Chapter 1: Introduction

Table of Contents

1.0 Overview

energy balance

This chapter begins with an overview of energy balance and the global climate system. Various definitions of the tropics are presented. The role of the tropics in the global energy and momentum balance is presented. Atmospheric structure of temperature and humidity are discussed in terms of latitudinal variability. Pressure ranges and scales of atmospheric motion in the tropics are reviewed. Seasonal and geographic distribution and the diurnal cycle of surface temperature and the influencing factors are examined in detail. Finally, we review tropical air masses and climates.

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1.0 Overview »
Learning Objectives

At the end of this chapter, you should be able to:

  • Understand the concept of radiation balance between intake and loss of energy by the earth and atmosphere
  • Describe the role of the tropics in the global climate system
  • List the mechanisms that allow for exchange of energy between the surface and the atmosphere
  • Recall at least three methods of defining the tropics
  • List the components of the surface energy budget
  • Understand the atmosphere and ocean transport in the global energy balance
  • Describe how latent heat and tropical deep convection fit into the global energy budget
  • Relate the tropical tropopause height to higher latitude tropopause height
  • Recall the basic structure of the trade wind inversion and its impact on east-west cloud distribution over the tropical oceans
  • Describe the range of surface humidity in the tropics relative to higher latitudes
  • Describe the geographic and seasonal distribution of surface temperature and the factors that influence that distribution
  • Recall the typical diurnal distribution of temperature and factors that influence that distribution
  • Describe the meridional transport of angular momentum and the role of the tropics
  • Recall at least three tropical air masses and locations
  • Recall at least three tropical climate regions

1.1 What is Tropical Meteorology?

Tropical or midlatitude rain, which of these photographs do you think is typical for each zone?

midlatitude raintropical rain

The correct answer is b.

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. Driving these events are important energy sources and sinks like surplus radiation, latent heat, sensible heat, evapotranspiration, and ocean heat storage.

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.

Tropical cumulus clouds are a conduit for the release of latent heat,1 their main source of energy. Tropical cyclones are the most spectacular manifestation of this energy exchange. Local and mesoscale systems play a greater role in tropical weather than synoptic systems. There is strong interaction among these various scales of cumulus convection, mesoscale, and large-scale circulations. Such interactions are important for tropical weather and climate prediction but are not fully understood, hence the need for ongoing research.

We must also remember that meteorology is an observational science, where knowledge is advanced by improvements in the observation network. Tropical meteorology has benefited especially from advances in satellite remote sensing. Satellites have become the primary source of tropical weather and climate observations because of the sparseness of the standard surface and upper air network. Thus, remote sensing technologies and their applications are an integral part of tropical meteorology.

A study of tropical meteorology must also consider tropical atmosphere-ocean coupling. Tropical sea surface temperature (SST) variations influence diabatic heating distribution and atmospheric motion helps to drive upper ocean circulation. In addition, diabatic heating associated with tropical precipitating systems can induce waves2,3 that lead to remote responses in the tropics and extratropics. Intraseasonal variability of tropical rainfall is dominated by the 40-50 day oscillation known as the Madden-Julian Oscillation (MJO)4,5 while the interannual scale is dominated by the El Niño-Southern Oscillation (ENSO).6,7 The past two decades have seen a rapid improvement in our understanding of tropical variability; based on theoretical dynamics, observations, and numerical models. However, operational tropical prediction is still relatively immature compared with the midlatitudes.

Large areas of the tropics are influenced by monsoon circulations,8 which result from the temperature contrast that results from the different temperature responses to heating of land and ocean. Warming of the land relative to the ocean leads to “onshore” flow, cumulus convection, and warming of the troposphere by latent heat release; the reverse circulation occurs during the winter. Entire economies in parts of the tropics depend on an adequate supply of monsoon rains. Knowledge of monsoon variability, including the onset and retreat of the monsoon, breaks in the monsoon, and rainfall amounts is critical information for billions of people in Asia and Africa. Too little or too much rainfall can have devastating consequences.

The ability of tropical societies to reduce their vulnerability to environmental disasters is tied to the quality and application of weather and climate information and prediction. For example, knowledge of the spatial and temporal variability of precipitation is needed to manage agriculture, water resources, public health, and renewable energy. Such information is gained from analysis of the observations and predictive models.9 Advances in tropical prediction are due to the combination of physical insight, skillful computer model representations of the physical laws governing the weather and climate, and advances in computing technology. Numerical (computer) models also facilitate examination of physical processes that influence weather and climate.

Knowledge of the tropical atmosphere has increased significantly during recent decades, especially because of satellite observations, field programs, and modern numerical models. The textbook will cover those and other topics such as energy distribution; surface-atmosphere interactions; waves in the equatorial region; cumulus convection and its interactions with mesoscale circulations and large-scale motions; the variability of circulations across scales from turbulence to decadal; the vertical transport of heat, moisture, and momentum; and tropical cyclones.


1.2 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.

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:

missing(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 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 transfer from the sun to the earth is nearly all by radiation (some negligible mass is associated with the solar wind). Figure 1.1 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.

annual incoming solar radiation
Fig. 1.1. Latitudinal variations in the annual incoming solar radiation (insolation) density and distance to the surface.

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 relates the energy emitted per unit area (from all wavelengths) to the absolute temperature (K) by,

Stefan-Boltzmann equation(2)

where the Stefan-Boltzmann constant, σ = 5.67 x 10-8 W m-2 K-4

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.

Wien's Law equation(3)

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 Fig. 1.2. 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.

Solar and Thermal blackbody curves
Fig. 1.2. Conceptual depiction of the energy emitted per unit area as represented by the area under the blackbody curve for solar radiation (yellow) and terrestrial radiation (red).

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

energy balance at top of atmosphere(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.

global annual mean energy balance
Fig. 1.3. Global annual mean energy balance for March 2000 to May 2004 (W m-2).10

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 (Fig. 1.3). Of these processes, conduction contributes the least to warming the tropospheric column.

Slight variations in energy budget percentages will occur depending on the period for which the averages are taken and the data used to estimate the contributions. The primary data sources for global radiation budgets are satellite measurements.

In studying the atmosphere, a distinction is made between weather and climate. Weather describes short-term state of the atmosphere at a given time and place; for example, today, tomorrow, or the near future. Short-term weather fluctuations are caused by internal atmospheric instabilities rather than changes in solar output.

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.

The heat engine of the earth system is driven by the tropical latitudes. Malkus (1962)11 refers to the tropics as the “firebox” of the heat engine (Fig. 1.4). 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 animation of Fig 1.4a.

temp gradient between tropics and poles
Cross-section showing the meridional circulation
Fig. 1.4. (a) Schematic of the temperature gradient between the warm tropics and the cold poles and the resulting pressure gradient force (at the top of the atmosphere) and atmospheric motion. (b) Cross-section showing the meridional circulation between the tropics and subtropics driven by heating in the equatorial regions and upward transport by deep convection in the tropics and sinking motion in the subtropics. Adapted from Simpson (1992)12

Riehl and Malkus (1958)1 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 Fig. 1.5. 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.

Satellite IR image at 1200 UTC on 18 Aug. 2001
Fig. 1.5. Satellite IR image at 1200 UTC on 18 Aug. 2001; the yellow line outlines the region of deep convection cloud systems responsible for most of the upward vertical transport of energy in the tropical atmosphere.

1.3 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, (Fig. 1.6).
  • The region of surplus radiation where annual solar input minus terrestrial output is positive, ± 35 to 40° latitude, (Fig. 1.7).
  • The region of net upward motion and surface low pressure (Fig. 1.8): 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 1735,13 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.
  • The region in which winds blow primarily from the east (approximately ± 30° latitude), except for the regional monsoon (Fig. 1.9). 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.
  • The region where the annual range of temperature is less than or equal to the average daily range. We will examine the surface temperature distribution in Section 1.6 and focus sections at the end of the chapter.
  • 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. Temperature and precipitation characteristics of tropical climates are presented in Section 1.9.

Riehl (1979)14 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.

tropics defined by solar declination angle
Solar elevation angle at local noon
Fig. 1.6. (a) The Tropics defined by solar declination angle and (b) solar elevation angle at local noon.
tropics defined by net gain in radiation compared with deficits at the poles
Global annual radiation budget by latitude
Fig. 1.7. (a) The tropics defined by net gain in radiation compared with deficits at the poles. (b) Global annual radiation budget (W m-2).
 The Tropics defined by upward motion, low pressure, surface winds, and net surface heating
Fig. 1.8. The Tropics defined by upward motion, low pressure, surface winds, and net surface heating.
 tropics defined by zonal winds
Fig. 1.9. The tropics defined by zonal winds, predominantly easterly winds between 30°S and 30°N (dark blue and purple shades); the mean zonal winds for January–December 1979 to 2009.

1.4 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 (Fig. 1.10a).

 Conceptual model relating net diabatic heating and large-scale atmospheric dynamics
 simplified conceptual model relating energy, density, and large-scale ocean dynamics.
Fig. 1.10. (a) Conceptual model relating net diabatic heating and large-scale atmospheric dynamics and (b) simplified conceptual model relating energy, density, and large-scale ocean dynamics.

The large-scale oceanic dynamics are driven by 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 (Fig. 1.10b). Salinity changes are caused by processes such as evaporation, precipitation, and runoff. Heat is transported by convection, turbulent mixing, downwelling, or upwelling.

1.4 Energy Balance and the Role of the Tropics »
Box 1-1 Averaging Procedures

A variety of energy sources contribute to the general motion of the atmosphere and ocean. In order to understand the interactions and relative contributions of each source, we separate quantities into a mean and eddy (perturbation) over a given time or length scale. This concept is applied to variables such as moisture and momentum. For example, fluid motion can be considered as waves that are the sum of the mean and eddy components (Fig. 1B1.1).

fluid motion as a wave
Fig. 1B1.1. Conceptual diagram of motion and its mean and eddy components.

Averaging around a latitudinal belt can provide meaningful climate information because daily-averaged insolation is dependent on latitude and independent of longitude. Most climatologies are based o2n time averages such as monthly, seasonal, or annual averages. The effect of weather events appears as eddies or perturbations from the time mean. Different notation is used for each type of averaging, e.g.,

equation

Using this concept of mean and eddies, we can examine variations on a latitude circle that are associated with quasi-stationary features that show up in the time averages. Features that appear in the time averages can be perturbed spatially and be calculated as the deviation of the time mean from its zonal average,

equation

Note that, by definition, averages of deviations in the same mode (space or time) are equal to zero. For example, the time average of the time deviations of X is zero or

equation

Among the most commonly used averages are:

  • Time average,
equation

where ∆t is period of record

  • Latitudinal or zonal average over longitude, λ,
equation

where L is the circumference of the latitude circle or other length scale

  • Area average,
equation

where A is the area
 

  • Vertical average between two pressure levels,
equation

A more complex average is the transport or flux of quantity, X, across a latitude circle between pressure levels, p0 and pt, and averaged over time, ∆t (Fig. 1B1.2).

Vertical average between two pressure levels
Fig. 1B1.2. Schematic of transport across a latitude circle and between pressure levels.
equation

Length of latitude circle, L = 2πacosΦ, a is the radius of the earth, Φ is latitude, and g is acceleration due to gravity.

These concepts are the basis for calculating the energy budget of the globe, specific regions, or latitudinal zones. For example, Section 1.4.2 describes the meridional transport of energy, which is calculated from zonal and annual averages.

1.4 Energy Balance and the Role of the Tropics »
1.4.1 Surface Energy Budget

Following the formula in Eqn. (4), the net radiation at the surface can be written as:

equation(5)

where the subscripts s and a refer to the surface and atmosphere, respectively. Τ is the temperature (K), αs is the surface albedo. Arrows indicate direction of energy transfer. The atmosphere is largely transparent to solar radiation while absorbing and emitting longwave radiation (Fig. 1.3). 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”.

The surface energy budget relates the net radiation to the sensible heat, potential energy, latent heat, kinetic energy, storage, and advection:

Radiación neta        (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; Δf is the horizontal flux or advection; and G is the heat transferred in and out of storage (subsurface layers). The first two terms in (6) are the sensible and latent heat, respectively. Generally, the kinetic energy can be neglected because it is miniscule compared with the other energy components. Table 1 shows the mean energy in the northern hemisphere as calculated by Oort (1971).15

Table 1.1. Kinds and amount of energy in the northern hemisphere.
(adapted from original data listed in Oort 1971)15
Kind of Energy Formula Amount x 1021 Joules
Latent Heat Lq 0.931
Sensible CpT 33.94
Potential gz 8.37
Kinetic ½ (υ2 + ν2) 0.0017

Under steady state conditions, such as would be assumed for the annual average (6) can be reduced and partitioned into:

ocean surface heat budget equation       (7)
land surface energy budget equation       (8)

Note that several small terms have been left out of (6), 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.

1.4 Energy Balance and the Role of the Tropics »
1.4.2 Meridional Energy Transport by Atmosphere and Ocean

Satellite measurements and in situ observations are used to create an energy budget, which allows us to to calculate the kinds of energy transported in each latitudinal zone. For example, Fig. 1.11 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.

Zonal annual mean of the atmospheric energy budget
Fig. 1.11. Zonal annual mean of the atmospheric energy budget (W m-2): net radiative cooling of the atmosphere (Ra), latent heat flux (evaporation, LH), sensible heat (SH) at the surface into the atmosphere, and net energy into the atmospheric column (sum of the three) averaged over longitude and the annual cycle for 1985-1988. LH and SH data are from Sellers (1965).16 Adapted from Zhang and Rossow (1997)17

Atmospheric and oceanic contributions to the meridional transport of energy are shown in Fig. 1.12. 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),18 and ocean circulation transportation derived from Global Ocean Data Assimilation System (GODAS).19

The atmospheric transport of energy dominates except at low tropical latitudes where the ocean current transport is maximized (Fig. 1.12d). These currents account for about 25% of the total meridional transport of energy. The poleward transport by the ocean from the tropics into the winter hemisphere exceeds 4 petawatts (1015 Watts) in the tropics (red and deep blue areas of Fig.1.12c). The annual mean poleward transport by the ocean peaks at 11°S and 15°N (blue line in Fig. 1.12d), while the atmospheric transport peaks near 45°N (red line).

 annual mean meridional energy transport
Fig. 1.12. The ERBE period zonal mean annual cycle of the meridional energy transport by (a) the atmosphere and ocean as inferred from ERBE RT, NRA dAE/dt, and GODAS dOE/dt; (b) the atmosphere based on NRA; and (c) by the ocean as implied by ERBE + NRA FS and GODAS dOE/dt. Stippling and hatching in (a)–(c) represent regions and times of year in which the standard deviation of the monthly mean values among estimates, some of which include the CERES period (see text), exceeds 0.5 and 1.0 PW, respectively. (d) The median annual mean transport by latitude for the total (gray), atmosphere (red), and ocean (blue) accompanied with the associated twice the standard deviation from the mean (shaded). Units are petawatts (PW= 1015 W). Adapted from Fasullo and Trenberth (2008)20

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 (Fig. 1.13). The Hadley cell is stronger during winter than it is in the summer; this means that the northern hemisphere MMC depicted in Fig. 1.13 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).

Longitudinal gradients in energy distribution are due to topography and oceanic warm pools. The latter are created by wind stress and associated ocean currents.

annual averaged northward energy flux
Fig. 1.13. Annual averaged northward energy flux by the atmosphere (data from Oort 1971, AMS ).15 Units are 1015 W.

1.4 Energy Balance and the Role of the Tropics »
1.4.3 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.

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 (Fig. 1.14). In regions of deep convection, precipitation generally exceeds evaporation (blue and magenta regions within about 10° of the equator in Fig. 1.14). In the subtropics (from about 10°-40° latitude, centered on yellow and orange areas), evaporation exceeds precipitation. 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 (Fig. 1.8). Water vapor brought by the trade winds into the equatorial low pressure areas feeds the deep cumulonimbus cloud systems.

 Mean outgoing long-wave radiation
annual mean net surface solar radiation
Fig. 1.14. (a) Mean outgoing long-wave radiation; maxima in deep convective clouds occur over the tropical continents and maritime continent, within ±10° of the equator. (b) Annual mean net surface solar radiation. Units are W m-2.
external linkmovie Annual cycle of cold cloud tops, http://cimss.ssec.wisc.edu/wxwise/gifs/LWALL.mpg

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. Note that the flow of latent heat is northward across the equator because the mean position of the ITCZ is north of the equator; especially evident across the eastern and central Pacific (note blue and magenta regions in Fig. 1.14). You will learn more about the ITCZ position in the chapter on moisture and precipitation.

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.

1.4 Energy Balance and the Role of the Tropics »
1.4.4 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 (Fig. 1.12d).

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. 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.

Figure 1.15 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. Various scales of variability in tropical weather and climate, including coupled atmosphere-ocean phenomena, are explored in Chapter 4, Tropical VariabilityChapter 4, Tropical Variability.

 Schematic model of the climate system
Fig. 1.15. Schematic model of the climate system that includes critical processes and interactions in the atmosphere, hydrosphere, biosphere, and cryosphere.

1.5 Atmospheric Structure

1.5 Atmospheric Structure »
1.5.1 Temperature Profiles

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 (Fig.1.16).

 (a)Standard atmospheric temperature profile and (b) height of the tropopause in the tropics, midlatitudes, and poles
Fig 1.16. Standard atmospheric temperature profile (left) and height of the tropopause in the tropics, midlatitudes, and poles (right).

The lowest layer of the troposphere is known as the planetary boundary layer or PBL (Fig. 1.17). 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. 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 (Fig. 1.17). Note that a well-defined boundary layer is not always present. The boundary layer will be explored in a subsequent chapter.

 temperature profile
Fig. 1.17. Temperature profile in the troposphere and lower stratosphere. Inset graph shows the atmospheric boundary layer and an inversion at the top of the boundary layer. Sounding taken from St. Helena Island in the tropical south Atlantic.

1.5 Atmospheric Structure »
1.5.2 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 (Fig. 1.18a). Subsidence dries and warms the layer above the boundary layer and creates an 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.

Schematic of mean sea level pressure and air flow
conceptual model of the vertical profile of the trade wind inversion
Fig. 1.18. (a) Schematic of mean sea level pressure and air flow and its relationship to stability in the troposphere for the tropical north Atlantic and (b) conceptual model of the vertical profile of the trade wind inversion, from west to east across the equatorial oceans.

The trade wind inversion weakens towards the west with increasing SSTs and increasing instability of the troposphere. Figure 1.18b is a schematic of the trend from east to west across the tropical oceans. The vertical structure of the tropical troposphere, including the trade wind inversion, will be explored in more detail in a separate chapter.

1.5 Atmospheric Structure »
1.5.3 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 (Fig. 1.19).

distribution of surface water vapor percentageannual mean water vapor content profile
Fig. 1.19. (a) The distribution of surface water vapor percentage by latitude and (b) annual mean water vapor content (specific humidity) profile. Data in (b) from Oort (1983)21

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:

Clausius-Clapeyron Equation (9)

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:

equation

In general water vapor decreases with height (Fig. 1.19a) but is much more variable in space and time than temperature. The contrast between the relatively smooth temperature and the humidity profiles is evident in Fig. 1.20. The soundings are from St. Helena Island in the tropical south Atlantic (left) and Penang in near equatorial Malaysia (right). 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. In contrast, Penang is under the influence of the tropical warm pool where deep convection is frequent (Fig. 1.14a) and the relative humidity remains fairly high in many layers of the troposphere.

relative humidity and temperature profiles
Fig. 1.20. Relative humidity and temperature profiles in the troposphere and lower stratosphere for the tropical South Atlantic (marked by cool SSTs, low stratus clouds, and subsidence from the subtropical high) and the near equatorial Maritime Continent (the tropical warm pool area).

Sometimes, advection of dry air can change the profile of moisture, even in the humid tropics. Figure 1.21 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.

relative humidity and temperature profiles
Fig. 1.21. Profiles of mean relative humidity over the West Indies (Caribbean Sea) from Jordan (1958) in black, Non-SAL (solid blue) and SAL environment (dashed blue). Courtesy of Dunion and Velden (2004)22

1.5 Atmospheric Structure »
1.5.4 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 1979Typhoon Tip in October 1979, but such extremely low pressure tropical cyclones are rare. We will explore more about pressure gradients and circulations in the chapter on global circulationglobal circulations.

1.6 Temperature

1.6 Temperature »
1.6.1 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 (Fig. 1.1). 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 Fig. 1.22. 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.

annual cycle of temperature for stations in southern hemisphere
Fig. 1.22. The annual cycle of temperature for stations in the southern hemisphere (negative sign indicates southern latitudes).

Latitude as a control of mean temperature and the annual temperature range is shown clearly in Fig. 1.22. 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.

Review Fig. 1.22 carefully. Can you find the discrepancy in the latitudinal temperature trend? What do you think may be causing that discrepancy?

Feedback:

The discrepancy in the trend is that Gaborone at 25°S has a greater temperature range than Durban at 29°S. Clearly, latitude is not the only influence on surface temperature. In this case, the location of Durban on the eastern coast moderates its temperature range more than the inland location of Gaborone. Oceanic and continental properties are among other factors that affect surface temperature. Box 1-2 describes various factors that affect the annual temperature cycle.


Figure 1.23 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.

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.

Jan mean temperature 2m
Jul mean temperature 2m
Fig. 1.23. Mean temperature at 2 m for (a) January and (b) July.
Annual range of monthly surface temperature
Fig. 1.24. Annual range of monthly surface temperature (°C); data from the Japanese Reanalysis Project (JRA25), 1979-2004.

In general, the tropics have the lowest annual temperature range for the globe, less than 1.5°C in the near equatorial regions (Fig. 1.24). 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. The distribution shown in Fig. 1.24 is as expected from the influences described in Box 1-2.

1.6 Temperature »
Box 1-2 Major influences on Annual Surface Temperature Distribution

length of day and insolation
Fig. 1B2.1. Variation of length of day and insolation with latitude. Thick lines mark the equinox and the summer and winter solstices (northern hemisphere).

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 (Fig. 1B2.1). The amount of daylight in the tropics has a small range between solstices.
  • 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 (Fig. 1B2.1, Fig. 1.1).
  • 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 (Fig. 1.5).

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 (Fig. 1.24).

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

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.

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.

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.

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. Extremes of annual temperature in the tropics are presented in a focus sectionfocus section at the end of the chapter.

1.6 Temperature »
1.6.2 Diurnal Temperature Variability in the Tropics

diurnal cycle of surface temperature and the net energy rate
Fig. 1.25. Schematic of the typical diurnal cycle of surface temperature (red) and the net energy rate due to incoming solar (black) and outgoing longwave radiation (blue).

The diurnal variation in temperature has its maximum at the surface, following the daily cycle of surplus heating. Figure 1.25 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.

Can you think of events that may change the typical curve of the diurnal cycle?

Feedback:

One of the biggest influences on the diurnal cycle of temperature is the presence of clouds, which generally reduce the solar heating at the surface and the net incoming longwave radiation received at the surface during the night. The timing of the cloudiness will affect the shape of the temperature curve. The arrival of a cold or warm front can change the shape of the curve. A cold front arriving during the morning causes temperatures to fall during what would typically be a period of increasing temperatures; resulting in a morning temperature maximum. The topography of a station will also affect its diurnal temperature range. Sea/land breezes and mountain/valley breezes can bring air of differing temperatures to a region.


Like the annual temperature range, the diurnal temperature range varies with length of daylight hours. The midlatitude and the poles generally have a larger amplitude range than the tropics. The mean diurnal temperature range is also affected by:

  • 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.

  • 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.

  • Wind speed

During windy conditions air with different temperatures is more easily mixed and helps to moderate the temperature range.

  • 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.

  • 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. 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 (Fig. 1.26). When those temperature differences are large, the pressure gradients can lead to local wind maxima developing at low altitudes.

The combined effect of these influences causes oceanic and coastal regions to have small diurnal temperature ranges while deserts have large diurnal temperature range. The Sahara Desert experiences some of the most extreme diurnal temperature ranges and holds the record for the highest surface temperature (Focus 1).

The effect of elevation and altitude on the range of temperature between day (upper) and night
Fig. 1.26. The effect of elevation and altitude on the range of temperature between day (left) and night (right).

1.7 Moisture and Precipitation

In the tropics, precipitation and the amount of water vapor integrated over a column of air are closely related.23,24 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 (Fig. 1.19b) fuels the tall convective clouds that produce the 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. Competing theories exist for thermodynamic versus dynamic control of tropical surface winds and precipitation fields.25 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.

precipitaion patterns by latitude
Fig. 1.27. Simplified classification of precipitation patterns by latitude.

A simple classification of precipitation by latitude is shown in Fig. 1.27. 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.

1.8 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 × × r = mr2ω.

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.

The absolute velocity is the sum of the angular velocity of the earth and the relative zonal wind speed (Fig. 1.28). 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. 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

angular momentum Equation (10)

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.

Angular momentum component about the axis of rotation of the earth
Conceptual model of the angular momentum balance maintained by transport in the tropics and midlatitudes
Fig. 1.28. (a) Angular momentum component about the axis of rotation of the earth. (b) Conceptual model of the angular momentum balance maintained by transport in the tropics and midlatitudes.

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 (see animation of Fig. 1.28b).

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.

1.8 Role of the Tropics in Momentum Balance »
1.8.1 Spatial and Temporal Scales in the Tropics

Space and time-scales of dynamical atmospheric processes
Fig. 1.29. Space and time-scales of dynamical atmospheric processes.

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. Figure 1.29 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.

1.9 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 (Fig. 1.30). 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).

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.

air mass classifications
Fig. 1.30. Air mass classifications around the globe. Tropical air masses are in light red.

Along with the temperature (Fig. 1.23), the climate of a region is determined by its precipitation characteristics (Fig. 1.27). A more complicated climate classification emerges when the effects of elevation and complex topography are added (Fig. 1.31). The climate classification shown in Fig. 1.31 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.

climate zones classification
Fig. 1.31. The climate zones based on the Köppen classification.

Within the region that we define as the tropics are the following climates:

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 (Fig. 1.31).

  • 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.

  • 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.

  • 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.

  • 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. At high elevation in the tropics, the climate is similar to sub-polar regions; highland climates are found in the tropical regions of Africa, the Americas, and Asia.

Learn more about climate classification in the COMET Module, Introduction to Climatology, http://www.meted.ucar.edu/afwa/climo/intro/main.htm.

Focus Areas

Focus Areas »
Focus 1: Temperature Extremes in the Tropics

Table F1.1 Extreme temperatures recorded in the tropics.
Continent Highest Temp. Place Elevation (m/ft) Date
Africa 57.8 °C/ 136 °F El Azizia, Libya 111.9/ 367 13 Sep 1922
Australia 53.3 °C/ 128 °F* Cloncurry, Queensland 189.6/ 622 16 Jan 1889
Oceania 42.2 °C/ 108 °F Tuguegarao, Philippines 21.9/ 72 29 Apr 1912
Continent Lowest Temp. Place Elevation (m/ft) Date
Africa -23.9 °C/-11°F Ifrane, Morocco 1634.9/ 5364 11 Feb 1935
Oceania -24.4 °C/-12 °F Mauna Kea Observatory, Hawaii 4198/ 13,773 17 May 1979
Adapted from NOAA National Climatic Data Center, NCDC. Sources: The first, "Climates of the World", is an NCDC publication that lists global average temperature and precipitation information for particular locations, with highlighted global extremes. The second publication is the updated "Weather and Climate Extremes" (TEC-0099) published by the US Army Corp of Engineers.
* Note: This temperature was measured using the techniques available at the time of recording, which are different to the standard techniques currently used in Australia. The most likely Australian high-temperature record using standard equipment is an observation of 50.7°C (123°F) recorded at Oodnadatta in 1960.

Focus Areas »
Focus 2: Regional Influences on Tropical Temperature Variability

Here we focus on individual station climographs to illustrate how temperature is influenced by latitude, surface characteristics, prevailing winds (e.g., downstream from mountain or ocean) and oceanic flow (warm and cold currents), clouds and precipitation.

monthly mean temperature
monthly mean temperature
Fig. 1F2.1. Monthly mean temperature for selected stations in northern Africa (upper) and southern African (lower).

The dominance of latitude can be seen in climographs for northern and southern Africa (Fig. 1F2.1). Temperature range is about 3 °C at the near equatorial stations, Mogadishu, Yaounde, and Mombassa and greater than 20 °C at higher tropical latitudes. As noted in Section 1.6.1, although latitude dominates, it is not the only factor. Aswan is inland, gets little precipitation, and therefore has a much warmer summer than Cairo although their winter minima are only a few degrees apart. While Africa straddles the equator, the temperature ranges in the northern hemisphere and southern hemisphere are not mirror images at the similar latitudes. The southern stations have a smaller temperature range than those in the northern hemisphere; for example, compare Gaborone and Aswan. The main cause is the smaller land mass in southern Africa and higher elevation at Gaborone. Note that Walvis Bay and Gaborone are close in latitude but have dramatically different temperature ranges and Walvis Bay is several degrees cooler. Continental Gaborone is warmer and heats and cools at a faster rate than coastal Walvis Bay, which is next to the cold Benguela current.

The impact of continentality is evident when we compare stations at similar latitudes. Inland Timbuktu is warmer and has a greater range than coastal Dakar (Fig. 1F2.2). Although Belem and Manaus are close in latitude and both are along the Amazon River, the temperature graph at coastal Belem is nearly flat while inland Manaus has a range of about 5°C.

monthly mean temperature
monthly mean temperature
Fig. 1F2.2. Monthly mean temperature for selected stations in West Africa (upper) and South America (lower).

Prevailing atmospheric and oceanic flow help explain the differences between stations shown in Fig. 1F2.3. Beira in southeastern Africa is influenced by the warm Mozambique Current while Walvis Bay is next to the cold Benguela current flowing from Antarctica. Salvador on the northeast Brazil coast is under the influence of warm trade winds and warmer ocean currents compared with Lima, Peru where the cold Humboldt Current dominates the climate.

monthly mean temperature
monthly mean temperature
monthly mean temperature
Fig. 1F2.3. Monthly mean temperature for selected stations in southern Africa (upper) and South America (middle) and equatorial Africa (lower).

Even near the equator the ocean temperature factors into the temperature cycle, although the impact is less than stations at higher latitude. Libreville is slightly cooler than east coast Mogadishu because of upwelling along the west coast of Africa.

monthly mean temperature
Fig. 1F2.4. Monthly mean temperature for selected stations in Ecuador, South America.

Elevation is the critical difference between the temperature in Quito and Guayaquil (Fig. 1F2.4). High altitude Quito is more than 10 degrees cooler than sea-level Guayaquil.

 

monthly mean temp
annual mean temp
Fig. 1F2.5. Monthly mean temperature for Kozhikode (Calicut), India and Belize City, Belize.

Clouds and precipitation will reflect sunlight and lead to a decrease in mean surface temperature. That influence is strong in the monsoon regions. Compare Kozhikode (Calicut), India, which experiences the monsoon, to Belize City, which is mainly in the trade wind regime (Fig. 1F2.5). For Kozhikode, the temperature is at its maximum in May, just prior to the monsoon cloudiness and change in the prevailing wind; it decreases with the monsoon rains to a minimum in July then increases slightly thereafter. Meanwhile in Belize City, the rainfall is fairly constant with only a slight reduction during winter and the temperature gradient is small with a broad maximum from June to September. Note that although higher in latitude than Kozhikode, Belize City is coastal and downstream from the Caribbean Sea, which moderates the temperature.

Summary

In this chapter we demonstrated the importance of the tropics in the global energy balance and climate. Annually-averaged, the earth-atmosphere system is in radiation balance, but the tropics are a region of surplus heating while the poles are a region of net cooling. The atmosphere and the ocean circulations act to remove the heating imbalance. Surplus heating in the tropics drives the general circulation and energy in the tropical troposphere is transported upward in deep convective clouds. Surface to atmosphere energy transfer in the tropics is mostly by latent heat (evaporation from the tropical oceans and condensation in tropical cumulus). Most of the upward transport is concentrated in deep convection over the tropical continents and oceanic warm pool regions.

We examined different methods by which the tropics are defined. For this text, the tropics exist between subtropical high pressure regions whose latitudinal extent depends on the location of net surface heating.

As a result of the surplus heating and strong upward motion in tropical convection, the tropical tropopause is much higher than the tropopause at higher latitudes. The atmosphere is heated from below by latent heat, sensible heat, and radiation. The surface energy budget consists of those contributions as well as storage and horizontal advection of energy. The atmospheric energy budget is dominated by the sensible heat, latent heat, and potential energy while the kinetic energy is very small. The atmosphere is responsible for most the energy transport towards the poles but the ocean dominates at low latitudes. The peak in poleward energy transport occurs during the winter. The Hadley cells transport most of the energy from the tropics to the subtropics while waves or cyclones are responsible for most of the transport in the midlatitudes. Methods for calculating zonal, meridional, and time averages were described.

We examined the seasonal, geographic, and diurnal distribution of temperature, the most important climate variable. In general, surface temperature is controlled by latitude, continentality, prevailing atmospheric and oceanic flow, relief, slope aspect, clouds and precipitation cycles, and surface albedo.

Because of the relationship between saturation vapor pressure and temperature, like temperature, average water vapor content decreases with latitude and altitude in the troposphere although moisture profiles are more variable than temperature profiles. On average, the profiles of temperature and moisture are a study in contrast between the eastern and western tropical ocean. On the eastern side, subsidence in the subtropical highs and cool SSTs are the norm with upwelling of cool water where surface currents are driven by the trade winds. The result is an inversion, known as the trade wind inversion. Warmer waters are pushed westward; the inversion weakens towards the west; and finally disappears. The atmosphere over the western ocean is unstable and conducive to the development of deep cumulus clouds.

The tropics also play a significant role in maintaining angular momentum balance. In the tropical easterlies, the atmosphere gains angular momentum from the surface and that momentum is transferred upward and poleward by the Hadley cells.

Finally, we explored tropical air masses as well as tropical climates, their definitions, and spatial distribution across the global tropics.

Questions for Review

  • Explain the concept of radiation balance in the earth and atmosphere system.
  • List the types of energy transfer that occur between the surface and the atmosphere.
  • You were introduced to various methods of defining the tropics, please discuss at least three methods and give their approximate latitudinal boundaries.
  • Describe the components of the surface energy budget and rank them in terms of their relative amounts in the atmosphere.
  • Describe the relative contributions and difference between the meridional transport by the ocean and the atmosphere.
  • Compare the relative contributions of the Hadley cell and transient waves to atmospheric energy transport from the tropics to the poles in the northern hemisphere.
  • Explain how latent heat and tropical deep convection fit into the global energy balance.
  • Why is the tropopause highest in the tropics?
  • Describe the trade wind inversion and its impact on east-west cloud distribution over the tropical oceans.
  • Describe the general distribution of surface water vapor content from the tropics to the poles and explain the reason for the distribution.
  • What is the typical range of atmospheric surface pressure between the subtropics and the equator?
  • Describe the geographic and seasonal distribution of surface temperature and the factors that influence that distribution.
  • Describe the factors that influence the diurnal temperature cycle.
  • What is the role of the tropics in maintaining the angular momentum of the earth-atmosphere system?
  • List three tropical air masses, their characteristics, and regions where each type of air masses originates.
  • List at least three tropical climates, their temperature and precipitation characteristics, and a region where each may be found.

QUIZ

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Biographies

Herbert Riehl (1915-1997)

Herbert Riehl wrote the first “Tropical Meteorology” textbook and conducted some of the foundational research on tropical cyclones, easterly waves, and tropical meteorology. Interestingly, this widely regarded “father of tropical meteorology” was born in Munich, Germany, in 1915 and immigrated to the United States in 1933. He earned an M.S. in meteorology from New York University in 1942 and a Ph. D. from the University of Chicago in 1947. Dr. Riehl conducted research in Puerto Rico when the U.S. Army Air Corps and the University of Chicago created an Institute of Tropical Meteorology at the University of Puerto Rico in 1943 to facilitate studying equatorial meteorology. He directed the institute from 1945-1946. He was a faculty member at the University of Chicago from 1947-1960, after which he moved to Fort Collins, Colorado, to help found and head Colorado State University’s Department of Atmospheric Science in 1960. He stepped down as department head in 1968 but continued working there until 1972. In that year he moved back to Germany to direct the Institute of Meteorology and Geophysics at the Free University in Berlin. He returned to the United States in 1976 to join the National Center for Atmospheric Research and the Cooperative Institute for Research in Environmental Sciences in Boulder, Colorado as a senior scientist. He retired in 1989, but continued to consult in tropical meteorology research in North and South America.

In addition to his work on tropical meteorology, he also studied the global general circulation, jet streams, mid-latitude forecasting, the atmospheric water cycle, climate change, and air pollution. Riehl oversaw the Ph.D. work of Joanne and Robert Simpson, Noel LaSeur, Charles Jordan, Mike Alaka, Jose Colon, T.N. Krishnamurti, William Gray, James Rasmussen, and Russell Elsberry. He wrote the landmark textbook, “Tropical Meteorology” in 1954, “Climate and Weather in the Tropics” in 1979, and “The Hurricane and Its Impact” in 1981, the latter with Robert Simpson, as well as than 150 journal articles. He was awarded the Losey Award of the American Institute for Aeronautics and Astronautics in 1962 and the American Meteorological Society’s Carl-Gustaf Rossby Research Medal in 1979. The CSU Department of Atmospheric Science grants the annual Herbert Riehl Award to a graduate student who submits the best technical manuscript for publication in the refereed literature.

Joanne Simpson (1923-2010)

Joanne Simpson blazed trails in both meteorology and women’s progress when she became the first woman to earn a Ph.D. in meteorology. She sparked the field of cloud modeling when she developed the first mathematical model of clouds using only a slide rule. She helped explain the connection between the heat transported in “hot towers”, tall cumulus clouds in the tropics, and the forces driving the trade winds and poleward motion in the Hadley Cells, a theory that also shed light on the heat engine of hurricanes. She also conducted research into cloud seeding and weather modification as part of her study of clouds, and finally worked on a tropical satellite project that resulted in the first satellite radar profile of precipitation and latent heat estimates from tropical clouds. Dr. Simpson began studying meteorology while she was a student pilot during World War II.

Enthused by the subject, she sought a Ph. D. only to be warned by a faculty advisor that “No woman has ever earned a Ph. D. in meteorology. No woman ever will. Even if you did, no one would give you a job.” She went on to disprove this advice, not only earning a Ph.D. in meteorology from the University of Chicago in 1949, but even landing a job as an Assistant Professor of Physics at the Illinois Institute of Technology. Later she worked at Woods Hole Oceanographic Institution, UCLA, NOAA, the University of Virginia, and finally NASA’s Goddard Space Flight Center as Chief Scientist for meteorology. Among her most famous accomplishments was her work with Herbert Riehl in the 1950s, describing tropical atmospheric circulation through “hot towers”, which determine the prevailing winds in the surrounding latitudes and govern how heat and moisture move from the tropics to the mid-latitudes. At NASA, she spearheaded the Tropical Rainfall Measuring Mission (TRMM), which she considered her greatest achievement and which helped scientists understand tropical cyclone genesis, dust and smoke’s influence on rainfall, and how much latent heat is released by tropical clouds. She was also a mentor to countless scientists of both genders, either in person or through her writings, speeches, and actions.

Dr. Simpson was awarded the American Meteorological Society’s highest prize, the Carl-Gustav Rossby Research Medal in 1983, and became the AMS’s first female president in 1989. She was also made a member of the National Academy of Engineering and was the first woman to win the International Meteorological Organization Prize in 2002. Greg Holland, director of the Earth Systems Laboratory at the National Center for Atmospheric Research in Boulder, Colo., summed her up in her obituary in the Washington Post: "There is zero doubt that there has never been a more capable woman in meteorology, and she would also be in the top five of all meteorologists in history, no matter the gender.”

References

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2. Matsuno, T., 1966: Quasi-geostrophic motions in the equatorial area. J. Meteor. Soc. Japan, 44, 25-43.
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4. Madden, R., P. Julian, 1971: Detection of a 40-50 day oscillation in the zonal wind in the tropical Pacific. J. Atmos. Sci., 28, 702-708.
5. Madden, R., P. R. Julian, 1972: Description of global scale circulation cells in the Tropics with 40–50 day period. J. Atmos. Sci., 29, 1109-1123.
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