# Carbonate "clumped" isotope signatures in aragonitic scleractinian and calcitic gorgonian deep-sea corals

J. Kimball, R. E. Tripati, R. Dunbar

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## To cite this version:

J. Kimball, R. E. Tripati, R. Dunbar. Carbonate "clumped" isotope signatures in aragonitic scleractinian and calcitic gorgonian deep-sea corals. Biogeosciences Discussions, 2015, 12, pp.19115-19165. 10.5194/bgd-12-19115-2015. insu-03580284

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# Carbonate “clumped” isotope signatures in aragonitic scleractinian and calcitic gorgonian deep-sea corals

J. Kimball¹,², R. E. Tripati¹,³, and R. Dunbar²

¹Institute of the Environment and Sustainability, Department of Atmospheric and Oceanic Sciences, Department of Earth, Planetary, and Space Sciences, University of California, Los Angeles, CA, 90095, USA

²Department of Environmental Earth System Science, Stanford University, Stanford, CA 94025, USA

³European Institute of Marine Sciences (IUEM), Université de Brest, UMR 6538/6539, Rue Dumont D'Urville, and IFREMER, Plouzané, France

Received: 27 October 2015 – Accepted: 29 October 2015 – Published: 4 December 2015

Correspondence to: R. E. Tripati (robert.tripati@gmail.com)

Published by Copernicus Publications on behalf of the European Geosciences Union.

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Deep-sea corals are a potentially valuable archive of the temperature and ocean chemistry of intermediate and deep waters. Living in near constant temperature, salinity and pH, and having amongst the slowest calcification rates observed in carbonate-precipitating biological organisms, deep-sea corals can provide valuable constraints on processes driving mineral equilibrium and disequilibrium isotope signatures. Here we report new data to further develop “clumped” isotopes as a paleothermometer in deep-sea corals as well as to investigate mineral-specific, taxon-specific, and growth-rate related effects. Carbonate clumped isotope thermometry is based on measurements of the abundance of the doubly-substituted isotopologue $^{13}C^{18}O^{16}O_2$ in carbonate minerals, analyzed in CO$_2$ gas liberated on phosphoric acid digestion of carbonates and reported as $\Delta_{47}$ values. We analyzed $\Delta_{47}$ in live-collected aragonitic scleractinian (*Enallopsammia* sp.) and calcitic gorgonian (Isididae and Coralliidae) deep-sea corals, and compared results to published data for other aragonitic scleractinian taxa. Measured $\Delta_{47}$ values were compared to in situ temperatures and the relationship between $\Delta_{47}$ and temperature was determined for each group to investigate taxon-specific effects. We find that aragonitic scleractinian deep-sea corals exhibit higher values than calcitic gorgonian corals and the two groups of coral produce statistically different relationship between $\Delta_{47}$-temperature calibrations. These data are significant in the interpretation of all carbonate “clumped” isotope calibration data as they show that distinct $\Delta_{47}$-temperature calibrations can be observed in different materials recovered from the same environment and analyzed using the same instrumentation, phosphoric acid composition, digestion temperature and technique, CO$_2$ gas purification apparatus, and data handling.

There are three possible explanations for the origin of these different calibrations. The offset between the corals of different mineralogy is in the same direction as published theoretical predictions for the offset between calcite and aragonite, although the magnitude of the offset is different. One possibility is that the deep-sea coral results

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reflect that crystals may attain nominal mineral equilibrium clumped isotope signatures only under conditions of extremely slow growth. In that case, a possible explanation for the attainment of disequilibrium bulk isotope signatures and equilibrium clumped isotope signatures by deep-sea corals is that extraordinarily slow growth rates can promote the occurrence of isotopic reordering in the interfacial region of growing crystals. We also cannot rule out a component of a biological “vital-effect” influencing clumped isotope signatures in one or both orders of coral. Based on published experimental data and theoretical calculations, these biological “vital” effects could arise from kinetic isotope effects due to the source of carbon used for calcification, temperature- and pH-dependent rates of CO₂ hydration and/or hydroxylation, calcifying fluid pH, the activity of carbonic anhydrase, the residence time of dissolved inorganic carbon in the calcifying fluid, and calcification rate. A third possible explanation is the occurrence of variable acid digestion fractionation factors. Although a recent study has suggested that dolomite, calcite, and aragonite may have similar clumped isotope acid digestion fractionation factors, the influence of acid digestion kinetics on Δ₄₇ is a subject that warrants further investigation.

# 1 Introduction

“Clumped”-isotope paleothermometry is an emerging approach for determining carbonate mineral formation temperatures (Ghosh et al., 2006; Schauble et al., 2006; Eiler, 2007), which like the oxygen isotope thermometer before it, is founded on thermodynamic predictions of the distribution of isotopes. Crucially, instead of relying on an isotopic exchange reaction between different phases (eg. CaCO₃ and H₂O), clumped isotope thermometry relies on internal isotopic exchange between isotopes in a single phase (Schauble et al., 2006). This means that, in theory, all that is needed to determine mineral formation temperatures is the clumped isotope composition of the solid, and not the water from which it grew (Schauble et al., 2006; Eiler, 2007). When considering carbonate minerals, statistical thermodynamics predicts that heavy stable
isotopes of carbon (¹³C) and oxygen (¹⁸O) will increasingly bond or “clump” in a mineral as temperature decreases and will conform to isotopic equilibrium constants for reactions such as:

$$X^{12}C^{18}O^{16}O_2 + X^{13}C^{16}O_3 \leftrightarrow X^{12}C^{16}O_3 + X^{13}C^{18}O^{16}O_2$$

where X refers to cations such as Ca²⁺, Mg²⁺, Sr²⁺, Ba²⁺. In practice, the determination of ¹³C-¹⁸O bonds in carbonate minerals is accomplished by measurement of mass-47 CO₂ (predominantly ¹³C¹⁸O¹⁶O) liberated by phosphoric acid digestion (Ghosh et al., 2006). ¹³C-¹⁸O bonding is reported as a per mil enrichment from that which would be expected in the liberated CO₂ if the sample had a stochastic distribution of C and O isotopes among all isotopologues and is designated by the parameter Δ₄₇. Following the initial calibration of synthetic calcites and some coral taxa (Ghosh et al., 2006), several studies have focused on in-depth calibrations of biogenic carbonates that represent potential paleoclimate proxies. Proxy material calibrations thus far have included aragonitic scleractinian coral (Ghosh et al., 2006; Saenger et al., 2012; Tripati et al., 2015), aragonitic deep-sea corals (Ghosh et al., 2006; Thiagarajan et al., 2011), aragonitic otoliths (Ghosh et al., 2007), calcitic and aragonitic foraminifera (Tripati et al., 2010), aragonitic mollusks and brachiopods (Came et al., 2007, 2014; Eagle et al., 2013; Henkes et al., 2013) and land snails (Zaarur et al., 2011; Eagle et al., 2013), calcitic speleothems (Affek et al., 2008; Daeron et al., 2011), bioapatite (Eagle et al., 2010), and calcitic microbialites (Petryshyn et al., 2014).

It has been noted that calibration data on different biogenic carbonates generated in the same laboratory at Caltech and using similar analytical methods produce a relationship between temperature and Δ₄₇ values that was similar to the initial inorganic calcite calibration (Ghosh et al., 2006; Eagle et al., 2013). However there are differences in inorganic calcite calibrations produced in different laboratories (Ghosh et al., 2006; Dennis and Schrag et al., 2010; Zaarur et al., 2013; Tang et al., 2014; Wacker et al., 2014; Kluge et al., 2015), that are thought to either reflect methodological differences in clumped isotope measurements or differences in how synthetic carbonates

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were precipitated (Tripati et al., 2015). It is clear from studies in mollusks (Henkes et al., 2013; Eagle et al., 2013; Douglas et al., 2014; Petrizzo et al., 2014) and brachiopods (Henkes et al., 2013; Came et al., 2014) that calibrations of the same types of materials in different laboratories can yield different relationship between temperature and $\Delta_{47}$ values.

More recent studies have discussed some possible sources of methodological effects on clumped isotope calibrations relating to data handling and mass spectrometric effects (Wacker et al., 2014; Petrizzo et al., 2014; Defliese et al., 2015). It is also possible that phosphoric acid digestion technique and temperature may be playing a role (Wacker et al., 2014; Defliese et al., 2015), as is known to be the case for conventional $\delta^{18}$O measurements (Swart et al., 1991). Nonetheless carbonate standards and a deep-sea scleractinian coral were found to yield broadly comparable results in an inter-lab comparison study between four different laboratories (Dennis et al., 2011).

To date, there is relatively little published data attempting to resolve whether different biogenic calibrations may originate from methodological differences between laboratories by simply measuring multiple different types of biogenic carbonates interspersed with each other using the same methods. Therefore it remains difficult at this time to distinguish between methodological differences and genuine differences in clumped isotope compositions of different materials, for example biologically mediated fractionations or “vital-effects”. The little work that has been done includes *Porites* corals that appears to exhibit a kinetic isotope effect that drives their $\Delta_{47}$ values significantly out of equilibrium (Ghosh et al., 2006; Saenger et al., 2012), and a temperate coral *Oculina arbuscula* cultured at the same temperature and variable pH that exhibits large kinetic effects (Tripati et al., 2015).

Here we explore $^{13}$C-$^{18}$O bond abundances in deep-sea coral species from the orders Scleractinia and Gorgonacea, that precipitate aragonite and high-Mg calcite respectively. As these corals were recovered from the same environment, analyzed by the same methods and mass spectrometer during the same analytical period, and subject to the same data processing, they represent an opportunity to explore the potential
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role of mineralogy and biology in governing clumped isotope signatures. We discuss results in the context of recent studies, including of acid digestion fractionation (De- fliese et al., 2015), and theoretical models and experimental data that have been used to constrain differences in clumped isotope composition between different dissolved inorganic carbon (DIC) species (Hill et al., 2014; Tripati et al., 2015), as it has been suggested that in cases where a carbonate mineral may inherit an isotopic signature of DIC then factors affecting DIC speciation such as pH, salinity, and temperature may also influence carbonate clumped isotope signatures (Hill et al., 2014; Tripati et al., 2015).

1.1 Deep-sea corals in paleoceanography

Deep-sea corals represent a potentially valuable archive of intermediate and deep
ocean temperatures, and have been a target for models of stable isotope fractiona-
tion. These archives could give valuable insight into the natural variability of regions of
the ocean that play an active role in large-scale climate dynamics. With fossil deep-sea
coral recorded for at least the last 225 000 years (Robinson et al., 2007; Thiagarajan
et al., 2013) it is a proxy with potential to extend our observations of ocean physics and
climate into the Pleistocene with decadal to centennial resolution. It has been estab-
lished that deep-sea corals have significant skeletal “vital” effects (disequilibrium stable
isotope fractionations) that compromise the classic δ¹⁸O paleotemperature method
(Emiliani et al., 1978; McConnaughey, 1989, 2003; Adkins et al., 2003; Smith et al.,
2000; Hill et al., 2011; Kimball et al., 2014). Attempts to circumvent these kinetic ef-
fects have focused on the “lines” method that recovers an average temperature over
the lifetime of a coral (Smith et al., 2000; Hill et al., 2011; Kimball et al., 2014) and
a Rayleigh-based multi-element (e.g. Mg/Ca, Sr/Ca, Ba/Ca) paleotemperature method
(Gaetani et al., 2011).

Despite deep-sea coral species showing significant disequilibrium in δ¹³C and δ¹⁸O values, initial calibration work using clumped-isotopes in scleractinian corals (Thiagarajan et al., 2011) revealed a good agreement between aragonitic scleractinian deep-sea
coral Δ₄₇ and the inorganic calcite calibration of Ghosh et al. (Ghosh et al., 2006). Therefore at least some species that show significant “vital effects” on δ¹³C and δ¹⁸O yield apparent equilibrium Δ₄₇ values, an observation also made on foraminifera and coccoliths (Tripati et al., 2010). Recently scleractinian deep-sea coral Δ₄₇ measurements have been used in applied paleoceanographic reconstructions (Thiagarajan et al., 2014).

# 2 Samples and methods

## 2.1 Samples

Thirteen live-collected specimens of deep-sea coral were examined. One specimen, PV 703-7 was dead at the base and so a living branch was analyzed instead. Specimens were collected by deep-sea submersible diving on Warwick Seamount, Gulf of Alaska (DSRV Alvin, 2002), the Hawaiian Islands (DSRV Pisces V, 2007) and Line Islands (DSRV Pisces IV, 2005) (Fig. 1). The thirteen corals belong to the Gorgonacea (gorgonian) and Scleractinia (scleractinian) orders. Gorgonian corals represent the Isididae and Coralliidae families and scleractinian corals are *Enallopsammia rostrata*. Isididae samples were identified as *Keratoisis, Isidella, or Acanella* spp. (collectively referred to as bamboo corals) and the Coralliidae sample as *Corallium* sp., most likely *Corallium secundum* (Table 1). Identification at the time of collection represents the current and best available taxonomic understanding.

Temperature data from Warwick Seamount is averaged from two Seabird CTD casts taken at the dive location during the dive, while in situ temperature was measured on DSRV Pisces V and IV at the time of collection for Hawaiian and Line Island corals. Temperature ranges from Hawaiian Island coral were calculated from Hawaiian Ocean Time Series (HOTS) (data collected at Station ALOHA during the years 1990–2012). Although Station ALOHA is located about 328 miles east of our samples, a comparison between the HOTS vertical temperature profile matches closely to CTD profiles

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taken from the sample locations. Additionally, Station ALOHA was originally chosen as a monitoring station because it is believed to be representative of the North Pacific subtropical gyre and we take it to represent conditions experienced by Hawaiian corals. Temperature ranges from Warwick Seamount are estimated from CTD casts taken at approximately 48.02 N, -130.66 W from 11 cruises during 1972–1998 at different times of the year. Temperature data was extracted from National Oceanic Data Center at 634±5 m, 704±5 m, 720±5 m and 872±5 m to assess temperature variability at the depths the corals were collected at, in the same oceanic region. Similarly temperature ranges from the Line Islands are estimated from CTD casts extracted from National Oceanic Data Center taken at approximately 6° N, -160° W from 9–27 cruises (depending on depth) during 1972–1998.

## 2.2 Sample preparation

Disks were cut from near the base for all coral skeletons, except PV 703-7 which was dead at the base. For this specimen, a disk was cut from the living branch for sampling. Disks were cleaned by the simple method of sonication in nanopure water and air-drying at room temperature (~25°C) All specimens lacked visual organic contamination of any kind and following the findings of Thiagarjan et al. (2011) and Eagle et al. (2013), which found cleaning steps to be unnecessary in deep-sea corals and mollusks, respectively, no further cleaning was performed. For most corals, sample powders of ca. 50–70 mg were milled using a Merchantek micromill from the outer portions of disks which represents the most recently accreted part of the skeleton (Fig. 2). In one sample, PV-703-5, a sample was milled from both the center and outer edge of the disk. In a few cases (PIV 148-2, PV 703-2 PIV 146-6, PV 694-3, PV 694-13) samples were first milled from the outer edge and later, to obtain additional sample, whole disks were ground into powder using a mortor and pestle. Although milling was preferred, in some cases, disks were small enough that the entire disk was necessary to yield sufficient sample for repeated $\Delta_{47}$ measurements.
In order to explore intra-coral heterogeneity and sampling effects, specimen PIV 146-6 is sampled both along the outer edge by micromilling as well as by grinding and homogenizing an entire disk, while PV 703-5 was micromilled in the center and outer edge (Fig. 2). In this way, intra-specimen reproducibility and the effects of sampling were assessed.

## 2.3 Notation and reference frame

The parameter $\Delta_{47}$ is a measure of the enrichment in per mil of $^{13}$C$^{18}$O$^{16}$O in CO$_2$ relative to the predicted stochastic abundance:

$$ \Delta_{47} = (R_{47}/R_{47}^* - 1) - (R_{46}/R_{46}^* - 1) - (R_{45}/R_{45}^* - 1) \times 1000 $$

where $R_{47}$, $R_{46}$, $R_{45}$ are the measured abundance ratios of masses 47/44, 46/44 and 45/44 in the sample and $R_{47}^*$, $R_{46}^*$, $R_{45}^*$ are the predicted abundance ratios of the same masses in the sample if it had a stochastic distribution of C and O isotopes among CO$_2$ isotopologues.

Standard gases of different bulk compositions were equilibrated at different temperatures (25 or 1000 °C) and measured each day. Values of $\Delta_{47}$ are reported in the absolute reference frame (ARF) and calculated as described in another publication using equilibrated gases (Dennis et al., 2011). Published data are presented in the absolute reference frame with conversion of data from Thiagarajan et al. (2011) and Ghosh et al. (2006) to the ARF, as reported in Eagle et al. (2013).

## 2.4 Analytical measurement of $\Delta_{47}$

Isotopic measurements were conducted on a dual inlet Thermo Scientific MAT 253 mass spectrometer coupled to a custom-built semi-automated sample digestion and purification system located at the University of California, Los Angeles. The set-up is modeled on that used at Caltech and John Hopkins, which are described in Passey et al. (2010) and Henkes et al. (2013). Samples were run during two separate time

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periods (winter: January–March 2013, summer: May–July 2013) and samples of scle-
ractinian (aragonite) and gorgonian (high-Mg calcite) were run interchangeably during
both time periods. For each analysis (previously referred to as extraction) 8–10 mg of
material was used to yield sufficient gas to maintain a steady signal at extended count-
ing times. Also, recent results suggest that digestion of samples that are significantly
smaller might undergo secondary equilibration with water, resulting in elevated Δ<sub>47</sub>
values compared to > 7 mg size samples (Wacker et al., 2013).

Mass spectrometric configuration was set to measure ion beams corresponding to
*m/z* = 44,45,46 (amplified by 3 × 10⁸ to 1 × 10¹¹Ω resistors) and 47, 48, 49 (ampli-
fied by 10¹²Ω resistors) with the ion beam *m/z* = 44 fixed at ~ 16V. Each analysis
contains 8 measurements of 7 cycles between sample and reference gas with 26 s
of integration per cycle. The total integration time of 1456 s is sufficiently long that er-
rors should be able to approach shot noise error predictions (Thiagarajan et al., 2011;
Huntington et al., 2009). Internal precision of Δ₄₇ was ca. 0.005–0.012‰, 1σ (based
on 8 measurements of 7 cycles of sample) and reference gas comparison within a sin-
gle analysis, (Table 2, Suppl. 1) and external precision was ca. 0.004–0.008‰, 1σ
(determined from repeat analyses of standards, Table 3) To increase external preci-
sion, samples were measured 3–7 times each, which resulted in standard errors for
Δ₄₇ of 0.002–0.010‰ for the corals in this study (Table 2).

The semi-automated sample digestion system allows for sample to be introduced as either a carbonate powder or a gas. Carbonate powders are introduced via a Costech autosampler which drops samples in vacuo into a common ~ 103% phosphoric acid (ρ ≈ 1.91 mg/ml) bath held at 90 °C and allowed to react for 20 min. Evolved CO₂ from acid digestion is then passed through a cooled ethanol trap (–78 °C) and collected in a liquid nitrogen trap (–200 °C). CO₂ is liberated by warming with the same ethanol trap (–78 °C), and with a helium carrier gas, passes through a Porapak Q 120/80 mesh GC (gas chromatograph) column at –20 °C and silver wool which removes organic contaminants and scavenges sulfur compounds, respectively. After GC passage, CO₂ is again collected in liquid nitrogen and undergoes one final cryogenic purification step

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in vacuo before introduction to the mass spectrometer. Gaseous CO₂ can also be prepared on a vacuum line and introduced using quartz tubes. A tube cracker which leads into the second cooled ethanol trap prior to the GC step allows CO₂ to pass through the autoline in the same way as carbonate samples, minus the common acid bath.

In addition to carbonate powders, on most days, at least one gaseous “equilibrated CO₂” sample is analyzed for use in defining the absolute reference frame for Δ₄₇ measurements. Producing equilibrated CO₂ gases with varying bulk isotopic (δ¹³C and δ¹⁸O) composition is accomplished by utilizing a very depleted CO₂ (δ¹³C = -25‰, δ¹⁸O = -30‰) and equilibrating Oztech CO₂ (δ¹³C = -3.6‰ PDB, δ¹⁸O = -15.83‰ PDB) with isotopically “heavy” water (¹⁸O enriched and produced by boiling house DI). These two isotopically end member gases are then either heated at 1000°C for 2 h to produce a nearly stochastic distribution of isotopes among isotopologues, or equilibrated in a water bath held at 25°C. Equilibrated CO₂ is then cryogenically purified on a vacuum line and captured into quartz tubes. This procedure produces four isotopically unique gaseous CO₂ samples which define the empirical transfer function (ETF) used to covert Δ₄₇ values to the ARF scale (Dennis et al., 2011). The carbonate standards Carrara Marble, Carmel Chalk, TV01, and 102-GC-AZ01 were run in concert with samples and were shown to have values that were indistinguishable from those determined in the Caltech lab (Table 2).

An acid fractionation factor of 0.081‰, was applied to all data and to theoretical predictions to normalize to previously reported data at 25°C (Passey et al., 2010) that is nearly identical to what was recently reported in another study (Defliese et al., 2015). We also show results calculated using different acid digestion fractionation factors (Henkes et al., 2013; Wacker et al., 2013; Defliese et al., 2015).
## 3.1 Scleractinian coral $\Delta_{47}$

When compared with previously reported deep-sea coral data the aragonitic scleractinian coral results from this study agree very closely to those measured in Thiagarajan et al. (2011) and Ghosh et al. (2006) (Figs. 3 and 4). Measured and predicted coral isotopic measurements are presented in Table 3 and 4, respectively. Since specimen PV 703-7 was not completely alive when collected and could have had a more complicated life history, it was excluded from calibration analysis. When the scleractinian corals are combined with the data from Thiagarajan et al. (2011), a similar $\Delta_{47}-T$ relationship to that originally reported is found. In a plot of $\Delta_{47}$ vs. $10^6/T^2$ the linear regression of the 11 corals from Thiagarajan et al. (2011) (slope: $0.0643 \pm 0.008$, intercept: $-0.029 \pm 0.103$, $R^2 = 0.934$) is almost identical to that of the combined data set of 14 scleractinian corals (slope: $0.0631 \pm 0.008$, intercept: $-0.013 \pm 0.098$, $R^2 = 0.85$). As noted in Thiagarajan et al. (2011) this is closely similar to that of the Ghosh et al. (2006) calibration of inorganic calcite (slope: $0.0620$; intercept: $-0.0021$).

The temperature sensitivity of the combined scleractinian coral calibration (slope: $0.0631 \pm 0.008$, intercept: $-0.013 \pm 0.098$, $R^2 = 0.85$) was examined. At high (17.4°C) and low (2.3°C) temperatures, 95% confidence intervals of $\Delta_{47}$ give $\pm 0.0165‰$ and $\pm 0.0105‰$, which correspond to temperature uncertainties of ca. $\pm 4–5°C$ and $\pm 3–4°C$, respectively. At the average temperature of the data set (9.6°C) uncertainty can be as good as $\pm 0.0075‰$ or $\pm 2°C$. Uncertainty in $\Delta_{47}$ is converted to uncertainty in temperature in accordance with Huntington et al. (2009).

## 3.2 Gorgonian coral $\Delta_{47}$

Gorgonian deep-sea corals precipitate skeletal carbonate in the form of high-Mg calcite, with 5–10 mol% $MgCO_3$ (Noe and Dullo, 2006; Kimball et al., 2014). Compared to scleractinian deep-sea coral and all other previously reported biogenic carbonates, $\Delta_{47}$

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is depleted in gorgonian deep-sea corals. In a plot of $\Delta_{47}$ vs. $10^6/T^2$, the linear regression of the 9 gorgonian deep-sea corals (slope: $0.018 \pm 0.012$, intercept: $0.511 \pm 0.157$, $R^2 = 0.23$) analyzed in this study gives a significantly shallower slope and $\Delta_{47}$ offset of ca. $0.04–0.07‰$ in the temperature range of $3.2–11.2°C$ from scleractinian corals as well as the Ghosh et al. (2006) calibration, however is more similar to the inorganic calibration of Dennis and Schrag (2010) (Fig. 3). Temperature sensitivity of the gorgonian coral calibration was not examined due to poorness of fit ($R^2 = 0.23$).

## 3.3 Intra-specimen sampling

Intraspecimen measurements from sampling different portions of the skeleton gives insight into the $\Delta_{47}$ heterogeneity present in an individual coral. Specimen PIV 146-6 was sampled both along the outer edge by micromilling as well as by grinding and homogenizing an entire disk, while PV 703-5 was micromilled in the center and outer edge. In both corals, aliquots produced from a given sampling method are significantly different in bulk isotopes. PIV 146-6 has differing bulk isotopic compositions between the two aliquots of more than 1‰ in $\delta^{13}$C and 0.4‰ in $\delta^{18}$O. PV 703-5 has almost a 3‰ difference in $\delta^{13}$C and 0.5‰ in $\delta^{18}$O (Table 5). $\Delta_{47}$ of the aliquots are approximately 0.03 and 0.02 different from each other in PIV 146-6 and PV 703-5, respectively. For PIV 146-6, the average of all 7 measurements is $0.804 \pm 0.008$ and for the two aliquots the averages are $0.808 \pm 0.008$ and $0.823 \pm 0.01$. For PV 703-5 the average of all 6 measurements is $0.750 \pm 0.007$ and the averages for the two aliquots are $0.740 \pm 0.005$ and $0.760 \pm 0.010$.
# 4 Discussion

## 4.1 Trends in deep-sea coral data

The scleractinian corals analyzed in this study are solely *Enallopsammia* sp. In contrast, the deep-sea corals presented in Ghosh et al. (2006) and Thiagarajan et al. (2011) are scleractinian *Desmophyllum*, *Enallopsammia*, and *Caryophyllia* sp. All scleractinian species precipitate a skeleton composed of aragonite although growth habits vary between species. *Desmophyllum* sp., also known as cup corals, grow in a rosette shape, whereas *Caryophyllia* and *Enallopsammia* sp. have a dendritic growth habit. The *E. rostrata* specimens measured in Thiagarajan et al. (2011) allow comparison between deep-sea corals of the same species measured in this study.

The scleractinian deep-sea coral data from this study have values that overlap with those reported by Thiagarajan et al. (2011). In contrast, the gorgonian corals show significantly lower $\Delta_{47}$ values compared to the scleractinian corals for the same temperature range. An analysis of covariance (ANCOVA) between the scleractinian and gorgonian coral data gives a P value < 0.01 which indicates that the differences between the two groups are statistically significant. This 0.04–0.07‰ offset between scleractinian and gorgonian corals is observed in non-acid digestion corrected $\Delta_{47}$ values. Possible explanations for this offset are discussed below.

## 4.2 Comparison with other published $\Delta_{47}$ calibration data sets

Since its inception, numerous theoretical, inorganic and biogenic calibrations have been put forward to calibrate the $\Delta_{47}-T$ relationship in different carbonate materials. To date, at least seven synthetic calibrations have been published (Ghosh et al., 2006; Dennis and Schrag, 2010; Zaruur et al., 2013; Tang et al., 2014; Defliese et al., 2015; Kluge et al., 2015; Tripati et al., 2015). Numerous biogenic carbonate materials have been studied to assess their agreement with inorganic and theoretical studies (Ghosh et al., 2006; Thiagarajan et al., 2011; Saenger et al., 2012; Eagle et al., 2013;

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Henkes et al., 2013). The scleractinian corals in this study agree well with the inorganic calibrations of Ghosh et al. (2006) and Zaruur et al. (2014) and Tripati et al., 2015, while the gorgonian corals show agreement with the Dennis and Schrag (2010), Tang et al. (2014), and Kluge et al. (2015) inorganic calibrations (e.g., Fig. 4).

## 4.3 Comparison with theoretical predictions

Theoretical modeling of $\Delta_{63}$ for different carbonate minerals (Schauble et al., 2006; Hill et al., 2014; Tripati et al., 2015) have been combined with theoretical acid digestion fractionations to give predicted $\Delta_{47}$ values for a variety of carbonate minerals. These predictions generally agree with measured inorganic and biogenic studies however are more depleted in $\Delta_{47}$ for a given temperature. Further, offsets due to mineralogy, while predicted in these theoretical modeling studies, have not been resolved in measured samples. Compared to deep-sea coral $\Delta_{47}$, theoretical predictions are more depleted for aragonite, however the calcite prediction agrees well with the calcitic gorgonians from this study (Fig. 5). The offset of 0.03‰ between aragonite and calcite is on the lower end of what we observe. A recent theoretical modeling effort (Hill et al., 2014) has produced similar predicted $\Delta_{47}$ values for aragonite and calcite although with more depleted (less agreement) between predicted and observed aragonite $\Delta_{47}$ and a smaller offset between aragonite and calcite.

## 4.4 Acid digestion fractionation factor uncertainties

It is possible that uncertainties in the acid digestion fractionation factor for carbonates digested at 25 and 90 °C may account for some of the above offsets. Three values have been proposed for the acid digestion fractionation factor for reporting data at 90 °C and converting to 25 °C on the absolute reference frame. To explore the sensitivity of results to choice of acid digestion fractionation factor, Figs. 6–9 show calculations using a range of acid digestion fractionation factors of isotope ratios, and apparent deviations of measured isotopic ratios from isotopic equilibrium. Passey et al. (2010) cal-

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culates and used a value of 0.081‰, whereas Henkes et al. (2013) suggests a value of 0.092‰ ±0.012‰, which is within error of the Passey et al. (2010) value and the Defliese et al. (2015) value of 0.091‰. In a recent study Wacker et al. (2013) suggests there may be a mineral-specific AFF, which is also predicted in the theoretical study by Guo et al. (2009). Application of the Wacker et al. $\Delta_{25-90}^*$ correction to the aragonite and calcite $\Delta_{47}$ data can explain 0.009‰ of the 0.03–0.07‰ offset observed in our aragonite and calcite data. Based on theoretical predications for dolomite (Guo et al., 2009), a further offset might be expected due to Mg incorporation in calcite. Still, experimental (Wacker et al., 2013; Defliese et al., 2015) and theoretical (Guo et al., 2009) studies suggest acid digestion differences due to mineralogy may not be large enough to explain the ca. 0.04–0.07‰ offset that we observe.

Deep-sea coral $\Delta_{47}$ were corrected using $\Delta_{25-90}$ AFF of 0.081‰ (Passey et al., 2010), 0.092‰ (Henkes et al., 2013 – near identical to the value of 0.091‰ from Defliese et al., 2015) and 0.066‰ (Wacker et al., 2013, aragonite) and 0.075‰ (Wacker et al., 2013, calcite). Figure 6 shows the result of the AFF sensitivity analysis and indicates that although the absolute values of the $\Delta_{47}$ change with application of AFF values, the offset between aragonite and calcite is unaffected except for in the case of application of mineral specific AFF from Wacker et al. (2013). Application of Wacker et al. (2013) mineral specific AFF values results in a decrease of the offset between the aragonite and calcite deep-sea coral $\Delta_{47}$ of only 0.009‰, therefore not altering our conclusions.

However, we acknowledge the kinetics of acid digestion and effects on clumped isotope signatures on $\Delta_{47}$ values is relatively poorly constrained. It is possible there are variable acid digestion fractionation factors that are influencing these observations, possibly due to cation substitution into carbonates (Guo et al., 2009), mineral bulk isotope composition (Guo et al., 2009), the presence of organic matter, as well as other factors. Until systematic studies of acid digestion is undertaken to investigate each of these factors, the contribution of acid digestion fractionation to the observed differ-
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ences between aragonitic scleractinian and high-Mg calcitic gorgonian corals cannot
be accurately quantified.

**4.5 Analytical procedure, data comparison**

Since its inception, the analytical measurement of Δ₄₇ has improved in precision and accuracy due to advances in mass spectrometric analysis (e.g. development of autoline and higher counting times) and standardization of reporting procedures (Huntington et al., 2009; Dennis et al., 2011). Because the methods involved in producing and reporting Δ₄₇ have changed over time, the comparison of data between laboratories and even within laboratories can be an issue. For relevant comparisons between previously reported data and this study's data, we refer to the recalculations of Δ₄₇ performed in Eagle et al., (2013) using standard values to project into the absolute reference frame. Still, we can not exclude the possibility that analytical artifacts remain. For instance, Zaruur et al. (2013) noted that the shorter counting times and fewer replications of the trailblazing work of Ghosh et al. (2006) may contribute to any discrepancies between data sets. We assess interlaboratory comparability from standards, which indicate accuracy and precision comparable to what has been reported for other laboratories (Table 2; Dennis et al., 2011). Furthermore, we note that good agreement between the aragonitic scleractinian corals of this study and that of Thiagarajan et al. (2011) reinforces the comparability of our data to the wider body of measurements reported from other laboratories.

Interlaboratory comparability should not be a critical issue for the interpretation of
data presented herein since the offset between aragonitic and calcitic corals was ob-
served in data generated within our lab, while running samples interchangeably using
the same analytical procedure. Standards were also run concurrently during these time
periods (Table 3) and found to be in good agreement with accepted values. We there-
fore conclude it is unlikely that the offset is due to internal analytical variability.

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## 4.6 Mixing and intraspecimen $\Delta_{47}$ heterogeneity

It has been shown that both scleractinian and gorgonian deep-sea corals exhibit significant disequilibrium in $\delta^{13}$C and $\delta^{18}$O of several per mil within an individual coral. Because of the subtle saddle-shaped curvature in $\Delta_{47}$ relative to $\delta^{13}$C and $\delta^{18}$O (Eiler and Schauble, 2004; Thiagarajan et al., 2011; Defliese and Lohmann, 2015), $\Delta_{47}$ is sensitive to mixing of end member $\delta^{13}$C and $\delta^{18}$O values and will result in non-conservative mixing (Fig. 10). Deep-sea corals, with their several per mil range in $\delta^{13}$C and $\delta^{18}$O values will be particularly susceptible to this effect. In fact, spatial variability in $\delta^{13}$C and $\delta^{18}$O values within deep-sea coral skeletons varies on the scale of 10s to 100s of microns (Smith et al., 2000; Adkins et al., 2003; Roark et al., 2006; Kimball et al., 2014) making the possibility of sampling an area with homogenous $\delta^{13}$C and $\delta^{18}$O values for $\Delta_{47}$ measurements unfeasible.

Both Thiagarajan et al. (2011) and Henkes et al. (2013) have addressed mixing in biogenic carbonates by calculating the $\Delta_{47}$ effect from mixing samples of differing bulk isotope composition. The end members from the Thiagarajan et al. (2011) calculations are similar to what we observe ($\delta^{13}$C = 2‰, $\delta^{18}$O = 5‰ and $\delta^{13}$C = -10‰, $\delta^{18}$O = -2‰). Using these end member values, Thiagarajan et al. (2011) calculated a positive $\Delta_{47}$ increase as large as 0.02‰ for the resulting mixture, if 50% of each end member is used. They note that while this mixing effect is not consistent with their data, mixing and then re-equilibration could potentially result in the trend seen in their data. Since sampling an isotopically homogenous area of a deep-sea coral is unfeasible (especially given the sample size requirements for clumped isotopes) we must assume that this mixing effect is indeed occurring. However, the extreme example of end members and the 50% contribution from each would be an unlikely occurrence. More likely is mixing between skeletal carbonate of less extreme differences in bulk isotopic composition. This is also supported by replicate analyses of the same sample aliquot. Standard errors for $\delta^{13}$C and $\delta^{18}$O in different preparations of the same coral measured range

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from 0.019 to 0.204‰ and 0.018 to 0.095‰, respectively (excluding those specimens which included multiple aliquots). This is consistent with relative small mixing effects.

As a further investigation into intraspecimen heterogeneity effects on $\Delta_{47}$ however, portions of two coral samples that differed in bulk isotopic composition were analyzed. In two corals, PIV 146-6 (scleractinian) and PV 703-5 (gorgonian), $\delta^{13}$C, $\delta^{18}$O and $\Delta_{47}$ were measured from intraspecimen aliquots of carbonate produced from sampling portions of the skeleton which differed significantly in their bulk isotopic composition (Table 5). Although $\Delta_{47}$ of the aliquots are approximately 0.02‰ different from each other in both corals, this is similar to that reported in other studies between a few replicates (Huntington et al., 2009; Thiagarajan et al., 2011). Additionally there is not a consistent trend observed between $\delta^{18}$O and $\Delta_{47}$, as for coral PIV 146-6, the aliquot with the more positive $\delta^{18}$O also has the higher $\Delta_{47}$, but the opposite is observed for coral PV 703-5. We therefore interpret the difference in $\Delta_{47}$ values between the aliquots as noise in the measurements, which further reinforces the need to run many replicates.

## 4.7 Diffusion

Both Thiagarajan et al. (2011) and Saenger et al. (2013) have suggested the diffusion of CO$_2$ across the calicoblastic membrane in corals as a possible mechanism for fractionation of $\Delta_{47}$. A generalized schematic of the coral calcifying region is shown in Fig. 11. Thiagarajan et al. (2011) calculated 0.7 and 1.6‰ decreases in $\delta^{18}$O and $\delta^{13}$C, respectively and a 0.036‰ increase in $\Delta_{47}$ for liquid phase diffusion of CO$_2$ (Fig. 10). For this $\Delta_{47}$ fractionation to be heritable by the mineral, however, the mineral must preserve the isotopic state of ordering of the DIC rather than obtain internal equilibrium within the mineral carbonate itself. Additionally carbonate mineral growth would have to occur before isotope clumping could reach equilibrium. Furthermore this fractionation only applies to DIC produced from diffused CO$_2$ and therefore the ratio of DIC from diffused DIC to “DIC derived from seawater ‘leaking’” directly to the calcifying
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environment would also have to be known. Both Thiagarajan et al. (2011) and Saenger et al. (2013) found fractionation associated with CO₂ diffusion across the calicoblastic layer or, additionally proposed by Saenger et al. (2013), across a boundary layer developed between the growing mineral and calcifying fluid, to be an unlikely explanation of the trends observed in their data.

It is possible the observed pattern could be explained if scleractinian deep-sea corals contain a larger fraction of diffused CO₂ compared to gorgonian coral, if this CO₂ then underwent relatively little isotopic equilibration in the calcifying fluid (Fig. 10). In cross plots of δ¹⁸O and Δ₄₇ residuals, scleractinian corals do exhibit elevated Δ₄₇ and depleted δ¹⁸O values, relative to Zaruur et al. (2013), Dennis and Schrag (2010) and theoretical aragonite (Schuable et al., 2006; Guo et al., 2009) in accordance with calculated diffusion effects. However, since scleractinian deep-sea corals generally conform to the majority of reported biogenic and synthetic calibrations, this would have be a widespread effect occurring in other materials reported as well.

## 4.8 Mineralogy and Δ₄₇

Despite Epstein et al.'s (1953) original finding that the nacreous layer (aragonite) and prismatic layer (calcite) of mollusks showed no oxygen isotopic offset, it has since been recognized that aragonite is enriched in ¹⁸O and ¹³C relative to calcite both in theory and experiment (Tarutani et al., 1969; Kim and O'Neil, 1997, 2007; Mavtomatis, 2012). An analogous mineralogical offset is predicted for Δ₄₇ (with aragonite predicted to have greater abundance of ¹³C-¹⁸O bonds than calcite that formed at the same temperature) (Schable et al., 2006; Guo et al., 2013; Hill et al., 2014; Tripati et al., 2015).

The corals in this study differ in mineralogy, with scleractinian corals precipitating a fully aragonitic skeleton and gorgonian corals secreting carbonate in the form of high-Mg calcite. To date, a combination of aragonitic and calcitic biogenic materials have been analyzed. Calcitic specimens have been represented by certain species of mollusks (Eagle et al., 2013; Henkes et al., 2013), brachiopods (Came et al., 2007;
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Henkes et al., 2013), and foraminifera (Tripati et al., 2010). Aragonitic specimens have included deep-sea (Thiagarajan et al., 2011) and shallow (Ghosh et al., 2006; Saenger et al., 2013) scleractinian corals, mollusks (Eagle et al., 2013; Henkes et al., 2013), brachiopods (Came et al., 2007) and land snails (Zaarur et al., 2011), all of which have not revealed statistically significant mineralogical offsets from the synthetic calibrations of calcite (Ghosh et al., 2006; Dennis et al., 2010; Zaruur et al., 2013). More careful study, however of the XRD data of Ghosh et al. (2006) has revealed aragonite was also present in their precipitates (Zaruur et al., 2013).

Despite not observing this offset between calcite and aragonite $\Delta_{47}$ in synthetic or biogenic carbonates, it has been predicted in theoretical studies. By combining a previous theoretical model of $^{13}$C-$^{18}$O clumping in carbonate minerals (Schauble et al., 2006) with their theoretical model of kinetic isotope effects associated with phosphoric acid digestion, Guo et al. (2009) proposed predicted $\Delta_{47}-T$ relationships for a number of carbonate minerals. For the temperature range 0–30 °C, $\Delta_{47,aragonite-calcite} \approx 0.02‰$ and $\Delta_{47,aragonite-dolomite} \approx 0.04‰$ with the direction of fractionation in agreement with the offset that we observe (Fig. 5). A similar pattern was reported by Hill et al. (2014) and Tripati et al. (2015). If these models of mineral-specific $\Delta_{47}-T$ relationships is correct, then this could potentially explain a substantial portion of the offset we observe between the aragonite and high-Mg calcite corals in this study. We discuss possible reasons for an observed mineralogical effects in deep-sea corals and not in other types of carbonates in more detail below.

## 4.9 Calcification in deep-sea corals

Non-equilibrium partitioning of oxygen and carbon isotopes have been observed in numerous calcifying organisms, and linked to calcification processes (McConnaughy, 1989; Spero et al., 1997; Zeebe, 1999; Adkins et al., 2003: Watson, 2004; DePaolo, 2011; Gabitov et al., 2012; Watkins et al., 2013). Clumped isotopes represent another tool that can be used to probe mechanisms for disequilibrium isotopic signals in biological carbonates, and to determine when equilibrium precipitation has occurred,

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through the coupling of theoretical calculations and inorganic precipitation experiments with culturing studies (Thiagarajan et al., 2011; Tripati et al., 2015). Figure 10 shows a framework for how multiple paired stable isotope measurements can be used to trace the origin of disequilibrium effects.

In corals, precipitation of carbonate is believed to occur in a space between the cali-
coblastic layer and the hard skeleton (Fig. 11). It is likely that calcifying fluid chemistry
(pH, salinity) as well as growth rate and DIC equilibrium can all be potential sources
of non-equilibrium signals preserved in the solid. The chemical composition and DIC
sources from which the carbonate mineral forms have been a main focus for under-
standing non-equilibrium δ¹³C and δ¹⁸O values seen in corals. Biomineralization in
shallow symbiotic scleractinian corals has been well-studied over the past few decades
(see Cohen and McConnaughy, 2003 for a review). Early studies of shallow symbiotic
scleractinian corals revealed time independent δ¹⁸O and δ¹³C offsets from theoreti-
cal carbonate-water fractionation curves. This was addressed through species-specific
calibration studies, which allowed offsets to be corrected for, revealing the underlying
environmental signals (Weber and Woodhead, 1972; Dunbar and Wellington, 1981). In
contrast, linear trends between δ¹⁸O and δ¹³C with depletions of several per mil, are
observed in an individual deep-sea coral. Living in near constant temperature, salinity
and pH environments, these large disequilibrium signals are attributed to their biomin-
eralization mechanism (i.e. “vital effects”). Different mechanistic models aimed at ex-
plaining non-equilibrium δ¹³C and δ¹⁸O values seen in deep-sea corals have been
proposed to account for these so-called “vital effects”.

The first of these models, proposed by McConnaughey (1989a) calls on a kinetic
fractionation during the hydration and hydroxylation steps of DIC speciation in which
precipitation outpaces the ability of the system to reach isotopic equilibrium (“Kinetic
model” in Fig. 11). The range of δ¹³C and δ¹⁸O values is therefore a reflection of
a spectrum of varying amounts of equilibrium being attained amongst DIC species.
The master variable in this case being time, McConnaughy (1989b), using calcification
rates, calculated that calcification could outpace the time it took for isotopic equilibrium
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to be reached. Henkes et al. (2013) further noted that at lower temperatures, days to weeks are required for oxygen isotopic equilibrium to be reached in DIC.

The second model proposed by Adkins et al. (2003) extensively examines potential sources of DIC to the extracellular calcifying fluid and introduces the concept of a variable pH environment as the master controller on skeletal $\delta^{13}$C and $\delta^{18}$O (“pH model” in Fig. 11). For oxygen isotope processes, they draw on work characterizing the oxygen isotopic fractionation between DIC species (i.e. CO$_2$, H$_2$CO$_3$, HCO$_3^-$, CO$_3^{2-}$) (Usdowski et al., 2009; Zeebe, 2007). Their model suggests that the mineral preserves the oxygen isotopic ratio of HCO$_3^-$/CO$_3^{2-}$. Because each DIC species has a unique oxygen fractionation relative to water, the oxygen isotope fractionation between the sum of DIC (DIC$_{sum}$) species and water will vary depending on the proportion of the DIC species. For seawater, at high pH (pH > 9), DIC will be dominated by CO$_3^{2-}$, which has the lowest fractionation factor and will therefore result in a more depleted $\delta^{18}$O value of the DIC$_{sum}$ from which the mineral precipitates. At lower (pH 6–9) fluid pH, as the HCO$_3^-$/CO$_3^{2-}$ ratio of the fluid rises, so does the precipitated mineral’s $\delta^{18}$O value. $\delta^{13}$C is not expected to be sensitive to pH effects since the exchange of carbon between DIC species is very rapid. $\delta^{13}$C, therefore, is proposed to be a function of two sources of CO$_2$, membrane diffused and seawater leaked, and the proportion of them, a function of pH gradients across the membrane. The “break” that they uncover at the most depleted isotopic values in which $\delta^{18}$O continues to decrease but $\delta^{13}$C plateaus, is interpreted as the point at which the maximum flux of diffused CO$_2$ across the membrane is reached (during rapid calcification). Qualitatively their model captures the isotopic trends that they see in their corals and although numerical modeling is able to reproduce the “break” and range of $\delta^{18}$O and $\delta^{13}$C that they report, it fails to reproduce the numerical value of the $\delta^{18}$O vs. $\delta^{13}$C slopes that they observe. $\delta^{18}$O vs. $\delta^{13}$C slopes have been observed in other studies and appear to be a conservative property of biomineralization in deep-sea corals (Smith et al., 2000; Adkins et al., 2003; Hill et al., 2010; Kimball et al., 2014).

Additional biomineralization models have been proposed in shallow symbiotic corals which incorporate the presence and activity of zooxanthellae (Cohen and Mc-
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Connaughy, 2003; Rollian-Bard et al., 2003) and deep-sea corals, based on trace ele-
ments (Gaetani et al., 2010). However only those models that examine effects on δ¹⁸O
and δ¹³C are considered with respect to the data presented herein. Additionally, as of
yet, a model that addresses the distinct high-Mg calcite mineralogy of gorgonians has
not been proposed.

More generally, models for equilibrium/disequilibrium stable isotope signatures have
been proposed that examine processes that influence the ion-by-ion growth of crys-
tals (Nielsen et al., 2012; Watkins et al., 2013; Tripati et al., 2015), as well as diffu-
sional/reordering processes that occur near the calcifying fluid/crystal interface (Wat-
son and Liang, 1995; Watson, 2004; Gabitov et al., 2012; Tripati et al., 2015).

4.10 Δ₄₇ in deep-sea coral skeletal carbonate: biomineralization model implications

The Δ₄₇ of a mineral is predicted to conform to the K<sub>eq</sub> of a homogenous isotope ex-
change reaction (Reaction R1) and be independent of the fluid composition from which
it forms, as well as the bulk isotopic composition of the crystal. Recent experimental and
theoretical studies have suggested that mineral equilibrium with respect to Δ₄₇ might
not always be attained however (Saenger et al., 2013; Tripati et al., 2015). Because of
this, Δ₄₇ can be utilized as an additional geochemical skeletal signal which could give
more insight into the mechanisms involved in skeletal precipitation in deep-sea corals
and other biological carbonates.

In a recent study (Tripati et al., 2015), the impact of DIC speciation on mineral Δ₄₇ signatures was examined with respect to Δ₄₇. Contrary to initial theoretical modeling by Hill et al. (2014) which calculated differences of ~ 0.03‰ in the Δ₄₇ values of DIC species at equilibrium, Tripati et al. (2015), using experimental results, found that HCO₃⁻Δ₆₃ could be as much as 0.05 to 0.07‰ higher than CO₃²⁻. They further propose that this fractionation could be partially or fully inheritable by the solid. Central to their mechanism of preservation of disequilibrium mineral Δ₆₃ signals is the concept of differing mineral isotopic archival effects occurring during “slow” and “fast” crystal growth,
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with “slow” and “fast” being defined as particular to a specific temperature, salinity and pH growth environment, rather than conforming to a particular value. Depending on the proportion of HCO₃⁻ and CO₃²⁻ incorporated into the growing crystal and the ability of the crystal lattice to conform to Kₐ₇ql (Reaction R1), disequilibrium in Δ₆₃ and therefore Δ₄₇ could be preserved. Only during “slow” crystal growth, they argue, is the mineral able to fully conform to Kₐ₇ql (Reaction R1). In this scenario, disequilibrium in Δ₆₃ and bulk isotopes are not correlated and it would therefore be possible to attain equilibrium with respect to Δ₆₃, but not in bulk isotopes.

Deep-sea corals, being one of the slowest calcifiers, are a likely potential candidate for attainment of Δ₆₃ mineral equilibrium, if the mechanism proposed by Tripati et al. (2015) is correct. Bulk isotopes, as previously mentioned, vary greatly within an individual deep-sea coral, while Δ₄₇ values are reproducible within instrumental error. This indicates that while bulk isotopes are controlled by processes such as in the biomineralization models outlined above, Δ₄₇ values are independent and possibly conform to equilibrium values. While a complete quantitative biophysical model has yet to be proposed, one possibility is that the range in bulk isotopes is controlled by processes related to the chemical environment (eg. calcifying solution pH, kinetics) while the slow growth rates in deep-sea corals allow isotopic reordering at the crystal-solution interface resulting in attainment of nominal mineral equilibrium for Δ₄₇ (Fig. 11).

Adkins et al. (2003), however, finds that in the more dense (interpreted as faster calcification) growth bands, δ¹⁸O is depleted compared to the less dense (more slowly growing) bands. In the Tripati et al. (2015) model, they predict that during rapid calcification both δ¹⁸O and Δ₄₇ would be more positive. Relative to the calcification rates required for inheritance of these signals however, perhaps even at the most elevated calcification rates, deep-sea corals, are still able to reach mineral isotopic equilibrium with respect to Δ₄₇, while bulk δ¹⁸O and δ¹³C is still under the control of processes associated with the calcifying fluid environment.

Cross plots of δ¹⁸O and Δ₄₇ residuals (Figs. 7, 8, and 9) can be used to assess conformation to proposed biomineralization models. For the McConnaughy et al. (2003)
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model of coral biomineralization, a range of depletions in $\delta^{18}$O is expected, with coral $\delta^{18}$O approaching or potentially reaching mineral equilibrium $\delta^{18}$O, depending on the calcification rate. Similarly in the Adkins et al. (2003) model, a range of pH would result in a range of depletions in recorded mineral $\delta^{18}$O, assuming that the proportion of $HCO_3^-$/$CO_3^{2-}$ incorporated into the skeleton is independent of calcification rate (Fig. 10). Although neither model addresses $\Delta_{63}$, in both cases, all else being equal, and assuming a mechanism for inheritance of DIC $\Delta_{63}$ into mineral $\Delta_{63}$, a decrease in $\delta^{18}$O would be tied to a decrease in $\Delta_{63}$. For the McConnaughy et al. (2003) model, an increase in calcification rate resulting in only partial equilibration of the DIC-water system would result in depleted $\delta^{18}$O and $\Delta_{63}$ values of $HCO_3^-$ and $CO_3^{2-}$. Similarly, in Adkins et al. (2003) model, varying solution pH would have the same effect on $\delta^{18}$O and $\Delta_{63}$ with both values decreasing (increasing) as pH increases (decreases). Comparing residual cross plots from various $\Delta_{47}$ calibrations however, fails to capture a positive relationship between $\delta^{18}$O and $\Delta_{47}$ and instead appears that $\delta^{18}$O and $\Delta_{47}$ are varying independently from each other. This suggests that either the models are not accurately portraying the mechanism by which variable $\delta^{18}$O values are inherited in the mineral or the $\Delta_{63}$ is not under the same controls as $\delta^{18}$O (i.e. not inheriting the $\Delta_{63}$ value of the fluid).

## 4.11 Other biomineralization effects

Unlike scleractinian zooxanthellate and azooxanthelate corals, gorgonian corals have been much less studied. In particular, biomineralization in gorgonian corals has yet to be addressed from a geochemical perspective. The fact that gorgonians precipitate high-Mg calcite however, in contrast to the aragonitic skeletons of scleractinians, is a simple indication of differing biomineralization mechanisms (Ries, 2010). Additionally, although it has been shown that deep-sea gorgonian corals exhibit $\delta^{18}$O and $\delta^{13}$C trends similar to their scleractinian counterparts, they have not been shown to exhibit the “break” that is observed in Adkins et al. (2003). When growing inorganic
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carbonate precipitates, mineralogy can be controlled through the Mg/Ca ratio of the precipitating fluid, with molar Mg/Ca > 2 producing aragonite and high-Mg calcite and molar Mg/Ca < 2 producing calcite. Current seas are ca. Mg/Ca = 5.2 (Ries, 2010) and the biomineralization mechanism responsible for precipitation of high-Mg calcite in gorgonians is unknown.

One possibility is that biomineralization is regulated in gorgonians through an organic template. The presence of an insoluble organic matrix in both isidid bamboo and *Coralium* sp. (Allemand et al., 1994; Ehrlick et al., 2006; Noe and Dullo, 2006) supports the possibility that an organic template may be directing mineral precipitation. The presence of an organic matrix is unique to gorgonian corals (has not been observed in scleractinians) and only preliminary studies have been undertaken to characterize the amino acid composition of the organic matrixes (Allemand et al., 1994; Ehrlick et al., 2006).

It seems plausible, that the organic template, in addition to directing lattice structure, could also be influencing isotopic fractionation within a growing mineral, as has recently been shown for Mg/Ca ratios in carbonates (Wang et al., 2009). Therefore, in addition to theoretical depletions expected in $\Delta_{63}$ due to mineralogy, the presence of an organic template might act as an additional control on bond ordering. Carbonate anhydrase has also been implicated in scleractinian hermatypic coral biomineralization models, and presence/absence and variable activities of carbonic anhydrase may another factor to be considered.

## 4.12 Suggestions for future work

In-depth experimental studies are needed to explore the role of multiple processes on acid digestion kinetics, such as Wacker et al. (2013) and Defliese et al. (2015). Possible factors influencing acid digestion fractionation factors that still need to be examined include (but are not limited to) organic matter content, the influence of cation substitution (e.g., high vs. low Mg-calcite), variations in bulk composition ($\delta^{13}$C, $\delta^{18}$O), the type of reaction vessel and clean-up method used, and phosphoric acid composition.

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Direct constraints on the impact of variable cation concentrations on the clumped isotopic composition of carbonate minerals from theory and synthetic experiments may be a useful tool to distinguish between whether these different deep-sea coral orders in fact are recording equilibrium mineral-specific offsets. Additionally, measurement of coral calcifying fluid pH and the use of labeled carbon sources during growth experiments would provide important constraints on whether the observed patterns reflect kinetic effects.

A unifying model which explains the observed patterns of isotopic and trace element variability that have been observed could also help to test the different explanations posed in this study and other publications, and could help us advance our understanding of biomineralization in scleractinian and gorgonian deep-sea corals. For clumped isotopes, such a model would need to include the factors that influence equilibrium fractionation, i.e., temperature and well-constrained mineral-specific fractionation factors. Additional factors that need to be included to accurately model kinetic isotope effects include temperature, the carbon source used for calcification (the relative fraction of seawater-derived DIC vs. CO₂ diffused across cellular membranes), DIC speciation as controlled by multiple factors (pH, temperature, salinity), the timescale for DIC equilibration vs. precipitation rate (affected by the residence time of DIC in the calcification space), as modulated by the occurrence and activity of the enzyme carbonic anhydrase.

## 5 Conclusions

Two groups of deep-sea corals with differing mineralogy (aragonite and calcite) were analyzed for δ¹³C, δ¹⁸O and Δ₄₇. Aragonitic scleractinian deep-sea coral exhibit a nearly coincident Δ₄₇–T relationship to previously reported scleractinian deep-sea corals (Thiagarjan et al., 2011) and can be used to predict temperatures to ±2 °C at the mean temperature and ±3–5 °C at the extreme ends of the data set. Calcitic gorgonian deep-sea corals show lower Δ₄₇ values. This offset between taxa is consistent with the-

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oretical predictions for mineralogic differences, albeit significantly larger in magnitude than expected. Therefore the offset between gorgonian and scleractinian deep-sea corals may reflect, at least in part, as a difference in mineralogy that is only observed in deep-sea corals and not in data published for other carbonates simply because of their very slow growth rates as compared to all other studied carbonates. Another equally likely explanation is that differences in carbon sources and biomineralization pathways in coral taxa result in distinctive clumped isotope vital effects in gorgonian and scleractinian deep-sea coral. Finally, we can not preclude that offsets may reflect variable acid digestion kinetics for the different types of materials that we have surveyed.

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Data availability

All raw data for samples and standards and equilibrated gas lines are in Supplement 1.

The Supplement related to this article is available online at
doi:10.5194/bgd-12-19115-2015-supplement.

Acknowledgements. This work was supported by a NSF Graduate Fellowship, NSF OCE-1437166, DOE BES grant DE-FG02-13ER16402, and the “Laboratoire d’Excellence” LabexMER (ANR-10-LABX-19) and co-funded by a grant from the French government under the program “Investissements d’Avenir”. We thank the Tripati Lab Group for their assistance with analyses and discussion of this work, and Ben Elliott for assistance with drafting figures.

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<table><caption>**Table 1.** Thirteen scleractinian and gorgonian deep-sea corals were collected live by submersible diving with depth and temperature measured at the time of collection. Temperature ranges are mean with SD and come from Hawaii Ocean Time Series (Karl and Lukas, 1996) and cruises extracted from the National Ocean Database (Boyer et al., 2013).</caption><thead><tr><th>ID</th><th>Order</th><th>Coral</th><th>Location</th><th>Latitude</th><th>Longitude</th><th>Depth (m)</th><th>Temp (°C)</th><th>Temperature range (°C)</th></tr></thead><tbody><tr><td>ALV 3806-1</td><td>Gorgonacea</td><td>Isididae</td><td>Warwick, AK</td><td>48°04' N</td><td>132°48' W</td><td>872</td><td>3.2</td><td>3.4 ± 0.1</td></tr><tr><td>ALV 3808-1</td><td>Gorgonacea</td><td>Isididae</td><td>Warwick, AK</td><td>48°04' N</td><td>132°48' W</td><td>758</td><td>3.5</td><td>3.5 ± 0.1</td></tr><tr><td>ALV 3808-3</td><td>Gorgonacea</td><td>Isididae</td><td>Warwick, AK</td><td>48°04' N</td><td>132°48' W</td><td>720</td><td>3.5</td><td>3.7 ± 0.1</td></tr><tr><td>ALV 3808-4</td><td>Gorgonacea</td><td>Isididae</td><td>Warwick, AK</td><td>48°04' N</td><td>132°48' W</td><td>704</td><td>3.6</td><td>3.7 ± 0.1</td></tr><tr><td>ALV 3808-5</td><td>Gorgonacea</td><td>Isididae</td><td>Warwick, AK</td><td>48°04' N</td><td>132°48' W</td><td>634</td><td>3.6</td><td>3.8 ± 0.2</td></tr><tr><td>PV 703-5</td><td>Gorgonacea</td><td>Coralliidae</td><td>Twin Banks, NWHI</td><td>23°07' N</td><td>163°08' W</td><td>942</td><td>4</td><td>4.2 ± 0.1</td></tr><tr><td>PV 592-1</td><td>Gorgonacea</td><td>Isididae</td><td>Big Island, HI</td><td>19°48' N</td><td>156°07' W</td><td>386</td><td>9.4</td><td>10.3 ± 0.2</td></tr><tr><td>PV 694-13</td><td>Gorgonacea</td><td>Isididae</td><td>East French Frigate Shoals, NWHI</td><td>23°54' N</td><td>165°23' W</td><td>356</td><td>11.2</td><td>10.6 ± 0.6</td></tr><tr><td>PV 694-3</td><td>Gorgonacea</td><td>Isididae</td><td>East French Frigate Shoals, NWHI</td><td>23°54' N</td><td>165°23' W</td><td>351</td><td>11.2</td><td>11.1 ± 0.7</td></tr><tr><td>PV 703-2</td><td>Scleractinia</td><td>E. rostrata</td><td>Twin Banks, NWHI</td><td>23°07' N</td><td>163°08' W</td><td>1108</td><td>3.7</td><td>3.7 ± 0.1</td></tr><tr><td>PIV 146-6</td><td>Scleractinia</td><td>E. rostrata</td><td>Kingman Reef</td><td>06°26.0' N</td><td>162°27.5' W</td><td>788</td><td>4</td><td>5.5 ± 0.2</td></tr><tr><td>PIV 148-2</td><td>Scleractinia</td><td>E. rostrata</td><td>Palmyra Atoll</td><td>05°50.784' N</td><td>162°06.741' W</td><td>588</td><td>6.5</td><td>6.9 ± 0.4</td></tr><tr><td>PV 703-7</td><td>Scleractinia</td><td>E. rostrata</td><td>Twin Banks, NWHI</td><td>23°07' N</td><td>163°08' W</td><td>534</td><td>6.7</td><td>6.7 ± 0.3</td></tr></tbody></table>

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<table><caption>**Table 2.** Coral specimens analyzed for δ<sup>13</sup>C, δ<sup>18</sup>O and Δ<sub>47</sub>. Δ<sub>47</sub> is reported relative to the absolute reference frame (ARF). Standard error, 1σ, is reported with number of replicate measurements (n). Individual analyses are reported in Supplement Table 1.</caption><thead><tr><th rowspan="2">ID</th><th rowspan="2">Order</th><th rowspan="2">mineral</th><th rowspan="2">n</th><th>δ<sup>13</sup>C</th><th>δ<sup>18</sup>O</th><th>Δ<sub>47</sub></th><th>Δ<sub>47</sub></th></tr><tr><th>(‰ VPDB)</th><th>(‰ VPDB)</th><th>(‰ ARF; adj. 25 °C)</th><th>(‰ ARF; 90 °C results)</th></tr></thead><tbody><tr><td>ALV 3806-1</td><td>Gorgonacea</td><td>high-Mg calcite</td><td>8</td><td>-3.74<br>(±0.091)</td><td>0.76<br>(±0.109)</td><td>0.744 (±0.004)</td><td>0.663 (±0.004)</td></tr><tr><td>ALV 3808-1</td><td>Gorgonacea</td><td>high-Mg calcite</td><td>3</td><td>-3.15<br>(±0.024)</td><td>0.53<br>(±0.064)</td><td>0.731 (±0.004)</td><td>0.650 (±0.004)</td></tr><tr><td>ALV 3808-3</td><td>Gorgonacea</td><td>high-Mg calcite</td><td>2</td><td>-4.58<br>(±0.04)</td><td>0.08<br>(±0.025)</td><td>0.720 (±0.002)</td><td>0.639 (±0.002)</td></tr><tr><td>ALV 3808-4</td><td>Gorgonacea</td><td>high-Mg calcite</td><td>6</td><td>-3.15<br>(±0.208)</td><td>0.66<br>(±0.095)</td><td>0.730 (±0.01)</td><td>0.649 (±0.01)</td></tr><tr><td>ALV 3808-5</td><td>Gorgonacea</td><td>high-Mg calcite</td><td>3</td><td>-5.87<br>(±0.131)</td><td>-0.21<br>(±0.043)</td><td>0.738 (±0.004)</td><td>0.657 (±0.004)</td></tr><tr><td>PV 703-5</td><td>Gorgonacea</td><td>high-Mg calcite</td><td>6</td><td>-6.35<br>(±0.543)</td><td>-1.09<br>(±0.345)</td><td>0.743 (±0.007)</td><td>0.662 (±0.007)</td></tr><tr><td>PV 592-1</td><td>Gorgonacea</td><td>high-Mg calcite</td><td>2</td><td>-2.42<br>(±0.068)</td><td>-0.18<br>(±0.029)</td><td>0.727 (±0.004)</td><td>0.646 (±0.004)</td></tr><tr><td>PV 694-13</td><td>Gorgonacea</td><td>high-Mg calcite</td><td>4</td><td>-0.71<br>(±0.019)</td><td>0.10<br>(±0.018)</td><td>0.722 (±0.004)</td><td>0.641 (±0.004)</td></tr><tr><td>PV 694-3</td><td>Gorgonacea</td><td>high-Mg calcite</td><td>5</td><td>-4.19<br>(±0.108)</td><td>-0.05<br>(±0.154)</td><td>0.707 (±0.008)</td><td>0.626 (±0.008)</td></tr><tr><td>PV 703-2</td><td>Scleractinia</td><td>aragonite</td><td>8</td><td>-1.91<br>(±0.19)</td><td>2.11<br>(±0.338)</td><td>0.79 (±0.008)</td><td>0.709 (±0.008)</td></tr><tr><td>PIV 146-6</td><td>Scleractinia</td><td>aragonite</td><td>7</td><td>-2.90<br>(±0.166)</td><td>1.40<br>(±0.083)</td><td>0.791 (±0.006)</td><td>0.710 (±0.006)</td></tr><tr><td>PIV 148-2</td><td>Scleractinia</td><td>aragonite</td><td>7</td><td>-3.90<br>(±0.351)</td><td>0.70<br>(±0.044)</td><td>0.796 (±0.004)</td><td>0.715 (±0.004)</td></tr><tr><td>PV 703-7</td><td>Scleractinia</td><td>aragonite</td><td>3</td><td>-3.10<br>(±0.065)</td><td>0.40<br>(±0.038)</td><td>0.746 (±0.002)</td><td>0.665 (±0.002)</td></tr></tbody></table>
<table><caption>**Table 3.** Average standard values with standard error reported for the two periods of time when samples were measured compared to accepted values. $\Delta_{47}$ values are in ‰ relative to the absolute reference frame.</caption><thead><tr><th>Standard</th><th>n</th><th>Δ<sub>47</sub> UCLA<br>(‰ ARF; adj. 25 °C)</th><th>Accepted Δ<sub>47</sub> (Caltech)<br>(‰ ARF; adj. 25 °C)</th></tr></thead><tbody><tr><td>Cararra marble<sup>a</sup></td><td>8</td><td>0.395 (±0.008)</td><td>0.395</td></tr><tr><td>Cararra marble<sup>b</sup></td><td>15</td><td>0.390 (±0.005)</td><td>0.395</td></tr><tr><td>Carmel Chalk<sup>a</sup></td><td>7</td><td>0.688 (±0.004)</td><td>0.697</td></tr><tr><td>Carmel Chalk<sup>b</sup></td><td>13</td><td>0.695 (±0.004)</td><td>0.697</td></tr><tr><td>102-GC-AZ01<sup>a</sup></td><td>5</td><td>0.729 (±0.004)</td><td>0.713</td></tr><tr><td>102-GC-AZ01<sup>b</sup></td><td>3</td><td>0.699 (±0.006)</td><td>0.713</td></tr><tr><td>TV01/TV03<sup>b</sup></td><td>7</td><td>0.720 (±0.007)</td><td>0.713</td></tr></tbody></table>

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<footnote><sup>a</sup> winter.</footnote>

<footnote><sup>b</sup> summer 2013.</footnote>
<table><caption>**Table 4.** Coral specimens analyzed for $\delta^{13}$C, $\delta^{18}$O and $\Delta_{47}$. Average $\Delta_{47}$ for all replicate measurements is reported relative to the absolute (ARF) reference frame. $\delta_{47}$ is reported relative to intra-laboratory working gas (WG).</caption><thead><tr><th rowspan="2">ID</th><th>δ<sup>18</sup>O<sub>coral</sub></th><th>δ<sup>18</sup>O<sub>water</sub></th><th>δ<sup>18</sup>O predicted</th><th rowspan="2">δ<sub>47</sub> (‰ WG)<sup>c</sup></th><th>Δ47</th><th>Δ<sub>47</sub> predicted</th></tr><tr><th>(‰ VPDB)</th><th>(‰ VPDB)<sup>a</sup></th><th>(‰ VPDB)<sup>b</sup></th><th>(‰ ARF; adj. 25 °C)</th><th>(‰ ARF; adj. 25 °C)<sup>d</sup></th></tr></thead><tbody><tr><td>ALV 3806-1</td><td>0.78 (±0.123)</td><td>-0.2</td><td>2.6</td><td>14.79</td><td>0.744 (±0.004)</td><td>0.833</td></tr><tr><td>ALV 3808-1</td><td>0.53 (±0.064)</td><td>-0.2</td><td>2.53</td><td>15.12</td><td>0.731 (±0.004)</td><td>0.831</td></tr><tr><td>ALV 3808-3</td><td>0.08 (±0.025)</td><td>-0.2</td><td>2.53</td><td>13.2</td><td>0.720 (±0.002)</td><td>0.831</td></tr><tr><td>ALV 3808-4</td><td>0.66 (±0.095)</td><td>-0.2</td><td>2.51</td><td>15.24</td><td>0.730 (±0.01)</td><td>0.831</td></tr><tr><td>ALV 3808-5</td><td>-0.21 (±0.043)</td><td>-0.2</td><td>2.51</td><td>11.66</td><td>0.738 (±0.004)</td><td>0.831</td></tr><tr><td>PV 703-5</td><td>-1.70 (±0.165)</td><td>-0.05</td><td>2.26</td><td>10.11</td><td>0.743 (±0.007)</td><td>0.828</td></tr><tr><td>PV 592-1</td><td>-0.18 (±0.029)</td><td>0.1</td><td>0.87</td><td>15.1</td><td>0.727 (±0.004)</td><td>0.797</td></tr><tr><td>PV 694-13</td><td>0.10 (±0.018)</td><td>0.1</td><td>0.45</td><td>16.97</td><td>0.722 (±0.004)</td><td>0.786</td></tr><tr><td>PV 694-3</td><td>-0.05 (±0.154)</td><td>0.1</td><td>0.45</td><td>17.03</td><td>0.707 (±0.008)</td><td>0.786</td></tr><tr><td>PV 703-2</td><td>2.00 (±0.318)</td><td>-0.1</td><td>3.13</td><td>18</td><td>0.790 (±0.008)</td><td>0.830</td></tr><tr><td>PIV 146-6</td><td>1.40 (±0.083)</td><td>-0.1</td><td>3.06</td><td>16.289</td><td>0.791 (±0.006)</td><td>0.828</td></tr><tr><td>PIV 148-2</td><td>0.70 (±0.044)</td><td>0.1</td><td>2.28</td><td>14.62</td><td>0.796 (±0.004)</td><td>0.813</td></tr><tr><td>PV 703-7</td><td>0.40 (±0.038)</td><td>0.1</td><td>1.6</td><td>14.94</td><td>0.746 (±0.002)</td><td>0.796</td></tr></tbody></table>

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<footnote>a $\delta^{18}$O<sub>water</sub> values are from WOCE and HOTS.</footnote>

<footnote>b $\delta^{18}$O predicted values are calculated using the equations reported in Kim and O'Neil (1997) for calcite and Kim et al. (2007) for aragonite.</footnote>

<footnote>c $\delta_{47}$ is relative to intralaboratory working gas with a known isotopic composition.</footnote>

<footnote>d $\Delta_{47}$ predicted value is calculated from the inorganic calibration of Ghosh et al. (2006); similar values are predicted from other inorganic calibrations including Zaarur et al. (2013) and Tripati et al. (2015).</footnote>
<table><caption>**Table 5.** Two coral specimens were analyzed using different portions of their skeletons. PIV 146-6 was sampled by micromilling the outer portion of a disk (outer) as well as grinding an entire disk (whole). PV 703-3 was micromilled in both cases, but was milled in both the center and outer portion of a disk. All data and averages are reported as well as standard error, 1σ.</caption><thead><tr><th>ID</th><th>n</th><th>δ<sup>13</sup>C (‰ VPDB)</th><th>Δ<sub>47</sub> predicted<br>δ<sup>18</sup>O (‰ VPDB)</th><th>(‰ ARF; adj. 25 °C)<sup>d</sup></th></tr></thead><tbody><tr><td>PIV 146-6 (outer)</td><td>4</td><td>-2.67 (±0.161)</td><td>1.33 (±0.02)</td><td>0.796 (±0.007)</td></tr><tr><td>PIV 146-6 (whole)</td><td>3</td><td>-3.37 (±0.086)</td><td>1.77 (±0.042)</td><td>0.820 (±0.010)</td></tr><tr><td>PIV 146-6 (average)</td><td>7</td><td>-2.90 (±0.166)</td><td>1.40 (±0.083)</td><td>0.808 (±0.008)</td></tr><tr><td>PV 703-5 (center)</td><td>3</td><td>-7.76 (±0.038)</td><td>-1.70 (±0.165)</td><td>0.760 (±0.010)</td></tr><tr><td>PV 703-5 (outer)</td><td>3</td><td>-4.95 (±0.116)</td><td>-0.47 (±0.025)</td><td>0.740 (±0.005)</td></tr><tr><td>PV 703-5 (average)</td><td>6</td><td>-6.35 (±0.543)</td><td>-1.09 (±0.345)</td><td>0.750 (±0.007)</td></tr></tbody></table>

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<figure><img src="image_195.png" /><figcaption>**Figure 1.** Sampling locations of three expeditions from which coral specimens were collected by deep-sea submersible diving. Locations approximate Warwick Seamount, Gulf of Alaska (DSRV Alvin, 2004), Hawaiian and Northwest Hawaiian Islands (DSRV Pisces V, 2005 and 2007) and Line Islands (DSRV Pisces IV, 2006).</figcaption></figure>

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<figure><img src="image_201.png" /><figcaption>**Figure 2. (a)** A bamboo coral from Warwick Seamount, ALV 3808-4. The bottommost internode is cut and the disk used for milling along the outer, most recently accreted portion. **(b)** A portion of a radial disk from PV 703-5 in which samples were milled from the center and outer portion.</figcaption></figure>

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<figure><img src="image_212.png" /><figcaption>**Figure 3.** Scleractinian and gorgonian deep-sea coral compared to deep-sea corals reported in Thiagarajan et al. (2011) and average Δ$_{47}$ values recalculated relative to ARF from Eagle et al. (2013). Sample 703-7 is not included in the calibration as it was not fully alive when collected.</figcaption></figure>
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<figure><img src="image_218.png" /><figcaption>**Figure 4.** Scleractinian corals from this study and Thiagarajan et al. (2011) are combined (slope: 0.643, intercept: 0.288) and along with gorgonian deep-sea corals are compared to reported calibrations of Eagle et al. (2013) and Ghosh et al. (2006).</figcaption></figure>
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<figure><img src="image_224.png" /><figcaption>**Figure 5.** Scleractinian and gorgonian corals from this study and Thiagarajan et al. (2011) are compared to the theoretical predictions for $\Delta_{63}$ for aragonite, calcite and dolomite. Acid fractionation values were added to convert to $\Delta_{47}$.</figcaption></figure>
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<figure><img src="image_230.png" /><figcaption>**Figure 6.** Sensitivity analysis of varying acid digestion fractionation factor ($\Delta_{25-90}$) on results.</figcaption></figure>
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<figure><img src="image_236.png" /><figcaption>**Figure 7.** $\Delta_{47}$ residuals are calculated from average measured $\Delta_{47}$ values for each coral relative to the inorganic calibrations of Ghosh et al. (2006), Zaruur et al. (2013) and Dennis and Schrag (2010). $\delta^{18}\text{O}$ residuals are calculated from average measured $\delta^{18}\text{O}$ values relative to the calibrations of Kim and O'Neil (1997) and Kim et al. (2007) for calcite and aragonite, respectively.</figcaption></figure>
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<figure><img src="image_242.png" /><figcaption>**Figure 8.** $\Delta_{47}$ residuals are calculated from average measured $\Delta_{47}$ values for each coral relative to the biogenic calibrations of Eagle et al. (2013). $\delta^{18}\text{O}$ residuals are calculated from average measured $\delta^{18}\text{O}$ values relative to the calibrations of Kim and O'Neil (1997) and Kim et al. (2007) for calcite and aragonite, respectively.</figcaption></figure>
<figure><img src="image_243.png" /><figcaption>**Figure 9.** $\Delta_{47}$ residuals are calculated from average measured $\Delta_{47}$ values for each coral relative to the theoretical calibration of Schauble et al. (2006) paired with theoretical acid digestion fractionations from Guo et al. (2009). $\delta^{18}\text{O}$ residuals are calculated from average measured $\delta^{18}\text{O}$ values relative to the calibrations of Kim and O'Neil (1997) and Kim et al. (2007) for calcite and aragonite, respectively.</figcaption></figure>

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<figure><img src="image_249.png" /><figcaption>**Figure 10.** Schematic illustrating general trajectories for different processes on stable isotope signatures ($\Delta_{63}$ or $\Delta_{47}$ values and $\delta^{18}$O), relative to an arbitrary point. Equilibrium temperature dependence is shown (Ghosh et al., 2006). Note that mixing of DIC from different sources results in curved trajectories (in $\delta^{18}$O space, mixing is linear whereas in $\Delta_{47}$ space can be non-linear) (Eiler and Schauble, 2004; Thiagarajan et al., 2011; Defliese and Lohmann, 2015). Diffusion results in enrichment of clumped species and depletion in $\delta^{18}$O (Thiagarajan et al., 2011). CO$_2$ hydration reactions are strongly temperature pH-dependent and also result in enrichment of clumped species and depletion in $\delta^{18}$O (Affek, 2013; Tang et al., 2014; Tripati et al., 2015). At any given temperature, equilibrium DIC species have differing clumped and oxygen isotope signatures, potentially also giving rise to pH effect (Hill et al., 2014; Tripati et al., 2015). Solid-state diffusion processes (Passey and Henkes, 2012; Henkes et al., 2014), including hypothesized reordering in the interfacial region (Tripati et al., 2015) only influence clumped isotope signatures and not bulk isotopic signatures (e.g., $\delta^{18}$O).</figcaption></figure>

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**Figure 11.** Schematic of the coral calcifying region (modified from McConnaughey, 1989; Adkins et al., 2003) illustrating processes that are hypothesized to result in stable isotope signatures observed in corals. Calcium pumping across the cell wall establishes a pH gradient between seawater and the extracellular calcifying fluid. Within the calcifying fluid, carbon can be sourced from CO₂ that is diffusively transported across the cell wall, and/or through seawater leakage. The hydration and/or hydroxylation of CO₂ are slow reactions, particularly at low temperature and at high pH, driving disequilibrium in δ¹⁸O and in clumped isotope signatures. The enzyme carbonic anhydrase will catalyze this reaction, if present and sufficiently active. At a given temperature, both HCO₃⁻ and CO₃²⁻ ions have distinct clumped isotope and δ¹⁸O signatures, and therefore changes in extracellular calcifying fluid pH may also affect the isotopic composition of minerals, as explored through various models including the pH, surface entrapment, and interfacial models mentioned in the text. Surface kinetic and interfacial processes and properties (e.g., surface speciation, attachment/detachment rates for isotopologues, the occurrence of defects, the nature of protein and saccharide macromolecules) also affect crystal growth and chemistry.

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