METEOR 4100 - Introduction.pdf

METEOR 4100 Tropical Meteorology (Introduction) Jophet D. Flores Instructor  Outline What is tropical meteorology? Temperature Energy and the global climate Seasonal and Geographic Dis...

METEOR 4100 Tropical Meteorology (Introduction) Jophet D. Flores Instructor  Outline What is tropical meteorology? Temperature Energy and the global climate Seasonal and Geographic Distribution of Temperature Defining the tropics Major Influence of Annual Surface Energy balance and the role of tropics Temperature Distribution Surface Energy Budget Diurnal Temperature Variability in the Tropics Meridional Energy Transport by Atmosphere and Ocean Moisture and precipitation Latent Heat and Deep Convective Cloud Distribution Role of tropics in the momentum balance Surface-Air Interactions Spatial and Temporal Scales in the Tropics Atmospheric structure Tropical air masses and climates Temperature Profiles The Trade Wind Inversion Atmospheric Humidity Pressure Ranges What is tropical meteorology? Tropical meteorology is the study of the tropical atmosphere including thunderstorms and lightning, tropical cyclones (hurricanes, typhoons), monsoons, dust storms, El Niños, squall lines, equatorial Rossby waves, Madden-Julian Oscillation events, trade wind inversions, easterly jets, snow and ice, and much, much more. What is tropical meteorology? Driving these events are important energy sources and sinks like surplus radiation, latent heat, sensible heat, evapotranspiration, and ocean heat storage. What is tropical meteorology? The meteorology of the tropics is different from the meteorology of higher latitudes in various ways. The Coriolis force is weak or non-existent and pressure gradients are very weak (except in tropical cyclones). Temperature contrasts are minimal so that air masses are fairly homogeneous and weather disturbances are initiated by modest differences in wind velocity gradients or heating. In contrast, midlatitude weather is dominated by synoptic cyclones that form in response to strong gradients of air temperature and density. What is tropical meteorology? The mid-latitude cyclone is a synoptic scale low pressure system that has cyclonic (counter-clockwise in northern hemisphere) flow that is found in the middle latitudes (i.e., 30° N - 55° N) IT IS NOT A HURRICANE OR TROPICAL STORM There is a location (tropics vs. mid-latitudes) and size difference between hurricane and mid- latitude cyclone Typical size of mid-latitude cyclone = 1500-5000 km in diameter Typical size of a hurricane or tropical storm = 200- 1000 km in diameter Energy and the Global Climate The sun is the primary source of energy for the earth system (atmosphere, hydrosphere, biosphere, cryosphere, and lithosphere). The earth systems are constantly exchanging matter and energy through physical and bio-geochemical cycles. The physical cycles, including atmospheric and oceanic motion, are driven by solar energy which is why we need to understand the distribution of solar energy around the globe. Energy and the Global Climate A fundamental principle of meteorology is that of the principle of a conserved quantity, one that is conserved except for external sources and sinks. The first law of thermodynamics states that energy is conserved. For a closed system, any heat added or removed must be equal to the change in internal energy plus the work done, which can be expressed as: 𝑑𝑄 = 𝑑𝑈 + 𝑑𝑊 (1) where 𝑑𝑄 is the amount of heat added or removed, 𝑑𝑈 is the change in internal energy (stored energy), and 𝑑𝑊 is the work done (energy used to do work). Energy and the Global Climate Energy is transferred by: radiation (no mass exchange, no medium required, radiation moves at the speed of light); conduction (no mass exchanged, heat transferred by vibration and collision among atoms and molecules); and convection (mass exchanged, fluid parcels with different amounts of energy change places, the net movement of mass is not necessary for energy to be transferred). Energy and the Global Climate Energy transfer from the sun to the earth is nearly all by radiation (some negligible mass is associated with the solar wind). The figure shows the latitudinal distribution of incoming solar radiation and its effect on energy density received at the surface. The maximum at the equator and minimum at the poles occur because of the solar beam spreading over a greater area at higher latitudes and attenuation, or beam depletion, by the atmosphere. Energy and the Global Climate Since most of our energy comes from solar radiation, it is helpful to briefly review two basic laws that govern electromagnetic radiation (Stefan Boltzmann Law and Wien’s Law). Stefan Boltzmann Law relates the energy emitted per unit area (from all wavelengths) to the absolute temperature (K) by, 𝐸 = 𝜎𝑇 4 (2) where the Stefan-Boltzmann constant, 𝜎 = 5.67 × 10−8 𝑊𝑚−2 𝐾 −4 Energy and the Global Climate Wien’s Law tells us that the wavelength, 𝜆𝑚𝑎𝑥 (𝜇𝑚), of maximum blackbody emission is inversely proportional to its absolute temperature, Τ(𝐾). The hotter the object, the shorter the peak wavelength at which it emits. 2897.9 𝜆𝑚𝑎𝑥 = (3) 𝑇 Energy and the Global Climate Therefore, incoming solar radiation has more energy per unit area and a shorter peak wavelength than radiation from the earth-atmosphere system, as depicted in the figure. The peak of the solar radiation is at ~0.5 𝜇𝑚 (shortwave) in the visible wavelength range while the terrestrial radiation peak is ~10 𝜇𝑚 (longwave) in the infrared range. Energy and the Global Climate However, it is not the incoming solar energy or insolation that determines climate but the net radiation, the balance between the incoming and outgoing radiation from the earth-atmosphere system. The annually-averaged energy balance at the top of the atmosphere is 𝐹𝑠𝑤 1 − 𝛼𝑝 = 𝜀𝜎𝑇𝑒4 (4) Incoming solar radiation (shortwave) = Outgoing terrestrial radiation (longwave) where FSW is the incoming solar radiation, αp is the planetary reflectivity or albedo, ε is the emissivity of the atmosphere, σ is the Stefan-Boltzmann constant, Te is the effective temperature (K), the temperature required to balance the solar energy absorbed Energy and the Global Climate The atmosphere is largely transparent to solar radiation; only about 20% is absorbed, whereas most of the solar radiation that reaches the earth's surface is absorbed. On average 70% of the solar radiation entering the top of the atmosphere is absorbed by either the atmosphere or the surface. The surface warms the atmosphere from below by emission of long-wave radiation, sensible heat (conduction and dry convection), and latent heat release through evaporation and moist convection. Of these processes, conduction contributes the least to warming the tropospheric column. Energy and the Global Climate The fluid motions in the earth system, primarily the ocean and atmosphere, act to compensate for the radiative imbalance between the warm equator and the cold poles. Thus, it is appropriate to consider fluid motion as a heat engine that helps to equilibrate the earth system. Energy and the Global Climate The heat engine of the earth system is driven by the tropical latitudes. Malkus (1962) refers to the tropics as the “firebox” of the heat engine. Surplus heating in the tropics and net cooling at the poles creates a horizontal temperature gradient in the atmosphere. The consequent horizontal pressure gradients set the atmosphere into motion. This relationship is depicted in the figure. Energy and the Global Climate Riehl and Malkus (1958) hypothesized that the upward transport of energy into the upper troposphere is concentrated in deep convective weather systems, such as the intermittent band of systems outlined by the yellow line in the figure. They estimated that between 1500 and 5000 deep cumuli were needed to balance the global heat budget. The upward motion in these convective cores, referred to as “hot towers”, is balanced by downward motion in the space between the clouds. Defining the Tropics The region of the earth known as the tropics straddles the equator. However, its latitudinal limits are defined variously as: The region where the angle of declination can be 90°; whose outer limits are identified as the Tropic of Cancer and the Tropic of Capricorn, ±23.5° latitude. Defining the Tropics The region of the earth known as the tropics straddles the equator. However, its latitudinal limits are defined variously as: The region of surplus radiation where annual solar input minus terrestrial output is positive, ± 35 to 40° latitude. Defining the Tropics The region of the earth known as the tropics straddles the equator. However, its latitudinal limits are defined variously as: The region of net upward motion and surface low pressure: positive net radiation sets air in motion leading to general upward motion and low pressure at the surface surrounded by sinking air and high pressure at the subtropics. This circulation is referred to as the Hadley cell in honor of George Hadley who, in 1735, proposed that excess radiation in the tropics would lead to upward motion and corresponding subsidence at the poles. Later studies showed that his circulation model was incomplete as it did not account for the midlatitude westerlies and the indirect circulations known as the Ferrel cells. Defining the Tropics The region of the earth known as the tropics straddles the equator. However, its latitudinal limits are defined variously as: The region in which winds blow primarily from the east (approximately ± 30° latitude), except for the regional monsoon. The easterly trade winds flow out of the subtropical high into the equatorial trough. They converge at the Intertropical Convergence Zone (ITCZ), which is usually identified as an intermittent band of clouds in the low pressure belt or equatorial trough. Defining the Tropics The region of the earth known as the tropics straddles the equator. However, its latitudinal limits are defined variously as: The region where the annual range of temperature is less than or equal to the average daily range. Defining the Tropics The region of the earth known as the tropics straddles the equator. However, its latitudinal limits are defined variously as: The region that is better described by a wet and dry season than the four seasons of higher latitudes because annual rainfall varies much more from place to place than annual temperature. Defining the Tropics Riehl (1979) described the tropics as “that part of the world where atmospheric processes differ decidedly and sufficiently from those in higher latitudes, so that one is justified writing a separate book on tropical weather and climate alone”. In this text, the tropics encompass the region of relatively low surface pressure located between high pressures belts in the subtropics. This definition emphasizes the dynamic nature of atmospheric circulations as a response, primarily, to solar heating of the earth, and, secondarily, to other factors such as surface properties. Energy Balance and the Role of the Tropics In order to maintain the energy balance in the earth-atmosphere system and within latitudinal zones, energy is transported by the ocean and the atmosphere. For example, differential heating of the atmosphere influences the temperature patterns, which lead to pressure gradients and winds driven by the pressure gradient force. The winds, in turn, advect air of differing temperatures. Energy Balance and the Role of the Tropics The large-scale oceanic dynamics are driven primarily by wind stress at the upper surface as well as net radiation and sensible heat fluxes, and density variations that are due to changes in salinity. Salinity changes are caused by processes such as evaporation, precipitation, and runoff. Heat is transported by convection, turbulent mixing, downwelling, or upwelling. Energy Balance and the Role of the Tropics (Surface Energy Budget) The net radiation at the surface can be written as: 𝑅𝑠 = 𝐹𝑠𝑤 1 − 𝛼𝑠 − 𝜀𝜎𝑇𝑠4 + 𝜀𝜎𝑇𝑎4 (5) where the subscripts s and a refer to the surface and atmosphere, respectively. Τ is the temperature (K), αs is the surface albedo. The signs indicate direction of energy transfer. The atmosphere is largely transparent to solar radiation while absorbing and emitting longwave radiation. The earth’s surface is warmer than it would be without an atmosphere because of solar radiation and downward longwave radiation from the atmosphere; a condition known as the “Greenhouse effect”. Energy Balance and the Role of the Tropics (Surface Energy Budget) The surface energy budget relates the net radiation to the sensible heat, potential energy, latent heat, kinetic energy, storage, and advection: 𝑅𝑠 = 𝑐𝑝 𝑇 + 𝐿𝑞 + 𝑔𝑧 + 𝑘 + 𝐺 + ∆𝑓 (6) 𝑁𝑒𝑡 𝑟𝑎𝑑𝑖𝑎𝑡𝑖𝑜𝑛 = 𝑠𝑒𝑛𝑠𝑖𝑏𝑙𝑒 + 𝑙𝑎𝑡𝑒𝑛𝑡 + 𝑝𝑜𝑡𝑒𝑛𝑡𝑖𝑎𝑙 + 𝑘𝑖𝑛𝑒𝑡𝑖𝑐 + 𝑠𝑡𝑜𝑟𝑎𝑔𝑒 + ℎ𝑜𝑟𝑖𝑧𝑜𝑛𝑡𝑎𝑙 𝑎𝑑𝑣𝑒𝑐𝑡𝑖𝑜𝑛 where Rs is the net radiation at the surface; cp is specific heat at constant pressure; Τ is temperature; L is latent heat of vaporization; q is the specific humidity; g is acceleration due to gravity; z is altitude; k is the atmospheric kinetic energy, ½ (υ2 + ν2) where υ, ν are horizontal wind components; G is the heat transferred in and out of storage (subsurface layers); and Δf is the horizontal flux or advection. Energy Balance and the Role of the Tropics (Surface Energy Budget) The first two terms in equation (6) Kind of Formula Amount x 1021 Energy Joules are the sensible and latent heat, respectively. Latent 𝐿𝑞 0.931 Generally, the kinetic energy can be neglected because it is Sensible 𝑐𝑝 𝑇 33.94 miniscule compared with the other Potential 𝑔𝑧 8.37 energy components. The table shows the mean energy Kinetic 1 2 𝑢 + 𝑣2 0.0017 in the northern hemisphere as 2 calculated by Oort (1971). Energy Balance and the Role of the Tropics (Surface Energy Budget) Under steady state conditions, such as would be assumed for the annual average, equation (6) can be reduced and partitioned into: 𝑂𝑐𝑒𝑎𝑛, 𝑅𝑠 = 𝑐𝑝 𝑇 + 𝐿𝑞 + ∆𝑓 (7) 𝐿𝑎𝑛𝑑, 𝑅𝑠 = 𝑐𝑝 𝑇 + 𝐿𝑞 (8) Note that several small terms have been left out of the equation, such as latent heat of fusion for melting ice and snow, conversion of kinetic energy of winds and waves to thermal energy, energy used for photosynthesis, geothermal energy, and heat from fossil fuel burning. Energy Balance and the Role of the Tropics (Meridional Energy Transport by Atmosphere and Ocean) Satellite measurements and in situ observations are used to create an energy budget, which allows us to calculate the kinds of energy transported in each latitudinal zone. For example, the figure shows that latent heat flux is the dominant energy input from the surface to the tropical atmosphere, with prominent peaks near 15° latitude (greater in the southern hemisphere). The sensible heat flux is much smaller, varies little at tropical latitudes, peaks over the northern hemisphere near 30°N, and is negative around the poles. Net radiative cooling varies little at tropical latitudes. The net energy flux into the atmosphere peaks in the subtropics with a relative minimum around the equator. Energy Balance and the Role of the Tropics (Meridional Energy Transport by Atmosphere and Ocean) Atmospheric and oceanic contributions to the meridional transport of energy are shown in the figure. Measurements were taken by the Earth Radiation Budget Experiment (ERBE; February 1985–April 1989) and the Clouds and Earth’s Radiant Energy System (CERES; March 2000–May 2004) satellites, National Centers for Environmental Prediction–National Center for Atmospheric Research (NCEP–NCAR) Reanalysis (NRA), and ocean circulation transportation derived from Global Ocean Data Assimilation System (GODAS). Energy Balance and the Role of the Tropics (Meridional Energy Transport by Atmosphere and Ocean) Meridional energy transport is divided between mean meridional transport and eddy motion (waves). Between the equator and about 15°N, a large portion of the energy transport is by the mean meridional circulation (MMC), also known as the Hadley cell. The Hadley cell is stronger during winter than it is in the summer; this means that the northern hemisphere MMC depicted in the figure transports much of its energy for the year in January through March. During the northern hemispheric winter, the Hadley cell is fairly intense compared with the summer circulation. In comparison, midlatitude transport is mostly by eddies (midlatitude cyclones). Energy Balance and the Role of the Tropics (Latent Heat and Deep Convective Cloud Distribution) Latent heating is the most important method of surface to atmosphere transfer in the tropics and the dominant source of energy for tropical circulation. Latent heat is added to the atmosphere during condensation and precipitation and removed during evaporation. The tropical oceans, which occupy most of the equatorial belt, supply most of the moisture for convection. The strength and location of convection depends on various factors including surface fluxes of sensible heat, which may alter the stability of the atmosphere. Energy Balance and the Role of the Tropics (Latent Heat and Deep Convective Cloud Distribution) Deep convective clouds are usually identified by their cold cloud tops which emit low values of outgoing longwave radiation (OLR). Minima in OLR, a proxy for deep convective clouds, are typically found over the tropical continents and the warm pools of the tropical western Pacific and Indian oceans. In regions of deep convection, precipitation generally exceeds evaporation (blue and magenta regions within about 10° of the equator). In the subtropics (from about 10°-40° latitude, centered on yellow and orange areas), evaporation exceeds precipitation. Energy Balance and the Role of the Tropics (Latent Heat and Deep Convective Cloud Distribution) In order to maintain latent heat balance, water vapor must be exported from regions of excess evaporation to regions of excess precipitation. Most of the transport of water vapor is done by the mean meridional circulation, the Hadley cell. Water vapor brought by the trade winds into the equatorial low pressure areas feeds the deep cumulonimbus cloud systems. Energy Balance and the Role of the Tropics (Latent Heat and Deep Convective Cloud Distribution) Deep convective clouds are the primary mechanism (or conduit) for transferring the sun's heat into the tropical atmosphere. That energy drives upward motion in the tropical atmosphere and the larger scale motions of the general circulation. Energy Balance and the Role of the Tropics (Latent Heat and Deep Convective Cloud Distribution) Spatial and temporal change in the regions of convection and maximum rising motion can lead to dramatic changes in the global climate. For example, the east-west shifts in the region of deep convection over the equatorial Pacific during an El Niño. The physical characteristics of clouds influence the distribution of radiative heating and cooling in the troposphere. Thus cloud processes that determine the cloud radiative forcing are significant components of the large- scale circulation in the tropics. Energy Balance and the Role of the Tropics (Surface-Air Interactions) An integral part of the transport of energy within and from the tropics is the interaction between the surface and the atmosphere. Since the ocean occupies such a large area of the surface, ocean-atmosphere interactions are a dominant component of that energy exchange. Because water has a larger heat capacity, the oceanic response to seasonal variations of solar heating is smaller than over land. The oceans are then able to store heat during the summer and release it during the winter. The peak in oceanic transport of heat to the poles occurs at low latitude. Energy Balance and the Role of the Tropics (Surface-Air Interactions) As noted earlier, the ocean to atmosphere transfer is primarily of heat and moisture through evaporation. On the other hand, winds transfer momentum to the ocean by moving the waters around. A tropical cyclone is an impressive example of these interactions; its energy is gained from the warm ocean waters through latent heat transfer while its winds stir the ocean surface, transfer momentum downward, and leave a cool anomaly in its wake. Furthermore, regions of heating and strong convection are observed over the tropical continents and warm ocean basins. Energy Balance and the Role of the Tropics (Surface-Air Interactions) Various weather and climate phenomena are generated in response to these heating maxima. One prominent example is the Madden-Julian Oscillation (MJO), a 30-60 day oscillation of surface pressure and winds that influences tropical weather from small-scale convection to planetary-scale circulations. The MJO is a coupled atmosphere-ocean phenomenon that is commonly identified by a broad area of active cloud and rainfall propagating eastward around the equator. The oceanic signature of the MJO is evident in the SSTs, depth of the isothermal ocean surface layer, and latent heat exchange. Energy Balance and the Role of the Tropics (Surface-Air Interactions) The figure on the right illustrates some critical surface-atmosphere interactions such as winds and waves, evaporation, heat, and salinity exchanges. Climate and weather models try to account for these interactions.  Atmospheric Structure (Temperature Profile) The troposphere is heated from below by latent heat, longwave radiation, and sensible heat. The tropics experience surplus heating and vertical expansion of the troposphere in response to that heating. In addition, deep tropical clouds transfer latent heat high in the atmosphere. The result is that the tropopause is highest in the tropics. The average height of the tropical tropopause is almost 7 km higher than the average tropopause height at the poles Atmospheric Structure (Temperature Profile)  Atmospheric Structure (Temperature Profile) The lowest layer of the troposphere is known as the planetary boundary layer or PBL. This layer of the atmosphere is in contact with the surface and experiences the effects of friction. It is the layer through which heat, moisture, and momentum are exchanged between the atmosphere and the surface. Atmospheric Structure (Temperature Profile)  Atmospheric Structure (Temperature Profile) Motion in the boundary layer is turbulent. Because it feels the effects of the surface, the PBL experiences large diurnal changes in temperature, winds, and depth. It varies from a few 100 m over tropical oceans to 6 km over the hot, dry Sahara. It is often capped by an inversion. Note that a well-defined boundary layer is not always present.  Atmospheric Structure (The Trade Wind Inversion) One of the most prominent features of the tropical boundary layer is the trade wind inversion, which is strongest over the eastern regions of the tropical oceans. These areas are marked by upwelling of cooler, deep water and cool SSTs. Subtropical ridges suppress the marine boundary layer in the eastern tropical oceans. Subsidence dries and warms the layer above the boundary layer and creates an inversion.  Atmospheric Structure (The Trade Wind Inversion)  Atmospheric Structure (The Trade Wind Inversion) As a result of the strong inversion and cool SSTs in this region, moisture content increases within the marine boundary layer and, with saturation, clouds form over a wide area of the eastern tropical oceans. Typically stratus is near the coast, stratocumulus is offshore, and trade-wind cumuli are over the relatively warm ocean to the west.  Atmospheric Structure (Atmospheric Humidity) As expected from the Clausius-Clapeyron equation, which describes the relationship between temperature and the saturation vapor pressure at equilibrium, surface water vapor content is highest on average in the tropics and lowest at the poles.  Atmospheric Structure (Atmospheric Humidity) As temperature increases more molecules are needed to achieve equilibrium between vapor and liquid. Conversely, fewer molecules are needed for equilibrium as the temperature decreases. The relationship may be expressed as: 𝑒 𝐿 1 1 ln 𝑒 𝑠 = 𝑅𝑣 −𝑇 (9) 𝑠0 𝑣 𝑇0 where T is the temperature in K, To is 273K, es is the saturation water vapor pressure in hPa, eso is the saturation water vapor pressure at temperature To (6.11 hPa), Lv is the latent heat of vaporization (2.453 × 106 J kg-1), and Rv is the water vapor gas constant (461 J kg-1 K-1). We can rewrite (9) to calculate the saturation vapor pressure at T as: 𝐿𝑣 1 1 𝑒𝑠 = 6.11 𝑒𝑥𝑝 − 𝑅𝑣 𝑇0 𝑇  Atmospheric Structure (Atmospheric Humidity) In general water vapor decreases with height but is much more variable in space and time than temperature.  Atmospheric Structure (Atmospheric Humidity) The contrast between the relatively smooth temperature and the humidity profiles is evident in the figure. The soundings are from St. Helena Island in the tropical south Atlantic (left) and Penang in near equatorial Malaysia (right).  Atmospheric Structure (Atmospheric Humidity) St. Helena Island is in a subtropical high pressure area with subsidence above a moist boundary layer. The profile therefore shows high relative humidity in the boundary layer, with a dramatic decrease in relative humidity near 800 hPa. The relative humidity decreases in the temperature inversion layer.  Atmospheric Structure (Atmospheric Humidity) In contrast, Penang is under the influence of the tropical warm pool where deep convection is frequent and the relative humidity remains fairly high in many layers of the troposphere.  Atmospheric Structure (Atmospheric Humidity) Sometimes, advection of dry air can change the profile of moisture, even in the humid tropics. The figure shows the large decrease in relative humidity over the Caribbean due to the dry Sahara Air Layer (SAL). Between 800 and 600 hPa, the difference between the SAL and non-SAL environment is greater than 30%. The figure also shows that the mean tropical sounding calculated by Jordan (1958) is moister in the lower troposphere and drier in the mid- troposphere compared with the mean non-SAL environment for 1998 and 2000.  Atmospheric Structure (Pressure Ranges) Horizontal pressure gradients are fairly weak across the tropics. Sea level pressure in the Hadley cell ranges from strong high pressure in the subtropics (about 1024 hPa) to the extremely low pressure in tropical cyclones (about 980 hPa in weak tropical cyclones). A record minimum of 870 hPa was observed in Typhoon Tip in October 1979, but such extremely low pressure tropical cyclones are rare.  Seasonal and Geographic Distribution of Temperature The primary influence on the mean annual temperature is latitude. The period of daylight and the solar declination angle vary with the latitude. The amount of energy received per unit area decreases towards the poles. The equator always has 12 hours of daylight. Places between the Tropics of Cancer and Capricorn experience the overhead beam and less attenuation by the atmosphere while for higher latitudes more insolation is attenuated because of the longer distance through the atmosphere.  Seasonal and Geographic Distribution of Temperature Even small differences in the solar declination angle can affect the temperature range and the annual average temperature as illustrated by the seasonal cycles shown in the figure. The higher latitudes have a greater temperature range with a peak in July (northern hemisphere) and January (southern hemisphere). The temperature lags the solstice as the atmosphere responds to the heating of the surface. The ocean heats up more slowly than land because water has a higher specific heat capacity.  Seasonal and Geographic Distribution of Temperature Latitude as a control of mean temperature and the annual temperature range is shown clearly in the figure. In general, the mean temperature declines from equator to poles and the range increases. For the low latitude stations, such as Pontianak, the monthly mean stays above 25°C and the range between minimum and maximum monthly mean temperature is 2°C or less. Moving to higher latitudes, the minimum monthly temperature is getting cooler for longer periods but is warmer or similar to the near equatorial station temperature during summer. Seasonal and Geographic Distribution of Temperature  Seasonal and Geographic Distribution of Temperature The figure shows the mean temperature for the two seasonal extremes, January and July and the animation shows the full annual cycle. Maxima are found over the continental regions of the tropics. The pattern is nearly zonal, matching the latitude, over the oceans while near the coasts and over land there is greater meridional variation. Minima at tropical latitudes are found at high elevation such as the Andes and East Africa. The northern hemisphere, with its greater land mass, has warmer summers and cooler winters than the southern hemisphere. Meridional temperature gradients are larger in winter than summer.  Seasonal and Geographic Distribution of Temperature The influence of the prevailing ocean currents is seen along the eastern and western ocean basins and nearby coastal regions. For example, off southeastern Africa, the temperature is 30°C or higher while off the southwestern coast, at the same latitude, the temperature is 20°C or lower. Longititudinal temperature contrasts are relatively small (< 6°C), and are correlated with continents, which are colder than oceans in winter and warmer during the summer.  Seasonal and Geographic Distribution of Temperature In general, the tropics have the lowest annual temperature range for the globe, less than 1.5°C in the near equatorial regions. The mid-continental regions of Africa and Australia have the highest range within the subtropics. The maximum annual temperature range in Africa occurs in northwestern Africa, south of the Atlas Mountains. Australia experiences the highest mean annual temperature range of the tropical continental regions.  Major influences on Annual Surface Temperature Distribution Latitude Continentality Relief Prevailing atmospheric and oceanic flow Clouds and Precipitation Albedo  Major influences on Annual Surface Temperature Distribution (How?) Latitude Period of daylight varies from 12 h at the equator to the extremes of 0 and 24 h at the pole that is tilted away from or towards the sun, respectively. The amount of daylight in the tropics has a small range between solstices.  Major influences on Annual Surface Temperature Distribution (How?) Latitude Less attenuation of the incident solar radiation occurs over the tropics compared to the poles. The incident solar beam has greater depletion with longer distance through the atmosphere  Major influences on Annual Surface Temperature Distribution (How?) Latitude The solar elevation angle is highest in the tropics, which has the highest heating per unit area. With decreasing declination, the solar beam spreads out and the heating density decreases.  Major influences on Annual Surface Temperature Distribution (How?) Continentality Continental regions have a larger annual temperature range than the oceans because the specific heat capacity of water is greater than that of land. Land heats and cools faster than water. Therefore, the northern hemisphere with its larger land mass has warmer summers and colder winters than the southern hemisphere and coastal areas have smaller temperature range than mid- continental regions at the same latitudes.  Major influences on Annual Surface Temperature Distribution (How?) Relief Higher elevations are colder. Leeward sides of mountain ranges are warmer and drier than windward sides. Rain shadows develop on the leeward side due to adiabatic warming with descending flow while the windward side is characterized by rising motion, condensation, and precipitation which keeps the surface cool. Slopes facing equatorward are generally warmer than slopes facing poleward.  Major influences on Annual Surface Temperature Distribution (How?) Prevailing atmospheric and oceanic flow Regions are warmed or cooled by currents flowing from the equator or the poles, respectively. Prevailing winds from the ocean moderate the temperature range while prevailing flow from the land leads to a greater range of temperatures.  Major influences on Annual Surface Temperature Distribution (How?) Clouds and Precipitation The seasonal variation of clouds and precipitation has a strong effect on annual temperature variation in some tropical regions. The temperature decreases during the rainy season. For example, over southwest India, the temperature is at its maximum in early May, decreases with the monsoon rains to a July minimum and reaches a secondary maximum in September.  Major influences on Annual Surface Temperature Distribution (How?) Albedo It seems intuitive to assume that surfaces with high albedo or reflectivity will absorb less sunlight and thus have cooler annual mean temperature. For example, snow-covered regions near the poles and at high elevation are colder because of the high reflectivity of snow. Such places will become warmer if the snow is replaced with a lower albedo surface. However, latitude and atmospheric circulation patterns are more dominant than albedo in determining annual temperature range. So, subtropical deserts have high albedo and high annual mean temperature.  Major influences on Annual Surface Temperature Distribution The highest average annual surface temperatures are in the inland regions of the subtropical deserts, whereas the lowest are found in inland of Antarctica, Greenland, and Russia. Diurnal Temperature Variability in the Tropics The diurnal variation in temperature has its maximum at the surface, following the daily cycle of surplus heating. The figure is an idealized graph of the daily temperature with a minimum at sunrise and a maximum in the afternoon. The daily maximum generally lags the solar maximum as the heated surface is warming the surrounding air.  Major influences on Diurnal Surface Temperature Distribution Humidity Cloudiness Wind speed Albedo Elevation  Major influences on Diurnal Surface Temperature Distribution (How?) Humidity Temperature variation is more moderate in humid environments because water vapor is a good absorber and emitter of longwave radiation. Water vapor also absorbs in the near infrared part of the solar radiation, which reduces the energy reaching the surface during the daytime. Therefore, daily maximum temperatures are lower in humid environments and higher in dry environments. Furthermore, where the land surface is dry and the air is dry, the conductive capacity is reduced and the diurnal temperature range is greater. The near surface air responds to the rapid heating and cooling of the surface. Thus, the Sahara desert, for example, has a large diurnal temperature range while the Indonesian rainforest has a small diurnal temperature range.  Major influences on Diurnal Surface Temperature Distribution (How?) Cloudiness Clouds are good absorbers and emitters of long-wave radiation and good reflectors of sunlight (shortwave radiation) therefore cloudiness leads to cooler days and warmer nights and a smaller diurnal temperature range.  Major influences on Diurnal Surface Temperature Distribution (How?) Wind Speed During windy conditions air with different temperatures is more easily mixed and helps to moderate the temperature range.  Major influences on Diurnal Surface Temperature Distribution (How?) Albedo The diurnal temperature range is influenced by the albedo in the same manner as the annual cycle. Highly reflective surfaces are cooler than surfaces with low reflectivity.  Major influences on Diurnal Surface Temperature Distribution (How?) Elevation Mountain areas are warmed earlier than the valleys below and cool more rapidly after sunset. The difference in heating causes valley (upslope) and mountain (downslope) breezes, respectively. The aspect of the slopes also affects the diurnal temperature cycle. The part of the slope facing the rays of the sun will be warmer than the sheltered side.  Major influences on Diurnal Surface Temperature Distribution (How?) Elevation In most cases, west-facing slopes will be warmer because the sun is in the west during the hottest part of the day (unless prevailing flow causes clouds and precipitation). High terrain surface warms and cools more rapidly than air at the same altitude. When those temperature differences are large, the pressure gradients can lead to local wind maxima developing at low altitudes. Temperature Extremes in the Tropics Regional Influences on Tropical Temperature Variability Regional Influences on Tropical Temperature Variability Regional Influences on Tropical Temperature Variability  Moisture and Precipitation In the tropics, precipitation and the amount of water vapor integrated over a column of air are closely related. Condensation and precipitation require high relative humidity, which is supplied by evaporation from the surface or horizontal transport by the winds. The high water vapor content fuels the tall convective clouds that produce the precipitation.  Moisture and Precipitation However, unlike temperature and pressure, which are fairly homogeneous across the tropics (except in intense convective systems such as tropical cyclones), tropical precipitation is mostly convective, episodic, and variable in nature. With little or no Coriolis force in the tropics, horizontal motion and associated gradients of moisture are not governed by the balance between Coriolis and pressure gradient force that is characteristic of midlatitude weather. Rather perturbations in the wind field, e.g., zones of low-level convergence (negative divergence) have a greater influence on horizontal gradients of moisture, temperature, and pressure. Furthermore, horizontal gradients in moisture between dry and wet regions over land can create a dynamical response.  Moisture and Precipitation Competing theories exist for thermodynamic versus dynamic control of tropical surface winds and precipitation fields. For example, in quasi-equilibrium theory it is assumed that convection is in statistical equilibrium with perturbations in the large- scale flow which creates a thermally stratified troposphere that is moist adiabatic. The large amount of moisture in the tropical boundary layer is the dominant factor and the velocity fields dictate when and how the convection will organize.  Moisture and Precipitation A simple classification of precipitation by latitude is shown in the figure. As expected, the equatorial low and sub-polar low regions have precipitation year round with dry subtropical highs. In general, tropical zones between the equatorial trough and dry subtropical highs have wet summers and dry winters. It is interesting to note that tropical meteorology includes all precipitation types, including frozen precipitation in clouds and on the ground at high elevation.  Role of the Tropics in Momentum Balance In addition to maintaining the global energy balance, tropical circulations are also critical for maintaining the global angular momentum balance. The absolute angular momentum = mass × the angular or rotation velocity × the perpendicular distance from the axis of rotation, in equation form as m × rω × r = mr2ω.  Role of the Tropics in Momentum Balance The conservation of absolute angular momentum means that as the distance from the axis of rotation changes then the absolute angular velocity also changes to maintain momentum balance. When we consider the angular momentum balance for the atmosphere and earth combined, we must also account for the transfer of momentum between the earth and the atmosphere as well as the transfer of momentum within the atmosphere.  Role of the Tropics in Momentum Balance The absolute velocity is the sum of the angular velocity of the earth and the relative zonal wind speed. The earth revolves at the rate of once every 23 hours, 56 minutes and 4.1 seconds. So, the angular velocity of the earth, Ω, is 2π/86164.1 rad s- 1 or 7.292 x 10-5 rad s-1.  Role of the Tropics in Momentum Balance Since Ω is constant, then the relative zonal wind speed will change when the distance from the axis of rotation changes, i.e., with latitude or altitude. Since the atmospheric depth is thin relative to the radius of the earth, we can focus on the latitudinal radius. The latitude has a strong influence on the absolute angular momentum, which can be calculated, per unit mass of the atmosphere, as 𝑀 = Ω 𝑎 cos ∅ + 𝑢 𝑎 cos ∅ where a is radius of the earth, Ω is the angular velocity of the earth, u is the zonal wind speed, and Φ is the latitude.  Role of the Tropics in Momentum Balance When a parcel moves from the equator towards the poles, it retains the same angular momentum unless it exchanges angular momentum with other air parcels or with the surface. As the parcel moves poleward, the distance to the axis of rotation decreases, so its eastward velocity will increase to maintain a constant total angular momentum. In the region of easterly winds where the atmosphere is moving more slowly than the earth’s surface, the atmosphere gains momentum from the earth. In the region of the westerlies, the atmosphere is rotating faster than the surface, and it gives up westerly angular momentum to the surface. Role of the Tropics in Momentum Balance  Role of the Tropics in Momentum Balance Within the Hadley cells, parcels moving upward and poleward will accelerate eastward to conserve angular momentum. Note that shifts in atmospheric pressure patterns, especially near mountains, and wind velocity can change the rate of rotation of the solid earth. For example, when pressure patterns shift during El Niño, the earth’s rotation adjusts to maintain momentum balance.  Spatial and Temporal Scales in the Tropics Motion and momentum transfer in the atmosphere are occurring at various scales simultaneously. Instabilities in the atmosphere and ocean are created by gradients of temperature, winds, humidity, and SSTs. Weather and climate phenomena are responses to these instabilities. Scales of atmospheric motion range from the short length and time scales of friction and turbulent motion to the decadal and planetary- scale circulations.  Spatial and Temporal Scales in the Tropics The figure illustrates some dynamical processes of the tropical atmosphere and their typical space and time scales. Note that most occupy a range of scales and that smaller scale features can occur within a larger scale circulation. For example, tropical cyclones consist of numerous thunderstorms; tropical cyclones can be spawned within the intraseasonal MJO; and both are modulated by the inter-annual ENSO.  Tropical Air Masses and Climates As a consequence of the latitudinal variations in temperature and moisture and underlying surface characteristics, different types of air masses form around the globe. The first letter of each air mass identifies its moisture characteristics (maritime or continental) and the second its temperature characteristics (Tropical, Polar, Arctic or Antarctic).  Tropical Air Masses and Climates Maritime air masses (mE and mT) dominate much of the tropics because of the vast area of ocean. Continental tropical (cT) air masses originate over North Africa, Australia, and North America; cT air masses are relatively dry, which means that they heat and cool faster than maritime air masses.  Tropical Air Masses and Climates Along with the temperature, the climate of a region is determined by its precipitation characteristics. A more complicated climate classification emerges when the effects of elevation and complex topography are added.  Tropical Air Masses and Climates The climate classification shown in the figure was developed by Wladimir Köppen (1918) and revised by Rudolf Geiger. It categorizes zones according to temperature extremes; precipitation amount and type as well as seasonality of precipitation. The three-letter classification identifies, respectively, the main climate, the precipitation, and the temperature. Note that these climate zones do not have sharp boundaries nor are they static. There is gradual transition from one zone to another with subregions defined by local physiography.  Tropical moist climates (Group A, average temperature of each month above 18°C) Af – tropical wet or tropical rainforest Abundant rainfall year round (> 60 mm per month); annual temperature range is less than 3°C; diurnal temperature range, about 10°C on average, is greater than the annual temperature range; cloudiness and high humidity moderate maximum temperatures. This climate is found in the near equatorial region.  Tropical moist climates (Group A, average temperature of each month above 18°C) Am- tropical monsoon Annual rainfall similar to tropical wet but interrupted by one or two months of little or no precipitation. Such climates are found along the southwest coast of India and the Indochinese peninsula, for example.  Tropical moist climates (Group A, average temperature of each month above 18°C) Aw – tropical wet with dry winter Distinct dry season with less rainfall than other tropical moist climates (one month with precipitation < 60 mm); amount and timing of rainfall can vary widely from one year to the next or even within a season; cooler temperatures in winter, especially overnight because of dry conditions. Tropical wet-dry regions are, generally, just poleward of tropical wet and monsoon regions.  Dry climates (Group B, subgroup h with average monthly temperature above 18°C) BWh – arid and hot Little or no precipitation; precipitation is highly irregular in timing and frequency, a station may receive its annual precipitation in one day; large diurnal temperature range, from scorching maximum above 45°C to cool minimum of 25°C and below; found mainly in the subtropics, centered along the Tropics of Cancer and Capricorn and extending between 15° and 30° latitude. Dry climates dominate northern Africa, the Arabian Peninsula, Australia, and the west coasts of continents where cold currents are a big influence.  Dry climates (Group B, subgroup h with average monthly temperature above 18°C) BSh – semi-arid and hot The transition from tropical moist to arid regions where precipitation can fluctuate widely from one year to the next; such variability has had devastating consequences in regions like the Sahel, just south of the Sahara. When rains are late, not far enough north, or reduced in amount, crop failure has led to widespread famine.  Moist climate with mild winter (Group C) Cfa – humid, subtropical This climate is found in a few areas of the tropics, e.g., parts of southern China, Florida, and the east coast of Mexico, where elevation and/or atmospheric circulation patterns create similar conditions to midlatitude areas; mild winters with average temperature from -3°C to 18°C during coldest month; distinct summer and winter with enough precipitation to be classified as humid; winter weather is changeable because of the passage of midlatitude cyclones; summer temperatures are warmer than in tropical moist climates.  Moist climate with mild winter (Group C) Cwa – dry winter Adjacent to tropical wet and dry areas; found in parts of Africa, South America, and Asia that are high in elevation and thus too cool to be considered tropical; this climate is similar in temperature to humid, subtropical but with dry winters.  Highland climates (Group H) As noted earlier, as elevation increases both temperature and moisture content decrease. 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