Marine Calcification:
An Alkali Earth
Metal Isotope Perspective
Anton Eisenhauer1, Başak Kısakürek1, and Florian Böhm1
1811-5209/09/0005-0365$2.50
DOI: 10.2113/gselements.5.6.365
C
alcium and magnesium isotope fractionation, recorded in shells of
marine organisms that biomineralize calcium carbonate, is helping
scientists to understand the transport of trace elements from seawater
to the site of calcification, as well as trends in seawater composition throughout
time. This knowledge would be difficult to obtain otherwise, and is important,
especially now, for assessing the threat of ocean acidification to shell-producing
marine organisms.
Reef-building corals like
these branching
Acropora and Diploria
colonies from the Red
Sea are major sinks for
the alkali earth
elements Ca and Sr.
The measurement of
trace element ratios
(e.g. Sr/Ca) and their
CALCIUM ISOTOPE
isotope systems
FRACTIONATION IN provides
information
MARINE CALCIFYING
about trace metal
transport, environORGANISMS
mental conditions,
Recent advances in thermal and climate change.
The isotope composition of any material is usually reported
in the δ notation (for a definition, see Bullen and Eisenhauer
2009 this issue). Isotope fractionation of the alkali
earth elements (e.g. Ca, Mg, Sr) is defi ned as the isotope
difference between, for example, calcite (δ44/40 CaCC) and a
growth solution (δ44/40 CaGS), for example seawater; thus,
Δ44/40 CaCC-GS = δ44/40 CaCC – δ44/40 CaGS. The δ value is a material characteristic, whereas the Δ value describes the fractionation process eventually dependent on temperature
(T) and precipitation rate (R). Usually Δ44/40 CaCC-GS is negative because the light isotopes tend to become enriched in
the mineral precipitate. External reducibility of the Ca and
Mg data reported here is in the order of about 0.1 to 0.2‰.
1 Leibniz-Institut für Meereswissenschaften, IFM-GEOMAR
Wischhofstraße 1-3, D-24148 Kiel, Germany
E-mail: [email protected]
E LEMENTS , V OL . 5,
PP.
365–368
Δ44/40C aCC-GS (‰ )
and plasma mass spectrometry have allowed the precise
measurement of small differences
in Ca isotope composition in
KEYWORDS : calcium isotope, magnesium isotope, biocalcification,
CaCO3. As a result, recent isotope
inorganic carbonate, shell
studies of Ca biomineralization
have been able to test whether Ca
isotopes can provide an environINTRODUCTION
mental indicator—for example, of temperature (T)—using
marine CaCO3 as an archive. It was hoped that developThe alkali earths calcium (Ca) and magnesium (Mg) play
a central role in a variety of geological, environmental, ment of this tool would provide information complementary to that gained from more traditional marine T data,
and biological processes. Changes in the Ca concentration
such as Mg/Ca in foraminifera and Sr/Ca in corals.
of seawater, together with changes in alkalinity and pH,
have been suggested as a major driving force for the onset
For both calcite and aragonite—two of the six polymorphs
of Ca carbonate (CaCO3) biomineralization at the
of CaCO3 —precipitation experiments have shown that
Precambrian–Cambrian boundary. Ca and Mg concentraΔ44/40 Ca CC-GS correlates with T (~0.02‰/°C) (F IG . 1).
tions have also been implicated in driving the alternation
However, aragonite displays an offset of about –0.5‰ relabetween aragonite and calcite seas observed throughout
tive to calcite. In contrast, Ca isotope fractionation in
Earth’s history. Our ability to confi rm these and other
biogenic CaCO3 depends on more factors than simply polytheories hinges on being able to decipher the marine
morphism. Aragonitic sclerosponges and pteropods secrete
CaCO3 record. This depends on being able to recognize
skeletons with a fractionation factor similar to that for
markers that can differentiate between inorganic CaCO3
precipitation and physiologically controlled calcification
0
by organisms. Another important piece of the puzzle is to
understand the interplay of Ca and Mg flux in and out of
the oceans, and its role as an indicator of long-term changes
in continental weathering, mid-ocean ridge spreading
-1
rates, and climate. Ca and Mg stable isotope fractionation
provides new understanding of CaCO3 formation processes
that are needed to interpret the marine record.
-2
inorganic calcite (1)
coccolithophores (2)
corals (3)
biogenic aragonite (4)
inorganic aragonite (5)
most planktonic foraminifera (6)
G. sacculifer (7)
except
N. pachyderma (8)
-3
{
-4
0
10
20
30
Temperature (ºC )
Temperature-dependent Ca isotope fractionation in
(1) inorganic calcite (Marriott et al. 2004), (2) calcite
of coccolithophores (Langer et al. 2007), (3) aragonite skeletons of
scleractinian corals (Böhm et al. 2006), (4) aragonite shells of sclerosponges and pteropods (Gussone et al. 2005), (5) inorganic aragonite
(Gussone et al. 2003), (6) calcite of most planktonic foraminifera
(Gussone et al. 2003; Griffith et al. 2008a), (7) planktonic foraminifer
Globigerinoides sacculifer (Nägler et al. 2000), and (8) planktonic
foraminifer Neogloboquadrina pachyderma (sin.) (Hippler et al. 2009)
365
FIGURE 1
D ECEMBER 2009
inorganic aragonite, whereas aragonitic scleractinian
corals, calcitic coccolithophores, and calcitic planktonic
foraminifera have comparable degrees of isotope fractionation, lying between the values for inorganic calcite and
aragonite (FIG. 1). Two species of planktonic foraminifera,
Globigerinoides sacculifer and Neogloboquadrina pachyderma
(sin.), show greater dependence of Δ44/40 CaCC-GS on T (FIG. 1).
Given their different habitats (tropical versus polar surface
ocean), both species show surprisingly similar T sensitivity
(∼0.24‰/°C). This was observed in laboratory cultures and
in planktonic foraminifera collected from the ocean with
tow samplers. However, other studies involving these same
species failed to reproduce the large T dependence (Griffith
et al. 2008a), suggesting an influence of salinity and T
thresholds for the various species (Hippler et al. 2009). The
latter observation indicates a complex and unpredictable
physiological control on Ca uptake by calcifying organisms
prior to CaCO3 formation, challenging its use as a paleotemperature proxy.
dence time (FIG. 2) cannot be explained by imbalances in
the ocean Ca budget. In this regard, Heuser et al. (2005)
suggested that the formation of dolomite [MgCa(CO3)2]
leads to isotope fractionation as Ca is replaced by Mg in
marine CaCO3. The enrichment of dolomite in heavy Ca
isotopes enriches seawater in lighter Ca isotopes. In
contrast, diminishing dolomitization in the ocean enriches
seawater in heavier Ca isotopes, possibly explaining the
gradual increase in the Ca isotope composition of seawater
from the Miocene Climatic Optimum (time interval
between 17 and 15 Ma with high atmospheric and seawater
temperatures; also characterized by large continental shelfs
and sub- to anoxic conditions in ocean water) to the
present.
A sudden increase of the Ca isotope values at the beginning
of the Late Paleozoic (350 Ma; FIG. 2) correlates with a shift
from the Early Paleozoic calcitic sea to the Late Paleozoic
THE PHANEROZOIC CALCIUM
ISOTOPE RECORD
Calcium is provided to the ocean by rivers that drain
mineral-weathering areas on the continents and by hydrothermal activity in the oceans. An additional source of Ca
flux in the geological past was dolomitization, in which
Ca in CaCO3 was replaced by Mg. In the modern ocean,
the most important sinks for Ca are biogenic precipitation
and sedimentary deposition of CaCO3. Less important is
the formation of CaCO3 during low-T alteration of ocean crust.
0.0
Δ44/40C a CC-GS (‰ )
A
Phanerozoic history of seawater Ca isotope composition
(δ44/40Ca) recorded in marine fossils relative to the Ca
standard SRM915a. Data from Farkaš et al. (2007b) (brachiopods and
belemnites) and Heuser et al. (2005) (planktonic foraminifera) are
averaged in 10-million-year bins. The record shows a general increase
with time from low values in the Early Paleozoic (540 to 360 Ma) to high
values in the Early Cretaceous (146 to 100 Ma) and Neogene (23 to
0 Ma). A sharp rise during the Carboniferous (360 to 300 Ma) is well
documented in the data. A second rise occurred during the Late
Jurassic (160 to 145 Ma), but its precise onset is not known because
data are not available for the earlier Jurassic.
FIGURE 2
2σ
0.0
2σ
-0.4
-0.4
-0.8
-0.8
-1.2
T=21±1ºC
T = 40ºC
T = 25ºC
T = 5ºC
stirred
2
unstirred
a
-1.6
1
3
4
5 -1
lo g (R ) (μmol/m2/h)
E LEMENTS
0
1
b
2
3
4
-1.6
5
lo g (R ) (μmol/m2/h)
Growth rate (R)–dependent Ca isotope fractionation for
inorganically precipitated calcite grown from (A) wellstirred CaCl2–SrCl2–NH4Cl solutions with ionic strength = 0.035 molal
(M) and Ca concentration = 10 mmol/l (Tang et al. 2008b) and
FIGURE 3
-1.2
B
Δ44/40C a CC-GS (‰ )
Oceanic Ca isotope composition changes only if the input
and output fluxes vary in their δ44/40 Ca composition or
become imbalanced in magnitude on a timescale of 0.5
million years. The actual isotope compositions of Ca sinks
and sources are similar, so the modern ocean is close to or
at steady state (Schmitt et al. 2003). However, the situation
must have been different in the geological past. For
example, De La Rocha and DePaolo (2000) showed that
δ44/40 Ca varied considerably during the last 80 My. During
the Early Paleogene, relatively high δ44/40Ca values reflected
the fact that the Ca input flux was ∼80% of the output flux
for a period of ∼25 My, causing the Ca concentration in
seawater to decrease. Further studies confi rmed that
substantial fluctuations in the δ44/40Ca of seawater occurred
during the Neogene (cf Griffith et al. 2008b) and earlier in
Earth history. Variations much longer than the Ca resi-
(B) unstirred and stirred CaCl2–NH4Cl solutions with ionic strength =
0.45 to 0.85 M and Ca concentration = 15 mmol/l (open symbols)
and 150 mmol/l (closed symbols) (Lemarchand et al. 2004).
366
D ECEMBER 2009
aragonitic sea. Ca isotope fractionation is offset by about
0.5‰ between calcite and aragonite, so the shift from a
calcite- to an aragonite-dominated ocean would be expected
to cause a shift in the seawater Ca isotope composition.
-0.4
(1) inorganic calcite
Δ44/40Ca CC-GS (‰)
-0.6
CONTROLS ON ISOTOPE FRACTIONATION
OF CALCIUM IN INORGANIC CALCITE
MAGNESIUM ISOTOPES IN CARBONATES
Although no resolvable temperature sensitivity of Mg
isotope fractionation has been detected, differences in Mg
isotope fractionation are apparent between aragonite and
high- and low-Mg calcite (FIG. 5). In contrast to the observations on Ca isotopes, aragonite skeletons of corals and
sclerosponges show the least Mg isotope fractionation.
High-Mg calcite of marine organisms as well as inorganic
precipitates display a relatively constant degree of Mg
isotope fractionation, whereas low-Mg tests of planktonic
foraminifera reveal the largest degree of Mg isotope fractionation (FIG. 5). In contrast, coccolithophores (<1 mol%
of MgCO3) and benthic foraminifera with high Mg content
have distinct Mg isotope signatures that are heavier than
those of planktonic foraminifera. The major marine calciE LEMENTS
-0.8
(2) brachiopods
-1.0
(3) planktonic
(8) soft coral
foraminifera
P. fulvum
(4) blue mussel
M. edulis
(7) benthic foraminifer
Amphistegina spp.
2σ
-1.2
(6) coccolithophore
E. huxleyi
(5) oyster
Ostrea
-1.4
-1.6
0.0
0.1
0.2
0.3
0.4
DSr
Ca isotope fractionation versus Sr partition coefficient
(DSr) from (1) inorganic calcite (gray diamonds; Tang
et al. 2008b); low-Mg calcite of (2) brachiopods (brown squares;
Steuber and Buhl 2006; Farkaš et al. 2007a), (3) planktonic
foraminifera Globigerinoides sacculifer (see also FIG. 1; light blue
triangle; Gussone et al. 2004), (4) mussel Mytillus edulis (blue
square), (5) oysters Ostrea (purple squares; Steuber and Buhl 2006),
and (6) coccolithophores Emiliania huxleyi (green diamond; Langer
et al. 2006, 2007); and high-Mg calcite of (7) benthic foraminifera
Amphistegina spp. (red triangle), (8) soft coral Parerythropodium
fulvum (orange circle), and (9) sclerosponge Acanthochaetetes wellsi
(yellow circle; Gussone et al. 2005).
FIGURE 4
0
(2) coccolithophores
(<1 mol% MgCO3)
(1) Aragonite
corals and sclerosponges
-1
Δ26/24Mg CC-GS (‰)
Inorganic CaCO3 precipitation experiments have shown
that Ca isotope fractionation is controlled by both precipitation rate (R) and T. Using a CO2 -diffusion technique in
well-stirred solutions with Ca and Sr concentrations close
to those of modern seawater, Tang et al. (2008a, b) showed
that, at constant T, Δ44/40 CaCC-GS decreased as R increased
(FIG. 3A). They also observed a secondary control by T,
where Δ44/40 CaCC-GS increased with increasing T at constant
R (F IG. 3A). The inverse relationship between R and
Δ44/40 CaCC-GS is best explained by the surface entrapment
model, in which isotope fractionation of Ca develops under
disequilibrium conditions during precipitation by two
counteracting processes: (1) preferential incorporation of
light Ca at the crystal surface at greater R and (2) reequilibration of Ca isotopes in the surface layer by ion diffusion
at lesser R.Using a different calcite precipitation technique,
Lemarchand et al. (2004) observed that Δ44/40 Ca CC-GS
increases with increasing R at constant T (FIG. 3B). However,
at constant R, a smaller Δ44/40 CaCC-GS value was obtained in
unstirred solutions relative to stirred solutions (FIG. 3B).
The lowest Δ44/40 CaCC-GS value, about –1.5‰, was observed
in unstirred solutions at presumed equilibrium between
crystal growth and dissolution, when R was close to zero.
In contrast, with increasing R, Δ44/40 CaCC-GS comes close to
zero. The contrasting R dependence of Δ44/40 CaCC-GS, seen
in FIGURE 3, might be related to differences in the experimental and chemical setup. Furthermore, the stirring rate
of the growth solution is thought to control aqueous ion
diffusion and hence might influence the precipitation environments. A particularly exciting result of the inorganic
precipitation experiments of Tang et al. (2008a, b, c) is that
partitioning of Sr between the growth solution and calcite
is correlated with Ca isotope fractionation (FIG. 4). Although
a possible influence of pH cannot yet be excluded, the
relationship between Δ44/40 CaCC-GS and the partition coefficient DSr [i.e. DSr = (Sr/Ca)carbonate/(Sr/Ca)GS] is independent
of T, R, (Sr/Ca)GS, ionic strength, and isomorphic substitution of Ca by Mg (FIG. 4). This implies that element partitioning and isotope fractionation are controlled by similar
processes, independent of the distinct physical and chemical characteristics of elements and their isotopes (Tang et
al. 2008b). Recently obtained data from oysters, brachiopods, foraminifera, and blue mussels plot close to the inorganic regression line, whereas those from coccolithophores,
soft corals, and sclerosponges are offset, indicating that
physiological effects overprint inorganic control.
(9) sponge
A. wellsi
(6) benthic foraminifer
Amphistegina spp.
(~5 mol% MgCO3)
-2
(4) calcitic speleothems
(0.5-7.5 mol% MgCO3)
(7) soft coral P. fulvum
(~13 mol% MgCO3)
-3
(5) inorganic calcite
(0.4-3.7 mol% MgCO3)
-4
(3) planktonic foraminifera
(<1 mol% MgCO3)
-5
0
5
10
15
20
25
30
35
Temperature (°C)
Mg isotope fractionation plotted versus T in (1) aragonitic
corals and sclerosponges (dark blue line; Wombacher et
al. 2006); low-Mg calcite of (2) cultured coccolithophores E. huxleyi
and C. braarudii (filled and empty green diamonds, respectively) and
(3) planktonic foraminifera (light blue line; Chang et al. 2004; Pogge
von Strandmann 2008); and high-Mg calcite of (4) speleothems
(yellow squares; Galy et al. 2002), (5) inorganic precipitates (brown
squares), (6) cultured benthic foraminifer Amphistegina spp. (red triangles), and (7) cultured soft coral Parerythropodium fulvum (orange
circles).
FIGURE 5
fying organisms (coccolithophores, foraminifera) discriminate Mg relative to Ca in order to precipitate calcite. The
low Δ26/24Mg values in some calcifying organisms (up to
about –4.5‰ for Δ26/24Mg) may therefore be a result of Mg
transport across channels and pumps embedded in their
cell membranes. In this regard, calcifying organisms may
not follow the pathway that consumes the least energy to
precipitate calcite, that is, increasing pH by removing H+
ions from their calcifying vesicles (Zeebe and Sanyal 2002).
367
D ECEMBER 2009
Other criteria, such as greater thermodynamic stability of
low-Mg calcite relative to high-Mg calcite, might be more
important. Preliminary results indicate that the Mg isotope
composition of inorganically precipitated calcite produced
both naturally and in the laboratory is controlled by equilibrium fractionation, whereas the Mg isotope composition
of planktonic foraminifera is controlled by kinetic fractionation (Kısakürek et al. 2009). The Mg isotope ratios of
the benthic foraminifer Amphistegina spp. and of coccolithophores are more easily reconciled with multistage
rather than single-stage mass-fractionation pathways.
However, consistent errors due to instrumental mass bias and/
or fluid composition cannot yet be ruled out, calling for comparative studies in the future.
FUTURE DIRECTIONS
In the future, Ca and Mg isotope studies will be extended
by the knowledge of Sr isotope fractionation, and these
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