Download O 3

Chapter 8
The atmospheric environment
The
atmosphere
Little mixing
between the
troposphere and
stratosphere
ozone
Increasing temp
here because
of proximity to
O3 layer
Troposphere
Well-mixed
Figure 8-1. The U.S. Standard Atmosphere, 1976. Note the various temperature
reversals, which act as thermal lids on the lower parts of the atmosphere. In the
troposphere, gases are well mixed. From Neiburger et al. (1982).
Atmospheric gas composition
Nitrogen, oxygen, argon, neon, xenon constant on 1000 yr timescale
Oxygen varies on geologic timescales
Carbon dioxide, nitrous oxide, methane increasing
Near surface ozone increasing
Stratospheric ozone (ozone layer) decreasing
Very upper atmosphere has gradients in each gas due to gravity effect on the
different molecular weights
Solar Radiation and Atmospheric Heating
Solar constant – measure of the amount of energy passing through a unit
surface area perpendicular to the direction of the solar radiation
Albedo – amount of energy reflected back to space from the surface and the
Atmosphere
Insolation – amount of energy reaching the earth’s surface
Incoming solar radiation is short-wavelength
Outgoing solar radiation is long-wavelength
Any gas with multiple bonds will absorb some long-wavelength radiation and
turn it into heat…..greenhouse gas
Water vapor, carbon dioxide, nitrous oxide, methane, CFCs
Venus…
It is estimated that the surface temperature on Venus would actually be below
0°F without the Greenhouse effect. However, because of the greenhouse
effect, the average surface temperature is 467°C or 872°F.
Atmospheric Composition at
Surface Level
Major Components (by volume)
CO2
96.5%
N2
3.5%
Minor Component (ppm)
SO2
150
Ar
70
H2O
20
CO
17
He
12
Ne
7
Perfect Radiator: any substance that emits the maximum
amount of electromagnetic energy at all wavelengths.
Total amount of energy emitted is a function of temperature
and described by the Stefan-Boltzmann law:
E = sT4
The wavelength of the maximum emitted energy varies
inversely with the temperature and is described by the Wien
displacement law:
lM = aT-1
Short wavelength radiation: general term for radiation
coming from the sun.
Long wavelength radiation: general term for radiation
coming from the earth.
There is a latitudinal disequilibrium of heat on the planet
Figure 8-4. Incoming (shortwave) and outgoing (longwave) radiation as a
function of latitude. The crossover occurs at ~40o. At lower latitudes there is a heat
excess, at higher latitudes a heat deficit.
Yet the heat flux of the planet is just about in steady state. Heat redistributed due to
atmospheric (and oceanic) circulation
Global Redistribution of Heat through circulation
without the spin of the Earth
Coriolis Effect
Fc = (2Wsinf)v
Where W is the angular
velocity of the earth’s rotation
in radians (7.29 x 10-5 rad s-1),
f is the latitude, and v is the
velocity of the moving mass.
Note that at the equator the
Coriolis force would be zero,
and at the poles the Coriolis
force would be at its
maximum value.
Global Atmospheric Circulation with the spin of the Earth
Polar Cell
Ferrel Cell
Ferrel Cell
Reykjavik, Iceland 64°8’
Boston, MA latitude 42°23’ N
Death Valley, CA 36°34’N
George town, Bahamas 23°51’ N
Khartoum, Sudan 15°62’N
Macapá, Brazil 0°02’N
Wellington, New Zealand 41°26’S
Capitán Arturo Prat, Antarctica 62°33’
Hydrostatic Equation: Dp = -rgDh
where Dp is the change in pressure, r is the density of the fluid, g is the
acceleration due to gravity, and Dh is the change in height.
Dry adiabatic lapse rate: the rate at which an air
parcel cools if lifted in the atmosphere or warms if
forced to lower levels, as long as no condensation
occurs in the air parcel. (= ~9.8 K km-1.)
Absolute humidity: the amount of water vapor
actually present in the air.
Relative humidity: the amount of water vapor in
the air divided by the amount of water vapor the
air can hold at any particular temperature,
expressed in percent.
Wet adiabatic lapse rate: the rate at which an air
parcel cools when condensation occurs. It is a
function of temperature and pressure.
Environmental lapse rate: the observed rate at
which temperature changes in a column of air.
Inversion: the reversal of the normal temperature pattern
Radiation inversion: caused by
radiational cooling of the land surface
and a decrease in the temperature of
the atmosphere at low levels.
Subtropical inversion: caused by
sinking air at high pressure center.
*Remember, when air descends its
temperature increases.
Frontal inversion: caused by the
relative movement of warm air over
cold air.
Air Pollution
Primary pollutants – direct products of combustion or evaporation
VOCs, CO, CO2, SOx, NOx
Secondary pollutants – products of atmospheric reactions involving primary pollutants
Ozone, compounds produced by photochemical oxidation (PANs: Peroxyacytyl
nitrate) VOCs and NOx’s are important reactants in forming secondary pollutants
Ta ble 8-2. Classification of air pollutants
Major Class
Sub class
Inorganic gases
O xides of nitrogen
N 2 O, NO, N O 2
O xides of sulfur
SO 2 , SO 3
O xides of carbon
CO, CO 2
Other ino rganics
O 3 , H 2S, HF, N H 3, Cl 2 , Rn
Hydro carbons
M ethane ( CH 4), butane (C4 H 10 ), octane (C8 H 18),
benzene (C6 H 6), acetylene (C2 H 2), ethylene (C2 H 4)
Aldehydes and ketones
For maldehyde, acetone
Other organ ics
Chlorofluo rocarbons, PA Hs, alcohols, organic acids
Solids
Fu me, dust, smoke, ash, carbon soot, lead, asbestos
Liquids
M ist, spray, oil, grease, acids
Organic gases
Particulates
Examp les
Aerosols
Solid particles or liquid droplets ranging in size up to 20um in radius. Can
either be put into the air directly or created in atm
SO2(g) + H2O  H2SO4(l)
Sulfuric acid aerosol
Ta ble 8-3. Sour ces of aerosols an d c ontri butions of natural versus anthr opog enic
sourc es*
Natural
(1 0 12 g y -1)
An throp ogenic
(1 0 12 g y -1)
Soil and rock dust
3000 - 4000
?
Se a salt
1700 - 4700
Sou rc e
Biogenic
100 - 500
Bio mass burning (soot)
6 - 11
Volca nic
15 - 90
Dire ct e missions - fuel, inc inerators, industry
36 - 154
15 - 90
Gaseous e missions
Su lfa te fro m bioge nic DMS
Su lfa te fro m volc anic SO 2
51
18 - 27
Su lfa te fro m fossil fuel
105
Nitrate fro m NO x
62
128
A mmoniu m fro m NH 3
28
37
Biogenic hydroca rbons
20 - 150
Anthropogenic hydroca rbo ns
Tota l
100
5000 - 9619
421 - 614
* Modified from Be rne r and Be rner (1996)
Factoid: Following large volcanic eruptions, the sulfuric acid aerosol in the atm
increases the earth’s albedo leading to temporary global cooling
Ta ble 8-3. Sources of aerosols an d contri butions of natural versus anthropog enic
sources*
Natural
(1 0 12 g y -1)
An throp ogenic
(1 0 12 g y -1)
Soil and rock dust
3000 - 4000
?
Sea salt
1700 - 4700
Biogenic
100 - 500
Sou rce
Bio mass burning (soot)
6 - 11
Volcanic
15 - 90
Direct emissions - fuel, incinerators, industry
36 - 154
15 - 90
Gaseous emissions
Su lfate fro m biogenic DMS
Su lfate fro m volcanic SO 2
51
18 - 27
Su lfate fro m fossil fuel
105
Nitrate fro m NO x
62
128
A mmoniu m fro m NH 3
28
37
Biogenic hydrocarbons
20 - 150
Anthropogenic hydrocarbo ns
Total
* Modified from Berner and Berner (1996)
100
5000 - 9619
421 - 614
Smog
Smoke + fog
More prevalent during inversions
Photochemical smogs – maximum during midday
Ta ble 8-4. Types of s mogs an d their characteristics
Characteristic
Industrial
Photoch emical
First occurrence
Londo n
Los Angeles
Principal pollutants
SO x
O 3 , NO x, H C, CO, free radicals
Principal sou rces
Industrial and h ousehold fuel co mbustion
M otor vehicle fuel comb ustion
Effect on hu mans
Lung and throat irritatio n
Eye and respiratory irritation
Effect on co mpounds
Reducing
O xidi zing
Time of worst events
Winter mo nths in the early morning
Su mmer months aroun d midday
Industrial smog – mostly acid aerosols, corrode buildings and retinas
Photochemical smogs – mostly formation of ozone, NOx, and PAN, respiratory
distress
The Great London Smog of 1952
Greenhouse Gases: CO2, CH4, N2O & CFCs that absorb
long-wave, out-going radiation causing them to vibrate
and generate heat.
Positive and negative feedbacks of greenhouse effect and
climate change
Hotter world = more water vapor = more heat trapping = positive feedback
Hotter world = more cloud cover = less insolation = negative feedback
Hotter world = less snow = less albedo = more net insolation = postive feedback
Hotter world = faster decomp = more CO2 = more heat trap = positive feedback
More CO2 + temp = more photosynthesis = less CO2 = negative feedback
Geologic record indicates periods of warm wet earth and cold dry earth
The relative effect of an individual gas on the greenhouse
effect over time depends upon:
Molecular scale radiative forcing (how well it converts radiation to heat)
Atmospheric concentration
Rate of increase in the atmosphere
Atmospheric residence time
Ta ble 8-5. Data for greenhouse gases*
Gas
Concen tration
1990 (ppmv)
CO 2
35 4
Positive
radiative
forcing
(W m-2)
% total
radiative
forcing
Relative
instantan eous
Lifeti me radiative forcing
(y)
(molecular basis)
Global
W arming
Potential
(100 y)
1.5
61
50 - 200
1
1
12
43
21
310
CH 4
1.72
0.42
17
H 2 O strat
-
0.14
6
N2 O
0.310
0.1
4
120
250
CF C-11
0.00028
0.062
2.5
65
15,000
3,400
CF C-12
0.00048 4
0.14
6
130
19,000
7,100
Other CFCs
0.085
3.5
Total
2.45
100.0
CF C substitutes
HFC-23
264
650
HFC-152a
1.5
140
50,000
6,500
3,200
7,400
CF 4
C6 F1 4
*Fro m Berner and Berner (1996), IPCC (1996), van Loon and Duffy (2000)
CO2
Imbalance = (fossil fuel) – (increase in CO2) – (ocean storage) + (deforestation)
The ‘missing’ carbon sink. The global CO2 budget is out of balance. We would
predict more increase in atm CO2 than what is actually observed.
Figure 8-8. Mean monthly concentrations of CO2 at Mauna Loa, Hawaii. From Berner and Berner (1996).
Read Case Study 8-2
Global C box model
Net reaction for atm CO2 uptake in ocean
CO2 + CO32-  H2O + 2HCO3-
Fast
slow
Figure 8-9. The carbon cycle. Reservoir concentrations are in 1015 g (Gt) carbon. Fluxes are in
Gt C y-1. From Berner and Berner (1996).
Methane
2nd most important greenhouse gas
Sinks for methane:
1) chemical oxidation in the troposphere
2) stratospheric oxidation
3) microbe uptake
Figure 8-10. Variation in methane abundance from 1841 to 1996. The fitted curve is a sixth-order polynomial. Data from
Etheridge et al. (1994) and IPCC (1966).
Methane
Ta ble 8-7. Sourc es an d sinks for at mos phe ric methane*
Source or Sink
CH 4 (10 1 2 g C y -1)
% tota l
Sources
Natu ra l
How could 14C help tell us
The source of atm methane?
Wetla nds
86
22.5
Te rmites
15
3.9
Ocea ns
8
2.1
La kes
4
1.0
Me thane hydrates
4
1.0
117
30.5
Ene rgy production/use
69
18.0
Ente ric fer menta tion
63
16.4
Rice
45
11.8
Ani mal wastes
20
5.2
La ndfills
29
7.6
Bio mass burning
21
5.5
Do mestic sewa ge
19
5.0
266
69.5
Total natural
Anthropoge nic
Total anthropoge nic
Total fo r sou rc es
383
Sin ks
At mosphe ric re moval
353
88.2
Re moval by soils
23
5.8
At mosphe re
24
6.0
Total fo r sinks
400
* Data fro m Berne r and Be rne r (1996), IP CC (1 992)
Nitrous Oxide…its no laughing matter
N 2O
Sources: denitrification, nitrification, biomass burning & fertilizer production
No N2O sinks in the troposphere
The third largest contributor to global warming behind CO2 and CH4.
Also responsible for stratospheric ozone destruction.
Climate Change and the Geologic Record
Ice cores
Stable isotopes
Direct gas measurement of bubbles
Paleotemp from
isotopes
Figure 8-11. Variation in temperature, CO2, and CH4
concentrations in Antarctica during the past 240,000 years.
From Lorius et al. (1993).
Climate Change and the Geologic Record
Sediment record
Stable isotopes (oxygen and/ carbon)
Oxygen isotopes in carbonates
from a sediment core in the Western
Pacific
Several glacial / interglacial periods
Figure 8-12. Surface temperature of the Pacific Ocean based on
oxygen isotope ratios. From THE BLUE PLANET, 2nd Edition by
B. J. Skinner, S. C. Porter and D. B. Botkin. Copyright © 1999.
This material is used by permission of John Wiley & Sons, Inc.
Ozone
Good ozone – stratosphere
Bad ozone – troposphere
Ozone production requires energy from photons
O2 + hv  O* + O*
O* + O2 + M  O3
(M is a catalyst …e.g. N; DHR0 = -106.5 exothermic)
Net reaction …3O2 + hv  2O3
Ozone destruction also involves photons
O3 + hv O2 + O*
O* + O3 O2 + O2
Why good ozone is good.
Figure 8-13. Absorption cross sections for oxygen and ozone in the 100 to 300 nm
wavelengths. Also shown is the solar flux density and the wavelengths of biologically
harmful radiation (UV-B and UV-C). From vanLoon and Duffy (2000).
The polar vortex is a persistent, large-scale cyclone located near the
Earth's poles, in the middle and upper troposphere and the stratosphere. It
surrounds the polar highs and is part of the polar front.
Nitric acid in polar stratospheric
clouds reacts with CFCs to form
chlorine, which catalyzes the
photochemical destruction of
ozone. Chlorine concentrations
build up during the winter polar
night, and the consequent ozone
destruction is greatest when the
sunlight returns in spring
(September/October). These
clouds can only form at
temperatures below about -80°C,
so the warmer Arctic region does
not have an ozone hole.
Maximal ozone will form where form where gas molecule density and uv photon
denisty are optimal.
Ozone layer
Figure 8-14. Altitude versus variations in photon and molecular densities.
The optimum altitude for ozone formation occurs where these curves cross.
Stratospheric ozone distribution
Higher over poles (stratospheric transport)
Higher in summer vs. winter
The ozone hole
Hole varies in size due
to meteorological factors
Additions of N2O,
CFCs, and bromine
Compounds caused the
decline in ozone
Figure 8-15. Seasonal variation of ozone concentrations (in Dobson units) at Halley Bay,
Antarctica, for two different time periods. From Solomon (1990).
Ozone destroying reactions
N2O
N2O + O*  2NO
NO + O3  NO2 + O2
NO2 + O  NO + O2
O + O3  2O2
CFCl3
CFCl3 + hv  CFCl2’ + Cl’
Cl’ + O3  ClO’ + O2
ClO’ + O  Cl’ + O2
O + O3  2O2
For each of these reactions
the Cl’ and the NO return to their
original state. They are catalysts
only and do not participate in the
reaction
Calculating reaction rates for various ozone destroying chemicals
Ta ble 8-8. Kin etic data for variou s re actan ts in th e catalytic des tru ction of ozon e at 235 K*
X + O3
X
Concen tration
( molecu les cm-3)
A
(cm3 molecules -1 s -1)
Ea
(kJ mol-1)
k 235
(cm3 molecules -1 s-1 )
O
1.0 x 10 9
H
2.0 x 10 1 5
1.4 x 10 -10
3.9
1.9 x 10 -11
OH
1.0 x 10 6
1.6 x 10 -12
7.8
3.0 x 10 -14
NO
5.0 x 10 8
1.8 x 10 -12
11.4
5.3 x 10 -15
Cl
Very small
2.8 x 10 -12
21
6.0 x 10 -17
XO + O
XO Calc reaction
Concen tration rate forA Cl’ + O  ClO’
E
Example 8-3:
+ O3 at k235K.
3
( molecu les cm )
(cm molecules s )
(kJ mol )
(cm molecules s )
Cl’ conc. = 5.0 E11. O3 conc = 2.0 E12. Reaction rate = k [Cl][O3]
a
-3
3
-1
-1
235
-1
3
-1
O2
5.0 x 10 1 6
8.0 x 10 -12
17.1
1.3 x 10 -15
HO
1.0 x 10 6
2.3 x 10 -11
0
2.3 x 10 -11
7
-11
-11
19
-12
-12
7
-11
-11
-1
-Calc k using Arrhenius eqn
HO
2.5 x 10
2.2 x 10
-0.1
2.3 x 10
k =NOAe-Ea/RT 5.0 x 10
9.3 x 10
0
9.3 x 10
-12
3
-1
-1
–(21
KJ
mol-1)(8.314
kJ
mol-1K-1)(235K)
k =ClO
(2.8 E 2.0
cmx 10 molecules
s
)e
4.7 x 10
0.4
3.8 x 10
3
-1 s-1
k=*The6.0
e-17 cm
con centrations
of themolecules
species are for an altitude
of 30 k m with the exception of ClO which is for an
2
2
-Rate calc
altitude of 35 km. The co ncentration of o zone at 30 k m is 2.0 x 1 0 12 mo lecules cm-3. Fro m van Loon and
Du ffy (20 00)
rate = (6.0 E-17 cm3 molecules-1 s-1)(5.0 E11 molecules cm-3)(2.0 E12 molecules cm-3)
rate = 6.0 E7 molecules cm-3 s-1
Tropospheric ozone
Bad ozone
Photochemical smog
OH radicals or NO is a catalyst for the production of troposhperic ozone
NO (nitric oxide) released during fuel combustion
NO converted to NO2 by a host of reactions
NO2 + hv  NO + O
O + O2 + M  O3 + M …remember M is a catalytic particle
Of which, O3 is one
Figure 8-16. Variation in abundances of various species, on a 24-hour cycle, produced during a
photochemical smog event. From vanLoon and Duffy (2000).
Radon – 222Rn
Produced from the 238U decay chain
Problematic when bedrock contains uranium
and 214Po progeny are particle active, particles inhaled, lodged in lungs…
subsequent alpha decay damages lung tissue
218Po
Consider Rn levels inside:
Generalized steady state equation for an inside pollutant
Ri = kexCi – kexCo
Ci = Co + Ri/kex
Ci = inside conc
Co = outside conc
kex = exchange coef
Ri= production rate of pollutant
Since Rn is radioactive, the expression is modified to acct. for decay
Ri= kex Ai + lAi – kexA0
Indoor activity of Rn is:
Ai= (Ri + kexA0 ) / (kex + l)
Example 8-4: Radon release from soils to a basement at a rate of 0.01 Bq L-1 h-1
Outdoor air has Rn acitivity of 4.0 E-3 Bq L-1 h-1. Assume air exchange coeff of 10 h-1
What is the steady state indoor actvitiy of Rn?
Ai= (Ri + kexA0 ) / (kex + l)
Plug and chug…answer is Ai = 5.0 E-3 Bq L-1 h-1
Radon flux depends on any factors that change gas diffusion
-soil moisture
-temp (solubility of Rn)
-freezing (caps Rn)
-barometric pressure
Rn is elevated in groundwaters and can be used as a tracer for gw inputs to
Surface waters
Rainwater Chemistry
Compounds found in rainwater come from seawater, terrestrial or pollution sources
T a ble 8 - 1 0 . S o u rces o f in di vi du a l io n s in ra in wa t er *
Origin
Ion
Ma rine inputs
Te rrestrial inputs
Pollution inputs
Na+
Sea salt
Soil d ust
Biomass burning
Mg 2+
Sea salt
Soil d ust
Biomass burning
K+
Sea salt
Bioge nic aerosols
Soil d ust
Biomass burning
Fertilize r
Ca2 +
Sea salt
Soil d ust
Ce ment ma nufacture
Fuel burning
Biomass burning
H+
Gas rea ction
Gas re action
Fuel burning
Cl -
Sea salt
None
Industrial H Cl
Sea salt
DM S from biologica l
dec ay
DM S, H 2S, etc ., fro m
biologica l de cay
Volc anoes
Soil d ust
Biomass burning
N 2 plus lightning
NO 2 fro m biologic al de cay
N 2 plus lightning
Auto e missions
Fossil fu els
Biomass burning
Fertilize r
NH 3 from biologica l
activity
NH 3 fro m bac terial deca y
NH 3 fertilize rs
Hu man , ani mal waste
de co mposition
(Co mbustion)
Bioge nic a erosols adsorbe d
on sea salt
Soil d ust
Biomass burning
Fertilize r
HC O3
CO 2 in air
C O 2 in air
Soil d ust
None
Si O 2, Al, Fe
None
Soil d ust
Land clea ring
2
SO 4

NO 3

NH 4
3
PO 4

Cl- in rain is assumed to
be from a seawater source
Cl- and other ions in from
seawater are assumed to
have a constant proportion
Rain sample can be ‘corrected’
for seawater contribution
*Fro m Berner a nd Berner (199 6)
Excess ion X = total ion X – [(ratio of ion X to Cl- in seawater) (Cl- conc)]
Ta ble 8-11. W eigh t r atios of major ions in s ea wate r r elati ve to C l-- or N a++ *
Ion
W eight ratio to Cl -
Weight ratio to Na+
Cl-
1.00
1.80
Na +
0.56
1.00
M g 2+
0.07
0.12
SO 4
0.14
0.25
Ca2+
0.02
0.04
K+
0.02
0.04
2
*So urce of data for ratio calculations, Wilson (1975)
Ta ble 8-12. P ri mary as s oci ation s for r ain w ater*
Origin
Association
Marin e
Cl - Na - M g - SO 4
Soil
Al - Fe - Si - Ca - (K, M g, Na)
Biological
N O 3 - NH 4 - SO 4 - K
Bio mass burning
N O 3 - NH 4 - P - K - SO 4 - (Ca, Na, Mg)
Industrial pollution
SO 4 - NO 3 - Cl
Fertilizers
K - PO 4 - NH 4 - NO 3
*Fro m Berner and Berner (1996)
Figure 8-17. Average Cl- concentration (mg L-1) of rainwater for the United States from July 1955 to
June 1956. From Berner and Berner (1996).
Marine influence on rainwater chemistry
Low ratios reflect
dust inputs from sodium-rich
rocks
Seawater
Cl- /Na+ = 1.8
Figure 8-18. Average Cl-/Na+ weight ratio of rainwater for the United States from July 1955 to June
1956. From Berner and Berner (1996).
Marine influence on rainwater chemistry
Gaseous species
SO2(g) + 2OH(g)  H2SO4(aq)  2 H+ + SO42- gas phase
SO2(g) + H2O2(aq)  H2SO4(aq)  2H+ + SO42- liquid droplets
NO2(g) + OH(g)  HNO3(aq)  H+ + NO3-
NH3(g) + H2O  NH4OH(aq)  NH4+ + OH-
Acid deposition
and what else?
Figure 8-19. Global SO2 produced by the
burning of fossil fuel, 1940 to 1986, in Tg SO2 S y-1 (1 Tg = 106 metric tons = 1012 g). From
Berner and Berner (1996).
Figure 8-20. Global NOx produced by the burning
of fossil fuel, 1970 to 1986, in Tg NOx - N y-1.
From Berner and Berner (1996).
Figure 8-21. Generalized isoconcentration contours for SO42- (in mg L-1) for atmospheric precipitation over the
contiguous United States in 1995. Source of data is the NADP. From Langmuir (1997).
Figure 8-22. Generalized isoconcentration contours for NO3- (in mg L-1) for atmospheric precipitation over
the contiguous United States in 1995. Source of data is the NADP. From Langmuir (1997).
Two most important species for acid rain
are nitrate and sulfate
Figure 8-23. Average pH for precipitation in 1955-1956 and 1972-1973 for the northeastern United
States and Canada and in 1980 for the contiguous United States and Canada. From Langmuir (1997).
Example 8-7: Calc pH for a stream receiving acid rain.
Calc moles of sulfate and nitrate (from Table 8-13)
sulfate = 2.165 E-5 mol L-1
nitrate = 2.355 E-5 mol L-1
Calculate H+ produced based on what you know about the normality of sulfuric and
Nitric acid. One mole H+ per mole nitrate, two moles H+ per mole sulfate.
Moles H+ = 6.685 E-5 mol L-1
pH = -log [H+] = 4.17
The Nitrogen Cycle
Hog Production in USA
(1 dot= 10,000 Hogs and Pigs)
Chemistry and sources of atmospheric particulates (aerosols)
Primarily tropospheric
transport
Figure 8-24. Sources of atmospheric particulates. Arrows with dashed lines indicate that there is a
gaseous emission associated with the source.
Mineral dust
Fine particles
Aeolian transport
Sahara dust
Trace element delivery to remote oceans (e.g. Antarctica)
Sea Salts
Bursting of bubbles
Pure sea salt aerosols have a predictable ratio of the major ions in seawater
Cl/Na , S/Na , N/Na
Sulfates
Sulfate aerosols can be in the form of (NH4)2SO4, or H2SO4 primarily
Sources: Anthropogenic - combustion
Natural – DMS (dimethyl sulfide), volcanoes (SO2 and H2S)
Carbon-derived particle
Black carbon (soot) – incomplete combustion
Soot from coal (fly ash) = high K, Fe, Mn, Zn
Soot from oil = high V
Organic aerosols- VOCs
Bioaerosols – spores, pollen, and volatile bio-organic compounds (Blue mountains)
Dry deposition – dust settling
Rate determined by radius of particle (Stokes Law)
Wet deposition – washout
Precip (rain or snow)
Condensation
Aerosols serve as condensation nuclei for the formation of
clouds
Source tracking for aerosol deposition
Air mass trajectories
Using atmospheric circulation models to reconstruct the history of an
air mass.
Aerosol : Crust Enrichment Factor
Determination of the amount of additional element that has been added to a
particulate above that amount which you would expect based upon its crustal
Concentration
Assume that the particulate conc. of Al, Fe, Si, Ti, or Sc are solely from the
crustal contribution.
Calc EFcrust the crustal Enrichment Factor
X
RE particulate
EFcrust=
X = conc. of element X
RE = conc. of reference element
X
RE crust
Calculating the noncrustal conc. of element X
Xnoncrustal= Xtotal - REparticulate
X
RE
crust
Ta ble 8-14. Ele men tal co m pos ition of th e c on tin e ntal cru s t*
Concen tration (pp m)
Element
Concentration (pp m)
Upper crust
Bulk crust
Al
80,400
84,100
Se
0.05
0.05
Fe
35,000
70,700
Mo
1.5
1.0
Sc
11
30
Ag
0.050
0.080
Ti
3 000
5400
Cd
0.098
0.098
V
60
230
Sn
5.5
2.5
Cr
35
185
Sb
0.2
0.2
Mn
600
1400
W
2.0
1.0
Co
10
29
Au
0.0018
0.003
Ni
20
105
Pb
20
8.0
Cu
25
75
Th
10.7
3.5
Zn
71
80
U
2.8
0.91
As
1.5
Element
Upp er crust
Bulk Crust
1.0
*Data fro m Taylor and McLen nan (19 85). Both bulk co ntinental crust an d upper contin ental
crust have been used to calculate crust ratios. In some cases the results may differ
significantly. For example, the Pb/Al ratio for th e up per crust is 2.5 x 1 0 -4 while for the bulk
crust the ratio is 9.5 x 10 -5. Given a samp le that has a Pb /Al ratio of 2.5 x 10 -4, th e up per crust
gives an EF of 1 while the bulk crust gives an EF of 2.6. The difference is great eno ugh that
different conclusions might b e drawn regarding the source of the Pb (natu ral for upp er crust
nor malization and anth ropogenic fo r bulk crust normalization). Cr would sho w an even
greater difference, but in the opposite sense. The Cr/Al ratio for the upper crust is 4.4 x 10-4
while for the bulk crust the ratio is 2.2 x 10 -3. Given a sample with a Cr/Al ratio of 2.2 x 10 -3,
the bulk crust normalization gives EF = 1 while the upper crust nor malization gives EF = 5
suggesting that there is an anthropogenic contribution to the Cr content of the sample.
Example 8-9: Aerosol sample has 1000ppm Al and 7ppm Cr. Calculate the
crustal enrichment factor and the noncrustal concentration of Cr
X
RE particulate
7
1,000
EFcrust=
particulate
= 3.2
=
X
RE crust
185
84,100
Xnoncrustal= Xtotal - REparticulate
Crnoncrustal
= 7 - 1000
Crnoncrustal = 4.8 ppm
X
RE
185
84,100
crust
crust
Elemental, molecular and isotopic signatures are used
to source aerosols from crust, marine, or pollution sources
Elemental
X = Xcrust + Xmarine + Xpollution
Using Al, Na, and Se, for reference elements for crust, marine, and pollution sources
respectively...
X = Al
X
Al
+ Na
crust
X
Na
+ Se
marine
X
Se
pollution
To solve for a particular component, divide both sides of equation by the ref element
for that component. Ex: Crust
X = Al
Al
Al
X
Al
crust
+ Na
Al
X
Na
+ Se
marine
Al
X
Se
pollution
Pairwise plotting of the elemental ratios (X/Al, X/Na, X/Se) shows the contributions
of the different sources graphically
100% crustal
source
100% marine
source
100% crustal
source
100% pollut
source
Molecular signatures
‘Fingerprints’ mostly for VOCs
CPI (carbon preference index) = #odd carbon chains / #even carbon chains
petroleum= CPI = 1, natural compounds = CPI < 1
Chain length ratios, aromaticity, trace compounds, biomarkers
Isotopic signatures
Lead revisited- multiple lead isotopes that can be ratio-ed to each other and used
to derive other ratios.
Example 8-10: Aerosol sample:
what’s the 208Pb / 206Pb ratio?
206Pb/204Pb
= 18.004, 208Pb/204Pb = 38.08,
(208Pb/204Pb)(206Pb/204Pb) = 38.08/18.004 = 2.115
3 endmember mixing applied to lead isotopes
Need two markers and 3 equations
1)
(207Pb/206Pb) meas = (207Pb/206Pb)a fa + (207Pb/206Pb)b fb + (207Pb/206Pb)c fc
2)
(206Pb/208Pb) meas = (206Pb/208Pb)a fa + (206Pb/208Pb)b fb + (206Pb/208Pb)c fc
3)
f a + f b + fc = 1
See Example 8-11
Isotopic signatures
Example: 14C
Nuke bomb testing added a lot of 14C to the atm. This excess is decaying and
the atm is not in equilibrium with respect to 14C
Percent modern carbon (pmc) = 0.95 x 14C content of a standard
Percent biogenic carbon = ( 14C of sample / 14C atm at time of sampling) x 100
{all 14C measurements in pmc}
Example 8-12: Aerosol containing formic acid. 14C in sample = 88.7,
14C in atm = 110.5
%biogenic = (14Cmeas/14C atm) x 100
= (88.7/110.5) x 100
= 80%
Happy Thanksgiving!!