The calcium isotopic composition of carbonate hardground cements: A new record of 1 2 changes in ocean chemistry? 3 4 5 Andrea M. Erhardt, Alexandra V. Turchyn, Harold J. Bradbury, J.A.D. Dickson 6 5 7 8 9 10 Abstract 11 12 13 44 14 Reconstructing changes in the calcium isotopic composition ( Ca) of the ocean over Earth 15 10 history has been challenging. This difficulty is due to the large range of calcium isotope 16 fractionation factors during mineral precipitation and the potential for overwriting the initial 17 44 44 18 Ca of minerals during shallow marine diagenesis. We present a new Ca record 19 measured in carbonate hardground cements, an inorganic carbonate-mineral precipitate that 20 rapidly forms at or near the sediment-water interface. The range in the 44Ca for any 21 15 particular carbonate hardground cements is between 0.05 and 0.56. In some cases, the 22 44 23 progressive increase in the Ca during precipitation can be observed, consistent with 24 precipitation in a closed-system. Our data show an average calcium isotope fractionation 25 during carbonate hardground cement precipitation that is -0.57±0.27, similar to the calcium 26 isotope fractionation factor for inorganic calcite precipitates in previous laboratory and 27 20 modelling studies, and closer to what is considered a kinetic end member calcium isotope 28 29 fractionation than growth at equilibrium. This is consistent with the rapid carbonate mineral 30 precipitation expected for carbonate hardground cements. Our 44Ca record over the 31 Phanerozoic is similar to other calcium-bearing mineral records over the same time interval, 32 with average 44Ca becoming lower going back in time by about 0.5 to 0.7 . Our results 33 44 34 25 add further support for the evolution of seawater Ca over time, and we discuss the possible 35 causes of these changes with suggestions for future studies. 36 37 38 Introduction 39 40 41 42 30 The growing interest in understanding the biogeochemical calcium cycle over geological time 43 44 stems from the hope that, by reconstructing the calcium cycle, we will gain insight into the 45 global carbon cycle (De La Rocha and DePaolo, 2000; Farka et al., 2007a; Blättler and 46 47 Higgins, 2017). In principle, the dominant source of calcium to the global ocean, rivers 48 49 through carbonate and silicate weathering, and the dominant sink for calcium from the global 50 51 35 ocean, the deposition and lithification of carbonate-bearing sediments, are the same processes 52 53 moving carbon through, to, and from Earths surface environment (Fantle, 2010; Fantle and 54 55 Tipper, 2014; Blättler and Higgins, 2017). If we could accurately reconstruct changes in 56 these fluxes across Earths surface environment over time, would we gain insight into the 57 58 evolution of the carbon cycle that has, so far, eluded us? 59 60 40 61 62 63 64 65

The tool that has been most used to reconstruct the calcium cycle over time is the 1 2 measurement of the calcium isotopic composition of calcium bearing minerals; particularly 3 4 carbonate minerals (Farka et al., 2007a; Blättler et al., 2012). The calcium isotopic 5 44 40 6 composition is reported as the ratio of the heavy Ca isotope to the lighter Ca isotope 7 8 45 versus a standard in delta notation, with units of parts per thousand, or permil. 9 10 11 (Eq1) 12 13 14 15 16 There are three reference scales (standard in Eq 1) that have been used throughout the 17 18 calcium isotope community: seawater (SW), a carbonate standard (915A), and bulk silicate 19 20 earth (BSE). Each has their advantages and limitations. The carbonate reference scale, 21 22 50 915A, is the most commonly used, suffers from lack of availability for use in many labs as it 23 is no longer commercially available. The BSE standard similarly suffers from the fact that 24 25 silicate earth varies in its 44Ca, and some have argued that seawater may vary in a similar 26 27 manner, making an international agreement on the value of the reference 44Ca/40Ca to use 28 29 elusive (Kang et al., 2017; Chakrabarti et al., 2018). It has become common practice to report 30 31 55 calcium isotope data versus all three reference scales for comparison among various studies 32 33 and approaches. 34 35 36 The use of 44Ca to reconstruct the calcium cycle, and by proxy the carbon cycle, over 37 38 geological time has been challenging. The 44Ca of all minerals precipitating in the ocean is 39 44 40 40 60 lower than the Ca of the ocean, that is minerals precipitating take in more of the Ca 41 44 42 isotope, leaving Ca behind (Fantle and DePaolo, 2005, 2007; Kasemann et al., 2005, 2014; 43 44 Nielsen et al., 2012). This kinetic isotope effect varies in its magnitude, depending largely on 45 the rate of precipitation of the mineral in question and the mineral, or polymorph of calcium 46 47 carbonate, that is precipitating (Gussone et al., 2003; Tang et al., 2008; Blättler et al., 2012). 48 49 65 If you measure only one type of calcium-bearing mineral over time, then any change in the 50 44 44 51 Ca of this mineral could reflect either a change in the Ca of the ocean or a change in 52 53 the nature and conditions of precipitation of this mineral over Earth history, influencing the 54 55 calcium isotope fractionation on mineral precipitation. This is further complicated by the fact 56 that many calcium carbonate minerals dissolve, precipitate, and recrystallise during 57 58 70 sedimentary diagenesis. This recrystallisation can shift the 44Ca of the mineral, and not 59 60 always in the same direction depending on the degree of fluid-buffered versus sediment- 61 62 63 64 65

buffered (or open-system versus closed-system) conditions (Fantle and DePaolo, 2005, 2007; 1 2 Fantle, 2015; Higgins et al., 2018; Ahm et al., 2019). 3 4 5 6 75 This problem has been largely circumvented by measuring calcium isotopes in a single 7 8 mineral, often a biomineral, thus avoiding measurements of bulk carbonate sediments, which 9 may record a range of different primary carbonate mineralogies and where the nature of 10 11 diagenesis may be difficult to disentangle (Higgins et al., 2018; Ahm et al., 2019; Tostevin et 12 13 al., this issue). Many studies have explored the calcium isotope fractionation associated with 14 15 80 biogenic carbonate mineral precipitation, including corals (Gothmann et al., 2016), 16 17 foraminifera (Heuser et al., 2005; Hippler et al., 2006; Sime et al., 2007), brachiopods 18 19 (Farka et al., 2007b; Brazier et al., 2015), belemnites (Farka et al., 2007a,b; Blättler et al., 20 2012), and rudists (Immenhauser, 2005). When these records show long-term temporal 21 22 changes in the calcium isotope composition of the single biomineral, the conclusion is that 23 24 85 either the calcium isotopic composition of the ocean has changed (Blättler et al., 2012) or that 25 26 the calcium isotope fractionation factor on biomineral precipitation has changed (Gothmann 27 28 et al., 2016). 29 30 31 In theory, if one biomineral is measured, and the calcium isotope fractionation factor for that 32 33 90 mineral is known, then a reasonable estimate of seawater 44Ca could be made (Blättler et al., 34 35 2012). Although seawater calcium should have a single 44Ca, as calcium is homogenized in 36 37 the ocean and has a long residence time, the seawater 44Ca as reconstructed from single 38 39 biominerals and calcium-in-barite, over the Phanerozoic, has significant scatter (Figure 1). 40 44 41 This range in the Ca at any given time point has been attributed to difference in the 42 43 95 calcium-isotope fractionation factor during mineral precipitation or stabilisation, which has 44 45 been attributed, at times, to local temperature variability (Tang et al., 2008), changes in the 46 original carbonate-mineral polymorph that precipitated (Gussone et al., 2005), varying rates 47 48 of carbonate mineral precipitation (DePaolo, 2011), and varying degrees of recrystallization 49 50 during diagenesis (Higgins et al., 2018). Additionally, all of these biomineral-based proxies 51 52 100 for the calcium cycle rely on bootstrapping the record to more recent, better characterized 53 54 species. 55 56 57 58 59 60 61 62 63 64 65

1 3.0 44 2 Records of seawater d Ca Marine Barite

3 Forams 4 Brachiopods 5 2.5 6 Belemnites 7 Coral 8 9 2.0

10 ) e l

11 a c S 12 A

5 1.5

13 1 9 14 ( a C

15 4 4 16 d 17 1.0 18 19 20 21 0.5 22 23 24 Cenozoic Mesozoic Paleozoic 25 0.0 0 50 100 150 200 250 300 350 400 450 500 26 27 Age (Ma) 28 29 30 31 Figure 1. Compiled record of calcium isotopes in biogenic carbonate and marine 32 barite. The raw measurements have been corrected using the known calcium isotope 33 fractionation factors for each biomineral, listed in Table 4, to yield what is thought to 34 44 35 be paleo-seawater Ca, allowing the records to be compared to one another as each 36 biomineral has a different calcium isotope fractionation factor. Records include 37 marine barite (purple circles, Griffith et al., 2008); forams (yellow diamonds, Heuser 38 et al., 2005; Hippler et al., 2006; Nägler et al., 2000; Sime et al., 2007), brachiopods 39 40 (light blue circles, Farka et al., 2007a) belemnites (green squares, Blättler et al., 41 2012; Farka et al., 2007a,b), and coral (red triangles, Gothmann et al., 2016). 42 Measurement error, as reported, is typically ~0.1 (2 ). Exact uncertainties as 43 reported in the original publications are shown in Figure 2. 44 45 46 47 105 48 49 50 51 52 53 54 55 56 110 57 58 59 60 61 62 63 64 65

1 2 3 Measured d44Ca of biominerals 44 1.8 Figure 2. Individual Ca records 4 Forams for various minerals shown in 5 1.4 6 composite in Figure 1, here not 7 1.0 corrected for calcium isotope 8 fractionation (i.e. raw data, on the 9 0.6 915A reference scale). Here, only 10 11 0.2 the last 140 Million years are 12 plotted. Error bars represent 13 -0.2 reported 2SD error if available. 14 1.5 Marine Barite 15 1.0 ) e 16 l a

c 0.5 S

17

A What is apparent from the 18 5 0.0 1 9

( 44

19 a -0.5 various records of Ca over the C

20 4 4 21 d -1.0 115 Phanerozoic is that there is some 22 23 -1.5 broad shape or curvature to the 24 0.8 25 Belemnites record (Figure 1, 2). In the 0.6 26 calcium isotope fractionation- 27 0.4 28 corrected compilation of 29 0.2 30 120 biominerals in Figure 1, there is 31 0.0 32 an increase in the reconstructed 33 -0.2 1.4 44 34 Coral seawater Ca from the

35 ) e l 1.0 a

36 c Mesozoic to today. When these S

37 A 5

1 0.6 corrections are removed and each 38 9 (

a

39 C 4

4 0.2 125 biomineral is considered 40 d 41 independently, this increase in 42 -0.2 1.4 43 Brachiopods 44Ca becomes even more 44 1.2 45 apparent (Figure 2). This 46 1.0 47 increase in 44Ca has been seen 48 0.8 49 130 in other calcium-bearing records 50 0.6 51 0 20 40 60 80 100 120 140 not presented here such as 52 Age (Ma) 53 phosphates (Schmitt et al., 2003; 54 Soudry et al., 2004, 2006) and evaporite minerals (Blättler and Higgins, 2014). The 55 56 ambiguity associated with the calcium isotopic fractionation factors of such a range of 57 58 135 minerals necessarily makes absolute comparisons challenging. However, the fact that raw 59 60 comparisons show similar temporal trends (Figure 2) suggests that there is a signal within 61 62 63 64 65

these datasets, that each mineral is responding to a similar change possibly in the 44Ca of the 1 2 ocean. It remains an open question what the drivers of a long-term evolution of the 44Ca of 3 4 the ocean might be. 5 6 140 7 44 8 We add to this discussion the analysis of the Ca of carbonate hardground cements as a 9 10 further calcium-bearing mineral carrier that we have analysed from nine time periods over the 11 12 course of the Phanerozoic. Carbonate hardground cements allow for the constraint of some 13 of the variables plaguing other biogenic carbonate minerals records, including the potential 14 15 145 skewing of the record given the high degree of biological control exerted by some organisms 16 17 in making their calcium carbonate skeletons (Blättler et al., 2012, Gothmann et al., 2016). 18 19 Carbonate hardgrounds are rocks that have rapidly lithified at the sediment/water interface, 20 21 resulting from a combination of high carbonate saturation state (Rameil et al., 2012; Christ et 22 23 al., 2015), low sediment accumulation rates (Shinn, 1969; Kennedy and Garrison, 1975), or 24 150 elevated fluid flow through the subsurface promoting rapid precipitation (Dravis, 1979, 25 26 Lighty, 1985). The timescale for lithification is variable, though short on a geologic time 27 28 scale (10s to 1000s years, Christ at al., 2015). Forming at shallow water depths, carbonate 29 30 hardgrounds are found in warm environments in the modern ocean, such as Pacific atolls 31 32 (Carpenter et al., 1991), the Arabian Gulf (Khalaf et al., 1987), and the Caribbean (Malone et 33 34 155 al., 2001). When sea level falls, carbonate hardgrounds may be subject to meteoric water 35 influence. They are often identified in the geological record from the subsequent boring and 36 37 encrustation of marine organisms into the lithified rock and their abundance is clustered in 38 39 the Mesozoic and early Paleozoic (Christ et al., 2015). 40 41 42 43 160 In theory, carbonate hardground cements should precipitate at the sediment-water interface 44 45 and the assumption is that the fluid should be minimally evolved from seawater (Christ et al., 46 2015). However, pore fluids evolve rapidly in their chemical composition from the sediment- 47 48 water interface and the 44Ca can be particularly variable, given how much more calcium is 49 50 in the mineral phase relative to the fluid phase (therefore small changes in mineral dissolution 51 52 165 and precipitation can greatly impact the 44Ca of the pore fluid) (Fantle and DePaolo, 2007; 53 54 Fantle, 2015). As carbonate hardground cements are effectively inorganic carbonate 55 56 precipitates that are rapid on geological timescales, they may preserve a seawater signal 57 58 better than biominerals or other calcium-bearing phases that are subject to varying vital 59 effects or unclear timing of precipitation. Our results will be compared to other existing 60 61 62 63 64 65

170 records and considered in the context of seawater, porewater, and the carbon cycle evolution 1 2 over the Phanerozoic. 3 4 5 6 Methods 7 8 Samples from carbonate hardgrounds were obtained from the collections of the London 9 175 Museum of Natural History, Wooster University, the University of Colorado, Eötvös Loránd 10 11 University, University of Saskatchewan, and the University of Cambridge. Samples are from 12 13 the modern, Holocene, Cretaceous, Jurassic, Triassic, Permian, Carboniferous, Ordovician, 14 15 and Cambrian. These samples were selected based on evidence of early lithification at the 16 17 sediment/water interface through the presence of marine boring and encrusting organisms. 18 19 180 All samples are thought to be either primary or minimally altered calcite. This presumption is 20 based off of both diagenetic and petrographic considerations. For more uniform comparison, 21 22 any aragonite or previously aragonite cements were visually identified through botryoidal or 23 24 needle crystal morphology and excluded. Additionally, geochemical evidence of primary or 25 26 altered aragonite, through low magnesium concentrations or evidence of alternation, resulted 27 28 185 in sample exclusion. These samples have been previously analysed for composition and 29 18 30 diagenetic alteration using magnesium, strontium, and manganese concentrations, REE, O, 31 13 87 86 32 C, and Sr/ Sr ratios (Erhardt et al., in review). Table 1 summarizes the ages and 33 depositional environments of these samples. 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 62 63 64 65 1 2

3 4 5 6 7 Age (Ma) Stage Period Location Reference Depositional Environment Cement type 8 Enewetak Atoll, Marshalls Carpenter et al., Shallow water modern Atoll, Carbonate reef grainstone- fibrous high 0 Modern Neogene 9 Islands 1991 classic beachrock Mg calcite cement 10 Meerssen Member, the van der Ham et From the type section for the Bladed low-Mg calcite with blocky pore- 66 Maastrichtian 1 Cretaceous 11 Netherlands al., 2007 Mastrictian. Seagrass present filling calcite 12 Immenhauser et Shallow-water cycle tops, likely 66-72 Maastrichtian 2 Cretaceous Qishn Formation, Oman Microcrystalline calcite 13 al., 2004 intertidal 14 100 Cenomanian Cretaceous Upper Greensands, Devon, UK Gallois, 2005 Subtidal Microcrystalline calcite 15 Császár et al., 16 Tölgyhát Limestone, Deepwater sediment, pelagic 'red' Microcrystalline early calcite with blocky 168-170 Bajocian Jurassic 2012 17 Gerecse Mts. Of Hungary limestone pore-filling calcite 18 Platform carbonate- Radiaxial fibrous 19 From a cycle top - often calcrete 331-347 Visean Carboniferous Eskett Limestones, England Dean et al., 2011 calcite, secondary blocky pore filling 20 (subaerially exposed) 21 calcite cement Intertidal and subtidal carbonate Fossiliferous micrite- 22 461-468 Dariwillian Ordovician Kanosh Formation, Utah McDowell, 1986 23 sedimentation microcrystalline calcite cement 24 485-497 Furongian Cambrian 1 Furongian, Banff National Park Westrop, 1989 Subtidal, storm dominated shelf Microcrystalline and drusy calcite 25 495-505 Drumian Cambrian 2 Marjuman, Banff National Park Aitken, 1979 Peritidal Microcrystaline calcite 26 190 Table 1- Samples Used in this Study 27 28 29 30 31 32 33 34 35 36 37 38 39 40 41 42 43 44 45 46 47 48 49

1 2 One key objective was to isolate cements from the primary biominerals that make up the 3 4 carbonate hardground. Early cements were identified by their distinct chemical composition, 5 6 including concentrations of magnesium, calcium, strontium, manganese, and rare earth 7 8 195 elements (Erhardt et al., in review) from both the primary clasts and any interior-pore 9 precipitation. Samples for calcium isotope analysis were collected using a tungsten needle 10 11 with 1m point under an optical microscope or microdrilled, attempting to collect only the 12 13 authigenic carbonate phase (Dickson et al., 2008). Approximately ~1-2 mg of sample was 14 15 collected to support multiple geochemical analysis, though only 10 g was needed for each 16 17 200 calcium analysis. The procedural blank from sample dissolution and separation on the 18 19 Dionex ICS5000+ HPIC, as determined independently by ICP-OES, is 96 ng of calcium 20 when 7 ml of eluent was collected from the Dionex. During the collection of 4.4 mg of 21 22 calcium using the Dionex, this represents ~2% of the collected calcium. 23 24 25 42 26 205 These carbonate samples were dissolved in distilled 2% nitric acid, spiked with Ca and 27 48 28 Ca, and the calcium was isolated from other cations using a Thermo Fisher Scientific 29 30 Dionex ICS5000+ HPIC following the methods of Bradbury and Turchyn (2018). This 31 sample, approximately 4 g of calcium, was loaded onto rhenium filaments with phosphoric 32 33 acid as an activator. The calcium isotope ratios were analyzed on a Thermo Fisher Scientific 34 35 210 Triton Plus Multicollector Thermal Ionization Mass Spectrometer (TIMS) at the University 36 37 of Cambridge. The NIST carbonate standard 915B was analyzed every 4-5 samples. Based on 38 39 the long-term reproducibility of the 915B standard, an error of 0.09 (2) was assigned to 40 41 all analyses. 42 43 44 215 When we present previously published data, the conversion between 44/42Ca and 44/40Ca 45 46 was made using the conversion in Sime et al. (2005) (Equation 2) . Results used from Blättler 47 48 et al. (2012) were converted to account for differences from precipitation temperature. Given 49 50 the overall small magnitude of that conversion relative to the magnitude of change in 44Ca 51 52 observed in this study, further temperature conversions were not applied other samples in this 53 54 220 study. 55 56 Equation 2: 57 58 59 60 61 62 63 64 65

1

2

3 4 5 6 7 8 9 10 11 225 Results 12 44 13 The Ca of carbonate hardground cements ranges from 0.47 to 1.53 ( 915A) with one 14 15 outlier from the late Cretaceous at +1.83 ( versus 915A Table 1, Figure 3). Within any 16 44 17 time interval sampled, the range in the Ca of the carbonate cement is relatively narrow, 18 19 between 0.05 and 0.56 (among data points of the same time interval), disregarding the 20 230 outlier from the late Cretaceous. Within each time period we report multiple analyses of 21 22 chronically adjacent hardgrounds, along with the sampling of both the earliest and likely 23 24 secondary cements when this fine degree of sampling was possible. All 44Ca data from our 25 26 carbonate hardground cements lie below the reconstructed 44Ca for seawater over the 27 28 Phanerozoic (Figure 3), as expected during carbonate mineral precipitation where the 40Ca is 29 30 235 preferentially taken into the carbonate mineral lattice. 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 62 63 64 65

1 44Ca 2 44Ca 44Ca ( 3 Age Location ( ) ( ) 915A) 4 5 0 Modern Enewetak Atoll, Marshalls Islands 0.42 -0.52 1.42 6 0 Modern Enewetak Atoll, Marshalls Islands 0.43 -0.51 1.43 7 8 0 Modern Enewetak Atoll, Marshalls Islands 0.53 -0.41 1.53 9 0 Modern Enewetak Atoll, Marshalls Islands 0.46 -0.48 1.46 10 11 66 Maastrichtian 1 Meerssen Member, the Netherlands 0.23 -0.71 1.23 12 66 Maastrichtian 1 Meerssen Member, the Netherlands 0.34 -0.60 1.34 13 14 66-72 Maastrichtian 2 Quishn Formation, Oman -0.03 -0.97 0.97 15 66-72 Maastrichtian 2 Quishn Formation, Oman 0.02 -0.92 1.02 16 Quishn Formation, Oman 17 66-72 Maastrichtian 2 0.13 -0.81 1.13 18 66-72 Maastrichtian 2 Quishn Formation, Oman 0.21 -0.73 1.21 19 Quishn Formation, Oman 20 66-72 Maastrichtian 2 0.83 -0.11 1.83 21 Upper Greensands, Dunscome Cliff, 22 100 Cenomanian Devon, UK 0.07 -0.87 1.07 23 Upper Greensands, Dunscome Cliff, 24 100 Cenomanian Devon, UK 0.11 -0.83 1.11 25 Upper Greensands, Dunscome Cliff, 26 100 Cenomanian Devon, UK 0.12 -0.82 1.12 27 Tölgyhát Limestone, Gerecse Mts. Of 28 168-170 Jurassic Hungary- Composite -0.19 -1.13 0.81 29 168-170 Jurassic Gerecse Mts. Of Hungary- Composite -0.20 -1.14 0.80 30 31 168-170 Jurassic Geresce Sequence 1- earliest cement -0.53 -1.47 0.47 32 168-170 Jurassic Geresce Seq. 1- intermediate cement -0.42 -1.36 0.58 33 34 168-170 Jurassic Geresce Sequence 1- later cement -0.05 -0.99 0.95 35 168-170 Jurassic Geresce Sequence 2- earliest cement -0.46 -1.40 0.54 36 37 168-170 Jurassic Geresce Sequence 2- later cement 0.03 -0.91 1.03 38 Sixth/Seventh Limestones, Eskett 39 331-347 Mississipian Limestones, England -0.38 -1.32 0.62 40 Sixth/Seventh Limestones, Eskett 41 331-347 Mississipian Limestones, England -0.35 -1.29 0.65 42 Sixth/Seventh Limestones, Eskett 43 331-347 Mississipian Limestones, England -0.32 -1.26 0.68 44 461-468 Ordovician Kanosh Formation, Utah -0.20 -1.14 0.80 45 46 461-468 Ordovician Kanosh Formation, Utah -0.15 -1.09 0.85 47 485-497 Cambrian 1 Furongian, Banff National Park -0.44 -1.38 0.56 48 49 495-505 Cambrian 2 Marjuman, Banff National Park -0.07 -1.01 0.93 50 495-505 Cambrian 2 Marjuman, Banff National Park -0.07 -1.01 0.93 51 52 495-505 Cambrian 2 Marjuman, Banff National Park -0.04 -0.98 0.96 53 495-505 Cambrian 2 Marjuman, Banff National Park 0.03 -0.91 1.03 54 55 Table 2 - Our data, including the age, Period or Epoch, location where the sample comes 56 from, its calcium isotope composition versus bulk silicate earth, seawater, and vs. 915A. 57 58 59 60 61 62 63 64 65

1 2 3 2.5 4 Records of seawater d44Ca 5 6 7 2.0 8 ) e l

9 a c 1.5 S 10 A 5

11 1 9 (

12 a 1.0 C 4

13 4 14 d 15 0.5 16 17 18 0.0 19 0 50 100 150 200 250 300 350 400 450 500 20 240 Age Ma 21 22 Figure 3. Results from this study (red squares) of 44Ca in carbonate hardground 23 cements. The grey dots represent the complied existing data as seawater (see also 24 Figure 1), with the dark dashed line representing the average fit through those results, 25 44 26 as a best estimate curve of the Ca of the ocean over time from which the carbonate 27 hardground cements precipitate. Measurement error for our samples is ± 0.09 (2 ). 28 29 30 Discussion 31 44 32 Range in the δ Ca of the carbonate hardground cements 33 34 35 At any given point in time, we observe a range of 44Ca measured in the carbonate 36 37 hardground cement samples of between 0.54 and 1.03 ( Table 2, Figure 3). In theory, if 38 39 245 the carbonate hardground cements were precipitating in a fully closed system, then the initial 40 44 41 carbonate precipitate would be offset from the Ca of the pore fluid by some amount 42 43 determined by the polymorph of carbonate mineral precipitated and the rate of mineral 44 precipitation. This removal of 40Ca would leave 44Ca behind in the pore fluid, and the next 45 46 carbonate mineral to precipitate would derive from this isotopically evolved fluid, differing in 47 48 250 its 44Ca from the initial carbonate mineral precipitate. In modern environments where 49 50 carbonate minerals are precipitating, this distillation (i.e. Rayleigh fractionation) of calcium 51 52 isotopes has been seen in the sedimentary pore fluids, and has been used to reconstruct the 53 54 rate of carbonate mineral precipitation and its calcium isotope fractionation (Teichert et al., 55 2009; Bradbury and Turchyn, 2018). During the subsequent recrystallization of carbonate 56 57 255 minerals, the pore fluids evolve to a lower 44Ca, reflecting the addition of 40Ca-rich from 58 59 carbonate mineral phases (it is thought that there is no calcium isotope fractionation on 60 61 62 63 64 65

dissolution) (Fantle and DePaolo, 2005, 2007; Fantle, 2015). The range of 44Ca that has 1 2 been measured in modern sedimentary pore fluids is approximately 2 due to this range of 3 4 processes (Fantle and DePaolo 2005, 2007; Turchyn and DePaolo, 2011; Fantle, 2015; 5 6 260 Bradbury and Turchyn, 2018). 7 8 9 In the case of carbonate hardground cements, we expect the pore fluids to be affected less by 10 44 11 recrystallization or mineral dissolution lowering the Ca of the pore fluid, and more by the 12 13 amount of carbonate cement precipitation and the degree of open-versus-closed system 14 15 conditions during cement formation. This is because of the rapid nature of carbonate 16 265 hardground cement precipitation near to the sediment-water interface. If the system is 17 18 closed and carbonate mineral precipitation is occurring, we should see the isotopic 19 20 distillation discussed above; this would be reflected in an increase in the 44Ca across the 21 22 cement. However, if the system is more open largely to exchange with seawater or other 23 44 24 fluids, this distillation may not be seen, and a more constant Ca found in the precipitating 25 26 270 cements. 27 28 29 We see two examples of closed-system carbonate mineral precipitation in our carbonate 30 31 hardground cements. In one Jurassic sample we were able to identify and sample multiple 32 cement phases using a tungsten needle. This sample shows three distinct phases of 33 34 precipitation, transitioning from a micrite to blocky calcite phases. There is an increase in the 35 36 275 44Ca of the cement with growth (Figure 4). In this case, the early phase may be a rock- 37 38 buffered pore fluid (i.e. Pruss et al., 2018) with the following cements from burial or later 39 40 diagenetic processes. In theory, as the cement grows onto the previously precipitated 41 42 carbonate mineral, the initial mineral precipitate is not isotopically reset. 43 44 45 46 47 48 49 50 51 52 53 54 55 56 57 58 59 60 61 62 63 64 65

1 A. B. 2 44 d Ca 915A (‰) 3 44 1) Ca 915A (‰) 4 : 84; B- 6; 4 (D) 5 , - 4C89- 6; ' 2

6 & 1 & " 7 0.470.47 % + 0 / .

8 0.58 $ - 0.58 , % % 9 0.950.95 # ! 10 " 11 ! 100 m 12 " ! )+ ! )* ! )( ! )' ! )& ! )% ! )$ ! )# ! )" ! 1000 m 13 µ 34- 56789.: 84; .< - 6; 4., - =57>? .@; ? - 7979A 14 15 16 17 Figure 4. A) Three layers of the Jurassic carbonate hardground cement show an increase 18 in the 44Ca of subsequent layers of carbonate precipitation with increasing 44Ca. The 19 inset picture shows the optical image of this sample. B) Rayleigh fractionation of 20 21 porewater and carbonate during simple carbonate precipitation in a closed system with a 22 constant calcium isotope fractionation factor of = 0.9985 (DePaolo, 2011). 23 24 44 25 280 If we assume both the initial Ca of the fluid, a constant calcium isotope fractionation 26 44 27 factor, and closed system precipitation, the Ca of the cement could be used to track the 28 fraction of calcium remaining in the porewater during cement precipitation (Figure 4B). 29 30 However, this would assume that we know both what the 44Ca of the initial pore fluid was at 31 32 this time, the calcium isotope fractionation on mineral precipitation and, more critically, that 33 34 285 there was no local mineral dissolution (or recrystallisation). All three of these assumptions 35 36 are not well constrained. In marine sediments dominated by sedimentary carbonate mineral 37 38 precipitation, although carbonate mineral precipitation exceeds the rate of dissolution by as 39 much as 10 or 20 times, there is still some mineral dissolution as suggested by pore fluid 40 41 models (Fantle and DePaolo, 2005, 2007; Fantle, 2015; Huber et al., 2017; Bradbury and 42 43 290 Turchyn, 2018). 44 45 46 We can make a simple assumption that in modern sedimentary pore fluids where there is 47 44 48 active carbonate precipitation occurring, the Ca of the pore fluid is either equal to or 49 50 higher than seawater (but critically not less than). This is consistent with modern sedimentary 51 52 pore fluids in sediment with active sedimentary carbonate precipitation. While there is no 53 295 current data of pore water 44Ca in the upper centimeters of the sediment column where 54 55 hardgrounds form, the initial pore fluid 44Ca that has been reported at 3 or 5 meters below 56 57 the surface is often at or higher than seawater 44Ca (Teichert et al., 2007; Turchyn and 58 59 DePaolo 2011; Bradbury and Turchyn, 2018) with the exception of samples from Site 807A 60 61 62 63 64 65

(Fantle and DePaolo 2007). This means that in any given time interval, the lowest 44Ca we 1 2 300 measure in the carbonate hardground cement may represent the earliest precipitate, and 3 44 4 therefore the one that precipitated from the least evolved pore fluid, closest to seawater Ca. 5 6 Making this assumption gives an estimate of the calcium isotope fractionation factor on 7 8 carbonate hardground cement precipitation and a bound on the trajectory of changes in the 9 44Ca of the fluid over time. 10 11 12 13 305 We use this assumption to estimate the calcium isotope fractionation factor for the various 14 carbonate hardground cements (Table 3), which averages -0.57 ± 0.27 for the nine time 15 16 intervals studied. There is a range in calcium isotope fractionation for the carbonate 17 18 hardground cements, with the lowest calcium isotope fractionation in the Cretaceous 19 20 carbonate hardground cement (-0.22) and the highest in the Jurass ic (-1.11). There are 21 22 310 several reasons why these numbers should be taken as very rough estimates rather than firm 23 44 24 calcium isotope fractionation factors. For example, the reconstruction of seawater Ca 25 (Figure 3) is highly uncertain, and furthermore it is likely that the initial fluids from which 26 27 the carbonate hardground cements precipitated are evolved in some way from any paleo- 28 29 seawater 44Ca. Furthermore, given the physical difficulty of only sampling the cement, it is 30 31 315 probable that some of the host carbonate biominerals were included in our analysis. Hence, 32 44 33 the measured Ca may not only record cement from pore fluids. Accidental sampling of the 34 44 35 biomineral with the cement would skew our data lower in its Ca on the whole, and thus the 36 calculated calcium isotope fractionation during carbonate hardground cement precipitation 37 38 (Table 3) would be larger than in reality. 39 40 41 320 With these caveats on our calculated calcium isotope fractionation factor during carbonate 42 43 hardground cement formation, we can compare our average (-0.57±0.27) to the calcium 44 45 isotope fractionation factor for other calcium-bearing carbonate minerals. We note that in 46 47 general the calcium isotope fractionation for inorganic carbonate precipitates is less than for 48 49 biogenically precipitated minerals (Table 4), and our carbonate hardground cement data has a 50 51 325 similar calcium isotope fractionation as inorganic calcite. It is also consistent with the 52 estimated calcium isotope fractionation calculated at the kinetic limit for inorganic 53 54 carbonate mineral precipitates (Nielsen et al., 2012). We suggest that the calcium isotope 55 56 fractionation on carbonate hardground cement precipitation is kinetically controlled and not 57 58 growing at equilibrium with the pore fluids, and unlikely to be recrystallised in fluid-buffered 59 60 330 conditions. 61 62 63 64 65

Interpolated Calcium 1 2 Lowest seawater isotope 44 44 3 cement Ca Ca ( fractionation 4 Age of hardground ( 915A) 915A) factor () 5 6 0 Modern 1.42 1.84 - 0.42 7 66 Maastrichtian 1 1.23 1.45 - 0.22 8 66-72 Maastrichtian 2 0.97 1.48 - 0.51 9 10 100 Cenomanian 1.07 1.69 - 0.62 11 168-170 Jurassic 0.47 1.58 - 1.11 12 13 33`-347 Carboniferous 0.62 1.44 - 0.82 14 461-468 Ordovician 0.8 1.29 - 0.49 15 16 485-497 Cambrian 1 0.56 1.24 - 0.68 17 495-505 Cambrian 2 0.93 1.24 - 0.31 18 19 Average -0.57 0.27 20 21 22 Table 3. Lowest measured 44Ca in the carbonate hardground cement, the interpolated 23 44 24 seawater Ca, and the calculated calcium isotope fractionation factor during cement 25 precipitation from each time period studied. 26 27 28 29 We compare our hardground cements with other Phanerozoic carbonate cements. Previously 30 44 31 published Ca of carbonate cements arent from hardgrounds and therefore may have more 32 ambiguity regarding the timing of precipitation, however they serve as a check for our 33 34 calcium isotope fractionation factor. Steuber and Buhl (2006) measured Cenomanian 35 36 335 diagenetic calcite with an average 44Ca of 1.1 +/- 0.05 2SD (converted to 44Ca on the 37 38 915A reference scale), which is identical to our Cenomanian sample of 1.1 +/- 0.06 2SD. 39 44 40 Additionally, the Steuber and Buhl (2006) Ca for Carboniferous cement was 1.14 +/- 41 44 42 0.15 2SD, which is higher than our Mississippian Ca of 0.62 +/- 0.05 2SD. Our 43 44 44 44 Mississippian carbonate hardground cement Ca are among the lowest Ca that we have 45 340 measured (Figure 3). Other studies have analysed calcite cements filling voids in Cretaceous 46 47 (Toronian) ocean basalts at 1.3 1.5 (Weinzierl et al., 2018) and Ediacaran age cements 48 49 that were originally aragonite at ~0.6 (Pruss et al., 2018). These 44Ca for other carbonate 50 51 52 53 Carbonate Phase/ Approx. calcium isotope fractionation 54 Source 55 Seawater Proxy from seawater () 56 57 58 Inorganic calcite -0.57 +/- 0.27 This Study 59 60 61 62 63 64 65

1 -0.4 to -1.6 dependent on rate Tang et al., 2008 (laboratory study) 2 3 4 -0.6 to -0.7 Nielsen et al., 2012 (modelled result) 5 6 -0.6 to -0.8 Marriott et al., 2004 (laboratory study) 7 8 9 Brachiopod -0.85 Farka et al., 2007a 10 11 12 -0.94 Heuser et al., 2005 13 14 Forams -1.4 to -1.1 Sime et al., 2005 15 16 17 -1.2 to -0.8 (Orbulina universa) Gussone et al. 2003 18 19 20 -1.12 (modern seawater) 21 Gothmann et al., 2016 22 ∆ (Coral-SW) = -0.06 ± 0.01*[Ca] - 0.6 ±- 23 Coral 0.2 24 25 -1.2 at 21° C 26 Böhm et al., 2006 27 -1.0 at 29° C 28 29 Belemnites -1.40 Farka et al., 2007b 30 31 Inorganic -1.8 at 10°C and -1.5 at 30°C Gussone et al., 2003 32 aragonite 33 34 35 Marine barite 2.01 +/- 0.15 Griffith et al., 2008 36 37 Table 4. Compilation of calcium isotope fractionation from fluid for inorganic and 38 biogenic carbonate minerals from modern seawater. 39 40 cements compare favourably with our results for the 44Ca of carbonate hardground cements. 41 42 43 345 This range of calcium isotope fractionation during the precipitation of different biogenic and 44 45 non-biogenic minerals remains a focus of research (Table 4). The calcium isotope 46 47 fractionation during carbonate mineral precipitation is largely driven by environmental 48 49 conditions, such as temperature and ion concentration, that impact the rate of mineral 50 precipitation. The controls on the rate of carbonate mineral precipitation have been 51 52 350 investigated in detail by multiple workers and are only briefly summarised here (Lemarchand 53 54 et al., 2004; Gussone et al., 2005; Tang et al., 2008; DePaolo, 2011). Ultimately the rate of 55 56 mineral precipitation is related to the concentration of calcium and dissolved inorganic 57 58 carbon, which determines whether the precipitation is controlled by the concentration of ions 59 60 at the surface or the need to transport these ions to the surface of the growing mineral 61 62 63 64 65

355 (surface-controlled versus transport-controlled) (DePaolo, 2011). When a carbonate 1 2 mineral is growing under surface-controlled conditions then it is closer in theory to 3 4 thermodynamic equilibrium and therefore the calcium isotope fractionation is lower, while in 5 6 transport-controlled there is a larger calcium isotope fractionation (Fantle and DePaolo, 7 8 2007). These scenarios were modelled by DePaolo (2011) to empirically determine the 9 360 relationship between precipitation rate vs. calcium isotope fractionation. This work was 10 11 continued in Nielsen et al. (2012) where they developed a model to simultaneously consider 12 13 both growth rate and calcium isotope fractionation as a function of solution stoichiometry and 14 15 oversaturation. Overall, they found that the relationship of the concentration of calcium 16 17 versus carbonate ions influences the calcium isotope fractionation factor; significantly higher 18 19 365 concentrations of calcium versus that of carbonate ions drives the system toward isotopic 20 equilibrium (Nielsen et al., 2012). However, increasing supersaturation overall, that is 21 22 increasing both the concentration of calcium and carbonate ions, drives the system toward 23 24 kinetic end-members, increasing the precipitation rate and increasing the calcium isotope 25 26 fractionation factor (Nielsen et al., 2012). Using this model, inorganic carbonate mineral 27 28 370 precipitation, particularly calcite, from seawater gave a calcium isotope fractionation factor 29 30 between -0.6 and -0.7 (Nielsen et al., 2012), similar to both laboratory studies and our 31 results from inorganic calcite hardground cements. 32 33 34 The rate of carbonate hardground cement precipitation is likely much slower than laboratory 35 36 experiments, i.e. less than 10-9 mol/m2/s (Tang et al., 2008), and much higher than the rates in 37 38 375 deep-sea sediments of ~10-17 mol/m2/s (Fantle and DePaolo, 2007). This is based in the 39 40 anecdotal evidence of the formation of carbonate hardgrounds over 10s 100s of years 41 42 (Christ et al., 2015). This gives support to the idea that the carbonate hardground cements are 43 44 precipitating near a kinetic limit and the calcium isotope fractionation factor could be similar 45 to laboratory experiments and those predicted from modeling (Nielsen et al., 2012). In 46 47 380 contrast, biomineral precipitation dominantly occurs within microenvironments, where 48 49 locally the saturation state can be greater than surface seawater, which will impact the rate of 50 51 mineral precipitation (Bentov and Erez, 2005; Sime et al., 2005). As a result, the calcium 52 53 isotopic fractionation of biogenic minerals from seawater is typically greater than inorganic 54 55 precipitates and can scale differently with changes in biomineral growth rate (Table 4, 56 385 Blättler et al., 2012). 57 58 59 Change in the calcium isotope composition of the ocean over time 60 61 62 63 64 65

Our data add to the growing weight of evidence suggesting that the calcium isotopic 1 2 composition of the ocean can change over geological time. While we cant resolve possible 3 4 changes in the calcium isotope fractionation factor changing over geological time, we do 5 44 6 390 observe the same increase in the Ca over the late Mesozoic and Cenozoic that has now 7 8 been reported for biominerals and calcium-in-barite (Figure 1, 2). Going forward, it will be 9 44 10 key to resolve both the magnitude and the timing of this change in seawater Ca. As we 11 better understand the actual calcium isotope fractionation factor during mineral precipitation, 12 13 and measure a wider range of minerals, we will be able to better tease out the magnitude, 14 15 395 timing, and rate of change in the 44Ca of seawater. 16 17 18 19 There are several observations we can make now that will help frame the future research and 20 44 44 21 discussion on this change in seawater Ca. First, it is notable that the change in the Ca 22 23 observed in our data and seen in other calcium-bearing mineral proxies is on a longer time 24 25 400 scale than one would expect given the short (~0.6-1 Myr) residence time of calcium in the 26 ocean (Fantle and Tipper, 2014). Indeed, during short-term perturbations to the carbon - and 27 28 linked calcium - cycles, there have been previously reported perturbations to the 44Ca that 29 30 are not fully linked to changes in depositional mineralogy (Payne et al., 2010; Griffith et al., 31 32 2015). However, the change observed over the Cenozoic and into the Mesozoic is over closer 33 34 405 to 50-60 Million years, much longer than would be expected given the short residence time of 35 36 marine calcium. 37 38 39 Any driver of an increase in the 44Ca of the ocean must either remove more of the 40Ca 40 41 isotope (or less of the 44Ca isotope) or add more of the 44Ca isotope (or less of the 40Ca 42 43 410 isotope). The calcium biogeochemical cycle is dominated by one source (river weathering) 44 45 and one sink (carbonate mineral deposition), with a smaller contribution from hydrothermal 46 47 circulation and chemical reaction with clay minerals and the off-axis oceanic crust. Due to 48 49 the dominance of one source and one sink, it has been thought to be difficult to 50 fundamentally change the 44Ca of the ocean, as the 44Ca of the source and sink are similar, 51 52 415 and fundamentally the surface calcium cycle is dominated by the recycling of carbonate 53 54 minerals and rocks throughout Earths surface environments (Blättler et al., 2012; Gothman 55 56 et al., 2016; Blättler and Higgins, 2017). 57 58 59 60 61 62 63 64 65

One possibility is a change in chemical weathering over the Mesozoic and Cenozoic. As 1 2 420 calcium in rivers is dominated by carbonate weathering, which has much more 40Ca in it 3 4 (rivers have low 44Ca), a river-driven increase in the 44Ca of the ocean would imply a shift 5 6 to more silicate-sources calcium or a reduction in carbonate weathering (Sime et al., 2007; 7 44 8 Griffith et al., 2008.) Driving changes in the Ca of the ocean by changes in chemical 9 10 weathering has been previously explored by Fantle and Tipper (2014), where they showed 11 425 that a switch from rivers weathering 100% carbonate minerals to weathering 100% silicate 12 13 minerals would only result in a ~0.4 shift in marine 44Ca, making it an ineffective lever to 14 15 change marine 44Ca. Tectonic events over the late Mesozoic and Cenozoic, particularly the 16 17 uplift of the Himalayas (60-40 Ma) and Andes (50-30 Ma), would in theory expose more 18 44 19 silicate minerals for chemical weathering increasing the Ca of riverine calcium. However, 20 21 430 the uplift of the Alps, a dominantly carbonate-rock mountain range also occurred over this 22 44 23 time interval which might counteract any change in river Ca from the weathering of 24 25 silicate terranes. Furthermore, it is unclear whether there is enough calcium in silicate rocks 26 relative to the large amount in carbonate rocks to drastically change the 44Ca of the river 27 28 input. 29 30 435 31 32 Another driver of a change in the 44Ca of the ocean could be the switch between calcite- 33 34 dominated and aragonite-dominated oceans (Sime et al., 2007, Farka et al. 2007a, Blättler et 35 36 al., 2012). In the Mesozoic, the dominant carbonate polymorph precipitated was calcite 37 38 (Stanley and Hardie, 1998), in oceans that had higher overall calcium concentrations and 39 440 lower magnesium and sulfate concentrations than the modern ocean. In contrast, aragonite is 40 41 the thermodynamically preferred carbonate polymorph in the modern ocean (Friedman et al., 42 43 1965), particularly for marine invertebrates (Zhuravlev and Wood, 2009) and inorganic 44 45 carbonate minerals (Christ et al., 2015). If all other environmental variables are equal 46 40 47 (temperature and rate of mineral precipitation), aragonite has more Ca in it, or a larger 48 49 445 calcium isotope fractionation factor from seawater than calcite, and has been shown across 50 multiple carbonate-precipitating organisms (~0.6 more depleted than calcite Table 4 51 52 Blättler et al., 2012). In principle, if more carbonate-as-aragonite is removed from the ocean, 53 54 more 40Ca will be removed as well, driving the 44Ca of the ocean to higher, or more 44Ca- 55 56 enriched values. Large shifts in the abundance of aragonitic vs. calcitic precipitating 57 58 450 organisms have been observed throughout the Phanerozoic, that might support this driver of 59 44 60 ocean Ca (Zhuravlev and Wood, 2009). The problem with this explanation is that 61 62 63 64 65

carbonate minerals will recrystallise during diagenesis, with the aragonite recrystalising to 1 2 calcite, which is a more stable carbonate mineral polymorph. In certain unique conditions of 3 4 carbonate recrystallisation, the aragonite-turned-calcite will retain the lower 44Ca from its 5 6 455 precipitation, however most of the time the calcite ultimately formed will acquire a new 7 44 8 Ca during recrystallisation and be higher than the initially precipitated aragonite (Higgins 9 10 et al., 2018; Pruss et al., 2018). Thus it isnt clear that precipitating more carbonate-as- 11 aragonite would ultimately result in removing more 40Ca from the ocean. 12 13 14 15 460 However, the 44Ca of carbonate minerals deposited in shallow marine environments is 16 17 fundamentally more 44Ca-enriched, or isotopically heavy than the 44Ca of carbonate 18 19 minerals deposited in deep-marine environments. The reason for this is thought to be that 20 21 high fluid flow through shallow carbonate platforms and carbonate-rich sediments tends to 22 44 23 increase the Ca of the carbonate minerals towards those in equilibrium with seawater 24 25 465 (Fantle and DePaolo 2007; Higgins et al., 2018). Carbonate platforms are fluid-buffered 26 systems, i.e. open systems with sufficient and repeated exposure to the overlying ocean to 27 28 overprint the initial mineral 44Ca. While certainly not all shallow marine environments are 29 30 fluid buffered, shallow marine environments like have greater fluid flow through submarine 31 32 groundwater discharge, changes in eustatic sea level, evaporation-driven density gradients, 33 34 470 etc. (Higgins et al., 2018). This results in carbonate minerals in shallow marine environments 35 44 36 that are, on the whole higher, and closer to seawater Ca. 37 38 39 In contrast, the 44Ca of carbonate minerals in deep marine environments tend to preserve 40 41 their original water-column acquired 44Ca, or at a minimum remain around 1 to 1.4 lower 42 43 475 than seawater. Fluid flow through deep marine sediments is restricted, partially due to the 44 45 increased clay compositions of these sediments and lower external fluid-forcing mechanisms; 46 47 this creates a sediment-buffered system. A shift, therefore, from carbonate mineral burial 48 49 primarily in shallow, fluid buffered environments, to carbonate mineral burial partitioned 50 40 51 between shallow and deep environments could result in preferentially more Ca being 52 44 53 480 removed from the ocean, driving the Ca of the ocean higher. 54 55 56 Exactly such a scenario happened in the middle Mesozoic, with the Marine Mesozoic 57 58 Revolution (Ridgwell, 2005). The Marine Mesozoic Revolution was the colonising of the 59 60 open ocean by pelagic calcifying organisms, pushing carbonate mineral production into open 61 62 63 64 65

485 water columns, where before it had been confined to near continental shelf and slope 1 2 environments. The Marine Mesozoic Revolution has been suggested to have changed the 3 4 regulation of the carbon cycle (Ridgwell and Zeebe, 2005) by creating a pelagic carbonate 5 6 sink, which enabled Earths system to better respond to perturbations to the carbon cycle 7 40 8 (Payne et al., 2010). This pelagic carbonate sink would have had more Ca in it, and this 9 44 10 490 would have allowed the Ca of the ocean to increase through the expansion of this deep 11 carbonate sink (initially suggested in Blättler et al., 2012). Future work focusing on this 12 13 interval with the calcium isotopic composition of the ocean in mind would both verify the 14 15 dynamics of this event and constrain the sensitivity of the ocean calcium reservoir to shifts in 16 17 sink locations, highlighting the applicability of the calcium isotope proxy. 18 19 495 20 21 Conclusions 22 We have measured the 44Ca in carbonate hardground cements in samples spanning the 23 24 Phanerozoic. When compared to the existing datasets of biogenic minerals, similar trends of 25 26 lower 44Ca further back in time are observed. It appears increasingly likely that independent 27 28 500 coral, foraminifera, marine barite, and inorganic carbonate precipitates are reflecting changes 29 30 in the isotopic composition of the calcium ocean reservoir. The offset between a 31 44 44 32 reconstructed seawater Ca and the lowest Ca in the carbonate hardground cements is 33 34 0.57±0.27 , consistent with the calcium isotope fractionation for inorganic carbonate 35 36 precipitates in laboratory studies (Tang et al., 2008) and predicted from models (Nielson et 37 505 al., 2012). Looking forward, the analysis of additional carbonate hardground cements may 38 39 help constrain both pore water dynamics and seawater calcium isotopic composition during 40 41 times of change in the Phanerozoic. Future directions include the modelling of the amount of 42 44 43 Ca that would need to be sequestered over time in fluid vs sediment buffered systems to 44 44 45 change the ocean Ca and higher resolution records involving single biominerals to better 46 44 47 510 constrain the timing and nature of changes in marine Ca. Carbonate hardground cements 48 49 offer a promising new avenue for studying earliest marine diagenesis and the evolution of 50 seawater chemistry. 51 52 53 54 55 515 Acknowledgements 56 57 We gratefully thank Paul Taylor, Mark Wilson, David Budd, Orsolya Gyri, Brian Pratt and 58 59 Stephen Lokier for providing samples for this study. We are thankful to the two anonymous 60 reviewers, along with the editor, who greatly improved the manuscript. Funding for this 61 62 63 64 65

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