Partial diagenetic overprint of Late Jurassic belemnites from New Zealand: Implications for the preservation potential of δ7Li values in calcite fossils

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Partial diagenetic overprint of Late Jurassic belemnites from New Zealand: Implications for the preservation potential of d7Li values in calcite fossils Clemens V. Ullmann a,b,⇑, Hamish J. Campbell c, Robert Frei a, Stephen P. Hesselbo b, Philip A.E. Pogge von Strandmann b, Christoph Korte a a

Department of Geography and Geology & Nordic Center for Earth Evolution (NordCEE), University of Copenhagen, Øster Voldgade 10, DK-1350 Copenhagen, Denmark b University of Oxford, Department of Earth Sciences, South Parks Road, Oxford OX1 3AN, UK c GNS Science, 1 Fairway Drive, Avalon 5010, New Zealand Received 9 November 2012; accepted in revised form 23 June 2013; available online 3 July 2013

Abstract The preservation potential and trends of alteration of many isotopic systems (e.g. Li, Mg, Ca) that are measured in fossil carbonates are little explored, yet extensive paleoenvironmental interpretations have been made on the basis of these records. Here we present a geochemical dataset for a Late Jurassic (153 Ma) belemnite (Belemnopsis sp.) from New Zealand that has been partially overprinted by alteration. We report the physical pathways and settings of alteration, the resulting elemental and isotopic trends including d7Li values and Li/Ca ratios, and assess whether remnants of the primary shell composition have been preserved or can be extrapolated from the measured values. The d18O and d13C values as well as Sr/Ca and Mn/Ca ratios were analysed along two profiles. In addition, 6 samples were analysed for 87Sr/86Sr, Sr/Ca and Mn/Ca ratios. Five samples from the same specimen and 2 from the surrounding sediment were analysed for d7Li values, Li/Ca, Sr/Ca and Mn/Ca ratios and are compared to results for 6 other Late Jurassic belemnite rostra (Belemnopsis sp. andHibolithes sp.) from the same region. The 87Sr/86Sr ratios are lower (less radiogenic) in the most altered part of the rostrum, whereas d7Li values become more positive with progressive alteration. The direction and magnitude of the trends in the geochemical record indicate that one main phase of alteration that occurred in the Late Cretaceous caused most of the diagenetic signature in the calcite. Despite relatively deep burial, down to 4 km, and thus elevated temperatures, this diagenetic signature has subsequently been preserved even for the highly mobile element lithium, suggesting that primary lithium-isotope values can be maintained over geological timescales, at least in thick macrofossil shells. Our best d7Li estimate for pristine Late Jurassic (155–148 Ma) belemnites is +27 ± 1&, which points to a Late Jurassic seawater d7Li of 29–32&, compatible with the modern value of 31&. Ó 2013 Elsevier Ltd. All rights reserved.

1. INTRODUCTION Stable isotope signatures of ancient low-magnesium-calcite (LMC) shells (e.g. belemnites, brachiopods and bivalve molluscs) can be utilized to reconstruct palaeoenviron-

ments, but this requires that the sample material is well preserved. Diagenetic fluid reactions and re-crystallization potentially modify the primary geochemical signatures of the shells (e.g. Longinelli, 1969; Al-Aasm and Veizer, 1986a,b), but the degree of alteration can be recognized

⇑ Corresponding author at: Department of Geography and Geology & Nordic Center for Earth Evolution (NordCEE), University of Copenhagen, Øster Voldgade 10, DK-1350 Copenhagen, Denmark. Tel.: +45 353 24 174; fax: +45 353 22 440. E-mail address: [email protected] (C.V. Ullmann).

0016-7037/$ - see front matter Ó 2013 Elsevier Ltd. All rights reserved. http://dx.doi.org/10.1016/j.gca.2013.06.029

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by cathodoluminescence microscopy, scanning electron microscopy (e.g. Slen and Karstang, 1989; Podlaha et al., 1998; O’Neil et al., 2003; Gro¨cke et al., 2007) and/ or analysis of trace element patterns (Veizer, 1974; Brand and Veizer, 1980; Al-Aasm and Veizer, 1986a). Lithium has recently become of higher interest for palaeoenvironmental reconstructions and Li/Ca and 6Li/7Li ratios are practical tracers. The residence time of Li in seawater is >1 Ma (Li, 1982) which is roughly 103 times longer than the time required for its mixing in the water masses (Misra and Froelich, 2012). As a result, both, Li/ Ca ratios and d7Li values are spatially uniform in the oceans (Millot et al., 2004 and references therein) but vary on geological timescales due to long-term shifts in the Li fluxes in and out of the oceans. This suggests that Li/Ca ratios can be used for past seawater reconstruction (Delaney et al., 1985) and some studies raise the possibility that Li/ Ca ratios offer the potential for palaeotemperature reconstruction using some biogenic carbonates (Delaney et al., 1989; Marriott et al., 2004a). Li isotope fractionation in inorganic and organically mediated carbonate precipitation is insensitive to temperature (Tomascak et al., 2003; Hall and Chan, 2004; Marriott et al., 2004a), making fossil carbonates a promising tool to estimate past seawater Li isotopic composition. This isotopic ratio has been of increasing interest for the characterization of the global weathering intensity (Huh et al., 2001; Kisaku¨rek et al., 2005; Hathorne and James, 2006; Pogge von Strandmann et al., 2010). Using foraminifera from deep sea cores, Delaney and Boyle (1986) made inferences on past Li/Ca of the oceans since the Cretaceous, and recently, Misra and Froelich (2012) published a first seawater d7Li curve for the entire Cenozoic reconstructed from foraminiferal tests in deep sea cores. A high potential of diagenetic alteration on the Li signature in carbonates has been hypothesized by Bailey et al. (2003) who found inconsistencies in the Li/Ca records of Toarcian (183–174 Ma) belemnites from Yorkshire (UK) and Germany, but comprehensive studies on carbonate diagenesis including Li are still missing. Belemnites, Mesozoic predators that went extinct at the Cretaceous–Paleogene boundary 66 Ma ago (Iba et al., 2011), are ideal to investigate diagenetic effects. Their massive calcite rostra allow spatial control of high-resolution sampling and provide enough material for multiple analyses of trace element and isotope ratios. This advantage is used in the present study, and several geochemical tracers (d13C, d18O, 87Sr/86Sr, Mn/Ca and Sr/Ca ratios) have been measured from a partly diagenetically overprinted belemnite rostrum to identify pristine and secondary modified parts, to reconstruct the alteration history, and to discover alteration paths. A partially altered fossil belemnite, specimen NZ0256B from New Zealand, has been adopted for this purpose. In addition, d7Li values have then been analysed on pristine parts and on zones with different degrees of alteration in order to assess potential diagenetic impacts on Li, and to evaluate the preservation potential for Li in belemnites. This high resolution elemental/isotopic study on a single, partly altered, belemnite rostrum is essential, because potential species dependent effects on d7Li and Li/Ca are not yet quantified and possible differing mecha-

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nisms of Li exchange during alteration, i.e. diffusion or dissolution and reprecipitation, can potentially be teased apart by the approach adopted here. Lithium isotope measurements on six additional Late Jurassic belemnites from the same region in New Zealand have been measured. The geochemical data from these rostra were then used to assess the consistency of the trends observed in the single rostrum. 2. MATERIALS AND METHODS 2.1. Sample locations Belemnite rostra of the genera Hibolithes and Belemnopsis from the Late Jurassic in New Zealand (northwestern North Island) were provided by GNS Science (Lower Hutt, New Zealand; see Table 1 for further details). The fossil rostra were found in an almost complete and unfractured state, lacking only parts of the alveolar area (see Fig. 1 for nomenclature). They are of Heterian to Puaroan age corresponding to the Kimmeridgian Ataxioceras hypselocyclum ammonite biozone (155 Ma) to the Tithonian Micracanthoceras ponti ammonite biozone (148 Ma) on the international timescale (Table 1; Stevens, 1997; Cooper, 2004; Gradstein et al., 2012). The fossil localities occur in a palaeogeographic setting on the southeastern margin of Gondwana close to southeast Australia and/or northeast Antarctica (Adams et al., 2007; Pole, 2009). Here, volcanic derived sediments were deposited in an island arc setting at high southern palaeolatitudes (Gray and Norton, 1988). These sedimentary rocks form a thick succession of Late Permian to earliest Cretaceous age (Murihiku Supergroup) and crop out along part of the west coast of the North Island, and in the Nelson, Southland and Otago regions of New Zealand (Mortimer, 2004). 2.2. Belemnite palaeobiology Many aspects of belemnite palaeobiology, depth habitats and their mode of life are poorly constrained (Malkoc and Mutterlose, 2010; Zakharov et al., 2011 and references therein). A fully marine habitat is inferred and their lifestyle is interpreted as either nekto-benthic, i.e. actively swimming close to the bottom of the water column, or nektonic, i.e. swimming within the water column (e.g. Pirrie et al., 2004; Wierzbowski and Joachimski, 2007; Rexfort and Mutterlose, 2009). Some evidence from both anatomical as well as geochemical data points to a nektonic lifestyle of Hibolithes with preferred water-depths of 50–100 m (Rexfort and Mutterlose, 2009). The life-span of these marine predators is estimated to be a few years (Rexfort and Mutterlose, 2006, 2009) by using Sepia officinalis as a modern analogue for the extinct belemnites. Several studies have shown systematic oxygen isotope variability in rostra that can be interpreted as seasonal signature (e.g. Stevens and Clayton, 1971; Zakharov et al., 2011) leading to the same estimates of lifespans for belemnites. Detailed descriptions of belemnite taxonomy and lithology of the studied sections are given by Meesook and Grant-Mackie (1995) and Challinor (1996, 1999, 2001, 2003).

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Table 1 Locations, lithostratigraphic units and ages of belemnite rostra taken for analysis. Biostratigraphic correlation to the Tethyan ammonite zones is from Stevens (1997) and numerical ages from Gradstein et al. (2012). Sample

Species

Fossil locality

Longitude

Latitude

Lithostratigraphic unit

Stage

Tethyan

Age

NZ0230

Hibolithes sp.

174.74796°E

37.42588°S

NZ0266A

174.86032°E

38.05676°S

174.98771°E

38.10355°S

NZ0256A,B

Hibolithes arkelli Hibolithes arkelli Belemnopsis sp.

R13/ f6650 R15/ f8628 R15/ f0001 R15/ f8012

175.00895°E

38.10937°S

Middle Tithonian Middle Tithonian Early Tithonian Kimmeridgian

Micracanthoceras ponti Semiformiceras fallauxi Semiformiceras semiforme Aulacostephanus eudoxus

147.72– 148.08 148.08– 149.87 149.87– 150.41 153.55– 153.96

NZ0261E

Belemnopsis sp.

174.78354°E

38.13128°S

Kimmeridgian

NZ0231

Belemnopsis sp.

R15/ f8560 R16/ f0230

174.76775°E

38.15416°S

Upper Puti Siltstone Lower Puti Siltstone Kinohaku Siltstone Waikiekie Tuffaceous Sandstone Waikutakuta Siltstone Ohineruru Formation

Crussoliceras divisum Ataxioceras hypselocyclum

154.47– 154.84 154.84– 155.60

NZ0242A

2.3. Methods The belemnite rostra were broken into several cm-sized pieces for further handling (see Figs. 1 and 2 for the altered specimen NZ0256B). Textural and geochemical screening methods were then applied to the fragments of the specimens. No chemical abrasion was performed for additional sample cleaning. Methodology and employed analytical machinery are summarized in Table 2 and explained in detail in Sections 2.3.1–2.3.6. 2.3.1. Scanning electron microscopy (SEM) Textural features of belemnite rostra were observed on fragments of >40 belemnite rostra from the studied sections including specimen NZ0256B and the other six specimens, for which geochemical analyses are reported. Fragments of about 1  1 mm size were broken off the rostra, mounted on an SEM stub and coated with gold. Images were taken

Fig. 1. Transverse section through a schematized belemnite rostrum with the position of the sampled fragment of NZ0256B indicated in pale grey.

Kimmeridgian

using the FEI Quanta 250 SEM of the Geological Museum in Copenhagen at 15 kV and in high vacuum mode. 2.3.2. Cathodoluminescence Two polished thick sections of the partly altered belemnite rostrum NZ0256B were investigated and photographed using the cold cathode Technosyn 8200 MK II Cathodoluminescence microscope at the GeoCenter Northern Bavaria, University of Erlangen. Operating conditions were 10– 15 kV and 350–500 mA. 2.3.3. Oxygen and carbon isotopes Carbonate samples from specimen NZ0256B were drilled from the rim to the apical line of the rostrum by progressive grinding along a dorsal section and a ventral section (Fig. 2), using a handheld drill with diamond coated steel bit ( 1 mm in diameter). The resulting linear resolution of the transects during progressive grinding is better than 200 lm/sample. The ventral profile was taken in the alveolar region, about one centimeter anterior to the former protoconch; the dorsal section in the stem region about two centimeters posterior of the former protoconch (Fig. 2). Analytical procedures of Spo¨tl and Vennemann (2003) were adopted for mass spectrometry. About 390 to 870 lg of prepared powders were transferred into 3.5 ml glass vials, sealed with rubber septa and the headspace was flushed with clean helium for 240 s to remove atmospheric contaminants. The samples were subsequently reacted with 0.05 ml of >100% H3PO4, equilibrated for >100 min at 70.0 °C, and the resultant carbon dioxide was analysed for d13C and d18O using the Micromass Iso Prime Isotope Ratio Mass Spectrometer at the University of Copenhagen. Weight dependent effects on the raw data were corrected by measuring a set of samples of the Copenhagen laboratory reference material (Carrara Marble: “LEO”), covering the weight ranges of the sample sets. The accuracy of the method was assessed by measuring the international standards NBS-18 and NBS-19, and the laboratory reference material from the Freie Universita¨t Berlin, covering a d13C and d18O range of 5.75& to +2.42& (PDB) and 23.2& to 2.03& (PDB), respectively. Our measured values deviate

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Fig. 2. (A) Transverse cut through the belemnite rostrum NZ0256B. The conical structure is the former phragmocone, now filled with siltstone with carbonate content of 20%. The shaded areas with roman numbers indicate the approximate extent of the discussed zones of differing degree of alteration. (B) Sampling areas for profiles of combined d13C, d18O and element ratios (See also Table. 5, Fig. 5). The arrows indicate the direction, in which the results are shown in Fig. 5. (C) Sample locations and results of 87 Sr/86Sr analyses (see Table 6 and Fig. 6 for details). (D) Sample locations and results of d7Li analyses. Sample powders from 1a and 1b were combined for a single analysis (see also Table 7 and Fig. 7).

by 0.05& or less from the adopted values for the in-house (Copenhagen and Berlin) and the accepted values for international reference materials apart from d18O values in NBS-18, for which a larger offset of 0.15& is observed (Table 3). The external reproducibility of the measurements over the protracted period of a year as determined by the standard deviation of the Copenhagen in-house reference material (LEO) is 0.08& for d13C and 0.18& for d18O (2 sd, n = 649). Uncertainties of measurements of reference materials are generally similar to the reported reproducibility (Table 3). 2.3.4. Elemental analyses Sr/Ca and Mn/Ca ratios were quantified using a Perkin Elmer Optima 7000 DV ICP-OES at the University of

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Copenhagen on the reacted carbonate aliquots that remained from the H3PO4 treatment for the stable isotope analyses (Coleman et al., 1989). Solutions were transferred into 15 ml centrifuge tubes and matrix matched with 2% nitric acid to a nominal calcium concentration of 25 lg/g. Measurements were calibrated against a set of synthetic solutions prepared from single element solutions. Limits of quantification (10 sd of baseline) were on average 1.3 lmol/mol for Sr/Ca and 5.5 lmol/mol for Mn/Ca at this concentration level of calcium. Potential contamination of the phosphoric acid was determined by checking the acid of each batch used for the reaction of the carbonates for the stable isotope preparation. The repeatability of the method was assessed by multiple measurements of the Copenhagen in-house reference materials LEO (Carrara marble; n = 25) and EGG (hen’s egg shell; n = 19), and the external reproducibility by the international reference materials JDo-1 (dolostone; n = 190) and JLs-1 (limestone; n = 100). The external reproducibility for Sr/Ca ratios of the reference materials was better than 2.4% (2 sd). The repeatability is 1.8% (2 sd) for EGG and 1.6% for LEO. The external reproducibility of the Mn/Ca ratios is better than 2.8% (2 sd) for JDo-1. The external reproducibility for JLs-1 is limited to 8.1% (2 sd) because of its low Mn/Ca ratio of 29 lmol/mol. The repeatability of the Mn/Ca ratio of LEO is 2.7% (2 sd) at a similar Mn/Ca ratio of 35 lmol/mol. Mn/Ca ratios of EGG were consistently below limits of quantification. The accuracy of the measurements was checked with JDo-1 and JLs-1 and all measured values are within 2% of the values published by Imai et al. (1996) (see Table 4). For the samples prepared for d7Li values, Li/Ca, Sr/Ca and Mn/Ca ratios were determined from a split of the dissolved carbonates on a Perkin Elmer Elan Quadrupole ICP-MS of the University of Oxford (refer to Section 2.3.6 for details on element purification and isotope ratio determination). Samples were matrix matched to 10 lg/g Ca, and were calibrated against a set of synthetic solutions made up from single element solutions. Accuracy and precision were assessed by repeated analyses of seawater and the international reference material JLs-1, as well as inhouse standards. Sample reproducibility of Li/Ca, Sr/Ca and Mn/Ca is 7% (2 sd, n = 7). 2.3.4.1. 87Sr/86Sr ratios. Six carbonate samples (1.4–7.4 mg carbonate) from the specimen NZ0256B were dissolved in 0.5 M HCl and, after drying on a hotplate at 80 °C, re-dissolved in 0.2 ml of 3 M nitric acid. The Sr was then purified using Sr-Spec resin (Horwitz et al., 1992). Eppendorfe pipette tips (1 ml), equipped with a filter, served as ion exchange columns. Purified Sr was eluted with de-ionized water and, after the addition of about 25 ll of 0.1 M H3PO4 and a subsequent drying on a hotplate at 80 °C, loaded on single rhenium filaments in 5 ll of Ta2O5– H3PO4–HF matrix. 87Sr/86Sr ratios were measured on the Sector VG 54 TIMS with 8 Faraday cups in dynamic multi-collection mode at 1250 to 1400 °C at the University of Copenhagen. Signal intensities were recorded for masses 85, 86, 87 and 88. 87Rb interferences on the 87Sr signals were monitored by 85Rb signal intensities and, if necessary,

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Table 2 Analytical procedures and equipment. Analytical step

Employed machinery

Machine type

Reference materials

Sample preparation Microtexture control Microtexture control d13C, d18O values Element ratios of C and O isotope samples 87 Sr/86Sr ratios Element ratios of 87Sr/86Sr samples d7Li values Element ratios of d7Li samples

Handheld drill Scanning electron microscope Cathodoluminescence microscope Gas source IRMS ICP-OES TIMS ICP-OES MC-ICP-MS Q-ICP-MS

Messner emtronic MIK 20 FEI Quanta 250 Technosyn 8200 MK II Micromass IsoPrime Perkin Elmer Optima 7000 DV Sector VG 54 Perkin Elmer Optima 7000 DV Nu Plasma HR MC-ICP-MS Perkin Elmer Elan Q-ICP-MS

– – – NBS 18, NBS 19a JLs-1, JDo-1 NISTSRM-987 JLs-1, JDo-1 Seawater JLs-1

a

In-house materials of the Freie Universita¨t Berlin FUB CAM, FUB KKS, FUB LM were measured for additional control; see Table 3 for results.

Table 3 Results for analyses of carbonate reference materials in 2011 with 2 sd uncertainty. Bold values are adopted values for materials KU LEO (Carrara Marble), FUB LM (Laaser Marble), FUB CAM (Carrara Marble) and FUB KKS (Kaiserstuhl Carbonatite) and accepted values for NBS-18 and NBS-19. Material

d13C

d18O

d13C

d18O

n

& PDB & PDB & PDB 2 sd & PDB 2 sd KU LEO FUB CAM FUB LM FUB KKS NBS-18 NBS-19

1.96 2.42 1.51 5.75 5.01 1.95

1.93 2.03 5.17 22.40 23.20 2.20

1.94 2.40 1.46 5.74 4.97 1.94

0.08 0.08 0.09 0.06 0.06 0.12

1.91 2.05 5.18 22.43 23.04 2.18

0.18 0.18 0.12 0.29 0.15 0.23

15 33 31 33 10 30

Table 4 Element ratios in international reference materials (JDO-1, JLs-1) and laboratory reference materials (LEO, EGG). Bold values are calculated from concentrations given in Imai et al. (1996) and taken for estimation of the measurement accuracy. Material

Sr/Ca

Mn/Ca Sr/Ca 2 sd Mn/Ca 2 sd n (lmol/mol) (lmol/mol)

JDo-1 JLs-1 LEO EGG

218.6 342.7

152.9 30.0

222.5 343.2 180.6 1169

4.8 8.1 2.9 21

150.5 29.5 35.0 d.l.

4.2 2.4 0.9 –

190 100 25 19

corrected for by the natural 87Rb/85Rb ratio of 0.3857. 87 Rb intensities were systematically below 150 ppm of the 87 total beam intensities. These levels of interferences make the 87Sr/86Sr ratios insensitive to Rb interference corrections beyond the sixth digit of the values. Instrumental mass bias was monitored by the ratio 86Sr/88Sr for which respective corrections on the 87Sr/86Sr values were performed using 0.1194 as a natural value. For all samples, 12 blocks with 10 cycles of about 40 s were measured. Ratios deviating more than 2 sd from the mean were rejected. Signal intensities were kept at 1–3 V for mass 88 optimize signal to noise ratio and analysis times. The repeatability of the method was determined by a whole procedure replication experiment using a marly chalk which gave the value of 0.707844 ± 0.000018 (2 sd, n = 18). The uncertainty of

this value represents the sum of the contributions of the primary heterogeneity of the sample material and all lab procedures, but not the long-term machine drift, as all analyses were performed within 48 h. Accuracy and reproducibility of the measurements were checked by multiple measurements of the reference material NISTSRM-987 giving a mean 87Sr/86Sr ratio of 0.710243 ± 0.000022 (2 sd, n = 15) which is within uncertainty of the certified value of 0.71034 ± 0.00026 and the value of 0.710248 adopted by McArthur et al. (2001). 2.3.5. d7Li values and element ratios Fourteen samples were prepared for Li isotope analyses: five from zones of differing degrees of alteration from the specimen NZ0256B, two from the sediment in the alveolar infill of this belemnite, and the remaining seven samples from additional six rostra. Aliquots between 75 and 130 mg of carbonate powders were drilled from the rostra using a handheld drill with tungsten bit. For all specimens other than NZ0256B material was drilled from clean, translucent parts of the rostra, and apical line as well as rims were avoided. Contamination of the samples with detrital Li is excluded, because samples were prepared from pure calcite from the interior part of the rostra. Samples were dissolved with 2 M HCl and dried down on a hotplate at 130 °C. In order to assess the isotopic composition of Li released from the sedimentary matrix by acid leaching, 12 mg of the calcareous siltstone were drilled out of the alveolar infill of the specimen NZ0256B. The extraction was done in a two step process with the first step releasing Li from the carbonate phase and weakly bound Li from the silicates, while during the second step stronger bound Li was released from the silicates. In the first step, the powdered material was leached with 0.2 M HCl at room temperature for 10 min and the resulting solution was centrifuged and the supernatant transferred into a Teflon beaker. The remaining silicate material was then leached with concentrated HCl at 130 °C for 2 h, the resulting solution centrifuged and the supernatant transferred to a Teflon beaker. Both solutions were then dried on a hotplate at 130 °C. For all samples the resulting dry residues were taken up by 0.2 M HCl in an approximate ratio of 2 ll per milligram calcium chloride and 100 ll of solution were taken for Li purification. Purification was performed in two steps using

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0.2 M HCl as an elutent and AG 50 W X12 as ion exchange resin (Pogge von Strandmann et al., 2011). In the first step Savillex Teflon columns with about 8.5 cm resin bed were used and for the second step smaller Teflon tubes with high aspect ratio and about 1 ml of resin were taken. Splits were collected before and after the Li window, and analysed to assess the loss of Li during the purification process. Concentrations of purified Li and splits were estimated using beam intensities on a Nu Plasma HR MC-ICP-MS at the University of Oxford compared to a reference solution of known concentration. Li concentrations in the splits consistently amounted to less than 0.3& of Li found in the samples. Li concentrations of the samples were adjusted to 20 ng/g (15 pA intensity on mass 7 for a 75 ll/min uptake rate). Measurements were done against the international Li isotope reference L-SVEC (Flesch et al., 1973), a lithium carbonate provided by the IAEA, using a standard sample bracketing protocol. Baseline intensities were measured after every second L-SVEC reference and subtracted from the raw intensities. Means of triplicate analyses with the double standard deviation are reported. The accuracy and reproducibility of the method were determined by multiple treatments and measurements of modern seawater, which gave a value of 31.1 ± 0.6& (2 sd, n = 9). This value is in complete agreement with the value of 31.0 ± 0.1& (2 err, n = 31) reported by Millot et al. (2004). 3. RESULTS 3.1. SEM General textural features of belemnite rostra are shown in Fig. 3A–C. Fractures of well-preserved belemnite rostra result in bent crystal surfaces with step-like appearance (Fig. 3A); observed thinning of single crystals is caused by concentric zones of void space around the apical line (Fig. 3B). A few per cent of such holes that are perhaps produced by the decay of original organic matter within the rostrum (Slen, 1989) have been found to contain iron sulphide (McArthur et al., 2007). The high alteration potential of the apical line and adjacent growth bands that are often found to be diagenetically overprinted (e.g. McArthur et al., 2000, 2007; Rosales et al., 2001; Bailey et al., 2003; Benito and Reolid, 2012; Li et al., 2012) is indicated by the fine-grained nature of the calcite crystal close to the apical line, which results in a large surface area and potential permeability for fluids (Fig. 3C). Dissolution and re-crystallization of the rostral calcite of specimen NZ0256B studied here is clearly visible in the form of sets of aligned voids and irregularly shaped bulges (Fig. 3D). 3.2. Cathodoluminescence (CL) Orange to red luminescence is observed in the section perpendicular to the apical line, in concentric bands (strongest in the first thicker band around the apical zone) and along fracture lines (Fig. 4A). In the section parallel to the apical line orange to red luminescence is visible in a zone adjacent to the sedimentary infill in the alveolar region and in several bands parallel to the growth bands (Fig. 4B).

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Fig. 3. SEM images of belemnite shell material. (A) Pristine fracture pattern in a Tithonian belemnite from Port Waikato, New Zealand for comparison. (B) Growth-ring-parallel zones of aligned cavities cutting through single crystals of calcite in a Kimmeridgian belemnite from Kawhia, New Zealand (this structure was not detected on the fragments of the altered specimen investigated by SEM but is indicated by the dark lines in Fig. 1; see also discussion). (C) Radial structure of calcite crystals growing outwards from the apical line in a Tithonian belemnite from Port Waikato, New Zealand. (D) Voids in the carbonate are aligned along preferential zones of dissolution and recrystallisation is visible in the form of bulges in the calcite (upper left). Fragments pictured in A–C are from belemnite rostra that have not been analysed for geochemical composition. The fragment pictured in D is derived from the rostrum NZ0256B.

The features of both sections can be correlated and put into a schematic representation of luminescent bands (Fig. 4C) that appears to traverse the rostrum parallel to the growth bands (Fig. 4B). 3.3. Isotopic signatures and trace elements The dorsal and ventral stable isotope/elemental sections of specimen NZ0256B show common characteristics, although Mn and d13C trends are generally less pronounced in the dorsal section (Fig. 5, Table 5). Four zones could be delineated in both transects. Starting from the central part of the rostrum in zone I d18O and the Sr/Ca ratios increase distinctly from the lowest values observed in both sections whereas Mn/Ca decreases (somewhat more steeply within the luminescent band labelled “a” in Fig. 4C); the d13C values in this interval are relatively stable. Further away from the apical line/alveolar infill in zone II, the Mn/Ca ratios increase as a pronounced peak in the ventral section and less clear in the dorsal section. These peaks in Mn/Ca ratios coincide with negative spikes in the Sr/Ca ratios and the d13C and d18O values, and with a set of luminescent bands labelled as “b” in Fig. 4C. Zone III is characterized by a series of values with relatively high Sr/Ca and low Mn/Ca ratios, and heavy, but slightly decreasing d13C and d18O values. The outermost zone IV is characterised by a Mn enrichment (ventral: 4.6 mmol/mol), including the highest

86

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value measured (dorsal: outermost part not sampled). In general, Mn enrichments (Fig. 5) coincide with the luminescence, first in bands around the apical zone, and between the peak in zone II and the brightest line in Fig. 4B. It is also notable that there is no visible geochemical fingerprint of the less prominent luminescent bands in the elemental record. 3.3.1. 87Sr/86Sr ratios 87 Sr/86Sr ratios were determined for five different subsamples of specimen NZ0256B and also for the sedimentary infill in the area of the former phragmocone (Fig. 2). The results of the belemnite calcite vary from 0.706724 ± 0.000010, close to the apical line, to 0.706928 ± 0.000011 in the rim and apical line areas (Figs. 2 and 6, Table 6); a trend that is broadly anti-correlated with the Sr/Ca ratios. The 87Sr/86Sr ratio of 0.707953 ± 0.000011 of the sedimentary infill is much more radiogenic. 3.4. Lithium The d7Li values of the different areas in specimen NZ0256B vary from +23.6 ± 0.5 to +40.1 ± 0.0& (Fig. 7, Table 7). The heaviest values were measured for Zone I, an intermediate value of 30.2 ± 0.8& was observed for Zone II, and the lightest values were measured for Zone III. The lowest Li/Ca ratios were detected for Zone III; intermediate and highest Li concentrations occur in the sample from Zone II and in the sample of the apical line, respectively (Fig. 7A). The Li isotopes and Li concentrations of the leachates from the two alveolar infill samples are 4.6 ± 0.3& and 3 lg/g, and 4.9 ± 0.9& and 29 lg/ g, respectively, for the leachings in 0.2 M HCl and concentrated HCl. The d7Li values of the 6 additional rostra are

A: ventral, anterior III

4

1.2

3 0.8

2 0.4

1

4 2 0 -2 -4 -6 -8

13

δ C

Mn/Ca

-10

δ18O

Sr/Ca

-12 alveolar infill

rim

B: dorsal, posterior Ib

II

III

IV

0.6 0.5

1.2

0.4

0.8

0.3 0.2

0.4

0.1

0.0

0

0.0

δ13C, δ18O ‰ PDB

1.6

5

IV

Sr/Ca (mmol/mol)

II

Mn/Ca (mmol/mol)

Ia

δ13C, δ18O ‰ PDB

Sr/Ca (mmol/mol)

1.6

Mn/Ca (mmol/mol)

Fig. 4. Cathodoluminescence images of (A) the dorsal section and (B) the ventral section of sample NZ0256B from the core to the rim. The patchy luminescence to the left is due to silicate in the alveolar infill. (C) Enrichments in manganese are visible as orange or red luminescence and are shown generalized from both dorsal and ventral sections. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

0.0

4 2 0 -2 -4 -6 -8

δ13C

Mn/Ca

-10

δ18O

Sr/Ca

-12 apical line

rim

Fig. 5. Trace element and stable isotope profiles through the rostrum NZ0256B. Both profiles are shown from the core (left) to the rim (right) of the rostrum. (A) ventral profile; length 5 mm, (B) dorsal profile, length 6 mm (see also Fig. 2 for locations). The width of the coloured bands represents the 2 sd uncertainties of the values as determined by the external reproducibility of the analytical methods. Note the difference in absolute values of Mn/Ca ratios between both profiles. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

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87

Table 5 Stable isotope and trace element results for the two profiles through the rostrum. Samples 1–42: dorsal profile; samples 43–69: ventral profile. Traces are from the rim towards the apical line and phragmocone, respectively. Uncertainties are 0.08& (2 sd) for d13C values, 0.18& (2 sd) for d18O values, 2.4% (2 rsd) for Sr/Ca and 2.8% (2 rsd) for Mn/Ca ratios. Sample

Sub

d13C (& PDB)

NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B

1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24 25 26 27 28 29 30 31 32 33 34 35 36 37 38 39 40 41 42 43 44 45 46 47 48 49 50 51 52 53 54 55 56 57 58 59 60 61

0.14 0.17 0.03 0.25 0.80 0.53 0.29 0.41 0.55 0.63 0.50 0.67 1.00 1.81 2.12 2.26 2.06 1.35 0.01 0.12 0.43 0.05 0.53 0.17 0.26 0.59 0.92 0.58 0.28 0.28 0.24 0.19 0.45 0.40 0.41 0.35 0.03 0.02 0.18 0.29 0.39 0.10 2.66 0.24 0.94 0.91 0.41 0.13 0.19 0.02 0.19 0.08 1.01 1.55 1.65 0.41 1.84 1.65 1.41 0.90 0.66

d18O (& PDB) 2.73 3.31 3.25 3.26 2.94 2.88 2.93 2.77 2.72 2.27 2.37 2.52 2.87 2.07 1.87 2.63 3.03 3.08 4.42 3.20 2.64 3.45 2.87 3.19 3.40 2.95 2.99 3.42 3.51 3.38 3.80 4.67 5.07 5.49 6.30 7.13 8.18 9.28 10.12 10.66 10.94 11.95 3.29 2.49 2.88 2.64 2.66 2.32 2.02 2.27 1.67 2.51 1.72 1.88 3.33 7.40 8.15 7.32 6.74 5.96 5.67

Sr/Ca (mmol/mol) 1.12 1.15 1.19 1.26 1.40 1.40 1.40 1.44 1.47 1.41 1.36 1.36 1.31 1.30 1.32 1.28 1.20 1.18 1.12 1.18 1.23 1.20 1.27 1.23 1.20 1.31 1.35 1.24 1.27 1.28 1.24 1.16 1.17 1.11 1.05 0.97 0.90 0.82 0.76 0.73 0.73 0.67 1.59 1.44 1.31 1.34 1.42 1.46 1.47 1.41 1.42 1.40 1.48 1.53 1.50 1.31 1.24 1.20 1.20 1.24 1.30

Mn/Ca (mmol/mol) 0.18 0.19 0.27 0.24 0.13 0.11 0.13 0.11 0.13 0.12 0.10 0.12 0.12 0.09 0.10 0.10 0.14 0.15 0.21 0.14 0.11 0.15 0.11 0.12 0.14 0.15 0.08 0.16 0.17 0.17 0.15 0.17 0.16 0.22 0.18 0.20 0.21 0.24 0.31 0.41 0.43 0.53 4.58 0.35 0.40 0.34 0.12 0.08 0.09 0.12 0.07 0.22 0.03 0.32 0.15 0.66 2.38 1.25 0.63 0.16 0.21 (continued on next page)

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Table 5 (continued) Sample

Sub

NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B

62 63 64 65 66 67 68 69

0.7080

d13C (& PDB)

d18O (& PDB)

0.66 0.21 0.09 0.40 0.56 0.73 0.43 1.26

6.23 6.74 6.30 8.02 8.55 9.23 10.05 10.96

Sr/86Sr 87

1.27 1.24 1.24 1.16 1.05 1.01 0.91 0.90

0.13 0.13 0.14 0.21 0.22 0.55 0.67 2.61

4.1. Alteration

0.7076 0.7072 New Zealand belemnites CE (Gröcke et al., 2003) 0.7068 Phanerozoic minimum 0.7064 0.0

Mn/Ca (mmol/mol)

4. DISCUSSION

F

2 sd

Sr/Ca (mmol/mol)

0.5

D B

A

1.0 1.5 Sr/Ca (mmol/mol)

2.0

Fig. 6. Strontium isotope results plotted against Sr/Ca ratios in the samples of NZ0256B. Phanerozoic minimum 87Sr/86Sr ratios (McArthur et al., 2001) and bracketing ratios for coeval belemnites from New Zealand (Gro¨cke et al., 2003) are shown for reference (see also Fig. 2 for locations). 2 sd uncertainty of the Sr/Ca ratios and 87Sr/86Sr ratios is
Table 6 Results for strontium isotope samples with double standard error. (A) close to apical line, (B) close to alveolar infill, (C–E) intermediate layers, (F) alveolar infill. See also Fig. 2 for sample positions. Uncertainties are 2.4% (2 rsd) for Sr/Ca measurements and 2.8% (2 rsd) for Mn/Ca ratios. Specimen

Sample Sr/Ca Mn/Ca (mmol/mol) (mmol/mol)

87

NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B

A B C D E F

0.706724 0.706744 0.706928 0.706893 0.706926 0.707953

0.95 0.82 1.25 1.22 1.31 0.58

0.55 0.49 0.15 0.23 0.29 15.22

Sr/86Sr 2err 0.000010 0.000010 0.000011 0.000010 0.000008 0.000011

between +25.8& and +31.5& and plot between the heaviest and the intermediate values of specimen NZ0256B (Fig. 7, Table 7). The Li/Ca ratios of the 6 additional belemnites vary from 9 to 97 lmol/mol with lower Li/Ca ratios in Hibolithes compared to Belemnopsis (Tables 1 and 7). The d7Li values of all analysed samples with Mn/ Ca ratios <0.18 mmol/mol, which is a commonly employed limit for well preserved belemnite calcite (e.g. Price and Mutterlose, 2004), yield an average of 27.0 ± 1.1& (2 sd, n = 5); samples with higher Mn/Ca ratios show significantly more variability and trend towards higher d7Li values (Fig. 7B, Table 7).

4.1.1. Geochemical trends of alteration The chemical and isotopic composition of biogenic carbonates are not in equilibrium with diagenetic fluids and during alteration some geochemical proxies evolve along predictable trends (e.g. Brand and Veizer, 1981). Manganese is generally enriched in carbonates that have been re-crystallized, because Mn has distribution coefficients that are consistently >1 during co-precipitation with calcite despite a strong precipitation rate dependence of the absolute value on the distribution coefficient (Mucci, 1988; Pingitore et al., 1988). Mn concentrations are therefore often used to identify altered calcite fossils (e.g. Veizer, 1974; Brand and Veizer, 1980; Korte et al., 2003). Different Mn/ Ca thresholds have been suggested to identify pristine belemnites, for example <0.09 mmol/mol (Malkoc and Mutterlose, 2010) or <0.18 mmol/mol (Price and Mutterlose, 2004; Price and Rogov, 2009; Nunn and Price, 2010). Some indication exists that primary enrichments of Mn in shell calcite can occur (see data from McArthur et al., 2000; Freitas et al., 2006; Korte and Hesselbo, 2011). Such enrichments may be related to higher concentrations of dissolved Mn in dysoxic seawater, but the possible extent of this process is not yet well constrained. Cathodoluminescence (CL) can be used as a tool to qualitatively trace Mn enrichments (e.g. Slen and Karstang, 1989; Benito and Reolid, 2012), because luminescence in fossil calcites is primarily caused by enrichments of Mn2+ (Machel et al., 1991). CL photographs can therefore be used to detect altered shell material with high spatial resolution. Cautious interpretation of CL is advisable, because unaltered, luminescent shell material has been observed in modern brachiopods (Barbin and Gaspard, 1995) and cases of non-luminescent, altered shell material have been reported (Rush and Chafetz, 1990; Qing and Veizer, 1994). The concentration of Sr in biogenic LMC is usually higher than in non-biogenic calcite (Carpenter and Lohmann, 1992) and decreasing Sr concentrations in fossil calcites are a ubiquitous feature of diagenesis (e.g. Brand and Veizer, 1980, 1981). This is especially evident for belemnites that usually have higher Sr concentrations than other coeval calcite fossils (see Fig. 4 of Korte and Hesselbo, 2011). A depletion of Sr during diagenesis occurs, because the distribution coefficient of Sr between inorganic calcite and water is low (usually <0.1), possibly due to comparatively low precipitation rates of abiogenic calcites (Carpen-

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A 45 Ia

40 2 sd δ7Li ‰

Ib

35 II

30

III

25 III

20 0

20

40 60 Li/Ca (µmol/mol)

B 45

δ7Li ‰

40

Ia

2 sd

80

100

Ib

35 likely altered 30

III

II

25 III

20 0.0

0.2 0.4 Mn/Ca (mmol/mol)

C 45

Ia

δ7Li ‰

40 2 sd 35

Ib

20 0.5

monospecific fossil groups (Veizer, 1974; Steuber and Veizer, 2002; Korte and Hesselbo, 2011; Li et al., 2012). This variability has been related to vital effects such as growth rate (Lorens, 1981; Shen et al., 2001; see discussion in Korte and Hesselbo, 2011) and changes in the chemical composition of seawater (Steuber and Veizer, 2002; Lear et al., 2003). The application of absolute Sr/Ca ratios as markers for diagenesis therefore requires knowledge about the primary composition of fossil calcites from closely related species of the same age. Carbon and oxygen in diagenetically affected calcite generally tend to be depleted in the heavy isotope (see Veizer, 1983a and Marshall, 1992 for reviews) and d13C values are usually more resistant to diagenesis than d18O values (e.g. Banner and Hanson, 1990). This is because diagenetic fluids have a much lower C/O ratio than carbonates and part of the diagenetic carbon may be derived from the dissolution of ambient carbonates with a carbon isotope signal similar to the fossil shell material (Veizer, 1983a). 87 Sr/86Sr ratios of well-preserved marine biogenic LMC reflect the original Sr isotope composition of the seawater in which it was precipitated (e.g. Veizer et al., 1999). The 87Sr/86Sr ratios in marine settings are uniform due to the long residence time of Sr of >106 years (Li, 1982) in the oceans. Resolvable heterogeneity and/or offsets in the calcite of supposedly stenohaline organisms from coeval seawater 87Sr/86Sr are indicative of diagenesis, because freshwater derived Sr only in rare cases measurably affects the marine 87 Sr/86Sr ratio at salinity levels P30 psu (e.g. McArthur, 1994). Stable isotope fractionation can be excluded as a source of variability in the 87Sr/86Sr ratios, when these are normalized to an 86Sr/88Sr ratio of 0.1194 (Nier, 1938; Elderfield, 1986, see Section 2.3.5 for methodology).

II

30 25

0.6

89

III Hibolithes Belemnopsis NZ0256B

III

1.0 1.5 Sr/Ca (mmol/mol)

2.0

Fig. 7. Li isotope results plotted against element/Ca ratios in the samples of the partly altered rostrum NZ0256B (green diamonds) with zone numbers, and other Late Jurassic samples of Belemnopsis and Hibolithes (dark green triangles). For descriptions of the zones see Section 3.3.; Fig. 5. (A) d7Li values plotted against Li/Ca ratios. (B) d7Li values plotted against Mn/Ca ratios. The stippled vertical line at 0.18 mmol/mol delineates the commonly used acceptable upper limit for good preservation. (C) d7Li values plotted against Sr/Ca ratios. Note that error bars are only shown when larger than symbol size. Reproducibility of d7Li measurements is shown in the upper left corner of the plots. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

ter and Lohmann, 1992; Tang et al., 2008; DePaolo, 2011). Qualitative differences of Sr/Ca ratios in single specimens can therefore often be used to assess the preservation state of the shell material. The use of absolute Sr/Ca ratios for this purpose is complicated by large temporal and spatial differences in the primary Sr/Ca ratios in different and even

4.1.2. Alteration trends in specimen NZ0256B In many of the samples of specimen NZ0256B Mn/Ca ratios (up to 4.6 mmol/mol; Fig. 5, Table 5) exceed the suggested upper limits for good preservation and only 6 of 69 samples meet the strict criterion of Mn/Ca <0.09 mmol/ mol. Luminescence is evident in concentric bands around the apical line and in the sedimentary infill in the alveolar region (Fig. 4). The zones showing the brightest luminosity (a and b in Fig. 4C) can be correlated with Mn enrichments in zone I and the Mn spike in zone II. The strong enrichment of Mn at the rim of the rostrum (Fig. 5A) is not visible in the CL mosaics because of lacking spatial coverage. Often, thin luminescent bands occur where no clear spikes in Mn/Ca ratios are identified. This might be, because Mn enrichments are related to very narrow bands that could not be resolved by the sampling applied. Because diagenetically incorporated Mn was supplied by fluids, accumulations of this element are assumed to take place in areas of higher permeability. The declining trend of Mn/Ca perpendicular to the apical line and the Mn enrichment in zone II (Fig. 5) are thus likely caused by the fluid penetrating into the rostrum along crystal boundaries and, moreover, along zones with larger cavities (visible in Fig. 3B). In zones I and II of the studied belemnite, the Sr/Ca ratios and the d18O values correlate negatively with the Mn/ Ca ratios, and this indicates that the depletion of Sr is

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Table 7 Results for lithium isotope samples. For sample locations on sample NZ0256B refer also to Section 3.3, Fig. 2. Uncertainty of the element ratios is 7% (2 rsd). Specimen

Sample

NZ0230 NZ0266A NZ0242A NZ0256A NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0256B NZ0261E NZ0261E NZ0231 a b

Sample location

1 2 3 4 5 5 repeat anal. 6 (0.2 M HCl) 7 (conc. HCl) a b

Zone II Zone Ib Zone Ia Zone III Zone III Zone III Alveolar infill Alveolar infill

Sr/Ca (mmol/mol)

Mn/Ca (mmol/mol)

Li/Ca (lmol/mol)

d7Li (& L-SVEC)

2 sd

1.52 1.67 1.79 1.37 1.14 0.84 0.95 1.34 1.13

0.15 0.15 0.12 0.25 0.29 0.50 0.28 0.18 0.23

9 16 21 20 28 19 38 17 18

26.1 27.3 27.3 25.8 30.2 39.0 40.1 27.3 23.6 23.6 4.6 4.9 31.5 31.1 27.1

0.7 0.2 0.1 0.2 0.8 0.2 0.0 0.8 0.5 0.5 0.3 0.9 0.5 1.2 0.2

A B

1.36 1.33 1.01

0.24 0.34 0.09

97 73 46

Li concentration = 3 lg per gram sediment. Li concentration = 29 lg per gram sediment.

caused by diagenetic alteration (Brand and Veizer, 1980, 1981). The Sr/Ca and d18O values of zone I in the dorsal section show an excellent linear correlation (r2 of 0.995; Fig. 8). This linear correlation can be explained by a single phase of re-crystallisation at likely elevated temperatures, involving a possibly 18O-depleted fluid that maintained low Sr/Ca ratios. The negative correlation for oxygen isotopes and Sr/Ca with the Mn/Ca ratios, on the other hand, is non-linear (Fig. 8). Such non-linear trends are expected because of the compatibility of Mn in calcite (Mucci, 1988; Pingitore et al., 1988). In a (semi-) closed system the preferential uptake of Mn into calcite causes a rapid depletion of Mn in the diagenetic fluid during progressive interaction with calcite (Veizer, 1983a; Al-Aasm and Veizer, 1986a; Rimstidt et al., 1998; Rosales et al., 2001; and see Fig. 5 for resulting spatial trends). From the four zones

0.6 0.5

1.2

0.4 0.8 0.4

0.3 y = 0.073 x + 1.5 r2 = 0.995

0.0 -14

-12

-10 -8 -6 δ18O ‰ PDB

0.2

Mn/Ca (mmol/mol)

Sr/Ca (mmol/mol)

1.6

Sr/Ca 0.1 Mn/Ca 0.0 -4 -2

Fig. 8. Mn and Sr cross-plots with d18O for zone I of the dorsal section of sample NZ0256B. The excellent linear correlation of d18O with Sr/Ca is taken as an indication that dissolution and reprecipitation of calcite happened in one stage. The non-linear behaviour of Mn/Ca is related to the progressive depletion of manganese in the diagenetic fluid during interaction with the fossil calcite (see Section 4.1.2 for further explanation). Error bars are smaller than symbol size in all cases.

with differing degrees of alteration described in Section 3.3, zone III (Fig. 5) displays the lowest Mn/Ca ratios. The Sr/ Ca ratios of 1.28 to 1.52 mmol/mol (n = 20) for this zone fall entirely in the range of Sr/Ca ratios reported for Kimmeridgian and Tithonian (157–145 Ma) Belemnopsis and Hibolithes of New Zealand (Podlaha et al., 1998) for which Sr/Ca ratios of 0.98 to 1.61 mmol/mol (n = 18) were found. d13C values in the two profiles through specimen NZ0256B are dissimilar throughout zone I and most of zone II (Fig. 5) and the ventral section is consistently more depleted in 13C than the dorsal one. Furthermore, in this part of the profiles the d13C values of the dorsal section vary by only 0.9& (Fig. 5). Note that well-preserved belemnites are characterized by distinct fluctuations of carbon isotopes of several permil reflecting palaeoenvironmental changes and probably metabolic effects (e.g. Dutton et al., 2007; McArthur et al., 2007; Wierzbowski and Joachimski, 2009). A positive d13C excursion across zone III in both profiles with similar absolute values together with the lowest measured Mn/Ca ratios in this zone suggest an area with a nearly primary d13C signature (Fig. 5, Table 5). The range of observed d13C values from 0.41& to +2.26& (n = 20) is larger than the +0.11& to +1.00& reported by Stevens and Clayton (1971) (n = 5) for Belemnopsis from the same fossil locality but the mean value of our results (+0.68&) is very close to their mean (+0.49&). Larger scatter is to be expected from the larger number of samples and the much smaller sample sizes taken for preparation (1/10 of the sample size analysed by Stevens and Clayton, 1971). Lighter d13C values in the altered part of the ventral section suggest that most of the signal of external isotopically light carbon supplied from the fluid was taken up in the anterior part of the rostrum. The fraction of carbon in the fluid derived from partial dissolution of the rostrum progressively increased as the fluid migrated to the posterior part. This is consistent with alteration in a semi-enclosed system (see Veizer, 1983a). The sample, for which elemental ratios indicate best preservation (Table 6) has a 87Sr/86Sr ratio of

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0.706928 ± 0.000011 (sample C, Table 6, Figs. 2 and 6). This ratio is in complete agreement with the range of values reported by Gro¨cke et al. (2003) for Kimmeridgian belemnites from New Zealand that are stratigraphically closest to specimen NZ0256B (Fig. 6). Diagenetic alteration in the studied belemnite is the only plausible explanation for those 87 Sr/86Sr values that differ significantly from the Phanerozoic reference curve. This is further evidenced by the high Sr isotope variability within the fossil that cannot be acquired in an inferred fully marine setting (Figs. 2 and 6). It is notable that the 87Sr/86Sr ratios in the most altered parts of the rostrum near the apical line and close to the alveolar infill – reaching 0.706724 ± 0.000022 – are lower than in the better preserved parts (Figs. 2 and 6, Table 6). Diagenetic trends towards less radiogenic 87Sr/86Sr ratios were hypothesized also for Jurassic belemnites from New Zealand (Podlaha et al., 1998) and from Canada (Gro¨cke et al., 2007), and similar results have been reported for brachiopods (Brand, 1991). The low values observed in the present study are not only less radiogenic than the inferred seawater 87Sr/86Sr during the lifetime of the belemnite, but are about 0.00012 less radiogenic (corresponding to >10 sd. of the measurement uncertainty) than the Phanerozoic minimum of about 0.70685 in the Oxfordian around 162 Ma ago (Fig. 6; see also McArthur et al., 2001 and Jenkyns et al., 2002). Diagenetic fluids are often enriched in 87Sr because of the decay of 87Rb from detrital silicates (Veizer, 1983b), but a shift to lower 87Sr/86Sr ratios in diagenetically altered materials is also possible if Sr is supplied from young igneous rocks, such as basalts (Richter and DePaolo, 1988; Gao and Land, 1991). The low 87Sr/86Sr ratios in the altered parts of the studied rostrum are most probably a product of less radiogenic Sr derived from the rock matrix of the Waikiekie Tuffaceous Sandstone from which specimen NZ0256B was collected, and 80% of this sandstone comprises sediment derived from magmatic rocks of a volcanic arc setting with low Sr isotope values of about 0.704 (Roser et al., 2002). To summarize, the combination of element ratios and CL, d13C and d18O values as well as 87Sr/86Sr (Figs. 2, 4 and 5) provide detailed information of diagenetic patterns in belemnites, and permit the identification of clearly altered parts and even diagenetic pathways. From the multi-proxy-assessment of preservation state it can be concluded that geochemical data derived from the least altered zone III of specimen NZ0256B are most likely to provide meaningful constraints on palaeo-environmental conditions. 4.1.3. Constraints on the diagenetic process The observed geochemical trends in specimen NZ0256B indicate that the diagenetic fluid entered the rostrum from the anterior part along zones of higher permeability causing especially pronounced alteration in the area adjacent to the apical line. Parts of the rostrum exhibiting only minor signs of alteration are present, as attested by relatively high Sr/ Ca ratios, heavy oxygen isotope ratios and 87Sr/86Sr ratios compatible with values for coeval seawater (Fig. 6). These observations argue for partial preservation of the original calcite and its primary geochemical signature. The relatively

91

low Mn/Ca ratios in some clearly diagenetically overprinted areas (Fig. 5) point towards a generally low water–rock interaction ratio and to a low fluid flow, supplying too little Mn to attain equilibrium concentrations in all parts of the altered calcite during re-crystallisation. Dissolution and re-precipitation of calcite can best explain the observed co-variation of the measured geochemical proxies in the altered parts of the rostrum. The high degree of correlation of Sr/Ca, Mn/Ca and d18O values in zone I (Fig. 8) makes it likely that the effects of diagenesis on the carbonate were largely caused by a single event rather than by subsequent pulses. The low d18O values (down to 11.9&) in the same parts of the rostrum (Fig. 5) suggest elevated temperatures at the time of alteration and/or the involvement of meteoric, 18O depleted waters (e.g. Veizer, 1983a). A possible scenario is that the major phase of diagenesis took place during the Late Cretaceous (100 Ma), when rifting of the Tasman Sea commenced (Kamp and Liddell, 2000). The extensional regime of this rifting, involving up to 4 km of Cretaceous cover sediments (Kamp and Liddell, 2000), would have provided the necessary pathways for fluids that were relatively hot (>60 °C), depending on the former geothermal gradient. A diagenetic event in the Late Cretaceous is also compatible with the alteration towards unradiogenic 87Sr/86Sr ratios near the apical line and close to the alveolar infill, since the radiogenic ingrowth of 87Rb was probably not sufficient yet to raise the 87Sr/86Sr ratios of the volcaniclastic Sr source above those of the primary fossil carbonate. The much more radiogenic 87Sr/86Sr of the carbonate in the alveolar infill (>0.7079) is most probably a signature of a later, less pervasive phase of fluid movement that is not captured by the carbonate of the rostrum, at least not within the resolution achieved in this study. 4.2. Alteration of d7Li values and implication for d7Li in ancient samples The constraints on the diagenetic setting described above provide a framework for the interpretation of the highly variable d7Li values of the belemnite rostrum NZ056B (Fig. 7). The observed variation could have been controlled by the following different processes: (i) Primary Li isotope fractionation between liquid and calcite during precipitation of the rostrum. (ii) Continuous diffusion processes, Li is potentially highly mobile in the calcite lattice, because it probably resides in interstitial positions rather than on the Ca2+ position (Okumura and Kitano, 1986; Marriott et al., 2004a). (iii) Alteration of material with a primary Li isotope signature that is relatively constant during the dissolution and re-precipitation of rostral calcite also might lead to the observed variability. Each of these possibilities is discussed below. (i) Primary fractionation would suggest that biological processes can amount to a range of d7Li values of 16.5& in a single calcitic rostrum. This possibility is contradicted by analytical results on planktic and benthic foraminifera and corals, as well as inorganic

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carbonate, that generally show only little variability of d7Li values (Marriott et al., 2004a,b; Hathorne and James, 2006; Rollion-Bard et al., 2009; Misra and Froelich, 2012). Clear indication for alteration of other geochemical proxies in parts of specimen NZ0256B makes it very unlikely that d7Li values should have been entirely unaffected by diagenetic overprinting. (ii) Continuous diffusion would be expected to lead to kinetic isotope fractionation effects across the sample, and specifically an enrichment of the heavy isotope in the residual phase due to its lower average diffusion speed (Richter et al., 2006). The measured isotopic ratios in specimen NZ0256B would then need to be compatible with a diffusion gradient leading to the addition of Li with an isotopically light signature in zone III (Figs. 2 and 7), either from the apical zone or from the rim. This scenario is hard to justify for several reasons. Li/Ca ratios are higher in the isotopically heavy area of the rostrum (Fig. 7) and this is incompatible with the addition of isotopically light Li into zone III. Also d7Li values in specimen NZ0256B covary with geochemical proxies for alteration. Specifically, a positive linear correlation with Mn/Ca ratios (r2 = 0.47) and a negative linear correlation with Sr/Ca ratios (r2 = 0.65) are observed. These trends are also indicated when including the results for the other rostra (Fig. 7) but are less clear probably because of differences in the fluid chemistry related to different lithologies (see Table 1) and differing primary Sr/Ca ratios of the rostra. (iii) We consider dissolution-reprecipitation to be the most likely process that partially altered the primary Li isotope signatures of the rostra. The constraints on the physical and chemical composition of the diagenetic fluid predict that extensive interaction with the rock matrix at possibly elevated temperatures took place. Alteration of silicates leads to an enrichment of 7Li in the fluid phase (Chan and Kastner, 2000; Huh et al., 2001; Kisaku¨rek et al., 2005; Pogge von Strandmann et al., 2006). The consequent heavy isotopic signature in pore waters (Chan and Kastner, 2000) is then recorded by the diagenetic carbonate in the same fashion as the less radiogenic Sr leached from the matrix is taken up by the recrystallized calcite. Heavier d7Li values are compatible with highand low-temperature diagenetic settings (Scholz et al., 2010) and pore waters with very positive d7Li values of up to +45& (Chan and Kastner, 2000; Scholz et al., 2010), similar to heaviest values reported in this study have been observed. Li derived from dissolution of the aragonitic phragmocone most likely played no role during alteration as aragonite dissolution would have been an early feature of diagenesis. Furthermore aragonite is found to incorporate even less and preferentially lighter Li than calcite (Marriott et al., 2004a). This is in contrast to the observed heavy Li isotope signatures of the apical line and zone adjacent to the alveolar infill

(Fig. 7). The leachates of the alveolar infill have a very light Li isotope signature ( 4.6 ± 0.3& and 4.9 ± 0.9&). While we cannot exclude, that a small effect of kinetic Li isotope fractionation during sample leaching with HCl contributes to the low d7Li values of the Li derived from the sediment samples, we note that our measurements fall in the range of d7Li values measured for marine sediments (Tomascak, 2004). The difference of 45& between the adjacent calcite and the rock matrix is much larger than expected in equilibrium, suggesting that isotopic disequilibrium between the fossil calcite and the rock matrix could be maintained. The enrichment of Li during recrystallization is hard to explain, because current calibrations for Li distribution factors predict lower Li uptake into the carbonate at elevated temperature (Marriott et al., 2004a). These calibrations are performed using data in the range of 0 to 30 °C at ambient pressures (Hall and Chan, 2004; Marriott et al., 2004a,b) and it is currently unclear if extrapolation to the predicted elevated temperatures and pressures experienced during alteration of the samples is permissible. The distribution factor for Li rises with salinity (Marriott et al., 2004a) so that a highly saline solution could potentially produce a trend to higher Li concentrations during diagenesis. The Gulf of Mexico, where pore waters are highly saline and have high Li concentrations of >400 lM and very positive d7Li values of up to +45& (Scholz et al., 2010), might pose an example for such a setting. Adopting the favoured interpretation of diagenetic control on the Li isotopic composition of the rostrum via dissolution and reprecipitation has strong implications on the use of Li isotope ratios from fossil carbonates in deep time. This suggests that the Li isotopic signature of fossil calcite can be preserved over geological timescales, at least in some thick-shelled organisms, even after deep burial and prolonged exposure to elevated temperatures unless overprinted by recrystallization. The mean d7Li value from all belemnites with Mn/Ca ratios below 0.18 mmol/mol of +27.1 ± 1.1& (Fig. 7) is in the range of values measured for Cenozoic foraminifera (Misra and Froelich, 2012). Very heavy primary d7Li values on the other hand would necessitate the global Li cycle to have operated at much more extreme conditions than seen in the Cenozoic. We thus speculate that a d7Li of +27& is representative of Late Jurassic (155–148 Ma) belemnites. Assuming fractionation factors of 2–5& reported for calcite (Marriott et al., 2004a; Hathorne and James, 2006; Misra and Froelich, 2012), seawater d7Li was about +29& to +32& in the Late Jurassic, which is similar to Cenozoic values and in complete agreement with the modern value of +31&. 5. CONCLUSIONS For this study isotopic and elemental ratios of multiple elements were analysed on a single belemnite rostrum. The findings are compatible with results of earlier studies showing that post-depositional alteration of the rostrum leads to a decrease in d18O values and Sr/Ca ratios, whereas

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Mn/Ca ratios increase. The impact of the diagenetic fluid on the elemental ratios is strongly controlled by the permeability structure of the rostrum and by the chemical composition of the fluid. The 87Sr/86Sr ratio in calcite fossils can be used as a tool for constraining the sources of the elements in diagenetic fluids and in some cases for an estimate of the timing and setting of alteration. Diagenetic mobilization of Li in calcite and all other geochemical tracers measured here seems to be controlled by dissolution and re-precipitation rather than by diffusional processes. The Li isotopic signature in well-preserved parts of belemnite rostra appears to be robust for inferences on the primary composition, even for Jurassic materials that experienced deep burial of several kilometres. An average d7Li value of +27 ± 1& is derived for well-preserved Late Jurassic (155–148 Ma) belemnite material from New Zealand, suggesting that seawater had a d7Li very similar to the modern value during that time interval. ACKNOWLEDGEMENTS We acknowledge John Simes for assistance during sample collection at GNS Science and Neville Hudson and Donald A.B. MacFarlan for logistic help and assistance during sampling in the field. Michael Joachimski is acknowledged for help with cathodoluminescence microscopy at the University of Erlangen. Toni Larsen, Toby Leeper, Sten Lennart and Bo Petersen are thanked for support during preparation and analyses at the University of Copenhagen. We thank the associate editor Dr. Claudine Stirling, Dr. Sambuddha Misra and two anonymous reviewers for extensive suggestions and comments that helped to greatly improve the manuscript. This project was funded by the Danish Council for Independent Research–Natural Sciences (Project 09-072715) and by NERC Research Fellowship NE/I020571/1.

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