Moisture and Stability of the Atmosphere

Moisture and the Diurnal Temperature Cycle

Background on Atmospheric Moisture

We are constantly effected by the moisture in our atmosphere. One of the obvious effects of atmospheric moisture is something we see and feel, precipitation. Other effects also relate to our comfort level, like humidity. Water in out atmosphere exists in three different phases. Clouds, snow, rain, fog, and haze are all made of some form of water.

The moisture content of the atmosphere is measured in many ways (all discussed in lecture):

Absolute Humidity- The ratio of the mass of water vapor to the volume occupied by a mixture of water vapor and dry air.
Specific Humidity- The mass of water vapor per unit mass of air, including the water vapor.
Mixing Ratio- mass of water vapor/mass of dry air.
Saturation Mixing Ratio- mass of water vapor when a parcel is saturated/mass of dry air in the parcel.
Vapor Pressure- Pressure of water vapor constituent of the atmosphere.
Saturation Vapor Pressure- The pressure of water vapor constituent when the atmosphere is saturated.
Relative Humidity- Vapor Pressure/ Saturation vapor pressure.
Dew Point Temperature- The temperature at which air with the current amount of vapor in it will become saturated.

We have to remember that there are only two ways of increasing the relative humidity:

Cooling the air so it becomes closer to the due point temperature
Adding water vapor to the air

An air parcel with a large moisture content usually signifies the potential for that parcel to produce a great amount of precipitation. Air with a mixing ratio of 13 g/kg will likely rain a greater amount of water than air with a mixing ratio of 6 g/kg. It is also noteworthy that warmer air can also hold more moisture that colder air. For example, if we had one parcel with a temperature of 27°F and a dew point temperature of 21°F and a second parcel with a temperature of 82°F and a dew point temperature of 55°F, the second parcel would contain more water than the first, however, the first would have a larger relative humidity and be more likely to form precipitation.

Moisture and the Diurnal Temperature Cycle

Moisture in the atmosphere can affect the temperature itself through alternations in radiative properties. We already discussed that water has a large heat capacity, energy need to raise the temperature of 1 gram of water 1°C. It turns out that the effect of water vapor heat capacity in the atmosphere is significant and moisture usually has a moderating effect, causing nights to be warmer and also causing days to be cooler.

Moisture also moderates annual temperatures, but not quite as clearly. Proximity to water has a much larger effect on the amplitude of the seasonal temperature cycle than the actual moisture content in the atmosphere.

So we know that moisture in different parts of the country is variable. Well just like the way moisture can vary in this horizontal plane, moisture is variable in the vertical. The vertical and horizontal distribution of moisture play a large role in location of precipitation. Moisture helps the atmosphere become more conducive to convection by reducing the rate at which a parcel of air will cool as it rises. This is due to condensation, which releases latent heat. So, as the parcel rises, it naturally cools and expands as it becomes lower in pressure. However, if a parcel of air is condensing water vapor, forming clouds, the latent heat release will warm the parcel. This warming of the parcel due to the release of latent heat is not enough to prevent the parcel from cooling as it rises, however, it will reduce the rate at which it cools.

Vertical Profiles of the Atmosphere

Vertical profiles of the atmosphere are taken at 0000 UTC (7:00 am) and 1200 UTC (7:00 pm) at about 80 stations across the country and many more stations across the globe. Weather balloons rise to over 50,000 feet and take measurements of several meteorological variables using the radiosonde instrument attached to the balloon. These variables include:

Temperature
Dew Point Temperature
Wind
Pressure

From these variables we can find the following:

Saturation Mixing Ratio
Mixing Ratio

From these profiles we can identify the cloud layers, the dry layers, the inversions and much more.

To begin with we will examine the vertical profile (these are also called soundings) of the atmosphere over Omaha, NE from 0000 UTC 23 May, 2004 (link at the right). This profile shows us the temperature and dew point from the surface to 100mb (which happens to be at 16,420 meters above ground).

Skew-T Log-P

So what does this graph tell us?

At the surface (the lowest level): T ~ 27°C and T_d ~ 20°C
When the balloon reached 850 mb (1402 meters above the ground): T ~ 16°C and T_d ~ 15°C

Based on this we can say that the air at 850 mb was closer to saturation than the air at the surface.

Between 800 and 750 mb, T = T_d, the air is saturated and we can assume there is a cloud at this level.
The tropopause is located at 200 mb (~12km)
From 500mb (5710m, ~5.7km) to 400mb (7370m, ~7.4km) the Temperature drops about 13°C this give that layer a lapse rate of 13°C / 1.7km

This last point brings us to our next question...Why does the temperature decrease with height?

Lapse Rates

A lapse rate is the rate at which the temperature decreases with height. The lapse rate we just described in the sounding was an environmental lapse rate. The environmental lapse rate is variable, it changes every day and at different heights.

Other lapse rates refer to parcels moving up and down in the atmosphere. We know that pressure decreases with height. If we were to take a parcel that was 1000mb and raise it to 700mb it would expand, as it did the temperature would decrease. However, if we were to bring that same parcel back down from 700mb to 1000mb the temperature would increase because the parcel would be compressed.

Before we discuss the next lapse rates we need to address Adiabatic Processes. An adiabatic process is a process in which no heat is added to a system. In terms of a parcel rising/sinking, this means that no heat is exchanged between the parcel and the surrounding environment. In otherwords, the parcel is cooling/warming purely because the pressure is changing. Recall the the first law of thermodynamics

The rate that a parcel cools with height depends the moisture within that parcel.

If a parcel is unsaturaded it rises/sinks adiabatically, so we can say it follows the Dry Adiabatic Lapse Rate

If a parcel is saturated water vapor condenses as it rises and releases latent heat. This means that the parcel will not cool as rapidly as an unsaturated parcel, it will cool at a rate that combines the DALR and warming from latent heat release. The parcel will follow the Moist Adiabatic Lapse Rate.

So now we know something about the path a parcel would take if it were lifted, if it was dry it would follow the DARL, if it was saturated it would follow the MARL. But how do you know when to switch from the dry to the moist lapse rate?

In order to find the height at which a parcel makes the switch from following the Dry ALR to the Moist ALR, we need to know the mixing ratio. How do we get the mixing ratio?

So lets draw the path a parcel would take if it was lifted from the surface...

Start you parcel with the same characteristics as the surface, if your parcel is unsaturated draw a line upward following the dry adiabatic lapse rate from the surface temperature until it hits the surface mixing ratio. This is the level at which the parcel will reach saturation, it is called the Lifting Condensation Level or LCL. After this the parcel will follow the moist adiabatic lapse rate as it rises. Please following figure to help you.