N. pachyderma (sin.)
2. Proxies for ocean carbonate chemistry and their limitations 1 Carbonate chemistry
2.1.1 Ba/Ca to infer alkalinity
Ba is a nutrient-like tracer similar to Cd and δ13C, because biological activity extracts these elements from surface waters and gravitation transfers them toward the seafloor in sinking particles. On its way from the North Atlantic to the deep North Pacific, deep water is progressively enriched in Ba. The close correlation between Ba and alkalinity in seawater (Chan et al., 1977; Lea and Boyle, 1989) is mechanistically not well understood (Bishop, 1988; Chan et al., 1977; Chow and Goldberg, 1960; Lea, 1993), but is thought to be related to the simultaneous release of alkalinity through CaCO3 dissolution and regeneration of Ba at the seafloor. However, Lea (1993) suggested that changes in the thermohaline circulation redistribute Ba and alkalinity similarly, thereby allowing reconstruction of past alkalinity distributions from benthic foraminiferal Ba/Ca.
The main limitation of Ba as a paleoproxy is due to the short oceanic residence time on the order of 10,000 years (Broecker and Peng, 1982; Chan et al., 1977). The Ba-alkalinity correlation is not perfectly applicable on a time scale longer than this period. However, the fact that Ba is incorporated into foraminifera shells in direct proportion to the seawater concentration (Lea and Boyle, 1989; Lea and Spero, 1992; Lea and Spero, 1994) allows us to estimate paleo-Ba concentrations from foraminifera deposited in sediments. Independent estimates of seawater carbonate chemistry would offer an opportunity to verify whether the present-day slope of the Ba-alkalinity relationship is applicable to the past as well. A multiproxy approach would provide the best means of calculating alkalinity for various time
Table A1. Reconstructing past ocean carbonate chemistry: proxies, limitations and estimates. Please note that the table merely presents proxies that were discussed during the workshop. Proxy target parameter interfering parameters estimate for LGM limitations references Ba/Ca seawater [Ba2+ ] and alkalinity none on planktonics; pressure on benthic DBa
+20-25 ± 10 µmol kg-1 (Circumpolar Deep Water) short residence time, mechanistic linkage between [Ba2+ ] and alkalinity not understood
Lea and Boyle 1989; Lea 1993; Lea and Spero 1992, 1994 δ11 B pHspecies-specific, symbionts, maybe temperature and early diagenesis
+0.2 ± 0.1 units (surface water) +0.3 ± 0.1 units (deep ocean) material and time consuming, only mixed benthics
Sanyal et al. 1995, 1996, 2001 Carbonate ion effect on δ13 C [CO32- ]depth habitat (G. sacculifer and G. ruber)
+55 ± 63 µmol kg-1 (Indian Ocean) Spero et al. 1997,1999; Bijm 1999; Russell and Spero 2000 U/Ca [CO32- ]species-specific, temperature? +80-100 ± 60 µmol kg-1Mn contamination Russell et al. 1994; Russell, 2001 SO42- /CO32- [CO32- ]Mg, temperature equivalent to +0.2 pH units (deep ocean ) BaSO4 contamination Erez et al., 2001 Size- normalized weight
deep [CO32- ] dissolution offset between pore water and bottom water saturation, growth conditions no data local calibrations required, reproducibility of weight measurements: ± 6%
Lohmann 1995; Lohmann et al. 1999; Rosenthal et al. 2000; Broecker et al. 2001; Broecker and Clark 2001 size fraction index deep [CO32- ] dissolution no data (± 5 µmol kg-1 ) non calcite material trapped in shells, Corg rain, continuous breakup after burial?
Broecker and Clark, 1999 Reflectence/ lightness + weight
corrosion qualitative estimates only Helmke and Bauch, 2002
scales which can then be compared with the Ba concentrations recorded in foraminiferal shells.
2.1.2 δ11B to infer pH
Inferring seawater pH from stable boron isotopic compositions is based on the isotopic fractionation between dissolved boron in seawater and boron in CaCO3. The uncharged boron species B(OH)3 is enriched in 11B by ~20‰ over the charged species B(OH)4-. Of these two dominant aqueous species, it is the charged borate that is incorporated into carbonate minerals. As the fraction of B(OH)4- andB(OH)3 changes with pH, so must their respective isotopic compositions. The isotopic composition of boron in CaCO3 therefore is enriched with the heavier isotope 11B with increasing pH (Hemming and Hanson, 1992; Sanyal et al., 2000).
The boron isotopic composition in carbonates is also highly sensitive to local variations in pH. The calcification process itself and microenvironments (i.e. associations with symbionts
15 20 25 30 35 40
7.5 8 8.5 9 9.5
inorg. precipitation cultured O. universa B(OH)4
-Holocene G. sacculifer Holocene O. universa cultured G. sacculifer
δ11 B (‰)
pH
modern oceanic range
T = 25°C & S = 35‰
(pK = 8.60; α = 19.4‰)
Figure A1. Present state of the δ11B proxy calibration. Red: theoretical curve according to Kakihana et al. (1977); black: inorganic precipitation results (Sanyal et al., 2000); blue:
Orbulina universa from culture experiments (closed circles, Sanyal et al., 1996) and core-top sediment samples (open circles, Sanyal et al., 1997); green: Globigerinoides sacculifer from culture experiments (closed triangles, Sanyal et al., 2001) and core-top sediment samples (open triangles, Sanyal et al., 1995).
or precipitation within extrapallial fluids) must therefore be considered. While the planktonic foraminifer G. sacculifer and benthic foraminifera appear to incorporate δ11B with little or no fractionation compared to the theoretical curve, O. universa shows an offset from those foraminifera by ~3.3 ‰ (Sanyal et al., 2001). The offset was suggested to be due to a vital effect, although its nature could not be explained: both planktonic species are spinose and symbiont-bearing and should therefore react similarly.
The effect of symbiont photosynthesis has recently been investigated in a diffusion-reaction model by Zeebe at al. (subm.). They calculated a constant but significant offset between δ11B in planktonic foraminiferal calcite and the isotopic signature of B(OH)4- in the seawater medium. Recent laboratory culture data (Hönisch et al., subm.) are in good agreement with the model results.
Stable boron isotopic analyses, using negative thermal ionisation mass spectrometry (nTIMS), have several complications. First of all, the technique requires several hours of permanent operator assistance and numerous replicate analyses until accurate values can be obtained. Second, to achieve a reproducible result, approximately 4 ng boron are required per analysis. Since foraminifera contain 5-15 ppm B (Hemming et al., 1998), up to 10 mg foraminiferal calcite are needed per sample (when considering weight loss during cleaning and multiple replicate analyses). Especially for the investigation of deep water chemistry the second point is crucial, as the abundance of benthic foraminifera is too low to routinely allow single-species analyses. Sanyal et al. (1995) therefore combined several species for their deepwater record despite possible differences in habitat (epifaunal/infaunal) characterized by a range of pH conditions, and species-specific offsets like the ones found for planktonic foraminifera (Sanyal et al., 2001). These factors (Sanyal et al., 1997; Sanyal et al., 1996) may have biased the obtained value, which suggested a 0.3 pH units increase for last glacial deepwater (Sanyal et al., 1995).
In order to solve the problems named above, it is desirable to reduce the amount of material required for analyses, to speed up measurements and to generally expedite the analytical procedure so that the investigation of past ocean acidity can be realised extensively in the future.
2.1.3 Deconvolution of the carbonate ion effect to infer [CO32-]
This approach is based on the deconvolution of foraminiferal δ13C records to calculate the change in surface [CO32-] and δ13CΣCO2 through time: The stable carbon and oxygen isotopic compositions of planktonic foraminifera decrease with increasing carbonate ion concentration (Bijma et al., 1999; Spero et al., 1999). Among the investigated planktonic foraminifera, G. sacculifer and G. ruber share the same habitat but the slope in δ13C vs.
[CO32-] is twice as large in G. ruber as in G. sacculifer. This species-specific difference is used to distinguish between the effect of [CO32-] and a simultaneous change in δ13CΣCO2. Application to the sediment record leads to the estimate of +55 ± 63 µmol kg-1[CO32-] for the Indian Ocean during the last glacial (Spero et al., 1999). Unfortunately, this method is restricted to tropical surface waters, where G. sacculifer and G. ruber occur.
2.1.4 U/Ca to infer [CO32-]
Laboratory experiments revealed that U/Ca in planktonic foraminifera shells is inversely related to [CO32-] (Russell, 2001). The symbiont-barren G. bulloides incorporates approximately twice U/Ca than the symbiont-bearing O. universa at the same [CO32-]. No consistent temperature effect on the record has been found above 19°C. Application of the U/Ca relationship to Caribbean cores suggested that glacial [CO32-] was 80-100 ± 60 µmol kg
-1 higher than during the Holocene.
Although the approach is generally promising, the study of several sediment cores revealed that contamination by Mn-carbonates places a significant diagenetic overprint on the incorporated U/Ca which may limit the general applicability of this proxy to sediments above the redox front.
2.1.5 SO42-/ CO32- to infer [CO32-]
In laboratory culture experiments, Erez et al. (2001) observed a constant distribution coefficient between SO42-/CO32- in the shells of benthic and planktonic foraminifera and SO4
2-/CO32- ratio in seawater. Since the seawater SO42- inventory is not expected to have changed on glacial-interglacial time scales, [CO32-] can be reconstructed. In situ calibrations of this proxy in the Gulf of Eilat gave similar results to those of the laboratory experiments.
However, in the Little Bahama Bank a temperature effect was revealed which may have been caused by changes in the Mg ion content, apparently affecting the SO42- content of foraminiferal shells. Correction of this temperature effect leads to the empirical negative
correlation between seawater [CO32-] and foraminiferal SO42-/CO32- as found in laboratory culture experiments. Preliminary comparisons of SO42-/CO32- from Holocene and glacial benthic foraminifera show variability in pH similar in magnitude to that estimated independently from δ11B (approximately 0.2 pH units increase in the glacial deep Pacific).
The advantage of this proxy is the very small sample size required for routine measurements using a Magnetic Sector ICP-MS. Hence it is practical to be used for benthic foraminifera from deep sea sediments. However, the proxy is still under development and not much is known about its limitations. One possible interference may be the contamination with extraneous phases like barite (BaSO4).
2.2 Carbonate preservation 2.2.1 Size-normalized shell weight
The average mass of planktonic foraminifera is primarily determined by their size, but there is a measurable secondary relationship of shell mass to water depth (Lohmann, 1995;
Lohmann et al., 1999; Rosenthal et al., 2000). Due to dissolution, the size-normalized mass of nearly all species is lower in deeper water than it is in shallow water, and the decrease is continuous over a wide range of carbonate saturation states, even well above the calcite lysocline. Based on shells of the three species G. sacculifer, Pulleniatina obliquiloculata and Neogloboquadrina dutertrei, Broecker and Clark (2001a) determined an average weight-loss slope of 0.3 ± 0.05 µg (µmol kg-1)-1 decrease in pressure-corrected deep sea carbonate ion concentration. This relationship allows estimates of changes in seawater carbonate content from the size-normalized mass of planktonic foraminifera.
To use the relationship as a paleocarbonate ion proxy, this method requires that the offset between pore and bottom water saturation is constant. However, numerous investigations (Archer et al., 1989; Berelson et al., 1990; Berelson et al., 1994; Hales and Emerson, 1996; Hales and Emerson, 1997a; Jahnke et al., 1994; Jahnke et al., 1997) have applied microelectrodes and benthic flux chambers to validate the theory of respiration-driven dissolution in-situ (Table A2, see also section 2.3.1). They conclude that 40-60 % of the calcite dissolution above the saturation horizon can be attributed to metabolic processes. The amount of organic matter reaching the seafloor varies between sites and depends on depth.
Assuming increased productivity on glacial time scales, the magnitude of this effect might have been even stronger. Application of this proxy should therefore be restricted to locations where strong changes in paleoproductivity are not expected.
Table A2. In situ investigation of sedimentary carbonate dissolution. Reference Location Depth (m) ∆ (µmol l-1 ) or ΩδrMD1 (%) MD2 (% Hales et al. (1994) North Atlantic 2100-5400 Ω = -0.2 – +1.75 35 – 67 Jahnke et al. (1994) Eastern North Atlantic 3100 ∆ = +270no Archer et al. (1989) Equatorial Atlantic 3800-5000 ∆ = -27 – +9 some Hales and Emerson (1997) Western eq. Atlantic (Ceara Rise) 3300-4700 Ω = +0.8 – +1.2 > 20 Martin and Sayles (1996) Western eq. Atlantic (Ceara Rise) 3200-4700 ∆ = -28 – +13 0.8 – 1.4 36 – 6637 – 92 Berelson et al. (1990) Central eq. Pacific 4400-5000 ∆ = -6 – +12a no data Berelson et al. (1994) Central eq. Pacific 3380-4560 Ω = +0.78 – +0.87 no data60 – 100 Cai et al. (1995) California continental margin 4100 Ω = +0.69 1.0 no data Jahnke et al. (1997) California continental margin 800-3700 ∆ = -18 – -93.6 – 7.3 no data Hales and Emerson (1996) Western eq. Pacific (Ontong Java Plateau)2300-3000 Ω = +0.75 – +0.91 no data Jahnke et al. (1994) Western eq. Pacific (Ontong Java Plateau)3000, 4400 ∆ = -6, -37 no data ∆ = [CO32- ]in situ - [CO32- ]c measures the saturation state, where [CO32- ]c . Alternatively, Ω = [CO32- ]in situ - [CO32- ]c. The rain ratio Corg/CaCO denoted by δr. MD1 denotes the fraction of carbonate above the saturation horizon that is dissolved. MD2 denotes the contribution of metabolic processes to carbonate dissolution below the saturation horizon. a H. Jansens’s calculation using GEOSECS data, the values are not mentioned in the paper.
10 20 30 40 50 60
0 100 200 300 400 500 600 700
shell weight (µg)
CO32- (µmol kg-1)
(6) (4)
(5)
(2) (4)
(2)
y = 27.3 + 0.025x R2 = 0.39
10 20 30 40 50 60 70 80
0 100 200 300 400 500 600 700
shell weight (µg)
CO32- (µmol kg-1)
y = 29.5 + 0.051x R2 = 0.67 y = 23.19 + 0.11x R2 = 0.55
Figure A2. Increased foraminiferal shell weight under higher [CO32-] during shell growth: a) Linear fit regression for G. sacculifer at 29°C, real size range: 493-575 µm. Numbers in brackets represent the number of shells per average. Data compiled from several laboratory culture experiments. b) Linear fit regression for O. universa at 22°C, real size range: 500-600 µm. The shaded area represents the range of ambient [CO32-]. Data modified after Bijma et al. (1999).
Complication also arises from observations on an initial increase in the average shell weight of a freshly sedimented foraminiferal population: At the initial stages of dissolution (probably still above the lysocline) the thinshelled, light-weighted individuals disintegrate first. The loss of these shells from the sedimentary community is consistent with an observed increase in foraminiferal δ18O and δ13C which can be attributed to the lighter isotopic composition of thinshelled individuals (Erez, 1979a; Erez, 1979b). Lohmann‘s data indeed show the maximum weight per individual slightly above the lysocline. It should also be noted that the dissolution rate of foraminifera is not a linear process. This is due to the exponential increase of the specific surface area of a foraminifera with dissolution (Honjo and Erez, 1978) and to the non-linear nature of calcite-dissolution as a function of undersaturation (Keir, 1980).
An intriguing factor is also the dependency of the proxy on the target parameter itself, i.e. deep [CO32-] does not exclusively determine the preservation state of foraminiferal shells in the sediment but shell weight also depends on surface [CO32-] experienced during shell precipitation. Figure A2 demonstrates this dependency for the species G. sacculifer and O.
universa. For a glacial increase in surface water [CO32-] on the order of 55-100 µmol kg-1 (as determined from U/Ca and carbonate ion effect, Table A1), the shell weight of G. sacculifer
thus increases by approximately 1.6–2.9 µg – independent of concomitant changes in deep water saturation. A detailed examination of available culture and sediment trap data is required to better estimate the magnitude of the observed growth effects and to determine how temperature affects the carbonate dependent growth variability at a certain locality over glacial/interglacial time scales.
2.2.2 Reflectance/ lightness of foraminiferal shells
A qualitative estimate of carbonate corrosion prior to foraminiferal test fragmentation is the combination of weight and light reflectance measurements of planktonic foraminiferal tests of the polar species Neogloboquadrina pachyderma (sin.). The method was developed by Helmke and Bauch (2002) and is restricted to regions and time intervals where carbonate preservation is generally good. Carbonate corrosion leads to changes in the surface structure of the calcite crystals and has a profound influence on the reflectivity of foraminiferal tests.
An inverse relationship between light reflectence and weight was found. Application to Nordic Sea sediments revealed better preservation during glacial periods, which is consistent with higher deep sea [CO32-] for this time scale. However, the method is yet far from being used for quantitative estimates.
2.3 Estimating coccolithophorid paleoproductivity
Coccolithophorids are major contributors to the biogenic carbonate content in deep-sea sediments (Archer et al., 2000; Milliman, 1993; Westbroek et al., 1993). Recently, there has been increased interest in utilizing the elemental and isotopic chemistry of coccoliths. The chemistry of coccolith carbonate may record different information than that of foraminiferal carbonate because coccolithophorids, unlike foraminifera, are primary producers. Knowledge about their paleoproductivity is of major importance for e.g. rain ratio estimates and δ13Calkenone-based paleobarometer reconstructions (for review: Laws et al., 2001).
One limitation in the use of coccolith carbonate for geochemical studies has been their very small size and therefore the inability to seperate monospecific coccolith assemblages for analysis. New techniques now permit separation of fractions whose carbonate is highly dominated (>70 % and often >90 %) by a single coccolith species (Stoll and Ziveri, in press).
As with foraminifera, calcite produced by different species of coccolithophorids has different minor element partitioning and oxygen and carbon isotope fractionations. Specific examination now opens a new field for paleoceanography.
2.3.1 Coccolith Sr/Ca and stable carbon and oxygen isotopes to infer growth rate and cell size
The Sr/Ca ratio of coccoliths has been recently proposed as a potential indicator of past coccolithophorid growth rates. The hypothesis is based on correlations between Sr/Ca in polyspecific coccolith samples and primary productivity, alkenone-estimated growth rates, and CaCO3 rain rates in deep sediment traps (e.g. Stoll and Schrag, 2000) across the Equatorial Pacific upwelling region. Subsequently, a number of culture studies have investigated controls over Sr/Ca ratios in coccoliths of several species. For identical temperature and media composition, Sr partitioning is linearly related to rates of calcite production/cell (Stoll et al., 2001). Higher calcification per cell at higher growth rates observed in light-limited cultures of Gephyrocapsa oceanica, Calcidiscus leptoporus and Emiliania huxleyi cultures (Paasche, in press; Stoll et al., in press) may suggest that active uptake and calcification become increasingly important at higher growth rates. If coccolith Sr/Ca is a reliable indicator of coccolithophorid productivity, it provides an index of past productivity directly recorded by a primary producer. Furthermore, productivity estimates from coccolith Sr/Ca do not rely on conventional determinations of sediment accumulation rates which are often imprecise.
With regard to oxygen and carbon isotope fractionations, culture studies indicate different nonequilibrium effects for different species of coccolithophorids (Dudley et al., 1986; Ziveri et al., 2000; Ziveri et al., in prep.). These nonequilibrium effects appear to reflect changing ecological and physiological responses of the organisms. In light- and nutrient-replete cultures, the non-equilibrium effects in δ18O correlate highly with cell division rates across a range of species. At similar calcification temperature and media composition, the growth rates of the most common living species, E. huxleyi and G. oceanica, are strong and δ18O is 3 ‰ offset with respect to equilibrium composition. In contrast, species with low growth rates such as Umbilicosphaera sibogae var. foliosa have a δ18O fractionation effect of
~-2 ‰. Systematic relationships were also found between the carbon and oxygen isotopic composition of the coccolith calcite for each species and the surface area/volume ratio of the cells, which determines the diffusive flux of CO2 available to the cell (Ziveri et al., in prep.).
Clearly, more work is needed to test the validity of this proxy in constraining coccolithophorid growth rates. Nevertheless, we are encouraged that qualitative or quantitative determination of past variations in species-specific algal growth rates may be possible.
2.4 Model results
2.4.1 Lysocline reconstruction
In a modeling study, Jansen et al. (subm.) investigate the relationship between lysocline and saturation horizon. The relative positions of these properties in a modern sediment profile are described in Figure A3. The model results demonstrate that a combination of changing CaCO3 and Corg production in combination with an increase in the remineralization depth of organic carbon may have decoupled the lysocline and saturation horizon during the last glacial maximum (LGM). Instead of the conservative estimate of 20 µatm, changes in the marine carbonate pump thus may have been responsible for a greater portion of the observed glacial/interglacial atmospheric pCO2 shift on the order of 30-50 µatm.
Figure A3. a) Calcium carbonate content in a modern sediment profile. Lysocline and calcite compensation depth (CCD) are set to the depth where the sediment calcite content drops below 95% and 10%, respectively. Near to the saturation horizon (SH), the rain rate of CaCO3 exceeds the dissolution rate, explaining the position above the lysocline. The transition zone between lysocline and CCD is indicated by the yellow band. b) Sedimentary calcite saturation profile. The solid blue line denotes the calcite saturation in bottom water, while the broken blue line refers to pore water saturation. Note that above
-~3.6 km, the sediment is more corrosive than the bottom water due to organic carbon remineralization. Below that depth, calcite dissolution becomes progressively stronger, resulting in a pore water saturation state that is higher than in bottom waters.
Geological records suggest that during the last glacial, the Atlantic lysocline was 0.3-1 km shallower (Crowley, 1983; Curry and Lohmann, 1986) than today, while it was about 0.8 km deeper in the Pacific Ocean (Farrell and Prell, 1989). Assuming that the lysocline has not changed its position relative to the saturation horizon, these changes roughly correspond to a decrease in atmospheric pCO2 by approximately 20 µatm (Broecker et al., 2001). However, ice core observations suggest a glacial/interglacial shift in atmospheric pCO2 of 80 µatm (Neftel et al., 1982; Petit et al., 1999). To bring the two records into line, additional reduction in atmospheric pCO2 can be brought about by decoupling the lysocline from the saturation horizon, due to respiration-driven carbonate dissolution in the upper 10 cm of sediments, as observed e.g. by Hales and Emerson (1996). By changing the amount of organic carbon arriving at the seafloor, the amount of carbonate dissolution above the saturation horizon can change dramatically, shoaling the lysocline relative to the saturation horizon.
In contrast to the model used by Archer and Maier-Reimer (1994), Jansen et al.
(subm.) do not consider the dissolution of CaCO3 in the water column, as until now, the underlying mechanism of this proposed feature has not been found (Jansen and Wolf-Gladrow, 2001; Jansen et al., 2002; Milliman et al., 1999). Although this might result in an overestimation of the decoupling, it does not affect glacial/interglacial changes in the whole ocean carbonate inventory. Archer and Maier-Reimer (1994) tested scenarios for a glacial ocean where Corg production was three times as high and CaCO3 production 60 % lower than at present. They concluded that such a variation, operated by a shift from calcareous to siliceous organisms during glacial times, might have driven atmospheric pCO2 to glacial values. More recently, Sigman et al. (1998) argued that an increase in respiration-driven calcite dissolution has no significant effect on the decoupling of lysocline and saturation horizon as increased shallow water dissolution of carbonates would deepen the lysocline due to mass balance considerations. However, it is questionable whether carbonate production and dissolution are balanced at all (Milliman, 1993).
In contrast to Sigman et al. (1998), Jansen et al. (subm.) demonstrate that a decoupling of lysocline and saturation horizon is possible. However, respiration-driven dissolution can only significantly influence the atmospheric pCO2 when the export ratio depends on primary production. Otherwise, the amount of organic carbon reaching deep-sea sediments would be too small. With fixed export ratios, variability in CaCO3 production has a greater influence on atmospheric pCO2 than variability in organic carbon production has.
Organic carbon productivity has been assumed to have increased during the LGM in the range of up to +100 % relative to today (Berger et al., 1989; Kumar et al., 1995; Paytan et
al., 1996), while CaCO3 productivity ranged between –60 % and +60 % relative to modern times (Archer et al., 2000; Broecker and Henderson, 1998; Kumar et al., 1995). These estimates yield Corg: CaCO3 rain ratios that are comparable to modern rain ratios in high productivity areas. Thus, glacial pCO2 levels of ~ 230-250 µatm pCO2 are achieved within the assumed rain ratio ranges. The result of Archer and Maier-Reimer (1994), who found that the glacial to interglacial shift in atmospheric pCO2 is completely explainable by a decoupling mechanism could not be reproduced by Jansen et al. (subm.). Rather, their model indicates that ~40-60 % of the glacial pCO2 reduction may be attributed to changes in the marine carbonate pump.