Most of our knowledge of past seawater temperature history is based on
δ18O values of calcium carbonate fossil shells. However, the
determination of past temperatures using this proxy requires the knowledge
of past seawater δ18O values, which is generally poorly
constrained. Other paleothermometers using carbonate archives, such as Mg/Ca
ratios and clumped isotopes (Δ47), have been developed to allow for paleotemperatures to be estimated independently and to allow past ocean δ18O values to be calculated using various groups of calcifying
organisms. Articulated brachiopod shells are some of the most commonly used
archives in studies of past oceanic geochemistry and temperature. They are
abundant in the fossil record since the Cambrian, and for decades, their low
Mg–calcite mineralogy has been considered relatively resistant to diagenetic
alteration. Here, we investigate the potential of brachiopod shells as
recorders of seawater temperatures and seawater δ18O values
using new brachiopod shell geochemical data by testing multiple well-established or suggested paleothermometers applied to carbonate archives.
Modern articulated brachiopod shells covering a wide range of temperatures
(-1.9 to 25.5 ∘C), depths (5 to 3431 m) and salinities (33.4 to
37.0 PSU) were analysed for their stable isotope compositions (δ13C, δ18O and Δ47) and their elemental ratios
(Mg/Ca, Sr/Ca, Na/Ca and Li/Ca). Our data allowed us to propose a revised
oxygen isotope fractionation equation between modern-brachiopod shell
calcite and seawater:
T=-5.0(±0.2)(δ18Oc-δ18Osw)+19.4(±0.4),
where δ18Oc is in ‰ VPDB, δ18Osw is in ‰ VSMOW, and T is in ∘C. Our results strongly support the use of clumped isotopes as an
alternative temperature proxy but confirm significant offsets relative to
the canonical relationship established for other biogenic and abiogenic
calcium carbonate minerals. Brachiopod shell Mg/Ca ratios show no
relationship with seawater temperatures, indicating that this ratio is a
poor recorder of past changes in temperatures, an observation at variance
with several previous studies. Despite significant correlations with
brachiopod living temperature, brachiopod shell Sr/Ca, Na/Ca and Li/Ca
values indicate the influence of environmental and biological factors
unrelated to temperature, which undermines their potential as alternative
temperature proxies. Kinetic effects (growth rates) could explain most of the
deviation of brachiopod shell calcite from expected isotopic equilibrium
with seawater and part of the distribution of Sr/Ca, Na/Ca and Li/Ca ratios.
Agence Nationale de la RechercheANR-18-CE31-0020Introduction
Oxygen isotope ratios (δ18O) of brachiopod shell calcite
constitute the most extensive record of marine temperatures over the
Phanerozoic eon (Prokoph et al.,
2008; Veizer and Prokoph, 2015). Nevertheless, our limited knowledge of
δ18O values of past seas and oceans prevents a confident
interpretation of this record, as carbonate δ18O values
(δ18Oc) are dependent on both shell growth temperature and
living-water δ18O (δ18Osw; Epstein
et al., 1953; Kim and O'Neil, 1997; Kim et al., 2007; Brand et al., 2013,
2019). Other paleotemperature proxies, such as trace element ratios and
clumped isotopes (Δ47), may be the key to building a more
confident record of Phanerozoic ocean temperatures. Owing to the
determination over the last decades of reliable oxygen isotope fractionation
equations between calcite and seawater, such independent temperature
estimates can be combined with δ18Oc measurements to
reconstruct δ18Osw values and hence to characterize oceanic
or local hydrography.
Alternative carbonate archive thermometers have been developed for different
calcifying organisms and have been proposed in the literature as potentially
applicable to brachiopod shells. Element / Ca ratios such as Mg/Ca, Sr/Ca and
Li/Ca are the most common in the literature. Mg/Ca thermometry of
foraminiferal calcite (Nürnberg, 1995; Anand et
al., 2003) is a widely used proxy that is the basis for the reconstruction of
changes in δ18Osw and ice volumes over the Cenozoic era (Lear,
2000; Billups and Schrag, 2003; Miller et al., 2020). Building on the
pioneering work of Lowenstam (1961), Brand et al. (2013, 2019) documented a positive correlation between seawater temperature
and the MgCO3 content of brachiopod shells sampled at a worldwide
scale. Plus, following the conclusions of
Jiménez-López et al. (2004), Brand et
al. (2013) developed an isotopic fractionation equation that considers the
amount of Mg2+ substituted for Ca2+ in the crystal lattice of
the low-Mg calcitic shell which was abandoned in their more recent works
(Brand et al., 2019). The MgCO3–temperature relationship is described by a curve named the
Global Brachiopod Mg Line (GBMgL), yet Brand et al. (2013, 2019)
acknowledge species-specific deviations from this concept that fall well above
the said line. Plus, brachiopod shell Mg content has been found to have
a strong taxonomic trend, with shells of thecideid and craniid brachiopods being made of
high-Mg calcite in contrast to the low-Mg calcite of rhynchonellide and
terebratulide brachiopod shells (Brand
et al., 2003; Ullmann et al., 2017). This undermines the brachiopod shell
MgCO3–temperature relationships which uses a dataset that consists of thecideid,
rhynchonellide and terebratulide (Brand et al.,
2013, 2019).
While Sr/Ca ratios are mostly used to assess the shell preservation of
fossil specimens (Brand and Veizer, 1980), this ratio also
constitutes a potential marine paleothermometer applicable to brachiopod
shells, as is the case for corals (McCulloch
et al., 1994; Shen et al., 1996; Marshall and McCulloch, 2002; Swart et al.,
2002; Ayling et al., 2006; DeLong et al., 2011). Previous studies already
suggested Sr/Ca ratios as marine paleothermometers for brachiopod shells
owing to their correlations with oxygen isotope ratios and seawater
temperature (Lowenstam,
1961; Mii and Grossman, 1994; Pérez-Huerta et al., 2008; Butler et al.,
2015; Ullmann et al., 2017). Similarly, the brachiopod Li/Ca ratio was also
suggested as a potential temperature proxy in the seminal study of
Delaney et al. (1989). Subsequent studies have also
confirmed that Li/Ca may constitute a potentially accurate thermometer in
brachiopods (Dellinger
et al., 2018; Rollion-Bard et al., 2019; Washington et al., 2020) and in
corals (Montagna et
al., 2014; Marchitto et al., 2018). For this latter group, however, the Li/Mg
ratio has been considered to be a much more precise thermometer than Li/Ca (Montagna et al.,
2014; Marchitto et al., 2018) despite some identified caveats such as the
scrambling role of organic matter (Cuny-Guirriec et al., 2019).
On top of element / Ca proxies, clumped isotope (Δ47) thermometry
has been developed more recently in relation to different carbonate materials
(Ghosh et al., 2006;
Zaarur et al., 2013). This parameter measures the anomaly of
13C–18O bonds within the carbonate lattice relative to their
stochastic abundance. Δ47 values are strongly correlated to
crystallization temperature and are independent from δ18Osw.
This proxy has been applied to different marine calcifying organisms to
estimate δ18Osw in the deep past
(Petersen
et al., 2016; Bergmann et al., 2018; Henkes et al., 2018; Wierzbowski et
al., 2018; Price et al., 2020; Vickers et al., 2020, 2021; de Winter et al.,
2021; Meckler et al., 2022). Although promising, brachiopod shell Δ47 composition was only used in a couple of studies to estimate past
temperatures (Henkes et al., 2013; Came
et al., 2014). Indeed, the few existing modern-brachiopod Δ47–temperature data are in good agreement with the canonical Δ47–temperature relationship established for other biogenic and
abiogenic calcium carbonate minerals (Henkes et al., 2013; Came et al.,
2014). Recent compilation of Δ47–T calibration data from
different laboratories using the new carbonate-based standardization
(Bernasconi
et al., 2018, 2021) concluded with a unique Δ47–T relationship
for synthetic and natural carbonates (Anderson et al., 2021). The
establishment of such a unique calibration equation strengthens the good
agreement observed between natural and synthetic Δ47–T
relationships and would suggest that this calibration is valid when applied
to brachiopod shells, especially as empirical calibrations are in very good
agreement with theoretical predictions (Jautzy et
al., 2020). However, Δ47 measurements from brachiopod shells
are absent from this compilation, and other studies suggest that they may be
subjected to significant deviations from Δ47–T relationships that are
likely related to differences in growth rates
(Bajnai et al., 2018, 2020).
Beside, a number of studies suggested that living temperature alone may not
explain the variability of those different parameters within brachiopod
shells. First, biogenic carbonate precipitation rate depends on
environmental and biological factors (temperature and salinity tolerance,
age, reproduction period, ontogeny, etc.), which causes biases in
the paleoenvironment record extracted from the shell
(Peck et al., 1997;
Schöne, 2008). Brachiopod shell calcite is secreted by the outer
epithelial cells of the mantle (Williams,
1968; Simonet Roda et al., 2019a, b, 2022). Calcite elements are in close
relationship with organic matrices which control the mineral growth (Williams,
1968; Gaspard et al., 2008; Gaspard and Nouet, 2016; Simonet Roda et al.,
2019a, b, 2022). Finally, nanoparticles are transported via ion transport
through the cell membrane, highlighting calcite precipitation as an
amorphous calcium carbonate precursor stage (Griesshaber et al.,
2009; Schmahl et al., 2012). The chemistry of the calcifying environment of
brachiopod shell calcite is poorly constrained (Immenhauser et al., 2016)
Testing thermodynamic equilibrium at the mineralization site is therefore,
to date, out of reach. Consequently, the chemistry of brachiopod shells is
usually compared to that of the surrounding water in order to approach the
processes involved in biomineralization from a chemical point of view.
Numerous studies highlighted that brachiopod shell calcite does not
precipitate sensu stricto in isotopic equilibrium with the surrounding water but, at best,
mimics thermodynamic equilibrium (Carpenter
and Lohmann, 1995; Auclair et al., 2003; Parkinson et al., 2005; Yamamoto et
al., 2010a, b; Cusack et al., 2012; Takayanagi et al., 2012, 2013, 2015;
Bajnai et al., 2018, 2020; Romanin et al., 2018; Rollion-Bard et al., 2019).
Brand et al. (2019) characterize their oxygen isotope fractionation equation
as representing a brachiopod based-equilibrium. This term is opposed to
the thermodynamic equilibrium, which is best described by very-slow-growing
calcites (Coplen, 2007;
Daëron et al., 2019), thus ensuring isotopic equilibration of the
dissolved inorganic carbon (DIC) species with water
(Watkins et al., 2013, 2014).
Trace element incorporation into calcite is also affected by kinetic
effects. In synthetic carbonates, the relative abundance of elements such as
Sr, Na and Li rises with higher growth rates (Lorens,
1981; Busenberg and Plummer, 1985; Tesoriero and Pankow, 1996; Gabitov et
al., 2011, 2014). Such kinetic trends were evidenced at an intra-individual
level among brachiopods by the negative correlation of element / Ca ratios
with stable isotope values (δ13C, δ18O; Ullmann et al., 2017;
Rollion-Bard et al., 2019).
In this study, we assess the potential of various geochemical proxies
(δ18O, Δ47, Sr/Ca; Mg/Ca, Li/Ca, Na/Ca) as
recorders of seawater temperature by analysing a new set of modern
articulated brachiopod shells collected from various depths and latitudes
during institutional oceanographic cruises and covering a broad range of
water temperatures comprised between -1.9 and 25.5 ∘C. We discuss the validity and robustness of our oxygen isotope
fractionation equation established with measured δ18Oc of
calcite and with seawater temperature and δ18Ow estimated from
oceanographic data. The dependence of elemental ratios on seawater
temperature is discussed in light of this new available dataset of
modern brachiopods. Finally, we highlight kinetic effects as non-negligible
sources of isotopic and trace element variability.
Material and methodsSample collection and environmental parameters
The studied material consists of 37 articulated brachiopod shells that were
collected in situ during institutional oceanographic cruises or other scientific
missions (see Supplement for details). The sampling location was
documented with depth and geographic coordinates for most samples.
Temperature and salinity, which were not measured in situ for most samples, were
estimated independently using the NOAA World Ocean Atlas 2018
(Locarnini et al.,
2018; Zweng et al., 2018) or, whenever available, using long-term local
records in the literature (Table 1). For both temperature and salinity, mean
annual values (MAT), as well as higher and lower monthly averages, were
considered. Similarly, δ18Osw values were usually not
measured in situ but were calculated from the salinity data with the appropriate
regional δ18Osw–S relationship from
LeGrande and Schmidt (2006). The oceanographic parameters
used in this study are listed in Table 1.
Brachiopod samples with taxonomic identification, sampling location
and environmental parameters. Most environmental parameters are derived from
the World Ocean Atlas 2018 (Locarnini et al.,
2018; Zweng et al., 2018). For a few samples, environmental parameters are
derived from local literature, as listed below. Almost all δ18Osw values are calculated using the regional δ18Osw–salinity relationship published by LeGrande
and Schmidt (2006), unless direct measurements covering seasonal variations
were available in the literature.
1 Goodwin and Cornelisen (2012); 2 Woods et
al. (2014); 3 Meredith et al. (2013);
4 Jacobson (1983).
Sample preparation
Encrusting organisms covering the shells were first mechanically removed
using stainless steel dental tools, and each shell was then placed in diluted
bleach (NaClO 5 %) for 5 to 10 min in an ultrasonic bath to remove
organic matter and other contaminants. As preliminary tests have revealed
that the more delicate shells (e.g. Macandrevia africana; Cooper, 1975) broke apart during the
ultrasonic bath, immersion in diluted bleach for a few hours without the
ultrasonic bath was subsequently preferred for a few specimens. In all cases,
the specimens were rinsed with deionized water and oven dried at
50 ∘C for a few hours.
The sampling of brachiopod shell calcite was performed following the
recommendations of Romanin et al. (2018). The umbo, edges and muscle scar area were avoided because they
record major kinetic or metabolic effects (Carpenter
and Lohmann, 1995; Auclair et al., 2003; Parkinson et al., 2005; Yamamoto et
al., 2010b; Ullmann et al., 2017; Romanin et al., 2018), and sampling was
focused on the middle part of the shell. Considering the complex structure
of brachiopod shells (Williams,
1968; Gaspard et al., 2018; Simonet Roda et al., 2022) relative to the
precision of our sampling method (dental tools and engraving bit), we were
unable to prepare pure samples for each layer of the brachiopod shell
(acicular primary layer, fibrous secondary layer and columnar tertiary
layer). In this study, we differentiate the outer layers, i.e. a mix of the
primary acicular layer and some amount of the outer secondary fibrous layer,
which record significant kinetic effects (Carpenter
and Lohmann, 1995; Auclair et al., 2003; Parkinson et al., 2005; Bajnai et
al., 2018; Romanin et al., 2018; Rollion-Bard et al., 2019), from the inner
layers, i.e. the inner secondary fibrous layer and/or the tertiary
columnar layer when present. Outer shell layers were removed manually with
dental tools, and resulting powders were kept to compare their elemental
ratios, δ18O and δ13C with those of the inner shell layers
(Carpenter and
Lohmann, 1995; Auclair et al., 2003; Pérez-Huerta et al., 2008). After
removal of the outer layers, the inner layers were sampled using an
engraving bit fitted to a DREMEL Micro™ drill adjusted to the lowest
possible speed. For each sample, the powder was used for all selected
geochemical analyses which correspond to δ18O, δ13C, Δ47 and element concentration. As clumped isotope
analyses require large amounts of material (>10 mg is necessary
to operate multiple replicate analyses), a large area (a few
cm2) was sampled in the middle part of the shell, allowing
several mg of calcitic powder to be collected. It is worth noting that
this large sampling area may correspond to several months or years of shell
growth, depending on the species.
This sampling protocol was applied to most of our samples, except for the
smallest and most fragile shells that quickly broke apart during sampling.
Consequently, in the case of the small shells of Terebratulina latifrons, Frenulina sanguinolenta and some specimens of
Fallax neocaledonensis, the outer layers were kept, and only the area covering the umbo up to the
muscle scar was avoided for sampling. The weakness and thinness of M. africana
precluded a total removal of the outer layers; however, the umbo, muscle
scars and edges were removed from the bulk shell. Fragments of these more
fragile shells were ground to a fine powder in an agate mortar.
Carbon and oxygen stable isotopes
Stable isotope compositions of the sampled powders were determined using a MultiPrep™ Autosampler coupled to a dual-inlet GV
IsoPrime® mass spectrometer. For each sample, an aliquot of
about 400 µg of calcium carbonate was reacted with anhydrous
oversaturated phosphoric acid at 90 ∘C for 20 min. Oxygen
isotope ratios of calcium carbonate were computed assuming an acid
fractionation factor 1000lnα(CO2–CaCO3) of 8.1 between
carbon dioxide and calcite (Swart et
al., 1991). All sample measurements were duplicated and adjusted to the
international references NIST NBS18 (δ18OVPDB=-23.2 ‰; δ13CVPDB=-5.01 ‰) and NBS19 (δ18OVPDB=-2.20 ‰; δ13CVPDB=+1.95 ‰) and an internal standard of Carrara Marble
(δ18OVPDB=-1.84 ‰; δ13CVPDB=+2.03 ‰). Since 2019,
reproducibility of the Carrara Marble in-house standard has been ±0.096 ‰ for δ18O (2σ, n=1062)
and ±0.066 ‰ for δ13C (2σ,
n=1062). These stable isotope data are completed with δ18O
and δ13C data obtained from clumped isotope measurements, for
which the analytical procedures are reported in detail in Sect. 2.5.
Elemental ratios
Elemental concentrations were obtained by dissolving 2 to 20 mg of carbonate
powder in 10 mL HNO3 (2 %). For all samples, pairs of aliquots were
prepared with 10× and 100× dilutions, depending on the considered
trace and major elements, along with a fixed amount of Sc and In that were
added to correct concentrations from instrument drift. Solutions were
analysed using an inductively coupled plasma-optical emission spectrometer
(iCAP 7000 ICP-OES) and a quadrupole ICP-mass spectrometer (i-CAP-Q ICP-MS)
for minor and trace elements, respectively. The 100× diluted aliquots were
used to calculate Ca concentration, while the 10× diluted aliquots served for
Mg, Sr, Na and Li concentrations. The calculated element / Ca ratios are
reported in mmol / mol for Mg/Ca, Sr/Ca and Na/Ca and in µmol / mol for
Li/Ca. The reproducibility of measurements was assessed through the analysis
of the carbonate standard CCH1 (Roelandts and Duchesne,
1988).
Clumped isotopes
Carbonate samples were converted to CO2 by phosphoric acid reaction at
90 ∘C in a common, stirred acid bath for 15 min. The initial
phosphoric acid concentration was 103 % (1.91 g cm-3), and each batch
of acid was used for 7 d. After cryogenic removal of water, the evolved
CO2 was helium-flushed at 25 mL min-1 through a purification column
packed with Porapak Q (50/80 mesh, 1 m length, 2.1 mm ID) and held at
-20 ∘C, then it was quantitatively recollected by cryogenic trapping
and transferred into an IsoPrime 100™ dual-inlet mass spectrometer
equipped with six Faraday collectors (m/z 44–49). Each analysis took about
2.5 h, during which analyte gas and working reference gas were allowed
to flow from matching 10 mL reservoirs into the source through deactivated
fused-silica capillaries (65 cm length, 110 µm ID). Every 20 min,
gas pressures were adjusted to achieve m/z= 44 current of 80 nA, with
differences between analyte gas and working gas generally being below 0.1 nA.
Pressure-dependent background current corrections were measured 12 times for
each analysis. All background measurements from a given session were then
used to determine a mass-specific relationship linking background intensity
(Zm), total m/z= 44 intensity (I44) and time (t) as follows:
Zm=aI44+P(t),
with P being a polynomial of degree 2 to 4.
Background-corrected ion current ratios (δ45 to δ49) were
converted to δ13C, δ18O and “raw” Δ47 values as described by Daëron et al. (2016) using the IUPAC oxygen-17 correction parameters.
The isotopic composition (δ13C, δ18O) of our
working reference gas was computed based on the nominal isotopic composition
of carbonate standard ETH-3 (Bernasconi et al.,
2018) and an oxygen-18 acid fractionation factor of 1.00813
(Kim et al., 2007). Raw Δ47 values were
then converted to the I-CDES (Intercarb-Carbon Dioxide Equilibrium Scale) Δ47 reference frame by comparison
with four ETH carbonate standards
(ETH
1–4; Bernasconi et al., 2021) using a pooled-regression approach
(Daëron, 2021). Full
analytical errors are derived from the external reproducibility of unknowns
and standards (Nf= 89) and conservatively account for the uncertainties
in raw Δ47 measurements, as well as those associated with the
conversion to the absolute Δ47 reference frame.
Results
All geochemical results obtained from modern brachiopods (δ13C,
δ18O, Δ47, Mg/Ca, Sr/Ca, Na/Ca and Li/Ca) are
reported in Table 2. Linear regression models of Δ18Oc-w,
Δ47, Mg/Ca, Sr/Ca, Na/Ca and Li/Ca, alongside mean annual temperature
(MAT), are reported in Fig. 1 and Table 3.
Geochemical data of modern-brachiopod samples. Stable isotope
values and element / Ca ratios.
Geochemical data and linear regression model for each geochemical
parameter tested for its relationship with brachiopod living temperature.
Parameters associated with each regression model are listed in Table 3.
Carbon and oxygen stable isotopes
The whole dataset of δ13C and δ18O values (n=73) from
modern-brachiopod shells ranges from -2.2 ‰ to
3.2 ‰ and from -1.2 ‰ to
4.0 ‰, respectively. There is a robust negative
linear correlation (R2>0.73, p slope < 0.001) between Δ18Oc-w (=δ18Oc-δ18Ow) and MAT (Fig. 1; Table 3). The correlation
observed in bulk shell samples is indistinguishable from that in inner-layer
samples (Fig. 1a). On the contrary, the correlation obtained from outer-layer samples shows a significant offset in relation to both inner-layer and bulk-layer samples. This is in line with the differences between the outer and
inner layers observed within the same specimen, which are in most cases
above analytical uncertainties and range from -3.7 ‰ to
-0.2 ‰ (mean =-1.4 ± 0.4 ‰; 2σ; n=17) and -0.8 ‰ to 0.5 ‰ (mean =-0.4±0.3 ‰; 2σ; n=17) for δ13C and δ18O, respectively. Our new set of data from
inner- and bulk-layers samples provides the following oxygen isotope
fractionation equation:
T=-5.0(±0.2)(δ18Ocδ18Osw)+19.4(±0.4),
with T being the temperature in ∘C and Δ18Oc-w (=δ18Oc-δ18Ow) being the oxygen isotope
fractionation between brachiopod calcite and seawater, excluding data from
the outer-shell layers. As opposed to the ordinary least-squares regression
used in Fig. 1 and Table 3, this equation is established using a York
regression (York et al., 2004) using seasonal variation for
temperature uncertainties and propagated uncertainties from δ18Oc measurements and δ18Ow estimates for
Δ18Oc-w uncertainties.
Trace element <inline-formula><mml:math id="M638" display="inline"><mml:mo>/</mml:mo></mml:math></inline-formula> Ca ratios
The Mg/Ca, Sr/Ca, Li/Ca and Na/Ca ratios of modern-brachiopod shells range
from 3.57 to 25.30 mmol / mol, 0.45 to 2.05 mmol / mol, 3.28 to 58.47 µmol / mol and 1.86 to 18.28 mmol / mol, respectively (N=47). Sr/Ca ratios show
significant but weak negative correlations (0.28<R2<0.44; p slope < 0.01) with temperature. Li/Ca ratios show
significant and good negative correlations (0.61<R2<0.86; p slope < 0.001) with temperature, with the strongest
correlation (R2= 0.86) being reached in outer-layer
samples. Na/Ca ratios show significant negative correlations (0.38<R2<0.60; p slope < 0.001) with temperature.
The concentrations of these three elements (Sr, Li and Na) in brachiopod shells
are strongly positively correlated with each other (0.66<R2<0.93 for the whole dataset), especially when
considering only inner-layers sample (0.84<R2<0.96). No significant correlation (p slope > 0.2) of
Mg/Ca with temperature is observed in the dataset (Fig. 1; Table 3).
Measured element / Ca ratios are systematically higher in the outer layers
than in the inner layers of brachiopod shells (Student's t test: p< 0.01; Fig. 1c–e).
Clumped isotopes
Δ47 values obtained from the selected brachiopod shells range from
0.59±0.01 ‰ to 0.69±0.01 ‰ I-CDES and are
strongly correlated with ambient temperature (Table 3; Fig. 1b).
Temperatures inferred from clumped isotope data were calculated using the
INTERCARB calibration determined by Anderson et al. (2021), who
reprocessed the data coming from various laboratories using the same
carbonate standards for data correction. Clumped isotope temperatures are
significantly lower than estimated MAT, with a mean deviation from
environmental temperatures of -2.4±1.3∘C (95 % CI (confidence interval);
n=19).
Interpretation and discussionBrachiopod shell paleothermometersValidity and robustness of our <inline-formula><mml:math id="M678" display="inline"><mml:mrow class="chem"><mml:msup><mml:mi/><mml:mn mathvariant="normal">18</mml:mn></mml:msup><mml:mi mathvariant="normal">O</mml:mi><mml:msup><mml:mo>/</mml:mo><mml:mn mathvariant="normal">16</mml:mn></mml:msup><mml:mi mathvariant="normal">O</mml:mi></mml:mrow></mml:math></inline-formula>-based fractionation equation
The oxygen isotope fractionation equation derived from our data deviates
substantially from that determined by Brand et al. (2019), which is also derived from articulated brachiopod shells (Fig. 2). The two equations follow
similar trends and are only offset by ∼ 0.6 ‰ at tropical temperatures (20–30 ∘C; Fig. 2a), with the equation from Brand et al. (2019) predicting a lower
fractionation factor. At temperate and polar temperatures (20 to
0 ∘C), our new equation highlights a higher sensitivity of oxygen
fractionation to temperature than that of Brand et al. (2019; Fig. 2).
The two equations overlap at around 10 ∘C and deviate
substantially around ∼ 0 ∘C, with a Δ18Oc-w difference of up to ∼ 1 ‰
between the equation of Brand et al. (2019) and our equation (Fig. 2a), a
pattern also observed in the geochemical datasets (Fig. 2b). Examining the
geochemical dataset revealed that the large offset between the two equations
at low temperatures partly results from the use of a second-order
regression model by Brand et al. (2019), as the regression curve overlaps
with the highest Δ18Oc-w data points at temperatures of
∼ 1 ∘C. We preferred to use a linear regression that was more
appropriate to data in the low-temperature (<50∘C)
linear domain of the hyperbolic curve (α=f(1/T2)) predicted by equilibrium thermodynamics. Still, the main discrepancies
between the equation proposed herein and that of Brand et al. (2019) are
mainly driven by a difference in the geochemical dataset. The case of the
polar brachiopods Liothyrella uva from the island of Rothera (Antarctica) highlights the likely
reasons for this discrepancy at low temperatures. Indeed, δ18O
values reported in both studies are the same (Brand et al., 2019),
indicating that the discrepancy lies in the environmental parameters used to
derive the equation. Brand et al. (2019) used published environmental
parameters measured locally in late February for this site, while we
approached this estimation using regional annual averages. This difference
in integration time of the environmental parameters is important because
the month of February records the annual maximum and minimum in seawater
temperature and δ18Osw values, respectively
(Meredith et al., 2013). This particular
brachiopod species has a higher growth rate during the Austral winter than
during the Austral summer (Peck et
al., 1997), further increasing the likely seasonal bias. These discrepancies
are individually relatively small (± 1–2 ∘C; ±0.5 ‰ VSMOW), but put together, they easily explain the
1 ‰ offset between both equations in the low-temperature
range.
Plot comparison of (a) regression lines for different
temperature–oxygen isotope fractionations for brachiopod calcite (this study;
Brand et al., 2019), laboratory-precipitated calcite (Kim and O'Neil, 1997)
and slow-growing cave calcite (Daëron et al., 2019); and (b) corresponding
datasets for brachiopod calcite (modified after Brand et al., 2019).
Uncertainty envelope of 95 % CI of the York least-squares regression
(York et al., 2004).
We use a published independent dataset of modern-brachiopod δ18O values (Bajnai et al., 2018) to test our equation
against other published equations in their ability to predict environmental
temperatures when δ18Osw is known (Fig. 3). This dataset
has a MAT range of 0 to 29 ∘C and uses two sets of δ18Osw for most data, one derived from direct measurements and one
derived from δ18Osw–salinity relationships
(LeGrande and Schmidt, 2006). Plus, Mg/Ca values are
available for all samples, allowing us to apply the equation of Brand et al. (2013) to the dataset. Isotopic temperatures were calculated and then normalized
as the temperature deviation from environmental temperature (ΔT in
∘C) with full propagation of uncertainties. Of all the equations
tested, the equations derived from measurements of modern brachiopods (Brand
et al., 2013, 2019; this study) yield the most accurate temperature
predictions. Mean ΔT are statistically indistinguishable from 0,
using salinity-based δ18Osw estimate results for all four
equations. Interestingly, the accuracy of isotopic temperature estimates is
poorer when measured δ18Osw values are used. The equations
of Brand et al. (2013, 2019) now significantly underestimate growing
temperatures (Fig. 3; p(meanΔT=0)<0.01),
while our equation yields mean ΔT values that are statistically
indistinguishable from 0 (p(meanΔT=0)=0.11). The
equation of Kim and O'Neil (1997), derived from synthetic
carbonates, and that of Epstein et al. (1953),
derived from molluscs, both underestimate the growing
temperature of brachiopod shells (Fig. 3; p(meanΔT=0)<0.001) in
both δ18Osw hypotheses. On the other end of the spectrum,
the equation of Daëron et al. (2019), derived from slow-growing calcite,
overestimates the growing temperature of brachiopod shells (Fig. 3, p(meanΔT=0)<0.001) in both δ18Osw
hypotheses. Although it lessens the accuracy of δ18O
temperatures in this dataset, the use of locally measured δ18Osw values over the salinity-based estimates improved their
precision (4.0<2σ<5.3∘C, n=13
rather than 5.4<2σ<6.4∘C, n=18).
Offsets between the δ18O temperatures and MAT from
the modern-brachiopod dataset of Bajnai et al. (2018) using different
fractionation equations. Circles: temperature offset for each sample;
vertical transparent bars: temperature offset 95 % confidence interval for
each sample; horizontal bar: temperature offset arithmetic mean and
associated 95 % confidence interval of the mean. (a) Temperature offset for
isotopic temperatures calculated with δ18Osw values
calculated from the LeGrande and Schmidt (2006) database; n=18. (b) Temperature offset for isotopic temperatures calculated with locally
measured δ18Osw values; n=13. OLS refers to ordinary least-squares regression; YLS refers to York least-squares regression (York
et al., 2004).
Independent from the fractionation equation or the δ18Osw
value used, temperature estimates are associated with large scatter of more
than 4 ∘C (2σ) around mean ΔT. We suggest that
this scatter is related primarily to the variability of brachiopod shell
δ18O values observed within a population (Fig. 2; Brand et
al., 2019) due to isotopic fractionation mechanisms other than temperature-dependent fractionation, namely kinetic effects, metabolic effects and pH
effects. Using the MgCO3-corrected equation of Brand et al. (2013)
instead of that of Brand et al. (2019), both derived from very similar datasets, has
the only effect of slightly improving the precision of the estimates,
decreasing the scatter by ∼ 0.6 ∘C around the mean
deviation from environmental temperatures. This effect is about 1 order of
magnitude below the remaining dispersal of data.
For the different equations
derived from modern-brachiopod data, we push further the comparison of the deviation between isotopic temperature calculated with the different equations and environmental temperatures and look at the distribution of these offsets (ΔT) against seawater temperature (Fig. 4).
This highlights a significant correlation between ΔT and seawater temperature when applying both the equations of Brand et al. (2013, 2019) to the comparative dataset of Bajnai et al. (2018; puncorrelated<0.001; Fig. 4). On contrary, there is no
significant correlation between ΔT and seawater temperature when
applying both our new equations to the comparative dataset of Bajnai et al. (2018; puncorrelated>0.3). This comparison highlights that
the application of the equations of Brand et al. (2013, 2019) does not
completely account for temperature as an explicative variable for brachiopod
Δ18Oc-w in the dataset of Bajnai et al. (2018).
Distribution of isotopic temperature deviation from living
temperature against living temperature using the equations of (a) this
study, (b) Brand et al. (2013) and (c) Brand et al. (2019) in relation to the dataset of
Bajnai et al. (2018).
In the following, we further assess the impact of trace element
incorporation in brachiopod calcite on the fractionation factor between
calcite and water using our new data. Brachiopod calcite can be viewed as a
three-component molar mixture of CaCO3, MgCO3 and SrCO3, and
the fractionation factor of this mixture (Δbrach.calcite–water)
of divalent metal carbonates is equivalent to the sum of the mineral–water
fractionation factors weighted from the mineral molar fractions. This bulk
isotopic fractionation factor is thus expressed as follows:
Δ(brach.calcite–water)=∑i=1nXi⋅Δ(mineral–water),
with the n (=3) minerals noted as i and X being the molar fraction of each mineral
constituting the brachiopod calcite mixture. As, to our knowledge,
a MgCO3–water oxygen isotope fractionation equation has not yet been
established from synthetic carbonates, we used the theoretical fractionation
factors (α(x-y)) of Chacko and Deines (2008) to compare the mixture and pure calcite fractionation factors.
Δ(brach.calcite–water) values, calculated for brachiopod inner calcite
layers with the same range of CaCO3, MgCO3 and SrCO3 contents as that
recorded by our elemental data, do not exceed 0.18 ‰
compared to the calcite end member at environmental temperatures (mean = 0.08 ‰; n=24). This difference is mainly explained by
Mg2+ incorporation, as (1) MgCO3 is more abundant than SrCO3 in
brachiopod calcite; and (2) 1000lnα(MgCO3–water) is higher than
1000lnα(CaCO3–water) by about 13 ‰ in this
temperature range, whereas 1000lnα(SrCO3–water) is lower than
1000lnα(CaCO3–water) by only ∼ 1 ‰
(Chacko and Deines,
2008). This basic calculation illustrates that Mg2+ incorporation into
brachiopod shells with low Mg calcite can affect calcite δ18O by no more
than 0.2 ‰, which is about twice the range of analytical
uncertainties at 2σ. As a consequence, the incorporation of
Mg2+ has a weak effect on the δ18O of brachiopod calcite.
The MgCO3 correction proposed by Brand et
al. (2013) is not really needed for paleoclimatic and paleoceanographic
studies, which most likely justifies why it was later abandoned (Brand et al., 2019).
Brachiopod shell clumped isotopes: an alternative paleothermometer?
Although there is a strong correlation between our new Δ47
values from brachiopod shells and growing temperatures, Δ47
values at a given temperature are generally higher than what is expected
from the state-of-the-art equation of Anderson et al. (2021; Fig. 5).
This observation is in line with the results of Bajnai et al. (2018) gathered in the CDES reference frame. To discuss modern-brachiopod
Δ47 values in relation to a larger dataset, the data from
Bajnai et al. (2018) were adjusted to the CDES 90 reference
frame for better comparison with our dataset in the I-CDES reference frame (Bernasconi
et al., 2021). We acknowledge that the different standardization protocols
used for the two datasets could hamper the strength of this comparison.
However, this comparison is supported by the good agreement between sample
143 from Bajnai et al. (2018; Δ47=0.670±0.008 ‰ CDES 90) and its replicate sample Mv143b
(Δ47=0.664±0.006 ‰ I-CDES; Bajnai et al., 2020; Fiebig et
al., 2021). The Δ47–temperature relationship derived from this
combined brachiopod dataset yields the following equation:
Δ47(I-CDES90∘C)=0.042±0.002×106T2+0.124±0.016,
with Δ47(I-CDES90∘C) in ‰ and T being the temperature in K. This equation, while still nearly within the error
of the equation of Anderson et al. (2021; Fig. 5b), highlights
systematically higher brachiopod Δ47 values relative to what is
expected from the canonical clumped isotope equation.
Comparison of Δ47 values from modern brachiopods with the
state-of-the-art Δ47 calibration of Anderson et al. (2021).
Envelope and error bars are at 2σ. (a) Comparison of York least-squares
regressions (York et al., 2004) derived from two modern-brachiopod datasets (Bajnai et al., 2018; this study) with the equation of
Anderson et al. (2021). (b) Comparison of York least-squares regression
derived from the combined modern-brachiopod dataset with the equation of
Anderson et al. (2021). (c) Comparison of modern-brachiopod Δ47 temperatures calculated using the equation of Anderson et al. (2021) with seawater temperatures.
A similar underestimation of Δ47 temperatures using the
equation of Anderson et al. (2021) was recently reported on culture-grown
bivalves (de Winter et al., 2022). The
authors argued that the Δ47-1–T2 relationship is
non-linear, implying that equations including warm (>100∘C) data points, such as that of Anderson et al. (2021),
systematically underestimate temperatures in the 30 to -2∘C
range (de Winter et al., 2022). To
test whether the equation of Anderson et al. (2021) is responsible for the
cold bias reported here, we compare our data with three other equations
established in the marine temperature range within the I-CDES reference
frame (Meinicke
et al., 2021; Huyghe et al., 2022; Peral et al., 2022). While applying the
equation of Meinicke et al. (2021) slightly reduced the mean Δ47–temperature offset relative to brachiopod living temperature
(-2.4±1.2∘C; 2 SE, n=37) compared to the equation of Anderson
et al. (2021) (-3.9±1.2∘C; 2 SE, n=37), applying of
the equation of Peral et al. (2022; -3.8±1.2∘C; 2 SE,
n=37) or of Huyghe et al. (2022; -3.0±1.2∘C; 2 SE,
n=37) gave similar offsets (see File S3 in the Supplement for a detailed comparison).
This comparison shows that equations derived from samples precipitated
exclusively in the marine temperature range give very similar offsets and
does not support the hypothesis that the offset is due to very warm data
points (>30∘C) influencing the slope of the linear
regression of Anderson et al. (2021). We also argue that the difference
between the equations of Peral et al. (2018, 2022) and Meinicke et al. (2020, 2021), both based on foraminiferal calcite, can be explained by
different ways of estimating calcification temperature rather than by different
observations, as discussed by Meinicke et al. (2020). At least a fourth of
brachiopod calcite samples show statistically significant deviation from
previously published equations (9 to 14 of the 37 samples considered here,
depending on the equation). We conclude that the Δ47 offsets
observed in brachiopod calcite reflect primary deviations from previous
observations that should be addressed to prevent temperature biases when
using Δ47 values of fossil brachiopods as paleotemperature
proxies.
Our new data suggest that, at a given temperature, brachiopod shell Δ47 values are higher than what has been observed among foraminifera
(Breitenbach
et al., 2018; Peral et al., 2018, 2022; Daëron et al., 2016; Meinicke et
al., 2020, 2021), laboratory-precipitated calcites (Jautzy
et al., 2020; Anderson et al., 2021; Fiebig et al., 2021), slow-growing cave
calcites (Daëron
et al., 2019; Anderson et al., 2021; Fiebig et al., 2021) and bivalve
calcite (Huyghe et al., 2022). This
observation is more relevant in the Bajnai et al. (2018) dataset, with a mean
Δ47 temperature deviation from seawater temperatures of -5.6±1.6∘C (2 SE, n=18), than in our new dataset, with a mean
Δ47 temperature deviation from seawater temperatures of -2.4±1.3∘C (2 SE, n=19; Fig. 5a). However, we argue that
the difference between the two datasets of modern-brachiopod Δ47 may largely result from distinct taxonomic assemblages and/or
sampled locations in both datasets. Indeed, our new Δ47 data
from New Zealand brachiopods agree very well with previous data from the
same brachiopod species and similar locations (Bajnai et al.,
2018; see File S3 in the Supplement). The Bajnai et al. (2018) dataset is dominated by
brachiopods from the suborder Terebratellidina (12/18), a group for which
the combined dataset shows the most deviation from the Anderson et al. (2021) equation at low (-2 to 5 ∘C) temperatures but most importantly in the
middle (8–15 ∘C) temperature range (Fig. 5c). This group is
associated with a mean Δ47 temperature deviation from seawater
temperatures of -5.6±1.7∘C (2 SE, n=19) when applying
the Anderson et al. (2021) equation, while the other most represented group
in the combined dataset, the suborder Terebratulidina, better fits the
canonical equation (Fig. 5c) with a mean Δ47 temperature
deviation from seawater temperatures of -1.8±1.4∘C (2 SE,
n=13). This apparent taxonomic difference among brachiopod groups suggests
that the deviation observed for brachiopod shell Δ47 values
relative to other calcifying organisms relates to processes other than
ambient temperatures, controlling bound ordering within the calcite crystal
of brachiopod shells (Bajnai et al., 2018). Such processes
still need to be confidently identified (Sect. 4.2) but likely lie in
the different ways the shell is secreted under different growth rate dynamics
and/or different crystal arrangements. Available brachiopod Δ47
values suggest that, while the equation of Anderson et al. (2021) could be
confidently applied to brachiopods of the suborder Terebratulidina (Fig. 5), this would not be the case for species of the suborder Terebratellidina.
For the other brachiopod groups, data are still too scarce to propose an
enlightened conclusion.
<inline-formula><mml:math id="M852" display="inline"><mml:mrow class="chem"><mml:mi mathvariant="normal">Mg</mml:mi><mml:mo>/</mml:mo><mml:mi mathvariant="normal">Ca</mml:mi></mml:mrow></mml:math></inline-formula> thermometer for brachiopod shells
Our Mg/Ca data obtained from the inner layers of brachiopod shells do not
show any significant correlation with growth temperature (p=0.94; Fig. 1c; Table 3), at odds with the strong correlation between the Mg/Ca molar
ratio of the brachiopod shell secondary layer and the growth temperature found
by Brand et al. (2013, 2019). However, these authors acknowledged that some species do not
follow the trend they illustrated as the GBMgL and suggested that such
deviations resulted from either lower growth rates (Brand et al., 2013) or higher growth rates (Brand et al., 2019) of the
considered species relative to other brachiopods. This issue alone
illustrates how, from the largest dataset of modern brachiopods available so
far (i.e. Brand's database), one cannot firmly determine whether or not
Mg/Ca correlates with seawater temperatures. We emphasize that our dataset
lacks brachiopod shell samples from equatorial surface marine waters
(27 ∘C <T< 32 ∘C), which record the
highest MgCO3 molar content in Brand's database (Brand et al.,
2013, 2019). Interestingly, the highest MgCO3 contents reported by
these authors were measured on brachiopods belonging to the order
Thecideida. This taxonomic singularity is of prime importance because (1) Thecideide brachiopod shells are essentially made of a shell fabric unique
to its group (Simonet Roda et al., 2022), and
secondary shell fabric (fibrous layers, which form the majority of
Rhynchonellida and Terebratulida shells) is restricted to small areas of the
shell, such as the teeth and the inner sockets (Williams,
1968, 1973; Baker, 2006); (2) Thecideide shell is formed from high-Mg calcite
(Ullmann et al., 2017). The exclusion of Thecideida
data from the database published by Brand et al. (2013) produces a
substantially weaker (R2= 0.23) yet significant (p<0.01) correlation between shell Mg/Ca and temperature. In
addition, magnesium has also been shown to be more concentrated in major
growth bands (Gaspard
et al., 2018; Müller et al., 2022) than in other parts of the secondary
shell layer in articulated brachiopods, which could explain the high Mg/Ca
values observed from slow-growth and long-lifespan brachiopods
(Brand et al., 2019), especially
when considering samples obtained from a large sampling area covering
multiple growth lines. Given our new data and all these considerations, the
brachiopod Mg/Ca should be regarded as an unreliable proxy for annual or
seasonal temperature reconstructions, at least in the -2 to 25 ∘C
range (Fig. 1; Brand et al.,
2013, 2019).
An alternative or complementary way to explain the scattering of
temperature–Mg/Ca paired data may arise in the variability of marine Mg/Ca.
All the elements considered in this study (Mg, Sr, Li and Na) have much
larger residence times (10, 3.5, 1.8 and 44 Myr, respectively) than water
molecules in the oceans, meaning that, at first order, they are uniformly
distributed in the main bodies of the oceans
(Lécuyer, 2016). A recent study, however,
has shown that sizable variations in element / Ca may occur in the vicinity of
continental margins under the influence of oceanic currents at the
interface with groundwaters or the atmosphere or within water masses
isolated from the global thermohaline circulation (e.g. Black Sea,
Mediterranean Sea, Red Sea; Lebrato et al., 2020).
For example, despite the large residence time of Sr, non-homogeneous marine
strontium isotopic ratios (87Sr/86Sr) have been measured in
coastal environments (southern Okinawa Trough, South China Sea and Kaoping
Canyon) as the result of seawater mixing with groundwaters
(Huang et al., 2011;
El Meknassi et al., 2020). Given the poor constraint on local coastal water
chemistry in the deep time, brachiopod samples from offshore environments
should be preferred for trace-element-based paleotemperature
reconstructions, as was suggested for marine biogenic apatite by
Balter and
Lécuyer (2010).
Potential of other trace element <inline-formula><mml:math id="M876" display="inline"><mml:mo>/</mml:mo></mml:math></inline-formula> Ca proxy
Our new brachiopod Sr/Ca data show a significant but weak negative
correlation with ambient temperature (pslope<0.001;
R2<0.40). This correlation is supported by a weak
positive correlation between Sr/Ca and Δ18Oc-w of the
inner layers (pslope<0.01; R2=0.31), as we
also expect Δ18O to decrease with increasing temperatures (i.e.
Sect. 4.1). This contradicts the pioneering findings of
Lowenstam (1961), which suggested a negative
correlation between brachiopod shell Sr/Ca and Δ18Oc-w.
The brachiopod database of Brand et al. (2013) shows no significant
correlation between Sr/Ca and temperatures (pslope=0.20). Given
the diverging conclusions from one dataset to another, Sr/Ca values of
brachiopod shells do not appear to be controlled by growing temperatures at
the first order.
Li/Ca in brachiopod shells is significantly correlated to temperature
(Fig. 1e; Table 3), a covariation initially reported by
Delaney et al. (1989) and confirmed by more recent studies
(Dellinger
et al., 2018; Rollion-Bard et al., 2019; Washington et al., 2020). Most of
the data presented here follow the Li/Ca–temperature trend derived from
literature data (Washington et
al., 2020), with several samples (n=9), exclusively collected from the
inner layer, showing significantly lower Li/Ca. Those low Li/Ca values lie
below the Li/Ca–temperature relationship derived from inorganic calcite (Marriott et al., 2004). Similarly, low Li/Ca values from brachiopod shells were reported by previous studies (Delaney et al., 1989;
Washington et al., 2020). Interestingly, most of those low Li/Ca values
belong to samples from the genus Tichosina (Delaney et al., 1989;
Washington et al., 2020, Table 2), so one could consider some peculiar
species-dependent trace element partitioning driven by metabolic or kinetic
processes. Caution is warranted, however, because very low Li/Ca values are
also apparent in the other tropical genus studied here, Stenosarina Cooper 1977, as well
as in one specimen of Aerothyris kerguelenensis among the nine specimens studied there.
Other environmental parameters or biological processes may control the
element / Ca incorporation into calcite. Sr/Ca, Li/Ca and Na/Ca show
significant negative correlations with the local salinity (0.37<R2<0.76; p slope < 0.01), with similar
scattering as that for the temperature relationships. Mg/Ca ratios, again, show no
significant correlation with this parameter (p slope = 0.7). A wider look at
the geochemical dataset highlights very strong correlations of Sr/Ca, Li/Ca,
and Na/Ca ratios in brachiopod shells with one another (0.68<R2<0.92; p slope < 0.01) over the whole
dataset and particularly from inner-layer samples (0.87<R2<0.97; p slope < 0.01). Correlation in
our dataset is strongest between Sr/Ca and Na/Ca and weakest between Sr/Ca
and Li/Ca. Similar correlation trends also appear at the intra-specimen
level (Romanin et
al., 2018; Rollion-Bard et al., 2019). Such strong correlations, at both
intra-individual and inter-specific scales, may suggest that the first-order
variability of those different element concentrations within brachiopod
shells may be explained by similar processes, either environmental or
biological. Ullmann et al. (2017) reported a distinct geochemical signature
between brachiopods species, especially within the Mg/Ca–Sr/Ca space. Our
inner-shell-layer dataset presents a similar distribution within the low-Mg/Ca range (<10 mmol / mol), where Sr/Ca ratios of Terebratulidina
shells are distinctly lower than those of Terebratellidina shells. Owing to
the strong correlation between Sr/Ca, Li/Ca and Na/Ca, our dataset
differentiates between Terebratulidina shells with low element / Ca ratios (0.4<Sr/Ca<0.7 mmol / mol; 0<Li/Ca<10µmol / mol; 1<Na/Ca<4 mmol / mol) and Terebratellidina
shells with high element / Ca ratios (0.9<Sr/Ca<1.2 mmol / mol;
35<Li/Ca<60µmol / mol; 8<Na/Ca<12 mmol / mol), with the exception of the Antarctic brachiopod Liothyrella uva (Broderip,
1833; Terebratulidina) which shows the highest Sr/Ca, Na/Ca and Li/Ca values
(Fig. 6a–b). With this exception and that reported by
Ullmann et al. (2017), a taxonomic relationship is
not satisfactory, especially considering the significant yet noisy
relationship between element / Ca and environmental parameters such as
temperature and salinity within our new dataset. Indeed, we also observe a
strong relationship between brachiopod shell element / Ca values and
geographical location (Fig. 6c–d). Tropical shells (Guadeloupe and New
Caledonia – only Terebratulidina) have low element / Ca values, the middle
latitudes (New Zealand and Crozet Islands – dominated by Terebratellidina)
have intermediate to high element / Ca values, and polar brachiopods
(Antarctica and Norway – mixed assemblage) have high element / Ca values.
Despite this first-order distribution, it appears that, within the same
environment, species difference appears to dominate, as is the case in
Doubtful Sound (New Zealand), where Terebratella sanguinea (Leach, 1814; Terebratellidina) shells
have higher element / Ca values than Liothyrella neozelanica (Thomson, 1918; Terebratulidina) shells.
Distribution of trace element content of the inner shell layers
among the different brachiopod taxonomic groups (same symbols as in Fig. 5c). The data show a strong correlation between Sr/Ca and Li/Ca(a) and
Sr/Ca and Na/Ca(b), as well as strong latitudinal partitioning of Li/Ca(c)
and Sr/Ca(d) values.
Our new data suggest that temperature is not the prime parameter that can
describe trace element incorporation into articulated brachiopod shell
calcite from a dataset of widely distributed and taxonomically diverse
brachiopods. This conclusion does not exclude the possibility of any
temperature control on trace element incorporation in brachiopod shell
calcite, which may be relevant at the species level, as proposed by
Butler et al. (2015) and observed in other calcifying organisms such as bivalves (Freitas et al., 2006) and
foraminifera (Anand et al., 2003). However, the influence of other
controls on trace element incorporation in brachiopod shell calcite (growth
rate, environment fluid chemistry, etc.; Jurikova et al., 2020)
is strong enough to blur the likely temperature control so that the trace
element / Ca ratios studied here are not suited to being used as
paleothermometers.
Precipitation of brachiopod shell calcite out of equilibrium with
seawater
Our new data confirm that brachiopod shell calcite does not precipitate in
isotopic equilibrium with seawater, as highlighted by the significant
deviation from thermodynamic isotopic equilibrium relationships constrained
from very-slow-growing calcite (Daëron
et al., 2019; Anderson et al., 2021, Figs. 2–5). It is worth noting that
most brachiopod shell δ18O data (Brand et al., 2019; this
study) plot in between thermodynamic equilibrium isotopic fractionation
(Daëron et al., 2019) and the equation of Kim and
O'Neil (1997; Fig. 2). The deviation of brachiopod shell calcite δ18O from thermodynamic equilibrium (Δ18Oeq=δ18Obrach.calcite-δ18Oequilibrium)
is obvious when we compare our dataset to the equation of Daëron et al. (2019) in the clumped isotope–oxygen isotope fractionation space (Fig. 7a),
with most of our dataset showing higher Δ47 and lower δ18O than predicted (Fig. 7b), a feature previously highlighted by
Bajnai et al. (2018). As a result, brachiopod calcite shows
apparent Δ47 and δ18O temperatures that are, respectively,
lower and higher than those predicted by the thermodynamic equilibrium. This
deviation from equilibrium values has a Δ47–Δ18Oeq slope of 0.019±0.005 in our new brachiopod dataset,
indistinguishable from the slope of 0.017±0.003 reported by Bajnai et
al. (2018) using δ18Osw values from the δ18Osw database (LeGrande and Schmidt, 2006).
Non-equilibrium fractionation and kinetic effects in brachiopod
shell calcite as evidenced by clumped isotopes and oxygen isotopes. (a) Plot
of brachiopod inner shell layers in the Δ47–1000lnα(calcite–water) space compared with the equilibrium line (blue line)
defined by very-slow-growing abiogenic calcite (Daëron et al., 2019;
Anderson et al., 2021). The red highlight shows the results from A. kerguelenensis (Crozet Islands) and
their expected geochemical values on the equilibrium line. Note the negative
correlation between Δ47 and 1000lnα(calcite–water) evidenced in this subset. (b)δ18Oc
and Δ47 offset from their expected equilibrium values
calculated from environmental parameters using the equations of Daëron
et al. (2019) and Anderson et al. (2021), respectively. The trends followed
by these offsets within our new dataset are similar to those reported from
other brachiopod samples by Bajnai et al. (2018).
Brachiopod shell Δ18O data (Brand
et al., 2003, 2013, 2019; this study) also show a large geochemical
variability among brachiopods sampled from similar marine environments (T,
S), with a Δ18O range often higher than 1 ‰
(Fig. 3; Brand et al., 2019). Isotopic fractionation equations (Brand et
al., 2013, 2019; this study) usually rely on Δ18Oc-w
values plotted against mean annual temperatures. Thus, it would be tempting
to attribute such scattering of data to seasonal variations in both
temperature and δ18Osw, but this would only apply to
specimens that lived above or in the upper part of the thermocline. Such
isotopic scattering is indeed expected if we consider datasets based on
multiple sampling along the growth axis of an individual (Yamamoto et al.,
2011; Takayanagi et al., 2015), such as in the case of Brand et al. (2013,
2019). However, seasonal scattering in data resulting from intra-individual
geochemical variability is unlikely for our measurements, which are performed on a
large sampling area that should correspond to multiple months or years of
growth.
For example, we note a large isotopic variability in both δ18O
and Δ47 among the nine specimens of Aerothyris kerguelenensis from the Crozet Islands even
though the specimens were recovered from depths in the 105 to 355 m range
with a narrow range of temperature (3–5 ∘C) and salinity
(33.8–34.2). These nine samples, all collected from the shell inner layers of
different individuals, record a Δ18O variability of
2 ‰, which is equivalent to a temperature range of 8 to
10 ∘C if interpreted as such, and a Δ47 variability
of 0.02 ‰, which is equivalent to a temperature range of about
5 ∘C. This subsample shows a significant Δ47–Δ18Oeq slope of 0.027±0.007 (2σ;
pslope<0.01; R2=0.95, n=5), as well as a
strong positive correlation between Δ18Oc-w and δ13C (pslope<0.01; R2=0.69, n=9).
These covariations most likely result from kinetic effects as defined by
McConnaughey (1989), here recorded at the intra-specific level.
Our new data replicate the observations made by Bajnai et al. (2018) concerning kinetic fractionation in brachiopod shells. In
consequence, we follow their conclusions and consider it to be the case that different growth
rates of the specimens are responsible for the second-order isotopic
variability and the deviation from the stable isotope and clumped isotope
equilibrium observed in brachiopod datasets, both at an inter-specific and
intra-specific level.
Neither minor nor trace elements are incorporated into brachiopod shell
calcite in equilibrium with seawater relative to abiogenic calcite
(Rollion-Bard et al.,
2019; Simonet Roda et al., 2022). Kinetic effects are known to influence
trace element incorporation into calcite (Lorens,
1981; Busenberg and Plummer, 1985; Tesoriero and Pankow, 1996; Gabitov et
al., 2011, 2014) and may be a process, among others
(Jurikova et al., 2020), that explains trace element
incorporation into brachiopod shell calcite. Owing to the very limited
knowledge of brachiopod shell growth rates (for only 5 out of 17 species in our
dataset), we are unable to directly test this hypothesis. However, if
kinetic effects have such an influence on element incorporation into brachiopod
shell calcite, we should observe strong negative correlations of Sr/Ca, Na/Ca and Li/Ca ratios with Δ18Oeq. We observe such significant
correlations in the new dataset in both the inner- and outer-layer subsets
(0.40<R2<0.55, p slope < 0.01),
with a significant slope change between the two subsets. The variation in
element / Ca related to a similar variation in Δ18Oeq, as
suggested by linear regression models, is about twice as high in the inner-layer subset than in the outer-layer subset. This difference may be
explained by the different pathways brachiopods use to form their different
shell layers (acicular, fibrous and prismatic/columnar;
Simonet Roda et al., 2022). However, while looking
at Aerothyris kerguelenensis, which showed the most important kinetic effects according to stable
isotope data, there is no significant correlation of element / Ca ratios with
Δ18Oeq. Element / Ca ratios appear to be independent from
isotope kinetic effects at the species level despite strong correlations of
element / Ca ratios and Δ18Oeq in the multi-specific
dataset.
In line with the significant correlation between Δ18Oeq and element / Ca values in brachiopod shell inner layers, the environmental
and taxonomic distributions of isotopic offsets (Δ18Oeq,
Δ47 offset) in our dataset are similar to the distribution of
element / Ca values. We observe more deviation from the equilibrium in high-latitude environments than in tropical environments, and we generally observe
more deviation from isotopic equilibrium in Terebratellidina than
in Terebratulidina in the whole dataset, although the difference is not as
striking as that documented for trace elements. As the geographical
distribution of the two suborders in the new dataset is highly biased, this
trend cannot be confidently interpreted as arising from different shell
formation processes in the different groups. However, several examples of
Terebratellidine and Terebratulidine living in the same environment show a
similar pattern. In Doubtful Sound (New Zealand), Terebratellasanguinea (Terebratellidina)
deviates more from equilibrium isotopic composition than Liothyrella neozelanica (Terebratulidina; this study; Bajnai et al., 2018). In Friday Harbor (Washington, USA),
Terebratalia transversa (Sowerby, 1846; Terebratellidina) records lower and more variable δ18Oc than Terabratulina sp. (Terebratulidina; Ullmann et
al., 2017). These authors also reported high isotopic variability associated
with the strong correlation of δ18O with δ13C among
Terebratellidine but not among Terebratulidine.
From the relationship between the different parameters studied here, we
suggest that kinetic effects in brachiopod shells may be a major source of
deviation from calcite precipitation in equilibrium with seawater. Kinetic
effects can be, on first approximation, controlled by shell growth rates
(McConnaughey, 1989; Bajnai et al., 2018). Therefore, both
biological and environmental parameters should control the extent of the
kinetic effects. At first order, one can expect species growing larger
shells to record more important kinetic effects. Moreover, shell growth
rates in calcifying organisms increase when environmental conditions are
close to optimal living range and decrease when the conditions strongly
derive from that optimum, a pattern that is dependent on taxon-related
preferences (Schöne, 2008). However, the poor knowledge of
brachiopod shell growth rate, relative to other marine calcifying organisms,
prevents an in-depth appreciation of its exact role. Kinetic effects are
still insufficient to describe all of the deviation from equilibrium observed
here, but our dataset does not allow for a discussion of other potential factors
(Jurikova et al., 2020).
Conclusions
Our multi-proxy (δ18O, δ13C, Δ47,
Mg/Ca, Sr/Ca, Na/Ca and Li/Ca) geochemical analyses of 37 modern brachiopods
from different locations, spanning living temperatures from -1.9 to 25.5 ∘C, allowed us to investigate the temperature
relationship of several geochemical markers and to evaluate or re-evaluate
their potential as paleothermometers for brachiopod shell archives. δ18O and Δ47 are the most reliable paleothermometers,
although we propose a revised δ18O paleotemperature equation
for brachiopod shell calcite, and Δ47 values highlight a
significant deviation from the canonical temperature relationship. The
element / Ca ratios investigated here should not be used to infer
environmental temperatures from brachiopod calcite owing to either the
absence (Mg/Ca) or a poor yet significant (Sr/Ca, Li/Ca and Na/Ca) correlation
with living temperature. Our new dataset confirms previous evidence for the interpretation of non-equilibrium isotopic fractionation interpreted as significant kinetic
effects recorded in the brachiopod shell inner layers. In general, the
precipitation of brachiopod shells that takes place out of isotopic equilibrium
leads to (1) an underestimation of living temperature from Δ47
and (2) an overestimation of living temperature from oxygen isotopes, which,
once the proxies are combined, could lead to (3) an underestimation of living-water δ18Osw relative to what would be expected from
equilibrium precipitation. Kinetic effects and, by extension, shell growth
rates could explain the second order of isotopic variability in brachiopod
shells and can be a larger source of isotopic variability than environmental
parameters (temperature, δ18Osw) at the local scale. Thus,
we advise that this variability be considered as a significant source of
uncertainties when applying paleotemperature proxies. Although we reject
their use as a paleothermometer, we do not exclude any temperature control
on trace element (Mg, Sr, Li and Na) incorporation into brachiopod shell
calcite. Rather, our findings indicate that other processes, such as kinetic
effects (growth rate) or shell microstructure and their associated
mineralization pathways, have a significant control over trace element
incorporation into brachiopod shell calcite, resulting in the high
scrambling observed in the element / Ca–temperature relationship.
Code availability
No code has been developed for the purpose of this study. Statistical analysis was done using RStudio build 492. The R package IsoplotR version 4.4 (https://github.com/pvermees/IsoplotR, Vermeesch, 2022) was used for its York regression function. Clumped isotope data were processed using the D47crunch Python package (https://doi.org/10.5281/zenodo.6300900, Daëron, 2022).
Data availability
The new dataset generated for this study is provided in the Supplement.
The WOA 2018 dataset used in this study is available on the website of the National Centers for Environmental Information: National Oceanic and Atmospheric Administration (https://www.ncei.noaa.gov/access/world-ocean-atlas-2018/, last access: 31 March 2023).
A version of the dataset to be viewed with Ocean Data View (ODV) is available at https://odv.awi.de/data/ocean/world-ocean-atlas-2018/ (Ocean Data View, 2023).
The dataset of Bajnai et al. (2018) used in this study is available from the Supplement of the original publication (10.1038/s41598-017-17353-7).
The supplement related to this article is available online at: https://doi.org/10.5194/bg-20-1381-2023-supplement.
Author contributions
TL and CL designed the study in close collaboration with GS and DG. DG
obtained the samples from various contributors. TL performed shell sampling
in the lab. AVL and TL performed the stable isotope analysis and data
treatment. TL and FAG performed the elemental analysis and data treatment.
TL and MD performed the clumped isotope analysis and data treatment. TL
performed the statistical analysis of the data and wrote the paper with
input from all the authors.
Competing interests
The contact author has declared that none of the authors has any competing interests.
Disclaimer
Publisher's note: Copernicus Publications remains neutral with regard to jurisdictional claims in published maps and institutional affiliations.
Acknowledgements
The authors gratefully thank the leaders and teams of the oceanographic vessels of the
following expeditions: the Walda cruise in the southeast Atlantic Ocean (Michel Segonzac, IFREMER, Brest); sampling from Norway coasts (courtesy of the
Trondheim Biological Station); MD 30 BIOMASS around the Crozet Islands
(Patrick Arnaud, Marseille, France; and Nadia Ameziane, MNHN, who provided the
material); Musorstom 4, SMIB 3 and 6, and Norfolk 1 around Loyauté and Lifou islands in New Caledonia (Bernard Laurin, Université de Bourgogne, Dijon; and
Bertrand Richer de Forges); Karubenthos 2 around the French Caribbean island of Guadeloupe
(Philippe Bouchet with Pierre Lozouet, Laure Corbari and Philippe Maestrati, with
the help of Jerôme Mainguy, MNHN, Paris, France) with a supplementary sampling
(Dominique Lamy, Université des Antilles et de la Guyane); additional material from
New Zealand (Daphne Lee, Marine Biology Lab, University of Otago; David I. MacKinnon, University of Canterbury, Christchurch, New Zealand; and Anthony E. Aldridge; CEAMARC off the French
Antarctic area (Nadia Ameziane and Marc Eléaume, MNHN, Paris,
France); and sampling from Palmer Peninsula, Antarctic (courtesy of Thomas Desvignes, USA).
We thank Adrian Immenhauser and two anonymous reviewers for their
constructive comments that helped improve the quality of this paper.
Financial support
This research has been supported by the Agence Nationale de la Recherche (grant no. ANR-18-CE31-0020).
Review statement
This paper was edited by Tina Treude and reviewed by Adrian Immenhauser and two anonymous referees.
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