam top exits the top of the boundary, reflectivity decreases at point A.

The beam top starts exiting the top of the boundary fineline and the returned reflectivity decreases more rapidly as the percentage of target coverage drops. Knowing the height of the beam top is more crucial than knowing the beam center height. One does not know where in that beam volume the fineline top is located.

• Therefore, use the beam top elevation to estimate the depth of a boundary.

Therefore the best method for detecting the top of a linear feature like a fineline is knowing how high the beam top is when the reflectivity starts to fall. For that reason, we have developed an applet for the web that is located on the front web page for this session.

Unlike reflectivity where you seeing direct evidence of ascending air manifesting as finelines. Here, you are only observing convergence/divergence. Calculating the integrated values of convergence below and up to the level of the radar beam is the only real way of knowing the sign of the vertical velocity. If you see convergence in the velocity display, most likely there is ascent occurring at that level. In a classic frontal-like circulation, the level of non-divergence is where the maximum ascent occurs. Therefore, ascending air should exceed the height of any convergence you see.

• However, detecting convergence is difficult due to the large default velocity increments of ± 10 kts.

Most weakly forced boundaries exhibit less than 20kts of cross-boundary wind differential. The current velocity increments would never adequately capture most boundaries, especially when the wind directions are mostly tangential.

• Solution, when possible reduce the velocity increments in the low range at the UCP.

Look in the Selection of product parameters menu in the UCP then pull of the velocity submenu. You can reduce the increment levels in there when you’re experiencing difficulty in discerning boundaries. Don’t forget to change it back when severe storms start forming.

A boundary passing over a radar maybe completely invisible in the base velocity if the winds on both sides of the boundary are tangential. Try to draw various windfields that might not show in a clear air base velocity display.

• However, velocity echoes display at lower threshold reflectivity.

The precipitation mode reflectivity displays do not show the lowest reflectivity that the radar observes. The base velocity shows velocity for all observable echoes. Therefore, you can observe boundaries on occasion out to greater ranges in the base velocity when the radar is in precipitation mode.

This is a simple but very useful benchmark for knowing when a boundary ascent zone reaches the LCL. For all the capabilities of radar, it simply cannot observe cumulus clouds. And it’s always a good idea to have cumulus to develop before a storm develops. The image on the right is a highly magnified view of a boundary in GOES-8 visible imagery. There are two east-west cumulus lines that indicate boundary lifting through the LCL.

• Organized TCU indicates deeper boundary lifting and more likely initiation.

This is not so easy to discern. There is no benchmark for observing when cumulus break through the LFC. Towering cumulus are also difficult to define from satellite data alone. One good candidate for a benchmark is the 3.9 micron reflectivity or albedo data during the day and the fog product or nocturnal 3.9 micron albedo product at night. These products (especially the daylight ones) are very sensitive to the onset of glaciation. In rapid scan imagery, the timing of glaciation in cumulus can be an important benchmark to anticipating storm development in a short term forecast.

• Moisture/temp gradients on low-level water vapor imagery indicates deep boundary.

This idea is uncertain. However, the deeper the air on the cool or dry side of a boundary, the more likely, the low-level water vapor sounder imagery will observe a temperature/moisture gradient.

Boundary # 1 is just emerging from a group of thunderstorms as an arc of stratiform clouds. This is not surprising judging from the lack of cumulus in the air preceding this boundary. Although there isn’t enough information to make any claims about the stability of the boundary yet. In any case, the forced lifting is reaching the LCL of the environment, whatever depth that might be.

Boundary #2 is not easy to define. Upon looping the visible imagery, notice the cumulus clouds moving east and dissipating along a north to south line just south of the label. There is a boundary along that line of dissipation. Go on to the next two slides for an closer look.

The thin green line crossing NW to SE across the radar site representes a velocity cross section in the next page.

Also note that the shape of the boundary head consists of an initial rise in depth followed by a depression. This matches the shape of outflow boundaries produced by numerical and theoretical results.

• The depth of the windshift might be related to the depth of the convergence.

This is a similar problem to that of the 88-D base velocity. Convergence can be estimated by the shift of the winds as the boundary passes. But to really calculate the vertical velocity profile, the convergence between two wind profiles on either side of the boundary needs to be calculated and then integrated. For visual inspection, the deeper the initial windshift, the more likely the convergence is deeper.

In the example on the right, the wind shifts immediately for the first two elevation gates (1 and 2 kft) above the radar in one volume scan. The 3kft gate wind remains unchanged for another 20 minutes (or 5nm behind the boundary). Finally, the 4kft wind (4th level) remains unchanged for 28 minutes before going to ‘No Data’. Most likely, the boundary interface has reached close to the 4kft range in 28 minutes.

It is difficult to say which wind profiles should be compared to get a picture of what the convergence is between the two sides of the boundary. In this case, say the width of the fineline is about 3nm. The wind profile at 1808 UTC would be inside the fineline. The 1803 profile sampled the air previous to the boundary. The 1813 profile could be close to the back edge of the fineline while the 1818 UTC should be clear of the back edge of the fineline. Assuming the winds directly in front of and behind the fineline are characteristic of the pre and post-boundary winds, and the convergence depth is equal to the depth of the windshift, then the convergence depth is between 2 and 3kft above radar level.

• The level of maximum ascent will likely be at the top of the convergence zone.

Cautions:

• Large VAD range will smooth out true wind gradients.

If the VAD range is very high, both sides of a boundary will be sam pled and a wind direction estimate will be attempted using data from disparate windfields. The default VAD range is 33km. A boundary moving at 20km/hr means there will be disparate wind fields for over an hour. If possible, bring in the VAD range as close as possible to the cone of silence at times of boundary passages.

The VAD range is the range out from the radar at which horizontal winds are calculated.

• Poor vertical resolution will allow for only .

• Horizontal variations in boundary depth not accounted.

Just like in the surface convergence estimations, wind profiles can only be used to estimate broad estimations of boundary depth convergence when you can safely ignore along-boundary variations in wind.

This illustrates the importance of being cognizant of how the VWP realistically depicts the environmental winds in regions of strong windshear. Decreasing the VAD range not only helps to resolve shallow windshear layers but it can also more accurately measure the low-level velocity gradients of boundary passages.

Estimating the depth of the radar fineline may underestimate the depth of the ascending air column especially for higher slices. It is not always the case that echo causing particulates reach as high as the ascending air column.

• From wind profiles and velocity:
• Depth of convergence can be estimated
• Ascending air should be higher than the convergence depth.

It is not clear how much higher the ascent reaches above the convergence layer.

• From satellite:

• Presence of cumulus indicates ascending air reaching LCL.

The uses of satellite will probably be greater than the other sensor sources when you actually forecast convection. It is the only tool that explicitly observes growing cumulus convection.

The next several frames point this caution out in even more detail.

• Strength of convergence doesn’t always mean deep convergence. A boundary may have strong convergence down low and be capped well below the LFC.

This is a problem especially when the LFC is high and/or if there’s a significant capping layer below it. Drylines are notorious for exhibiting strong sfc convergence and yet shallow circulations.

One thing to remember is that when looking at the vertical continuity equation, the vertical velocity at some level depends on the integrated values of convergence/divergence below. Vertical velocity is highest at the level of non-divergence where convergence resides below and divergence above. In most cases, the deeper the convergence, the higher the level of maximum vertical velocity will become.

One could easily imagine a boundary as a two dimensional interface between two regions of differing winds. In this example, there are no irregularities in the wind field on either side of the boundary as one travels up or down the boundary.

• Then the only component of air doing the converging is that oriented perpendicular to the boundary (red vectors).

• In this case, convergence can be calculated by calculating differential velocity parallel to the baseline (yellow line). Which baseline distance?

Convergence doesn’t need to be calculated along X and Y cartesian coordinates. There is radial convergence and in this case, convergence calculated along a rotated coordinate system. The only axis the red vector winds are converging upon is the axis of the boundary. Therefore, the baseline that we calculate convergence is the only baseline down which the winds change which are the ones perpendicular to the axis of the boundary. You can almost think of convergence calculated in natural coordinates where n is the distance along the yellow baselines. Then dv/dn is the way to calculate convergence.

The problem is, over what length (n) do you calculate convergence? The best length is that where the winds change. The next frames show a typical boundary and typical convergences calculated using typical station spacings and a realistic convergence over the right baseline length.

The dv/dn = convergence where n= 140km and is along the long yellow baseline. Since we only know the winds from the two METAR stations at the end of the baseline, convergence can only be calculated over this huge distance.

Now consider dense station coverage of 46 km. Dv/dn along the shorter line left increases to 7.1 X 10^-5 s^-1.

Now adding the mesonetwork of stations, we have a much shorter station spacing. In this example, convergence between the two endpoints of the 46km baseline is up an order of magnitude from the METAR based convergence.

But is that where the air is starting to converge in reality? Let’s go to the next frame.

The clear air echoes in the boundary fineline are the result of insects and plant debris caught in ascending air. In that case, the width of the convergence line is much less than the station spacing of even the mesonetwork.

• After assuming the winds at each end of the yellow line are also true at either end of the purple line segment.

Some of the stations away from the baseline but close to either side of the boundary suggest the winds are pretty homogeneous until right next to either side of the boundary fineline.

• The convergence increases to 6.86 X 10^-3 s^-1, 3 orders higher than with METARs alone.

• Convergence should be estimated across the width of the boundary.

• Estimating one- dimensional convergence assumes a uniform airflow on either side. In other words, a front with no undulations or intersections with other boundaries.

• Convergence will be higher at intersections, mergers, collisions with other sources of ascent.

Don’t forget the inverse as well. Convergence across a boundary will be lower where there is along-boundary divergence.

See the example on the right side. The width of the ascending air on this boundary is taken to be the width of the small cumulus line. Using the ‘Distance Bearing’ function in D2D, the boundary width is about 4mi. There are a couple things to consider when attempting to estimate the width of the boundary

• The ascending zone maybe sloped causing an unrealistically large width of the boundary.

• The cumulus clouds mark the upper part of the ascending zone and may not be the same width as the region of convergence down near the surface.

• Then estimate the winds directly on either side of the ascent zone.

This is a challenge here because of the lack of stations. However, observing the wind tendencies at each station as the boundary passes is a possible space/time method of estimating the pre and post boundary winds. In this case, the pre-boundary winds were westerly at 7kts. The post-boundary winds were northerly 15kts.

• Finally take the difference in the wind component normal to the boundary divided by the width.

First, you need to know the boundary orientation. The long blue line from the ‘Distance Bearing’ tool marks the rough orientation of the boundary. The line is oriented 100°. Therefore the line normal to the boundary is oriented from 010°. The component of the pre and post boundary winds should be taken along 010°. Then the difference of those components should be divided by the width of the boundary in meters.

As in the shear considerations, these theories were derived from density current boundaries.

• A boundary-relative headwind (tailwind) increases (decreases) the depth of the ascending air column.

Based on numerical modeling of an outflow boundary, the boundary height increases when a boundary moves into a headwind.

• Boundaries with a headwind move more slowly than those with a tailwind.

• From Liu and Moncrieff, 1996 MWR, 124, pp 2282.

From these theoretical and numerical experiments, the ascending air column will be deeper for boundaries experiencing a significant headwind than for those following into a tailwind. It stands to reason that low-level convergence with headwind (tailwind) boundaries is higher (lower).

• Data from METARs and mesonets will underestimate convergence in boundaries by 2 to 3 orders of magnitude.
• Model analysis will underestimate convergence similarly.

All operational model resolutions are too poor to capture the true width of a boundary. Also, stormscale boundaries are typically unobserved by models.

Typical boundaries are 3-10 km wide.

• Actual convergence will be occuring on these scales.

• Convergence will be higher at boundary intersections with other boundaries, circulations, rolls.

In this case, two dimensional convergence becomes important since winds vary both along and normal to the boundary axis. • The deeper the convergence zone relative to the LFC, the more likely initiation will occur.

There’s a growing body of research including theoretical and observational that may promise in forecasting convective initiation on boundaries without having to directly observe the depth of a convergence zone or fineline. Simply knowing the properties of the windfield in the reference frame of the moving boundary may give some important considerations as to how deep a boundary ascent zone maybe which leads ultimately to the propensity for convective initiation.

Factors of influence on estimating the depth of a boundary ascent zone:

• Boundary-relative low-level flow.
• Boundary-relative steering-layer flow
• Boundary-orthogonal shear.

Each of these topics will be considered in more detail in this section. There are addendums and quid pro quos in this section as with the last one. These concepts will not explicitly tell you where convection will form. Also, boundaries exhibiting unfavorable values of the above parameters can also initiate convection. Also, we do not know many of the characteristics we should be observing. Some or all of these theories may altered or even discarded as more evidence becomes available.

Related to deep layer shear considerations but this applies to any boundary.

When a boundary does not have any density current characteristics, the shear considerations do not apply in the sense in which they were presented. In some sense shear is still important since boundary-normal shear can be too strong potentially ripping apart an ascending zone of air before it can reach the LFC.

However there is another way to look at the potential depth of an ascent zone. Given an ascending air column reaching toward the LFC, what kind of flow aloft will favor the ascending air column to remain over the boundary for the longest time? And what kind of flow aloft should be considered useful?

Observations by Wilson and Megenhardt, 1997.

• Boundaries in FL were most active if cell motion -boundary motion 5 m/s.
• Before storms, the 2 – 4 km mean wind = cell motion.

Several observational studies including this one sampled the mean flow in a 2- 4 km layer and found the boundaries were most convectively active when the boundary-normal flow was less than 5 m/s. The 2-4 km layer was chosen since it correlated best with new storm motion. Boundary-relative flows greater than 5 m/s exhibited less storm activity and those storms that did form on the boundary would quickly be left behind in less stable air and dissipate.

This study was done during a field project around Cape Kennedy during the summer. Therefore the storms were mostly driven by surface-based phenomenon and there were small amounts of shear. Whether these findings can apply to stronger synoptically forced situations still remains a question worthy of testing at any of your locations.

Note that not all storms will dissipate if the boundary-relative 2-4 km flow is high. Convection can persist or even thrive well after they leave the original boundary if there is still sufficient uncapped instability.

The first frame shows cumulus equally distributed around the expanding outflow boundary.

The second – fourth frames show that the cumulus on the northeast side of the boundary are able to ride with the boundary. These parts of the expanding circular boundary have a low boundary-relative steering-layer flow. The other cumulus quickly advect over cool air left by the outflow and dissipate.

Supposing that the atmosphere is such that a front is needed to initiate convection, the top cumulus cloud would have the best chance to develop since it has the maximum residence time above the boundary.

Density current boundaries include strong fronts, seabreeze boundaries and outflow boundaries.

• Deepest ascent occurs when boundary orthogonal environmental shear balances cold pool vorticity (Rotunno et al. 1988)

• Or when the boundary speed (C) cancels out positive shear magnitude ( U_L). C/ U_L ~ 1.

When the boundary-normal (or orthogonal) component of the shear is oriented forward away from the cold side of the boundary (positive shear) and it has the same magnitude as the forward speed of the boundary, the conditions are optimal for the strongest, most vertically oriented ascent zone. That is C/ Delta U_L ~ 1. This is graphically shown in the upper-left schematic.

What often occurs is shown in the lower right figure where the shear vector is directed front to rear of the boundary (negative shear). In that case, no balance occurs in C/ UL and the ascent zone becomes much shallower.

A question may arise, how can boundary speed C with units of m/s be compared with Delta U_L with units of s^-1? A possible answer lies in the assumed properties of a density current behind the leading edge. The strongest flow behind the boundary exists near the surface. The flow weakens with height until reaching near zero near the top edge. Actually there are numerous shearing eddies along the top edge of the cold pool. Then the velocity difference behind the boundary directed back from the leading edge with a magnitude equal to the forward speed of the boundary. Therefore, Rotunno et al. 1988 proposed that the forward motion of the boundary, C, can be used to compare to the magnitude of the shear (Delta U ) in the layer of air in front of the boundary.

Much has still yet to be understood regarding what shear layer should be used to anticipate convective initiation. None of these theories take into account the thermodynamic profile of the environmental air. Does the depth of the shear layer need to match some thermodynamically based parameter? Is one shear layer important for initiation considerations versus storm maintenance? Or is the shear argument relevant at all? Recent findings suggest there is no relationship between cold pool strength and low-level shear in the maintenance of long-lived severe squall lines. Refer to this work in review by Evans and Doswell, 00. There has not been enough observational evidence to support or refute the shear theories for convective initiation. However, each NWS office has the tools to evaluate whether boundary orthogonal shear is relevant.

The depth of the leading edge of the outflow boundary (shaded in blue) corresponds to the depth of the ascending zone. The probabilities of new convection are proportional to the depth of the ascending zone in each scenario. There are no thermodynamic considerations in this schematic. The wind profile is given as a vertical profile of horizontal wind vectors pointing toward which the wind is blowing. The low-level flow is depicted by the lowest wind vector while the shear is visualized by the change in wind vectors with height. All the wind vectors are ground-relative. The depth of the shear layer depicted by the wind profile is about 4 km AGL. It is intended that the depth of the outflow is lower than the depth of the shear layer.

This figure is adapted from Moncrieff and Liu, 1999.

The second frame in this loop is a schematic considering the effects of mean layer flow on probabilities for initiation. In these considerations, the boundary depth itself is not known to be affected by shear. Instead the concern is the residence time for developing cumulus. Therefore, a boundary with strong headwinds, may not be able to produce deep convection because the steering layer flow prevents an upright deep ascending column.

Here is a brief mini exercise where you issue short term forecasts based on the theories on how boundary depth is related to surface convergence, shear and boundary-normal steering layer flow.

Presented here is a sounding representative of the forecast area to be presented in the next page. Note in this case, there is not much shear present in the sounding but there is a steady southerly flow in the lowest 5 km, moderate CAPE and little CIN.

The next page is a loop of satellite imagery and the goal is to draw low, medium and high probability contours of where you expect the most convection to develop on an outflow boundary to be generated by a thunderstorm nearest to the red pointer.

In the teletraining version of this session, you will draw probability contours of low, medium, and high chances of new convection around the target storm knowing that a circular outflow boundary is about to develop. Let’s assume the outflow boundary will initially move out at 7 m/s (14 kts)in all directions from its source. Do your thunderstorm probabilities for a period of two hours.

• The boundary motion direction is defined as the direction from where the boundary has moved orthogonal to the boundary orientation.
• Then boundary-orthogonal flow and shear can be visualized.

This is very similar to the coordinate transformations done with storm-relative flow parameters. Remember that for a storm-relative framework, the storm is the origin but there is no coordinate rotation. For a boundary, you can visualize the boundary axis as a an original axis rotated to the orientation of the boundary. Then the origin of the new coordinate would lie on the boundary axis at the intersection of the boundary motion vector (S_b)

Note that sometimes, boundary-relative flow and boundary-orthogonal flow is used interchangeably. To reduce confusion, boundary-relative flow implies the orthogonal component of the flow relative to boundary motion. Boundary-orthogonal flow also implies the same orthogonal component of flow relative to boundary motion.

Applying shear/cold pool theories to your area.

Slide 1: What is the most potentially active boundary orientation and movement given an environmental: • Shear profile? • Low-level wind? • Mean steering wind?

The next several slides are meant to show two ways to visualize the shear, low-level wind and steering layer wind in the reference frame of a boundary. These are forecasting aids to help you apply or test the considerations about how these parameters affect convective initiation in your area.

We can use a hodograph to immediately visualize the flow in the boundary coordinate system. This hodograph was taken from 29 June 1998 in Omaha, NE.

Delta V _0-4km is the 0 – 4 km velocity difference or shear.
V_low is the low-level flow vector.
V_cell is the cell motion observed for that day.

We can actually plot the boundary motion vector and the boundary axis at the tip of the vector to show that the hodograph can be plotted in the boundary frame of reference.

Slide 2:

Adding the boundary motion • Visualize Boundary- relative flow and boundary-orthogonal shear on a hodograph.

• Plot the boundary motion vector on the hodograph.

To plot the motion vector of the boundary on the hodograph, we chose the vector direction as from which the boundary is moving. The boundary speed is taken as the displacement distance of the boundary normal to the axis of the boundary over a period of time of say 30 minutes. The result is a boundary motion vector called Vfront

• Then draw the frontal axis perpendicular to the vector.

In the frame of reference of the boundary, the orientation of the boundary can serve as a new X-axis. The red line depicts the orientation of the boundary at the physical point where we estimated the boundary motion vector.

Applying shear/cold pool theories to your area.

Slide 3:

Now the wind and shear in a boundary-normal sense can be visualized.

• Boundary-relative low-level flow is marked by the long yellow vector.

The hodograph shows that a southeastward moving boundary will create a strong headwind in the low-levels. That should be considered a positive for good low-level convergence.

• Boundary-relative cell motion is marked by the white line.

The white line is actually a white arrow eminating from the boundary to the Storm motion point. The ground-relative cell motion is fairly quick but when taken relative to the boundary, the cell is just barely making it ahead of the boundary.

• Boundary-orthogonal shear (Delta U_{b 0-4km}) is marked by the cyan line.

The boundary orthogonal shear vector is fairly strong and positive, a positive sign for initiation of convection. Note that the full shear vector was moved from its origin out to the boundary itself. Remember that shear is galilean invariant meaning that it is unaffected by any frame of reference. Also remember that the shear/cold pool theories only consider the boundary-orthogonal component of the shear vector.

Applying shear/cold pool theories to your area.

Visualizing boundary-relative flow on a hypothetical expanding circular boundary.

Now we present a hypothetical circular boundary (like an outflow) conceptual figure where you have all possible boundary orientations. Just like on a hodograph, you may plot the mean wind or cell motion vector, the shear vector and the low-level wind vector.

• Determine V_cell from mean flow or actual cell motion.

Vcell is the cell motion vector and in this case, it is from the west at 30 (any units).

• Determine which boundary motion will yield the smallest relative cell motion.

Here, boundary motion is defined similarly as with the hodograph situation. Take a direction from the origin out to the cyan circle, the intersection point is the boundary motion vector. Note the boundary motion vector that yields the smallest boundary-relative V_cell. In this case, an eastward moving boundary of 20 (any units) yields the smallest value. The smallest V_cell still exceeds the boundary speed by 10.

• Observe actual boundary motion and compare

If you have an actual boundary motion, you can plot the direction of motion on this hodograph. Then you can immediately compare the optimal boundary orientation to the actual boundary orientation and compare. If your boundary was moving out of the northeast, then your V_cell would be -35, or rapidly front to rear of the boundary.

You can use this relation to calculate boundary-relative cell speed. U_b = V_cell cos(theta) -V_front where U_b = boundary-relative cell speed V_cell = Cell motion or mean steering layer velocity. = angular difference between cell motion direction and boundary orientation. V_front = Speed of the boundary.

Applying shear/cold pool theories to your area.

Visualizing boundary-orthogonal shear on a hypothetical expanding circular boundary.

To calculate boundary- orthogonal shear…

• Subtract the surface flow vector from a velocity vector 2 -4 km AGL from a hodograph or by using Product Maker on model analysis.

This step is for estimating the shear vector of the environment on the most unstable side of a boundary. There is still a question of how deep a layer the shear should be calculated. This would be a good question a local study could answer. As a first application, we choose the 0-3km shear. At the lower end, you may wish to choose a representative wind in the lowest 100m.

• Use equation 2] to calculate the boundary orthogonal shear (S_b).

This equation, U _b = V cos (theta) is similar to the one to calculate boundary-normal flow except for shear, boundary motion is irrelevant. The angle (theta) is still the angle between the direction from which the boundary is moving and the shear vector direction. What if the boundary is not moving? Then the angle for the boundary is taken here as the angle perpendicular to the boundary axis taken from the most stable to the least stable side.

For example, a stationary boundary oriented SW to NE with the most stable air located north of the boundary would have a boundary orientation of 315°. This is just a convention but it does help calculate the boundary-normal shear vector.

• See if V_front / U _b ~ 1 results in the most action.

For the example on the right, a shear vector of 180 degrees at 10 kts is displayed as a red arrow at multiple points around the hypothetical circular outflow moving away from a point also at 10 kts. According the Rotunno et al. 1988, the optimal boundary motion vector would be one where the shear vector is pointed in a positive direction (away from the stable side) and the magnitude equals the boundary speed. In this example, an east-west boundary tangent to the top of the hypothetical circle moving north at 10 kts is the optimal boundary for initiating convection according to theory.

We have applied this hypothetical circular outflow boundary to calculate boundary-normal flow and shear to find out which boundary orientation would be most optimal for initiating convection. In the two past examples, the circular boundary was expanding outward at 10 units/time (e.g., m/s, kts). What happens with a slower or faster boundary speed? Well, for boundary-normal steering layer flow, equation 1 factors in the boundary speed and from that you may calculate the most optimal flow. For boundary-normal shear, only boundary orientation is important so that you may calculate the component of shear normal to the boundary. However when you are looking for the optimal condition where shear balances out boundary speed, V_front / Delta U _b ~ 1, then boundary speed comes into importance.

For wind direction and speed, simply put the direction into Cell Dir. and speed, place it into Cell Speed. Any wind or cell motion is fine for this calculation. For boundary speed, you can place any observed or estimated number into Boundary Speed. The calculator will show you the boundary-normal motion vector for various boundary orientations.

Example case study: Using the Calculator.

1. Using the ‘distance bearing’ tool, orient a segment parallel to the target boundary.
   Upon clicking the ‘distance bearing’ tool, numerous line segments appear on a plan view map. After making these lines ‘editable’, take one of them and oriented it parallel to the target boundary. The line segment will have distance between the two endpoints and the orientation as a legend.
2. Track the boundary motion with the ‘distance speed’ tool and normal to the boundary orientation that you measured in part a.
   Now take the ‘distance speed’ tool to track the boundary in a loop. The direction of motion should be perpendicular to that of the boundary orientation and the direction of motion should be from which the boundary has moved.

Example of measuring boundary orientation and motion.

GOES-8 visible loop 5 hours after sounding time.

We go back to the same example that we used to estimate the boundary width when calculating low-level convergence. The next frame shows a loop of 1km GOES-8 visible imagery where the boundary motion can be tracked and its orientation measured. We measure the orientation of the boundary axis to be 100 degrees (blue distance bearing line).

The boundary is moving south, perpendicular to the axis of the boundary. Making sure we use the ‘distance speed’ utility to track the boundary motion perpendicular to its axis, we find it is moving toward 190 degrees at 25 kt. By convention, the boundary is moving from 010 degrees at 25 kt.

Remember that the ‘Distance Speed’ Tool marks the direction toward instead of from which your object is moving.

Example of measuring mean wind and shear.

12 UTC Sounding hodograph at KOMA.

• What is the environmental cell motion?

• Track incipient TCU on satellite, new cells on radar.
• Estimate mean wind in the cell bearing layer.

The hodograph on the right was taken from a sounding 5 hours before the previous satellite loop. It shows a 0-6 km mean wind of 280 deg at 22kt. There were no incipient cells to track to see how well they agree with the mean wind. What is the environmental shear?

• Depth of shear layer? There are conflicting answers.

The 0-3km shear layer is similar to the mean wind.

Applying observed boundary motion and 0-6 km mean wind to the boundary applet.

The boundary-normal 0-6km wind displays around the perimeter of the hypothetical circular outflow whose speed is moving outward at 13 m/s. Note, however, that we measured the boundary moving from 010 degrees. Therefore, the boundary-normal 0-6km wind is taken where the real boundary motion and axis(thick blue line) lies tangent to the hypothetical circle.

The 0-6 km mean boundary-normal wind is slightly greater than the -10.9 m/s value listed nearby. The applet marks the boundary-normal flows that are optimal for initiating convection in red. The values for this boundary indicate strongly negative boundary-normal mean wind and therefore, we would assign a low probability of this boundary becoming convectively active in the immediate future.

The GOES-8 loop a couple frames back supports this forecast as there are no growing cumulus along the boundary.

2hr nowcast Discussion.

• The boundary is moving too quickly to permit incipient cells to remain in the mesoscale updraft.
• The 0-3 km boundary-normal shear vector is oriented rear-to-front of the updraft but C/(Delta) U = 13/4 which is greater than 1 indicating the boundary is overpowering the shear.
• Satellite data shows no developing cumulus ahead of the boundary.

• Nowcast at this time would not favor initiation in the next couple hours near this segment of the boundary.

What are they?

All boundaries exhibit horizontal variations in their properties that can focus initiating convection.

Studies of Boundary intersections, mergers, collisions:

•  
• Intersections with boundary-layer rolls.
• Seabreeze/outflow boundary collisions.
• Intersecting outflow boundaries.
• Dryline triple points.

There have been numerous studies since the 1970’s that observed the tendency for convection to initiate on boundary intersections, mergers and collisions. These features have often been used to highlight areas where convection is expected to develop.

• Variations in stability.

There is no airmass with spatially uniform instability. Boundaries with poor potential to initiate convection based on flow and shear considerations can still be prolific thunderstorm producers given low enough CIN and LFC.

•  
• Wilson et al. 1992 MWR, 120, pp 1785,
• Fankhauser et al.1995, MWR, 123, pp 291.

• Seabreeze/outflow boundary collisions.

•  
• Kingsmill, 1995, MWR, 123, pp 2913.

• Intersecting outflow boundaries.

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• Wilson and Shreiber, 1986, MWR, 115, pp 2516
• Browning et al. 1997, WAF, 12, pp 915.

• Dryline triple points, intersections.

•  
• Hane et al. 1997, MWR, 125, pp 231.

There are many more papers that exist and some of them are listed in the boundary bibliography on the main boundary homepage.

Boundary Intersections, Mergers and collisions.

Boundary Intersections, Mergers and collisions

Probabilities of initiating convection

Boundary intersections with rolls.

• Initiation is favored where roll updrafts intersect another boundary.

•  
• Especially for large rolls or rolls with the largest cumulus street.

Here is a schematic figure which shows how the updraft axis of the rolls becomes superimposed on the ascending zone of the boundary to create enhanced regions of updraft. Conversely, the downdraft axis of the rolls will negate or cancel out the boundary ascending air to limit the potential for convection in these areas.

Most average roll spacing is only about 4km, the scale of cumulus clouds. There is not much room for enhanced cumulus to grow at the intersection points along a boundary. Occasionally, large or enhanced rolls exist with an aspect ratio of greater than 7. These rolls exhibit greater depths of circulations and will be even more likely areas of initiation where they intersect boundaries.

example of rolls interacting with a boundary

Also note that the rolls actually are lifted up and north of the cold front for a short distance.

Variations in Stability.

• Large variations in stability exist in very short distances.

•  
• CAPE (CIN) can double (half) in value for soundings within ascending (descending) regions of roll circulations.

There may not be much difference in surface dewpoints between the updraft and downdraft axis’ of Horizontal Convective Rolls (HCRs). However the depth of the surface moisture changes rapidly. Therefore Mixed Layer CAPE will change dramatically between the updraft and the downdraft parts of HCRs. Satellite is an excellent tool for judging relative variations in stability by observing:

•  
• Areas of concentrated cumulus in imager data.
• Areas of enhanced moisture/instability in derived product imagery.

For short term forecasting of convection (0-2 hr), there is no better way to monitor small spatial variations in atmospheric instability than by observing horizontal changes in the behavior of cumulus clouds by GOES satellite. Boundaries that lie or move into pre-existing fields of enhanced cumulus are very likely to initiate convection. Likewise, HCRs with enhanced cumulus are the ones to pay close attention to, especially if they intersect with boundaries. Again, boundaries with unfavorable values of boundary-normal shear and flow for initiating convection may do so prolifically if they interact with a pre-existing field of enhanced cumulus.

In fact, monitoring rapid changes in cumulus fields is a more important reason to invoke Rapid Scan Operations than waiting until there are mature severe storms.

Summary.

For intersections, the convection may appear to lag the intersection point in cases where either or both boundaries are closing the gap before the intersection. Look for initiation about 15-30 minutes following the intersection for any one point.

For a boundary moving into a field of enhanced cumulus, look for these cumulus to initiate into storms between 10 and 40 minutes after boundary passage.

Collisions favor convective initiation at the point of collision about 15 to 30 minutes after the collision.

Stationary boundaries will tend to develop storms within 10km of their locations.

These rules of thumb were developed by Wilson and Schreiber, 1986 for typical summer convection. Given different CIN, LFC and instability, and strength of a boundary, these rules of thumb may change dramatically.

• Initiation is favored

• CAPE > 0, low CIN
• Region of enhanced cumulus
• Deep boundary relative to the LFC height
• Strong low-level convergence
• Low boundary-relative steering layer flow
• Moderate positive shear?

• Initiation is discouraged.

•  
• CAPE>0, higher CIN
• No cumulus in area.
• Shallow boundary relative to LFC.
• Weak low-level convergence.
• High boundary-relative steering flow.
• Negative shear.

Remember that these variables may contradict each other. For example, a boundary may have excellent shear and flow parameters but there maybe a cap too strong to allow initiation. Conversely, a boundary with poor shear and flow parameters may exist in an environment where any forced ascent will cause numerous thunderstorms.

As in warning environments, nowcasting convection requires a careful consideration of all parameters, some of which were covered in this section.

An example of horizontal variations in stability, boundary collisions and multicell propagation.

Upper air plots 12 UTC

Our area of concentration will be in Minnesota and the Eastern Dakotas.

The white box on the 925 mb plot represents the forecast area.

The low-levels show a jet arcing northeastward through Southern Minnesota and a stationary east-west frontal boundary north of International Falls.

There is a weak mid-level shortwave trough passing through Wisconsin and Minnesota. The forecast region is behind the trough axis.

A northwesterly jet extends from Montana to Nebraska. The forecast region seems to be in a diffluent region. The direction in the upper air flow is uncertain in the forecast region.

July 13, 1999.

An example of horizontal variations in stability, boundary collisions and multicell propagation.

Watervapor loop and 350-250mb PV

The next page shows RUC-II 20 UTC soundings corresponding to some of the station ID’s in the last frame of this loop.

The CAPE and CIN were calculated using a mixed 50mb surface layer.

The teletraining version contains some more information and two rounds where you make the nowcast.

We’ll focus the forecast area (FA) to the right of the black line visible on the 925 mb plot in North and South Carolina.

The low-level pattern indicates a large anticyclone northeast of the FA and a easterly wave type trough to the southwest. An upper-tropospheric anticyclone is located southwest of the FA. Thus, low-level easterly flow is replaced by mid and upper-level southwesterlies.

The moisture values plotted here are actual dewpoints and not depressions. The FA is rich with low-level moisture but at 700mb and above, the airmass is dry.

The MHX hodograph is somewhat different in that the low-level flow is westerly and the shear in the lowest one km is northerly. However, ESE shear exists between one and three km followed by a shear reversal above to WSW. The differing wind direction in the lowest km may be a function of the coastal landbreeze. Later in the day, a seabreeze develops likely leading to a differing wind profile.

Estimating storm motion based on these hodographs is very problematic. The 0 – 6 km mean wind at MHX is 004° 2 m/s and at GSO, 164° 3 m/s. From 0 – H_lfc(2km)the mean wind is easterly about 6 m/s. The wind remains easterly up to 3 km. Therefore, cumulus clouds reaching the LFC should be embedded in a steady easterly flow and therefore incipient cell motion should be from the east at 6 m/s. However, mature cell motion is more difficult to discern. Following the mean 0 – 6 km wind, mature cells should remain nearly stationary.

We go on to the first satellite loop. The GOES-8 is in RSO mode.

For the instructors in this exercise:
1. have the participants analyze surface boundaries and areas of best instability (using different colors). Refer to the teletraining session for the actual exercise. This loop is here to help familiarize yourself with the case.