DOCSLIB.ORG

Atmosphere and Climate

Total Page：16

File Type：pdf, Size：1020Kb

Download full-text PDF Read full-text
Atmosphere and Climate Troposphere The troposphere is the lowest atmospheric layer reaching an altitude of about 20 km. Density, pressure and temperature decline with altitude. The troposphere is largely convective, which translates into a uniform chemical composition. Critical component is H2O vapor from evaporation and transpiration processes. Water vapor decreases with altitude because of condensation and cloud formation. Temperature profile of the Troposphere Q T W Heat transfer in air: k A kair 0.026 thermalconductivity t z mK Heat transfer takes place by energy transfer through cross sectional area A! Vertical convective motion takes place by moving volume of air in a reversible adiabatic process (no heat exchange with surrounding environment Q=0!) This is subject to energy conservation according to the first law of thermodynamics: dU Q W Q PdV Heat exchange between air packet and surrounding is: kJ dQ C dT P dV C 1.010 v v Specific heat or heat capacity of K kg an air volume element at constant kJ volume or constant pressure. CP 1.005 K kg Adiabatic conditions: Q 0 Cv dT P dV For adiabatic conditions Q 0 Cv dT P dV dU Internal energy for V const : dU C dT C v v dT dU dQ specific heat for V const : Cv for P const :CP dT V dT P Gas Cp J/mole·K Cv J/mole·K He 20.79 12.52 Ar 20.79 12.45 Universal gas constant N 29.12 20.8 2 R = 8.314 J/mole·K O2 29.37 20.98 CO 29.04 20.74 R CP Cv CO2 36.62 28.17 H2O vapor 37.47 28.03 Conversion: 1 mole ≡ A g air 29.07 20.95 To derive the temperature profile along the vertical atmosphere axis as function of altitude T(z) requires differentiating the ideal gas equation P·V=R·T for Q=0: PdV V dP RdT CP Cv dT R dT CP dT Cv dT CP dT P dV P dV V dP CP dT V dP for adiabatic element dP Reformulating the hydrostatic equation: g g 9.81m / s2 dz z V dP V g dz m g dz C dT m g dz z P Lapse rate defines the rate at dT m g which temperature T decreasesR C CT dz C with increasing altitude z. P v P (For the troposphere a negative value) T Adiabatic lapse rate: =-9.7 K/km for dry air. The adiabatic lapse rate describes the temperature change with altitude for absolutely dry air, it serves as a normalization parameter since air is never dry. CP increases with humidity, CP=1.88 KJ/km K for H2O vapor. An increase in humidity decreases the temperature gradient. The actual lapse rate L depends on humidity and is around 6-7 K/km! dT g V L dz CP K 6 7 km Higher humidity increases CP and changes the temperature gradient that -L < -Γ or L > Γ. Flight height of 10 km, T=-50oC=223K If the density ρ in an air volume element V is smaller than the density ρs of surrounding environment, the buoyancy causes acceleration upward, if density of air volume is larger, element will stabilize. dT V g smaller L dz CP Unstable situation, convection dT V g larger L dz CP stable situation, no convection wet g dTair V dz L dz Convective equilibrium CP z dT dry g dT V dz dz dz dry CP convection m 0.828 0.232 29 amu dry 3 radiation V V V m m 0.7228 0.1832 0.118 27.7 amu wet 3 T V V V m Clouds Rising Cloud physics and weather see next chapter The stratosphere The stratosphere is the second layer of the atmosphere above the troposphere ranging from about 20 km to 50 km altitude. The density declines in this layer from 100 g/m3 to 3 g/m3, the pressure drops exponentially from 100mb to 1mb. The temperature remains constant at the lower layer of the stratosphere and then rises with increasing altitude. At the top of the stratosphere the thin air may reach temperatures close to 0°C (273 K). Temperature profile of Stratosphere ~20-50km altitude f Temperature is dominated by radiative energy balance f Incoming flux must be balanced by outgoing flux. The factor f is the fraction of transmitted radiation. F T 4 f T 4 f in E S 4 4 Fout 1 f TE f TS f 4 4 4 4 TE f TS 1 f TE f TS 4 4 4 TE 1 f TE 2 f TS T T 4 2 T 4 T E E S S 21/ 4 TE = 255 K TS = 215 K This suggests that temperature in stratosphere remains constant. This is only partially correct since ozone layer in stratosphere causes further absorption and an increase of the stratosphere temperature up to 270K. Chemical composition of Stratosphere Temperature is low and constant for ~10 km, that region is frequently called tropopause. Towards higher altitude the temperature increases due to the absorption of cosmic radiation, reaching a maximum of about 0oC (273 K) at top of stratosphere. This is associated with the absorption of UV radiation by Oxygen and Ozone causing stratosphere heating in its upper layers. UV range Nuclear test program deposition A total of 2057 nuclear weapon explosions since 1945 51017 Bq=500 PBq Mesosphere Mesosphere is the atmospheric layer above stratosphere. It is characterized by steadily decreasing density. High altitude (~80km) location for rare kind of clouds (noctilucent clouds) in polar zone, frequency of occurrence seems to increase (signature for climate change, consequence of space shuttle exhaust?) NLC: first observed 1885 (two years after the Krakatau eruption), increasing appearance of noctilucent clouds is interpreted as a consequence of a decrease in temperature at the altitude where the clouds form parallel with an increase in water vapor by complex NLC formation requires a combination of photochemical processes. The warming very low temperatures, a source of water in the in the lower atmospheric layers vapor, and condensation points (meteoritic causes the cloud layer to cool dust, volcano ashes, aerosols?) increasing the probability of formation. Mesosphere is too high to be reached by balloon or aircraft and is therefore poorly understood! New satellites such as AIM mission are launched for more detailed exploration. Temperature declines towards its lowest value in the atmosphere of about T=-80 oC (193K). Limited absorption of the solar radiation flux and possible cooling by de-excitation of CO2 vibrational excitation modes by collision with oxygen atoms, which are in equilibrium at higher density in lower atmospheric regions. 960 Q 2 e T collision rate This leads to the discussion of possible enhanced cooling with increasing CO2 budget! Aerosol impact on stratosphere temperature Example: Mount Pinatubo eruption in June 13, 1991 causing emission of particles into higher atmosphere with direct impact on optical depth (·d) and temperature T. Optical depth for 1020 nm range F d e d 40km; 0 d 0.001 v d 0.06 F0 F F e0 d 99% v ev d 94% F0 F0 Impact on high altitude temperature Heavy metal enrichment Large numbers of meteorites are absorbed by During the early phase of the planet atmosphere, they typically evaporate in the formation, the accretion rate was in mesosphere, enriching the layer with about 40 the range of more than millions /sec tons/day of heavy metal containing dust! from meteors to planetesimals. FREQUENCY OF IMPACTORS: Pea-size meteoroids - 10 per hour 2.7 Walnut-size - 1 per hour N 37 D Grapefruit-size - 1 every 10 hours Basketball-size - 1 per month N: number of meteorites 50-m rock - 1 per 100 years D: diameter of meteorites 1-km asteroid - 1 per 100,000 years 2-km asteroid - 1 per 500,000 years Meteorites fragment in higher atmosphere layers and the fragments evaporate in the mesosphere. Meteoritic material is enriched in heavy elements (iron meteorites) which form dust particles in the high altitude layers. .
Load more

Recommended publications
  • The California Low-Level Coastal Jet and Nearshore Stratocumulus
    Calhoun: The NPS Institutional Archive Faculty and Researcher Publications Faculty and Researcher Publications 2001 The California Low-Level Coastal Jet and Nearshore Stratocumulus Kalogiros, John http://hdl.handle.net/10945/34433 1.1 THE CALIFORNIA LOW-LEVEL COASTAL JET AND NEARSHORE STRATOCUMULUS John Kalogiros and Qing Wang* Naval Postgraduate School, Monterey, California 1. INTRODUCTION* and the warm air above land (thermal wind) and the frictional effect within the atmospheric This observational study focus on the boundary layer (Zemba and Friehe 1987, Burk and interaction between the coastal wind field and the Thompson 1996). The horizontal temperature evolution of the coastal stratocumulus clouds. The gradient follows the orientation of the coastline data were collected during an experiment in the (317° on the average in central California). Thus, summer of 1999 near the California coast with the the thermal wind has a northern and an eastern Twin Otter research aircraft operated by the component and local maximum of both Center for Interdisciplinary Remote Piloted Aircraft components of the wind is expected close to the Study (CIRPAS) at the Naval Postgraduate School top of the boundary layer. Topographical features (NPS). The wind components were measured with may intensify this wind jet at significant convex a DGPS/radome system, which provides a very bends of the coastline like Cape Mendocino, Pt accurate (0.1 ms-1) estimate of the wind (Kalogiros Arena and Pt Sur. At such changes of the and Wang 2001). The instrumentation of the coastline geometry combined with mountain aircraft included fast sensors for the measurement barriers (channeling effect) the northerly flow can of air temperature, humidity, shortwave and become supercritical.
    [Show full text]
  • Using the Weak Temperature Gradient Approximation to Understand Self-Aggregation 1
    Multiple Equilibria in a Cloud Resolving Model: Using the Weak Temperature Gradient Approximation to Understand Self-Aggregation 1 Sharon Sessions, Satomi Sugaya, David Raymond, Adam Sobel New Mexico Tech SWAP 2011 nmtlogo 1This work is supported by the National Science Foundation Sessions (NMT) Multiple Equilibria in a CRM May2011 1/21 Self-Aggregation Bretherton et al (2005) precipitation 1 WVP = g rt dp R nmtlogo Day 1 Day 50 Sessions (NMT) Multiple Equilibria in a CRM May2011 2/21 Possible insight from limited domain WTG simulations Bretherton et al (2005) Day 1 Day 50 nmtlogo Sessions (NMT) Multiple Equilibria in a CRM May2011 3/21 Weak Temperature Gradients in the Tropics Charney 1963 scaling analysis Extratropics (small Rossby number) F δθ ∼ r θ Ro Tropics (Rossby number not small) δθ ∼ Fr θ 2 −3 Froude number: Fr = U /gH ∼ 10 , measure of stratification −5 Rossby number: Ro = U/fL ∼ 10 /f , characterizes rotation nmtlogo Sessions (NMT) Multiple Equilibria in a CRM May2011 4/21 Weak Temperature Gradients in the Tropics Charney 1963 scaling analysis Extratropics (small Rossby number) F δθ ∼ r θ Ro Tropics (Rossby number not small) δθ ∼ Fr θ 2 −3 Froude number: Fr = U /gH ∼ 10 , measure of stratification −5 Rossby number: Ro = U/fL ∼ 10 /f , characterizes rotation In the tropics, gravity waves rapidly redistribute buoyancy anomalies nmtlogo Sessions (NMT) Multiple Equilibria in a CRM May2011 4/21 Weak Temperature Gradient (WTG) Approximation Sobel & Bretherton 2000 Single column model (SCM) representing one grid cell in a global circulation
    [Show full text]
  • NWS Unified Surface Analysis Manual
    Unified Surface Analysis Manual Weather Prediction Center Ocean Prediction Center National Hurricane Center Honolulu Forecast Office November 21, 2013 Table of Contents Chapter 1: Surface Analysis – Its History at the Analysis Centers…………….3 Chapter 2: Datasets available for creation of the Unified Analysis………...…..5 Chapter 3: The Unified Surface Analysis and related features.……….……….19 Chapter 4: Creation/Merging of the Unified Surface Analysis………….……..24 Chapter 5: Bibliography………………………………………………….…….30 Appendix A: Unified Graphics Legend showing Ocean Center symbols.….…33 2 Chapter 1: Surface Analysis – Its History at the Analysis Centers 1. INTRODUCTION Since 1942, surface analyses produced by several different offices within the U.S. Weather Bureau (USWB) and the National Oceanic and Atmospheric Administration’s (NOAA’s) National Weather Service (NWS) were generally based on the Norwegian Cyclone Model (Bjerknes 1919) over land, and in recent decades, the Shapiro-Keyser Model over the mid-latitudes of the ocean. The graphic below shows a typical evolution according to both models of cyclone development. Conceptual models of cyclone evolution showing lower-tropospheric (e.g., 850-hPa) geopotential height and fronts (top), and lower-tropospheric potential temperature (bottom). (a) Norwegian cyclone model: (I) incipient frontal cyclone, (II) and (III) narrowing warm sector, (IV) occlusion; (b) Shapiro–Keyser cyclone model: (I) incipient frontal cyclone, (II) frontal fracture, (III) frontal T-bone and bent-back front, (IV) frontal T-bone and warm seclusion. Panel (b) is adapted from Shapiro and Keyser (1990) , their FIG. 10.27 ) to enhance the zonal elongation of the cyclone and fronts and to reflect the continued existence of the frontal T-bone in stage IV.
    [Show full text]
  • Thermal Conductivity of Materials
    • DECEMBER 2019 Thermal Conductivity of Materials Thermal Conductivity in Heat Transfer – Lesson 2 What Is Thermal Conductivity? Based on our life experiences, we know that some materials (like metals) conduct heat at a much faster rate than other materials (like glass). Why is this? As engineers, how do we quantify this? • Thermal conductivity of a material is a measure of its intrinsic ability to conduct heat. Metals conduct heat much faster to our hands, Plastic is a bad conductor of heat, so we can touch an which is why we use oven mitts when taking a pie iron without burning our hands. out of the oven. 2 What Is Thermal Conductivity? Hot Cold • Let’s recall Fourier’s law from the previous lesson: 푞 = −푘∇푇 Heat flow direction where 푞 is the heat flux, ∇푇 is the temperature gradient and 푘 is the thermal conductivity. • If we have a unit temperature gradient across the material, i.e.,∇푇 = 1, then 푘 = −푞. The negative sign indicates that heat flows in the direction of the negative gradient of temperature, i.e., from higher temperature to lower temperature. Thus, thermal conductivity of a material can be defined as the heat flux transmitted through a material due to a unit temperature gradient under steady-state conditions. It is a material property (independent of the geometry of the object in which conduction is occurring). 3 Measuring Thermal Conductivity • In the International System of Units (SI system), thermal conductivity is measured in watts per -1 -1 meter-Kelvin (W m K ). Unit 푞 W ∙ m−1 푞 = −푘∇푇 푘 = − (W ∙ m−1 ∙ K−1) ∇푇 K • In imperial units, thermal conductivity is measured in British thermal unit per hour-feet-degree- Fahrenheit (BTU h-1ft-1 oF-1).
    [Show full text]
  • Atmospheric Stability Atmospheric Lapse Rate
    ATMOSPHERIC STABILITY ATMOSPHERIC LAPSE RATE The atmospheric lapse rate ( ) refers to the change of an atmospheric variable with a change of altitude, the variable being temperature unless specified otherwise (such as pressure, density or humidity). While usually applied to Earth's atmosphere, the concept of lapse rate can be extended to atmospheres (if any) that exist on other planets. Lapse rates are usually expressed as the amount of temperature change associated with a specified amount of altitude change, such as 9.8 °Kelvin (K) per kilometer, 0.0098 °K per meter or the equivalent 5.4 °F per 1000 feet. If the atmospheric air cools with increasing altitude, the lapse rate may be expressed as a negative number. If the air heats with increasing altitude, the lapse rate may be expressed as a positive number. Understanding of lapse rates is important in micro-scale air pollution dispersion analysis, as well as urban noise pollution modeling, forest fire-fighting and certain aviation applications. The lapse rate is most often denoted by the Greek capital letter Gamma ( or Γ ) but not always. For example, the U.S. Standard Atmosphere uses L to denote lapse rates. A few others use the Greek lower case letter gamma ( ). Types of lapse rates There are three types of lapse rates that are used to express the rate of temperature change with a change in altitude, namely the dry adiabatic lapse rate, the wet adiabatic lapse rate and the environmental lapse rate. Dry adiabatic lapse rate Since the atmospheric pressure decreases with altitude, the volume of an air parcel expands as it rises.
    [Show full text]
  • Thermal Equilibrium of the Atmosphere with a Convective Adjustment
    jULY 1964 SYUKURO MANABE AND ROBERT F. STRICKLER 361 Thermal Equilibrium of the Atmosphere with a Convective Adjustment SYUKURO MANABE AND ROBERT F. STRICKLER General Circulation Research Laboratory, U.S. Weather Bureau, Washington, D. C. (Manuscript received 19 December 1963, in revised form 13 April 1964) ABSTRACT The states of thermal equilibrium (incorporating an adjustment of super-adiabatic stratification) as well as that of pure radiative equilibrium of the atmosphere are computed as the asymptotic steady state ap­ proached in an initial value problem. Recent measurements of absorptivities obtained for a wide range of pressure are used, and the scheme of computation is sufficiently general to include the effect of several layers of clouds. The atmosphere in thermal equilibrium has an isothermal lower stratosphere and an inversion in the upper stratosphere which are features observed in middle latitudes. The role of various gaseous absorbers (i.e., water vapor, carbon dioxide, and ozone), as well as the role of the clouds, is investigated by computing thermal equilibrium with and without one or two of these elements. The existence of ozone has very little effect on the equilibrium temperature of the earth's surface but a very important effect on the temperature throughout the stratosphere; the absorption of solar radiation by ozone in the upper and middle strato­ sphere, in addition to maintaining the warm temperature in that region, appears also to be necessary for the maintenance of the isothermal layer or slight inversion just above the tropopause. The thermal equilib­ rium state in the absence of solar insolation is computed by setting the temperature of the earth's surface at the observed polar value.
    [Show full text]
  • ATM 316 - the “Thermal Wind” Fall, 2016 – Fovell
    ATM 316 - The \thermal wind" Fall, 2016 { Fovell Recap and isobaric coordinates We have seen that for the synoptic time and space scales, the three leading terms in the horizontal equations of motion are du 1 @p − fv = − (1) dt ρ @x dv 1 @p + fu = − ; (2) dt ρ @y where f = 2Ω sin φ. The two largest terms are the Coriolis and pressure gradient forces (PGF) which combined represent geostrophic balance. We can define geostrophic winds ug and vg that exactly satisfy geostrophic balance, as 1 @p −fv ≡ − (3) g ρ @x 1 @p fu ≡ − ; (4) g ρ @y and thus we can also write du = f(v − v ) (5) dt g dv = −f(u − u ): (6) dt g In other words, on the synoptic scale, accelerations result from departures from geostrophic balance. z R p z Q δ δx p+δp x Figure 1: Isobaric coordinates. It is convenient to shift into an isobaric coordinate system, replacing height z by pressure p. Remember that pressure gradients on constant height surfaces are height gradients on constant pressure surfaces (such as the 500 mb chart). In Fig. 1, the points Q and R reside on the same isobar p. Ignoring the y direction for simplicity, then p = p(x; z) and the chain rule says @p @p δp = δx + δz: @x @z 1 Here, since the points reside on the same isobar, δp = 0. Therefore, we can rearrange the remainder to find δz @p = − @x : δx @p @z Using the hydrostatic equation on the denominator of the RHS, and cleaning up the notation, we find 1 @p @z − = −g : (7) ρ @x @x In other words, we have related the PGF (per unit mass) on constant height surfaces to a \height gradient force" (again per unit mass) on constant pressure surfaces.
    [Show full text]
  • Thickness and Thermal Wind
    ESCI 241 – Meteorology Lesson 12 – Geopotential, Thickness, and Thermal Wind Dr. DeCaria GEOPOTENTIAL l The acceleration due to gravity is not constant. It varies from place to place, with the largest variation due to latitude. o What we call gravity is actually the combination of the gravitational acceleration and the centrifugal acceleration due to the rotation of the Earth. o Gravity at the North Pole is approximately 9.83 m/s2, while at the Equator it is about 9.78 m/s2. l Though small, the variation in gravity must be accounted for. We do this via the concept of geopotential. l A surface of constant geopotential represents a surface along which all objects of the same mass have the same potential energy (the potential energy is just mF). l If gravity were constant, a geopotential surface would also have the same altitude everywhere. Since gravity is not constant, a geopotential surface will have varying altitude. l Geopotential is defined as z F º ò gdz, (1) 0 or in differential form as dF = gdz. (2) l Geopotential height is defined as F 1 z Z º = ò gdz (3) g0 g0 0 2 where g0 is a constant called standard gravity, and has a value of 9.80665 m/s . l If the change in gravity with height is ignored, geopotential height and geometric height are related via g Z = z. (4) g 0 o If the local gravity is stronger than standard gravity, then Z > z. o If the local gravity is weaker than standard gravity, then Z < z.
    [Show full text]
  • Temperature Gradient
    Temperature Gradient Derivation of dT=dr In Modern Physics it is shown that the pressure of a photon gas in thermodynamic equilib- rium (the radiation pressure PR) is given by the equation 1 P = aT 4; (1) R 3 with a the radiation constant, which is related to the Stefan-Boltzman constant, σ, and the speed of light in vacuum, c, via the equation 4σ a = = 7:565731 × 10−16 J m−3 K−4: (2) c Hence, the radiation pressure gradient is dP 4 dT R = aT 3 : (3) dr 3 dr Referring to our our discussion of opacity, we can write dF = −hκiρF; (4) dr where F = F (r) is the net outward flux (integrated over all wavelengths) of photons at a distance r from the center of the star, and hκi some properly taken average value of the opacity. The flux F is related to the luminosity through the equation L(r) F (r) = : (5) 4πr2 Considering that pressure is force per unit area, and force is a rate of change of linear momentum, and photons carry linear momentum equal to their energy divided by c, PR and F obey the equation 1 P = F (r): (6) R c Differentiation with respect to r gives dP 1 dF R = : (7) dr c dr Combining equations (3), (4), (5) and (7) then gives 1 L(r) 4 dT − hκiρ = aT 3 ; (8) c 4πr2 3 dr { 2 { which finally leads to dT 3hκiρ 1 1 = − L(r): (9) dr 16πac T 3 r2 Equation (9) gives the local (i.e.
    [Show full text]
  • Air Pressure and Wind
    Air Pressure We know that standard atmospheric pressure is 14.7 pounds per square inch. We also know that air pressure decreases as we rise in the atmosphere. 1013.25 mb = 101.325 kPa = 29.92 inches Hg = 14.7 pounds per in 2 = 760 mm of Hg = 34 feet of water Air pressure can simply be measured with a barometer by measuring how the level of a liquid changes due to different weather conditions. In order that we don't have columns of liquid many feet tall, it is best to use a column of mercury, a dense liquid. The aneroid barometer measures air pressure without the use of liquid by using a partially evacuated chamber. This bellows-like chamber responds to air pressure so it can be used to measure atmospheric pressure. Air pressure records: 1084 mb in Siberia (1968) 870 mb in a Pacific Typhoon An Ideal Ga s behaves in such a way that the relationship between pressure (P), temperature (T), and volume (V) are well characterized. The equation that relates the three variables, the Ideal Gas Law , is PV = nRT with n being the number of moles of gas, and R being a constant. If we keep the mass of the gas constant, then we can simplify the equation to (PV)/T = constant. That means that: For a constant P, T increases, V increases. For a constant V, T increases, P increases. For a constant T, P increases, V decreases. Since air is a gas, it responds to changes in temperature, elevation, and latitude (owing to a non-spherical Earth).
    [Show full text]
  • Aerosol Indirect Effects on the Temperature-Precipitation Scaling
    Atmos. Chem. Phys. Discuss., https://doi.org/10.5194/acp-2018-1334 Manuscript under review for journal Atmos. Chem. Phys. Discussion started: 28 January 2019 c Author(s) 2019. CC BY 4.0 License. Aerosol indirect effects on the temperature-precipitation scaling Nicolas Da Silva1, Sylvain Mailler1,2, and Philippe Drobinski1 1LMD/IPSL, École polytechnique, Université Paris Saclay, ENS, PSL Research University, Sorbonne Universités, UPMC Univ Paris 06, CNRS, Palaiseau, France 2ENPC, Champs-sur-Marne, France Correspondence: Nicolas Da Silva ([email protected]) Abstract. Indirect effects of aerosols were found to weaken convective precipitation through reduced precipitable water and convective instability. The present study aims at quantifying the relative importance of these two processes in the reduction of summer precipitation using the temperature-precipitation scaling. Based on a numerical sensitivity experiment conducted over central 5 Europe aiming to isolate indirect effects, all others effects being equal, the results show that the scaling of hourly convective precipitation with temperature follows the Clausius-Clapeyron (CC) relationship whereas the decrease of convective precipita- tion does not scale with the CC law since it is mostly attributable to increased stability with increased aerosols concentrations rather than to decreased precipitable water content. This effect is larger at low temperatures for which clouds are statistically more frequent and optically thicker. At these temperatures, the increase of stability is mostly linked to the stronger reduction 10 of temperature in the lower troposphere compared to the upper troposphere which results in lower lapse rates. 1 Introduction The temperature-precipitation relationship has often been studied because it has been hypothesised to give an insight of the change of precipitation in a warming climate.
    [Show full text]
  • Appendix a Fundamentals of Heat Transfer
    Appendix A Fundamentals of Heat Transfer A.1 Heat Conduction When a temperature gradient exists in a body, energy will transfer from the high- temperature region to the low-temperature region by conduction. The general equa- tions of heat conduction is given by [1] ∂T ∇·(k∇T) +˙q = ρc (A.1) in ∂t where T is temperature; q˙in is energy generated per unit volume; k, ρ, and c are thermal conductivity, density and specific heat, respectively. In fire engineering, one-dimensional (1D) heat transfer is usually considered. The equation of 1D conduction is given by (ignore energy generation, thus q˙in = 0) ∂2T 1 ∂T = (A.2) ∂x2 α ∂t where α = kρc is thermal diffusivity. Fourier’s law states that the quantity of heat transferred per unit time per unit area is proportional to the temperature gradient, thus ∂T q˙ =−k (A.3) ∂x To solve Eq. A.2, the following boundary conditions are usually used [2, 3] 1. Initial condition, T(x, 0) = T∞ (A.4) 2. Dirichlet condition, T(0, t) = Tg(t) (A.5) © Springer-Verlag Berlin Heidelberg 2015 113 C. Zhang, Reliability of Steel Columns Protected by Intumescent Coatings Subjected to Natural Fires, Springer Theses, DOI 10.1007/978-3-662-46379-6 114 Appendix A: Fundamentals of Heat Transfer 3. Neumann condition, ∂T − k (0, t) =˙q = h[T (t) − T(0, t)] (A.6) ∂x 0 g where, x denotes opposite normal direction to the surface; T∞, Tg are environment temperature and fire temperature, respectively; q˙0 is heat flux per unit time transferred from fire environment to the surface; and h is coefficient for convection or radiation.
    [Show full text]