Lecture 10: Ocean Carbonate Chemistry:

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Lecture 10 Ocean Carbonate Chemistry
Ocean Distributions Ocean
Distributions Controls on Distributions
What is the distribution of CO2 added to the
ocean?
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CO2 rocks HCO3- clays
CO2
River Flux
Gas Exchange
Atm
Ocn
CO2 ? H2CO3 ? HCO3- ? CO32-
Upwelling/ Mixing
H2O CH2O O2
Ca2 CaCO3
CO2
BorgC
BCaCO3
Biological Pump
Controls pH of ocean Sediment diagenesis
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Air-Sea CO2 Disequilibrium
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Influences on pCO2
Ko Solubility of CO2 K1, K2 Dissociation
constants Function of Temperature, Salinity
Depends on biology and gas exchange
Depends on biology only
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What happens to the CO2 that dissolves in water?
CO2 is taken up by ocean biology to produce a
flux of organic mater to the deep sea
(BorgC) CO2 H2O CH2O O2 Some carbon is
taken up to make a particulate flux of CaCO3
(BCaCO3) Ca2 2HCO3- CaCO3(s) CO2
H2O The biologically driven flux is called the
Biological Pump. The sediment record of BorgC
and BCaCO3 are used to unravel paleoproductivity.
The flux of BorgC to sediments drives an
extensive set of oxidation-reduction reactions
that are part of sediment diagenesis. Carbonate
chemistry controls the pH of seawater which is a
master Variable for many geochemical processes.
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Ocean Distributions versus depth, versus ocean
Atlantic
Pacific
Points 1. Uniform surface concentrations 2.
Surface depletion - Deep enrichment 3. DIC lt
Alk 4. DDIC gt DAlk
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Ocean Distributions of PO4, Ct, Alk, O2 versus
Depth and Ocean
The main features are 1. uniform surface
values 2. increase with depth 3. Deep ocean
values increase from the Atlantic to the
Pacific 4. DIC lt Alk and DDIC gt
DAlk 5. Profile of pH is similar in
shape to O2. 6. Profile of PCO2 (not shown)
mirrors O2.
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Carbonate ion (CO32-) and pH decrease from
Atlantic to Pacific
x 10-3 mol kg-1 x 10-6 mol
kg-1 Alk DIC CO32- pH Surface
Water 2.300 1.950 242 8.30 North
Atlantic 2.350 2.190 109 8.03 Deep
Water Antarctic 2.390 2.280 84 7.89 Deep
Water North Pacific 2.420 2.370 57 7.71 Deep
water
Deep Atlantic to Deep Pacific DAlk
0.070 DDIC 0.180 So DAlk/DDIC 0.40
CO32- decreases from surface to deep Atlantic to
deep Pacific. CO32- Alk - DIC
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Controls on Ocean Distributions
A) Photosynthesis/Respiration Organic matter
(approximated as CH2O for this example) is
produced and consumed as follows CH2O O2 ?
CO2 H2O Then CO2 H2O ?
H2CO3 H2CO3 ? H HCO3- HCO3- ? H
CO32- As CO2 is produced during respiration we
should observe pH ? DIC ? Alk ? PCO2 ? The
trends will be the opposite for
photosynthesis. B) CaCO3 dissolution/precipitatio
n CaCO3(s) ? Ca2 CO3 2- Also written
as CaCO3(s) CO2 H2O ? Ca2 2
HCO3- As CaCO3(s) dissolves, CO32- is added to
solution. We should observe pH ? DIC ?
Alk ? PCO2 ?
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Composition of Sinking Particles and Predicted
Changes
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New estimates of CaCO3/orgC export ratio from the
euphotic zone (100m) Approach a box model
for surface ocean mass balances for alkalinity
and nitrate Result Global average ratio of
CaCO3 / org C 0.06 0.03 Maximum values are
in the equatorial regions. Details See
Sarmiento et al (2002) A new estimate of the
CaCO3 to organic carbon export ratio. Global
Biogeochemical Cycles 16, doi10.1029/2002GB001919
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Ocean Alkalinity versus Total CO2 in the
Ocean (Broecker and Peng, 1982)
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Carbonate System Calculations
pH and CT
Alkalinity and PCO2
A useful shorthand is the alpha notation, where
the alpha (a) express the fraction each carbonate
species is of the total DIC. These a values are a
function of pH only for a given set of acidity
constants. Thus H2CO3 ao CT HCO3- a1
CT CO32- a2 CT The derivations of the
equations are as follows ao H2CO3 / CT
H2CO3 / (H2CO3 HCO3 CO3) 1 / ( 1
HCO3 / H2CO3 CO3/H2CO3) 1 / ( 1 K1/H
K1K2/H2) H2 / ( H2 HK1 K1K2) The
values for a1 and a2 can be derived in a similar
manner. a1 HK1 / (H2 H K1 K1K2) a2
K1K2 / ( H2 H K1 K1K2) For example Assume pH
8, CT 10-3, pK1' 6.0 and pK2'
9.0 H2CO3 10-5 mol kg-1 (note the answer is
in concentration because we used K') HCO3-
10-3 mol kg-1 CO32- 10-4 mol kg-1
Alk HCO3 2 CO3 OH - H For this
problem neglect H and OH (a good assumption ),
then CT a1 2 CT a2 CT (a1 2a2) We can
use this equation if we have a closed system and
we know 2 of the 3 variables (Alk, CT or pH). For
an open system we can express CT in terms of PCO2
as follows We know that H2CO3 CT ao ( you
can also use this equation if you know pH and
PCO2) But H2CO3 can be expressed in terms of the
Henry's Law KH PCO2 CT ao So CT KH
PCO2 / ao Now Alk (KH PCO2 / ao ) ( a1
2a2) Alk KH PCO2 ( (a1 2 a2 ) / ao ) Alk
KH PCO2 ( HK1 2 K1K2 / H2 ) Assume Alk
10-3 PCO2 10-3.5 pK1' 6.0 pK2' 9.0 Then
pH 8.3
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What controls the pH of seawater? pH in
seawater is controlled by alkalinity and DIC and
can be calculated from these two parameters as
shown below. Alk ? HCO3- 2 CO32- Alk ? CT a1
2 CT a2 Alk CT (HK1' 2 K1' K2' ) / (H2
H K1' K1'K2') Rearranging, we can
calculate pH from Alk and CT. (H) ?-K1'
(Alk-CT) (K1')2 (Alk-CT)2 - 4 Alk K1' K2'
(Alk - CT) ? / 2 Alk So the question boils
down to what controls alkalinity and total
CO2. Internal variations of pH in the ocean and
controlled by internal variations in DIC and
alkalinity which are controlled by
photosynthesis, respiration and CaCO3
dissolution and precipitation. The long term
controls on alkalinity and DIC are the balance
between the sources and sinks and these are the
weathering (sources) and burial (sinks) of
silicate and carbonate rocks and organic
matter.
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Revelle Factor
The Revelle buffer factor defines how much CO2
can be absorbed by homogeneous reaction with
seawater. B CT / PCO2 (?PCO2/?CT)alk CT
(?PCO2/?H)alk PCO2 (?CT/?H)alk After
substitution B CT / (H2CO3 CO32-) For
typical seawater with pH 8, Alk 10-2.7 and CT
10-2.7 H2CO3 10-4.7 and CO32- 10-3.8 then
B 11.2
Field data from GEOSECS Sundquist et al.,
(1979) A value of 10 tells you that a change of
10 in atm CO2 is required to produce a 1 change
in total CO2 content of seawater, By this
mechanism the oceans can absorb about half
of the increase in atmospheric CO2
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Photosynthesis/respiration (shown as apparent
oxygen utilization or AOU O2,sat O2,obs) and
CaCO3 dissolution/precipitation vectors (from
Park, 1969)