Author: nwjunker

Heavy rainfall forecasting training manual

Published in 2008

Table of contents

The information in this training package is best viewed using Internet Explorer as some animation files do not work properly using Mozilla

1)

Convection and Mesoscale Convective Systems

A)

Scale Characteristics of MCSs

B)

On the evolution, shape and cell movement of MCSs

C)

Propagation and movement

i)

The original vector method

ii)

The revised vector method

iii)

An example of a backbuilding MCS

iv)

When is a forward propagating MCS likely

D)

How long will an MCS last before dissipation

E)

Exercise I. Forecasting convective movement

F)

Exercise 2. Forecasting convective movement

G)

An ingredients based methodology for predicting convective rainfall.

i)

Rainfall rates, vertical moisture flux and precipitation efficiency

2)

Original Maddox et al. MCS arch-types associated with flash flooding

A)

Synoptic type heavy rainfall events

a)

A forecast exercise, using pattern recognition and ensemble guidance to predict an heavy rainfall event

B)

Frontal type heavy rainfall events

a)

Regions of weak inertial stability

C)

Mesohigh type heavy rainfall events

a)

Use of Climatological thickness

b)

“saturated thickness” concept

D)

A few thoughts about elevated convection

E)

Synoptic Scale Circulation Systems

F)

Latent heat and the generation of mesoscale convective vortices

a)

A mesoscale convective vortex exercise

3)

Western U.S. heavy rains. Overview

A)

Maddox type I

B)

East slope of the front range events

C

Maddox type II

D

Chappell deformation type heavy rainfall event

E)

Maddox Type IV events

F)

An example of a warm season western flash flood

4)

West coast cold season heavy rainfall events.

A)

Role of tropical convection and atmospheric rivers

B)

Southern California pattern

C)

Using climatoglical anomalies along the west coast

a)

A case study

D)

Leeside heavy rainfall events-spillover events

5)

Hurricanes and heavy rainfall

A)

Rainfall at Landfall

a)

Kraft rule of thumb

b)

TRaP rainfall

B)

Rainfall primarily west of track storms

C)

Rainfall primarily east of track storms

D)

East Pacific Storms

a)

Effects over southwestern U.S.

b)

Effects over Texas and Oklahoma

E)

Hurricanes and the Appalachians

a)

More on hurricanes and the Appalachians, an exercise

b)

Hugo and Fran, a quick comparison

F

Weakening tropical storms and depressions

G)

References

A big thanks to those who helped collect data for cases used in this manual (Rich Grumm and Mike Bodner in particular) and to the many authors of the journal and preprint articles that were used to illustrate concepts discussed within the manual. Also, thanks to the American Meteorological Society and the National Weather Digest which which were both sources of material used within various parts of the manual. Finally, thanks to Ed Danaher and Jim Hoke for supporting the project.

Wes Junker, HPC Hydrometeorological Testbed/IM Systems Group

Convection and Mesoscale convective systems (MCSs)

A number of researchers have noted the importance of convection and especially, mesoscale convective systems in producing warm season precipitation. Heideman and Fritsch (1984) estimated that about half of warm season precipitation over the United States is produced directly by mesoscale systems or phenomena. Fritsch et al., in a study of long lasting, circular shaped, convective systems, using satellite imagery found that such systems accounted for approximately 30 to 70% of warm season precipitation between the Rocky Mountains and Mississippi River.
More recently, Schumacher and Johnson (2005) examined the synoptic and mesoscale structures of extreme rainfall events over the eastern two thirds of the U.S. using radar imagery. They defined “extreme rainfall” as the 24 hour precipitation amount that exceeds the 50-yr recurrence interval amount at that point. They noted that 65.5% of extreme events were associated with MCSs and that 26.7% were associated with synoptic scale systems. However, the latter figure may be misleading since they noted that for synoptic systems, over 90 percent of the extreme rainfall events were associated with convection. In other words, 97.2% of all “extreme” rainfall events were associated with convection.

From Shumacher and Johnson, 2005

Therefore, understanding what governs the development of convection and convective systems is essential for forecasting heavy to extreme rainfall events.

Heavy rainfall typically occurs with moist deep convection. Doswell et al. (1996) have said that in order to produce moist, deep convection three ingredients are needed: 1) the environmental lapse rate must be conditionally unstable, 2) there must be enough lifting so that a parcel will reach its level of free convection, 3) there must be enough moisture present that a rising parcel’s associated moist adiabat has a level of free convection.
“Nothing can replace an examination of an individual sounding” according to Doswell (1982). Soundings are the best tool to look at the depth of the moisture, the CAPE, CIN, LFC, LCL and Equilibrium level and for estimating what the cloud base will be. We’ll assume that forecasters understand how to read a sounding. Besides using CAPE calculated from soundings and model soundings to assess instability, forecasters at HPC also often look at Lifted and K indices. The latter gives an indication of the depth of the moisture. It is also important to remember that the stability can be changed by diurnal heating, differential advection of moisture and temperature, vertical motion and or differential vertical motion.
While there is a clear association between large scale lift and convection, synoptic scale lift usually is not sufficient to release convective instability. Instead, it plays a role in steepening the lapse rate and decreasing the convective inhibition or cap. By contrast, mesoscale lifting produces near surface low level convergence that forces the moist boundary layer air upward, when the moisture and low level convergence is sufficient for a parcel to reach its LFC, convection occurs.
Typical sources of mesoscale lift include: fronts, orographically induced convergence zones, sea breeze boundaries, dry lines and outflow boundaries and even land use boundaries. Therefore, it is important for a forecaster to analyze surface, upper air data and radar and satellite imagery closely for possible mesoscale features that might provide lifting.
Moisture helps modulate the cap and whether the parcel can reach its LFC. Available moisture also helps determine how much moisture will be available for a thunderstorm to process. The higher the influx of moisture into a convective system, the more potential it has to produce high rainfall rates. However, there are other factors that also modulate how much rainfall a system will produce. These factors will be discussed later in the manual.

Convective systems (MCSs) come in many shapes and forms and models often have a hard time forecasting them and their associated precipitation. To correctly predict the precipitation associated with an MCS, a model has to correctly forecast: 1) the timing and location of the onset of convection (the initiation of convection), 2) the convective system evolution (including mode of development, the size of the convective system, how it will be effected by propagation and duration of the system), 3) the rainfall rates associated with the system which may vary across the MCS and may fluctuate with time, 4) the impact on other downstream weather elements such as clouds and temperatures which may in turn impact on the vertical temperature structure and therefore the stability of the atmosphere.
Maddox (personal communication) has stated that if a forecaster or model messes up forecasting the initiation of convection (when and where it forms), there is little hope of forecasting the resultant convective system or its associated weather correctly. There are a number of reasons why mesoscale models have problems predicting initiation of convection. One obvious problem is the lack of sufficient observations to accurately depict the initial state of the atmosphere. This is especially true of measurement of the horizontal and vertical distribution of moisture, both of which modulate the stability of the atmosphere. A second problem is that a number of atmospheric processes occur at scales below the scale that mesoscale models are capable of resolving. The most obvious example is convection. All convective parameterization schemes have weaknesses (Gallus 1999). Their prime purpose is to keep the mesoscale model from suffering from grid scale convection which can feedback erroneously into the model mass and wind fields. These parameterization schemes have problems simulating outflow boundaries and do not include sophisticated enough microphysical processes to correctly simulate the rear inflow jet.
During past decade, a number of researchers have undertaken studies in hopes of improving warm season precipitation forecasts. Stensrud (1999b) illustrated that there was a need to improve the initialization of such mesoscale features as cold pools. However, Gallus and Segal (2001) showed that various methods to improve mesoscale initiation did not consistently improve QPF. At first glance, using ensembles to assess probabilities offers for various precipitation thresholds. However, studies have shown that the early SREF ensemble runs had insufficient spread (Hamill and Colucci 1998), probably because they were based on perturbing initial conditions which during the weak flow of summer did not typical grow fast enough (Stenrud et al. (2000) to encompass all the possible solutions. In the latter article, the authors discussed that the physics and initial conditions have to be varied in ensemble forecasting. NCEP introduced physics uncertainties into the SREF in 2004 (Du et al. 2004) by varying the convective parameterization schemes. However, the members still tend to cluster around the convective scheme rather than truly dispersing.
A study by Jankov et al. (2005) looked into varying model physics in additional ways to study the impact of various physics packages on the WRF QPF. They found that no one model configuration was clearly superior in all cases and that varying how convection was handled led to the most significant differences in the forecasts. They also noted that the rainfall rate was typically highest when no convective parameterization scheme was used. Threat scores for all model formulations was low for the higher thresholds of precipitation. For the foreseeable future, forecasters will continue to be an important element in predicting heavy convective rainfall events.

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Scale Characteristics of MCSs.

Kane et.al (1987) found that the composite precipitation patterns associated with various types of MCSs (based on the Maddox archetypes) were very similar but that the scale of the events differed. They also found that synoptic and frontal type systems were typically larger than mesohigh type events. Why are there differences in scale between these different MCC archtypes? A study of MCSs over the Southeast by Geerts (1998) found that MCSs were more common in summer but were of larger scale and lasted longer during winter.

This suggests that the scale and strength of the forcing may be one factor that determines the scale of an MCS. Augustine and Caracena (1994) have shown that the very largest scale, longest lived MCSs are associated with stronger frontogenetical forcing than smaller scale shorter lived MCSs

Their work supports the idea that one of the factors that governs the size of the MCS is the scale of the forcing and the strength of the low-level frontogenesis.
However, the scale of the event also may be dependent on the available moisture (the mean relative humidity and the precipitable water). Junker et al. (1999) studied the rainfall associated with the MCSs during the 1993 warm season and found that the scale of the area covered by 3 inches or greater rainfall depended on the precipitable water and thickness at which the rain fell. Almost all of the larger scale events occurred with mean 1000-500 relative humidity values of greater than 70 percent. With lower relative humidity than 70 percent, if a 3 inch area was observed, it was very small.

The core of the heavy rainfall associated with MCSs and MCCs is also usually small making forecasting its location a major forecast problem (Kane et al. (1987), McAnelly and Cotton 1986), Junker et al, 1999). Kane et al. studied 74 mesoscale convective systems that met the size and shape requirements to be defined as MCCs. They noted that the maximum rainfall for the 74 cases averaged 104 mm and that the median was 89 mm. However, they also noted that the size of the 77 mm or greater amounts was usually very small compared to the size of the MCC. Therefore, when they developed a normalized composite of their cases, they found probability of rainfall amounts greater than 77 mm (3 inches) was limited to a little more than 10 percent. This suggests that even if a forecaster were correct in predicting the development, evolution and track of an MCS, it would still be very difficult to predict the core of greater than 77 mm rainfall. This argues for a probabilistic approach to predicting heavy rainfall.

The figure below shows the cumulative rainfall that caused the Fort Collins flash flood. The rainfall during the Fort Collins flood illustrates the small scale that extreme rainfall events can have. The 4 inch or greater rainfall fit into a 10 by 10 km grid. The small scale associated with many extreme convective events argues that in the foreseeable future, numerical models will show little, if any, skill in accurately forecasting them explicitly.

However, high resolution, non-hydrostatic ensemble runs with explicit convection may offer potential in assessing the probability of a small scale heavy rainfall event occurring within a general area providing the ensemble suite is unbiased.

A figure from McAnelly and Cotton (1989) shows that the area of precipitation thresholds associated with MCSs decay almost logarithmically.

From McAnelly and Cotton 1989

HPC verification (figures above) show a similar decay as the thresholds increase but this time the scale on the left has been converted to a logarithmic one. Note the almost log linear decay of precipitation thresholds and of the HPC forecast accuracy (Threat Scores).

Model guidance often does a pretty good job of predicting the scale of the event when strong forcing is present due to strong shortwave that is accompanied by a strong low level jet and warm advection pattern. However, the models do poorer when the dynamics are weak and the scale is small. A graph showing the threat scores for four different types of convective systems (Watson 1999) illustrates this point.

Higher threat scores were associated with synoptic and cool type low events than during frontal or mesohigh events which typically have a smaller scale than either the synoptic or cool low type event. The small scale of the heavier amounts makes it difficult to provide meaningful deterministic or probabilistic forecasts for a specific point. Note that the verification of 12-h MOS PQPF for 2.00 inch or greater amounts shows only about a 3% improvement over climatology for the 24-36 hr period (from Shirey).

A more rational approach to forecasting extreme precipitation thresholds that are associated with mesoscale convective systems might be to forecast and verify them within a pre-defined area. For example a forecast 5 inch area that verified somewhere within 150 km of the forecast might be considered a hit.

However, HPC currently forecasts precipitation thresholds associated with MCSs including the rare high end thresholds deterministically. The following pages offer suggestions on predicting the evolution and movement of MCS in hopes of improving precipitation forecasts of convective systems. However, HPC excessive rainfall potential outlooks are now probabilistic.

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Unfortunately the current generation of operational models often cannot handle propagation effects since the NMC models convective parameterization schemes and microphysical packages do not realistically simulate downdrafts. Next, propagation will be discussed because of its importance in determining the mode of development (archtype) and movement of MCSs.

On the evolution, shape and movement of MCSs

A number of researchers have documented that MCSs evolve in a number of ways (Bluestein 1997, Houze et al. 1990. Loehrer and Johnson 1995 have suggested there are at least four ways in which convection can organize into a linear MCS (see below).

From Hilgendorf and Johnson 1998

Parker and Johnson (2000) have identified three linear MCS archetypes and indicated that the type that develops is dependent on the system’s relative wind field. One type is the trailing stratiform type (TS) which has convection along its leading edge and stratiform precipitation along its rear edge.

Note the deep storm relative winds (line-perpendicular component) for the TS type compared to the LS type

This type of MCS develops when the system relative flow is from the front to the rear of the system. Note that on the right hand side of the figure that the line perpendicular flow (the arrows) is fairly strong from the surface to the mid-levels. Such a system typically propagates forward causing it to move faster than the mean flow which is the reason the system relative flow is from the front to the rear of the system.
A second type (LS) has stratiform precipitation along its leading edge with its convection along the rear edge of the storm. This type of system is typically much slower moving than the trailing stratiform archtype and has the system relative flow at upper levels working from the rear to the front of the system. New cells forming along the rear of the system tend to retard the system’s movement so the MCS moves slower than the mean flow.
A third type is the parallel stratiform type (PS). This latter type tends to translate with the speed of the associated mid-level trough with the mean flow generally aligned with the system. Most of the Maddox et al. synoptic type flash floods are associated with this latter type of system. Usually, the heavy rain producing PS type systems are associated with a slow-moving mid level trough.
Shumacher and Johnson identified two other archtypes. The first, the trailing line-adjoining stratiform system produced more extreme rainfall than any other type of MCS. Its cell motion is approximately parallel to the convective line which is parallel to the low level frontal boundary and on the cool side of the boundary. The overall structure of the MCS is influenced by the shear. The low-level winds are often perpendicular to the boundary and convective line while the mid level (cloud bearing) winds are parallel to the boundary. This type of MCS is usually associated with a stationary front, pre-existing outflow boundary or slow moving warm front.

It often occurs when the frontal boundary is oriented in a northwest to southeast direction. In this type case there is usually a trough or even closed low to the east which helps keep the mean mid level flow parallel to the low-level frontal boundary. A strong southwesterly low level jet is often impinging upon this boundary. Most HPC forecasters would think of this type of MCS as a Maddox frontal type.

From Shumacher and Johnson 2006

The second type they identified was the backbuilding/quasi-stationary archtype. They identified it as a separate archetype than the leading stratiform type but it is in many respects similar. Both types typically have propagation due to new cell growth act in the opposite direction as the movement of the individual cells which approximately move with the mean wind.

From Shumacher and Johnson 2006

The two main differences between the original Parker Johnson leading stratiform archtype and the back-building MCSs is that during the latter , its band of convection is typically oriented more parallel to the mid level flow than the leading stratiform which tend to be more perpendicular. My view is that the original Parker and Johnson leading stratiform type is rare. The back-building archtype is probably more common than the Parker and Johnson trailing stratiform type and produces more extreme rainfall events. It often develops a stratiform shield on its down-shear side. Generally, MCSs having stratiform precipitation on the east or northeast side are more likely to produce extreme rainfall than MCSs with their stratiform precipitation on the west side of the system.

Doswell and his colleagues (1996) have discussed how the shape and movement of an MCS can help determine the rainfall that falls at any point. They developed the diagram below to explain how the shape of an MCS and its orientation with respect to its direction of movement can modulate the precipitation. The light blue dots represent the location where the theoretical rainfall rates would be observed for the each MCS. The arrows represent the direction of movement of the convective system and the red areas represent the brightest radar returns. The shaded grey area is the stratiform precipitation that often develops with various linear MCSs. When the convection is oriented normal to system’s direction of movement (the left hand side of the figure), a relatively brief period of heavy rainfall is expected. Conversely, when an MCS is oriented parallel to its direction of movement, a more prolonged period of intense rainfall is likely.

The alignment of the low level convergence upstream with respect to the 850-300-hPa mean flow also appears to modulate precipitation (Junker et al. 1996). Low level convergence upstream from your location allows promotes the development of new cells providing the airmass is sufficiently moist and unstable. The Junker et al. figure below shows the length of the low level convergence that was parallel to the 850-300-hPa mean wind upstream from where the max precipitation was observed (vertical axis) and the analyzed precipitable water at the location where the heaviest rain was observed (bottom axis) just prior to Midwest rainfall events during the June-September period of 1993. The zeros plotted on the diagram represent events in which no 3 inch area was reported while the fours represent events in which there was a 3 inch or greater area of at least 3600 sq. nautical miles. The findings above suggest that the larger scale extreme rainfall events are more likely when a boundary is parallel to the mean flow of the cloud bearing layer. Such an alignment increases the likelihood that cells will train along the same axis.

Which is the most untrue statement about a trailing stratiform

A) they are usually aligned normal to the mean flow

B) they are more likely to be associated with training of cells than a parallel stratiform MCS

C) they move faster than the mean flow within the cloud bearing layer

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How long will an MCS last before dissipation

One of the most difficult forecast problems associated with forecasting the precipitation associated with an MCS is deciding how long it will last. Will it quickly weaken and dissipate or will it last into your forecast period. In the shorter time ranges, monitoring satellite and radar imagery to see whether the associated cloud tops are warming or whether the radar echoes are weakening may offer some clues about a system’s future. In the past forecasters have often relied on climatology to try to anticipate when the convection associated with an MCS will weaken and die. For example, many MCS weaken over the Plains and Mississippi Valley during the 1200-1800 UTC period. However, that presupposes a knowledge of the regional climatology of MCSs which may not exist and also assumes that all MCS over a region behave similarly.

Recent work done by Evans and Doswell (2001) and Gale et al. ( 2002) suggest that the storm-relative inflow may also be important in determining whether a MCS will be long lasting or not. Both the surface system relative inflow and elevated system relative inflow tend to decrease as an MCS nears dissipation, due in part because of the decrease in the speed of movement.

From Gale et al. 2002

The figure above show box and whisker diagrams of the surface system-relative inflow (SSRI) versus time until dissipation for MCSs north (left figure) and south (right figure) of a boundary. The SSRI for long lasting MCSs is much stronger for the long lasting MCSs versus the shorter lasting ones, especially south of the boundary. The same holds for the elevated system relative inflow (ESRI) as shown below, especially those in the warm sector.

From Gale et al. 2002

From Gale et al. 2002

System relative inflow is strongly related to system movement and the mean wind of the cloud bearing layer. Therefore, rapid moving systems should be associated with strong system-relative inflow. Therefore, it is not surprising that faster moving MCSs typically last longer than slower ones since they should be associated with stronger system relative flow.

From Gale et al. 2002

From Gale et al. 2002

Gale et al. also noted that MCSs that were affected by the low-level jet at some time during their life cycle often weakened once they were no longer being influenced by the low level jet. Systems with strong system-relative low-level inflow on the forward side of the system and weak storm relative flow at mid levels last longer than systems having weaker low-level storm-relative inflow. The Gale et al. work also suggested that equivalent potential temperature advection might be a potential predictor of MCS dissipation since its advection decreased prior to MCS dissipation.

Rotunno et al. (1988) has theorized that the balance between the line-normal shear (the shear perpendicular to the linear convective system) and the evaporation induced cold pool circulation determines the structure of long-lived convective systems. Weisman (1992) has suggested that the evolution of the rear inflow jet and whether it has descended to the surface or remained elevated could also play a role in how long a convective system lasts before dissipation. Weisman and Rotunno (2004) have suggested that surface based shear distributed to greater than 5-km is detrimental to producing strong convective systems and that under these conditions that 2-D simulations show little sensitivity to changes in the upper-level shear. However, the cold pool-shear (low level) balance ratio theory of Rotunno et al. has not been supported by observations (see figure below) and appears to play little role in determining how quickly an MCS will dissipate. Note on the figure below how the mean cold-pool shear balance ratio is almost the same for the MCS that did not dissipate for 12 to 15 hours as for those that dissipated within three hours.

From Gale et al. 2002

Two articles by Coniglio and his co-authors (2001, 2006) indicate that adding upper-level shear to a wind profile with weak to moderate low-level shear increases the vertical displacement of parcels despite the decrease in the vertical velocity along the cold pool interface. This research suggests that looking at the wind profile through the entire troposphere needs to be considered when trying to analyze how long a linear MCS will last.

From Coniglio et al. 2006

There were differences in the structures of the storm-relative hodographs between mature MCSs and decaying ones. There was significant unidirection shear associated with mature MCSs with a mean steering level near 8-km which the article notes is often tied to pre-existing upper level jets. The upper-level shear decreased significantly as MCSs decayed. Coniglio et al. note that the decrease in upper level shear can often be related to the propagation of the system away from the upper level jet.

Typically, MCSs weaken and dissipate when they move into a region of reduced low-level moisture support and weaker forcing (Cotton et al. 1989). Each of these factors can often be related to weakening of the low-level jet either through diurnal weakening or by the MCS moving away from it. However, the former can also be a result of the weakening of the system-relative moisture inflow. Another factor in weakening may be that the system moves into an area where the airmass is less unstable and may actually be stable. These unfavorable conditions lead to the weakening and eventual decay of the system.

The models typically do a much better job handling the temporal evolution of convective systems that are associated with strong mid-level forcing and do a poorer job in handling systems that have weaker forcing. However, even for the stronger systems, the models may still have trouble handling an MCSs life cycle since they often have problems in predicting the stability (ie. CAPE) and also lack the sophisticated cloud microphysical processes needed to correctly simulate the development of the rear inflow jet. Therefore, a model may underplay the speed at which the MCS moves and therefore underplays the system relative inflow.

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Propagation and movement
Propagation is influenced by a number of factors. One of the most important of these factors is where low level moisture convergence and instability is located relative to the initial convection or convective system. New cells form where there is moisture, instability and lift. The cold pool and resulting outflow boundary from the convection associated with the MCS plays and important role in determining where the low level convergence and instability will be located.
The system-relative inflow into the MCS also helps in determining the region of strongest moisture convergence and ascent. Finally, non-hydrostatic pressure gradients and buoyancy gradients may play aid in generating the rear inflow jet which in turn plays a role where the strongest low-level convergence is located relative to the system.
The following paragraphs will attempt to provide insight into how to operationally forecast the approximate direction and speed of propagation that will be associated with an MCS, and how the resulting propagation can be used to help forecast the movement of the MCS. System movement depends on cell movement and propagation. Individual cells usually move at around 90% of the mean wind with a slight deviation to the right. Finally, propagation is affected by what flank of the storm the new cells are forming on and by the rate of development of the new cells.
Chappell (1986) has noted that it is convenient to represent the movement of a convective system as the sum of the vector representing the movement wind flow and another vector representing the effects of propogation.

Original Corfidi vector method

Fankhauser (1964) noted that the movement of individual convective cells could be approximated by the mean wind of the cloud bearing layer. More recently, Corfidi et al (1996) suggested that during MCS initiation the movement of the convective cells can be approximated by the mean wind between 850- and 300-hPa as defined by,
Vcell=(V850+V700+V500+V300)/4
Corfidi et al. created two scatter diagrams showing that there is a high correlation between the 850-300 hPa mean wind speed and direction at the time of convective initiation.

However, it is important to note that for convection with unusually low cloud tops using the 850-300 hPa mean wind may be a poor approximation for cell motion since the cloud layer may be considerably lower, and the clouds move with the winds within the cloud bearing layer. In addition, for potentially deeper clouds (equilibrium levels higher than 300-hPa), you may need to use 850-200 hPa layer for mean wind calculations to represent the movement of individual cells.
When you are trying to predict the motion of individual cells, remember to try to anticipate whether you will be dealing with low topped convection or not. The equilibrium level provides the top limit to the cloud bearing layer while the LCL can be used to estimate the bottom of the cloud bearing layer. The next step is to find a way to approximate the propagation vector. The scatter diagram shows that the direction of propagation is anti-parallel (in the opposite direction) to the low level jet (see figure). However, there were two significant outliers. The big question then becomes how to determine the magnitude of the vector. Corfidi et al. (1996, henceforth called (CMF) reasoned that the stronger the low level jet, the stronger low level convergence should be along the flank of the MCS that is being impinged upon by the jet. They therefore assumed that the propagation vector was equal to and of opposite sign as the low level jet. Therefore, Vprop= -Vllj

They then tested the following formula to see whether it provided a good estimate of the motion of the most active elements of the convection system.
VMCS=Vcell-Vllj
Where VMCS stands for MCS motion, Vcell stands for motion of the individual cells which is approximated by the mean wind between 300 hPa and 850 hPa, and Vllj is the vector describing the low level jet.
CMF also found fairly high correlations for both speed and direction using the formula (see figure below). This formula has sometimes been called the original Corfidi vector method or the CMF vector method. One important point to make is that most of cases of the original data set were large scale MCCs similar to those found by Maddox et al. Therefore, many of the cases were slow moving, backward propagating MCSs. Also, the low-level jet they found during various cases was often not at 850-hPa. It is a mistake to assume that the low-level jet will always be located at 850-hPa when using the technique. Therefore, it is important to look at a sounding or forecast sounding to define the level of the low-level jet before assigning the propagation vector. If the low level jet has the same general direction and magnitude as the 850 winds, then the approximation on the HPC workstation can be used if you expect a backward propagating system. The original technique works best when moist, unstable air is located upstream from the initial area of convection, while cooler, more stable air is found downstream from where the initial thunderstorms associated with the MCS are expected to develop.

The Revised Corfidi Vector method

More recently, Corfidi (2003) has noted that the wind profiles associated with quasi-stationary convective systems was sometimes very similar to the vertical wind profile found for the Trailing Stratiform (Parker and Johnson TS type) MCSs (derechoes) studied by Johns and Hurt (1987). Therefore, the original CMF method underestimated the speed of movement for a number of these forward propagating MCSs. The original technique does not work as well when: 1) there is moist unstable air downstream from where the initial thunderstorms develop or are expected to develop, 2) when the mid level flow is relatively strong, 3) when there is a dry layer at mid levels.
Corfidi (2003) has suggested that the original vector technique would sometimes seriously under-estimate the motion of a squall line because it failed to account for propagation that occurs in the down-shear or downwind direction. A forecaster needs to consider the direction of greatest system-relative convergence and the distribution of surface-based instability when applying the vector technique. The latter is often determined by the evolution of the cold pool and associated outflow boundary or gust front. Corfidi offers a hypothetical example of the evolution of a gust front associated with a fairly uniform vertical wind profile (see figure below). In such a case the leading edge of the gust front will move with the mean wind speed.

Along the leading edge of the gust front, a forward propagating convective system is typically found. However, farther west along the gust front, outflow boundary may become aligned parallel to the mean flow allowing training of cells. Essentially, the vector method says that if the air mass is unstable downstream of the initial convection, then a forward propagating MCS

is a good bet when the mid-level wind field is relatively strong, especially if there is a dry layer to promote the downward transfer of momentum. The resulting motion of the MCS is the sum of the mean wind and a system-relative propagation vector which is dependent on the mean wind and the low-level jet. For example, if the convection is moving east at 40 kts, the system is experiencing inflow of similar magnitude assuming there is no low level wind. Relative to the west-to-east moving convective system, easterly winds are feeding into it. If the airmass is moist and unstable on the east side of the system, then the system relative inflow will produce moisture convergence and convection along the leading edge of the convection.
The inflow due to the system motion acts the same way that the low-level jet worked in the original CMF vector method. System movement imparts a system-relative inflow of equal magnitude to the system movement but in the opposite direction. To a west-to-east moving system, this is analogous to having an east wind on the east side of the system. Remember in the original CMF vector method, the propagation vector was equal in magnitude but opposite in direction to the low level jet. In the revised technique, the system-relative low-level wind acts similarly. Therefore, the MCS movement vector from the original MCF method is added to the vector that describes the mean wind. Essentially, in the revised vector method you double the mean wind and then subtract the low level jet from that sum.
VMCS=2(Vcell)-Vllj
When using the vector method it is important to remember that the decision on whether to use the original CMF method or the revised method should be dependent on the type of MCS that is anticipated. Most of the time the original vector method should work best for back-building convective system or for systems in which the propagation due to new cells forming on the rearward flank acts to slow system movement.

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Propagation and movement (continued)

The N-AWIPS workstations have versions of the Corfidi method based on model forecasts. The old version is available on NTRANS if you click on the GFS or NAM and then click on QPF. The c-vectors are located where it says PW, PMSL, C-VEC. The revised and original methods can be found using NMAP2, click on grid and then the model you want and then QPF. There are a few caveats that need to be kept in mind when using the NTRANS and NMAP versions of the C-vector technique. The first is that the method assumes the mean flow of the cloud bearing layer is 850-300 which may not be true. Secondly, the method assumes that the low level jet is located at 850 hPa, which is an approximation. Finally, the method uses a low level jet which is taken at the same location as the mean flow. Corfidi’s research used a low level jet a little south of the location of where the convection developed. Therefore, the method sometimes may underestimate the impact of the low level jet. While the revised method is an improvement to the original method when dealing with forward propagating MCSs, it sometimes moves them too quickly. Still the method often offers a decent estimation of which MCSs are the most likely to be slow moving.

An example of a backbuilding/quasi-stationary convective system

During the Des Moines flash flood that knocked out the water supply during the 1993 Great Midwest flood is a good example of a back-building of quasi-stationary convective system. Note that at 21 UTC the first cell started to develop in Iowa and that new cells formed to the west between 2100 UTC July 8 and 0200 UTC July 9.

From Junker and Schneider, 1997

However, with time its character changed and it evolved in an MCS with considerable training as the initial MCS over Iowa merged with another MCS moving eastward from Nebraska. MCSs often change character during their evolution.

From Junker and Schneider, 1997

1000-850 mb layer mean moisture flux (vectors)moisture flux magnitude (dashed) and moisture flux divergence (-4 x10-7s-1 are shaded), the red dot represents the location where convection started

In this case, the strongest instability was to the southwest of where the initial convection developed and the low level moisture convergence was located to the west of the initial cell. Such a combination favors quasi-stationary convection.

From Junker and Schneider, 1997

Mean sea-level pressure analysis and observed moisture convergence (highest values are shaded), the red dot is the location where the first cell formed.

When new convection keeps developing on the west or upstream side of the system, the system moves slower than the mean flow and often leads to stratiform clouds forming on the downstream side of the convection.
For short range forecasts, it is important to look at what is happening to the low level convergence. If it strengthening to the upstream of the initial convection, the system has a good chance of having new cells form to the west of the initial ones.

Mean sea-level pressure, 850-300 hPa mean wind (flag=50 kts, barb=10 kts, half-barb= 5 kts) and 982-hPa equivalent potential temperature. Red dot is where the initial convective cell developed.

In this case the 850-300 hPa mean wind was angled slightly away from the boundary. Chappell (1986) has noted that such an orientation allows the initial convection to move away from the boundary as new cells form along it. When the mean flow is absolutely parallel to the boundary, the cold pool associated with the outflow sometimes may disrupt the boundary especially when the mean relative humidity is initially marginal for heavy rain and when the CAPE is high. As the boundary sinks southward, new cells advecting with the mean wind are more likely to track across new areas that have not already received rain rather than across areas that have already received rain.

Which type of MCS is the least likely to produce a significant flash flood

A) A linear MCS that ends up having a trailing stratiform area

B) a parallel stratiform MCS type system

C) one that evolves into a MCS with a leading stratiform area

At the very shortest time ranges, extrapolation of radar and satellite imagery often works reasonably well for forecasting MCS movement. The observed orientation of the boundary and how cells along it are forming with respect to the initial convection can be used as input into short range (0-6 hr) QPF. However, after an hour or two, extrapolation of the current trends rarely works well because systems do not move or evolve linearly. Complex non-linear processes determine how MCSs move and also play a role in determining where new cells will form in relation to the system.

How much rainfall an MCS will produce is dependent on its shape and size, rainfall rates and movement. The latter is determined by the rate at which cells are advecting with the cloud bearing winds and the rate at which new cells form along some flank of the storm (propagation). The importance of propagation on the evolution and movement of MCSs will be discussed in the next section.

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850-hPa winds (full barb=10 kt and half barb=5 kt), Θe (solid, every 4oK), lifted index (values less than –8 are shaded), front (thick line) and surface trough (dashed line) at 00 UTC.

Propagation and movement (cont.)

When using the vector method it is important to remember that the decision on whether to use the original CMF method or the revised method should be dependent on the type of MCS that is anticipated. Most of the time, the original vector method should work best for back-building convective system or for systems in which the propagation due to new cells forming on the rearward flank acts to slow system movement.
Forward propagating systems are more likely when:
1) When the 850-300 hPa mean winds are moderate to strong through a deep layer.
2) When a progressive shortwave is translating to the east.
3) When the air mass upstream of the developing convection is strongly capped but the air mass downstream is unstable and less strongly capped.
4) When the low level theta-e ridge is located coincident and downstream of the system.
5) When the low level convergence is located on the downstream flank of the convection or is expected to remain on the downstream flank. If the mean flow is moderately strong and there is an elevated dry layer, usually any convection that develops will move fast enough to keep strong moisture convergence going along the leading edge of the convection.
Unfortunately, MCSs and the effects of propagation are much more complicated than presented above. Convective systems sometimes exhibit forward and backward propagation at the same time (see figure below). This is most common when wind fields are unidirectional because then the cold pool can become elongated. Corfidi offered this hypothetical example showing how different types of MCS can form along the same gust front.

In this case, the elongation of the cold pool allows its northwestern portion to become parallel to the mean flow even as the forward section of the cold pool boundary remains perpendicular to the mean flow. The portion of the boundary that is perpendicular to the mean flow then becomes a forward propagating MCS while the portion that is parallel to the mean flow becomes a focus for training of cells and slow movement (Corfidi, 2003)

Knowledge of the effects of propagation and being able to anticipate whether a convective system will most likely be a forward or backward propagating system can be used to help assess model QPF. Currently, when numerical models predict slow movement and very heavy rainfall in association with an potential convective system, a forecaster is left to wonder whether the model is suffering from grid-scale feedback or not. If the pattern and model forecast soundings suggest that a rapid moving forward propagating TS type MCS is most likely, then a forecaster can have some confidence in cutting back the numerical model QPF and also have confidence in moving the MCS and its associated precipitation shield more quickly than forecast. Conversely, when the diagnosis favors a system where propagation effects should lead to back-building or training of cells, the forecaster might wish to beef up the model QPF.

The loop above shows an example of how the outflow boundary of a forward propagating system can lead subsequent quasi-stationary convection as the cold pool aligns with the mean flow. Note on the loop that a fast moving MCS dives southeastward across Wisconsin into Minnesota. Later in the loop, new convection forms in northwestern Wisconsin where training of cells occurs and over 3 inches of rain were reported. The outflow boundary enhanced the front across the region. The strong low level jet and differential thermal advection associated with it, then produced another outbreak of convection during the night.

For an exercise on which vector method click here.

For a section that discusses factors that govern how long an MCS might last, click here

850 analysis valid 00Z, 24 July 2005

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Exercise 1: Forecasting convective movement

In this exercise you’ll be asked to predict where a convective system is likely to develop and then predict the speed and direction in which it is most likely to move. The exercise is meant to help forecasters learn clues that might aid them in forecasting what type of MCS is most likely and help decide which vector method will most likely provide the best guess at where the MCS will track.

A NAM 2100 UTC MSLP and 1000-500 hPa thickness forecast valid at 2100 UTC and an HPC MSLP analysis verifying at the same time as shown below. The red and blue dot are location where NAM model soundings were secured to aid in the forecast

NAM MSLP forecast valid at 2100 UTC, the red and blue dot locations is where model sounding below were chosen.

2100 UTC HPC surface map

From the 2100 and 0000 UTC observed and forecast maps, try to determine where you think the front is located and note any differences in the pressure and low level convergence patterns that might have an impact on where convection might develop.

NAM MSLP forecast valid at 0000 UTC

0000 UTC HPC surface map

2100 UTC boundary layer winds, moisture convergence and Θe

0000 UTC boundary layer winds, moisture convergence and Θe

Do you notice any areas where convergence might differ. Think why any difference might be important?

The observed surface pressure pattern and wind fields at 2100 and 0000 UTC suggest that the pattern of low-level moisture convergence forecast by the NAM might be in error. For example, the observed wind field suggests low-level convergence is taking place along the front in northeastern Wyoming.

Forecast sounding valid 2100 UTC at the blue dot

Forecast sounding valid 2100 UTC at the red dot

2100 UTC forecast sounding at the location of the blue dot in southeast SD on the earlier maps. A few important details about the sounding: CAPE=4367 J/kg, PW=1.16”, LCL=801mb, Mean RH=35%, Mean LRH=51%

2100 UTC forecast sounding at the location of the red dot along the ND/SD border on the earlier maps. A few important details about the sounding: CAPE=3239 J/kg, PW=.93”, LCL=745mb, Mean RH=38%, Mean LRH=30%

Using the sounding on the left (blue dot location), If you were trying to determine the mean winds of the cloud bearing layer, what layer would probably be the smartest to use?
The relatively low RH suggests that there will be decent cold pool generation by any convection that develops. There also is good potential for the downward transport of momentum to bring some of the stronger winds at mid levels down to the surface.

A) 850 hPa-300 hPa

B) 850 hPa-200 hPa

C) 800 hPa-200 hPa

Using the sounding on the right (where the red dot was located), which of the following would not necessarily point you towards using the revised vector method to predict MCS movement across South Dakota?

A) The high CAPE of 3239 J/kg

B) The relatively strong unidirectional mean winds (approximately 40 kts)

C) The low mean relative humidity and low relative humidity of the subcloud layer

At 2100 UTC the model is forecasting one vorticity max over the central portion of the Dakotas, another over northern Wyoming and third, much stronger shortwave, extending from southern Canada southward into northern Idaho (see figures below). Based on the forecast cape, vorticity, and boundary layer convergence where would you expect convection to form within the next 3 hours? Will it be in southern South Dakota, eastern Wyoming or eastern Montana? Where the initial convection forms is important in determining the archtype of MCS that you are likely to be dealing with.
One of the big problems in forecasting the evolution and movement of an MCS is trying to predict where convective initiation will occur. The models often have problems predicting stability and can make mistakes in their predictions of the low level convergence, especially in the warm season.

The model initialization may also not have the some important boundaries very well defined. Note the lack of low level convergence with the front in western North and South Dakota.

NAM CAPE and CIN forecast V.T. 2100 UTC

In what way might the low level convergence near the front change the stability from what the model is forecasting.

A) CAPE forecasts might be lower and CIN might be lower in southeastern Montana, northeastern Wyoming and western South Dakota than predicted by the NAM.

B) CAPE forecasts might be higher and CIN might be lower in southeastern Montana, northeastern Wyoming and western South Dakota than predicted by the NAM.

NAM 500 height and vorticity forecast V.T. 21Z

NAM 500 height and vorticity forecast V.T. 00Z

Observed 500 mb heights V.T. 00Z

Click on the radar imagery and compare the radar features shown on the loop to the evolution of the NAM forecasts of vorticity between 21Z and 00Z. Does the vorticity look real? Note that the radar echoes are decreasing as the convection moves into the region where the NAM is showing the strongest low level convergence and instability. The convection in southeastern Montana is moving eastward and looks like it may still be increasing. Your next task will be to decide which vector method you think would work best for different areas of convection.

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Exercise 1 continued

3 hr accumulated precipitation valid at 2100 UTC

3 hr accumulated precipitation valid at 0000 UTC

The red dot is the location of the soundings below

Radar image valid 0000 UTC

IR Satellite image valid 0000 UTC

Radar and satellite Imagery are indicating that an MCS is developing over eastern Montana and that a few cells are also developing over central South Dakota

3 hr NAM accumulated precipitation valid at:

0300 UTCZ

0900 UTC

1200 UTC

The NAM forecasts very little precipitation with the blob moving eastward from MT and instead forecasts its development over eastern SD into MN. First, concentrate on the MCS developing in MT and try to forecast its movement from the following model output.

0000 UTC forecast sounding at the location of the red dot along the ND/SD border on the earlier maps. A few important details about the sounding: CAPE=1338 J/kg, PW=1.14”, LCL=690mb, Mean RH=49%, Mean LRH=37%

0300 UTC forecast sounding at the location of the red dot along the ND/SD border on the earlier maps. A few important details about the sounding: CAPE=1320 J/kg, PW=.95”, LCL=727mb, Mean RH=31%, Mean LRH=39%

NAM CAPE and CIN valid 0300 UTC

Using the sounding above and forecast fields, decide on which vector method to use from the maps below, the original or revised vector methods. Then, use the following Corfidi vector wind forecasts to forecast the movement of an MCS that is moving from Montana into the Dakotas. The new vector method winds are on the left hand side and the original vector method winds are on the right.

NAM 0300 UTC revised method winds

NAM 0300 UTC original CMF method winds

NAM 0600 UTC revised method winds

Now that you picked which method to use. Forecast the direction and speed of the MCS as it moves across the Dakotas.

Which of the following do you think will be closest to the actual speed and movement of the MCS .

A) 270o (east) at 15 knots

B) 270o (east) at 30 knots

C) 270o (east) at 40 knots

0200UTC radar

1000 UTC radar

The MCS moved at about 45 Knots to the east.

Below is a list of questions you should ask when trying to decide whether of forward propagating MCS is more likely than of backward propagating one.

1.) Are moderate to strong mean winds present in the cloud bearing layer.
2.) Is there a dry layer present in the mid-levels or in the subcloud layer?
3.) is there an area of favorable CAPE coincident or downstream of the initial convection.
4.) Is a 850 hPa theta-e ridge downstream or coincident with the convection.
5.) is the low-level jet and axis of strongest low level moisture flux coincident or downstream of the convection.
6.) Is the MCS moving with a mid-level trough.
7.) Is there and area of low-level convergence located along and downstream of the convection.

Adapted from Moore’s COMET presentation

The more of these questions that you can answer with a yes, the better the probability of a forward propagating MCS and the better the revised vector method will work.

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NAM 850 wind, Øe and Øe advection valid 0300 UTC

NAM 0600 UTC CMF original method winds

MCS movement exercise 2

Use the following maps and forecast soundings to :
1) decide which Parker et al. archtype of MCS to predict:
to forecast the MCS’s movement using the appropriate vector method
Decide whether to increase or decrease the precipitation that the models are forecasting with the system.

1800 UTC

2100 UTC

0000UTC

0000 UTC 500 hPa analysis

NAM 2100 UTC CAPE and CIN forecast

It’s very unstable along and south of the front with 3000-4000 J/kg CAPE across Oklahoma. The loop below shows an infrared satellite loop for the area of concern.

Double click the left button to play. To view the last image of the sequence, pull the drag bar all the way to the right. The loops plays from 1215UTC to 0015 UTC

NAM 2100 UTC 850-hPa wind, Øe and Øe advection forecast

NAM 0000 UTC 850-hPa wind, Øe and Øe advection forecast

Where the convection first develops is often important in determining the archtype of convection that develop. Based on the satellite loop along with the other charts think of some questions that you should ask when trying to forecast where the first convection might fire?

NAM 0300 UTC 850-hPa wind and total precipitable water forecast advection forecast

Concerning convection initiation , which statement is the least likely to be true.

A) A good place for convection to develop is just north of the 850-hPa Øe ridge in the southern portion of the region of Øe advection.

B) The high forecast CAPE and precipitable water values suggest convection will first form south of the analyzed front.

C) The axis from southeastern Kansas to Arkansas might be more favored than northeastern Oklahoma for initiation

0230 UTC radar

0215 UTC Infrared imagery

0330 UTC radar

0315 UTC Infrared imagery

The first convection develops over across southeast Kansas and Arkansas along and north of the boundary. Once you’ve determined where the convection will start, the next step is to determine what type of system will develop.

NAM 0300 UTC 850-hPa wind and total precipitable water forecast advection forecast.

Note where the initial convection broke out in Kansas on the 0315 UTC and 0330 UTC imagery above and then look at the forecasts of precipitable water and 850-hPA winds. One clue to where the first convection might occur was the apparent development of altocumulus clouds or cumulus clouds over southeastern Kansas.
Note the location of the axis of the low-level jet compared to the initial convection. Where does the axis of low level jet impinge upon the frontal zone, is it east or west of the initial convection?

NAM 0300 UTC 850-hPa wind, Øe advection and moisture convergence forecast

NAM 0600 UTC 850-hPa wind, Øe advection and moisture convergence forecast

From the forecast fields above which MCS type do you think is more likely.

A) A forward propagating MCS

B) A training line-adjoining stratiform MCS

C) A parallel stratiform system

Which of the two Corfidi techniques will probably work best?

NAM 0300 UTC 850-300 hPa mean wind forecast

Valid 0300 UTC

Valid 0600 UTC

The NAM forecast of the maximum precipitation is obviously going to be too far southwest and it will also be too light.

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To see what actually happened. Click here.

NAM 0000 UTC CAPE and CIN forecast

NAM 0600 UTC 850-hPa wind and total precipitable water forecast advection forecast.

A) The original CMF method

B) The revised technique

Exercise 2 recap

Note how slow moving the MCS moves between 0315Z and 0715Z. The revised method predicted 20 to 25 kt movement to the east-southeast while the original method predicted 10 to 15 kts.

NAM 0300 UTC original CMF vectors

NAM 0300 UTC revised method vectors

0315 UTC

0415 UTC

0715 UTC

Note how the MCS radar signature took on a configuration with the brightest radar echoes on the western edge of the MCS and that new cells kept forming over Kansas where the initial convection started for a prolonged period. The original vector method used the most active beta elements to track the speed of the MCS. During this case even though individual cells were moving southeastward at around 25 kts, the MCS only moved around 10 knots. Also, the brightest clouds were on the edge of the convection where the strongest instability and low level convergence were located. It had the look of a leading stratiform type convective system or a cross between a training line-adjoining stratiform and back building type of MCS.

Backward propagating (upshear propagating) systems often produce a heavier core of rainfall than the models. In this case, both the NAM and GFS underpredicted the rainfall associated with the MCS. The GFS had a better forecast of the axis of heaviest rainfall but it only predicted .50 inches. The 1.00 inch or greater area extended from southeastern Kansas into northwestern Arkansas.

12-36 hr accumulated precipitation forecast from the NAM

12-36 hr accumulated precipitation forecast from the GFS

The GFS often forecasts its rainfall axis to the north of the NAM. Observations of the HPC forecasters is that if it has a bias, it is to be too far north with its axis. However, for this case, it’s more northerly axis was correct.

24 hr accumulated precipitation for the same period

The NAM does not have a consistent northward or southward bias. However, the model does have bias of not predicting heavy enough rainfall with convection because it tends to overpredict light rainfall which releases the instability too early. The NAM also sometimes will predict the rainfall axis too far south when strong instability (high CAPE) is present along a front where there is also strong CIN. That appears to be what happened in this case.

Neither model predicted heavy enough rainfall during this event. The two small observed 2.00 greater areas on the analysis illustrate how small scale the core of heaviest rainfall associated with an MCS can be.

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An ingredients based methodology for forecasting precipitation associated with MCS’s

Funk (1991) documented some of the techniques that are used within HPC to forecast quantitative precipitation associated with convective rainfall events. He emphasized the importance of conceptual models such those developed by Maddox et al. (1979) and others as important tools in the forecast process. More recently Doswell et al. (1996) have suggested using an ingredients based methodology to forecast flash floods and extreme rainfall.
In their article Doswell and his colleagues note that there are many different type of flash flood producing storms. In particular, they document that supercells, squall lines, MCSs of many shapes and even non-convective systems can all produce flash flooding. Therefore, pattern recognition and the use of conceptual models should not be used in a vacuum. Forecasters need to understand why various patterns are favorable of extreme rainfall to keep from using a particular pattern when a basic ingredient is missing that will prevent that system from producing heavy rainfall for a sufficient period of time to produce flash flooding. Also, applying an ingredients based methodology should also allow a forecaster to predict heavy rainfall from events that do not quite fit any conceptual model but still produce extreme rainfall.
Therefore, the ingredients based methodology espoused by the Doswell et al. article will be reviewed. Chappell noted (Doswell et a. 1996) that the heaviest rain falls were high rainfall rates are located for the longest time. Factors that govern rainfall rates and the duration of the rainfall therefore need to be explored.
Rainfall rates are dependent on the vertical moisture flux that is being fed into the cloud. Therefore, high rainfall rates require that a high moisture content (measured by mixing ratios or precipitable water) be present along with strong ascent. A continued feed of moisture is needed to sustain high rainfall rates for a sustained period. MCSs associated with sustained heavy rainfall rates are usually located just north of an axis of strong low-level horizontal moisture flux.
The two figures above are the 850-hPa moisture flux (left) and 850-hPa moisture flux convergence just prior to the development of MCSs that produced more than 100 mm of rainfall during the night (Junker et al. 1999). The red dots are the location where the subsequent rainfall maximum was observed. The horizontal and vertical lines represent the degrees latitude and longitude (every 2 degrees) from the center of the heaviest rainfall. Strong moisture flux convergence is located along the northern edge of where the 850-hPa moisture flux convergence decreases. The combination of strong horizontal moisture flux and low-level moisture convergence implies strong vertical moisture flux is taking place.
However, rainfall rates are also dependent on the precipitation efficiency (PE) of the system defined as the ratio of the mass of water falling as precipitation to the total mass of the moisture flux that is ingested into the convective system during its life cycle. The biggest single factor the lowers the precipitation efficiency is evaporation (Doswell et al. 1996). They suggest that the key observable factor that governs evaporation is the environmental relative humidity. Convective cells during days when the mean relative humidity is relatively low should experience lower precipitation efficiency than days when the mean relative humidity is high. Fankhauser (1988) found that systems with lower cloud bases were more efficient than ones with higher cloud bases.
More recently, Work by Market and Allen (2003) suggests that the apparent relationship between cloud base height and PE may be due to the mean relative humidity (RH) of the layer from the surface to the LCL (subcloud layer). When the mean RH in the subcloud layer is low, usually the cloud base will be higher than if the subcloud layer RH is high (see figure). The higher cloud base means that precipitation will be falling through a deeper unsaturated layer. The also found that shear had a negative impact on precipitation efficiency.

Rainfall rates

Precipitation Efficiency

From Market and Allen 2003

There are also microphysical processes that may impact precipitation efficiency. The air masses origins and whether the cloud is undergoing warm or cold rain processes also plays a role in precipitation efficiency. As Braham (1959) has noted, the all water (warm rain mechanism) starts as soon as a parcel of cloud is formed and continues throughout of the life of the cloud while cold rain process only operate after the cloud top temperature is colder than about –10 to 15oC. The origins of a cloud are important because maritime air masses have a much larger number of large hygroscopic nuclei composed of various salts than a continental air mass. These large hygroscope nuclei are more efficient at producing droplets than the smaller nuclei found over continental regions. Braham notes that there is an order of magnitude decrease in the number of these particles as maritime air moves from the Gulf of Mexico to Missouri during the summer months.

Paradoxically, during convective events where warm rain processes are more common, the raindrop size, while large, is limited as raindrops reach a critical size through collision coalescence and they break up into smaller drops. By contrast, raindrops that once where associated with hail at some level tend to have a larger variation in droplet size and have bigger drops associated with them. This is the reason why different ZR relationships are used for continental convection and tropical storms and hurricanes. The following figures illustrate how the reflectivity is impacted by droplet size and density for the two different types of air masses.

Florida

Arizona

From Kelsch lecture

Note that the droplets over Florida have a lower range of mean drop diameter than the Arizona figures while also having a higher drop density per cubic meter than the Arizona figures. The figures suggest that the relationship between droplet density and rainfall intensity is better for the more maritime air mass than the extreme continental air mass.

The scale of the mesoscale precipitation system also probably affects PE as small isolated convective systems are more likely to have dry air entrained into their core than larger systems. This might explain the negative correlation between Convective Inhibition (CIN) and PE found by Market and Allen (2003).

While the environmental lapse rate must be conditionally unstable for free convection to occur, high CAPE does not necessarily favor high PE. The max updraft associated with convection can be estimated from CAPE using the formula below..
Convection associated with extremely high CAPE may produce such intense updrafts that some of the mass (moisture) may be ejected out of the top of the cell only to advect downstream where it may evaporate. Therefore, the shape or positive area of the sounding may actually to some extent modulate precipitation efficiency. The two soundings below have identical CAPE. However, the sounding on the left with the rather fat, positive would produce a much stronger updraft than the CAPE on the right, especially in the lower portion of the cell. The stronger updraft would carry more precipitation upward and reduce the amount falling into the lower portion of the updraft. This would minimize the effect of moisture loading which normally weakens the updraft.
The long “skinny” positive area on sounding ‘B” would have a slower updraft acceleration and a taller thunderstorm. The thunderstorm associated with the skinnier looking sounding would also have less precipitation carried into the higher portions of the thunderstorm and would lose less mass (ice crystals) from the top. It therefore would probably be more efficient. Also the weaker updraft would allow for a longer time period in which collision-coalescence and warm rain processes could act.

A forecasters should not get the false impression that a thunderstorm having high CAPE and a fat look to the positive area of the sounding like the area shown on sounding A will not produce extreme rainfall and flash flooding. If sufficient moisture is available and is fed into the precipitation system, the prodigious amount of moisture processed in the updraft can lead to extreme rainfall and flash flooding even if the system has a low PE. The important factor is not the size or shape of the positive area but instead is whether the system can produce intense enough rainfall rates for a long enough period to produce run off problems. An environment having a very high CAPE and deep moisture (ie. high precipitable water) may produce intense rainfall rates despite having less than optimal PE.

Kelsch in a COMET lecture has noted that deep warm layer (the layer from the cloud base to the wet bulb freezing level) of at least 3-4 km tends to lead to increased precipitation efficiency. He suggests that such a deep layer increases the residence time within the cloud where collision-coalescence and warm rain processes are occurring. However, Market and Allen were unable to establish a clear correlation between Precipitation Efficiency and the depth of the warm layer.

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Original Maddox et al. MCS archtypes associated with flash flooding

Maddox et al. (1979 and 1980) studied 150 heavy rainfall events and identified three primary types of mesoscale convective systems as often producing excessive rain. Two (Meso-High and Frontal types) were primarily mesoscale phenomena; the other (Synoptic) was driven by synoptic forcing but still often occurred on the Meso-a scale instead of the synoptic scale. A fourth type, the cyclonic circulation system, was identified by Spayd and Scofield (1982).

Numerical models still often do not forecast explicitly forecast these mesoscale type of extreme rainfall events, but often show the synoptic pattern which is associated with their development. This is where conceptual models come into play enabling a forecaster to identify which synoptic features are typically present when mesoscale heavy rainfall events develop. However, as noted earlier, when using pattern recognition or conceptual models it is imperative to still apply an ingredients based methodology to forecasts as try to aniticpate whether the convective system will be a forward or backward propagating one.

A number of features were common to many of the excessive rainfall events studied by Maddox and his colleagues. They state:
1). Most flash floods were associated with convective storms.
2). Storms occurred where surface dew points were high.
3). Relatively high moisture contents were present through a deep tropospheric layer.
4). Weak to moderate vertical wind shear was present through the cloud depth.
5). Convective storms and/or cells repeatedly moved over the same area.
6). A weak, mid-tropospheric, Meso-a scale trough helped to trigger and focus the storms.
7). The storm position was very near the mid-tropospheric, large-scale ridge position.
8). Storms often occurred during the nighttime hours. Remember the bulk of the Maddox MCCs were located in the Plains

Synoptic type heavy rainfall events.
Synoptic type heavy rain events, Maddox et al. (1979), comprised about 20 percent of the flash flood events studied by Maddox and his colleagues. Synoptic type events are most common during the spring and early summer and again in fall and early winter. In synoptic type flash floods, a strong 500-hPa trough is usually moving slowly eastward or northeastward. A weak upper-level shortwave often lifts northeastward ahead of the main trough and weakens. Most of the associated height falls shift more northeastward then eastward. Therefore, any surface low lifts northeastward and the trailing front slows or stalls as it becomes parallel to the upper flow. Fronts attending these events are usually oriented southwest to northeast, but Maddox noted a few that extended west to east.

The typical surface pattern associated with synoptic type flash floods is shown above (top left). The yellow shading is where surface dewpoints are 70oF or higher. The boxes represent the flash flood threat area. Note that the heavy rains usually occur in the warm sector but that the area may be located near an old frontal boundary or dewpoint discontinuity. The front is usually slow moving and is aligned parallel to the mean flow which promotes training of cells.
Typical 500-hPa stream line pattern and 250-hPa wind maxima position during synoptic events is illustrated on the top left panel. A strong 500-hPa trough is usually located to the west. Often an initial shortwave lifts northeast and another one drops into the trough holding the southern portion of the trough almost stationary. The second shortwave often induces a weak frontal wave on the front.

As early a 1972, Smith and Younkin (1972) showed that relationship of heavy rain to the jet stream with a digging polar jet. Maddox (1979) states that the cases Smith and Younkin investigated were probably synoptic type heavy rain events. The 1972 study never mentioned the concept of jet streaks. However, the favored position they found for the heaviest rain however is in a favorable location that would allow strong vertical motion to develop. Their position would allow the exit region of the jet max swinging around the base of the trough and the entrance region of the jet streak developing on the eastern side of a trough to enhance the vertical motions. A similar process has been described for heavy snowstorms by Uccellini and Kocin (1987).

The threat area (red boxes) is often along the right entrance region of an upper level jet streak and is sometime in the region where the entrance region of the polar jet and exit region of the subtropical jet are juxtaposed (Beleville and Stewart, 1983).

Strong winds are typically present at 850 mb and the winds often are confluent. There is often tongue of 850 dewpoints of 15C or greater with the threat area being located just north of the dewpoint maximum. In essence the area of maximum rainfall potential is located just north of a Θe ridge. The figure above (right hand side) is a composite of the wind, temperature and dewpoints depressions at various levels. The number 2 plotted below the 15 at 850 mb indicates that the average 850 dewpoint temperature was 13oC. However, the temperature and dewpoint plotted at the surface on the figure are given in oF. Precipitable water values are typically around 1.46” or about 181% percent of normal. Remember that precipitable water is the integrated moisture through the column and is not the moisture at any one level. On the workstations, model forecasts of precipitable water is often plotted with 850-hPa winds.
There are several important points to make about synoptic type events. First, while the winds aloft are typically stronger than found for other type of event, the mean winds in the cloud bearing layer are almost always parallel to frontal boundary. Cells within the front track northward or northeastward with the mean flow only to be replaced by new cells that form in axis of low level convergence associated with the front. The movement of the band of convection is often linked to the movement of the forcing associated with the mid level trough.

The Maddox synoptic type flash flood is usually associated with a subset of the Parker and Johnson Parallel stratiform (PS) type of MCSs. The size of the circle with the “x” inside it on the figure above is the strength of the winds parallel to the axis of convection or line. The arrows at right are the line perpendicular winds.

The period Mar-May 1995 was a very active period of heavy to excessive rainfall from Texas eastward into western Alabama. During this period the track of 500-hPA lows was generally from the southeast across Colorado and then northeastward towards the Great Lakes region. Such a track usually means that the strongest height fall are north of the Gulf coast region. The track also implies that the mean trough remains well west of the high precipitation area.

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The synoptic archtype can occur during any time of year but is most common during the transition seasons from late winter to early summer and from fall into early winter. Synoptic type events are usually associated with strong 500-hPa troughs that are lifting to the northeast. There is a distinct maximum to 4 inch or greater rainfall events across the Gulf Coast region from late winter into the spring (figure below left).

A forecast exercise, using pattern recognition and ensemble guidance to predict a heavy rainfall event.

First we’ll use the forecast funnel approach espoused by Snellman (1977) and look at the synoptic scale pattern and compare it to the synoptic type flash flood pattern discussion on the previous page. Well then look at various ensemble forecasting system products to see how they can possibly be used during a potential rainfall event.

12 hr NAM surface pressure and boundary layer wind and temperature v.t. 1200 UTC Oct. 26, 2006. The thick red lines is a warm frontal position.

24 hr NAM surface pressure and boundary layer wind and temperature v.t. 0000 UTC Oct. 27, 2006. The thick red lines is a warm frontal position.

A fairly cold airmass was located over the east and the NAM and GFS predicted the development of a low to the lee of the Rocky Mountains in western Oklahoma. The remains of an old frontal boundary was located over Louisiana. This boundary was predicted to move very little during the 12 hours between the two forecast maps shown above. South of the warm front, surface dewpoints were forecast to be above 70oF to the south of the warm front across eastern Texas and over the western Gulf of Mexico. There is also a cold front moving across west Texas with a surface trough to its east. Having a prefrontal trough is a common variation of synoptic type event as it moves east of the Rocky Mountains. Compare the pattern to the surface pattern for a synoptic scale heavy rainfall event (below left).

Schematic diagram of surface pattern for Maddox et al. (1979) heavy rainfall events.

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12hr GFS 500-hPa heights and vorticity valid 1200 UTC 26 Oct 2006.

27hr GFS 500-hPa heights and vorticity valid 0300 UTC 27 Oct 2006.

27hr GFS 250 heights and isotachs valid 0300 UTC 27 Oct 2006.

The forecast upper level pattern is also a close match for a synoptic type heavy rainfall event. A strong, positive tilted trough is moving eastward towards the area. The 250-hPa isotach pattern shows a double jet structure with the right entrance region of the jet being near the threat area which extends across eastern Texas into Arkansas and Louisiana.

24 hr GFS 850-hPa wind and PW forecast valid at 0000 UCT 27 Oct. 2006.

The forecast 850-hPa winds and precipitable water also support a synoptic event. During the night, the 850-hPa winds are forecast to strengthen to 45 to 50 kts across Louisiana. The over 2.00” PW values predicted across Louisiana are significantly higher than the average 1.46” found by Maddox for such events.

Standardized anomalies are a useful but not a stand-alone forecast tool. They should be used as and adjunct to an ingredients based approach to assessing the pattern and its precipitation potential and that many heavy rainfall events are associated with convection and mesoscale and storm scale processes. However, they can give an idea of how unusual the pattern is. Below is the percent of the time that the distribution will fall within 2.00, 3.00 and 4.00 standard deviations (SD) based on the shape of the curved representing the distribution. While moisture flux and PW curves are skewed, they’re distribution within various standard deviations remains closer to the bell shaped than to the distribution based on Chebyshev’s inequality. Therefore, a 2 SD anomaly is a relatively uncommon event and a SD of greater than 4 is a rare event.

Normalized anomaly	any shape	Bell curve
2 S.D	75%	95.4%
3 S.D	89%	99.7%
4 S.D	94%	99.99%

Model forecast of various fields can be assessed to see how unusual such high values are for that field. Ensemble mean forecasts can be very useful at times. When an ensemble mean shows very high values of precipitable water, a strong low-level jet and moisture flux, it signals that the majority of members are predicting similar timing and evolution of those fields since ensemble means typically smooth features. This smoothing usually decreases the maximum value and often broadens the field. For example, verification of ensemble mean precipitation forecasts exhibit a high bias for low precipitation thresholds and a low bias for higher amounts (1.00” or greater amounts). For more information on using ensemble forecasts see
http://www.hpc.ncep.noaa.gov/ensembletraining/ and (http://meted.ucar.edu/nwp/pcu1/ensemble/print.htm).
The ensemble mean forecast of PW and its normalized anomaly for the Oct. 2007 events are shown below. The ensemble mean forecast showing PWs above 2.00 indicates that there is a lot of agreement among individual members that the PW will exceed 2.00 inches. The mean suggests that the majority of the members are predicting PW values significantly higher than the average value found by Maddox for synoptic type events.

15-hr SREF ensemble mean forecast of PW (brown, inches) and standardized anomaly of PW (dashed red) valid 1200 UTC 26 Oct. 2006.

27-hr SREF ensemble mean forecast of PW (brown line, in inches) and standardized anomaly of PW (dashed red) valid 0000 UTC 27 Oct. 2006.

27-hr SREF ensemble mean forecast of 850-hPa magnitude of moisture flux (scale at left, (g/kg)(ms-1), standardized anomaly of magnitude of moisture flux (dashed red) and 850-hPa wind (barbs).

The standardized anomalies, especially of the magnitude of the moisture flux suggested this could be a significant event. Based on the greater than 4 SD anomaly, over 99 percent of the time that the observed moisture flux values during the same time of year will be lower than the forecast values. They anomalies are suggesting this will not be a run-of-the-mill rainfall event.

12-36 hr GFS QPF valid 1200 UTC 27 Oct. 2006.

Also, the SREFS ensemble mean forecasts of normalized anomalies suggested that the heavier amounts predicted by the GFS were more likely than the NAM maximum amounts. However, the high anomalies do not really aid in pinpointing the precise region where the maximum rainfall. The location is usually determined by low-level boundaries that help focus the low level convergence.

That both the GFS and NAM forecast heavy rainfall across the region from eastern Texas across Louisiana into Texas is no surprise. In a sense, the two forecasts provide two members of an ensemble. Ensemble products provide an estimate of the uncertainty of the event and can provide a rough idea of the probability of an event occurring. In this case, the operational models indicated there is a high likelihood of a heavy rainfall event but differ on how heavy and where the maximum might be located. Ensemble forecasts from the SREF and GEFS are additional tools that can and should be used to assess the likelihood of this being a significant rainfall event.

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Schematic diagram of upper level pattern during Maddox et al. synoptic type events. Green lines are 500-hPa streamlines, winds are at 500-hPA, light blue is location of 250-hPa wind maxima.

15-hr SREF ensemble mean forecast of 850-hPa magnitude of moisture flux (scale at left, (g/kg)(ms-1), standardized anomaly of magnitude of moisture flux (dashed red) and 850-hPa wind (barbs).

12-36 hr NAM QPF valid 1200 UTC 27 Oct. 2006.

24 hr 3 inch spaghetti diagrams valid 1200 UTC 17 Oct. 2006

SREF members, eta members are red/yellow, rsm members are blue/purple and wrf members are green.

GEFS members, only initial conditions are perturbed.

Based on the two sets of ensemble products and your knowledge of the GEFS and SREF ensemble guidance which is the least true.

A)

The ensembles members suggest that the NAM maximum rainfall may be too light.

B)

The ensemble members suggest that the highest threat for over 3-inches of rain extends along an axis from central and northern Louisiana into central or northern Mississippi.

C)

The SREF ensemble members will typically show more spread than the GEFS members.

D)

The ensemble members suggest there is a high probability of over 3 inches during the event and that there is some potential for over 5 inches.

The ensemble members definitely suggest heavy rainfall is likely across the Gulf states. However, it’s important to remember that the GFS ensemble members only have their initial conditions perturbed so they may not have sufficient diversity. Also, the lower resolution of the GFS based ensembles members compared to the operational GFS may also make it more difficult for them to predict the higher thresholds of rainfall than the GFS. Also, they will likely suffer the same biases as the GFS in terms of the location of the maximum rainfall.

A number of researchers have suggested that when trying to forecast mesoscale events and precipitation that not only do you need to perturb the initial conditions but also need to perturb various parameterization schemes (Mullen et al.1999) in order for ensemble systems to have a sufficient range of solutions to properly capture the uncertainty in a forecast situation. The SREFS makes an attempt to capture at least a part of this uncertainty by also perturbing the physics. UCAR offers a matrix that provides information on how the GEFS and SREFS are perturbed, see see http://meted.ucar.edu/nwp/pcu2/ens_matrix/.

So far, we’ve mainly focused on the synoptic scale and to a certain extent have neglected an ingredients based approach. Let’s complete the picture and also discuss how any MCS that forms might align compared the model forecast of the precipitation axis.

12-hr NAM forecast of , boundary layer theta-e, winds and moisture convergence v.t. 1200 UTC 26 Oct. 2006. .

Observed LCH sounding for 1200 UTC 26 Oct. 2006.

One of the most worrisome aspects of the forecast is whether the surface boundary predicted to be across Louisiana at the beginning of the period will lift northward or not. The low predicted low level convergence over eastern Texas and southwestern Louisiana argue that convection may already be going on at 1200 UTC 26 Oct. While no forecast sounding is available for this case, the observed sounding for Lake Charles for the same time argues that convection probably had already started in southwestern Louisiana. The sounding also shows little shear between 850-hPa and 300-hPa suggesting any convection that organizes into a system may be slow moving. Note that the predicted CMF vectors (see below) suggest that any MCS will be slow moving and may move more eastward than northeastward.

The relatively low LCL suggested by the sounding also argues for a fairly low cloud base. The high relatively humidity and high PWs argue for heavy rainfall rates.

24 hr NAM mslp, boundary layer wind and temperature v.t. 00Z 27 Oct, 2006

The axis of the observed precipitation maximum was southwest of where the axis was predicted by any GEFS member (see below). Also, the axis was more east-west oriented than forecast. The maximum was south of all but two SREFS members, both of the members being ETA members with the BMJ convective scheme.

24 hr observed accumulated precipitation valid 1200 UTC 17 Oct 2006. Magenta is over 6 inches, white over 10.

24 hr SREF member 3-inch contours valid 1200 UTC 17 Oct. 2006.

24 hr GEFS member 3-inch contours valid 1200 UTC 17 Oct. 2006.

A 2006 in-house study by Barthold indicated that the location of the GFS rainfall maximum was much more likely to be predicted too far north versus too far south. This is consistent with one of the known biases of the Arakawa Schubert scheme (A/S) that is mentioned in the UCAR Model Matrix which notes that bulls-eyes from the A/S scheme have their convection form too late and too far north along the sloped frontal surface rather than near to the warm air where the actual convection usually develops. The SREF ensemble members offered a better idea of the threat area having it extend from southwestern Louisiana into southeastern Arkansas and northern Mississippi. However, its members will often cluster based on their convective schemes with the A/S and Kain-Fritsch members in one group and the Eta members that use the BMJ scheme in another so even the SREF members may not have enough diversity.

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For more on using standardized anomalies as a forecasting tool, click here.

Maddox Frontal Type heavy rainfall events

Comprised 25% of Maddox’s cases. The primary distinguishing characteristic for this type of event is an east-west frontal boundary with abundant moisture to its south. Precipitable water values in the Maddox et all study averaged around 1.60 inches or about 160% of normal on the warm side of the frontal boundary. In the frontal events studied, the warm moist air was forced to overrun the front so that heaviest rain typically fell on the cool side of the front. This type of event primarily occurs at night, possibly because of the nocturnal maximum in the LLJ documented by Bonner (1966). Wind speeds generally increased little from 850 mb through 300 mb although significant veering was present. The upper-level winds were usually found to parallel the front allowing convective cells to train over an area.

850 mb

The greatest threat of excessive rain was skewed slightly to the right of a warm tongue of moist unstable air. 850 mb dew points were usually in the mid to upper teens within this “tongue” and K-indices were around 38. Since the best moisture inflow is often found along the western or southwestern portion of a Meso-scale Convective System (MCS), cells often redevelop along the western flank. This redevelopment can allow the most active convection to remain stationary or even propagate backwards. Schofield and Robinson (1990) have also presented a conceptual model for backwards propagating MCS.

Note that that the moist unstable air with higher low level Θe is located upstream from the initial convection. Junker et at. noted that the low level moisture convergence is also located upstream (usually to the west) during frontal type of events during the 1993 Midwest floods. The Maddox frontal type system is essentially of Parker and Johnson leading stratiform type event. New cell development causes the MCS to move slower than the mean flow.

Adapted from Maddox et al. 1980

Maddox and his colleagues noted that frontal heavy rainfall events often occurred near the 500 mb ridge where the mid level flow is fairly weak. The weak mean flow associated with many frontal type events makes it easier for propagation effects to offset cell motion. Also, the lack of shear in the cloud bearing layer, raises the potential for training of cells.
Doswell and his colleagues (1996) have suggested that the reason that MCSs and many flash floods develop near the 500h ridge axis may be be related to its synoptic scale subsidence suppressing convection which allows diabatic heating processes to steepen lapse rates and also allows for the accumulation of moisture in the area that later is used to fuel the MCS. They speculate that if convection developed earlier in the day it would process the moisture making less moisture and instability available to fuel the sustained convection needed for MCS development.

Blanchard et al. (1998) have suggested that a position near the axis of an upper level ridge is also an area of weak inertial stability. Therefore, when convection does develop near the ridge that the weak inertial stability associated with the anticyclonic curvature might make it easier to develop upper level divergence. Therefore, convection might be more likely to organize into an MCS where the upper level divergence is enhanced by the weak inertial stability. The region to the east or south of an upper level jet where the horizontal shear is anti-cyclonic is another area of weak inertial stability.

Adapted from Maddox et al.

Adapted from Maddox et al. 1980

Frontal type events often develop into training line-adjoining stratiform convective type systems with cells aligned parallel to the surface front but to its north. The line of thunderstorms are oriented parallel to the mean flow so cells train across an area.

From Shumacher and Johnson 2006

Frontal type events are most common during the warm season when dynamics are weak. However, hybrid synoptic/frontal systems can occur across the south during the winter.

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Mesohigh Heavy Rain Events

MESOHIGH 850 mb

Approximately half the time, these mesohigh events occur east of a slow moving frontal band where it is intersected by the outflow boundary. Surface dewpoint are usually above 70 deg F south of both the front and outflow boundary. The typical surface, 850 and 500 mb patterns for mesohigh events is shown below

850 mb dew points south of the boundary are generally in the mid teens leading to K-indices in the upper 30s. The greatest threat area of excessive rainfall is usually found in the southern or western portions of the mesohigh. Essentially, mesohigh events and frontal events are the same. The scale of development is the main difference. An old heavy precipitation forecasting admonition is bewaring of east-west frontal or thermal boundaries, they are dangerous.

Mesohigh events have a propensity to develop near the axis of the upper level ridge where the winds at mid and upper levels are fairly weak and the shear is relatively weak.

What makes the Frontal and Meso-high type patterns identified by Maddox et al. particularly favorable for heavy rains. Let’s look at them from an ingredients based methodology. The article notes that the airmass is unstable with anomalously high precipitable water (PW) for both types. The combination should produce adequate vertical moisture flux into the cloud to produce heavy rains. Both composite shows an axis of higher 850 winds that are also coincident with high 850 dewpoints suggesting strong there is maximum of horizontal moisture flux feeding into the thermal boundary. This is a plus for trying to sustain the vertical moisture flux into the system.
The low level jet is often located to the southwest of the boundary which helps maintain low level convergence on the southwest side of any convective system than develops. The mean flow is westerly so the MCS often develops into a Parker and Johnson leading stratiform type system. Rain cooled air reinforces the boundary as the initial cell moves away from it only to be replaced by a new cell developing along it. The rain cooled air on the east side or northeast side of the boundary stabilized the air mass east of the boundary but moisture unstable air remains along its southwestern flank. The new cells forming along the southwest flank means that the propagation due to new cell formation is opposing the movement of the system due to advection. Therefore, resultant MCS moves slower than the mean flow which helps prolong the period that high rainfall rates will be affecting the area.
The likelihood of multiple cell training across an area increases as new cells generate which also tends to increase the rainfall efficiency of the system by increasing the relative humidity and decreasing the negative effects of entrainment. It’s important when you think that the pattern favors a Maddox type heavy rainfall event to still evaluate the pattern using an ingredients based methodology. Try to anticipate which is the most likely Parker and Johnson arch-type is the most likely to develop. For example, is new convection likely to develop upstream or downstream from the initial convection that develops.

An example that fits the Maddox et al. archtype and using climatological thickness to forecast where an MCS might form

500-hPA height and vorticity forecast

MSL and 1000-500 thickness forecast

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00 UTC 17 July

00 UTC 17 July

Now let’s try to apply what you’ve learned from studying about the Maddox arch-types. However, We’ll also try to apply climatological thickness to helping solve where the axis of heavy rain will fall. During the 1960’s one of the forecasters conducted a study to try to find out the average value of 1000-500-hPa thickness values that 1.00 inch or more rainfall was observed in a 12 hour period during the various months of the year (Bohl and Junker, 1987). HPC have used the climatological thickness to try to help forecast where convection might develop. Climatological thickness works best during the summer when dynamics are relatively weak. Let’s try to use the climatological thickness to estimate the most likely region where heavy precipitation might develop using the fields shown above. Note that a shortwave is approaching the 500 hPa ridge axis. Such a 500 hPa pattern is a common one for Maddox frontal and meso-high events. The graph below shows the climatological thickness for the northern plains and portions of the upper Midwest. The mean thickness for region 1 using the chart below is 575. However, the median is 576 and the mode 579 for the region suggesting that the axis of heavy rain will occur in northern Nebraska extreme northwestern Iowa.

A)

Click A to go to a with the climatolgical thickness for various regions of the country.

NAM forecast of precipitable water and 850 hPa wind

1.25

1.50

1.25

1.50

Valid 00 UTC 17 July

Valid 03 UTC 17 July

Forecasters at HPC have found that the amount of available moisture in the atmosphere plays a strong role in determining the whether heaviest convection will occur along or north of the front to whether outflow boundaries from the initial convection will help focus convection south of the front. The forecasters also have noted that when the mean relative humidity reaches 70% or more the lifting is present, convection usually develops. This relationship is the basis for using the saturation thickness table. If for example the average PW values south of the front are about 1.15 to 1.25 inches, then saturation would most likely occur right along the front as represented on the bottom figure on the right hand side. If on the other hand, the PW values were lower, let’s say 0.80”, then for hypothetical case below, then north of the boundary.

From Funk 1999

If you were using the saturated thickness concept to predict the axis of heaviest rainfall for the 17 July case, the new forecast axis would be:

A).

north or northeast of the location found using the climatological thickness.

B).

Almost identical to the one found using climatological thickness

C).

South or southeast of the location suggested by climatological thickness.

Please click on the correct answer, A, B, or C.

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Junker et al. (1999) have noted that in reality that convective systems do form on the left side of the line where the mean RH is below 70% but that their scale tends to be smaller. In that 1999 study, MCSs were classified by the scale of their 3 inch or greater coverage. The figure above plots thickness versus PW with the diagonal black line given the saturated thickness values for PWs taken from the Funk table. All of the two largest scale category rainfall events (the most extreme ones) occurred along or to the right of the saturated thickness line. In other words, all occurred with a mean relative humidity that was equal to or greater than 70%. More recently, Kelsch noted that a number of severe flash flood cases occurred well to the right of the saturated thickness line. The small arrows on the figure above represent the thickness and PW that was present for various flash floods. The plot of the flash flood cases suggests that when a convective event occurs with at high precipitable water values relative to its thickness that the event should be monitored closely for its potential to produce heavy to extreme rainfall.

A few ideas on forecasting elevated convection

MSL and 1000-500 thickness forecast

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The figure at left shows the location of the axis of heavy rain according to the saturated and climatological thickness methods (transparent white area). The area suggested by either thickness method suggested that the rainfall event would be an elevated one to the north of the low level boundary and the boundary layer moisture convergence. The rainfall during the 24 hours was concentrated in this same area.

00 UTC 17 July

03 UTC 17 July

The saturated and climatological thickness suggested the precipitation would be well north of where the NAM was forecasting a weak surface boundary of some sort (see forecasts below)

NAM forecast of boundary layer wind, moisture convergence and dewpoint

The climatogical and saturated thickness methods provide one way of determining how close to the front to forecast the axis of convection in the plains. Another aid in determining whether convection will be capped or not near the front is to use the 700 hPa temperatures. The figure on the left shows the 700-hPa temperature that were present at the onset of the development of convective rainfall events that produced at least 2 inches or rainfall. The majority occurred at temperatures below 12oC though a few occurred with temperatures as warm as 14oC.

A air mass will usually be capped at temperatures above 12oC except in the high plains. An even better indicator of the cap is CIN (convective inhibition). Essentially, CIN is a measure of the energy needed to overcome the cap and initiate convection. When strong CIN is present along the surface the boundary and the CIN is forecast to remain strong, elevated convection may develop to the north of the surface front in the area of strong warm advection providing there is a sufficiently deep elevated unstable layer.

Moore et al. (2003) studied twenty-one warm season heavy rainfall events that developed north of a surface boundary. They found that the heavy-rain region of an elevated MCS is location on average about 160-km north of a quasi-stationary frontal band along a region of low level convergence that is elongated to the west (see the left hand figure below). In his composites the low-level front is located across Oklahoma and straddles the Arkansas and Missouri border. They note that large values of CIN (greater than 110 J kg-1 inhibit convection near the front. Therefore, expect elevated convection in the Plains when strong CIN is present. The exact values found by Moore are an average value for the CIN calculated in one manner and are observed CIN values. The CIN on the workstation may not be calculated in exactly the same manner and model forecast CIN may be biased. Therefore, it is important to learn any biases that may be inherent in the model forecast CIN that can be displayed on the workstation. However in general, when there is high values of CIN present along a front, convection will be capped or will develop as elevated convection on the cold side of the front.

Adopted from Moore 2003, Schematic diagrams that summarize the typical conditions associated with warm season elevated thunderstorms attended by heavy rainfall:

a) low-level plan view and In (a), dashed lines are representative Θe values decreasing to the north, dashed-cross lines represent 925-850-hPa moisture convergence maxima, the shaded area is a region of maximum Θe advection, the broad stippled arrow denotes the LLJ, the encircled X represents the MCS centroid location, and the front is indicated using standard notation.

Convective inhibition (CIN) computed using the mean parcel from the lowest 100 hPa. The dot represents the MCS centroid location.

(b) middle-upper-level plan view. In (b), dashed lines are isotachs associated with the upper-level jet, solid lines are representative height lines at hPa, the stippled arrow denotes the 700-hPa jet, and the shaded area indicates where the mean surface-to-500-hPa relative humidity exceeds 70%.

Maximum-Θe CAPE (J kg-1) computed using the parcel with the maximum Θe in the first 300-hPa of the sounding.

Moore and his colleagues also noted that elevated convection was more common along the right entrance region of an upper-level jet in a region of 250-hPa divergence and 850-hPa convergence. The elevated convection is also located to the north of the highest CAPE in a region where there is a strong gradient of CAPE. They presented a cross-sectional view taken parallel to the low-level jet to illustrate the typical 3-dimensional circulation pattern associated with most elevated MCSs. Note that the parcel is being lifted along an isentropic surface that is sloped with the slope largely determined by the strength of the frontal boundary and if looking at a plane view, its thermal gradient. They note that the distance north of the front is partly governed by the CAPE, CIN and slope of the isentropic surface.

Dashed lines represent typical Θe values, the large stippled arrow represents the ascending LLJ, the ageostrophic direct thermal circulation associated with the upper-level jet is represented by the closed solid line with the large arrows on it, the dash-dotted oval represents the direct thermal circulation associated with the low-level frontogenetical forcing. The area of aloft enclosed by dotted lines indicates the location of upper-level divergence. The enclosed solid lined area aloft is the location of the upper-level jet streak. Note the horizontal distance between the MCS and upper –level jet maximum is not to scale.

Schumacher and Johnson identified that the training line-adjoining stratiform MCS that was discussed in an earlier section as the most frequent type of MCS associated with extreme rainfall. They noted that such an MCS typically was typically elevated. They note that the key to such events causing extreme rainfall is the orientation of the convective line with respect to the flow of the cloud bearing layer. When they are parallel to each other, extreme rainfall is much more likely.

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SYNOPTIC SCALE CYCLONIC CIRCULATIONS
As mentioned by Spayd and Schofield (1983), and Elsner et al, this type of excessive rainfall producer is usually associated with slow moving warm top convection. Generally, the models show a fairly strong 500 mb vorticity maximum with a very round look to it. Usually a weak surface wave develops in response to this feature. Often the convection takes on a comma shape with the heaviest rain usually occurring at the southern edge of the comma head, north of northeast of the surface low and about 2-3 degrees of latitude to the north of northeast of the 850 mb low. The squares shown on the figures below represent the typical threat area for extreme rainfall.

Composite of the 500-hPa vorticity for warm-top convection embedded in a synoptic-scale cyclonic circulation system. Vorticity isopleths are dashed lines, streamlines are solid lines, and low dewpoint depression areas are scalloped.

Composite of the 850-hPa chart for warm-top convection embedded in a synoptic scale circulation system. Wind barbs and dewpoint contours are annotated.

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65oF

70oF

A composite of the surface chart for warm-top convection embedded in a synoptic scale circulation. Dewpoints are dashed every 5oF. Pressure is depicted by solid lines every 2 hPa.

Composite of soundings for the warm-top convection embedded in a synoptic scale cyclonic circulation. The surface temperature and dewpoint is in o F, At upper levels temperatures (o C) and dewpoint depressions are given. The average precipitable water and surface to 500-hPa mean relative humidity are shown.

The average precipitable water is lower for this type of heavy rainfall producing system than is typical for the various arch-types discussed by Maddox. The convection often has a much lower cloud top temperature than usually observed for most summer-season convective events. The most important characteristic of satellite cloud pattern is that it remains stationary for an extended period though it also often takes on a comma type shape. The circulation associated with these events usually extends from the surface to at least 850-hPa and sometimes to 500-hPa.
The high surface to 500-hPa relative humidity implies that the convection associated with these type systems will have low cloud bases implying the precipitation efficiency during these events is high. The axis of heavy rainfall usually occurs along a deformation zone to the north of a surface low.

The case below is a good example of the rainfall associated with a cyclonic circulation system.

12 UTC 12 AUG 2000

00 UTC 13 AUG 2000

Surface maps

12 UTC 12 AUG 2000

00 UTC 13 AUG 2000

925-hPa maps

The sounding at OKX (Upton) New York exhibited a distinct easterly low-level jet of 25 knots at around 925 mb at 12 UCT 12 August and 00 UTC 13 August. Most cyclonic circulation systems that produce heavy rainfall probably are similar to winter storms that produce heavy mesoscale bands of precipitation often in the form of snow except that the area of heavy precipitation is usually smaller in scale than the winter events. Nicosia and Grumm (1999) have noted that there is a zone of differential moisture advection where mid-level drying occurred above the region dominated by the moisture laden easterly low-level jet associated with the cold conveyor belt.
where lower branch of the cold conveyor belt (in this case the easterly low level jet)

Mid-level frontogenetic region

Cloud shield

Cold Conveyor belt

Dry Tongue Jet

Zone of stability reduction reduction

Surface low

Warm conveyor belt

Figure adapted from Nicosia and Grumm (1999)

The figure above is adapted from their model. They note that there is a synergistic relationship between frontogenesis and EPV reduction and differential moisture advection (Θe advection). They point out that for the two snow cases that they studied, the mesoscale bands of heavy snow occured in the zone of where the stability and EPV were being reduced and the southern edge of region of mid-level frontogenesis. The same concept holds for heavy synoptic type rainfall events except they typically occur during the warm season when airmasses are often already conditionally unstable. Therefore, the precipitation is usually a result of upright convection.

12 UTC 12 AUG 2000

00 UTC 13 AUG 2000

850-hPa maps

Note that during this event that the 850-hPa circulation was strengthening. Such deepening of the low or mid-level circulation is typical of a cyclonic circulation system that produces heavy rainfall from thunderstorms having relatively warm cloud top temperatures.

Infrared image valid 1315 UTC 12 August 2000UTC

Corresponding 0.5 degree radar image

Sounding valid 1200 UTC

Sparta, New Jersey received 12 inches of rain during this case and several other locations in northwest New Jersey reported as much 10 inches. However, the convection did not look that impressive on the infrared imagery. Cloud top temperatures were only around –30oC fairly warm for a summer season heavy rainfall event. The equilibrium level is typically low during these type systems and for this case was at 400-hPA.

12 UTC 12 AUG 2000

00 UTC 13 AUG 2000

500-hPa maps

Terrain map

Radar rainfall estimates

Slow movement of the 500-hPa circulation is another common aspect of these cyclonic circulation associated heavy rainfall events and certainly was evident for the Sparta New Jersey case (see 500-hPa maps above). Slow system movement allows the mesoscale forcing to remain relatively stationary. In the case, the axis of low level jet, frontogenesis remained relatively stationary. In this case, another big plus was lift provided by the terrain. The easterly low-level jet impinging upon the hilly terrain of northwestern New Jersey provided a stationary source of lifting. The rainfall maximum (see above) was located in northwestern NJ in proximity to the mountains.

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Latent heat and the generation of mesoscale convective vortices

The effect of latent heat on the general circulation pattern is to superimpose an anticyclonic perturbation at the upper-levels and a cyclone at the low and mid-levels just above a shallow layer of cool air near the surface (Kreitzberg, C.W and D.J. Perkey, 1977, Maddox and Fritsch, 1980, Fritsch and Maddox 1981 Perkey and Fritch 1981). Maddox (1980) found that the convection associated with quasi-circular systems usually started to form during the afternoon but that the convection usually did not organize into a system until early evening. He noted that most of these systems grew to maximum size on the satellite imagery after midnight and persistent into the morning hours.
The effects of latent heat can play an important role in the development and enhancement of mesoscale circulations (Keyser and Johnson, 1984). The development of a cyclonic low-level circulation acts to enhance the low-level southerly jet on the east side of the convective system. This leads to an increase in the amount of moisture feeding into the complex. The release of latent heat also helps to superimpose an upper level anticyclone over the convection. The following two figures give a general idea of the affects of latent heat and the typical mesoscale circulation that often develops. The impact of convection it to develop and anticyclone at upper-levels and generate cyclonic vorticity in the mid-levels.

From Frisch et al, 1994

The cyclonic low and mid-level circulation that develops with a strong convective complex sometimes shows up on the satellite imagery once the parent convective complex dissipates. Typically, these convectively induced mesoscale cyclonic vortices (MCV) show up first on the visible imagery. Cumulus and towering clouds usually define the circulation center once diurnal heating gets underway. A banded structure usually becomes defined between 1800 and 2100 UTC, during the time of maximum heating (Bartels and Maddox, 1989). The figure below left shows a typical looking MCV on the visible imagery and the figure on the left shows the system relative winds at 600-hPa.

From Davis et al. 2004

MCVs tend to from in regions of weak shear and light mid and upper level winds. The figure below shows the location of MCVs with respect to the synoptic scale mid-level ridge. Note that number of the cases occur either downstream from the ridge or to its south. Bartels and Maddox found that 10 of the 24 meso-vortices that they studies were associated with subsequent MSCs or a series of MCSs. Therefore, such vortices need to be monitored closely.

From Bartels and Maddox 1991

The maximum vertical motion and precipitation associated with the MCS typically occurs near the center or slightly downshear (Davis and Weisman 1994). Often, during the day, the convection is removed from the center and occurs along outflow boundaries. However, at night it tends to shift towards the circulation center and typically occur near or just to the northeast of it. The ability of an MCV to initiate convection appears to be related at least in part to shear. Weak shear confined to the lowest layers of the atmosphere seem to the be the optimal conditions for subsequent convective development (Davis and Weisman 1994)

A mesoscale convective vortex exercise

The circle on the 500-hPA analysis below is the location where a mesoscale convectively induced vorticity center was located at 12-UTC.

250-hPa analysis valid 12 UTC 6 May

850-hPa analysis valid 12 UTC 6 May 2000

700-hPa analysis valid 12 UTC 6 May 2000

From the four maps above, which is the truest statement about the MCV.

A)

The 850 wind field across eastern OK and MO is strong enough to suggest there will be too much shear for the system to maintain itself into the evening.

B)

The lack of shear in the mid and upper levels suggest that the MCV may be able to maintain itself through the day so it should be watched closely during the afternoon and evening for possible convective development.

0000 UTC visible satellite imagery

0000Z UTC infrared satellite imagery

Obviously from the pictures above, the MCV did maintain its identity. The lack of shear at mid and upper levels allow it to maintain the circulation. Having relatively strong low level winds is not that detrimental to MCVs.

500-hPa analysis valid 00 UTC 7 May 2000

250-hPa analysis valid 00 UTC 7 May 2000

700-hPa analysis valid 00 UTC 7 May 2000
May 2000

Analyzing the 579 decameter height on the 500-hPa plot (see above left) suggests that the vorticity center still is well defined. The shear between 700 hPa and 250-hPa remains fairly weak at mid and upper level data. Note that the winds that are implied across eastern Missouri are all the 20 to 25 knot range from 700-hPa to 250 hPa. The big question becomes whether new convection will form during the evening.

Look at the 00 UTC sounding for SGF and decide which are the most important and representative fields on the sounding of the surrounding environment.

A)

The saturated and unstable layer between 800 and 650 hPA?

B)

The 123.3 J/kg CAPE?

C)

The 40 knot southwesterly winds at around 850 hPa?

D)

The vertical wind structure?

Click here for answer

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500-hPa analysis valid 12 UTC 6 May 2000

Mesoscale Convective Vortex exercise (continued)

850 hPa analysis valid at 00 UTC 7 May 2000

Surface analysis valid at 00 UTC 7 May 2000

The surface plot is fairly benign looking. The surface dewpoints are in the low 60s with very little in the way of a boundary to focus lifting. The NAM and GFS had very innocuous precipitation forecasts. The NAM predicted a little more than .50 inch just to the southwest of STL. The GFS had .50 inch or more, mostly in Illinois. Most forecasters would probably forecast some precipitation with the MCV and might beef up the models some.

0215 UTC 7 May 2000 infrared imagery

Convection started to develop at around 02 UTC near the center of the circulation. Remember the NAM and GFS had some convective precipitation forecast in this general area but did not predict any amounts greater than 1.00 inch. Often the convection associated with MCVs produce their rainfall maximum at night.

The 06 hr forecast sounding from the 00 UTC RUC indicated that the CAPE would be around 1000 J/kg at Union MO and that there would be a potent low level jet of 45 kts (see below).

The RUC sounding above shows a number of factors that suggest that any MCS that forms and persists from 0215 UTC to 0600 UTC will probably produce high rainfall rates. Which of the following would support such a forecast?

A)

The low LCL that suggests a low cloud base

B)

The low values of CIN

C)

The 12795 ft or 3.88 km freezing level that suggests a deep layer where warm rain collision coalescence processes can take place.

D)

The high mean RH and deep layer on the sounding that is near saturation.

E)

The 1.78 inch PW which is higher than the average value Maddox found associated with frontal or mesohigh type flash floods.

F)

The 45 kt low-level jet which along with the high PW which suggests the MCS will have no trouble replenishing the moisture that fall as rain.

G

All of the above combined together suggest high rainfall rates are a good bet.

The 6-h RUC sounding forecast did a good job of indicating there was potential for convection that could produce intense rain rates.

0545 UTC 7 May 2000 infrared imagery

Radar reflectivity for 0600 UTC 7 May 2000 (from Schumacher and Johnson 2005)

Once convection developed, a low-level boundary became apparent along with a persistent area of moisture convergence.

0600 UTC meso-surface analysis (from Glass et al. 2001)

0600 UTC RUC analysis of 950-hPa moisture convergence and wind (from Glass et al. 2001)

1000 UTC meso-surface analysis (from Glass et al. 2001

1000 UTC RUC analysis of 950-hPa moisture convergence and wind (from Glass et al. 2001

Radar reflectivity from 0900 UTC 07 May 2000 (from Schumacher and Johnson 2005)

The observed rainfall rates at Union Missouri were in the 2-3 inch per hour range from 0630 UTC to 1030 UTC. If you were a forecasting in St Louis looking at a time cross section showing reflectivity, there was a big clue that heavy rainfall might be occurring, the high reflectivity area on the cross section was located below the freezing level. Therefore, it was unlikely that hail was causing the high reflectivity or that the NEXRAD algorithms would be overestimated rainfall rates.

Time series of accumulated precipitation gray bars and 30 minute rainfall observations (black bars) for Union Missouri. (From Glass et al. 2001)

Precipitation analysis inched for the 24 hr period ending 1200 UTC 7 May 2000 (from Glass 2001).

16 inched of rain fell just west of union which received more than 12. Not all MCVs produce such copious rainfall, in fact many dissipate without producing a second round of convection.

In Summary

MCVs are more likely to form and persist in regions of weak mid to upper level shear. A highly sheared environment is not favorable for their development or maintenance. If the surrounding airmass is moist and unstable there is a good possibility that a new round of convection will develop, such systems are especially dangerous if a low level jet develops and feeds moisture into upstream flank (usually southwestern flank) of the developing MCS (Trier and Davis 2002).

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Convection continued to develop along the western edge of the boundary through around 1200 UTC. The southwest winds kept the moisture feeding into the boundary and new cells kept forming even as the old cells move east and weakened. This is a case where the MCV helped generate convection but then the MCV kept moving eastward.

Western U.S. convective heavy rain events

Predicting the rainfall needed for flash flooding in the western U.S. during the warm season is especially difficult since rains of only 2 to 3 cm (around and inch) can produce flash flooding over a small, steeply sloped drainage basin (Maddox et al. 1978). Data coverage is limited due to the how sparse population over much of the west, especially the intermountain region. Also, radar coverage in the west is not as dense as over the east because of beam blockage problems related to the complex terrain. Even where there is radar coverage, the beam may end up overshooting the strongest echoes because the radar is often located on elevated terrain. Therefore, a combination of sensors need to be used to monitor moisture surges and the development of convection. GOES sounder data is available at (http://cimss.ssec.wisc.edu/goes/realtime/grtmain.html)
to monitor moisture surges. Note that if you compare the sounder product (below right) to the NAM initial analysis of PW over the Gulf of Mexico at the same time (below left), there appears to be a difference just off the coast of Louisiana. A similar comparison can be made for the gulf of California region during the Monsoon season. Another useful link

NESDIS satellite rainfall estimates also can provide valuable data to supplement radar imagery to monitor and track convective cells and systems in the west during the summer.

The two figures below are from Chappell’s COMET notes. They show the locations of flash floods over the western U.S and the type of convective system that produced the event. The number of storms near probably skewed by population since if is unlikely that a flash flood would be recorded if it did not impact someone.

Time of onset of heavy precipitation during western flash flood events (from Maddox et al 1978)

Western events have a strong diurnal maximum during the 6 hour period (2:00PM to 8:00PM LST) and generally occur during the monsoon season of late July through September.

In his COMET notes Chappell identified various types of western flash flood events They differed somewhat from the ones identified by Maddox et al. (1980). Where the composites differ, both sets will be presented.

Maddox type I (western)

This type event is associated with a shortwave lifting into a mean trough. The shortwave often takes on a negative tilt that acts to increase the upper level divergence over the area where the convection and heavy rain develops. The heaviest rain also usually occurs along or north of a weak surface boundary just west of the 500 ridge axis. The approach of the upper level trough to the ridge axis and associated shortening of wavelength also probably serves to increase the upper level divergence. Maddox and his colleagues caution that there are a number of variations of the pattern and that the amplitude of the ridge and trough do not have to be as strong as in the depiction below. This type system was the most common they found during their study.
The generalized 500-hPa and surface patterns for the type I event is offered below. The figure (at left) includes 500-hPa streamlines, area with the 500-hPa dewpoint depression that is less than or equal to 6oC (scalloped area).

Maddox and his colleagues noted that the interaction between weak shortwaves and surface fronts or troughs was common during these events but that the fronts often did not show up on the NCEP surface analysis. They suggest forecasters do their own analyses.
The table above presents the mean conditions associated with this type of event. Deep moisture is present from the surface to 300-hPa. Winds at all levels are relatively light (below 40 kts). Precipitable water averaged .94 inches (184% of the average monthly value) for these events and k indices were 40 attesting to the high relative humidity at 700-hPa.
The elevated terrain over the Rocky Mountains and west helps build a mean ridge over the west during summer. Whenever a trough starts impinging upon this ridge during the monsoon season is a prime period favoring isolated flash flooding in the intermountain region. The low mean precipitable water across this region usually means that heavy rainfall events in the intermountain region have a small scale and sometimes occur from rather scattered convection with most of the convection producing no significant flash flooding while a few isolated cells or clusters of cells produce the flood problems.
A 500-hPa pattern very similar to the one typical of a Maddox type I event often causes heavy to excessive rainfall to the east of north-south oriented mountain ranges. This type of event will be discussed next.

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East slopes of mountain range type events

Orographic lifting may trigger convective storms along the slopes of mountains given the correct flow regime (Maddox et al. 1998). Four such examples have been studied extensively, the Big Thompson (Colorado, the Rapid City (South Dakota), the Cheyenne (Wyoming) and Madison County (Virginia) flash floods (Maddox et al. 1978, Chappell and Rodgers 1988). The upper-level pattern was very similar for all four events. A full latitude trough was located along or just off the west coast of the U.S. A much smaller scale shortwave was noted to be lifting northward on the east side of trough during each case.

From Pontrelli et al. 1999

The figures from an article by Pontrelli et al (1999) below show the 500-hPA patterns associated with the four flash floods. The dashed line is the weak shortwave and the red line is the axis of the upper level ridge. Note that latter has a negative tilt in each case. The negative tilting 500h ridge was associated helped induce the development of a surface high that supplied low-level easterly flow into the mountains during each case. North of the low moving or stationary front, each also had a ribbon of moist air where there also was an easterly low level jet. The shaded area on the figures below show the ribbons of high dewpoint air at the onset of each storm.

From Pontrelli et al. 1999

All four events were characterized by having low cloud bases relative to the terrain and low centroid tops to the convection with most of the condensate forming and falling below the freezing level. The mid– and upper-level flow during each of the events was relatively light (Pontrelli et al 1999).

From Pontrelli et al. 1999

Pontrelli and his colleagues offered an 1800 UTC sounding for IAD during the Madison county event. The sounding illustrates how deep the moisture was across the region. In all likelihood, the relative humidity in the layer below 800-hPa on the IAD sounding would have higher relative humidity with less of a dewpoint depression than farther to the northwest where the cooler air would have been deeper. Such a modification to the IAD sounding would suggest a very low LCL and cloud base. The 1745 UTC reflectivity cross section of the Madison county storm clearly shows a low centroid of high reflectivity.

From Pontrelli et al. 1999

More that 600 mm of rainfall were estimated during the Madison County rainfall event. 250 mm were reported by several observers in a 2 hour period attesting to the intense rainfall rates that occurred.

Top number is temperature in oC except at the surface and the bottom number is the dewpoint depression

There are a few differences differences between the events along the front range of the Rocky mountain and similar events along the east slopes of the Appalachians is average precipitable water associated with the events. The figure at the left provides the average profile of the moisture and instability associated with the Rocky mountain events. The precipitable water for the western events is usually significantly lower than during Appalachian events. The scale of the events is often smaller. However, in precipitable water values are typically much higher than normal during these west coast events.

Figures adapted from Chappell COMET notes

The figures above show the seasonal frequency of east slope events along the front range of the Rocky mountains and the local time of occurrence. In the west these events are strongly tied to the monsoon and typically occur during the afternoon and evening. Along the Appalachians such events have been observed from June-early November

The idealized 500-hPA and surface patterns below show the typical threat area relative to the mountains and the typical look of the synoptic pattern. Note the similarity with Maddox type I western events.
In Summary, the following synoptic and mesoscale characteristic are common for warm season extreme rainfall events along the front range of north-south mountain ranges.
A weak 500-hPa shortwave swings northward or northeastward towards the threat area.
Light southeast to south-southeast winds are present in the mid-and upper levels.
A slow moving or stationary front lying just to the south of the threat area.
High moisture content though a deep layer
West and south of the threat area, afternoon heating in combination with cold advection to the east of the threat area helps intensify thickness and pressure gradients
A narrow band of moist and unstable air located just north of the front. A mechanism to provide a cap that limits development until the air is lifted by terrain.
The convection and heavy rains usually forms to the east of the highest terrain.
Convection may slowly move but keeps reforming on the terrain along the western flank of the convection.

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The deep moisture and warm layer on the sounding also have another important implication. The most efficient was to product downdrafts is through evaporation and melting. When most of the condensate is occurring in a very moisture atmosphere with a deep warm layer that is above freezing, the potential of cooling trough evaporation and melting are minimized which limits downdrafts. A weaker cold pool leads to less movement to the boundary giving a system more potential for new cells to form in place.

Maddox et al. type II events

This type of event accounted for about 20% of the western U.S. flash flood cases in the Maddox et al. study. These events are typically associated with a 500-hPa short-wave trough that is moving southward down the eastern side of the long-wave trough. The typical 500-hPa and surface pattern are condensed in the figure below. Note that the threat area (the shaded area) is usually found in advance of the 500-hPa short-wave. During these cases there is almost always a 500-hPa trough of closed low just off the coast of Baja. There is also usually a weak surface low over southern California. These two features help pull monsoonal moisture northward in advance of the southward digging short-wave. The remains of an old frontal boundary is sometimes, but not always, located in advance of the shortwave and is available to help provide low level convergence.

The table (above right) presents the mean vertical profile for some fo the significant levels. Wind aloft are more westerly than for either Maddox type I or IV events. However, speeds are generally light and are often below 40 knots throughout the vertical profile. Surface dewpoints are often in the upper 50s but can be even higher. The dewpoint depressions are low throughout the profile.
The airmass is typically quite moist and unstable with PWs, K indices and LIs averaging an inch (145% percent or normal), 39, -5, respectively.
Maddox and his colleagues note that the Sept. 14, 1974 El Dorado, Nevada flash flood could be classified as a type II event. However, though it shows a digging 500-hPa shortwave, the shortwave transitions into a weak 500-hPa low. Therefore, the pattern looks quite a bit different than the composite.

Chappell deformation type heavy rainfall event

This type of event typically occurs when monsoonal moisture has been pulled northward into the southwest, often but not always in association with an weak upper low to the west or southwest. In that respect this type of event is similar to the Maddox et al. type II event. Like the former, the forcing associated with the southern end of the trough then interacts with the monsoonal moisture. The idealized 500-h-Pa pattern associated with this type of event is provided below (left). Another similarity shared with type II events is a 500-hPA low or trough located near of just off the coast of Baja. The threat area below for flash flooding is near where the entrance region of the upper level jet would be located. The idealized 700-hPa pattern suggests that there is convergence at 700-hPa and that frontogenesis is taking place or that a weak frontal boundary is associated with the trough.

Idealized 500-hPa height , wind and vorticity pattern associated with deformation type flash flood events.

Idealized 700-hPa height , wind and dewpoint pattern associated with deformation type flash flood events.

A weak surface low is often located over Baja or southern California. A weak frontal boundary is also often evident. Precipitable water over the threat is around 1.15 inches (160%) of normal. However, even higher precipitable water is often found across southern Arizona and New Mexico where surface dewpoints are often in the 60s.

Idealized surface pattern for the deformation type events.

Mean vertical profile for the events. Top number is temperature bottom number is dewpoint depression in oC except for values at the surface which is in oF with the bottom number being dewpoint.

Winds are southwesterly with the profile being very similar to that of the Maddox type II composite. The main difference between the Maddox composite and the Chappell deformation type event, is in the latter, the 500-hPa trough is more progressive than in the Maddox II type composite. Otherwise the events are almost identical.

It is not critical that the pattern exactly fit one the Chappell or Maddox composites.
Here are a few important questions to ask about the pattern.
1) is high (1.00 or greater) precipitable water available nearby that can be tapped?
2) will the airmass be unstable?
3) will the wind shear be on the light side?
4) Will a mechanism such as a shortwave be in place to provide forcing to weaken the cap?

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Maddox type IV event

Another flash flood pattern occurs during the monsoon season when a shortwave, usually and easterly wave moves under or around the southern and southwestern edge of the subtropical ridge. This type of event is strongly tied to the southwest monsoon season and is almost exclusively observed in Arizona, New Mexico, southern California, southern Utah and southern Nevada. The surface pattern with these events is usually ill-defined. One configuration of the general idealized 500-hPA (left) and 700-hPa (right) patterns is shown below.

Both figures are adapted from Chappell’s COMET notes.

There is often a weak mid-level low near or just west of Baja that combines with ridge or 500-hPa low near or just east of the 4-corners to produce deep southeasterly low level flow. A weak spoke of vorticity often rotates around the weak upper low with an extension of it that moves into the southwest U.S. The threat area is usually located in advance of this vorticity lobe. Often satellite imagery, especially water vapor imagery, provides the best clues of the existence of the weak 500 low and any vorticity lobes associated with it. Also, satellite images sometimes show very intense thunderstorm complexes that develop first over Mexico or the Gulf of California and then subsequently propagate northward or northwestward into the U.S. during the late night or early evening hours. At 700-hPa there is often a thermal ridge located near the threat area with some warm advection implied across the threat area.

The surface pattern is often indistinct but there is often a thermal low over mexico with a thermal trough extending across southern Arizona into California. Precipitable water averages 1.32” but surges can lead to values much high than that. The GOES sounder Precipitable water product and AMSU total precipitable water loops are great resources for monitoring moisture surges. A site that will provide a loop of the latter is located at http://amsu.cira.colostate.edu/TPW/global.htm.
Winds at all levels are very weak.

The original Maddox et al. article offered a slightly different composite than Chappell for type IV events. His composite has a easterly wave moving under the 500-hPa trough with any upper low farther south than the Chappell pattern. The main similarity is that the threat area is in advance of the westerly moving impulse in an area of deep moisture and weak mid and upper level flow. The area of low dewpoint depressions scalloped in the figure.

Surface dewpoints above 60oF extend across Arizona into southern California.

An example of a warm season western flash flood

Now let’s look at an actual flash flood case in the west. First look at the various archtypes. Note that there is a heat low over depicted over southern California and to its east, there is a large area of 60oF plus

12 UTC ETA 850-hPa winds, dewpoints superimposed on a precipitable water image

12 UTC surface analysis, surface dewpoints are dashed

3

6

9

6

3

6

9

6

12 UTC 8 Jul 1999 500-hPa height and absolute vorticity analysis

15 UTC 8 Jul 1999 500-hPa height and absolute vorticity analysis

Surface dewpoints extending across Arizona and southern California that is very similar to the area shown on the Maddox et al., type IV archtype. There is also a weak vorticity lobe (the area labeled with the 9) extending across southern California at 500-hPa and an inverted looking trough suggesting that some type of impulse is working westward to the south of the upper level ridge. The soundings at 1200 UTC for Desert Rock, Nevada (DRA) and San Diego (NKX) indicate there is a deep, moist unstable layer and that the precipitable water values are well above normal (1.68 for DRA (around 168% of normal) and 2.01 inches for NKX).

15 UTC 8 Jul 1999 300-hPA, heights winds and temperatures.

The soundings suggest more moisture than shown on the composite type IV from the Chappell COMET notes. The deep moisture and almost MAUL (moist, absolutely unstable layer) and weak winds aloft suggest that the cells may move slower than is typical for cells over the southwest. The rather moist sounding suggests that potential downdrafts and outflow will be weaker than normal. The weak flow suggests advection of the cells will be weak.

The winds on the 15 UTC 300-hPa analysis are fairly typical of the pattern associated with this type of event. The flow is often anticyclonic at upper-levels. This ensures weak inertial stability which makes it easy for upper- level divergence to form with any cell that develops.

500-hpa heights and water vapor imagery for 18 UTC 8 July 1999

Radar precipitation estimates for the event

Water vapor imagery (top left) indicates the convection had started to develop by 18 UTC. The QPE for the date is shown above. There were a number of areas of enhanced convection but that the serious flash flooding was located near Las Vegas rather than where the radar estimates were showing the heaviest rainfall accumulations farther to the northeast. The heavy rainfall amounts suggested by the radar imagery farther to the northeast may have been overdone because of the distance of the echoes from the radar site. At such a distance, the beam may be sampling a higher part of the cloud where hail contamination is more likely. Another possibility is that there was very heavy rainfall across the area to the north but that few people live in the area so any flooding would go unnoticed. The lack of good radar coverage is a forecast problem in the west and is a reason to combine radar information with satellite estimates when assessing rainfall over the intermountain region.
Floodwaters in the Las Vegas metropolitan area damaged or destroyed 369 homes with damage to public property amounting to $20,500,000 and caused two deaths.

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West Coast cold season heavy rainfall events

Extreme rainfall events along the west coast typically occur during the cold season from late fall to through spring. The figures below show the number of 4-inch per days by state during a 10-year period (1991-2000) for the west coast states during the various seasons based on the HPC precipitation analyses. Note that only 2 events were observed during the June-Aug period.

A

B

D

Number of 4 inch events by state during A) fall , B) winter, C) spring, D) summer.

The wet season in the mountains from southern California northward into British Columbia is during the wet season (Neiman et al. 2007). The same study noted that most heavy rainfall events across this region were typically associated with atmospheric rivers (plumes of moisture) with the rainfall and that the magnitude of the rainfall seemed to be related to the magnitude of the integrated water vapor within the plume. The distribution of rainfall is linked closely to the terrain with a one group of maxima along the coastal ranges and another over the Cascade and Sierra Nevada ranges

Climatological profiles of wind speed, specific humidity and moisture flux for the various seasons. Figure adapted from Pandey et al. 1999.

The figures above indicate the seasonal differences in the moisture, winds and moisture flux along the west coast. They imply that the cold season extreme rainfall maximum appears to be primarily due to the stronger winds and horizontal moisture flux in the cold season versus the warm season rather than due to large differences in the absolute moisture in the column (Pandey et al. 1999).

Pattern recognition remains an important component of forecasting extreme rainfall events along the west coast. The 500-hPa patterns that are associated with landfalling moisture plumes are very similar no matter which latitude along the coast they occur at. The two figures below are composite 500-hPa anomaly patterns associated with landfalling atmospheric rivers. The figure at left illustrates the pattern for when the river intersects the blue square. The right

From Neiman et al. 2006

hand figure is the pattern associated with “rivers” that intersect at the red box and produce heavy rainfall in California. Note that each figure has a negative anomaly to the west northwest of location where the river intersects the coast and a positive anomaly to the south and east. The magnitude of the two anomalies differ with the pattern conducive to the northern intersection with the coast having a strong positive anomaly to the south and east than the pattern associated with the southern intersection. The configuration of both sets of anomalies ensures that deep southwesterly flow is present along and off the west coast. This deep southwesterly flow taps the moisture in the subtropics and helps produce the pineapple express. One of the best ways to monitor the moisture within any atmospheric river/pineapple express is by using satellite imagery and the CIRRA link (http://amsu.cira.colostate.edu/gpstpw). The product not only provides the total precipitable water (TPW) but when you scroll down indicates the percent of normal of the TPW.

West coast events, tropical convection and atmospheric rivers

The Pacific jet exhibits fluctuates in response to tropical convection which in turn influences the large scale pattern. The placement of the Pacific jet plays a significant role in determining where and when major west coast rainfall events may occur. The changes in the seas surface temperatures associated with ENSO that influences the pattern along the west coast. The warmer sea surface temperatures in the East Pacific during El Nino years usually allows convection to extend east of the date line. The latent heating associated with this convection helps promote an eastward extension of the East Asian Jet and its associated exit region. During some El Nino years, especially the stronger ones, the exit region of the jet and its associated region of strong cyclogenesis becomes located just upstream from California. The southward shift of the jet and westerlies leads to a southward shift of the storm track (Bell and Higgins, 2005). Extreme rainfall events along the length of the west coast are most common during neutral years (Higgins et al. 2000) so they can occur during any season.
1980 was a weak El Nino, its pattern is fairly typical of the Nino associated pattern for heavy rains in southern California and Arizona. The upper level jet is suppressed southward, in this case, the jet cut under a blocking high. Not how the exit region of the jet extends eastward towards the coast and how height off the coast of California are much below normal.

200-hPa analysis of geopotential heights and isotachs for 1200 UTC 13 Feb 1980.

500-hPa analysis of geopotential heights and of the standardized height anomalies for 1200 UTC 13 Feb 1980.

The block and suppressed jet helps produce a series of storms over the region. By contrast, during La Nina years, the cold SST over the east Pacific keep the convection confined to the western Pacific which leads to the exit region of the jet being much farther to the west which leads to the mean upper level ridge being located farther to the west and to above normal heights across California which pushes the jet northward towards Washington and Oregon (From Bell and Higgins 2005)

Convection associated with the Madden-Julian Oscillation play a role in influencing where extreme precipitation falls along the west coast (Bond and Vecchi 2003, Higgins et al. 2000, Mo and Higgins (1998). CPC has developed a schematic that roughly shows how the synoptic pattern, upper-level jet and moisture plume tends to respond to tropical forcing (see below). They note that as convection become enhanced near Indonesia that a strong blocking ridge develops and that as the convection moves eastward in response to the MJO that the subtropical jet starts undercutting the ridge which retrogrades and a trough amplifies to the south and east of the ridge. Concurrently, a moisture plume or atmospheric river begins to elongate northeastward and eventually impinges on the coast.

Mo and Higgens noted that drier than normal conditions were typical over California when the blob of convection associated with the MJO moved eastward to 120E. One way to assess the convection is by looking at negative anomalies of outward longwave radiation (OLR). Wet periods are more likely to occur develop as the convection shifts eastward with onset usually occurring when the convection approaches 150E.

Bell and Higgins have noted that the evolution of the pattern over southwestern U.S. during the December 18, 2004-January. The area of green (lower values of outward longwave radiation) indicate on Fig. a indicates there is enhanced convection of Indonesia. At around the same time, note there was a trough and enhanced upper level jet just to its north and a block just off the coast of western Canada very similar to the pattern that typically occurs 7-10 days prior to a multi-day period of enhanced rainfall along the west coast. Note how the mean ridge then retrogresses westward and allows a trough to form off the coast of California. In this case, the first heavy rainfall event occurred earlier

Mean 200-hPa heights and wind speeds v.t. 22 Dec. 2004

v.t. 28-29 Dec. 2004

v.t. 7-11 January 2004

than would be expected by their schematic and occurred when the enhanced convection was still over Indonesia when the first major rainfall event occurred over southern California during the Dec 27-29 period. The early onset suggests caution is needed in following the timing of the schematic too closely. The MJO convection does not move steadily but in fits and starts. Chaos also plays a role in how patterns evolve. Not all heavy rainfall events are associated with the MJO.

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C

Accumulated rainfall in inches from 6 different storms during a 9-day period (Feb 13-21, 1980) in southern California.

Polar diagrams of floods on the Sauk, Snoqualmie, and Chehalia Rivers during a) Oct-Dec., b) Jan-Mar. for 1979-2000. The magnitudes of the floods in terms of streamflow are proportional to the radius (the outer ring represents the all time record peak on each river): the timing of these floods with respect to the phase of the MJO is shown by the azimuth (see labels on the perimeter of each diagram). From Bond and Vecchi, 2003.

The Bond and Vechii research shown above suggests that the MJO may be an important player in flood events during Oct.-Dec. period in western Washington as 81% of the floods during that period occurred in MJO phases from very late in phase 5-through early phase 8 with the bulk of the events falling in phase 7. From the Figure, Bond and Vechii avowed that there was evidence of MJO modulation of the Oct-Dec floods with a statistical significance of 99%.
The relationship to floods on the western Washington rivers to the MJO is less clear during Jan-March. The phases associated with flooding in Jan-Mar. period also differed from the phases associated with the fall and early winter floods. It is unclear how much help the MJO can be at anticipating the onset of extreme rainfall along the west coast in the medium range.

Southern California pattern

The composite mean rainfall for a number of heavy rainfall events in southern California is presented below. The composite mean 500-hPa anomaly pattern for these events is also shown. The negative anomaly center is farther to the south than for the cases when the plume is impacting northern California, almost directly west of San Francisco. Such a position leads provides a a fairly strong southwesterly to south-southwesterly gradient across the mountains of southern California. normal to the mountains.

Composite of 24-hr rainfall for heavy rainfall events in the mountains near Los Angeles. From Bell et al., 2005

Composite mean 250-hPa winds pattern for the beginning of the period of heavy rainfall. From Bell et al., 2005.

Composite mean 850-hPa winds pattern for the beginning of the period of heavy rainfall. Outer magenta area is where the anomaly of the v-component (southerly) of the 850 winds have standard deviation (S.D.) of a greater than 1.8 and the inner oval is where the departure is greater than 2.1 S.D.

The Bell et al study suggested that the left exit region of an upper level jet streak might be enhancing the upper level divergence across southern California while also contributing to stronger than normal v-component to the 850-winds due to the ageostrophic response to the upper branch of the jet streak’s transverse circulation.
The most extreme events flood cases often result from multiple storms with one or two being of them producing extreme rainfall. February of 1980 in southern California and Arizona is a case in point. The steady onslaught of storms brought many streams record volumes of run-off in southern California. The 500-hPa analysis (bottom right) shows one

An example of a southern California event.

The NCEP operational models often do a good job forecasting the basic synoptic pattern associated with many of the heavy rainfall events along the west coast especially in the shorter time ranges.

Analysis of 500-hPa heights and standardized anomalies v.t. 1200 UTC 27 Dec. 2004.

12hr GFS 500-hPa and vorticity v.t. 0000 UTC 28 Dec. 2004.

30hr GFS 500-hPa and vorticity v.t. 18 UTC 28 Dec. 2004.

The GFS forecast of the 500-hPa pattern indicated that the intense 500 low with a greater than 3 S.D. height departure would be slow moving during the period. The 500-hPa forecast fits the composite for 4 inch or greater rainfall, What about other fields? Note how similar the forecast 250-hPa isotach (below) field is to the one in the composite for the southern California heavy events.

30-hr GFS 250-hPa heights, isotachs, ageostropic winds (arrows) and divergence v.t. 1800 UTC 28 Dec. 2004

30-hr GFS 850-hPa winds and PW v.t. 1800 UTC 28 Dec. 2004.

The predicted 850 winds in the vicinity of southern California is actually stronger than on the composite mean for the southern California events and the PW field shown above is very similar to the composite mean suggesting that above normal moisture flux was also present at 850-hPa. The pattern predicted by the models is very conducive to getting 5 inches of rain in a 24-hr period along the coast or over the east-west oriented mountains of southern California.

Legend (feet)

30hr NAM boundary layer moisture transport (arrows), Orographic omega (color filled are positive values), and 80-km depiction of model terrain (brown lines, in meters.

10-km resolution depiction of terrain.

Now look at the 10-km find the areas where the changes in elevation are greatest. The areas with the strongest gradients on the 10-km depiction are considerably different than shown on the 80-km eta depiction. Now try to visualize where the strong orographic lifting will be based on the boundary layer moisture flux arrows predicted by the ETA and gradients using the higher resolution terrain. The yellow arrows point to the location of the tightest gradient and the location where the terrain induced lifting is greatest. The PRISM precipitation climatology has rainfall maximum over both regions where the gradient is tight. Neither rainfall maximum is right on the coast. Admittedly, the PRISM relies on gradients of terrain to infer

Rainfall maximum

10-km PRISM climatology

precipitation in data scarce regions. Western region uses PRISM to help downscale their precipitation analyses so it is usually wise to model precipitation maximum fairly close to the PRISM maxima if the flow is typical for significant rainfall event.

Neiman et al. (2002) have developed a conceptual representation of the distribution of rainfall in California’s coastal mountains for blocked and unblocked flow. Blocked flow is essentially flow in which there is considerably veering of the winds with height at the lowest levels and situations when there also is a marked terrain parallel barrier jet.

(c)

Conceptual representation of orographic rainfall distribution in California’s coastal mountains, and the impact of terrain-blocked flow on this distribution: (a) plan view of unblocked and blocked flow, and (b) and © show the cross-section perspective of unblocked and blocked flow respectively, with representative coastal profiles of wind velocity (flags and barbs) and the correlation coefficient (based on the magnitude of the upslope flow at the coast vs. the rain rate in the coastal mountains) shown on the left. The variable h in (b) and © is the scale height of the mountain barrier. The spacing between the rain streaks in (b) and © is proportional to rain intensity. The small circle in © represents the location of terrain-parallel barrier jet. From Neiman et al. 2002.

The research by Neiman and his colleagues suggests when the airmass is stable in the low layers with veering winds with height that the precipitation is less concentrated moderate to heavy rainfall extends upstream farther from the mountains then cases without veering winds with heights or the stable layer. Their research was conducted for the coastal mountains but there is reason to suspect that similar blocking can also occur along the Cascade and Sierra ranges (Doyle 1997, Steenburgh and Mass 1996).

When might the modeling the precipitation to the PRISM data distribution be misleading and lead you to underpredict the rainfall in the valleys of California.

A) With the post frontal rainfall

B) a system with strong veering with heights and warm advection.

C) With strong frontogenesis with some instability.

D) with strong and deep southwesterly flow when an atmospheric river is impinging into the coast.

For the answer, see next page

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Composite mean 500-hPa height anomaly pattern for the beginning of the period of heavy rainfall on southern California. From Bell et al., 2005

Southern California case continued

All the answers but “A” and possibly “D” are good answers. During post frontal situations along the west coast, the flow often is fairly uniform with little veering and blocking, also cold mid-level temperatures often lead to instability. Answer D is tricky as deep southwesterly flow is conducive to enhanced orographic precipitation. However, a barrier jet can form with the jet being parallel to the mountains. When this occurs, some blocking of flow is likely which can lead to rainfall occurring upstream of the mountains. Strong veering with heights is favorable for blocking (answer B). In these cases the heaviest rainfall still may be near the mountain tops but there is often moderate to heavy rainfall upstream from the mountains (see the figure on the previous page). Strong frontogenesis (answer C) in conjunction with instability often allows heavy rainfall to spread downstream from mountain barriers, in this case the coastal mountains allowing moderate to heavy rainfall to occur in the valley.

The distribution of rainfall along the coastal mountains, Sierra and Cascade ranges sometimes are also modulated by narrow but intense bands of associated with the front (Reynolds, 1996). The complex structure that sometimes occurs during these storms is illustrated by the two IR satellite images of the Dec. 28, 2004 heavy rainfall event. The images also suggest how difficult it can be to predict correctly in which 6-hr period the heaviest rainfall will occur as there sometime are periods of subsidence between bands. For example, at 1500 UTC 27 December the band right on the coast shows more enhancement than the next band of brighter clouds to its west but by 2100 UTC the initial band has weakened while the clouds in the second band have become enhanced and are obviously associated with convection.

GOES IR imagery v.t. 1500 UTC 27 Dec. 2004

GOES IR imagery v.t. 2100 UTC 27 Dec. 2004

The area of apparent subsidence between the two bands also shows signs of cooling at 2100 UTC. However, the extensive convection within the band off the coast also suggested that the heavy rainfall might not be limited to near the crest of the mountains but could also fall right on the coast as this convection shifts into the coast. The GFS and NAM forecasts of the timing of the band of heaviest precipitation differs (see below). Resolving the differences between the timing is often problematic. In the shortest time ranges (0-12 hrs) satellite imagery can sometimes be useful. However, at longer time ranges, the best bet is often to rely on the consensus of all the operational models and ensemble members to time the bands.

24-hr ETA 6-hr precipitation v.t. 1200 UTC 28 Dec. 2004.

30-hr ETA 6-hr precipitation v.t. 1200 UTC 28 Dec. 2004.

24-hr GFS 6-hr precipitation v.t. 1200 UTC 28 Dec. 2004.

30-hr GFS 6-hr precipitation v.t. 1800 UTC 28 Dec. 2004.

The 6-hourly RFC observed rainfall on the figures is an example of how the heaviest precipitation often evolves non-linearly. The first maximum that occurs during the 6-hr period ending at 1800 UTC appears to be related to orographic lifting from the southerly winds in advance of the front and the frontal band itself. Note that over 3 inches of rain was observed during the 6-hr period over the mountains near Los Angeles and that this maximum is underpredicted by either model at least when the precipitation is depicted at the 80-km resolution of the workstation. This again argues for looking at precipitation forecasts in complex terrain at the highest resolution possible.

0600 UTC 28 Dec

1200 UTC 28 Dec

1800 UTC 28 Dec

0000 UTC 29 Dec

0600 UTC 29 Dec

1200 UTC 29 Dec

ETA heights and vorticity v.t. 0900 UTC 29 Dec.

A second maximum was fairly well forecast and was associated with the initial mid level trough opening up. By this time, the model was predicting unstable LIs (not shown). Having a break between the two heaviest period of heavy rainfall is not all that uncommon in a case where a vorticity maximum and the cold pool aloft shift over the region behind the main baroclinic band.

Using climatological anomalies as a forecast tool along the west coast.

Hart and Grumm (2001) have used standardized anomalies to rank synoptic scale events from a climatology developed from the NCEP reanalysis. The standardized anomalies essentially scale the anomaly by the its standard deviation from the mean. The beauty of using standardized anomalies (SDs) is that the standard deviation provides information about the frequency of occurrence of whatever forecast parameter than you are looking at. As noted by Grumm and Hart (2001), a departure from normal of 2.5 SDs implies that the observed value occurs between 5% and 16% of the time. For fields like 500-hPa or 700-hPa heights the actually confidence limit is probably close to that of a normal distribution of 5%. Neither precipitable water (PW) nor the magnitude of 850-hPa magnitude of the moisture flux (MF) have a normal distribution, but even they appear to have confidence limits near that of a normal distribution. Note how over 99 percent of the time, the PW and 850-hPa magnitude of moisture flux is below 3 SDs.

From Junker et al. 2007

Magnitude of moisture flux

Precipitable water

Histogram showing the percentage of the distribution of PW and 850-hPa magnitude of moisture flux.

The 700-hPa height and standardized anomaly pattern for three different multi-day extreme rainfall events are shown below. Note that each has a greater than 3 SD height anomaly off California associated with it. The greater than 3 SD departure from normal suggests that 99% of the time during this time of year, the 700-hPa geopotential height anomaly is less than the observed anomalies during these three events.

-3 SD

-3 SD

-3 SD

700-hPa geopotential height and standardized anomaly of the height field (shaded) valid 0000 UTC 9 Jan 1995.

700-hPa geopotential height and standardized anomaly of the height field (shaded) valid 1200 UTC 9 Mar 1995.

700-hPa geopotential height and standardized anomaly of the height field (shaded) valid 1200 UTC 31 Dec 1997.

Heavy west coast precipitation events during winter usually are associated with a fairly strong negative anomaly with a weaker positive height anomaly to the southeast. All three of these multiday extreme rainfall events were associated with such a pattern which pretty much assures a strong southwesterly geopotential height gradient. The strong gradient between the two anomalies and the southward displacement of the jet helped pull and plume of moisture or atmospheric river northeastward into the coast. Note that during both the March 1995 and Dec. 31 events, that precipitable water values exceeded 3 standard deviations.

3 SD

4 SD

1200 UTC 9 Mar 1995 PW (contour) and Normalized PW anomaly (shaded, unit=1 SD.

1200 UTC 31 Dec 1997 PW (contour) and Normalized PW anomaly (shaded, unit=1 SD.

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5 SD

5 SD

850-hPa moisture flux (arrows) and magnitude of the moisture flux (solid) and anomaly of the magnitude of the moisture flux (shaded, unit= 1SD) valid 0000 UTC 9 Jan 1995.

850-hPa moisture flux (arrows) and magnitude of the moisture flux (solid) and anomaly of the magnitude of the moisture flux (shaded, unit= 1SD) valid 1200 UTC 9 Mar 1995.

All three events were associated with plumes of moisture that impinged upon the coast. The role of these plumes or atmospheric rivers is even more evident when anomalies of the magnitude of 850-hPa moisture flux is examined. Let’s first look at the moisture flux anomalies associated with the January and March 1995 events (see below). The shaded areas are anomalously high MF anomalies with the white area having a 5 SD departure from normal.

The 5 SD moisture flux anomaly impinging upon the coast suggests that the both 1995 rainfall events might be relatively rare as more than 99.999% of the time the moisture flux will be lower than on the two maps above. Significant flooding and flash flooding also occurred over the Sacramento, Napa, and Sonoma Valleys as a result of rainfall during the eight-day period, January 7-15, 1995 (NOAA 1995), prompting thirty-four counties to be declared federal disaster areas. The March 9-10, 1995 caused considerable flooding on the San Joaquin basin (Reynolds 1996). The combined damage from the January and March 1995 floods exceeded 305 million dollars in the Sacramento Valley and 193 million in the San Joaquin Valley (Army Corps of Engineers 1999).
Let’s explore the 850-hPa moisture flux and MF anomalies during the period from 0000 31 Dec. 1996 through 0000 UTC 2 Jan. 1997 and explore the relationship of the maximum rainfall to the axis of strongest 850-hPa MF. During the entire period, the standardized MF anomalies were at least 4 SD and usually 5 SD. Again the normalized anomaly pattern signalized that there was the potential for a relatively rare, extreme rare precipitation event over California. This particular event, caused nearly two billion dollars in property damage and forced over 100,000 people from their homes (Baird and Robles 1997).

850-hPa moisture flux (arrow, scale shown at bottom of image), magnitude of moisture flux (interval=0.05 ms-1) and normalized anomaly of moisture flux (magnitude of the anomaly is shaded with the scale on the right hand side of the figure, interval=1 SD) for a) 0000 UTC 31 Dec. 1996, b) 1200 UTC 31 Dec. 1996, c) 0000 UTC 1 Jan. 1997, and d) 0000 UTC 2 Jan. 1997.

During upslope events, the anomalously moisture flux impinging upon the coast and mountains helps determine the location of where the heaviest rainfall will occur. The axis of strongest moisture flux was located north of San Francisco during the 24 hour periods ending at 1200 UTC 31 Dec. 1996 and 1200 UTC 01 Jan. 1997. By 0000 UTC 2 Jan. 1997 the axis had shifted south of San Francisco, only then did very heaviest rainfall shift south of the city (see figures below).

24 hr HPC precipitation analysis for, a) 1200 UTC 31 Dec 1996, b) 1200 UTC 1 Jan 1997, and c) 1200 UTC 2 Jan 1997. From Junker et al. 2007

All three of these multi-day rainfall events was associated with a “pineapple express” or atmospheric river where there was anomalously high PW and MF. Ralph et al. (2006) seven floods on the Russian River that were associated with landfalling atmospheric rivers attesting to their importance in producing excessive rainfall.

a

c

b

d

Composite mean standardized 850-hPa moisture flux anomalies of moisture flux field at the beginning of the period of heavy rainfall, a) for 4 inch or greater events and b) for 2-4 inch events; at the end of the period, c) for 4 inch or greater events, and d) for 2-4 inch events (from Junker et al. 2007).

Composite mean standardized 500-hPa departure from normal anomalies at the beginning of the period of heavy rainfall, a) for 4 inch or greater events and b) for 2-4 inch events; at the end of the period, c) for 4 inch or greater events, and d) for 2-4 inch events. (from Junker et al. 2007).

There are differences in the moisture flux and PW anomaly fields between events that produce heavy rainfall compared to those that produce more modest amounts (see above). Note that 500-hPa departures from normal indicate that the shortwaves associated with the heaviest events are slower moving than the more moderate ones. The differences in the anomaly fields for moderate and heavy support the idea that anomalies can be used to identify major rainfall events. The similarity of the 500-hPa, 700-hPa and 850-hPa moisture flux patterns at locations from the Pacific Northwest (Lackmann and Gyakum) and southern California (Bell and Pereira 2005) to the anomaly patterns shown above for northern California suggest that standardized anomalies may also be a useful forecast tool along the entire west coast.

An example of how you might use standardized anomalies to make a forecast.

Pattern recognition remains an important component of predicting extreme rainfall events in the mountains along the Pacific coast of North America. The 500-hPa geopotental height and associated anomaly pattern is very similar for all heavy rainfall events with the primary difference being associated with the location of negative anomaly associated with the trough and the positive anomaly to the southeast. Note how similar the anomaly pattern associated with the heavy precipitation events in the Pacific Northwest (below right) to the pattern associated with heavy events in northern California (below right) at the beginning of the period of heavy precipitation. These similarities suggest that once a forecaster is able to recognize the ingredients associated with heavy rainfall in one place along the west coast, he or she should be able to translate that knowledge to enable them to identify possible extreme events in other places along the west coast.

From Lackmann and Gyakum 1999

Pacific Northwest composite

Northern California composite

Composite mean 500-hPa departure from normal (positive departures are solid, negative departures are dashed, contour interval=3 dam). Shaded areas are the 95% and 99% probably that the composite belongs to a distinct population from climatology.

Composite mean 500-hPa normalized departures from normal. Unit is 0.3 standard deviations. Negative values are below normal heights, positive are above normal heights. From Junker et al. 2007.

Now we’ll attempt to apply pattern typing and the use of normalized anomalies to forecast a rainfall event over the Pacific Northwest. The first step is to compare the forecast geopotential height field to the composite for heavy events over that area. Note the similarities between the GFS forecast 500-hPa geopotential height field to the composite (below). Both exhibit a fairly strong trough in the height field over the Gulf of Alaska and distinct ridging over California.

534

546

96

582

570

558

16

20

Composite mean 500-hPa height (solid) and sea level pressure (dahsed , every 4 hPa) at the beginning of the period of heavy precipitation over the Pacific Northwest.

12-hr GFS forecast of 500-hPa geopotential height and vorticity valid 1200 UTC 6 Nov. 2006

76

96

20

12-hr GFS forecast mslp and 100-500-hPa thickness (dashed) valid 1200 UTC 6 Nov. 2006

Also note the basic look of the forecast mslp pattern with both the surface low position in the Gulf of Alaska and the ridge axis extended from California west-southwest well into the Pacific and compare it to the composite (dashed line, above left). The axis of the area of strongest surface pressure gradient is also similar. However, there is a notable difference, the gradient is much stronger in the forecast than in the Lackmann et al. composite.

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Example continued

In their article, they defined a heavy rainfall day as one in which Seattle (SEA), Olympia (OLM) and Stampede Pass (SMP), WA and Astoria (AST), OR all received 0.50 inches of rain. Heavy but certainly not extreme precipitation. The surface pressure gradient suggests that the precipitation might end up being considerably heavier than that. The long fetch implied by the pressure gradient also argues that a strong low level jet may help pull a plume of tropical moisture east-northeastward.

12-hr GFS forecast of PW and 850-hPA winds valid 1200 UTC 6 Nov. 2006

24-hr GFS forecast of PW and 850-hPA winds valid 0000 UTC 7 Nov. 2006

The GFS forecasts of PW and 850-hPa winds indicate there is an impressive fetch of deep moisture and that there is definitely and atmospheric river associated with the event. One way to check if the model appears correct is by accessing the CIRA blended TPW product at
http://amsu.cira.colostate.edu/TPW/global.htm. During this case, the blended product supported the model forecast PWs. The “pineapple express” or atmospheric river was going to be a player in the forecast.

An experienced forecaster would probably recognize that there was above normal PWs within this river but he or she still might not have an objective way of knowing how often such a plume occurs or how uncommon such high PWs are. The individual would have an even tougher time assessing a less commonly used diagnostic tool like moisture flux. Standardized anomalies offer a way of assessing how often the values of a particular meteorological field that are being forecast by the numerical models is present providing that the model has no bias in predicting that particular field.

SREF ensemble mean PW (brown lines, in inches) and standardized PW anomaly (dashed red) valid 1200 UTC 6 Nov 2006

SREF ensemble mean PW (brown lines, in inches) and standardized PW anomaly (dashed red) valid 0000 UTC 7 Nov 2006

SREF ensemble mean PW (brown lines, in inches) and standardized PW anomaly (dashed red) valid 1200 UTC 7 Nov 2006

The SREF ensemble mean forecast of of a standardized PW anomaly of greater than 3 standard deviations (SD) is impressive since ensemble means average values so small differences in the location of the PW axis between the members tend to produce a broader moisture band and with lower maximum precipitable water than the individual members. Therefore, the high standardized anomalies suggest that there was strong agreement between the SREF members as to the location and movement of the band. The 3 SD anomaly or greater PW anomalies also indicate that over 99 percent of time the observed precipitable water is lower than is being forecast suggesting that the PWs being forecast are in the top 1% of events occurring during the same time of year.

SREF 850-hPa moisture flux (g/kg) ms-1 and standardized magnitude of moisture flux anomaly valid 1200 UTC 6 Nov. 2006.

SREF 850-hPa moisture flux (g/kg) ms-1 and standardized magnitude of moisture flux anomaly valid 0000 UTC 7 Nov. 2006.

SREF 850-hPa moisture flux (g/kg) ms-1 and standardized magnitude of moisture flux anomaly valid 1200 UTC 7 Nov. 2006.

The SREF ensemble mean 850-hPa moisture flux anomalies are even more impressive reaching over 5 standard deviations. Such a high standardized anomaly suggests that November observed moisture flux has rarely, if ever, exceeded the forecast values since 1950. Unless the SREF ensemble mean forecast of moisture flux has a high bias, this has the potential to possibly be a rare event. Such a slow moving moisture flux axis with such anomalously high moisture argues that this event might have the potential to be a historic one.

Knowing that the SREF predicted moisture flux anomalies are suggesting that a rare heavy event might be in the offing, how might the forecaster modify the model forecasts below?

12-36 hr QPF from the NAM model, scale in inches at left. Depicted on a 32-km grid.

12-36 hr QPF from the GFS model. Depicted on a 32-km grid.

When depicted on a 32-km grid, the NAM and GFS predicted a significant rainfall event, however, the areal distribution of the rainfall differed between the models. In a case like this one, it’s important to look at the NAM precipitation output at the highest resolution possible. The model with the better horizontal resolution of the terrain predicted less rain then the coarser resolution GFS. Both models predicted a relatively uncommon event but neither predicted amounts that suggested record rainfall might occur. However, the highly anomalous moisture flux and PWs suggested this could indeed be a near record or record event.

C

A

D

12-36 hr QPF from the HPC forecaster. Depicted on a 32-km grid.

One aid to amending model forecasts in this type cases is to model rainfall maximum to the prism climatology areas where enhanced rainfall is likely and by making the changes based on how long the axis of highly anomalous moisture flux is situated over those regions. To a certain extend, the HPC forecast looks like it employed some form of this method.

Four points are identified on the 12-36-hr HPC QPF. Looking at the SREF forecasts of moisture flux and the displayed of the normalized anomalies. Try to identify whether you would decrease or increase the rainfall over each of the 4 points.

Let’s first look at point A. Note that point A is the heaviest location on the PRISM climatology. The SREF forecasts indicate that the moisture flux anomalies during the first 12 hours of the period are at least 4 standard deviations (sigma) and then decrease to 2-sigma by the end of the period. The GFS forecast only predicted between 3 and 4 inches across this region. The SREF moisture flux anomalies certainly argue for opting for the heavier forecast by the NAM over the Olympic mountains but also argue for more along the north coast of Oregon. The strength, anomalously high values of the moisture flux for an extended period and climatology might suggest that the rainfall at point “A” would be at least as high as point “C”.
The HPC prediction of the rainfall at point “B” compared the forecasts by either the NAM or GFS is a major improvement and is well supported by the moisture flux anomalies and climatology. Moisture flux anomalies are forecast to be above 3 sigma for 24 hours and are predicted to be above 5 sigma for over 12 hours. The anomalies at both point “B and “D” are indicative of a rare event. The PRISM climatology indicates that on average, point B will get more precipitation than point “D” (see below left). The combination suggests that point “B” might be a candidate for as much rainfall as point “D” despite the model forecasts to the contrary since the primary forcing is due to orography.

B

D

B

D

Reforecast product PQPF for 4.00 inches during the 24-hr period ending at 1200 UTC 7 Nov. 2006.

The analog PQPF product based on retrospective forecasts (Hamill and Whitaker 2006) offers an aid for determining which areas are more likely to get the heaviest 24-hour rainfall. These product provide calibrated probabilities for various thresholds based on a 25-year set of probabilities. Note that the 4-inch probabilities were higher over the Olympic mountains than over the Washington Cascades near point “C”. The same holds for the coastal ranges near the coast compared to the Cascade Range near point “D”. The PQPF for various thresholds are available on the HPC workstations under NMAP2, GRIDS, analog PQPF. These products are also available on the Internet at http://www.cdc.noaa.gov/reforecast/narr/.

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B

The Hamill and Whitaker reforecast method also can be used to provide a guidance product that tells the forecaster the precipitation amount that was only exceeded in 10 percent of the cases identified for that particular day. The graphic below left is such a product for the 24 hrs ending at 1200 UTC 7 Nov. 2006.

24-h analog ensemble QPF-90th percentile (upper decile) (in) product valid 1200 UTC 7 Nov. 2006.

The greater than 3 SD and 5 SD departure from normal for PW and 850-hPa moisture flux predicted by the SREF ensemble mean suggests that this event will at least equal the 24-h QPF from upper 10% of the reforecast project analogs suggesting a heavier rainfall maximum over the Olympic Mountains and coastal ranges of northwest Oregon than predicted by the NAM or GFS.

One downside of the technique is that the forecasts are predicated on a low resolution ensemble forecast system which may not be a skillful as the SREFS. Therefore, the high precipitation amounts on the decile product can extend over too large an area at times in the shorter time range when in reality the SREF output is suggesting that the spread is quite small and that the heavy amounts will be confined to those areas with strong moisture transport and strong orographic lifting.

24-h accumulated precipitation (32 km grid) valid 1200 UTC 7 Nov. 2006.

24-h accumulated precipitation (4 km grid) valid 1200 UTC 7 Nov. 2006.

Two different precipitation analyses are shown for the event, one a mapped to 32-km, the other to 4-km. During this period, Lee’s Camp in Oregon received 14.3 inches and June Lake in Washington received over 15. Record flooding was reported on 10 rivers and 16 counties were declared to be in a state of emergency by the governor or Washington. The standardized anomalies correctly identified the potential for a rare event. For a summary of the effects of the event on the Oregon and Washington see http://www.ocs.orst.edu/page_links/whats_new/november_flooding.html.

Leeside cool season spillover events

There is usually a significant “rain shadow” to the lee of major mountain ranges like the Sierra Nevada range of California and the Cascade range of Oregon and Washington. However, there are cases when the “shadowing” effect is minimized. Wallmann and Milne (2007) (henceforth WM) note that there are a lot of misconceptions about when the “rains shadow effect” will be strong or weak. They site the following taken from NWS Forecast discussions. “South winds across the Sierra Nevada will limit the rain-shadow effect and result in more precipitation for Reno,” “moisture depth will be sufficient for precipitation to spill over into western Nevada,”strong southwest winds will result in a strong rainshadow effect.” None of these “rules of thumb” works particularly well since they are not based on scientific understanding of the blocking and the effect it has on the “rain shadow.” This section which will attempt to provide a better understanding of the “rain shadow” draws heavily on the work done by Wallmann and Milne (2007) and on Wallman’s 2007 presentation at HPC.

Research by Colle (2004) identified a number of factors that affect the distribution of orographic precipitation. He noted that the distribution was “highly dependent” on how the terrain-induced gravity wave modifies the circulation aloft. His research suggests that the amount of precipitation that “spills over” the mountain crest is a function of stability, mountain height and width, wind shear, and freezing level.
Colle’s research indicates that the amount of precipitation that spills over the mountains increases and the stability decreases and as the wind speed increases and also indicates that a higher proportion of the condensate fall on the windward slopes when the freezing level is high versus when it is low if all other factors are equal. When the freezing level is high, most of the condensate is in the form a rain which falls quicker which positions the rainfall farther upstream from the mountain crest than when the freezing level is low.

When trying to decide how much precipitation will slop over, probably the most important factor is the stability. When the air mass is stable, blocking occurs even when there is strong winds into the terrain. Also, strong southwesterly flow towards the mountains can lead to the development of a southerly barrier jet. Not much precipitation makes it to the lee side except for brief showers. WM offered the following sounding as typical of a blocking case and accompanying radar image. They note that the stable isothermal layer from 750-hPa to 675-hPa and less unstable layer above it is very indicative of cases with strong subsidence resulting in a enhanced “rain shadow”.

A similar stable layer near the crest level has been noted in downsloping wind events to the lee of mountains (Colle and Mass 1998).

0600 UTC 31 Dec. 2004 KREV sounding

0600 UTC 31 Dec. 2004 radar reflectivity. Red lines are Nevada state line.

The WM paper also offers the following sounding as an example of an unstable one where there was significant spillover of precipitation to the lee of the mountains. Note that it is only barely unstable in the classical sense in the layer between the two lines. They note that there will be some spillover for weakly stable events but that they will not have the precipitation extend as far downstream as the unstable cases. For weakly stable cases, how far it extends downstream from the mountains crest is dependent on the wind. The figure below shows the simulation of rainfall with respect to the crest of a mountain for different wind speeds when the static stability is weak. Note that there is little spillover with weak winds with the maximum rainfall on the lower slopes of the terrain. As wind speeds increased, the maximum shifted towards the crest and the distance the rainfall extended downstream from the crest increased.

Rainfall distribution with respect to the mountains crest (vertical dashed line) for various wind speeds under weakly stable conditions. The gray line with an arrow head is the location where the maximum increase in precipitation was found as the wind was increased to the next level. (From Colle 2004).

Wallmann cautions that even when the air mass is unstable there usually needs to be some kind of synoptic or mesoscale forcing to realize the instability to get precipitation to extend well east of the mountains when the winds are below 30 knots. Without forcing, precipitation east of the mountains will usually weaken and be scattered. He also for jet streaks, frontogenesis and isentropic lift to assess the synoptic and mesoscale forcing and cautions against using Q-vectors as they lack the mesoscale resolution needed to distinguish the scale of the lifting. My own feeling is that looking at the isentropic lift to the lee of the mountains often has the same problem. However, Wallmann that in a moist, weakling stable airmass, its presence can help nudge the “rain shadow” downstream from the mountains.

GFS forecast of 250-hPa isotachs (color fill), 700-hPa winds and temperatures v.t. 1800 UTC 31 Dec. 2005. (Images from Wallmann)

GFS forecast of 700-hPA frontogenesis (shaded, warm colors are frontogenetic) and 850-500-hPa moisture convergence (dashed show convergence) v.t. 1800 UTC 31 Dec. 2005.

A strong jet streak on the west side of the shortwave approaching the coast argues that it will remain strong as it pushes eastward and might even amplify. The entrance region of a jet streak is developing and is associated with frontogenesis and low and mid-level convergence (dashed area) and noted that by 1800 UTC the spillover would be ongoing if the model was correct.

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0.50 deg. Reflectivity for 1500 UTC 31 Dec. 2005.

GFS forecast sounding valid 1800 UTC 31 Dec. 2005.

The radar shows that plenty of rain made it east of the mountains by 1500 UTC. Note that the GFS sounding is saturated and pretty much follows the moist adiabatic. Such a sounding looks like it is common for spillover events. Note the lack of any inversion just above the level of the mountain crest. Wallmann suggests that a forecaster look at a cross section with saturated moisture potential vorticity on it to find when the instability or CSI might be present. Model sounding from NSHARP can help. When the forecast sounding looks similar to the one above and strong forcing is present east of the Sierra ridge top, forecast spillover.

12-36 hr ETA precipitation forecast valid 1200 UTC 31 Dec. 2005

24 hr precipitation analysis forecast valid 1200 UTC 31 Dec. 2005

NAM 36hr 850-hPa wind and moisture flux valid 1200 UTC 31 Dec. 2005.

The was associated with an atmospheric river with moisture flux/transport anomalies of greater than 5 S.D. The high snow levels of 8500 ft. led to heavy rains across the Tahoe basin where rainfall amounts of 4 to 8 inches were reported over a 2 day period. 2 to 4 inches fell in Reno. The heavy rainfall led to widespread flooding along the Susan, Pit, Truckee and Carson Rivers.

The NAM predicted spillover during the Dec. 31, 2005 event but it does not always correctly predict spillover as can be seen below. Spillover events remain a difficult forecast problem.

24-h accumulated precipitation valid 1200 UTC 20 Oct. 2006.

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12-36 hr ETA QPF valid 1200 UTC 20 Oct. 2004.

Hurricanes and extreme rainfall.

Hurricanes, tropical systems and depressions pose a special forecast problem. Tropical storms and hurricanes or their remains are capable of producing extremely heavy rainfall and often produce serious flooding and flash flooding. A presentation by Rusty Pfost documented some of the more impressive rainfall associated with tropical cyclones. He notes that rainfall rates of up to 6 inches per hour have been reported and that 43 inches of rain was reported in 24 hours at Alvin, Texas as a result of the remains of Tropical Storm Claudette. Flooding resulting from the rainfall associated with tropical systems accounted for 292 deaths during the 1970-1999 which was 59% of the total deaths from tropical systems (Rappaport et al. 1999).
HPC quantitative precipitation forecasts are to a certain extent dependent of the track of the system provided by the Tropical Prediction Center. Forecasts are usually modeled to the NHC track with amounts to a certain extent dependent on the size of the storm and its speed. However, the amount of rainfall that a storm produces does not appear to be related to the intensity of the storm. The Rappaport et al. study suggests that there is no apparent relationship between the intensity of the tropical cyclone at landfall and the number of rainfall induced deaths that occur because of the storm.
The heaviest rainfall associated with tropical systems usually falls along or near the coast and typically occurs slightly to the right of the track of the storm. A 46 storm composite of the average 48 hr rainfall is offered below (Goodyear 1968). The same study indicates that there is about a 50% chance that the period of heaviest rainfall will occur in the 12-h period starting 6 hours prior to landfall and ending 6 hours after landfall. The 6 hour period just prior to landfall is the period most likely to receive the maximum 6 hourly accumulations during the storm.

Another study, done more recently by Rusty Pfost, for hurricanes with a landfall in Florida, found that the maximum rainfall was usually a little to the east of the storm. However, Pfost found a few exceptions (see figure below). The maximum rainfall usually falls within the eyewall or with the spiral bands and occur where the frictional convergence is maximized on the east side of the storm. Unfortunately, some storms are highly asymmetric and have more convection on the west side of the storm than the east. Others may be highly sheared with most of their convection located well east of the center.

Forecasters in the in the 60s and 70s often used a rule of thumb (the Kraft rule of thumb technique) to estimate the rainfall associated with tropical systems.

Maximum rainfall (inches)=100/speed (in knots)

Northern Gulf coast storms

Florida landfall storms

Adapted form PFost Powerpoint Presentation

The two graphs above argue that forecasters should not wed themselves to the Kraft technique. The second fastest moving northern Gulf of Mexico landfall storm produced the third highest precipitation maximum of all northern Gulf storms. On both graphs above, a curve has been drawn to try to best fit the points. The slope of both curves is very flat except at speeds below 10 knots where the slope steepens. The lack of a well defined relationship between the speed of the cyclone and its associated precipitation maximum should come as no surprise since the Kraft method fails to take into account the following: 1)the size of the storm, 2) the amount of convection associated with it, 3) that some storms are asymmetric due to shear and 4) the impact of nearby fronts. The graphs suggest that the maximum storm total rainfall will be in the 5 to 10 inch range for storms moving between 6 and 30 knots. The rainfall maxima for storms moving slower than 6 knots was usually above 15 inches.

The Satellite Tropical Rainfall Potential (TRaP) technique provides another more scientifically based method of estimating the rainfall potential of a tropical cyclone as it makes landfall (Kidder et al. 2005). Equation used to determine the rainfall using the TRaP technique is provided in the figure below.

Images from Scofield, Kuligowski

Where Ravg is the hourly rainfall rates measured by the satellite sensors, D is the measure length along the axis of the storms movement and V is the velocity of the storm. The speed and direction that TRaP uses for its rainfall predictions comes from the NHC forecast tracks.

TRaP rainfall forecast

Eta model forecast

Multi-sensor observed rainfall analysis

The technique’s precipitation projections were quite a bit better than the Eta’s. If a storm changes little in character and the NHC track is fairly accurate, the TRaP rainfall should provide a pretty good estimate of the system’s rainfall potential at landfall.

TRaP assumes that the rainfall rates measured by the satellite sensors are accurate, that the precipitation rates and distribution of precipitation within and around the storm will remain the same with the same. It fails to take into account changes due to eyewall replacement, changes that are taking place in the environment around the storm that may impact it. For example, dry air being entrained into the storm or changes in the shear, both of which can impact the structure of the storm.

Roth (personal communication) has noted that there are a number of factors that modulate the distribution of rainfall associated with tropical cyclones. Storm size, morphology, track, interaction with synoptic and mesoscale features and even the time of day can impact the precipitation associated with the storm. He offered the figure at left to illustrate how complex the precipitation pattern can be. The heaviest rainfall remained on the east side of the storm as it moves across southern Georgia.

As it approached the Appalachains, strong flow into mountains produced very heavy rainfall along their east slopes and a minimum on the west side of the mountains (this minimum was very pronounced over Virginia). However, as the remains turned northeastward another maximum formed along a weak front associated with the entrance region of an upper level jet streak. Here, the maximum ended up to the west of the storm track.

As illustrated by the figure above, how to forecast the distribution of rainfall associated with a tropical system becomes more difficult once it moves inland and turns towards the north or northeast and gains latitude. The chances of entraining dry air into some quadrant of the storm increases as it moves away from the coast, which increases its chances of becoming asymmetric. Also, as a storm turns northward, it often starts to interact with a mid-level trough. When this occurs, the axis of heaviest precipitation then often falls to the west of the track of the surface center. The proximity to a mid-level trough and increase in latitude raise the chances that the storm might interact with a frontal boundary.

From Bosart and Dean 1991

Bosart and Dean (1991) have offered a schematic of the mid-tropospheric streamlines associated with possible tropical storm related flooding. The southern area is associated with the forcing of the storm itself, while the northern area is usually associated with moisture interacting with a frontal boundary. Strong frontogenesis is often occurring within this area. Where the heaviest rainfall will be located relative to the storm is strongly related to its structure.
The next few pages will attempt to explain how the storms structure can influence its area distribution of rainfall. Remember that convection and heavy rainfall related to a tropical system but not directly due to it can develop well in advance of the storm along a boundary, the entrance region of a jet streak or along the upwind side of a mountain range (Cote 2007).

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Precipitation primarily left of Track storms

A study by Atallah and his colleagues (2007) composited 14 cyclones that had the bulk of their rainfall occur to the west of the track of the cyclone circulation. Six of the storms made landfall along the Gulf of Mexico west of the Appalachians. Their composite argues strongly that storms that have most of their rainfall distributed on the west side of storm’s track interact with upper level trough and potential vorticity maximum in a synergistic way.

1000-500 hPa thickness (dashed black lines, contour interval=60 m), 1000-hPa geopotential height (solid black lines, contour interval=30 m) and 850-200 hPa wind shear (shaded, contoured every 5 m s-1, starting at 20 m s-1) , for left of track cases. Thick black line is the trough axis. (from Atallah et al, submitted to Mon. Wea. Rev. 2006)

Storms undergoing a transition to becoming extratropical typically have their heaviest rainfall to the west of the track. The 850-200 hPa shear suggests that there was a strong jet streak present to the north or northwest of the cyclone and implies that the storm was located near the right entrance region of a jet streak where frontogenesis is favored. The composite also suggests that the axis of the upper-level trough shifts from having a slight positive tilt or neutral one to having a negative tilt as the storm undergoes extratropical transition (ET).

Agnes illustrates one way in which a tropical storm or hurricane may interact with a trough. If the upper-level trough has a pronounced negative tilt, the hurricane circulation may phase with the trough and be absorbed by the deepening 500-hPa trough. When this occurs, the surface low may re-intensify through baroclinic processes. The surface low also may almost stall or loop as it is pulled into the 500-hPa circulation. Note how the gradient of the isotherms tightens between 0000 UTC and 1200 UTC at 500-hPa and the surface.

From Bosart and Dean, 1991

Strong frontogenesis was occurring with the heaviest rainfall located along and to the west of the surface trough associated with the front. The front on the north side of the storm often becomes stationary for a period. This adds to the systems potential to produce an extended period of heavy rainfall. The figure below shows the movement of the front. Note how slowly it moves over Pennsylvania.

From Bosart and Dean, 1991

Such strong phasing between a tropical system and a strengthening negative tilted mid-level trough to its west is unusual. The 200-hPa analyses above both have red ovals which roughly depict the area where the right rear quadrant of the north-south oriented jet streak is juxtaposed with the left exit region of another jet streak to the south. Uccellini and Kocin (1987) have documented that similar jet streak configurations are conducive to the deepening of east coast snowstorms.

850-700-hPa PV (shaded in cool colors every 0.15 PVU starting at .45 PVU where 1 PVU=10-6Kkg-1 m2 s-1) and 300-200-hPa PV (shaded in warm colors every 1 PVU starting at 2 PVU) and winds (barbs, knots convention). Thick dashed contours indicate areas of positive PV advection in the 300-200 hPa layer. The thick black line represents the axis of cross sections that except for at T+24 are not shown. The arrows define the axis of the upper trough. (from Atallah et al, submitted to Mon. Wea. Rev. 2006)

The potential vorticity (PV) structure for the LOT storms evolves very much like an east coast snowstorm. The upper level PV maximum approached a low level center with each center initially having their own area of PV advection. However, as the distance between them decreases, the advection areas merge together, generally to the north and northwest of the tropical cyclone center.

Kocin and Uccellini (2004) have noted a similar evolution of the PV fields during many East coast snowstorms and note the synergistic interactions between the upper level dynamical processes (the PV and jet streak evolution) and the low level diabatic processes in the development of east coast snowstorms. A similar interaction and feedback takes place in tropical system undergoing extratropical transition (ET).

Cross section of PVU (shaded, starting at .4 PVU as given by the color bar), isentropes (thin solid lines, contoured every 3oK), relative humidity (dashed line contoured every 10%) for line onT+24 shown above . The thick solid line indicates the position of the warm core associated with the tropical system. from Atallah et al, submitted to Mon. Wea. Rev. 2006

Hurricane Floyd is a classic case where ET is taking place, and not surprisingly, the bulk of its rainfall occurred to the left of the track.

Storms interacting with fronts tend to be some of the wetter storms and can produce heavy rain even as though the remains of the storm may be fairly fast moving. Neither the Kraft rule of thumb or precipitation forecast based on TRaP will work well with transitioning storms. The operational models sometimes handle the precipitation well but at other times do not do a good job. They often have problems in handling the latent heating and associated ridging (Atallah et al, 2006).

250-hPa winds and heights valid 0000 UTC 16 Sept 1999, the red dot indicates the positions of Floyd

250-hPa winds and heights valid 1200 UTC 16 Sept 1999, the red dot indicates the positions of Floyd

At 0000 UTC, as Floyd moved northward (the red dot on the figure shows its position), a strong positively tilting upper level trough was moving eastward in the vicinity of the Great Lakes and associated upper level jet streak was located to its east. Note that as the upper trough shifted eastward and the distance between the trough and Floyd decreased, the geopotential heights across eastern New England actually rose and the flow across the region became anticyclonic suggesting that the latent heating associated with Floyd was feeding back and was altering the pattern. The half wavelength between the base of the trough and this shortwave ridge was shortening which increased the upper level divergence. This is a fairly common characteristic of storms undergoing extratropical transition.
Another important thing to note about the pattern is the location of Floyd along the right entrance region of the upper level jet streak, one that has a significant along stream variation. The strong along stream variation suggests two things: 1) that the lower branch of the associated transverse circulation will be acting in concert with the low level circulation associated with Floyd to enhance any thermal gradient across the region, 2) will act to increase the upper level divergence across the region.

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925-hPa height and winds (black), temperature (blue lines, oK), valid 0000 UTC 16 Sept 1999

925-hPa height and winds (black), temperature (blue lines, oK), valid 1200 UTC 16 Sept 1999

A pool of cooler drier air is usually associated with mid- and upper level troughs, especially the stronger ones. Note the almost cold air damming look to the isotherms over Virginia and the Carolinas. Clearly a front was present and showed up on the 925-hPa 0000 UTC Sept. 16 analysis. As the Floyd interacted with the trough and moved north, the thermal gradient tightened to its north and east. Strong frontogenesis and isentropic lift was taking place along the thermal boundary

The 4-panel GOES-8 loop shows water vapor imagery (top right) and three goes sounding channels. The two bottom panels of the loop are sensitive to mid- (left) and low-level (right) moisture. Note how the clouds increase to the west and northwest of the storm and that the heavy rainfall across North Carolina begins well in advance of Floyd.

Loop courtesy of Rod Scofield

925-hPa winds, frontogenesis (orange contours) and 700-hPa omega (upward motion is red and white area) valid 1200 UTC 16 Sept

Composite radar reflectivity valid 1200 UTC 16 Sept. From Atallah and Bosart 2003.

Image from Kelsch lecture

The radar echoes and region of rainfall elongated along the frontal band. Such an elongation of the precipitation pattern is typical when a front is present. The axis of precipitation of left of track (LOT) storms often is aligned parallel to the movement of the storm which increases the period of heavy rainfall. Bosart (personal commuciation) believes that the LOT cases on average produce a larger total volume of rainfall than right of track (ROT) cases. The idea is supported by two six-case composites (shown below) that one of his students, Josh Darr, assembled for his M.S. thesis at SUNY Albany.

Strong extra-transition 6 case composite cases, precipitation contoured in mm.

Composite of non-transition cases, precipitation contoured in mm.

Precipitation primarily right of the track (ROT).storms.

The composites of storms with the bulk of their precipitation to the right of the track from the Atallah et al. study are shown below. There are several notable differences in the ROT composite from the pattern typically associated with LOT storms. The axis of the 850-200-hPa shear is typically oriented more westerly than the more north-south oriented axis associated with the approaching trough during LOT cases. Also, the location of the upper level jet streak is much farther removed for ROT cases than for LOT cases. The baroclinic zone as shown by the leading edge of the 1000-500 hPa thickness is also much farther removed from the storm. Also, implied 1000-hPa circulation is much weaker, especially on the northern and west side of the system where there is a much weaker gradient implying that the easterly low level winds are quite a big weaker for ROT storms than LOT. storms.

1000-500 hPa thickness (dashed black lines, contour interval=60 m), 1000-hPa geopotential height (solid black lines, contour interval=30 m) and 850-200 hPa wind shear (shaded, contoured every 5 m s-1, starting at 20 m s-1) , for right of track cases. Thick black line is the trough axis. (from Atallah et al, submitted to Mon. Wea. Rev. 2006)

The more zonal look to the pattern helps keep any northern stream PV anomaly well removed from the low level PV center associated with the tropical system. The right hand composites lack the negative tilt.

850-700-hPa PV (shaded in cool colors every 0.15 PVU starting at .45 PVU where 1 PVU=10-6Kkg-1 m2 s-1) and 300-200-hPa PV (shaded in warm colors every 1 PVU starting at 2 PVU) and winds (barbs, knots convention). Thick dashed contours indicate areas of positive PV advection in the 300-200 hPa layer. The thick black line represents the axis of cross sections that except for at T+24 are not shown. (from Atallah et al, submitted to Mon. Wea. Rev. 2006)

Don’t be fooled by the composite’s location of either LOT or ROT cases. 11 of the 16 cases for the ROT composite made landfall over the Gulf of Mexico and 5 made landfall along the east coast, usually near the Georgia-South Carolina border. 8 of the storms transitioned from a ROT storm to a LOT storm.

The track and speed of Hurricane Bertha were very similar to that of Floyd making it an interesting storm to study since it rainfall amounts and rainfall distribution were different than occurred during Bertha. The axis of heaviest rainfall during Bertha was to the right of the track across the southern Middle Atlantic States but shifted to the west side from New Jersey northward.

Bertha and Floyd were both Category 2 hurricanes at landfall. However, the circulation associated with Bertha was smaller. An even bigger difference between the systems was the thermal structure around the storm.

250-hPa winds and heights valid 0000 UTC 13 July 1996, the red dot indicates the positions of Bertha

925-hPa height and winds (black), temperature (blue lines, oK), valid 0000 UTC 13 July 1999

During Bertha the 292oK isotherm dipped southward around Bertha which helped make the isotherms almost parallel to the winds to the north and northeast of Bertha (see figure above). At 00 UTC 13 July, no frontal boundary was present near the storm. The lack of warm advection suggested that the heavy rainfall would be confined to near the center or to feeder bands on its east side. Also, the jet was well removed from the storm with generally westerly flow across the storm, the latter suggests that the storm may have been affected by shear. Also, the orientation of the upper level jet was more west to east than north to south when compared to Floyd. The implied shear would tend to keep the convection on the east side of the system and at that time, the heaviest rainfall was located to the east of storm. The structure of the storm is an important component in predicting the rainfall with a storm.

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Rainfall maximum to right of track composite (ROT)

Rainfall maximum to left of track composite(LOT)

Cross section of PVU (shaded, starting at .4 PVU as given by the color bar), isentropes (thin solid lines, contoured every 3oK), relative humidity (dashed line contoured every 10%) for line onT+24 shown above . The thick solid line indicates the position of the warm core associated with the tropical system.

The Atallah et al. composite cross section for ROT versus LOT shows several marked differences. The area of high relative humidity on the ROT cases is deeper than on the LOT case composite suggesting upright convection might be the dominant type of precipitation. The warm core (thick black line) on the ROT composite is not as well defined at lower layers as on the LOT composite and there is actually a weak cold pool at the lowest layers. The authors caution that the cold pool may be a function of the lack of resolution of the reanalysis data. However, the same cold pool was present in the NARR data during the period when the bulk of the rainfall was located to the right of Bertha’s track so the cold pool is probably present during at least a portion of the ROT cases.

Back to Bertha, the thermal structure of the system slowly changed during the next 15 hours and by 15 UTC July 13, a thermal boundary developed which promoted low level warm advection (the area enclosed by the thick back line). By then, the 925-hPa absolute vorticity (red lines) looks like it is being captured by a weak negative tilted trough (axis is defined by magenta line). The upper trough and low level thermal boundary are both much weaker than during Floyd.

925-hPa height and winds (black), temperature (blue lines, oK), valid 1500 UTC 13 July 1999

250-hPa winds and heights valid 1500 UTC 13 July 1999

In summary, the structure of a storm seems to determine whether the heaviest rainfall will occur to the left or right of the track. Storms interacting with upper level troughs, jet streaks and baroclinic zones tend to have their precipitation transition to the left side of the track. There is a deep frontal boundary clearly defined in the isentrope field on the LOT composite. No frontal structure is evident on the ROT composite.

Another possible useful tool is the cyclone phase space diagram developed by Hart (2003). The diagram provides a tool to objectively diagnose whether the system is warm or cold core and whether the thermal structure will promote asymmetry. A sample of the product is shown below. The ordinate measures the symmetry of the system while the abscissa measures the thermal structure and whether the system has warm or cold core structure. The location of the storm at various times is given at the upper right hand corner of the figure.

From Evans and Hart 2003

Note when Floyd was labeled with the number 16 it crossed the line between being a symmetric and asymmetric warm core system and that during the next 12 hours the storm became increasingly asymmetric. This was also about the time when the bulk of the precipitation transitioned to the west side of the track. The phase space diagrams provide a quick way to investigate the structure of a storm using model forecasts of the storm. When the diagram starts indicating that a storm is starting to become asymmetric, look closely at model diagnostic fields to see whether the storm is starting to transition. Also look for fronts and jet streaks to that might help focus the precipitation to the left of the storms track.

Click here to look at the The Florida State University website with phase space diagrams

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East Pacific storms

Effects over Southwest U.S.

Hurricanes and tropical systems that move into Mexico or track northeastward off the coast of Baja can produce wide ranges a precipitation. Some storms that move into Mexico are very wet across the southwest while others produce only spotty rainfall. The difference appears to be related to whether the basic synoptic pattern is favorable for pulling moisture northward into the southwest. Note how similar the 500-hPa pattern prior to Kathleen to the 500-hPa patterns that are typically associated with monsoonal surges into the southwest. A 500-hPa low off the coast and a strong ridge to the east. Such a configuration produces deep southerly flow. Another common factor is surface or 1000-hPa trough that extends northward across California. The rapid movement of the storm and abnormally warm ocean temperatures associated with an El Nino allowed Kathleen maintain its intensity more than

500-hPa geopotential heights and normalized departures from normal (color shaded) valid 00 UTC 9 Sept 1976. The red dot is the position of Kathleen.

most storms with a similar track. Kathleen is one of five Pacific storms that have maintained strength enough to produce gale force winds over California. Almost 12 inches of rain fell over the mountains of southern California. Over half the town of Ocotillo was destroyed by flash flooding. Kathleen was one of the few southwest storms that produced significant flooding without interacting with an upper level jet streak. However, some storms can produce significant rainfall without being in the classic right rear entrance region of a jet streak (For example Octave).

The coverage and amount of precipitation seems to be modulated by the speed of movement of the trough and by the strength of the southerly or southwesterly flow at 700 hPa, especially for those storms that dissipate and then indirectly produce sustained heavy rainfall. Tropical Storm Octave is a case in point. Octave moved northward and became a tropical storm on October 01, 1983. At 00 UTC 01 Oct., Octave was still located southwest of Baja, south of the bottom end of the 250-hPa analysis shown below.

250-hPa geopotential heights, winds and isotachs v.t. 00 UTC 30 Sept 1983

250-hPa geopotential heights, winds and isotachs v.t. 00 UTC 01 Oct 1983

The strong jet streak on the west side of the trough and the lack of any upstream kicker suggest that the upper trough will be very slow moving and that the upper low may actually dig southward.

700-hPa geopotential heights, winds and isotachs v.t. 12 UTC 30 Sept 1983

700-hPa geopotential heights, winds and isotachs v.t. 00 UTC 02 Sept 1983. The red dot is the approximate location of Octave

A wind max at 700-hPa remained aimed at Arizona for 2 days with both analyses above suggesting speed convergence at 700-hPa. The winds at 850 were also from the south but were quite a bit weaker than at 700-hPa. Such deep southerly flow pulled moisture northward with an area of 1.40 inch or greater PWs extending into southern Arizona. The pattern was such that even without Octave, there probably would have been pockets of heavy rainfall across Arizona. Octave’s moisture served to enhance the rainfall.

The majority of east pacific storms that produce heavy rainfall across the southwest are associated with upper troughs and most are associated with an entrance region of an upper level jet maximum that helps to enhance the upper level divergence across the region. Norma (1970), Norman (1978), Lester (1992) and Olivia (1982) all were storms associated with and upper-level trough and the right rear quadrant of an upper level jet streak. Lester is a case in point.

250 hPa geopotential heights, winds and isotachs v.t. 00 UTC 23 August 1992

700 hPa geopotential heights, winds and isotachs v.t. 00 UTC 23 August 1992. Red dot is NHC position of Lester.

Precipitable water (mm) v.t. 00 UTC 23 August 1992

MSL pressures lowered across Nevada and southern California in response to the upper trough digging towards Nevada and by 00 UTC 23 August, a surface low was located over southern Nevada. The low level southerly winds developed across western Arizona (not shown) ahead of Lester which helped pull a surge of anomalously high PW moved into Arizona.

Heavy rainfall occurred right as the storm made landfall over Baja. The location of the right entrance region over Arizona and the along stream variation of the winds within the jet, argue that strong upper level divergence should be going on across Arizona.

250 hPa geopotential heights, winds and isotachs v.t. 12 UTC 23 August 1992

700 hPa geopotential heights, winds and isotachs v.t. 12 UTC 23 August 1992. Red dot is NHC position of Lester.

24-hour precipitation verifying at 12 UTC 23 Aug. 1992

Lester storm total precipitation

250 hPa geopotential heights, winds and isotachs v.t. 12 UTC 24 August 1992

700 hPa geopotential heights, winds and isotachs v.t. 12 UTC 24 August 1992. Dashed red line is track of the remains of Lester.

Despite a surge of strong southerly winds across New Mexico in advance of Lester, the storm produced most of its rainfall in the southwest across Arizona and southwestern Colorado. Dickinson and his colleagues (2002) have noted that Lester underwent extratropical transition. Both Lester and Norman (1978) produced the bulk of their rainfall to the west suggesting that some eastern Pacific storms may undergo extratropical transition and have their precipitation shield shift to the northwestern side of the storm.

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1000-hPa heights and Precipitable water anomalies (color shaded) valid 00 UTC 9 Sept 1976

Portions of southeast Arizona were inundated and record peak flows were recorded on several large rivers. Clifton and Marana just north of Tuscon were devastated. Property damage was estimated to be $500,000. A report written in 1985 indicated that the flooding associated with Octave and the strong upper trough that it interacted with was the most damaging flood ever recorded in Arizona (Smith et al, 1985).

Let’s explore another case, one that produced much less rain than Lester even though the storms hit Mexico at about the same strength. Hurricane Isis (the red dot) lifts slowly northeast between 12 UTC 2 Sept and 00 UTC 3 Sept, the next 12 hours the storm weakens as it lifts northward. A trough/upper low that is developing off the California coast help pull Isis slowly north. However, the associated jet streak is very weak and Isis is located along the right front quadrant of this weak jet streak. More importantly during most of 24 hour period ending at 12 UTC 4 Sept., there is considerable shear with the easterly low level flow across Arizona and Mexico. Winds at 700 and 850 hPa

00 UTC 3 Sept 1998 250-hPa geopotential heights and winds, red dot is Hurricane Isis, the magenta dot is the 12 UTC 2 Sept position.

00 UTC 3 Sept 1998 700-hPa geopotential heights and winds, red dot is Hurricane Isis, the magenta do it the 12 UTC 2 Sept position.

00 UTC 4 Sept 1998 700-hPa geopotential heights and winds, red dot is Hurricane Isis at 12 UTC 3 Sept position.

00 UTC 4 Sept 1998 250-hPa geopotential heights and winds .

24-hr hPC precipitation analysis valid 12 UTC 4 Sept. Remember only amounts greater than 0.50 inches are analyzed

are easterly across Arizona and New Mexico at 00 UTC 4 Sept. Such flow is not as conducive to rainfall as deep southerly flow.

00 UTC 5 Sept 1998 700-hPa geopotential heights and winds.

00 UTC 5 Sept 1998 250-hPa geopotential heights and winds .

24-hr hPC precipitation analysis valid 12 UTC 4 Sept.

By 00 UTC 5 Sept., the ridge had shifted eastward and deep southerly flow was present across Arizona and the area of moderate to heavy rainfall consequently expanded. However, the rainfall associated with Isis was not nearly as impressive as with Lester, a storm that interacted with the right rear quadrant of an upper level jet streak and a storm that showed signs of extratropical transition. The contrast between the rainfall associated with Isis compared to other east Pacific storms that were discussed earlier in the section argues that deep southerly flow and a favorable interaction with a jet streak both act to increase the potential for heavy rains across the southwest.

Remnants of hurricanes in the East Pacific
can help produce heavy rainfall across Oklahoma and Texas.
Two examples are presented below, Norma and Tico. These two example illustrate that convection indirectly associated with Pacific hurricanes can develop prior to a storm’s landfall in Mexico.

Infrared image v.t. 0000 UTC 12 Oct

At 0000 UTC Norma was still located off the coast of Mexico, convection had already developed across parts of northern Texas and Oklahoma along and north of an old frontal boundary. This boundary clearly shows up on the 1000-hPa temperature field below. The storm is located towards the entrance region of an upper level jet streak that starts near baja and extends northeastward into Texas.

250-hPa height, and winds valid 0000 UTC 12 Oct. 1981

850-hPa height, and winds and 1000-hPa temperature (green, interval=2oK) valid 0000 UTC 12 Oct. 1981

250-hPa height, and winds valid 0000 UTC 13 Oct. 1981

850-hPa height, and winds and 1000-hPa temperature (green, interval=2oK) valid 0600 UTC 13 Oct. 1981

The upper-level jet streak strengthened during the next 24 hours probably in response to the diabatic heating associated with the convection. A weak upper level trough was also embedded in the flow. It lifted northward during the night and worked in concert with the jet streak to increase the upper level divergence across the region which in turn probably acted to increase low level jet across Texas.

The implied relationship between the track and precipitation may be misleading as the track used was not a low level center but was a mid-level vorticity center. No true low level circulation made it across the mountains.

A pre-existing front was also present prior to Tico and it too helped focus heavy rainfall across northern Texas and Oklahoma (see NARR analysis below).

850-hPa geopotential height, winds and 1000-hPa temperatures valid 0000 UTC 18 Oct. 1983

250-hPa height, and winds valid 0000 UTC 18 Oct. 1983

250-hPa height, and winds valid 1200 UTC 18 Oct. 1983

250-hPa height, and winds valid 0000 UTC 19 Oct. 1983

The entrance region of an upper level jet streak appear to play a role in enhancing the dynamics during each case. Note how the jet streak strengthens as a northern stream trough presses eastward and helps tighten the gradient.

850-hPa geopotential height, winds and 1000-hPa temperatures valid 1200 UTC 18 Oct. 1983

850-hPa geopotential height, winds and 1000-hPa temperatures valid 0000 UTC 19 Oct. 1983

250-hPa height, and winds valid 1200 UTC 19 Oct. 1983

Tico actually had a remnant low move across Mexico into Texas. However, the bulk of the rainfall fell along and north of the front across the region. When pondering whether heavy rainfall will develop in response to an east Pacific tropical system, look to a preexisting front and for some mechanism to strengthen the upper level jet streak that is helping enhance the low level convergence and upper level divergence in the vicinity of the front.

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850-hPa geopotential height, winds and 1000-hPa temperatures valid 0000 UTC 19 Oct. 1983

Hurricanes and the Appalachians

The mountains in the East can play a significant role in focusing lifting along the windward slopes of the rising terrain. Storms crossing the mountains or interacting with them that come from west or east of the mountains can produce heavy rainfall along the east slopes of the mountains.
Camille is the poster child of storms that produce catastrophic rainfall when they interact with the Appalachians. More than 150 people died as a result of flash flooding and mud slides caused by the heavy rainfall associated with Camille making it the worst natural disaster in the history of Virginia (Schwartz 1970). Over 27 inches of rainfall was reported during Camille and much of it fell in a 6 hour period. Remnants of Camille definitely helped focus rainfall but terrain was not the only factor that led to the heavy rainfall.

From Schwartz 1970

Simplified terrain, warm color are higher terrain

From Schwartz 1970

Mean Sea Level pressure chart for the period 2200 to 0300 EST on August 19-20, 1969. The dots are hourly positions of the low starting at 7PM EST.

The maximum in rainfall occurred north of the track of the surface low in a region where low level southeasterly, and then easterly, flow was intercepted by the mountains. However, when you compare the orientation of the mountains to the orientation of the axis of heavy rainfall, there was a difference. Other focusing mechanisms must have been present to account for the differences in the orientation. One of the difficulties predicting rainfall near the mountains in the Mid Atlantic Region and northeast is that there also is often a front, jet streak or upper level trough that will also be contributing to the lifting and ultimate rainfall pattern.

A closer look at Camille helps explain why the precipitation axis was oriented east-west while the terrain had a more southwest to northeast orientation. At 7:00 PM EST the remnant circulation was located over eastern Kentucky with the low being well defined on the satellite imagery (not shown). The remnant low tracked almost due east across Virginia with the bulk of its precipitation falling north of the low. The low was forced eastward under the trough that was dominating the Northeast and was forced eastward along the right entrance region of an upper level jet streak (see below)

200 hPa analysis valid at 12 UTC 20 August 1969

From Chien and Smith 1977

Sea level pressure (contour interval=4 mb) and surface temperature (contour interval=5oC) valid 00 UTC 20 Aug 1969

Sea level pressure (contour interval=4 mb) and surface temperature (contour interval=5oC) valid 12 UTC 20 Aug 1969

As the low shifted eastward, there was a significant tightening of the surface temperature gradient across PA to the east of the surface low. In all likelihood, the jet streak and the remnant low were acting synergistically to produce frontogenesis. Chen and Dell’Osso (1987) have noted that diabatic heating associated with heavy rainfall has an impact on strength of the upper level jet and its associated transverse circulations (see below). In their study, diabatic heating increased the strength of the ageostrophic circulation associated with the upper branch of the jet even as the latent heating increased the strength of the winds within the jet streak itself. Latent heat also increases the low level convergence and frontogenetical forcing along any thermal boundary. The tightening of the low level thermal gradient across Virginia suggests this synergistic interaction took place which probably contributed to re-energizing the storm. The jet streak and terrain both probably enhanced the lifting. Atallah and Bosart (2003) have noted that the effects of latent heat on the upper level structure in the vicinity of a tropical system may at times have significant errors.

From Chen and Dell’Osso, 1987

The jet streak and terrain may explain a large portion of why rainfall amounts were much heavier east of the mountains than west of them. However, the differences in precipitation may have also resulted from the typically nocturnal fluctuation that often occurs with the remains of old storms. Convection associated with tropical systems typically flares up during the night. The important thing to remember is not to give up on a storm’s potential when it still has a well defined circulation, as it crosses the mountains, precipitation may enhance significantly.

From Chien and Smith 1977

Hurricane Danny is an interesting storm to contrast with Camille. It too came ashore in the central Gulf of Mexico and then tracked around the periphery of the subtropical ridge and produced heavy rainfall on the east sides of the Appalachians.

Like Camille, Danny produced heavier rainfall east of the mountains than west. Where Danny differed was in the strength of the upper level jet across the Mid Atlantic region. This jet streak actually weakened as Danny approached (see below). Danny also had a secondary precipitation maximum along the east slopes of the Appalachians south of the track of the storm where there was strong speed convergence present at 850-hPa (see figure below).

850 hPa and 1000-hPa temperatures v.t. 12 UTC 17 Aug 1985

250-hPa heights and winds v.t. 12 UTC Aug 1985

850 hPa and 1000-hPa temperatures v.t. 00 UTC 18 Aug 1985

250-hPa heights and winds v.t. 00 UTC 18 Aug 1985

850 hPa and 1000-hPa temperatures v.t. 12 UTC 18 Aug 1985

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200 hPa analysis valid at 00 UTC 20 August 1969.

Danny produced a rainfall maximum along the east slopes of the mountains just north of the storm’s track. It also produced a stripe of heavy rainfall with another maximum near Richmond in an area where the gradient in the isotherms increased. Even without a strong jet streak, if a storm holds together and has a decent 850-hPa circulation while crossing the mountains, it still may produce heavy rainfall.

More on tropical systems and the Appalachians, an exercise

Tropical storms and hurricanes also cross the mountains when moving from west to east. Most of the time there is an upper low to the west or southwest of these type of storms which allows deep easterly flow to develop to the north and northwest of the storm. This storm is forecast to make landfall in the Carolinas and then is forecast to track northwestward. Below is an ETA/NAM 500-hPa geopotential height forecast of the pattern associated with the approaching hurricane. The hurricane is just showing up east of Florida at 24 hours.

00-hr model 500-h forecast v.t. 00 UTC

12-hr model 500-h forecast v.t. 12 UTC

Forecast hurricane location

24-hr model 500-h forecast v.t. 00 UTC

36-hr model 500-h forecast v.t. UTC

The airmass is predicted to be unstable across the region at the time of maximum heating. The PW is around 1.80 across central North Carolina and Virginia

18 hr CAPE v.t. 18 UTC

24 hr CAPE v.t. 00 UTC

During the 12-h to 36-h period on the maps above, do you think forecasters in North Carolina and Virginia should be worrying about heavy rainfall during 24-h period ending at 36-h? Click on the answer you like best.

A)

The subsidence ahead of the hurricane and the lack of a front should keep a cap in place across the Carolinas and Virginia limiting convection and rainfall

B)

The weak vorticity maximum that swings from South Carolina to West Virginia suggests there could be scattered showers with locally heavy rainfall so I’d monitor the situation in case a special weather statement is needed. I’d also look at the situation in more detail.

C)

Slow moving upper lows are dangerous but it looks like the upper low is probably too far from North Carolina and Virginia during the 24 hour period to be a major concern for those two states. However there could be some scattered showers with locally heavy rainfall.

Let’s look at the case in a little more detail. The position of the upper low is one that is favorable for producing deep easterly flow into the mountains. 30 kt winds are predicted at 850-hPa across portions of Virginia with a strong easterly component.

24-hr model 850 wind and PW (dotted line. inches) forecast v.t. 00 UTC

The faint, dotted line that extends across central North Carolina into Virginia near Roanoke is the 1.80” PW contour. The highest PW and CAPE is forecast to be across eastern and central Virginia (see below). The ETA and GFS forecasts for the period are shown below. The models differed on their forecasts of the rainfall potential and on where the heaviest rainfall would occur.

ETA QPF for period

GFS QPF for period

The GFS suggested the maximum might be along the east slopes of the Appalachians. What about the potential for upslope? Whenever deep easterly flow is present, there is potential for the upslope flow to enhance the vertical motion. One tool to look at is the upward motion due to the vertical transport of motion due to terrain that is available on the NMAP workstations. (see figures below). In this case, the eta was predicting an extended period with strong orographic lifting across portions of western North Carolina and Virginia.

18 hr v.t. 18 UTC boundary layer winds, terrain (thin lines) and area of strong orographic omega (dark line)

24 hr v.t. 00 UTC boundary layer winds, terrain (thin lines) and area of strong orographic omega (dark line)

30 hr v.t. 06 UTC boundary layer winds, terrain (thin lines) and area of strong orographic omega (dark line)

The combination of having a 500-hPa upper low to the southwest of the mountains of western North Carolina, a very moist airmass with positive CAPES and anomalously strong 850-hPa winds directed into the mountains should provide a signal that there is potential for convective rainfall with the potential for the convection to remain focused along the east slopes of the mountains for an extended period if the eta model orographic omega was in the ball park. The predicted high PW and high relative humidity also suggested the potential for fairly high rainfall rates. Evaluation of the pattern suggests that the GFS had the right idea about the location of the maximum but that there is potential for heavier rainfall than it is forecasting.

18 hr 700-hPa height, vertical motion and relative humidity (dark shading greater than 60%, light shading greater than 80%.

24 hr 700-hPa height, vertical motion and relative humidity (dark shading greater than 60%, light shading greater than 80%.

36 hr 700-hPa height, vertical motion and relative humidity (dark shading greater than 60%, light shading greater than 80%.

Look again at the model QPF and think how you would modify it. Choose the answer that best describe the changes to the models QPF and the actions you would make on shift. (Click on an answer and you’ll be linked to the verifying rainfall analysis)

A)

I would forecast a rainfall maximum near the axis of where the GFS predicted its maximum but would beef up its amounts near the mountains and would keep any two inch rainfall away from the coast. I would also issue a slight risk of excessive rainfall near the mountains.

B)

I would forecast two rainfall maximum of almost equal magnitude one near where the eta has the eastern portion of its two inch contour and another where the GFS has its maximum and would issue a slight risk of exceeding flash flood guidance values across both areas.

C)

I would forecast two rainfall maximum, a heavier one near where the GFS predicted its maximum and a second smaller, lighter maximum near the coast with a relative minimum in between. I would also issue a slight risk of exceeding flash flood guidance near the mountains.

The hurricane off the southeast coast was Fran. In some respects it and Hugo were similar. Both produced flash flooding in the Appalachians well in advance of landfall. Both had upper lows located to the west or southwest as they approached the coast. However, the scale of the heavy rainfall was much smaller with Hugo than Fran. It also produced quite a bit less rainfall than Fran even though both hurricanes were fairly large ones. A few differences between these two storms will be discussed on the next page.

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Hurricane’s Hugo and Fran, a quick comparison

Though Hugo and Fran were fairly similar in size, they produced markedly dissimilar precipitation footprints in part because of the differences in how fast they moved. Fran took two and a half days to track from the Carolina coast to a position near Buffalo, New York. By contrast, Hugo took less than a day to move almost the same distance (see tracks below). This certainly helps explains the differences of rainfall across the Piedmont of the Carolinas during the two storms.

Hurricane Hugo

Hurricane Fran

2201 UTC 21 Sept 1989

1632 UTC 5 Sept 1996

Hugo’s track.

Fran’s track

Hugo storm total rainfall

Fran’s storm total rainfall

The structure of the storms as they interacted with the mountains is another distinction between the two storms. 500-hPa and 850-hPa geopotential heights are shown for Fran below (the top three panels). Fran still maintained a circulation at 850-hPa that was conducive to maintaining orographic lifting across the mountains of southern Pennsylvania on 00 UTC 7 September.

The NARR reanalysis data below is from The Pennsylvania State University

500-hPa height and vorticity analysis valid 12 UTC 6 Sept. 1996

500-hPa height and vorticity analysis valid 00 UTC 7 Sept. 1996

850-hPa geopotental height and RH analysis valid 00 UTC 7 Sept. 1996

500-hPa analysis valid 12 UTC
22 Sept. 1999

500-hPa analysis valid 00 UTC
23 Sept. 1999

850-hPA geopotential height and RH analysis valid 00 UTC 23 Sept. 1999

Hugo’s 500– and 850-hPa patterns are shown on the three bottom panels. Hugo was actually stronger than Fran at landfall but quickly lost its closed circulation at 500– and 850-hPa as it interacted with the strong trough digging into the Great Lakes region. The lack of strong easterly flow by 00 UTC 23 Sept. limited the vertical motion as the flow quickly shifted to having a westerly component across the mountains which causes downsloping flow east of the mountains. Hugo is a cautionary tale showing that a storms interaction with synoptic scale trough does not always result in a large areas of heavy rainfall associated with isentropic lift and frontogenesis. In this case the shearing of the circulation center, the rapid movement, and the lack of a strong frontal boundary nearby kept Hugo from producing widespread heavy rainfall.

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Parker, M. D. and R. H. Johnson, 2000: Organizational modes of midlatitude mesoscale convective systems. Mon. Wea. Rev., 128, 3413-3436.
Rappaport, E.N., M. Fuchs, and M. Lorentson, 1999: The threat to life in inland areas of the United States from Atlantic tropical cyclones. Preprints, 23rd Conf. on Hurricanes and Tropical Meteorology, Dallas, TX, Amer. Meteor. Soc., 339-342.
Reynolds, D. W., 1996: Similarities of three major precipitation events in California during 1995. Preprints, 15th Conference Weather Analysis and Forecasting, August 19-23, Norfolk, VA, J127-J130.
Schwartz, F. K., 1970: The Unprecedented rains in Virginia associated with the remnants of Hurricane Camille. Monthly Weather Review: Vol. 98, No. 11, pp. 851–859.
Schumacher, R. S. and R. H. Johnson, 2005: Organization and environmental properties of extreme-rain-producing mesoscale convective systems. Mon. Wea. Rev., 133, 961-976.
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Weakening tropical storms and depression

Tropical storms and depressions often can produce heavy rainfall, sometimes days after landfall and at times with a distinct nocturnal maximum. Grice and Ward (1983) conducted a 10-year study of tropical depressions and hurricanes that moved inland into Texas in hopes of separating the “wet ones” from the relatively dry ones. They found that 5 of the heaviest 24-hour precipitation totals occurred one or more days after landfall. They found for the “wet” and “dry” cases, the lower and mid-tropospheric were often surprisingly similar. The biggest difference appeared to be in the upper levels and that the wet cases were usually associated with upper-level tropospheric waves and or jet streaks.

From Grice and Ward, 1983

A case in point was Tropical Storm Amelia. Like many weak tropical storms, the low level circulation associated with Amelia weakened once the storm came inland. However, at least a portion of Amelia’s 500-hPa circulation held together an contributed to heavy rainfall on two consecutive nights in Texas. Am unofficial report of over 45 inches of rain was observed in the Texas hill country over a 2 day. period.

The surface circulation weakens and only a weak cyclonic shear axis was present by 0000 UTC 1 August. Such weakening is not that unusual for weak tropical systems once they move inland, especially into Texas.

Courtesy of Dave Roth

IR image valid 0000 UTC 1 August 1978 (Caracena and Fritsch 1983)

The convection associated with the remains of the storm also weakened. In such cases, Caracena and Maddox have cautioned that forecasters should.
Anticipate the dissipation of thermally induced showers as boundary layer temperatures decrease.
Anticipate the development of deep convection at night somewhere near the core (within about 200 km radius of a weakening tropicsl storm and or an associated shortwave vorticity maximum

3) Look for possible mesoscale features that might help focus convection (ie. Terrain, outflow boundaries, low level jet or area of low level cyclonic shear).

4) Locate the most unstable are flowing into the core region. The horizontal distribution of unstable versus stable air is a key in determining whether propagation effects might lead to a quasi-stationary convective system.

IR image valid 1230 UTC 1 August 1978 (Caracena and Fritsch 1983)

Terrain (warmer colors are higher terrain)

Note that by 1230 UTC 1 August that thunderstorms had developed over the Hill Country. By that time, the surface circulation had died. However, there was a 500-hPa low to the southwest that helped supply deep upslope flow into the escarpment (see below). Low-level southeasterly flow was also present. When a storm is dying start thinking of it as a potential mesoscale convective system producer.

500-hPa analysis valid 1200 UTC 1 August 1978

200-hPa-hPa winds and divergence for 0000 UTC 2 August 1978. The trough axis is shown by the dashed line (Caracena and Fritsch 1983).

The 500-hPa low from the remains of Amelia merged with the low over Mexico by 0000Z 2 August (not shown). Grice and Ward noted that the proximity to an upper level trough seemed to be favorable for producing wet storms. Note the approach of a trough at 200-hPa and that the convection developed in the region where there was strong divergence anticyclonic curvature in the 200-hPa wind field.

IR image valid 0000 UTC 2 August 1978 (Caracena and Fritsch 1983)

IR image valid 1200 UTC 2 August 1978 (Caracena and Fritsch 1983)

During the night the convection contracted towards the core of the system. The cloud top temperatures also warmed which can be misleading as extreme rainfall rates still can occur when the bases of the clouds lower and mean relatively humidity becomes high.

The moisture associated with Amelia produced another MCS the next day over northern Texas. A frontal boundary that dropped south into northern Texas was the focus of this latter MCS. Again, look for mesoscale forcing when dealing with the remains of a tropical system.

The remains of Elena

Sometimes the remains of a tropical system will act very similar to a mesoscale convective vortex (see http://www.hpc.ncep.noaa.gov/research/mcs_web_test_test_files/Page1207.htm for a discussion of MCVs). These systems can be quite wet even after the surface low has weakened. They are more liable to act as MCVs when the mid– and upper level shear is weak. Note that the 500-hPa circulation remains removed from the westerlies and that no surface low is evident but that tropical moisture has spread northward along the western side of the ridge over the southeast.

500-hPa height and vorticity analysis v.t. 1800 UTC 3 Sept. 1985.

PMSL (contour interval=2 hPa) and precipitable water ( mm) analysis v.t. 1800 UTC 3 Sept. 1985.

1800 UTC 3 Sept. 1985 GOES vis imagery.

1800 UTC 3 Sept. 1985 GOES IR imagery.

A distinct swirl and apparent circulation was present on the satellite imagery but the convection was rather spotty and found around the periphery of the system rather than near its center. This is rather typical of a mesoscale convective vortex. Almost no convection was present by 0400 UTC 4 Sept.

0400 UTC 4 Sept. 1985 GOES IR imagery

If you treat the remains as an MCV it becomes easier to decide whether to another round of heavy rainfall might be expected. The first step to take is to look at the low-level wind field to determine where the source region for any subsequent convection might be. During the day, MCVs often have convection form along the periphery but at night convection often develops directly below the MCV center (Trier et al 2000).

At 0400 UTC there was only a few weak cells associated with the remains of the storm.

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Elena Continued

0300 UTC 4 Sept 1985 500-h and vorticity

0300 UTC 4 Sept 1985 850-hPa wind and 2-m dewpoints

0300 UTC 4 Sept 1985 850-hPa wind and CAPE

Despite the lack of convection there was reason to believe that a new outbreak might occur. In addition to having a distinct circulation apparent on the satellite imagery, there also was a fairly strong 500-hPa vorticity center apparent at 0300 UTC. A strong vorticity center even without a cold pool near the surface can produce convection.

In this case there is also a cold pool associated with the remains of convection. The most unstable air is located to the southwest of this cold pool which acts to promote new cell growth on the west side of the initial convection. The basic mesoscale pattern is very similar to the one offered by Fritsch et al. in the schematic below.

From Fritsch et al. 1994

Conceptual model of MCV that is moving slow enough that air is overtaking the system due to a low-level jet to the west or southwest. The dashed lines are potential temperature, solid lines are potential vorticity. Note that where there is rain cooled air the isentropes bulge upward and where there is warming due to latent heat the potential temperature lines bulge downward. The positive PV anomaly is located where the potential temperature gradient is greatest between the two.

Isentropic lift should occur where the low level jet is forced to ascend along the isentropes associated with the cold pool. Complicating the picture is that forcing for vertical motion can also be produced by the interaction of the cyclonic vortex and PV anomaly with the ambient shear (Raymond 1992, Trier et al. 2000) and lifting can happen without a pronounced cold pool.
Schumacher and Johnson (2005) have suggested that backbuilding convection can occur from elevated pockets of convergence within a conditionally unstable layer with sufficient relative humidity to allow a moist absolutely unstable layer (MAUL) to develop. Ascent initially is produced by differential vorticity advection or vorticity advection by the thermal wind. In early stages of development ascent is produced on the downshear (downstream) side of the PV anomaly. However, with time the vortex becomes tilted because of the vertical shear. Because of the tilt that occurs due to shear, the maximum vertical displacements often shift in an anticyclonic manner toward the mid-level circulation center (Trier et al. 2000).

950-hPA MF (color fill), wind and 900-700 PV (red lines, contour interval=0.5 PVU) valid 0600 UTC Sept. 4, 1985

850-hPA MF (color fill), wind and 900-700 PV (red lines, contour interval=0.5 PVU) valid 0600 UTC Sept. 4, 1985

Let’s look at the 0600 UTC 4 Sept. NARR moisture flux, winds at 950-hPa and 850-hPa. One thing to note is that the low level jet during Elena was not located at 850-hPa but ascended from around 950-hPa over Texas and southern Arkansas to 850-hPa over Missouri. .

0600 UTC 4 Sept cross section showing PV (color filled, interval=0.50 PVA), potential temperature (red lines, interval=2oK), omega (cyan lines). Upward motion is negative and at 0600 UTC was very weak.

The cross section below was taken along the black line on the figure above. The cross section is oriented almost parallel to the low-level jet and the time of the cross section was prior to the nocturnal outbreak of convection. Note that the cross section suggests there was a weak cold pool present where the isentropes sloped upward towards the PV maximum. This cross section is very similar to the idealized one offered by Fritsch et al. shown earlier on this page.

NARR 0600 UTC 925-500 hPa layer mean RH (color fill, legend at left), Θ (red lines), winds (kts)

NARR 0600 UTC sounding, Cape=914, Cin=-36, PW=2.04 at 34.14oN 92.98oW.

At 0600 UTC weak lifting was present to the southwest of the cold pool where the airmass was conditionally unstable and very moist. The high PW and high layer mean RH has the potential to produce a moist absolutely unstable layer if the lifting acts long enough. The combination of a decent vortex signature with strong low level jet typically produces the necessary lifting. Trier et al. have suggested that an absolutely unstable layer often develops with MCVs as parcels originate in the higher thetae air along the periphery of the vortex where there is less cloudiness and higher CAPE. These parcels are then are lifted and become saturated. Being elevated, the parcels therefore are unlikely to be impacted by the nocturnal cooling that takes place in the boundary layer along the periphery of the system where there are fewer clouds. This nocturnal cooling around the periphery of the system tends to increase the CIN around the MCS while the CIN is decreased near the system center due to the development of the MAUL.

NARR 0900 UTC sounding, Cape=120, Cin=-22, PW=2.16 at 34.7oN 92.44oW.

NARR 0900 UTC hodograph

The 0900 UTC NARR f00 sounding shows a saturated layer just above 700-hPa that looks like it might a weak MAUL. By this time, the nocturnal convection had developed. Operational model soundings may have a hard time forecasting a true moist absolutely unstable layer depending on its convection scheme. For example, the BMJ reference profiles make it hard to get a saturated layer if the scheme is activated.
Schumacher (personal communication) has suggested that a hodograph has a very distinctive look most times when an MCV develops new convection near the center. The part of the hodograph that represents the low levels has the look of the top of a coat-hanger.

950-hPA MF (color fill), wind and 900-700 PV (red lines, contour interval=0.5 PVU) valid 0900 UTC Sept. 4, 1985

0900 UTC 4 Sept cross section showing PV (color filled, interval=0.50 PVA), potential temperature (red lines, interval=2oK), omega (cyan lines). Upward motion is negative and at 0600 UTC was very weak.

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Elena continued

Once convection develops, potential vorticity (PV) is increased in the low levels and decreased near the cirrus anvil. Comparison between the PV on the cross sections at 0600 UTC and 0900 UTC 0600 UTC above clearly show this change while also suggesting that the moisture flux at 950-hPa increasing during this period. A compact MCS developed just n of the area of strongest moisture transport near the southern end of the layered PV maximum. It produced over 7 inches of rain.

GOES infrared imagery valid 1000 UTC 4 Sept. 1985.

IR imagery valid 1800 UTC 4 Sept. 1985

IR imagery valid 0400 UTC 5 Sept. 1985.

The convection weakened and by 0400 UTC 5 Sept. cloud debris remained. The NARR analysis below indicates there was still a vorticity maximum and a 30 kt or greater low-level jet. Interestingly enough, despite these features the NARR

0900 UTC 5 Sept NARR 500-hPa heights and vorticity analysis.

925 mb MF (color fill), wind (kts), 900-700 PV (red line) valid 1200 UTC 5 Sept 1985.

GOES IR satellite imagery valid 1000 UTC 5 Sept. 1985.

analysis showed little in the way of upward motion at 0900 or 1200 UTC even though another round of convection and heavy rainfall developed which produced another small area of greater than 7 inch rainfall.

Another example, Erin

More recently, Tropical Storm Erin showed a similar diurnal cycle to its convection though it maintained some convection throughout the day as the storm tracked across Texas into Oklahoma. However, it lost its surface center but still produced heavy rainfall over southwestern Missouri. The interaction of low level jet with the mid level vorticity associated with the remains of the storm again appeared to play a critical role in producing the resulting heavy rainfall event (Schumacher and Johnson 2008). The interaction between mid level vorticity and the low level jet also appears to play a role in MCVs that produce heavy rainfall.

GOES IR imagery of the remains of Erin valid 1800 UTC 18 August 2007 (left), 0000 UTC 19 August (center) and 0600 UTC 19 August (right)

One complicating factor when dealing with weak or decaying tropical systems is that numerical models can have problems forecasting the magnitude and placement of the precipitation and also have problems maintaining the mid-level vorticity center. Model forecasts of precipitation often differ significantly depending on the convective parameterization scheme that is used. In the case of Erin, the NAM,

12-36 hr accumulated precipitation (in mm) for the period ending 1200 UTC 19 August 2007, NAM forecast (top left), GFS forecast (bottom left). Verifying precipitation (top right).

underpredicted the precipitation and at times lost the low-level circulation center. It’s PV maximum was much more poorly forecast than the PV maximum forecast by the GFS. The lack of rainfall and the associated latent hear release negatively impacted on the low level circulation center and low level jet which subsequently helped feed the next round of convection which re-strengthened the PV center each night. Note how the wetter GFS had a stronger PV maximum and better track of the PV maximum even though it was slow moving it northeastward (see below).

The bottom line is don’t give up on the remains of a tropical system when it still has a well defined mid-level center, is embedded in an area of high relatively humidity and precipitable water, especially if a decent low level jet is present to interact with the PV maximum

The remains of Erin strengthened due to the diabatic feedback during the diurnal flare-up of convection near the center. Watonga reported sustained winds of 52 mph and had sustained tropical storm force winds for the better part of three hours.

1004 UTC 19 August radar

Which of following are usually associated with MCVs or decaying tropical systems that have convection redevelop near the circulation center.
stronger low level winds than the mid and upper level winds.
High relative humidity and near saturated conditions
An area of negative saturated EPV (equivalent potential vorticity)
A hodograph with a coat-hanger bend
A well defined mid-level vorticity center

The correct answer is all are typically present though the models may lows one or more of them during it’s forecast cycle if it mishandles the convection.

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