Kinetic effects during the experimental transition of aragonite to calcite in aqueous solution: Insights from clumped and oxygen isotope signatures

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ScienceDirect Geochimica et Cosmochimica Acta 248 (2019) 210–230 www.elsevier.com/locate/gca

Kinetic effects during the experimental transition of aragonite to calcite in aqueous solution: Insights from clumped and oxygen isotope signatures Yangrui Guo a,b, Wenfeng Deng a,c,⇑, Gangjian Wei a a

State Key Laboratory of Isotope Geochemistry, Guangzhou Institute of Geochemistry, Chinese Academy of Sciences, Guangzhou 510640, China b College of Earth and Planetary Sciences, University of Chinese Academy of Sciences, Beijing 100049, China c Radiogenic Isotope Facility, School of Earth and Environmental Sciences, The University of Queensland, Brisbane, Queensland 4072, Australia Received 22 May 2018; accepted in revised form 5 January 2019; available online 14 January 2019

Abstract Paleoenvironmental reconstructions based on the clumped isotopes (D47) and traditional isotope (d13C or d18O) techniques are often problematic for carbonates that have undergone diagenetic alteration. One of the most common types of diagenesis is the transition between polymorphs, such as the replacement of aragonite by calcite. The isotope fractionation during such transitions in aqueous solutions remains unclear. We conducted a series of aragonite-to-calcite transition experiments in aqueous solutions of varying salinity and experiment durations at 25 °C and 90 °C to examine the variations of d13C, d18O, and D47 values for carbonates at varying degrees of the transition. The results confirm the retarding effect of Mg2+ and the catalytic effect of Na+ and Ca2+ on the transition, which are consistent with previous findings. Compared with the results from the transitions at 25 °C, the 90 °C experiments show a greater transition to calcite and a more depleted oxygen isotope composition. However, both clumped and carbon isotope values show no significant variation in the various experiments, suggesting that they are unaffected by mineralogical transition. Based on published equilibrium calibrations for d18O and D47, the results demonstrate that kinetic effects in isotope systems are controlled primarily by the rate of polymorph transition, but there is a significant kinetic difference between the clumped isotope bond (13CA18O) reordering and the 18O exchange within the dissolved inorganic carbon pool. This kinetic difference results in partial equilibration in d18O with no significant reordering in D47 toward their equilibriums, which can be accounted for by the difference of equilibration rates between the oxygen isotopes bound to 12C and those bound to 13C. This study provides a clear observation of the response of isotope systems to carbonate polymorph transitions and sheds light on their reliability as paleoenvironmental proxies. Ó 2019 Elsevier Ltd. All rights reserved. Keywords: Clumped isotopes; Carbonate polymorphism; Isotope kinetic effect; Oxygen isotope; Carbon isotope; Bond reordering; Dissolved inorganic carbon

1. INTRODUCTION

⇑ Corresponding author at: State Key Laboratory of Isotope

Geochemistry, Guangzhou Institute of Geochemistry, Chinese Academy of Sciences, Guangzhou 510640, China. E-mail addresses: [email protected], [email protected] (W. Deng). https://doi.org/10.1016/j.gca.2019.01.012 0016-7037/Ó 2019 Elsevier Ltd. All rights reserved.

The temperature dependence of oxygen isotope partitioning between water and carbonates forms the basis of a paleo-thermometer that has become one of the most important tools for reconstructing past seawater temperatures (Urey, 1947; Epstein et al., 1953; Emiliani, 1955; Shackleton, 1967; Kim and O’Neil, 1997). However, this

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carbonate–oxygen isotope paleo-thermometer does not provide direct temperature information because it also relies on the oxygen isotope composition of the water (d18Owater) in which the carbonate precipitated. Consequently, it is not applicable in situations where the d18Owater is unknown. Fortunately, the emerging clumped isotope thermometry method is potentially able to solve this problem by determining the carbonate formation temperature directly, without needing to also know d18Owater (Ghosh et al., 2006a; Eiler, 2011). This novel thermometry technique is based on measurements of the extent to which 13 C and 18O chemically bond to each other within the carbonate crystal lattice (Schauble et al., 2006; Eiler, 2011). The degree of the 13CA18O bond ordering (or clumping) is controlled solely by temperature and is conventionally quantified by the D47 value as follows:  D47 ¼

  46   45  R47 R R  1   1   1  1000‰ R47 R46 R45

where the R47, R46, and R45 are the measured ratios of mass 47, 46, and 45 relative to mass 44, respectively. The R47 , R46 , and R45 are the corresponding mass ratios where all isotopes in the sample gas are in stochastic distribution (Eiler and Schauble, 2004; Huntington et al., 2009). Over the past decade, this powerful thermometry tool has been applied to various topics in geological science, including the evolution of seawater temperature during the Phanerozoic (Came et al., 2007; Dennis et al., 2013; Douglas et al., 2014; Tripati et al., 2014; Rodrı´guez-Sanz et al., 2017; Henkes et al., 2018), the elevation history of Earth’s surface (Ghosh et al., 2006b; Huntington et al., 2010; Lechler et al., 2013; Garzione et al., 2014; Hough et al., 2014; Snell et al., 2014; Kar et al., 2016; Methner et al., 2016), and carbonate diagenesis in various settings (Huntington et al., 2011; Dale et al., 2014; Bradbury et al., 2015; Shenton et al., 2015; Defliese and Lohmann, 2016; Stolper and Eiler, 2016; Winkelstern and Lohmann, 2016; Mu¨ller et al., 2017c; Ritter et al., 2017; Ahm et al., 2018; Staudigel et al., 2018; Stolper et al., 2018). It is crucial to ensure that the carbonate isotope composition represents that of the primary mineral, as diagenetic alteration resulting from various post-depositional processes, such as meteoric, burial, and metamorphic diagenesis (Swart, 2015), can influence isotope signatures. Good preservation of the original aragonite has historically been used as an indicator for unaltered geochemical archives (Stahl and Jordan, 1969). Aragonite is highly susceptible to post-deposition diagenetic alteration and readily undergoes a process known as carbonate polymorphic transition (Bischoff and Fyfe, 1968; James, 1974; James and Choquette, 1984; Budd, 1988). Aragonite is thermodynamically unstable in subaerial environments and transforms to calcite under certain geological conditions (Plummer and Busenberg, 1982; Budd, 1988; Morse et al., 2007; Radha and Navrotsky, 2013). This can even occur during sample pretreatment in experimental environments (Waite and Swart, 2015; Staudigel and Swart, 2016). As the original isotope composition is potentially modified during transition, it is important to under-

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stand the degree to which the original isotope signatures are preserved after the transition. The aragonite-to-calcite transition occurs at high temperatures under anhydrous condition (Yoshioka and Kitano, 1985) and affects D47 values (Staudigel and Swart, 2016), mainly due to a solid-state reordering of the 13 CA18O bond within the carbonate crystal lattice under high temperature (>100 °C) (Dennis and Schrag, 2010; Passey and Henkes, 2012; Stolper and Eiler, 2015). However, the transition can also occur in aqueous solutions via a dissolution-reprecipitation process usually at lower temperatures (Taft, 1967; Bischoff, 1969; Yoshioka et al., 1986; Perdikouri et al., 2011), which can also lead to reordering of the 13CA18O bond within the dissolved inorganic carbon (DIC) pool derived from the aragonite (Eiler, 2011). The state of the bond reordering approaches the equilibrium specific to the temperature of the reprecipitated calcite during the transition. Hydrothermal experiments with aragonite (conducted at 100 °C) have shown a marked change in D47 values following the replacement of aragonite with calcite (Mu¨ller et al., 2017c; Ritter et al., 2017). In these studies, D47 variations were attributed to clumped isotope solid-state reordering similar to the changes that occur under high-temperature conditions. However, experiments studying the transition of aragonite to calcite and alteration of D47 under aqueous conditions at lower temperatures (<100 °C) have not been conducted. It is not known if D47 alteration will occur in the ‘solutionstate’ transition at temperatures below the threshold for solid-state reordering. As the reprecipitated calcite will resist further mineralogical alteration (Defliese and Lohmann, 2016), it is important to understand what meaningful information can be interpreted from the calcite D47. When applying both D47 and d18O to study diagenesis associated with the transition of aragonite-to-calcite, there is an implicit assumption that the reprecipitated calcite is in isotopic equilibrium with the diagenetic fluids (Huntington et al., 2011). The d18Owater of diagenetic fluids can then be directly estimated based on carbonate d18O values, the temperatures constrained by clumped isotope thermometry, and the equilibrium carbonate–water fractionation factor experimentally determined by slow calcite precipitation (Kim and O’Neil, 1997). This assumption is valid in environments where the carbonate polymorphic transition proceeds sufficiently slowly to allow isotopic exchange at equilibrium. However, there is currently no convincing evidence to support this assumption of isotopic equilibrium. In addition, diagenetic fluids usually contain different concentrations of dissolved salts to pure water (Zeebe and Wolf-Gladrow, 2001). As ions such as Na+ and Ca2+ facilitate the aragonite-to-calcite transition, their presence may increase transition rates (Taft, 1967; Bischoff and Fyfe, 1968; Yoshioka et al., 1986), resulting in potential kinetic effects on isotopic composition. The potential influence of dissolved salts on both D47 and d18O has recently been investigated in carbonate precipitation, which suggested a significant effect due to high concentrations of CaCl2 (Kluge and John, 2015). It is thus important to consider the kinetic effects that may occur during the rapid reprecipitation of calcite from aragonite in presence of dissolved salts.

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In this study, we conducted transition experiments at two temperatures (25 °C and 90 °C) and a range of durations, using coral aragonite and solutions of different salt additions including NaCl, CaCl2 and MgCl2. At the conclusion of the transition experiments, we measured the mineralogical variations, clumped isotope composition, and bulk isotope compositions (d13C and d18O) of the carbonates and d18O of the solutions. The main purposes of this study are to better understand the extent to which the altered carbonates preserve their isotopic signals during the polymorphic transition, and to identify the potential kinetic effects on the D47 and d18O isotope systems during the solution-state transition. 2. MATERIALS AND METHODS 2.1. Starting materials We used coral skeletal aragonite pieces to supply the large quantity (1 g) of aragonite powder required for each experimental run. Biogenic aragonites were taken from a Porites sp. colony growing on a fringing reef in Sanya Bay at the southern edge of Hainan Island, China. All soft tissue was removed, and the coral pieces were washed with distilled water several times and air-dried. The materials were crushed, ground to powder with a mortar and pestle, and sieved to obtain the <75 lm grain size fraction. X-ray diffraction analysis revealed that the resulting powder contained 5 wt% calcite. This trace amount of calcite would likely serve as seed crystals for stimulating rapid calcite recrystallization. 2.2. Transition experiments The aragonite-to-calcite transition experiments were performed in several air-tight plastic containers with different treatments, similar to the experiments of Yoshioka et al. (1986). To simulate diagenetic alteration with different aqueous fluids and investigate the potential effects of dissolved salts on isotope fractionation, we used several different solutions: deionized water (prepared by an UltrapureTM 18 MX system); water with NaCl, CaCl2, or MgCl2(H2O)6; and water with all these reagents or various combinations of two of the reagents. The concentrations of the solutions were similar to that of modern seawater (Table 1). Solutions with high concentrations of dissolved salts were included in the experiments because natural diagenetic processes usually involve saline fluids with concentrations higher than those of seawater (Bein and Land, 1983; Budd et al., 2013). As salt additives may influence the transition process even at very low concentrations (Meyer, 1984; Gutjahr et al., 1996), the chemicals used for this study were all reagent-grade (available from CodowÒ) to minimize the effects of any other dissolved ions. For each experiment, 1 g of aragonite powder was placed in a polypropylene bottle and 100 ml of solution was added with no air headspace. This reaction container was tightly sealed by a cap with a rubber septum inside and quickly immersed in a thermostatic water bath oscillator which constantly stirs the solution at a rate of 100 rpm. The experiments were set at 25 °C and 90 °C maintained by

a temperature controlled water circulation system with thermal equilibrium reached in a few minutes. The experiment duration time for most treatments was 265 h. Treatments at 90 °C in deionized water, water–NaCl, and water–CaCl2 were analyzed in time series due to their high mineral transition rates (see the experimental conditions listed in Table 1). The pH and salinity were measured at the beginning and end of every experimental run, using a Thermo Orion 4-star Plus pH/conductivity meter with an uncertainty of 0.02 in both cases. After the experiment, the container was removed from the thermostatic bath and rapidly cooled in cold water to avoid possible isotopic re-equilibration. The solid powder adherent on the filter was dried at room temperature in a vacuum drying oven and collected into a sample vial ready for subsequent analyses. The solution was passed through a 0.45 lm nylon filter and partly preserved in a capped bottle containing one drop of saturated HgCl2 solution to prevent isotopic fractionation from bacterial metabolism. The remnant solution was analyzed for its calcium concentration to estimate possible dissolution of the carbonate after the experiments. The results suggested that the dissolution was unlikely to affect the carbonate isotopic composition. We thus assumed that the transition reactions involved mineral dissolution and reprecipitation without any net change in the overall mass of carbonate minerals (see Table S1 in the Supplemental Materials). 2.3. Mineralogical observation X-ray powder diffraction (XRD) measurements were performed at the Key Laboratory of Mineralogy and Metallogeny at the Guangzhou Institute of Geochemistry, Chinese Academy of Sciences, Guangzhou, China, using a Bruker D8 Advance instrument with Ni-filtered Cu Ka radiation. Samples of 0.5 g of powder were loaded into a plastic sample holder and the surface of the powder was smoothed with a piece of flat glass. The working voltage and current were set at 40 kV and 40 mA, respectively. The diffraction spectral pattern was measured at a 2h step scan of 0.01° between 10° and 60°. The XRD patterns were directly obtained from a software (Evaluation) supplied with the XRD instrument. The relative weight percentages of aragonite and calcite were determined using the reference intensity ratio (RIR) method (Chung, 1974; Snyder, 1992) similar to the one used by Dominguez-Villar et al. (2017). This was achieved by using Jade 5.0 software interfaced with the Powder Diffraction File (PDF), a mineral database wherein the published RIRs specific to calcite and aragonite can be found. The uncertainty of this quantitative analysis was 5% (1r) as assessed from repeated analyses by in-home standards. The morphologies of samples before and after the transition experiment were observed with a Hitachi (SU8010) fieldemission scanning electron microscope (SEM). 2.4. Isotope analysis of solid samples Clumped isotope and bulk isotope measurements were conducted at the State Key Laboratory of Isotope Geochemistry (SKLaBIG) at the Guangzhou Institute of Geo-

Table 1 Experimental conditions and isotopic results. The uncertainty in the calcite phase quantification is 5%. The d18Osolution is measured oxygen isotope composition in solutions with precision better than 0.1‰. Both the d18Osolid and d13Csolid refer to the isotope composition of the carbonate samples. The D47, RF-AC is the clumped isotope value after acid correction (see Section 2.4.5). The uncertainties for all the isotope values of carbonates are given as one standard deviation (1r) of replicate analyses. The predicted oxygen isotope compositions for carbonate products of the transition experiments are calculated using the calcite–water isotope calibration of Kim and O’Neil (1997), employing the measured d18Osolution values in this study, and assuming the newly-formed calcite and remaining aragonite follow a linear mixing regime. The predicted clumped isotope composition is based on the laboratory-determined D47–T calibration of Kluge et al. (2015) and the mixing model of Defliese and Lohmann (2015). Experiment no.

Temp (°C)

Duration (h)

No salt added

Salt concentration (g/L)

Corresponding ion concentration (mol/L)

Mg

Na

Ca

Mg

90 90 90 90 90 90 90 90 25

23 49 93 170 265 265 265 265 265

_ _ _ _ _ _ _ _ _

_ _ _ _ _ _ _ _ _

_ _ _ _ _ _ _ _ _

Na (NaCl) added NA-1 NA-2 NA-3 NA-4 NA-5 NA-6 NA-7 NA-8

90 90 90 90 90 25 90 25

23 49 93 170 265 265 265 265

_ _ _ _ _ _ _ _

10 10 10 10 10 10 80 80

_ _ _ _ _ _ _ _

Ca (CaCl2) added CA-1 CA-2 CA-4 CA-5 CA-6 CA-7 CA-8

90 90 90 90 25 90 25

23 49 170 265 265 265 265

_ _ _ _ _ _ _

_ _ _ _ _ _ _

0.4 0.4 0.4 0.4 0.4 2.0 2.0

Mg (MgCl26H2O) added MG-1 MG-2 MG-3 MG-4

90 90 25 25

265 265 265 265

1.3 0.1 1.3 0.1

_ _ _ _

_ _ _ _

d18Osolution VSMOW (‰)

9.51 9.55 9.40 9.40 9.22 9.22 9.23 9.24 9.80

0.16 0.15 0.17 0.18 0.18 0.18 0.19 0.19 0.13

5 18 40 54 56 54 58 62 4

6.0 5.6 5.8 5.9 6.4 5.3 6.5 5.9 5.4

9.40 9.40 9.36 9.32 9.20 9.72 8.77 9.20

31.0 30.6 30.2 30.6 29.1 29.2 128 144

75 82 85 82 82 5 93 4

8.34 8.25 8.38 8.20 8.53 7.88 8.02

1.81 1.76 1.78 1.63 1.70 7.05 7.22

9.02 9.38 9.49 9.62

7.33 0.89 7.33 0.78

d13Csolid

1r

d18OsolidVPDB (‰)

1r

D47 RFAC (‰)

1r

n

d18Opredicted (‰)

D47predicted (‰)

1.87 1.87 1.87 1.88 1.87 1.88 1.86 1.86 1.86

0.01 0.01 0.01 0.02 0.02 0.01 0.00 0.01 0.01

4.97 5.41 6.16 6.94 7.88 7.30 7.59 7.67 4.48

0.02 0.02 0.04 0.04 0.15 0.05 0.04 0.01 0.01

0.718 0.722 0.721 0.724 0.709 0.697 0.702 0.708 0.726

0.014 0.011 0.013 0.014 0.008 0.010 0.009 0.008 0.008

3 3 3 3 3 3 3 3 3

5.15 6.93 10.13 12.26 12.85 11.99 13.22 13.41 4.51

0.722 0.701 0.664 0.639 0.635 0.638 0.631 0.624 0.722

5.9 5.5 5.8 5.9 5.5 5.7 5.8 5.9

1.86 1.87 1.87 1.88 1.87 1.84 1.89 1.86

0.00 0.01 0.01 0.01 0.01 0.02 0.01 0.01

6.83 8.63 9.23 8.79 8.91 4.48 10.87 4.49

0.04 0.06 0.04 0.01 0.04 0.02 0.03 0.07

0.734 0.724 0.725 0.722 0.716 0.720 0.720 0.715

0.013 0.012 0.016 0.022 0.018 0.007 0.016 0.008

5 3 4 4 5 3 3 3

15.30 16.03 16.77 16.37 16.07 4.58 17.78 4.54

0.602 0.589 0.582 0.588 0.588 0.723 0.569 0.723

61 79 86 87 10 92 8

6.0 5.6 6.2 5.7 5.9 5.7 5.5

1.89 1.88 1.87 1.87 1.87 1.87 1.87

0.02 0.02 0.01 0.01 0.01 0.01 0.02

5.43 6.71 7.20 7.72 4.47 7.53 4.45

0.06 0.06 0.14 0.06 0.06 0.04 0.03

0.724 0.735 0.729 0.725 0.729 0.735 0.720

0.003 0.004 0.012 0.003 0.016 0.009 0.004

3 3 3 3 3 3 3

13.36 15.69 17.10 16.86 4.78 17.67 4.67

0.626 0.594 0.582 0.580 0.721 0.570 0.722

4 4 5 4

5.4 5.7 5.9 5.8

1.84 1.86 1.85 1.85

0.01 0.01 0.01 0.01

5.61 5.42 4.46 4.47

0.03 0.03 0.14 0.05

0.694 0.705 0.717 0.732

0.012 0.025 0.006 0.014

3 4 3 3

4.92 4.99 4.59 4.53

0.721 0.721 0.723 0.723

VPDB

(‰)

Ca

0.43 0.43 0.43 0.43 0.43 0.43 3.48 3.48

0.01 0.01 0.01 0.01 0.01 0.05 0.05

0.053 0.004 0.053 0.004

Calcite content (%)

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DW-1 DW-2 DW-3 DW-4 DW-5-1 DW-5-2 DW-5-3 DW-5-4 DW-6

Na

Salinity (TDS, g/ L)

pH

213

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4 0.011 0.723 0.06 0.01 1.86

4.39

0.01 0.03 0.00 0.01

5.9 0.01 7.04

5

30.3 34.1 8.45 35.0 0.43

2.4.1. CO2 extraction and purification Mass 47 signal of CO2 derived from carbonate acid digestion can be used to determine the clumped isotope composition of the reactant carbonate (Ghosh et al., 2006a; Schauble et al., 2006). It is conventional to extract CO2 from carbonates through reaction with anhydrous phosphoric acid. According to the method of Burman et al. (2005), the acid used in this study was prepared to be 1.93 g/cm3 in density corresponding to 104% acid concentration. Between 6 and 8 mg of powdered carbonate sample or standard was loaded in a quartz cup (1 cm3 in volume) and placed in a custom-built reaction vessel in which the prepared acid was placed at the bottom without touching the carbonate powder. The reaction vessel was directly connected to a custom-made stainless-steel vacuum line (Fig. S1). Prior to acid digestion, the reaction vessel was pumped down to <10–3 to 10–2 mbar and the reaction temperature was stabilized to 90 °C (±0.2) using a water bath. This acid digestion setup follows the method descried in Kluge and John (2015) and is similar to a common acid bath. Each reaction required 3 ml new acid and a precleaned vessel for each sample. Starting the reaction was achieved by using magnets between the vessel walls to push the quartz cup into the phosphoric acid. The CO2 gas produced from the carbonate was simultaneously collected in a liquid nitrogen (LN2) trap (198 °C). After the reaction was completed (confirmed by the complete cessation of gas bubbling), the reaction vessel was isolated and the CO2 remaining in the LN2 trap was ready for further purification. The procedures for CO2 cleaning were based on those in previous studies (Eiler and Schauble, 2004; Ghosh et al., 2006a; Dennis and Schrag, 2010; Kluge et al., 2015). The first LN2 trap was replaced by an ethanol slush (100 ± 5 °C) to release the CO2 and isolate the water. The CO2 was then allowed to pass through a U-shaped PoraPakTM Q (PPQ) trap held at 20 °C without a carrier gas and was collected in a second LN2 trap. The PPQ trap removed contaminations (e.g., heavy hydrocarbons or chlorinated hydrocarbons) that would interfere with the mass 47 (Eiler and Schauble, 2004; Dennis and Schrag, 2010). The time taken for the gas passing through the PPQ trap is dependent on sample size and was set to 25 min in our setup to ensure complete gas recovery. Before finally transferring the gas into a coldfinger (or being sealed into a quartz tube), non-condensable gas was pumped away. The remaining CO2 in the second LN2 trap was then released by the ethanol slush and collected into the coldfinger ready for further analysis.

Starting materials Coral aragonite Deionized water

Salts combined NC90 NM90 CM90 NCM90

90 90 90 90

265 265 265 265

_ 1.3 1.3 1.3

10 10 _ 10

0.4 _ 0.4 0.4

0.053 0.053 0.053

0.43 0.43

0.01 0.01

0.01

8.39 9.05 8.92 8.74

92 4 4 4

5.7 6.0 5.8 5.8

1.88 1.87 1.86 1.86

8.26 5.87 5.51 5.73

0.04 0.05 0.02 0.01

0.727 0.710 0.699 0.707

0.014 0.029 0.006 0.009

3 3 3 3

17.72 4.91 4.96 4.99

0.570 0.721 0.721 0.721

chemistry, Chinese Academy of Sciences, between October 2017 and January 2018, using a dual-inlet isotope ratio mass spectrometer (253 Plus, Thermo ScientificTM) by the following method.

2.4.2. Standard gas preparation As with previous clumped isotope studies (Huntington et al., 2009; Dennis et al., 2011; Mu¨ller et al., 2017a), the 253 Plus used in this study shows a slight non-linear effect between actual and measured R47 values, which can affect

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the D47 value of a sample gas if its bulk composition is significantly different to that of the reference gas (Eiler and Schauble, 2004). Therefore, a suite of CO2 gas standards with the same clumped isotope composition but different bulk compositions is needed to establish a calibration line to correct for this effect. The standard gases can be prepared by the heated gas method or equilibrated gas method. In the heated gas method, CO2 is sealed in a quartz tube, heated to 1000 °C in a muffle furnace for 2 h, and then rapidly quenched to room temperature to maintain its 13 CA18O bond-ordering state. In the equilibrated gas method, CO2 is isotopically equilibrated at 25 °C through a H2OACO2 reaction for 2 days. The raw clumped isotope compositions of the gas standards used in this study are listed in the supplemental materials (Tables S2 and S3). In addition, two in-home standards (BACS, tropical Porites coral aragonite; ISCS, Iceland spar calcite) and two Carrara marble standards (NBS-19 and C1, available from IAEA) were used to monitor instrument drift. Prior to analysis, the carbonate and gas standards were digested or cleaned in the same way as the sample gas, and freshly measured to prevent re-equilibration. 2.4.3. Instrumentation The 253 Plus installed at SKLaBIG has a Faraday cup collector configuration for collecting signals of mass 44– 49. An additional half-mass detector (cup 47.5) is included to provide simultaneous baseline monitoring. The signals for mass 44–46 are measured with standard amplifications of 3  108, 3  1010, and 1  1011 X, respectively, and the signals for mass 47–49 (including mass 47.5) are measured with 1  1013 X amplification. The original nickel capillaries from Thermo ScientificTM were replaced by 3-foot-long Ò nickel capillaries manufactured by VICI (Part No. TEFNI.505) due to the clumped isotope redistribution reac13 18 16 tion (e.g., C O O + 12C16O16O M 13C16O16O 12 18 16 + C O O) that occurred inside the capillaries as gases passed through (Passey et al., 2010; Wacker et al., 2013). The gases were introduced via a dual-inlet system. Bellows from both the reference and sample sides were automatically adjusted to achieve 10 V for mass 44. Peak centering, background measurement and pressure adjustment were performed before each acquisition. Six acquisitions, each consisting of 10 cycles of 26 s of integration time and a 12 s changeover delay time, were used for one analysis. The internal precision (1 standard error, SE) was usually 0.01‰ for d13C and d18O, and 0.009‰ for D47. The precision for D47 at the total integration time of 1560 s was close to its corresponding shot noise level as calculated by Merritt and Hayes (1994), suggesting there were no significant precision-limiting artifacts besides ion detection and counting statistics. The reference gas was sourced from pure CO2 tank gas (Linde Group) with d13CVPDB = 26.78‰, d18OVPDB = 8.95‰, and was normalized to two IAEA standards (NBS19 and NBS18) and checked daily against in-home standards as mentioned in Section 2.4.2. 2.4.4. Background treatment Generally, backgrounds are measured when no gas is presented in the ion source, while scattered ion back-

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grounds are monitored during measurement. Due to their pressure dependence, the second backgrounds have been generally referred to as pressure baseline (PBL). They have been considered to affect the linearity of the instrument (He et al., 2012). Their origin remains unclear; however, they can be corrected for by measuring the off-peak background signals and substituting these to represent the true background during on-peak measurements (He et al., 2012; Bernasconi et al., 2013; Rosenheim et al., 2013). Failure to correct this background has been shown to impact not only the internal precision but also the accuracy of the CO2 D47 measurement (Rosenheim et al., 2013; Fiebig et al., 2016). In this study, the intensity of the mass 47.5 cup did not directly represent the true PBL received in the mass 47 cup but had to be multiplied by a factor (f), and this is taken into account when detecting the true mass 47 signal (e.g., I47,true = I47,measured  f  I47.5 (Radke et al., 2017)). The f value can be modified from Isodat, a software available from the 253 Plus, and determined by calibration with gas standards of known clumped isotope compositions. Compared with the traditional method used to solve this linearity issue (Eiler and Schauble, 2004; Huntington et al., 2009), our background treatment did not have a significant difference on the final reported D47 value (Figs. S2– S4). Therefore, all the D47 measurements reported in this study were treated with the 47.5 cup-based PBL correction. 2.4.5. Data processing Raw data files were extracted from Isodat and processed with Easotope, a community software specific for clumped isotope data reduction (John and Bowen, 2016). As different isotopic parameters can potentially introduce significant errors in D47 calculations (Schauer et al., 2016), we used a value of 0.528 for k (Barkan and Luz, 2005) and revised isotope abundances for VPDB 13C/12C (Zhang et al., 1990), VSMOW 18O/16O (Baertschi, 1976) and VSMOW 17 O/16O (Assonov and Brenninkmeijer, 2003), following Dae¨ron et al. (2016). As both d18O and d13C values for solid samples are measured along with the clumped isotopes, the bulk isotope compositions were processed with the same isotopic parameters and exported from Easotope. As CO2 produced from carbonate acid digestion involved oxygen isotope fractionation, the phosphoric acid fractionation factors for both 25 °C or 90 °C for both calcite and aragonite were taken from Kim et al. (2007a). For samples with mixtures of calcite and aragonite, oxygen isotope compositions were calculated proportionally using the corresponding acid fractionation factors; e.g., d18OMixture = (1  f )  (d18Oaragonite) + f d18Ocalcite, where the f is the content ratio of calcite as determined by XRD analysis (Table 1). The non-linearity correction of the raw D47 values was calculated using the heated gas data (Huntington et al., 2009) and then transferred to an absolute reference frame (ARF) by the empirical transfer function (ETF) (Dennis et al., 2011). The ETF was established using both heated and equilibrated gases standards to correct for the scale compression or stretching due to fragmentation and recombination of isotopologues in the ion source (Dennis et al., 2011). A detailed correction procedure for the absolute ref-

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erence frame was taken from Dennis et al. (2011). In this study, there were four periods in which different heated gas lines were used (P1–P4, see Table S4). This was due to occasional source bake-out, inlet part repair, or the use of different 47.5 cup f values. Carbonate standard measurements were obtained over the last three periods, but during the P4, only equilibrated gas measurements were available for the calculation of the accepted D47 values on ARF scale (D47-RF). Therefore, the D47-RF values for all other carbonate standards as well as samples in both P2 and P3 were determined by two secondary transfer functions (Table S5). These transfer functions were established through a linear regression between the linearity-corrected data and the accepted ‘‘true” values of the standards (Dennis et al., 2011). As the pioneering clumped isotope analyses used an acid digestion temperature of 25 °C (Ghosh et al., 2006a) rather than 90 °C, as used in this study, an acid fractionation factor (AFF) was necessary to normalize our data for comparison with existing results. We used an AFF value of 0.068‰ (±0.013, SD), which was determined through digestion of three carbonate standards (NBS19, C1, and BACS) at 25 °C in a McCrea-type reaction vessel (Ghosh et al., 2006a), following the same purification method as used with on other samples in this study and reporting the D47 values on the ARF scale (Table S6). The AFF is close to 0.069‰, the theoretically predicated value for calcite digested at 90 °C (Guo et al., 2009). As the AFF is considered to be independent of carbonate mineralogical differences under the current analyti-

cal resolution (Wacker et al., 2013; Defliese et al., 2015), the same AFF was applied to all measured carbonates. However, it is uncertain whether the AFF is mineralogydependent with respect to all different carbonates (e.g., calcite vs. dolomite) (Bonifacie et al., 2017; Mu¨ller et al., 2017b) and this issue requires further research. Using the data processing methods described above, the average D47 values on the ARF scale with acid correction for four carbonate standards (BACS, ISCS, C1, and NBS19) were 0.721 ± 0.013‰ (1SD, n = 20), 0.569 ± 0.013‰ (1SD, n = 20), 0.387 ± 0.015‰ (1SD, n = 28) and 0.380 ± 0.014‰ (1SD, n = 9), respectively. The values obtained for the two Carrara marble standards were in agreement with published values (Dennis et al., 2011). Raw data relevant to carbonate samples and standards are included in the supplemental material (Tables S7 and S8). 2.5. Isotope analysis of solutions d18Osolution was measured according to the classic CO2– H2O equilibration technique (Cohn and Urey, 1938; Epstein and Mayeda, 1953) and using a dual-inlet mass spectrometer (GV IsoPrime) coupled with an online sample preparation system (MultiprepTM) at SKLaBIG. Details of the analytical procedures were described by Xie et al. (2011). Results for all solution samples were normalized to VSMOW scale using the IAEA standards (GISP, VSMOW2, and SLAP2). The analytical precision for d18Osolution was 0.10‰ based on repeated measurements

Fig. 1. SEM images of selected samples and corresponding XRD patterns. Panel A: coral aragonite with trace amounts of calcite, as used as starting materials for the experiments. Panel B: sample (DW-2) treated with deionized water for 49 h at 90 °C shows corroded crystal surfaces, indicating partial dissolution of the initial aragonite and >18% inversion to calcite. Panel C: sample (CA-6) treated with CaCl2 for 265 h at 90 °C shows rhombic calcite assemblages, confirmed by a high calcite peak in the XRD pattern. Note that the imperfections of these recrystallized calcites are probably due to rapid growth during the transformation reaction.

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Fig. 2. Isotope and mineralogy results from transition experiments at 25 °C after 265 h duration, plotted relative to the concentration of added salts. (A) Calcite content, (B) d13C, (C) d18O, (D) D47. Error bars represent uncertainites (2r) in post-experiment measurements. Shaded blue areas represent the uncertainty (2r) of the initial value. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

of an in-home standard run with samples in an analytical sequence. 3. RESULTS The pH ranged from 7.88 to 9.72 for all final experimental solutions, indicating different degrees of carbonate saturation (Table1). The salinity of the solutions ranged from 0.78 to 144 (TDS, g/l), generally corresponding to the concentrations of the salt additives. The oxygen isotope compositions in the final solutions of all treatments showed

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Fig. 3. Isotope and mineralogy results from transition experiments at 90 °C for different durations, plotted relative to the concentration of added salts. (A) Calcite content, (B) d13C, (C) d18O, (D) D47. Error bars represent uncertainties (2r) in post-experiment measurements. Shaded blue areas represent the uncertainty (2r) of the initial value. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

no significant differences from the initial compositions (Table 1). The SEM photographs of selected solid samples showed significant modifications in mineral structure with rod-like aragonite transforming to rhombic calcite crystal. This corresponded with the observed XRD patterns (Fig. 1), thereby verifying the carbonate polymorphic transition. The XRD analysis revealed various amounts of calcite, ranging from 4% to 93%. Calcite content showed no significant change after a duration of 265 h in the low-

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Fig. 4. Calcite content and d18O results plotted relative to duration time for high-temperature (90 °C) transition experiments with different treatments. (A) Calcite content; (B) d18O. Error bars represent uncertainties (2r) in post-experiment measurements. Shaded blue areas represent the uncertainty (2r) of the initial value. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

temperature (25 °C) experiments (Fig. 2A), but was significantly higher in 90 °C experiments with the addition of Na+ or Ca2+ (Fig. 3A). However, addition of Mg2+ to the solutions resulted in no change from the initial mineralogy (Fig. 3A). Treatments with additional Na+ or Ca2+ resulted in a final calcite content of >60% after 23 h, indicating higher transition rates compared with those of the deionized water treatments (Fig. 4A). Plots of the isotopic values against the solution salt concentrations for the 25 °C experiments (Fig. 2) showed no significant difference from those of the initial aragonite. However, the d18O values of 90 °C experiments displayed a considerable change from the initial value throughout the duration of the experiments and were inversely correlated with the change in mineralogy (Fig. 4B). The effect of different combinations of salts was shown in Fig. 5, where the addition of both Na+ and Ca2+ significantly influenced the rate of the polymorphic transition and

Fig. 5. Isotope results plotted relative to mineralogy for hightemperature (90 °C) transition experiments with different combinations of added salts. (A) d13C, (B) d18O, C) D47. Error bars represent uncertainties (2r) in post-experiment measurements. Shaded blue areas represent the uncertainty (2r) of the initial value. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

accompanying d18O values. Addition of Mg2+ resulted in no significant reprecipitation of calcite, but the d18O values were still depleted from the initial value by 1‰ (Fig. 5B). The D47 and d13C values of samples after all experimental treatments were close to the initial values within their analytical precisions, respectively (Table 1). 4. DISCUSSION 4.1. Controls on the aragonite-to-calcite transition Aragonite is thermodynamically metastable at Earth’s surface. It is prone to diagenetic alteration and is usually replaced by calcite (Swart, 2015). As this polymorphic transition may influence the geochemical information used for paleo-environment research, a better understanding of the mechanisms behind the aragonite-to-calcite transition is required. Based on the published studies, it is generally agreed that the transition is best described by a dissolution–reprecipitation mechanism, where aragonite dissolves and calcite reprecipitates from the solution (Ogino et al., 1987; Finch and Allison, 2003; Perdikouri et al., 2011; Casella et al., 2017). A further consideration is the role of various dissolved salts in the solution, as they

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can influence the transition. For example, the presence of Mg2+ retards the rate of transition of aragonite to calcite, whereas both Ca2+ and Na+ are known to increase the rate (Taft, 1967; Bischoff and Fyfe, 1968; Bischoff, 1969; Berner, 1975; Morse et al., 1980). In this study, we investigated the role of temperature as a primary factor in determining the transition rate and found it can magnify the effect induced by salt additions. Within the precision of the XRD method (5%), no difference in calcite contents could be determined for the 25 °C experiments, even with the addition of Na+ or Ca2+ to the solutions (Fig. 2A). However, the 90 °C experiments resulted in high final calcite contents, suggesting increased transition rate of aragonite to calcite (Fig. 3A). This dependence of transition rate on temperature is clearly evident from our results and is consistent with previous work (Taft, 1967; Bischoff, 1969). However, the influence of higher temperature in the transition rate can be suppressed by the presence of Mg2+, even in combination with Na+ and Ca2+ (Fig. 5), resulting in little or no extra calcite in solid samples. This is because the Mg2+ adsorbs onto the crystal surface and blocks the crystal site to prevent any further calcite formation (Berner, 1975; Mucci and Morse, 1983). As the basic lattice-building ion within the carbonate crystal, Ca2+ plays an important role in boosting the degree of supersaturation with respect to both aragonite and calcite within solutions. As the solubility of aragonite is normally higher than that of calcite (Plummer and Busenberg, 1982), a pure solution will be undersaturated with respect to aragonite prior to any calcite dissolution. As dissolution occurs, Ca2+ and DIC ions dissolved from the aragonite will result in the solution becoming increasingly saturated with respect to calcite, resulting in calcite reprecipitation. This initiates the transition reaction, the rate of which increases with increased Ca2+ concentration as a result of the common ion effect principle (Harris, 2007). This is confirmed by our experimental results where the addition of Ca2+ in solutions always led to faster rates of calcite reprecipitation than found for the deionized water treatments for identical transition times (Figs. 3A and 4A). Our results indicate that Na+ accelerated the transition rate to a similar extent as Ca2+. However, a combination of both Na+ and Ca2+ does not appear to further enhance the transition rate (Fig. 5). This could possibly be attributed to analytical error, but is more plausibly an indication of a competitive relationship between these two ions. Although Na+ is similar in size to Ca2+, the two ions have different charges, and it remains unclear whether Na+ can readily substitute for Ca2+ in carbonate (Mitsuguchi et al., 2010). Moreover, Na+ enhances the transition rate but does not have a comparable influence to Ca2+ upon the Sr/Ca ratio of solid samples (Fig. S5). This suggests that Na+ does not necessarily competitively replace divalent ions in the calcite crystal lattice but likely acts as a medium to facilitate the calcite crystallization during the transition. 4.2. Interpretation of carbon isotope variations As the aragonite-to-calcite transition reaction produces calcite reprecipitated from a pool of DIC dissolved from

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the original aragonite, d13C variations can provide insights into the potential kinetic effects on isotopic fractionation in 2 a calcite–DIC (mainly HCO 3 and CO3 ) system. Carbon isotope fractionation between calcium carbonate and gaseous CO2 is known to be temperature dependent (Emrich et al., 1970; Mook et al., 1974; Romanek et al., 1992). However, in our transition experiment, no CO2 gas exchange occurred, as the experiment was carried out in a closed system. We can therefore rule out the possible influence on the d13C due to isotopic exchange with ambient CO2 gas or CO2 degassing. Rather, the primary control on the d13C results is isotopic fractionation between the DIC pool and the newly-formed calcite. If the calcite–DIC fractionation factor for carbon isotopes is constant or independent of temperature, the kinetic effect discriminating between heavy and light carbon isotopes will not be significant. Therefore the newly-formed calcite will have similar d13C values as the DIC or the precursor carbonate. Our results have supported this explanation, as the d13C values from all experiments were comparable with that of the original aragonite (Table 1), suggesting that no significant fractionation occurred after the transition. Furthermore, Calcite precipitation experiments under different conditions conducted by Levitt et al. (2018) and Romanek et al. (1992) have showed little to no temperature dependence for the carbon isotope fractionation between the DIC and the calcite. Similar results have been reported by Ritter et al. (2017) for the experimental diagenetic alteration of biogenic aragonite, as the d13C in their solid samples remained unaltered in all experiments. The authors suggested that the d13C in the DIC pool was dominated by the carbon contained in the solid phase, which was incorporated into the reprecipitated calcite without significant d13C fractionation. Consequently, our d13C results suggest that no additional sources of carbon for the DIC pool to exchange with leads to no change in the d13C of the DIC pool. Meanwhile, kinetic effects on d13C fractionation between the calcite and DIC are insignificant under the conditions of our transition experiments. 4.3. d18O disequilibrium during the transition To explore the d18O disequilibrium during the transition reaction, it’s necessary to understand the equilibrium conditions for oxygen isotope partitioning between the reaction phases. We used the calcite–water oxygen isotope calibration from Kim and O’Neil (1997) as the isotopic equilibration reference. This calibration determines the fractionation factor for a given temperature, enabling the prediction of the reprecipitated calcite d18O, provided the d18O value of the solution is also known. We assumed any mixing between the newly-formed calcite and the remaining aragonite during transition is linear and accounted for this mixing when predicting the d18O value of the mixed carbonates (Table 1). As the calibration from Kim and O’Neil (1997) was determined experimentally by slowly precipitating calcite, we assume that any further isotope kinetic fractionation during their calcite growth was negligible. Therefore, any deviation from the prediction can be understood as a kinetic effect or disequilibrium fractionation.

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Fig. 6. Deviations of d18O and D47 from the expected equilibrium values and their connection to the transition rate. In panels A and B, both measured and expected equilibrium d18O and D47 values are presented relative to mineralogy. Note that the equilibrium values are taken from Table 1 based on the carbonate d18O–T calibration of Kim and O’Neil (1997), D47–T calibration of Kluge et al. (2015), and mixing model of Defliese and Lohmann (2015). The difference between the two lines in panel B indicates a significant kinetic offset for d18O (T-test, pvalue = 0.009), but not for D47, for which no significant kinetic offset is observed (T-test, p-value = 0.88). In panels C and D, deviations of measured d18O and D47 from the expected equilibrium values are presented relative to the transition rate. The transition rates are estimated using the ratio of the calcite content to elapsed time. The yellow dashed line represents the mean value of treatments with added Ca2+ and the blue dashed line represents the mean value of treatments with added Na+. In panel B, 95% confidence interval (95%) of the fitting line is indicated by dashed lines. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

Based the above assumption, we explored the kinetic effects by plotting the predicted equilibrium values against the measured d18O values (Fig. 6A). The predicted d18O values decreased linearly with progress of the transition reaction. As the results of the 25 °C experiment showed no apparent calcite reprecipitation, only the results from the 90 °C experiment were considered. These results showed that the d18O values were significantly positive compared with the predicted values, indicating apparent kinetic effects associated with the transition. We also plotted the deviations between the measured and predicted d18O values against the transition rates to show the dependence of the d18O kinetic effect upon the transition reaction rate (Fig. 6B). The deviation increases with transition rate from 0 to 0.5%/h (dotted box in Fig. 6B) and remains constant after the rate exceeds a value of 0.5%/h. Note that the diminished kinetic offset observed at the lowest conversion rates does not necessarily indicate approach to an equilibrium state, as there is no significant calcite reprecipitation in the experiments. A positive kinetic offset was observed in experiments with additions of Ca2+ compared with those with additions of Na+ when the transition rate exceeds the value of 0.5%/h (Fig. 6B). As the magnitude of the Na+/ Ca2+-induced kinetic effect was lower than that induced

from the transition rate, we suggest that the disequilibrium during the transition reaction is controlled mainly by the transition rate with a further influence from the addition of dissolved salts (especially Ca2+), which is discussed in detail below. 4.3.1. Dependence of the d18O kinetic effect on transition rate The d18O kinetic effect due to the high calcite growth rate has been documented in numerous studies of T–d18O calibration (Kim and O’Neil, 1997; Coplen, 2007; Kim et al., 2007b; Dietzel et al., 2009; Gabitov et al., 2012; Watkins et al., 2014). In most cases, higher growth rates lead to lighter d18O values in calcite precipitated from a DIC pool where the d18O is either at equilibrium or at disequilibrium. This growth rate dependence in d18O kinetic effects can derived from two scenarios (Devriendt et al., 2017): (1) kinetic fractionation between DIC species and water, and (2) kinetic fractionation between calcite and DIC species. In our study, calcite was precipitated from the solution in which the DIC species were dissolved from the original aragonite. If higher growth rates result in insufficient time for isotopic exchange between the DIC and water (inset in Fig. 6B), the resulting d18O in the calcite or the mixed

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carbonate will deviate from the predicted equilibrium value. Specifically, given the limited amount of time, DIC with light isotopes will preferentially exchange with the water (McConnaughey, 1989), leading to the remaining DIC becoming enriched in heavy isotopes, ultimately resulting in higher d18O in the final calcite after the transition. As the 18O exchange within the DIC pool occurs through CO2(aq) hydration/dehydration (Affek, 2013), the lower CO2 (aq) concentration in a closed system also likely limits the 18O equilibrium between water and both HCO 3 and CO2 3 . Higher growth rate during the transition leads to higher difference in d18O values between the insufficient exchange DIC pool and the predicted final equilibrium state. The pH variation during the transition can be considered as another potential effect to account for the d18O variation in this study. It is noted that there is a significant difference in d18O equilibrium values between the DIC spe2 cies (HCO 3 and CO3 ) for a given temperature (Usdowski and Hoefs, 1993; Beck et al., 2005). Higher pH values (10) corresponding to higher CO2 3 proportions in the DIC will result in depleted d18O in the DIC pool and then in the final precipitated carbonates (Zeebe and Wolf-Gladrow, 2001). The pH values of final solutions in this study are generally higher than those of Kim and O’Neil (1997), which were

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designed to precipitate carbonate at equilibrium. Therefore, this pH effect will lead to lower than expected d18O in carbonates after the transition, which will be more significant in high pH solutions. However, since most d18O values we observed are higher than expected (Fig. 6B) opposite to the pH effect, the net disequilibrium fractionation is still dominated by partial isotopic exchange between the DIC and water due to the high transition rate. With regard to the kinetic fractionation between calcite and DIC species, we consider that the observed kinetic effects may arise not only during the isotopic exchange between DIC and solution but also during the transfer of DIC from solution to the growing crystal surface (Watkins et al., 2014; Devriendt et al., 2017). Therefore, the disequilibrium within the DIC pool potentially compounds additional kinetic effects related to processes controlling the isotope fractionation during crystal growth from the surface to lattice incorporation sites (Watkins et al., 2014; Devriendt et al., 2017). A diffusion process has been suggested to control the preferential incorporation of light isotopes into the inner part of the crystal (Turner, 1982). Preferential deprotonation of 16O-enriched HCO 3 and incorporation of 16O-enriched CO2 have been pro3 posed to result in lower d18O in carbonate minerals (Kim et al., 2006). It has also been found that ion entrapment

Fig. 7. Schematic of aragonite-to-calcite transition region illustrating hypothesized processes to account for the observed isotopic variations. Based on the dissolution–reprecipitation mechanism (Ogino et al., 1987; Finch and Allison, 2003; Perdikouri et al., 2011; Casella et al., 2017), aragonite starts to dissolve when the ion product (Q) is lower than the aragonite solubility product (Ksp,A). The reprecipitation of calcite starts 2 when Q is higher than the calcite solubility product (Ksp,C). DIC species (mainly HCO 3 and CO3 ) undergo isotopic exchange with the solution via CO2 hydration/dehydration or dehydroxylation/dehydroxylation. In the DIC pool, 18O fractionation is dependent on the molecule exchange rate for reactions involving oxygen atoms; clumped isotope fractionation is dependent on the 13C–18O reordering rate within each DIC species via self-exchange reactions. In a closed system where no other source of CO2 is introduced and the initial DIC in the solution is low, 13C fractionation in the dissolved DIC is insignificant. A disequilibrium scenario for both d18O and D47 can be expected, given that the precipitation rate (Rreprecipitaion) is higher than the molecule exchange rate (Rexchange) and the 13C–18O bond reordering rate (Rreordering) is lower than the Rexchange. Note that further kinetic fractionation will potentially take place in the solution–mineral boundary layer through several potential processes (e.g., ion detachment or attachment, mass-dependent diffusion, and surface chemistry and properties), although its direction and magnitude are currently unknown.

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at the crystal surface can potentially lead to a lower isotope composition throughout the whole crystal lattice in the case of high growth rates (Watson, 2004; Watson and Mu¨ller, 2009), and ion attachment at the crystal surface can result in higher concentrations of the light isotopes within fastgrowing minerals (DePaolo, 2011). Note that in this study, the specific calcite growth rate (lmol m2 h1) was not explicitly determined but only roughly approximated by the transition rate (%/h). This makes it difficult to estimate the magnitude of kinetic effects based on published models. However, these findings suggest that these type of kinetic effects can potentially result in depleted d18O in the calcite, opposite to the d18O enrichment in reprecipiated carbonates due to the incomplete isotopic exchange between DIC and water in the case of our transition experiments. Therefore, we suggest that the high transition rate has a significant control on the overall kinetic isotope effects. It leads to different degrees of partial exchange of 18O between the DIC species and water, which is responsible for the observed d18O disequilibrium. 4.3.2. Effect of salt addition on d18O kinetic effect As well as the transition rate, addition of dissolved salts plays important role in d18O disequilibrium. As shown in Fig. 6B, there is a significant positive kinetic offset evident in experiments with the added Ca2+ relative to those with the added Na+, even when transition rates are similar. The offset is also apparent for Na+ experiments relative to deionized water experiments. However, d18O values from the deionized water treatments appear to approach those of Na+ experiments as the transition rate increases (inset in Fig. 7B). A deeper exploration of the kinetic offset evident in the Ca2+- and Na+- added experiments follows. The influence of dissolved salts on oxygen isotope fractionation in aqueous solution has long been recognized, particularly in measurements using the CO2AH2O equilibration technique (Taube, 1954; Sofer and Gat, 1972; Truesdell, 1974; Horita et al., 1993b; Driesner and Seward, 2000; Le´cuyer et al., 2009; Kim et al., 2012). This so-called oxygen isotope salt effect is believed to result from preferential bonding of heavy or light isotopes in a hydration sphere of dissolved ions and localized changes to the structure of the water around the solute. Truesdell (1974) introduced the concept of ‘structure region’ of water to explain the salt effect on the oxygen isotope composition of water, where heavy isotopes are concentrated in the structure region, and light isotopes in the structurebroken regions. The balance between these two different hydration regions determines the ultimate isotopic fractionation regime between the water in the hydration sphere and the remaining bulk solution (O’Neil and Truesdell, 1991). In the present study, this isotope salt effect was observed during calcite reprecipitation from a DIC pool in which different salts were introduced. We therefore discuss the salt effect in the context of the relationship between the ionic hydration and the rate of oxygen isotope exchange between DIC and water. The oxygen isotope salt effect has been observed in several previous studies for solutions containing NaCl or CaCl2 (Truesdell, 1974; O’Neil and Truesdell, 1991;

Horita et al., 1993a; Horita et al., 1993b; Le´cuyer et al., 2009; Kim et al., 2012). For example, the addition of Ca2+ to the solution leads to enrichment of 18O in the strongly-bound hydration sphere and depletion of 18O in the remaining bulk solution or unhydrated water (Truesdell, 1974; O’Neil and Truesdell, 1991; Driesner and Seward, 2000). No significant oxygen isotope salt effect has been observed with the addition of Na+ (Horita et al., 1993b; Kim et al., 2012), except at higher salinities (up to 250 g/L), where a significant oxygen isotope salt effect has been observed (Le´cuyer et al., 2009). Although some discrepancies exist among these various works regarding the magnitude of d18O fractionation, the influence of both temperature and salt concentration has been confirmed. The exact fractionation value specific to salt addition is not the focus in this study. Rather, the major concern is relative fractionation difference between the two different salt additions, and we tentatively estimate the salt effects based on existing studies in which the temperature and salt composition are similar to those of our experiment. In the case of Truesdell (1974), the d18O fractionation difference between Ca2+ and Na+ treatments was determined to be <0.2‰ at 90 °C with salt concentrations in the range of 37.2 to 80 g/L for Ca2+ and 23 to 92 g/L for Na+. According to Horita et al. (1993b), the fractionation difference was likely to be <0.4‰ at 100 °C with salt concentrations in the range of 20 to 40 g/L for Ca2+ and 11.5 to 92 g/L for Na+. These estimated salt effect differences between Ca2+ and Na+ are too small to account for the observed offset in Fig. 6B. Furthermore, our 18Osolution results showed no significant difference between the salt solutions and the initial pure water (Table 1). This suggests that the observed salt effect is not detectable via CO2AH2O equilibration under our experiment conditions but is sensitively captured via carbonate precipitation. If the reprecipitation of calcite relied completely on the unhydrated water, the salt effect (or the salt-induced kinetic offset) would not be observed, suggesting an important role for the hydration sphere. The possible explanation of how the isotopic fractionation takes place in the hydration sphere is that the 18 O from the hydration sphere surrounding the Ca2+ ions has limited time to exchange with the surrounding DIC as the calcite precipitates. This scenario was first suggested by Kluge and John (2015) where they explained that the 18 O can be preferentially incorporated into the mineral as the water is stripped off the hydration sphere. An alternative explanation is that the rate of 18O exchange between the bulk water and the DIC is somehow slower when additional Ca2+ is presented. However, the reason for this remains unclear. 4.4. Disequilibrium of clumped isotopes during the transition During the aragonite-to-calcite transition, it can be expected that some chemical bonds will break and/or form in DIC species derived from aragonite dissolution. This can lead to reordering of the 13CA18O bonds with the new solution environment. Given that the clumped isotope signature is implicitly assumed to be dependent solely on the temperature of mineral formation (Ghosh et al., 2006a; Schauble

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et al., 2006; Eiler, 2011), the 13CA18O bond reordering will proceed towards a new internal thermodynamic equilibrium. As such, the equilibrated D47 value in newly-formed calcite can be predicted for a known temperature by a D47–T calibration, determined either experimentally (Ghosh et al., 2006a; Dennis and Schrag, 2010; Zaarur et al., 2013; Kluge et al., 2015; Kelson et al., 2017) or empirically (Tripati et al., 2010; Henkes et al., 2013; Wacker et al., 2014; Kele et al., 2015). In this study, we predicted the D47 value using the calibration from Kluge et al. (2015) for two reasons: (1) the temperature they used for carbonate precipitation is between 23 °C and 250 °C, which encompasses the temperature range used in this study; and (2) their analytical procedures for sample purification are similar to those used in this study. Although different D47 predictions can be obtained using the different calibrations due to unsolved issues of inter-laboratory inconsistencies (Dennis et al., 2011; Kelson et al., 2017), the main signature of the isotope kinetic isotope effect is likely independent of the choice of calibration. Using a recent published calibration of Bonifacie et al. (2017), the patterns of D47 kinetic fractionation are similar to those based on the Kluge calibration except a slight difference (0.03‰) in magnitude (Fig. S6), which does not affect the conclusion in this study. In the same manner as exploring the d18O disequilibrium, we interpreted the deviation of measured D47 from the predicted values as a clumped isotope kinetic effect (Fig. 6C). The most remarkable feature is that the D47 values of all treatments are similar to the value of the original aragonite. The mean D47 value of 0.718‰ for these samples is within the range of typical fast-growing aragonitic surface corals (Saenger et al., 2012). Their standard deviation of 0.011‰ (1r) is comparable to the long-term reproducibility of BACS (see Section 2.4.5), which is the carbonate standard regularly measured along with these samples in this study. This indicates that any variation in D47 that occurs during the transition reaction can be attributed to the analytical uncertainty. In addition, the significant correlation between the D47 kinetic effect and transition rate (Fig. 6D) suggests that the calcite growth rate surpasses the limit required for 13CA18O bonds in the DIC to reorder towards equilibrium. Similar to the d18O disequilibrium in Fig. 6B, the D47 kinetic effect remains constant when the transition rate exceeds the value of 0.5%/h but there is no additional kinetic fractionation difference between the experiments using Ca2+ and Na+ solution (Fig. 6D). It suggests that the highest kinetic fractionation is controlled mainly by the high transition rate and is limited by the difference in D47 values between the equilibrated prediction and the original aragonite. The disequilibrium in carbonate D47 can be used to infer the d18O disequilibrium within the DIC pool (Kluge and Affek, 2012; Kluge et al., 2014). As such, the link between the D47 and d18O in the solution system is crucial for understanding the isotopic equilibrium state in carbonate minerals. Affek (2013) observed a kinetic similarity between D47 and d18O in the CO2AH2O isotope exchange reaction. Clog et al. (2015) performed similar experiments and also observed the similar exchange rates for both D47 and d18O. These studies have suggested that oxygen isotope

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exchange between CO2(g) and water can also lead to reordering of 13CA18O bonds in CO2(g). However, this similarity may be due to the rate limitation during the CO2 dissolution (Clog et al., 2015; Staudigel and Swart, 2018). In other words, if the exchange rate of CO2 with the water is fast enough, it is possible to see the kinetic difference between D47 and d18O. Furthermore, in a DIC pool dominated HCO–3 and CO2– 3 , it is unknown if and how this kinetic similarity still holds. The present results showed significant variations in d18O values, whereas the clumped isotope reordering was observed to be sluggish. This raises questions regarding the rate difference between the isotope clumping within the isotopologues and 18O exchange within the DIC pool, especially in the case of rapid calcite precipitation. There are several possible reasons for the observed D47 variations: (1) analytical artifacts from contamination of the original carbonate material, (2) non-linear mixing effects due to mixing of carbonates with different compositions, and (3) solution chemistry (e.g., salinity and pH). 4.4.1. Potential sources of D47 variations The first possible explanation for the deviation between predicted and actual D47 values is contamination of original coral aragonite. It is possible that Porites coral preserves some organic matter in the skeleton after deposition (Wang et al., 2015; Erler et al., 2016). As the coral material was homogenized prior to the experiments, any organic matter would be also homogenized, which would lead to consistent contamination in each experiment aliquot. After the gas purification, this organic matter would break down in the ion source of the mass spectrometer, generating mass 47 fragments, which results in higher than expected D47 (Eiler and Schauble, 2004). According to the T–D47 equilibrium calibration of Kluge et al. (2015), if all aragonite is reprecipitated as calcite without kinetic effects, the expected D47 value of the final calcite sample will be 0.547‰ at 90 °C. Compared with the initial value of 0.723‰, this large enrichment seems unlikely to be the result of organic contamination. As the expected D47 value decreases as calcite content increases (see Fig. 5C), the resulting D47 values due to contamination will not remain constant. Furthermore, previous work using coral sample without oxidative cleaning by H2O2 showed no significant alteration in the clumped isotope composition (Saenger et al., 2012; Spooner et al., 2016). Overall, we conclude that the organic contamination is not a significant cause of the D47 variations observed in the present study. As the D47 calculation is in part dependent on the d13C and d18O values, mixing of phases with the same D47 values but different bulk isotope compositions will result in a mixed phase with a D47 value different to the weighted sum of the phases’ D47 values (Eiler and Schauble, 2004; Thiagarajan et al., 2011; Defliese and Lohmann, 2015; Kimball et al., 2016). Different degrees of completion of the aragonite-to-calcite in aqueous solution will likely produce larger variations in d13C and d18O in the final mixed carbonate, with mixing effects potentially affecting clumped isotope compositions. In this study, we observed no significant variation in d13C after different transition experiments but there were large variations in d18O. We quantified the

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maximum mixing effect based on predicted d18O and D47 in the reprecipitated calcite at 90 °C and the measured d18O and D47 in the original aragonite. Using the calcite–water isotope calibration of Kim and O’Neil (1997) and the measured d18Osolution values, we calculated the predicted d18O in the reprecipitated calcite to be 19.6‰ to 18.4‰ (VPDB) for experiments at 90 °C. The non-linear mixing model of Defliese and Lohmann (2015) gives a mixing effect of <0.002‰ for D47 in the case of 50% of each end-member. As this value is much lower than the analytical precision, we rule out this possibility to account for the clumped isotope results. Several studies have discussed the influence of solution salinity and pH on D47 during carbonate precipitation (Thiagarajan et al., 2011; Saenger et al., 2012; Tang et al., 2014; Tripati et al., 2015; Watkins and Hunt, 2015; Kelson et al., 2017). The clumped isotope results of this study could be attributed to these effects, given the significant variation in salinity and pH in this study. Salinity effects are likely related to the concentration of different saline ions and have been investigated previously (Kluge and John, 2015; Kluge et al., 2015). It has been suggested that high concentrations of NaCl (375 g/L) and moderate concentrations of MgCl2(H2O)6 (150 g/L) within the solutions have little impact on D47 (Kluge and John, 2015). Addition of CaCl2 (200 g/L) to the solution resulted in a positive correlation between CaCl2 concentration and D47. This correlation is likely independent of temperature (Kluge and John, 2015), with the observed D47 offset being +0.03‰ per 40 g Ca2+ in 1L of solution. However, considering the highest Ca2+ concentration in this study (2 g/ L, Table 1) and using the correlation of Kluge and John (2015), the corresponding magnitude of this salt effect is too small to account for our D47 offsets. Furthermore, we found no significant correlation when plotting deviations between the measured and predicted D47 values against the salinity (Fig. S7A7), which also rules out the potential salinity effect on D47. The effect of pH on D47 results from its influence on DIC speciation, and it is also commonly referred to as the DIC speciation effect (Tripati et al., 2015). Generally, HCO 3 is the dominant species for pH values of 6–10 and CO2 3 dominates at higher pH, although the speciation is also slightly affected by salinity and temperature (Zeebe and WolfGladrow, 2001). At equilibrium, these species exhibit different degrees of 13CA18O bond ordering (quantified as D63) (Hill et al., 2014; Tripati et al., 2015). The theoretically determined D63(HCO is significantly higher than 3) D63(CO2 3 ), with a difference of 0.0255‰ at 90 °C (Hill et al., 2014). During calcite precipitation, DIC undergoes deprotonation to form calcite and 13CA18O bonds are reordered toward equilibrium throughout the crystal lattice. However, rapid crystal growth limit the time available for the clumped isotope equilibration, leading to a disequilibrium signal in part or all of the clumped isotope signature of the DIC species as they are incorporated into the calcite lattice (Hill et al., 2014). As such, calcite growing from solution under different pH conditions will exhibit D47 variations if the growth rate is sufficiently rapid. In the present study, the pH values are in the range of 7.88 to 9.72 leading

to large variation in the proportion of HCO 3 relative to CO2 3 in the DIC pool. Moreover, these pH values are generally higher than those from Kluge and John (2015) where they established their D47–T calibration (Kluge et al., 2015) used in our study. If the DIC pool is isotopically at equilibrium and the calcite growth during the transition outpaces the ability of the 13CA18O bonds to completely reorder at equilibrium throughout the crystal lattice, samples reprecipitated under higher pH will exhibit depletion in D47 compared with equilibrium predictions (Hill et al., 2014; Tripati et al., 2015). Assuming the transition rate is proportional to the calcite growth rate for each experiment, the pH effect can be roughly tested by plotting deviations between the measured and predicted D47 values against pH. Note that the pH of the solution was not strictly monitored during each experiment. Even so, we observed a weak negative correlation (Fig. S7B). However, the D47 values we observed are mainly higher than expected (Fig. 6D) opposite to the expected effect of pH on D47 in DIC pool. The DIC pool is still likely out of equilibrium as there is not sufficient time for it to reach equilibrium. The pH effect on D47 in DIC is thus not a likely cause of the observed D47 offset in our experiments. 4.4.2. Kinetic difference between 13CA18O reordering and 18O exchange within the DIC pool As none of the existing mechanisms outlined above are able to explain the D47 variations in the present results, another explanation is explored as follows. As shown in Fig. 7, several hypothesized processes explain why the clumped isotope signature remained unchanged in the case of 18O isotope exchange within the DIC pool. Firstly, the fractionation of D47 and d18O in the transition environment is considered to occur within three phases: solution, mineral, and solution–mineral boundary layer. In the solution, the isotope exchange within the DIC pool is achieved through CO2 (aq) hydration/dehydration, as well as hydroxylation/dehydroxylation under high-pH conditions (McConnaughey, 1989; Affek, 2013). In the mineral phase, 13 C–18O bond ordering can be influenced by solid-state exchange of oxygen and carbon isotopes within the mineral lattice (Dennis and Schrag, 2010; Passey and Henkes, 2012). However, given the duration (265 h) and temperature (90 °C) of the present experiments, this solid-state reordering effect is unlikely to be significant (Dennis and Schrag, 2010; Passey and Henkes, 2012; Brenner et al., 2018). In the solution–mineral boundary layer, massdependent fractionation of oxygen isotopes has been widely recognized (Watson and Mu¨ller, 2009) and several surface kinetic models have been proposed to predict the conditions under which isotopes attain equilibrium or disequilibrium (Watson, 2004; DePaolo, 2011; Gabitov et al., 2012; Watkins et al., 2013; Watkins et al., 2014; Tripati et al., 2015; Watkins and Hunt, 2015). It is presently unknown how these models can explain the D47 and d18O signatures in our transition experiments, but it is worth considering the potential contribution from the boundary layer to the overall kinetic effects on both D47 and d18O. It is most likely that isotope exchange between the DIC species and the solution plays an important role. If the rate

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Fig. 8. Comparison of model-based equilibration paths with measured values for d18O and D47. Measured isotopic values from all 90 °C experiments are plotted as red circles with ± 1r error bars. Vector field (blue lines) and equilibrium paths (green and black lines) are generated by the model of Staudigel and Swart (2018), given the isotopic compositions of the initial aragonite (red dot) and the final calcite (black circle) along with two different K1318 12 values (0.982 or 1). The equilibrium d O composition is calculated by assuming a constant d18Osolution value of 5.8‰ (relative to VSMOW) which is averaged from all d18Osolution measuring values in this study and the calcite–water isotope calibration of Kim and O’Neil (1997) at 90 °C. The equilibrium D47 composition is calculated by the D47–T calibration of Kluge et al. (2015) at 90 °C. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

of molecule exchange reactions (e.g., CO2 hydration or HCO dehydration) is sufficiently rapid that 18O 3 approaches a new thermodynamic equilibrium, we would expect no kinetic fractionation in d18O of the solution phase, as 18O has reached its mass balance in the DIC– H2O system. However, this may not be the case for the clumped isotopes. As in the DIC species, 13CA18O bond reordering reflects an internal redistribution or rearrange2 ment of isotopes within the HCO ions, which 3 or CO3 is likely affected by oxygen isotope replacement between water and DIC species. If the 18O bound to 13C shows a different exchange rate compared with the one bound to 12C it will be reasonable to expect a rate difference between the two isotope systems (Staudigel and Swart, 2018). Although the timescales for reaching CO2–water equilibrium are similar for both D47 and d18O (Affek, 2013; Clog et al., 2015), it will be more important for us to know whether the rate of 13 CA18O bond reordering is different from the rate of 18O exchange within a DIC pool which is dominated by the 2 HCO 3 or CO3 species. A recent study of Staudigel and Swart (2018) investigated the rates of DIC equilibration for the two isotope systems using barium carbonate precipitation. They observed a decoupled behavior for D47 vs. d18O during the equilibration and developed a model based on their experimental results. Previous studies from speleothem carbonates have identified a higher disequilibrium

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sensitivity in D47 relative to d18O during calcite precipitation (Kluge and Affek, 2012; Affek and Zaarur, 2014; Affek et al., 2014; Kluge et al., 2014). These results indicate that the rate difference, as mentioned above, commonly exists between the two isotope systems. As suggested by Staudigel and Swart (2018), the difference in equilibration rates can be explained by the fact that the exchange rate in DIC species bound to 12C is slightly higher than those bound to 13C. This kinetic difference (K13-12) has been experimentally determined to be a constant, measured as 0.982, which is generally insensitive to temperature (Staudigel and Swart, 2018). We therefore applied the Staudigel and Swart (2018)’s model using the initial aragonite and 90 °C equilibrated calcite compositions from our study to simulate possible equilibration paths. As shown in Fig. 8, assuming a K13-12 value of 1 results in a linear path suggesting no decoupled behavior for both isotopes. However, using the K13-12 value of 0.982 results in an equilibration path, which generally agrees with measured values of 90 °C experiments in this study. The model can also predict the direction and magnitude of the rate change of D47 relative to d18O for any initial point within Fig. 8 (denoted as the direction and the length of blue lines, respectively). The agreement between the measured values and model predication suggests that the 13CA18O bond reordering is the slow step in the molecule exchange reaction between DIC and water in our transition experiment. As such, molecule exchange in the DIC pool leads to partial disequilibrium for d18O and complete disequilibrium for D47 during the aragonite-to-calcite transition. In a scenario where the rate of calcite precipitation outpaces molecule exchange and there is no further isotope fractionation in the surface–mineral boundary layer, we would expect the reprecipitated calcite to inherit the disequilibrium signals (D47 and d18O) from the DIC pool. If the rate of calcite precipitation is sufficiently low that the DIC pool reaches equilibrium with respect to 18O, but not for complete reordering of 13 CA18O bonds, then we may expect equilibrated d18O but partial disequilibrated D47 signals in the reprecipitated calcite. In this study, both D47 and d18O results directly support the first scenario. The second scenario, however, could be tested in future work if the transition experiment can be controlled at lower rates, possibly by adjusting the solution chemistry or the ratio of water to solid phases. Such data will be also critical to further verify the reliability of using oxygen isotope equilibrium to infer equilibrium of the clumped isotopes, which is the common assumption in T– D47 calibration studies (Ghosh et al., 2006a; Zaarur et al., 2013). 5. CONCLUSIONS In this study, aragonitic carbonate was experimentally transformed into calcite in aqueous solutions with different salt additions at two temperatures (25 °C and 90 °C). Mineralogical and isotopic analyses were conducted for the carbonates, which had undergone various degrees of the transition. The main conclusions are as follows.

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(1) The presence of Mg2+ ions in solution hinders the aragonite-to-calcite transition, whereas both Na+ and Ca2+ enhance the transition reaction. These findings are consistent with previous studies. (2) Significant variation was observed in d18O values, indicating different degrees of isotope exchange within the DIC pool. The calcite reprecipitated from the DIC pool at high transition rates recorded this partial equilibrated state. This disequilibrium effect can be exaggerated by the addition of Ca2+ to the solution, and is likely related to the hydration sphere of the bivalent cations in solution. (3) Under the experimental conditions of this study, both D47 and d13C remain constant during the transition reaction, suggesting that the temperature signatures recorded by D47 and the carbon isotope composition in the original aragonite are faithfully preserved in the reprecipitated calcite. (4) Kinetic effects influence both the 13CA18O bond reordering and molecule exchange within the DIC pool during the aragonite-to-calcite transition. The main control factor is the high transition rate, which leads to the disequilibrium for both D47 and d18O in the DIC pool prior to reprecipitation. The key finding is the different equilibration rates in the two isotopic systems during the transition toward equilibrium in which the D47 is completely out of equilibrium while the d18O can be partly equilibrated. The difference in the equilibration rate of D47 relative to d18O is also not constant on the path towards complete equilibrium. This likely provides important insights to studies relied on an assumption of constant disequilibrium state for both D47 and d18O in carbonates.

ACKNOWLEDGEMENTS This work is supported by the National Key Research and Development Program of China (2016YFA0601204), the National Natural Science Foundation of China (41722301, 41673115, and 41173004), the State Key Laboratory of Isotope Geochemistry (SKLIG-RC14-02), the Guangzhou Institute of Geochemistry, Chinese Academy of Sciences (135PY201605) and the Technology Research and Development Project of the Chinese Academy of Sciences (yg2010021). We are grateful to discussions with John Eiler, Hagit Affek, Tobias Kluge, Brad Rosenheim, Jens Fiebig, and Xu Wang about analytical methods involving clumped isotope of CO2. We thank Edwin Schauble, Philip Staudigel and three anonymous reviewers who provided constructive feedback that improved this manuscript appreciably. We wish to thank Ti Zeng for the field work and providing the coral samples used in this study, and Jingming Wei for her help with XRD analysis. The English of the manuscript was improved by Stallard Scientific Editing. This is contribution No. IS-2635 from GIGCAS. The data for this paper are available as the tables in the main text and the online supplementary materials.

APPENDIX A. SUPPLEMENTARY MATERIAL Supplementary data to this article can be found online at https://doi.org/10.1016/j.gca.2019.01.012.

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