Assessing Strong Updrafts

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From a weather forecaster's perspective, predicting thunderstorms and severe weather is always challenging (for a variety of reasons that we'll gradually uncover in this course). Seasoned forecasters develop their own forecasting routines and favorite tools that help to approach the problem in the most consistent way possible. In time, you'll develop your own specific approach and favorite tools for forecasting thunderstorms, but your routine should certainly include:

• getting a firm handle on the big picture (synoptic-scale weather pattern) at the surface and aloft
• getting a sense for overall instability and the potential for strong thunderstorm updrafts
• assessing the magnitude and role of vertical wind shear

While these items are listed as separate bullet points, the reality is that they're all intertwined. Aspects of the big picture impact the potential for strong updrafts as well as the magnitude and role of vertical wind shear, all of which are crucial pieces of any forecast for thunderstorms. In order to see how the big picture impacts the potential for strong updrafts and the magnitude of vertical shear (one of the overarching goals of this lesson), we have to cover a few basics first.

We'll start by covering how forecasters assess the potential for strong updrafts, and doing so requires tackling a few definitions. This discussion requires a good basic knowledge of skew-T diagrams and associated concepts, which you've studied previously. If you're rusty on these topics, I strongly recommend that you spend some time reviewing skew-T basics from Lesson 6 of METEO 101. The basics that you learned previously about how to read information and move parcels on skew-Ts will be absolutely critical in this discussion and our deeper look at skew-Ts later on, so don't skimp on any review time you might need!

With that caveat out of the way, let's start with a few definitions.

The Level of Free Convection

Recall from your previous studies that the lifting condensation level (LCL) is the level where net condensation begins in a lifted parcel. If the parcel continues to rise above the LCL (there's no guarantee it will do so), it now cools at a reduced rate -- the moist adiabatic lapse rate -- marked by the thin blue curve in the image below. If something keeps forcing the parcel to rise (lifting from low-level convergence is one possibility), eventually, the parcel reaches an altitude where its temperature equals the temperature of the environment. Assuming the air parcel rises slightly above this altitude, it becomes positively buoyant, and accelerates upward, setting the stage for deep, moist convection. In light of this convective scenario, meteorologists refer to the altitude where the air parcel first becomes positively buoyant above the LCL as the Level of Free Convection (LFC). In this context, the adjective, "free", means that the positively buoyant parcel will rise freely through a deep layer of the troposphere. No further lifting by an external force is required.

If a parcel is forced to rise moist adiabatically above the LCL, eventually it reaches the LFC, the point above which the parcel becomes positively buoyant and will accelerate upward on its own.

Credit: Penn State University

You may be asking yourself, "is there always an LFC?" The answer is a resounding, "no." For example, take a look at this sounding from Pittsburgh, Pennsylvania at 12Z on January 6, 2016. A couple of things should jump out at you immediately: First, the lower troposphere is overwhelmingly stable and dry. It was warmer at 700 mb than it was at the surface at this time! Could a parcel lifted from the surface ever become positively buoyant (warmer than its environment) through a deep layer? Absolutely not! Check out this annotated sounding showing the path a parcel would take from the surface to its LCL (not far above the surface), and then lifted moist adiabatically. The parcel is always to the left of the temperature sounding, so it's colder. In other words, even with Herculean lifting, a parcel will never become positively buoyant. It has no LFC.

Even though this particular example came from a location in a cold, dry, Arctic air mass in the winter, the environment may not have an LFC at any time of year -- even in summer, when it's warm and humid. The presence of an LFC (or lack thereof) and its altitude depend largely on lapse rates and low-level temperatures and dew points (we'll explore these issues more shortly).

Equilibrium Level and CAPE

Once a parcel becomes positively buoyant above its LFC (assuming an LFC exists and the parcel makes it to that level), where does the positive buoyancy stop? The answer is called the "equilibrium level." Formally, the Equilibrium Level (EL) is the altitude above the Level of Free Convection where the temperature of a positively buoyant parcel again equals the temperature of its environment (the EL often occurs near the tropopause).

With the LFC and EL safely tucked under our learning belts, we're ready to assess the potential for strong updrafts. On skew-Ts (plotted from radiosonde measurements or model forecasts), CAPE, which stands for Convective Available Potential Energy, is simply the area between the temperature sounding and the local moist adiabat that a lifted air parcel follows between the Level of Free Convection (LFC) and the Equilibrium Level (EL). The positive area on the idealized skew-T near the top of this page (shaded in green) represents CAPE. Of course, CAPE is zero whenever there isn't any surface-based LFC (it's way too stable for air parcels lifted from the ground to become positively buoyant).

So, what does CAPE mean in a practical sense? Let's start with the idea of  "Potential Energy". Quite simply, the word "Potential" refers to the possibility that air parcels lifted from the surface make it to the LFC. What about "Energy?" If air parcels lifted from the surface are able to reach the LFC, they become positively buoyant and accelerate upward through a relatively deep layer of the troposphere thanks to the temperature difference between a parcel and its surroundings, which paves the way for deep, moist convection. In light of this process, CAPE (positive area on a skew-T) is a proxy for the total possible amount of kinetic energy that an air parcel can gain between the LFC and the equilibrium level because of its positive buoyancy. The parcel's positive buoyancy is determined by the size of the temperature difference between a parcel and its surroundings, which governs the magnitude of the parcel's upward acceleration, a relationship which you can explore in the interactive tool below.

Interpreting and Using CAPE

Whenever you're working with CAPE, you should always be aware of the proper units. For the record, the units of CAPE are Joules (a unit of energy) per kilogram (J/kg). How can we interpret values of CAPE? In general, you should treat values of CAPE between 0 and 1000 Joules per kilogram as small. When you see CAPE values higher than 2500 Joules per kilogram, think large. But, I wouldn't get carried away with small and large values of of CAPE, because severe thunderstorms (and tornadoes) can and do occur with small values of CAPE (only a few hundred Joules per kilogram). On the other hand, sometimes environments with large values of CAPE (well over 2500 Joules per kilogram) fail to yield a single thunderstorm!

You'll often see CAPE described as "an overall measure of instability in the troposphere." But, treating CAPE this way has some problems. When CAPE is really high and thunderstorms fail to materialize (air parcels lifted from the surface never make it to the LFC), equating CAPE with instability is, at the very least, misleading because some parcels were nudged upward and didn't continue rising as they would in an unstable situation (they weren't forced up far enough to reach the LFC). Plus, the general public tends to equate instability with thunderstorms, so it's wise to avoid using instability to describe CAPE. When you see tables describing CAPE values on the Internet, I wouldn't put much stock in them.

So how should you think about CAPE? I like to treat CAPE as a measure of the potential for strong updrafts. If air parcels lifted from the surface reach the LFC in an environment with CAPE (especially moderate to high CAPE), they accelerate upward, acquiring kinetic energy and forming strong updrafts in developing thunderstorms. If air parcels don't make it to the LFC in an environment with high CAPE, there certainly was a potential for strong updrafts, but they never materialized.

The moral of this story is that there just isn't any universal way to interpret values of CAPE. While CAPE helps forecasters assess the potential for strong updrafts, specific values of CAPE do not guarantee that thunderstorm updrafts will actually form, and cannot be connected to specific updraft speeds. For more of a quantitative look at the connection between CAPE and updraft speeds, check out the materials in the Explore Further section below.

Indeed, to interpret CAPE you must take into account climatology, the season, and the prevailing weather pattern. To see what I mean, consider that values of CAPE along the West Coast are, on average, much smaller than average values over the Middle West. Moreover, CAPE is usually smaller, on average, during winter than it is during spring and early summer. 

Ultimately, what determines whether or not strong updrafts will actually materialize if CAPE is present? Let me introduce "Convective Inhibition."

Convective Inhibition

Convective Inhibition (CIN) is a proxy for the amount of energy needed to lift a parcel to its LFC. So, if CIN is great, and lift rather weak, thunderstorms probably won't happen because parcels won't make it to the LFC and accelerate upward. On an idealized skew-T (see below), CIN is the area between the temperature sounding and the dry adiabat / moist adiabat followed by a lifted parcel on its way to its LFC. CIN is represented by the negative area (in red).

A sounding showing the spatial relationship between the LFC and negative area. Also known as Convective Inhibition (CIN), this area is directly proportional to the amount of energy required to lift an air parcel to its LFC.

Credit: Penn State University

In the idealized skew-T above, the temperature inversion near 850 mb, and the stable layer (small lapse rates) just above it are responsible for a large chunk of the CIN (the negative area shaded in red). To give you an idea of how to interpret CIN values, keep in mind that because CIN is a "negative area," its values are negative, and the more negative the number, the greater the CIN. In general, you can rank CIN values between 0 and minus 25 Joules per kilogram as weak inhibition. CIN values between minus 25 and minus 50 Joules per kilogram typically qualify as moderate. When you see CIN values of at least minus 50 Joules per kilogram or more, think large inhibition.

As it turns out, CIN can be reduced. How's that? In short, the synoptic-scale weather pattern helps to prime local environments for deep, moist convection by reducing CIN primarily via the following three processes:

• low-level heating
• low-level moistening
• synoptic-scale lift

We'll be exploring how synoptic-scale lifting can reduce CIN throughout this lesson, but the role of low-level heating and moistening in reducing CIN are fairly intuitive. For starters, check out this interactive tool showing how low-level heating reduces CIN (pink shading indicates CIN). In a nutshell, heating of the ground and the overlying layer of air moves the lower portion of the temperature sounding toward increasing temperatures (to the right). Meanwhile, the lapse rate in the gradually deepening boundary layer trends toward dry adiabatic, reducing CIN. Also note how the LFC lowers and CAPE (positive area) increases with time in response to low-level heating.

How does low-level moistening reduce CIN? This interactive tool illustrates the consequences of low-level moistening for CIN. In essence, moistening moves the lower portion of the dew-point sounding to the right toward higher dew points. Increasing moisture causes the LCL to lower because an increase in moisture means that air parcels need not rise as far to achieve net condensation. Moreover, note that the LFC also lowers and CAPE (positive area) increases. And, of course, the lapse rate in the well-mixed boundary layer remains dry adiabatic.

These first two cases should not surprise you because increasing surface temperature and dew points ultimately translates to an increase in energy that's available for deep, moist convection. But, before we move on, I want to make an important point: If you're thinking that CIN must be zero for thunderstorms to initiate, wipe this notion from your mind. It's incorrect. Some CIN usually exists when thunderstorms erupt (but its magnitude is fairly small). Overcoming existing CIN is a major theme that we'll cover throughout this lesson.

To access real-time model analyses of CAPE and CIN, SPC's Mesoscale Analysis Page provides a great resource. Under the "Thermodynamics" menu, you'll find several "varieties" of CAPE. The most basic form, that we covered on this page, is "Surface-Based" CAPE. We'll cover some of the other types later in the course (and one in the Explore Further section below).

While forecasters use CAPE as a tool to assess the potential for strong updrafts, they look at environmental lapse rates in the lower half of the troposphere to get a more direct sense for the instability that exists. Let's investigate further.