Lecture 10: Ocean Carbonate Chemistry:
Ocean Distributions
Ocean Distributions
Controls on Distributions
What is the distribution of CO2 added to the ocean?
See Section 4.4 Emerson and Hedges
Sarmiento and Gruber (2002) Sinks for Anthropogenic Carbon
Physics Today August 2002 30-36
CO2 + rocks = HCO3- + clays
CO2
Gas Exchange
Atm
River Flux
Ocn
CO2 → H2CO3 → HCO3- → CO32-
Upwelling/
Mixing
+ H2O = CH2O + O2
+ Ca2+ = CaCO3
CO2
BorgC
BCaCO3
Biological Pump
Controls:
pH of ocean
Sediment diagenesis
Influences on pCO2
Ko: Solubility of CO2
K1, K2: Dissociation constants
Function of Temperature, Salinity
Depends on biology only
Depends on biology
and gas exchange
Ocean Distributions – versus depth, versus ocean
Atlantic
Pacific
Points:
See Key et al (2004)
GBC
1. Uniform surface
concentrations
2. Surface depletion Deep enrichment
3. DIC < Alk
Q?
4. DDIC > DAlk
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+ + HCO3HCO3-  H+ + CO32As CO2 is produced during respiration we should observe:
pH  DIC  Alk  PCO2 
The trends will be the opposite for photosynthesis.
B) CaCO3
dissolution/precipitation
CaCO3(s)  Ca2+ + CO3 2Also written as:
CaCO3(s) + CO2 + H2O  Ca2+ + 2 HCO3As CaCO3(s) dissolves, CO32- is added to solution. We should observe:
pH  DIC  Alk  PCO2 
Influence of Nitrogen Uptake/Remineralization on Alkalinity
NO3- assimilation by phytoplankton
106 CO2 + 138 H2O + 16 NO3- → (CH2O)106(NH3)16 + 16 OH- + 138 O2
NH4 assimilation by phytoplankton
106 CO2 + 106 H2O + 16 NH4+ → (CH2O)106(NH3)16 + 16 H+ + 106 O2
NO3- uptake is balanced by
OH- production
Alk ↑
NH4+ uptake leads to
H+ generation
Alk ↓
Alk = HCO3- + 2 CO32- + OH- - H+
See Brewer and Goldman (1976) L&O
Goldman and Brewer (1980) L&O
Experimental Culture
Ocean Distributions of, DIC, Alk, O2 and PO4 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 < Alk
DDIC > DAlk
5. Profile of pH is similar in
shape to O2.
6. Profile of PCO2 (not shown)
mirrors O2.
Inter-Ocean Comparison
Carbonate ion (CO32-) and pH decrease from Atlantic to Pacific
x 10-3 mol kg-1
x 10-6 mol kg-1
S = 35
T = 25C
Alk
2.300
DIC
1.950
CO32246
pH
8.12
North Atlantic
Deep Water
2.350
2.190
128
7.75
Antarctic
Deep Water
2.390
2.280
101
7.63
Deep Atlantic
to Deep Pacific
DAlk = 0.070
DDIC = 0.180
7.46
So
DAlk/DDIC = 0.40
Surface Water
North Pacific
Deep water
2.420
Q? CO2Sys/CO2Calc
2.370
72
CO32- decreases from
surface to deep Atlantic
to deep Pacific. These CO32- are from CO2Sys.
Can Approximate as CO32- ≈ Alk - DIC
Composition of Sinking Particles and Predicted Changes
Ocean Alkalinity versus Total CO2 in the Ocean
(Broecker and Peng, 1982)
DDIC/DAlk ≈ 1.5/1
Work Backwards
DAlk / DDIC ≈ 0.66
= 2/3
= 2 mol Org C / 1 mol CaCO3
Emerson and Hedges Color Plate
What is composition of sinking particles?
Data from annual sediment traps deployments
5 g POC g m-2 y-1 / 12 g mol-1 = 0.42 mol C m-2 y-1
40 g CaCO3 g m-2 y-1 / 105 g mol-1 = 0.38 mol C m-2 y-1
Org C / CaCO3 ~ 1.1
From Klaas and Archer (2002) GBC
PIC/POC in sediment trap samples
POC and CaCO3 Export Fluxes
This Study
Previous Studies
POC (Gt a−1)
Global export
9.6 ± 3.6
11.1–12.9 [Laws et al., 2000]b
9.2 [Aumont et al., 2003]c
8.6 [Heinze et al., 2003]c
8.7–10.0 [Gnanadesikan et al., 2004]c
9.6 [Schlitzer, 2004]d
5.8–6.6 [Moore et al., 2004]c
CaCO3 (GtC a−1)
Global export
0.52 ± 0.15
0.9–1.1 [Lee, 2001]b
1.8 [Heinze et al., 1999]c
1.64 [Heinze et al., 2003]c
0.68–0.78 [Gnanadesikan et al., 2004]c
0.38 [Moore et al., 2004]c
0.84 [Jin et al., 2006]c
0.5–4.7 [Berelson et al., 2007]b
POC/CaCO3 = 9.6 / 0.52 = 18.5
Based on Global Model results of
Sarmiento et al (2992) GBC; Dunne et al (2007) GBC
Revelle Factor
The Revelle buffer factor defines how much CO2 can be absorbed by
homogeneous reaction with seawater. B = dPCO2/PCO2 / dDIC/ DIC
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., Science (1979)
dPCO2/PCO2 = B dDIC/DIC
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
B↑ as T↓ as CT↑
Revelle Factor Numerical Example (using CO2Sys)
CO2 + CO32- = HCO3-
CO2
350ppm + 10% = 385ppm
Atm
Ocn
CO2 → H2CO3 → HCO3- → CO32-
at constant alkalinity
DIC
11.3 mM
1640.5 mM
183.7
1837
+1.2 (10.6%)
+27.7 (1.7%)
-11.1 (-6.0%)
+17.9 (+0.97%)
12.5
1668.2
174.2
1854.9
The total increase in DIC of +17.9 mM is mostly due to a big change
in HCO3- (+27.7 mM) countering a decrease in CO32- (-11.1 mM).
Most of the CO2 added to the ocean reacts with CO32- to make HCO3-.
The final increase in H2CO3 is a small (+1.2 mM) portion of the total.
Air-Sea CO2 Disequilibrium
Emerson and Hedges Plate 8
Effect of El Nino on ∆pCO2 fields
High resolution pCO2 measurements in the Pacific since Eq. Pac-92
Eq Pac-92 process study
PCO2sw
Always greater
than atmospheric
El Nino Index
Cosca et al. in press
Expression of Air -Sea CO2 Flux
Magnitude
Mechanism
Apply over larger space time domain
k-transfer velocity
S – Solubility
From Sc # & wind speed
From SST & Salinity
F = k s (pCO2w- pCO2a) = K ∆ pCO2
pCO2w
From measurements
and proxies
pCO2a
From CMDL
CCGG network
Global Map of Piston Velocity (k in m yr-1) times CO2 solubility (mol m-3) = K
from satellite observations (Nightingale and Liss, 2004 from Boutin).
∆pCO2 fields
Overall trends known:
* Outgassing at low latitudes (e.g. equatorial)
* Influx at high latitudes (e.g. circumpolar)
* Spring blooms draw down pCO2 (N. Atl)
* El Niños decrease efflux
JGOFS Gas Exchange Highlight #4 ∆pCO2 fields:Takahashi climatology
Monthly changes in pCO2w
Fluxes: JGOFS- Global monthly fluxes
Combining pCO2 fields with k: F = k s (pCO2w- pCO2a)
On first order flux and ∆pCO2 maps do not look that different
CO2 Fluxes: Status
Do different parameterizations between gas exchange and wind matter?
Global uptakes
Liss and Merlivat-83:
Wanninkhof-92:
Wanninkhof&McGillis-98:
1 Pg C yr-1
1.85 Pg C yr-1
2.33 Pg C yr-1
Zemmelink-03:
2.45 Pg C yr-1
Yes!
Global average k (=21.4 cm/hr):
2.3 Pg C yr-1
We might not know exact parameterization with forcing but forcing is
clearly important
Compare with net flux of 1.3 PgCy-1 (1.9 - 0.6)
in Sarmiento and Gruber (2002), Figure 1
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.
Photosynthesis/respiration (shown as apparent oxygen utilization or AOU = O2,sat – O2,obs)
and CaCO3 dissolution/precipitation vectors (from Park, 1969)
CH2O + O2 → CO2 + H2O
as O2↓ AOU ↑ CO2 ↑