PALEOCEANOGRAPHY, PHYSICAL AND CHEMICAL PROXIES | Dissolution of Deep-Sea Carbonates

Dissolution of Deep-Sea Carbonates S Barker, Cardiff University, Cardiff, UK ã 2013 Elsevier B.V. All rights reserved. This article is reproduced from the previous edition, volume 2, pp. 1710–1722, ã 2007, Elsevier BV.

Introduction Marine carbonates represent the major component of deep-sea sediments (Figure 1; Table 1). The distribution of carbonates throughout the world’s oceans reflects the interplay between production in the surface ocean and dissolution, either in the water column, or at the seafloor before final burial. Since the modern rate of carbonate production is several times greater than the global burial rate, dissolution is an important term. The dissolution of carbonates in the deep sea is ultimately controlled by the saturation state of mean ocean water, which itself is a function of the balance between the rate of supply (through weathering) and burial of carbonate at the sea floor. In this way, the system can be thought of as self-regulating. The importance of carbonate dissolution within the global carbon cycle can be appreciated by its relation to mean ocean saturation and atmospheric CO2. As a rule of thumb, increased dissolution will result from increased levels of CO2 and a corresponding decrease in ocean saturation state. Past changes in atmospheric CO2 have been linked to changes in carbonate dissolution intensity. Similarly, much of the CO2 currently emitted by fossil fuel burning will ultimately be sequestered by an increase in deep-sea carbonate dissolution.

Global Distribution of Deep-Sea Carbonates Surface Production Almost all the carbonate in deep-sea sediments is manufactured by marine organisms living in the top few hundred meters of the water column. Various species of phytoplankton and zoo plankton produce calcium carbonate (CaCO3) shells or platelets, that fall to the seafloor after death. Some organisms build tests made of calcite (e.g., coccolithophores and Foraminifera) while others (e.g., pteropods) form aragonite shells. Differences in the carbonate mineral and form result in different distributions with water depth due to their varying solubilities (see the Section ‘Mechanisms of Carbonate Dissolution’). The primary control on the distribution of deep-sea carbonates is therefore the pattern of biological production in the surface ocean (Figure 1). Carbonate-rich sediments (oozes) are to be found beneath highly productive regions such as coastal or equatorial upwelling zones (a notable exception being the Southern Ocean), whereas oligotrophic areas are associated with sediments lower in carbonate content.

Dissolution in the Deep Sea One of the most striking features of ocean sedimentation is the dramatic decrease in the carbonate content of seafloor sediments with increasing water depth. Topographic highs such as mid-ocean ridges tend to be draped in carbonate-rich oozes,

whereas abyssal depth sediments are almost devoid of carbonates (Figure 2). The transition between these two realms is a zone, several hundred meters wide, over which the CaCO3 content drops from perhaps 85–95% to almost zero. The top of this zone (where CaCO3 content starts to decrease rapidly) is generally termed the sedimentary lysocline (this term has also been applied to the entire transitional zone). The depth below which %CaCO3 is effectively zero (where the flux of carbonate to the seafloor is matched by the rate of dissolution) is termed the carbonate compensation depth (the calcite compensation depth, CCD, is deeper than the aragonite compensation depth, ACD, due to the greater solubility of aragonite; see the Section ‘Mechanisms of Carbonate Dissolution’). Understanding why this transition occurs and the overall role of carbonate dissolution within the global carbon cycle requires an appreciation of certain thermodynamic and kinetic considerations.

Thermodynamics of the Carbonate System in Seawater Dissolved Inorganic Carbon, DIC The dissolution of carbonate in seawater is intimately related to the marine carbon dioxide (CO2) system. CO2 dissolved in seawater exists in three inorganic forms: CO2 (aq.) (aqueous CO2), HCO3
(bicarbonate ion), and CO32
(carbonate ion). A fourth form, H2CO3 (carbonic acid), is present only in very low concentrations and is generally considered in combination with CO2 (aq.) and denoted as CO2: ½CO2  ¼ ½CO2 ðaq:Þ þ ½H2 CO3 

[1]

Here square brackets denote the stoichiometric concentrations (molarity or mol kg
1). The total dissolved inorganic carbon (DIC or S CO2) in seawater is therefore defined as 2
DIC  SCO2 ¼ ½CO2  þ ½HCO
3  þ ½CO3 

[2]

Under equilibrium conditions (no net exchange of CO2 between air and sea), the concentration of CO2 in the surface ocean is related to the fugacity (or partial pressure) of CO2 in the atmosphere according to Henry’s law (the fugacity of a gas is similar to its partial pressure but allows for its nonideality): K

CO2 ðgÞ ¼0 CO2 ðaq:Þ Table 1

[3]

Global distribution of the major sedimentary types

Component

Atlantic

Pacific

Indian

World ocean

Calcareous ooze Siliceous ooze Red clay

68 6.7 25.3

36.2 14.7 49.1

54.3 20.4 25.3

47.7 14.2 38.1

Data from Thurman HV and Trujillo AP (2002) Essentials of Oceanography, 7th ed. Upper Saddle River: Prentice-Hall.

859

860

PALEOCEANOGRAPHY, PHYSICAL AND CHEMICAL PROXIES | Dissolution of Deep-Sea Carbonates

90

180 150 120

90

60

30

0

30

60

90

120 150 180

70 60 50 30 10 0 10 30 50 60 70 90 Neritic Continental

Pelagic Calcareous ooze

Abyssal clay

Siliceous ooze Diatom Radiolarian

Figure 1 Global distribution of dominant sedimentary types. Reprinted from Thurman HV and Trujillo AP (2002) Essentials of Oceanography, 7th ed. Upper Saddle River: Prentice-Hall.

where K0 is the solubility coefficient of CO2 in seawater: K0 ¼ ½CO2 =f CO2

[4]

The dissolved carbonate species are related by the following equilibria: K

CO2 ðaq:Þ þ H2 O ¼H H2 CO3 K1

CO2 þ H2 O ¼ Hþ þ HCO
3 K2 HCO
3 ¼

þ

H þ

CO2
3

[5] [6] [7]

where K1 and K2 are the first and second dissociation constants of carbonic acid (which do not differentiate between CO2 (aq.) and H2CO3). Thus, the system may be simplified to K1

K2

þ 2
þ CO2 þ H2 O Ð HCO
3 þ H Ð CO3 þ 2H

leads to the thought that the oceans must play an important role in controlling atmospheric CO2 levels. It can also been seen from Figure 3 that [HCO3
] is relatively constant for average oceanic pH values. This leads to a valuable rule of thumb; the concentration of carbonate ion, [CO32
], is inversely related to [CO2]. This generalized statement can be derived numerically, by combining Eqns [9] and [10]: 2 2
K1* =K2* ¼ ½HCO
3  =½CO2 =½CO3 

K1*,

K2*,

Since and are effectively constant for a given T, S, and P, Eqn [11] reduces to ½CO2 =½CO2
3   constant


K2* ¼ ½Hþ ½HCO2
3 =½HCO3 

[9] [10]

K1* and K2* are dependent on temperature (T ), salinity (S), and pressure (P) (Table 2, Figure 3). Expressions for K0,1,2 have been determined repeatedly and the reader is referred to DOE (1994) for current recommended usage and references. The partitioning of DIC between CO2, HCO3
, and CO32
can be usefully illustrated with a Bjerrum plot (Figure 3). It can be seen that, for typical ocean water, HCO3
dominates (90%), followed by CO32
. CO2 represents only a few percent of the total dissolved inorganic carbon in seawater. This biasing away from CO2 explains why the oceanic carbon reservoir is so much (some 50–60 times) larger than the atmosphere and

[12]

From Eqn [11] we may also derive the pH reaction

[8]

K1 =K2


CO2 þ CO2
3 þ H2 O Ð 2HCO3

It is most common to use the stoichiometric equilibrium constants, K1* and K2*, which are defined as K1* ¼ ½Hþ ½HCO
3 =½CO2 

[11]

[HCO3
]

[13]

Alkalinity, TA The total alkalinity (TA) of seawater may be defined as the charge difference between the major conservative (concentration unaffected by changes in pH, pressure, or temperature) cations and anions: TA ¼ S conservation cations
S conservative anions TA  ð½ Naþ þ 2½Mg2þ þ 2½Ca2þ þ ½Kþ Þ
ð½Cl
 þ 2½SO2
4 

[14]

The small excess positive charge ( 2 as compared with  600 mmol kg
1 total positive charge concentration) is balanced by contributions from carbonate, borate, and water alkalinities (Figure 4):

PALEOCEANOGRAPHY, PHYSICAL AND CHEMICAL PROXIES | Dissolution of Deep-Sea Carbonates

West N Atlantic

Water depth (km)

3

Central equatorial Pacific

East N Atlantic

861

Western equatorial Pacific

4

5

6

0

20

40 60 80 100 0 %CaCO3

20

40 60 80 100 0 %CaCO3

20

40 60 80 100 0 %CaCO3

20

40 60 80 100 %CaCO3

Figure 2 %CaCO3 in core-top sediments from the Atlantic and Pacific Oceans. Carbonate is preserved to significantly deeper depths in the Atlantic. Reprinted from Broecker WS and Peng T-H (1982) Tracers in the sea. Palisades: Eldigio Press.

Table 2 Influence of salinity, temperature, and pressure on the dissociation constants of carbonic acid and the solubility products of calcite and aragonite. 1 dbar  1 m T (C)

S

25 2 25 2

P (dbar)

35 35 34 35

pK1*

pK2* a

0 0 0 3000

5.8563 6.0860a 5.8594a 5.9397c

pKsp* (calcite) a

b

8.9249 9.3433a 8.9342a 9.2536c

6.3692 6.3667b 6.3854b 6.1043c

pKsp* (aragonite) 6.1880b 6.1656b 6.2034b 5.9189c

a

DOE (1994). Mucci (1983). c Pressure correction from Millero FJ (1995) Thermodynamics of the carbon dioxide system in the oceans. Geochimica et Cosmochimica Acta, 59: 661–677. b

Fractional contribution

1

[HCO3− ]

[CO2]

important contributor to TA, which leads to a useful approximation:

pK2*

pK1*

   2
 TA  CA ¼ HCO
3 þ 2 CO3

[CO2− ] 3

Since [CO2] is only a minor component of SCO2,    2
 SCO2  HCO
3 þ CO3

0.1

0.01

[16]

[17]

which leads to 

0.001 4

5

6

7

8

9

10

  TA
SCO2 CO2
3

[18]

11

pH Figure 3 Bjerrum plot for the carbonate system in seawater for various 2
P, T, S. Blue, red, and green represent [CO2], [HCO
3 ], and [CO3 ], respectively. Heavy curves are for S ¼ 35, T ¼ 25 C, P ¼ 0 dbar. Narrow curves are for S ¼ 35, T ¼ 0 C, P ¼ 0 dbar, dashed curves are S ¼ 35, T ¼ 0 C, P ¼ 3000 dbar. Note that the pK* values (¼
log 10 {K*}) for 2

K1* and K2* are equal to pH when [HCO
3 ] ¼ [CO2] and [CO3 ] ¼ [HCO3 ], respectively. Shaded area represents typical ocean pH values.

TA  PA
2

þ ¼ ½HCO
3 þ 2½CO3 þ ½BðOHÞ4  þ ½OH 
½H 

[15]

where PA is the practical alkalinity (alkalinity for most practical purposes). In fact, carbonate alkalinity (CA) is by far the most

Balance between DIC and TA Alkalinity is delivered to the oceans by continental weathering and hydrothermal exchange and is lost mainly through sedimentary CaCO3 burial and reef formation. Weathering or burial of CaCO3 adds or removes TA and DIC in the ratio 2:1. Therefore, an increase in one or other of these terms will cause a change in [CO32
] (Eqn [18]) and therefore pCO2 (see the Section ‘Changes in Carbonate Dissolution through Time’). Similarly, the manufacture or dissolution of CaCO3 within the marine system causes a decrease or increase in TA and DIC in a 2:1 ratio. This means that the formation of CaCO3 actually causes an increase in pCO2 (Figure 5).

862

PALEOCEANOGRAPHY, PHYSICAL AND CHEMICAL PROXIES | Dissolution of Deep-Sea Carbonates

Determining Carbonate Parameters

650 600

[K+] 2[Ca2+]

550

] 2[SO2− 4

([HCO−3]+ 2[CO2− 3 ]+ [B(OH)−4 ])

It is not possible to measure the individual species of the carbonate system directly. However, the four parameters, DIC, TA, fCO2, and pH, can be determined. From a knowledge of any two of these parameters, along with temperature, salinity, and pressure, it is possible to calculate all six (SCO2, [CO32
], [HCO3
], [CO2], pH, TA).

2[Mg2+]

500

mmol kg–1

450

Mechanisms of Carbonate Dissolution Thermodynamics 100

40

Efforts to understand the mechanisms of carbonate dissolution have involved many laboratory and field-based studies on both synthetic and biogenic carbonates. However, there remains considerable uncertainty regarding the specific parameters involved. Equilibrium solution of calcium carbonate in seawater may be expressed as

20

CaCO3 ðsÞ ¼ Ca2þ þ CO2
3

80 60 Cl−

Na+

Ksp

[19]

The stoichiometric solubility product, Ksp * is then defined as 0 Cations

* Ksp ¼ ½Ca2þ sat ½CO2
3 sat

Anions

Figure 4 Charge balance for the major ions in seawater. The small excess positive charge of the conservative cations over conservative 2

anions is balanced (mainly) by [HCO
3 ], [CO3 ] and [B(OH)4 ] with
þ minor contributions from [OH ] and [H ]. Reprinted from Zeebe RE and Wolf-Gladrow D (2001) CO2 in Seawater: Equilibrium, Kinetics, Isotopes. Amsterdam: Elsevier.

where the subscript ‘sat’ refers to the concentration at saturation. The tendency for CaCO3 to dissolve in seawater depends on the degree of saturation (O) of the surrounding seawater with respect to the specific mineral phase in question (calcite or aragonite): * O ¼ ½Ca2þ ½CO2
3 =K sp

DIC vs. TA @ 25C, S = 35

2
O ¼ ½CO2
3 =½CO3 sat

CaCO3 dissolution (supply)

TA(μmol kg–1)

250 2260 300 350 pCO2 2220

2
2
D½CO2
3  ¼ ½CO3 
½CO3 sat * depends Ksp

1980

[22]

When O ¼ 1, the system is at equilibrium and should be stable. O < 1 indicates that the solution is undersaturated with respect to carbonate and we should expect dissolution to occur. When O > 1, the solution is said to be supersaturated and, thermodynamically, we may expect precipitation of carbonate. However, due to complexities such as the composition of seawater, spontaneous precipitation is rare (see Morse and He, 1993). A widely used alternative for expressing the degree of saturation is the D[CO3
] notation (Figure 8), where

200 2280

CaCO3 precipitation (burial) 2200 1900 1920 1940 1960

[21]

where [Ca2þ] and [CO32
] are the concentrations in solution. Since the concentration of Ca2þ in seawater is relatively constant, we can make the approximation

2300

2240

[20]

2000

DIC (μmol kg–1) Figure 5 DIC versus TA for the formation and dissolution of CaCO3 in seawater. Contours are constant pCO2. Formation of CaCO3 from its dissolved constituents causes a decrease in TA and DIC in the ratio 2 : 1 and a corresponding increase in pCO2. The supply of CaCO3 to the oceans from weathering (or its burial as sediments) also increases (or decreases) TA and DIC in the ratio 2 : 1.

[23]

on the particular mineral phase of The value of interest and varies with temperature, salinity, and pressure such that solubility increases with decreasing temperature and increasing pressure (Figure 6; Table 2). For our concerns, pressure has the most important influence given the great depth of the ocean. The pressure dependence of carbonate solubility relates to the difference in volume occupied by the solid phase of CaCO3 compared with the dissolved ions. The change in partial molar volume, DV, between solid and dissolved phases, where

PALEOCEANOGRAPHY, PHYSICAL AND CHEMICAL PROXIES | Dissolution of Deep-Sea Carbonates

Kinetics

0 A

Water depth (m)

1000 Saturation calcite Saturation aragonite Atlantic Pacific

2000

If solid CaCO3 is in contact with undersaturated water, it will tend to dissolve. The rate of dissolution ultimately depends on the degree of undersaturation and therefore both [CO32
] and * Ksp need to be known accurately. The most commonly employed form of kinetic rate equation for the dissolution of CaCO3 (Morse and Berner, 1972) is given by R ¼ kð1
OÞn

B

C 4000 D 80

120

160

200

[CO2− 3 ] Figure 6 Water-column profiles of in situ [CO2
3 ] from the Atlantic and Pacific Oceans plotted with the saturation [CO2
3 ] profiles for aragonite and calcite. Saturation horizons for aragonite and calcite in the Pacific and Atlantic are labeled A–D, respectively. In situ data are from WOCE (sections A05 and P15N, available from http://cdiac.ornl.gov). Solubility coefficients are from Mucci (1983). Pressure correction from Millero (1995).

D V ¼ V Ca þ VCO3
VCaCO3

[25]


1

3000

5000 40

863

[24]

is negative for calcite and aragonite, that is, the volume occupied by Ca2þ and CO32
is smaller than when they are combined to form CaCO3. This explains why solubility increases with increasing water depth. Aragonite is more soluble than calcite for any given T, S, and P. It is for this reason that the saturation horizon for aragonite is shallower than for calcite (Figure 6, Table 2). Experimental determinations of KSP* at various T and S and 1 atm pressure suggest an uncertainty of  5% for both calcite and aragonite (Mucci, 1983). Difficulties in determining accurate values for DV (a range of  35–45 cm3 mol
1 for calcite, e.g., Ingle (1975)) lead to a total uncertainty in KSP* of  11% at 4 km water depth. Accordingly, our ability to define the vertical profiles of [CO32
]sat (Figure 6) are similarly uncertain. Including the error associated with determining in situ [CO32
], our uncertainty in calculating O becomes 12%. In all, this means that we cannot define the depth of the saturation horizon in the oceans to better than 0.90 km (see Emerson and Hedges (2003) for a fuller description). From purely thermodynamic considerations, we should expect the CaCO3 content of marine sediments to drop from a maximum value to zero precisely where the water column becomes undersaturated. But as stated earlier, the actual transition from carbonate-rich to carbonate-poor sediments occurs over a zone spanning several hundred meters with undissolved carbonate being found considerably deeper than the watercolumn saturation horizon. It is therefore necessary to consider the rate of CaCO3 dissolution.

where R is the rate (usually % day ), k is the rate constant (same units as R), and n is the reaction order. Determination of k and n has proved to be extremely nontrivial due to the difficulties in extrapolating between laboratory results and the real ocean and the fact that CaCO3 dissolution in seawater is the sum of several different chemical reaction pathways and involves the effects of impurities and surface coatings as well as multiple carbonate phases. A notable uncertainty in our understanding of carbonate dissolution kinetics is highlighted by the disparity between dissolution rates obtained in the laboratory with those predicted from field observations (which are smaller by 2–3 orders of magnitude). The most extensive laboratory-based measurements of carbonate dissolution were made by Keir (1980) using a variety of biogenic and synthetic carbonates at near-equilibrium conditions. Keir obtained a reaction order of n ¼ 4.5 for calcite and a very wide range of rate constants depending on the form of calcite tested, in many cases k being inversely correlated with grain size. Later work by Hales and Emerson (1997) suggested that uncertainties in Keir’s calculation of O had led to a particularly high rate order. They reinterpreted his results in terms of first-order kinetics. Using a value of n  1–2 reduced the associated range of rate constants and reduced discrepancies between various field observations. However, differences in rate constants determined in the laboratory compared with in situ field experiments remain unexplained.

Shape of the Sedimentary Lysocline The transition from carbonate-rich to carbonate-poor sediments as illustrated in Figure 2 is the ultimate expression of carbonate dissolution in the deep sea. The shape and position of the transition zone is sensitive to deep-water carbonate chemistry, the flux of CaCO3 and refractory material, and also the flux of organic carbon and its oxidation to CO2 (see also Archer (1991)). Interbasin differences in the depth of the lysocline are mainly due to deep-water chemistry variations. For example, dissolution occurs at much shallower depths in the Pacific than in the Atlantic Ocean (Figures 2 and 6). This is a result of mixing water masses with differing [CO32
] (North Atlantic Deep Water, NADW, has relatively high [CO32
] compared with deep water produced in the Southern Ocean) as well as the constant addition of CO2 (and corresponding decrease in [CO32
]) to deep waters by the oxidation of raining organic carbon as they progress along the ocean conveyor. The fact that carbonate may be preserved below the depth at which the water column becomes undersaturated (i.e., where O < 1) implies that dissolution is sufficiently slow (relative to the flux arriving at the sea floor) to allow burial. Also, diffusion of ocean water through the sediments must be slow enough to

864

PALEOCEANOGRAPHY, PHYSICAL AND CHEMICAL PROXIES | Dissolution of Deep-Sea Carbonates

allow pore waters to reach equilibrium saturation (by addition of CO32
through dissolution). Higher fluxes of CaCO3 to the sediment will tend to deepen the lysocline relative to the saturation horizon. This is apparent under certain upwelling areas (Figure 1). In regions where productivity is low, the CCD will tend to approach the saturation horizon. The proportion of noncarbonate material (e.g., clays) will also affect the shape of the transition zone. For example, the generally thicker lysocline of the Atlantic relative to the Pacific Ocean is thought to be a consequence of a higher flux of terrigenous material to Atlantic sediments (Archer, 1996). The effect is related to the sensitivity of %CaCO3 to dissolution (see the Section ‘Quantifying the Extent of Dissolution’); a lower initial CaCO3 content will mean greater sensitivity of % CaCO3 to a given degree of dissolution (Archer, 1991).

Supralysoclinal Dissolution Besides carbonate preservation below the saturation horizon, significant dissolution may occur above this depth. This phenomenon is the result of pore waters becoming undersaturated with respect to CaCO3 through the addition of metabolic CO2 by oxidation of organic carbon within the sediment (cf. eqn [13]). An early modeling study by Emerson and Bender (1981) suggested that release of respiratory CO2 within pore waters could cause up to 50% dissolution of CaCO3 even at the saturation horizon although the actual number is probably somewhat lower, perhaps 20–40%. In situ microelectrode measurements of O2 and pH confirm that dissolution of CaCO3 within pore waters must be occurring in response to the release of CO2 (e.g., Hales and Emerson, (1997); Figure 7). Further studies, using in situ benthic flux measurements have generally confirmed this conclusion although supersaturated settings Oxygen (μmol kg–1) 30 50 70 90 110130 −1

with a high initial %CaCO3 have proved to be equivocal. R. A. Jahnke and D. B. Jahnke (2004) have suggested that the low alkalinity fluxes observed at these sites may be the result of inorganic CaCO3 precipitation in the supersaturated sediment surface, followed by CO2-driven redissolution deeper in the sediments. Hales (2003) argues that the observations can be reconciled using revised estimates for the rates of respiration and dissolution. A consequence of pore-water dissolution is that loss of CaCO3 will tend to be more intense under highly productive regions (such as continental margins) for a given bottom water saturation. This is somewhat opposed to the notion that increased CaCO3 production will tend to deepen the lysocline. However, Archer (1991) has shown that it is the ratio of Corg to CaCO3 reaching the sediments (the benthic rain ratio) which makes the most difference to the proportion of CaCO3 dissolved. Increasing the benthic rain ratio causes a shoaling of the sedimentary lysocline, whereas increasing the absolute flux of CaCO3 will generally cause a deepening of the lysocline (Figure 8).

Quantifying the Extent of Dissolution Dissolution Proxies The importance of deep-sea carbonate dissolution within the global carbon cycle is reflected by the enumerable attempts to reconstruct the history of dissolution on various timescales. Furthermore, because dissolution can affect other paleoceanographic proxies employed to reconstruct climatic variations, it is evermore important to be able to quantitatively assess the extent of dissolution at any given location at any given time. At this time there is no perfect dissolution proxy; each suffers from

ΔpH −0.04 −0.02

−0.1

0

0

0

0 1 1

No dissolution

2

2

Depth (cm)

3 Depth (cm)

ΔpH −0.05

4 5 6

3 4 5

7

6

8 7

9 (a)

(b)

(c)

Figure 7 Pore-water profiles of (a) oxygen; (b) DpH; (c) modeled DpH from 2,235 m water depth in the west equatorial Pacific (Ontong Java Plateau) from Hales and Emerson (1996). Measurements were made in situ using microelectrodes. Decreasing O2 reflects the oxidation of organic matter and the production of respiration CO2. Heavy curves in (c) were computed assuming that respiration CO2 does not cause dissolution of carbonate. Light curve is a best fit to the data. The observed pH values demand that dissolution occurs. Reprinted from Hales B and Emerson S (1996) Calcite dissolution in sediments of the Ontong-Java Plateau: In situ measurements of pore water O2 and pH. Global Biogeochemical Cycle, 10: 527–541.

PALEOCEANOGRAPHY, PHYSICAL AND CHEMICAL PROXIES | Dissolution of Deep-Sea Carbonates

865

Fraction of CaCO3 in sediment 0

0.9 0

0.9 0

0.9 0

0.9 0

0.9

40

ΔCO3

org = 0.0 cal

org = 1.5 cal

org = 1.0 cal

org = 0.67 cal

org = 0.5 cal

6 12

0

6

18

12

–40

30

24 30

18 24

30

18

12 18 24

12

24

6

6 6

36 42

36

54

60 60

60

Figure 8 Results from Archer’s (1991) sedimentary lysocline model. Curve numbers represent various CaCO3 fluxes (micro-mol.cm
2.yr
1). Increasing the benthic rain ration (Corg/CaCO3) causes a shoaling of the lysocline, whereas increasing the absolute flux of CaCO3 generally causes a deepening (these results are from Archer’s oxic/anoxic model. Note results are plotted versus D[CO2
3 ] (cf. eqn [23]). Reprinted from Archer D (1991) Modeling the calcite lysocline. Journal of Geophysical Research-Oceans 96: 17037–17050.

particular difficulties, most commonly a sensitivity to changes in environmental conditions at the sea surface. It is wise therefore to consider more than one proxy at any given location. A further complication involves the offset in [CO32
] between bottom waters and pore waters. For example, in sediments above the water-column saturation horizon, dissolution can only occur under the influence of undersaturated pore waters. To infer changes in bottom-water [CO32
] would therefore require, at the very least, a knowledge of the contemporary benthic flux ratio of Corg/CaCO3. (see Carbon Cycle Proxies (d11B, d13Ccalcite, d13Corganic, Shell Weights, B/Ca, U/Ca, Zn/Ca, Ba/Ca)).

other workers have considered the ratio of benthic to planktonic Foraminifera (benthic or bottom-dwelling species tending to be more resistant to dissolution) or pteropods to Foraminifera as well as coccolith assemblages. A major complication to this approach is the environmental control on initial species composition. Planktonic foraminifera species assemblages are very sensitive to the prevailing environmental conditions at the sea surface and are in fact widely used to reconstruct changes in, for example, sea-surface temperatures (SSTs) through time. This complication makes interpretation in terms of dissolution subject to significant uncertainty.

Fragmentation indices Carbonate content (%CaCO3) The most apparent effect of dissolution in deep-sea sediments is the decrease in carbonate content with increasing water depth (Figure 2). However, the relationship between %CaCO3 and the extent of dissolution is not simple and as a result %CaCO3 is generally not considered a reliable indicator of dissolution. As described above, the carbonate content of a particular sediment is a function of the production and dissolution of CaCO3 and dilution by noncarbonate material. Additionally, %CaCO3 is fairly insensitive to the degree of dissolution until this becomes significant. For example, dissolving 50% of the carbonate from a sediment originally containing 90% CaCO3 would only decrease this value by 8% (Figure 9).

Changes in species composition Dissolution causes the thinning and breakup of Foraminifera tests and coccoliths. Different species are more or less susceptible to the effects of dissolution depending on the initial thickness of their tests as well as more cryptic differences such as crystal habit or chemical composition. These differences lead to changes in the species composition (faunal or floral assemblage) as dissolution proceeds and can be correlated with the extent of dissolution. Many variations of this type of proxy have been developed. As well as Berger’s (1970) classic solubility ranking of planktonic Foraminifera species (Table 3),

The extent of breakup or fragmentation of Foraminifera tests during dissolution can be used to assess the overall preservation state of the sediment. This approach has been applied extensively to reconstruct dissolution intensity in a wide variety of settings. A potential problem with fragmentation indices is their nonlinearity; since a single shell can break into a number of fragments, a simple ratio of fragments to total entities tends to be oversensitive to changes at the onset of dissolution and much less sensitive as dissolution proceeds (Figure 10). This matter was addressed by Le and Shackleton (1992), who suggested taking the average number of fragments per whole shell into account. %Fragment ¼ 100  ð#fragments=xÞ= fð#fragments=xÞ þ #whole shellsg

[26]

where x is the average number of fragments that a typical Foraminifera shell will break into (Figure 10). Another variant is to compare the number of fragments to whole shells of a single species (e.g., Mekik et al. (2002)).

Grain-size indices Grain-size indices are based on the fact that progressive dissolution causes a decrease in the average grain size of a given sediment packet as entities such as Foraminifera tests break up. A typical example of this type of indicator is the proportion of coarse fraction to total sediment (e.g., % >150 mm). Grain-size

866

PALEOCEANOGRAPHY, PHYSICAL AND CHEMICAL PROXIES | Dissolution of Deep-Sea Carbonates

ΔCO32− = +5

9

ΔCO2− 3 = –3

9

ΔCO2− 3 = –6

8.7

9

μmol 8 kg

9

9 I 90% Calcite

μmol kg

I

I

0.3 I 23% Calcite

I

100

4 I 80% Calcite 1 I 50% Calcite

Calcite content of sed. (%)

ΔCO3– = –9

I

μmol kg

5

Lysocline

μmol kg

80 60 40 20 0

0

20 40 60 80 Calcite dissolved (%)

100

Figure 9 Illustration of the insensitivity of %CaCO3 to dissolution until the fraction dissolved exceeds  50%. Note in this case initial carbonate content is 90%. A lower initial value would result in increased sensitivity to dissolution. Straight arrows represent fluxes of carbonate and noncarbonate to the sediment and burial after dissolution (indicated by wavy arrows). Reprinted from Broecker WS and Peng T-H (1982) Tracers in the Sea. Palisades: Eldigio Press.

Table 3 Berger’s (1970) solubility index of planktonic Foraminifera (rank 1 is most soluble) Low resistance

High resistance

1. Globigerinoides ruber 2. Orbulina universa 3. Globigerinella siphonifera 4. Globigerina rubescens 5. Globigerinoides sacculifer 6. Globigerinoides tenellus 7. Globigerinoides conglobatus 8. Globigerina bulloides 9. Globigerina quinqueloba 10. Globigerinita glutinata 11. Candela nitida

12. Globorotalia hirsuta 13. Globorotalia truncatulinoides 14. Globorotalia inflata 15. Globorotalia cultrata 16. Globorotalia dutertrei 17. Globigerina pachyderma, s.l. 18. Pulleniatina obliquiloculata 19. Globorotalia crassiformis 20. Sphaeroidinella dehiscens 21. Globorotalia tumida 22. Turborotalia humulis

100

From Berger WH (1970) Planktonic foraminifera: Selective solution and the lysocline. Marine Geology 8: 111–138.

% Fragments

80

60

40

20

0 1000

800

600

400

200

0

# Whole shells

indices are sensitive to the initial particle size distribution, for example, coccoliths versus Foraminifera, and are therefore subject to environmental change. They may also suffer from a lack of sensitivity at the onset of dissolution.

Figure 10 Dissolution causes progressive fragmentation of foraminiferal shells. Each shell will produce several fragments, leading to nonlinearity of a straightforward measure of %fragmentation (red curve). Le and Shackleton (1992) accounted for this by including a term for average number of fragments per whole shell into their fragmentation index (cf. Eqn [26], blue curve).

Foraminiferal shell weights The individual weights of Foraminifera tests within a narrow size range have been correlated to bottom water [CO32
] and used to infer changes in this parameter (related to CaCO3 dissolution, cf. Eqn [21]) through time (e.g., Broecker and Clark (2001)). Again, this proxy is subject to variations in initial shell weight, which may result from changing conditions during growth (e.g., Barker and Elderfield (2002)).

Observational indices Changes in the ultrastructure of Foraminifera tests, caused by progressive dissolution, can be observed using scanning electron microscopy (SEM) (Figure 11). Various studies have employed this technique in an attempt to reconstruct the history of dissolution in different locations although difficulties

PALEOCEANOGRAPHY, PHYSICAL AND CHEMICAL PROXIES | Dissolution of Deep-Sea Carbonates

R1-1

R1-2

R1-3

R1-4

R2-1

R2-2

R2-3

R2-4

R3-1

R3-2

R3-3

R3-4

867

Figure 11 Scanning electron microscopy (SEM) can be used to assess the preservation state of carbonate sediments. In this case, numbers 1–4 show progressive dissolution of tests of Globigerina bulloides. Reprinted from Dittert et al. (2000).

may arise from changes in initial shell-wall structure as well as the inherent subjectivity of the method (this is also a problem with many of the other proxies discussed here).

Crystallinity A relatively unexplored indicator of carbonate dissolution is the so-called crystallinity of foraminiferal calcite. Crystallinity, defined as the peak width (at half maximum height) of the (104) calcite X-ray diffraction peak, is essentially a measure of how perfect the calcite crystal lattice is. As dissolution proceeds, more poorly crystallized calcite is thought to be removed, causing a narrowing of the (104) diffraction peak. Crystallinity has been applied to core-top sediments and may, in the future, prove to be a reliable dissolution indicator, providing environmental conditions do not play a major role (Bassinot et al., 2004).

Composite indices In order to circumvent the various difficulties associated with many of the individual proxies for carbonate dissolution, several workers have adopted so-called composite dissolution indices (CDIs). Peterson and Prell (1985) used a combination of six different measurements: %CaCO3, a size index (>63 mm), three fragmentation indices (% whole Foraminifera, ratio of benthic to planktonic Foraminifera, and % whole Globorotalia menardii) and % radiolarians (ratio of radiolarians to radiolarians þ whole planktonic Foraminifera).

Chemical indices As well as indicators based on the physical effects of dissolution, various chemical proxies have been developed with the aim of directly reconstructing deep-sea carbonate system

parameters. For example, elemental ratios, such as Zn/Ca and Cd/Ca measured on benthic Foraminifera, have been applied to assess changes in deep ocean [CO32
] (e.g., Marchitto et al. (2005)) whereas the boron isotopic composition (d11B) of the same organisms has the potential for quantifying changes in surface and deep-water pH (e.g., Sanyal et al. (1995)).

Effect of Dissolution on Other Paleoceanographic Proxies Many of the proxies employed by paleoceanographers utilize foraminiferal calcite. It has been widely shown that many of these proxies are actually susceptible to the effects of dissolution. For example, planktonic foraminiferal Mg/Ca ratios (used as a proxy for temperature) are known to decrease with progressive dissolution of the test. The actual mechanism is poorly understood but probably reflects the chemical heterogeneity of foraminiferal calcite. The effects of dissolution on Mg/Caderived temperatures can be significant and as such it is important to assess the state of sediment preservation before interpretation of results. Other foram-based proxies that may be affected by dissolution include d11B, d18O, and faunal assemblages (see above).

Changes in Carbonate Dissolution through Time From the relations described in Sections ‘Thermodynamics of the carbonate system in seawater’ and ‘Mechanisms of Carbonate Dissolution,’ it is possible to understand the importance of carbonate dissolution within the broader context of the global carbon cycle, particularly with respect to atmospheric CO2. Our understanding of atmospheric CO2 changes

868

PALEOCEANOGRAPHY, PHYSICAL AND CHEMICAL PROXIES | Dissolution of Deep-Sea Carbonates

in the past as well as the consequences of future changes depends on our understanding of their interaction with marine carbonate chemistry and CaCO3 dissolution.

Carbonate Budgets and Carbonate Compensation The marine carbonate cycle involves the balance between the production of CaCO3 (mainly through biogenic processes) from its dissolved constituents, the delivery of those dissolved ingredients to the ocean from continental weathering or hydrothermal exchange, and the burial of solid CaCO3 as sediment or reefs. The actual rates of carbonate production, burial, and the influx from weathering are sources of significant uncertainty, as discussed by Milliman and Droxler (1996) (Table 4). However, it is clear that the global production rate of carbonate is significantly greater than the rate of supply from weathering (by a factor of  4 or more). This overproduction of carbonate is (approximately) balanced by dissolution; the depth of the lysocline is such that it intersects and dissolves approximately the correct proportion of sinking CaCO3 to maintain the balance between supply and burial. As mentioned in the Section ‘Thermodynamics of the Carbonate System in Seawater,’ CaCO3 weathering (or burial) adds (or removes) alkalinity and DIC in a ratio of 2 : 1. From Eqn [18], we see, for example, that an increase in CaCO3 weathering will cause an increase in [CO32
] and deepening of the lysocline such that a smaller proportion of raining CaCO3 will be dissolved, thereby regaining the balance between supply and burial. This process is known as carbonate compensation and has a timescale of adjustment of the order 4–5 ka (Broecker and Peng, 1982). Thus, changes in CaCO3 dissolution through time may be linked with changes in weathering, the rate of carbonate production, or changes in the locus of CaCO3 deposition (see below). (The present discussion concerns timescales of thousands of years. On longer (Myr) timescales, the control of weathering on CO2 is rather different, being dependent on the rate of silicate weathering alone and not on the imbalance between weathering and burial rates (see Berner (2004) for a full description).) Importantly, these processes are also linked to atmospheric CO2 by the inverse relation between CO2 and [CO32
] (Eqn [12]).

Table 4 Global estimates for the rates of supply, production, and burial of CaCO3(units are 1012 mol CaCO3 per year). Uncertainties are probably less than a factor of 1 for most values

Total

Inputs

Production

Outputs (burial)

Rivers (13)

Coral reefs, shelves (25) Pelagic (inc. slopes) (65)

Neritic (15) Deep-sea (17)

90

32

Hydrothermal (3–12) Groundwater ( 5) 21–30

Data compiled from Milliman JD and Droxler AW (1996) Neritic and pelagic carbonate sedimentation in the marine environment: Ignorance is not bliss. Geolgische Rundschau 85: 496–504.

Glacial–Interglacial Changes in Atmospheric CO2 Ice-core records tell us that atmospheric CO2 was approximately 30% lower during the Last Glacial Maximum (LGM) than the preindustrial Holocene (see CO2 Studies). They also tell us that CO2 has varied in tandem with the late Pleistocene glacial cycles, each cycle comprising a stepwise decline from an interglacial high of about 280 ppmv to a glacial low of about 190 ppmv followed by a more rapid increase to interglacial values during glacial terminations (Petit et al., 1999). The reason for lowered atmospheric CO2 during glacial times has been debated for more than quarter of a century and is still the subject of enquiry. However, since the oceans contain some 50–60 times more carbon than the atmosphere (see the Section ‘Thermodynamics of the Carbonate System in Seawater’), it is assumed that they must play a major role in the drawdown of CO2 during glacial times. Currently favored mechanisms involve variations in vertical stratification and air–sea exchange in the Southern Ocean, enhanced biological activity under the influence of Fe fertilization, and changes in the dominant plankton type (see Sigman and Boyle (2000) for a review and additional references). However, any mechanism will ultimately involve interaction with the marine carbonate cycle. A few examples of mechanisms involving changing carbonate dissolution patterns are described below. Records of %CaCO3 from the deep Pacific Ocean show cycles of increased glacial CaCO3 content and decreased % CaCO3 during interglacials (Figure 12). There has been debate over whether the Pacific CaCO3 cycles represent changes in dissolution intensity or surface ocean CaCO3 production although dissolution is the generally favored explanation. The observation of a deeper calcite lysocline during glacial times (Figure 12) suggests that deep ocean [CO32
] was higher and could help to explain the decrease in CO2. In an attempt to explain the glacial decrease in CO2, Berger (1982) suggested that rising sea levels associated with glacial terminations would increase the accommodation space for coral reef communities. Since coral reefs and shallow carbonate deposition account for almost half of all modern CaCO3 burial (Table 4; Milliman and Droxler, 1996), an increase in this mode of deposition at the end of glacial times would cause global CaCO3 burial to exceed supply, leading to a decrease in whole ocean [CO32
] and an increase in atmospheric CO2. This would in turn lead to a shoaling of the lysocline through compensation and a shift in the locus of carbonate deposition from the deep sea to more shallow waters. However, the mean change in lysocline depth for the glacial period is not sufficient to explain the 30% reduction in CO2 via this mechanism. CaCO3 dissolution actually increased in the deep Atlantic as a consequence of low [CO32
] Antarctic Bottom Water (AABW) invading the glacial deep Atlantic (Barker et al., 2004). In fact, an estimate of global deep-sea CaCO3 accumulation during the last glacial period suggests that it was probably very similar to the modern day (Catubig et al., 1998). Furthermore, since the deglacial rise in CO2 preceded the rise in sea level, it cannot have been driven mainly by changes in shallow-water accommodation space. Another suggestion was made by Archer and Maier-Reimer (1994). This involved the influence of pore-water CaCO3 dissolution and the benthic rain ratio of Corg to CaCO3 (see

PALEOCEANOGRAPHY, PHYSICAL AND CHEMICAL PROXIES | Dissolution of Deep-Sea Carbonates

0–20% Isotope

1

CaCO3

3

20–40%

5 B 3

90% isopleth

Weight percent CaCO3in sediment

10% isopleth, CCrD

7

40–60% Brunhes 9 11

B 5

B 7

0.2

0.3

B 9

60–80% 13

B 11

869

80–100%

15

17

B 13

B 15

21 B 17 M 1

Increasing dissolution

Water depth (km)

4.2

4.4

4.6

4.8

5.0 0.0

0.1

0.4

0.5

0.6

0.7

0.8

Myr BP Figure 12 Reconstructed sedimentary lysocline in the equatorial Pacific made by combining several records of %CaCO3 from different water depths (Farrell and Prell, 1989). Glacial periods are characterized by a deepened and narrower transition zone suggesting better carbonate preservation. Reprinted from Farrell BF and Prell W (1989) Climatic change and CaCO3 preservation; an 800,000 year bathymetric reconstruction from the Central Equatorial Pacific Ocean. Paleoceanography 4: 447–466.

the Section ‘Shape of the Sedimentary Lysocline’). By increasing the production ratio of organic carbon to inorganic carbonate (e.g., by increasing the ratio of diatom to coccolithophore production), Archer and Maier-Reimer argued that increased supralysoclinal dissolution would cause carbonate supply to exceed burial, thereby raising [CO32
] and lowering CO2. In this scenario, the depth of the sedimentary lysocline would not need to change by very much since it would shoal relative to a deepened saturation horizon. Although the ‘rain-ratio’ hypothesis escapes some of the problems associated with the coral reef hypothesis, it has not yet passed the rigors of validation (e.g., there are difficulties associated with the width of the predicted transition zone (Sigman et al., 1998)). Thus the search for a viable explanation for the glacial lowering of atmospheric CO2 continues. In all likelihood, it will comprise a combination of several mechanisms rather than a single process.

inventory of fossil fuel reserves of  5000 Gt C (Sundquist, 1985). Modeling studies (Archer et al., 1997, 1998) suggest that on timescales of hundreds of years, 70–80% of anthropogenic CO2 will dissolve into the oceans. Neutralization by CaCO3 dissolution will occur on timescales of thousands of years and will eventually account for 60–70% of total CO2 emissions.

See also: Ice Core Methods: CO2 Studies; Overview. Paleoceanography: Paleoceanography An Overview. Paleoceanography, Biological Proxies: Benthic Foraminifera; Coccolithophores; Planktic Foraminifera. Paleoceanography, Physical and Chemical Proxies: Carbon Cycle Proxies (d11B, d13Ccalcite, d13Corganic, Shell Weights, B/Ca, U/Ca, Zn/Ca, Ba/Ca).

References Carbonate Dissolution in the Future: Sequestration of Anthropogenic CO2 Atmospheric CO2 has increased by around 90 ppmv (30%) since the start of the industrial revolution. However, this increase reflects only about half of the total CO2 added; most of the rest has been taken up by the oceans. The oceans provide a buffer to increasing CO2 by reaction with CO32
, thereby lowering the saturation state. This has consequences for marine life in the surface ocean but will also result in deep-sea CaCO3 dissolution. In fact, much of the CO2 released by fossil fuel burning will ultimately be neutralized by CaCO3 dissolution in the deep sea. Estimates of the inventory of deep-sea CaCO3 available for CO2 neutralization are of the order 2000–5000 Gt C (Archer et al., 1997). These can be compared to a global

Archer D (1991) Modeling the calcite lysocline. Journal of Geophysical ResearchOceans 96: 17037–17050. Archer D (1996) A data-driven model of the global calcite lysocline. Global Biogeochemical Cycle 10: 511–526. Archer D, Kheshgi H, and Maier-Reimer E (1997) Multiple timescales for neutralization of fossil fuel CO2. Geophysical Research Letter 24: 405–408. Archer D, Kheshgi H, and Maier-Reimer E (1998) Dynamics of fossil fuel CO2 neutralization by marine CaCO3. Global Biogeochemical Cycle 12: 259–276. Archer D and Maier-Reimer E (1994) Effect of deep-sea sedimentary calcite preservation on atmospheric CO2 concentration. Nature 367: 260–263. Barker S and Elderfield H (2002) Foraminiferal calcification response to glacial– interglacial changes in atmospheric CO2. Science 297: 833–836. Barker S, Kiefer T, and Elderfield H (2004) Temporal changes in North Atlantic circulation constrained by planktonic foraminiferal shell weights. Paleoceanography 19: http://dx.doi.org/10.1029/2004PA001004. Bassinot F, Melieres F, Gehlen M, Levi C, and Labeyrie L (2004) Crystallinity of foraminifer shells: A proxy to reconstruct past bottom water CO3- changes? Geochemistry Geophysics Geosystems 5: http://dx.doi.org/10.1029/ 2003GC000668.

870

PALEOCEANOGRAPHY, PHYSICAL AND CHEMICAL PROXIES | Dissolution of Deep-Sea Carbonates

Berger WH (1970) Planktonic foraminifera: Selective solution and the lysocline. Marine Geology 8: 111–138. Berger WH (1982) Increase of carbon-dioxide in the atmosphere during Deglaciation – the coral-reef hypothesis. Naturwissenschaften 69: 87–88. Berner A (2004) The Phanerozoic carbon cycle: CO2 and O2, p. 158. New York: Oxford University Press. Broecker WS and Clark E (2001) Glacial to Holocene redistribution of carbonate ion in the deep sea. Science 294: 2152–2155. Broecker WS and Peng T-H (1982) Tracers in the Sea. Palisades: Eldigio Press. Catubig NR, Archer DE, Francois R, deMenocal P, Howard W, and Yu EF (1998) Global deep-sea burial rate of calcium carbonate during the last glacial maximum. Paleoceanography 13: 298–310. Dittert N and Henrich R (2000) Carbonate dissolution in the South Atlantic Ocean: Evidence from ultrastructure breakdown in Globigerina bulloides. Deep-Sea Research Part I 47: 603–620. DOE (1994) Dickson AG and Goyet C (eds.) In: Handbook of Methods for the Analysis of the Various Parameters of the Carbon dioxide System in Sea Water, version 2, (ORNL/CDIAC-74). Emerson S and Bender M (1981) Carbon fluxes at the sediment-water interface of the deep sea: Calcium carbonate preservation. Journal Marine Research 39: 139–162. Emerson S and Hedges JI (2003) Sediment diagenesis and benthic flux. In: Elderfield H (ed.) Treatise on Geochemistry: The Oceans and Marine Geochemistry, vol. 6, pp. 293–319. Amsterdam: Elsevier. Farrell BF and Prell W (1989) Climatic change and CaCO3 preservation; an 800,000 year bathymetric reconstruction from the Central Equatorial Pacific Ocean. Paleoceanography 4: 447–466. Hales B (2003) Respiration, dissolution, and the Iysocline. Paleoceanography 18. Hales B and Emerson S (1996) Calcite dissolution in sediments of the Ontong-Java Plateau: In situ measurements of pore water O-2 and pH. Global Biogeochemical Cycle 10: 527–541. Hales B and Emerson S (1997) Evidence in support of first-order dissolution kinetics of calcite in seawater. Earth and Planetary Science Letter 148: 317–327. Ingle SE (1975) Solubility of calcite in the ocean. Marine Chemistry 3: 301–319. Jahnke RA and Jahnke DB (2004) Calcium carbonate dissolution in deep sea sediments: Reconciling microelectrode, pore water and benthic flux chamber results. Geochimica et Cosmochimica Acta 68: 47–59. Keir RS (1980) Dissolution kinetics of biogenic calcium carbonates in seawater. Geochimica et Cosmochimica Acta 44: 241–252. Le J and Shackleton NJ (1992) Carbonate dissolution fluctuations in the western Equatorial Pacific during the late Quaternary. Paleoceanography 7: 21–42. Marchitto TM, Lynch-Stieglitz J, and Hemming SR (2005) Deep Pacific CaCO3 compensation and glacial–interglacial atmospheric CO2. Earth and Planetary Science Letter 231: 317–336.

Mekik FA, Loubere PW, and Archer DE (2002) Organic carbon flux and organic carbon to calcite flux ratio recorded in deep-sea carbonates: Demonstration and a new proxy. Global Biogeochemical Cycle 16: (art. no.-1052). Millero FJ (1995) Thermodynamics of the carbon dioxide system in the oceans. Geochimica et Cosmochimica Acta 59: 661–677. Milliman JD and Droxler AW (1996) Neritic and pelagic carbonate sedimentation in the marine environment: Ignorance is not bliss. Geolgische Rundschau 85: 496–504. Morse JW and Berner RA (1972) Dissolution kinetics of calcium-carbonate in sea-water 2: Kinetic origin for Lysocline. American Journal Science 272: 840. Morse JW and He SL (1993) Influences of T, S and P(Co2) on the pseudohomogeneous precipitation of Caco3 from seawater-implications for Whiting formation. Marine Chemistry 41: 291–297. Mucci A (1983) The solubility of Calcite and Aragonite in seawater at various salinities, temperatures, and one atmosphere total pressure. American Journal of Science 283: 780–799. Peterson LC and Prell WL (1985) Carbonate preservation and rates of climatic change: An 800 kyr record from the Indian Ocean. In: Sundquist E and Broecker WS (eds.) The Carbon Cycle and Atmospheric CO2: Natural Variations Archean to Present. Geophysical Monograph Series, vol. 32, pp. 251–269. Washington, DC: AGU. Petit JR, Jouzel J, Raynaud D, et al. (1999) Climate and atmospheric history of the past 420,000 years from the Vostok ice core Antarctica. Nature 399: 429–436. Sanyal A, Hemming NG, Hanson GN, and Broecker WS (1995) Evidence for a higher pH in the Glacial ocean from boron isotopes in foraminifera. Nature 373: 234–236. Sigman DM and Boyle EA (2000) Glacial/interglacial variations in atmospheric carbon dioxide. Nature 407: 859–869. Sigman DM, McCorkle DC, and Martin WR (1998) The calcite lysocline as a constraint on glacial/interglacial low-latitude production changes. Global Biogeochemical Cycle 12: 409–427. Sundquist E (1985) Geological perspectives on carbon dioxide and the carbon cycle. In: Sundquist E and Broecker WS (eds.) The Carbon Cycle and Atmospheric CO2: Natural Variations, Archean to Present, pp. 5–59. Washington, D. C: AGU. Thurman HV and Trujillo AP (2002) Essentials of Oceanography, 7th ed. Upper Saddle River: Prentice-Hall. Zeebe RE and Wolf-Gladrow D (2001) CO2 in Seawater: Equilibrium, Kinetics, Isotopes. Amsterdam: Elsevier.

Relevant Website http://cdiac.ornl.gov – Carbon Dioxide Information Analysis Center.