Evaluation of boron isotope ratio as a pH proxy in the deep sea coral Desmophyllum dianthus: Evidence of physiological pH adjustment

Earth and Planetary Science Letters 349–350 (2012) 251–260

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Letter’s

Evaluation of boron isotope ratio as a pH proxy in the deep sea coral Desmophyllum dianthus: Evidence of physiological pH adjustment E. Anagnostou a,n, K.-F. Huang b, C.-F. You c, E.L. Sikes a, R.M. Sherrell a,d a

Institute of Marine and Coastal Sciences, Rutgers University, New Brunswick, NJ 08901, USA Department of Geology and Geophysics, Woods Hole Oceanographic Institution, MA 02543, USA c Department of Earth Sciences and Earth Dynamic System Research Center, National Cheng Kung University, Tainan, Taiwan d Department of Earth and Planetary Sciences, Rutgers University, Piscataway, NJ 08854, USA b

a r t i c l e i n f o

abstract

Article history: Received 1 December 2011 Received in revised form 22 June 2012 Accepted 5 July 2012 Editor: G. Henderson Available online 16 August 2012

The boron isotope ratio (d11B) of foraminifers and tropical corals has been proposed to record seawater pH. To test the veracity and practicality of this potential paleo-pH proxy in deep sea corals, samples of skeletal material from twelve archived modern Desmophyllum dianthus (D. dianthus) corals from a depth range of 274–1470 m in the Atlantic, Pacific, and Southern Oceans, ambient pH range 7.57–8.05, were analyzed for d11B. The d11B values for these corals, spanning a range from 23.56 to 27.88, are found to be related to seawater borate d11B by the linear regression: d11Bcoral ¼ (0.76 7 0.28) d11Bborate þ (14.67 74.19) (1 standard error (SE)). The D. dianthus d11B values are greater than those measured in tropical corals, and suggest substantial physiological modification of pH in the calcifying space by a value that is an inverse function of seawater pH. This mechanism partially compensates for the range of ocean pH and aragonite saturation at which this species grows, enhancing aragonite precipitation and suggesting an adaptation mechanism to low pH environments in intermediate and deep waters. Consistent with the findings of Trotter et al. (2011) for tropical surface corals, the data suggest an inverse correlation between the magnitude of a biologically driven pH offset recorded in the coral skeleton, and the seawater pH, described by the equation: DpH ¼ pH recorded by coral  seawater pH ¼  (0.75 7 0.12) pHw þ (6.88 7 0.93) (1 SE). Error analysis based on 95% confidence interval(CI) and the standard deviation of the regression residuals suggests that the uncertainty of seawater pH reconstructed from d11Bcoral is 7 0.07 to 0.12 pH units. This study demonstrates the applicability of d11B in D. dianthus to record ambient seawater pH and holds promise for reconstructing oceanic pH distribution and history using fossil corals. & 2012 Elsevier B.V. All rights reserved.

Keywords: deep sea coral pH boron isotopes calibration acidification

1. Introduction It is now well documented that atmospheric carbon dioxide (pCO2) varied with the growth and decay of polar ice sheets over the last 800 kyr (Luthi et al., 2008; Petit et al., 1999). In general, during periods of increasing atmospheric pCO2, surface ocean pH decreases ¨ as a result of CO2 uptake (Honisch and Hemming, 2005). The driving mechanisms that cause atmospheric pCO2 variations, and how they relate to the Earth’s climate system, are not well constrained. Since the Industrial Revolution, however, atmospheric CO2 has increased beyond the range of natural Pleistocene variations, displaying a sharp rise in the last  50 years (Keeling, 1961; Luthi et al., 2008; Petit et al., 1999) mainly as a result of fossil fuel burning (Le Que´re

n Corresponding author. Now at: OES, National Oceanography Centre Southampton, University of Southampton Waterfront Campus, Southampton SO14 3ZH, UK. Tel.: þ 44 23 8059 5075. E-mail address: [email protected] (E. Anagnostou).

0012-821X/$ - see front matter & 2012 Elsevier B.V. All rights reserved. http://dx.doi.org/10.1016/j.epsl.2012.07.006

et al., 2009). Such an increase has resulted in substantial decrease in oceanic pH (Sabine et al., 2004), which has raised much concern for the physiological and ecological response of marine calcifying organisms (Gattuso and Hansson, 2011; Hoegh-Guldberg et al., 2007). Insights into these issues can be gained by developing paleo-proxies for aspects of the ocean carbonate system. One that has received much recent focus is the boron isotope (d11B) proxy for seawater pH, as preserved in various biogenic carbonate archives ¨ (e.g. Hemming and Honisch, 2007). An important step in the development of the d11B proxy is its calibration against pH using specimens of modern biocalcifying organisms that grew at locations where seawater pH is well quantified (e.g. Rae et al., 2011). The biogenic carbonate d11B proxy has been explored and described in extensive detail elsewhere (Douville et al., 2010; ¨ Hemming and Hanson, 1992; Honisch and Hemming, 2005; Klochko et al., 2006; Rae et al., 2011; Sanyal et al., 1996; Spivack et al., 1993; Trotter et al., 2011; Zeebe et al., 2001, 2003). At typical modern seawater pH, boron (B) exists mostly

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in the forms of boric acid (B(OH)3) and borate ion (B(OH)4 ) with a distinct isotopic fractionation between the two species such that boric acid is enriched in the heavier isotope (11B) by 27.2% compared to borate ion (Klochko et al., 2006). The relative abundance of these two B species is pH dependent while total seawater B varies only on long time scales given its oceanic residence time of 10–14 million years (Lemarchand et al., 2000; Paris et al., 2010). Because the isotopic composition of total boron in seawater must be maintained as the relative abundance of the species changes with pH, the isotope ratio of each species must also change. Therefore the isotopic composition of each species for recent times in earth history depends entirely on pH. On the basis of the similarity of the boron isotopic composition of all marine carbonates to that of seawater borate, boron is believed to be incorporated into biogenic and inorganic carbonates overwhelmingly as the borate species (Hemming and Hanson, 1992). Incorporation in this form results in the d11B of marine carbonates tracking the d11B of borate ion and thus being sensitive to pH variations, with a typical  1% increase in d11B for a 0.1 unit increase in seawater pH. Boron isotope measurements in foraminiferal calcite and tropical coral aragonite have been used extensively as proxies of surface oceanic pH or pCO2 ¨ ¨ (Douville et al., 2010; Honisch and Hemming, 2005; Honisch et al., 2009; Liu et al., 2009; Palmer and Pearson, 2003; Pelejero et al., 2005; Sanyal and Bijma, 1999; Wei et al., 2009). However, there have been only a few efforts to reconstruct the evolution of ¨ intermediate to deep ocean pH using this proxy (Honisch et al., 2008; Rae et al., 2011; Sanyal et al., 1995; Yu et al., 2010). Successful demonstration of the utility of the boron isotope pH proxy in a new carbonate archive, Deep Sea Corals (DSC), would open a new window on deep ocean pH records. The boron isotope proxy applied to DSC offers several advantages compared to foraminifera: DSC are not only cosmopolitan in geographic distribution (Roberts et al., 2006), and of sufficient mass for precise U/Th dating (Cheng et al., 2000; Edwards et al., ¨ 2003; Goldstein et al., 2001; Schroder-Ritzrau et al., 2003), but they can also potentially provide single-individual snapshot records of climate-related oceanic variables with subdecadal resolution that only ice cores can rival (e.g. Adkins et al., 1998). In addition, most species are aragonitic and relatively rich in B (70–100 ppm; Douville et al., 2010), compared to 10–20 ppm for calcitic forams (Rae et al., 2011) allowing analyses of d11B in small samples with high precision. Deep sea corals do not harbor symbiotic algae, hence they are not subject to physiological processes related to zooxanthellae activity that can contribute to pH-sensitive vital effects in other corals (Al-Horani et al., 2003) and in foraminifera (Zeebe et al., 2003). Although the majority of DSC grow in waters that are oversaturated with respect to aragonite (Guinotte et al., 2006), some can survive in conditions unfavorable to aragonite precipitation (Oarag o1, where Oarag ¼ [Ca2 þ ][CO23  ]/K*arag, and K*arag is the stoichiometric solubility product for aragonite). Therefore it is suspected that deep corals have mechanisms to elevate the aragonite saturation of the extracellular calcifying fluid (ECF) to induce carbonate precipitation, consistent with models for biocalcification in tropical corals (Adkins et al., 2003; Cohen and McConnaughey, 2003; McConnaughey, 1989a; Sinclair and Risk, 2006). Further modification of the pH of the physiologically controlled internal fluid reservoirs could be caused by respiration or other aspects of polyp metabolism or by variations in calcification rate (Adkins et al., 2003; Goreau, 1977; Krief et al., 2010; McConnaughey, 1989b). Regardless, foraminifera modeling studies have shown that the presence of such biological factors does not necessarily compromise the d11B pH proxy if their effect can be adequately characterized and is constant over a range of relevant pH (Zeebe et al., 2003). Generalizing from this case, physiological vital effects that

shift the internal pH by a factor that is either constant or a function of external ambient seawater pH may be correctable and therefore preserve the signal of seawater pH in the d11B of the skeleton. Here we present a calibration of the d11B proxy in globally distributed modern Desmophyllum dianthus (D. dianthus) coral specimens, against ambient pH at their growth locations. The relatively large size of these corals provides enough aragonite to allow the measurement of multiple geochemical proxies. A multiproxy approach on individual corals growing at different locations in intermediate and deep waters could thereby reveal concomitant variations in chemical and physical properties in a single carbonate archive, a rare luxury. Our previous work on this species has demonstrated modern calibrations of P/Ca, Ba/Ca, and U/Ca ratios as proxies of seawater phosphate, barium, and carbonate ion (Anagnostou et al., 2011). The addition of a pH proxy would augment this suite of geochemical tools and, because of fundamental biogeochemical relationships among nutrients and carbonate system parameters, provide checks on paleoceanographic interpretations.

2. Materials and methods 2.1. Sample preparation and analytical approach Corals were sampled from the National Museum of Natural History (Smithsonian Institution, Washington, D.C.) and from the National Institute for Water and Atmosphere (NIWA), Greta Point, Wellington, NZ (Fig. 1). Coral septa were separated mechanically and cleaned with a power dental saw to remove any surface crust material. Subsequently, the coral pieces were rinsed and ultrasonicated with distilled water several times, and any remaining tissue was removed with a brush. Next, the corals were dried in a laminar flow clean bench and crushed to o2 mm pieces. Informed by recent studies suggesting internal skeletal inhomogeneity in the DSC Lophelia pertusa (L. pertusa) (Blamart et al., 2007), sufficiently large skeletal samples (17–108 mg; Table 1) were excised to integrate any possible micro-structural heterogeneity in D. dianthus d11B. Although we cannot prove that such heterogeneity was averaged in a completely consistent manner, the goal of this initial study is to test the potential of a d11B-pH proxy method that is characterized by good subsampling

Fig. 1. Locations of the D. dianthus specimens used in this study. Grey circles represent samples for which reliable hydrographic data were not available and thus were not used in proxy calibration in proxy calibration.

Table 1 Deep sea coral (D. dianthus) specimens used in this study, hydrographic data and calculated carbonate system variables for the specimen locations, and boron isotope data with replicate analyses 2  d11B range about the mean (n ¼ 2) (%)

DpH

17.1 32.1 34.8 28.2 26.8 107.9 61.8 58.3 78.1 28.7 20.9 36.1 41.1 18.7 61.1 79.4 35.8

26.01 28.13 26.15 26.99 26.63 25.89 26.02 24.95 26.81 27.48 23.56 26.32 24.96 26.71 25.87 25.91 25.49

0.18 0.15 0.02 0.05 0.09 0.00 0.02 0.24 0.24 0.24 0.01 0.09 0.13 0.11 0.30 0.04 0.08

0.71 0.91 0.80 0.86 1.04 1.06 0.86 1.20 0.85 0.96 1.08 0.83

Replicate sampling/analysis

Weight (mg)

d11B (%)

2  d11B range about the mean (n ¼ 2) (%)

d11Boffset of d11Baverage all replicates (%) replicated (%)

47409b 19249b 83583b 47413b 94069b 62309b 48473b 80358b 48739b 47407b 84820b 84818b

14.7 10.6 24.4 17.5 12.8

26.82 26.08 25.04 26.89 27.89 25.82 25.97 25.61 27.64 24.34 26.74 26.39

0.21 0.24 0.20 0.19 0.22 0.03 0.28 0.05 0.06 0.10 0.07 0.12

0.02 0.19 0.09 0.08 0.41 0.20 0.17 0.40 0.49 0.62 0.02 0.52

80358 48739 48473 48740 47409 19249 62309 83583 47413 94069 78630 Z9725 47407 84820 84818 19168 82065

a b

Depth Lat (m) (1N)

358 825 1107 1470 673 274 522 464 421 710 312 276 549 806 400 636 586

Long (1W)

Date Weight collected (mg)

48.0 7.9 1967 47.6 7.3 1973 47.7 8.1 1973 48.7 10.9 1973  54.5 39.4 1966 34.0 119.5 1889 40.4 67.7 1979 32.9 127.8 1986  50.6  167.6 1964  30.5 178.7 1993 46.8 130.8 1982  45.2  171.6 1999  54.8 129.8 1964 0.3 91.6 1986 0.1 90.3 1986  51.9 73.7 7888  54.8 129.8 1964

a b b b b b b

Date WOCE hydrographic line

Station Salinity Temp (1C)

pHw total

pH Oarag Total alk error (mmol/kg)

TCO2 (mmol/kg)

pKB

error

0.02 0.02 0.02 0.00 0.01 0.05 0.04 0.08 0.02 0.02 0.05 0.01

1997 1997 1997 1997 2005 2001 1997 2004 1996 2003 2008 1996

33 32 32 32 19 171 73 165 66 191 138 77

8.05 7.98 7.97 7.99 7.89 7.76 7.97 7.57 8.04 7.95 7.58 8.00

0.01 0.01 0.01 0.00 0.01 0.05 0.04 0.08 0.02 0.00 0.05 0.01

2129 2172 2182 2169 2260 2196 2179 2284 2124 2146 2248 2130

8.75 8.74 8.74 8.76 8.85 8.79 8.80 8.81 8.81 8.78 8.83 8.78

DpH

26.69 25.95 24.98 26.83 27.62 25.92 26.06 25.81 27.88 24.65 26.72 26.13

A24 A24 A24 A24 A16S P02 A22 P02 P15S P06W P12 P15S

35.54 35.53 35.55 35.12 34.66 34.04 35.05 34.20 34.25 34.49 33.91 34.51

11.0 9.5 8.4 5.0 1.7 8.1 6.1 6.3 5.9 7.3 5.4 8.1

2.1 1.6 1.5 1.3 1.0 1.0 1.4 0.6 1.6 1.4 0.6 1.7

2336 2344 2348 2324 2347 2264 2317 2293 2284 2284 2259 2288

d11B(OH)4 (%)

16.49 15.96 15.89 15.84 14.61 14.25 15.43 13.45 15.85 15.44 13.52 15.77

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d11B (%)

Coral ID

Subsampled from the same solid aliquot. Replicate aliquot from the same dissolved sample.

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Table 2 Neptune MC-ICPMS operating conditions RF power

1200 W

Ar cooling gas flow rate Ar sampling gas flow rate Ar auxiliary gas flow rate Interface cones Nebulizer type Inlet system

15.0 L min  1 1.1 L min  1 0.7 L min  1 Ni-Sampler, X-Skimmer PFA microflow, 100 mL min  1 (Elemental Scientific) Stable inlet system with dual-quartz spray chamber Low H3 (11B) and L3 (10B) 1.049 s

Resolution Faraday cup Integration time

reproducibility and ease of sample processing and d11B measurements (Table 1). Boron purification followed the micro-sublimation method established by Wang et al. (2010), and isotopic analyses were performed using multi-collector inductively coupled plasma mass spectrometry (MC-ICPMS) (Neptune, Thermo Scientific) at the National Cheng Kung University, Taiwan, as per previously published methods (Foster, 2008). Typical operating conditions and introduction system details are summarized in Table 2. Instrumental mass bias was adequately corrected using a standard-sample bracketing technique (Foster, 2008; Wang et al., 2010), using NIST SRM 951 boric acid. Standard concentrations were typically 20 ppb B, which gave 0.4–0.5 V of 11B signal intensity. Every sample was analyzed in duplicate and an average value with the range about the mean is reported (Table 1). A number of coral specimens were sub-sampled, crushed and processed as independent samples; offsets between these replicate pairs ranged 0.08–0.41% (5 replicated samples; Table 1). For the 62309 coral, the crushed skeleton was homogenized and split into duplicate subsamples. The remaining 6 duplicated samples were separated aliquots from the dissolved coral solution prior to sublimation (marked as b in Table 1). Differences between duplicate pairs in these cases ranged 0.02–0.62% (Table 1). Thus, overall, replicate subsamples agree within 0.3% about the mean. Given the poor wash out and increased memory effect characteristic of B compared to other elements (Al-Ammar et al., 2000), samples were diluted to 20–40 ppb B prior to analyses. Samples, standards and washes were introduced in 0.1 M HNO3. The wash out using this protocol reduces signal to less than 2.5% of initial in 110 s, following introduction of 50 ppb B samples. Blanks monitored throughout analytical runs that included a variety of samples (20–40 ppb B) and NIST SRM 951 boric acid standard (typically 20 ppb B) were reproducibly o1 ppb B or 0.01 V for 11B.

2.3. Carbonate system data The twelve deep-sea coral specimens used for the calibration analyses (Fig. 1) are assumed to have been collected alive since residual animal tissue was observed on all corals except specimens 94069 and 48739. The corals were growing in areas with relatively well-constrained seawater properties, although the quality of these data depends on availability of reliable alkalinity and dissolved inorganic carbon (DIC) concentrations from the CDIAC ocean carbon system database (http://cdiac.ornl.gov). Using the Matlab version of carbonate system calculation software CO2sys Ver. 1.1 (van Heuven et al., 2011) with constants from Mehrbach et al. (1973) refit by Dickson and Millero (1987), available alkalinity and DIC concentrations were used to calculate ambient pH. The total pH scale is used throughout the manuscript. The hydrographic stations corresponding to each coral location were selected based on their proximity to the coral collection coordinates, taking into consideration the quality of the carbonate parameters measured. Of the seventeen samples analyzed, five were collected at locations for which proximal hydrographic data with appropriate carbonate system variables could not be found; these were not included in the calibration. Correction for anthropogenic CO2 invasion, by subtracting the total regional cumulative anthropogenic contribution to DIC and recalculating pH (Sabine et al., 2004), resulted in o0.01 pH error at stations where such estimations were available. This effect was comparable to the measurement error and was therefore neglected. The quality of a proxy calibration depends on the uncertainty in the oceanographic variable as well as in the proxy measurement; we estimated the former in several different ways. First, there is uncertainty inherent in the calculation of pH from alkalinity and DIC, but it is small. Typical uncertainties in the laboratory determination of these variables propagate to a pH uncertainty of o0.01 pH units (See EPOCA Guide to Best Practices for Ocean Acidification Research and Date Reporting; http://www. epoca-project.eu/index.php/guide-to-best-practices-for-ocean-aci dification-research-and-data-reporting.html). Very likely a larger uncertainty is imposed by the agreement in space and time between the hydrographic data used and the coral sample site. We used the GLODAP Gridded database (http://cdiac.ornl.gov/ oceans/glodap/Glop_grid_OV.html) to estimate ranges of alkalinity and DIC within a grid box of 1–8o from the coral location, depending on data availability, and depth range within 70 m of the coral depth. These carbonate system parameters were combined with temperature, salinity, and inorganic phosphate data from the selected hydrographic station (Table 1) to calculate using CO2sys the range of pH of seawater on the total scale (pHw) within the grid box, and from that a conservatively estimated uncertainty in pHw.

3. Results and discussion 2.2. Accuracy and precision 3.1. Proxy development To overcome the lack of accepted reference standards for B isotope composition, the JCp-1 carbonate standard was analyzed repeatedly, since this standard has been tested extensively for d11B by MC-ICPMS, N-TIMS and P-TIMS (Aggarwal et al., 2009; Douville et al., 2010; Wang et al., 2010). Full procedural replicates for analyses of JCp-1 give a long-term precision of 24.41 70.30% 2SD (n¼ 10), in agreement with the values of 24.2270.28% and 24.570.2% as previously reported (Wang et al., 2010; Douville et al., 2010, respectively). Additionally, repeated analyses of an inhouse boron high purity solution standard (Alfa-Aeser) exhibited a high level of reproducibility (  5.41 70.19%, 2SD; n ¼16).

The B isotopic composition of the D. dianthus corals appears to vary with pH across a wide range of oceanographic locations (Table 1). The corals analyzed in this study were selected from a broad geographic and depth range, so that each specimen grew at a unique combination of temperature, salinity, and pressure. Therefore, direct quantitative comparison of coral d11B to the seawater borate d11B-pH curves (typically defined for 25 1C, 1 atm, and S ¼35) is problematic. Nevertheless, for visual presentation purposes only, we plot the new D. dianthus data against ambient pH at each of the coral sampling locations (Fig. 2), with

E. Anagnostou et al. / Earth and Planetary Science Letters 349–350 (2012) 251–260

255

32

δ11 B (‰)

30 28

Desmophyllum dianthus (this study)

26

Acropora nobilis (Honisch et al. 2004)

24

Acropora sp. (Reynaud et al. 2004)

22

Porites cylindrica (Honisch et al. 2004)

20 18

Porites (Krief et al. 2010)

16

Stylophora pistillata (Krief et al. 2010)

14

Cibicidoides sp. (Rae et al. 2011)

12 10 7.1

7.3

7.5

7.7 pHw

7.9

8.1

8.3

Fig. 2. Coral d11B plotted against ambient seawater pH. D. dianthus data (solid black circles; this work) are plotted with error bars representing range about the mean of replicate analyses and estimated uncertainty in pH based on hydrographic data (Table 1). Published d11B-pH data from tropical coral culture and benthic foram studies are also plotted, with horizontal error bars representing precision of pH control in incubation experiments, as well as the seawater borate curve (red line) as described by Klochko et al. (2006). We note that this figure should be used only for a visual representation of the d11B magnitude of each species and not for regression or comparison with the seawater borate curve, as each deep coral grew at conditions which are different from the seawater borate curve, determined at 25 1C, 1 atm, and S ¼35 (Klochko et al., 2006). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

the Klochko et al. (2006) seawater curve shown for context, along with a selection of literature d11B data for tropical corals and benthic (calcitic) foraminifera included for comparison. Little can be done quantitatively with this graph because the seawater curve is not appropriate to the variable depths and temperatures of the deep coral samples. It should be noted that a significant error in the vibrational spectrum term of borate used by Kakihana et al. (1977) has been identified recently (Klochko et al., 2006; Rustad and Bylaska, 2007). The value of aB ¼1.0194 (Kakihana et al., 1977) is therefore not a correct equilibrium fractionation factor for seawater, but has come to be considered an approximate fitting parameter that can be used to describe the curvature of the d11B vs. pH curve of some marine carbonates (Fig. 2; ¨ Honisch et al., 2004, 2007; Krief et al., 2010; Reynaud et al., 2004; Sanyal et al., 1996, 2001). To allow an appropriate evaluation of the relationship of the deep coral data to the isotopic composition of ambient seawater, the measured skeletal d11B was plotted (Fig. 3) against seawater borate d11B (as per Rae et al. (2011)) calculated iteratively using ambient hydrographic variables (Table 1) and the following equation: 11

11

11

11

pH ¼ pK nB logððd BSW d Bborate Þ=ðd BSW aB d Bborate ðaB 1Þ1000ÞÞ

ð1Þ where pH is in situ as calculated by CO2sys and tabulated in Table 1, pK nB is the equilibrium constant for the dissociation of boric acid to borate at ambient temperature and pressure conditions (computed by CO2sys using Culberson and Pytkowicz (1968) and Dickson (1990); salinity correction contributes r0.003 pH uncertainty and was thus considered negligible), aB is the fractionation factor of 1.0272 for the exchange of 10B and 11B between the two major aqueous species (Klochko et al., 2006), d11Bsw is the isotopic composition of seawater (39.6%; Foster et al., 2010), and d11Bborate is the isotopic composition of borate in seawater.

For comparison, the 1:1 line with seawater borate d11B is also shown, indicating where the coral B isotope composition should fall if borate were simply incorporated at ambient pH and with no opportunities for fractionation. Indeed, recent data for the calcitic benthic foraminifera Cibicidoides sp. (Rae et al., 2011) shows a regression indistinguishable from the 1:1 seawater line. Also shown are data from published tropical coral culture studies in which pH was artificially manipulated in incubation experiments (Fig. 3). To evaluate the potential influence of various hydrographic variables on d11B of D. dianthus, the offsets of coral d11B from the d11B of borate predicted using the Klochko fractionation factor were regressed against ambient salinity, pressure, temperature, and carbonate ion concentration to test for possible secondary dependencies. No obvious relationships were found (all correlations resulted in R2 o0.1). Therefore, although there may be a relevant temperature dependence of the Klochko aB value, despite the lack ¨ of consensus on the numerical value of that dependence (Honisch et al., 2008; Klochko et al., 2006; Zeebe, 2005), no temperature correction is supported by the current D. dianthus d11B offsets given the lack of correlation with temperature or any other hydrographic parameter. Therefore, we applied no temperature correction to aB. A similar approach was taken by Rae et al. (2011), in their study of benthic foraminifera, with the same result. Were a temperature correction to aB significant, it is worth noting that thermodynamic principles dictate that aB would be higher at lower temperatures compared to the published value determined at 25 1C (Klochko et al., 2006). Since the corals used in this study all grew in temperatures o25 1C, a significant temperature dependence would manifest in Fig. 3 as a flattening of the D. dianthus data points, to a regression slope o0.76. In this sense, the slope calculated for our data may be, without a temperature correction, an upper limit estimate. While all tropical corals plot above the Klochko et al.’s (2006) seawater borate d11B (Fig. 3), the D. dianthus d11B values are

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30

y(D.dianthus) = 0.76x + 14.67 R² = 0.43

28 26 δ11B coral (‰)

D.dianthus (this study): slope 0.76

24

Acropora sp. (Reynaud et al. 2004): slope 0.71 Porites (Krief et al. 2010): slope 0.76

22

S.pistillata (Krief et al. 2010): slope 0.71

20

A.nobilis (Honisch et al. 2004): slope 0.81 P.cylindrica (Honisch et al. 2004): slope 0.76

18

Cladocora (Trotter et al. 2011): slope 0.75

16

Cibicidoides (Rae et al. 2011): slope 1.00

14 12

14

16 δ11B B(OH)4 - ( ‰)

18

20

Fig. 3. Coral d11B (blue squares, overall mean for each sample) vs. calculated d11B of seawater borate using Eq. (1) for ambient pH, temperature and salinity conditions for each coral location. Error bars represent uncertainty of coral replicate analyses as described in Table 1. For comparison, a number of studies in tropical corals and benthic foraminifera are included. The lines represent least squares linear regression fits, the slopes of which are provided in the legend, while the dashed curves around the D. dianthus data represent the 95% CI about the regression. The dashed line through the Cibicidoides sp. is the 1:1 line with seawater borate ion, and the dotted line through the D. dianthus is parallel to the 1:1 line, offset by 11%. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

heavier than the d11B of tropical surface corals Acropora sp., ¨ Porites sp. and Stylophora pistillata (S. pistillata) (Honisch et al., 2004; Krief et al., 2010; Reynaud et al., 2004), the temperate coral Cladocora caespitosa (C. caespitosa) (Trotter et al., 2011), and the planktonic foraminifer Orbulina universa at pH¼8.19 (Kasemann et al., 2009). The benthic foraminifera Cibicidoides sp. and Planulina are the only organisms identified to date that incorporate boron close to the seawater borate d11B, showing no obvious offset from the seawater borate curve (Klochko et al., 2006; Rae et al., 2011). The D. dianthus data of Fig. 3 can be interpreted as a simple regression or by imposing the assumption that the offset from the 1:1 line is independent of pH and thus that the slope ¼1.0. A least squares linear regression fit is described as

d11 Bcoral ¼ ð0:76 70:28Þd11 Bborate þð14:67 74:19Þ ð1SE, R2 ¼ 0:43Þ ð2Þ The slope of the regression is similar to that for P. cylindrica ¨ (Honisch et al., 2004) and Porites (Krief et al., 2010). If instead we force a regression parallel to the 1:1 line with the seawater borate, the intercept equals 11.0 ( 70.3 1SE, R2 ¼0.39), within the uncertainty in both the slope and intercept in the simple least squares linear regression scenario.

3.2. Explaining the magnitude of the boron isotopic composition of corals The differences in absolute d11B range among the species plotted in Fig. 3 suggest distinct mechanisms of calcification and/ or B isotope incorporation. Although the bulk analysis approach followed in this study cannot offer a detailed understanding of the mechanism of boron inclusion during biomineralization in D. dianthus, the heavy d11Bcoral compared to seawater borate d11B suggests that one or more of several potential processes may be operating to drive the coral d11B higher than that of the ambient seawater borate, and higher than that of tropical corals.

The d11B offset of tropical corals and planktonic foraminifera from the seawater borate curve has been suggested to result from some combination of (1) possible uptake and incorporation of some of the isotopically heavier boric acid species during mineralization with potential subsequent isotopic redistribution through intermediate boron-complexes (Klochko et al., 2009; Rollion-Bard et al., 2011a); (2) manipulation of ECF pH assisted by (a) photosynthesis (Cohen and McConnaughey, 2003) which consumes protons and may increase alkalinity (Holcomb et al., 2010; Langdon and Atkinson, 2005), causing the coral to record ¨ higher pH than ambient (Honisch et al., 2004; Krief et al., 2010; Reynaud et al., 2004), or (b) the action of Ca-ATPase, carbonic anhydrase, proton pumps and other enzymes with a combined result of ECF super-saturation and aragonite precipitation; and (3) other vital effects, e.g. growth rate and temperature dependent metabolic processes (e.g. Trotter et al., 2011; McCulloch et al., 2012). Currently, the importance of seawater boric acid to boron incorporation in D. dianthus is unknown. Although threefoldcoordinated boron could be present in the coral aragonite, as suggested for L. pertusa (Rollion-Bard et al., 2011a) and other marine carbonates (Klochko et al., 2009), there is no firm evidence to suggest that it originates from seawater boric acid. Instead, if present, trigonal boron in coral skeletons may be a transformation product of seawater borate ion during coral calcification (Klochko et al., 2009). Nevertheless, the hypothesis of boric acid incorporation in DSC needs to be tested further using independent techniques, and examined in corals growing at low pH where boric acid in seawater is more abundant. The d11B offset for DSCs cannot be related to photosynthesis, because these deep-dwelling organisms lack symbiotic algae. Taxon-specific vital effects including polyp metabolism and physiological modification of the ECF pH are, however, mechanisms that may contribute to the observed offset. While metabolic vital effects cannot be ruled out, it is not clear how polyp metabolism might affect D. dianthus d11B. Certainly polyp respiration is not a likely candidate for the observed results, as respiratory CO2 would lower internal pH, not raise it. In addition,

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9.0

Apparent coral ECF pH

8.9

8.8

8.7

8.6

8.5

8.4 7.5

7.6

7.7

7.8 pHw

y = -0.75x + 6.88 R² = 0.80

1.2

1.0

y(McCulloch) = -0.72x + 6.50 R² = 0.79

0.8

D. dianthus (this study)

0.6 Cladocora (Trotter et al. 2011)

7.3 11‰

11B

(‰)

35

y(Trotter) = -0.70x + 6.24 R² = 0.95

7.5

7.7 pHw

7.9

8.1

Fig. 6. Coral reconstructed pH minus seawater pH (DpH) vs. seawater pH (pHw; total scale). The linear regression fit is indistinguishable from a logarithmic fit (as expected from the curvature of the seawater borate curve) for the range of seawater pH in this study. For comparison, published data for C. caespitosa (Trotter et al., 2011) are shown as closed circles, and for D. dianthus (McCulloch et al., 2012) as closed diamonds. Dotted lines represent 95% CI about the regression fit to the D. dianthus data.

pH ~0.7

30 pH ~1 25 11‰

15 10 7.5

8.1

0.4

40

20

8.0

1.4

Dianthus (McCulloch et al. 2012)

45

7.9

Fig. 5. Calculated pH recorded in corals assuming only borate ion is incorporated, and that no further fractionation occurs in the ECF during calcification. The calculated pH is derived using the seawater borate curve of Eq. (1) with the Klochko et al. (2006) fractionation factor for ambient temperature and salinity.

ΔpH (Coral recorded - Seawater)

respiratory CO2 is not expected to contribute significantly to DIC in the ECF, as isotopic studies have shown that D. dianthus calcifies using seawater, not respiratory, DIC (Adkins et al., 2002). Thus discounting a respiration driven vital effect at the calcifying site, the results for D. dianthus are best explained by physiological control of the pH of the ECF, thereby affecting the bulk coral d11B measurements. It is reasonable that corals need to increase the ECF pH to enhance calcification (Erez et al., 2011), and evidence for this process is mounting from recent pH measurements under the calcifying tissue of zooxanthelate or facultative zooxanthelate tropical and temperate corals (Al-Horani et al., 2003; Ries, 2011; Venn et al., 2011). This mechanism is especially relevant to D. dianthus, which survives in waters with Oarag o1 (Table 1). It is known that a process of pH adjustment would be subject to enzymatic activity in the coral; key enzymes studied to date include carbonic anhydrase (Furla et al., 2000) and Ca-ATPase (Cohen and McConnaughey, 2003; McConnaughey and Whelan, 1997). However, the mechanisms by which physiology and biochemistry might affect boron isotope fractionation and boron incorporation in coral aragonite remain uncertain (Rollion-Bard et al., 2011a, 2011b). The slope of the D. dianthus regression in Fig. 3 can thus be interpreted as the degree of ECF pH control. Perfect control of the ECF to constant pH (and constant omega, assuming DIC is constant as well), across the range of seawater pH, would result in a flat response of d11B in the coral. Varying degrees of imperfect compensation would result in slopes between 0 and 1. What is striking about Fig. 3 is that all corals display a slope of 0.7570.5, suggesting that the degree to which ECF pH is modified as ambient pH decreases is common to all corals, even if the absolute offset of ECF relative to seawater pH is taxondependent. We point out here that, graphically and mathematically, the curvature of the seawater borate curve necessarily dictates than an observed constant offset in d11B (Fig. 3 dotted line through D. dianthus data) requires a pH offset that is more pronounced at lower pH values, as shown for a hypothetical case in Fig. 4. Thus a constant offset observed in a d11Bcoral vs. pHw plot (which is convenient but in no way required) is tantamount to evidence supporting an ECF pH modification that is dependent on ambient pH. A slope of o1, as shown by all of the coral data (Fig. 3), is consistent with enhanced ECF pH adjustment relative to the constant d11B offset case (slope¼1).

257

7.8

8.1

8.4

8.7

9.0

9.3

pHw Fig. 4. Seawater d11B vs. pH curve (Klochko et al., 2006) and a hypothetical parallel curve drawn at 11% offset from the seawater curve. This is a graphical representation of the important point that an invariant d11B offset connotes varying degrees of pH offset, with higher offset at lower seawater pH.

We calculated the apparent ECF pH using the seawater borate fractionation factor and the coral d11B, to investigate the relationship between ECF pH and ambient seawater pH. Such an approach assumes that the coral d11B reflects the coral d11B of borate in the coral ECF and that no further fractionation occurs during incorporation in the newly forming aragonite. The data suggest that corals modify their ECF pH by a variable offset that is inversely related to ambient pH, yielding a narrow range of ECF pH values (Fig. 5). The ECF pH falls in the range of 8.68–8.93 for all samples. At the calculated average ECF pH of 8.8 for our DSC samples

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(based on aB ¼1.0272; Klochko et al., 2006), and at DIC assumed unaltered from that of ambient seawater, the Oarag is within the range of 7–9 for all of the D. dianthus specimens. These results suggest that physiological pH adjustment of the calcifying fluid by 0.7–1.2 pH units, with greater adjustment at lower ambient pH, could reasonably explain the results for D. dianthus. Following a model of elevation of the ECF pH, Trotter et al. (2011) suggested that a calibration of the biological offset in pH as determined from d11B in the surface coral C. caespitosa, could be used iteratively to calculate ambient seawater pH at which a coral grew. We explore a similar approach for D. dianthus (Fig. 6), which results in the regression line described as

calcification. Nevertheless, the low abundance of D. dianthus in the Pacific and Indian Oceans could point towards additional factors influencing deep sea coral survival, including food availability, oxygen levels, and temperature (Atkinson and Cuet, 2008 ¨ and references there in; Form and Riebesell, 2011; Portner, 2008; Tanaka et al., 2007; Tribollet, 2008). Therefore a careful consideration of coral adaptation strategies in low pH environments should include an investigation of the effects of other environmental variables on DSC behavior and biogeography, especially in regions where the organisms may need to expend maximum metabolic energy to overcome undersaturated conditions.

DpH ¼ ð0:75 70:12ÞpHw þ ð6:88 7 0:93Þ ð1SE, R2 ¼ 0:80Þ

3.3. Potential for paleo-reconstructions

ð3Þ

where DpH is the difference between calculated pH of the ECF and measured pH of the seawater. It should be noted that if corals were able to maintain the pH of the ECF at a constant value independent of ambient pH, the slope of the DpH vs. pH regression would be equal to one and a pH proxy would be impossible. Here, however, we observe that the biological pH increase in the ECF is not sufficient to completely counteract the ambient seawater pH decrease, indicating that D. dianthus can only partially compensate for low Oarag in ambient seawater. This observation is in agreement with studies of temperate and tropical corals whose exposure to seawater acidification resulted in a moderate increase in the pH offset of the ECF from seawater pH, compared to corals grown under normal conditions (Cohen et al., 2009; Ries, 2011). While this paper was in revision, a few additional data points for d11B of D. dianthus specimens became available (McCulloch et al., 2012). Although the sample size and ambient pH range are smaller, we note that the DpH vs. pH regression in this study is indistinguishable, within error, from our regression, and we plot the additional data in Fig. 6 for comparison to ours. Even with the small sample number, the agreement is striking given different approaches to subsampling and independent analytical techniques. This convergence lends strength to our conclusions regarding the potential of D. dianthus as a viable recorder of ocean pH. The DpH vs. pHw regression for D. dianthus results in a slope similar to that suggested for the faster growing C. caespitosa, Porites sp., and S. pistillata corals (Trotter et al., 2011 and references therein). The DpH values for D. dianthus, however, are higher throughout a similar ambient pH range, suggesting a taxon-specific mechanism that achieves higher ECF pH than that observed for tropical corals. These results are of great interest for understanding fundamental mechanisms of biocalcification and their response to seawater pH. Under conditions of increased oceanic CO2, corals may have to exert additional energy to counter a large pH gradient between seawater and calcifying space, posing a challenge to the energy budget of these organisms (Cohen and Holcomb, 2010). Adaption may involve allocation of energy expenses, e.g. by allowing ECF pH to be lower than the more optimal ECF pH possible at higher seawater pH, as observed for D. dianthus (Fig. 5), by reproducing infrequently, or by calcifying at a slower rate (Holcomb et al., 2012 and references therein). Coral studies on the response of coral calcification rates to ocean acidification are equivocal (e.g. summary by Andersson et al. (2011)). The DSC species L. pertusa (Maier et al., 2009) and Oculina sp. (Fine and Tchernov, 2007; Ries et al., 2009) clearly exhibit a tight relationship between calcification rates and pH, with lower calcification rates observed at lower seawater pH. The slow calcification rate of D. dianthus (e.g. Cheng et al., 2000) may be linked to the high observed ECF high throughout the modern ocean pH range, as even moderate rates of ion pumping may be able to compensate for slow proton production during

Our data suggest that seawater pH could be reconstructed from coral d11B using Eq. (1) after seawater borate d11B is estimated from regression (2). The reconstruction uncertainty based on 95% CI around the regression equals 0.7% ( 0.07 pH) for d11B of borate  15%, increasing to 1% (  0.10 pH) for d11B of borate o14% and 415.5% at the extremes of the regression. This uncertainty is similar to that calculated based on the standard deviation of the residuals of the coral d11B data from the regression fit (2) and Fig. 3 (1.16%,  0.12 pH). Alternatively, the DpH vs. pH Eq. (3) could be used to reconstruct seawater pH. The reconstruction uncertainty based on the regression error envelope (95% CI) is 0.05 pH, and it increases to 0.10 for seawater pHo7.8. The individual DpH uncertainties for our samples, propagating the error on both pHw and pH of ECF, are between 0.00 and 0.08, dominated by the uncertainty in the hydrographic pH. The standard deviation of the residuals in DpH vs. pH regression fit equals 0.08 pH, which is within the range of the uncertainty based on the error envelope (95% CI). Although Eq. (3) results in a somewhat lower reconstructed pH error, we suggest the use of Eq. (2) for proxy application, as it does not involve ECF pH and so does not imply knowledge of the carbonate system or the details of biologically controlled processes at the site of calcification. Limited sample availability at locations where seawater pH is low, the bulk analytical approach followed in this study, and uncertainty in hydrographic pH could account for the scatter in d11B vs. pH plots observed here. A larger number of corals analyzed from intermediate to low pH oceanic regimes are needed to refine the slope in the DpH regression, and verify the offset between seawater and coral d11B. Additionally, microsampling might be required to obtain tighter regressions, minimizing the possible inclusion of variable fractions of COCs and fibrous aragonite among different coral specimens (Rollion-Bard et al., 2011a). Microsampling by laser ablation was found necessary for other proxy elements in this species (Anagnostou et al., 2011), driven by the need to avoid contamination associated with surface phases (Mason et al., 2011), potential diagenetic alteration, and compositionally unique COCs; the influence of these skeletal features on coral d11B is not well known at this point. Uncertainty in hydrographic pH could be overcome through long-term culture studies under tightly controlled conditions, or through careful carbonate system measurements accompanying new specimen collections.

4. Conclusions This study demonstrates the feasibility of using d11B in the scleractininan deep sea coral D. dianthus to reconstruct seawater pH. The d11B of 12 globally distributed specimens of D. dianthus fall in the range of 23.56–27.88, and plot above the seawater borate d11B vs. pH curve (Klochko et al., 2006) and above the d11B of tropical corals measured in previous studies. In addition, a

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regression of d11B for this species vs. d11B of seawater borate gives a slope of 0.76, very similar to the slope determined for tropical corals, but offset to higher d11Bcoral values. This behavior is explained by physiological adjustment of the pH of the extracellular calcifying fluid, to a degree that is inversely related to ambient seawater pH, and spanning a range of pH 8.67–8.93 for seawater pH of 7.57–8.05. While we cannot explain why ECF pH is apparently higher in this species than in tropical corals or foraminifera, the results suggest strong control of internal pH by this species, which could reflect adaptation to the low pH environments in which the species is currently found. Based on this initial study, the D. dianthus d11B-pH proxy could be used to reconstruct pH with an uncertainty of  0.07 pH for d11Bborate of15%, increasing to 0.1 pH for d11Bborate o 14% and 415.5% (based on 95% CI about the regression, Eq. (2)). The uncertainty calculated using the standard deviation of the scatter of the coral d11B data about the regression (2) equals  0.12 pH units. With further development to refine the sampling approach, address possible boric acid incorporation, and investigate further the potential artifacts associated with diagenesis and surface contaminants, especially in fossil corals, B isotope measurements in D. dianthus could offer sufficient reconstruction precision on ocean pH to provide a valuable complement to other proxies of ocean carbonate system and biogeochemistry in the past.

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