|UFDC Home||myUFDC Home | Help|
This item has the following downloads:
QUANTIFICATION OF THE EXTENT OF DIAGENESIS IN BIOGENIC APATITE
OF CENOZOIC SHARK CENTRA
JOANN LABS HOCHSTEIN
A DISSERTATION PRESENTED TO THE GRADUATE SCHOOL
OF THE UNIVERSITY OF FLORIDA IN PARTIAL FULFILLMENT
OF THE REQUIREMENTS FOR THE DEGREE OF
DOCTOR OF PHILOSOPHY
UNIVERSITY OF FLORIDA
Joann Labs Hochstein
This dissertation is dedicated to my parents, Robert and Denise, and my Husband, Jason
for all their love and support.
Firstly, I would like to thank my advisor, Bruce MacFadden, for his insights, expert
guidance, support, patience, and continual enthusiasm for this project. I also wish to
thank my committee members, Ellen Martin, Phil Neuhoff, Neil Opdyke, and John
Krigbaum, for their patience and support throughout this project. George Kamenov,
Jason Curtis, and Penny Higgins gave invaluable insight during sample preparation and
analysis. I am grateful for Clifford Jeremiah for providing the inspiration of this project
with the donation of Otodus obliquus specimens analyzed in this study. I would also like
to thank Michael Gottfried, Gordon Hubbell, Dirk Nolf, O. Sakamoto, and Sabine
Wintner for allowing me to borrow and sample specimens.
I would like to thank all my friends, especially Helen Evans and Steve Volpe, for
their continual support and comic relief. The staff of the Geological Sciences
Department, Ron Ozbun, Jody Gordon, and Mary Ploch, has been very helpful during my
study at the University of Florida. I would like to thank my parents and sister for their
constant support and patience. Finally, I would like to thank my husband, Jason, for the
encouragement to accomplish my dreams and his unconditional love.
This research was supported by Geological Society of America Grant number
2009018 and National Science Foundation grant EAR 0418042.
TABLE OF CONTENTS
A C K N O W L E D G M E N T S ................................................................................................. iv
L IST O F T A B L E S .............. ..... ..................................... .......... .. .. ............. .. vii
LIST O F FIG U R E S ......................................................... ......... .. ............. viii
A B ST R A C T .......... ..... ...................................................................................... x
1 IN T R O D U C T IO N ............................................................................. .............. ...
2 DIAGENSIS AND VARIATION OF THE OXYGEN ISOTOPIC SIGNATURES
IN VERTEBRAL CENTRA FROM Otodus obliquus............... ............... 3
Intro du action ..................................................................................................... .... .. 3
M materials and M methods ................................................................. ....................... 8
G ross X -ray A nalyses ................... .......... .... .......... .. ...... ........... ... ...... 9
Fourier Transform Infrared Spectroscopy Preparation and Analysis..................9
Stable Isotope A nalyses................................................ ............ ............... 10
R results and D discussion ................................... .. ...... ... ........ .... ............ 12
Fossil and Recent Shark Centra Mineralization and Diagenesis......................12
Physical Increments and Variation.................. ...... ...............15
Stable Isotope (6180) Signal Archived in Eocene Otodus obliquus centra........17
C o n c lu sio n s..................................................... ................ 2 1
3 QUANTIFICATION OF DIAGENESIS IN CENOZOIC SHARKS:
ELEMENTAL AND MINERALOGICAL CHANGES ....................... .................23
Introdu action ................. .............. ................... ............. ................. 2 3
B one Chem istry and D iagenesis...................................... ........................ 24
R are Earth Elem ents ........................................................................... 26
M materials and M ethods ......................................................................... ............... 30
Fourier Transform Infrared Spectroscopy (FT-IR) ..........................................31
Elem ental A analysis (ICPM S) ........................................ ......................... 33
R results ...........................................................................................33
M ineralogical Changes ......................................................... ............... 33
E lem ental C concentration ........................................................... ....................34
D iscu ssion ........................................................................................ .. 4 1
M ineralogical Characterization of Centra ....................................................41
Implications for Diagenetic and Biological Signal Reconstruction ....................42
C o n clu sio n s..................................................... ................ 4 6
4 OXYGEN ISOTOPIC AND RARE EARTH ELEMENTAL ANALYSIS OF
MODERN LAMNID SHARKs: DETERMINATION OF LIFE HISTORY? ...........48
Intro du action ............. ......... .. .. ......... .. .. ............................................. 4 8
R are Earth Elem ents ........................... ........... ...................... ...............50
B background ............. .. .. .... .. ..........................................................................51
Great White (Carcharodon carcharias) ..................... ..................... 51
Longfin Mako (Isuruspaucus) ............................ ........................52
Shortfin M ako (Isurus oxyrinchus) ........................................ ............... 53
M materials and M methods ....................................................................... ..................54
X-radiograph Analyses ............... ........................... ............... 54
Oxygen Isotopic Preparation and Analyses......... ............................. 55
Bomb Carbon Dating Preparation and Analysis ...........................................55
Inductively Coupled Plasma Mass Spectroscopy (ICPMS) .............................57
Results and D discussion .................................... ..... .. ...... .............. 58
O ntogenic A ge D eterm nations ........................................ ....................... 58
R are Earth Elem ents .......................................................................... 61
C o n c lu sio n s..................................................... ................ 6 3
5 SUMMARY AND CONCLUSIONS......................................................................68
L IST O F R EFE R E N C E S ............................................................................. ............. 72
B IO G R A PH IC A L SK E TCH ..................................................................... ..................82
LIST OF TABLES
2-1. Comparison of crystallinity index (CI) and carbonate content (C/P) from FT-IR
spectrum for samples treated with acetic acid and samples that were not treated
with acetic acid and from Eocene and modern samples ............................... 13
2-2. Stable isotopic data for three vertebral centra of Otodus obliquus, UF 162732,
from the Early Eocene of Morocco.................................................20
3-1. Some possible substitutions in the apatite crystal structure. .....................................25
3-2. Lamnid shark specimens used in this study ...... ......... ....................................... 31
3-3. Elemental and mineralogical data of nine shark vertebral centra. Elemental
concentration are in ppm .. ............................. .... .......................................36
4-1. Lamnid specimens used in this study. ............................................ ............... 53
4-2. Ontogenic age estimates based growth ring counts (GR), oxygen isotopic (6180)
cyclicity, and bom b carbon (A13C)................................................ ............... 59
4-3. Elemental data (in ppm), and oxygen isotopic and bomb carbon dating ages ..........64
LIST OF FIGURES
2-1. Modern white shark, Carcharodon carcharias (UF211352, left), and Eocene
Otodus obliquus (UF162732, right) showing well defined growth increments.........5
2-2. Infrared spectrum of O. obliquus (UF162732A, sample j102-53) between 500 and
7 0 0 c m 1 .......................................................... ................ .. 7
2-3. Fourier transform infrared spectra of (a) Otodus obliquus centrum (UF162732C,
sample j102-58) (scale on y-axis 0 to 1.4) and (b) Isurus paucus centrum
(UF211353, sample j102-319) (scale on y-axis 0 to 1.2) ........................................ 14
2-4. Graph representing decrease in the amount of carbonate with an increase in
crystallinity ............................................................................................ 15
2-5. Contact prints of the three Otodus obliquus (UF162732) x-rays. The "." symbols
indicate dark growth bands on centra............................................ ...............16
2-6. Centrum of Otodus obliquus, UF 162732A, from the early Eocene of Morocco
showing exact sampling locations (grooves, top) and plot of variation in 6180c
(b o tto m )............................................................................................... 19
3-1. Nine vertebral centra used in this study........................................................ ........ 32
3-2. FT-IR spectra of all nine shark vertebral centra illustrating the differences from
modern (C. carcharias) to fossil biogenic apatites. ............................................35
3-3. FT-IR spectra from 400 to 850 cm-1, illustrating the y4 PO43- band differences
between modern (solid line) to fossil (dashed line) shark centra...........................35
3-4. Carbonate content (C/P) vs. crystallinity index (CI) of the nine shark vertebral
3-5. Isocon plots (Grant, 1982) depicting variations in fossil elemental concentrations
and ratios due to diagenesis........................................................... ............... 38
3-6. PAAS normalized REE of the nine vertebral centra divided into four diagenetic
groups ................ ......... ............................. ........................... 39
3-7. Compilation of observed (La/Yb)N vs. (La/Sm)N in biogenic apatites of various
ag es an d ty p es................................................... ................ 4 0
4-1. Scanned contact print ofBTO433 centra ....................................... ............... 49
4-2. BTO433 bomb carbon data plotted vs. two reference curves.................................57
4-3. Oxygen isotopic data (VPDB) for the shark centra analyzed...............................61
4-4. Post Archean Australian Shale normalized rare earth element plots for the eight
sharks analyzed. .................................................... ................. 65
4-5. Depth estimates for the eight lamnid sharks. Arrow indicates direction of
increasing w after depth............ .. .................. ........ ... ... .. ........ .... 66
Abstract of Dissertation Presented to the Graduate School
of the University of Florida in Partial Fulfillment of the
Requirements for the Degree of Doctor of Philosophy
QUANTIFICATION OF THE EXTENT OF DIAGENESIS IN BIOGENIC APATITE
OF CENOZOIC SHARK CENTRA
Joann Labs Hochstein
Chair: Bruce J. MacFadden
Major Department: Geological Sciences
Diagenesis of bone in the fossil record is pervasive; however, the extent of this
process varies with depositional environment. Diagenesis is any chemical or physical
change that occurs below 200C. This study quantifies the extent of diagenesis in shark
vertebral centra through analysis of a suite of physical and chemical properties including
crystallinity index, carbonate content, isotopes, and major, minor, and trace elemental
concentrations. The sharks used in this study (Family Lamnidae) range in geographic
location and geological age from the Cretaceous to Recent. Although shark skeletons are
initially cartilaginous, the cartilage of the vertebral centra is replaced with carbonate
hydroxyapatite during growth of the individual. Understanding chemical changes to
biogenic apatite informs of the extent diagenesis has altered the biological signal
preserved in vertebrate bones.
Modern lamnid vertebral centra establish a modem analog for comparison to fossil
lamnid sharks. Rare earth element (REE) compositions, A14C, and 6180, give indications
of timing of when these sharks migrate, changes in their eating habits, changes in water
depth, and determination of ontogenic age and growth rates.
Fossil shark centra used in this study have undergone diagenesis; therefore how
have the processes of diagenesis affected the original signal recorded in these centra?
Shale-normalized REE patterns indicate that diagenesis has erased the original signal;
however because diagenesis occurred at or near the seawater/sediment interface, a
seawater REE signal may still be preserved in lamnid shark centra (related to the time of
deposition and location). Therefore, with caution, geochemical data from biogenic
apatite of fossil marine vertebrates, such as lamnid sharks, may be used to understand
paleoceanography and paleoenvironment. Also, the centra from Otodus obliquus
demonstrate that the biological oxygen isotopic signal is not completely erased by
diagenesis. Therefore, biological signals and diagenetic signals of fossil lamnid sharks
can be utilized to understand paleobiology, paleoclimatology, and paleoceanography.
Diagenesis is a fundamental and pervasive aspect of the fossil record. Diagenesis
is defined here as all physical and chemical changes that occur to vertebrate fossils
(bones and teeth) after deposition and below 200C. For vertebrate bones, rapid physical
and chemical changes are known to occur within relatively short geological timescales.
For example, Trueman (2004) demonstrated that after death and decay of a modem
animal, their bones will uptake rare earth elements (REE) within 25 years. The changes
that occur in bones during diagenesis can affect crystallinity, organic content,
mineralogy, density, major, minor, and trace element concentrations, and stable isotopic
compositions. Any scientist interested in interpreting the fossil record must understand
how diagenesis has affected the evidence. Despite the importance of understanding how
diagenesis affects biogenic apatite, this process is poorly understood and frequently
considered to be an impediment to progress. In many cases the effects of diagenesis are
explained as being of minor importance, or insignificant, to the interpretation of
geochemical data archived in the fossil record.
In this context, the fundamental questions that I seek to answer in this research
* How does diagenesis affect the geochemistry of fossil bone?
* How can these changes be quantified?
* Can life histories of sharks (e.g., age determination and ecology) be determined by
the chemical signal preserved in modem and fossil shark vertebral centra?
Although initially cartilaginous, sharks deposit carbonate hydroxyapatite in their
vertebral centra throughout their life, and therefore their centra are prone to fossilization
(except for the ubiquitous teeth, the rest of the skeleton remains cartilaginous and usually
does not fossilize). Lamnid shark centra were chosen to understand the affects of
diagenesis on fossil bone for the following reasons:
1. Vertebral centra in sharks are incrementally calcified during an individual's
lifetime. Calcified bands deposited in concentric rings are either, dark and compact
representing periods of slow growth whereas light, less dense bands correspond to
periods of faster growth. According to Ridewood (1921) and Moss (1977), there is
no resorption or redeposition of bone during life as is the case in other vertebrates,
e.g., the limb bones of reptiles (e.g., Francillon-Vieillot et al. 1990).
2. Lamnid sharks (Family Lamnidae, sensu Gottfried et al., 1996) are widely
distributed in space, time, and different marine sedimentary (phosphate, carbonate,
and plastic) environments throughout the Cenozoic.
3. Lamnid sharks are large, including the largest shark ever to have lived, the Mio-
Pliocene Carcharodon megalodon (Gottfried et al. 1996). Consequently, their
vertebral centra are physically large and therefore can be more easily sampled for
4. Modem lamnids such as, Carcharodon carcharias (the Great White Shark) and
Isurus makoss), are available for analysis of an unaltered end-member.
This dissertation will discuss the use of oxygen isotopic data as a tool for
incremental growth studies in fossil sharks, and the chemical and mineralogical alteration
of lamnid shark centra caused by diagenesis. Also, discussed is the use of growth ring
counts, bomb carbon, and oxygen isotopic data in modern analogs (i.e., modem great
white and mako centra) as ontogenic age determination tools.
DIAGENSIS AND VARIATION OF THE OXYGEN ISOTOPIC SIGNATURES IN
VERTEBRAL CENTRA FROM Otodus obliquus
Sharks have been an abundant part of the marine fossil record since they first
appeared in the Devonian. Sharks are usually represented in the fossil record by their
durable teeth; most of the remainder of the skeletal tissue is composed of hyaline
cartilage, and thus not prone to fossilization. However, vertebral centra, and in rare
instances the jaws, both undergo secondary calcification during ontogeny resulting in
more durable skeletal elements composed of carbonate hydroxyapatite that are frequently
preserved in the fossil record (Ridewood, 1921; Goodrich, 1930; Applegate, 1967; Moss,
1977; Compango, 1999). Shark centra grow incrementally, laying down a dark band that
represents time of slow growth and a light band that represents times of faster growth.
Most shark species deposit a set of bands (one dark, which represents winter and one
light, which represents summer) annually (called annuli), but this may vary depending on
the species, physical environment (including temperature and water depth), food
availability, and stress (Branstetter et al., 1987). Clearly, the periodicity of the growth
bands cannot be established in modem or fossil sharks just by counting the growth rings.
Therefore, the challenge for paleobiological interpretation is how to interpret the
periodicity of the growth bands. Stable isotopic analysis provides the potential to
independently determine whether growth band increments represent annual growth, and
therefore the ontogenetic age of an individual. In most modern fish 6180 in body fluids is
close to that of ambient water, and since P043- and CO32- are cogenetic oxygen-bearing
phases in isotopic equilibrium with the same oxygen reservoir at the same temperature, a
linear correlation should exist between the oxygen isotopic composition of the phosphate
and carbonate, 6180p and 6180, respectively (lacumin et al., 1996). The results below
show that isotopic 6180, are archived in mineralized bone of fossil shark centra, even
when it is diagenetically altered. These data may not represent the actual amplitude of
temperature change, but demonstrate seasonal cycles that can be used to corroborate age
determinations based on counting physical growth bands. The preservation of
incremental growth layers in fossil vertebrate skeletal tissues provides the opportunity to
assess growth rates of individual species and the evolution of developmental strategies in
ancestral and descendant species. In addition to the physical archives of bone growth
preserved in the fossil record, recent studies have also applied stable isotope analyses to
understand periodic growth and related parameters of diet and seasonality preserved in
fossil bone and teeth (Longinelli and Nuti, 1973; Kolodny et al. 1983; Cerling and Sharp
1996; Bocherens et al., 1996; MacFadden et al., 1999; and Vennemann et al., 2001).
Results from associated vertebral centra of the lamnid shark Otodus obliquus
from the early Eocene (Ypresian) of Morocco (Fig. 2-1) are presented in this chapter.
This is a cosmopolitan species and is well represented in the highly fossiliferous
phosphate mines of the Oued Zem, central Morocco (Arambourg, 1952). The current
study was undertaken to determine:
If 50-million-year-old shark centra preserve an archive of incremental growth and
isotopic data that can be interpreted in a meaningful ontogenetic and phylogenetic
The extent to which secondary calcification during ontogeny and/or diagenesis
after death obscures or removes the biologically significant incremental growth
and stable isotopic variation preserved in the centra.
If there is intravertebral variation in the physical incremental growth preserved in
Figure 2-1. Modern white shark, Carcharodon carcharias (UF211352, left), and Eocene
Otodus obliquus (UF162732, right) showing well defined growth increments.
The Eocene shark Otodus obliquus was chosen for this study for several reasons.
This species is represented by excellent specimens of intact, fossilized portions of the
vertebral column in association with teeth. Otodus obliquus is conservatively classified
within the Lamnidea, the family that includes the modern mako (Isurus oxyrinchus),
white (Carcharodon carcharias), and extinct shark species, including the Carcharodon
megalodon (Gottfried and Fordyce, 2002). Therefore, this study is a necessary
foundation for further studies of the evolution of development and body size in extinct
Isotopic studies of bone are rare due to the potential for diagenetic alteration as a
result of large surface area and small crystal size. During fossilization, replacement and
recrystallization occurs within the crystal lattice, which changes the original composition
of carbonate hydroxyapatite to carbonate fluorapatite (sometimes referred to as
"francolite") and eventually to fluorapatite. Transformation of carbonate hydroxyapatite
to fluorapatite occurs with the loss of CO2 and OH- and addition of F-, which causes an
increase in crystallinity (Barrick, 1998; Wang and Cerling, 1994; Shemesh et al., 1983).
Fourier transform infrared (FT-IR) spectroscopy has been effectively used to evaluate
mineral characteristics of fossils. Infrared spectroscopy measures the absorption of
infrared radiation by the sample at the vibrational frequencies of its component molecular
bonds, allowing characterization of its structural sites. In addition, the magnitude of IR
absorption is proportional to the concentration of a molecular species in the sample
(Sibilia et al., 1988). The crystalline structure of bone can be determined by calculating
the crystallinity index (CI) from the extent of phosphate peak splitting at 565-605 cm-1 in
an FT-IR spectrum (Figure 2-2). The 605 cm1 peak intensity increases with respect to
the 565 cm- peak intensity with an increasing degree of fluorination. Ultimately, the CI
is influenced by the size distribution of crystallites and the degree of the substitutional
order-disorder within the crystal lattice (Shemesh, 1990). Apatites with larger, more
ordered crystals show greater separation of these peaks and a higher CI, while in poorly
crystallized apatites the peaks are closer together and therefore have a lower CI (Weiner
and Bar-Yosef, 1990; Wright and Schwarcz, 1996).
In the apatite lattice carbonate can substitute in two sites, OH-1 and P04-2
(Shemesh, 1990; Lee-Thorp and van der Merwe, 1991; Rink and Schwarcz, 1995),
indicated by the superscripts A and B in the formula:
Ca5 [(P04)3-B(CO3)B] [(OH)-A(C03)A]
which results in two sets of absorption bands in an FT-IR spectrum, corresponding to
A(1545-1450-890 cm-1) and B (1465-1412-873 cm-1). Carbonate content can be
estimated from the ratio of the absorbance of the CO3 and P04 peaks (C/P) in the FT-IR
spectrum (Shemesh, 1990; and Wright and Schwarcz, 1996). The amount of carbonate
present in apatite affects the CI, due to type B carbonate substitution for P04, which
produces smaller crystals with greater strain; therefore, highly carbonated apatites show
little peak splitting and have lower crystal indices. Fourier transform infrared spectrum
will indicate the presence of non-apatite mineral structures, such as fluorine. Francolite
(carbonate fluorapatite) has a characteristic peak at 1096 cm-1, therefore the presence of
fluorine can be determined (Wright and Schwarcz, 1996).
700 650 600 550 500
Figure 2-2. Infrared spectrum of 0. obliquus (UF162732A, sample j102-53) between 500
and 700 cm- 1. The crystallinity index, (CI) is calculated by (A + B)/C, where
A, B, and C represent the peak height from the baseline.
It has been found that recrystallization during diagenesis does not necessarily affect
the isotopic composition (Barrick, 1998). When bone recrystallizes in a closed system,
there is no alteration of the isotopic value (Stuart-Williams et al. 1996). Studies have
shown (see Lee-Thorp and van der Merwe, 1991; Wright and Schwarcz, 1996) that when
a sample is treated with buffered 1M acetic acid any apatite that has been diagenetically
enriched in CO32- and secondary carbonate minerals can be removed (Lee-Thorp and van
der Merwe, 1991; Wright and Schwarcz, 1996). It is known that when CO3-2 substitution
increases apatite solubility, rendering it more susceptible to diagenesis (Krueger, 1991;
Lee-Thorp and van der Merwe, 1991; Wright and Schwarcz, 1996). As a consequence,
6180, values may not be a reliable paleothermometer in fossil bone. However, fossils that
have been affected by diagenesis the oxygen isotope composition may have some
biological information preserved, as will be discussed below.
Materials and Methods
Three associated fossil shark precaudal centra were analyzed from a single
individual catalogued in the Vertebrate Paleontology Collection, Florida Museum of
Natural History (FLMNH), University of Florida (UF) 162732. These centra are
identified as Otodus obliquus based on association with a diagnostic dentition. This
specimen was collected from the Early Eocene (Ypresian) unit within the phosphate
mines at Oued Zem, central Morocco. For comparison with the fossil shark, four modern
shark centra representing Carcharodon carcharias (great white, UF Environmental
Archaeology specimen 31648 and UF Vertebrate Paleontology specimen 211351), Isurus
oxyrinchus (shortfin mako, UF Environmental Archaeology specimen 47943), and Isurus
paucus (longfin mako, UF Vertebrate Paleontology specimen 211353) were analyzed.
Gross X-ray Analyses
Traditional x-radiographic ("x-rays") photography can reveal physical differences
in bone density, such as those representing incremental growth bands. X-rays of the
centra were taken at the C.A. Pound Human Identification Laboratory at UF. The x-rays
are set at 78 kV for 2 minutes. The x-rays are used to make contact prints, which are a
reversed pattern of the x-ray (i.e. dark lines on the x-ray are the light lines on the contact
print). The dark and light alternating growth rings are easily seen on the contact prints
and are marked to indicate the location of the samples used for oxygen isotope analysis.
These contact prints were scanned digitally and modified for presentation using Adobe
Fourier Transform Infrared Spectroscopy Preparation and Analysis
For the intended study presented here, Fourier transform infrared spectroscopy has
its advantages over x-ray diffraction (XRD), including: (1) only a small amount of
sample is required (<1 mg); (2) preparation is easier and produces more accurate results;
and (3) carbonate content can be assessed from FT-IR.
Four -2 mg microsamples were drilled with a low speed Foredom drill from each
of the three centra. For the FT-IR analyses, two samples came from the center, and two
from the edge of each centrum. One-half of the samples were treated with the same
procedures as those analyzed for oxygen isotope composition (Table 2-1) to compare
with the results from samples not treated with acetic acid and hydrogen peroxide. The
other half of the samples were left untreated to serve as a control. The samples were
weighed out to 0.8 mg, and combined with 150 mg of spectral grade KBr, and ground
together in a ball mill. The KBr dye was put under vacuum for 5 minutes and
compressed (under vacuum) at 20,000 psi for another 8 minutes. The vacuum was
removed and the KBr dye remained under 20,000 psi for another 2 minutes, which
generated a 13 mm pellet. Infrared spectra were obtained between 4000 and 400 cm-1 on
a FT-IR Nicolet 20 SXB Bench in the Major Analytical Instrument Center in the UF
Material Science and Engineering Department. Interferences from KBr were cancelled
by subtracting a standard KBr spectra from the sample spectra. The crystallinity index,
(CI) measures the degree of P043- band splitting and is defined by:
CI= (A605 + A565)/(A595)
where Ax is the absorbance at wave number x (Shemesh, 1990), assuming a straight
baseline between 700 and 500 cm1 (Fig. 2-2). An estimate of the carbonate content is
given by the absorption ratio of the height of the carbonate peak at 1428 cm-1 to the
height of the phosphate peak at 1042 cm- of the FT-IR spectrum (Featherstone et al.,
1984; Lee-Thorp and van der Merwe, 1991; Wright and Schwarcz, 1996; Stuart-Williams
et al., 1996), that is:
Stable Isotope Analyses
For each of the three centra, from 24 to 28 microsamples of ~5 mg each were
drilled with a low speed Foredom drill across the growth axis starting from the center and
ending at the external margin. As far as practicable, the goal was to sample each annulus
twice, i.e., once in the dark portion and another in the light portion of the mineralized
bone. Sample powders were treated with standard isotope preparation techniques (e.g.,
MacFadden et al., 1999) used to analyze teeth. This included first washing with H202
overnight to remove organic contaminants and then with weak (0.1 N) acetic acid
overnight to remove mineral (principally CaCO3) contaminants, and then dried using
methanol. About 2 mg of each treated sample powder were then measured into
individual vials and placed in the automated Multiprep device for introduction into the
VG Prism mass spectrometer in the Stable Isotope Laboratory in the UF Department of
Geological Sciences. The sample runs were calibrated to internal laboratory and NBS 19
standards. The carbon and oxygen isotopic results are reported in the standard "6"
convention: 6-value = [(Rsample/Rstandard)-l] x 1,000 (parts per mil, %o), where R = 13C/12C
or 180/160, and standard is VPDB (Vienna Pee Dee Belemnite).
After the isotopic data were run and plotted against distance from origin of the
centra, they were the entered into a time series analysis program, AnalyseriesT version
1.2. This is necessary because each of the three centra are of slightly different sizes and
the growth lines do not match exactly, i.e., as measured by the distance from center.
Analyseries is traditionally used in paleoclimatological interpretation to construct age-
depth relations for sedimentary records. In Analyseries the method for establishing an
age-scale on a sedimentary record is to use a comparable well-dated signal as a reference
signal and then to optimize some measurement of the similarity between the two series,
while changing the depth scale of the first one to the age-scale of the second (Labeyrie
and Yiou, 1996). Analyseries then generates a pointer file, which allows plotting of the
isotope data from the two patterns on the same scale. The oxygen isotope values versus
distance from the origin were used to generate the pointer files and the carbon isotope
data (otherwise not discussed in this paper) was used as a check for the quality of the
match of the oxygen isotope data.
Results and Discussion
Fossil and Recent Shark Centra Mineralization and Diagenesis
FT-IR spectra indicate differences in shape and crystallinity indices between
modern shark centra and Eocene 0. obliquus centra (Fig. 2-3 and Table 2-1). Modern
shark centra (UF47943 and UF31648, FLMNH Environmental Collection) are
characterized by low CI (2.79 -2.84), low C/P values (0.32 0.34), and the absence of the
1096 cm-1 peak. The 0. obliquus centra (UF162732) have high CI (4.44 4.83), low
carbonate content (0.15 0.20), and a pronounced 1096 cm-1 peak, which indicates an
increase in crystallinity, a decrease in amount of carbonate, and the formation of a new
mineral phase after burial (francolite). There was no significant variation between the
samples treated with acetic acid compared to those samples not treated with acetic acid.
Fourier transform infrared spectroscopy allows for the evaluation of the mineral
characteristics of modern and fossil shark centra and enables detection of any diagenetic
changes to the mineralogy. The differences between the crystallinity indices of modern
shark centra and fossil 0. obliquus (Table 2-1) suggest that the three 0. obliquus centra
have been recrystallized, presumably due to diagenetic processes involving the growth of
larger crystals at the expense of smaller ones. The FT-IR spectra of the 0. obliquus
centra are indicative of two chemical changes occurring during recrystallization: (1) type
B carbonate substitution for P04-2 decreases with increasing crystallinity, which is seen in
reduced C/P values and increased CI (Fig. 2-4); and (2) an increase in fluorine content
with increasing crystallinity, which is indicated by a distinct peak at 1096 cm-1 (Fig. 2-3).
These changes in the FT-IR spectra signify a transformation from dahllite (carbonate
Table 2-1. Comparison of crystallinity index (CI) and carbonate content (C/P) from FT-
IR spectrum for samples treated with acetic acid and samples that were not
treated with acetic acid and from Eocene and modern samples.
Sample # CI C/P
Eocene 0. obliquus, UF 162732A
j102-51 4.62 0.18
j102-52 4.67 0.18
j102-53 4.62 0.19
Eocene O. obliquus, UF 162732B
j102-54 4.83 0.16
j102-56 4.51 0.20
j102-26* 4.58 0.17
j102-47* 4.53 0.15
Eocene O. obliquus, UF 162732C
j102-57 4.55 0.17
j102-58 4.57 0.17
j102-94* 4.44 0.15
j102-106* 4.45 0.16
Modem C. carcharias, UF 47943
j102-60 2.82 0.34
Modem C. carcharias, UF 211351
j102-317 2.80 0.32
Modem/. oxyrhinchus, UF 31648
j102-318 2.79 0.34
Modems. paucus, UF 211353
j102-319 2.84 0.33
*Samples that have been treated with acetic acid.
< 0.7 P043-
1900 1800 1700 1600 1500 1400 1300 1200 1100 1000 900 800 700 600 500 400
m 0.7- CO32-
1900 1800 1700 1600 1500 1400 1300 1200 1100 1000 900 800 700 600 500 400
Figure 2-3. Fourier transform infrared spectra of (a) Otodus obliquus centrum
(UF162732C, sample j102-58) (scale on y-axis 0 to 1.4) and (b) Isurus paucus
centrum (UF211353, sample j102-319) (scale on y-axis 0 to 1.2). The arrow
shows peak 1096 cm-1, which is diagnostic of francolite (F-apatite)and is only
present in the fossil (UF 162732C) specimen and not the modern specimen
hydroxyapatite) to francolite (carbonate fluorapatite). The lack of significant differences
in the FT-IR spectrum of the 0. obliquus samples treated with acetic acid and those that
were not treated suggest that these samples were not diagenetically enriched in CO3-2 and
no secondary carbonate minerals were present. The absence of a 710 cm-1 peak on the
FT-IR spectrum, which represents the presence of calcite, indicates that no secondary
carbonate minerals were present. Comparison of the FT-IR spectra from various
locations along the growth axis of a single 0. obliquus centra and between the three
centra show insignificant variation, which indicates that the chemical changes that occur
during diagenesis are uniform along the growth axis and between the three centra.
o3 o UF162732B
A UF162732B (T)
(0.2 UF162732C (T)
O U o Modern
2 2.5 3 3.5 4 4.5 5
Figure 2-4. Graph representing decrease in the amount of carbonate with an increase in
Physical Increments and Variation
Physical growth couplets annulii) are evident in the gross morphology and x-
radiographs of each of the fossil Otodus obliquus centra. Contact prints of x-rays of the
three 0. obliquus centra (Fig. 2-5) were used to enhance the visibility of the growth rings.
The counts on the three contact prints indicate a total of 19 growth couplets for each
centrum. A growth couplet is defined as a band pair, composed of one opaque (darker)
band and one lighter band. The location of the 19 growth couplets is consistent among
the three centra and fit a characteristic growth function (von Bertalanffy, 1938; and von
Bertalanffy, 1960), i.e., with the most rapid growth during early ontogeny and
incrementally decreasing growth rate during later ontogeny. Given the fact that all three
centra come from the same individual, the observation of 19 equivalent band couplets is
both expected and corroborates these as growth related phenomena. Assuming one
growth couplet is deposited annually as proposed for other lamnids (Cailliet at al. 1983;
Campana et al. 2002; Labs-Hochstein, submitted) this shark has a minimum age of 19
Figure 2-5. Contact prints of the three Otodus obliquus (UF 162732) x-rays. The "*"
symbols indicate dark growth bands on centra.
Stable Isotope (6SO ) Signal Archived in Eocene Otodus obliquus centra
A series of microsamples taken along the growth axis of three centra (UF
162732A, UF162732B, and UF162732C) of Otodus obliquus reveals a systematic pattern
of change in 6180 values (Fig. 2-6; Table 2-2). There appears to be a systematic,
sinusoidal, variation of 6180 values representing at least eight isotopic cycles. The
extremes of these cycles are interpreted to represent warm seasonal signals, with lower
6180 values between approximately -3.5 to -4.0%o and cold seasonal signal with higher
6180 values between about -2.0 to -3.3%o.
There are several points of discussion concerning these data. Firstly, two adjacent
microsamples were taken, so far as possible, to correspond with dark-light band couplets
within an annulus. The results indicate that more positive 6180 values correspond to the
darker bands, confirming the prediction that these represent a colder signal. Conversely,
the more negative 6180 values were taken in the lighter band, indicating a warmer signal.
This correspondence, plus the non-random pattern of isotopic variation therefore suggests
that these isotopic data are archiving a record of paleoclimate cycles (although perhaps
not the actual amplitude of paleotemperature change, see below).
Secondly, the oxygen isotopic data of all three centra have eight annual growth
band couplets. However, there is no consistency in the amplitude of the bands. In
different centra the same growth bands have different isotopic values, and the extremes of
each annual cycle are not recorded uniformly in each of the three centra. This may be
due to several reasons, such as; (1) the growth band is not sampled in the exact same
place on each of the three centra; (2) there may be an ossification order such that the
centra closer to the head ossify before the centra towards the tail or vice versa
(Ridewood, 1921); (3) the seasonal variation may not have been as strong at various
times during the sharks life; and/or (4) possible dampening/amplification of the original
signal by diagenesis.
Finally, there are fewer annuli indicated from the isotopes (8) than those observed
from the physical growth couplets (19). Otodus obliquus could have been
homoeothermic (maintain a body temperature above its surroundings). All modern day
lamnids have the ability to be homoeothermic (Campango, 2002) and it is not known
when lamnids evolved this characteristic. However since porebeagles, makos, and great
whites are all homoeothermic it suggest that the common ancestor may also have been
homoeothermic. Using molecular clocks calibrated for sharks Martin (1996)
demonstrated that all three genera of lamnids diverged at nearly identical times during the
Paleocene or early Eocene. Another possibility is that all modern lamnids independently
evolved homeothermy. Only seven out of the more than 360 known modern species of
sharks are known to be homoeothermic, the five species of modem lamnids and two
species of modem thresher sharks (Campango, 2002). Therefore homeothermy is rare in
modern sharks and most likely evolved in lamnids through a common ancestor. The
possibility exists that the oxygen isotopic data is showing the shifts in body temperature
in 0. obliquus as it came in contact with varying temperatures of water and not the actual
water temperature. The oxygen isotopic signal gives an age of half what the growth ring
counts estimates suggesting that this shark may not have been migrating into waters with
enough variation in the oxygen isotopic signal on an annual basis or it was not encounter
waters that changed its body temperature on an annual basis if 0. obliquus was
-2.5 \ I --UF162732C
10 15 20 25 30 35 40 45 50 55
Distance from Center (mm)
Figure 2-6. Centrum of Otodus obliquus, UF 162732A, from the early Eocene of
Morocco showing exact sampling locations (grooves, top) and plot of
variation in 6180c (bottom), the three plots represent the correlation between
each of the centra produced by using AnalyseriesTM. The gray lines connect
the dark growth bands that correspond to the possible annual signal.
Table 2-2. Stable isotopic data for three vertebral centra of Otodus obliquus, UF 162732,
from the Early Eocene of Morocco.
mm 618~ 0/oo mm 618O 0/oo mm 68Ce /oo
Because of their cartilaginous skeletons, which characteristically do not fossilize,
the secondarily calcified centra provide a unique opportunity to assess incremental
growth and age determination in fossil sharks. The FT-IR spectra indicate that the three
0. obliquus centra have undergone diagenesis. Modem skeletal tissue is composed of
carbonate hydroxyapatite while according to the FT-IR spectra the 0. obliquus centra
have francolite (carbonate fluorapatite) and have lost carbonate and organic relative to
modern shark centra. All three centra have undergone similar diagenesis indicating that
in this case the diagenesis was uniform throughout this specimen. There was no
difference between the acetic acid treated samples and those that were not treated,
indicating that these specimens did not have any secondary carbonate present and the
acetic acid does not affect the structural carbonate.
Despite this degree of alteration, two important conclusions derived from this study
are that: (1) the incremental growth banding, i.e., annuli, retain their original physical
structure; and (2) the non-random pattern of oxygen isotopic variation therefore suggests
that these isotopic data are archiving a record of paleoclimate cycles. With regard to the
latter observation, I caution that the oxygen isotopic variation may represent a signal
damped/amplified by diagenesis or the possibility exists that 0. obliquus was
homoeothermic and therefore the shifts in oxygen isotopes represent changes in body
temperature as it came in contact water of different temperatures. Accordingly, I do not
advocate using these data to attempt calculations of paleotemperatures in the early
Eocene seas. The oxygen isotopic signal gives an age of about half of what the growth
ring counts estimates suggest (assuming one growth ring pair per year). One explanation
of this observation is that this shark may not have been migrating into waters with
enough oxygen isotopic variation on an annual basis. Another explanation, if 0. obliquus
was warm-bodied, is it was not encountering waters that changed its body temperature on
an annual. These findings are similar to modern day lamnids, where by larger sharks
oxygen isotope estimates yield ages equal to approximately 1/2 of the ring counts (Labs-
The analysis of Otodus obliquus centra from the early Eocene of Morocco
potentially have broad ramifications for understanding the evolution of developmental
strategies in fossil sharks. This application can potentially answer some unresolved
questions about the developmental mechanisms that resulted in huge body size in fossil
lamnid sharks (Gottfried et al., 1996) such as Miocene Carcharodon megalodon, which is
a close relative of 0. obliquus. In all such future studies, however, analytical techniques
that assess diagenesis, such as FT-IR, should be used in combination with isotopic studies
to produce the most insightful analysis of fossil shark paleobiology.
QUANTIFICATION OF DIAGENESIS IN CENOZOIC SHARKS: ELEMENTAL AND
Fossilized vertebrate skeletal tissues, including teeth and bone, have recently
received considerable attention as geochemical archives of paleoecological and
paleoenvironmental information (Piper, 1974; Kolodny et al., 1983; Elderfield and
Pagett, 1986; Kolodny and Luz, 1991; Lecuyer et al., 1993; Picard et al., 1998; Shields
and Stille, 2001; Picard et al., 2002; MacFadden et al., 2004; Puceat et al., 2004). In these
studies, fossil tooth enamel has been the preferred material for analysis because it is
compact, relatively non-porous mineral and consists of >95% hydroxyapatite. However,
isotopic data have been used to eludicate paleobiological information (e.g., to reconstruct
dinosaur physiology; Barrick and Showers, 1994). 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 (i.e., composed of soft supporting tissue that does not fossilize), the
cartilage is replaced in the vertebral centra by 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
Branstetter et al., 1987). These growth rings are easily observed in both modern and
fossilized shark centra. During a related research project investigating stable isotopic
signatures archived in fossil shark centra (MacFadden et al. 2004), we became interested
in the extent of diagenesis and how it may have affected 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, carbonate
content, and major, minor, trace elemental concentrations. The sharks are all from the
group known as the superfamily Lamnoidea (Capetta 1987), that includes the modern
great white (Carcharodon carcharias), and six closely related, extinct species that range
in geologic age from the Cretaceous to the Pliestocene. 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.
Bone Chemistry and Diagenesis
Stable isotopes and rare earth elements (REE) of biogenic apatites have been used
for paleoclimate 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; Lecuyer et al.,
1993; Picard et al., 1998; Shields and Stille, 2001; Picard et al., 2002; MacFadden et al.,
2004; Puceat 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 the biogenic apatite (Reiche et al., 2003); therefore, the original
chemical signatures of biogenic apatites may be modified through diagenesis, resulting in
the interpretation of erroneous biological signals (Puceat et al., 2004).
Modem bone is composed of carbonate hydroxyapatite, Calo(PO4)6(CO3)x(OH)2-
2x, 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 small
crystallite size and high surface area of the bone hydroxyapatite (Trueman, 1999). Many
substitutions are possible for both the anions and cations in biogenic hydroxyapatite
(Table 3-1; Nathan, 1981). In modern biogenic apatites, carbonate (C032-) can substitute
for either OH- (A site) or P043- (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 distorts the crystal lattice increasing the solubility of biogenic
apatite (Nelson, 1981; Nelson et al. 1983).
During fossilization, the biogenic apatite alters to a more stable, less reactive form
of apatite (carbonate fluorapatite some times referred to as "francolite") by losing
carbonate and hydroxyl ions and gaining fluoride (Nathan and Sass, 1983; Newesely,
1989; Greene et al., 2004). The loss of carbonate decreases the defect densities within
the hydroxyapatite lattice resulting in an increase in crystallite size and order and
decreaed solubility relative to carbonate hydroxyapatite (Greene et al., 2004).
Table 3-1. Some possible substitutions in the apatite crystal structure.
Constituent ion Substituting ion
Ca Na K Sr Mn Mg Zn ,Ba Sc3,
Y3+, REEs, U4
P043- CO32-, SO42-, Cr042-, CO3sF3-,C03 OH4-, SiO44-
OH- F-, C1-, Br, 02-
Through the processes of diagenesis, trace element concentrations can either
increase or decrease relative to unaltered bone concentration. This phenomenon is well-
documented and described in the literature (Elderfield and Pagett, 1986; Wright et al.,
1987; Williams, 1988; Grandjean and Albarede, 1989; Koeppekastrop and DeCarlo,
1992; Grandjean-Lecuyer et al., 1993; Denys et al., 1996; Hubert et al., 1996; Laenen et
al., 1997; Reynard et al., 1999; Trueman, 1999; Staron et al., 2001). Trace elements are
most likely incorporated into bone apatite during early diagenesis through the process of
substitution. Trace element signatures acquired during the initial stages of diagenesis
appear to be stable and resistant to later modification (Bernat, 1975; Grandjean and
Albarede, 1989; Grandjean-Lecuyer et al., 1993). Because of similar ionic size, REE3
ions are easily substituted into the Ca2+ site by means of coupled substitution (Whittacker
and Munts, 1970). Adsoprtion of trace elements onto the surface may also be a
qunatatively significant mode of uptakeof biogenic apatite (Reynard et al., 1999).
Surface adsorbed species are typically only weakly bound to the mieral surfaces, and
therefore ions adsorbed are susceptible to exchange as long as the crystal surface remains
exposed (Reynard et al., 1999). However, if during diagenesis the inter-crystalline
porosity is closed, then individual crystallite surfaces will be protected from further
exchange. Ultimately, the final trace element composition of the biogenic apatite is
controlled by the concentration of trace elements in the system, apatite-fluid partition
coefficienst, chemistry of the burial microenvironment, bone microstructure, and length
of exposure (Trueman, 1999).
Rare Earth Elements
Although the REE typically exist in the 3+ oxidation state two exceptions are
Cerium and Europium. Cerium can undergo oxidation in seawater from the solvated 3+
state to the relatively insoluble 4+ state (de Baar et al., 1985a). 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) and under
reducing conditions Ce3+ may be released back into the water column or pore waters.
Therefore when Ce is depleted (i.e., under oxic conditions) in the water column a
negative Ce anomaly (Ce anom.) is present and vice versa (German and Elderfield, 1990).
Europium may undergo oxidation from the Eu3+ to Eu2, which is significant in the
oceans because of its preferential ability of Eu2+ to substitute for Ca2+ in apatites
(Elderfield, 1988). de Barr et al. (1985a) illustrate that in both the Atlantic and Pacific
Oceans all REE, with the exception of Ce, increase with water depth. Concentrations of
Ce decrease with water depth and therefore the negative Ce anomaly observed increases
with depth (de Baar et al., 1985a). Thus, when the seawater signal is preserved in
biogenic apatites the Ce anomaly may be a useful relative indicator of water depth at
which fossils have been deposited.
Fossil biogenic apatites contain several tens to several hundreds parts per million
(ppm) of REE, whereas REE maximum concentration 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). Bernat (1975) reported high REE
concentrations in ichthyoliths from the upper-most 4 cm of sediment of ocean cores.
These ichthyoliths exhibit bulk REE patterns similar to overlying waters. Furthermore,
Bernat (1975) analyzed ichthyoliths from 4 to 600 cm of the sediments, and showed that
their REE compositions were similar to the ichthyoliths from the upper-most 4 cm of the
same cores. These results suggest that in this case ichthyoliths inherit a REE composition
directly from seawater at the sediment/seawater interface during early diagenesis, with
little or no fractionation. Direct uptake of REE in biogenic apatites from pore waters
and/or seawater raises serious problems. Assuming that REE are directly taken up
through advection of pore waters, about a ton of pore water would be needed to give the
biogenic apatite enough REE to fit observed concentrations (several tens to several
hundreds parts per million, ppm). Such a water/rock ratio would require an exceedingly
high flux of pore water through the sedimentary column, and reasonably this cannot be
considered for the cause of the enrichment of biogenic phosphates (Grandjean et al.,
1987; Grandjean and Albarede, 1989; Grandjean-Lecuyer et al., 1993). 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 specific surfaces, which easily absorb 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, oxy-hydroxide flocs) and organic
(pellets, organic debris) phases. The decay of the REE rich carriers at the
sediment/seawater interface, associated with that of biogenic apatite, and the resulting
reducing conditions eventually cause the dissolution of Fe-Mn oxy-hydroxides and favor
transfer of REE from large volumes of seawater to recrystallized biogenic apatite in a
rather short period of time (Grandjean et al., 1989). This extension of Bernat's (1975)
model implies that both that the ultimate source of phosphatic REE is seawater, and there
is an early diagenetic transfer to the phosphate through a short-lived phase made of
oxihydroxides and organic detritus. Upon completion of the early diagenetic processes
and once most oxihydroxides have been dissolved, phosphate remains the major non-
detrital REE repository in sediments, so its REE concentrations must reflect to a large
extent the flux of seawater derived REE exclusive of the detrital particulate accumulation
(Grandjean et al., 1989).
Variations in host lithology of marine biogenic apatites may influence the REE
contents due to differences in permeability, the flux of REE from diagenetic fluids
expelled from sediments, and organic and oxy-hydroxide contents (Grandjean-Lecuyer et
al., 1993; Lecuyer et al., 2004). Terrestrially derived sediments have shale-normalized
REE patterns that carry a shale-like signal (i.e., flatten REE pattern) and no Ceanom. is
expected from common fine-grained detrital material (Grandjean et al., 1987). REE in
pore waters are derived from the surrounding sediment particles and development of
large pore water concentration gradients will allow fluxes of REE from sediments to
seawater (i.e., diagenetic fluids expelled from sediments into seawater; Elderfield and
Sholkovitz, 1987). Therefore, REE contents of biogenic apatites deposited in terrestrial
derived sediments (clays and sands) may have flattened shale-normalized REE 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
carbonatess and phosphorites) with little to no terrestrial input have diagenetic fluids
reflecting the composition of the overlying water column. Biogenic apatite deposited in
marine precipitated sediments should show a seawater REE pattern since the diagenetic
signature would be the same as the overlying water column (Lecuyer et al., 2004).
Therefore, the REE signature in fossil biogenic apatites results from a mass balance
between the flux of REE from decaying organic and oxy-hydroxides (primary carriers
with seawater signature), the flux of REE from diagenetic fluids expelled from sediments
(diagenetic signature),and the flux of REE from rivers (detrital signature) (Grandjean and
Materials and Methods
This study analyzes the composition of nine shark centra (Fig. 3-1), two modern
great white, Carcharodon carcharias, and seven fossil specimens ranging in age from
Cretaceous to Pliocene (Table 3-2). These specimens were selected because they all are
within the monophyletic superfamily Lamnoidea, the group that includes great white
sharks, makos (Isurus), and their close fossil relatives. Likewise, these sharks were
selected because they span an age range from Cretaceous to Recent and are widely
distributed geographically. A broad geographic distribution of fossils should illuminate
the effects of different extents and environments of diagenetic alteration. The vertebral
centra were chosen because they are the primary ossified skeletal tissue that fossilizes in
sharks (i.e., other than teeth).
Two analytical techniques were used to determine the chemical and mineralogical
properties of the nine shark centra: (1) Fourier Transform Infrared Spectroscopy (FT-IR).
FT-IR, which is used here to determine crystallinity. FT-IR has advantages over x-ray
diffraction (XRD), because only a small amount of sample is required (<1 mg),
preparation is easier and produces more accurate results; and additionally, carbonate
content can be assessed from FT-IR. (2) Inductively Coupled Plasma Mass Spectrometry
(ICPMS), which allows for the determination of the major, minor, and trace elements in
modern and fossil shark centra. ICPMS is a comprehensive technique that is extremely
sensitive (detection limits in the ppb range for many elements in aqueous solution). The
high level of relative accuracy (1 to 2%) coupled with sensitivity allows analysis at
concentrations ranging over more than nine orders of magnitude (Montaser, 1998).
Table 3-2. Lamnid shark specimens used in this study.
ecie M I Lcait Age Sediment & Depositional
Species Museum ID Locality Eim
Carchar n BT0433 E coast, South Africa
Carcharon UF211351 Islamorada, Florida
Pliocene Shallow bay sandstone
Isurus hastalis UF211358 Pisco Fm., Peru ene a a a
(Brand et al., 2004)
Carcharodon 120A Saitama Prefecture, Miocene Nearshore sandy siltstone
megalodon Japan (Hayashi et al., 2003)
Carcharodon OU22261 Kokoamu Greensand, Oligocene Shelf glauconitic sand
angustidens New Zealand (Ayress, 1993)
Carcharodon Brussels Sand, Oligocene Near shore shelf sandstone
au s EF809A Belm (Hooyberghs, 1990; and
SHerman et al. 2000)
Eocene Shelf phosphorite (Lancelot
Otodus obliquus UF162732B Oued Zem, Morocco E e Sf p
and Seibold, 1978)
Eocene Shelf phosphorite (Lancelot
Otodus obliquus UF162732D Oued Zem, Morocco and Seibo, 1978)
and Seibold, 1978)
Cretaceous Shallow epicontinental sea
reotxyrina UF211357 Niobrara Fm., Kansas outer shelf chalky shale
mantelli (Hattin, 1981)
Fourier Transform Infrared Spectroscopy (FT-IR)
Three -1 mg samples were drilled with a low-speed Foredom drill from each of
the nine centra. Samples were taken along the growth axes, i.e., one from the center, one
from the middle, and one from the edge of each centrum. Potassium Bromide (KBr)
pellets were prepared using the method discussed in MacFadden et al. (2004). Infrared
spectra were obtained between 4000 and 400 cm'1 on a FT-IR Nicolet 20 SXB Bench
housed at the Major Analytical Instrument Center in the UF Material Science and
Engineering Department. Interferences from KBr were cancelled by subtracting a
standard KBr spectrum from the sample spectra. The size distribution of crystallites and
the degree of substitution order-disorder within the crystal lattice of biogenic apatite can
be determined by calculating the crystallinity index (CI) from the extent of phosphate
peak splitting at 565-605 cm-1 in an FT-IR spectrum. Apatites with larger, more ordered
crystals show greater separation of these peaks and a higher CI (Shemesh, 1990; Wright
and Schwarcz, 1996). An estimate of the carbonate content is given by the absorption
ratio of the height of the carbonate peak at 1428 cm-1 to the height of the phosphate peak
at 1042 cm-1 of the FT-IR spectrum (Featherstone et al., 1984; Lee-Thorp and van der
Merwe, 1991; Stuart-Williams et al., 1996; Wright and Schwarcz, 1996).
Figure 3-1. Nine vertebral centra used in this study. A. Carcharodon carcharias
(BT0433) B. Carcharodon carcharias (UF211351) C. Isurus hastalis D.
Carcharodon megalodon E. Carcharodon angustidens F. Carcharodon
auriculatus G. Otodus obliquus H. Otodus obliquus I. Creotxyrhina
Elemental Analysis (ICPMS)
Approximately 6 mg of bulk sample was drilled with a slow speed Foredom drill
from each of the nine centra. 5 mg of each sample were weighed out and placed into 3
mL Savillex vials, dissolved in 1 mL of 3M HNO3, and heated overnight. Samples were
allowed to cool and then dried. Next 2 mL of 1% HNO3 was added, heated overnight, and
allowed to cool. Samples were analyzed on an Element 2 High Resolution Inductively
Coupled Plasma Mass Spectrometer (HR-ICP-MS) at the Center for Trace Element
Analysis at the University of Southern Mississippi. All samples were corrected by
subtracting the blank, corrected for instrumental drift based on internal machine standards
that were analyzed during the run (initial quantification based on comparing the corrected
ion counts of the samples with ion count for the standards), and corrected ion counts to a
constant response to the known amount.
All REE values were shale-normalized to PAAS (Post-Archean Australian Shale
Standard) in order to clearly illuminate enrichment-depletion trends relative to average
crust (e.g., Grandjean et al., 1988; Elderfield et al., 1990; Grandjean-Lecuyer et al., 1993;
Reynard et al., 1999; Trueman and Tuross, 2001). The Ce anomaly (Ceanom) was
calculated from Ceanom=Log[3CeN/(2LaN+NdN)] (Elderfield and Greaves, 1982).
FT-IR spectra of the modern and fossil shark centra are shown in Fig. 3-2. Both
the modem 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 H20 bands (which
usually mask the OH- band at 3567 cm-1) and the presence of organic represented by the
three amide group bands amidee I 1660 cm-1, amide II 1550 cm-1, and amide III 1236 cm
1, mean values). The FT-IR spectra for fossil specimen are characterized by reduced H20
bands, lack of OH- band, and absence of one or more of the amide group bands (Fig.3-2).
There are three intense phosphate (P043-) absorption bands: the main absorbance peak is
recorded at 1041 cm-1 and a doublet at 605 cm-1 and ~ 568 cm-1, consistent with
previous studies (Shemesh, 1990). In the modem specimens, the 605 cm-1 absorption
band has a smaller intensity than the 568 cm-1 band and but the reverse holds for the
fossil specimens (Fig. 3-3). The modern specimens have an average crystallinity index
(CI) of 2.83, while in the fossil specimens CI is increased to 3.19-5.39 (Table 3-3; Fig. 3-
4). B-type carbonate substitution (replacement of P043- by CO32-; Shemesh, 1990; Dahm
and Risnes, 1999) is represented by a set of absorption bands at 1460 cm-1, 1428 cm-1,
and 870 cm-1 (average values). Carbonate content (C/P) is much greater in modern
specimens (0.35 and 0.43) than fossil specimens (range from 0.10 to 0.29) (Table 3-3;
Fig. 3-4). The lack of the 713 cm-1 absorbance band in all samples indicates that there is
no authigenic calcite present, and the 1092 cm-1 band (average value), which is found
only in the fossil specimens, demonstrates the presence of fluorine (Fig. 3-2).
The effects of diagenesis on elemental concentrations can be assessed by
comparing the modem unaltered centra with the altered fossil centra. This is illustrated
by the isocon plots (Grant, 1982) in Fig. 3-5. The linear trends (labeled isocon in Fig. 3-
5) are the average of the two modem shark centra elemental compositions and represent
no elemental loss or gain during diagenesis. In modem sharks the major and minor
elements vary (Table 3-3), therefore in the isocon plot of the major and minor elements
Figure 3-2. FT-IR spectra of all nine shark vertebral centra illustrating the differences
from modern (C. carcharias) to fossil biogenic apatites.
605 cm-1 A
800 750 700 650 600 550 500 450 400
Figure 3-3. FT-IR spectra from 400 to 850 cm-1, illustrating the y4 P043- band
differences between modem (solid line) to fossil (dashed line) shark centra.
2.50 3.00 3.50 4.00
Crystalllnlty Index (CI)
* O. obliquus
1 O. obliquus
A C. carcharias
+ C. mantelli
X I. hastalis
A C auriculatus
o C. angustidens
o C. megalodon
4.50 5.00 5.50 600
Figure 3-4. Carbonate content (C/P) vs. crystallinity index (CI) of the nine shark vertebral
Table 3-3. Elemental and mineralogical data of nine shark vertebral centra. Elemental
concentration are in ppm. Crystallinity index (CI) and carbonate content
(C/P) were calculated from FT-IR spectra.
Specimen Ca P Mg Na Zn Fe K Sr Al Si Mn Ba Pb B
C. carcharias 732027 59982.3 1752.9 4053.4 904.8 177.2 1133.6 352.5 266.0 28.8 4.8 1.5 9.1 9.1
C. carcharias 131144.2 1156974 4239.8 3446.9 2318.8 523.9 314.3 644.8 189.6 1663.7 12.4 1.9 47.0 23.8
I hastials 179681.8 132370.8 5234.1 18385.2 449.6 3408.4 2242.6 1031.2 2999.2 118.2 494.6 30.6 13.7 46.9
C. megalodon 129279.7 78068.0 1521.4 1540.9 575.5 2790.4 239.2 1716.6 2592.8 543.9 667.9 444.4 9.9 49.8
C. auriculatus 88019.48 61402.8 2027.8 1879.8 1417.0 957.9 183,9 519.9 373.3 953.7 31.7 26.2 29.7 24.5
C angustidens 101257.9 61242.3 489.8 1600.8 18.8 883.8 588.5 512.9 568,7 117.6 11.0 24.6 2.6 17.9
O. obliquus 159115.7 116988.6 1029.3 3260.4 942.4 133.0 98.9 653.5 149.5 0.0 3.8 62.9 8.1 17.5
O. obliquus 186700.9 130473.9 1149.3 37458.4 853.6 113.1 86.0 791.9 106.4 0.0 3.2 76.9 7.9 21.6
C. mantelli 89909.63 58346.8 648.6 1151.9 516.8 1109.5 145.9 1326.5 553.6 0.0 10.2 283.7 4.7 19.9
Specimen Y La Ce Nd Sm Gd Dy Yb EREE U La/Yb Ca/P Ce I CI C/P
C carcharias 0.56 0.23 0.09 0.14 0.02 0.04 0.04 0.03 1.15 0.28 7.7 1.2 -0.73 2.82 0.43
C. carcharias 0.62 0.30 0.57 0.34 0.07 0.07 0.09 0.04 2.11 0.27 7.7 1.1 -0.11 2.83 0.34
I hastials 9.94 15,04 22.65 9.93 1.85 1.55 1.31 0.97 63.2 6.66 15.5 1.4 -0.09 4.30 0.14
C. megalodon 16.14 23.78 41.26 16.22 3.37 3.29 2.59 1.03 107.7 1.59 23.1 1.7 -0.05 5.39 0.10
C. ariculatus 75.95 104.39 128.85 75.73 13.37 14.71 10.25 3.8 427.1 24.84 27.5 1.4 -0.21 3.41 0.24
C. angustidens 21.11 12.46 13.29 10.85 2.16 2.36 2.19 1.01 65.4 32.9 12.3 1.7 -0.18 3.60 0.30
O. obliquus 31.64 14.19 3.79 6.57 1,26 1.92 2.07 1,86 63.3 48.05 7.7 1.4 -0,85 4.55 0.17
O. obliquus 32.89 3.76 6.88 1.29 1.98 2.19 2.00 66.4 57.92 7.6 1.4 -0.85 4.55 0 17
C. mantelli 98.90 66.39 91.75 39.43 7.40 8.63 9.99 6.34 328.8 27.26 10.5 1.5 -0.14 4.02 0.21
(Fig. 3-5A) the elemental concentrations range in modern sharks for each element is
indicated by a vertical line. Most fossil sharks have comparable Ca, P, Zn, Si, Pb, and B
concentrations to the modem sharks (Fig. 3-5A). The minor elements Mg, Na, K, Sr, and
Mn are more variable, with some fossil specimens exhibiting concentrations similar to
modern sharks while Fe, Al, and Ba have no fossils with similar concentrations as
modern sharks (Fig. 3-5A). All fossil specimens are enriched in Ba relative to modern,
while most fossils are enriched relative to modern in B, Mn, Al, Fe, and Sr (Fig. 3-5A).
All fossil sharks are depleted relative to modem in Pb, while most fossils are depleted
relative to modern in K, Si, Zn, Mg, and Na (Fig. 3-5A). It can be seen in Fig. 3-5B that
Y, U, and REE's are enriched in fossil bone relative to modern. Fig. 3-5C shows the
Ba/Ca ratio for the fossils are higher than in modern sharks, because diagenesis has
significantly enriched Ba (one to two orders of magnitude greater in fossils than modem
shark centra) in all the fossil centra used in this study. The fossil centra Ca/P ratios are
slightly higher (0.2-0.5 higher) than modern sharks (Fig. 3-5C).
The shale-normalized REE concentrations of the shark centra are enriched relative
to seawater by about 105-107 (Fig. 3-6). The modem specimen C. carcharias
(UF211351), which was caught off the east coast of Florida, has a shale-normalized REE
pattern similar to seawater but is slightly flattened and depleted in Yb, a small negative
Ceanom. (-0.11), a higher concentration of Ce than the other C. carcharias (BT0433), and
(La/Sm)N and (La/Yb)N ratios that overlap those found in coastal waters (Fig. 3-6 and
Fig. 3-7). C. carcharias (BT0433), which was caught off the east coast of South Africa,
has a seawater-like shale-normalized REE pattern (enriched with HREE), has depleted Ce
concentrations, a much larger Ceanom. (-0.73) than the other C. carcharias (UF211351),
Ba rn sr Z
B \ 0
1 10 100 1000 10000 100000 100t
Modem Shark Centra Concentrations (ppm)
Nd & GdU$
A Yb 8S
+ + x x
0 A o 0
Modem Shark Centra Concentrations (ppm)
* 0. obliquus
o 0. obliquus
+ C. mantelli
A C. auriculatus
* C. angustidens
0.00001 U0.001 0.001 U.01 U.1 1 1U
Modem Shark Centra Concentrations (ppm)
. Isocon plots (Grant, 1982) depicting variations in fossil elemental
concentrations and ratios due to diagenesis. The isocon lines are the average
of the elemental concentrations in the two modern shark centra. A. Isocon
plot of the major and minor elements, with the red vertical lines represent the
variation between the elemental concentrations in the two modern shark
centra. B. Isocon plot of Y and REE. C. Isocon plot of Ba/Ca, Sr/ Ca, and
U.UUUU ... .... .. .. -.
and (La/Sm)N and (La/Yb)N ratios that overlap with oceanic waters (Fig. 3-6 and Fig. 3-
Fossil specimens (Fig. 3-6) may be divided into four groups based on their REE
composition. The first group, which includes the two 0. obliquus specimens, have shale-
normalized REE patterns similar to the modern C. carcharias (BT0433) and seawater
(HREE enriched) (Fig. 3-6A), show a minimum at Sm (disregarding Ce),
La Ce Nd Sm Gd Dy Yb
La Ce Nd Sm Gd Dy Yb
Figure 3-6. PAAS normalized REE of the nine vertebral centra divided into four
diagenetic groups. A. First group has seawater-like REEN pattern. B. Second
and third groups have mixed shale- and seawater-like REEN patterns. C.
Fourth group has shale-like REEN patterns. Seawater is average seawater
(Elderfield and Greaves, 1982) and multiplied by 106 and modern C.
carcharias, UF211351 and BT0433, are multiplied by 101.
0.1 C. carcharias(UF2 1351)
001 C N S
La Ce Nd Sm Gd Dy Yb
VI r j (r Coastal waters
0.1- AA* A A A
@ 00 :00
Continental waters w
0.1 1 10
Figure 3-7. Compilation of observed (La/Yb)N vs. (La/Sm)N in biogenic apatites of
various ages and types, and fresh and oceanic waters (based on Reynard et al.,
1999). *, Jurassic fish and reptile teeth (Picard et al., 2002); A, Tertiary and
Mesozoic fish teeth (Grandjean et al., 1988; Grandjean and Albarede, 1989);
Cretaceous reptile and dinosaur bones (Samoilov and Benjamini, 1996) and
Jurassic coprolites (Kemp and Trueman, 2002); 0, North Atlantic seawater
from 0-4000 m depth (Elderfield and Greaves, 1982), Southern Ocean from 0-
4000 m depth (German et al., 1990 ) and North Pacific surface and deepwater
(Piepgrass and Jacobsen ,1992); *, bottom waters (German et al., 1991); 0,
coastal waters (Elderfield and Sholkovitz, 1987) and coastal waters (Hoyle et
al., 1984); -, Scotland river waters (Hoyle et al., 1984); @, anoxic waters
(Elderfield and Sholkovitz, 1987); and current research, A and 0, C.
carcharias, X, I. hastalis, 0, C. angustidens, A, C. auriculatus, o, C.
megalodon, 1 and 0, O. obliquus, and +, C. mantelli. Oceanic, Coastal, and
Continental boxes are based on values given by Reynard et al. (1999). Blue
circles indicate samples with seawater-like REEN patterns, green circles
indicate samples with shale-like REEN patterns, and red circles indicate
samples with mixed seawater- and shale-like REEN patterns.
large Ceanom. (greater than -0.3), and an (La/Yb)N ratio that falls within coastal waters. 0.
obliquus is the only sample that has a (La/Sm)N ratio greater than oceanic waters. The
second group, which includes C. angustidens and C. mantelli, have shale-normalized
REE patterns similar to the modern shark pattern of C. carcharias (UF211351) and close
to seawater but with some flattening (Fig. 3-6B), show a minimum at Nd (disregarding
Ce), (La/Yb)N and (La/Sm)N ratios that fall within or just outside of coastal waters (Fig.
3-7), and small negative Ceanom. (less than -0.3). The third group, which includes C.
auriculatus, shows a minimum in HREE, shale-normalized REE pattern between
seawater and the modern C. carcharias (UF211351) pattern (Fig. 3-6B), (La/Yb)N ratio
above marine and continental waters, (La/Sm)N ratio within coastal and oceanic waters
(Fig. 3-7), and a small negative Ceanom. (less than -0.3). The fourth group, which includes
C. megalodon and I. hastalis, show a minimum at Yb, a flat shale-normalized pattern or a
maximum in the heavy-middle REE (Fig3- 6C), a high (La/Yb)N ratios (above seawater
and continental waters), and (La/Sm)N ratio within coastal and oceanic waters (Fig. 3-7).
Mineralogical Characterization of Centra
Fossil specimens have lost most, if not all, of their organic content through
diagenetic processes and contain less absorbed (3430 cm-1 band) and structural H20
(3330 cm-1 band; Holcomb and Young, 1980; Michel et al., 1995) than modern
specimens. The weak intensity of the absorption band near 1660 cm-1, corresponding to
vCONH of the amide group amidee I), and the absence of the two other amide bands
amidee II and III) signify a significant loss of organic (Reiche et al., 2003) in the fossils.
In contrast, the FT-IR spectra of the modern specimens show intense amide I, II, and III
bands and thus indicate the presence of organic matter (Reiche at al., 2003).
When comparing the two modem specimens, the CIs are similar, but the carbonate
content differs, with the South African C. carcharias shark (BT0433) having a higher
carbonate content than the one from Florida. This may be due to due to natural
variability within a shark species or overlapping absorption peaks that have both
carbonate substitution in the A-site for OH- and B-site for PO43- (i.e., peaks 1460 cm-1
and 870 cm-1) are not included in the estimation of carbonate content. Since, the
carbonate content estimation is only normalized by a phosphate peak, the use of A-site
substituted carbonate would not be an accurate representation of carbonate content.
Usually in modern specimens the OH- band at 3567 cm-1 is masked by the water bands
and is completely absent in fossil specimens (Fig. 3-2), making it difficult for
normalization. Lower C/P and higher CI in fossil specimens (Fig. 3-4) indicate
diagenetic loss of carbonate during recrystallization and possibly dissolution of the
Implications for Diagenetic and Biological Signal Reconstruction
The shale normalized REE patterns of the two modern shark specimens (Fig. 3-6)
have similar patterns to seawater, however, C. carcharias (BT0433) from South Africa
has a more negative Ceanom. and greater HREE enrichment than UF211351 from Florida.
The larger negative Ceanom. in C. carcharias (BT0433) may indicate that sharks from the
coast of Africa live and/or spent most of their time at greater depths than sharks from the
East coast of North America. Alternatively, larger great white sharks may feed and/or
spend a majority of their time at shallower depths than smaller great whites (UF211351
was 12'9" in length when captured and BT0433 was 6'9" in length when captured).
Further study of modem great white shark behavior is needed to fully understand the bulk
chemistry variations within this species.
The first fossil group, which includes the two 0. obliquus specimens, have
seawater-like patterns indicating that REE enrichment involved quantitative uptake of
REE without fractionation. In other words, there was no absorption taking place, no
indications of a terrestrially derived diagenetic signature (flattening of the shale-
normalized REE patterns; Grandjean et al., 1987), or no fractionation of REE (i.e. no
MREE enrichment or bell shape pattern; Reynard et al., 1999). The La/Yb ratio of about
7 and a larger negative Ceanom. points to deposition in an environment with a deepwater
influence (Grandjean, 1987; Grandjean et al., 1988; German and Elderfield, 1990;
Pipegras and Jacobsen, 1992). Results from DSDP Leg 41 off the Moroccan coast have
identified the onset of abundant chert deposition during the early Tertiary reflecting the
input of cold bottom water (Lancelot and Seibold, 1978). Cappetta (1981) showed that
Ypresian fish associations from the Ouled Abdoun basin (Oued Zem, location of O.
obliquus, is located in the eastern part of the Ouled Abdoun basin) were indicative of
greater depth than previous periods. All this evidence indicates either upwelling of deep
water onto the continental shelf or progressive deepening of the troughs and channels in
which the phosphorite was deposited (Grandjean, 1987) and suggests an influence of
deepwater during the diagenesis of 0. obliquus. The original REE signal of the 0.
obliquus centra have been replaced during diagenesis at/or near the sediment/seawater
interface with the seawater signal present at time of deposition of the centra. Therefore,
the two 0. obliquus centra preserved a seawater signal at or near the seawater/sediment
boundary and will be useful in reconstructing paleoceanographic environments.
The second diagenetic grouping, which includes C. angustidens and C. mantelli,
have shale-normalized REE patterns similar to the modem C. carcharias (UF211351)
and close to seawater but with some flattening (Fig. 3-6B). C. angustidens was deposited
in glauconitic shelf sand (Kokoamu Greensand) at a water depth of 50-100 m. The
presence of glauconite indicates slow sedimentation and low inputs of detritus from land.
New Zealand, as a whole, was low-lying and almost fully submerged at this time and
therefore the terrestrial input was minor (Ayrees, 1993). Minor flattening of the C.
angustidens shale-normalized REE pattern relative to seawater supports that the
diagenetic signature, which is produced by large pore water concentration gradients that
allow fluxes of REE from sediments to seawater, has had a small influence from
terrestrial derived sediments. C. mantelli was deposited in the Smoky Hill Chalk
Member of the Niobrara Formation in a water depth between 30-180 m (Hattin, 1981).
The sediments of the Smoky Hill Chalk Member are from the mid to outer shelf of the
Cretaceous epicontinental seaway. Once again, slight flattening of C. mantelli shale-
normalized REE pattern relative to seawater supports a small influence from a
terrestrially derived diagenetic signature. Neither C. angustidens nor C. mantelli have
any indications of REE fractionation ("bell-shaped" shale-normalized REE pattern). C.
angustidens and C. mantelli centra have the potential to be used for a general
paleoceanographic and paleoenvironmental reconstruction, but with caution, because they
do have a mixture of seawater signal and diagenetic signature, which will decrease the
Ceanom. as well as flatten the shale normalized REE pattern.
The third group, which includes C. auriculatus, has a shale-normalized REE
pattern between seawater and the modern C. carcharias (UF211351) (Fig. 3-6B). The C.
auriculatus centrum was deposited in the Brussels Sand in a shelf/nearshore environment
(Hooyberghs, 1990; Herman et al., 2000). The shale-normalized REE pattern, unlike the
second fossil group and C. carcharias (UF211351), has a high at Gd and is more depleted
in Dy and Yb. The flattening of the shale-normalized REE pattern, the (La/Yb)N ratio
greater than seawater, and the depletion of Dy and Yb in the C. auriculatus centrum
indicates a diagenetic signature that is more strongly derived from terrestrial sediments
than that of the second fossil group. The C. auriculatus centrum has the potential to be
used for a general paleoenvironmental reconstruction, but with caution, because of the
terrestrially influenced diagenetic signature, which will decrease the Ceanom. and flatten
the shale-normalized REE pattern.
The fourth group, which includes C. megalodon and I. hastalis have a flat shale-
normalized pattern or a maximum in the heavy-middle REE (Fig3- 6C). The C.
megalodon centrum was deposited in sandy siltstone near the Kanto Mountains in Japan
(Hayashi et al., 2003). The shale-normalized pattern of C. megalodon is almost
completely flat indicating a diagenetic signature that is extensively derived from
terrestrial sediments. The I. hastilas centrum was deposited in a shallow bay sandstone
and has a flat shale-normalized REE pattern, which indicates a diagenetic signature
derived primarily from terrestrial sediments and/or major influence from river water.
This is supported by fresh water diatoms in the sediments of the Pisco Formation
suggesting the influx of river water entering the basin (Brand et al., 2004). The high
(La/Yb)N ratios found in C. megalodon and I. hastalis can be explained by the extensive
terrigenous influence (Grandjean et al., 1987). C. megalodon and I. hastalis do not show
any indications of late diagenesis ("bell-shaped" shale-normalized REE pattern). These
two centra indicate the quantitative intake of REEs during early diagenesis with a strong
influence of continental sedimentary supply in a near-shore environment; therefore they
are not useful in paleoceanographic reconstruction (global controls) but the record of the
diagenetic signature indicates the sedimentary environment (local controls) can be
For the fossil shark centra used in this study the major control of the shale-
normalized REE pattern would be the diagenetic signature and how the terrestrially
derived sediments generate concentration gradients between pore waters and seawater.
The more prominent the terrestrial sediments the flatter and more shale-like the shale-
normalized REE signal become. Each depositional environment must be assessed in
order to determine which fossils can be used in paleoceanographic studies versus
depositional environment reconstructions.
While certain variables independently provide information about diagenesis, the
simultaneous use of FT-IR and elemental concentrations gives a much better picture of
depositional environment and the extent of diagenesis. There is no doubt that diagenetic
alteration has affected the REE composition of these seven fossil shark centra. Through
the processes of diagenesis, the centra have been imprinted with an REE seawater and
diagenetic REE pattern at the sediment/seawater interface. The REE seawater signature
was incorporated into the biogenic apatite via a transfer from a short-lived phase made of
oxy-hydroxides and organic detritus (Grandjean et al., 1987). The diagenetic signature is
caused by the development of pore water concentration gradients, which allow fluxes of
REE from sediments to seawater. The amount of terrigenous input and therefore the REE
composition of diagenetic signature controls whether the shale-normalized REE patterns
and Ce anomalies are representative of the original seawater signal for these seven fossil
shark centra. This is clearly seen for I. hastalis, and C. megalodon, where these two
specimens have a strong terrestrially influence diagenetic signature, which disturbs the
oceanic signal and Ceanom.. In contrast, the two 0. obliquus centra preserve a REE
seawater signal at the time of deposition and have no indication of a diagenetic signature
derived from terrestrial sediments. The remaining three centra (C. angustidens, C.
megalodon, and C. auriculatus) have diagenetic signatures that have some influence from
terrestrial sediments, which is evident by the slightly flattened shale-normalized REE
patterns. These later kinds of samples would require extreme care when interpreting Ce
anomalies because it is difficult to determine how much the Ceanom. has been reduce by
the diagenetic signature. Hence, even if there is no negative Ceanom. that does not
necessarily indicate an anoxic environment but may represent a strong continental
influence in the depositional environment. In summary, geochemical data from biogenic
apatite of fossil marine vertebrates, like lamnid sharks, have the potential to be used to
understand diagenesis, depositional environments (local controls), and/or
paleoceanography (global controls).
OXYGEN ISOTOPIC AND RARE EARTH ELEMENTAL ANALYSIS OF MODERN
LAMNID SHARKS: DETERMINATION OF LIFE HISTORY?
Lamnid sharks, Family Lamnidae (great white sharks and their relatives), are of
great interest not only to the scientific community but the public as well. Scientists have
spent a great deal of time trying to study and understand the life history of great whites
(Carcharodon carcharias) and their relatives. Because great whites do not survive well
in captivity, tagging and recapture studies and captured sharks from fishermen have been
the main source of study.
Sharks deposit light and dark bands on their vertebral centra throughout their lives
(Fig. 4-1). It is known that in most sharks the darker, denser portions are deposited
during slower growth times (e.g., winter) and lighter portions are deposited during more
rapid growth (e.g., summer). The problem is that the growth rate is affected by the
physical environment (including temperature and water depth), food availability, and
stress (Branstetter et al., 1987). Therefore, it cannot be assumed that a band pair (one
light and one dark band) reflects a single year (called annulus).
Wintner and Cliff (1999) estimated ages of great white sharks from the East coast
of South Africa by counting growth rings in the centra. The vertebrae of 61 females and
53 males were x-rayed and counts were made from the x-rays. X-rays enhance the
visibility of growth rings in shark centra and have been used successfully to accurately
determine ontogenic age of several species (Cailliet et al., 1983; Yudin and Cailliet,
1990; Ferreira and Vooren, 1991). Of particular interest from Wintner and Cliffs (1999)
study was the one shark that was injected with oxytetracycline (OTC) on October 10,
1994, and was recaptured on May 27, 1997 (specimen BTO433). BTO433 was tagged at
140 cm and grew 69 cm within that two year, seven month and 27 day period. The OTC
indicated annual growth ring deposition in most of the centra from BTO433; however,
this could not be confirmed from growth ring counts of the entire sample (Wintner and
Dark growth band
Figure 4-1. Scanned contact print of BTO433 centra. (Top) Dark growth bands, OTC
mark and birth mark have all been indicated. Dark growth rings on the
contact print show up as white bands, which can be seen in the lower image of
Bomb Carbon and Oxygen Isotopes
Atmospheric testing of atomic bombs in the 1950s and 1960s resulted in a rapid
increase in radiocarbon (14C) in the world's oceans (Druffel and Linick, 1978). The
period of radiocarbon increase was almost synchronous in marine carbonates such as
corals, bivalves, and fish otoliths around the world (Kalish, 1993; Campana, 1997; Baker
and Wilson, 2001), providing a date marker in calcified structures exhibiting incremental
growth. More recently, analysis of bomb radiocarbon has been used to validate age
estimates derived from vertebral centra of sharks (because sharks do not contain otoliths
that grow incrementally; Campana et al., 2002). Campana et al. (2002) used bomb
carbon dating as an age validation method for long-lived sharks. They compared
radiocarbon assays in young, known-age porbeagles (Lamna) collected in the 1960s with
corresponding growth bands in old porbeagles collected later. With this method Campana
et al. (2002) confirmed the validity of porbeagle vertebral growth band counts as accurate
Because shark vertebrae grow incrementally, the oxygen isotopic signal preserved
should reflect the seawater conditions at the time of formation. Oxygen isotopes vary
with temperature and salinity (Hoefs, 1988), consequently variations preserved in the
vertebral centra could indicate migration into various water bodies and/or water depths.
Also, it might be possible to determine the frequency of the migration and, if annual, the
oxygen isotopic signals could be used to estimate ontogenic age of the individual.
Rare Earth Elements
The rare earth elements (REE) consist of fifteen elements which form a series from
the lightest REE, lanthanum (La), to the heaviest, lutecium (Lu). With the exception of
multiple oxidation states for Ce and Eu, the other REEs have trivalent oxidation state in
most natural waters. Cerium may undergo oxidation in seawater from the solvated Ce3+
state to insoluble Ce4 consequently Ce fractionates relative to other REE (German and
Elderfield, 1990). Europium may undergo reduction from the Eu3+ to Eu2 which
substitutes readily for Ca2+ inCa- bearing minerals such as apatites (Elderfield, 1988).
The residence time of REE's in seawater is 102-103 years and is therefore shorter than the
mixing time of the oceans (1600 years) making these elements useful tracers of
oceanographic events and processes (Elderfield and Greaves, 1982; Bertram and
Elderfield, 1993; Nozaki et al., 1999; Lacan and Jeandel, 2001). Because most REE
concentrations increase with water depth (Elderfield and Greaves, 1982; deBaar et al.,
1985a; deBaar et al, 1985b; Dubinin, 2004) they may act as a proxy to indicate the
relative water depth at which these individual sharks are living
This study tests whether the chemistry of vertebral centra can be used to improve
our understanding of the life history of sharks. The possibility exists to elucidate shark
migration, the relative depth of habitat, and determination of individual age using rare
earth elemental compositions, bomb carbon, and oxygen isotopes.
Five species of lamnid sharks live today, Carcharodon carcharias (great white),
Isurus paucus (longfin mako), Isurus oxyrinchus (shortfin mako), Lamna ditropis
(salmon shark), and Lamna nasus (porbeagle). This study focuses only on the largest
three members of Lamnidae, the great white, longfin mako, and shortfin mako (Table 4-
1), due to ease in sampling larger vertebral centra.
Great White (Carcharodon carcharias)
The great white is the largest extant lamnid shark and has one of the broadest
distributions of all modern sharks. The great white is cosmopolitan in cold temperate to
tropical seas, living primarily in coastal and offshore habitats of continental and insular
shelves. However large individuals have been recorded off oceanic islands. The known
depth range of great whites is from the surface to at least 6,150 ft (1.875 km). Great
whites maintain a body temperature up to 270F (150C) warmer than surrounding waters
by an adaptation to its circulatory system that does not allow the heat generated within
the body to escape through the gills (Compango, 2002). Great whites prefer waters with
a sea surface temperature between 59-720F (15-220C). Most individuals are 12-16 ft
(3.7-4.9 m) long with a maximum length of 20 ft (6.1 m). At birth a great white is
between 3'11" to 4'3" in length. Juvenile sharks feed on bottom-dwelling teleost fish,
small sharks, and rays, while adult sharks feed on sharks, rays, teleost fish, seals, sea
lions, dolphins, whale blubber (scavenged), squid, seabirds, marine turtles, crabs, and
snails (Campagno, 2002).
Longfin Mako (Isurus paucus)
The longfin mako was first described in 1966 and is one of the least-known
lamnids. It is also the second largest member of Lamnidae, after the great white.
Longfin makos have an appearance similar to shortfin makos but have a slimmer body,
larger eyes, and larger pectoral fins. Most specimens are about 7 ft (2.2 m) long and the
maximum known length is 14 ft (4.3 m), which is based on a male specimen taken from
15 mi (24 km) off Pompano Beach, FL, in February 1984 (Compango, 2002). At birth a
longfin mako is between 3' to 3' 11" (92 to 120 cm) in length. Longfin makos are widely
distributed in tropical to warm temperate seas. They are fairly common in the western
Atlantic (Gulf Stream waters, northern Cuba to southeast Florida) and possibly in the
central Pacific (near Phoenix Island and north of Hawaii), although rare, longfin makos
have been recorded off northwestern Africa and the Iberian Peninsula, from the northern
Gulf of Mexico to the Grand Banks, Bahamas and off New South Wales Australia
(Compango, 2002). Most specimens of longfin makos are caught on long-lines in deep
tropical waters, from depths of 360-720 ft (110-220 m). The long broad pectoral fins
suggest that longfin makos are slower and less active than shortfin makos. This inference
is also supported by the fact that longfin makos have the same heat-retaining
modifications to the circulatory system as other lamnids; however longfin makos are
unique among the members of its family in that this species is not warm bodied. The diet
of a longfin mako consists of schooling fish and pelagic squid (Compagno, 2002).
Table 4-1. Lamnid specimens used in this study.
Taxon Specimen Location of Capture of Length
Cn BTO433 Capetown, S. Africa 1 6'11"
Cr UF211351 Islamorada, Florida 1 12'9"
Cn UF211352 Marathon, Florida 1 12'5 /4"
Cr UF31648 Florida 1 N/A
Isuruspaucus UF211355 Miami, Florida 3 8
Pompano Beach, 8
Isuruspaucus UF211354 o o 3 84"
Isuruspaucus UF211353 o ao a 1 14
Isus UF47943 Florida 1 N/A
Shortfin Mako (Isurus oxyrinchus)
The shortfin mako is the fastest swimming shark and has a global distribution in
tropical and temperate waters. Shortfin makos are common in coastal and oceanic
regions of tropical and temperate seas but seldom occur in waters less than 610F (160C).
They range from California to Chile in the Pacific Ocean and from the Grand Banks of
the Bahamas to Brazil, including the Gulf of Mexico and the Caribbean Sea in the
Atlantic Ocean. In the eastern Atlantic, shortfin makos range from Norway to South
Africa, including the Mediterranean, and is found throughout the Indian Ocean from
South Africa to Australia. In the western Pacific it can be found from Japan to New
Zealand and in the central Pacific it occurs from the Aleutian Islands to the Society
Islands (Compango, 2002). The known depth range is from the surface down to at least
1,300 ft (400 m). Shortfin makos tend to follow movements of warm water in extreme
northern and southern parts of its range. Tagging studies off the northeastern US show a
seasonal pattern of abundance along the western margin of the Gulf Stream, moving
inshore and into higher latitude waters as the stream shifts northward from April-October,
possibly wintering in the Sargasso sea from November to March (Compango, 2002). At
birth a shortfin mako is 2 to 2'3" long with most individuals 6-8 ft (1.8-2.5 m) long and a
maximum recorded length of 12.8 ft (3.9 m). Shortfin makos maintain body temperatures
12.5-180F (7-100C) warmer than the ambient water and are capable of rapid acceleration
and bursts of speed when hooked or in pursuit of prey. Adults have been clocked at 31
mph (50 km/hr) (Compango, 2002).
Materials and Methods
X-radiographs were used to enhance the visibility of the growth rings. X-
radiographs of whole centra were taken at the C.A. Pound Human Identification
Laboratory at the University of Florida. The x-rays are set at 78 kV for 2 minutes. The
x-rays are used to make contact prints, which are a reversed pattern of the x-ray (i.e., dark
lines on the x-ray are the light lines on the contact print). The dark and light alternating
growth rings are easily seen on the contact prints (Fig. 4-1). These contact prints were
digitally scanned and Adobe PhotoshopT was used to enhance the images. Growth ring
counts were made from the scanned contact prints and interpretation of the vertebral
growth bands was made using published criteria for porbeagles and Pacific shortfin
makos. Campana et al. (2002) has proven with the use of bomb carbon, annual growth
band pairs form in porbeagles, and Cailliet at al. (1983) demonstrated that a single band
pair formed each year in Pacific shortfin mako based on growth ring counts.
Oxygen Isotopic Preparation and Analyses
Oxygen isotopes can be used for the interpretation of temperature. For each of the
centra, about 1 mg was drilled with a foredoom slow speed drill approximately every 1
mm across the growth axis starting from the birthmark and ending at the external margin.
An interval of 1 mm insured that each primary growth band was sampled. Sample
powders were treated with standard isotope preparation techniques (e.g., MacFadden et
al., 1999). At least 0.75 mg of each treated sample powder was then measured into
individual metal boats and placed in the carousel of the isocarb device for introduction
into the VG Prism mass spectrometer in the Stable Isotope Laboratory in the UF
Department of Geological Sciences. The sample runs were calibrated to internal
laboratory and NBS 19 standards. The oxygen isotopic results are reported in the
standard "6" convention: 6 (parts per mil, %o) = (Rsampe/Rstandard)-I) x 1,000), where R =
13/13C or 180/160, and the standard is VPDB. Ontogenic ages based on the oxygen
isotopic data were estimated by counting the number of peaks or valleys present long the
growth axis of the centra.
Bomb Carbon Dating Preparation and Analysis
Atmospheric testing of atomic bombs in the 1950's and 1960's produced a time-
specific radiocarbon marker, which allows for material formed between the 1950's and
the present to be dated (Kalish, 1993; Weidman and Jones, 1993). To determine if bomb
carbon dating would work on great white centra, only one shark centra, BTO433, was
chosen for dating. BTO433 was chosen because the capture date is known and this shark
was injected with oxytetracycline (OTC) on October 30th 1994. The OTC mark gives a
second reference for the accuracy of the bomb carbon dates generated.
About 60 mg of sample was extracted from BTO433 first formed growth band
(corresponding to the first year of growth) and the last growth band. The external surface
of the centrum was removed in order to minimize surface contamination. The sample was
weighed to the nearest 0.01 mg in preparation for assay with Accelerator Mass
Spectrometry (AMS). The sample was assayed at the Keck Carbon Cycle AMS
Laboratory, UC Irvine for 613C (to determine the carbon source) and A14C (measure of
radiocarbon), with A14C calculated per Stuiver and Polach (1977).
To assign dates of formation to an unknown sample, it is necessary that the A14C of
the unknown sample be compared with a A14C of a known-aged material. BTO433 A14C
data were compared to the A14C values of Pagrus auratus otolith collected off the East
coast of North Island, New Zealand (Kalish, 1993). An age for BTO433 based off the
Campana et al. (2002) reference curve was assigned by correcting the 1997 A14C value of
BTO433 to the1997 porbeagle reference curve value generated by Campana et al. (2002),
then the same correction was applied to the first growth band A14C value and that value
was compared to the porbeagle curve to determine a year for deposition of the first
growth band (Fig. 4-2). This was done to determine if the second method would produce
similar results to the first, because reference curves are not always available for the area
in which the specimen was recovered and this would allow for the use of previously
produced reference curves. It is important to have a local reference curve since the
carbon was not uniformly distributed at the time of atomic bomb testing, especially
between the Northern and Southern Hemisphere as most of the bomb testing was done in
the Northern Hemisphere.
1952 1956 1960 1964 1968 1972 1976 1980 1984 1988 1992 1996 2000 2004
Figure 4-2. BT0433 bomb carbon data plotted vs. two reference curves (a) fish otolith
from New Zealand (Kalish, 1983) and (b) porbeagle (Campana, 2002).
BT0433 is plotted on the fish otolith curve and BT0433B is plotted on the
Inductively Coupled Plasma Mass Spectroscopy (ICPMS)
Inductively coupled plasma mass spectroscopy (ICPMS) allows for the
quantification of elemental abundances within the shark centra. Approximately 5-10 mg
of bulk sample was drilled from each of the twelve centra. 5 mg of each sample were
weighed into 3 mL Savillex vials and dissolved in 1 mL of 3M HNO3 and heated
overnight. Samples were allowed to cool and then dried. Then 2 mL of 1% HNO3 was
added, heated overnight, and allowed to cool. Samples were analyzed on an Element 2
High Resolution Inductively Coupled Plasma Mass Spectrometer (HR-ICP-MS) at the
UF Department of Geological Sciences. All samples were corrected by subtracting the
blank, corrected for instrumental drift based on internal machine standards that were
analyzed during the run, and correcting ion counts to a constant response to the known
All REE values were shale-normalized to PAAS (Post-Archean Shale Standard) in
order to eliminate the odd-even effect of the natural abundances of the REE. The Ce
anomaly (Ceanom) was calculated with the Elderfield and Greaves (1982) formula,
Results and Discussion
Ontogenic Age Determinations
The ontogenic ages determined from oxygen isotopic cyclicity, growth ring counts
and the A14C of the eight sharks are presented in Table 4-2. The ontogenic ages
estimated based on the oxygen isotopic cyclicity (Fig. 4-3) give ages that are too young
with the exception of UF31648 and BTO433 (Table 4-2) when compared to growth ring
counts. For UF211351, UF211352, and UF47943 the oxygen isotopic signal gave an
ontogenic ages of 5+ years, 7+ years, and 4+ years, respectively, which are all several
years too young compared to growth ring counts. For specimens UF211354 and
UF211353 the oxygen isotopic data gave ages of 6+ and 8+, respectively, which are
about half of the ages estimated from the growth ring counts. The three UF211355 have
age estimates of 3+, which is 8+ years too young according to the growth ring counts.
One problem with the oxygen isotopic data is that sampling every growth ring is
challenging. As sharks get older, their growth rate decreases and therefore the size of the
growth rings become narrower making it difficult to sample only one growth ring at a
time. Also, the location of the centra in the sharks vertebral column (i.e., towards the
head or towards the tail) will affect the age estimate for both oxygen isotopic analysis and
growth ring counts. Sampling the exact same location for multiple centra within an
individual is difficult to accomplish because not all the vertebrae grow at the exact same
time. This accounts for the offset of the isotopic data in sharks UF211354 and UF211355
Table 4-2. Ontogenic age estimates based growth ring counts (GR), oxygen isotopic
(6180) cyclicity, and bomb carbon (A13C). Ages estimated from growth ring
counts should be considered minimum ages since the last growth band may
not have been formed completely at time of capture.
Taxon Specimen GR counts 610 age A14C age
Carcharodon carcharias BTO433 3+ 3+ 3
Carcharodon carcharias UF211352 9+ 7+
Carcharodon carcharias UF211351 7+ 5+
Carcharodon carcharias UF47943 4+ 3+
Isurus paucus UF211355A 11+ 3+
Isurus paucus UF211355B 11+ 3+
Isurus paucus UF211355C 11+ 4+
Isurus paucus UF211354C 11+ 4+
Isurus paucus UF211354B 13+ 5+
Isurus paucus UF211354A 12+ 5+
Isurus paucus UF211353 16+ 8+
Isurus oxyrhincus UF31648 3+ 3+
of the three centra analyzed (Fig. 4-3). Choosing the largest centra and/or centra that are
closer to the head than the tail would probably alleviate some of these effects. Another
possibility is that the sharks studied did not migrate into bodies of water with enough
temperature variation to alter the oxygen isotopic signal. Also, great whites and shortfin
makos are warm-bodied and maintain a body temperature above the surrounding water
temperature, which will affect the oxygen isotopic signal, (i.e., the isotopic variation
might not reflect the actual water temperature change, but the changes in body
temperature as the sharks encounter colder or warmer waters). It has been shown that
muscle temperature of lamnid sharks changes in response to changes in ambient water
temperature (Carey et al., 1982; Tricas and McCosker, 1984; and Carey et al. 1985).
Oxygen isotopic signals may reflect changes in body temperature; these changes may
accentuate or dampen the oxygen isotopic signal preserved in the centra. The shark
centra analyzed in this study do not seem to have a dampened oxygen isotopic signal (i.e.,
oxygen isotopic values along the growth axis have similar absolute values) indicating that
great whites and shortfin makos do not maintain a constant body temperature (Fig. 4-3).
Unlike the shortfin mako and great white, the longfin mako is not warm-bodied and
therefore does not maintain a body temperature above that of its surroundings and
consequently the oxygen isotopic signal should reflect the temperature of the surrounding
BT0433 age is estimated by the A14C values, which are plotted against the two
reference curves in Fig. 4-2. BT0433 A14C values were plotted against the first reference
curve, otoliths from fish off the coast of New Zealand (Kalish, 1993), and no correction
was necessary because the 1997 age plotted on the curve. The second method entailed
subtracting the 1997 A14C reference curve value generated from the porbeagle centra
(Campana, 2002) from the BT0433 1997 A14C value and then subtracting that same
amount from the BT0433 first growth ring A14C value. Both methods produce the same
ontogenic age of 3 years. More work needs to be done in order to determine if
subtracting the difference of the A14C values of the known age of the specimen to the
reference curve works for sharks that are older. However, the fact that both
methods(growth ring counting and bomb carbon dating) produced an ontogenic age of 3
years for BT0433 indicates that at least at young ages great whites deposit growth bands
on an annual basis. Also, if two growth band pairs form per year, this shark would only
be 11/2 years old and that is impossible since 2.6 years passed before recapture of BT0433
after the OTC injection (Wintner and Cliff, 1999). However one individual cannot be
used to provide a definitive growth rate for great whites, further radiocarbon assays of
lamnid shark vertebrae is needed to accomplish this goal.
8 9 10 11 12 13 14 15 16 17 18 19 20 21 22
-6.5 Great White
11 12 13 14 15 16 17 18 19 20 21 22 23 24 25 26 27 28
7 8 9 10 11 12 13 14 15 16 17 18 19
1 Great White I
9 10 11 12 13 14 15 16 17 18 19 20
5 6 7 8 9 10 11 12 13 14 15
5 6 7 8 9 10 11 12 13 14
.5 Longfin Mako
5 6 7 8 9 1011 1213141516171819202122232425
Figure 4-3. Oxygen isotopic data (VPDB) for the shark centra analyzed. UF211354 and
UF31155 both have three centra per specimen.
Rare Earth Elements
The elemental concentrations are given in Table 4-3 and the shale-normalized REE
patterns are shown in Fig. 4-4. The shale-normalized REE concentrations are enriched
relative to seawater by about 104 -105. For all the sharks captured off the coast of
Florida, Eu is enriched relative to seawater, Eu was not analyzed for the specimen caught
0 11 12 13 14 15 16 17 18 19 20 21 22 23 24 25 26 27 28
off the coast of South Africa. Eu is known to replace Ca in apatite more readily than
other REE since Eu2+ has the same oxidation state and a similar ionic radius to Ca2+
(Elderfield, 1988), therefore expectations are for marine biogenic apatite to be enriched in
Eu relative to seawater. Specimens UF211353, UF31648, UF47943, and UF211352 all
have depletion in Gd relative to seawater. The degree of enrichment of Eu and depletion
of Gd varies between specimens. BT0433 and UF211352 have shale normalized REE
patterns similar to average seawater, ignoring the Eu enrichment (Fig. 4-4). UF211351,
UF211354, and UF211355 have shale-normalized REE patterns between seawater and
coastal waters (Fig. 4-4), ignoring the Eu enrichment. UF211353, UF31648, and
UF47943 have shale normalized REE patterns similar to coastal waters (Fig. 4-4),
ignoring the Eu enrichment.
The REE patterns differ with location along the Florida coast, indicating that REE
shale-normalized patterns may serve as a provence tool (i.e., location determination). In
this study the two centra, UF31648 and UF47943, belong to sharks without capture
location data but exhibit REE distributions similar to that of UF211353. To use bulk
samples as a provenace tool numerous sharks from an area should be analyzed, which is
evident from the differences between UF211354 and UF211353. Both of these sharks
were caught off the coast of Pompano Beach, Florida but they show different shale-
normalized REE patterns, because they probably spent time in different waters/depths
throughout their lives. The REE pattern throughout life will be averaged when bulk
samples are used; therefore changes in the REE shale-normalized patterns along the
growth axis of the centra (i.e., migration patterns) should be shown when microsampling
The REE content variation in seawater correlates to the circumcontinental zonality
(Dubinin and Rozanov, 2001; Tachikawa et al., 1999) due to an exchange between the
dissolved REE and absorbed complex of terrigenous suspended matter. The decrease in
terrigenous particulate content toward the pelagic area leads to the increase of the
dissolved REE concentration (Dubinin, 2004). Dissolved REE content increases with
water depth by several times in the Atlantic Ocean (de Barr et al., 1985; Sholkovitz,
1994; Dubinin, 2004), therefore REE concentrations could be a proxy for relative water
depth for these shark centra. The plot of La+Sm+Yb (ppm) versus (Sm/Yb)N (Fig. 4-5)
indicates that (a) UF211351, UF211353,UF211355, UF47943, and UF31648 lived at
similar depth closest to the surface; and (b) BTO433, UF211352, and UF211354 also
lived at similar depth, but deeper than the previously mentioned group. UF211353,
UF211354, and UF211355 should plot at similar depths (i.e., have similar REE
concentrations) since it is known that most longfin makos live and feed between 360-720
ft and expectations are that the REE values would not vary much within this depth range
since it is only a difference of 360 ft. Shortfin makos have a larger depth range in which
they feed and inhabit than longfin makos, however according to these data UF47943
lived or ate at a fairly shallow depth relative to the other sharks analyzed.
This is the first study to analyze the chemical and light stable isotopic composition of
lamnid shark centra. Because lamnid sharks cannot be kept in captivity, REE
concentrations, oxygen isotopes, and bomb carbon are new approaches to gaining
knowledge about these animals. Bomb carbon dating is an effective way to accurately
date shark centra, and this has been tested with an individual with a known capture date.
Table 4-3. Elemental data (in ppm), and oxygen isotopic and bomb carbon dating ages. N/A represents elements that were not
analyzed and (-) indicates that the concentrations were below detection limits of the ICPMS.
Taxon Specimen Na Mg Al Mn Fe Ni Cu Zn Sr Y Ba Pb U P
C. carcharas BT0433 4053.43 1752.92 1 4.80 177.22 N/A N/A 904.81 352.49 0.56 1.47 9.11 0.28 59982.29
C. carcharas UF211352 8391.58 3196.68 82.32 6.95 35.19 1.54 2.13 255.64 1428.63 0.05 3.20 1.04 2.42 144448.89
C. carcharas UF211351 6559.68 2678.65 30.78 6.43 408.77 7.91 6.65 29.51 1120.88 0.03 1.76 0.54 0.14 116058.87
C. carcharas UF47943 5012.73 2822.44 75.03 4.82 19.52 5.72 8.62 36.04 1181.54 0.05 3.47 2.47 0.11 119935.33
I. paucus UF211355A 6012.29 3053.64 88.06 9.80 42.78 0.66 1.15 52.40 1541.16 0.06 6.62 0.19 0.10 131921.99
I. paucus UF211355B 5001.04 1868.17 52.46 10.60 16.28 0.75 0.93 31.45 1545.61 0.04 8.27 0.15 0.11 135493.72
I. paucus UF211355C 5992.55 3050.59 83.15 9.25 14.71 0.86 1.81 46.09 1478.26 0.04 6.67 0.27 0.09 130841.17
I. paucus UF211354C 4915.74 2651.50 46.06 8.94 79.74 0.84 1.09 35.57 1417.09 0.08 6.78 0.64 0.21 131198.65
I. paucus UF211354B 4548.74 2327.36 21.16 7.01 41.85 0.38 0.60 28.17 1226.32 0.04 6.05 0.32 0.18 117790.59
I. paucus UF211354A 4782.93 2627.17 49.99 7.77 54.53 0.76 0.95 32.09 1337.93 0.08 7.43 0.67 0.21 125275.09
I. paucus UF211353 7522.55 2775.05 11.37 10.52 16.66 0.45 1.05 54.17 1596.44 0.01 4.52 24.26 0.17 137821.84
I. oxyrhincus UF31648 5869.75 3483.01 20.78 7.62 14.73 1.61 5.87 54.54 1563.24 0.09 5.87 0.57 0.21 142857.41
Taxon Specimen K Ca La Ce Pr Nd Sm Eu Gd Dy Yb Lu anom.
C. carcharas BT0433 1133.63 73202.7 0.23 0.09 N/A 0.14 0.02 N/A 0.04 0.04 0.03 N/A -0.68
C. carcharas UF211352 322.38 302628.1 0.168 0.028 0.033 0.135 0.025 0.014 0.007 0.024 0.016 0.003 -1.08
C. carcharas UF211351 261.62 246712.3 0.044 0.068 0.01 0.035 0.009 0.006 0.005 0.008 0.005 0.002 -0.11
C. carcharas UF47943 185.32 250650.4 0.055 0.138 0.011 0.044 0.011 0.013 0.008 0.005 0.001 0.1
I. paucus UF211355A 187.57 280707.6 0.081 0.097 0.018 0.078 0.014 0.022 0.029 0.014 0.008 0.003 -0.25
I. paucus UF211355B 187.16 8 0.086 0.132 0.021 0.087 0.018 0.023 0.012 0.016 0.009 0.002 -0.15
I. paucus UF211355C 217.55 277130.3 0.185 0.17 0.041 0.16 0.029 0.026 0.018 0.027 0.009 0.002 -0.35
I. paucus UF211354C 131.12 280654.2 0.476 1.45 0.13 0.62 0.145 0.047 0.156 0.183 0.103 0.019 0.1
I. paucus UF211354B 105.88 252319.5 0.2223 0.444 0.057 0.245 0.051 0.024 0.036 0.055 0.034 0.007 -0.05
I. paucus UF211354A 109.88 270683.3 0.271 0.773 0.073 0.336 0.07 0.033 0.1 0.089 0.058 0.009 0.08
I. paucus UF211353 639.50 291824.4 0.05 0.071 0.012 0.047 0.01 0.017 0.01 0.006 0.001 -0.18
I. oxyrhincus UF31648 192.68 30000.6 0.055 0.131 0.013 0.051 0.013 0.018 0.007 0.006 0.001 0.05
* BT0433 Caught off coast of S. Africa
-F UF211352 Caught off coast of Marathon, FL
- UF211351 Caught off coast of Islamorada, FL .-...""
X" ...... .....
---- Seawater (x 06)
---- Connecticut Coastal Waters (x104)
La Ce Nd Sm Eu Gd Dy Yb Lu
Caught off coast of Miami, FL
.. ... .... -.... .....
-+-UF211355B --a-Seawater (x10s)
--UF211355C *x-- Connecticut Coastal Waters (x10)
La Ce Nd Sm Eu Gd Dy Yb Lu
La Ce Nd Sm Eu Gd Dy Yb Lu
Figure 4-4. Post Archean Australian Shale normalized rare earth element plots for the
eight sharks analyzed compared with average seawater (Elderfield and
Sholkovitz, 1987) and Connecticut coastal waters (Elderfield et al., 1990).
Caught off coast of J --UF211354B
Pompano Beach, FL -o.UF21135A
Unkown capture location --_UF47943
A m BT0433
A E UF211351
E A 0UF47943
C A AUF211355
+E0.1 A *UF31648
S Increasing water depth
0.01 .. ...
0.1 1 10
Figure 4-5. Depth estimates for the eight lamnid sharks. Arrow indicates direction of
increasing water depth.
Both methods of bomb carbon age estimation (correcting to a known curve and plotting
against a known curve) seem to work for at least the one great white analyzed here.
BTO433 has annual growth ring deposition giving it an age of 3+ years, which coincides
with the growth ring counts from the contact prints of the centra.
Oxygen isotopes give an indication of changes in body temperature for shortfin
makos and great white sharks, and water temperature for longfin makos throughout their
lives. Ontogenic age estimated based on the oxygen isotopic data did not give similar
ages for most sharks when compared to the growth ring counts. Either the eight sharks
studied were not encountering waters that shifted their body temperature enough to
change the oxygen isotopic signal seasonally, or during sampling some growth bands
were missed and/or more than one band was sampled averaging the oxygen isotopic data.
Annual growth deposition in lamnid sharks is supported by bomb carbon dating from
Camapana (2002) and from the great white BTO433 from this study. Also, UF211353 a
14' longfin mako, which is twice the usual size, would only be 8 years old if two growth
band pairs were deposited annually, indicating that the longfin mako, which is the slow
moving non-warm bodied lamnid, would grow at a faster rate than the warm-bodied fast
moving shortfin mako.
REE shale normalized patterns have the potential to be utilized for assessing the
general location of sharks that have unknown capture locations. Elemental analysis was
done on bulk samples, which would average the REE values along the growth axis.
Sharks with similar REE patterns most likely lived in similar waters. UF31648 and
UF47943 were captured off the coast of Florida and no other location data were given.
When compared to the other centra analyzed from the coast of Florida UF31648 and
UF47943 REE shale-normalized patterns are almost identical to UF211353, indicating
that these sharks may have spent a significant amount of time in similar waters. Also
with microsampling the changes of the shale-normalized REE patterns along the growth
axis of the centra will allow for the determination of migration timing. The use of
La+Sm+Yb vs. (La/Sm)N demonstrates the possibility of estimating relative water depth
for sharks. According to this interpretation (Fig. 4-5), the great whites in this study have
a large depth distribution and longfin makos in this study have a fairly restricted depth
distribution, which is supported by that fact that great whites have been seen from the
surface down to 6,150 ft and longfin makos are usually captured depths between 360-720
ft. More shark centra must be analyzed in order to determine whether REE in shark
centra accurately represent water depth. This study demonstrates the potential of using
elemental and isotopic analysis to learn more about the life histories of sharks from the
chemical make-up of their vertebral centra.
SUMMARY AND CONCLUSIONS
Diagenesis is pervasive in fossil bones and ancient sedimentary environments even
though in paleontology this process is all too often ignored. Understanding how
diagenesis affects the biological signal preserved in fossils is imperative to correctly
interpret analytical findings. The study presented here gives several procedures
(including FT-IR, ICMPS, and stable isotope analysis) to help unravel the diagenetic
story. A modem analog must always be understood before the fossil data can be
interpreted. Therefore, if the end memebr composition of unaltered specimens is not
known there is no way to interpret the signal preserved in the fossil (all fossil undergo
diagenesis just to varying degrees).
Otodus obliquus centra from the Eocene of Morocco demonstrate that a biological
oxygen isotopic signal remains preserved along the growth axis even though diagenesis
has taken place. The oxygen isotopic signal does not seem to be annual in this case,
because if annual growth ring deposition is assumed as in other lamnids, then the oxygen
isotopic age is eight years too young. Given that the oxygen isotopes vary with
temperature, this distribution relative to the growth rings may indicate that these sharks
were not migrating annually into waters with temperature shifts detectable by the oxygen
isotopes. Another possibility is that Otodus obliquus, like modern lamnids, may be
warm-bodied. If this were the case, then the oxygen isotopic data represent body
temperature and not water temperature and 0. obliquus may be migrating on an annual
basis but it may not be encountering waters that raise or lower its body temperature. The
oxygen isotopic analysis of Otodus obliquus centra have prospective broad ramifications
for understanding the evolution of growth rates and developmental strategies in fossil
sharks; however, analytical techniques that assess diagenesis should be used in
combination with isotopic studies in order to produce the most insightful analysis of
fossil shark paleobiology.
Diagenesis can be quantified by the use of multiple variable analyses. The use of
ICPMS and FT-IR data provide a clearer picture of the depositional environment and the
extent of diagenesis. Seven fossil lamnid shark centra from all over the world were
analyzed for elemental and mineralogical composition. All seven fossil centra are
diagenetically altered, which is evident from the FT-IR spectra indicating the presence of
fluorine and the decrease in carbonate content, and ICPMS data which show an
enrichment in REE, Y, and U. However, through diagenesis the centra have been
imprinted with the seawater signal at/near the sediment/water interface. The type of
diagenetic fluids expelled from sediments into the water column determines how
representative the shale-normalized REE patterns and Ce anomalies are to the original
seawater signal. The more terrigenous sediments present in the depositional environment
the flatter and more shale-like the normalized REE patterns of the centra. Therefore
samples that contain a terrestrial influence would require extreme care when interpreting
Ce anomalies because it is difficult to determine how much the Ce anomaly has been
reduced by continental input. In these cases, the chemistry of depositional environment
is determined rather than the chemistry of the seawater. These seven shark centra
illustrate that geochemical data from biogenic apatites of fossil marine vertebrates have
the potential to be used to understand diagenesis, depositional environments, and/or
Modem lamnid sharks, which include great whites, longfin makos, and shortfin
makos, have vertebral centra composed of carbonate hydroxyapatite similar to other
vertebrates. The oxygen isotopic signals preserved along the growth axis of great whites
and shortfin makos represent changes in body temperature as they encounter varying
water temperatures, while in the longfin mako the actual water temperature is
represented. This is because great whites and shortfin makos are warm-bodied and
longfin makos are cold-bodied (Campana, 2002). Unfortunately, the oxygen isotopic
signal recorded in the eight modem sharks studied did not demonstrate annual cyclicity.
Growth ring counts made on all specimens did indicate annual growth, but the oxygen
isotopic data usually gave an age of about half of the actual age. Annual growth ring
deposition in lamnid sharks was supported by bomb carbon dating of lamnid sharks
(BT0433) and a comparison of growth rates for three different lamnid species. Rare
earth elemental data illustrate that shale-normalized rare earth elements can be used as to
determine habitat for sharks with an unknown capture site off the coast of Florida and can
be used for relative depth estimation of the eight sharks in this study.
These studies will potentially serve as a general model for other researchers
interested in assessing the extent of diagenesis of their fossils in a particular study area.
These studies therefore have broad applicability to paleobiologists, paleoclimatologists,
paleoceanographers, and archaeologists.
Plans to continue this study include geochemical modeling of the diagenetic system
using thermodynamics and kinetics to better understand the stability of biogenic apatites
and chemical pathways for diagenesis. Also a comparison of the accuracy of laser
ablation techniques with conventional dissolution and dilution methods currently used to
quantify the elemental concentrations on the ICPMS. The precision of laser ablation will
allow for fine scale resolution along the growth axis of the shark's centra and will
minimize destruction to the specimen. Future work will include applying the techniques
presented in this dissertation to the terrestrial system in order to determine how
diagenesis affects the orthodentine of Edentates (sloths and armadillos). Edentates,
unlike other mammals, do not have enamel on their teeth and are therefore more prone to
diagenesis. Finally, future work will be to expand the techniques used here to other
groups of fishes (including modem and fossil) and possibly mososaurs and plesiosaurs
(which preserve incremental growth in their vertebrae).
LIST OF REFERENCES
Applegate, S.P. 1967. A survey of shark hard parts In .\,l/iuA ,\skAte/, and Rays (ed P.
Gilbert et al.). Baltimore: Johns Hopkins Univ. Press. pp. 37-67.
Arambourg, C. 1952. Les vertebres fossiles des gisements de phosphates (Maroc-Algerie-
Tunisie). Service Geologie Maroc, Notes et Memoires 92,1-372.
Ayress, M.A. 1993. Ostracod biostratigraphy and palaeoecology of the Kokoamu
Greensand and Otekaike Limestone (late Oligocene to early Miocene), North Otago
and South Canterbury, New Zealand. International Symposia on Ostracoda 11,
Baker, M.S., and Wilson, C.A. 2001. Use of bomb radiocarbon to validate otoliths
section ages of red snapper Lutjanus campechanus from the Northern Gulf of
Mexico. Limnol. Oceanogr. 46, 1819-1824.
Barrick, R.E., and Showers W. J. 1994. Thermophysiology of Tyrannosaurus rex:
Evidence from oxygen isotopes. Science 265, 222-224.
Barrick, R.E., 1998. Isotope paleobiology of vertebrates: ecology, physiology, and
diagenesis. In: Isotope Paleobiology and Paleoecology (eds R.D. Norris and R.M.
Corfield). The Paleontological Society Papers 4,101-137.
Bernat, R.T. 1975. Les isotopes de l'uranium et du thorium et les terres rares dans
l'environnement marin. Cah. ORSTORM Ser. Geol. 7, 68-83.
Bertram, C.J. and Elderfield, H. 1993. The geochemical balance of the rare earth
elements and neodymium isotopes in the oceans. Geochim. Cosmochim.acta 75,
Bocherens, H., Koch, P.L., Mariotti, A., Geraads, D., and Jaeger, J.-J. 1996. Isotopic
biogeochemistry (13C, 180) of mammalian enamel from African Pleistocene
hominid sites. Palaios 11, 306-318.
Brand, L.R., Esperante, R., Chadwick, A.V., Poma Porras, O., and Alomia, M. 2004.
Fossil whale preservation implies high diatom accumulation rate in the Miocene-
Pliocene Pisco Formation of Peru. Geology 32, 165-168.
Branstetter, S., Musick, J. A., Colvocoresses, J.A. 1987. A comparison of the age and
growth of the Tiger Shark, Galeocerdo cuveri, from off Virginia and from the
northwestern Gulf of Mexico. Fishery Bull. 85, 269-279.
Cailliet, G.M., Martin, L.K., Kusher, D., Wolf, P., and Welden, B.A. 1983. Techniques
for enhancing vertebral bands in age estimation of California elasmobranches. In
Proceedings of the international workshop on age determination of oceanic pelagic
fishes: tunas, billfishes, and sharks (eds E.D. Prince and L.M. Pulos). U.S. Dep.
Commer., NOAA Tech. Rep. NMFS 8. pp157-165.
Campana, S.E. 1997. Use of radiocarbon from nuclear fallout as a dated marker in
otoliths of haddock, Meganogrammus aeglefinus. Mar. Ecol. Prog. Ser. 150, 49-56.
Camapna, S.E., Natanson, L.J., and Myklevoll, S. 2002. Bomb dating and age
determination of large pelagic shraks. Can.. J. Aquat. Sci. 59, 450-455.
Cappetta, H. 1987. Handbook ofPaleoichthyology, Chliui iIn hiie' II. Gustav Fisher
Cappetta, H. 1981. Additional a la faune des Selaciens fossils du Maroc: sur la presence
des genres Heptranchias, Alopias et Odontorhytis dans l'Ypresien des Ouled
Abdoun. Geobios 14, 53-575.
Carlson, S. 1990. Vertebrate dental structures. In .,keletaIl biomineralization: Patterns,
Processes and Evolutionary Trends Volume 1 (ed. J. S. Carter). Van Nostrand
Reinhold, New York. pp. 531-556.
Carey, F.G., Casey, J.G., Pratt, H.L., Urquhart, D., and McCosker, J.E. 1985.
Temperature, heat production, and heat exchange in lamnid sharks. So. Calif. Acad.
Sci. Mem. 9, 92-108.
Carey, F.G., Kanwisher, J.W.,Brazier, O.,Gabrielson, G., Casey, J.G., and Pratt, H.L.
1983. Temperature and activities of a white shark, Carcharodon carcharias.
Copeia 2, 254-260.
Cerling, T. E., and Sharp, Z. 1996. Stable carbon and oxygen isotope analysis of fossil
tooth enamel using laser ablation. Palaeogeo. Palaeoclim. Palaeoecol. 126,173-
Compango, L.J.V. 2002. Sharks of the world: an annotated and illustrated catalog of
shark species known to date. FAO Species Catalog for Fishery Purposes 1, 96-125.
Compango, L.J.V. 1999. Chapter 3: Endoskeleton In .,\h/i, k, Skates, andRays: The
Biology ofElasmobranch Fishes (ed W.C. Hamlett).. Baltimore, Johns Hopkins
Univ. Press. pp 69-92.
deBaar, H.J.W., Bacon, M.P., Brewer, P.G., and Bruland, K.W. 1985a. Rare earth
elements in the Pacific and Atlantic Oceans. Geochim. Cosmochim. Acta 49, 1943-
deBaar, H.J.W., Brewer, P.G., and Bacon, M.P. 1985b. Anomalies in rare earth
distribution in seawater: Gadolinium and terbium. Geochim. Cosmochim Acta 49,
DeNiro, M.J., and Epstein, S. 1978. Influence of diet on the distribution of carbon
isotopes in animals. Geochim. Cosmochim. Acta 42, 495-506.
Denys, C., C. T. Williams, Y. Dauphin, P. Andrews, and T. Femandez-Jalvo. 1996.
Diagenetical changes in Pleistocene small mammal bones from Olduvai Bed 1.
Palaeogeo. Palaeoclim. Palaeoecol. 126, 121-134.
Druffel, E.M., and Linick, T.W. 1978. radiocarbon in annual coral rings of Florida.
Geophys. Res. Lett. 5, 913-916.
Dubinin, A.V. 2004. Geochemistry of rare earth elements in the ocean. Lith. Miner.
Resour. 39, 289-307.
Dubinin, A.V., and Rozanov, A.G. 2001. Geochemistry of rare earth elements and
thorium in sediments and ferromanganese nodules of the atlantic ocean. Lith.
Miner. Resour. 36, 368-279.
Elderfield, H. 1988. The oceanic chemistry of rare-earth elements. Phil. Trans. R. Soc.
Lond. A 325, 105-126.
Elderfield, H., Upstill-Goddard, R., and Sholkovitz, E.R. 1990. The rare earth elements in
rivers, estuaries, and coastal seas and their significance to the composition of ocean
waters. Geochim. Cosmochim. Acta 54, 971-991.
Elderfield, H., and Sholkovitz, E.R. 1987. Rare earth elements in the pore waters of
reducing near shore sediments. Earth Planet. Sci. Lett. 82, 280-288.
Elderfield, H., and Pagett, R. 1986. Rare earth elements in ichthyoliths: variations with
redox conditions and depositional environment. Sci. TotalEnviron. 49, 175-197.
Elderfield, H., and Greaves, M.J. 1982. The rare earth elements in seawater. Nature 296,
Featherstone, J.D., B, S. Pearson, and R. Z. LeGeros. 1984. An infrared method for
quantification of carbonate in carbonate apatites. Caries Res. 18, 63-66.
Ferreira, B.P., and Vooren, C.M. 1991. Age, growth, and structure of vertebra in the
school shark Galeorhinus galeus (Linnaeus, 1758) from southern Brazil. Fish Bull.
Fleet, A.J. 1984. Aqueous and sedimentary geochemistry of the rare earth elements. In
Rare Earth Element Geochemistry (ed. P. Henderson). Amsterdam: Elsevier. pp.
Francillon-Vieillot, H., de Buffrenil, V., Casternet, J., Geraudie, J., Meunier, F. J., Sire,
Y. J., Zylberberg, L. & de Ricqles, A. 1990: Microstructure and mineralization of
vertebrate skeletal tissues. In .\keleItil Biomineralization: Patterns, Processes and
Evolutionary Trends I 9ed J.G. Carter): New York: Van Nostrand Reinhold. pp
German, C.R., and Elderfield, H. 1990. Application of the Ceanom. as a paleoredox
indicator: the ground rules. Paleocean. 5, 823-833.
German, C.R., Holliday, B.P., and Elderfield, H. 1991. Redox cycling of rare earth
elements in suboxic zone of the Black Sea. Geochim. Cosmochim. Acta 53, 3179-
Goodrich, E.S. 1930. Studies on the Structure and Development of Vertebrates. London:
Gottfried, M.D., Fordyce, R.W. 2002. An associated specimen of Carcharodon
angustidens (Condrichthyes, Lamnidae) from the late Oligocene of New Zealand,
with comments on Carcharodon interrelationships. J. Vert. Paleontol. 21, 730-739.
Gottfried, M.D., Compagno, L.J.V., Bowman, S.C. 1996. Size and skeketal anatomy of
the giant "megatooth" shark Carcharodon megalodon. In Great White .h\lii k The
Biology of Carcharodon carcharias (ed A. P. Klimley and D. G. Ainley). San
Diego: Academic Press. pp. 55-66.
Grandjean-Lecuyer, P., Feist, R., and Albarede, 1993. Rare earth elements in old biogenic
apatite. Geochim. Cosmochim. Acta 57, 2507-2514.
Grandjean, P, and Albarede, F. 1989. Ion probe measurement of rare earth elements in
biogenic phosphates. Geochim. Cosmochim. Acta 53, 3179-3183.
Grandjean, P.H., Capetta, H., and Albarede, F. 1988. The REE and Nd of 40-70 Ma old
fish debris from the West-African platform. Geophy. Res. Lett. 15, 389-394.
Grandjean, P., Cappetta, H., Michard, A., and Albarede, F. 1987. The assessment of REE
patterns and 143Nd/144Nd ratios in fish remains. Earth Planet. Sci. Lett. 84, 181-196.
Grant, J.A. 1982. The isocon diagram A simple solution to Gresen's equation for
metasomatic alteration. Econ. Geol. 81, 1976-1982.
Greene, E.F., Tauch, S., Webb, E., and Amarasiriwardena, D. 2004. Application of
diffuse reflectance infrared Fourier transform spectroscopy (DRIFT) for the
identification of potential diagenesis and crystallinity changes in teeth. Microchem.
J. 76, 141-149.
Hattin, D. 1981. Petrology of Smoky Hill Member, Niobrara Chalk (Upper Cretaceous),
in type area, Western Kansas. AAPG Bull. 65, 831-849.
Hayashi, H., Kurihara, Y., Horiuchi, S., Iwashita, T., and Yanagisawa, Y. 2003.
Planktonic foraminferal biostratigraphy of the Miocene Sequence in the Iwadono
Hills, Central Japan: an integrated approach. Palaios 18, 176-191.
Henderson, P. Marlow, C.A., Molleson, T.I., and Williams, C.T. 1983. Patterns of
chemical change during bone fossilization. Nature 306, 358-360.
Herman, J., Steurbaut, E., and Vandenberghe, N. 2000. The boundary between the middle
Eocene Brussel Sand and the Lede Sand formations in the Zaventem-
Nederokkerzeel area (Northeast of Brussels, Belgium). Geol. Belgica 3, 231-255.
Hoefs, J. 1988. Stable Isotope Geochemistry, 3rd ed. Springer-Verlag, Berlin.
Holcomb, D.W., and Younf, R.A. 1980. Thermal decomposition of tooth enamel. Calcif
Tissue Int. 31, 189-201.
Hooyberghs, H.J.F. 1990. New palaeoecological studies in benthic foraminifera from the
Brussels Sands Formation (Lutetian, Middle Eocene) in Belgium. Bull. De la
Socidtd belge de Geol., 3, 337-354.
Hoyel, J., Elderfield, H., Gledhill, A., and Greaves, M. 1984. The behavior of the rare
earth elements during mixing of river and seawaters. Geochim. Cosmochim. Acta
Hubert, J.F., Panish, P.T., Chure, D.J., and Prostak, K.S. 1996. Chemistry,
microstructure, petrology, and diagenetic model of Jurassic dinosaur bones,
Dinosaur National Monument, Utah. J. Sed. Res. 66, 531-547.
lacumin, P., Bocherens, H., Mariotti, A., and Longinelli, A. 1996. Oxygen isotope
analysis of co-existing carbonate and phosphate in biogenic apatite: a way to
monitor diagenetic alteration of bone phosphate? Earth Planet. Sci.Lett. 142,1-6.
Kalish, J.M. 1993. Pre- and post-bomb radiocarbon in fish otoliths. Earth Planet. Sci.
Lett. 114, 549-554.
Kemp, R.A., and Trueman, C.N. 2002. Rare earth elements in Solnhofen biogenic apatite:
Geochemical clues to the palaeoenvironment. Sed. Geol. 155, 109-127.
Koch, P. L., Halliday, A.N.,Walter, L.M., Stearley, R.F., Huston, T.J., and Smith, G.R.
1992. Sr isotopic composition of hydroxyapatite from recent and fossil salmon: the
record of lifetime migration and diagenesis. Earth Planet. Sci. Lett. 108, 277-287.
Koeppenkastrop, D., and De Carlo, E.H. 1992. Sorption of rare earth elements from
seawater onto synthetic mineral particles: an experimental approach. Chem. Geol.
Kolodny Y., Luz B, Sander M., and Clemens W. A. 1996. Dinosaur bones: fossils or
pseudomorphs? The pitfalls of physiology reconstruction from apatitic fossils.
Palaeogeog. Palaeoclimat. Palaeoecol. 126, 151-160.
Kolodny, Y., and Luz, B. 1991. Oxygen isotopes in phosphates of fossil fish- Devonian
to Recent. In Stable Isotope Geochemistry: A Tribute to Samuel Epstein (eds. H.P.
Taylor, Jr., J.R. O'Neil, and I.R. Kaplan). The Geochemical Society, Special
Publication 3, 105-119.
Kolodny, Y., Luz, B., and Navon, O. 1983. Oxygen isotope variation in phosphate of
biogenic apatites, I. Fish bone apatite-rechecking the rules of the game. Earth
Planet. Sci. Lett. 64, 398-404.
Krueger, H. W. 1991. Exchange of carbon with biological apatite. J. Archaeol. Sci. 18,
Labeyrie, P.D.L., and Yiou, P. 1996. Macintosh program performs time-series analysis.
Eos 77, 379.
Labs-Hochstein, J. submitted. Oxygen isotopic analysis, and rare earth elements of
lamnid shark vertebral centra: determination of life history? Earth Plant. Sci. Lett.
Labs-Hochstein, J., and MacFadden, B.J. accepted. Quantification of diagenesis in
Cenozoic sharks: elemental and mineralogical changes. Geochim. Cosmochim.
Lacan, F., and Jeandel, C. 2001. tracing Papua New Guinea imprint on the central
Equatorial Pacific Ocean using neodymium isotopic composition and rare earth
element patterns. Earth Planet. Sci. Lett. 186, 497-512.
Laenen, B. Hertogen, J., and Vandenberghe, N. 1997. The variation of the trace-element
content of fossil biogenic apatite through eustatic sea-level cycles. Palaeogeog.
Palaeoclimat. Palaeoecol. 132, 325-342.
Lancelot, Y., and Seibold, E. 1978. The evolution of the Central Northeastern Atlantic:
Summary of results of DSDP Leg 41. Initial Report DSDP 41, 1215-1245.
Lecuyer, C., Reynard, B., and Grandjean, P. 2004. Rare earth element evolution of
Phanerozoic seawater recorded in biogenic apatites. Chem. Geol. 204, 63-102.
Lecuyer, C., Grandjean, P., O'Neil, J.R., Cappetta, H., and Martineau, F. 1993. Thermal
oceanic excursions at the Cretaceous-Tertiary boundary: the 6180 record of
phosphatic fish debris (northern Morocco). Palaeogeog. Palaeoclimat. Palaeoecol.
Lee-Thorp, J.A., and N. J. van der Merwe. 1991. Aspects of the chemistry of modem and
fossil biological apatites. J. Archaeol. Sci. 18, 343-354.
Longinelli, A. Ed. 1996. Biogenic phosphates as palaeoenvironmental indicators.
Palaeogeog. Palaeoclimat. Palaeoecol. 126, 31-44.
Longinelli, A., and S. Nuti. 1973. Oxygen isotope measurements of phosphate from fish
teeth and bones. Earth Planet. Sci. Lett. 20,337-340.
MacFadden B. J., Labs-Hochstein J., Quitmyer I., and Jones D. S. 2004. Incremental
growth and diagenesis of skeletal parts of the lamnoid shark Otodus obliquus from
the early Eocene (Ypresian) of Morocco. Palaeogeog., Palaeoclimat., Palaeoecol.
MacFadden, B.J., Cerling, T.E., Harris, J.M., and Prado, J. 1999. Ancient latitudinal
gradients of C3/C4 grasses interpreted from stable isotopes of New World
Pleistocene horse (Equus) teeth. Global Ecol. Biogeo. Lett. 8,137-149.
Manly, B.F.J. 1995. Multivariate Statistical Methods: A primer. Second edition.
Chapman and Hall, London.
Martin, A.P. 1996. Systematics of the Lamnidae and the origination time of Carcharodon
carcharias inferred from the comparative analysis of mitochondrial DNA
sequences In Great White .\lii kA The Biology of Carcharodon carcharias (eds
A.P. Kimley and d.G. Ainley). Academic Press, New York. pp49-54.
Michel, V., Ildefonse, P., and Morin, G. 1995. Chemical and structural changes in Cervus
elaphus tooth enamels during fossilization (Lazaret cave): a combined IR and XRD
Rietveld analysis. Applied Geochem. 10, 145-159.
Montaser, A.E. 1998. Inductively coupled plasma mass spectrometry. New York: Wiley.
Moss, M.L. 1977. Skeletal tissues in sharks. Amer. Zool. 17, 335-342.
Nathan, Y., and Sass, E. 1983. Stability relations of apatites and calcium carbonates.
Chem. Geol. 34, 103-111.
Nelson, D.G.A. 1981. The influence of carbonate on the atomic structure and reactivity of
hydroxyapatite. J Dental Res. 60,1621-1629.
Nelson, D.G.A., Featherstone, J.D.B., Ducan, J.F., and Cutress, T.W. 1983. Effect of
carbonate and fluoride on the dissolution behavior of synthetic apatites. Caries Res.
Newsley, H. 1989. Fossil bone apatite. Applied Geochem. 4, 233-245.
Nozaki, Y., Alibo, D.S., Amakawa, H., Gamo, T., and Hasumoto, H. 1999. Dissolved
rare earth elements and hydrography in the Sulu Sea. Geochim. Cosmochim. Acta
Picard, S., Lecuyer, C., Barrat, J., Garcia, J., Cromart, G., Sheppard., S.M. 2002. Rare
earth element content of Jurassic fish and reptile teeth and their potential relation
to seawater composition (Anglo-Paris Basin, France and England). Chem. Geol.
Picard, S., Garcia, J.P., Lecyuer, Sheppard, S.M.F. Cappetta, H., and Emig. C.C. 1998.
6180 values of coexisting brachiopods and fish: temperature differences and
estimates of paleo-water depths. Geology 26, 975-978.
Piepgrass, D.J., and Jacobsen, S.B. 1992. The behavior of rare earth elements in
seawater: precise determination of variations in the North Pacific water column.
Geochim. Cosmochim. Acta 56, 1851-1862.
Piper, D.Z. 1974. Rare earth elements in the sedimentary cycle: a summary. Chem. Geol.
Price, T.D.E. 1989. The Chemistry of Prehistoric Human Bone. Cambridge: Cambridge
Puceat,E., Reynard, B., and Lecuyer, C., 2004. Can crystallinity be used to determine the
degree of chemical alteration of biogenic apatites? Chem. Geol. 205, 83-97.
Reiche, I., Favre-Quattropani, L., Vignaud, C., Bocherens, H., Charlet, L., and Menu, M.
2003. A multi-analytical study of bone diagenesis: the Neolithic site of bercy
(Paris, France). Meas. Sci. Technol. 14, 1608-1619.
Reynard, B., Lecuyer, C., and Grandjean, P. 1999. Crystal-chemical controls on rare-
earth element concentrations in fossil biogenic apatites and implications for
paleoenvironmental reconstructions. Chem. Geol. 155, 233-241.
Ridewood, W.G. 1921. On the calcification of the vertebral centra in sharks and rays.
Phil. Trans. Royal Soc. London. Ser. B. 210, 311-407.
Rink, W.J., and Schwarcz, H.P. 1995. Tests for diagenesis in tooth enamel: ESR dating
signals and carbonate contents. J. Archaeol. Sci. 22, 251-255.
Samoilov, V.S., and Benjamini, C. 1996. Geochemical features of dinosaur remains from
the Gobi Desert, South Mongolia. Palaios 11, 519-531.
Shemesh, A. 1990. Crystallinity and diagenesis of sedimentary apatites. Geochim.
Cosmochim. Acta 54, 2433-2438.
Shemesh, A., Kolodny, Y., and Luz, B. 1983. Oxygen isotope variation in biogenic
apatites. II. Phosphate rock. Earth Plant. Sci. Lett. 64, 405-416.
Shields, G., and Stille, P. 2001. Diagenetic constraints on the use of cerium anomalies as
palaeoseawater redox proxies: an isotopic and REE study of Cambrian
phosphorites. Chem. Geol. 175, 29-48.
Sholkovitz, E.R., Landing, W.M, and Lewis, B.L. 1994. Ocean particle chemistry: The
fractionation of rare earth ellemnts between suspended particles and seawater.
Geochim. Cosmochim. Acta 58, 1567-1580.
Sibilia, J.P., Hammond, W.B. and Szobota, J.S. 1988. Molecular spectroscopy In A
Guide to Materials Characteristization and Chemical Analysis (ed. J.P. Sibilia).
New York: VCH Publishers. pp 13-19.
Starton, R.M., Grandstaff, B., Gallagher, W.B., and Grandstaff, D.E. 2001. REE
signatures in vertebrate fossils from Sewell, NJ: Implications for location of the K-
T boundary. Palaios 16, 255-265.
Stuart-Williams, H.Q., Schwarcz, H.P.White, C.D., and Spence, M.W. 1996. The isotopic
composition and diagenesis of human bone from Teotiuacan and Oaxaca, Mexico.
Palaeogeog. Palaeoclimat. Palaeoecol. 126, 1-14.
Stuiver,M., and Polach, H.A. 1977. Reporting od C-14 data. Radiocarbon 19, 355-363.
Tachikawa, K., Jeandel, C. Vangriesheim, A., and Dupre, B. 1999. Distribution of rare
earth elements and neodymium isotopes in suspended particles of the tropical
Atlantic Ocean (EUMELI site). Deep Sea Res. 46, 733-755.
Tricas, T.C., and McCosker, J.E. 1984. Predatory behavior of the white shark
(Carcharodon carcharias), with notes on its biology. Proc. Calif. Acad. Sci. 43,
Trueman, C.N. 1999. Rare earth element geochemistry and taphonomy of terrestrial
vertebrate assemblages. Palaios 14:555-568.
Trueman, C.N. 1996. Variation of REE patterns in dinosaur bones from North West
Montana: implications for taphonomy and preservation. Geoscientist 6, 27-30.
Trueman, C.N, Behrensmeyer, A. K.; Tuross, N.; and Weiner, S. 2004. Mineralogical and
compositional changes in bones exposed on soil surfaces in Amboseli National
Park, Kenya: diagenetic mechanisms and the role of sediment pore fluids. J. Arch.
Sci. 31, 721-740.
Trueman, C.N., and Tuross, N. 2001. Trace elements in recent and fossil bone apatite. In
Phosphates- Geochemical, geobiological, and materials importance (M.L. Kohn.,
Rakovan. J., and Hughes. J.M. eds.). Reviews in Mineralogy and Geochemistry. pp
Trueman, C.N., and Benton, M.J. 1997. A geochemical method to trace the taphonomic
history of reworked bones in sedimentary settings. Geology 25, 263-266.
Veeh, H.H., 1982. Concordant 230Th and 231 Pa ages of marine phosphorites. Earth.
Planet. Sci. Lett. 76, 45-56.
Vennemann, T. W., Hegner, E., Cliff, G., and G. W. Benz. 2001. Isotopic composition of
recent shark teeth as a proxy for environmental conditions. Geochim. Cosmochim.
Von Bertalanffy, L. 1960. Principles and theory of growth. In Fundamental Aspects of
Normal andMalignant Gi ,i i il (ed W.W. Nowinski).. Amsterdam: Elsevier. Pp.
Von Bertalanffy, L. 1938. A quantitative theory of organic growth (Inquiries on growth
laws. II). Human Biol. 10, 181-213.
Wang, Y., and Cerling T.E. 1994. A model of fossil tooth and bone diagenesis:
Implications for paleodiet reconstruction from stable isotopes. Palaeogeog.,
Palaeoclimat., Palaeoecol. 107, 281-289.
Weiner, S., and Proce, P.A. 1986. Disaggregation of bone into crystals. Calcif Tissue
Intern. 39, 365-375.
Weidman, C.R., and Jones, G.A. 1993. A shell-derived time history of bomb C-14 on
Georges Bank and its Labrador Sea implications. J. Geophys. Res. 98, 14577-
Whittacker, E.J.W., and Muntus, R. 1970. Ionic radii for use in geochemistry. Geochim.
Cosmochim. Acta 34, 945-956.
Williams, C. T. 1989. Trace elements in fossil bone. Applied Geochem. 4, 247-248.
Williams, C.T. 1988. Alteration of chemical composition of fossil bones by soil processes
and groundwater., In Trace elements in environmental history (eds. G. Grupe and
B. Herrmann). Berlin: Springer-Verlag. pp. 27-40
Wintner, S.P., and Cliff, G. 1999. Age and growth determination of the white shark,
Carcharodon carcharias, from the east coast of South Africa. Fish. Bull. 97,153-
Wright, L.E. and Schwarcz, H.P. 1996. Infrared and isotopic evidence for diagenesis of
bone apatite at Dos Pilas, Guatemala: Palaeodietary implications. J. Archaeol. Sci.
Wright, J., Schrader, H.M. and Holser, W.T. 1987. Paleoredox variations in ancient
oceans recorded by rare earth elements in fossil apatite. Geochim. Cosmochim. Acta
Yudin, K.G., and Cailliet, G.M. 1990. Age and growth of the gray smoothound, Mustelus
californicus limbatus from the east coast of South Africa. Fish. Bull. 94, 135-144.
Joann Labs Hochstein was born in Rockville Center, NY, on January 13, 1977.
She graduated from Commack High School on Long Island in June 1995. In August
1995 she attended South Dakota School of Mines and Technology in Rapid City, SD,
where her research was focused on vertebrate paleontology. In 1999 she graduated with
honors with her Bachelor of Science degree in geology. After, Joann attended the
University of Florida, focusing her research on the high resolution study of the variations
in paleointensity of the Earth's magnetic field. In May 2001 she received her Master of
Science degree in geological sciences. In August 2001 Joann entered the Ph.D. program
at the University of Florida. Her research focused on understanding how diagenesis
affects the chemistry of fossils. Joann will complete her Ph.D. at the University of
Florida in 2005.