Quantification of diagenesis in Cenozoic sharks: Elemental and mineralogical changes

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Abstract

Diagenesis of bone during fossilization is pervasive, however, the extent of this process varies with depositional environment. This study quantifies diagenesis of shark vertebral centra through analysis of a suite of physical and chemical characters including crystallinty index (CI), carbonate content, and elemental concentrations. Although shark skeletons are initially cartilaginous, the soft cartilage of the vertebral centra is replaced with carbonate hydroxyapatite during growth. Nine vertebral centra are analyzed from lamnoid (Lamnoidea) sharks ranging in age from the cretaceous to recent using Fourier transform infrared spectroscopy (FT-IR) and inductively coupled plasma mass spectrometry (ICPMS). The variables CI, carbonate content, rare earth element (REE) concentrations, Ca/P, Ba/Ca, Sr/Ba, (La/Yb)N, (La/Y)N, (La/Yb)N vs. (La/Sm)N, La/Yb, and Ce anomalies elucidate the diagenetic and depositional environments of the seven fossil vertebral centra. The two extant centra demonstrate the initial, unaltered end-member conditions for these variables. Two fossil vertebral centra (Carcharodon megalodon and Isurus hastalis) demonstrate a strong terrestrial influence during diagenesis (distinctive flattening of shale-normalized REE patterns) that masked the seawater signal. Three centra (Carcharodon auriculatus, Carcharodon angustidens, and Creotxyrhina mantelli) have indications of some terrestrial influx evident by some flattening of the REE patterns relative to seawater. The terrestrial influence in these five shark centra (C. megalodon, I. hastalis, C. auriculatus, C. angustidens and C. mantelli) are interpreted to represent a primarily nearshore habitat for these species. In contrast, the two Otodus obliquus centra have REE patterns that represent the original seawater signal and have no indications of terrigenous input. These results indicate that fossil shark vertebral centra have the potential to understand diagenesis and reconstruct paleooceanographic environments.

Introduction

Fossilized vertebrate skeletal tissues, including teeth and bone, have recently received considerable attention as geochemical archives of paleoecological and paleoenvironmental information. In these studies, fossil tooth enamel has been the preferred material for analysis because of the compact, relatively non-porous mineral consisting of >95% hydroxyapatite. In contrast, some interesting studies of broad relevance have been presented using isotopic data from fossil bone (e.g., to reconstruct dinosaur physiology; Barrick and Showers, 1994). However, these studies have come under close scrutiny (Kolodny et al., 1996) because porous bone is more prone to diagenesis than teeth (Wang and Cerling, 1994).

There are certain situations in which fossil bone is either the only skeletal material available for study (e.g., in those vertebrates that lack teeth, such as most birds) or is preferred because certain skeletal elements archive incremental growth. One example of an archive of incremental growth is shark vertebral centra. Although shark skeletons are initially cartilaginous (a soft supporting tissue that does not fossilize), the cartilage is replaced in the vertebral centra with carbonate hydroxyapatite during the growth of the individual. This growth is periodic, and incremental rings are called annuli because of their presumed annular cyclicity, although this is not always the case, (e.g., Branstetter et al., 1987). These growth rings can be easily seen in both modern and fossilized shark centra. During a related research project investigating stable isotopic signatures archived in fossil shark centra (e.g., MacFadden et al., 2004), we became interested in the extent of diagenesis and how it potentially affects the geochemistry of fossil bone.

The purpose of this study is to quantify diagenesis of shark bone through analysis of a suite of physical and chemical characters including crystallinty index (CI), carbonate content, and major, minor, trace elemental concentrations. The sharks are all from the group known as the superfamily Lamnoidea (Cappetta, 1987) which includes the modern great white (Carcharodon carcharias) and six closely related extinct species ranging in age from Cretaceous to Pleistocene. The modern shark species are included in this study to provide an unaltered “end-member” in which initial physical parameters and elemental concentrations can be determined. Lamnoid sharks were selected because these sharks are widely distributed in space and time. A broad geographic distribution of fossils should illuminate the effects of different degrees and environments of diagenesis. The vertebral centra were chosen because they are the primary ossified skeletal tissue that fossilizes in sharks (i.e., other than teeth).

Stable isotopes and rare earth elements (REE) of biogenic apatites have been used for paleoclimatic reconstruction, to trace ocean currents and water masses, to quantify redox conditions, for incremental growth studies, and to reconstruct diet (Piper, 1974, Kolodny et al., 1983, Elderfield and Pagett, 1986, Kolodny and Luz, 1991, Lécuyer et al., 1993, Picard et al., 1998, Picard et al., 2002, Shields and Stille, 2001, MacFadden et al., 2004, Pucéat et al., 2004). Partial or complete dissolution, precipitation, recrystallization, and ion uptake by adsorption and diffusion may lead to changes in chemical composition and lattice structure of biogenic apatite (Reiche et al., 2003). As such, the original chemical signatures of biogenic apatites may be modified through diagenesis, resulting in the interpretation of erroneous biological signals (Pucéat et al., 2004).

Modern bone is composed of hydroxyapatite (Ca5(PO4)6(OH)2) that has small crystallites, large surface area (200 m2/g; Weiner and Price, 1986) and high organic content (∼35%, principally collagen and water; Williams, 1989, Carlson, 1990, Koch et al., 1992). The high reactivity of biogenic apatite is due to the small size and high surface area of the bone hydroxyapatite crystallites (Trueman, 1999). Many substitutions are possible for both the anions and cations in biogenic hydroxyapatite (Table 1; Nathan, 1981). In modern biogenic apatites, carbonate (CO32−) can substitute for either OH (A site) or PO43− (B site) but primarily substitutes for the latter (Shemesh, 1990, Lee-Thorp and van der Merwe, 1991, Rink and Schwarcz, 1995). Substitution of carbonate for phosphate can distort the crystal lattice and further decrease the stability of biogenic apatite (Nelson, 1981, Nelson et al., 1983).

During fossilization the carbonate hydroyxapatite alters to a more stable form of apatite (francolite or carbonate fluorapatite) by losing carbonate and hydroxyl ions and gaining fluoride (Nathan and Sass, 1983, Newsley, 1989, Greene et al., 2004). The loss of carbonate in francolite decreases the defects of the hydroxyapatite lattice, resulting in increased crystal size and stability relative to carbonate hydroxyapatite (Greene et al., 2004).

Through the processes of diagenesis trace element concentrations either increase or decrease relative to those of unaltered bone (Elderfield and Pagett, 1986, Wright et al., 1987, Williams, 1988, Grandjean and Albaréde, 1989, Koeppenkastrop and De Carlo, 1992, Grandjean-Lécuyer et al., 1993, Denys et al., 1996, Hubert et al., 1996, Laenen et al., 1997, Reynard et al., 1999, Trueman, 1999, Starton et al., 2001). Trace elements are most likely incorporated into bone apatite during early diagenesis through substitution. After this initial “recrystallization,” trace element signatures appear to be stable and resistant to later diagenesis (Bernat, 1975, Grandjean and Albaréde, 1989, Grandjean-Lécuyer et al., 1993). REE3+ ions are similar in size to Ca2+, consequently they readily substitute into the Ca site (Whittacker and Muntus, 1970). Because REE are typically trivalent, introduction of these cations into the Ca site occurs by coupled substitution, for example, REE3+ + Na+  2Ca2+.

Adsorption also occurs during diagenesis of biogenic apatite. Binding is usually weak and reversible and therefore ions adsorbed are susceptible to exchange as long as the crystal surface remains exposed. However, if the inter-crystalline porosity is closed during diagenesis, individual crystallite surfaces will be closed to further exchange. Finally, cations (such as Sr2+) may be incorporated into fossil bone via growth of authigenic apatite (Reynard et al., 1999, Trueman and Tuross, 2001). Ultimately, the final trace element composition of the biogenic apatite is controlled by the concentration of trace elements in the system, the apatite–fluid partition coefficient, the chemistry of the burial microenvironment, bone microstructure, and the length of exposure (Trueman, 1999).

REE research on biogenic apatites is currently focused on the possible uses of the REE signal, determination of the signal source (i.e. environmental, biological, or diagenetic), and the processes of incorporation into the fossil apatite (Henderson et al., 1983, Fleet, 1984, Grandjean et al., 1987, Wright et al., 1987, Williams, 1988, Trueman, 1996, Trueman and Benton, 1997, Laenen et al., 1997, Reynard et al., 1999, Shields and Stille, 2001, Picard et al., 2002, Trueman and Tuross, 2001). REE in seawater typically exist in the 3+ oxidation state. One exception is cerium, which can undergo oxidation in seawater from the solvated 3+ state to the relatively insoluble Ce4+ state (deBaar et al., 1985). Under oxic conditions, Ce4+ is readily removed from seawater onto particle surface coatings or into authigenic minerals (Sholkovitz et al., 1994, Koeppenkastrop and De Carlo, 1992), resulting in a negative Ce anomaly (Ceanom.). In contrast, under reducing conditions Ce3+ is typically released back into the water column or pore waters (German and Elderfield, 1990). The oceanic distribution and typical seawater pattern of REE is largely controlled by adsorptive scavenging by settling particles. deBaar et al. (1985) illustrated that in both the Atlantic and Pacific oceans, all REE except Ce increase with water depth.

Bernat (1975) who described icthyolith (fish teeth) in which REE concentrations in from the upper-most 600 cm of ocean-core sediments have a bulk REE pattern similar to overlying waters. These results indicate that biogenic apatites incorporate a REE composition at the sediment/seawater interface during early diagenesis (with little or no fractionation) and are not prone to late diagenetic exchange of REE. Two main theories of REE incorporation into the biogenic apatite at the sediment seawater interface have been proposed. Firstly, some have argued that REE incorporated into biogenic apatite must be transported by and introduced into fossils directly from seawater or porewater (Henderson et al., 1983, Williams, 1988, Trueman, 1996, Trueman and Benton, 1997). However, direct uptake of REE in biogenic apatites from pore waters and/or seawater raises serious problems. Fossil biogenic apatites contain several tens to several hundreds parts per million (ppm) of REE, whereas maximum REE concentrations in pore water and seawater are in the range of parts per billion (ppb) and part per trillion (ppt), respectively (Elderfield and Greaves, 1982, Elderfield and Sholkovitz, 1987). Assuming that the REE are taken up directly through diffusion of pore waters, approximately one ton (103 kg) of pore water would be required to give the biogenic apatite enough REE to fit observed concentrations (several tens to several hundreds ppm; Grandjean et al., 1987, Grandjean and Albaréde, 1989, Grandjean-Lécuyer et al., 1993).

Secondly, Grandjean et al. (1987) proposed quantitative uptake of non-detrital REE locally released at the sediment/seawater interface to explain biogenic apatite enrichment. Abundant debris with large surfaces, which easily adsorb large amounts of REE from seawater, are dispersed in the oceans and are known to settle to the ocean floor. Such a rain of REE-rich carriers has been identified in sediment traps (Murphy and Dymond, 1984) and comprises a variety of inorganic (detrital minerals, oxyhydroxides) and organic (pellets, organic debris) phases. The decay of the REE-rich carriers at the sediment/seawater interface associated with biogenic apatites and the resulting reducing conditions eventually cause the dissolution of Fe–Mn oxyhydroxides, which transfer their REE to the recrystallized biogenic apatite. The transfer of REE from oxyhydroxides to biogenic apatites occurs within a rather short period of time (Grandjean and Albaréde, 1989). Upon completion of the early diagenetic processes and once most oxyhydroxides have dissolved, apatite remains the major non-detrital REE repository in sediments (Grandjean and Albaréde, 1989). This extension of Bernat’s (1975) model indicates that for biogenic apatites: (1) the primary source of REE is seawater, (2) early diagenetic transfer of REE to biogenic apatites occurs through a short-lived phase consisting of oxyhydroxides and organic detritus, and (3) REE are incorporated within a short period of time and do not undergo late diagenetic exchange.

Variations in host sediments can typically influence the REE compositions of biogenic apatites due to differences in permeability, the REE flux from diagenetic fluids expelled from sediments (diagenetic signal), and organic and oxyhydroxide contents (Grandjean-Lécuyer et al., 1993, Lécuyer et al., 2004). Porewater REE are derived from the surrounding sediments. The higher concentrations of REE in porewaters relative to seawater will allow fluxes of REE from sediments to seawater (i.e., diagenetic fluids that undergo diffusion from sediments into seawater; Elderfield and Sholkovitz, 1987). Terrestrially derived sediments characteristically have shale-normalized REE (REEN) that are relatively flat and no Ceanom. (particularly from common fine-grained detrital material; Grandjean et al., 1987). Therefore, the REE contents of biogenic apatites deposited in terrestrially derived sediments (clays and sands) typically have flattened REEN patterns that are intermediate between those of seawater and those of shale (Grandjean et al., 1988, Elderfield et al., 1990). Sediments that precipitate directly from seawater (carbonates and phosphorites) with little or no terrestrial input have diagenetic fluids reflecting the composition of the overlying water column. Biogenic apatite deposited in these types of sediments typically show a seawater REE pattern because the diagenetic signature in the sediments is the same as the overlying water column (Lécuyer et al., 2004). Therefore, the REE signature in fossil biogenic apatites results from a mass balance between the flux of REE from: (1) decaying organic and oxyhydroxides (primary carriers with seawater signature), (2) diagenetic fluids that undergo diffusion from sediments to seawater (diagenetic signature), and (3) rivers (detrital signature; Grandjean and Albaréde, 1989).

There is evidence for REE fractionatation from seawater during incorporation into biogenic apatites. Reynard et al. (1999) generated a model in which partition coefficients of REE between apatites and water for substituted ions were extrapolated from mineral/melt partition data. These were compared with available experimental partition coefficients for REE adsorption. Reynard et al. (1999) demonstrated that bell-shaped REE patterns (middle REE enrichment) in fossil apatites are due to fractionation with seawater or continental fluids at low temperatures. Therefore, in these cases, the fluid composition can only be determined if the fractionations are known, in which case the fluid composition can be backed out.

Of the numerous geochemical variables that are potentially available to quantify diagenesis, CI, carbonate content, and REE concentrations and ratios seem to be most promising. Crystallinity index, carbonate content, and REE are consistent in modern shark centra and alter during diagenesis; therefore, modern shark centra provide an initial end member. In contrast, as demonstrated below, minor elemental concentrations are highly variable in modern shark centra and are therefore considered less useful in assessing diagenesis.

Section snippets

Materials and methods

This study analyzes the geochemistry of nine shark centra (Fig. 1): two modern great whites (Carcharodon carcharias) and seven fossil specimens ranging in age from Cretaceous to Pliocene (Table 2). The chemical and mineralogical properties of the nine shark centra were determined by Fourier Transform Infrared Spectroscopy (FT-IR) and Inductively Coupled Plasma Mass Spectrometry (ICPMS). FT-IR is used here to determine crystallinity and estimate carbonate content. FT-IR has advantages over X-ray

Mineralogical changes

FT-IR spectra of the modern and fossil shark centra are shown in Fig. 2. Both the modern and fossil samples have the same characteristic absorption bands as the FT-IR spectra of synthetic apatites containing CO32− at both A-and B-sites (Bonel, 1972). The FT-IR spectra for modern specimens are characterized by large H2O bands (which usually mask the OH band at 3567 cm−1) and the presence of organics represented by the three amide group bands (amide I 1660 cm−1, amide II 1550 cm−1, and amide III

Mineralogical characterization of centra

In contrast to the modern specimens, the fossils have lost most, if not all, of their organic content through diagenesis and therefore contain less absorbed (3430 cm−1 band) and structural H2O (3330 cm−1 band; Holcomb and Younf, 1980, Michel et al., 1995). The weak intensity of the absorption band near 1660 cm−1, corresponding to νCONH of the amide group (amide I), and the absence of the two other amide bands (amide II and III) signify a significant loss of organics in the fossils (Reiche et al.,

Conclusions

While certain variables independently provide information about diagenesis, the simultaneous use of crystalinity index, carbonate content, and elemental concentrations has the potential to more fully elucidate depositional environment and diagenesis. Diagenesis resulted in these seven fossil shark centra the incorporation of a combination of seawater signal, diagenetic signal (diagenetic fluids expelled from sediments), and detrital signal (river water) at the sediment/seawater interface. These

Acknowledgments

We thank M. Gottfried, G. Hubbell, C. Jeremiah, S. Witner, D. Nolf, and O. Sakamoto for allowing us to borrow and sample the shark centra. We thank G. Kamenov, A. Heatherington, C. Bohn, and A. Shriller for assistance in the laboratory. C. Lécuyer, T. Lyons, and the other two reviewers provided helpful comments that improved this manuscript. This research was supported by NSF Grant EAR 0418042. This is University of Florida Contribution to Paleobiology 569.

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