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2011 Stories of the Proterozoic earth: Paleoproterozoic Chuanlinggou arcitarchs and Neoproterozoic ¹⁷O- depleted Barite Yongbo Peng Louisiana State University and Agricultural and Mechanical College, [email protected]
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Recommended Citation Peng, Yongbo, "Stories of the Proterozoic earth: Paleoproterozoic Chuanlinggou arcitarchs and Neoproterozoic ¹⁷O-depleted Barite" (2011). LSU Doctoral Dissertations. 2552. https://digitalcommons.lsu.edu/gradschool_dissertations/2552
This Dissertation is brought to you for free and open access by the Graduate School at LSU Digital Commons. It has been accepted for inclusion in LSU Doctoral Dissertations by an authorized graduate school editor of LSU Digital Commons. For more information, please [email protected]. STORIES OF THE PROTEROZOIC EARTH: PALEOPROTEROZOIC CHUANLINGGOU ACRITARCHS AND NEOPROTEROZOIC 17 O-DEPLETED BARITE
A Dissertation Submitted to the Graduate Faculty of Louisiana State University and Agricultural and Mechanical College in the partial fulfillment of the requirements for the degree of Doctor of Philosophy
in
The Department of Geology and Geophysics
by Yongbo Peng B.S., Yantai, Ludong University, 2003 M.S., Nanjing Institute of Geology and Palaeontology, Chinese Academy of Sciences, 2007
May 2010
This dissertation is dedicated to my grandfather.
ii ACKNOWLEDGMENTS
First of all, I would like to express my deepest appreciation to my major advisor Dr.
Huiming Bao, for his patience, friendship, and generosity. His enthusiasm for science, boundless energy, and never-ending ideas are examples that I will follow for my whole life. I am also grateful to each member of my graduate advisory committee: Dr. Gary
Byerly, Dr. Annette Summers Engel, Dr. Laurie Anderson, Dr. Ayman Okeil, and Dr.
Mostafa Elseifi for their help and instructive discussion and suggestions concerning my projects.
I am indebted to Dr. Xunlai Yuan and Dr. Chuming Zhou of the Nanjing Institute of Geology and Palaeontology, Chinese Academy of Sciences and Dr. Shuhai Xiao of
Department of Geosciences, Virginia Tech. Without their encouragement and financial support, I would not have finished this dissertation.
I would like to thank Dr. Duofu Chen and Dr. Dong Feng of the Guangzhou
Institute of Geochemistry, Chinese Academy of Sciences for providing some of the barite samples and for coordinating strontium isotope analysis.
Next, I would like to thank our laboratory members, Margarita Khachaturyan,
Issaku Kohl, Tao Sun, Xiaoqian, Li and Changjie, Liu. During the past five years we have been like a family and I have been most fortunate to have the company, assistance, and encouragement.
I also would like to thank the faculty, staff, and students of Department of Geology and Geophysics at LSU for their friendship and help with my studies.
I would like to offer special gratitude to my wife Guiying Su for her love and my beautiful girl Emily Peng for so much good luck she brings to me.
iii I would like to thank my parents for giving birth to me and supporting me spiritually all the way through my life.
This research was supported by National Natural Science Foundation of China
(40828002 to HB and 40625006 to XY), Chinese Academy of Sciences (KZCX2-YW-
153 and KZCX-SW-141to XY), Chinese Ministry of Science and Technology
(2006CB806400 and 2003CB716805 to XY), State Key Laboratory of Ore Deposit
Geochemistry (Chinese Academy of Sciences) (200903 to HB), Louisiana State
University (Faculty Research Grant program to HB), and National Science Foundation
(USA) (EAR-0952057 to HB), Shreveport and New Orleans Geologic Society scholarship and Harvey Broyles Endowment Fund Scholarship ( 2010-2011, to YP).
iv TABLE OF CONTENTS
ACKNOWLEDGMENTS...... iii
LIST OF TABLES...... vii
LIST OF FIGURES...... viii
ABSTRACT...... x
CHAPTER 1 INTRODUCTION TO STUDY...... 1 1.1 My Intellectual Journey and the Dissertation Process...... 1 1.2 Dissertation Objectives...... 9 1.3 References...... 11
CHAPTER 2 NEW MORPHOLOGICAL OBSERVATIONS FOR PALEOPROTEROZOIC ACRITARCHS FROM THE CHUANLINGGOU FORMATION, NORTH CHINA……………………………………...... ………………….……………14 2.1 Introduction………………...... ………….....………….……………………...14 2.2 Material and Methods ……………...... ……………………………...... ………....15 2.3 Results………………...... ……..19 2.4 Discussion……………………………………………………………….………..28 2.4.1 A Morphological Model for Chuanlinggou Ovoidal Acritarchs...... 28 2.4.2 Comparison to Mesoproterozoic Acritarchs from Roper Group and Ruyang Group...... 29 2.4.3 Is the Chuanlinggou Ovoidal Group the Earliest Eukaryote Body Fossil...... 31 2.5 Conclusions ……...... ……………………….…………………..…37 2.6 References...... 38
CHAPTER 3 NEOPROTEROZOIC 17 O DEPLETED BARITE FROM TWO MARINOAN CAP DOLOSTONE SECTIONS, SOUTH CHINA...... 43 3.1 Introduction…………………………………………………………………...... 43 3.2 Sedimentological and Petrographic Contexts of the 17 O-depleted Barite ...... 45 3.3 Methods ……………...... ………….....…..50 3.3.1 Isotope Analytical Methods...... 50 3.3.2 Modeling Methods...... 51 3.4 Results...... 56 3.4.1 Isotope Results...... 56 3.4.2 Modeling Results...... 57 3.5 Discussion...... 69 3.5.1 The Effect of Diagenesis...... 69 2− 3.5.2 An Isotopically Variable Sulfate Pool Reflecting a Low [SO 4 ] Ocean...... 69 3.5.3 Barite Snapshots of an Isotopically Variable Sulfate Pool...... 70
v 3.5.4 A Depositional Model for the 17 O-depleted Barite...... 73 3.6 Conclusions …………………………...... ……………………..…75 3.7 References...... 75
APPENDIX A. LETTERS OF PERMISSION...... 80
APPENDIX B. TRIPLE OXYGEN, SULFUR, AND STRONTIUM ISOTOPE DATA FOR BARITE ASSOCIATED WITH MARINOAN CAP DOLOSTONES FROM SONGLIN AND BAIZHU, SOUTH CHINA...... 81
VITA……………………………………………………………………………………86
vi LIST OF TABLES
Table 2-1 Comparing morphological features between Mesoperterozoic acritarchs from Roper Group or Ruyang Group and Paleoproterozoic Chuanlinggou acritarchs from this study...... 33
Table 3-1 Values used in the model...... 53
Table 3-2 Parameters and constraints for 28 control model runs...... 54
Table 3-3 Ranges and mean values of the ∆17 O, δ18 O, and δ34 S for B1 and B2 barite at Songlin and Baizhu...... 62
vii LIST OF FIGURES
Figure 2-1 Location and stratigraphy of the Changcheng Group, Jixian, northern China...... 16
Figure 2-2 Field photograph of sample locality near Jixian...... 18
Figure 2-3 Comparison between images of optical microscopy and of SEM on the same individual microfossil …...... ………………….....………...…..20
Figure 2-4 The ovoidal acritarchs from the Chuanlingguo Formation, northern China …………...... …….……………....…..21
Figure 2-5 The ovoidal acritarchs from the Chuanlingguo Formation, northern China ……………...... …………….………..….…...23
Figure 2-6 The spheroidal acritarchs from the Chuanlingguo Formation, northern China...... 25
Figure 2-7 Lengths of the longer axis of ovoidal acritarchs from Chuanlinggou Formation...... 26
Figure 2-8 A schematic of the fold pattern of an ovoidal acritarch and a schematic of the arc lengths along the longer axis (a) and the shorter axis (b)...... 34
Figure 2-9 TEM photomicrograph of extant algae ultrathin sections...... 36
Figure 3-1 Simplified stratigraphic column and paleogeography of the Yangtze Block during late Neoproterozoic to Cambrian times ...... 46
Figure 3-2 Outcrop photos of barite in cap dolostones at Songlin...... 47
Figure 3-3 Outcrop photos of cap dolostones and barite at Baizhu ...... 48
Figure 3-4 Microphotographs of polished slabs and thin sections of barite...... 49
Figure 3-5 Relationships between the 17 O and δ34 S for all B1 and B2 barite samples from Songlin (A) and from Baizhu (B)...... 59
Figure 3-6 Relationships between the δ 18 O and δ34 S for all B1 and B2 barite samples from Songlin (A) and from Baizhu (B) ...... 60
Figure 3-7 Relationships between the δ 18 O and 17 O for all B1 and B2 barite samples from Songlin (A) and from Baizhu (B)...... 61
Figure 3-8 Drill-sampled slabs and their corresponding δ 18 O, 17 O, and δ34 S data...... 63
viii Figure 3-9 Relationships between δ18 O and 17 O for crystal-fan-specific barite samples from Songlin and Baizhu...... 64
Figure 3-10 Relationships between 17 O and δ34 S for crystal-fan-specific barite samples from Songlin and Baizhu...... 65
Figure 3-11 Relationships between δ18 O and δ34 S for crystal-fan-specific barite samples from Songlin and Baizhu...... 66
Figure 3-12 Model results for runs R1-R6 plotted as a function of time from 0 −2.7 million years for variations in the 17 O, δ18 O, and δ34 S of marine sulfate....67
Figure 3-13 Model results for runs R1-R6 plotted for variations in the 17 O, δ18 O, and δ34 S of marine sulfate...... 68
Figure 3-14 Sketches showing a sequence of post-glacial depositional events on the Yangtze Blockcof South China ...... 71
ix ABSTRACT
This dissertation consists of two independent topics, both critical to understanding
Earth history during the Proterozoic Eon.
(1) Paleoproterozoic Chuanlinggou Acritarchs : The eukaryotic affinity of acritarchs from the Chuanlinggou Formation has been questioned because of lack of indisputable morphological evidence. In this study, Chuanlinggou acritarchs were examined by jointly using optical microscopy, scanning electron microscopy, and transmission electron microscopy. In most cases, an ovoidal group of acritarchs represents a half-vesicle following a complete longitudinal rupture, which is a morphological model different from a whole envelope with medial splits, as proposed by earlier studies. This ovoidal group displays a bipolar morphology, longitudinal rupture, and occasionally striated wall structures that are consistent with a eukaryotic affinity. Thus, the Chuanlinggou ovoidal acritarchs probably extend the eukaryotic body fossil record into the Paleoproterozoic, circa 200 million years earlier than the morphologically more complex acritarchs from the Mesoproterozoic Roper Group.
(2) Neoproterozoic 17 O-depleted Barite : Distinct, non-mass-dependent 17 O depletion was reported in barite from Marinoan cap dolostones, and has been interpreted as an indication of an extremely high-pCO 2 atmosphere. Understanding the origins of this barite — and particularly the source of sulfate — is critical to interpreting the anomalous
17 O signature and its implications for sulfur and oxygen cycles after Marinoan glaciation.
In this study, together with field, petrographic, and Sr isotope data, the expanded dataset
(1) confirms large variability in ∆17 O, δ18 O, and δ34 S of barite; (2) demonstrates a hyperbolic relationship between the ∆17 O and δ34 S; (3) reveals that individual barite crystal fans and fans of the same generation possess well-clustered sets of δ18 O, 17 O,
x and δ34 S values; and (4) shows that barite crystal fans of different layers bear different sets of 17 O, δ18 O, and δ34 S values. The study suggests that 17 O-depleted barite crystals were formed under supersaturation when Ba 2+ from sulfate-free deepwater came to mix with sulfate-bearing shallow water. The large variability in sulfur and triple-oxygen isotope composition and the high 87 Sr/ 86Sr ratios indicate that the two sites were sufficiently close to continents so that the isotopic composition of sulfate was easily influenced by changes in riverine flux.
xi CHAPTER 1
INTRODUCTON TO STUDY
1.1. My Intellectual Journey and the Dissertation Process
My undergraduate major was in biology, and my master’s thesis focused on the fossil record of biological evolution. Prior to joining the Louisiana State University (LSU)
Doctor of Philosophy (Ph.D.) program, I was fascinated with the emergence of animals, an important milestone in biological evolution. In particular, I was interested in the carbonaceous compression fossils Protoarenicola baiguashanensis , Pararenicola huaiyuanensis , and Sinosabellidites huainanensis from the early Neoproterozoic
Liulaobei and Jiuliqiao Formations (700-800 Ma) in Anhui, North China. These were previously interpreted as worm-like metazoans by Wang (1982) and Zheng (1980). If correct, then these specimens would be some of the earliest-known bilaterian animals.
To test this hypothesis, I worked with Lin Dong at Virginia Tech to reexamine these fossils. Together, we collected samples from Anhui province, northern China, in the summer of 2005. While Dong had responsibility for morphological observations, I used a maceration method to extract the carbonaceous compression fossils from rock and studied the ultrastructure by scanning electron microscopy (SEM) and transmission electron microscopy (TEM). However, the fossils were heavily carbonized. After examining the morphological features of a large population of these carbonaceous fossils, we concluded that previous interpretations of P. baiguashanensis and P. huaiyuanensis as worm-like bilaterian animals were weakly surpported. Instead, they can be alternatively interpreted as erect epibenthic organisms, possibly coenocytic algae. This study was published in
Palaeogeography, Palaeoclimatology, Palaeoecology in February of 2008 (Dong et al.,
2008), in which I was one of the co-authors.
1 My initial motivation in joining Dr. Bao’s group at LSU and working towards a
Ph.D. was to expand on my study of well-preserved acritarchs in a shale ( ca. 1,700 Ma) in North China, the work of my master’s thesis at the Chinese Academy of Sciences
(Peng et al., 2007). One group of those acritarch fossils displays some of the unambiguous morphological features supporting its eukaryotic affinity. The Eukarya are more complex than the other two domains (Bacteria and Archaea) in that Eukarya possesses a membrane-bound nuclei and a dynamic cytoskeletal and endomembrane systems.
The timing of the emergence of eukaryotic organisms in the geological record is important because it provides an independent line of evidence for the evolutionary history of life. Although still being debated, the earliest body fossil evidence for eukaryotes is believed to go back as early as ca . 1,500 Ma, from the Roper Group, North Australia
(Javaux et al., 2001). Acritarchs from the Chuanlinggou Formation ( ca. 1,700 Ma) of the
Jixian section (ca . 18,000 Ma to 800 Ma), northern China, were first reported by Xing and Liu (1973), and some were later described as eukaryotes by Yan (1982). A eukaryotic affinity of these acritarchs, however, remains unclear because of a lack of indisputable morphological evidence, though Chuanlinggou acritarchs are some of the oldest on the Earth and still have ultrastructure.
My master’s thesis had limited samples, and detailed morphological reconstruction was unattainable. During my first two years at LSU, I worked on the project, which was published in Precambrian Research (Peng et al., 2009). Upon completion of this acritarch project, I could have gone ahead to apply the same methods to study acritarch fossils in the older Changzhouguo Formation that was known to have similar shale beds occurrences. I did not think that would be interesting enough, although, because the same 2 ovoidal acritarchs were recently found in the Changzhougou Formation, underling the
Chuanlinggou Formation (Lamb et al., 2009).
Instead, I forged ahead on two new research paths. I tried to do a three-dimensional geometric reconstruction of these compressed individual ovoidal acritarchs to further test our model. Three-dimensional geometric reconstruction is an entirely new scientific field for computer and mathematical sciences. I attempted to resolve the problem by collaborating with Dr. Ye Duan in the Department of Computer Science at the University of Missouri in Columbia and Dr. Scott Baldridge in the Department of Mathematics at
LSU. Because too much information was lost during the folding and pressing processes of the vesicle, however, we found it difficult to achieve a unique solution for a 3-D image.
The second research path I took was related to the background of Earth’s atmospheric and oceanic condition in which acritarchs first evolved. The evolution of the
Earth’s atmosphere is largely a story of oxygen in the atmosphere. The accumulation of oxygen in the atmosphere is controlled by the burial of organic carbon and pyrite sulfur in the sediments, as hypothesized originally by Ebelmen (1845) and re-emphasized by
Garrels and Perry (1974) and Holland (1973). The amount of buried organic carbon,
13 however, did not change significantly over the Proterozoic Era because the δ Ccarb remained essentially constant, despite large fluctuations at the beginning and end of the
Proterozoic. To reconcile the two lines of evidence, Catling et al. (2005) hypothesized that an increased burial of sulfides over this period. If true, then the sulfur isotope composition (the δ34 S) of the seawater sulfate should have increased steadily during this period. If the δ34 S does not display such a trend, then it would support that hydrogen escape played a substantial role in the irreversible oxygenation of the Earth’s surface
3 environment in the Proterozoic. This in turn suggests that methane concentration was still high in the Proterozoic atmosphere.
At this time, there are few sulfur isotope data for seawater sulfate available to test of the validity of this hypothesis. Therefore, I decided to use high-resolution sulfur isotope record of trace amounts of carbonate-associated sulfate (CAS) within limestones or dolostones in the Jixian section (ca. 1,800 to 800 Ma), northern China, Catling’s (2005) hypothesis. The Jixian section is one of the best preserved and least altered Proterozoic sections in the world. However, the validity of using CAS needed be checked first because of the possibility of sulfide oxidation during post-depositional exposure or via the CAS extraction procedure. Together with Dr. Bao, I collected about 230 samples from the Jixian section in the summer of 2006. Considering the possibility of sulfide oxidation during post-depositional exposure or the CAS extraction procedure, I decided that it was essential to do a careful examination of the validity of using CAS in those pyrite-bearing carbonates.
During CAS extraction, pyrite contained in carbonate rocks can be oxidized to sulfate, which will affect the primary isotopic signatures of CAS containing carbonate rocks. The study of Marenco et al. (2008) shows that samples containing 1 wt% pyrite can result in up to −4‰ depletion in the resulting sulfur isotope ratios, simply from the aqueous oxidation of coexisting pyrite. The generation of secondary, post/syn-diagenetic
CAS, however, has not been investigated. This secondary sulfate could become locked in the carbonate rock. We hypothesize that sulfate generated from syn/post-diagenetic in- situ oxidation of pyrite is present in variable amounts within the carbonate host rocks proximal to the pyrite grains. Issaku Kohl and I have been working with Dr. Amitava Roy at the Center for Advanced Microstructures and Devices (CAMD) at LSU to test this 4 hypothesis. We are using sulfur K-edge X-ray absorption near edge structure, which will allow for the identification of elevated sulfate concentrations proximal to pyrite grains. If there was variation in peak intensity or sulfur species in observed proximal to pyrite grains, then secondary diagenetic sulfate would be verified. We prepared polished rock chips from pyrite-bearing carbonate, carbonaceous shale rocks, and standard pyrite more than two years ago. The X-ray beam at CAMD is still not set up yet at this time.
While waiting for the CAS project to go ahead, I was intrigued by the lucinid bivalves and endosymbiosis. The association between sulfur-oxidizing (thiotrophic) bacteria and the lucinid bivalve clade is particularly intriguing because the inferred antiquity of the relationship (>400 million years) seems at odds with the relatively loose ecologic linkage of living members. Because only half of genus-level lucinid taxa are extant and the δ13 C of shell carbonate exhibits no systematic difference between symbiotic and non-symbiotic bivalves (Jones et al., 1988; CoBabe, 1991), a new morphologically independent proxy to determine whether fossil taxa possessed thiotrophic endosymbionts has been needed. The δ34 S of CAS in bivalve shells may hold promise because biogenic carbonate incorporates sulfate into its crystal structure during
2- 34 biomineralization. Incorporation of microbially derived SO 4 (with a more negative δ S value due to its reduced sulfur origin) into the lucinid-shell crystal lattice would,
34 2- therefore, impart a distinctly lower δ SCAS value than that from seawater SO 4 and would be distinguishable from CAS values of co-occurring heterotrophic bivalves.
34 18 In the summer of 2007, I measured CAS contents, δ SCAS and δ OCAS values of 15 sets of lucinid, and co-occurring infaunal and epifaunal heterotrophic bivalve shells collected by Dr. Laurie Anderson and Dr. Annette Engel from LSU. These were from modern and Cenozoic shallow marine sites. The data show that the modern bivalve shells
5 had variable CAS content, from 100 to 2,600 p.p.m. Epifauna often had the highest
34 18 concentrations relative to the other ecological groups. The δ SCAS and δ OCAS clustered at values corresponding to modern seawater sulfate, but with significant scatter. There was no systematic isotopic compositional difference among all of the bivalves in the same habitat or among the same lucinid, infaunal, or epifaunal groups across different sites. The fossil bivalve shells tended to preserve lower CAS concentrations, and the isotope compositions further deviated from seawater values. These data suggest that (1) pore-water sulfate in shallow sediments is highly heterogeneous in its concentration and isotope composition, probably due to active microbial sulfate reduction, bioturbation, and water-pumping by bivalves and other infaunal filter feeders; (2) CAS is derived from ambient pore water or pumped-in seawater for infauna or epifauna, as well as for lucinids; and (3) CAS concentration and isotope compositions are vulnerable to later diagenetic processes. This project has been finished, but the writing of the manuscript is delayed by
34 the need to repeat some δ SCAS analyses.
In late 2006 and early 2007, the Oxy-Anion Stable Isotope Consortium (OASIC) at
LSU, was largely taken over by an unexpected discovery — non-mass-dependently depleted 17 O in sulfate of barite in cap carbonates from the Marinoan “snowball Earth” episode. By the time the work was published (Bao et al., 2008), we knew that the highest anomaly should be preserved in massive Marinoan diamictites. There was an urgent need to test this hypothesis. One of my friends, Dr. Dong Feng in the Department of
Oceanography and Coastal Sciences at LSU, told me he had barite samples from a carbonate lens in Marinoan diamictites from Songlin, South China. He and I collected more samples of barite from the “carbonate lens” in the spring and summer of 2008.
However, we found that the so-called “lens” is in fact an artifact of a small fault in the
6 area based on the field survey in the summer of 2008. This means that the “carbonate lens” and barite within is the same as those reported in Bao et al. (2008).
Thereafter, my research focus shifted to understanding why there was a significant variation in the magnitude of the 17 O anomaly, even among barite samples from the same hand specimen, as identified by Bao et al. (2008). In Chapter 3 of my dissertation origins of this barite are determined by combining sedimentary and geochemical evidence.
Understanding the origins of this barite and its anomalous 17 O signature is critical to interpreting the anomalous 17 O signature and its implication on sulfur and oxygen cycles in the aftermath of the Marinoan glacial.
My dissertation focus is on two cap dolostone sections, the Songlin and Baizhu, where barite is well developed. I used a high-resolution sampling approach by which barite can be analyzed with spatial resolution at millimeters for its ∆17 O, δ18 O, δ34 S, and
87 Sr/ 86 Sr. Zhou, Bao, myself, and other colleagues are in the meantime trying to piece together the sedimentological and tectonic context and the event sequence associated with the deposition of the 17 O-anomalous barite. My multiple isotope data for barite turned out to be crucial to pinning down the exact sequence of events. We proposed that the cap dolostones were uplifted due to isostatic rebound and experienced karstic dissolution in both shallow platform and transitional facies of the Yangtze Block. The 17 O-depleted barite minerals precipitated when the transgression flooded the karstified cap dolostones shortly after the initial meltdown of the Marinoan snowball Earth. This study was published by Geology in October of 2010 (Zhou et al., 2010).
While focusing on barite from the two sites in South China, I was seeking to expand my search for barite or sulfate hosts in potential Marinoan continental diamictites
7 at one site in central Guizhou, South China. I found malachite (Cu 2(CO 3)(OH) 2) in the
Nantuo purple diamictites with associated mineral barite (BaSO 4), cuprite (Cu 2O), and
34 chalcocite (Cu 2S). Preliminary data show that (1) barite and chalcocite have similar δ S values of about −12.0‰ (VCDT); and (2) δ18 O value of barite is as low as –20.3‰
(VSMOW). These findings are important because the sulfate in barite and cuprite may have recorded the stable oxygen isotopic compositions of melted ice water of the snowball Earth, a piece of information that is highly informative but has been lacking to date.
This is still an on-going project, and I do not think it is developed enough to be included as one of my dissertation chapters. I hope that part of the work will become something I can explore in a postdoctoral position. For example, given that the textures of diamictites are so complex and individual grains are so small, it is not applicable for conventional techniques to analyze stable oxygen isotope compositions of cuprite. The high precision in situ oxygen isotope measurements by secondary ion mass spectrometry
(SIMS) become an ideal choice to address this problem. I have submitted a proposal to the NASA postdoctoral fellowship program to pursue this possibility. The specific objectives of this project are to (1) determine the temperature dependent fractionation of oxygen isotopes between water and Cu 2O; (2) develop an isotopically homogenous barite and cuprite standard for SIMS analysis; and (3) in situ analysis of stable oxygen isotope compositions of barite and cuprite in diamictites. This study will help in understanding the conditions of the atmosphere, hydrosphere, and biosphere after retreat of the
Marinoan glaciers. In particular, cuprite will become a new tool used to monitor ancient environments. Additionally, this study will serve as an example for in situ analyzing
8 stable isotope compositions of barite and cuprite by SIMS. The proposal has already passed the initial round of competition.
In addition to my interest in barite sulfate of the Neoproterozoic age, I have also explored sulfate minerals of other time periods. In the summer of 2008, I collected barite from the Wumishan Formation ( ca. 1,000 Ma) from Jixian, North China, and alunite
(KAl 3(SO 4)2(OH) 6) from the Mesoproterozoic Yanshan Formation, Shuangxiwu Group,
South China. I gathered these samples in hopes that a temporal evolution of sulfate triple oxygen isotope composition can be constructed. Preliminary data of barite from the
Wumishan Formation show that ∆17 O, δ34 S and δ18 O range from −0.18‰ to −0.32‰, from 8.4‰ to 10.4‰, and from 22.8‰ to 23.7‰, respectively. The preliminary data of alunite from the Yanshan Formation show that ∆17 O, δ34 S and δ18 O range from −0.06‰ to −0.08‰, from 11.6‰ to 13.1‰, and from 1.1‰ to 2.0‰, respectively. These data can serve as the seeds from which to develop my future research projects.
1.2 Dissertation Objectives
After describing my intellectual journey during the past five years at LSU, I hope I have depicted the landscape in which this dissertation sits. The atmosphere, biosphere, hydrosphere, and lithosphere of the Earth affect each other and co-evolve over time.
Understanding the evolution of each component is equally important to understanding the history of the Earth’s system. The Proterozoic Eon spans from 2,500 Ma to 542 Ma, which comprises over 40% of all the Earth’s history. It is the most intriguing period in
Earth’s history because many important events involving the atmosphere, biosphere, hydrosphere, and lithosphere occurred, including global glaciation and the rise of atmospheric oxygen in the Paleoproterozoic and Neoproterozoic Eras (Hoffman et al.,
1998; Catling and Claire, 2005; Kopp et al., 2005); the origin of eukaryotes in the
9 Paleoproterozoic (Javaux et al., 2003); and animals in the Neoproterozoic Era (Yin et al.,
2007). These events resulted in irreversible changes in the Earth’s system and paved the way for the Phanerozoic evolution of life.
The Proterozoic Era began and ended with episodes of global glaciation (Hoffman et al., 1998; Hoffman and Schrag, 2002; Kopp et al., 2005). Highly positive or negative
δ13 C excursions were reported for carbonates deposited before and after the glaciations in many parts of the world (Karhu and Holland, 1996; Melezhik et al., 1999; Bekker et al.,
2001; Hoffman and Schrag, 2002; Hurtgen et al., 2005). This indicates dramatic variations in organic and inorganic carbon budgets. The net result of the Paleoproterozoic and Neoproterozoic glaciations was probably a dramatic increase in atmospheric O 2 concentration (Rye and Holland, 1998; Catling and Claire, 2005). Since aerobic respiration and synthesis of membrane lipids (sterols and unsaturated fatty acids) by eukaryotic organisms require certain levels of molecular oxygen (Ourisson et al., 1987;
Catling et al., 2005), the link between the rise of atmospheric O 2 concentration in the
Paleoproterozoic and the appearance of sophisticated morphologies of eukaryotes at ca .1,700 Ma is appealing.
In view of the fact that the aerobic respiration of animals may require a higher level of molecular oxygen (Catling et al., 2005; Sreenivas and Murakami, 2005), the emergence of animals was likely the consequence of the rise of atmospheric O 2 concentration after the Neoproterozoic Marinoan “snowball Earth” event (Hoffman et al.,
1998; Knoll, 2003). This dissertation, therefore, focuses on two topics — the origin of eukaryotes in the Paleoproterozoic Era and a “snowball Earth” event in the
Neoproterozoic Era. The former is an important milestone in biological evolution, while
10 the latter marks a critical shift in the Earth’s surface environment at the dawn of animal evolution.
1.3 References
Bao, H.M., Lyons, J.R., and Zhou, C.M., 2008, Triple oxygen isotope evidence for elevated CO 2 levels after a Neoproterozoic glaciation: Nature, v. 453, p. 504-506.
Bekker, A., Kaufman, A.J., Karhu, J.A., Beukes, N.J., Swart, Q.D., Coetzee, L.L., and Eriksson, K.A., 2001, Chemostratigraphy of the Paleoproterozoic Duitschland Formation, South Africa: implications for coupled climate change and carbon cycling: American Journal of Science, v. 301, p. 261-285.
Catling, D.C., and Claire, M.W., 2005, How Earth's atmosphere evolved to an oxic state: A status report: Earth and Planetary Science Letters, v. 237, p. 1-20.
Catling, D.C., Glein, C.R., Zahnle, K.J., and McKay, C.P., 2005, Why O 2 is required by complex life on habitable planets and the concept of planetary "oxygenation time": Astrobiology, v. 5, p. 415-438.
CoBabe, E., 1991, Detection of chemosymbiosis in the fossil record: the use of stable isotopy on the organic matrix of lucinid bivalves, in Dudley, E., ed., The Unity of Evolutionary Biology:Proceedings of the Fourth International Congress of Systematic and Evolutionary Biology: Portland, OR, Bioscoroides Press, p. 669-673.
Dong, L., Xiao, S.H., Shen, B., Yuan, X.L., Yan, X.Q., and Peng, Y.B., 2008, Restudy of the worm-like carbonaceous compression fossils Protoarenicola, Pararenicola, and Sinosabellidites from early Neoproterozoic successions in North China: Palaeogeography Palaeoclimatology Palaeoecology, v. 258, p. 138-161.
Ebelmen, J.J., 1845, Sur les produits de lad´ecomposition des especes min´erales de la famille des silicates: Annales Mines v. 7, p. 3-66.
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12 Runnegar, B., 1991, Precambrian oxygen levels estimated from the biochemistry and physiology of early eukaryotes: Palaeogeography Palaeoclimatology Palaeoecology, v. 97, p. 97-111.
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13 CHAPTER 2
NEW MORPHOLOGICAL OBSERVATIONS FOR PALEOPROTEROZOIC ACRITARCHS FROM THE CHUANLINGGOU FORMATION, NORTH CHINA *
2.1 Introduction
The origin of eukaryotes is one of the most important milestones in evolutionary history. C 28 -C30 steranes extracted from Archean shale (Brocks et al., 1999) and sequence comparisons of proteins (Hedges et al., 2001) place the minimum timing for the origin of eukaryotes at ca. 2,700 Ma. Convincing eukaryotic body fossils, however, only go back as far as ca. 1,500 Ma, represented by microfossils such as Tappania sp., Satka favosa , and Valeria lophostriata from the Roper Group, northern Australia (Javaux et al., 2001;
Javaux et al., 2004).
Acritarchs are organic-walled microfossils of uncertain affinity, mostly interpreted as metabolically inert resting stages of algae (Strother, 1996). Acritarchs from the
Chuanlinggou Formation ( ca. 1,700 Ma), Changcheng Group, North China, are some the oldest acritarchs found on Earth. Xing and Liu (1973) first documented the Chuanlinggou acritarchs, and Yan (1982) later described some of the acritarchs as eukaryotes. Their eukaryotic affinity remained dubious, however, for absence of indisputable morphological evidence (Knoll et al., 2006). The Chuanlinggou Formation was re- sampled in this study and contained abundant well-preserved acritarchs. Their morphological features of these acritarchs were examined by jointly using optical microscopy, scanning electron microscopy (SEM), and transmission electron microscopy
(TEM). Although morphologically simpler than acritarchs found in the Mesoproterozoic
Roper Group, the Paleoproterozoic Chuanlinggou acritarchs possess many of the key morphological features of eukaryotes.
* Modified from Peng , Y., Bao, H., and Yuan, X., 2009, New morphological observations for Paleoproterozoic acritarchs from the Chuanlinggou Formation, North China: Precambrian Research, 168(3-4), 223-232 . Reprinted by permission of Precambrian Research (Appendix A). 14 2.2 Material and Methods
Acritarchs were collected from the lower member of the Chuanlinggou Formation of the Changcheng Group in Jixian County (40°12' N, 117°29'E), ca. 100 km east of
Beijing, northern China (Chen et al., 1980) (Figure 2-1 and Figure 2-2A). The
Changcheng Group ( ca. 2,700 m thick), consisting of the Changzhougou, Chuanlinggou,
Tuanshanzi, Dahongyu, and Gaoyuzhuang Formations, in ascending order (Figure 2-1), is largely un-metamorphosed and overlies nonconformably an Archean metamorphic complex (Chen et al., 1980). Residing at the lower part of the Changcheng Group, the
Chuanlinggou Formation consists of black shale with interbedded siltstones (Figure 2-
2A-B), representing a nearshore or shallow-marine environment (Yan and Liu, 1998).
The underlying Changzhougou Formation consists mainly of coarse sandstones and conglomerates while the overlying Tuanshanzi Formation is dominantly argillaceous dolostones (Figure 2-1). The sequence of the Changcheng Group displays a typical transition from a nearshore high-energy environment to an offshore, low-energy setting during the early period of basin development (Yan and Liu, 1998). A U-Pb SHRIMP detrital zircon age of ca. 1,750-1,800 Ma from the bottom of the underlying
Changzhougou Formation represents the maximum age for the onset of the Changcheng
Group sedimentation (Wan et al., 2003). A single grain U-Pb volcanic zircon age of
1,683 ± 67 Ma for the middle of the overlying Tuanshanzi Formation (Li and Lu, 1995), and in conjunction with a few Pb-Pb ages within the Chuanlinggou Formation, e.g., 1,705
±42 Ma (illite, Lu et al. 1989), 1,757±113Ma (shale, Li et al. 1984), and 1,785±19 (illite,
Kusky and Li, 2003), suggests an age of 1,700 ± 50 Ma for the Chuanlinggou Formation
(Figure 2-1).
Individual acritarchs were extracted from siltstone and shale samples using a
15
Figure 2-1. Location and stratigraphy of the Changcheng Group, Jixian, northern China, modified from Chen et al. (2008).
16 standard palynological techniques (Vidal, 1988; Grey, 1999). Ten petrographic thin sections of the acritarch-bearing rocks were cut parallel to bedding to examine in situ preservation of acritarchs. Rock sample surfaces were cleaned before subsequent wet- chemistry treatment. About 100 g samples were crushed to < 5 mm and washed three times with distilled water. Samples were then transferred to a 500 ml plastic screw- topped container and treated with a 48% hydrofluoric acid (HF) solution until very little residual sample was left. The residues normally settled to the bottom of the container. We decanted the solution, added distilled water, allowed residues to settle out, and decanted again. This washing process was repeated until decanted solution was neutral in pH.
Residue was transferred to a 500 ml glass beakers and ca. 250 ml of 32% hydrochloric acid was added to the sample. The sample in the beaker was gently boiled until the greyish, gel-like fluorides had dissolved and the solution was allowed to cool. Again, solution was decanted, and the residue was washed until it reached a neutral pH. Finally, microfossils were obtained by filtration through a sieve (pore size =20 m).
We noted that shale in the lower part of the shale-silt/sandstone assemblage of the
Chuanlinggou Formation is especially rich in acritarch fossils. Extracted fossils were pipetted onto a glass slide. Individual acritarchs were picked up under a stereoscope and were first observed by optical microscopy and then coated with 22 nm layer of gold palladium for SEM observation using a JEOL JEM-1230 microscope. For TEM, individual microfossils were entombed on filter membrane and dehydrated in a series of ethanol solutions, infiltrated with a mixture of propylene oxide/ethanol, followed by propylene oxide/ethanol resin. After dehydration, the membrane with microfossils was polymerized by Epon 812 in an oven at 37 ˚C for 12 h, 45 ˚C for 12 h, and 65 ˚C for 12 h.
17
Figure 2-2. (A) Field photograph of sample locality near Jixian (40°12' N, 117°29'E), hammer for scale. (B) Black shale with interbedded siltstones from the Chuanlinggou Formation.
18 Ultra-thin sections (40-50 nm thick) were cut by LEICA ULTRACUT UCT with a diamond knife and observed under a Leica M216A at Nanjing Institute of Geology and
Paleontology and a JEOL 100CX at Louisiana State University (LSU). For modern algae, we used the same procedures but first fixed the algae in 1% osmium tetroxide (phosphate buffered, pH= 7.2) for 1 hour before dehydration.
A aspect ratios of the apparent longer to shorter arc of 33 well-preserved ovoidal acritarchs were measured (Adobe Photoshop CS3 software) while considering folding and overlapping.
2.3 Results
A total of ca. 800 individual acritarchs were observed for this study. On the basis of their color (orange-brown), the Chuanlinggou acritarchs appear to be an early mature on
Thermal Alternation Index (Batten, 1996), suggesting that the Chuanlinggou acritarchs have undergone minimal thermal alternation.
The acritarchs of Chuanlinggou Formation have been divided into ca. 80 species since first reported by Xing, Yan, and their colleagues (Xing and Liu, 1973; Yan, 1982;
Yan and Liu, 1998). Comparison between images of optical microscopy and of SEM on the same individual microfossils indicates that some morphological features previously used for assigning species are questionable, however. For example, SEM shows that the
“reticulate” structure observed by optical microscopy is actually the result of erosion and pressure by mineral grains (Figure 2-3A-C), and the striations, different openings, inner bodies, or multi-layer membrane (Figure 2-4D-E) observed by optical microscopy were caused by various types or degrees of folding on a vesicle. Overall, our observations identified two basic acritarch morphologies: one ovoidal (Figure 2-4 and 2-5) and one spheroidal (Figure 2-6).
19
Figure 2-3. (A) Optical photomicrograph, (B), and (C) SEM photomicrograph) Compare between images of optical microscopy and of SEM on the same individual microfossil shows reticulate structure on surface of vesicle result from erosion and pressure by mineral grain.
20 Figure 2-4. The ovoidal acritarchs from the Chuanlingguo Formation, northern China. (A-B) Optical photomicrographs showing multiple fold lineations parallel to the longer axis of the vesicles. (C-E) SEM photomicrographs of the (B) showing smooth rupture lines and smooth surface of the vesicle; close-up views (D) and (E) of the two ends of the (C) arguing against the possibility of forming an ellipsoidal body by the ‘sheet’ itself and therefore supporting that it is most likely part of the vesicle. (F) Thin-section photomicrograph of an in-situ occurrence of the ovoidal acritarchs in the Chuanlinggou shale. (G-I) SEM photographs of striated ornamentation consisting of ridges spaced at 0.2 to 0.3 µm on the inner surface of the vesicle; close-up views (H) of red rectangle part of (G); close-up views (I) of red rectangle part of (H). (J-L) TEM photomicrographs of an ultrathin-section perpendicular to the longer axis of the ovoidal acritarchs (J), the ultrastructure of a wall with alternating electron-dense and electron-tenuous bands spaced at 0.2 to 0.3 µm (K), and trilaminar structure (TLS) of a wall that has two thin (0.01 µm) electron-dense layers (arrows) sandwiching a thick (0.1 to 0.3 µm in thickness) electron- tenuous homogeneous layer (L).
21 22 Figure 2-5. The ovoidal acritarchs from the Chuanlingguo Formation, northern China. (A-D) SEM photomicrographs of an acritarch displaying a longitudinal rupture with a straight and smooth rupture line in its vesicle (A) and close-up views of the split (B), (D) and (C). (E) Optical photomicrograph of an acritarch with two half vesicles attached. F-H: SEM photomicrographs of the (E). (I-L) A proposed model for an original olive- shaped acritarch that had two somewhat equal-volume half vesicles. The fold pattern (J) is based on a fossil individual (I) (light microphotograph). Cartoon (K) is the three- dimensional reconstruction of the fossil (I). The longitudinal striations depicted in (K) represent those observed both by SEM (Figure 4H) and TEM (Figure 4J). The model shows that the ovoidal acritarchs (L) consists of two parts that were often preserved as two individual ovoidal fossils when enrolled and compressed.
23
24
Figure 2-6. The spheroidal acritarchs from the Chuanlingguo Formation, northern China. (A) (Optical photomicrograph) and (B) (SEM photomicrograph) Showing smooth surfaces and irregular folds of the vesicle. (C) Thin-section photomicrograph of an in-situ occurrence of a spheroidal acritarch in the Chuanlinggou shale. (D) TEM photomicrograph of a folded and compressed vesicle of a spheroidal acritarch. (E) TEM photomicrograph of the trilaminar structure (TLS) of a spheroidal acritarch – the same as that of an ovoidal acritarch (Figure 1-4K).
25
Figure 2-7. Lengths of the longer axis of 128 ovoidal acritarchs from Chuanlinggou Formation observed under optical microscopes.
26 The ovoidal acritarchs are large in size with the longer axis measuring from 40 to
250 µm and an average value of 125 µm (n =128) observed by optical microscopy
(Figure 2- 7). The aspect ratios of the apparent longer arc (a) to the shorter one (b) are tightly clustered at 1.47 (s = 0.21) (Figure 2-8A-B). Most individuals show multiple folds on the surface layers that have fold lineations parallel to the longer axis of the vesicles
(Figure 2-4A-F). An important observation is that most ovoidal individuals exhibit elliptical and smooth edges that cannot by themselves form an enclosed three dimensional body when overlaps and folds are considered (Figure 2-4A-E and Figure 2-
5I-L). Both SEM and optical microscope examinations on the same specimens reveal that most of the ovoidal acritarchs vesicles are in fact roll-ups of a piece of organic ‘sheet’.
The edges of the ‘sheet’ are mostly smooth (Figure 2-4C-E), suggesting that the rupture pattern had an underlying structural constraint.
Further SEM examination on the vesicle walls of ca. 100 ovoidal acritarchs individuals shows that all the outer surfaces and most of the inner surfaces are smooth
(Figure 2-4C). A few ( ca. 3 out of 100) individuals’ inner walls, however, display regular striations with ridges parallel to the longer axis of the vesicles and are spaced at 0.2 to 0.3
µm (Figure 2-4G-I). Subsequent TEM observation of the wall ultrastructure reveals that regular or irregular electron-dense and electron-tenuous bands are found for six of all eight examined ovoidal individuals (Figure 2-4J-K). The bands are also spaced at 0.2 to
0.3 µm in ultra-thin sections that were cut perpendicular to the longer axis of the ovoidal vesicles, and are less well defined for non-perpendicular sections. Finally, TEM reveals that some segments of the wall display a typical trilaminar structure (TLS) in which a thick inner electron-tenuous layer (0.1 to 0.3 µm) is sandwiched between two thin ( ca.
0.01 µm) electron-dense surface layers (Figure 2-4L).
27 The other group of the Chuanlinggou Formation, the spheroidal acritarchs, have a smaller average size with a diameter ranging from 25 µm to 200 µm (Peng et al., 2007).
Despite the many irregular folds on the smooth surface of the vesicles, they are enclosed envelopes and do not possess either an identifiable rupture pattern (Figure 2-6A-C) or a consistent longitudinal rupture typical of the ovoidal acritarchs. The ultrastructure of the wall shows the same TLS as the ovoidal acritarchs when observed by TEM (Figure 2-6D-
E) (Peng et al., 2007).
2.4 Discussion
2.4.1 A Morphological Model for Chuanlinggou Ovoidal Acritarchs
The observed morphological features suggest that the original ovoidal acritarch was an olive-shaped body with two somewhat equal-volume half vesicles, as depicted in
Figure 2-5 (I-L), and most of the ovoidal fossil vesicles are in fact compressed roll-ups of these half vesicles. This morphological model is also supported by two rare fossil vesicles
(two out of 140 individuals) that have their two halves still attached and straight longitudinal rupture lines displayed (Figure 2-5A-H). Furthermore, the aspect ratios of the apparent longer arc (a) to the shorter one (b) indicate that the longitudinal rupture had a consistent pattern, most likely a complete medial split of an olive-shaped vesicle
(Figure 2-8A-B). In this morphological model, the almost exclusively longitudinal fold lineations or ruptures are probably constrained by the underlying ultrastructure of the vesicle wall: a series of longitudinally parallel striations revealed as ridges on inner surfaces or as electron-dense or electron-tenuous bands in ultra-sections. The bands are unlikely to be artifacts because the bands were only seen in ovoidal and not in spheroidal acritarchs that were treated with the same preparation techniques (Peng et al., 2007). This
28 model can be tested by future three-dimensional geometric reconstruction of these compressed individuals.
The ovoidal acritarchs from the Chuanlinggou Formation have been named many different genera, including Schizofusa , Schizospora , Shizovalvia , Thecatovalvia ,
Diplomembrane , Goniocystis , Leiovalia , Valvimorpha , Leioarachnitum , Scaphita ,
Trachyarachnitum (Yan, 1982, 1985, 1995). They were treated as a complete envelope by Yan (1982), not a half envelope as we proposed here. Schizofusa , Schizospora,
Diplomembrane , Goniocystis Leiovalia , Valvimorpha , Leioarachnitum , Scaphita ,
Trachyarachnitum are characterized by a single medial split along the long axis, but with different sizes of opening or degrees/numbers of folds among them. The vesicles of
Shizovalvia and Thecatovalvia display two separate parts with one or two splits.
According to our model, however, Schizofusa , Schizospora , Diplomembrane , Leiovalia ,
Valvimorpha , Leioarachnitum , Scaphita , Trachyarachnitum , and Goniocystis may be only half of a complete envelope resulting probably from the split of Shizovalvia and
Thecatovalvia . Butterfield and Chandler (1992) reported some groups of spheroidal acritarchs with medial splits in the ca. 1,250 Ma Agu Bay Formation, northern Canada, which are similar to Shizovalvia , one of Chuanlinggou ovoidal acritarchs, described by
Yan (1985).
2.4.2 Comparison to Mesoproterozoic Acritarchs from Roper Group and Ruyang Group
Microfossils from Roper or Ruyang Group are the oldest arcitarchs except for those from the Changcheng Group. The Chuanlinggou ovoidal acritarchs share a few morphological features with them, but in general are much less complex than those
Mesoproterozoic acritarchs (Table 2-1).
29 Tappania is characterized by asymmetric distribution of processes, a neck-like tubular extension, and a single-layered electron-dense and homogeneous wall ultrastructure (Yin et al., 2005). The wall of Satka favosa consists of polygonal plates up to 15 microns in maximum dimension that form a tessellated pattern as observed by SEM
(Javaux et al., 2004). Dictyosphaera is ornamented with reticulate vesicle surface and with interlocked polygonal plates (Yin et al., 2005). Neither processes nor plates are observed in Chuanlinggou acritarchs. Valeria lophostriata has a concentric striation pattern on the surface as observed by optical microscopy and a homogeneous electron- dense wall ultrastructure by TEM (Javaux et al., 2004); neither of which are seen in the
Chuanlinggou acritarchs. Valeria , however, has parallel ridges uniformly spaced one micron apart on the internal surface, similar to the striations seen in some of the ovoidal
Chuanlinggou acritarchs. Javaux and colleagues (Javaux et al., 2001; Javaux et al., 2004) first reported Valeria lophostriata having roll-ups of striated half vesicle in the ca. 1500
Ma Roper Group and in the ca. 1,650 Ma Mallapunyah Formation, northern Australia.
This roll-up feature is commonly observed among our Chuanlinggou ovoidal acritarchs, in which the rolling axis is parallel to the longer axis of the vesicles and the rolling probably has to do with the longitudinal rupture pattern. Finally,
Shuiyousphaeridium macroreticulatum is a spheroidal vesicle with a diameter ranging from 50 to 300 mm, having a reticulated surface and numerous regularly spaced cylindrical hollow processes that flare outward. SEM examination shows that the vesicle’s outer surface is covered with ridges that delimit polygonal fields ca. 2 mm across. Inner wall surfaces show closely packed and beveled hexagonal plates (Javaux et al., 2004). TEM images show that the ca. 1.5 mm-thick wall is multilayered, with the electron-dense, homogeneous layer of organic plates lying above a thin electrontenuous
30 layer (Javaux et al., 2004). Observations suggest it is morphologically much more complex than the Chuanlinggou acritarchs.
In terms of microfossil assemblages, Chuanlinggou acritarchs are much lower in diversity than Mesoproterozic acritarchs according to an empirical morphospace constructed for acritarch evolution from Paleoproterozic to Cambrian (Huntley et al.,
2006).
2.4.3 Is the Chuanlinggou Ovoidal Group the Earliest Eukaryote Body Fossil?
The origin of eukaryotic organisms was one of the most important evolutionary events in the history of life. It was suggested, by comparisons of subunit rRNA genes
(Pace, 1997) and whole genomes (House and Fitz-Gibbon, 2002), that eukaryotes diverged from the other two domains early in Earth history. Eukaryotes are distinctive from Bacteria and Archaea in possessing linear chromosomes, histones, mitochondria, dictyosomes and membrane-bound organelles, which are impossible to be preserved in the fossil record. Most eukaryotic cells are larger than 10-20 µm while prokaryotic cells are much smaller. It is unreliable, however, to assign Precambrian microfossils to eukaryotes using size as the only criterion due to the size overlap between prokaryotes and eukaryotes (Schopf, 1977). For example, some sulfur bacteria can be as large as 750
µm in diameter (Walterbury and Stanier, 1978; Moreira and Lopez-Garcia, 2002).
Although geochemical techniques now permit microchemical analysis of individual microfossils, there is little comparative information of biopolymers produced by extant organisms (Marshall et al., 2005; Marshall et al., 2006; Javaux and Marshall, 2007).
Fossil morphological features, e.g. shape, surface ornamentation, wall structure and ultrastructure, and excystment structures observed with the combination of optical microscopy, SEM, and TEM are still the best venue for determining biological affinity of
31 Precambrian microfossils, which has been discussed extensively by Javaux et al. (2003).
At present, most would not deny that microfossils from the Roper Group, northern
Australia ( ca. 1,500 Ma), such as Tappania sp., Satka favosa , and Valeria lophostriata
(Javaux et al., 2001; Javaux et al., 2004) are the oldest convincing eukaryotic body fossils.
The Chuanlinggou ovoidal acritarchs display two key features which are consistent with a eukaryote affinity. First, the Chuanlinggou acritarchs possess micron-scale regular striations on their walls. Micron-scale patterns of lineations, fields, and spines are characteristic of eukaryotic microfossils and are not known among prokaryotes. Some prokaryotic cells possess nano-scale concentric rings of juxtaposed filaments, fibrils, pili, or fimbriae (Boone and Castenholz, 2001). But these proteinaceous surface ornamentations are easily removed by degradation during diagenesis or by laboratory chemical treatment (Javaux et al., 2003). More importantly, the surface parallel striation pattern seen by SEM matches the banded ultrastructure observed by TEM in both dimension and direction, implying a common structural origin of the two patterns.
Second, the Chuanlinggou ovoidal acritarchs display regular longitudinal rupture, similar to those commonly seen among Phanerozoic or the extant green algae, e.g., spores of the extant green alga Spirogyra (Van Geel and Grenfell, 1996). The envelopes of prokaryotes (e.g. cyanobacteria) can sometimes split with uniform and smooth split margins (Walterbury and Stanier, 1978), but cyanobacteria neither separate clearly into two halves nor roll up in a regular pattern. This is because envelopes of prokaryotes are made of homogeneous peptidoglycans while the envelopes of most eukaryotes are constructed of carbohydrate polymer with construction of cytoskeleton and endomembrane system (Hoiczyk, 2000). It is also noted that, unlike the complete separation of the half vesicles among Chuanlinggou ovoidal acritarchs, the two half
32 Table 2-1. Comparing morphological features between Mesoperterozoic acritarchs from Roper Group (Javaux et al., 2001) 1 or Ruyang Group (Yin et al., 2005) 2 and Paleoproterozoic Chuanlinggou acritarchs from this study.
Groups Mesoproterozoic Paleoproterozoic
Characters acritarchs Chuanlinggou acritarchs
Striation Valeria lophostriata 1, 2 , The ovoidal group
Spiromorpha sp 2.
Shuiyousphaeridium
Ornamentation Plates macroreticulatum 2,
Satka favosa 1, None
Dictyosphaera
delicate 1, 2
Tappania sp. 1, 2 ,
Processes Shuiyousphaeridium None
macroreticulatum
Longitudinal rupture Valeria lophostriata The ovoidal group
TLS wall ultrastructure Leiosphaeridia crassa 1 The ovoidal group and
spheroidal group
33
Figure 2-8. (A) (1) A schematic of the fold pattern of an ovoidal acritarch; (2) A schematic of the arc lengths along the longer axis (a) and the shorter axis (b), estimated under optical microscope where the fold pattern is more apparent than under SEM. (B) The distribution of the aspect ratio (a/b) for 33 measured ovoidal fossils. The values are well clustered at 1.47 with a σ value of 0.21, despite the errors associated with interpreting a three-dimensional body using a largely two-dimensional image.
34 vesicles in the younger green algae remain solidly attached. It is speculated that
Chuanlinggou acritarchs display a primitive longitudinal rupture pattern that demanded less polarity control than do the recent green algae.
To achieve the demonstrated level of morphological sophistication, i.e., bipolar olive-shape body, longitudinal rupture, and regularly striated wall structures , the ovoidal
Chuanlinggou acritarchs probably had to possess a cytoskeleton system and endomembrane together with supporting genetic materials for transfer and construction, as was argued by many including Cavalier-Smith (2002) and Menzel (Menzel, 1996).
These are regarded as fundamental features that distinguish eukaryotes from prokaryotes
(Schmidt and Hall, 1998; Jansen, 1999). As Cavalier-Smith (2002) pointed out, “The most fundamental events in converting a bacterium into a eukaryote were not generalized changes in the genome or modes of gene expression, but two profound cellular changes:
(i) a radical change in membrane topology associated with the origin of coated vesicle budding and fusion and nuclear pore complexes and (ii) a changeover from a relatively passive exoskeleton (the microbial cell wall) to an endoskeleton of microtubules and microfilaments associated with the molecular motors dynein, kinesin and myosin.” Some bacteria do have cytoskeletal elements (Moller-Jensen and Lowe, 2005), which, however, was probably later transferred from eukaryotes (Schlieper et al., 2005).
Supporting evidence for a eukaryotic affinity also comes from the typical TLS
Chuanlinggou acritarchs having in their wall ultrastructure. Javaux et al. (2004) reported partial preservation of TLS structure in Leiosphaeridia crassa from Roper Group. Yan
(1985) compared the ovoidal Chuanlinggou acritarchs to modern dinophysales on the basis of morphology. Our Chuanlinggou TLS feature, however, is similar to those of the
35
Figure 2-9. TEM photomicrograph of extant algae ultrathin sections. (A) TEM photomicrograph of extant green algae ( Chlorella zofingiensis ). (B) TEM photomicrograph of wall ultrastructure of extant green algae ( Chlorella zofingiensis ) showing the trilaminar structure (TLS) in which a thick inner electron- tenuous layer (b in C) is sandwiched between two thin electron-dense surface layers (a and a’ in C) (arrows). (C) The diagram of wall ultrastructure. (D) TEM photomicrograph of extant dinoflagellate algae (Alexandrium tamarense ). (E) TEM photomicrograph of wall ultrastructure of extant dinoflagellate algae (Alexandrium tamarense ) showing the electron-tenuous homogeneous wall.
36 extant green alga Chlorella zofingiensis (Figure 2-9A-C) rather than the electron-tenuous.
However, it is rather tentative to claim similarity between Chuanlinggou acritarchs and green algae based on some similarity in their trilaminar wall ultrastructure. It has been difficult to prepare ultra-thin cross-sections of these small compressed fossils and the number of studies on wall ultrastructure in acritarchs remains small. Often, well- preserved microfossils are required (Talyzina and Moczydlowska, 2000). Since the 1960s, wall ultrastructure of acritarchs has been described carefully in only a few taxa (e.g.,
Tasmanites , Acanthodiacrodium , Chuaria, Leiosphaeridia , Tappania plana,
Shuiyousphaerium macroreticulatum, Satka favosa, and Valeria lophostriata et al.) (Jux,
1968; Kjellström, 1968; Loeblich, 1970; Jux, 1971; Martin and Kjellström, 1973; Jux,
1977; Peat, 1981; Colbath and Grenfell, 1995; Amard, 1997; Arouri et al., 1999; Arouri et al., 2000; Talyzina, 2000; Talyzina and Moczydlowska, 2000; Javaux et al., 2004).
Because of a lack of experimental studies, the TLS could also be either the result of diagenesis or differential biodegradation that may or may not mimic the original wall structure.
The millimeter-wide and centimeter-long coiled filaments, Grypania , dating back as early as Paleoproterozoic ( ca. 1,874 Ma), were regarded as possibly the earliest eukaryotes (Han and Runnegar, 1992; Schneider et al., 2002). Their eukaryotic affinities are debated, however, largely because of the lack of morphological details (Samuelsson and Butterfield, 2001). Additionally, the megascopic carbonaceous compressions found in the underlying Changzhougou Formation ( ca. 1,800 Ma) were recently concluded to be pseudo-fossils rather than body fossils (Lamb et al., 2007).
2.5 Conclusions
37 Combining observations from optical microscopy, SEM, and TEM, this study provides new lines of morphological evidence for eukaryotic affinity of an ovoidal group of the ca. 1,700 Ma Chuanlinggou acritarchs from North China. Based on the proposed morphological model, most of the ovoidal fossil vesicles are compressed roll-ups of the half vesicles. Although the Chuanlinggou acritarchs are lower in morphological diversity than known Mesoproterozoic acritarchs (e.g. ca. 1,500 Ma Roper Group), they do express, bipolar morphology, striations on the inner surface of vesicles, and bands in wall ultrastructure, and the regular longitudinal rupture pattern, which suggest they probably represent an earlier, albeit simpler group of eukaryote body fossils in the
Paleoproterozoic.
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40 Menzel, D., 1996, The role of the cytoskeleton in polarity and morphogenesis of algal cells: Current Opinion in Cell Biology, v. 8, p. 38-42.
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41 Talyzina, N.M., and Moczydlowska, M., 2000, Morphological and ultrastructural studies of some acritarchs from the Lower Cambrian Lukati Formation, Estonia: Review of Palaeobotany and Palynology, v. 112, p. 1-21.
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42 CHAPTER 3
NEOPROTEROZOIC 17 O-DEPLETED BARITE FROM TWO MARINOAN CAP DOLOSTONE SECTIONS, SOUTH CHINA
3.1. Introduction
The Neoproterozoic Era (1000-542 Ma) is of particular interest to this dissertation because of the occurrence of multiple ice ages (Kaufman et al., 1997; Hoffman et al.,
1998). The Marinoan glaciation ( ca. 635 Ma) is probably the most widely accepted glacial to have occurred on a global scale during this time period (Hoffman et al., 1998).
The Marinoan glaciation was suggested to have triggered another dramatic rise of atmospheric O 2 concentration and the subsequent appearance of macroscopic metazoans
(Hoffman and Schrag, 2002; Knoll, 2003; Yin et al., 2007).
The immediate aftermath for the dynamic Earth system — the interaction among the hydrosphere, atmosphere, and biosphere — is recorded in the carbonates on top of glacial diamictites. Most of the carbonates are dolostones, namely “cap dolostones.”
Recently, Bao et al. (2008) reported a distinctive, non-mass-dependent 17 O depletion in a sedimentary mineral — barite — from the Neoproterozoic cap carbonates in South China and West Africa. It was further proposed that the anomalous 17 O signal was transferred to
17 sulfate from tropospheric O 2 through oxidative weathering, and the highly O-depleted atmospheric O 2 was most likely the result of an extremely high-pCO 2 atmosphere in the immediate aftermath of the Marinoan glaciation (Bao et al., 2008). Bao et al. (2008) reveal a large variability in the magnitude of the 17 O depletion in barite samples from
South China, even for those from a single-hand specimen. Thus, a spatial heterogeneity is suspected, and microanalysis in a clear, petrographic context is warranted for the barite.
43 In addition to South China, barite was also found deposited on the top of the cap dolostones in the Taoudéni Basin in northwest Africa (Shields et al., 2007) and the
Mackenzie Mountains in northwestern Canada (James et al., 2001; Hoffman and Schrag,
2002). The occurrence of barite associated with cap carbonates may be a global phenomenon. Shields et al. (2007) advocate a sedimentary exhalative origin for the barite in Taoudéni Basin, associated with methane seepage after glaciation. Their model was based on the wide range of δ34 S and contemporaneous seawater 87 Sr/ 86 Sr ratio for their barite samples. The barite of South China, however, was deposited at the onset of a transgression after the initially deposited cap dolostones had been exposed and disrupted.
It was clearly deposited earlier than the calcite bearing a methane hydrate signature (Zhou et al., 2010). Therefore, understanding the origins of barite in South China’s cap dolostones can help to determine if the barite deposition, as well as the proposed sequence of events in the aftermath of Marinoan glaciation (Zhou et al., 2010), is of global significance.
A good model for the origin of barite should be able to (1) explain the heterogeneity of barite’s sulfur and triple-oxygen isotope composition; (2) identify the source of sulfate and barium; (3) explain its peculiar position in the sequence of events; and (4) offer clues to the local depositional environment, as well as global ocean chemistry in the aftermath of the Marinoan glaciation. To achieve this objective, we selected two sites in South China — Songlin in the Guizhou Province and Baizhu in the
Hubei Province (Figure 3-1) — where barite is well exposed in Marinoan cap dolostones.
After examining polished slabs and thin-sections of these baritic dolostones, we used a high-resolution sampling approach by which barite can be analyzed with spatial resolution at millimeters for its ∆17 O, δ18 O, and δ34 S. This means that multiple samples
44 can be obtained from individual barite crystal fans within the same barite layer or among different layers. Since 87 Sr/ 86 Sr ratios and δ34 S have been used to distinguish the various types of the barite deposits (Torres et al., 2003), 87 Sr/ 86 Sr ratios of selected barite samples were measured as additional evidence of the barium source and the depositional environment. A simple box model is used to explore the relationship between ∆17 O, δ18 O, and δ34 S of seawater sulfate enforced by riverine input sulfate after the melting of
Marinoan glaciers.
3.2. Sedimentological and Petrographic Contexts of the 17 O-depleted Barite
Sedimentary barite occurs in several horizons throughout the Neoproterozoic
Doushantuo Formation in South China. The barite studied in this paper was exclusively associated with the Marinoan cap dolostone at the base of the Doushantuo Formation in the Songlin (27°43' N, 106°39'E) and Baizhu (31°42' N, 110°56'E) areas (Figure 3-1).
Barite, occasionally silicified, occurs in two different growth patterns — (1) B1: larger, dendritic crystal fans within a dolomite matrix, with the height of individual fans ranging from 2 to 5 cm (Figure 3-2A&B, Figure 3-3C, and Figure 3-4A); and (2) B2: smaller, denser clusters of crystal fans, with the height of individual fans ranging from 0.1 cm to 2 cm (Figure 3-2C-F, Figure 3-3B, and Figure 3-4B-D). Field observations show that there are many layers of barite growth. B1 appears to precipitate consistently earlier than B2 in both the Songlin and Baizhu sections . In Songlin, the 2- to 3-m thick cap dolostones consist mainly of pink dolostone, capping the grey diamictite of the Nantuo Formation.
They are overlaid by purple, muddy dolostones and black shale of the Doushantuo
Formation. Here, the well-exposed outcrops show B1 barite blading in the dolostone matrix (Figure 3-2A&B) and B2 barite coating the surfaces of the disrupted cap dolostones (Figure 3-2C-F). In Baizhu, B2 barite occurs about 0.5 m above the B1 barite
45 Figure 3-1. Simplified stratigraphic column and paleogeography of the Yangtze Block during late Neoproterozoic to Cambrian times illustrating the location of the Songlin and Baizhu Sections (modified after Zhou et al., 2003).
46
Figure 3-2. Outcrop photos of barite in cap dolostones at Songlin. A and B: larger, dendritic crystal fans within the dolomite matrix (B1barite); C to F: smaller, denser crystal fans (B2 barite), often as encrustation on the surface of disrupted boulders (C) or walls of fractures of the cap dolostones (D) (Scale =1cm); E: a close-up view of B2 barite crystals on a weathered surface; F: multiple layer of B2 barite within weathered dolostones.
47
Figure 3-3. Outcrop photos of cap dolostones and barite at Baizhu. A: the position of two barite layers, marked by the position of hammer or hand in the cap dolostones; B: close- up view of B2 barite; C: close-up view of B1 barite.
48
Figure 3-4. Photographs of polished slabs or thin sections of barite layers. A: B1 barite crystal fans from Baizhu, growing on a cap dolostone clast; B: B2 barite crystal fans from Songlin, orange matrices are dolomite mixed with clay minerals; C: densely packed B2 barite crystal needles in dolomite matrix; D: densely packed B2 barite crystal fans.
49 layer, sandwiching a dolostone with wavy, laminated structures (Figure 3-3A). The
Songlin and Baizhu cap dolostones are believed to have been deposited in a shallow marine platform environment (Zhu et al., 2003). The barite was deposited at the onset of a regional transgression after the basal dolostones were deposited, exposed, and disrupted
(by dissolution), as outlined in Zhou et al. (2010).
3.3. Methods
3.3.1 Isotope Analytical Methods
We analyzed 160 micro-sampled barite samples from 20 hand specimens collected from several sections in Songlin and Baizhu. Hand specimens were first cut and polished, and thin-sections were cut parallel to the barite crystal grow direction. Different barite crystal fans or layers were identified before being micro-sampled by a dental drill. Each drilling hole sample is about 2 mm deep and 5 mm in diameter. A 2 M HCl solution was then added to about 100 mg powders to dissolve and remove any carbonates. The pretreatment procedure and triple oxygen isotope analysis for barite can be found in Bao et al. (2008). In brief, the acid-treated residues were further treated with a DDARP method (Bao, 2006) and a 2 M NaOH solution (Bao et al., 2007) to precipitate pure barite
(BaSO 4). Sample pretreatment and all oxygen-isotope measurements were carried out at
17 LSU. The O of barite was determined using O 2 generated from BaSO 4 in a CO 2 laser- fluorination system (Bao et al., 2000) and run on a MAT253 in a duel inlet model. The standard deviation associated with the 17 O is ± 0.03‰ based on multiple (N = 3) runs of the same O 2 gas on the MAT 253, and is ±0.05‰ for replicates of the same BaSO 4 via
18 laser-fluorination. The δ O of barite was measured using CO gas converted from BaSO 4 using a Thermal Conversion Elemental Analyzer (TCEA) at 1450°C, coupled with the
18 MAT253 in a continuous-flow mode. The standard deviation associated with δ OSO4 is
50 * 18 16 17 16 ±0.5‰. Here the δ = (Rsample /Rstandard − 1), in which R is the ratio of O/ O, O/ O, or
34 32 17 17 18 S/ S, and O ≡ δ' O – 0.52× δ' O, in which the δ'≡ ln (Rsample /RVSMOW ).
The δ34 S measurement of sulfate was conducted at the University of Maryland and at Colorado Plateau Stable Isotope Laboratory (CPSIL), where BaSO 4 was converted to
SO 2 using an Elemental Analyzer (EA) at 1050°C and run on a Micromass Isoprime
(Maryland) or a Thermo-Electron Delta Plus Advantage CPSIL mass spectrometer in a continuous-flow mode. The standard deviation associated with δ34 S measurement is
±0.3‰.
The δ34 S measurement of sulfate was conducted at the University of Maryland and at Colorado Plateau Stable Isotope Laboratory (CPSIL) where BaSO 4 was converted to
SO 2 using an Elemental Analyzer (EA) at 1050°C and run on a Micromass Isoprime
(Maryland) or a Thermo-Electron Delta Plus Advantage (CPSIL) mass spectrometer in a continuous-flow mode. The standard deviation associated with δ34 S measurement is
±0.3‰.
The 87 Sr/ 86 Sr ratios were measured for 10 barite samples drilled from hand specimens, following a standard procedure (e.g. Paytan et al. 1993) at the Chinese
Academy of Sciences (barite powders were pre-leached in 1M acetic acid at room temperature for 24 hours to remove carbonate minerals). Ratios of 87 Sr/ 86 Sr were determined on a Thermo-Fisher Triton mass spectrometry. External precision is
±0.000007 (2 standard deviations). During the period of analyzing these barite samples, the mean 87 Sr/ 86 Sr ratio of the standard NBS 987 was 0.710258.
3.3.2 Modeling Methods
The modeling method and values are revised from Turchyn and Schrag (2006).
Given that evaporite deposits in Precambrian rocks are rare (Grotzinger and Kasting,
51 1993; Zharkov, 2005; Warren, 2010), we assume all input of sulfate is from pyrite
34 17 weathering with a δ SSO4-rivers = 0‰ and ∆ O SO4-rivers = −4.2‰ (Bao et al., 2009). If we assume atmospheric O 2 concentration is 20% of the present (Holland, 2006), then the rate of pyrite oxidation is about half of the present (Williamson and Rimstidt, 1994). Thus, the
11 flux of sulfate from rivers (Frivers ) is set at about 3.23×10 mole/year. We set marine sulfate concentration at 635 Ma at 0.28 mM, about 1% of the modern value of 28 mM.
Microbial sulfate reduction rate is set at 8×10 11 mole/yr, about 10% of the present rate
(Canfield et al., 2000; Bolliger et al., 2001). Based on one set of data of barite (27‰) and associated pyrite (9‰) from Songlin and the model of Habicht et al. (2002), the sulfur isotope fractionation ( εS) between sulfate and pyrite is set at 18‰.
To model the sulfur cycle we use the equations for sulfate:
= + × − (1) in which
MSO4 amount of sulfate in the ocean,
flux of sulfate from rivers,
rate of sulfate reduction,
fraction of sulfide produced from sulfate reduction that is subsequently
reoxidized to form sulfate in water column;
18 For δ OSO4 , we have
d δ O M dt
= F δ O + f × F δ − F δ O − ε (2)
oxygen isotopic composition of marine sulfate, δ O
52 Table 3-1. Values used in the dynamic model.
Description Term Value Unit 17 Mass of sulfate in the ocean MSO4 3.78 ×10 mole 11 Flux of sulfate from rivers Frivers 3.23 ×10 mole/y ear 11 Sulfate reduction rate Freduction 8×10 mole/y ear Fraction of sulfide produced from sulfate f 83 % reduction which is subsequently re-oxidized 18 Oxygen isotopic composition of marine sulfate δ OSO4 -maine 13 ‰ 18 Oxygen isotopic composition of riverine sulfate δ OSO4 -rivers 0 ‰ Oxygen isotopic composition for sulfate δx 15 ‰ produced by sulfide re-oxidation associated with sulfate reduction
Oxygen isotopic fractionation associated with a εo 7 ‰ preferential selection for 16 O-enriched sulfate during sulfate reduction 34 Sulfur isotopic composition of sulfate in the δ SSO4-marine 20 ‰ oceans 34 Sulfur isotopic composition of sulfate in rivers δ SSO4 -rivers 0 ‰
Sulfur isotope fractionation between sulfate and εS 18 ‰ sulfide during sulfate reduction 17 17 ∆ O sulfate of sulfate in the oceans ∆ O SO4-marine −1 ‰ 17 17 ∆ O sulfate of sulfate in rivers ∆ O SO4-rivers −4.2 to −1 ‰ 17 17 ∆ O sulfate from re -oxidation of sulfides ∆ OSO4 -reoxidation 0 ‰ 17 17 ∆ O fractionation associated with sulfate ∆ OSO4-reduction 0 ‰ reduction
53 Table 3-2. Parameters and constraints for 28 controlled model runs. For R1-1 to R1-6, we 16 17 increase MSO4 from 3.78×10 moles to 7.53×10 Moles while other variables are fixed. 9 11 For R2-1 to R2-4, we increase F rivers from 3.23×10 Mole/ year to 3.23×10 moles while other variables are fixed. R3-1 to R3-4 differs from R2-1 to R2-4 only in initial MSO4 . For 16 17 R4-1 to R4-6, we increase MSO4 from 3.78×10 moles to 7.53×10 Moles with other variables fixed but different from R1 to R3. For R5-1 to R5-4 and R6-1 to R6-4, we 17 decrease the ∆ O SO4-river from −1.0 ‰ to −4.0 ‰ in which the initial MSO4 is different for R5 and R6.
Conditions
Runs Initial values Frivers 17 17 18 34 MSO4 ∆ O SO4-river ∆ O SO4-marine δ OSO4-marine δ SSO4-marine (mole/year) (Mole) (‰) (‰) (‰) (‰) R1 -1 3.78×10 16 -4.2 -1 13 20 0 R1-2 7.56 ×10 16 -4.2 -1 13 20 0 R1-3 1.74 ×10 17 -4.2 -1 13 20 0 R1-4 3. 02 ×10 17 -4.2 -1 13 20 0 R1-5 3. 78 ×10 17 -4.2 -1 13 20 0 R1-6 7.53 ×10 17 -4.2 -1 13 20 0 R2-1 3.78×10 16 -4.2 -1 13 20 3.23×10 9 R2-2 3.78×10 16 -4.2 -1 13 20 1.0 ×10 10 R2-3 3.78×10 16 -4.2 -1 13 20 2×10 10 R2-4 3.78×10 16 -4.2 -1 13 20 3.23×10 11 R3-1 3. 78 ×10 17 -4.2 -1 13 20 3.23×10 9 R3-2 3. 78 ×10 17 -4.2 -1 13 20 1.0 ×10 10 R3-3 3. 78 ×10 17 -4.2 -1 13 20 2×10 10 R3-4 3. 78 ×10 17 -4.2 -1 13 20 3.23×10 11 R4-1 3.78×10 16 -4.2 -0.1 23 46 3.23×10 11 R4-2 7.56 ×10 16 -4.2 -0.1 23 46 3.23×10 11 R4-3 1.74 ×10 17 -4.2 -0.1 23 46 3.23×10 11 R4-4 3. 02 ×10 17 -4.2 -0.1 23 46 3.23×10 11 R4-5 3. 78 ×10 17 -4.2 -0.1 23 46 3.23×10 11 R4-6 7.53 ×10 17 -4.2 -0.1 23 46 3.23×10 11 R5-1 3.78×10 16 -1.0 -1 13 20 3.23×10 11 R5-2 3.78×10 16 -2.0 -1 13 20 3.23×10 11 R5-3 3.78×10 16 -3.0 -1 13 20 3.23×10 11 R5-4 3.78×10 16 -4.0 -1 13 20 3.23×10 11 R6-1 3. 78 ×10 17 -1.0 -1 13 20 3.23×10 11 R6-2 3. 78 ×10 17 -2.0 -1 13 20 3.23×10 11 R6-3 3. 78 ×10 17 -3.0 -1 13 20 3.23×10 11 R6-4 3. 78 ×10 17 -4.0 -1 13 20 3.23×10 11
54 oxygen isotopic composition of riverine sulfate, δ O
oxygen isotopic composition for sulfate produced by sulfide re-
oxidation associated with sulfate reduction,
oxygen isotopic fractionation associated with a preferential selection
for 16 O-enriched sulfate during sulfate reduction;
For δ34 S, we have
= δ S − ( − × ) ×
(3) (δ S − )
sulfur isotopic composition of sulfate in the oceans, δ S
sulfur isotopic composition of sulfate in rivers, δ S
sulfur isotope fractionation between sulfate and sulfide during sulfate
reduction;
For ∆17 O, we have
∆ =
∗ ∆ O + × × ∆ O − ×
(4) ∆ O
∆17 O of sulfate in the oceans, ∆ O
∆17 O of sulfate in rivers, ∆ O
∆17 O of sulfate from re-oxidation of sulfides; this term also ∆ O
2- includes sulfate being introduced to the seawater via SO 3
APS back exchange within cells, f* the fraction of sulfide being re-oxidized to sulfate after sulfate
reduction, plus the fraction of sulfate introduced to the 55 2- seawater via SO 3 APS back exchange within cells, all with
respect to the sulfate being actually reduced.
17 34 18 Exploring temporal changes in the ∆ O SO4-marine , δ SSO4-marine , and δ OSO4-marine involves solving equations 1, 2, 3, and 4 using a time step of 20,000 years.
3.4. Results
3.4.1 Isotope Results
There is a large variability in the ∆17 O, δ18 O, and δ34 S among different barite crystal fans or layers from Songlin and Baizhu (Figure 3-5, 6, and 7 and Appendix B). In Songlin,
∆17 O, δ18 O, or δ34 S values of B1 and B2 barites are significantly different (Kolmogorov-
Smirnov test [KS -test], P <0.05, using an α value 0.05, N=50) (Figure 3-5A, Figure 3-6A, and Figure 3-7A). B1 barite generally has a higher ∆17 O, δ18 O, or δ34 S value than B2 barite (Table 3-3). There is a hyperbolic relationship between the ∆17 O and the δ34 S
(Figure 3-5A). In Baizhu, ∆17 O, δ18 O, or δ34 S values of B1 and B2 barites are also significantly different (KS-test, P <0.05, using an α value 0.05, N=37) (Figure 3-5B,
Figure 3-6B, and Figure 3-7B). Differing from Songlin barite, Baizhu’s B1barite has more negative ∆17 O values, higher δ18 O, but lower δ34 S values (Table 3-3).A hyperbolic relationship between ∆17 O and δ34 S is also found in Baizhu barite (Figure 3-5B).
We sampled 16 single barite crystal fans, each fan possessing well-clustered ∆17 O,
δ18 O, or δ34 S values. Each barite crystal fan had 3 to 11 micro-drilled samples, in which a standard deviation for the ∆17 O ranges from 0.02 to 0.09‰ (mean of standard deviation =
0.04‰), for δ18 O from 0.3 to 1.2‰ (mean of standard deviation = 0.7‰), and s for δ34 S from 0.2 to 1.9‰ (mean of standard deviation = 0.9‰) (Figure 3-8, 9, 10, and 11). In addition, individual barite crystal fans of the same layer bear similar ∆17 O and δ18 O values but not necessarily similar δ34 S (Figure 3-8B, C, D, and F and Figure 3-9, 10, and
56 11). Also, barite crystals of different layers have different δ18 O, δ34 S, or ∆17 O clusters
(Figure 3-8 A&E and Figure 3-9, 10, and 11). The 87 Sr/ 86 Sr ratios for barite from Songlin range from 0.71071 to 0.71457 (mean = 0.71256; N =10) (Appendix B).
3.4.2 Modeling Results
In run R1, when microbial sulfate reduction and sulfide reoxidation are active in the
17 18 marine system while the riverine sulfate influx is zero, the ∆ O SO4-marine and δ OSO4-marine
34 quickly increase to a steady-state value of 0‰ and +23.4 ‰, respectively, while δ SSO4- marine keeps increasing (Figure 3-12 A) over time. When MSO4 increases from 0.1% to 2%
17 18 of the present level in R1-1 to R1-6, the ∆ O SO4-marine and δ OSO4-marine take a longer time
34 to reach a steady state. It takes about 2.7 million years for δ SSO4-marine to increase from
20‰ to 46‰ when the marine sulfate reservoir is 3.78×10 17 Mole (i.e., 1% of the present).
When the marine sulfate reservoir was smaller, it would take a shorter time period to reach the same δ34 S value. The ∆17 O-δ34 S and δ18 O-δ34 S spaces follow a hyperbola, while the ∆17 O-δ18 O follows a positive linear correlation (Figure 3-13).
18 In runs R2 and R3, when the flux of riverine sulfate changes, the δ OSO4-marine and
17 ∆ O SO4-marine quickly get to steady state at values between 16‰ and 23.4 ‰, and −1.36‰ and 0‰, respectively. The steady-state values are controlled by Frivers and are independent of marine sulfate reservoirs (Figure 3-12 B). When the flux of riverine sulfate is larger
11 17 than about 2×10 Mole/year (i.e., 31% of the modern level), the ∆ O SO4-marine decreases to different steady-state values. The ∆17 O-δ34 S and the δ18 O-δ34 S spaces follow different hyperbolas, while the ∆17 O-δ18 O follows a negative linear correlation (Figure 3-13).
Conversely, when the flux of riverine sulfate is smaller than about 2×10 11 Mole/year, the
17 17 34 18 34 ∆ O SO4-marine increases to different steady-state values. The ∆ O-δ S, δ O-δ S, and
∆17 O-δ18 O spaces all follow hyperbolas (Figure 3-13). When the flux of riverine sulfate is
57 11 34 larger than about 1×10 Mole/year (i.e., 62% of the modern level), the δ SSO4-marine keeps decreasing over time. When the riverine flux is smaller than about 1×10 11 Mole/year, the
34 δ SSO4-marine keeps increasing.
In run R4, only MSO4 changes, and the flux of riverine sulfate is fixed. Initial values of the isotope composition of marine sulfate are different from other runs, however. The
17 18 ∆ OSO4-marine and δ OSO4-marine quickly decrease to a steady-state value of −1.36‰ and
34 16‰, respectively, while δ SSO4-marine slowly decreases to 9‰ (Figure 3-12 C). When
17 MSO4 increases from 0.1% to 2% of the modern level in runs R4-1 to R4-6, the ∆ O SO4-
18 marine and δ OSO4-marine take a longer time to reach a steady state. It takes about 2.7 million
34 years for the δ SSO4-marine to decrease from 46‰ to 20‰ when the marine sulfate reservoir is 3.78×10 17 Mole (i.e., 1% of the present). The ∆17 O-δ34 S and δ18 O-δ34 S spaces follow hyperbolas, while the ∆17 O-δ18 O follows a positive linear correlation (Figure 3-13).
17 In runs R5 and R6, when the value of ∆ O SO4-rivers varies while the flux of riverine
18 17 sulfate is fixed, the δ OSO4-marine and ∆ O SO4-marine quickly reach a steady-state value of
16‰ and between −1.36‰ and −0.3‰, respectively (Figure 3-12 D). The values are
17 controlled by the ∆ O SO4-rivers and are independent of marine sulfate reservoir size
17 17 (Figure 3-12 D). When the ∆ O SO4-rivers is more negative than −3‰, the ∆ O SO4-marine decreases to different steady-state values. The ∆17 O-δ34 S and δ18 O-δ34 S spaces follow hyperbolas, while the ∆17 O-δ18 O follows a negative linear correlation (Figure 3-13).
17 17 Conversely, when the ∆ O SO4-rivers is less negative than −3‰, the ∆ O SO4-marine increases to different steady values. The ∆17 O-δ34 S and δ18 O-δ34 S spaces follow hyperbolas, while the ∆17 O-δ18 O follows a positive linear correlation (Figure 3-13).
58 A 50.0
B1 -Songlin 45.0
B2 -Songlin 40.0
35.0 S 34 δ
30.0
25.0
20.0 -1.00 -0.90 -0.80 -0.70 -0.60 -0.50 -0.40 -0.30 -0.20 -0.10 0.00 ∆17 O
B 50.0 B1 -Baizhu B2 -Baizhu 45.0
40.0 S
35.0 34 δ
30.0
25.0
20.0 -1.00 -0.90 -0.80 -0.70 -0.60 -0.50 -0.40 -0.30 -0.20 -0.10 0.00 17 ∆ O
Figure 3-5. Relationships between the 17 O and δ34 S for all B1 and B2 barite samples from Songlin (A) and from Baizhu (B). The purple curves and green lines show the patterns displayed by the B1 and B2 barite, respectively.
59
A 50.0 B1 -Songlin B2 -Songlin 45.0
40.0 S
34 35.0 δ
30.0
25.0
20.0 12.0 14.0 16.0 18.0 20.0 22.0 24.0 δ18 O
50.0 B B1 -Baizhu B2 -Baizhu 45.0
40.0 S
34 35.0 δ
30.0
25.0
20.0 12.0 14.0 16.0 18.0 20.0 22.0 24.0 δ18 O
Figure 3-6. R elationships between the δ 18 O and δ34 S for all B1 and B2 barite samples from Songlin (A) and from Baizhu (B). The purple curves or line and green lines show the patterns displayed by the B1 and B2 barite, respectively.
60 δ18 O A 12.0 14.0 16.0 18.0 20.0 22.0 24.0 0.00
-0.10
-0.20
-0.30
-0.40 O
17 -0.50 ∆ -0.60
-0.70
-0.80 B1 -Songlin B2 -Songlin -0.90
-1.00 δ18 O B 12.0 14.0 16.0 18.0 20.0 22.0 24.0 0.00
-0.10
-0.20
-0.30
-0.40 O -0.50 17 ∆ -0.60
-0.70
-0.80 B1 -Baizhu -0.90 B2 -Baizhu -1.00
Figure 3-7. R elationships between the δ 18 O and 17 O for all B1 and B2 barite samples from Songlin (A) and from Baizhu (B). The purple and green lines show the patters displayed by the B1 and B2 barite, respectively.
61 Table 3-3. Ranges and mean values of the ∆17 O, δ18 O, and δ34 S of B1 and B2 barite at Songlin and Baizhu.
Sites and type of 17 18 34 ∆ O δ O δ S barite B1 −0.56‰ to −0.21‰ +14.1‰ to +20.6‰ +26.2‰ to +35.2‰ barite (mean = −0.34‰; N = (mean= +17.2‰; N (mean = +29.2‰; N Songlin 35) =35) =34) B2 −0.87 ‰ to −0.27‰ +13.3‰ to +16.9‰ +21.1‰ to +36.2‰ barite (mean = −0.59‰; N = (mean= +15.3‰; N (mean = +26.0‰; N 59) =59) =45) B1 −0.58‰ to −0.28‰ +17.6‰ to +23.5‰ +25.0‰ to +31.8‰ barite (mean = −0.40‰; N = (mean= +20.9‰; N (mean = +27.6‰; N Baizhu 43) =43) =37) B2 −0.45 ‰ to −0.09‰ +17.0‰ to +19.9‰ +22.4‰ to +45.5‰ barite (mean = −0.20‰; N = (mean= +18.7‰; N (mean = +36.6‰; N 23) =23) =16)
62
Figure 3-8. Drill-sampled slabs and their corresponding δ 18 O, 17 O, and δ34 S data. A (ZB-07-24-3), B (Jinhe-08-1-R), and C (SL-2-R) are barite from Songlin and D (ZB-07- 20), E (ZB-0714-2), and F (ZB-07-18-1-1) from Baizhu (Scale =1cm). A and E: sample #1 and #2 from the same individual barite crystal fan, other samples on the slab were from different barite layers; B and C: all samples on the slab were from the same individual barite crystal fan; D and F: samples on the slab were from the same layer of barite layer.
63 δ18 O 13.0 14.0 15.0 16.0 17.0 18.0 19.0 20.0 21.0 22.0 0.00
-0.10
-0.20
-0.30
-0.40 O
17 -0.50 ∆
-0.60
-0.70
-0.80
-0.90
-1.00
Jinhe -07 -4 Jihne -08 -1-R SL -2-R OB -4-1 ZB -07 -14 -1 ZB -07 -14 -2 ZB -07 -18 -1-1 ZB -07 -18 -1-2 ZB -07 -18 -2-1 ZB -07 -18 -2-2 ZB -07 -18 -3 ZB -07 -18 -4 ZB -07 -20 -1 ZB -07 -20 -2 ZB -07 -24 -5 ZB -07 -24 -3
Figure 3-9. Relationships between δ18 O and 17 O for crystal-fan-specific barite samples from Songlin (sample Jinhe-07-4, Jinhe-08-1-R, ZB-07-24-3, ZB-07-24-5, and SL-2-R) and Baizhu (all others). Multiple samples were drilled from a single crystal fan, except for sample Jinhe-07-4, ZB-07-24-3, ZB-07-14-1, and ZB-07-14-2, whose multiple data points were from different barite crystal fans.
64
Jinhe -08 -1-R 50.0
Sl -2-R
OB -4-1 45.0 ZB -07 -14 -2
ZB -07 -18 40.0 ZB -07 -18 -1-2
ZB -07 -18 -2-1
ZB -07 -18 -2-2 35.0 S 34 δ ZB -07 -20 -1
ZB -07 -20 -2 30.0 ZB -07 -24 -3
25.0
20.0 -1.00 -0.90 -0.80 -0.70 -0.60 -0.50 -0.40 -0.30 -0.20 -0.10 0.00 ∆17 O
Figure 3-10. Relationships between 17 O and δ34 S for crystal-fan-specific barite samples from Songlin (Jinhe-08-1-R, ZB-07-24-3, and SL-2-R) and Baizhu (all others). Multiple samples were drilled from a single crystal fan, except for sample ZB-07-14-2 and ZB-07-24-3, whose multiple data points were from different barite crystal fans.
65 50.0
45.0
40.0 S 34
δ 35.0
30.0
25.0
20.0 13.0 14.0 15.0 16.0 17.0 18.0 19.0 20.0 21.0 22.0 δ18 O
Jinhe -08 -1-R SL -2-R OB -4-1 ZB -07 -14 -2 ZB -07 -18 ZB -07 -18 -1-2 ZB -07 -18 -2-1 ZB -07 -18 -2-2 ZB -07 -20 -1 ZB -07 -20 -2 ZB -07 -24 -3
Figure 3-11. Relationships between δ18 O and δ34 S for crystal-fan-specific barite samples from Songlin (Jinhe-08-1-R, ZB-07-24-3, and SL-2-R) and Baizhu (all others). Multiple samples were drilled from a single crystal fan, except for sample ZB-07-14-2 and ZB-07-24-3, whose multiple data points were from different barite crystal fans.
66
Figure 3-12 Model results for runs R1-R6 plotted as a function of time from 0 −2.7 million years for variations in the 17 O, δ18 O, and δ34 S of marine sulfate. A: R1-1 to R1-6; B: R2-1 to R2-4 and R3-1 to R3-4; C: R4-1to R4-6; D: R5-1 to R5-4 and R6-1to R6-4. Please see Table 3 for details of model parameters.
67 A 45 R1 -1 R1 -2 R1 -3 R1 -4 40 R1 -5 R1 -6 35 R2 -1 R2 -2 R2 -3 R2 -4 30 R3 -1 R3 -2 25 R3 -3 R3 -4
S S R4 -1 R4 -2 20 34 δ R4 -3 R4 -4 15 R4 -5 R4 -6 R5 -1 R5 -2 10 R5 -3 R5 -4 5 R6 -1 R6 -2 R6 -3 R6 -4 0 -1.6 -1.4 -1.2 -1 -0.8 -0.6 -0.4 -0.2 0 Δ17 O
B 48 43 38 33 S S
34 28 δ 23 18 13 8 12 14 16 18 20 22 24
δ18 O
12 14 16 18 20 22 24 C 0 -0.2 -0.4 -0.6 -0.8 -1
O O -1.2 17 -1.4
Figure 3-13 Model results for runs R1 -R6 plotted for variations in the 17 O, δ18 O, and δ34 S of marine sulfate. Black dots are the starting points. Please see Table 3 for details of model parameters. Purple lines , purple curves, and green circles represent the pattern in Figure 3-5, 6, and 7, respectively.
68 3.5. Discussion
3.5.1 The Effect of Diagenesis
The stable isotope composition of barite may be affected by later fluid-rock interactions. Considering that later hydrothermal diagenesis can lead to silicification of barite in cap dolostones (Jiang et al., 2006), diagenetic effects need to be evaluated. Hydrothermal or meteoric-water alteration may change the initial δ18 O and
17 O values, but it would leave the δ34 S largely intact. While the δ18 O and ∆17 O are nearly uniform within the same crystal fan (Figure 3-8B, C, D, and F and Figure 3-9, 10, and 11), they can differ by 6‰ and 0.7‰, respectively, among fans of different layers
(Figure 3-8 A&E and Figure 3-9, 10, and 11). The preservation of these large heterogeneities at the millimeter scale indicates that, even if burial diagenesis or later alteration had some impact on barite’s isotope composition, most of the original signals or at least their differences are intact. Thus, the following discussion assumes the barite bears its original oxygen isotope signatures with little diagenetic alteration.
2− 3.5.2 An Isotopically Variable Sulfate Pool Reflecting a Low [SO 4 ] Ocean
While Bristow et al. (2009) suggests the Yangtze Gorges region of the Yangtze
Block as an inland lake environment during the deposition of most of the Doushantuo
Formation in the Neoproterozoic, most believe that it was a shallow platform in transition to open ocean basins (Zhu et al., 2003; Vernhet and Reijmer, 2010). Thus, as heavily influenced by continental runoff as the Songlin and Baizhu sites might have been, the water should be more or less connected to the open oceans and influenced by seawater sulfate. The duration for the multiple barite layers should be much less than 2.7 million years based on the age of 635.2 ± 0.6 Ma at the top of cap dolostones, where the 17 O- anomalous barite resides and the 632.5 ± 0.5 Ma for sediments is immediately above the
69 cap dolostones (Condon et al., 2005) (Figure 3-1). δ34 S of the multi-layeral barite fluctuated between 20‰ and 45‰ during this period, exceeding the range we see for the entire Phanerozoic Eon (540 million years) (Strauss, 1999). The rapidly changing δ34 S in barite can only be explained if the sulfate concentration in the water at that time was so exceedingly low that its isotopic composition was easily modified by sulfate fluxes into and out of the oceans. Our modeling results suggest that the isotope composition of the multi-layered barite could fluctuate in the way depicted in Figure 3-12 within 2.7 million years only if the marine sulfate reservoir was much less than 1% of the present one in the post-Marinoan oceans. The δ34 S of sulfate extracted from top of cap dolostones from
South China (McFadden et al., 2008), Northern Namibia (Hurtgen et al., 2006), and
Australia (Gorjan et al., 2000) show a range from about 20 ‰ to 35 ‰, which is in good agreement with barite of South China. This consistence suggests that sulfate of barite at
Songlin and Baizhu well presented the isotopic composition of contemporaneous seawater sulfate.
3.5.3 Barite Snapshots of an Isotopically Variable Sulfate Pool
The ∆17 O, δ18 O, and δ34 S and relationships among them provide clues to the origin and depositional environments of the 17 O-anomalous barite. We argue here that the main reason for the observed isotopic heterogeneity is the variable isotope composition of dissolved sulfate. Multiple barite depositional events at a time after the deposition and the subsequent disruption of the cap dolostones took many snapshots of the changing sulfate signatures in the post-glacial shallow oceans (Figure 3-14).
The dendritic crystal growth pattern (crystal fans) suggests that the barite was precipitated in situ on the seafloor or on surfaces of disrupted cap-dolostone boulders or
70
Figure 3-14. Sketches showing a sequence of post-glacial depositional events on the Yangtze Blockcof South China (modified after Shields, 2005). (1) Cap dolostones deposited in warm meltwaters by deglaciation; (2) Isostatic rebound and karstification of the cap dolostone; (3) a postglacial transgression induced multiple upwelling of Ba 2+ rich deep marine water to shallow platform leading multiple layer precipitation of barite.
71 clasts from a highly supersaturated solution (saturation index > 20) (Shikazono, 1994).
This conclusion is consistent with the well-clustered δ18 O, 17 O, or δ34 S values in individual barite crystal fans or among many barite fans of the same layer. The variable sulfur and triple-oxygen isotope composition for different layers of barite indicates the isotope composition of seawater sulfate sourcing the barite precipitation was changing with time.
The isotope composition of seawater sulfate is controlled by the flux and isotopic composition of the sulfate into and out of the ocean. The primary source of seawater
34 sulfate is the oxidative weathering of S-depleted sulfides by atmospheric O 2 and the dissolution of evaporite minerals. Removal of sulfate from the oceans is controlled by microbial sulfate reduction and the formation of sedimentary pyrite. Overall, seawater sulfate is enriched in 34 S relative to sulfate from riverine input. The magnitude of the enrichment depends on many factors, including sulfate availability, redox condition, and weathering intensity, among others (Bruchert et al., 2001; Canfield, 2001b, a; Habicht and Canfield, 2001; Habicht et al., 2002; Canfield, 2004; Brunner and Bernasconi, 2005;
Brunner et al., 2005; Canfield et al., 2006; Turchyn et al., 2009; Turchyn et al., 2010). In addition to the processes that affect the δ34 S of sulfate, the oxygen isotope composition of sulfate is also controlled by that of ambient water. The cell-internal and external
- 18 reoxidation of sulfides (HS or H 2S) can result in the δ O of seawater sulfate reaching a steady state with that of ocean water, instead of increasing continuously as is the case for
δ34 S in a closed system (Turchyn et al., 2010).
Our modeling results suggest that the isotope sulfate composition is easily modified by riverine input, microbial sulfate reduction, or sulfide reoxidation when sulfate concentration in the water at that time was exceedingly low (Figure 3-12). When riverine
72 sulfate input was dominant, the isotope composition of seawater sulfate had lower δ34 S and δ18 O and more negative ∆17 O values. When activity of microbial sulfate reduction and sulfide reoxidation were dominant, isotope composition of seawater sulfate had higher δ34 S and δ18 O but less negative ∆17 O values. This would generally result in positive correlations in ∆17 O-δ34 S, δ18 O-δ34 S, and ∆17 O-δ18 O spaces (Figure 3-13), which are consistent with what we observed in our barite data (Figure 3-9, 10, and 11).
Given the paleogeographic locations of Songlin and Baizhu (Figure 3-1) (Zhu et al.,
2003), riverine input of sulfate is expected to be significant. Sulfate from the oxidative weathering of sulfides possesses lower δ34 S and δ18 O (Turchyn et al., 2006; Turchyn et al.,
2009) and more negative ∆17 O values (Bao et al., 2008) than those of seawater sulfate.
This sulfate was delivered to the oceans by rivers and mixed with pre-existing seawater sulfate that had higher δ34 S, δ18 O, and less negative ∆17 O values, if any. The generally more negative barite 17 O and lower δ18 O and δ34 S values in Songlin compared to Baizhu suggests that Songlin was closer to continental runoff than Baizhu.
3.5.4 A Depositional Model for the 17 O-depleted Barite
Barite crystal fans rarely occur in the rock record of the Proterozoic Eon. The only ones that textually resemble the basal Doushantuo barite are barite deposits of the early
Archean age (3,230–3,260 Ma) (Bao et al., 2007). While the source of sulfate for Songlin and Baizhu barite was dissolved seawater sulfate, the source of barium had to be anoxic sulfate-free deep marine water. The Sr isotope compositions of Songlin barite, however, is higher than contemporaneous seawater (0.7068-0.7073) (Halverson et al., 2007), which suggest that the mineralizing fluids was locally affected by Sr source leached from sediments and from riverine input but was not dominated by hydrothermal sources which tend to have much lower 87 Sr/ 86 Sr ratios (White and Hofmann, 1982). The transport of
73 this Ba-bearing deep marine water to shallow platform sites can be achieved through upwelling or overturning of deep marine water or simply through cold seeps. The former is known to be associated with transgressions (Schröder and Grotzinger, 2007), while the latter is common on the modern seafloor (Torres et al., 1996; Torres et al., 2002). The wide and nearly continuous spatial distribution of barite in the Songlin and Baizhu sites, as well as the lack of other metal sulfide minerals or chimney buildups in the vicinity, argues against the cold-seep scenario. The basal Doushantuo barite is certainly not the biogenic barite whose grain size is on the micron scale (Paytan et al., 1993). Barite can also form beneath the sediment-water interface, where the diffusion-convection front of sulfate-bearing water from above meets the front of anoxic barium-bearing pore-water from below (Hanor, 2000). This in-sediment type of barite often occurs as irregular or porous concretions or as dispersed microcrystals in sediments (Dia et al., 1993; Fu et al.,
1994; Torres et al., 1996) and is not what we observed in the basal Doushantuo barite.
Thus, following the melting of the Marinoan glaciation, multiple episodes of upwelling of deep sulfate-free marine water brought Ba 2+ to shallow platform environments (Figure 3-14). These conditions might fit the “plumeworld” hypothesis that the cap dolostones deposited in low salinity meltwaters by a uniquely abrupt deglaciation at low latitudes (Shields, 2005). There was no visible barite deposition during the deposition of cap dolostones either because the upwelling did not occur or the sulfate concentration in shallow water was exceedingly low. The first episode of barite deposition occurred right at the onset of a regional transgression, after the karstification and disruption of the cap dolostones. Each wave of Ba 2+ input at the onset of the transgression created a supersaturated condition for BaSO 4, which resulted in one layer or layers of barite crystal fans being precipitated on pre-existing surfaces on the seafloor.
74 The minor differences in isotopic characteristics of barite between Songlin and Baizhu can be explained by the difference in their proximity to the continents.
3.6 Conclusions
(1) The crystal-fan/layer-specific clustering of ∆17 O, δ18 O, or δ34 S and the
strong correlation between the ∆17 O and δ34 S in both studied sections in South
China cannot be explained by any type of systematic experimental error.
(2) The non-mass-dependent 17 O-depleted barite in the basal Doushantuo
Formation was precipitated in multiple episodes, concentrated at the onset of a
regional transgression after the cap dolostones were deposited, exposed,
disrupted, and re-submerged. Transgression-related flooding of the shallow
platform by Ba 2+ -bearing deep water resulted in a rapid precipitation of barite
crystal fans on the seafloor and on pre-existing, disrupted cap dolostone surfaces.
(3) Both the Songlin and Baizhu sites were close to continental runoff.
Riverine sulfate contained the 17 O-depleted component that was the product of
oxidative weathering by tropospheric O 2 in the aftermath of the Marinoan
glaciation.
(4) The large variability in sulfur and triple-O isotope composition among
different layers of barite suggests that the isotope composition of sulfate is easily
modified by riverine input, microbial sulfate reduction, or sulfide reoxidation
when sulfate concentration in the water at that time was exceedingly low.
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79 APPENDIX A
LETTERS OF PERMISSION
80 APPENDIX B
TRIPLE OXYGEN, SULFUR, AND STRONTIUM ISOTOPE DATA FOR BARITE ASSOCIATED WITH MARINOAN CAP DOLOSTONES FROM SONGLIN AND BAIZHU, SOUTH CHINA.
δ18 O- ∆17 O- δ34 S- Barite Sample VSMOW VSMOW VCDT 87 Sr/ 86 Sr type Locality (‰) (‰) (‰)
Isotope analysis of bulk barite in Songlin and Bazihu, South China Jinhe-08-1-1 13.6 -0.63 23.3 − B2 Songlin Jinhe-08-1-1-1 14.3 -0.63 22.3 0.71188 B2 Songlin Jinhe-08-1-1-2 13.4 -0.60 23.7 0.71233 B2 Songlin Jinhe-08-2-1 14.1 -0.44 26.2 0.71457 B1 Songlin Jinhe-08-2-B 16.6 -0.43 27.1 − B1 Songlin Jinhe-08-3-B 16 -0.32 26.3 − B1 Songlin Jinhe-3t 15 -0.64 25.1 − B2 Songlin Jinhe-3t-1-1 14.1 -0.69 23.3 − B2 Songlin Jinhe-3t-1-1-1 15.2 -0.77 24.0 − B2 Songlin Jinhe-3t-1-2 14.2 -0.68 23.4 − B2 Songlin Jinhe-3t-2 14.7 -0.69 22.6 − B2 Songlin Jinhe-3t-3 14.9 -0.60 21.7 − B2 Songlin Jinhe-3t-4 13.3 -0.60 23.4 0.71384 B2 Songlin NC-11 16.2 -0.45 30.9 − B1 Songlin NC-11-1 15.6 -0.32 31.0 − B1 Songlin NC-11-2 14.8 -0.21 29.8 − B1 Songlin SL-07-24 17.5 -0.42 25.3 − B1 Songlin SL-07-25 13.9 -0.56 26.4 − B2 Songlin SL-07-26 19.3 -0.28 31.2 − B1 Songlin SL-1 15.4 -0.39 29.9 − B2 Songlin SL-1-1 15.4 -0.32 28.5 − B2 Songlin SL-1-2 15.8 -0.28 36.2 − B2 Songlin SL-1-2-1 16.9 -0.28 35.1 − B2 Songlin SL-1-2-2 16.4 -0.39 34.1 − B2 Songlin SL-1-2-3 15.2 -0.24 34.6 − B2 Songlin SL-1-3 15 -0.31 29.9 0.71443 B2 Songlin SL-2 17.9 -0.22 30.2 − B1 Songlin SL-2-1 18 -0.49 23.1 − B1 Songlin SL-2-2 17.7 -0.40 26.2 − B1 Songlin SL-2-3 16.7 -0.34 33.6 − B1 Songlin SL-3 13.6 -0.43 22.8 − B2 Songlin SL-3-1 14 -0.57 22.2 − B2 Songlin
81 SL-3-2 13.9 -0.50 24.8 − B2 Songlin SL-4 15.3 -0.29 35.2 − B2 Songlin SL-4-1-1 16.3 -0.76 23.6 − B2 Songlin SL-4-1-2 15.8 -0.31 35.2 − B2 Songlin SL-4-2-1 15.3 -0.26 33.3 − B2 Songlin SL-4-2-2 15.8 -0.20 35.6 − B2 Songlin SL-5 15.3 -0.33 27.3 − B1 Songlin SL-6-1 15.6 -0.36 28.3 − B1 Songlin SL-6-1-2-1 16.1 -0.32 28.9 − B1 Songlin SL-6-1-2-2 14.9 -0.22 28.8 − B1 Songlin SL-6-2 16.3 -0.28 30.4 − B1 Songlin ZB-07-14 17.6 -0.24 24.2 − B2 Baizhu ZB-07-17 17.6 -0.35 30.4 − B1 Baizhu ZB-07-18-1 21.3 -0.44 28.2 − B1 Baizhu ZB-07-18-2 19.7 -0.49 27.4 − B1 Baizhu ZB-07-18-3 18.9 -0.28 28.5 − B1 Baizhu ZB-07-20 19.1 -0.14 41.7 − B2 Baizhu ZB-07-22 20.1 -0.31 31.4 − B1 Baizhu ZB-07-24-1-1 17.3 -0.56 26.6 − B1 Songlin ZB-07-24-1-2 16 -0.39 23.6 0.71229 B1 Songlin ZB-07-24-B 14.6 -0.47 27.0 − B1 Songlin ZB-07-25-1 14.8 -0.86 23.5 − B2 Songlin ZB-07-25-2 14.1 -0.79 23.7 − B2 Songlin ZB-07-25-3 14 -0.72 21.5 0.71146 B2 Songlin ZB-07-25-4 13.9 -0.68 21.2 0.71384 B2 Songlin ZB-07-25-b-D 14.7 -0.70 21.1 − B2 Songlin ZB-07-26-1-1 18.2 -0.24 35.2 − B1 Songlin ZB-07-26-1-2 16.8 -0.24 27.0 − B1 Songlin ZB-07-26-1-3 18.2 -0.26 29.9 − B1 Songlin ZB-07-26-1-4 17.6 -0.33 33.8 − B1 Songlin ZB-07-26-1-5 16.4 -0.37 32.4 0.71272 B1 Songlin ZB-07-26-2-1 19.5 -0.33 27.6 − B1 Songlin ZB-07-26-2-2 20.6 -0.21 29.2 − B1 Songlin ZB-07-26-2-3 19.7 -0.30 28.3 − B1 Songlin ZB-07-26-2-4 19.1 -0.25 − 0.71071 B1 Songlin
82 Isotope analysis of high-resolution sampling barite in Songlin and Bazihu, South China Jinhe-07-4-1 16.4 -0.26 − − B2 Songlin Jinhe-07-4-2 16.7 -0.27 − − B2 Songlin Jinhe-07-4-3 17.3 -0.48 − − B2 Songlin Jinhe-07-4-4 17.2 -0.47 − − B2 Songlin Jinhe-07-4-5 16.6 -0.69 − − B2 Songlin Jinhe-07-4-6 16.7 -0.74 − − B2 Songlin Jinhe-07-4-7 15.9 -0.71 − − B2 Songlin Jinhe-07-4-8 16.5 -0.79 − − B2 Songlin Jinhe-07-4-9 16.1 -0.74 − − B2 Songlin Jinhe-07-4-10 16.3 -0.52 − − B2 Songlin Jinhe-08-1-1-R 16.4 -0.82 22.0 − B2 Songlin Jinhe-08-1-2-R 17.2 -0.87 21.7 − B2 Songlin Jinhe-08-1-3-R 17.2 -0.84 21.8 − B2 Songlin Jinhe-08-1-4-R 16.8 -0.79 22.1 − B2 Songlin Jinhe-08-1-5-R 15.4 -0.85 22.1 − B2 Songlin OB-4-1-1 20.0 -0.43 27.8 − B1 Baizhu OB-4-1-2 20.6 -0.44 28.1 − B1 Baizhu OB-4-1-3 21.6 -0.47 28.0 − B1 Baizhu OB-4-1-4 22.0 -0.50 27.6 − B1 Baizhu OB-4-1-5 21.5 -0.50 27.5 − B1 Baizhu SL-2-1-R 19.0 -0.44 31.2 − B1 Songlin SL-2-2-R 17.5 -0.39 31.5 − B1 Songlin SL-2-3-R 18.4 -0.35 31.7 − B1 Songlin SL-2-4-R 17.8 -0.30 31.6 − B1 Songlin SL-2-5-R 18.1 -0.31 30.8 − B1 Songlin SL-2-6-R 19.5 -0.39 31.8 − B1 Songlin ZB-07-14-1-1 17.3 -0.27 − − B2 Baizhu ZB-07-14-1-2 19.6 -0.13 − − B2 Baizhu ZB-07-14-1-3 19.2 -0.24 − − B2 Baizhu ZB-07-14-1-4 17.9 -0.15 − − B2 Baizhu ZB-07-14-1-5 18.9 -0.15 − − B2 Baizhu ZB-07-14-1-6 18.5 -0.15 − − B2 Baizhu ZB-07-14-1-7 18.1 -0.12 − − B2 Baizhu ZB-07-14-2-1 17.0 -0.45 22.4 − B2 Baizhu ZB-07-14-2-2 19.9 -0.24 28.8 − B2 Baizhu ZB-07-14-2-3 19.3 -0.26 34.5 − B2 Baizhu ZB-07-14-2-4 19.5 -0.09 36.5 − B2 Baizhu ZB-07-14-2-5 19.0 -0.09 36.0 − B2 Baizhu ZB-07-14-2-6 19.4 -0.16 34.6 − B2 Baizhu ZB-07-14-2-7 18.9 -0.11 32.7 − B2 Baizhu ZB-07-14-2-8 19.5 -0.13 31.7 − B2 Baizhu
83 ZB-07-18-1-1-1 19.1 -0.41 31.4 − B1 Baizhu ZB-07-18-1-1-2 21.9 -0.45 27.2 − B1 Baizhu ZB-07-18-1-1-3 21.9 -0.52 27.6 − B1 Baizhu ZB-07-18-1-1-4 22.0 -0.47 28.4 − B1 Baizhu ZB-07-18-1-1-5 21.7 -0.43 29.0 − B1 Baizhu ZB-07-18-1-1-6 22.6 -0.37 28.9 − B1 Baizhu ZB-07-18-1-1-7 21.2 -0.47 28.5 − B1 Baizhu ZB-07-18-1-1-8 23.3 -0.40 27.9 − B1 Baizhu ZB-07-18-1-1-9 22.1 -0.42 27.6 − B1 Baizhu ZB-07-18-1-1-10 23.5 -0.45 27.1 − B1 Baizhu ZB-07-18-1-1-11 23.2 -0.45 27.1 − B1 Baizhu ZB-07-18-1-2-1 18.5 -0.44 26.3 − B1 Baizhu ZB-07-18-1-2-2 20.6 -0.57 25.0 − B1 Baizhu ZB-07-18-1-2-3 21.4 -0.46 25.4 − B1 Baizhu ZB-07-18-1-2-4 20.6 -0.33 25.7 − B1 Baizhu ZB-07-18-1-2-5 20.9 -0.48 25.2 − B1 Baizhu ZB-07-18-1-2-6 20.9 -0.36 25.3 − B1 Baizhu ZB-07-18-2-1-1 21.2 -0.40 26.5 − B1 Baizhu ZB-07-18-2-1-2 21.0 -0.38 26.3 − B1 Baizhu ZB-07-18-2-1-3 21.4 -0.42 26.8 − B1 Baizhu ZB-07-18-2-1-4 21.4 -0.45 26.8 − B1 Baizhu ZB-07-18-2-1-5 20.7 -0.41 30.3 − B1 Baizhu ZB-07-18-2-1-6 21.1 -0.39 30.5 − B1 Baizhu ZB-07-18-2-2-1 19.7 -0.49 27.3 − B1 Baizhu ZB-07-18-2-2-2 21.9 -0.46 25.8 − B1 Baizhu ZB-07-18-2-2-3 20.9 -0.41 26.0 − B1 Baizhu ZB-07-18-2-2-4 20.4 -0.44 25.6 − B1 Baizhu ZB-07-18-3-1 19.8 -0.35 − − B1 Baizhu ZB-07-18-3-2 20.7 -0.34 − − B1 Baizhu ZB-07-18-3-3 21.9 -0.30 − − B1 Baizhu ZB-07-18-4-1 19.6 -0.56 − − B1 Baizhu ZB-07-18-4-2 19.4 -0.58 − − B1 Baizhu ZB-07-18-4-3 18.9 -0.54 − − B1 Baizhu ZB-07-20-1-1 18.6 -0.13 45.3 − B2 Baizhu ZB-07-20-1-2 18.0 -0.21 44.3 − B2 Baizhu ZB-07-20-1-3 18.6 -0.12 42.7 − B2 Baizhu ZB-07-20-2-1 19.2 -0.11 45.5 − B2 Baizhu ZB-07-20-2-2 18.4 -0.16 43.2 − B2 Baizhu ZB-07-20-2-3 19.4 -0.17 42.2 − B2 Baizhu ZB-07-24-2-1 14.2 -0.32 23.3 − B2 Songlin ZB-07-24-2-2 15.8 -0.79 24.3 − B2 Songlin ZB-07-24-3-1 16.0 -0.85 23.8 − B2 Songlin
84 ZB-07-24-3-2 16.4 -0.82 23.6 − B2 Songlin ZB-07-24-3-3 16.4 -0.55 25.2 − B2 Songlin ZB-07-24-3-4 14.2 -0.58 26.8 − B2 Songlin ZB-07-24-3-5 14.3 -0.59 26.9 − B2 Songlin ZB-07-24-3-6 14.0 -0.53 27.8 − B2 Songlin ZB-07-24-3-7 13.9 -0.39 29.0 − B2 Songlin ZB-07-24-5-1 15.2 -0.85 − − B2 Songlin ZB-07-24-5-2 16.6 -0.81 − − B2 Songlin ZB-07-24-5-3 16.0 -0.82 − − B2 Songlin ZB-07-24-5-4 15.4 -0.86 − − B2 Songlin
85 VITA
Yongbo Peng was born in Shandong Province in China in June, 1978. He received his Bachelor of Science degree in biological sciences from Ludong University in 2003 and a Master of Science degree in geology and paleontology from the Nanjing Institute of
Geology and Palaeontology in the Chinese Academy of Sciences in 2007. He was accepted as a doctoral student in geology and geophysics at Louisiana State University in the spring of 2006. He is currently completing his doctoral degree in the laboratory of Dr.
Huiming Bao.
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