Biogeosciences, 13, 64876505, 2016 www.biogeosciences.net/13/6487/2016/ doi:10.5194/bg-13-6487-2016 Author(s) 2016. CC Attribution 3.0 License.
Justine Kimball1,2, Robert Eagle1,3, and Robert Dunbar2
1Department of Atmospheric and Oceanic Sciences, Institute of the Environment and Sustainability, 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), Universit de Brest, UMR 6538/6539, Rue Dumont DUrville, and IFREMER, Plouzan, France
Correspondence to: Robert Eagle ([email protected])
Received: 27 October 2015 Published in Biogeosciences Discuss.: 4 December 2015 Revised: 11 August 2016 Accepted: 22 August 2016 Published: 12 December 2016
Abstract. 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 calcication 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-specic, taxon-specic, and growth-rate-related effects. Carbonate clumped isotope thermometry is based on measurements of the abundance of the doubly substituted isotopologue 13C18O16O2 in carbonate minerals, analyzed in CO2 gas liberated on phosphoric acid digestion of carbonates and reported as [Delta1]47 values. We analyzed [Delta1]47 in live-collected aragonitic scleractinian (Enallopsammia sp.) and high-Mg calcitic gorgonian (Isididae and Coralliidae) deep-sea corals and compared results to published data for other aragonitic scleractinian taxa. Measured [Delta1]47 values were compared to in situ temperatures, and the relationship between [Delta1]47 and temperature was determined for each group to investigate taxon-specic effects. We nd that aragonitic scleractinian deep-sea corals exhibit higher values than high-Mg calcitic gorgonian corals and the two groups of coral produce statistically different relationships between [Delta1]47temperature calibrations. These data are signicant in the interpretation of all carbonate clumped isotope calibration data as they show that distinct [Delta1]47temperature calibrations can be observed in different materials recovered from
Carbonate clumped isotope signatures in aragonitic scleractinian and calcitic gorgonian deep-sea corals
the same environment and analyzed using the same instrumentation, phosphoric acid composition, digestion temperature and technique, CO2 gas purication 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 reect high-Mg and aragonite crystals attaining nominal mineral equilibrium clumped isotope signatures due to 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 inuencing 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 calcication, temperature- and pH-dependent rates of CO2 hydration and/or hydroxylation, calcifying uid pH, the activity of carbonic anhydrase, the residence time of dissolved inorganic carbon in the calcifying uid, and calcication rate. A third possible explanation is the occurrence of variable acid digestion fractionation factors. Although a recent study has suggested that dolomite,
Published by Copernicus Publications on behalf of the European Geosciences Union.
6488 J. Kimball et al.: Carbonate clumped isotope signatures
calcite, and aragonite may have similar clumped isotope acid digestion fractionation factors, the inuence of acid digestion kinetics on [Delta1]47 is a subject that warrants further investigation.
1 Introduction
Clumped-isotope paleothermometry is an approach to determining carbonate mineral formation temperatures (Ghosh et al., 2006; Schauble et al., 2006; Eiler, 2007); like the oxygen isotope thermometer before it, it is founded on thermodynamic predictions of the distribution of isotopes. Crucially, instead of relying on an isotopic exchange reaction between different phases (e.g., 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
X12C18O16O2 + X13C16O3 ! gX12C16O3 + X13C18O16O2, (R1)
where X refers to cations such as Ca2+, Mg2+, Sr2+, and Ba2+. In practice, the determination of 13C18O bonds in carbonate minerals is accomplished by measurement of mass47 CO2 (predominantly 13C18O16O) liberated by phosphoric acid digestion (Ghosh et al., 2006). 13C18O 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 [Delta1]47. 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 zooxanthel-late corals (Ghosh et al., 2006; Saenger et al., 2012; Tripati et al., 2015), aragonitic scleractinian non-zooxanthellate 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), 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; Daron et al., 2011), bioapatite (Eagle et al., 2010), and calcitic micro-bialites (Petryshyn et al., 2015).
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 [Delta1]47 values that was similar to the initial inorganic calcite calibration (Ghosh et al., 2006; Eagle et al., 2013). However, there are differences in inorganic cal-cite 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., 2013; Kluge et al., 2015) that are thought to either reect methodological differences in clumped isotope measurements or differences in how synthetic carbonates 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 relationships between temperature and [Delta1]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., 2013; Petrizzo et al., 2014; Deiese et al., 2015). It is also possible that phosphoric acid digestion technique and temperature may be playing a role (Wacker et al., 2013; Deiese et al., 2015), as is known to be the case for conventional 18O 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 are relatively little published data attempting to resolve whether different biogenic calibrations may originate from methodological differences between laboratories by simply measuring multiple types of biogenic carbonates interspersed with each other using the same methods.Therefore, it remains difcult 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 appear to exhibit a kinetic isotope effect that drives their [Delta1]47 values signicantly 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 13C18O bond abundances in deep-sea coral species from the orders Scleractinia and Gorgonacea and 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 role of mineralogy and biology in governing clumped isotope signatures. We discuss results in the context of recent studies including of acid digestion fractionation (Deiese et al., 2015) and theoretical models and experimental data that have been used to constrain differences in clumped isotope com-
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J. Kimball et al.: Carbonate clumped isotope signatures 6489
position 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 effecting DIC speciation such as pH, salinity, and temperature may also inuence carbonate clumped isotope signatures (Hill et al., 2014; Tripati et al., 2015).
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 fractionation (Adkins et al., 2003; Gaetani et al., 2011). 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 the potential to extend our observations of ocean physics and climate into the Pleistocene with decadal to centennial resolution. It has been established that deep-sea corals have signicant skeletal vital effects (disequilibrium stable isotope fractionations) that compromise the classic 18O paleotemperature method (Emiliani et al., 1978;McConnaughey, 1989a, 2003; Adkins et al., 2003; Smith et al., 2000; Rollion-Bard et al., 2003, 2010; Lutringer et al., 2005; Hill et al., 2011; Kimball et al., 2014). Attempts to circumvent these kinetic effects have focused on the lines method that recovers an average temperature over the lifetime of a coral (Smith et al., 2000; Lutringer et al., 2005; Hill et al., 2011; Kimball et al., 2014) and a Rayleigh-based multielement (e.g., Mg / Ca, Sr / Ca, Ba / Ca) paleotemperature method (Gaetani et al., 2011).
Despite deep-sea coral species showing signicant disequilibrium in 13C and 18O values, sometimes termed vital effects, initial calibration work using clumped isotopes in scleractinian corals (Thiagarajan et al., 2011) revealed a good agreement between aragonitic scleractinian deep-sea coral [Delta1]47 and the inorganic calcite calibration of Ghosh et al. (2006). Therefore, at least some species that show significant vital effects on 13C and 18O yield apparent equilibrium [Delta1]47 values, an observation also made on foraminifera and coccoliths (Tripati et al., 2010). Recently scleractinian deep-sea coral [Delta1]47 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 mostly dead with patches of living tissue. Specimens were collected by deep-sea submersible diving on Warwick Seamount, Gulf of
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).
Alaska (DSRV Alvin, 2002); the Hawaiian Islands (DSRV Pisces V, 2007); and Line Islands (DSRV Pisces IV, 2005) (Fig. 1). The 13 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 identied 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). Identication at the time of collection represents the current and best available taxonomic understanding.
Temperature data from Warwick Seamount is averaged from two Seabird conductivitytemperaturedepth (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 19902012. Although Station ALOHA is located about 328 miles east of our samples, a comparison between the HOTS vertical temperature prole matches closely with CTD proles 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 Pacic 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 during 11 cruises during 19721998 at different times of the year. Temperature data was extracted from National Oceanic Data Center at 634 [notdef] 5 m, 704 [notdef] 5 m,
720 [notdef] 5 m, and 872 [notdef] 5 m to assess temperature variability at
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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 standard deviation and come from Hawaii Ocean Time Series (Karl and Lukas, 1996) and cruises extracted from the National Ocean Database (Boyer et al., 2013). NWHI: Northwest Hawaiian Islands.
ID Order Coral Location Latitude Longitude depth T T range(m) ( C) ( C)
ALV 3806-1 Gorgonacea Isididae Warwick, AK 48 04[prime] N 132 48[prime] W 872 3.2 3.4 [notdef] 0.1
ALV 3808-1 Gorgonacea Isididae Warwick, AK 48 04[prime] N 132 48[prime] W 758 3.5 3.5 [notdef] 0.1
ALV 3808-3 Gorgonacea Isididae Warwick, AK 48 04[prime] N 132 48[prime] W 720 3.5 3.7 [notdef] 0.1
ALV 3808-4 Gorgonacea Isididae Warwick, AK 48 04[prime] N 132 48[prime] W 704 3.6 3.7 [notdef] 0.1
ALV 3808-5 Gorgonacea Isididae Warwick, AK 48 04[prime] N 132 48[prime] W 634 3.6 3.8 [notdef] 0.2
PV 703-5 Gorgonacea Coralliidae Twin Banks, NWHI 23 07[prime] N 163 08[prime] W 942 4 4.2 [notdef] 0.1
PV 592-1 Gorgonacea Isididae Big Island, HI 19 48[prime] N 156 07[prime] W 386 9.4 10.3 [notdef] 0.2
PV 694-13 Gorgonacea Isididae East French Frigate 23 54[prime] N 165 23[prime] W 356 11.2 10.6 [notdef] 0.6
Shoals, NWHIPV 694-3 Gorgonacea Isididae East French Frigate 23 54[prime] N 165 23[prime] W 351 11.2 11.1 [notdef] 0.7
Shoals, NWHIPV 703-2 Scleractinia E.rostrata Twin Banks, NWHI 23 07[prime] N 163 08[prime] W 1108 3.7 3.7 [notdef] 0.1
PIV 146-6 Scleractinia E.rostrata Kingman Reef 06 26.0[prime] N 162 27.5[prime] W 788 4 5.5 [notdef] 0.2
PIV 148-2 Scleractinia E.rostrata Palmyra Atoll 05 50.784[prime] N 162 06.741[prime] W 588 6.5 6.9 [notdef] 0.4
PV 703-7 Scleractinia E.rostrata Twin Banks, NWHI 23 07[prime] N 163 08[prime] W 534 6.7 6.7 [notdef] 0.3
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 during 9
to 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 contamina
tion of any kind, and following the ndings of Thiagarajan et al. (2011) and Eagle et al. (2013), who 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. 5070 mg were milled using a Merchantek micromill from the outer portions of disks which represent 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 rst milled from the outer edge and later, to obtain additional sample, whole disks were ground into powder using a mortar and pestle. Although milling was preferred, in some cases, disks were small enough that the entire disk was necessary to yield the required weight of sample for replica [Delta1]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 homogeniz-
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.
ing 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 [Delta1]47 is a measure of the enrichment in per mil of 13C18O16O in CO2 relative to the predicted stochastic abundance:
[Delta1]47 = (R47/R 471)(R46/R 461)(R45/R 451)[notdef]1000,
where R47, R46, and R45 are the measured abundance ratios of masses 47/44, 46/44, and 45/44 in the sample and R 47,
R 46, and R 45 are the predicted abundance ratios of the same
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J. Kimball et al.: Carbonate clumped isotope signatures 6491
masses in the sample if it had a stochastic distribution of C and O isotopes among CO2 isotopologues.
Standard gases of different bulk compositions were equilibrated at different temperatures (25 or 1000 C) and measured each day. Values of [Delta1]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 [Delta1]47
Isotopic measurements were conducted on a dual inlet Thermo Scientic MAT 253 mass spectrometer coupled to a custom-built semiautomated sample digestion and purication system located at the University of California, Los Angeles. The setup is modeled on that used at Caltech and Johns Hopkins, which are described in Passey et al. (2010) and Henkes et al. (2013). Samples were run during two separate time periods (winter: January 2013March 2013; summer: May 2013July 2013) and samples of scleractinian (aragonite) and gorgonian (high-Mg calcite) were run interchangeably during both time periods. For each analysis (previously referred to as extraction) 810 mg of material was used to yield sufcient gas to maintain a steady signal at extended counting times. Also, recent results suggest that the digestion of samples that are signicantly smaller might undergo secondary equilibration with water, resulting in elevated [Delta1]47 values compared to > 7 mg size samples (Wacker et al., 2013).
Mass spectrometric conguration was set to measure ion beams corresponding to m/z = 44, 45, 46 (amplied by
3 [notdef] 108 to 1 [notdef] 1011[Omega1] resistors) and 47, 48, 49 (amplied by
1012 [Omega1] resistors) with the ion bean m/z = 44 xed at 16 V.
Each analysis contains eight measurements of seven cycles between sample and reference gas with 26 s of integration per cycle. The total integration time of 1456 s is sufciently long that errors should be able to approach shot noise error predictions (Thiagarajan et al., 2011; Huntington et al., 2009). Internal precision of [Delta1]47 was ca. 0.0050.012 , 1
(based on eight measurements of seven cycles of sample and reference gas comparison within a single analysis, (Table 2, Supplement Table S1) and external precision was ca. 0.0040.008 , 1 (determined from repeat analyses of standards, Table 3). To increase external precision, samples were measured three to seven times each, which resulted in standard errors for [Delta1]47 of 0.0020.010 for the corals in this study (Table 2).
The semiautomated sample digestion system allows for a 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 ( 1.91 mg mL1) 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 contam
inants and scavenges sulfur compounds, respectively. After GC passage, CO2 is again collected in liquid nitrogen and undergoes one nal cryogenic purication step 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 dening the absolute reference frame for [Delta1]47 measurements.Producing equilibrated CO2 gases with varying bulk isotopic ( 13C and 18O) composition is accomplished by utilizing a very depleted CO2 ( 13C = 25 , 18O = 3.6 Vienna
Standard Mean Ocean Water (VSMOW)) and equilibrating Oztech CO2 ( 13C = 3.6 Vienna Pee Dee Belem
nite (VPDB), 18O = 23.6 VSMOW) with isotopically
heavy water (18O enriched and produced by boiling house DI). These two isotopical 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 puried on a vacuum line and captured into quartz tubes. This procedure produces four isotopically unique gaseous CO2 samples which dene the empirical transfer function (ETF) used to covert [Delta1]47 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 3).
An acid fractionation factor of 0.092 was applied to all data and to theoretical predictions to normalize to previously reported data at 25 C (Henkes et al., 2013). In tables and gures, we also report results calculated using a value of0.082 (Passey et al., 2010; Deiese et al., 2015) as well as other acid digestion fractionation factors (e.g., Wacker et al., 2013).
3 Results
3.1 Scleractinian coral [Delta1]47
When compared with previously reported deep-sea coral data the aragonitic scleractinian coral results from this study agree very closely with those measured in Thiagarajan et
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6492 J. Kimball et al.: Carbonate clumped isotope signatures
Table 2. Coral specimens analyzed for 13C, 18O, and [Delta1]47. [Delta1]47 is reported relative to the absolute reference frame (ARF). Standard error, 1, is reported with number of replicate measurements (n). Individual analyses are reported in Table S1. Values are calculated using a 25 to 90 C AFF = 0.092 (denoted by 1; Henkes et al., 2013) and AFF = 0.082 (denoted by 2; Deiese et al., 2015, and Passey et al., 2010).
ID Order Mineral n 13C 18O [Delta1]47
( VPDB) ( VPDB) ( ARF; 90 C results) ( ARF; adj. 25 C1) ( ARF; adj. 25 C2)
ALV 3806-1 Gorgonacea high-Mg calcite 8 3.74 ([notdef]0.091) 0.76 ([notdef]0.109) 0.663 ([notdef]0.004) 0.744 ([notdef]0.004) 0.734 ([notdef]0.004)
ALV 3808-1 Gorgonacea high-Mg calcite 3 3.15 ([notdef]0.024) 0.53 ([notdef]0.064) 0.650 ([notdef]0.004) 0.731 ([notdef]0.004) 0.721 ([notdef]0.004)
ALV 3808-3 Gorgonacea high-Mg calcite 2 4.58 ([notdef]0.04) 0.08 ([notdef]0.025) 0.639 ([notdef]0.002) 0.720 ([notdef]0.002) 0.710 ([notdef]0.002)
ALV 3808-4 Gorgonacea high-Mg calcite 6 3.15 ([notdef]0.208) 0.66 ([notdef]0.095) 0.649 ([notdef]0.01) 0.730 ([notdef]0.01) 0.720 ([notdef]0.01)
ALV 3808-5 Gorgonacea high-Mg calcite 3 5.87 ([notdef]0.131) 0.21 ([notdef]0.043) 0.657 ([notdef]0.004) 0.738 ([notdef]0.004) 0.728 ([notdef]0.004)
PV 703- 5 Gorgonacea high-Mg calcite 6 6.35([notdef]0.543) 1.09 ([notdef]0.345) 0.662 ([notdef]0.007) 0.743 ([notdef]0.007) 0.733 ([notdef]0.007)
PV 592-1 Gorgonacea high-Mg calcite 2 2.42 ([notdef]0.068) 0.18 ([notdef]0.029) 0.646 ([notdef]0.004) 0.727 ([notdef]0.004) 0.717 ([notdef]0.004)
PV 694-13 Gorgonacea high-Mg calcite 4 0.71 ([notdef]0.019) 0.10 ([notdef]0.018) 0.641 ([notdef]0.004) 0.722 ([notdef]0.004) 0.712 ([notdef]0.004)
PV 694-3 Gorgonacea high-Mg calcite 5 4.19 ([notdef]0.108) 0.05 ([notdef]0.154) 0.626 ([notdef]0.008) 0.707 ([notdef]0.008) 0.697 ([notdef]0.008)
PV 703-2 Scleractinia aragonite 8 1.91 ([notdef]0.19) 2.11 ([notdef]0.338) 0.709 ([notdef]0.008) 0.790 ([notdef]0.008) 0.780 ([notdef]0.008)
PIV 146-6 Scleractinia aragonite 7 2.90 ([notdef]0.166) 1.40 ([notdef]0.083) 0.710 ([notdef]0.006) 0.791 ([notdef]0.006) 0.781 ([notdef]0.006)
PIV 148-2 Scleractinia aragonite 7 3.90 ([notdef]0.351) 0.70 ([notdef]0.044) 0.715 ([notdef]0.004) 0.796 ([notdef]0.004) 0.786 ([notdef]0.004)
PV 703-7 Scleractinia aragonite 3 3.10 ([notdef]0.065) 0.40 ([notdef]0.038) 0.665 ([notdef]0.002) 0.746 ([notdef]0.002) 0.736 ([notdef]0.002)
Thiagarajan et al. ( 2011)
Scleractinian deep-sea coral (not live-collected)
al. (2011) and Ghosh et al. (2006) (Figs. 3 and 4). Measured and predicted coral isotopic measurements are presented in Table 4. Since specimen PV 703-7 was mostly a dead coral except in small sections when collected and therefore could have had a more complicated life history, it was excluded from calibration analysis.For the purpose of calibration, we focus on data from live-collected specimens to compare with in situ temperature measurements to eliminate any uncertainty in growth temperatures or post-formation dissolution when considering proxy systematics. When the scleractinian corals are combined with the data from Thiagarajan et al. (2011), a similar [Delta1]47-T relationship to that originally reported is found. In a plot of [Delta1]47 vs. 106/T 2 the linear regression of the 11 corals from Thiagarajan et al. (2011) (slope: 0.0643 [notdef] 0.008; intercept: 0.029 [notdef] 0.103;
R2 = 0.934) is almost identical to that of the combined data
set of 14 scleractinian corals (slope: 0.0582 [notdef] 0.008; inter
cept: 0.0452 [notdef] 0.103; R2 = 0.81). 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: 0.0582 [notdef] 0.008; intercept:
0.0452 [notdef] 0.103; R2 = 0.81) was examined. At high (17.4 C)
and low (2.3 C) temperatures, 95 % condence intervals of [Delta1]47 give [notdef]0.0165 and [notdef]0.0105 , which correspond to
temperature uncertainties of ca. [notdef]45 and [notdef]34 C, respec
tively. At the average temperature of the data set (9.6 C),
uncertainty can be as good as [notdef]0.0075 or [notdef]2 C. Uncer
tainty in [Delta1]47 is converted to uncertainty in temperature in accordance with Huntington et al. (2009).
3.2 Gorgonian coral [Delta1]47
Gorgonian deep-sea corals precipitate skeletal carbonate in the form of high-Mg calcite, with 510 mol % MgCO3 (No and Dullo, 2006; Kimball et al., 2014). Compared to scle-
12.0 12.5 13.0 13.5
Scleractinian deep-sea corals
Gorgonian deep-sea corals (live-collected)
0.85
This study
Scleractinian deep-sea corals (live-collected)
0.80
47 ()
0.75
0.70
16 4
106/T2 (K); T in oC
Figure 3. Scleractinian and gorgonian deep-sea coral compared to deep-sea corals reported in Thiagarajan et al. (2011) and average [Delta1]47 values recalculated relative to ARF from Eagle et al. (2013).
Sample 703-7 is not included in the calibration and is not considered live-collected.
ractinian deep-sea coral and all other previously reported biogenic carbonates, [Delta1]47 is depleted in gorgonian deep-sea corals. In a plot of [Delta1]47 vs. 106/T 2, the linear regression through data derived from nine gorgonian deep-sea corals analyzed in this study gives a signicantly shallower slope (slope: 0.025 [notdef] 0.01; intercept: 0.403 [notdef] 0.129, R2 = 0.48)
and [Delta1]47 offset of ca. 0.040.07 in the temperature range of 3.211.2 C compared to scleractinian corals as well as compared to the Ghosh et al. (2006) calibration. It is, however, more similar to the inorganic calibration of Dennis and Schrag (2010) (Fig. 3). Temperature sensitivity of the gor-
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J. Kimball et al.: Carbonate clumped isotope signatures 6493
Figure 4
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, grows 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 signicantly lower [Delta1]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 signi-cant. This 0.040.07 offset between scleractinian and gorgonian corals is observed in non-acid-digestion-corrected [Delta1]47 values. Possible explanations for this offset are discussed below.
4.2 Comparison with other published [Delta1]47 calibration data sets
Since its inception, numerous theoretical, inorganic, and biogenic calibrations have been put forward to calibrate the [Delta1]47T relationship in different carbonate materials. To date, at least seven synthetic calibrations have been published (Ghosh et al., 2006; Dennis and Schrag, 2010; Zaarur et al., 2013; Tang et al., 2014; Deiese 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; Henkes et al., 2013). The scleractinian corals in this study agree well with the inorganic calibrations of Ghosh et al. (2006), Zaarur et al. (2013), 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 [Delta1]63 for different carbonate minerals (Schauble et al., 2006; Hill et al., 2014; Tripati et al., 2015) has been combined with theoretical acid digestion fractionations to give predicted [Delta1]47 values for a variety of carbonate minerals. These predictions generally agree with measured inorganic and biogenic studies; however, they are more depleted in [Delta1]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 [Delta1]47, theoretical predictions are more
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0.85
Ghosh et al. (2006) inorganic calibration Zaruur et al. (2013) inorganic calibration
Eagle et al. (2013) biogenic compilation calibration
0.80
47 ( )
0.75
0.70
0.65 11.5 12.0 12.5 13.0 13.5
106/T2 (K)
Figure 4. Scleractinian corals from this study and Thiagarajan et al. (2011) are combined (slope: 0.0582 [notdef] 0.008; intercept:
0.0452 [notdef] 0.103; R
2 =0.81) and, along with gorgonian deep-sea corals, are compared to other reported calibration studies.
gonian coral calibration was not examined due to poorness of t (R2 = 0.48).
3.3 Intra-specimen sampling
Intra-specimen measurements from sampling different portions of the skeleton gives insight into the [Delta1]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 signicantly different in bulk isotopes. PIV 146-6 has differing bulk isotopic compositions between the two aliquots of more than 1 in 13C and 0.4 in 18O. PV 703-5 has almost a 3 difference in 13C and 0.5 in 18O (Table 5). [Delta1]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 seven measurements is 0.791 [notdef] 0.008
and for the two aliquots the averages are 0.779 [notdef] 0.007 and
0.803 [notdef] 0.01. For PV 703-5 the average of all six measure
ments is 0.744 [notdef] 0.007, and the averages for the two aliquots
are 0.734 [notdef] 0.005 and 0.754 [notdef] 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
6494 J. Kimball et al.: Carbonate clumped isotope signatures
Table 3. Average standard values with standard error reported for the two periods of time when samples were measured compared to accepted values. [Delta1]47 values are in relative to the absolute reference frame. Values are calculated using a 25 to 90 C AFF = 0.092 (denoted by 1;
Henkes et al., 2013) and AFF = 0.082 (denoted by 2; Deiese et al., 2015 and Passey et al., 2010).
[Delta1]47 UCLA [Delta1]47 UCLA Accepted [Delta1]47 (Caltech)
Standard n ( ARF; adj. 25 C1) ( ARF; adj. 25 C2) ( ARF; adj. 25 C)
Carrara marblea 8 0.395 ([notdef]0.008) 0.385 ([notdef]0.008) 0.395
Carrara marbleb 15 0.390 ([notdef]0.005) 0.380 ([notdef]0.005) 0.395
Carmel chalka 7 0.688 ([notdef]0.004) 0.678 ([notdef]0.004) 0.697
Carmel chalkb 13 0.695 ([notdef]0.004) 0.685 ([notdef]0.004) 0.697
102-GC-AZ01a 5 0.729 ([notdef]0.004) 0.719 ([notdef]0.004) 0.713
102-GC-AZ01b 3 0.699 ([notdef]0.006) 0.689 ([notdef]0.006) 0.713
TV01/TV03b 7 0.720 ([notdef]0.007) 0.710 ([notdef]0.007) 0.713
a Winter. b Summer 2013.
Table 4. Coral specimens analyzed for 13C, 18O, and [Delta1]47. Average [Delta1]47 for all replicate measurements is reported relative to the absolute reference frame (ARF). 47 is reported relative to intra-laboratory working gas (WG).
ID 18Ocoral 18Owater 18O predicted 47 ( WG)c [Delta1]47 [Delta1]47 predicted( VPDB) ( VPDB)a ( VPDB)b ( ARF; adj. 25 C) ( ARF; adj. 25 C)d
ALV 3806-1 0.78 ([notdef]0.123) 0.2 2.6 14.79 0.744 ([notdef]0.004) 0.833
ALV 3808-1 0.53 ([notdef]0.064) 0.2 2.53 15.12 0.731 ([notdef]0.004) 0.831
ALV 3808-3 0.08 ([notdef]0.025) 0.2 2.53 13.2 0.720 ([notdef]0.002) 0.831
ALV 3808-4 0.66 ([notdef]0.095) 0.2 2.51 15.24 0.730 ([notdef]0.01) 0.831
ALV 3808-5 0.21 ([notdef]0.043) 0.2 2.51 11.66 0.738 ([notdef]0.004) 0.831
PV 703- 5 1.70 ([notdef]0.165) 0.05 2.26 10.11 0.743 ([notdef]0.007) 0.828
PV 592-1 0.18 ([notdef]0.029) 0.1 0.87 15.1 0.727 ([notdef]0.004) 0.797
PV 694-13 0.10 ([notdef]0.018) 0.1 0.45 16.97 0.722 ([notdef]0.004) 0.786
PV 694-3 0.05 ([notdef]0.154) 0.1 0.45 17.03 0.707 ([notdef]0.008) 0.786
PV 703-2 2.00 ([notdef]0.318) 0.1 3.13 18 0.790 ([notdef]0.008) 0.830
PIV 146-6 1.40 ([notdef]0.083) 0.1 3.06 16.289 0.791 ([notdef]0.006) 0.828
PIV 148-2 0.70 ([notdef]0.044) 0.1 2.28 14.62 0.796 ([notdef]0.004) 0.813
PV 703-7 0.40 ([notdef]0.038) 0.1 1.6 14.94 0.746 ([notdef]0.002) 0.796
a 18Owater values are from WOCE and HOTS. b 18O predicted values are calculated using the equations reported in Kim and ONeil (1997) for calcite and Kim et al. (2007) for aragonite. c 47 is relative to intra-laboratory working gas with a known isotopic composition. d [Delta1]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).
of results to the choice of acid digestion fractionation factor, Figs. 69 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) calculate and use a value of0.081 , whereas Henkes et al. (2013) suggest a value of0.092 [notdef] 0.012 , which is within error of the Passey et
al. (2010) value. Deiese et al. (2015) propose a value of0.082 [notdef] 0.014 , which is within error of both values,.
In a recent study Wacker et al. (2013) suggest there may be mineral-specic AFF, which is also predicted in the theoretical study by Guo et al. (2009). Application of the Wacker et al. (2013) [Delta1]2590 correction to the aragonite and calcite
[Delta1]47 data can explain 0.009 of the 0.030.07 offset ob-served in our aragonite and calcite data. Based on theoretical predictions for dolomite (Guo et al., 2009), a further offset might be expected due to Mg incorporation in calcite. Still,
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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 at the lower end of what we observe. A recent theoretical modeling effort (Hill et al., 2014) has produced similar predicted [Delta1]47 values for aragonite and calcite although with more depleted (less agreement) values between predicted and observed aragonite [Delta1]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 (AFF) for reporting data at 90 C and converting them to 25 C on the absolute reference frame. To explore the sensitivity
J. Kimball et al.: Carbonate clumped isotope signatures 6495
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, 1.
ID n 13C 18O [Delta1]47 predicted ( VPDB) ( VPDB) ( ARF; adj. 25 C)d
PIV 146-6 (outer) 4 2.67 ([notdef]0.161) 1.33 ([notdef]0.02) 0.796 ([notdef]0.007)
PIV 146-6 (whole) 3 3.37 ([notdef]0.086) 1.77 ([notdef]0.042) 0.820 ([notdef]0.010)
PIV 146-6 (average) 7 2.90 ([notdef]0.166) 1.40 ([notdef]0.083) 0.808 ([notdef]0.008)
PV 703-5 (center) 3 7.76 ([notdef]0.038) 1.70 ([notdef]0.165) 0.760 ([notdef]0.010)
PV 703-5 (outer) 3 4.95 ([notdef]0.116) 0.47 ([notdef]0.025) 0.740 ([notdef]0.005)
PV 703-5 (average) 6 6.35([notdef]0.543) 1.09 ([notdef]0.345) 0.750 ([notdef]0.007)
0.85
Theoretical aragonite (Schauble et al., 2006) + 0.280 Theoretical calcite (Schauble et al., 2006) + 0.278 Theoretical dolomite (Schauble et al., 2006) + 0.263
Wacker et al., 2013 (scleractinian) AFF 0.061
Henkes et al., 2013 (scleractinian) AFF 0.092
Passey et al., 2010 (scleractinian) AFF 0.081 (similar to Defliese et al., 2015)
Henkes et al., 2013 (gorgonian) AFF 0.092
Passey et al. , 2010 (gorgonian) AFF 0.081 (similar to Defliese et al., 2015)
0.85
Gorgonian deep-sea corals Scleractinian deep-sea corals
Wacker et al., 2013 (gorgonian) AFF 0.075
0.80
0.80
47 ( )
47 ( )
0.75
0.75
0.70
0.65 11.5 12.0 12.5 13.0 13.5
106/T2 (K)
Figure 6. Sensitivity analysis of varying acid digestion fractionation factor ([Delta1]2590) on results.
in the offset between the aragonite and calcite deep-sea coral [Delta1]47 of only 0.009 , therefore not altering our conclusions.
However, we acknowledge that the kinetics of acid digestion and the effects on clumped isotope signatures on [Delta1]47 values is relatively poorly constrained. It is possible there are variable acid digestion fractionation factors that are inuencing 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 are undertaken to investigate each of these factors, the contribution of acid digestion fractionation to the ob-served differences between aragonitic scleractinian and high-Mg calcitic gorgonian corals cannot be accurately quantied.
4.5 Analytical procedure, data comparison
Since its inception, the analytical measurement of [Delta1]47 has improved in precision and accuracy due to advances in mass
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106/T2 (K)
Figure 5. Scleractinian and gorgonian corals from this study and Thiagarajan et al. (2011) are compared to the theoretical predictions for [Delta1]63 for aragonite, calcite, and dolomite. Acid fractionation values were added to convert to [Delta1]47.
experimental (Wacker et al., 2013; Deiese et al., 2015) and theoretical (Guo et al., 2009) studies suggest that acid digestion differences due to mineralogy may not be large enough to explain the ca. 0.040.07 offset that we observe.
Deep-sea coral [Delta1]47 was corrected using [Delta1]2590 AFF
of 0.082 (Deiese et al., 2015) identical to the value from Passey et al., 2010 0.092 (Henkes et al., 2013) 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 [Delta1]47 change with the application of AFF values, the offset between aragonite and calcite is unaffected except in the case of the application of mineral-specic AFF from Wacker et al. (2013). The application of Wacker et al. (2013) mineral-specic AFF values results in a decrease
0.70 11.5 12.0 12.5 13.0 13.5
6496 J. Kimball et al.: Carbonate clumped isotope signatures
47 Ghosh residual (scleractinian corals)
offset between aragonitic and calcitic corals was observed 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 therefore conclude it is unlikely that the offset is due to internal analytical variability.
4.6 Mixing and intra-specimen [Delta1]47 heterogeneity
It has been shown that both scleractinian and gorgonian deep-sea corals exhibit signicant disequilibrium in 13C and 18O
of several per mil within an individual coral (eg. Smith et al., 2000; Adkins et al., 2003; Rollion-Bard et al., 2003, 2010;Lutringer et al., 2005; Kimball et al., 2014). Because of the subtle saddle-shaped curvature in [Delta1]47 relative to 13C and 18O (Eiler and Schauble, 2004; Thiagarajan et al., 2011;Deiese and Lohmann, 2015), [Delta1]47 is sensitive to the mixing of end member 13C and 18O values and will result in non-conservative mixing (Fig. 10). Deep-sea corals, with their several per mil range in 13C and 18O values will be particularly susceptible to this effect. In fact, both banding and spatial variability in 13C and 18O values within deep-sea coral skeletons vary on the scale of ten to hundreds of microns (Smith et al., 2000; Adkins et al., 2003; Rollion-Bard et al., 2003, 2010; Lutringer et al., 2005; Kimball et al., 2014). Any small-scale spatial variability makes the possibility of sampling within a skeletal band or an area with homogenous 13C and 18O values for [Delta1]47 measurements unfeasible, and therefore there is the potential for mixing effects in [Delta1]47 which we address here.
Both Thiagarajan et al. (2011) and Henkes et al. (2013) have addressed mixing in biogenic carbonates by calculating the [Delta1]47 effect from mixing samples of differing bulk isotope composition. The end members from the Thiagarajan et al. (2011) calculations are similar to the most extreme values that we observe ( 13C = 2 , 18O = 5 and
13C = 10 , 18O = 2 ). Using these end member
values, Thiagarajan et al. (2011) calculated a positive [Delta1]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 estimate the maximum artifacts in our data arising from mixing. The estimate for mixing artifacts from Thiagarajan et al. (2011) is based on observed variability in stable isotope values in scleractinian deep-sea coral, and represents an extreme example for this study, as they use 13C and 18O values of end members that differ by 12 and 7 , respectively. Such differences in scleractinian coral represent systematic isotopic depletions in centers of calcication compared to bers (Adkins et al., 2003), as well as variability associated with density bands;
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47 Ghosh residual (gorgonian corals)
47 Zaruur residual (scleractinian corals)
47 Dennis residual (scleractinian corals)
47 Dennis residual (scleractinian corals)
47 Ghosh residual (scleractinian corals)
0.10
47 measured- 47 predicted ()
0.05
0.00
-0.05
-0.10 -6 -4 -2 0 2 4
18Omeasured - 18O expected ()
Figure 7. [Delta1]47 residuals are calculated from average measured [Delta1]47 values for each coral relative to the inorganic calibrations of Ghosh et al. (2006), Zaarur et al. (2013), and Dennis and Schrag (2010). 18O residuals are calculated from average measured 18O values relative to the calibrations of Kim and ONeil (1997) and Kim et al. (2007) for calcite and aragonite, respectively.
spectrometric analysis (e.g., development of automated sample digestions and purication systems, abbreviated to auto-line, 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 [Delta1]47 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 studys data, we refer to the recalculations of [Delta1]47 performed in Eagle et al. (2013) using standard values to project onto the absolute reference frame. Still, we cannot exclude the possibility that analytical artifacts remain. For instance, Zaarur 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
J. Kimball et al.: Carbonate clumped isotope signatures 6497
no centers of calcication have been reported in gorgonian corals. The calculations from Thiagarajan et al. (2011) represent an estimate of the maximum artifact to arise from mixing for either taxon given the range of carbon and oxygen isotope values observed in scleractinian and gorgonian deep-sea coral. We also observed little variability in 13C and 18O
of different preparations of the same coral, with standard deviations of < 1 .
As a further investigation into intra-specimen heterogeneity, effects on [Delta1]47 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) 13C, 18O, and [Delta1]47 were measured from intra-specimen aliquots of carbonate produced from sampling portions of the skeleton which differed signicantly in their bulk isotopic composition (Table 5). Although [Delta1]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, no consistent trend is observed between 18O and [Delta1]47; as for coral PIV 146-6, the aliquot with the more positive 18O also has the higher [Delta1]47, but the opposite is ob-served for coral PV 703-5. We therefore interpret the difference in [Delta1]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. (2012) have suggested the diffusion of CO2 across the calicoblastic membrane in corals as a possible mechanism for the fractionation of [Delta1]47. A generalized schematic of a scleractinian coral-calcifying region is shown in Fig. 11. Thiagarajan et al. (2011) calculated 0.7 and 1.6 decreases in 18O and 13C, respectively, and a 0.036 increase in [Delta1]47 for liquid phase diffusion of CO2 (Fig. 10). For this [Delta1]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 CO2 and therefore the ratio of DIC from diffused DIC to DIC derived from seawater leaking directly to the calcifying environment would also have to be known.Both Thiagarajan et al. (2011) and Saenger et al. (2012) found fractionation associated with CO2 diffusion across the calicoblastic layer or, additionally proposed by Saenger et al. (2012), across a boundary layer developed between the growing mineral and calcifying uid to be an unlikely explanation of the trends observed in their data. Additionally, gorgonian corals are not known to have calicoblastic cells (No and Dullo, 2006), so this is relevant to scleractinian corals or
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0.10
47 Eagle biogenic compilation residual (scleractinian corals)
47 Eagle biogenic compilation residual (gorgonian corals)
47 Eagle mollusk residual (scleractinian corals)
47 Eagle mollusk residual (gorgonian corals)
47 measured- 47 predicted ()
0.05
0.00
-0.05
-0.10 -6 -4 -2 0 2 4
18Omeasured - 18O expected ()
Figure 8. [Delta1]47 residuals are calculated from average measured [Delta1]47 values for each coral relative to the biogenic calibrations of Eagle et al. (2013). 18O residuals are calculated from average measured 18O values relative to the calibrations of Kim and ONeil (1997) and Kim et al. (2007) for calcite and aragonite, respectively.
47 theoretical aragonite residual (scleractinian corals)
47 theoretical calcite residual (gorgonian corals)
47 theoretical dolomite residual (gorgonian corals)
0.10
47 measured- 47 predicted ()
0.05
0.00
-0.05
-0.10 -6 -4 -2 0 2 4
18Omeasured - 18O expected ()
Figure 9. [Delta1]47 residuals are calculated from average measured [Delta1]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). 18O residuals are calculated from average measured 18O values relative to the calibrations of Kim and ONeil (1997) and Kim et al. (2007) for calcite and aragonite, respectively.
6498 J. Kimball et al.: Carbonate clumped isotope signatures
O depletion
or enrichment
Seawater
Interfacial region
-T
Ca High [CO ]
Cell wall
Extracellular calcifying
fluid
Biogenic CaCO with organic macromolecules
mixing
O enrichment
pH/entrapment/interfacial models
CO Ca
Ca CO
CO Ca
+T
or depletion
Figure 10. Schematic illustrating general trajectories for different processes on stable isotope signatures ([Delta1]63 or [Delta1]47 values and 18O), 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 18O space, mixing is linear, whereas in [Delta1]47 space can be nonlinear) (Eiler and
Schauble, 2004; Thiagarajan et al., 2011; Deiese and Lohmann, 2015). Diffusion results in enrichment of clumped species and depletion in 18O (Thiagarajan et al., 2011). CO2 hydration reactions are strongly temperature pH-dependent and also result in enrichment of clumped species and depletion in 18O (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 inuence clumped isotope signatures and not bulk isotopic signatures (e.g., 18O).
if there is a similar functional cell playing a role in this for gorgonian corals.
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 uid (Fig. 10). In cross plots of 18O and [Delta1]47 residuals, scleractinian corals do exhibit elevated [Delta1]47 and depleted 18O values, relative to Zaarur et al. (2013), Dennis and Schrag (2010), and theoretical aragonite (Schauble 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 to be a widespread effect occurring in other materials reported as well.
4.8 Mineralogy and [Delta1]47
Despite Epstein et al. (1953) originally nding that the nacreous layer (aragonite) and prismatic layer (calcite) of mol-
Figure 11. Schematic of the scleractinian coral-calcifying region (modied from McConnaughey, 1989a; 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 uid. Within the calcifying uid, 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 18O and in clumped isotope signatures. The enzyme carbonic anhydrase will catalyze this reaction if present and sufciently active. At a given temperature, both HCO
3
and CO2
3 ions have distinct clumped isotope and 18O signatures, and therefore changes in extracellular calcifying uid 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 or detachment rates for isotopologues, the occurrence of defects, the nature of protein and saccharide macromolecules) also affect crystal growth and chemistry.
lusks showed no oxygen isotopic offset, it has since been recognized that aragonite is enriched in 18O and 13C rela
tive to calcite both in theory and experiment (Tarutani et al., 1969; Kim and ONeil, 1997; Kim et al., 2007; Mavromatis, 2013). An analogous mineralogical offset is predicted for [Delta1]47 (with aragonite predicted to have a greater abundance of 13C18O bonds than calcite that formed at the same temperature) (Schauble et al., 2006; Guo et al., 2009; 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;Henkes et al., 2013), and foraminifera (Tripati et al., 2010).
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Aragonitic specimens have included deep-sea (Thiagarajan et al., 2011) and shallow (Ghosh et al., 2006; Saenger et al., 2012) scleractinian corals, mollusks (Eagle et al., 2013;Henkes et al., 2013), and land snails (Zaarur et al., 2011), all of which have not revealed statistically signicant mineralogical offsets from the synthetic calibrations of calcite (Ghosh et al., 2006; Dennis et al., 2010; Zaarur et al., 2013).More careful study, however, of the X-ray diffraction (XRD) data of Ghosh et al. (2006) has revealed that aragonite was also present in their precipitates (Zaarur et al., 2013).
Despite not observing this offset between calcite and aragonite [Delta1]47 in synthetic or biogenic carbonates, it has been predicted in theoretical studies. By combining a previous theoretical model of 13C18O 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 [Delta1]47T relationships for a number of carbonate minerals.
For the temperature range 030 C, [Delta1]47, aragonite-calcite 0.02
and [Delta1]47, aragonite-dolomite 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-specic [Delta1]47T relationships are 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 observed mineralogical effects in deep-sea corals and not in other types of carbonates in more detail below.
4.9 Calcication in scleractinian deep-sea corals
Nonequilibrium partitioning of oxygen and carbon isotopes has been observed in numerous calcifying organisms and inorganic precipitation experiments and linked to calcication processes (e.g., McConnaughey, 1989a; Spero et al., 1997;Zeebe, 1999; Adkins et al., 2003: Watson, 2004; DePaolo, 2011; Gabitov et al., 2012; Watkins et al., 2014). 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, 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 scleractinian corals, precipitation of carbonate is believed to occur in a space between the calicoblastic layer and the hard skeleton (Fig. 11). It is likely that calcifying uid chemistry (pH, salinity) as well as growth rate and DIC equilibrium can all be potential sources of nonequilibrium signals preserved in the solid. The chemical composition and DIC sources from which the carbonate mineral forms have been a main focus for understanding nonequilibrium 13C and 18O
values observed in corals. Biomineralization in shallow symbiotic scleractinian corals has been well-studied over the past few decades (see Cohen and McConnaughey, 2003, for a review). Early studies of shallow symbiotic scleractinian corals revealed time-independent 18O and 13C offsets from theoretical carbonatewater fractionation curves. This was addressed through species-specic 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 18O and 13C 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 biomineralization mechanism (i.e., vital effects). Different mechanistic models aimed at explaining nonequilibrium 13C and 18O values seen in deep-sea corals have been proposed to account for these so-called vital effects.
The rst of these models, proposed by Mc-Connaughey (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 13C and 18O values is therefore a reection of a spectrum of varying amounts of equilibrium being attained amongst DIC species. The master variable in this case being time, McConnaughey (1989b), using calcication rates, calculated that calcication could outpace the time it took for isotopic equilibrium 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 uid and introduces the concept of a variable pH environment as the master controller on skeletal 13C and 18O (pH model in Fig. 11). For oxygen isotope processes, they draw on work characterizing the oxygen isotopic fractionation between DIC species (i.e., CO2,
H2CO3,, HCO3, CO23) (McCrea, 1950; Usdowski et al., 1991; Zeebe and Wolf-Gladrow, 2001; Beck et al., 2005;
Zeebe, 2007; Rollion-Bard et al., 2011). Their model suggests that the mineral preserves the oxygen isotopic ratio of HCO3 / CO23. Because each DIC species has a unique oxygen fractionation relative to water, the oxygen isotope fractionation between the sum of DIC (DICsum) 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 CO23, which has the lowest fractionation factor and will therefore result in a more depleted 18O value of the DICsum from which the mineral precipitates. At lower (pH 69) uid pH, as the HCO3 / CO23 ratio of the uid rises, so does the precipitated minerals 18O value. 13C is not expected to be sensitive to pH effects since the exchange of carbon be-
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tween DIC species is very rapid. 13C, therefore, is proposed to be a function of two sources of CO2, membrane-diffused and seawater-leaked, and the proportion of them, a function of pH gradients across the membrane. The break that they uncover in the most depleted isotopic values, in which 18O
continues to decrease but 13C plateaus, is interpreted as the point at which the maximum ux of diffused CO2 across the membrane is reached (during rapid calcication). 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 18O and 13C that they report, it fails to reproduce the numerical value of the 18O
vs. 13C slopes that they observe. 18O vs. 13C 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; Rollion-Bard et al., 2003, 2010; Lutringer et al., 2005; Hill et al., 2014; Kimball et al., 2014). Discrepancies between the biomineralization model of Adkins et al. (2003) and boron isotope data are yet to be resolved (Blamart et al., 2007; Rollion-Bard et al., 2010).
Additional biomineralization models have been proposed in shallow symbiotic corals which incorporate the presence and activity of zooxanthellae (Cohen and McConnaughey, 2003; Rollion-Bard et al., 2003; Allison et al., 2010) and for deep-sea corals (Rollion-Bard et al., 2010; Gaetani et al., 2010). However, only those models that examine effects on 18O and 13C are considered with respect to the data presented herein.
More generally, models for equilibriumdisequilibrium stable isotope signatures have been proposed that examine processes that inuence the ion-by-ion growth of crystals (Nielsen et al., 2012; Watkins et al., 2014; Tripati et al., 2015), as well as diffusionalreordering processes that occur near the calcifying uidcrystal interface (Watson and Liang, 1995; Watson, 2004; Gabitov et al., 2012; Tripati et al., 2015).
4.10 Calcication in gorgonian deep-sea corals
The growth structure, mineralogy, and biomineralization of gorgonian coral skeletons is distinctly different from scleractinian corals (No and Dullo, 2006). Additionally, the organic matrix-mediated calcication process is largely unexplored and unknown for most gorgonian families. The biomineralization models described for scleractinian corals, therefore, are not valid when interpreting potential vital-effect signals in gorgonian corals.
An exhaustive study of the growth structure and fabric of deep-sea isidid gorgonians in the same family as those investigated here, was done by No and Dullo (2006). Their ndings are based on the characterization of the isidid skeletal micro- and ultrastructure in combination with absolute age determinations. From their reconstructions of growth mode and fabric, they propose a biomineralization model. With no calicoblastic cells observed in gorgonians, they propose that
gorgonian produced from the endodermal epithelium acts as the structural framework, while a viscous slime surrounding the skeleton acts as the matrix, facilitating and regulating mineral growth. Their model, however, relies on physical observations and lacks geochemical data to inform their interpretations. So while offering some insight into the uniqueness of isidid biomineralization, the ndings are not useful in interpreting the geochemical data observed in this study.
Despite the lack of proposed biomineralization models in gorgonian corals, the trends in 13C and 18O have been explored and found to be similar in pattern to scleractinian corals, but with an offset that agrees with theoretical mineralogical fractionation predictions. The similarity in 13C, 18O, and 47 patterns suggests that while the particular biomineralization process is necessarily different in the two groups, there is a functional similarity that results in the conserved patterns in both gorgonian and scleractinian corals.
4.11 [Delta1]47 in deep-sea coral skeletal carbonate: biomineralization model implications
The [Delta1]47 of a mineral is predicted to conform to the Keq of a homogenous isotope exchange reaction (Reaction R1) and be independent of the uid composition from which is forms, as well as the bulk isotopic composition of the crystal. Recent experimental and theoretical studies have suggested that mineral equilibrium with respect to [Delta1]47 might not always be attained, however (Saenger et al., 2012; Tripati et al., 2015).Because of this, [Delta1]47 can be utilized as an additional geo-chemical 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 [Delta1]47 signatures was examined with respect to [Delta1]47. Contrary to initial theoretical modeling by Hill et al. (2014), who calculated differences of 0.03
in the [Delta1]47 values of DIC species at equilibrium, Tripati et al. (2015), using experimental results, found that HCO3[Delta1]63 could be as much as 0.05 to 0.07 higher than CO23. They further propose that this fractionation could be partially or fully inheritable by the solid. Central to their mechanism of the preservation of disequilibrium mineral [Delta1]63 signals is the concept of differing mineral isotopic archival effects occurring during slow and fast crystal growth, with slow and fast being dened as particular to a specic temperature, salinity, and pH growth environment, rather than conforming to a particular value. Depending on the proportion of HCO3 and CO23 incorporated into the growing crystal and the ability of the crystal lattice to conform to Keq (Reaction R1), disequilibrium in [Delta1]63 and therefore [Delta1]47 could be preserved.Only during slow crystal growth, they argue, is the mineral able to fully conform to Keq (Reaction R1). In this scenario, disequilibrium in [Delta1]63 and bulk isotopes are not correlated and it would therefore be possible to attain equilibrium with respect to [Delta1]63, but not in bulk isotopes (Tripati et al., 2015).
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Deep-sea corals, being one of the slowest calciers, are a likely potential candidate for the attainment of [Delta1]63 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 [Delta1]47 values are reproducible within instrumental error. This indicates that while bulk isotopes are controlled by processes such as in the biomineralization models outlined above, [Delta1]47 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 (e.g., calcifying solution pH, kinetics), while the slow growth rates in deep-sea corals allow isotopic reordering at the crystalsolution interface resulting in the attainment of nominal mineral equilibrium for [Delta1]47 (Fig. 11).
Adkins et al. (2003), however, nd that in the more dense (interpreted as faster calcication) growth bands, 18O is de
pleted compared to the less dense (more slowly growing) bands. In the Tripati et al. (2015) model, they predict that during rapid calcication both 18O and [Delta1]47 would be more positive. Relative to the calcication rates required for inheritance of these signals, however, perhaps even at the most elevated calcication rates, deep-sea corals, are still able to reach mineral isotopic equilibrium with respect to [Delta1]47, while bulk 18O and 13C is still under the control of processes associated with the calcifying uid environment.
Cross plots of 18O and [Delta1]47 residuals (Figs. 7, 8, and 9) can be used to assess conformation to proposed biomineralization models. For the McConnaughey et al. (2003) model of coral biomineralization, a range of depletions in 18O is
expected, with coral 18O approaching or potentially reaching mineral equilibrium 18O, depending on the calcication rate. Similarly in the Adkins et al. (2003) model, a range of pH would result in a range of depletions in recorded mineral 18O, assuming that the proportion of HCO3 / CO23 incorporated into the skeleton is independent of calcication rate (Fig. 10). Although neither model addresses [Delta1]63, in both cases, all else being equal, and assuming a mechanism for inheritance of DIC [Delta1]63 by mineral [Delta1]63, a decrease in 18O would be tied to a decrease in [Delta1]63. For the McConnaughey et al. (2003) model, an increase in calcication rate resulting in only partial equilibration of the DIC-water system would result in depleted 18O and [Delta1]63 values of HCO3 and CO23.
Similarly, in Adkins et al. (2003) model, varying solution pH would have the same effect on 18O and [Delta1]63 with both values decreasing (increasing) as pH increases (decreases).Comparing residual cross plots from various [Delta1]47 calibrations, however, fails to capture a positive relationship between 18O and [Delta1]47, and it instead appears that 18O and [Delta1]47 are varying independently of each other. This suggests that either the models are not accurately portraying the mechanism by which variable 18O values are inherited in the mineral or the [Delta1]63 is not under the same controls as 18O (i.e., not inheriting the [Delta1]63 value of the uid).
4.12 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 18O and 13C trends similar to their scleractinian counterparts, they have not been shown to exhibit the break that is observed in Adkins et al. (2003) (Kimball et al., 2014). When growing inorganic carbonate precipitates, mineralogy can be controlled through the Mg / Ca ratio of the precipitating uid, with molar Mg / Ca > 2 producing aragonite and high-Mg calcite and molar Mg / Ca < 2 producing calcite (reviewed by Ries, 2010). Current seas have 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 Coral-lium sp. (Allemand et al., 1994; Ehrlich et al., 2006; No 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; Ehrlich et al., 2006).
It seems plausible that the organic template, in addition to directing lattice structure, could also be inuencing 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 [Delta1]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 the presence or absence and variable activities of carbonic anhydrase may be another factor to be considered.
4.13 Suggestions for future work
In-depth experimental studies, such as Wacker et al. (2013) and Deiese et al. (2015), are needed to explore the role of multiple processes in acid digestion kinetics. Possible factors inuencing acid digestion fractionation factors that still need to be examined include (but are not limited to) organic matter content, the inuence of cation substitution (e.g., high vs. low Mg calcite), variations in bulk composition ( 13C, 18O), the type of reaction vessel and cleanup 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-specic offsets. Additionally, the measurement of coral-calcifying uid pH and the use of labeled carbon sources during growth experiments would provide important constraints on whether the observed patterns reect kinetic effects.
A unifying model which explains the observed patterns of isotopic and trace element variability that have been ob-served 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 inuence equilibrium fractionation, i.e., temperature and well-constrained mineral-specic fractionation factors. Additional factors that need to be included to accurately model kinetic isotope effects include temperature, the carbon source used for calcication (the relative fraction of seawater-derived DIC vs.CO2 diffused across cellular membranes), DIC speciation as controlled by multiple factors (pH, temperature, salinity), and the timescale for DIC equilibration vs. precipitation rate (affected by the residence time of DIC in the calcication 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 13C, 18O, and [Delta1]47. Aragonitic scleractinian deep-sea coral exhibit a nearly coincident [Delta1]47T relationship to previously reported scleractinian deep-sea corals (Thiagarjan et al., 2011) and can be used to predict temperatures to [notdef]2 C at the mean tem
perature and [notdef]35 C at the extreme ends of the data set.
High-Mg calcitic gorgonian deep-sea corals show lower [Delta1]47 values. This offset between taxa is consistent with theoretical predictions for mineralogic differences, albeit signicantly larger in magnitude than expected. Therefore, the offset between gorgonian and scleractinian deep-sea corals may be reected, 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 cannot preclude that offsets may reect variable acid digestion kinetics for the different types of materials that we have surveyed.
Information about the Supplement
All raw data for samples and standards and equilibrated gas lines are provided in the data tables in the Supplement.
The Supplement related to this article is available online at http://dx.doi.org/10.5194/bg-13-6487-2016-supplement
Web End =doi:10.5194/bg-13-6487-2016-supplement .
Acknowledgements. This work was supported by a NSF Graduate Fellowship, NSF OCE-1437166, DOE BES grant DE-FG02-13ER16402, and the Laboratoire dExcellence LabexMER (ANR-10-LABX-19), co-funded by a grant from the French government under the program Investissements dAvenir. We thank Aradhna Tripati and the Tripati Lab Group for their assistance with analyses and discussion of this work and Ben Elliott for assistance with drafting gures.
Edited by: A. ShemeshReviewed by: two anonymous referees
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Abstract
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 <sup>13</sup>C<sup>18</sup>O<sup>16</sup>O<sub>2</sub> in carbonate minerals, analyzed in CO<sub>2</sub> gas liberated on phosphoric acid digestion of carbonates and reported as Δ<sub>47</sub> values. We analyzed Δ<sub>47</sub> in live-collected aragonitic scleractinian (Enallopsammia sp.) and high-Mg calcitic gorgonian (Isididae and Coralliidae) deep-sea corals and compared results to published data for other aragonitic scleractinian taxa. Measured Δ<sub>47</sub> values were compared to in situ temperatures, and the relationship between Δ<sub>47</sub> and temperature was determined for each group to investigate taxon-specific effects. We find that aragonitic scleractinian deep-sea corals exhibit higher values than high-Mg calcitic gorgonian corals and the two groups of coral produce statistically different relationships between Δ<sub>47</sub>-temperature calibrations. These data are significant in the interpretation of all carbonate clumped isotope calibration data as they show that distinct Δ<sub>47</sub>-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<sub>2</sub> 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 reflect high-Mg and aragonite crystals attaining nominal mineral equilibrium clumped isotope signatures due to 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 CO<sub>2</sub> 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 Δ<sub>47</sub> is a subject that warrants further investigation.
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