Remote Sensing Tutorial Page 14-1
navigation image mapnext pagetable of contentsprevious page



Dr. William Nordberg This primer is dedicated to my first mentor at NASA Goddard Space Flight Center, Dr. William Nordberg (now deceased) who hired me at a critical time in my professional life and thus brought me into the field of remote sensing. Trained and practicing as a meteorologist, Dr. Nordberg (a native of Graz, Austria), took a lead role in the development of ERTS (Landsat) as he expanded his contributions from weather satellites to those doing Earth Observations for a much greater range of applications.Earlier, Bill, as all of his colleagues called him, had played important roles in the Nimbus and TIROS programs. Dr. Nordberg was one of the finest scientists and leaders I've ever served under. I am especially grateful for his tolerance of my demands for more field work - our arguments were vociferous but always controlled by a mutual friendship. By the time I accepted his offer to transfer to NASA (from a National Academy of Sciences Associateship), he had become Director of the Earth and Atmospheric Sciences Directorate. He took a special interest in transforming me from a ground geologist to a space geologist.

After this subsection was started, the writer while surfing the Internet in search of many of the illustrations used in the next four pages, found a fairly good Meteorology tutorial produced by staff of the Department of Meteorology at the Lyndon State College, in Vermont, web site. A number of illustrations from this site (identified by LSC just below image) are used in the subsection The reader may want to work through this now or at the end of the subsection. The subsection has the advantage of more detailed text that cover meteorological concepts and explains accompanying illustrations. However, the web site, divided into chapters, has an idiosyncracy -- there is no button or other mechanism to go from one chapter to the next (or any other chapter through a table of contents). To change chapters, go to the URL listed in the top of your browser, click right next to the "r" in chapter, delete the chapter number and replace it with the one you want (19 chapters in all), and click Enter. The new chapter will load. Repeat this as needed.


Meteorology (from the Greek meteõros meaning "high in the air") is defined as the science of atmospheric phemomena and processes. Weather denotes the short-termed (hours to a few days up to a week or so) behavior of the atmosphere, generally with the connotation of applying to local or regional parts of the Earth's globe, as it varies in the conditions used to describe weather - fair, rain, warm, windy, etc. Thus, each place is said to have a set of particular conditions, which tend to change over short time spans, that affected people describe as the weather for the day or perhaps as long as the next week. Climate refers to much longer time frames and describes the common characteristics of weather in broader parts of the Earth's globe. It applies to wider regions and depends on geographic location, physiographic conditions, time of year, and other factors. Climate is thus tied to 1) the larger, longer variation in typical or average weather in a region determined by the seasons - which in turn depend on the location of the Earth, with its tilted axis, as it rotates around the Sun in an annual cycle; and 2) the range of conditions expressed in day to day weather variations and extremes in such properties as temperatures, extent of cloud cover, and duration and types of rainfall/snowfall.

On Earth, the atmosphere is also known as "air" (with the Space Age we now know that some other planets and rarely even their moons have atmospheres; in this Primer we restrict our discussion to the Earth alone). The terrestrial atmosphere is part of the four "spheres": 1) the Lithosphere; 2) the Atmosphere; 3) the Hydrosphere; and 4) the Biosphere. These can, and do, interact and influence one another. The atmosphere itself is a mixture of gases of different compositions, plus water as vapor, liquid, or solid, and suspended particles of various natures. These gases follow various rules of thermodynamics (e.g., Charles and Boyles Laws). Thus, the physics of meteorology becomes the physics of gas behavior as described by temperature, pressure, circulation, density stratification, and energy utilization and transformation. The main (predominant) source of energy powering atmospheric behavior and change is solar radiation from the Sun. Also contribution to thermal inputs (or exchange) is terrestrial heat flow, in and out movement of heat from water reservoirs (mainly, the ocean), contributions from Man, animals, and plants, and chemical reactions within the atmosphere which can be exothermic or endothermic.

In the brief synoptic treatment of Meteorology that follows on these next pages, the primary sources of information and of illustrations has been taken from the Internet and from The Atmosphere: An Introduction to Meteorology, 4th Ed., by F.K. Lutgens and E.J. Tarbuck, Prentice Hall, 1994 and Physical Geography of the Global Environment, by H.J. de Bilj and P.O. Muller, John Wiley, 1993

To begin our survey, lets introduce the terminology that labels of the main structural components of the atmosphere as a function of height above the surface:

Subdivisions of the Earth's atmosphere.

We live within the troposphere. Jet aircraft are capable of flying in the upper troposphere and lowermost stratosphere (where the air is less dense [less drag on the plane and hence greater speed] and smoother, often cloud-free wind conditions) and experiment rocket planes have penetrated up to 90000 feet into the lower stratosphere. Knowledge of the air above this has come from balloons, sounding rockets, and orbiting satellites, as is considered in this Section.

A variant of this classification, similar to the one shown on page 14-1 that featured the temperature profile through the atmosphere, which includes some additional compositional information appears in this diagram

Profile of temperature and composition variations within the atmospheric layers.

The composition of the lower half of the atmosphere is dominated by two elements in their gaseous molecular forms, N2 and O2. This table lists all notable atmospheric constituents, their relative amount (in parts per billion), their sources, and their variations over time:


Table summarizing compositional properties of the atmosphere.

The two most notable numbers are N2 = 78.1% (7.81 x 108 divided by 109 x 100) and O2 = 20.1 %. The nitrogen was derived originally from escape from the Earth's interior. Oxygen has increased its percentage over time largely through photosynthetic processes, starting about 1.5 billion years ago. Argon (derived from radioactive decay of potassium in rocks) is the next most abundant component, about 1%, but can sometimes be matched by water vapor (0.1 to ~1.0%). Carbon dioxide (CO2) and Ozone (O3) Play small but important roles. Many of the other "trace" gases have either a biogenic origin, a volcanic origin, or result from human activity.

This simple pie chart diagram should help to simplify the generalization of atmospheric composition:

Simplifie diagram of atmosphere's composition.

The composition of air up to the outer limits of the stratosphere (out to about 70 km) is almost constant and homogeneous insofar as the three dominant species is concerned. Over time the addition of CO2 and possibly O2 slightly lowers the percentage of all other constituents (since they must be summed to 100% as a constant). In the upper atmosphere, which has a much less amount of all constituents (leading to lower densities), some of theses constituents have enough thermal energy to excape into outer space (hydrogen and helium being light are the principal gases that so leave).

Over the 4 plus billions of years in which Earth has had an atmosphere, its composition has changed significantly. The early atmosphere was largely ammonia and carbon dioxide plus some nitrogen, with little oxygen, but may have had more hydrogen.

The total mass of the global atmosphere is estimated to be ~5.1 x 1018 kilograms. Ninety percent of that mass is in the inner 10 km (6 miles). The average density of air at the Earth's surface is 1.2 kg/m3. The mean pressure of a column of air at sealevel, coming from gravitational attraction on all gases above it to the outer limits of the exosphere (where gas molecules diminish to the levels of outer space) is set at 1.013 bars, or more commonly expressed as 1013 millibars. A general pressure gradient for the atmosphere out to about 11 miles (18 km) is shown below:

Note that this and the next plot are exponential curves. At about 18000 feet (5.45 km), the pressure is 500 millibars, about half that at se level. This means that most of the mass is concentrated in the troposphere. (For pressures lower than about 400 mb, the amount of oxygen present is insufficient for humans to function for a long period without a supply of oxygen from tanks (climbing Mt Everst is possible without oxygen tanks but dangerous and requires a minimum stay about the 28000 ft level.) Air becomes progressively thinner with altitude, i.e., its density diminishes as shown in the graph below (which closely follows the pressure curve).

Atmospheric density variation with altitude out to about 10 km.

The parallel behavior of diminishing pressures and densities out to 500 km follows this same general exponential pattern:

Pressure and density changes out to 500 km altitude.

Measurement of this exponenential reduction in atmospheric density in the outer atmosphere (Mesosphere-Ionosphere) is beyond the reach of balloons but has been determined quantitatively by several means. Here is the results from several sounding rockets:

Uniform exponential drop in atmospheric pressure out to the Exosphere.

In several of the above graphs, the pattern of change in temperature has shifted from decrease with altitude to increase with several inflections. The causes for this are varied, as described here.

In the troposphere, the atmosphere follows the normal behavior of the physics of gases. From thermodynamics we learn that as an adiabatic (no external heat added or removed) expands, it cools. Conversely, compressing (reducing volume) of a gas, which requires an energy input, causes it to heat up. This is a fundamental concept that underlies much of what happens in observing meteorological phenomena in the inner air shell around the Earth.

The change of T-P states just described can be determined by applying the gas laws. Boyle's Law (for a constant temperature) is written as:

P1V1 = P2V2

Thus, if the volume is increased by 1/2, the pressure will decrease by that amount to maintain the equation's equality. That is, if V1 becomes 2 V2. then P2 must decrease by 0.5 (50%) to maintain equal multiplier products on both sides of the equation. Charles Law can be written V1/V2 = T1/T2. Again, to maintain equality, if a Volume is contracted (V1--->0.5V2), then the corresponding T2 must be 0.5T1; to be valid, the pressure must be maintained as constant. This results in cooling. This can also be expressed according to the Ideal Gas Law: PV = ρRT. where ρ is density and R is the gas constant of proportionality. (Note that T in this formula is absolute temperature in Kelvins K = C + 273). Using this formula, it is evident that if V increases (gas expansion), temperature likewise increases; if P decreases, the T falls - as indeed it does moving upwards through the troposphere.

The increasing thinness of the atmosphere above the Tropopause brings into play different physical phenomena that alter temperature patterns. In the Stratosphere, the presence of a very small amount of ozone (O3) has a significant influnece on temperatures since it absorbs heat from ultraviolet radiation. This temperature gradually increases until the stratopause (roughly equivalent to ozone being fully depleted) where a temperature inversion (inflection) brings about a resumption of adiabatic temperature drop down to about minus 90°C. This continues through the mesosphere. Another reversal leading to significant rise in temperature begins at the mesopause extending through the thermosphere. Even though its atmosphere is tenuous (very low densities), the presence of tiny amounts of oxygen and nitrogen is sufficient to absorb short wavelength solar radiation, leading to the positive temperature gradient. Levels as high as 1000° C are reached owing to the ability of moving atoms and molecules to attain higher speeds in the very low density atmosphere which means they are more energetic and hence this kinetic energy transforms into increased temperatures.

The atmosphere is kept in motion and develops its weather/climate characteristics primarily through energy input as heat (as a term this refers to quantities of thermal energy). Heat actions produce various temperatures (a measure of the degree of atomif/molecular motions through heat inputs and withdrawals). As we saw above, temperature variations give rise to differences in pressure (hence development of pressure gradients that drive gases into motions that include wind) and volumes. The dominant source of heat affecting the atmosphere come from solar irradiation or insolation (heat added from internal Earth flow adds a small amount to the land/ocean bottom surface). This insolation (= spectral irradiance, given in units of Watts/square meter) is primarily short wave radiation as shown in this diagram (see also page 9-2) which also indicates that some of that radiation is returned to and through the atmosphere as long wave radiation. The atmosphere has windows of transparency in the wavelengths shown.

Spectral distribution of incoming solar radiation and distribution of longer wavelength thermal radiation derived from the actions described below.

The average amount of solar energy flow as shortwave irradiance reaching the outermost edge of the atmosphere is defined as the solar constant. This has a value of 1370 W/m2 (note: Watts has time in its definition, so it describes a flow rate). The solar constant can also be expressed in terms of heat energy flow and is ~2.0 calories per square centimeters per minute.

The incoming radiation will react with a wide variety of "targets" From this one can calculate a partitioning of heat energy to various recipients such as the ground, water bodies, vegetation, air gases, clouds, etc. This can be shown diagrammatically as an Energy Budget or Heat Budget:

Heat Budget of the Earth.

The irradiation history can also be quantified in terms of irradiance and radiant emittance, in units of Watts/m2 as shown in this diagram:

The Solar energy balance in terms of radiant flux units (Watts per square meter)>.

In these examples, which show the same thing but in somewhat different partitioning modes,, the heat or energy flows must ultimately balance within the partitioning system in an equilibrium mode. But when the system is treated in terms of net effects leading to accumulation of heat in the atmosphere the outward flow into space will be less than the inward energy received from the Sun, so that warming results (see below).

The system utilizes conservation processes which are designed to attain a balance. The processes consist of inputs/outputs (flows) of radiant heat, sensible heat (that transferred by convection/conduction), latent heat (used or released in change of state processes such as water condensing into rain, freezing to ice or evaporating to cloud condensation), and ground (soil) heat.

Consider this pair of diagrams which is a variation of the first balance plot above, but specifically allots heating modes numerically to show the balance between incoming and outgoing radiation of a thermal nature:

Another version of the Energy Budget that partitions solar irradiance.
From Strahler and Strahler, Introducing Physical Geography, 1994

Drawing from their text, and condensing the explanation, let us account for the numbers in the diagram pair. In the left diagram, incoming solar radiation is arbitrarily set at 100 energy units (expressed as percent). 31% of the incoming short wave radiation is reflected by the ground, air molecules, clouds, and dust, representing an averaged albedo (reflected radiation). The atmosphere's dust and clouds absorb 21%. The remaining 48% is absorbed by land and ocean surfaces. The right diagram describes the various ways in which outgoing long wavelength radiation is ahsorbed by the main components of the system: land, water, air, and outer space. As emitted surface radiation 107 units are absorbed by the atmosphere and 6 units escape into space - this is a total of 113 units. Where does the extra percentage come from? The atmosphere counterradiates 97 units to the surface which must be added to the 48 units of incoming received and absorbed, giving a total of 145 units introduced to the surface as combined short and long wavelength radiation. On the right side, 10 units of sensible heat and 22 units of latent heat (mostly from ocean evaporation) combine with the 113 units of long wave radiation loss, yielding 145 units that balance the opposing flows. In the atmosphere itself, the partitioning of energy is somewhat different: The 63 units radiated to space are added to the 97 units counterradiated to the surface to give a total of 160 (percentage) units removed from the atmosphere. This is balanced in the long wave ledger by the 107 units released from the surface by radiant emission plus 21 units from outer space, and the 22 and 10 units from latent and sensible heat release respectively. The sum is 160 units as additions to the atmosphere.

However, atmospheric warming itself has its own numerical input that does not balance: This is based on the 145 units of incoming direct ground radiation and counterradiation (48 + 97 units) but of that number only 113 becomes outgoing radiation. The difference of 32 units accounts for a net gain in heat-producing energy in the atmosphere that yields the temperatures that characterize our planet's temperatures (whose ranges result from interactions as described above). This is analogous to the "Greenhouse Effect" described on the third page of this subsection.

In fact, one must also consider a somewhat different parameter - net radiation balance (not heat) - in which the outgoing radiation (397, in units of calories per square meter) is less than the incoming radiation (469 cal/m2) on an averaged global annual basis. This takes into account differences between heat response in the oceans and the land and in latitude locations.

Of great importance is the geographic distribution of net radiation. The excess of radiation at low latitudes results from radiation coming in faster than it goes out; the converse (slower in, faster out) takes over at the higher latitudes. This build-up of heat around the equator and depletion in polar regions is responsible for the poleward flow (transport) of heat energy to equalize the total energy distribution. That is one of the driving forces in the circulation patterns that mark Earth's atmosphere. This diagram shows that the low latitudes in the course of an annual seasonal cycle gain more heat whereas the higher latitudes loose heat.

Net gains and losses in radiation heating results during a year as as function of latitude location.

This figure summarizes the global net radiation distribution at the Earth's surface (in units of 1000 calories per square centimeter.

Distribution of net radiation globally on an annual basis.

From de Blij & Muller: Physical Geography of the Earth's Environment

The variations in net radiation leading to gains and losses should not be surprising. The above is a somewhat idealized picture that treats the atmosphere as a uniform entity rather than as a system in which inputs and outputs vary over the globe resulting in different conditions from place to place (such as albedo, cloud cover, etc.) that change with time. The same is true for the above heat budgets in the sense that the percentages vary with affecting conditions but should average out over time. It is the differences on a geographic and temporal basis that gives rise to the vicissitudes of weather and climate we observe on our planet.

This variation in heat content over the seasons affects the range of temperatures in marine waters, as shown below for several different oceans. In general, surface sea water experiences a minimum of temperature variations in low latitudes and a much high range of change as the polar regionsare approached.

Temperature range profiles from the equator to the poles.

The oceans act as major heat reservoirs that have significant effects on the roles that heat and radiation play in generating meteorological patterns. This generalized diagram shows the global variations in incoming solar radiation (irradiance) over the oceans.

Distribution of solar radiation annually over marine waters.

A fraction of that will escape the ocean water reservoir as sensible heat (top map) whereas another fraction enters the ocean as stored heat.

Amount of radiation returned to the atmosphere as sensible heat released from the ocean reservoir

Amount of heat retained in storage within the ocean reservoir.

The oceans experience differential heating, generally being warmer at the surface. Cooler water, being more dense, will sink. Since there are temperature differences within the oceans at various depths, these give rise to current movements. The general pattern of water transfer is shown here:

Generalize flow patterns of surface and deep currents in the oceans.

In the above illustration and the one below, the patterns are modified by the presence of land masses in "irregular" dispositions (locations) around the globe - the present configuration of continents results from the stage of plate tectonic movements now acting on the earth continental and oceanic lithospheres (crust and the uppermost mantle). Another factor: there is much more land in the northern hemisphere, so that 81% of the southern hemisphere is ocean water compared with 61% in the northern hemisphere; this suggests that marine thermal processes play a larger role south of the equator.

The surface currents result from cold water in higher latitudes moving equatorward, displacing the more heated (warm) waters that move poleward as a thermally driven circulation ensues. Movements are also affected by wind forces. A warm oceanic current transfers some of its heat to adjacent lands whereas the cooler current modifies land temperatures around coastal regions (best example, the West Coast of the U.S.). Best known of these surface currents to U.S. easterners is the Gulf Stream (a warming current); Benjamin Franklin was among the first to recognize its significance in weather control and pointed out that ships sailing from Europe westward took longer to reach North America since they often encountered the eastward flow of the Gulf Stream.

Generalized ocean surface currents; note their names.

Earth is characterized by its seasons, which are extended periods in which weather and climate have different average temperature, rainfall, and storm frequency and type conditions. In much of the earth beyond the tropical regions around the equator, the seasons also are marked by extended periods of active vegetation growth alternating with vegetation dormancy. The seasons more than any other factor influence human, animal, and vegetation activities over the course of a year. The reasons for seasons are mainly twofold: 1) the position of the Earth relative to the Sun during its annual 365.4 day revolution around the Sun; 2) the tilt of the Earth's axis of rotation. The first factor, shown here, has a much smaller effect than the second; distances from the Sun vary by less than 10% so that insolation intensity is not great but has some influence.

The Earth's orbit around the Sun; Earth is closer at perihelion.

The axis of Earth is tilted 23.5° from the plane of the ecliptic (in which is embedded its rotational path without any upward/downward motion during an annual cycle. This as we shall see shortly is the key to the changing seasons. Lets look first at the total path during that year, noting the direction of tilt.

The annual rotational positions of Earth at four times during the year; these are named as labeled and as described below.

At the Summer Solstice, the 23.5° tilt is leaning towards the Sun in the northern hemisphere; at the Winter Solstice the southern hemisphere is leaning 23.5° towards the Sun (it is winter in the northern hemisphere but summer in the southern hemisphere). Twice a year at the Vernal (norther Spring, southern Fall) Equinox and the Autumnal Equinox (northern Fall; southern Spring), the position of the Earth relative to the incoming Sun's rays is such that the length of the day is equal in both hemispheres.

To establish why the Earth's axis tilt is the prime cause of seasonal variation, start first with this diagram:

Effect of increased path length of solar rays moving from equator to poles.

In this simplified diagram, the sphere has a vertical rather than tilted axis (top to bottom). Note that at the equator, the length of the path of sunlight through an atmosphere of uniform thickness and identical density distribution is the shortest of all rays. That path length increases systematically moving to either pole. The shorter the sun ray path, the less is the interaction between the insolation and anything in the atmosphere that deflects, absorbs, or re-radiates the irradiance. Thus a greater proportion of radiation hits the surface or reacts in the troposphere. The result is increased warming (maximum at the equator). As the path lengthens poleward, the amount of radiation reaching the surface and lower atmosphere decreases, thus producing progressively less warming poleward. The mean (and maximum) temperatures will therefore be lessened going from low to high latitudes.

Now, examine the case where the axis is tilted 23.5° towards the Sun. The effect is to move the North American continent (or Eurasia) clockwise downward towards the equator and hence into regions of shorter pathlengths. Midlatitudes and even polar regions receive more surficial radiation and thus become warmer. This warming during this position of the pole and times of year close to this orientation leads to summer conditions. Vegetation commences annual growth as the conditions approach; subpolar and polar ice undergoes some degree of melting.

Radiation paths for an Earth whowe upper half (northern hemisphere) is tilted 23.5 degrees towards the Sun

Then consider the reverse case, a 23.5° tilt away from the Sun at the time of winter solstice:

The insolation pattern for the Earth in northern wintertime.

In this situation lands and oceans in the southern hemisphere move upward into the path patterns where the lengths through the atmosphere decrease and more warming radiation reaches these areas. This is the condition for southern hemisphere summer. Northern lands and seas have also moved upwards into the zones of greater path lengths and hence surface-troposphere temperature decreases. This is northern winter.

At the Equinoxes, the Earth's equator lies directly beneath the solar point. Between Summer and Winter solstices the equator is shifted a total of 2 times 23.5 or 47° over this six month period (the same occurs in the opposite direction between Winter and Summer. This oscillation accounts for the seasons. At intermediate times the heating patterns relative to the Earth's geography tend, on average, to fall between the extremes. In either hemisphere, this next diagram applies.

Variations between the observed angle to the Sun at Summer and at Winter.

In the summer case, the angle between a direct line to the Sun and the nearby horizon (Earth's surface meets the atmopheric base in the line of sight) is higher than in the winter case. This simply means that the Sun appears higher in the sky at Noon in Summer than in Winter. This is depicted thusly:

Top pair of panels: position of the Sun in the sky at Noon during a winter and a summer day; bottom pair: indicates that a lower sun angle (winter) spreads out the sun's rays (as though directed through a lens) over a wider area than during the summer.

The last diagram indicates two things: 1) a higher sun angle will also lengthen the time in 24 hours in which the Sun shines at a given place; and 2) the more concentrated solar beam leads to greater warming of an area. At the two Equinoxes, the length of day at the Equator is equal to the length of night (neglecting dawn and dusk effects due to light scattering). In the northern hemisphere's progression from Winter to Summer the duration of daylight goes from a minimum to a maximum (as it does for the southern hemisphere cycle 6 months apart). The further north one goes in, say, North America during any given day, the length of daylight (summer) or darkness (winter) increases as the season progresses. In Arctic regions, in mid-summer the Sun is always above the horizon ("perpetual" daylight) but in the winter months, polar regions may experience total diurnal darkness. (Same is true for the southern hemisphere; those doing studies in the Antarctic may leave the continent at such times, especially since then the weather is coldest.) In a hemisphere, the transition between Winter and Summer leads for a given region (at some latitude) to a gradually progressive warming; conversely, the trend from Summer to Winter is an overall cooling. (Low latitude regions, such as equatorial tropics, experience some changes in daily temperature variations [that will occur simply because of non-heating during the night hours], but with usually a minimal range and a higher average temperature spread over the full year).

Note too that the maximum and minimum temperatures do not necessarily coincide with the Summer and Winter Solstices per se. There is a lag effect that tends to delay the hottest days to the July months and the coolest days to the January months in the northern hemisphere.

As a transition to the second page on weather systems, which have both a seasonal and a daily character at individual locations, let's take a quick look at the temperature and pressure changes that typically mark a single day's activities. First, how does temperature vary at some given spot, say, in the eastern United States:

Diurnal variations in incoming-outgoing radiation and temperature changes.

From the lower chart, incoming solar radiation peaks around high noon. Outgoing radiation reaches its minimum around dawn. But, it also attains its maximum about 3 to 4 hours after noon, thus, not coincident with the radiation peak. This lag is the result of several factors, chief of which is the effects of thermal uplifting and winds that carry heat upwards and slow down the surface temperature rise until mixing in mid-afternoon produces the hottest time of the day.

Earthlings experience their weather conditions on the ground within a general layer of air called the boundary layer. This is the bottommost layer that has a thickness of 1 to 1.5 km (up to a mile) in the afternoon but which shrinks to about 100 meters (330 ft) deep in the night. The next, largely self-explanatory, four illustrations show temperature conditions, winds, and air movements at 4 times during daylight hours in a 24-hour thermal cycle. Note that maximum activity in the boundary layer (BL) involves upward flow of air heated at the surface in the afternoon. On a summer day, with local high pressure air and blue skies, clouds of the cumulus type (see next page) are the result.

Boundary Layer (BL) conditions at 6 AM.

BL conditions at 11 AM.

BL conditions at 3 PM.

BL conditions at 6 PM.

navigation image mapnext pageprevious page

Primary Author: Nicholas M. Short, Sr. email: