Remote Sensing Application PDF

What is Remote Sensing? Remote sensing is the science (and to some extent, art) of acquiring information about the Earth's surface without actually being in contact with it. This is done by sensing and recording reflected or emitted energy and processing, analyzing, and applying that information....

What is Remote Sensing? Remote sensing is the science (and to some extent, art) of acquiring information about the Earth's surface without actually being in contact with it. This is done by sensing and recording reflected or emitted energy and processing, analyzing, and applying that information." In much of remote sensing, the process involves an interaction between incident radiation and the targets of interest. This is exemplified by the use of imaging systems where the following seven elements are involved. Note, however, that remote sensing also involves the sensing of emitted energy and the use of non-imaging sensors.  Passive vs. Active Sensing The sun provides a very convenient source of energy for remote sensing. The sun's energy is either reflected, as it is for visible wavelengths, or absorbed and then reemitted, as it is for thermal infrared wavelengths. Remote sensing systems which measure energy that is naturally available are called passive sensors. Passive sensors can only be used to detect energy when the naturally occurring energy is available. For all reflected energy, this can only take place during the time when the sun is illuminating the Earth. There is no reflected energy available from the sun at night. Energy that is naturally emitted (such as thermal infrared) can be detected day or night, as long as the amount of energy is large enough to be recorded. Active sensors, on the other hand, provide their own energy source for illumination. The sensor emits radiation which is directed toward the target to be investigated. The radiation reflected from that target is detected and measured by the sensor. Advantages for active sensors include the ability to obtain measurements anytime, regardless of the time of day or season. Active sensors can be used for examining wavelengths that are not sufficiently provided by the sun, such as microwaves, or to better control the way a target is illuminated. However, active systems require the generation of a fairly large amount of energy to adequately illuminate targets. Some examples of active sensors are a laser fluorosensor and a synthetic  Sensors A sensor is a device that measures and records electromagnetic energy. Sensors can be divided into two groups. Passive sensors depend on an external source of energy, usually the Sun (although sometimes the Earth itself). The group of passive sensors cover the electromagnetic spectrum in the range from less than 1 picometre (gamma rays) to over 1 metre (micro and radio waves). The oldest and most common type of passive sensor is the (photographic) camera. Active sensors have their own source of energy. Measurements by active sensors are more controlled because they do not depend upon the (varying) illumination conditions. Active sensors include the laser altimeter (using infrared light) and radar. Figure 3 gives an overview of the types of the sensors that are introduced in this section.  Passive sensors Gamma-ray spectrometer: The gamma-ray spectrometer measures the amount of gamma rays emitted by the upper soil or rock layers due to radioactive decay, its main application is found in mineral exploration. Gamma rays have a very short wavelength on the order of picometres (pm)). Because of large atmospheric absorption of these waves this type of energy can only be measured up to a few hundred metres above the Earth’s surface. Aerial camera The camera system (lens and film) is mostly found in aircraft for aerial photography. Low orbiting satellites and NASA Space Shuttle missions also apply conventional camera techniques. The film types used in the camera enable electromagnetic energy in the range between 400 nm and 900 nm to be recorded. Its principal applications include medium and large scale(topographic) mapping and cadastral mapping. Video camera Video cameras are sometimes used to record image data. Most video sensors are only sensitive to the visible colours, although a few are able to record the near infrared part of the spectrum. Mostly, video images serve to provide low cost image data for qualitative purposes, for example, to provide additional visual information about an area captured with another sensor (e.g., laser scanner or radar). Multispectral scanner The multispectral scanner is an instrument that mainly measures the reflected sunlight in the optical domain. A scanner systematically ‘scans’ the Earth’s surface thereby measuring the energy reflected from the viewed area. This is done simultaneously for several avelength bands, hence the name ‘multispectral scanner’. Imaging spectrometer The principle of the imaging spectrometer is similar to that of the multispectral scanner, except that spectrometers measure only very narrow (5–10 nm) spectral bands. Imaging spectrometer data, therefore, can be used to determine the surface's mineral composition or the surface water's chlorophyll content. Thermal scanner Thermal scanners measure thermal data in the range of 10–14 μm. Wavelengths in this range are directly related to the temperature of an object. Cloud, land and sea surface temperature data are extremely useful for weather forecasting. For this reason, most remote sensing systems designed for meteorology include a thermal scanner. Thermal scanners can also be used to study the effects of drought (‘water stress’) on agricultural crops, and to monitor the temperature of cooling water discharged from thermal power plants. Radiometer EM energy with very long wavelengths (1–100 cm) is emitted from the soil and rocks on, or just below, the Earth's surface. The depth from which this energy is emitted depends on the properties, such as water content, of the specific material. Radiometers are used to detect this energy. The resulting data can be used in mineral exploration, soil mapping and soil moisture estimation Laser scanner Active sensors Laser scanners are mounted on aircraft and use a laser beam (infrared light) to measure the distance from the aircraft to points located on the ground. This distance measurement is then combined with exact information on the aircraft’s position to calculate the terrain elevation. Laser scanning is mainly used to produce detailed, high-resolution, Digital Terrain Models (DTM) for topographic mapping the production of detailed 3D models of city buildings, and for measuring tree heights in forestry. Radar altimeter Radar altimeters are used to measure the topographic profile parallel to the satellite orbit. They provide profiles (single lines of measurements) rather than ‘image’ data. Radar altimeters operate in the 1–6 cm domain and are able to determine height with a precision of 2–4 cm. Radar altimeters are useful for measuring relatively smooth surfaces such as oceans and for ‘small scale’ mapping of continental terrain models. Imaging radar Radar instruments operate in the 1–100 cm domain. Different wavelength bands are related to particular characteristics of the Earth’s surface. The radar backscatter is influenced by the illuminating signal and the illuminated surface characteristics (orientation, roughness, di-electric constant/moisture content). Since radar is an active sensor system and the applied wavelengths are able to penetrate clouds, it has ‘all-weather day- and-night’ acquisition capability. The combination of two radar images of the same area can provide information about terrain heights. Combining two radar images acquired at different moments can be used to precisely assess changes in height or vertical deformations (SAR Interferometry).  Platforms In remote sensing, the sensor is mounted on a platform. In general, remote sensing sensors are attached to moving platforms such as aircraft and satellites. Static platforms are occasionally used in an experimental context. Airborne observations are carried out using aircraft with specific modifications to carry sensors. Sometimes Ultra Light Vehicles (ULVs), balloons, Airship or kites are used for airborne remote sensing. Airborne observations are possible from 100 m up to 30–40 km height. For spaceborne remote sensing, satellites are used. Satellites are launched into space with rockets. Satellites for Earth Observation are positioned in orbits between 150–36,000 km altitude. The specific orbit depends on the objectives of the mission, e.g., continuous observation of large areas or detailed observation of smaller areas. Different types of orbits are required to achieve continuous monitoring (meteorology), global mapping (land cover mapping), or selective imaging (urban areas). For remote sensing purposes, the following orbit characteristics are relevant : altitude, which is the distance (in km) from the satellite to the mean surface level of the Earth. Typically, remote sensing satellites orbit either at 600– 800 km (polar orbit) or at 36,000 km (geo-stationary orbit) distance from the Earth. The distance influences to a large extent which area is viewed and at which detail. inclination angle, which is the angle (in degrees) between the orbit and the equator. The inclination angle of the orbit determines, together with field of view of the sensor, which latitudes can be observed.  Period is the time (in minutes) required to complete one full orbit. A polar satellite orbits at 800 km altitude and has a period of 90 minutes. A ground speed of 28,000 km/hour is almost 8 km/s. Compare this figure with the speed of an aircraft, which is around 400 km/hour. The speed of the platform has implications for the type of images that can be acquired (time for ‘exposure’). repeat cycle is the time (in days) between two successive identical orbits. The revisit time, the time between two subsequent images of the same area, is determined by the repeat cycle together with the pointing capability of the sensor. Pointing capability refers to the possibility of the sensor-platform to ‘look’ sideways. The following orbit types are most common for remote sensing missions: Polar, or near polar, orbit. These are orbits with inclination angle between 80 and 100 degrees and enable observation of the whole globe. The satellite is typically placed in orbit at 600–800 km altitude. Sun-synchronous orbit. An orbit chosen in such a way that the satellite always passes overhead at the same local solar time is called sun-synchronous. Most sun synchronous orbits cross the equator at mid-morning (around 10:30 h). At that moment the Sun angle is low and the resultant shadows reveal terrain relief. Sun- synchronous orbits allow a satellite to record images at two fixed times during one 24-hour period: one during the day and one at night. Examples of near polar sun- synchronous satellites are Landsat, SPOT and IRS. Geostationary orbit. This refers to orbits in which the satellite is placed above the equator (inclination angle is 0 degree) at a distance of some 36,000 km. At this distance, the period of the satellite is equal to the period of the Earth. The result is that the satellite is at a fixed position relative to the Earth. Geostationary orbits are used for meteorological and telecommunication satellites. Today’s meteorological weather satellite systems use a combination of geostationary satellites and polar orbiters. The geo-stationary satellites offer a continuous view, while the polar orbiters offer a higher resolution. The data of spaceborne sensors need to be sent to the ground for further analysis and processing. All Earth Observation satellites apply satellite communication technology for downlink of the data. If the satellite is outside the range of a receiving station the data can be temporarily stored by a tape recorder in the satellite and transmitted later.  Principle of Remote Sensing An electromagnetic signal is recorded by a detector after it interacts with a target containing molecules, particulates, and/or surfaces. If T and S denote the target and signal, respectively, then we may write symbolically 𝑆 = 𝐹(𝑇) The inverse gives 𝑇 = 𝐹 −1 (𝑆) The fundamental obstacle in all remote sensing inversion problems is the uniqueness of the solution. The nonuniqueness arises because the medium under investigation may be composed of a number of unknown parameters, a combination of which may lead to the same radiation signature.  Remote Sensing Using Transmitted Sunlight The remote sensing of aerosols and ozone from ground- based radiometers utilizes the direct beam of solar radiation transmitted through the cloudless atmosphere. The solar intensity of a given wavelength measured at the ground at a given time of the year can be written in the form 𝑟0 2 𝐼መ𝜆 = 𝐼⊙ 𝜆 exp[−𝜏 𝜆 𝑚(𝜃0 )] 𝑟 Where, 𝑟 and 𝑟0 are the actual and mean sun-earth distance 𝐼⊙ is the solar intensity at TOA 𝑚 𝜃0 = 1/ cos 𝜃0 is the mass factor 𝜃0 is the solar zenith angle Geometry of a ground-based radiometer for the measurement of a The total optical depth is the sum of the individual direct solar beam. A sun photometer can have a number of different values contributed from aerosols (A), Rayleigh filters to record direct solar radiation in component wavelengths molecules (R), ozone (3) and nitrogen dioxide (2), respectively, as follows: 𝜏 𝜆 = 𝜏𝐴 𝜆 + 𝜏𝑅 𝜆 + 𝜏3 𝜆 + 𝜏2 𝜆 Determination of Aerosol Optical Depth and Size Distribution Observational methods to determine the dust loading of the atmosphere were developed during the 1920s by Linke and Angstrom. In essence, aerosol total optical depth, sometimes also referred to as turbidity, is derived from direct spectral solar intensity measured on the ground. Wavelengths in the visible spectrum are normally employed because absorption due to water vapor can be neglected and the effect of ozone absorption is small. 𝜏𝐴 𝜆 = 𝜏 𝜆 − 𝜏𝑅 𝜆 − 𝜏3 𝜆 − 𝜏2 𝜆 𝑟0 2 We have, 𝐼መ𝜆 = 𝐼⊙ 𝜆 exp −𝜏 𝜆 𝑚 𝜃0 , 𝑟 1 𝑟0 2 give 𝜏 𝜆 = ln 𝐼 ∗ − ln 𝐼 , where 𝐼∗ = 𝐼⊙ 𝑚 𝜃0 𝑟 In the determination of aerosol optical depth, the optical depths associated with NO2 and O3 molecules are usually computed from the following parameterizations: 𝑒 − 𝑐+𝑑𝜆+𝜆 𝜏𝑅 𝜆 = 𝑎 + 𝑏𝐻 𝜆 𝑝/𝑝𝑠 , 𝜏2 𝜆 = 𝑘2 𝜆 𝐶 𝑁𝑂2 , 𝜏3 𝜆 = 𝑘3 𝜆 𝐶 𝑂3 Where, 𝑎 = 0.00864, 𝑏 = 6.5 × 10−6 , 𝑐 = 3.916, 𝑑 = 0.074, and 𝑒 = 0.050, H is the radiometer height (in km), 𝑝 is the atmospheric pressure at the radiometer, 𝑝𝑠 is 1013.25 mb, 𝑘2 is the absorption coefficient of NO2 and 𝑘3 is the absorption coefficient of O3, C is the concentration Further, some diffuse light that enters the radiometer can be taken into account by empirical adjustments. Thus, once these optical depths have been evaluated and the position of the sun is known, a measurement of the direct solar intensity can be used to infer the aerosol optical depth. In addition to the optical depth, the aerosol size distribution can also be retrieved. We shall first demonstrate the principle of using two wavelengths for this purpose. The aerosol optical depth corresponding to the entire atmospheric column can be expressed in terms of the extinction coefficient in the form 𝑧∞ 𝜏𝐴 𝜆 = න 𝛽𝑒 𝜆, 𝑧 𝑑𝑧 0 Let n(z,a) is the aerosol height-dependent size distribution. Then the extinction coefficient is 𝑎2 𝛽𝑒 𝜆, 𝑧 = න 𝜎𝑒 𝑎, 𝜆 𝑛 𝑧, 𝑎 𝑑𝑎 𝑎1 Where 𝜎𝑒 represents the extinction coefficient for the individual particles. The aerosol size distribution can ∗ be defined by the Junge distribution in the form𝑛 𝑧, 𝑎 = 𝐶(𝑧, 𝑎)𝑎−(𝑣 +1) , where C is scaling function and 2 ≤ 𝑣 ∗ ≤ 4, the size of aerosol can typically cover the ranges from 0.01 to 10 micro-meter.  ∗ Using the Junge size distribution, the aerosol optical depth is given by 𝜏𝐴 𝜆 = 𝑘𝜆(−𝑣 +2) Where, 𝑘 is constant. When 𝑣 ∗ = 3.3, 𝑘 is known as Angstrom turbidity coefficient. If the aerosol optical depth is measured at two wavelengths, then we have (−𝑣 ∗ +2) 𝜏𝐴 (𝜆1 ) 𝜆1 (−𝑣 ∗ +2) 𝑧Ƹ = = = 𝑦 𝜏𝐴 (𝜆2 ) 𝜆2 Which give the shaping factor as 𝑣 ∗ = 2 − ln 𝑧Ƹ /𝑦 We may also retrieve the aerosol size distribution directly if a number of aerosol optical depths are measured with a multispectral instrument known as a sunphotometer A sunphotometer tracks the sun and measures the intensity of a solar beam in several spectral channels ∞ Let us define the column aerosol size distribution in the form 𝑛𝑐 𝑎 = ‫׬‬0 𝑛 𝑎, 𝑧 𝑑𝑧 = 𝑓 𝑎 ℎ(𝑎) ∗ where ℎ 𝑎 = 𝑎−(𝑣 +1) and 𝑓 𝑎 is a slowly varying function. 𝑎 𝑎 𝜎𝑒 So, 𝜏𝐴 𝜆 = ‫ 𝑎׬‬2 𝑓 𝑎 ℎ 𝑎 𝜋𝑎2 𝑄𝑒 𝑚, 𝑑𝑎, since = 𝑄𝑒 according to Mie-theory. 1 𝜆 𝜋𝑎2 𝑎 For inversion method let 𝑔 = 𝜏𝐴 𝜆 and 𝐾𝜆 𝑎 = ℎ 𝑎 𝜋𝑎2 𝑄𝑒 𝑚, , Thus 𝜆 𝑎2 𝑔𝜆 = න 𝑓 𝑎 𝐾𝜆 𝑎 𝑑𝑎 𝑎1 This is the well-known Fredholm equation of the first kind in which 𝐾𝜆 𝑎 , the weighting function, is the kernel, and 𝑓 𝑎 is the function to be recovered from a set of 𝑔𝜆.  Determination of Total Ozone Concentration Dobson (1957) proposed a method for the estimation of total ozone concentration from a ground-based instrument. This method uses the Beer–Bouguer–Lambert law 𝑟0 2 𝐼መ𝜆 = 𝐼⊙ 𝜆 exp[−𝜏 𝜆 𝑚(𝜃0 )] 𝑟 for the transfer of UV radiation to retrieve ozone concentration. For many layers with different optical path 𝑟0 2 length 𝐼መ𝜆 = 𝐼⊙ 𝜆 exp[− σ𝑖 𝜏𝑖 (𝜆)𝑚𝑖 (𝜃0 )] 𝑟 The solar zenith angle Z is in reference to the height of about 20 km corresponding to the maximum ozone concentration. Let us define the total ozone concentration in the vertical column as Let, the absorption coefficient be denoted by 𝑘 𝜆 , and select a pair of wavelengths (𝜆1 , 𝜆2 ) in ozone absorption bands such that 𝑘(𝜆1 ) > 𝑘(𝜆2 ). Thus, መ 𝑖) 𝐼(𝜆 log10 𝐼መ = −𝑘 ∗ 𝜆𝑖 Ω sec 𝑍 − 𝜏𝐴∗ 𝜆𝑖 𝑚 − 𝜏𝑅∗ 𝜆𝑖 𝑚, where 𝑖 = 1,2. ⊙ (𝜆𝑖 ) This gives 1 መ 𝐼(𝜆 ) 𝐼 ⊙ 𝜆1 𝑁 = log10 𝐼(𝜆 መ − log 10 = −Ω sec 𝑍 ∆𝑘 − 𝑚∆𝜏𝐴 − 𝑚∆𝜏𝑅 ) 2 𝐼 ⊙ 𝜆2 ∗ ∗ Where ∆𝑘 = 𝑘 ∗ 𝜆1 − 𝑘 ∗ 𝜆2 and ∆𝜏𝐴,𝑅 = 𝜏𝐴,𝑅 𝜆1 − 𝜏𝐴,𝑅 𝜆2 Two pairs of wavelengths have been selected to minimize the aerosol effect because ∆𝜏𝐴 is the most uncertain term due to aerosol scattering. In the standard procedure developed by the World Meteorological Organization (WMO), these pairs are (0.3055, 0.3254 μm) and (0.3176, 0.3398 μm). An instrument using these pairs is referred to as a Dobson spectrometer. This gives 𝑁(1) −𝑁(2) Ω= − 𝑏, 𝑎 sec 𝑍 where the parameterized coefficients a and b are determined from known ozone absorption coefficients and the Rayleigh scattering theory in the forms 𝑎 = ∆𝑘 (2) − ∆𝑘 (1) = 1.388 (atm-1cm-1) and b=0.009ps The superscripts (1) and (2) denote the two pairs of wavelengths and ps is the surface pressure in units of atmospheres. Total ozone concentration has traditionally been measured in milli atm-cm, called Dobson units (DU). A DU is a vertical thickness of atmosphere in thousandths of a centimeter that is occupied by O3 when concentrated into a uniform layer of pure gas at the standard temperature and pressure. The total column ozone concentration normally ranges from 200 to 400 DU. Total ozone has been measured by the Dobson spectrometer at some 80 ground stations around the world Limb Extinction Technique Analogous to ground-based sunphotometer measurements of transmitted sunlight, measurements can also be made from space by scanning the sun’s disk as the sun rises and sets relative to the motion of the spacecraft, referred to as solar occultation. let the intensity measured at the center of the scan be denoted by 𝐼⊙ , and the intensity determined from a scan at a lower altitude ℎ𝑖 be 𝐼(ℎ𝑖 ) The transmittance in this case is given by 𝐼 𝜆, ℎ𝑖 𝑇 𝜆 = = exp −𝜏 𝜆, ℎ𝑖 , 𝐼⊙ 𝜆 Where the optical path is defined by the tangent height ℎ𝑖 as follows ∞ 𝐼 𝜆, ℎ𝑖 𝜏 𝜆, ℎ𝑖 = ln = න 𝛽𝑒 𝜆, 𝑥 𝑑𝑥 𝐼⊙ 𝜆 −∞ where the extinction coefficient is generally contributed by the extinction (scattering and absorption) of aerosols, Rayleigh molecules, O3, and NO2 𝛽𝑒 𝜆, 𝑥 = 𝛽𝑒,𝐴 𝜆, 𝑥 + 𝛽𝑒,𝑅 𝜆, 𝑥 + 𝛽𝑒,3 𝜆, 𝑥 + 𝛽𝑒,2 𝜆, 𝑥 We may divide the atmosphere into an appropriate number of layers (e.g., 80) so that 𝜏 𝜆, ℎ𝑖 = 2 ෍ 𝛽𝑒 𝜆, 𝑧𝑗 ∆𝑥𝑖𝑗 𝑗 where ∆𝑥𝑖𝑗 is the path length in the j’th layer represented by 𝑧𝑗 associated with the direct solar beam passing through the tangent height ℎ𝑖 The limb extinction technique is specifically useful for the determination of aerosols and other minor gases in the stratosphere. It was explored by the Stratospheric Aerosol Measurement (SAM) experiment aboard the Nimbus 7 satellite in 1978. Subsequently, the Stratospheric Aerosol and Gas Experiment I (SAGE I) and SAGE II were conducted in 1979 and 1984, respectively. SAGE II contains seven channels centered at 0.385, 0.448, 0.453, 0.525, 0.6, 0.94, and 1.02 μm. The 0.94 and 0.6 μm channels were used to infer the water vapor and ozone amounts, while the difference between 0.448 and 0.453 μm was used to determine the NO2 concentration. The solar occultation technique is extremely sensitive to the presence of high level clouds. The cirrus results derived from SAGE have complemented those from other satellite remote sensing methods, particularly in view of the fact that it is difficult to determine thin cirrus with optical depths less than about 0.5 based on reflected sunlight and/or emitted infrared radiation from nadir looking radiometers  Remote Sensing Using Reflected Sunlight Relation of scattering Θ, zenith (θ, θ’), and azimuthal angles (φ, φ’) in a spherical atmosphere. In the discussion of multiple scattering of a light beam, the notation is usually used to denote the scattering angle, while θ is used for the emergent angles.  The sunlight (denoted as In) reflected from a position-vector on the earth and the atmosphere is detected by a satellite (denoted as Out). The position of the sun is defined by the solar zenith angle θ0, while the satellite position is defined by the emergent zenith angle θ. The relative positions of the sun and the satellite are given by the azimuthal angle difference Δ𝜙. From spherical geometry the angle between the incoming and outgoing light beams counterclockwise from the incoming beam, known as the scattering angle , is defined by Any radiometer on board a satellite will have a finite field-of-view on a horizontal plane that collects radiation from the earth and the atmosphere, referred to as resolution, which is dependent on the specific instrument designed and its scan angle.  Any radiometer on board a satellite will have a finite field-of-view on a horizontal plane that collects radiation from the earth and the atmosphere, referred to as resolution, which is dependent on the specific instrument designed and its scan angle. The earth makes one complete resolution about the sun (2π radian) in one tropical year (365.2422 days). Thus, the right ascension of the sun changes at an average rate of about 1°𝑑𝑎𝑦 −1. If the inclination of the satellite is correctly chosen, the right ascension of its ascending node can be made to precess at the same rate. An orbit that is synchronized with the sun is called a sunsynchronous orbit or polar orbit. The point at which a satellite crosses the earth’s equatorial plane from south to north is known as the ascending node; The point passed as it crosses the plane from north to south is known as the descending node. The point directly beneath the satellite is called the subsatellite point or nadir. For a satellite with a height of about 870 km, its inclination angle needs to be about 99° for its orbit to be sunsynchronous.  Morning satellites ascend (or descend) between 06 and 12 h LST and descend (or ascend) between 18 and 24 h LST. Afternoon satellites ascend (or descend) between 12 and 18 h LST, and descend (or ascend) between 00 and 06 h LST. Most meteorological satellite instruments are designed such that the area viewed on one orbit touches or overlaps the area viewed on previous and successive orbits. If a satellite is moved farther from the earth, it would experience a weaker gravitational field and the centrifugal acceleration required to keep the satellite in orbit would be smaller. The rotation period of the satellite would also become longer. It is possible to select a distance at which the rotation period is exactly equal to the 1 day rotation period of the earth in such a manner that the satellite moves in a counterclockwise fashion. This orbit is called a geosynchronous or geostationary orbit. To achieve this distance, the satellite must be about 36,000 km above the earth’s surface Geostationary satellites remain essentially stationary above a point on the equator and are classified by the longitude of their subsatellite points. Presently, five geostationary satellites orbit the earth to gather weather data, including two operated by the United States, one by the European Space Agency (METEOSAT), one by Japan (GMS), and one by India (INSAT). The current U.S. system consists of two Geostationary Operational Environmental Satellites (GOES).  Satellite Remote Sensing of Ozone The basic principle involved in the estimation of ozone concentration utilizing reflected sunlight is to select a pair of wavelengths in the ozone absorption band. Wavelengths near the long-wavelength end of the band at which absorption is relatively weak are chosen so that most of the photons reaching the satellite instrument have passed through the ozone layer and backscattered from within the troposphere. The two wavelengths are separated by about 200 A˚ so that the scattering effect is about the same at each wavelength, while absorption for one of these wavelengths is stronger than the other. A pair such as (3125, 3312 A), for example, has been selected in the Nimbus 4 satellite experiment. If Z denotes the solar zenith angle at the level of maximum ozone concentration (about 20 km) at the subsatellite point, then the total attenuation path of the backscattered photons through the ozone layer is proportional to 1 + sec Z. Let 𝐹⊙ and I be the incident solar irradiance and the measured backscattered intensity at TOA, respectively.  መ 1) 𝐼(𝜆 𝐼⊙ 𝜆1 𝑁= log10 መ − log10 Can be written as 𝐼(𝜆2 ) 𝐼⊙ 𝜆2 ෡ 𝜆1 , 𝜆2 = log10 𝐹⊙ (𝜆1) − log10 𝐹⊙ (𝜆2). 𝑁 መ 𝐼(𝜆1 ) መ 𝐼(𝜆2 ) Determination of the total ozone concentration can be made by comparing the observed 𝑁 ෡ with values computed for a series of different standard ozone profile by means of the searching method. Three basic methods are as follows: (a) A set of table containing the computed quantities 𝐼 Ω, 𝜇0 , 0 , 𝑇 Ω, 𝜇0 and 𝑟(Ω) ҧ for different values of 𝜇0 and Ω are prepared a priori. (b) The effective surface albedo is determined by utilizing the radiometric measurement at a wavelength outside the ozone absorption band, say 𝜆3 (3800 𝐴° ). Surface albedo is measured 𝐼መ3 𝜇0 ,𝑟𝑠 −𝐼3 (𝜇0 ,0) with 𝑟𝑠 𝜆3 = The assumption is made that rs is independent of 𝑇 𝜇0 +𝑟[ҧ 𝐼መ3 𝜇0 ,𝑟𝑠 −𝐼3 (𝜇0 ,0)] wavelength so that it can be used for the pair of wavelengths (λ1, λ2). However, an empirical adjustment may also be performed from known surface albedo measurements. (c) With the surface albedo known, computations are then carried out to generate N(λ1, λ2) versus ෡ can then total ozone concentration. Best estimates of from the observed intensities via 𝑁 be made by an optimized search method. The Backscatter Ultraviolet Spectrometer (BUV) on board Nimbus 4 operated for seven years, from 1970 to 1977. The Total Ozone Mapping Spectrometer (TOMS) on Nimbus 7 produced daily global maps of total ozone at 50–150 km resolution in 1978. The Solar Backscatter Ultraviolet Radiometer (SBUV) on the NOAA satellites has provided daily ozone data over the globe on a routine basis since 1980.  Satellite Remote Sensing of Land Surfaces The remote sensing of vegetation and surface properties by satellites is a subject that is of interest to numerous disciplines, including meteorology, hydrology, geography, geology, biology, ecology, and electrical engineering. Usually, this is done by measuring the surface albedo with various techniques For water surfaces, the albedo ranges from about 6 to 9%, except for cases involving the low solar angle that is associated with the high latitudes of the winter hemisphere. The albedo can range from 10 to 40% for various land surfaces. For example, deserts and sand dunes have albedos of about 30–40%, whereas those for meadows and forests are about 10%. The albedos of snow and ice are greater than 40%. The albedos of some vegetation surfaces vary greatly with solar wavelength as noted below. If we select a wavelength in the visible spectrum (0.4–0.7 μm) such that the product of the surface albedo 𝑟𝑠 and the atmospheric bidirectional reflectance distribution function BRDF (generally less than 0.1) is much smaller than 1, then we can express the surface albedo in the form 𝑟𝑠 ≅ 𝑎𝑅෠ − 𝑏 The coefficients a and b are the so-called atmospheric correction terms involving the scattering contributions of aerosols and molecules, which can be empirically determined for specific applications. Thus, a measurement of BRDF 𝑅෠ at TOA determined by the sun–satellite geometry will provide a surface albedo value. Illustration of the step function transition in the surface spectral albedo of two types of vegetation at 0.7 μm, compared to the continuous spectral albedo for soil surfaces. These results were derived from measurements made in the Southern Great Plains of Oklahoma surrounding the Central Facility of the Department of Energy’s Atmospheric Radiation Measurement site (data taken from Li et al., 2002)  Remote Sensing Using Emitted Infrared Radiation In a non-scattering atmosphere that is in local thermodynamic equilibrium, the basic equation that governs the transfer of emitted thermal infrared (IR) radiance at a given wavenumber, ν, can be described by The optical depth is defined by with ρa the density of the absorbing gases and kν the absorption coefficient. The solution for upward radiance is given by the following integral equation: where τ∗ is the optical depth at the surface and Iν(τ∗) denotes the emitted surface radiance generally assumed to be isotropic. The first term on the right-hand side represents the surface emission contribution attenuated to the level τ, while the second term denotes the emission contribution of the atmosphere between τ and τ∗. For upwelling direction 𝜇 = 1. The emitted radiance at the surface Iν(τ∗) = εν Bν(Ts), where εν is the surface emissivity and Ts is the surface temperature. Moreover, in remote sensing the exponential terms are generally expressed in terms of the transmittance and weighting function defined in the following. The monochromatic transmittance is defined by For application to atmospheric remote sensing, the height or pressure coordinate is usually employed. Height and pressure are related via the hydrostatic equation, dp = -ρgdz, where ρ is the air density and g is the gravitational acceleration. The mixing ratio for a specific gas with density ρa is defined as q = ρa/ρ. An instrument can distinguish only a finite band width (Ψ(𝜈,ҧ 𝜈)), where Ψ denotes the instrumental response (or slit) function and 𝜈ҧ is the mean wavenumber. The measured radiance from a spectrometer over a wavenumber interval (ν1, ν2) in the normalization form is given by The effective spectral interval for the response function is usually small enough that the variation of the Planck function is insignificant. We can then replace its value by 𝐵𝑣ത (𝑇) without introducing noticeable errors. Where, when the instrumental response function is accounted for, the spectral transmittance is defined by Equation above for I(0) is fundamental to remote sensing of the atmosphere and the surface from orbiting meteorological satellites.  The potential information content in the thermal IR spectrum There are four regions over which water vapor, ozone, and carbon dioxide exhibit a significant absorption spectrum. Carbon dioxide absorbs IR radiation in the 15 μm band from about 600 to 800 cm-1. In addition, carbon dioxide also absorbs radiation in the 4.3 μm region that overlaps with solar radiation. Absorption due to ozone is primarily confined to the 9.6 μm band. Water vapor exhibits absorption lines over the entire infrared spectrum. The most pronounced absorption occurs in the 6.3 μm vibrational–rotational band and in the pure rotational band with wavenumbers less than about 500 cm-1. From about 800 to 1200 cm-1, referred to as the atmospheric Observed infrared spectrum displaying all the absorption window, absorption due to atmospheric gases shows a gases and their spectral location. This spectrum was obtained minimum, except in the 9.6 μm ozone band. from the scanning high-resolution interferometer sounder (S- There are also absorption bands for various greenhouse gases HIS), which measured the emitted thermal radiation between 3.3 and 18 μm, onboard the NASA ER-2 aircraft over the that can be used for their determination by remote sensing: Gulf of Mexico southeast of Louisiana on April 1, 2001 the CH4 7.6 μm band, the N2O 7.9 μm band, and some CFC lines in the window.  Surface Temperature Determination If observations are taken in the window region where the effect of the atmosphere is at a minimum, the upwelling radiance at TOA must be closely associated with emission from the surface. The above equation can be written as Where we have used 𝑖 instead of 𝜈,ҧ and defined average temperature 𝑇𝑎. The value of Ta generally varies by less than 1 K in the window region from about 10.5 to 12.5 μm, in which variability of the surface emissivity is insignificant. The atmospheric transmittance in the window region is primarily produced by the continuous absorption of water vapor and to a good approximation is given by where ki is the absorption coefficient of water vapor for a spectral band in the window and u is the water vapor path length. Where we have also expressed the observed radiances I1 and I2 in terms of the brightness temperatures Tb1 and Tb2. The objective of the split-window technique is to eliminate Ta, and to do so, we may expand the Planck function of temperature T by means of the Taylor series with respect to Ta in the form Applying this equation to the two channels with i = 1, 2 and eliminating (T - Ta) yields We can also have And, Sea surface temperatures (SSTs) have been routinely inferred from the operational NOAA satellites using one of the instruments aboard, the AVHRR. The 10.9 and 12.0 μm channels have been used to determine surface temperature, particularly over oceans. At night, the 3.7 μm channel has also been added to increase the accuracy of retrieval. Based on the split- window technique, surface temperature may be expressed by a general form of regression as follows: where a, b, and c are empirical coefficents derived from in situ observations obtained from drifting buoys, and Tb1 and Tb2 are the brightness temperatures involving a combination of AVHRR 10.9, 12.0, and 3.7 μm channels. The satellite-derived temperatures correspond to the temperature of a surface (skin temperature), whereas the buoy measurements are associated with a layer of water some meters deep. The regression approach of the multichannel technique is used partly to adjust (or tune) the satellite skin temperatures to the in situ bulk temperatures and partly to account for the water vapor absorption in the window, particularly in the moist tropical region. Since 1970, global SSTs have been operationally produced and archived. An important part of SST retrieval is the detection and elimination of clouds. Several threshold methods are employed for these purposes based on the bidirectional reflectance of the solar channel in daytime and the emission characteristics of IR channels with respect to the sea surface uniformity in nighttime. The split-window technique has also been used for the determination of surface temperature over land in conjunction with land–atmosphere interaction studies.