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Atmospheric Physics and Sustainable Aviation

CREATE Summer School

Atmospheric Physics and Sustainable Aviation (K. Strong) Slide 1

CREATE Summer School on Sustainable Aviation

Kimberly Strong

Department of Physics, University of Toronto

14 May 2013

Kimberly StrongDept of Physics, U of Toronto

Remote sounding of atmospheric composition from the ground, balloons, and satellites using UV-VIS-IR spectroscopy

Leader of the CANDAC/PEARL Arctic Middle Atmospheric Chemistry theme

Director of the NSERC CREATE Training Program in Arctic Atmospheric Science

PI, U of Toronto Atmospheric Observatory

ACE & Odin missions

Laboratory spectroscopy

Mars studies May 2011

Plan for the Day

Tuesday, May 14th

9:00am – 10:30am

Atmospheric Physics and Sustainable AviationPart 1: Introduction, Structure of the Atmosphere, Radiation, Chemistry, Clouds & Aerosols, General Circulation

10:30am – 10:45am Break with Refreshments

At h i Ph i d S t i bl A i ti


10:45am – 12:00pmAtmospheric Physics and Sustainable AviationPart 2: Stratospheric Ozone

12:00pm - 1:00pm Lunch

1:00pm – 2:30pmAtmospheric Physics and Sustainable AviationPart 3: Climate Change

2:30pm – 2:45pm Break with Refreshments

2:45pm – 4:00pmAtmospheric Physics and Sustainable AviationPart 4: Aviation and the Atmosphere

References: Atmospheric Physics Atmospheric Science, An Introductory Survey, 2nd Edition,

John M. Wallace & Peter V. Hobbs, Academic Press, 2006 http://www.elsevierdirect.com/v2/companion.jsp?ISBN=9780127329512

Atmosphere, Ocean, and Climate Dynamics, John Marshall & R. Alan Plumb, Academic Press, 2008 http://store.elsevier.com/product.jsp?isbn=9780125586917

Introduction to Atmospheric Physics, David G. Andrews, Cambridge University Press, 2000 http://www cambridge org/gb/knowledge/isbn/item1158689/?site locale=en GB


http://www.cambridge.org/gb/knowledge/isbn/item1158689/?site_locale=en_GB

Atmospheric Chemistry and Global Change, G. P. Brasseur, J. J. Orlando, G. S. Tyndall, Oxford University Press, 1999 http://ukcatalogue.oup.com/product/9780195105216.do#.UY5-2sqG4tY

Aeronomy of the Middle Atmosphere, 3rd Edition, Guy P. Brasseur & Susan Solomon, Springer, 2005 http://www.springer.com/environment/pollution+and+remediation/book/978-1-4020-3284-4

Climate Change 2007: The Physical Science Basis, Contrib. of Working Group I to the Fourth Assessment Report of the Intergovernmental Panel on Climate Change, 2007 http://www.ipcc.ch/publications_and_data/ar4/wg1/en/contents.html

References: Aviation & the Atmosphere European Scientific Assessment of the Atmospheric

Effects of Aircraft Emissions, Brasseur et al., Atmospheric Environment, 32 (13), 2329-2418, 1998 http://elib.dlr.de/10047/1/at-env-2329-1998.pdf

A Review of NASA's 'Atmospheric Effects of Stratospheric Aircraft' Project, US National Academy of Sciences, 1999 http://books.nap.edu/catalog.php?record_id=9604

Atmospheric Effects of Aviation: A Review of NASA's


pSubsonic Assessment Project, US National Academy of Sciences, 1999 http://www.nap.edu/catalog.php?record_id=6409

Aviation and the Global Atmosphere, IPCC, 1999 Full report: http://www.ipcc.ch/ipccreports/sres/aviation/index.php?idp=0

Summary for Policymakers: http://www.ipcc.ch/ipccreports/sres/aviation/index.php?idp=3

Transport Impacts on Atmosphere and Climate: The ATTICA Assessment Report, Atmospheric Environment, edited by R. Sausen, 44 (37), 4645-4816, 2010 http://www.sciencedirect.com/science/journal/13522310/44/37

http://www.pa.op.dlr.de/attica/


Part 1


Introduction

Structure of the Atmosphere

Atmospheric Radiation

Atmospheric Chemistry

Clouds and Aerosols

General Circulation of the Atmosphere

2

Aviation Emissions: Some Context

Concerns about aircraft emissions first focused on their contribution to local air quality Led to the introduction of the Clean Air Act in the USA and then to

International Standards for fuel venting and emissions of CO, NOx, unburned fuel, and smoke below 3000 ft (ICAO, 1981)

In the 1960s, the first civil supersonic aircraft were developed NO i i th t t th l


NOx emissions were seen as a threat to the ozone layer

Led to extensive research in the 1970s to try to quantify the impact

A proposed second generation civil supersonic aircraft (high speed civil transport – HSCT) again stimulated research into and assessments of the atmospheric effects of such a plane

The emphasis has now shifted to addressing the role and contribution of aviation emissions to the global issues of ozone loss and potential climate change

Brasseur et al., Atmos. Environment, 1998

Impacts of Aviation on the Atmosphere

Aircraft emissions affect ozone, aerosols and clouds, air quality, and


climate

IPCC 1999

Impacts of Aviation on Climate


https://www.atmosfair.de/en/air-travel-and-climate/ (based on IPCC 2007)

Introduction

A complete understanding of the atmosphere requires knowledge of:

Its evolution

Its physical structure

The processes determining its mass and composition

The distribution of density and composition


Motions within the atmosphere

Reasons for studying Earth’s atmosphere:

To understand its fundamental physics and chemistry

To monitor and understand the causes of change, both natural and anthropogenic, over short and long time scales

To use this understanding to predict the future state of the atmosphere

Aristotle’s Meteorologica, ~350 BC “The first difficulty is raised by what is called the air. What are we

to take its nature to be in the world surrounding the earth? And what is its position relatively to the other physical elements.”

“It [meteorology] studies also all the affections we may call common to air and water, and the kinds and parts of the earth and the affections of its parts. These throw light on the causes of winds and earthquakes and all the consequences the motions of


these kinds and parts involve.”

The hydrologic cycle: “Now the sun, moving as it does, sets up processes of change and becoming and decay, and by its agency the finest and sweetest water is every day carried up and is dissolved into vapour and rises to the upper region, where it is condensed again by the cold and so returns to the earth.”

“So summer and winter are due to the sun’s motion to and from the solstices, and water ascends and falls again for the same reason.”

http://ebooks.adelaide.edu.au/a/aristotle/meteorology/eBooks@Adelaide

Drivers for Atmospheric Science

Weather forecasting

Urban air quality

Acid rain

The ozone hole

Climate change

Improvement in the skill of weather forecasting (flow patterns at 5 km)

from 1981 to 2003

Atmospheric Physics and Sustainable Aviation (K. Strong) Slide 12Wallace and Hobbs, Figure 1.1

3

Atmospheric Science TodayCurrent Challenges

Tropospheric air quality globalization of

pollution

Stratospheric d l ti


ozone depletion recovery of the

ozone layer

Climate change global warming

and its impacts

Chemistry-climate interactions The Changing Atmosphere, ESA SP-1282, 2004 (after Isaksen, 2003)

How important is the atmosphere?


Part 1


Introduction




Clouds and Aerosols


Mean Composition of the Atmosphere

99.96% of the

atmosphere

Atmospheric Physics and Sustainable Aviation (K. Strong) Slide 16Brasseur et al., Atmos. Environment, 1998

The abundance of most gases in the atmosphere is quite small and so they are called trace gases.

Vertical Structure of the Atmosphere

The layers of the atmosphere are defined by the variation of temperature with altitude:

Troposphere:


p ptemperature decreases with altitude

Stratosphere: temperature increases with altitude

Weak vertical motions in the stratosphere

Strong overturning motion (convection) in the troposphere

Wallace and Hobbs, Figure 1.9

Vertical Structure of the Atmosphere


4

Atmospheric Processes Gas processes

Emission Photochemistry Gas-to-particle conversion Cloud removal

Radiative transfer UV/visible/near-IR/thermal-IR Scattering/absorption

Gas Aerosol Hydrometeor Snow, ice, water albedos

Aerosol processes Emission Nucleation/condensation Aerosol cloud coagulation

Meteorological processes Velocity Geopotential Pressure Water vapor Temperature Density


Jacobson, Fundamentals of Atmospheric Modeling 2nd Editionhttp://www.stanford.edu/group/efmh/FAMbook2dEd/index.html

Cloud processes Activation on aerosol Conden./evap./deposition/sublim. Hom./het./contact/evap. freezing Cloud, aerosol coagulation Precipitation/lightning Dissolution/chemistry

Aerosol, cloud coagulation Dissolution/chemistry/

crystallization Dry deposition/sedimentation Rainout/washout

Temperature Density Turbulence

Surface processes Temperatures and water content of

Soil Water Snow Sea ice Vegetation Roads Roofs

Surface energy/moisture fluxes Ocean-atmosphere exchange Ocean dynamics, chemistry

In the atmosphere, pressure and density decrease with altitude.

Pressure = force exerted per unit area, in N/m2 = Pa (Pascal)

Pressure at Earth’s surface: 101,325 Pa = 1013.25 hPa (SI unit)(= 1013.25 mb = 1 bar = 1 atm = 760 Torr) [1 mb = 1 hPa]

Pressure at 30 km: ~1,000 Pa

Atmospheric Pressure and Density


Density = amount of a material per unit volume

where m = mass in volume V

Mass density

= m / V(kg/m3)

Number densityno = N / V

(molecules/m3)where N = number of molecules in volume V

= M nowhere M is the molar mass (g/mole): M = m / n (n in moles)

Pressure and density are related quantities:

Near the surface, where the atmospheric pressure is largest the molecules will be “squeezed” most tightly together.

Higher up, where pressure is lower, the molecules are spread farther apart.

The Ideal Gas Law

T = atmospheric temperature (K)The Ideal Gas Law:


In terms of number density:

p = (n/V) R* T= (n NA/ V) (R*/NA) T= no k T

T = atmospheric temperaturek = Boltzmann’s constant (1.38 x 10-23 J/K)no = number of molecules per unit volumeNote: k = R* / Avogadro’s number NA

p V = mRT = nR*T p = (n/V) R* T

= (m/V) R T = R T

m = mass, n = number of molesR = gas constant (287.0 J/K/kg for dry air)R* = universal gas constant (8.3145 J/K/mole)= mass density (kg/unit volume)Note: R = R*/M where M = molar mass (“molecular weight”)

Mixing Ratios for Dry Air

Consider a small sample of the atmosphere, with volume V, pressure p, temperature T. Several gases, Gi, are present, each with Ni molecules and mass mi in the sample, mass density i and number density (no) i

Total number of molecules: Nd = Ni

Total number density: (no)d = (no)i

T t l N


Total mass: md = Ni mi

Total mass density: d = i

Volume mixing ratio (ppmv, ppbv, pptv):

Mass mixing ratio (kg/kg):

Mixing ratios are independent of temperature and pressure.

d

i

do

io

d

ii p

pn

n

NN

VMR

id

i

d

i

d

ii VMR

MM

mm

MMR

Units are Important


http://www.ldeo.columbia.edu/dees/ees/climate/slides/oz_prof.gif

S. Carn, Michigan Tech

Hydrostatic Balance

Atmospheric pressure at any altitude is due to the force per unit area exerted by the weight of all the air above that altitude.

Thus, atmospheric pressure decreases with altitude, resulting in a net upward force.

The atmosphere of a planet is subject to its gravitational field.

S th t d f i


So the net upward force is opposed by the downward gravitational force.

If these two forces are equal, the atmosphere is in hydrostatic balance.

meted.ucar.edu

5

The Hydrostatic Equation



http://www.physlink.com/fun/RickLondon.cfm

The Hydrostatic Equation

The forces acting upon on element of horizontal size A and vertical thickness z are [use Taylor expansionp(z+z) = p(z) + (p/z)z ]:

A zzp

A )z(pA )zz(pAp

p(z + z)

dA z + z

z

A


This net upward force is balanced by the downward gravitational force:

So we have:

And hence the Hydrostatic Equation:

(negative sign because p decreases as z increases; p<0, z>0 as z)

Force on upper face

Force onlower face

AzgV g

A z g A zzp

zgp

p(z)

Barometric Law and Scale Height We can combine the Hydrostatic Equation with the Ideal Gas Law

to get:

This is another form of the Hydrostatic Equation:

This equation is not, in general, integrable for an atmosphere because the temperature varies with the height z

dz g RTp

dz g dp T R p

pRTg

dzdp


because the temperature varies with the height, z.

If we assume an isothermal atmosphere (T = constant), then we can integrate to get:

where H = R T / g is the scale height it is a measure of how fast the pressure (or density) decreases with height (z for p by 1/e). for Earth’s near the surface: H ~ 8 km, for Venus: H ~ 12 km

Even when the atmosphere is not isothermal, “local scale height” is often used and is just defined at height z as H = R T(z) / g. At T = 210 K, H 6 km. At T = 290 K, H 8.5 km.

H

zz

RT

g

e )0(pe )0(p)z(p

Barometric Law

Vertical Profile of Atmospheric PressureFor an isothermal atmosphere (T = constant):

[ p = RT ]

where

( 0) 1013 25 hP

Hz

o

Hz

o

e )z(

e p)z(p


po = p(z=0) = 1013.25 hPa

(surface pressure)

o = (z=0) = 1.235 kg m-3

(global mean surface density)

Example: T = 270 K, H ~ 8 km.

At 8 km (height of Mt. Everest),

p = po e-1 ~ 370 hPaFrom Marshall and Plumb

First Law of Thermodynamics

We need the First Law of Thermodynamics to treat the movement of an “air parcel” in the atmosphere temperature structure.

Deals with the way in which a system can gain energy from or lose energy to its surroundings – it is just the conservation of energy principle applied to heat, work, and internal energy.


The energy of a system changes from an initial value Ui to a final value Uf due to heat Q and work W:

U = Uf – Ui = Q + W or for a small change: dU = dQ + dW when Q > 0, the system gains heat

when Q < 0, the system loses heat

when W > 0, work is done ON the system

when W < 0, work is done BY the system

Note: Some books usedu = dq – dwassuming unit mass and work done by the system

Specific Heat

If a small quantity of heat, dq, is given to a unit mass, resulting in the temperature increasing from T to T + dT, the specific heat is the ratio dq / dT.

Specific heat at constant volume:

dTdu

dTdu

dTdq

c VconstantVconstant

v

For mass m:dU / dT = m cv


using dq = du when V = constant, and Joule`s Law (internal energy depends only on temperature).

Specific heat at constant pressure:

cp cv because more heat is needed to do work as the system expands.

Can show that: cp = cv + R

p constantp dT

dqc

cv = 717 J/kg/K for dry aircp = 1004 J/kg/K for dry airR = 287 J/kg/K for dry air

6

Adiabatic Motion

Now, consider an air parcel that undergoes a change in its physical state (p, T, or V) is moved without exchanging heat with its surroundings (dQ = dq = 0), i.e., moved adiabatically

Such a parcel is assumed to be: Infinitesimally small

Thermally insulated from the environment

St i t th di t h


Staying at same pressure as the surrounding atmosphere

Moving slowly so that its kinetic energy is negligible

Is this realistic? If the atmosphere is vertically unstable, then it overturns quickly.

A very fast heat transfer mechanism would be needed to exchange a large amount of heat in or out of a parcel during its rising motion.

So on the timescale of convective overturning, we can treat the motion of an air parcel as adiabatic.

Dry Adiabatic Lapse Rate

We can derive the rate of change of temperature with altitude (the lapse rate, = – dT/dz) of a parcel of dry air using the Ideal Gas Law, the Hydrostatic Equation, and the First Law of Thermodynamics.

We will show that this is the line of stability between a


yconvective regime with much vertical mixing and a non-convective regime with convection suppressed. We will then look at how an atmosphere behaves convectively when a vertical temperature gradient is present.

The Ideal Gas Law again:

p V = m R Tor

p = R T( m = V / )


For adiabatic motion, the First Law of Thermodynamics becomes:dU = dQ + dW = 0 + dW dU – dW = 0

dU + p dV = 0 (using dW = F dx = – p dV)

Let’s look at each term in this equation:

dU + p dV = 0

dU = m cv dTF th Id l G L V R T

Note: Some books usedW = p dV assumingwork done by the system


Substitute these in to get: m cv dT + m R dT – (m/dp = 0

(cv + R) dT – dp/ = 0

Now, we can also use: cv + R = cp

So the change of temperature with pressure of

an air parcel moving adiabatically is:

v

for any air parcel From the Ideal Gas Law: p V = m R Tp dV + V dp = m R dT p dV = m R dT – (m/dp

pc1

dpdT


For a hydrostatic atmosphere, we can use:to get the change in temperature with altitude:

pc 1

)dz g (dT

dpdT

dzgdp

gdT


This gives us the dry adiabatic lapse rate:

What does this mean?

An air parcel moving in a hydrostatic atmosphere exhibits a fixed rate of change of temperature and density with altitude.

The air parcel will have the pressure of the surrounding atmosphere but will have its own temperature and density.

Near the Earth’s surface: d ~ 10 K/km (9.8 K/km)

pd c

gdzdT

Dry Adiabatic Lapse Ratep

d cg

dzdT


http://www.jonschrage.com/2xats113/2xE2L5/dalr.jpg

Convective Instability

Let’s consider three possible atmospheric temperature profiles.

(1) The atmospheric temperature profile changes more quicklywith altitude than the adiabatic profile. The air parcel starts at the surface and as it rises along the adiabat,

it becomes warmer than its surroundings because its temperature falls less rapidly than that of the surrounding atmosphere.


p y g p

Therefore it becomes lessdense and thus buoyant.

So it continues to rise at anaccelerating rate until it meets the atmospheric profile at some higher altitude.

This atmosphere is convectively unstable and will mix vertically because vertical motions are amplified.

Az

z + dz

TA (adiabatic profile)

TAtm (atmospheric profile)

> d

Indicates an unstable atmosphere(negative static stability)

7

Convective Stability

(2) The atmospheric temperature profile changes less quicklywith altitude than the adiabatic profile. The air parcel starts at the surface and as it rises along the adiabat,

it becomes cooler than its surroundings because its temperature falls more rapidly than that of the surrounding atmosphere.

Therefore it becomes moredense and tends to sink.

TAtm (atmospheric profile)


dense and tends to sink.

This atmosphere is convectively stable.

(3) The third possibility is thatthe atmosphere follows theadiabatic lapse rate. This atmosphere

is marginally stable.

Az

z + dz

TA (adiabatic profile)

< d

Indicates an stable atmosphere(positive static stability)

Atmospheric Stability - Dry Convection

Atmospheric Physics and Sustainable Aviation (K. Strong) Slide 38From Marshall and Plumb

Convection and Out-going Longwave Radiation (OLR)


Convection and Out-going Longwave Radiation (OLR)

OLR

total flux density (W/m2) in the infrared (emission from Earth and its atmosphere).


measured by downward-(nadir-) viewing satellites.

provides a measure of the temperature of the emitting region

From Marshall and Plumb

Out-going Longwave Radiation

OLR is high over deserts and cooler regions of the ocean as these are dry and cloud-free

OLR is low over polar regions as these are cold

OLR is also low over three wet equatorial regions, where radiation is emitted from the cold tops of deep convective clouds at high altitudes (10-15 km), not the warm surface


Where does convection occur? Almost everywhere.

Convection requires a warm surface relative to the air above.

This can happen by: Warming of the surface by sun

Cooling of the air above, e.g., winds bringing in cold air

At mid-latitudes: Shallow convection is frequent cumulus clouds

Deep convection is intermittent heavy rain, thunderstorms

Part 1


Introduction




Clouds and Aerosols


8

Radiation

The primary source of energy that drives the Earth’s climate system is the Sun.

The Sun’s energy comes to us mostly in the form of electromagnetic radiation.

Understanding climate requires understanding the nature of electromagnetic radiation and how it interacts with a


= 284Å = 304Å = 171Å = 195Å

Sun – SOHO observations 07/01/03, from http://sohowww.nascom.nasa.gov/

gplanetary atmosphere and surface.

Electromagnetic waveElectromagnetic waveMagnetic fieldElectric field

Defined by the solutions of Maxwell’s equations.

As they move, photons carry electric and magnetic fields that oscillate at a certain frequency, giving rise to an

The term electromagnetic radiation refers to a phenomenon that moves energy from one place to another, and carries with it an electric and a magnetic field.

Electromagnetic Radiation


EM waves are usually specified by:

E H direction of propagation

Wavelength () = distance between crests

Frequency (f or ) = number of oscillations per second = c/ Wavenumber ( or ) = number of crests per unit length = 1/

q y, g gelectromagnetic wave.

Can be thought of as particles (photons) that carry energy at speed

c (c = 2.998 x 108 m/s in vacuum).

Radiation as Wave Motion

The energy per unit area per unit time flowing perpendicularly into a surface is given by the Poynting vector :

units of W/m2

where

= vacuum permittivity

= electric field

HEcS o2

oE

E


= magnetic field

The energy transport is proportional to E2 becausethe electric and magnetic fieldsare related by a constant ofproportionality dependent onmedium in which propagation occurs.

H

H

S

direction of propagation

We measure the time average of the Poynting vector.

Radiometric Quantities

We will now define the fundamental quantities used to describe radiation and radiative transfer:

radiant flux density To do this we will also need

radiant exitance to define solid angle.

emittance

radiance = intensity


An aside – see Bohren and Clothiaux, Fundamentals of Atmospheric Radiation, page 204!

RADIOMETRY = acronym for Revulsive, Archaic, Diabolical, Invidious, Odious, Mystifying, Exotic Terminology Regenerating Yawns!

Solid angle

the area of the projection onto a unit sphere of an object, where lines are drawn from the centre of the sphere to every point on the surface of the object

equal to the ratio of the area A of a spherical surface intercepted to the square of the radius, r.

Solid Angle

z OBJECT (AREA A)

= A/r2 steradian (sr)


x

y

z

r = 1

UNIT SPHERE

OBJECT (AREA A)

r

PROJECTION ONTO UNIT SPHERE = SOLID ANGLE

AREA = 4 SPHERE OF RADIUS rAREA = 4 r

r

To obtain a differential element of solid angle, construct a sphere of radius r whose central point is denoted by O.

Differential area on the surface located at r from point O is:

Then the differential solid angle is:

Polar coordinates

Solid Angle

A

)d sinr)(dr(dA


where = cos = zenith angle

= 90 – elevation angle

= azimuth angle

[d(cos)=-sin; dropped - sign]

d dd d sinrA

dd 2

9

Radiant Flux Density

The electric and magnetic fields and the Poynting vector of an EM wave oscillate rapidly, so are difficult to measure instantaneously.

Usually measure the average magnitude over some time interval:

This is the radiant flux density (W m-2).

)S,H,E(

SF


The radiant flux density is redefined based on the direction of energy travel: radiant exitance (M) = radiant flux density emerging from an area

irradiance (E) = radiant flux density incident on an area

RADIANT EXITANCE, M IRRADIANCE, E

Intensity or radiance (I or L) = radiant flux density per unit solid angle (W m-2 sr-1).

It is the radiation flow in a particular direction at a particular point.

Strictly, the intensity represents the EM radiation leaving or incident upon an area perpendicular to the beam. For other directions, it must be weighted by cos .

Th di ti b l th d d t fi th

Intensity (or Radiance)

d

SI


The radiation may be wavelength dependent, so we prefix the energy-dependent terms by "monochromatic" or "spectral".

Symbols are subscripted accordingly, e.g., I, I, I refers to intensity per unit , , : I = I/d, , I=I/d, I /d Units: I in W m-2 sr-1 nm-1, I in W m-2 sr-1 Hz-1, I in W m-2 sr-1 (cm-1)-1

If the intensity is a function only of the direction but not position, then the field is homogeneous.

If the intensity is independent of all spatial and directional parameters, then the field is homogeneous and isotropic.

The radiant flux density, F, is the total or net energy flow in a particular direction, with

The monochromatic flux density (or the monochromatic irradiance) is defined by the normal component of integrated over the entire hemispheric solid angle:

Radiant Flux Density - Revisited

dt d dAFdE

d cosIF

etc. d/FF


In polar coordinates:

Note, for isotropic radiation(independent of direction):

Total flux density (or irradiance):

The total flux or radiant power (energy per time):

ddIF sincos),(2

0

2

0

0

dFF

AdAFf

I d d sincosIF2

0

2/

0

Independent of the type of material

Isotropic in nature (i e the same

Blackbodies

Blackbody energy distribution

A blackbody is a perfect emitter - it emits the maximum possible amount of radiation at each wavelength.

A blackbody is also a perfect absorber, absorbing at all wavelengths of radiation incident on it. Therefore, it looks black.


Isotropic in nature (i.e., the same in all directions)

Total energy is proportional to T4

Peak emission is given by

T2

Planck’s Blackbody Function

No real materials are perfect blackbodies. However, the radiation inside a cavity (whose walls are opaque to all radiation) is the radiation that would be emitted by a hypothetical blackbody at the same temperature. The cavity walls emit, absorb, and reflect radiation until equilibrium is reached.

Planck postulated that atoms oscillating in


Planck postulated that atoms oscillating in the walls of the cavity have discrete energies given by: E = n h where n = integer (quantum number), h = Planck's constant, = frequency

A quantum of energy emitted when an atom changes its energy state is then

E = h (n = 1).



Using these two assumptions, Planck derived the blackbody function, describing the radiance emitted by a blackbody:where

B = monochromatic radiance (W m-2 sr-1 m-1)

k = Boltzmann's constant

T = absolute temperature

1kThc

exp

hc2)T(B

52

c)T(B

51


T = absolute temperature

This can be written as:where

c1 = first radiation constant (1.191 10-16 W m2 sr-1)

c2 = second radiation constant (1.439 10-2 m K)

The radiance emitted by a blackbody depends only on and T.

B(T) increases with temperature

the of maximum B(T) decreases with temperature

1T

cexp

c)T(B

2

1

10


We can show that the equivalent form of Planck’s function as function of wavenumber is:

1T

cexp

c

1kThc

exp

hc2)T(B

2

31

32


Peak wavenumber for blackbody emission

dBdBdB :Note

Wien’s Displacement Law

The hotter the object, the shorter the of its maximum intensity (e.g., element on a stove).

This law can be used to determine the T of a blackbody from the position of the maximum monochromatic radiance


maximum monochromatic radiance.

The wavelength at which B(T) is a maximum is determined using:

This gives: m (T in K)

Or for wavenumbers:

Wallace and Hobbs, Figure 4.60)T(B

T9.2897

m

1m cm T2T962.1

Stefan-Boltzmann Law

The monochromatic radiant exitance is simply: M(T) = B(T)because the blackbody radiance is isotropic (independent of direction).

We can use this to determine the total radiant exitance from a blackbody:

415

Tc

d)T(Bd)T(M)T(M


or

where = Stefan-Boltzmann constant = 5.67 10-8 W m-2 K-4

If MBB (=F) is known, this can be used to calculate an equivalent blackbody temperature, or effective emission temperature, TE

420 0

Tc15

...d)T(Bd)T(M)T(M

4BB T)T(M

4T)T(B

The Blackbody Spectrum

Flash animation:

http://www.colorado.edu/physics/phet/simulations/blackbody/blackbody.swf

PhET Interactive Simulations, University of Colorado, http://phet.colorado.edu


Grey Bodies The emissivity of a material is used to quantify how closely it

approximates a blackbody:

For a blackbody, =1.

For other materials, called grey bodies, 0 < 1.

)T(BI

at radianceblackbody at radiance emitted

Wavelength-dependent


g y

Can also define: absorptivity absorbed radiance at / incident radiance at reflectivity R, reflected radiance at / incident radiance at transmittivity transmitted radiance at / incident radiance at

These three quantities describe the three possibilities for incident radiation. All have values between 0 and 1.

By the conservation of energy: + R + = 1.

Kirchoff's radiation Law: =

Any planetary surface looks reasonably “black” when considering thermal (i.e., infrared) radiation emitted from it.

Infrared Emittance

Some examples:


So we will not be too much in error if we treat planetary surfaces as blackbody surfaces.

In contrast, many polished surfaces have ~ 0 in the infrared (e.g., gold has = 0.01 despite its “brightness” in the visible).

11

particles: 0.0065-0.002 Wm-2

(mainly protons) and magnetic fields

SUN EARTH

solar wind

galactic cosmic rays

NRL LASCO coronagraph on SOHO

0.0000007 Wm-2

Atmospheric Physics and Sustainable Aviation (K. Strong) Slide 61PHY492F/PHY1498FLecture 4 Page 61

surface

photons: 1367 Wm-2

atmosphere

magnetosphere

sunspot

faculae

bow shock

plasmasphere

heliosphere

solar eruptions: flares,coronal mass ejection

Slide courtesy of Judith Lean, Naval Research Laboratory

Surface, troposphere

1.31367Total irradiance

Solar radiation

Terrestrial deposition

altitude

(km)

Solar Cycle change

(W m-2)

Energy

(W m-2)

Source

Properties of Solar Radiation


About 5000.0003Solar wind

0-900.000007Galactic cosmic ray

30-90 0.002Solar protons

Particles

50-5000.020.1UV 0 – 200 nm

0-500.1615.4UV 200-300 nm

p p

The solar radiation incident on the Earth’s atmosphere has the properties of blackbody radiation.

In the UV, visible, and IR regions, seen at low-medium resolution, the solar flux density is almost exactly like that from a blackbody aperture the size and shape of the Sun’s visible disk.

Temperature of this blackbody = 5780 K.

Can be used to calculate the spectral distribution of the radiation.

The Sun - A Hot Blackbody


Solar disc subtends an arc of 31.99 minutes at Earth.

This is very small, so the solar radiation is almost exactly a plane wave at the Earth.

Solar flux density can be described by:where is the solid angle subtended by the Sun (6.8 10-5 sr).

Intensity can be found using:

Thus, the solar flux density at Earth’s orbit is approx. 1367 W m-2.

This is the solar “constant”.

BM or IF BBss

1274 sr m W 102T)/(I

Solar and Terrestrial Radiation

Solar flux density or irradiance peaks at visible , near

0.48 m or 11,500 cm-1

falls off rapidly at IR known as shortwave

radiation


radiation

Earth’s irradiance peaks at IR , near 10 m

or 550 cm-1

emits no visible radiance

known as longwaveradiation

Absorption of Radiation by the Atmosphere

B

Blackbody spectra


Abs

orpt

ion

(%)

Absorption at the surface

The surface and atmosphere of a

planet will reflect and scatter some radiation

can see the Earth from “outside”

Solar radiation incident on a

planet

Absorption

Reflection

Miguel Olivares, New York City, Jan 5, 2003.

Absorption of Solar Radiation


can see the Earth from outside

The fraction of solar radiation reflected back into space from the

planet is called the planetary albedo.

General definition:(no units)

0 A 1, depending upon the surface or planet.

The fraction of solar radiation absorbed by a planet is (1 – A).

irradiancereflectiontodue exitance radiant

EM

A

12

Radiation Balance of Earth

Over the long-term must be an equilibrium between the solar radiation absorbed by a planet and the thermal radiation it emits.




We can write this balance as:

4einfrared

2Esun

2E

4einfrared

2Evisiblesunsun

2E

T R 4)A1( F R

T R 4 I R

Energy from the Sun Energy emitted


where we’ve used: Fsun = Isun sun and (1-A) = visible

and A is the albedo or the average amount of solar energy reflected.

intercepted by the spherical planet.

by the planet in the infrared.

4einfraredsun T 4)A1( F


Let’s plug in some numbers

K255 Wm13683.01FA1F

T

Wm4.2394 Wm1368

13.01

4FA1

TF

4

2-

4 s4 Earth

2-2-

s

infrared

4eEarth


K255 414

T 44 s

infrared

4 Earthe

The Greenhouse Effect

CO2, H2O, CH4, O3 naturally radiatively active gases

The fundamental reason for the existence of the greenhouse effect is that the temperature decreases with altitude in the troposphere.

Radiatively active gases, as well as clouds, absorb the radiation emitted by the warmer surface, while their emission of radiation to space occurs at cooler atmospheric temperatures


space occurs at cooler atmospheric temperatures.

The trapping of the radiation by radiatively active molecules produces an increase in the surface temperature of about 33C (assuming no change in albedo when atmosphere is removed…).

Recall:

This gives Te ~255 K (-18°C) (with Fsun=1367 W m-2, A=0.3, =1)

Using TS ~288 K (15°C), the atmosphere traps ~33°C.

4einfrared T 4)A1( F

A Simple Greenhouse Model - 1

Let’s add the atmosphere to our simple radiative equilibrium model – a Half-Dimensional Climate Model

Question: How?

Treat the atmosphere as a simple thin shell covering the planet.

“Unroll” the sphere to obtain a planar diagram.


Solar input to this plane: the average (over all orbital and global positions) input on a unit area of the planar surface:

(factor of 4 = ratio of disk area to surface area of a sphere).

4/So

Surface

AtmosphereASo/4

So/4

FS

FA

FA

Solar Thermal

Incident Solar Radiation

Total output from the sun: Q = 3.87 1026 W

Solar constant at Earth: So = Q / 4R2earth-sun 1367 W/m2

where Rearth-sun = 1.5 1011 m

A i i l fl d it / it t E th


Average incoming solar flux density / unit area at Earth:

= energy intercepted per unit time / surface area

= ( r2So ) / (4r2) = So/4 = 342 W/m2

13

Hypothesis:

Atmosphere is a planar layer of material that is

transparent to solar radiation

opaque to thermal radiation

Definitions:

So/4 is the average solar input (W/m2)


FA

Solar Thermal


o g p ( )

A is the planetary albedo provided entirely by ground reflection.

FS is the infrared flux density emitted by the surface: FS = TS

4

FA is the infrared flux density emitted by the atmosphere (as would beseen by a distant radiometer): FA = TA

4 Te4 (so TA Te)

= 1 (everything in the thermal region is considered to be black).

Surface

AtmosphereASo/4

So/4

FS

FA

Writing down the equation for vertical energy transfer abovethe atmosphere and stating that the atmosphere/planet system is in radiative equilibrium:

Aoo F4/SA4/S Planetary albedo

IR flux density emitted by the atmosphere


Average solar flux density


Below the atmosphere:

The surface temperature is therefore: TS = 21/4 Te =1.19 Te

SoAAo

SoAo

F4/SAF)F4/SA(

F4/SAF4/S

4e

4S

AS

T2T

F2F

Average solar flux density

Some Consequences

For Earth: Te = 255 K TS = 303 Ki.e., the surface temperature is higher than the effective radiating temperature of Earth

This is entirely due to the absorption of infrared terrestrial radiation by the atmosphere, which in turn re-radiates this back down to the surface, thus increasing the net downward radiative flux density at the surface


the surface

The flux density emitted by the atmosphere (FA) is very large, comparable to the solar flux density:

Actual mean TS is 288 K, so this simple model is an overestimation

Not all solar radiation incident at TOA reaches the surface

Not all radiation emitted by the surface is absorbed by the atmosphere

4/S)A1(F oA

Beer-Bouguer-Lambert Law

For a plane-parallel beam travelling through an absorbing di ll di h b h i i li d f ll

ds (changed from dx to match textbook)dA

I dII

d


medium, over small distances, the behaviour is linear and follows the Beer-Bouguer-Lambert Law (or Beer’s Law). See Perrin, Whose Absorption Law?, J. Optical Society of America,

38(1), 72-74, 1948.

Loss of the beam intensity dI is proportional to the amount of material traversed, ds, where is the density of the absorber:

ds k I-dI

extinction coefficient

Beer-Bouguer-Lambert Law

For the case of: monochromatic radiation

an absorber of uniform density

we can integrate Beer’s Law over a finite distance from 0 to s:

)0(Ie )0(Ie )0(Ids kexp)0(I)s(I s k

stransmission


This is the sort of thing we see in the laboratory as well as the atmosphere.

0

familiar (?) exponential decay law of transmission through uniform material

optical depth

Beer’s Law Applied to Atmosphere

Let’s integrate Beer’s Law from the top of the atmosphere (z=) down to any level z to determine what fraction of the incident radiation is absorbed or scattered, and what fraction transmitted.

We have ds = sec dz, so:

Integrate:z)z(I

dz sec k I -dI

Atmospheric paths are generally inhomogeneous varying in


where

= normal optical depth or optical thickness

= transmissivity

If no scattering, absorptivity is:

)(Ie )(Idz sec kexp)(I)z(I

dz sec k)(I)z(I

lnIdI

secz

z)z(I

)(I

θsec e1 1

inhomogeneous - varying in some parameter(s) such as temperature and pressure, so need to keep the integral.

14

k dx must be a “number” – meaning it has no units. There are several sets of units which can be used …

Most common units for the extinction coefficient, k: k in cm2 molecule-1 and in molecules cm-3 (“number density”)

or in SI units, with length dx in m: k in m2 kg-1 and in kg m-3

Cross-section (units of cm2 molecule-1) is also used to describe th t th f b ti i k h d it

Mass and Volume Extinction Coefficients


the strength of absorption, i.e., = k as we have used it. There is a lot of inconsistency in the use of and k! Be careful!

Whichever symbols are used, there are two absorption coefficients: The mass absorption coefficient or cross section has units of

m2 kg-1 or cm2 molecule-1

The volume absorption coefficient has units of (m-1 or cm-1) and equals the mass absorption coefficient density

Note that the product dx is the area density unit or column (e.g., molecules cm-2) – it represents the amount of material in the path

There are three atmospheric processes that can cause extinction of radiation.

Extinction of the beam:

- Absorption (ka)

- Simple scattering (ks)

Atmospheric Extinction Processes


- Resonant scattering (kr)

k = ka + ks + kr

total extinction coefficient, Note: each term has spectral dependence, not indicated here

Absorption

Spectroscopy describes the processes by which energy can be absorbed in planetary atmospheres.

These processes are related to the atomic and molecular properties of the gases in the atmosphere.

Rotation Types of transitions in the internal energy of the molecules in order of


Vibration

Electronic transitions

Photodissociation

Photoionization

energy of the molecules, in order of increasing energy

E0

E1

E2

h

absorption h emission

Interaction of the Radiation Field with Gases in the Atmosphere

Solar interactions Photoionization

– extreme UV strips electrons from atoms

Photodissociation – UV breaks apart molecules


Electronic (orbital)

Thermal IR interactions Electronic (orbital)

Vibration

Rotation

Born-Oppenheimer Approx:Energy of a gas moleculeE = Eelec + Evib + Erot + Etrans

Greenhouse Gases (GHGs)

forbidden

allowed

VIBRATIONAL MODES OF CO2

Greenhouse gases = gases with vib-rot absorption features at 5-50 m


allowed

allowed

• Major greenhouse gases: H2O, CO2, CH4, O3, N2O, CFCs,…

• Not greenhouse gases: N2, O2, Ar, …

Adapted from Daniel Jacob, Harvard

Simple scattering is the deflection of energy from its current direction into another direction (recall intensity has direction).

The energy is redirected rather than lost.

There is usually very little interaction with the kinetic energy reservoir and so little heating or cooling is associated with simple scattering.

- Need to pay attention to the angular

Simple Scattering


Consequences:

p y gproperties of the energy redistribution

- Energy is also scattered into the beam from other directions

Scattering of incident light wave by a particlePetty, Figure 7.1

15

Resonant scattering combines some of the properties of both absorption and single scattering.

Energy is first absorbed by the molecule, and then re-emitted some time later.

The re-emission may be at:

- same

- nearly the samedepending upon the quantum levels involved

Resonant Scattering


The re-radiation is usually isotropic.

Some degree of interaction between the process and the kinetic energy field.

The only atmospheric process by which radiation may be directly changed in frequency.

It is less important in the lower atmosphere.

y- very different

q

Ehf1

hf2

hf3

hf1 ~ hf2 + hf3

(f = frequency)

Scattering in Planetary Atmospheres

Scattering in the atmosphere generally occurs in three forms, by:

Molecules

Aerosols

Clouds

And there are


three regimesdetermined by the particle sizeand scatteringwavelength:

Rayleigh

Mie

GeometricWallace and Hobbs, Figure 4.11

Scattering by Larger Particles As particles becomes larger, the assumptions of Rayleigh

scattering become less justifiable Lorentz-Mie scattering (Lorentz 1890 and Mie 1908)

this requires solving Maxwell’s equations in all their glory!

Mie applied Maxwell's EM equations to the case of a plane EM wave incident on a sphere, and showed that the scattered


wave incident on a sphere, and showed that the scattered radiation for a sphere depends only on: viewing angle

complex index of refraction m = mr + i mi

size parameter 2r / , where r = radius of the sphere

Rayleigh scattering: < 0.1

Mie scattering: 0.1 < < 50

Geometric (optics) scattering: > 50

Atmospheric Effects of Scattering

Weak scattering just redirects the incoming solar beam slightly, allowing it to still reach the surface.

Stronger scattering increases the atmospheric pathlength, which increases the atmospheric absorption

e.g., Mars - dust storms fed by increased solar energy absorption

increases the back-scattering to space which therefore increases the atmospheric albedo


p e.g., Venus - clouds reflect (back-scatter) about 75% of solar radiation

Light that arrives at the top of the atmosphere from the Sun is unpolarized – this means that just as many photons have their planes of polarization in one direction as in any other direction. The degree of polarization of the light reaching the surface of the

Earth – or reaching some satellite instrument in space – is due to Rayleigh scattering in the atmosphere.

You can observe this by looking at the sky at 90° to the solar beam – rotate a polarizer and see how the transmitted light changes.

Why Is the Sky Blue?

We have seen that the intensity of Rayleigh-scattered light depends on the wavelength of the incident light and on the index of refraction of air molecules.

The intensity scattered by air molecules in a specific direction is proportional to 1/4.

A large portion of the solar energy is lies between the blue and red regions of the visible spectrum


The sky appears blue, when viewed away from the Sun’s disk.

red regions of the visible spectrum.

blue (~0.425 m) < red (~0.650 m)

Blue light scatters about 5.5 times more intensity than red light.

Why Does Twilight Look Red?

As the Sun approaches the horizon (at sunset or sunrise), the sunlight travels through a longer atmospheric pathlength, thus encountering more air molecules.

Therefore more and more blue light – and light of shorter wavelengths – is scattered out of the beam of light.

The Sun shows a deeper red color than at its zenith.


p

Since violet light ( ~ 0.405 m) has a shorter wavelength than blue, why doesn’t the sky appear violet?

The Sun is a deeper red color than at its zenith.

16

Why Does Twilight Look Red?

As the Sun approaches the horizon (at sunset or sunrise), the sunlight travels through a longer atmospheric pathlength, thus encountering more air molecules.

Therefore more and more blue light – and light of shorter wavelengths – is scattered out of the beam of light.

The Sun shows a deeper red color than at its zenith.


p

Since violet light ( ~ 0.405 m) has a shorter wavelength than blue, why doesn’t the sky appear violet?

Some of this violet light is absorbed by the upper atmosphere

There is more blue than violet in sunlight Our eyes are slightly more sensitive to blue than to violet

Atmospheric Source Processes

So far, we have discussed atmospheric processes that extinguish radiation

There are also three atmospheric processes that can cause augmentation of radiation i.e., they add energy to the beam.

Emission


The two scattering processes involve an increase in the intensity due to scattering into the direction of the beam from all other directions

Augmentation of the beam:

- Emission

- Simple scattering

- Resonant scattering

Emission

Emission is the process by which energy is transferred from the kinetic energy reservoir to the radiation field.

For a blackbody, there is no net change in kinetic energy, and so the absorption and emission must balance (Kirchoff’s Law). Therefore the absorption and emission coefficients must be equal.


The emitted intensity is thus:

Monochromatic blackbody function for T of gas

dx k )T(BdI ,a,emitted

Absorption coefficient = emission coefficient by Kirchoff`s Law

The Atmospheric Source Function

We can collect all of the source terms and include them in the equation for the change of intensity across some path dx:

where j refers to the summation of the three source terms

dx k Jdx k I -

dx jdxkI-dI

Source termsSink terms


j refers to the summation of the three source terms

J = j/k has units of intensity and makes the equation symmetric

This can be rearranged to give:

This is a fundamental equation for radiative transfer in planetary atmospheres - contains practically all the physics of the problem. Note: it applies to monochromatic intensity.

J I -dx k

dI

Schwarzchild’s Equation

Remote Sensing* / Sounding**

“Measurement at a distance”

Information is carried by electromagnetic radiation

Provides a method of obtaining information about the properties of the atmosphere without coming into physical contact with it. in contrast to extractive or in situ techniques

Advantagest b ti f th l b i b d


no perturbation of the sample being observed

sensitive to many gases and surfaces

can provide point, column or profile data

Disadvantages limited spatial resolution

interpretation of data can be difficult

*remote sensing generally applies to observations of the surface

**remote sounding generally applies to observations of the atmosphere

Remote sounding techniques can be classified by:

1) Radiation source passive – use natural radiation (solar, stellar, terrestrial)

active – use artificial sources of radiation (lasers, radar)

2) Type of interaction between radiation and the atmosphere absorption

Remote Sounding Approaches


p

emission

scattering

3) Spectral region ultraviolet

visible

infrared

microwave

17

Remote Sensing/Sounding Techniques

ACTIVE

IMAGING NON-IMAGING SOUNDING

RANGE POWER BACKSCATTER

PASSIVE

IMAGING NON-IMAGING SOUNDING

REFLECTEDSUNLIGHT

THERMALRADIATION

EXTINCTIONEMISSIONSCATTERING

SENSING / SOUNDINGSATELLITE REMOTE

TECHNIQUES


LASERPROFILER(VIS/IR)

RADARALTIMETER(m-WAVE)

SCATTER-OMETER(m-WAVE)

IMAGINGRADAR(m-WAVE)

LIDAR(VIS/IR)

SUNLIGHT RADIATION

AERIALPHOTO-GRAPHY

VIS/NEAR-IRSCANNER

THERMALINRAREDSCANNER

PASSIVEMICROWAVERADIOMETER

SCATTERING

RADIOMETER

BROAD-BAND(VIS/IR)

SELECTIVEFILTER(VIS/IR)

HETERO-DYNE(m-WAVE)

SPECTROMETER

FABRY-PEROT(VIS/IR)

FOURIERTRANSFORM(VIS/IR)

GRATING(UV/VIS/IR)

LASER

HETERO-DYNE(VIS/IR)

(SLAR)

SIDE-LOOKINGAIRBORNERADAR (SAR)

SYNTHETICAPERTURERADAR

RADAR(m-WAVE)

Satellite Remote Sounding of Temperature

Underlying principles:

Most of the radiation reaching a satellite in any spectral channel is emitted from near the level of unity optical depth for that channel.

Spectral channels with higher absorptivities are associated with higher altitudes of unity optical depth (and vice versa).


Applying Schwarzchild’s Equation

This applies to a nadir-viewing (downward-looking) or vertical-sounding satellite instrument measuring radiance in the [near] local vertical, in a number of discrete infrared channels.

1

)X(

X

0

'x kX k d B)X()0(I'dxe B ke)0(I)X(I

Surface emission Atmospheric emission


The contribution of the surface emission term is reduced if transmission is small, i.e., if the observation is NOT in an atmospheric window.

The atmospheric emission term includes information on the temperature through the blackbody function B(T), and on both the temperature and the concentration profile of the absorbing/emitting gas through the transmission and hence through the weighting function.

Applying Schwarzchild’s Equation

Channels on a real instrument each measure radiation from a range of altitudes, limiting the vertical resolution of the retrieved temperature profile.

For temperature retrievals, we want to observe the intensity for an absorbing gas which:(1) is well mixed in the atmosphere (has a constant mixing ratio), and (2) has well known absorption lines


and (2) has well known absorption lines.

Preferred regions (both gases are well mixed up to ~90 km):

(1) CO2 near 15 m

(2) O2 near 60 GHz (5 mm)

TOA Radiation Balance: Net Incoming Solar

Tropics: ~300 W/m2, with highest values over cloud-free regions over oceans (A~0.1) and lowest over deserts (A~0.2)

Annual mean


Polar regions: <100 W/m2, due to polar night and the summers offset by high solar zenith angles, clouds, and high albedo of ice and snow


TOA Radiation Balance: Outgoing Longwave

Equator-to-pole gradient is gentler

Regional variability in the tropics is greater

The equator-to-pole difference in surface air temperature produces a 2:1 difference in outgoing OLR

This is offset by the

Annual mean


This is offset by the higher altitude of cloud tops and the moist layer in the tropics

Low OLR over tropical continents is due to deep convective clouds with high, cold tops

Highest OLR is over deserts and cloud-free dry zones


18

TOA Radiation Balance: Net Radiation

Net downward radiation = incoming solar – outgoing longwave Positive = downward

Negative = upward

Excess of incoming solar at low latitudes

Excess of OLR athi h l tit d

Annual mean


high latitudes

OLR > incoming solar over hottestdesert regions


Global Energy Budget


Marshall and Plumb

Part 1


Introduction




Clouds and Aerosols


Residence Time

Under steady-state conditions, when the rate of production/injection of a constituent equals the rate of removal (by chemical and physical processes), we can define its residence time or lifetime:

= M / F

where M is the amount of constituent in the atmosphere (kg) and F is the rate of its removal from the atmosphere (kg/s).

Notes:


Notes:

Lifetimes of individual molecules may differ significantly from the lifetime of the constituent, e.g., due to local removal processes.

Lifetimes don’t represent how long it takes a constituent to react to an abrupt change in a source.

The concentration of a constituent with a short (long) lifetime will vary over short (long) spatial scales, so those with short lifetimes will have high (low) concentrations near (far from) sources, while those with long lifetimes will have more uniform concentrations.

Spatial & Temporal Scales of Variability



Trace Gas Sources Biogenic – production by biological processes, e.g., decay of

organic material, microbes, biomass burning, vegetation (emits volatile organic compounds – VOCs)

Solid Earth – volcanoes (many gases), rocks (He, Ar, Rn)

Oceans – reservoir for water-soluble gases (can be a source or sink), biological activity in the oceans (dimethyl sulphide – DMS, other sulphur-containing gases, methyl chloride – CH3Cl)


In situ formation – by chemical reactions in the atmosphere, often initiated by photolysis Homogeneous reactions – all reactants are the same phase

Heterogeneous reactions – reactants are in two or more phases

Anthropogenic – e.g., fossil-fuel burning, fertilization, evaporation of solvents, livestock, landfills, rice paddies

Note: Trace gases emitted by the biosphere, solid Earth, and oceans are usually in a reduced (low oxidation) state (e.g., hydrocarbons, ammonia, H2S) and are oxidized by in situ reactions in atmosphere.

19

Natural & Anthropogenic Sources (2000)


Wallace and Hobbs, Table 5.2

Transport of Trace Gases

Turbulent mixing in the planetary boundary layer (PBL or ABL) results in well-mixed gases in the lowest 1-2 km of the atmosphere during the day over land (a few 100 m during night over land)

Gases emitted from the surface will move into the free troposphere (if they are not returned to the surface or


consumed by in situ reactions in the PBL)

In the free troposphere, gases with long lifetimes will be transported by global circulation patterns Characteristics: restricted transport of tropospheric air across the

tropics, upward transport to the stratosphere occurs in the tropics, downward transport occurs at high latitudes

Gases will often remain within latitude bands

Northern Hemisphere: more affected by fossil fuel emissions

Southern Hemisphere: more affected by oceans and biomass burning

MOPITT Measurements of CO

First instrument to measure pollution from a satellite.

Launched Dec. 1999 on NASA’s


Animation created by American Museum of Natural History, October 2004. Three years of MOPITT data. http://www.acd.ucar.edu/mopitt/visualize.shtml

Terra satellite.

PI is Prof. Jim Drummond, U of Toronto /Dalhousie

Trace Gas Sinks

Transformation into other chemical species

Gas-to-particle conversion by chemical and physical processes

Wet deposition – the scavenging of gases and particles in the air by clouds and precipitation – this is major mechanism for cleaning the atmosphere

Dry deposition – collection of gases and particles by vegetation and other solid and liquid surfaces much slower than wet


and other solid and liquid surfaces – much slower than wet deposition but continuous

Deposition velocity of gas onto surface = (flux of gas to the surface) / (mean concentration of the gas just above the surface)

Oceans – flux depends on how unsaturated the ocean is for a gas Solubility is given by Henry’s Law: Cg = kH pg

where Cg = solubility of a gas in a liquid (mole/litre), kH = Henry’s Law constant, a temperature-dependent constant of proportionality (mole/litre/atm), and pg = partial pressure of gas (atm)

Tropospheric Trace Gases: OH

The hydroxyl radical (OH) is highly reactive – a key species

Produced by photolysis of O3, followed by reaction of O* with H2O

Primary sinks are oxidation of CO and CH4

Reacts quickly with most trace gases (an oxidant) “detergent of the

atmosphere”


atmosphere

Concentrations are low (<0.4 pptv)

Lifetime is short (~1 s)


Tropospheric Trace Gases: Nitrogen Species

NOx = NO (nitric oxide) + NO2 (nitrogen dioxide) Many sources, most emitted as NO, which establishes a daytime

equilibrium with NO2, main daytime sink is reaction with OH

Nitrate radical (NO3) Is the major oxidant at night when OH disappears (no photolysis)

Odd nitrogen = total reactive nitrogen (NOy) = NOx + all oxidation products of NO


products of NOx

Main sources are anthropogenic emissions, soils, lightning

Ammonia (NH3) Primary basic gas in the atmosphere, neutralizes acid species

Removed by conversion to ammonium(NH4)-containing aerosols followed by deposition

Nitrous oxide (N2O) Sources include bacteria in soils and oceans, fertilizers, livestock

Greenhouse gas, source of NOy in the stratosphere

20

The Nitrogen Cycle


http://www.waikato.ac.nz/wfass/subjects/geography/people/max/ConceptDiagrams/NitrogenCycle.jpg

Tropospheric Trace Gases: Organic Species

Organic compound contain carbon atoms

Hydrocarbons are organic compounds containing carbon and hydrogen Methane (CH4) is the most abundant, ~1.7 ppmv, ~9-year lifetime

Non-methane hydrocarbons (NMHCs) include alkanes (e.g., ethane, C2H6), alkenes (have a double bond, e.g., ethene,


, 2 6), ( , g , ,C2H4), and aromatics (e.g., benzene, C6H6)

Oxygenated hydrocarbons contain one or more oxygen atoms (e.g., acetone) and can be a source of hydrogen oxides

Tropospheric Trace Gases: CO

Carbon monoxide (CO) is produced by oxidation of CH4 and NHMCs, biomass burning, and fossil fuel combustion

The primary sink of CO is oxidation by OH – this is also the primary sink for OH, and so the distribution of OH is often determined by the distribution of CO

CO has a lifetime of ~2 months and so serves as a useful


tracer of atmospheric motions

It has a clear seasonal cycle in the extratropics: concentrations increase in winter when there is less OH and decrease rapidly in spring

Tropospheric Trace Gases: O3

Tropospheric ozone (O3) makes up about 10% of total ozone

Tropospheric ozone is created by photochemical reactions involving NOx, CO, and organic compounds

It is highly reactive and plays an important role in the oxidizing capacity of the troposphere


g p y p p

Mixing ratios have increased from pre-industrial values of 10-15 ppbv to 30-40 ppbv in 2000, due to increases in NOx

emissions from the use of fossil fuels

Tropospheric Trace Gases: Hydrogen Species

Hydrogen compounds are important oxidants

Odd hydrogen = HOx = H + OH + HO2 (x = 0, 1, 2)

H = atomic hydrogen – reacts quickly with O2 to form:

HO2 = hydroperoxyl radical

Other hydrogen compounds include H2, H2O2 (hydrogen peroxide) and H O


peroxide), and H2O

Rapid cycling occurs between OH and HO2

HOx and NOx can react to create ozone e.g., in upper troposphere where NOx is emitted by aircraft

Tropospheric Trace Gases: Sulphur Species

Sulphur is assimilated by living organisms and released as the end product of metabolism

Sulphur compounds can be in reduced and oxidized states

The more oxidized species have a high affinity for water and are more easily removed by wet processes

The most important are: Sulphur dioxide (SO )


Sulphur dioxide (SO2)

Hydrogen sulphide (H2S)

Dimethyl sulphide (CH3SCH3 or DMS)

Carbonyl sulphide (COS or OCS)

Carbon disulphide (CS2)

Primary sources of the reduced compounds are biogenic reactions in soils, plants, and oceans. They are oxidized to SO2 and SO4

2-.

Anthropogenic emissions are nearly all SO2, with 90% from NH.

Removal is by wet and dry deposition

21

The Sulphur Cycle


http://www.scienceclarified.com/Oi-Ph/Oxygen-Family.html

Rate Coefficients

Let’s define a few terms.

The concentration of a reactant is usually given in mol L-1, and is denoted by square brackets, as in [O3].

The reaction rate is the ratio of the change in concentration to the elapsed time = sources – sinks. The unit for rate is mol L-1 s-1. Depends on the concentration of the reacting species and a rate

constant


constant.

For example, rate = k[X]m[Y]n

The rate constant (or coefficient), k, describes the kinetics of a chemical reaction (and depends on temperature). Its units depend on the reaction.

The photolysis rate constant (or coefficient) (or j-value), j, is calculated from the product of the absorption cross section of the molecule, the quantum yield for the process, and the actinic flux, all integrated over the wavelength region of interest. Units are s-1.

Wallace and Hobbs Exercise 5.24

In the troposphere, the primary sink for CH4 is the reaction

CH4 + OH CH3 + H2O.

The rate coefficient for this reaction at a temperature typical of the troposphere is ~3.5 10-15 cm3 molecules-1 s-1.

If the average 24 h number density of OH molecules in the


If the average 24-h number density of OH molecules in the atmosphere is 1 106 molecules cm-3, what is the residence time of a CH4 molecule in the atmosphere?

Wallace and Hobbs Exercise 5.25Propane (C3H8) is a nonmethane hydrocarbon (NMHC) that reacts with OH via

C3H8 + OH C3H7 + H2O,

with a rate coefficient of 6.1 10-13 cm3 molecule-1 s-1 in the troposphere.

(a) Assuming the same OH number density as Exercise 5.24,


what is the residence time of C3H8 in the atmosphere?

(b) If the residence times of other NMHC are closer to that of C3H8 than to that of CH4, would you recommend that more attention be paid to the regulation of CH4 or to NMHC emissions?

(c) Why is CH4 the only hydrocarbon to enter the stratosphere in appreciable concentrations?

Wallace and Hobbs Exercise 5.27

The rate coefficient for the reaction

O + O3 2 O2

is k = 8.0 10-12 exp ( –2060 / T) cm3 molecule-1 s-1

(with temperature T in K).

If the temperature decreases from 20C to 30C what would


If the temperature decreases from –20 C to –30 C, what would be the percentage change in the rate of removal of O3 by this reaction?

Tropospheric Aerosols

Aerosols are suspended liquid and solid particles

They play a role in scattering and absorption of radiation, in the formation of cloud particles, and in atmospheric chemistry



22

Sources of Tropospheric Aerosols


Wallace and Hobbs, Table 5.3

Sinks of Tropospheric Aerosols

Particles are removed at about the same rate as they enter the atmosphere

Small particles (< ~0.2 m) coagulate to form larger particles, which can be removed

Precipitation processes remove 80-90% of the mass of particles through scavenging – aerosols serve as


p g g gcondensation nuclei, cloud drops collect particles by diffusion, and falling precipitation sweeps up more particles (impaction)

Dry deposition (gravitational settling) also removes aerosols (~10-20% of total mass removed)

Aerosol Sources and Sinks


http://www.ems.psu.edu/~lno/Meteo437/Aerosol.jpg

Particle Number Distributions

N(D) = number concentration of particles of diameter D

dN = number concentration of particles of diameter between D and D + dD

ContinentalMarine

Urban polluted


Note that N drops off rapidly as D increases. So the total number concentration (the Aitken or condensation nucleus (CN) count) is dominated by particles with D < 0.2 m, called Aitken or condensation nuclei


Particle Number Distributions

Straight line portions of the curves can be written:

ContinentalMarine

Urban polluted

CD)D(logd

dN

DlogC)D(logd

dNlog


where C = constant related to particle concentration and - = slope (usually 2-4)

A size distribution with = 3 is called a Jungedistribution

Note three regimes


Particle Surface and Volume Distributions Surface

distribution

Continental

Peak at D = 0.2 to 2 m due to coagulation and

evaporation: the accumulation mode

Peak at D > 1 m due


Volume distribution

ContinentalUrban

polluted

to dust and industrial

processes: the coarse particle mode

23

Lifetimes Smallest particles

(D < 0.01 m) have lifetimes < 1 day Removed by diffusion

to cloud particles, coagulation

Largest particles (D > 20 m) also

ContinentalMarine

Urban polluted

Surface distribution


(D > 20 m) also have lifetimes < 1 day Removed by

sedimentation, impaction, precip

Particles in accumulation mode have long lifetimes, Days in lower-middle

troposphere to months in upper troposphere


Air Pollution (Smog = smoke + fog)

If anthropogenic emissions of pollutants become large enough, this can adversely affect air quality, visibility, and human health.

Severe pollution episodes occur when rates of emission and/or formation of pollutants are much greater than the rates of dispersion, deposition, or chemical removal.


Combustion is the largest source of pollutants (CO, CO2, NOx, SO2) In ideal combustion, only CO2 and H2O are released – this is the

goal of modern vehicles which control the air-fuel ratio and have catalytic converters

Types of Smog London (or classic smog)

occurs mostly during cold winter days and consists primarily of a mixture of sulphur dioxide and particulate matter ("soot") usually derived from burning coal

particles swell under high relative humidity and serve as nuclei for fog droplets

SO2 dissolves in the droplets and is oxidized to form sulphuric acid


Los Angeles (or photochemical smog) occurs mostly during warm, sunny summertime days in strongly

polluted urban air that is subject to sunlight and stagnant meteorological conditions

characterized by a brown haze that reduces visibility and contains oxidants, such as ozone that cause respiratory problems, eye irritation and damage to plants

a mixture of pollutants: ozone, reactive hydrocarbons, NO, NO2, aldehydes, peroxyacetyl nitrate (C2H3NO5 = PAN), particulate matter

London Smog, 5-9 December 1952

Central Press/Hulton Archive/Getty Images


http://www.ems.psu.edu/~lno/Meteo437/Smoglond.jpg

Central Press/Hulton Archive/Getty Images

BBC: http://news.bbc.co.uk/2/hi/2545747.stm

Smog Formation


http://en.citizendium.org/wiki/File:Smog_formation.png

Evolution of Photochemical Smog

PAN = peroxyacetyl nitrate


24

Global Pollution: Arctic Haze

Arctic haze consists of pollution episodes due to the transport of contaminants from lower latitudes, primarily from December to April.


“Pathways of contaminants to the Arctic. Many POPs (persistent organic pollutants), heavy metals and other contaminants from emissions further south are accumulated in Arctic food chains and ultimately in indigenous peoples. This process is often referred to as long-range pollution or long-range transport of pollutants. While fear of these compounds sometimes has resulted in abandonment of traditional foods, this has also led to more unhealthy food habits acquired from non-indigenous peoples. Most indigenous peoples in smaller communities still supply a large share of their household foods from natural resources.”

AMAP 2002, ACIA 2004, Hugo Ahlenius, UNEP/GRID-Arendal http://maps.grida.no/go/graphic/pathways_of_contaminants_to_the_arctic

Stratospheric Chemistry

Concentrations of some gases change abruptly across the tropopause into the stratosphere e.g., water vapour decreases, ozone increases

These strong vertical gradients across the tropopause are because there is very little vertical mixing between the moist ozone-poor troposphere and the dry ozone-rich stratosphere


The stratosphere is convectively stable or neutral

There is no precipitation in the stratosphere to remove aerosols and trace gases, so constituents can remain there for long periods in stratified layers e.g., volcanic injections, aircraft emissions, chemicals transported

upwards by strong updrafts

We will look at stratospheric chemistry related to ozone.

Stratospheric Aerosols: Junge Layer

The Junge layer or stratospheric sulphate layer was discovered in the late 1950s

Consists of aerosol particles of diameter ~0.1-2 m at concentrations of ~0.1 cm-3

located at ~17-20 km

Composition is ~75% sulphuric acid and ~25% water


acid and ~25% water

Produced by oxidation ofSO2 to SO3, and reactionwith H2O to make H2SO4

H2SO4 vapour combines with water or condenses onto existing particles to make liquid H2SO4

Volcanoes are the primary source of stratospheric SO2 http://www.ems.psu.edu/~lno/Meteo437/Junge.jpg

Monthly Average Stratospheric AOD at 1 m

Aerosol optical depth (AOD) reflects the aerosol loading

AOD shows impact of PSCs (in Antarctica) and volcanoes


Impact of Mt. Pinatubo

Th ti f Mt Pi t b i

USGS/Cascades Volcano Observatoryhttp://vulcan.wr.usgs.gov/Volcanoes

/Philippines/Pinatubo/images.html


The eruption of Mt. Pinatubo in June 1991 caused a 30- to 40-fold increase in the surface area of particles available for enhancing chemical reactions

The aerosols remain in the stratosphere for 2-5 years

Record-low ozone levels were observed in 1992-1993

http://www.esrl.noaa.gov/csd/assessments/ozone/1998/faq4.html

Part 1


Introduction




Clouds and Aerosols


25

CloudsClouds - consist of water drops and ice crystals of radius ~10 m and are classified on the basis of size: cloud drops r = 10 m

drizzle r = 100 m

rain drops r = 1000 m (1 mm)


http://physics.uwstout.edu/wx/Notes/ch4notes.htm

How Are Clouds Formed? - 1

Warm, moist air rises adiabatically from Earth's surface.

As it rises it cools, raising the relative humidity of the air.

When the RH reaches 100%, the air is saturated with water.

This water vapour could stay in


This water vapour could stay in the atmosphere indefinitely. For a droplet to form by

condensation from the vapour, the surface tension must be overcome by a strong gradient in vapour pressure.

For “pure” water vapour, condensation does not occur at 100% RH.


How Are Clouds Formed? - 2

However, in the atmosphere, water does begin to condense when the air is saturated, and clouds do begin to form.

Why? Impurities in the air such as dust, smoke or air pollution attract and hold water molecules


together forming large droplets. Clouds form as the air reaches saturation and droplets begin to form around these impurities, called cloud condensation nuclei.

The type of cloud formed depends on factors such as air temperature, relative humidity, wind speed, and air pollution.


If the cloud continues rising and cools to <0°C, the water droplets

may freeze if ice nuclei are present.

How Cloud Drops Grow

A cloud drop has a typical diameter of ~ 10 m whereas a raindrop of sufficient size to precipitate is ~ 1 mm in diameter.

Question: How do cloud drops grow?

by diffusion of water vapour to the cloud particles and subsequent condensation

by collision and coalescence of particles



Why Do We Care About Clouds?

Clouds cover ~50% of the Earth.

Clouds are important because they:

influence Earth’s radiation budget

influence climate

influence daily weather

interfere with radiometric observations of the atmosphere and surface


Clouds influence the energetics of the atmosphere in two ways:

(1) By their role in the atmospheric water cycle.

latent heat is released on condensation and liquid water is removed from the atmosphere on precipitation.

(2) By scattering, absorption, and emission of solar and terrestrial radiation.

Reflection of solar radiation:Cooling effect

Blanketing of thermal infrared radiation:Warming effect

Why Do We Care About Clouds?


Evaporation Boundary layer

Precipitation

26

Clouds and Climate

Clouds are important to climate because they strongly modulate incoming solar and outgoing thermal radiation.

Clouds are the source of precipitation therefore they are a key element in the hydrologic cycle.

We need a better understanding of clouds and climate for better predictions of climate change, to guide policy in


p g , g p yameliorating or adjusting to change, and to provide better stewardship of our water resources.

Aerosols and Clouds?

Aerosol particles are important to climate: directly by scattering light

indirectly by serving as cloud condensation nuclei and by altering cloud properties

Clouds and Pollution - 1

Air pollution attracts and holds water molecules together, forming large droplets.

Clouds form as the air reaches saturation and droplets begin to form around these impurities.

The extent to which pollution affects level of cloudiness depends on the kind of pollution. A l li id lid ti l d d i th i


Aerosols are liquid or solid particles suspended in the air.

With increasing concentrations of trace gases of anthropogenic origin, the resulting clouds have different microphysical properties than clouds formed in clean atmospheric conditions.

Typically the number of particles increases and their sizes decrease in polluted air masses.

This can lead to different optical properties between clouds formed in polluted and clean conditions.

http://www.atm.helsinki.fi/aerosol/clouds.html

Clouds and Pollution - 2

Direct aerosol effect

aerosol particles scatter (cooling) and absorb (warming) solar and infrared radiation, thereby causing a direct radiative forcing

Indirect aerosol effects

1st aerosol particles alter cloud formation processes by increasing the concentrations


processes by increasing the concentrations of droplets and ice particles

2nd they decrease the precipitation efficiency of warm clouds and thereby cause an indirect radiative forcing

Aerosols have most likely made a significant negative contribution to the overall radiative forcing. However, their short atmospheric lifetimes means that they cannot be considered simply as a long-term offset to the warming influence of GHGs.

Cloud Formation in the Clean and Polluted Atmosphere

Clean Polluted

Atmospheric Physics and Sustainable Aviation (K. Strong) Slide 154Courtesy: Ulrike Lohmann

First Indirect Aerosol Effect

For the same liquid water content, the clean cloud (on left) reflects less sunlight because it has fewer and larger cloud droplets (smaller surface area) than the polluted cloud (on right).

Pollution: more small drops larger surface area more reflection cooling


Clean Polluted

Courtesy: Ulrike Lohmann

Second Indirect Aerosol EffectThe second indirect effect is due to additional effects of having more cloud condensation nuclei and hence more small droplets.

Smaller droplets less likely to collide lower precipitation efficiency increases the cloud cover and amount of liquid more reflection of solar radiation cooling


IPCC 2007

27

Radiative Effects of Clouds and Aerosols

Direct effect

BC

OC

Indirect cloud albedo effect

Courtesy: Ulrike Lohmann


Indirect cloud lifetime effect

Anthropogenic emissions

SO2

SO4--

OC

Cloud evaporationSemi-direct effect

Climate models suggest anthropogenic aerosol effect ranges from -1 to -4.4 W/m2

Past observations and simple models suggest 0 to -1.2 W/m2

Classification of Clouds - 1A little bit of history….

The French naturalist Lamarck (1744-1829) proposed a system for classifying clouds in 1802, but it did not receive wide acclaim.

One year later, Luke Howard, an English naturalist, developed a cloud classification system that did find general acceptance: sheet-like clouds as stratus (Latin for layers)

puffy clouds as cumulus (heap)


puffy clouds as cumulus (heap)

wispy clouds as cirrus (curl of hair)

rain clouds as nimbus (violent rain).

These are the four basic forms of clouds. Other clouds can be described by combination of these basic types. e.g., nimbostratus is a rain cloud that shows layering, whereas

cumulonimbus is a rain cloud that has pronounced vertical development

From a paper by S. Veerabuthiran - see: http://www.ias.ac.in/resonance/March2004/pdf/March2004p23-32.pdf

Classification of Clouds - 2

Clouds are generally classified based on characteristics such as altitude, appearance, or origin.

Altitude distinctions apply to those clouds that fit in various layers of the atmosphere as follows:

High clouds - have bases above 18,000 feet

Middle clouds - have bases between 7,000 and 18,000 feet


Low clouds - have bases below 7,000 feet

Fog - cloud in contact with the ground

Multi-level clouds... vertically thick spanning multiple layers

Orographic clouds - distinct clouds that form via interaction between wind and mountainous terrain features.

In appearance, clouds may be thick or thin, have well defined edges or be very diffuse, appear hairlike, cellular, towering, or in sheets, and be associated with fair weather or precipitation.

Classification of Clouds - 3 High clouds are primarily composed of ice crystals and include:

cirrus cirrocumulus cirrostratus Middle clouds contain ice crystals and/or water droplets and may

be associated with light precipitation: altocumulus altostratus

Low clouds are most often composed of water droplets, but can have ice crystals in colder climates:


y cumulus stratocumulus stratus fog

Multi-layer clouds are the heavy precipitation producers:

nimbostratus cumulonimbus

Orographic clouds are produced by the flow of air interacting with mountainous terrain:

cap clouds lenticular clouds

Special cloud types: Kelvin-Helmholtz instability waves aircraft contrails

Cirrus clouds - are high altitude wispy clouds. They are usually quite

thin and often have a hairlike or filament type of appearance.

Cirrus and Contrails

Atmospheric Physics and Sustainable Aviation (K. Strong) Slide 161http://vortex.plymouth.edu/clouds.html

Contrails (short for "condensation trails") can be formed from the vapor contained in the exhaust of a jet engine of an airplane when flying at high enough altitudes where cold temperatures cause the vapor to turn into ice crystals like cirrus clouds. These clouds look like lines in the sky.

Part 1


Introduction




Clouds and Aerosols


28

Scales of Motion

Molecular scale (<< 2 mm)Molecular diffusionMolecular viscosity

Synoptic scale 500-10,000 km)Pressure systemsWeather frontsTropical stormsHurricanesAntarctic ozone hole

Microscale (2 mm- 2 km)EddiesSmall plumes


Antarctic ozone hole

Jacobson, Fundamentals of Atmospheric Modeling 2nd Editionhttp://www.stanford.edu/group/efmh/FAMbook2dEd/index.html

pCar exhaustCumulus clouds

Mesoscale (2 - 2000 km)Gravity wavesThunderstormsTornadosLocal windsUrban air pollution

Planetary scale (>10,000 km)Global wind systemsRossby wavesStratospheric ozone lossGlobal warming

The Basics: Atmospheric Motion

Atmospheric motion is described by three velocity components:

drd

component meridionaldtd

Rdtdy

v

component zonaldtd

cosRdtdx

u

E

E


where

= latitude (positive in northern hemisphere)

= longitude (positive eastward)

r = height above sea level (radial direction)

component verticaldtdr

dtdz

w


The Basics: Atmospheric Motion

The horizontal velocity vector is V u i + v j (bold means a vector)where

i = unit vector in the zonal (x) direction

j = unit vector in the meridional (y) direction

Usually V >> w, i.e., the horizontal velocity is much greater than the


vertical velocity. So wind usually means horizontal velocity component.

A positive zonal velocity is westerly (from the west = eastward).

A negative zonal velocity is easterly (from the east = westward).

A positive meridional velocity is southerly (from south = northward).

A negative meridional velocity is northerly (from north = southward).

Eulerian and Lagrangian

There are two ways to describe the rate of change of a property of a fluid element

Eulerian derivative – describes the rate of change at a fixed point through which the fluid is moving /t = rate of change at a fixed point in rotating (x,y,z) space

Lagrangian derivative – describes the rate of change of an


g g gelement of fluid as that element is moving, rather than at a fixed point in space d/dt = rate of change following an air parcel as it moves along its

three-dimensional trajectory. Note, sometimes written as D/Dt.

Related by the Chain Rule:z

wy

vx

utdt

d

Material Derivative Suppose a fluid is associated with a velocity field V(x,t).

What is the rate of change of a fluid property, e.g., T?

elementfluidtheofpositionofchangeofrated

but

Tdtd

tT

dtdz

zT

dtdy

yT

dtdx

xT

tT

dtdT

t),t(T)t(z),t(y),t(xTT

x

x

x


So the total derivative (= material derivative) is:dtdz

w,dtdy

v,dtdx

u withdtd

)w,v,u(

element fluid the of position of change of ratedt

but

xV

TtT

dtdT

DtDT

V

Temperature advection = contribution to change in T due to fluid motion

Rate of change following the motion

Local rate of change at a fixed point

The Dynamics of Horizontal Flow

We can use Newton’s Second Law to describe the motion of an air parcel (in an inertial reference frame):

Three real forces act here: gravity, pressure gradient, and friction.

However, Earth is a rotating reference frame, i.e. accelerating,

i

iFm1

a


and so we need to introduce two apparent forces.

(1) Centrifugal force experienced by all objects in the rotating reference frame

pulls them outwards from the axis of rotation

(2) Coriolis force magnitude depends on the relative velocity of the object in the

plane perpendicular to the axis of rotation

also in the plane perpendicular to the axis of rotation and is perpendicular to the direction of motion

29

The Geostrophic Wind

When the Coriolis force and the pressure gradient force are balanced, it can be shown that the geostrophic wind blows parallel to isobars, with low pressure to the left (Northern hemisphere).

In both hemispheres, the geostrophic wind blows cyclonically around low pressure systems and anticyclonically around high pressure systems.

The tighter the spacing of isobars


The tighter the spacing of isobars(i.e., the larger p, since p constantand x or y smaller), the larger thepressure gradient force and thelarger the Coriolis force needed to balance it, hence the flow speeds up.

The pressure field determinesthe geostrophic wind field.


Geostrophic Balance

Flow is parallel to the isobars



Blue arrows show the sense of the circulation around high pressure systems (H) and low pressure systems (L), looking down on the South Pole (left) and North Pole (right).

Small arrows circling the poles show the sense of Earth’s rotation.

Cyclonic and Anticyclonic Circulation


Surface Winds and Pressure Contours

Winds are parallel to pressure contours. Note the direction of the

i d d


winds around high and low pressure systems in the northern and southern hemispheres.

The Primitive Equations

The primitive equations are the fundamental equations that describe large-scale atmospheric motion in both the horizontal and vertical.

The state of a fluid (atmosphere or ocean) at any time is defined by five variables:

Tp)wvu(V


These variables can be related by five independent equations:

The laws of motion applied to a fluid parcel, giving three equations in three directions, or two equations for motion in the horizontal (2-D) and vertical

The law of thermodynamics

The conservation of mass

T,p),w,v,u(V

Summary of the Primitive Equations

Horizontal equation of motion: Note that this has two

(x and y) components

Vertical equation of motion(Hydrostatic equation):

FVk

FVkV

f

fp1

dtd

pRT

porg

zp


Thermodynamic energy equation:

Continuity equation: or

Bottom boundary condition:

pcJ

pT

dtdT

0dtd

V Vp

sp

0ssss dp

zp

wpt

pVV

30

Using the Primitive Equations

Five equations with five dependent variables: u, v, w, (or p), T.

J and F are prescribed (parameterized) in terms of these variables.

The equations with time derivatives are prognostic equations.

The equations without time derivatives are diagnostic equationsand describe the relationships between the variables at any time t.


These equations can be numerically solved

Typically in global models, the variables are defined on a set of horizontal and vertical grid points and the equations are converted to Eulerian forms so that they can be calculated on this fixed set of grid points.

Some models use spherical harmonics to simplify the calculations – these are mutually orthogonal analytic functions (products of Legendre polynomials in latitude and sines and cosines in longitude) whose coefficients are used in the algebraic operations.


The atmospheric general circulation describes the global statistical properties of large-scale atmospheric motions.

We return to solar heating...

Annually averaged incoming solar radiation per unit area of Earth’s surface is greatest at the equator.


The tropics receive more energy from the Sun than they emit to space.

The polar regions emit more energy to space than they receive from the Sun.

Equilibrium requires some process to transport excess energy from the tropics to high latitudes, making the tropics cooler and the polar regions warmer than they would otherwise be.

This differential heating drives the atmospheric general circulation.

Kinetic Energy and Potential Energy

In small-scale convection, the kinetic energy generated by the upward buoyancy force on the rising plume is given to the vertical wind field (only vertical heating gradients are needed).

In large-scale atmospheric motions, the kinetic energy is given to the horizontal wind field (horizontal & vertical heating gradients needed). As warm (cold) air rises (sinks), the released potential energy does

work on the horizontal wind field by forcing it to flow across isobars


y gfrom higher to lower pressure.

The atmosphere is heated by radiative processes and by the release of latent heat of condensation in clouds.

As a result, potential energy is replenished by: (i) heating (cooling) in the tropics (polar regions), which maintains the

equator-to-pole temperature gradient on pressure surfaces, and

(ii) heating at lower levels and cooling at higher levels, which expands (compresses) air in the lower (upper) troposphere, lifting intermediate air and maintaining the height of the atmosphere’s centre of mass.

Earth’s Energy Budget

Atmospheric circulation must transport 61015 W



Features of the General Circulation

Tropospheric jet streams – westerly winds at mid-latitudes (~30) Blow from west to east year-round, strongest in winter

Mesospheric jet streams at ~60 km altitude in mid-latitudes Blow from the west in winter and from the east in summer

Baroclinic waves – derive energy from north-south temperature gradients at mid-latitudes, develop in response t i t biliti i th l l fl t t d


to instabilities in the large-scale flow, propagate eastward Dominated by extratropical cyclones in low-level flow

Cyclone = closed circulation in which the air spins in same direction as Earth’s rotation seen from above (CCW in NH, CW in SH)

Tropical cyclones derive energy from the release of latent heat of condensation in deep convective clouds Tend to be tighter, more axisymmetric, and more intense

than extratropical cylones, characterized by cloud-free eye at the centre

Zonal Temperatures and Winds


31

For an idealized aqua planet with no land-sea contrast, with the sun overhead on the equator:

The westerly wind belts dominate the extratropical circulation.

These are perturbed by the eastward-propagating baroclinic waves( ~ 4000 km, u ~ 10 m/s)

The trade winds dominate tropical circulation Northeasterly in NH southeasterly in SH

Features of the Surface Wind Field



Northeasterly in NH, southeasterly in SH

Due to the overturning Hadley cells

The extratropical westerlies and trade winds are separated by a tropical high pressure belt at 30 Weak surface winds, below the jet

streams at the tropopause

Weak sea-level pressure minimum at equator (trade winds converge)

Subpolar low pressure belt as extratropical lows migrate poleward

Surface Winds over the Oceans

In the real world, with land and ocean, the surface winds are usually stronger over the oceans due to less surface friction.

The winds over the Atlantic and Pacific exhibit an additional longitudinally dependent feature: the subtropical high-pressure belt has a series of distinct centres rather than being continuous. These are called subtropical

anticyclones centred over oceans


anticyclones, centred over oceans.


The ITCZ and Monsoon Circulation

The trade winds converge over the Atlantic and Pacific at about 7N – a belt called the Intertropical Convergence Zone (not on the equator due to land-sea geometry).

The surface winds over the tropical


The surface winds over the tropical Indian Ocean are dominated by the seasonally reversing monsoon circulation – begins as westward flow in winter hemisphere, crosses equator, and curves eastward, forming a belt of moist westerlies in the summer hemisphere. Driven by north-south land-sea

temperature differences Wallace and Hobbs, Figure 1.17

Surface Winds and Sea-Level Pressure Surface winds generally

blow parallel to the isobars, except in the tropics.

There is a systematic drift across the isobars from high to low p.

DJF


The Icelandic and Aleutian lows (mid-ocean cyclones) are associated with the subpolar low-pressure belt in winter.

NH oceanic subtropical cyclones are seen in summer.


JJA

P contour is 5 hPa. Blue: p > 1015 hPa. Yellow: p < 1000 hPa.

Atmospheric General Circulation Strong westerly

zonal flow at mid-latitudes (due to T decrease with latitude)

Surface winds are weak due to friction and thermal wind


and thermal wind balance

Surface winds are easterly in tropics and westerly at mid-latitudes; westerly angular momentum is transported from low to mid-latitudes to maintain equilibrium From Marshall and Plumb

Mean WindsZonal mean zonal winds are mostly eastward (westerly). Winds are stronger in the core of the subtropical jets. Tropical easterlies are weaker. Low-level easterlies are the trade winds.


Zonal mean meridional winds show the meridional overturning circulation (arrows).

There is upward motion on the summer sides of the equator due to convection, and downward motion on the winter side of the

equator. So the annual mean shows two weaker cells. The vertical motion is

accompanied by meridional flow, which makes the trade winds north-easterlies (NH

winter) and south-easterlies (SH winter).

32

Hadley Circulation

The west-to-east circulation in the upper troposphere dominates the large-scale atmospheric flow but does not cause the poleward transport of energy and angular momentum

The Hadley circulation provides this north-south flow (in the meridional plane)

Ch t i ti ( ll d)


Characteristics (annually averaged):

Upwelling near the equator

Poleward flow aloft

Subsidence in the subtropics

Equatorward return flow near the surface

This transports energy and angular momentum poleward within the tropics, but not in the extratropics (mid-latitudes)


Hadley Circulation



So What Happens at Higher Latitudes?

The Hadley circulation transports heat polewards in the tropics, but does not do so at high and mid-latitudes

An additional source of poleward heat transport is needed to balance the global energy budget

The answer lies in eddies, which arise when flow in the extratropics is broken down through baroclinic instability


p g y

The Coriolis acceleration is greater at mid-latitudes than at the equator – results in strongly eddying flow

Closed Circulation Cells

The baroclinic waves drive weak meridional circulation cells in the NH and SH – these are called Ferrel cells. Poleward Ekman drift at mid-latitudes (~45) due to friction, ascent

on the poleward side, and descent on the equatorward side.

This forces the Hadley cells into the tropics, so that descent in a Hadley cell occurs in the subtropics (~30).

A third type of meridional circulation cell


A third type of meridional circulation cell, the polar cells, also exist. Relatively warm, humid air at ~60C

rises, moves poleward in the upper troposphere, cools in the polar regions, descends as a cold, dry high pressure area, moves equatorward, and twists westward due to the Coriolis force to produce polar easterlies.

http://sparce.evac.ou.edu/q_and_a/air_circulation.htm

Thermally Direct and Indirect Circulations

The Hadley cells and polar cells are characterized by rising of warm air and sinking of cold air, with cross-isobar flow towards lower pressure, which releases potential energy and converts it into kinetic energy of the horizontal flow. These are called thermally direct circulations – they operate in the same sense as the global kinetic energy cycle.

The Ferrel cells operate in the opposite sense with rising of cold


The Ferrel cells operate in the opposite sense, with rising of cold air and sinking of warm air. These are called thermally indirectcirculations.

Globally, the loss of kinetic energy through frictional dissipation is offset by the release of potential energy through the sinking of cold, dense air and the rising of warm, less dense air.

Latitudinal Variations of Climate

The atmospheric circulation determines climate as well as winds:

Equatorial regions – convergence of trade winds; frequent and intense rainfall – rainforests

Subtropics (20-30) – warm, dry, descending branch of Hadley circulation; hot, dry climate – subtropical deserts

Poleward of 30 – baroclinic eddies dominate meteorology; local i d d il b t tl t l d th


winds vary daily but are mostly westerly; good weather associated with high pressure systems, stormy weather associated with low pressure systems

Solar forcing varies with season, shifting the Hadley circulation

There are longitudinal variations in climate

The oceans also transport heat and affect climate

33

Latitudinal and Vertical Variations in Wind and Climate


Motion on Smaller Scales Thermals are plumes of rising air, driven by solar heating of the

surface – generate cumulus clouds. Accompanied by subsidence.

Shallow convection occurs when the overturning circulation is confined to the boundary layer.

Deep convection occurs when plumes break through the temperature inversion in the boundary layer, reaching the tropopause (mostly in the tropics and warm, humid mid-latitude air).


p p ( y p )

Large-scale flow over surface topography causes boundary layer turbulence, generating waves and eddies as scales up to a few km.

Turbulence can also be generated by flow instabilities at higher altitudes, such as waves that develop in layers with strong vertical wind shear.

When such waves amplify and break, this wave breaking generates smaller-scale waves and eddies that become unstable. A succession of instabilities converts kinetic energy from the large-

scale wind field into a spectrum of small-scale motions.


Part 2


Part 2

Stratospheric Ozone

Ozone – Our Ultraviolet Shield

2011 Antarctic


ozoneas seen from spaceSource: NASA

ozonewatch.gsfc.nasa.gov

Stratospheric Ozone

http://en.wikipedia.org/wiki/Ozone

What is ozone?

A special form of oxygen (O3) that blocks harmful UV light from the Sun

A very reactive gas, present in small but significant quantities in the


WMO Ozone Assessment 2006

quantities in the atmosphere

Its concentration depends on altitude, with most ozone lying in a layer 20 km thick centred at 25-30 km

“To the Philosopher, the Physician, the Meteorologist, and the Chemist, there is perhaps no subject more attractive than that of Ozone.” Cornelius B. Fox, British chemist (1839-1884)

How Much Ozone Is There?

If all of the air in a column above us were compressed to surface pressure (0C and 1 atm) … …it would be a layer of air about 8 km thick.

If all of the ozone were separated out and compressed … … it would make a layer of ozone about 3 mm thick.

About the thickness


About the thicknessof 3 dimes!

Ozone is measured in Dobson Units (DU)

1 DU = 0.01 mm300 DU = 3 mm

34

Why Is Ozone Important?

Ozone absorbs solar radiation from 230-320 nm, including UV-B (290-320 nm) Only 1 part in 1030 of

250-nm solar radiation at TOA reaches the surface


Kerr and Fioletov, Atmos-Ocean 2008

surface

This warms the stratosphere (~10-50 km) Determines vertical

temperature structure

Ozone is also a GHG, absorbing IR radiation and heating the troposphere (~ 0-10 km)

Latitudinal Distribution of Stratospheric Ozone

Ozone concentrations vary with altitude, latitude, season, and meteorology

Production is a maximum in tropical stratosphere

Largest total columns are


gin spring Arctic in NH, and in spring mid-latitudes in SH due to meridional and downward transport (Brewer-Dobson circulation)

http://www.ccpo.odu.edu/~lizsmith/SEES/ozone/class/Chap_3/index.htm

Brewer-Dobson Circulation

Alan W. Brewer (1915–2007)UofT1962-1977See http://www.atmosp.physics.utoronto.ca/SPARC/News15/15_Norton.html

Gordon M.B. Dobson

(1889–1976)See

http://www.atm.ox.ac.uk/user/barnett/ozoneconf

erence/dobson.htm


http://www.ccpo.odu.edu/~lizsmith/SEES/ozone/class/Chap_6/index.htm

Stratospheric Ozone Chemistry

How is ozone created and destroyed?

(1) Chapman Cycle (1930) - oxygen-only reactions

Odd oxygen production:

Odd d t ti

32 M2

OOOOOhO

OOhO 23

Sydney Chapman(1888–1970)


Odd oxygen destruction:

(2) Catalytic Cycles (1970s) - destroy ozone

where reactive species X (= H, OH, NO, Cl, Br) is recycled

Crutzen, 1971Johnston, 1971Molina and Rowland, 1974

]slow OOO[

O2OO

2M

23

23

232

23

O2OOOXOXO

OXOOX

The Chapman Cycle PredictionsThe Chapman reactions:

Predict the correct shape for the vertical profile of ozone

Overestimate stratospheric [O3] by a factor of two in the tropics

Underestimate stratospheric [O3] at middle and high latitudes Predict too large a global production rate of ozone

Missing sinks:Chapman overpredicts b f t f 2

k1=jak =j


The maximum reflects ja (= k1), which is affected by:(1) decreasing [O2] with z

following (Barometric law)(2) increasing h with z

• Catalytic cycles• Brewer-Dobson

circulation

Alti

tude

Increasing photolysis

with altitude

by a factor of 2k3=jc

Based on www.chm.bris.ac.uk/~chdrg/Glowacki-AT207-L4.ppt

In 1985, a team of scientists from British Antarctic Survey reported that springtime stratospheric ozone column over their station at Halley Bay had

Discovery of Ozone Hole


at Halley Bay had decreased precipitously since 1970s.

Occurs in September-November.

No depletion was observed in other seasons. Farman et al., Nature, 1985

35

Antarctic Ozone DepletionBalloon-borne ozone profiles measured at South Pole in October

Measured vertical profiles show that the depletion of ozone is essentially total in the lowest region of the


the lowest region of the stratosphere between 10-20 km, which normally contains most of the total ozone column in polar spring. blue = 1967-1971 average red = lowest total ozone

recorded in 2001 green = lowest total ozone

recorded in 1986 NOAA CMDL , http://www.cmdl.noaa.gov/ozwv/ozsondes/spo/index.html

Antarctic Ozone Hole: 1979-2004


NASA/Goddard Space Flight Center Scientific Visualization Studio http://svs.gsfc.nasa.gov/vis/a000000/a003100/a003137/index.html

Polar Total Ozone Trends



Global Changes in Total Ozone



Vertical Total Ozone Trends

Measured and modelled verticalprofiles of ozone trends



Ozone Loss - Why Do We Care?

Decreasing ozone in stratosphere:

Increases UV radiation at Earth’s surface affects life

Leads to feedback processes: More UV radiation from the sun can lead to more ozone formation

in the troposphere

Ozone is a toxic gas for which air quality standards have been set


Tropospheric ozone is a greenhouse gas

More UV in troposphere will increase OH production affects other trace gases …

Changes the radiative balance in the stratosphere Less ozone less absorption of solar UV less IR heating

cooling of the stratosphere

36

Causes of Ozone Depletion

Emission of gases containing chlorine and bromine

chlorofluorocarbons (CFCs), halons, CH3Cl, CCl4, CH3Br, …

Transport of these gases

around the globe and into the stratosphere

Conversion to reactive halogen (Cl, Br) gases


halogen source gases are broken down by sunlight

Ozone-destroying chemical reactions

results in global chemical depletion of ozone (~3%)

Severe polar springtime depletion due to formation of polar stratospheric clouds and subsequent chemical reactions

Antarctic Ozone and ClO


From Rowland, 2006

Polar Ozone Depletion Processes

(1) Formation of the winter polar vortex (band of westerly winds) isolates cold dark air over the polar regions

(2) Low temperatures in the vortex, T<195 K polar stratospheric

clouds (PSCs) form in the lower stratosphere


stratosphere (liquid and solid HNO3, H2O, H2SO4)


Polar Ozone Depletion Processes

(3) Dehydration and denitrification remove water vapour and nitrogen oxides which would otherwise

react with and neutralize chlorine

(4) Release of CFCs, mixing, and transport to the polar regions enhanced levels of chlorine and other halogen species

(5) Heterogeneous reactions on the PSCs


convert inactive chlorine (HCl and ClONO2) to reactive Cl2

(6) Sunlight returns in the spring UV radiation breaks Cl2 apart to form Cl

(7) Catalytic chlorine cycles destroy ozone, while recycling Cl

This continues until the Sun causes a dynamical breakdown of the winter vortex and PSCs evaporate.

Formation of the Ozone Hole


Processes Affecting Stratospheric Ozone and Temperature

Chemical reaction rates

Stratospheric circulation

Stratospheric ozone

Stratospheric temperature UV

PSC formation


Brasseur, SPARC Lecture 2004, after Schnadt et al., Climate Dynamics 2002

Vertical propagation of planetary and gravity waves Anthropogenic emissions of

CO2, CFCs, CH4, N2O

Greenhouse gases

Troposphere-stratosphere exchange

Stratospheric chlorine, bromine, and nitrogen

oxidesCH4 oxidation

Stratosphericwater vapour

37

Radiative Forcing (1750-2000)



1985 - Vienna Convention for the Protection of the Ozone Layer

1987 - Montreal Protocol on Substances that Deplete the Ozone Layer (the “Ozone Treaty”)

Montreal Protocol & Its Amendments


Entered into force in 1989

Established controls on halogen source gases

Later strengthened by a series of Amendments

WMO Ozone Assessment 2006, 2010



The Impact of Emissions



Ozone Recovery

Gradual recovery of ozone is anticipated as stratospheric chlorine decreases.

ozone turnaround in the Arctic likely before 2020

Atmospheric Physics and Sustainable Aviation (K. Strong) Slide 221WMO Ozone Assessment 2010

vunerable to perturbations, such as volcanic aerosols

coupled to stratospheric cooling

extreme Arctic ozone loss is not predicted

Recovery of Stratospheric Ozone



38

Questions

What are the feedbacks between ozone and climate?

We need better knowledge of these

How will the polar stratosphere respond to climate perturbations?

Particularly while Cl and Br loading is high


How will changes in atmospheric circulation affect polar ozone?

Cooling (more depletion) or warming (less depletion)?

Will cooling of the stratosphere due to increasing greenhouse gases cause greater Arctic ozone loss?

Summary

Measurements and models led to the discovery and explanation of ozone depletion

The Montreal Protocol and later Amendments were negotiated to deal with this global issue

Compliance and the development of “ozone-friendly” substitutes appears to be working


substitutes appears to be working

Abundances of CFCs are decreasing

Expect ozone layer recovery later this century

This is a success story involving discovery, understanding, decisions, actions, and verification by many players

The links between stratospheric ozone and climate change remain an area of active investigation

For Further Information

WMO (World Meteorological Organization) Scientific Assessments of Ozone Depletion 2006 and 2010 http://www.wmo.int/pages/prog/arep/gaw/ozon

e_2006/ozone_asst_report.html

http://www.wmo.int/pages/prog/arep/gaw/ozone 2010/ozone asst report html


e_2010/ozone_asst_report.html

Our annual springtime campaigns at Eureka, Nunavut http://acebox.uwaterloo.ca/eureka/

http://www.candac.ca


Part 3


Part 3

The Climate System

Radiative Forcing

Observed Changes

Climate Model Predictions

Some DefinitionsWeather

the fluctuating state of the atmosphere around us, characterized by the temperature, wind, precipitation, clouds and other weather elements

Climate

the average weather in terms of the mean and its


the average weather in terms of the mean and its variability over a certain time-span and a certain area

“Climate is what we expect, weather is what we get.”

Mark Twain

Climate change

statistically significant variations of the mean state of the climate or of its variability, typically persisting for decades or longer

The Climate System - 1

The climate system is an interactive system forced or influenced by various external forcing mechanisms, the most important of which is the Sun.

The atmosphere is the most unstable and rapidly changing part of the system.

The climate of the Earth as a whole depends on factors that influence the radiative balance such as for example the


influence the radiative balance, such as for example, the atmospheric composition, solar radiation or volcanic eruptions.

39

Climate system:

Atmosphere

Hydrosphere

Cryosphere

L d f

Components of the Climate System


Land surface

Biosphere

Schematic view of the components of the global climate system, their processes and interactions IPCC 2007 (http://www.ipcc.ch/publications_and_data/ar4/wg1/en/faq-1-2.html)

Earth’s Energy Budget


Kiehl and Trenberth: Earth’s Annual Global Mean Energy Budget, Bull. Am. Met. Soc. 78, 197-208, 1997. Image from “Our Changing Planet”, CUP, 2007.

Earth’s Energy Budget with No GHGs


Numbers arepercentages

Earth’s Energy Budget With GHGs

This infrared loop, consisting of infrared radiation absorbed and reradiated by greenhouse

i h t

Numbers arepercentages


gases, is what keeps Earth’s surface warmer than its upper atmosphere.

This greenhouse effect is natural, and has existed for many millions of years.

The Scientific Basis: IPCC ReportsThe Intergovernmental Panel on Climate Change (IPCC)was established in 1988 by the WMO and UNEP to assess scientific, technical and socio-economic information relevant for the understanding of climate change, its potential impacts, and options for adaptation and mitigation.

Four reports have been published: 1990, 1995, 2001, 2007 The Fourth Assessment Report (AR4) was released in 2007


The Fourth Assessment Report (AR4) was released in 2007 Working Group I - The Physical Science Basis Working Group II - Impacts, Adaptation, and Vulnerability Working Group III - Mitigation of Climate Change

Oslo, 10 December 2007 - The IPCC and Al Gore were awarded of the Nobel Peace Prize "for their efforts to build up and disseminate greater knowledge about man-made climate change, and to lay the foundations for the measures that are needed to counteract such change".

Intergovernmental Panel on Climate Change

IPCC Homepage

www.ipcc.ch

IPCC Fourth Assessment Report: Climate Change 2007 (The Physical Science Basis)


Change 2007 (The Physical Science Basis)

Recommended Reading:

Summary for Policymakers (SPM), from Working Group I Report "The Physical Science Basis“, 18 pages

www.ipcc.ch/pdf/assessment-report/ar4/wg1/ar4-wg1-spm.pdf

40

Solar and Terrestrial Radiation

Solar flux density or irradiance peaks at visible , near 0.48 m or 11,500 cm-1

falls off rapidly at IR known as shortwave radiation

Earth’s irradiance peaks at IR , near 10 m or 550 cm-1


emits no visible radiance

known as longwave radiation

Solar Radiation Spectrum

5800 K Blackbody

Solar radiation peaks in the visible

Outside the atmosphere

Sea Level


ultraviolet infraredVIS

Thermal (IR) Radiation Spectrum

This figure shows the infrared radiation emitted by the Earth and its atmosphere.

Note the underlying blackbody emission and the absorption due to gases.

CO2 is uniformly mixed in the


atmosphere, while H2O and O3 vary in space and time.

Atmospheric Windows

What happens if you add a more of a gas which absorbs in an atmospheric window? (e.g., the 3-5 m and 8-13 m regions)

The effect can be large because there is no “competition” for the photons!

Gases that absorb in the atmospheric window are


ppotent greenhouse gases (e.g., CH4, N2O, CFCs).

Adding more CO2 or anothergas that absorbs in the sameregion as CO2 or H2O has asmaller effect because thereis already strong absorptionthere.

Greenhouse Gases (GHGs) Water vapour (H2O)

most common greenhouse gas increases as surface temperature rises

Carbon dioxide (CO2) released by plant and animal life, decay, and burning of fuels removed by plant photosynthesis and absorption by the oceans

Methane (CH4) very effective at trapping heat - powerful greenhouse gas


very effective at trapping heat powerful greenhouse gas wetlands, rice paddies, animal digestion, fossil fuel extraction,

decaying garbage Nitrous oxide (N2O)

soils and the oceans, some from burning fossil fuels and fertilizer use Ozone (O3)

most ground level ozone is from chemical reactions involving pollutants

Halocarbons chemicals containing bromine, chlorine, or fluorine, and carbon extremely powerful greenhouse gases

Greenhouse Gas Absorption

infraredUV

visibleAbsorption of radiationbetween the top of the atmosphere and the surface

O2 and O3 (in the stratosphere) absorb most of the harmful UV radiation


Most of the visible radiation from the sun passes through the atmosphere and reaches the surface

The efficient GHGs are the ones that absorb in the “atmospheric window” (8-13 um). Gases that absorb in the already-saturated regions of the spectrum are not efficient GHGs.

window

41

Infrared Absorption by CO2


1820s: Fourier 1850s:

Tyndall

We have long known about the Greenhouse Effect


1890s:Arrhenius

We have long known about the Greenhouse Effect.

1930s: Callendar

Fourier, Arrhenius, and Callendar


See http://www.aip.org/history/climate/co2.htm for more history.

The “Keeling Curve”

Charles David Keeling began Charles David Keeling began collecting data at Mauna Loa, at collecting data at Mauna Loa, at 3400 m elevation, in March, 1958.3400 m elevation, in March, 1958.

The Mauna Loa atmospheric CO

January 31, 2011: January 31, 2011: 391 391 ppmvppmv, ~40% higher , ~40% higher than in midthan in mid--1800s1800s


http://scrippsco2.ucsd.edu/http://www.mlo.noaa.gov/programs/coop/scripps/co2/co2.html

CO2

Observed Fact: The CO2 (and CH4, N2O,…) content of the Earth’s atmosphere is steadily increasing due to the burning of fossil fuels which releases the trapped carbon into the atmosphere.

The Mauna Loa atmospheric CO2

concentration measurements constitute the longest, continuous record of atmospheric CO2

available in the world.

IPCC 2007: Global CO2 increased from a pre-industrial value of ~280 ppm to 379 ppm in 2005. Range over the last 650,000 years is 180 to 300 ppm (from ice cores).(a) Direct measurements of atmospheric CO2, and of O2from 1990.

The Carbon Dioxide Record


(b) CO2 in Antarctic ice cores, with Mauna Loa data for comparison.

(c) CO2 in the Taylor Dome Antarctic ice core.

(d) CO2 in the Vostok Antarctic ice core.

(e) and (f) Geochemically inferred CO2.IPCC 2001, Figure 3.

GHGs: CO2, CH4, N2O

IPCC 2007 Figure SPM.1

Atmospheric concentrations of carbon dioxide, methane and nitrous oxide over the last 10,000 years (large panels) and since 1750 (inset panels).

CO2

CH4


Measurements are shown from ice cores (symbols with different colours for different studies) and atmospheric samples (red lines).

The corresponding radiative forcings are shown on the right hand axes of the large panels.

N2O

42

Ozone and Water Vapour Ozone

Stratospheric ozone loss over the past two decades has caused a cooling of 0.15 ± 0.1 W/m2 of the surface-troposphere system

Tropospheric ozone has increased by ~35% since pre-industrial times due to emissions of ozone-forming chemicals (nitrogen oxides, carbon monoxide, and hydrocarbons

This has caused a warming of 0.35 [+0.25 to +0.65] W/m2 (IPCC


This has caused a warming of 0.35 [ 0.25 to 0.65] W/m (IPCC 2007)

Water Vapour Can be difficult to discern trends

Average atmospheric water vapour content has increased since at least the 1980s over land and ocean as well as in the upper troposphere

This increase is broadly consistent with the extra water vapour that warmer air can hold

Global Warming Potentials

Global warming potential (GWP) is a measure of the relative radiative effect of a given substance compared to CO2 over a chosen time period.

Definition (IPCC 1998):

GWP = time-integrated warming effect due to an


instantaneous release of unit mass (1 kg) of a given green-house gas in today’s atmosphere, relative to the effect of CO2.

This concept was created to enable decision makers to evaluate options to evaluate future emissions of various greenhouse gases without having to perform complex model calculations.

Greenhouse Gases: Abundances, Lifetimes, and GWPs

Chemical species Formula Abundance ppt Life-time 100-yr GWP

1998 1750 (yr)

Carbon Dioxide CO2 (ppm) 365 280 5-200 1

Methane CH4 (ppb) 1745 700 8.4/12 23

Nitrous oxide N2O (ppb) 314 270 120/114 296

Perfluoromethane CF4 80 40 >50000 5700

Perfluoroethane C2F6 3.0 0 10000 11900

Sulphur hexafluoride SF6 4.2 0 3200 22200

HFC-23 CHF3 14 0 260 12000

HFC-134a CF3CH2F 7.5 0 13.8 1300

Important greenhouse halocarbons under Montreal Protocol and its Amendments


Important greenhouse halocarbons under Montreal Protocol and its Amendments

CFC-11 CFCl3 268 0 45 4600

CFC-12 CF2Cl2 533 0 100 10600

CFC-13 CF3Cl 4 0 640 14000

Carbon tetrachloride CCl4 102 0 35 1800

Methyl chloroform CH3CCl3 69 0 4.8 140

HCFC-22 CHF2Cl 132 0 11.9 1700

HCFC-142b CH3CF2Cl 11 0 19 2400

Halon-1211 CF2ClBr 3.8 0 11 1300

Halon-1301 CF3Br 2.5 0 65 6900

Other chemically active gases directly or indirectly affecting radiative forcing

Tropospheric ozone O3 (DU) 34 25 0.01-0.05 -

Tropospheric NOx NO + NO2 5-999 ? <0.01-0.03 -

Carbon monoxide CO (ppb) 80 ? 0.08 - 0.25 -

Stratospheric water H2O (ppm) 3-6 3-5 1-6 -

Observed Changes in Temperature

IPCC 2007: The 100-year linear trend (1906–2005) is

0.74°C, and is larger than earlier values because of record

warmth over the previous decade.

IPPC 2001: Global average surface temperature increased by

0.6°C for 1901-2000.


IPCC 1995: 0.15°C from pre-industrial values

Most of the increase in global temperature has occurred in

two distinct periods: 1910 to 1945 and since 1976

(0.15°C/decade).

Recent warming is greater over land than over oceans.

NASA Top 10NASA Top 10201020102005200519981998200320032002200220092009

tied

NASA GISS Surface Temperature Analysis (GISTEMP)


2009200920062006200720072004200420012001

• 2010, tied for warmest year globally on record (announced January 12, 2011)• 19 countries set record highs in 2010; only one set a record low• 2001-2010 was the warmest decade in recorded history• 2010 was the wettest year globally since 1900

Courtesy of Tom Pedersen, PICS CMOS Lecture, 2011

http://data.giss.nasa.gov/gistemp/

Surface Temperature Trends – 1000 yrs


IPCC 2001

43

Surface Temperature Trends – 400,000 yrs

Series of ice ages

Reasons for temperature swings:

Milankovitch cycles: periodic changes in


periodic changes in the flux of solar radiation received by Earth driven by changes in Earth’s orbit, amplified by changes in greenhouse gases?

100 200 300 400

Thousand years BP (before present)

E. Wolff, British Antarctic Survey

Regional Surface T Trends - 100 yrs


The size of each circle reflects the size of the trend that it represents. Red = positive trend. Blue = negative trend.

IPCC 2001, Figure 2.9

Modelling Surface Temperature Trends


NCAR Ensemble SimulationsFrom presentation by Rick Anthes, UCAR Meeting, Boulder, CO, 7 October 2003

Modelling Surface Temperature Trends

IPCC 2007 Fig SPM.4 Black line = decadal averages of observations for 1906-2005 relative to 1901-1950.

Blue shaded bands = 5-95% range for 19


gsimulations from 5 climate models using only the natural forcings due to solar activity and volcanoes.

Red shaded bands = 5-95% range for 58 simulations from 14 climate models using both natural and anthropogenic forcings.

Berkeley Earth Surface Temperature (BEST) Analysis

The Berkeley Earth analysis shows “that the rise in average world land temperature globe is approximately 1.5 degrees C in the past 250 years, and about 0.9 degrees in the past 50 years”.

http://berkeleyearth.org/


“Berkeley Earth also has carefully studied issues raised by skeptics,

such as possible biases from urban heating, data selection, poor station

quality, and data adjustment. We have demonstrated that these do

not unduly bias the results.”

Other Evidence of Climate Change (IPCC 2007)

Observations since 1961 show that the average temperature of the global ocean has increased to depths of at least 3000 m.

Mountain glaciers and snow cover have declined.

Global average sea level rose at 1.8 mm/year from 1961 to 2003.

Average Arctic temperatures increased at almost twice the global average rate in the past 100 years.

Satellite data since 1978 show that annual average Arctic sea ice


Satellite data since 1978 show that annual average Arctic sea ice extent has shrunk by 2.7% / decade (7.4% / decade in summer).

Long-term trends from 1900 to 2005 have been observed in precipitation amount over many large regions.

More intense and longer droughts have been observed over wider areas since the 1970s, particularly in tropics and subtropics.

There is observational evidence for an increase of intense tropical cyclone activity in the North Atlantic since about 1970, correlated with increases of tropical sea surface temperatures.

44

Causes of Climate Change

Increasing atmospheric concentrations of greenhouse gases

Increasing atmospheric concentrations of aerosols (microscopic airborne particles or droplets)

Changes in land surface properties


Variations in solar activity

Radiative Forcing The radiative energy budget of the Earth’s climate system

can be perturbed by such factors as: changes in the concentrations of greenhouse gases and aerosols

changes in the solar irradiance incident upon the planet, or

other changes that affect the radiative energy absorbed by the surface (e.g., changes in surface reflection properties).

A change in the net radiative energy available to the global


A change in the net radiative energy available to the global Earth-atmosphere system due to changes in forcing agents is called a radiative forcing (W m-2).

Positive radiative forcings warm the Earth’s surface

Negative radiative forcings cool the Earth’s surface

Radiative forcing is thus a measure of the relative global average impact on climate of different natural and anthropogenic causes.

Clouds: absorb and emit infrared radiation increase planetary albedo

high-altitude cirrus clouds cause warming low-level stratus clouds cause cooling

The Role of Clouds and Aerosols

Aerosols: absorb and scatter solar and infrared radiation change cloud formation processes


Overall clouds and aerosols tend to cool the Earth systemComplex feedback effects could lead to warming or cooling depending on specific cloud and aerosol changes in response to changes in climate forcing.

change cloud formation processes absorption causes warming scattering (reflection) causes cooling changes to cloud properties causes cooling

What About the Sun?


Global Radiative Forcing of Climate 1750-2005


IPCC 2007 Figure SPM.2

The climate sensitivity parameter (not wavelength here!) is defined as the global mean surface temperature response Ts to the radiative forcing F:

Ts = F

In one-dimensional radiative-convective models, wherein

Climate Sensitivity Parameter


the concept was first initiated, is about 0.3 K / (Wm-2).

If we assume 2.5 W m-2 as the total radiative forcing from increases in greenhouse gases since 1885, we get

Ts = 0.8°C [ vs. observed change of 0.74°C ]

Simulations with GCMs (global circulation models) give from 0.3 to 1.4 K / (Wm-2).

45

Key Questions About the Climate System


IPCC 2001, TS Figure 1

Human Influence on Climate - 1IPCC 1995: “the balance of evidence suggests that there is a discernible human influence on global climate”

IPCC 2001 (TAR): “In the light of new evidence and taking into account the remaining uncertainties, most of the observed warming over the last 50 years is likely to have been due to the increase in greenhouse gas concentrations.”

IPCC 2007 (AR4): “Most of the observed increase in globally


IPCC 2007 (AR4): Most of the observed increase in globally averaged temperatures since the mid-20th century is very likely due to the observed increase in anthropogenic greenhouse gas concentrations. This is an advance since the TAR’s conclusion that “most of the observed warming over the last 50 years is likely to have been due to the increase in greenhouse gas concentrations”. Discernible human influences now extend to other aspects of climate, including ocean warming, continental-average temperatures, temperature extremes and wind patterns.”

Human Influence on Climate - 2

The attribution of climate change to anthropogenic causes involves statistical analysis and the careful assessment of multiple lines of evidence to demonstrate, within a pre-specified margin of error, that the observed changes are:

unlikely to be due entirely to internal variability;


unlikely to be due entirely to internal variability;

consistent with the estimated responses to the given combination of anthropogenic and natural forcing; and

not consistent with alternative, physically plausible explanations of recent climate change that exclude important elements of the given combination of forcings.

Remaining Uncertainties

Simulating and attributing observed temperature changes at smaller scales. On these scales, natural climate variability is relatively larger,

making it harder to distinguish changes expected due to external forcings.

Simulating changes in atmospheric circulation. Anthropogenic forcing is likely to have contributed to changes in


p g g y gwind patterns, affecting extra-tropical storm tracks and temperature patterns in both hemispheres.

However, the observed changes in the Northern Hemisphere circulation are larger than simulated in response to 20th century forcing change.

Reconstructions of solar and volcanic forcing, which are based on proxy or limited observational data for all but the last two decades.

Cloud feedbacks remain a large source of uncertainty.

Climate Models

Used to simulate climate response to different scenarios

Based on physical laws represented by equations that are solved using a three-dimensional grid over the globe

Major components of the climate are represented in sub-models (atmosphere, ocean, land surface, cryosphere, biosphere) along with processes between them


biosphere), along with processes between them

Typical resolutions are a few 100 km in the horizontal and 1 km in the vertical

Equations typically solved every half hour

For processes that take place on much smaller spatial scales than the model grid, their average effects are approximated by using physically based relationships with the larger-scale variables (parameterization).

Development of Climate Models


Somerville and Hassol, Physics Today, October 2011

46

Climate Processes and Feedbacks

Water Vapour: positive feedback doubles the warming due to CO2

Clouds: net cloud feedback is still uncertain, as clouds can both absorb and reflect solar radiation (cooling the surface) and absorb and emit long wave radiation (warming the surface)

Stratosphere: radiative and dynamical processes affect the climate

Oceans: transport heat; exchange heat, water, CO2 with atmosphere


2

Cryosphere: reductions of sea ice and snow give positive feedback; melting can affect ocean salinity and circulation

Land surface: climate changes influence the land surface (e.g., soil moisture, albedo, roughness, vegetation); tropical deforestation reduces evaporation and increases surface temperature

Carbon cycle: CO2 cycles rapidly among the atmosphere, oceans and land, but removal of increased CO2 due to human activities from the atmosphere takes far longer

Positive and Negative Feedbacks in the Climate System

Increase CO2

in atmospherePlanet warms Increased

evaporationMore H2O inatmosphere

Positive Feedback


Increase CO2

in atmospherePlanet warms Ice sheets melt Planetary albedo

is reduced

•These feedbacks amplify the initial warming•The strength of the feedbacks varying from model to model

Courtesy of Dylan Jones

Positive and Negative Feedbacks in the Climate System

Negative Feedback

Increase CO2

in atmospherePlanet warms Increased

evaporationMore H2O inatmosphere

More low


altitude clouds

Planet cools

Courtesy of Dylan Jones

Projections of the Earth’s Future ClimateEmissions Scenarios of Special Report on Emissions Scenarios (SRES)

A1 - rapid economic growth, global population peaks in mid-century and declines thereafter, and rapid introduction of new efficient technologies A1FI - fossil intensive (“worst case”) A1T - non-fossil energy sources A1B - a balance across all sources

A2 - very heterogeneous world, self-reliance and preservation of local id titi ti l i i l ti f t d i


identities, continuously increasing population, fragmented economic development

B1 - convergent world with same global population as in A1, but with rapid change toward service and information economy, with reductions in material intensity and introduction of clean resource-efficient technologies (“best case”)

B2 - emphasis on local solutions to economic, social and environmental sustainability, continuously increasing global population at a rate lower than A2, intermediate levels of economic development, and less rapid and more diverse technological change than in A1 and B1

Step 1 - Projected Emissions

CO2 N2O

Atmospheric Physics and Sustainable Aviation (K. Strong) Slide 275IPCC 2001 TS, Figure 17

CH4 SO2

Step 2 - Resulting Atmospheric Concentrations

CO2CH4


IPCC 2001 TS, Figure 18

N2O

47

Step 3 - Resulting Radiative Forcing


Estimated historical anthropogenic radiative forcing to 2000 followed by radiative forcing for the six SRES scenarios. The values are based on the radiative forcing for a doubling of CO2 from seven AOGCMs.

IPCC 2001, Figure 9.13a

4.0°C (A1F1)

Step 4 - Resulting Temperatures

Multi-model global averages of surface warming, relative to 1980-1999.

Shading indicates the


1.8°C

grange of individual model annual averages (±1 standard deviation).

IPCC 2007, Figure SPM.5

Projected Temperature Changes


IPCC 2007, Figure 10.8

Summary of Climate Model Predictions

Globally averaged surface T will increase by 1.8 to 4.0°C from 1980–1999 to 2090–2099.

For the next two decades, a warming of about 0.2°C per decade is projected. Even if concentrations of GHGs and aerosols were kept constant at

2000 levels, a warming of ~0.1°C per decade would be expected.


There are many other predictions in the IPCC report regarding: Snow cover, sea ice, ice sheets, extreme weather, frequency and

intensity of hurricanes, precipitation, circulation in the oceans

Canada and Climate Change As a northern nation, Canada is expected to experience a

greater degree of warming than countries closer to the equator.

Canada's average annual temperature has been increasing over the past century, and it may be more than 4°C warmer by the end of the next century (compared to the projected global average increase of about 2°C).


p j g g ) For comparison, average global temperatures during the

last ice age were only about 4 to 8°C lower than they are today.

See the Arctic Climate Impact Assessment (2004) http://www.acia.uaf.edu/

Model Simulations for Canada

North American temperatures for 2000-2100, relative to the 1981-2000


average, simulated by Environment Canada’s Canadian Global Climate Model (CGCM3)

http://www.ec.gc.ca/sc-cs/default.asp?lang=En&n=0EC06FB9-1#top

48

Concluding ThoughtsThis is a very complex issue. Scientifically

The case is still being made By the time it is absolutely conclusive, it may be too late

Politically Politics is a very messy business The media likes an argument often a disservice to the


The media likes an argument - often a disservice to the science

Economically and Ethically How well do we need to understand the issue before

acting? What will we give up for a benefit that we cannot see? Should we employ the “Precautionary Principle”?

(“Scientific uncertainty should not be a reason to postpone measures to prevent harm.”)


Part 4


Part 4

Aviation and the Atmosphere

Aviation Emissions: Products

Combustion of fuel by aircraft jet engines produces:

Carbon dioxide (CO2) ~ 70%

Water vapour (H2O) ~30%

Nitrogen oxides (NO + NO2 = NOx) <1%

Carbon monoxide (CO) <1%


Carbon monoxide (CO) <1%

Sulphur oxides (SOx) <1%

Hydrocarbons, also known as volatile organic compounds (VOCs) <1% Some VOCs and particulates are hazardous air pollutants

Particulate matter (PM) <1%

Aviation Emissions: Sources Carbon dioxide is produced by complete combustion of fuel

Carbon in fuel combines with oxygen in to produce CO2

Water vapour is produced by complete combustion Hydrogen in fuel combines with oxygen to produce H2O

Nitrogen oxides are produced when air passes through high temperature/high pressure combustion

Nit d i th i bi t f NO


Nitrogen and oxygen in the air combine to form NOx

Hydrocarbons (VOCs), CO, and particulates are produced by incomplete combustion

Sulphur oxides are produced when small quantities of sulphur in the fuel combine with oxygen during combustion

Ozone (O3) is not emitted directly but is produced by the reaction of VOCs and NOx in heat and sunlight. Ozone is the primary constituent of smog.

Comparative Shipping Emissions


http://uk.reuters.com/article/2009/02/18/airlines-carbon-asia-idUKSP19731020090218

Aviation Fuel Usage and CO2 EmissionsTop panel:

Aviation fuel use since1940

Growth in air passenger traffic since 1970 in billions of revenue passenger kilometres (RPK)

Annual change in RPK Arrows indicate events that

affected aviation use: 1970s oil


affected aviation use: 1970s oil crises, Gulf war, Asian financial crisis, World Trade Center attack, and SARS health crisis

Bottom panel:

Growth in annual CO2 emissions for all anthropogenic activities and from aviation

Fraction of total anthropogenic CO2 emissions due to aviation

Lee et al., Atmos. Environment, 2009

49

Aviation Emissions: Roles

Substances that modify the chemical balance in the stratosphere and affect stratospheric ozone e.g., NOx, particulates, water vapour

Radiatively active substances (greenhouse gases – GHGs) e.g., CO2, water vapour

Emissions of chemical species that produce or destroy


p p yradiatively active substances such as ozone or methane e.g., CO, NOx

Emission of substances that trigger condensation trails (contrails), line-shaped cirrus clouds (contrail-cirrus), and cirrus clouds induced by the addition of heterogeneous ice nuclei from aviation (soot-cirrus) e.g., water vapour, soot

the formation of contrails is only associated with aviation

Aviation Emissions: Broad Impacts

Reduced air quality around airports

Changes in ozone concentrations due to NOx and particulate emissions at high altitudes

Short-term, regional changes in atmospheric radiation budgets due to emission of gases with short lifetimes (e.g., NOx, SOx, H2O) and particulates, and contrail formation


, , 2 ) p , These emissions can lead to radiative forcing that is regionally

located near the flight routes

Long-term, global climate change due to emissions of CO2, which is a long-lived and well-mixed greenhouse gas The effects of CO2 emissions from aviation are thus

indistinguishable from those emitted by any other source

Bruce Anderson, NASA, talk at Green Aviation Workshop, April 2009http://event.arc.nasa.gov/Green-Aviation/home/pdf/Anderson_GA_talk_apr26.pdf

Aviation Emissions: Impacts


Aviation Emissions: Impacts


Aviation Emissions: Importance Aviation accounts for about 3% of the EU's total GHG

emissions, but has had an 87% increase in CO2 since 1990

The transport sector's GHG emissions are increasing while other sectors are working to stabilize or reduce emissions Emissions from international air transport within the EU are

increasing faster than any other mode due to cheap air travel

In the UK flights produce ~7% of carbon emissions but this is


In the UK, flights produce 7% of carbon emissions, but this is projected to increase to as much as 25% over the coming decades

Emissions from aircraft have a greater impact than those from sources on the ground Two to four times the global warming potential of

ground-level CO2

As well as CO2 emissions, condensation trails (contrails) and aerosols from aircraft exhaust can alter cloud properties and deplete stratospheric ozone

http://news.bbc.co.uk/2/hi/science/nature/6955009.stm

Numbers from the International Air Transport Association (IATA)


http://www.iata.org/pressroom/facts_figures/fact_sheets/Pages/environment.aspx

50

“Any growth in the aviation industry will undermine efforts to tackle climate change.”

What do you think?

Strongly disagree


Strongly disagree

Disagree

Agree

Strongly agree

See BBC News Q&A: Aviation emissions, August 2007http://news.bbc.co.uk/2/hi/science/nature/6955009.stm

The Aviation Cycle



Aviation Emissions: Altitude

Aircraft emissions near the ground are primarily considered to be local air quality pollutants About 10% of aircraft emissions, except hydrocarbons and CO,

occur during airport ground level operations, landing, and takeoff

Aircraft emissions at altitude are primarily considered to be greenhouse gases


About 90% of aircraft emissions occur at altitude

For hydrocarbons and CO, about 30% is emitted at ground level and 70% at higher altitudes

No other significant anthropogenic sources at high altitudes

Aircraft are the only major source of emissions in the Arctic

Water vapour emitted at altitude is a greenhouse gas, and may produce contrails that also have a greenhouse effect

http://avstop.com/aviation_emissions/What_emissions_come_from_aviation.htm

Aviation and the Atmosphere

Anthropogenic emissions affect the global atmosphere through vertical and horizontal transport

Aircraft emissions are unique because they occur at altitude Effects can be very different from same emissions at the surface

The effects of aircraft emissions depend on the flight altitude and whether aircraft fly in the troposphere or stratosphere


y p p p

Air travel has grown by 5% per year since 1977, and continued growth is expected

Aviation fuel is 2-3% of total global fossil fuel use, and ~13% of fossil fuel used in transportation (IPCC 1999)

2006: global commercial aircraft fleet flew 31.26 million flights, burned 188.20 million metric tons of fuel and covered 38.68 billion km. Average flight covered 1237.2 km in 2.06 hours and produced 4.2 kg/km of carbon from CO2 (Wilkerson et al., ACP, 2010)

The Importance of Transport The impact of local aircraft emissions depends on how they

are transported: do they dilute and spread? are they quickly removed? do they undergo chemical reactions?

Aircraft fly close to the tropopause

Subsonic aircraft emissions are predominantly removed in the troposphere

St d ti l t t t k l i th t i h


Strong upward vertical transport takes place in the tropics, where the tropopause is high (above 15 km) and aircraft do not fly into the stratosphere

Since exchange and removal processes are very slow in the stratosphere, it is very sensitive to pollution Emissions directly into the stratosphere (only at mid and high

latitudes) are transported downwards into the troposphere and do not seem to undergo significant upwards transport that would perturb the whole stratosphereATTICA - Transport Climate Assessment, http://www.pa.op.dlr.de/attica/

Transport Processes

Atmospheric Physics and Sustainable Aviation (K. Strong) Slide 300ATTICA - Transport Climate Assessment, http://www.pa.op.dlr.de/attica/

51

Aviation and Air Quality

Concerns about aircraft emissions were originally focused on local air quality near airports

The transition during the 1960s and 1970s from turbojet to high-bypass ratio turbofan engines, for powering subsonic aircraft, resulted in improved fuel efficiency, lower noise levels and substantially reduced emissions of CO and


hydrocarbons through improved combustion efficiencies

NOx levels increased because of the higher engine internal temperatures and pressures. However, this increase has been minimised by the incorporation of advanced combustor technology designed to reduce NOx emissions

NOx emissions can be a source of ozone smog and the related negative health impacts


Tropospheric Chemistry

Atmospheric Physics and Sustainable Aviation (K. Strong) Slide 302Uherek et al., Atmos. Environment, 2010

Aviation and Ozone

Roles of ozone: air pollutant, greenhouse gas, ultraviolet shield

NOx is an ozone precursor so aircraft emissions perturb ozone Aircraft source of NOx ~1-2% of the total emissions

The lifetime of NOx increases with altitude, so ozone perturbations depend on the altitude of NOx injection and vary from regional in scale in the troposphere to global in scale in the stratosphere

Subsonic aircraft fly in the upper troposphere and lower


Subsonic aircraft – fly in the upper troposphere and lower stratosphere – the UTLS (9-13 km)

Supersonic aircraft – fly at higher altitudes in the stratosphere (17-20 km)

NOx emissions in the UTLS should increase ozone and decrease methane (slightly offset by decreases in ozone due to sulphur and water emissions)

NOx emissions at higher altitudes should decrease stratospheric ozone

Ozone formation and photolysis in the stratosphere:

Ozone formation in the troposphere: O2 undergoes reactions in which CH4, CO or NMHCs react with OH to form HO2, then:

Tropospheric Ozone Chemistry – 1


4, 2,

This atomic oxygen results in formation of ozone by Rxn (2)

Ozone is destroyed by recombinations with OH or HO2:


Tropospheric Ozone Chemistry – 2

Atmospheric Physics and Sustainable Aviation (K. Strong) Slide 305Lee et al., Atmos. Environment, 2010

Aviation and Tropospheric Ozone NOx concentrations increase by 30-40% in the main flight

corridors This enhances (4), decreases HOx, and increases ozone production

via (5) & (2)

The shift in the HOx partitioning to OH from (4) also reduces ozone loss via (6)

Both effects result in an ozone increase of ~3% in the NH troposphere


There are four regimes (R) for aviation-induced tropospheric ozone changes: RNOy, where ozone production is controlled by increased NOx and (4)

RHO2, where ozone loss is reduced due to a decrease in HO2 via (6)

RO3, where ozone loss is increased due to an increase in ozone concentrations via (5) and (6)

RPO3 close to the Earth’s surface, where the additional HO2 reacts with NO from other sources to form HNO3, which is easily washed out, reducing ozone production


52


Subsonic Aviation and Ozone Aircraft emissions of NOx are more effective at producing

ozone in the upper troposphere than an equivalent amount of emission at the surface (production of ozone per NO molecule is higher)

The radiative forcing effect of additional ozone molecules is larger at higher altitudes so ozone increases in the upper troposphere are more effective at than increases at surface


troposphere are more effective at than increases at surface

Due to these increases the calculated total ozone column in northern mid-latitudes is projected to grow by approximately 0.4 and 1.2% in 1992 and 2050, respectively

However, aircraft particulates, sulphate aerosols, and water tend to deplete ozone, partially offsetting the NOx-induced ozone increases

The largest increases in ozone concentration due to aircraft emissions are calculated to occur near the tropopause IPCC 1999

Aviation and Stratospheric Ozone Impact of subsonic NOx emissions

Current subsonic aviation emits NOx directly into the stratosphere only at mid and high latitudes, where the general circulation is directed downwards

Model studies suggest that subsonic aircraft emissions have a limited impact on stratospheric NOx and ozone, with ozone reductions below 0.05%

I f i i iNH total ozone column change as a function of EINOx in 2015 for a


Impact of supersonic emissions Models suggest a reduction in the

NH ozone column of 3.2% in 2015 for a NOx emission index of 15

Emissions can also lead to more polar stratospheric clouds, greater dehydration and denitrification, and enhanced heterogeneous ozone depletion during Antarctic spring


a function of EINOx in 2015 for a supersonic fleet of 500 HS

Supersonic Aviation and Ozone Early models suggested that NOx released by supersonic aircraft

would deplete stratospheric ozone substantially Assuming 500 aircraft cruising at 20 km altitude, predicted a reduction in

the global average ozone column of ~12% with a worst-case reduction of 25% near the flight corridor (CIAP, 1975)

With better understanding of ozone chemistry and more stringent NOx emissions, models now suggest that the response of ozone to NOx emissions by supersonic aircraft is negative (decrease) in


to NOx emissions by supersonic aircraft is negative (decrease) in the middle stratosphere and positive (increase) below Assuming 500 aircraft cruising at 20 km altitude, perturbation of the ozone

column could be slightly negative or positive (depending on latitude and season) but should not exceed a few percent

Impact on UV-B levels at surface should be less than 1-2%

The change in the occurrence frequency of PSCs resulting from the release in the polar stratosphere of water and NOx by future high-altitude aircraft could be a factor of 2 increase, but the corresponding impact on polar ozone remains uncertain


Summary of Ozone Processes


Climate Impacts of Transport Sector


Comparison of the climate impact on a 100 years time scale expressed in CO2

equivalent emissions [Tg CO2/yr] based on 100 year Global Warming Potentials (GWP100). Land transport is shown in red, aviation in blue and marine shipping in green. For NOx from aviation and for shipping values high and low impact estimations are shown.

ATTICA - Transport Climate Assessment, http://www.pa.op.dlr.de/attica/

53

Climate Effect of Emissions: Cause and Effect Chain

Atmospheric Physics and Sustainable Aviation (K. Strong) Slide 313Fuglestvedt et al., Atmos. Environment, 2010

Climate Impact of Subsonic Aviation Emissions of CO2 result in a positive RF (warming) Emissions of NOx result in the formation of tropospheric ozone

via atmospheric chemistry, with a positive RF (warming) Emissions of NOx destroy methane, with a negative RF

(cooling), accompanied by loss of tropospheric ozone Emissions of sulphate particles arising from sulphur in the fuel

lt i ti (di t) RF ( li )


result in a negative (direct) RF (cooling) Emissions of soot particles result in a (direct) positive RF

(warming) The formation of persistent linear contrails that may form in the

wake of an aircraft result in both positive and negative RF effects but overall, cause a positive RF effect (warming)

The formation of contrail-cirrus cloud results in both positive and negative RF effects but overall, is considered to cause a positive RF effect (warming) Lee et al., Atmos. Environment, 2010

Climate Impact of Supersonic Aviation Emissions of CO2 result in a positive RF (warming) Emissions of NOx result in destruction or increase of stratospheric ozone:

above ~20 km, models always predict ozone destruction, with a net ozone column decrease and a global negative RF (cooling)

Emissions of H2O result in a positive RF (warming): this is by far the dominant RF term for supersonic aircraft; an indirect (small) negative correction is obtained by additional HOx formation and ozone destruction

Emissions of NOx result in the destruction of methane with a negative RF


(cooling), although this is negligible in the stratosphere Emissions of sulphate particles result in negative RF (cooling): an indirect

effect occurs via heterogeneous chemistry whereby more NOx is lost to HNO3 and the ozone destruction/increase is reduced

In addition, the lower stratospheric ozone increase due to NOx emissions may change into destruction due to increasing Clx and Brx (since chlorine and bromine nitrates will decrease): the net effect on ozone is dependent upon altitude and transport

Emissions of soot result in positive RF (warming) The formation of persistent linear results in both positive and negative RF

effects but overall, causes a positive RF effect (warming) The formation of contrail-cirrus cloud causes a positive RF (warming)


Aviation and Climate – CO2

Current concentrations of CO2 and the associated radiative forcing are the result of emissions over the past century

The total amount of CO2 attributable to aviation is ~1 ppmv (1992 atmosphere), which is ~1% of the total anthropogenic increase

This percentage is lower than the percentage for emissions


p g p g(2-3%) because aircraft emissions have occurred in the last half-century

The accumulation of CO2 due to aviation over the next 50 years is projected to increase to 5-13 ppmv

IPCC 1999

Aviation’s Contribution to Canada’s Total GHG Emissions


Canada’s Action Plan to Reduce Greenhouse Gas Emissions from Aviation, Transport Canada, 2012, http://www.tc.gc.ca/eng/policy/acs-reduce-greenhouse-gas-aviation-menu-3007.htm

2006 Spatial Distribution of CO2Carbon Emissions from Aviation – 1

Carbon from CO2

(2-3% of total anthropogenic

CO2-C emissions)

(a) Column emissions per unit area showing

Atmospheric Physics and Sustainable Aviation (K. Strong) Slide 318Wilkerson et al., ACP, 2010

CO2-C = carbon from CO2 obtained by molecular weight ratio: MC/MCO2

area, showing regional effects

(b) Variation with latitude (peak in NH at 40N)

(c) Variation with longitude (peaks for US, Europe, East Asia)

54

2006 Spatial Distribution of CO2Carbon Emissions from Aviation – 2

Dominant flight corridors and emissions are at northern mid-latitudes (US, Europe, East Asia).


CO2-C = carbon from CO2 obtained by molecular weight ratio: MC/MCO2

Circulation patterns

prevent much exchange between NH and SH

spread emissions quickly across the northern mid-latitudes

Wilkerson et al., ACP, 2010

2006 Vertical Distribution of CO2-Carbon Emissions from Aviation

Atmospheric Physics and Sustainable Aviation (K. Strong) Slide 320Wilkerson et al., ACP, 2010

Emissions above 7 km are associated with cruising; below 7 km are associated with short flights, take-off, and landing

Peak emissions are at 10-12 km (UTLS) where contrails form Double peak is due to E-W and N-S flight clearance altitudes

74.6% of aviation fuel was burned above 7 km in 2006

Aviation and Climate – NOx, O3, CH4

In addition to increasing tropospheric ozone concentrations, aircraft NOx emissions decrease the concentrations of methane, which is also a greenhouse gas Reductions in methane cool the surface of the Earth

The methane concentration in 1992 is estimated to be about 2% less than that in an atmosphere without aircraft


Much smaller than the 2.5-fold increase since pre-industrial times

Changes in tropospheric ozone are mainly in the Northern Hemisphere, while those of methane are global in extent so that, even though the global average radiative forcings are of similar magnitude and opposite in sign, the latitudinal structure of the forcing is different so that the net regional radiative effects do not cancel

IPCC 1999

Aviation and Climate – H2O

Most subsonic aircraft water vapour emissions are released in the troposphere where they are rapidly removed by precipitation within 1 to 2 weeks

A smaller fraction of water vapour emissions is released in the lower stratosphere where it can build up to larger concentrations.


Because water vapour is a greenhouse gas, these increases tend to warm the Earth’s surface, though for subsonic aircraft this effect is smaller than those of other aircraft emissions such as CO2 and NOx

IPCC 1999

Aviation and Climate – Contrails Contrails cover ~0.1% of Earth’s surface on an annually

averaged basis with larger regional values (1992)

They form when water vapour and particles are released into cold ice-supersaturated air Their optical properties depend on the particles emitted or formed

in the aircraft plume and on atmospheric conditions

Contrails tend to warm the Earth’s surface similar to thin


Contrails tend to warm the Earth s surface, similar to thin high clouds The radiative effect of contrails depends on their optical properties

and global cover, both of which are uncertain

The contrail cover is projected to grow to 0.5% by 2050, faster than the rate of growth in aviation fuel consumption This faster growth in contrail cover is expected because air traffic

will increase mainly in the upper troposphere where contrails form preferentially, and may also occur as a result of improvements in aircraft fuel efficiency IPCC 1999

Aviation and Climate – Cirrus Clouds

Extensive cirrus clouds have been observed to develop after the formation of persistent contrails

Increases in cirrus cloud cover are positively correlated with aircraft emissions in a limited number of studies

About 30% of the Earth is covered with cirrus cloud

On average an increase in cirrus cloud cover tends to warm


On average an increase in cirrus cloud cover tends to warm the surface of the Earth (trapping of outgoing terrestrial radiation dominates reflection of solar)

An estimate for aircraft-induced cirrus cover for the late 1990s ranges from 0 to 0.2% of the surface of the Earth This may increase by a factor of 4 (0 to 0.8%) by 2050

IPCC 1999

Cirrus clouds and contrails; Source: Karlsruher Wolkenatlas © Bernhard Mührhttp://www.pa.op.dlr.de/attica/

55

Aviation and Climate – SOx and Soot

The aerosol mass concentrations in 1992 resulting from aircraft are small relative to those from surface sources

Although aerosol accumulation will grow with aviation fuel use, aerosol mass concentrations from aircraft in 2050 are projected to remain small compared to surface sources

Increases in soot tend to warm while increases in sulphate


ptend to cool the Earth’s surface

The direct radiative forcing of sulphate and soot aerosols from aircraft is small compared to those of other aircraft emissions

Because aerosols influence the formation of clouds, the accumulation of aerosols from aircraft may play a role in enhanced cloud formation and change the radiative properties of clouds

IPCC 1999

2006 NOx, SOx & Black Carbon Emissions


Column-integrated aircraft emissions of NOx (as NO2), sulphur from SO2, and black carbon show similar distributions to CO2-C

Wilkerson et al., ACP, 2010

Indirect RF Indirect climate effects result from

the fraction of aerosol (soot, sulphate, mineral dust) which leads to contrail and cloud formation

Figure shows modelled radiative forcing of all anthropogenic


forcing of all anthropogenic aerosols at the top of the atmosphere compared to the impact of aircraft soot Small changes in model input can

give slightly positive forcings for soot

The sign and magnitude of the effects are not yet well resolved at the moment.

Aviation Emissions & Radiative Forcing



Atmospheric Physics and Sustainable Aviation (K. Strong) Slide 329Lee et al., Atmos. Environment, 2009 Atmospheric Physics and Sustainable Aviation (K. Strong) Slide 330Lee et al., Atmos. Environment, 2009

56

Radiative Forcing


Projections for Aviation Radiative Forcing


Estimates of the globally and annually averaged total radiative forcing (without cirrus clouds) associated with aviation emissions under each of six scenarios for the growth of aviation over the time period 1990 to 2050.

IPCC 1999

Projections for Aviation Radiative Forcing


“The increased air traffic projected by 2020 and 2050 results in increased emissions and RF, despite the assumed technological improvements in the fleet.”

Aviation and Climate: Summary

Present day radiative forcing from aviation (2005) is estimated to be 55 mW m2 excluding cirrus cloud enhancement, which represents

3.5% of current anthropogenic forcing

78 mW m2 including cirrus cloud enhancement, which represents 4.9% of current forcing

Future radiative forcings may increase by factors of 3 4 over


Future radiative forcings may increase by factors of 3-4 over 2000 levels, in 2050.

The effects of aviation emissions of CO2 on global mean surface temperature last for many hundreds of years (like other sources), while non-CO2 aviation effects on temperature last for decades


Future Climate Impact of Aviation

Grewe & Stenke (ACP, 2008) developed a model “AirClim” to predict the climate impact of aircraft emissions for several aviation scenarios over varying time horizons.

AirClim calculates the global near-surface temperature change caused by CO2, NOx,water vapour, and contrails

The model includes atmospheric processes linking


p p gemissions to radiative forcing, and temperature change Accounts for the amount and location of the emissions

Considers uncertainty ranges for lifetimes and RF calculations

The model was used to assess the impact of subsonic aviation for 2000, 2100, and 2250.

Climate Impact – 1Impact of subsonic air traffic on (a) radiative forcing(b) near-surface temperature.

Red – due to CO2 emissions

Solid blue – due to impact of NOx emissions on O3


Hashed blue – due to impact of NOx emissions on O3 & CH4

Solid and hashed green –same as blue but for reduction of the NOx emission index by 40% from 2000 to 2050.(Note: 2250 is only taken into account to represent steady-state for non-CO2

greenhouse gases)Grewe & Stenke, ACP, 2008

57

For all years, RF from CO2 is larger than RF of NOx emission products

In 2000, T due to NOx emission products is larger than for CO2

This difference in the importance of CO2 and NOx emissions on RF and T is due larger efficacy of ozone (1.4) compared to CO2 (1)

In 2100, T due to CO2 is larger than

Climate Impact – 2


NOx emission products only for the 40% reduction of the NOx emission index between 2000 and 2050

This reduction will lower the importance of NOx emissions

Shows the different temporal evolution of radiative forcing and near surface temperature changes.

In 2250, the climate impact of CO2

emissions dominates that of NOxGrewe & Stenke, ACP, 2008

Growth in Air Travel Since 1977


http://www.boeing.com/commercial/cmo/pdf/2010_Farnborough_Presentation.pdf(presentation by Randy Tinseth, July 2010)

Growth in Air Travel

Passenger numbers are forecast to expand to 3.1 billion in 2013


Air freight volumes are forecast to increase slightly in 2013

http://www.iata.org/pressroom/facts_figures/Documents/economic-outlook-media-day-dec2012.pdf(presentation by Brian Pierce, IATA Global Media Day, December 2012)

Forecast Growth of Passenger Traffic

Between 2010 and 2014, Transport Canada forecasts 3.1% growth in domestic traffic


Canada’s Action Plan to Reduce Greenhouse Gas Emissions from Aviation, Transport Canada, 2012, http://www.tc.gc.ca/eng/policy/acs-reduce-greenhouse-gas-aviation-menu-3007.htm

Boeing’s Projected Growth in Air Travel


http://www.boeing.com/commercial/cmo/pdf/2010_Farnborough_Presentation.pdf(presentation by Randy Tinseth, July 2010)

RPK = revenue passenger kilometersProjected to increase by a factor of 6 between 1990 and 2029

Aviation Emissions & Fuel EfficiencyNote:

70% reduction in fuel per passenger kilometres since 1960s, significant


even since 1990

Saving of >4 billion tonnes of CO2 if technology were frozen in 1990

http://www.iata.org/pressroom/facts_figures/Documents/environment-media-day-dec2012.pdf(presentation by Paul Steele, IATA Global Media Day, December 2012)

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Projected Aviation Fuel Emissions


Aviation CO2 Emission Scenarios

Time series of aviation CO2 emission scenarios to 2100, based on IPCC/SRES (Special Report on Emissions Scenarios)

Atmospheric Physics and Sustainable Aviation (K. Strong) Slide 344Owen et al., Environ. Sci. Technol., 2010

Aviation CO2

Emission ScenariosTop panel: 2000 CO2

aviation emissions

Bottom panels: ratio of SRES 2050 emissions projections to 2000 emissions


In 2000, aviation contributes 677 Tg of CO2 = 12% of transport total

In 2100, aviation contributes 723-5067 Tg of CO2 = 7-24% of transport total

Owen et al., Environ. Sci. Technol., 2010

Target Aviation Emissions for CO2

In 2009, the aviation industry agreed to:

targets of 1.5% fuel efficiency to 2010


carbon neutral growth from 2020

50% reduction in net emissions by 2050


Target Aviation Emissions for CO2

Red line = no action

Green line = emissions needed to reach 2050 goal.

Shaded sections = contribution of “four pillars” to reaching target:


target:

Investment in technology

Better operations

Improved infrastructure

Economic measures


Key Science Questions1. The atmospheric distribution and the global budget of chemical

compounds which affect production and destruction of ozone

2. The chemical and photochemical processes that determine the partitioning between reactive species in the lower stratosphere and upper atmosphere, including heterogeneous mechanisms

3. The dynamical processes that determine the transport at diff t l f h i l d i l di


different scales of chemical compounds, including cross-tropopause exchanges and convective motions

4. The impact of soot and acids on particle formation and related changes in cloudiness and chemistry

5. Evaluations of the climatic impact of the changes in composition and cloudiness relative to the impact of CO2

6. The precise composition of aircraft exhaust as regard to particle formation processes


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Mitigation Options for Aviation

Ambitious technology targets for fuel burn and NOx emissions – requires technological breakthroughs

Reduction in contrail formation – requires reductions in the emission index of H2O, increase in the specific heat content of the fuel, or a decrease in propulsion efficiency Alternate option is operational avoidance – requires avoidance of


ice-supersaturated air

Reduction in NOx emissions – e.g., by manipulation of the combustor conditions or the cooling air Alternate option is varying cruise altitudes

Alternative fuels – e.g., biofuels, LH2 in cryoplanes Usage of biofuels in other sectors may yield more benefits than

utilization in the transport sector




The End!

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