Carbonate Solubility Dissolution mechanisms Dissolution rate expressions Saturation

Carbonate Solubility Dissolution mechanisms Dissolution rate expressions Saturation state in the ocean T, P, and CO 2 release In situ [CO 3– 2] and the saturation horizon Sedimentary evidence of carbonate dissolution Kinetic and/or thermodynamic controls

p. H scales Free Total Seawater Zeebe and Wolf-Gladrow

p. H scale comparisons p. K comparisons

Carbonate Solubility Ca. CO 3 <=> Ca+2 + CO 3– 2 Define satauration state as: � Ώ = IAP / Ksp With IAP (ion activity product) = [Ca+2][CO 3– 2] for the solution, and with the apparent solubility product: for a solution in equilibrium with solid Ca. CO 3 (T, P, S)

K’sp depends on mineralogy (calcite < aragonite < Mg-calcite, …) as well as T and P At equilibrium: Ώ = IAP / Ksp = [Ca+2][CO 3– 2] = 1 When Ώ > 1, solution is supersaturated When Ώ < 1, solution is undersaturated

Often, we take [Ca+2] as a constant function of salinity, and express Ώ in terms of the saturation carbonate ion concentration so, if [Ca+2] constant, Define DCO 3– 2 as ([CO 3– 2] in situ - [CO 3– 2]sat)

Controls on saturation state: Pressure and temperature effects on carbonate solubility Solubility higher in deeper and colder water Pressure, temperature, and respiration effects on [CO 3– 2] in situ [CO 3– 2] lower in deeper, colder, and “older” water Broecker and Peng

in situ saturometer Ben-Yakov et al. , 74

Profile of saturation carbonate ion concentration, based on in situ saturometer; strong increase in solubility with increasing depth. Broecker and Peng

Use observed K’sp in 0 -15 cm interval of in situ pore water profiles (Alk + p. H, = Ca) to estimate solubility. Sayles, 1985

Solubility vs depth for Sayles’ pore water data and lab solubility studies

Solubility vs depth comparison. Broecker and Peng

What is the relationship between degree of undersaturation and dissolution rate? Ca. CO 3 <=> Ca+2 + CO 3– 2 Define the mass-normalized dissolution rate: “k*” is the reaction rate (%/day) “n” is the reaction order “*” to account for % Ca. CO 3 in bulk sediments, and for surface area Can express dissolution rate in terms of carbonate ion concentration R = k* ([CO 3– 2]sat - [CO 3– 2] in situ )n Where ([CO 3– 2]sat is the saturation carbonate ion concentration (T, P)

An unresolved puzzle: Laboratory dissolution studies (Keir, 1980) and theoretical arguments (Morse) suggest that k* ~ 1000 / day and n = 4. 5 strongly non-linear fast dissolution really fast at low Ώ But ~ all porewater dissolution studies give k* ~ 1 - 10 / day (102 to 103 slower) and n = 3, or even 1 (Hales and Emerson, 1997)

Morse and Arvidson - High order kinetics consistent with control of dissolution by one of the surface processes (adsorption, migration, reaction, migration, desorption) Hales and Emerson, Sayles and Martin Shallow depth at which porewaters reach saturation implies low-order kinetics; otherwise final approach to saturation would take much longer (deeper).

Recalculate with new constants: n = 3. 2 Hales and Emerson Reassess solubiliy too: n = 1. 3

Why do we care? Influences of dissolution rate law include: vertical profile of dissolution in sediments impact of CO 2 release by benthic decomposition total dissolution rate shape of lysocline

Saturation state in the ocean Influence of T, P, and CO 2 release In situ [CO 3– 2] and the saturation horizon

Pressure, temperature, and respiration effects on [CO 3– 2] in situ. [CO 3– 2] lower in deeper, colder, and “older” water. Aging dominates.

How does [CO 3– 2] respond to changes in Alk or DIC? CT = [H 2 CO 3*] + [ HCO 3–] + [CO 3– 2] ~ [ HCO 3–] + [CO 3– 2] (an approximation) Alk = [OH–] + [HCO 3–] + 2[CO 3– 2] + [B(OH)4 -] – [H+] ~ [HCO 3–] + 2[CO 3– 2] (a. k. a. “carbonate alkalinity”) So (roughly): [CO 3– 2] ~ Alk – CT ↑ , [CO 3– 2] ↓ Alk ↑ , [CO 3– 2] ↑

Photosynthesis CO 2 + H 2 O => “CH 2 O” + O 2 D SCO 2 = -1 D Alk = 0 So D [CO 3– 2] = D Alk - D SCO 2 = 0 – (-1) = +1 [CO 3– 2] increases, Ώ increases. Respiration “CH 2 O” + O 2 => CO 2 + H 2 O D SCO 2 = +1 D Alk = 0 So D [CO 3– 2] = D Alk - D SCO 2 = 0 – 1 = -1 [CO 3– 2] decreases, Ώ decreases.

Zeebe and Wolf-Gladrow

Calcification Ca+2 + CO 3– 2 => Ca. CO 3 D SCO 2 = -1 D Alk = -2 So D [CO 3– 2] = -2 – (-1) = -1 [CO 3– 2] decreases, Ώ decreases. Dissolution: Ca. CO 3 => Ca+2 + CO 3– 2 D SCO 2 = +1 D Alk = +2 So D [CO 3– 2] = 2 – 1 = +1 [CO 3– 2] increases, Ώ increases.

% Ca. CO 3 vs. water depth “lysocline” – onset of dissolution “calcite compensation depth” – dissolution rate = rain rate

Takahashi and Broecker ’ 80; GEOSECS data

in situ Aragonite Calcite saturation

Another puzzle: NPac sediment traps – loss of carbonate above the aragonite saturation horizon. High aragonite/calcite ratio in sinking flux: 0. 2 to 20, most > 1 Betzer et al. , 1984

Carbonate flux in NAtl sediment traps – loss of carbonate above the aragonite saturation horizon. Thermodynamics? Highsolubility phase(s)? Magnesian-calcite) Artifacts? Milliman et al. , 1999

Sedimentary evidence of carbonate dissolution Kinetic and/or thermodynamic controls

Sediment evidence for dissolution?

Here, core-top aragonite distribution seems to match the depth of the aragonite saturation horizon.

Farrell and Prell

Adelseck 1977 – lab study of selective dissolution of planktonic foraminifera Assemblage change only after substantial carbonate dissolution; not a sensitive indicator

Berger et al. , 1982 – Selective dissolution in Eq. Pac Differential dissolution: a range of susceptibility Thermodynamics? (different solubilities? ) Kinetics? (different area/g, crystallinity? )