Table of Contents
Cover
PREFACE
Part I: Introduction
1 Chemostratigraphy as a Formal Stratigraphic Method
1.1. INTRODUCTION
1.2. BASIS AND DEVELOPMENT OF CHEMOSTRATIGRAPHY
1.3. CHEMOSTRATIGRAPHY AND CHRONOSTRATIGRAPHIC BOUNDARIES
1.4. CHEMOSTRATIGRAPHY AS FORMAL STRATIGRAPHIC METHOD
ACKNOWLEDGMENTS
REFERENCES
2 Glossary of Chemostratigraphy
2.1. INTRODUCTION
2.2. CHEMOSTRATIGRAPHY: TRANSCENDING BOUNDARIES AND EXPANDING POSSIBILITIES
2.3. GLOSSARY
2.4. SUMMARY
ACKNOWLEDGMENTS
REFERENCES
Part II: Precambrian
3 The Archean‐Proterozoic Boundary and the Great Oxidation Event
3.1. INTRODUCTION
3.2. SULFUR ISOTOPES
3.3. CHROMIUM ISOTOPES
3.4. MOLYBDENUM ISOTOPES AND CONCENTRATIONS
3.5. DISCUSSION AND CONCLUSIONS
ACKNOWLEDGMENTS
REFERENCES
4 Chronochemostratigraphy of Platform Sequences Across the Paleoproterozoic‐Mesoproterozoic Transition
4.1. INTRODUCTION
4.2. COLUMBIA SUPERCONTINENT AND POTENTIAL GSSP FOR PALEOPROTEROZOIC‐MESOPROTEROZOIC BOUNDARY
4.3. CHRONOSTRATIGRAPHIC CORRELATIONS FROM LATE PALEOPROTEROZOIC TO EARLY MESOPROTEROZOIC ERAS
4.4. CHEMOSTRATIGRAPHY
4.5. DETRITAL ZIRCON RECORD
4.6. CONCLUSIONS
ACKNOWLEDGMENTS
REFERENCES
5 Chemostratigraphy of the Mesoproterozoic‐Neoproterozoic Transition
5.1. INTRODUCTION
5.2. GEOCHRONOLOGIC AND CHEMOSTRATIGRAPHIC INFORMATION
5.3. AVAILABLE GEOCHEMICAL EVIDENCE OF MARINE OXYGEN LEVELS
5.4. MO ISOTOPE EVIDENCE OF OCEAN OXYGENATION
5.5. SUMMARY
5.6. METHODS
ACKNOWLEDGMENTS
REFERENCES
6 The Cryogenian‐Ediacaran Boundary in the Southern Amazon Craton
6.1. INTRODUCTION
6.2. GEOLOGICAL SETTING
6.3. THE MARINOAN GLACIAL DEPOSITS
6.4. THE CONTACT BETWEEN GLACIAL DEPOSITS AND CAP CARBONATE
6.5. THE MARINOAN CAP CARBONATE
6.6. PALEOBIOLOGY AND BIOSTRATIGRAPHY
6.7. CHEMOSTRATIGRAPHY AND GLOBAL CORRELATIONS
6.8. THE TIME SCALE OF CAP CARBONATE DEPOSITION AT THE CRYOGENIAN‐EDIACARAN BOUNDARY
6.9. FINAL REMARKS
ACKNOWLEDGMENTS
REFERENCES
7 The Ediacaran‐Cambrian Transition
7.1. A BIOGEOCHEMICAL PERSPECTIVE
7.2. CHEMOSTRATIGRAPHY
7.3. CLIMATE
7.4. ERNIETTAVILLE: EARLY BIOTURBATION?
7.5. METAZOANS TAKE THE STAGE
7.6. CONCLUSIONS
ACKNOWLEDGMENTS
REFERENCES
Part III: Paleozoic
8 δ
13
C Chemostratigraphy of the Ordovician‐Silurian Boundary Interval
8.1. INTRODUCTION
8.2. DOB’S LINN, SOUTHERN SCOTLAND
8.3. RÖSTÅNGA, SOUTHERN SWEDEN
8.4. WANGJIWAN, CHINA
8.5. TRURO ISLAND, CANADIAN ARCTIC
8.6. MIRNY CREEK, OMULEV MOUNTAINS, EASTERN SIBERIA
8.7. POSSIBLE RELATIONS BETWEEN δC CHEMOSTRATIGRAPHY, EUSTACY, AND EXTINCTION EVENTS
8.8. CONCLUSIONS
REFERENCES
9 Chemostratigraphy Across the Permian‐Triassic Boundary
9.1. INTRODUCTION
9.2. GEOLOGICAL SETTING
9.3. MATERIALS AND METHODS
9.4. CARBONATE MICROFACIES
9.5. CARBON AND OXYGEN ISOTOPE RESULTS
9.6. DISCUSSION
ACKNOWLEDGMENTS
REFERENCES
Part IV: Mesozoic
10 Chemostratigraphy Across the Triassic–Jurassic Boundary
10.1. INTRODUCTION
10.2. THE END‐TRIASSIC MASS EXTINCTION AND POTENTIAL CAUSES
10.3. INTENSELY STUDIED TRIASSIC‐JURASSIC BOUNDARY SUCCESSIONS
10.4. CHEMOSTRATIGRAPHY
10.5. CONCLUSIONS
ACKNOWLEDGMENTS
REFERENCES
11 Jurassic‐Cretaceous Carbon Isotope Geochemistry–Proxy for Paleoceanography and Tool for Stratigraphy
11.1. INTRODUCTION
11.2. CARBON ISOTOPE GEOCHEMISTRY: FROM A PALEOCLIMATIC TO A STRATIGRAPHIC TOOL
11.3. CARBON ISOTOPE STRATIGRAPHY: WIGGLES, EXCURSIONS, AND SPIKES
11.4. CARBONATE CARBON AND ORGANIC CARBON ISOTOPE STRATIGRAPHY
11.5. CARBON ISOTOPE STRATIGRAPHY: TERMINOLOGY
11.6. CARBON ISOTOPE STRATIGRAPHY: A PALEOCEANOGRAPHIC TOOL
11.7. C ISOTOPE STRATIGRAPHY AND THE JURASSIC‐CRETACEOUS BOUNDARY
ACKNOWLEDGMENTS
REFERENCES
12 Chemostratigraphy Across the Cretaceous‐Paleogene (K‐Pg) Boundary
12.1. INTRODUCTION
12.2. CAUSE FOR MASSIVE EXTINCTION AT THE CRETACEOUS‐PALEOGENE BOUNDARY: IMPACT OR VOLCANISM OR BOTH?
12.3. SELECTED CRETACEOUS‐PALEOGENE BOUNDARY SECTIONS
12.4. ELEMENTAL AND ISOTOPE CHEMOSTRATIGRAPHY
12.5. DISCUSSION AND CONCLUSIONS
ACKNOWLEDGMENTS
REFERENCES
Part V: Cenozoic
13 Cenozoic Chemostratigraphy
13.1. INTRODUCTION
13.2. CHRONOSTRATIGRAPHY OF THE CENOZOIC ERA
13.3. MAJOR GEOLOGICAL EVENTS AND SPATIOTEMPORAL SCALES
13.4. DISCUSSION
13.5. CONCLUSIONS
ACKNOWLEDGMENTS
REFERENCES
INDEX
End User License Agreement
List of Tables
Chapter 05
Table 5.1 [Mo], [Al], and δ
98/95
Mo of Carbonates from the Vazante Group.
Chapter 09
Table 9.1 Summary Statistics of Carbon and Oxygen Isotope Composition.
Chapter 12
Table 12.1 C, O isotopes and Ca isotopes, total organic carbon, Hg, Mo/Al, and Y/Ho ratios across the K/Pg boundary in a section at Højerup, Stevns Klint, Denmark.
Table 12.2 C, O and Ca isotopes, total organic carbon, Hg, Mo/Al, and Y/Ho ratios across the K/Pg boundary at Bottaccione, Gubbio, Italy.
Table 12.3 C, O and Ca isotopes, total organic carbon, Hg, Mo/Al, and Y/Ho ratios across the K/Pg boundary in the Um Sohryngkew section, Meghalaya, India.
Table 12.4 C‐, O‐, and Ca‐isotope analyses, Hg and TOC concentrations and Mo/Al and Y/Ho ratios for the Poty Quarry section, Paraíba Basin, Brazil.
Table 12.5 Rare earth elements (ppm) and PAAS‐normalized Y/Ho, Th/U, Ce/Ce*, Pr/Pr*, and Eu/Eu* for the Højerup section, Stevns Klint, Denmark.
Table 12.6 Rare earth elements (ppm) and PAAS‐normalized Y/Ho, Th/U, Ce/Ce*, Pr/Pr*, and Eu/Eu* for the Bottaccione section, Gubbio, Italy.
Table 12.7 Rare earth elements and PAAS‐normalized Y/Ho, Th/U, Ce/Ce*, Pr/Pr*, and Eu/Eu* for the Um Sohryngkew section, Meghalaya, India.
Table 12.8 Rare earth elements (ppm), PAAS‐normalize U/Th, Ce/Ce*, Pr/Pr* and Eu/Eu* for the Poty Quarry section, Brazil.
Table 12.9 Mo and U enrichment factors for some Højerup, Bottaccione and Um Sohryngkew samples in this study.
Table 12.10 Hg isotopes (‰ relative to NIST SRM 3133) in samples with Hg enrichments: spike I (Um Sohryngkew, Poty); spike II (Højerup, Bottaccione, Um Sohryngkew), spike III (Poty).
List of Illustrations
Chapter 02
Figure 2.1 Yearly increase of number of publications using geochemistry for stratigraphy.
Chapter 03
Figure 3.1 Summary of geological evidence of atmospheric oxygenation and glacial events in the later Archean and Paleoproterozoic. Glacial events named after the type Huronian glacials in Canada [
Papineau et al.,
2007]. BIF abundance curve according to
Isley and Abbott
[1999]; carbon isotope curve compiled from
Lindsay and Brasier
[2002],
Melezhik et al
. [2007], and
Maheshwari et al
. [2010]; and the other proxies from
Holland
[1994] and
Bekker et al
. [2006].
Figure 3.2 (a) δ
34
S and Δ
33
S global curves for the Archean and Paleoproterozoic. (b) Termination of the MIF sulfur signal (Δ
33
S) in the Transvaal Supergroup. Note that the transition, and thus the formation of a functional ozone layer, takes place in <4 m stratigraphically. Malm, Malmani Formation; Timeball, Timeball Hill Formation. Re‐Os age according to
Hannah et al
. [2004].
Figure 3.3 Cr isotope evolution of seawater in the Precambrian and present values. Shaded areas represent notable oxygenation events: OW1, Mesoarchean Oxygen Whiff; OW2, Neoarchean Oxygen Whiff; GOE, Great Oxidation Event; NOE, Neoproterozoic Oxidation Event. Gray line represents the envelope of the most fractionated data and approximates the evolution of atmospheric oxygen. Present‐day seawater values from
Paulukat et al
. [2016]. Other sources of data: Isua BIF:
Frei et al
. [2016], Ijzermyn iron formation (ca. 3 Ga):
Crowe et al
. [2013], Mesoproterozoic‐Tonian carbonates:
Gilleaudeau et al
. [2016], Otavi Group carbonates (Cryogenian):
Rodler et al
. [2016]; Ediacaran iron formations:
Frei et al
. [2013, 2017].
Figure 3.4 Mo isotope evolution in the Transvaal Supergroup between 2.64 and ca. 2.5 Ga. The curve in gray represents the δ
98/95
Mo evolution of marine carbonates, and the black curve the isotopic evolution of shales. U‐Pb ages according to Altermann and Nelson [1998 and references therein].
Figure 3.5 Summary of chemostratigraphic curves, BIF deposition, and glacial events (blue bars) near the Archean‐Proterozoic boundary. Three different proposals of Global Stratotype Section and Point (GSSP) for the boundary are shown. GSSP 1: current boundary at 2500 Ma. GSSP 2: base of the second and possibly near‐global Huronian glaciation [
Gradstein et al.,
2012]. GSSP 3: termination of the sulfur MIF (Δ
33
S) at ca. 2.33 Ga [
Luo et al.,
2016]. Sources of data: δ
53
Cr:
Frei et al
. [2009], δ
98/95
Mo of shales:
Wille et al
. [2007], Δ
33
S amplitude:
Williford et al
. [2011 and references therein],
Luo et al
. [2016]; δ
13
C, BIF abundance and glacial events: same as for Figure 3.1.
Chapter 04
Figure 4.1 World distribution of 2.1–1.8 Ga orogens and cratons which are the primary building blocks of Columbia/Nuna supercontinent shown in Figure 4.2. 1. Trans‐Hudson orogen; 2. Penokean orogen; 3. Taltson‐Thelon orogen; 4. Wopmay orogen; 5. Cape Smith‐New Quebec orogen; 6. Torngat orogen; 7. Foxe orogen; 8. Nagssugtoqidian orogen; 9. Makkovikian‐Ketilidian orogen; 10. Transamazonian orogen; 11. Itabuna‐Salvador‐Curaçá orogen; 12. Eburnean orogen; 13. Limpopo belt; 14. Moyar belt; 15. Capricorn orogen; 16. Trans–North China orogen; 17. Central Aldan belt; 18. Svecofennian orogen; 19. Kola‐Karelian orogen; 20. Transantarctic orogen. Paleoproterozoic to Mesoproterozoic basins occur in most cratons.
Figure 4.2 (a) Columbia/Nuna supercontinent configuration at 1.4 Ga [
Pehrsson et al
., 2015;
Evans et al
., 2016]. (b) Reconstruction of the Columbia supercontinent at the Paleoproterozoic‐Mesoproterozoic boundary showing the location of main depocenters and orogens. 1. Changcheng‐Jixian systems; 2. Mount Isa and McArthur superbasins; 3. Wernecke Supergroup; 4. Fort Simpson basin; 5. Coppermine basin; 6. Muskwa basin; 7. Athabasca Basin; 8. Thelon basin; 9. Baraboo sequence; 10. Sioux sequence; 11. Ortega sequence; 12. Kureika‐Anabar basin; 13. Teya‐Chapa basin; 14. Turukhansk basin; 15. Udzha and East Anabar basins; 16. Kotuy basin; 17. Uchur and Aimchan groups; 18. Espinhaço Supergroup; 19. Akanyaru and Kibara supergroups.
Figure 4.3 Chronostratigraphic framework of the preserved stratigraphic record of the Columbia supercontinent between the Statherian and Calymmian periods.
Figure 4.4 Simplified geological map of the Peninsular India showing the four main cratonic domains (Dharwar, Bastar, Aravalli‐Bundelkhand, and Singhbhum) and Proterozoic basins. MB, Marwar basin; VB, Vindhyan Basin; ChB, Chhattisgarh basin; CuB, Cuddapah Basin; IB, Indravati basin; PG, Pranhita‐Godavari basin; KBB, Kaladgi‐Bhima basin.
Figure 4.5 Stratigraphic sections and stable C isotope data for carbonate sequences of Laurentia (Modified after
Hahn et al
. [2013]) and North Australia [
Lindsay and Brasier
, 2000].
Figure 4.6 Stratigraphic sections and stable C isotope data for carbonate sequences of India [
Ray et al
., 2003] and North China Craton. Curves after 1.
Xiao et al
. [1997] as dot;
Chu et al
. [2007] as line; 2.
Hongwei et al
. [2011]; 3.
Guo et al
. [2013].
Figure 4.7 Isotope composition of δ
13
C‰ and δ
34
S‰ in carbonate sequences deposited from 1.7 and 1.4 Ga. (a) δ
13
C‰ for marine carbonates deposited from late Paleoproterozoic to early Mesoproterozoic transition in Australia, India, Laurentia, and North China. (b) δ
34
S‰ data of seawater sulfate [
Guo et al
., 2015].
Figure 4.8 Detrital zircon age histograms for the sedimentary basins formed between 1.8 and 1.4 Ga in Siberia [
Priyatkina et al
., 2016], Baltica [
Bogdanova et al
., 2015], Congo [
Fernadez‐Alonso et al
., 2012], India [
Collins et al
., 2015;
Lancaster et al
., 2015;
Sahoo et al
., 2018], North China [
Ying et al
., 2011;
Liu et al
., 2014;
Wang et al
., 2016], Laurentia [
Rainbird and Davis
, 2007;
Furlanetto et al
., 2013, 2016], Yangtze [
Greentree and Li
, 2008;
Zhao et al
., 2010;
Wang et al
., 2012;
Chen et al
., 2013], North Australia [
Page et al
., 2000], West Australia [
Martin et al
., 2008;
Spaggiari et al
., 2015], and São Francisco Craton [
Danderfer et al
., 2009, 2015;
Marques
, 2009;
Chemale et al
., 2012;
Santos et al
., 2013;
Franz et al
., 2014;
Guadagnin et al
., 2015a, 2015b]. See text for explanation.
Figure 4.9 Plot of
ε
Hf(t)
versus age of concordant detrital zircons from Statherian to Calymmian sequences of South Australia, India, Yangtze, North China, and Sao Francisco Craton. Depleted mantle (DM) curve after
Bodet and Scharer
[2000]. Average PPr (Paleoproterozoic), NAr (Neoarchean), MAr (Mesoarchean), and PAr (Paleoarchean) growth line for
176
Lu/
177
Hf = 0.013.
Chapter 05
Figure 5.1 Paleogeographic map of Rodinia at 1.0 Ga and paleogeographic location of Arequipa Massif (A), Amazonian Craton (AM), Avalonia (AV), Baltica (Russian craton) (B), Congo Craton (C), Coats Land‐Maudheim‐Grunehogna province of East Antarctica (CMG), Ellsworth‐Whitmore Mountains block (in Pangea position) (E), Florida (in pre–Pangea position within Gondwana and including the Carolina terrane) (F), Falkland‐Malvinas Plateau (F/MP), Kalahari Craton (K), Marie Byrd Land (MBL), New Guinea (NG), Rockall Plateau with adjacent northwest Scotland and northwest Ireland (R), Rio de la Plata Craton (RP), Siberia (Angara Craton) (S), São Francisco Craton (SF), Svalbard block (Barentia) (SV), West African Craton (WA), and hypothetical Texas plateau (TxP). Numbers correspond to carbonate successions: (1) Vazante Group, (2) Paranoá Group, (3) Turukhansk Group, (4) Atar Group, (5) Bylot Supergroup, and (6) São Caetano Complex.
Figure 5.2 Lithostratigraphy and C isotope chemostratigraphy of the Turukhansk Group, Turukhansk Uplift, Siberia.
87
Sr/
86
Sr inferred ages. Pb‐Pb ages. δ
53
Cr data.
Figure 5.3 Lithostratigraphy and C isotope chemostratigraphy of the Vazante Group, São Francisco Basin, Brasilia Deformed Belt, Brazil. Re‐Os ages. δ
53
Cr data. δ
98/95
Mo data (this study).
Figure 5.4 Lithostratigraphy and C isotope chemostratigraphy of the Paranoá Group. Detrital zircon U‐Pb ages. Pb‐Pb ages.
87
Sr/
86
Sr inferred ages. Biostratigraphic ages.
Figure 5.5 Lithostratigraphy and C isotope chemostratigraphy of the Atar Group. Re‐Os age.
Figure 5.6 Lithostratigraphic and C isotope chemostratigraphic correlation of several Mesoproterozoic‐Neoproterozoic carbonate successions. Note how the δ
13
C values decrease from ~4‰ at 1.1 Ga to ~2‰ during the latest Mesoproterozoic. These values increase again to ~4‰ just before the Mesoproterozoic‐Neoproterozoic boundary when they decrease to values below 0‰. The δ
13
C values increase again in the Neoproterozoic to values as high as ~4‰ and remain positive. Note the occurrence of an early Neoproterozoic discrete glacial event. This event was restricted to the São Francisco Craton as evidenced by the presence of glaciogenic deposits in the Vazante and Paranoá groups.
Figure 5.7 Lithostratigraphy and C isotope chemostratigraphy of the Bylot Supergroup. Re‐Os ages. U‐Pb ages.
87
Sr/
86
Sr inferred ages.
Figure 5.8 Lithostratigraphy and C isotope chemostratigraphy of the São Caetano Group, Transversal Zone, Borborema Province.
Figure 5.9 Evolution of the C, Sr, Cr, and Mo isotope composition of the seawater during the 2.5–0.5 Ga interval. Note how the increase in the variability of the seawater δ
13
C values in the late Mesoproterozoic coincides with an increase in the seawater δ
98/95
Mo and δ
53
Cr values [
Gilleaudeau et al.,
2016, this work]. This increase continues toward the Neoproterozoic, suggesting a continued increase in the ocean oxygen levels. The increase in the ocean oxygen levels implies an increase in oxidative weathering and thus an increase in atmospheric
p
O
2
. The increase in atmospheric and oceanic oxygen levels is supported by the increase in the sulfate concentrations in the global oceans [
Kah et al.,
2004]. This increase in biospheric oxygenation coincides with the increase in eukaryotic life as suggested by the increase in acritarch diversity [
Knoll et al.,
2006]. Note how the first appearance of
Bangiomorpha pubescens
at 1.045 Ga [
Gibson et al.,
2018] coincides with the highest δ
53
Cr values in the whole Mesoproterozoic. This suggests that the appearance of marine multicellular/sexual photosynthetic eukaryotic life also coincides with an increase in marine oxygenation. Normalized
87
Sr/
86
Sr curve after
Shields
[2007].
Chapter 06
Figure 6.1 Location and simplified geologic map of southwestern Amazon Craton with Neoproterozoic cap carbonate occurrences. The Cryogenian‐Ediacaran boundary is exposed in the main cap carbonate succession overlying crystalline and metasedimentary rocks of Precambrian basement in the southwest and south portion of the Amazon Craton.
Figure 6.2 Lithostratigraphy and major geologic events of the Neoproterozoic‐Cambrian deposits in the southern Amazon Craton and northern Paraguay Belt.
Figure 6.3 Geological‐structural setting of the southern Amazon Craton and the northern Paraguay Belt, near Planalto da Serra region. Spatial and geometrical distribution of Neoproterozoic and Cambrian units in map (a) and section (b) with primary and tectonic contacts established during post‐Ordovician brittle‐to‐brittle‐ductile deformation. The type section of the Serra Azul Formation [cf.,
Alvarenga et al.,
2007] is plotted in the map for better understanding of the structural‐geometric array of the region. Observed offshore deposits related to the Raizama Formation unconformably overlie the glaciogene diamictites of the Puga Formation.
Figure 6.4 Simplified stratigraphic section of Marinoan glaciogene deposits (Modified of
Silva et al
. [2015]). Two advance‐retreat cycles of coastal glaciers constitute the Puga succession (a). Core carried out by geological survey in the Rondonia Mineral Company (RMC) quarry showing the Cryogenian‐Ediacaran boundary, coincident with the diamictite‐dolostone contact (b). (c) Puga diamictite with disseminated clast.
Figure 6.5 The Cryogenian‐Ediacaran boundary in South America. (a and b) The deformed contact between glaciogene diamictites and cap dolostone, respectively, in Mirassol d'Oeste and Espigão d'Oeste regions. (c) Open fold in the base of cap dolostone onlapped by subhorizontal dolostone beds, exposed in Espigão d'Oeste region.
Figure 6.6 Neoproterozoic cap carbonate stratigraphic sections in the southern Amazon Craton. (a) Open pit of the RMC quarry, (b) outcrop at the Chupinguaia region, (c) open pit in the Calcario Tangará quarry, Tangará da Serra region, and (d) open pit in Terconi quarry, Mirassol d'Oeste region. The δ
13
C curves of the carbonate platform deposits in the Amazon Craton with values in ‰ were obtained in
Nogueira et al
. [2003, 2007],
Font et al
. [2006],
Riccomini et al
. [2007],
Alvarenga et al
. [2004, 2008],
Romero et al
. [2013], and
Soares et al
. [2013].
Figure 6.7 General aspects of cap carbonate in the Amazon Craton. (a) Microbialites. (b) Giant wave ripple. (c) Calcite crystal fans (pseudomorphs after aragonite) interbedded with undulated lamination. (d) Detail of (b), showing the climbing wave ripple cross‐lamination and well‐preserved crest of megaripple. (e) Fenestral porosity filled by bitumen. (f) Dolomitized crystal fans. (g) Macropeloids (scale bar = 2 cm). Bitumen impregnation detached all laminations of carbonates.
Figure 6.8 Representative fossils of the Marinoan cap carbonate in the Amazon Craton. (a) Typical microbial laminite. (b) Fenestral porosity filled by bitumen and euhedral dolomite. (c) Tubestone structures associated with microbial laminites. (d and e)
Leiosphaeridia
taxa from the intermediate part to top cap carbonate (scale bars = 20 μm).
Figure 6.9 Carbon isotope and
87
Sr/
86
Sr chemostratigraphy and lithostratigraphy of the Araras Group in the southern Amazon Craton (Modified of
Nogueira et al
. [2007]) compared with composite δ
13
C curve of the Otavi Group in Namibia [Halverson
et al
., 2005;
Macdonald et al
., 2009]. The
13
C isotopic trend of the Araras Group is correlated with the shelf values of Cryogenian Abeneb Subgroup. Approximately 500 m thick shelf deposits of Araras Group sections have been stretched to fit the Otavi Group section. The Maiberg anomaly is recorded in the base of Araras Group and marks the Cryogenian‐Ediacaran boundary. In this context, the Araras Group does not transpass the lower Ediacaran. Middle to late Ediacaran is not recorded in the southern Amazon Craton.
Figure 6.10 Insertion of the post‐Marinoan carbonates of the southern Amazon Craton in the global correlations for the Ediacaran key locations with associated isotopic, geochronologic, and biostratigraphic data. All the data are color coded for geographic location: Black, southern Amazon Craton, Central Brazil; orange, northwestern Canada; purple, South China; blue, Namibia; green, Oman; yellow, Australia; red, White Sea region of Russia; brown, Avalon Terrane of Newfoundland and the United Kingdom. The carbonates of the Araras Group and Espigão d'Oeste formations are close with lower Ediacaran age below of 614 Ma, corroborated by δ
13
C and
87
Sr/
86
Sr values. Enriched positive values of
87
Sr/
86
Sr in the upper portion of succession (dotted circle) represent samples probably contaminated with high content of siliciclastics. P, Puga Formation; M, Mirassol d'Oeste Formation; Guia, Guia Formation; SQ, Serra do Quilombo Formation; Nobres, Nobres Formation; Boc., Bocaina Formation; Tam., Tamengo Formation; Rz., Raizama Formation. For additional details and other abbreviation of units, see
Macdonald et al
. [2013].
Chapter 07
Figure 7.1 The agronomic revolution or Cambrian substrate revolution depicted in this modified illustration by Peter Trusler reveals profound changes in bioturbation across the Ediacaran (E)‐Cambrian (C) transition. The enigmatic Ediacaran biota (left) is rooted in ubiquitous microbial mats that carpeted the seafloor (Precambrian matgrounds) and largely sealed the anoxic sediments beneath from the free exchange of gases. The slow diffusion of sulfate into these mats provided an oxidant for microbial sulfate reduction, leading to the buildup of toxic H
2
S, a proposed critical nutrient for the Ediacaran biota. The horizontal burrows of animals within or beneath the mats and the activities of the Ediacaran organisms on its surface, which appear very late in the Ediacaran game, had little effect on sedimentary layering or the release of gases to seawater. In contrast, the deep dive of early Cambrian (right) animals into the sediments in search of food and shelter disrupted the mats and allowed the free exchange of gases across the sediment‐water interface, including O
2
. Ventilation and mixing of the sediments is implicated in the demise of the Ediacaran biota, if H
2
S was a critical physiological resource, as well as the buildup of sulfate in the oceans.
Figure 7.2 Generalized trend in the carbon isotope composition of marine carbonates through the TES and Fortunian stage of the Cambrian period. This compilation is based on data from a wide array of sources from successions in India [
Kaufman et al.,
2006], Morocco [
Maloof et al.,
2005], Namibia [
Kaufman et al
., 1991;
Saylor et al.,
1998;
Wood et al.,
2015], South China [
McFadden et al.,
2008;
Cui et al.,
2015, 2016a, 2016b, 2017], Oman [
Fike et al.,
2006], India [
Kaufman et al.,
2006;
Tewari and Sial
, 2007], Siberia [
Knoll et al
., 1995a;
Kaufman et al.,
1996;
Maloof et al.,
2010 and references therein;
Cui et al.,
2016c], South Australia [
Husson et al.,
2015], and the United States [
Corsetti and Kaufman
, 2003;
Hebert et al.,
2010;
Verdel et al.,
2011]. This compilation should be compared against those presented in
Xiao et al
. [2016] insofar as the upper reaches of the Shuram excursion are not pinned to the 551 Ma U‐Pb age at the Doushantuo‐Dengying contact (see text for explanation). Positions for ice ages are marked by ΔΔΔ. Note the uncertainties of the positions of the base of the TES and the end of the Fortunian, which relate to issues of correlating chemostratigraphic and biostratigraphic events between basins.
Figure 7.3 Generalized trend in the strontium isotope composition of well‐preserved high‐Sr marine limestones through the Ediacaran period modified from
Xiao et al
. [2016] with age constraints based on their “correlation 2” and assuming the Shuram excursion is pinned to the 551 Ma age for the Doushantuo‐Dengying boundary (see discussion in text and Fig. 7.2 caption). This compilation is based on data from successions in South China [
Sawaki et al
., 2010;
Cui et al.,
2015], Oman [
Burns et al.,
1994], South Australia [
Calver
, 2000], southern Siberia [
Melezhik et al.,
2009], and northern Siberia [
Cui et al.,
2016c]. The trend marked by the thick gray line indicates a plateau of ca. 0.7080 followed by a profound rise in
87
Sr/
86
Sr values coincident with the Shuram excursion up to as high as 0.7090 and likely associated with intense weathering up uplifted terrains worldwide. The trend then declines back to ~0.7080 very near to the Ediacaran‐Cambrian boundary.
Figure 7.4 (a) White to red colored Gaskiers diamictite overlain by a 50 cm thick white carbonate (left of the field assistant) on the shore of Conception Bay at Harbour Main, Newfoundland. Inset shows the brecciated and potentially karstified upper surface of the thin carbonate filled with green mudstone of the overlying Drook Formation. (b) Green meta‐basalt of the basal Catoctin Formation with diapirs (flame structures) of Fauquier Formation carbonate injected between the chilled pillow margins on Goose Creek near Aldie, Virginia, United States. Inset illustrates a hyaloclastic texture at contact between the Catoctin meta‐basalt and marble of the Fauquier cap carbonate, indicating that the sediments were water‐saturated during emplacement of the volcanic rocks. (c) A 20 m thick diamictite along the Tas‐Yuryakh River lying unconformably above Turkut Formation dolomites in the Olenek uplift, Arctic Siberia, Russia. The freshly exposed diamictite has a green‐gray sandy to clayey calcareous matrix with abundant cobble‐ to boulder‐sized clasts in a weakly stratified pile. The randomly oriented clasts (see inset) are primarily derived from the Turkut and underlying Khatyspyt formations, but they also consist of occasional green igneous rocks and metamorphic rocks of exotic origin. Subrounded clasts plucked from the surrounding carbonate‐rich matrix are notably faceted.
Figure 7.5 (a) Freshly exposed
Pteridinium
surfaces coated with yellow colored jarosite, a hydrous sulfate of potassium and iron that forms as a product of pyrite weathering, at Aarhauser on Farm Aar, near Aus, Namibia [
Hall et al.,
2013]. Inset shows a typical iron oxidize patina on an exposed and weathered surface of
Pteridinium
at the same locality. (b) “Mud chip” breccia associated with
Ernietta
‐bearing sandstone from Ernietta Hill on Farm Aar. The mud chips have the iron oxide patina indicated above, and a partially exposed erniettid is exposed on the surface (see red arrow). This suggests that the chips represent the surface connection between infaunal
Ernietta
bases and their epibenthic fronds. (c) Sock‐shaped sandstone concretions with flat upper surfaces likely to represent
Ernietta
specimens that have lost their tubular covering through exposure and weathering on Windy Peak, Farm Aar. These specimens are the same size and general shape to
Ernietta
preserved in situ with tubular structures (d) (yellow scale bar 1 cm).
Figure 7.6 Exceptionally preserved
Ernietta
and
Rangea
specimens discovered in sandy gutter casts on Farm Aar, southern Namibia. (a) The most complete
Ernietta
specimen known to date [
Ivantsov et al.,
2016], with white arrow pointing to the position of the connection between infaunal base and epibenthic frond (see Fig. 7.5b: coin is 2.26 cm in diameter). (b) Illustrated reconstruction of the
Ernietta
specimen depicted in (a) with tubular bilayer shown. (c) Illustrated reconstruction of
Rangea
with a tubular core and sixfold symmetry of vanes including at least three orders of fractal folding. (d) The most complete
Rangea
specimen known to date [
Vickers‐Rich et al.,
2013] showing sand‐filled base of the organism (image of same specimen in
Vickers‐Rich et al
. [2013] shows patina of yellow jarosite, a weathering product of pyrite).
Figure 7.7 (a) Conception‐style preservation [
Narbonne
, 2005] of
Charniodiscus spinosus and Cychrus procerus
in positive relief on a surface of the Mistaken Point Formation of Newfoundland. The fossil surface is overlain by a thin bed of volcanic ash (visible as dark layer on the upper right). (b) Spindle‐shaped
Fractofusus misrai
with primary and secondary branches on the same surface in Newfoundland with visible volcanic ash. Fossils in (a) and (b) are members of the Avalon Assemblage. These are representatives of the White Sea Assemblage. (c)
Dickinsonia costata
moving (see arrows) and resting traces. (d)
Kimberella quadrata
(right arrow) and associated trace fossil
Kimberichnus teruzzi
(left arrow). (e)
Aspidella
with holdfast and stalk, but no frond. (f)
Arborea arborea
.
Chapter 08
Figure 8.1 Diagram showing the common occurrence of stratigraphic gaps in the Ordovician‐Silurian boundary interval as illustrated by seven important successions in North America and Baltoscandia. All of these sequences have full or partial δ
13
C chemostratigraphic control. (1) The Monitor Range, Nevada, succession. (2) The Upper Mississippi Valley succession in Illinois. (3) The succession in Adams County, southern Ohio. (4) The Bruce Peninsula succession in southern Ontario, Canada. (5) The western Anticosti Island succession, Quebec, Canada. (6) The succession in the Siljan region, south‐central Sweden. (7) The northern Estonia succession.
Figure 8.2 Inferred Late Ordovician paleogeographic position of some important sections dealt with herein. Figured localities are as follows. (1) Dob’s Linn, Scotland. (2) Röstånga, S. Sweden. (3) Wangjiwan, Yangtze platform, China. (4) Truro Island, Canadian Arctic. (5) Mirny Creek, eastern Siberia. (6) Anticosti Island. (7) Canada; G, Monitor Range, Nevada.
Figure 8.3 δ
13
C
org
curve through the Ordovician‐Silurian boundary interval at the GSSP of the base of the Silurian at Dob’s Linn, Scotland. Note the range of the HICE through the
M. extraordinarius
and
M. persculptus
zones. Also note that the lower portion of the HICE exhibits a rather gradual, rather than abrupt, increase in δ
13
C
org
values.
Figure 8.4 δ
13
C
org
curve through the Ordovician‐Silurian boundary interval in the Röstånga‐1 drill core, southern Sweden. Note the abrupt increase in isotopic values at the base of the Hirnantian and the return to baseline values in the
M. persculptus
Zone in the topmost Hirnantian.
Figure 8.5 δ
13
C
org
and δ
13
C
carb
curves through the Ordovician‐Silurian boundary interval at the Wangjiwan (Riverside) section, China, which is very near the GSSP of the global Hirnantian stage. This is a stratigraphically very condensed succession, the total thickness of the Hirnantian stage being slightly less than one meter. However, the Hirnantian δ
13
C
carb
curve has a rather typical HICE appearance with a rapid increase in isotope values at the base of the Hirnantian; the δ
13
C
org
curve is more similar to the Dob’s Linn HICE curve with its essentially background‐size isotopic values in the lower Hirnantian.
Figure 8.6 δ
13
C
org
curve through the Ordovician‐Silurian boundary interval on Truro Island, Canadian Arctic. Note the characteristic shape of the HICE, which shows close similarity to the Röstånga curve (cf., Fig. 8.4 herein). As in the other successions discussed herein, the Ordovician‐Silurian boundary level is not marked by a prominent isotope excursion.
Figure 8.7 δ
13
C
carb
curve through the Ordovician‐Silurian boundary interval at Mirny Creek, eastern Siberia. The isotope curve through this important succession has good graptolite [e.g.,
Koren' et al.,
1988] and conodont [
Zhang and Barnes,
2007] biostratigraphic control through the upper Katian‐lower Rhuddanian interval. The Hirnantian segment of the isotope curve, which exhibits some similarity to the Truro Island and Röstanga‐1 curves, has little more than baseline values except for raised values at the beginning of the Hirnantian and a brief interval in middle portion of the
M. persculptus
Zone. There is also a notable negative curve trend in the lowermost Rhuddanian that is not obvious in the other successions dealt with herein, but the significance of this feature remains unclear.
Figure 8.8 Model of the inferred relations between chemostratigraphy, eustatic events, and major faunal extinction intervals in the Ordovician‐Silurian boundary interval. This model, here applied to the stratigraphically apparently continuous succession in the Röstånga‐1 drill core from southern Sweden, is similar to one proposed by
Bergström et al
. [2014; Fig. 19] for the Monitor Range succession in Nevada, which has a major stratigraphic gap at the systemic boundary. Although not identified in the Röstånga‐1 drill core, in many other successions worldwide, there are widespread, more or less prominent stratigraphic gaps at the base of the Hirnantian and in the middle‐upper part of Stage Slice Hi1 that are interpreted to reflect glaciations on Gondwana. These appear to correlate with significant global extinction events. As noted by
Bergström et al
. [2014], these extinction events are not at the end of the Ordovician as is commonly erroneously stated in the literature but in the early‐middle Hirnantian.
Chapter 09
Figure 9.1 A compilation of published bulk rock δ
13
C data from multiple sites in Iran (a) and the P‐Tr GSSP at Meishan in South China (b). The data points are placed on a dimensionless timeline to ensure a better comparison of trends in the δ
13
C data of both regions. The dimensionless timeline is based on conodont zones. (The biostratigraphic scheme and sources regarding sea‐level change and the extinction horizon are given in Supporting Information Text S1.) The trendline is based on a subsampling routine as outlined in
Schobben et al
. [2017]. The yellow error bars represent the value ranges of newly generated δ
13
C data for this study.
Figure 9.2 A survey of isotope stratigraphic markers as defined in previous studies on P‐Tr carbonate rock exposed in Iran (see the Supporting Information Text S1 and Table S1 for data and sources). The survey evaluates the methodology of material retrieval and classifies the isotope events. The isotope events are defined according to the following set of definitions. (a) Second‐order carbon isotope excursions:
S
is small (0.3–0.5‰),
M
is medium (0.5–1‰), and
L
is large (>1‰). If the amplitude is smaller than 0.3‰, we labeled the excursion as “close to the detection limit,” as the value is close to the analytical bias. (b) Monotonic carbon isotope trends. (c) Carbon isotope shifts of more than 0.3‰. (d) Bar plot that counts the magnitude of the identified second‐order excursions. (e) Bar plot that counts the material retrieval method for isotope analysis, cq: drilling or grounding of the whole rock. (f) Paleogeographic situation (Adapted from
Stampfli and Borel
[2002]. Reproduced with permission of Elsevier) of the sampling sites covered by the survey and the investigatory object of the current study.
Figure 9.3 Isotope sampling grid and results of the
argillaceous red, nodular, bioturbated wackestone
(Zl 11) (digits within the scale bars represent 2.5 mm). (a) Polished rock slab with sampling grid (circumference of individual drill sites is representative of actual dimensions). (b) Crossplot of δ
13
C versus δ
18
O. (c) Probability density distributions for δ
13
C and δ
18
O. (d) Spatial grid of δ
13
C and box plots (boxes depict the IQR, and bars within the boxes the median) for cumulative variation across the horizontal and vertical axis. (e) Spatial grid of δ
18
O and box plots for cumulative variation across the horizontal and vertical axis.
Figure 9.4 Isotope sampling grid and results of the
gray/yellow nodular sponge wackestone
(Zl 18a). See caption and legends of Figure 9.3 for explanations for panels a–e.
Figure 9.5 Isotope sampling grid and results of the
laminated bindstone
(Zl 23). See caption and legends of Figure 9.3 for explanations for panels a–e. The diamond in panel (d) and (e) demarcates an anomalous isotope value (see text).
Figure 9.6 Isotope sampling grid and results of the
oncoid wackestone‐floatstone
(Zl 35). See caption and legends of Figure 9.3 for explanations for panels a–e.
Figure 9.7 Microscopy images of rock thin sections (digits within the scale bars represent 50 μm).
Red, nodular, bioturbated wackestone
(Zl 11) with (a) ostracod test filled with sparite, (b) echinoderm fragment, (c) burrow with a geopetal infill, and (d) a dolomitized section of the matrix.
Gray/yellow nodular sponge wackestone
(Zl 18a) with (e) sponge spicules and (f) a transition from a fine micritic matrix to a more crystalline matrix, corresponding to the color transition from gray to yellow upsection (Fig. 9.4).
Laminated bindstone
(Zl 23) with (g) an ostracod test with a fine crystalline infill and (h) laminae formed by grain size and mineralogy differences and small discontinuity surfaces.
Figure 9.8 Microscopy images of rock thin sections (digits within the scale bars represent 50 μm).
Laminated bindstone
(Zl 23) with (a) a calcite vein crosscutting the internal lamination.
Oncoid floatstone
with (b) calcite veins interrupted and offset by stylolitization, (c) nuclei of oncoid partly consisting of dolomite and large‐sized sparite crystals, and (d) oncoids within clotted micritic matrix (the inset corresponds to panel c).
Figure 9.9 Bootstrap sampling exercise comprising multiple routines that probe for the effect of the sample size (i.e., the number of samples analyzed per limestone bed) on the median within‐bed isotope value. In other words, this approach mimics the effect of drilling individual subsamples from limestone beds (i.e., 1, 2, …, 25 drilled samples per stratigraphic horizon) on the obtained median carbon (a) and oxygen (b) isotope value. The cutoff values are based on whether a carbon isotope value deviates 0.3‰ from the bed‐specific median and estimates the probability that a randomly chosen within‐bed isotope value is therefore not representative for the whole rock value. The cutoff value is chosen to approximate the smallest wiggles used in carbon isotope chemostratigraphic studies (Section 9.1.3). A cutoff value of 1‰ was chosen for bed‐internal δ
18
O, or a 4 °C difference of the fluid, based on the equilibrium fractionation.
Chapter 10
Figure 10.1 Paleogeographic map of discussed and cited T–J boundary sections. Rocks of CAMP are marked in dark red, and the reconstructed CAMP area is colored pale and dark red and taken from
McHone
[2003]. 1: Fundy Basin [
Schoene et al
., 2006;
Blackburn et al
., 2013]; 2: Hartford Basin [
Whiteside et al
., 2010]; 3: Newark Basin [
Whiteside et al
., 2010;
Marzoli et al
., 2011;
Schaller et al
., 2011;
Blackburn et al
., 2013]; 4: Culpeper Basin [
Marzoli et al
., 2011]; 5: Argana Basin [
Deenen et al
., 2010;
Blackburn et al
., 2013]; 6: High Atlas Basin [
Marzoli et al
., 2004]; 7: Northern Calcareous Alps [
Kuerschner et al
., 2007;
Ruhl et al
., 2009, 2011;
Ruhl and Kürschner,
2011]; 8: Pelso Unit, Hungary [
Pálfy et al
., 2001, 2007]; 9: Western Carpathians [
Michalík et al
., 2007, 2010]; 10: Southern Alps [
Galli et al
., 2007;
van de Schootbrugge et al
., 2008;
Bachan et al
., 2012]; 11: Apennines [
van de Schootbrugge et al
., 2008;
Bachan et al
., 2012]; 12: Southern Germany [
van de Schootbrugge et al
., 2008;
Ruhl and Kürschner
, 2011]; 13: Polish Trough [
Pieńkowski et al
., 2012]; 14: Northern Germany [
van de Schootbrugge et al.,
2013]; 15: Danish Basin [
Lindström et al
., 2012]; 16: Southwest Britain [
Hesselbo et al
., 2002;
Korte et al
., 2009;
Clémence et al
., 2010;
Ruhl et al
., 2010]; 17: East Greenland [
McElwain et al
., 1999;
Hesselbo et al
., 2002]; 18: Queen Charlotte Islands [
Pálfy et al
., 2000;
Ward et al.,
2001;
Williford et al
., 2007;
Friedman et al
., 2008]; 19: New York Canyon, Nevada [
Guex et al
., 2004, 2012;
Ward et al
., 2004;
Schoene et al
., 2010;
Bartolini et al
., 2012]; 20: Utcubamba Valley, Peru [
Schaltegger et al
., 2008;
Schoene et al
., 2010]; 21: Arroyo Malo, Argentina [
Damborenea and Manceñido
, 2012;
Percival et al
., 2017].
Figure 10.2 (a) Comparison of Triassic‐Jurassic δ
13
C
carb
, δ
13
C
TOC
, and δ
13
C
wood
data from geographically distributed marine and terrestrial sections, with Hg/TOC ratios and inferred atmospheric
p
CO
2
. Atmospheric
p
CO
2
estimates are based on stomatal density analyses (Astartekløft, Greenland) and δ
13
C values of pedogenic carbonate sequences (Newark and Hartford basins). Note that this compilation is not exhaustive and that many more Triassic‐Jurassic boundary sections have been studied. (b) Highly resolved Triassic‐Jurassic boundary integrated stratigraphic framework based on key European successions studied for δ
13
C
TOC
, palynostratigraphy, and ammonite biostratigraphy.
Figure 10.3 Compilation of continental Triassic‐Jurassic sequences, which have been analyzed for δ
13
C
TOC
and δ
13
C
wood
and which record Central Atlantic Magmatic Province (CAMP) basalt emplacement.
Figure 10.4 Oxygen and carbon isotope values from pristine oysters originating from the earliest Jurassic successions at Lavernock Point, St Audrie’s Bay, and Watchet, United Kingdom, from
Korte et al
. [2009; including results from
van de Schootbrugge et al
., 2007] and bulk organic δ
13
C data from
Hesselbo et al
. [2002], plotted against stratigraphy.
Chapter 11
Figure 11.1 (a) Example of current‐reworked radiolarian sand, today preserved as siliceous limestone, with evidence for cross‐bedding. Pelagic Maiolica Formation, Lower Cretaceous. S. Alps. Italy. (b) Bioturbation features are used as indicator for winnowing in a pelagic setting. Bioturbated pelagic limestone reflects continuous sedimentation, and non‐bioturbated upper part was redeposited by currents. (c) Lower Cretaceous Maiolica Formation, Central Tethys Ocean (S. Alps, Italy). The picture illustrates the change from a white almost pure and up to several dm‐bedded nannofossil limestone (>90% CaCO
3
) to a darker thin‐bedded limestone. Transition is dated as lower Valanginian [
Weissert
, 1979]. (d) Limestone‐chert succession in pelagic limestones of earliest Cretaceous age (Maiolica Formation, S. Alps, N. Italy). Chert layers are interpreted as indicators of better surface water mixing and increased productivity and radiolarian flux into the pelagic sediment.
Chapter 12
Figure 12.1 Paleomap at 66 Ma showing paleogeography and a selection of published global terrestrial and nearshore marine K‐Pg boundary sites (yellow star) and the four K‐Pg boundary sections focused in this chapter (red star). Location of K‐Pg boundary sites has been compiled from
Smit
[1999],
Nichols and Johnson
[2008],
Schulte et al
. [2010],
Vajda and Bercovici
[2014], and
Punekar et al
. [2014]. The location of the Chicxulub impact site is shown as a yellow dot.
Figure 12.2 δ
13
C
carb
, δ
13
C
org
, δ
44/40
Ca, Hg/TOC, and Mo/Al variation patterns for (a) the Højerup section, Stevns Klint. (b) The Bottaccione (Gubbio) section.
Figure 12.3 δ
13
C
carb
, δ
13
C
org
, δ
44/40
Ca, Hg/TOC, and Mo/Al variation patterns for (a) an Um Sohryngkew River section. (b) A Poty quarry drill hole.
Figure 12.4 PAAS‐normalized REE patterns for (a) the Højerup section and (b) Bottaccione section.
Figure 12.5 PAAS‐normalized REE patterns for (a) an Um Sohryngkew River section and (b) drill core samples from a Poty drill hole.
Figure 12.6 (a) Cross‐plot of Ce/Ce* versus Pr/Pr* after
Bau and Dulski
[1996]. Samples in this study show true negative Ce anomaly, except those from the Um Sohryngkew section. (b) log (Ce/Ce*) versus Nd (ppm) plot (Modified from
Wang et al
. [2014]) in which all samples plot above the anoxic field. The K‐Pg layers exhibit much higher Nd values.
Figure 12.7 Mo
EF
versus U
EF
covariation for the Højerup, Bottaccione, Um Sohryngkew, and Poty K‐Pg boundary sections (Based on
Tribovillard et al
. [2012] and
Sosa‐Montes et al
. [2017]), in which Mo
EF
= [(Mo/Al)
sample
/(Mo/Al)
PAAS
] and U
EF
= [(U/Al)
sample
/(U/Al)
PAAS
]. The PAAS composition used is from
Taylor and McLennan
[1985]. The diagonal lines represent multiples of the Mo/U ratio of the present‐day seawater. The orange field represents a general pattern of Mo
EF
versus U
EF
covariation in unrestricted marine trend for modern eastern tropical Pacific (From
Tribovillard et al
. [2012], modified by
Sosa‐Montes et al
. [2017]), and the yellow field (From
Tribovillard et al
. [2012]) represents the particulate shuttle trend in which intense cycling of metal oxyhydroxide occurs within water column.
Figure 12.8 In a δ
202
Hg (MDF)‐Δ
201
Hg (MIF) plot, modified from
Sial et al
. [2016], samples from the K‐Pg layer (spike II) from the Højerup, Bottaccione, and Um Sohryngkew sections lie within the range for volcanogenic Hg. One sample from the K‐Pg layer (spike II), two samples from the spike I, and one from the spike III, from the Bidart section, added for comparison, lie within the range for chondrite/volcanogenic emission. One sample from the spike III of this section lies within the volcanogenic emission field, and another within the sediment, soil, and peat range. Ranges for volcanogenic and chondritic Hg are from
Bergquist and Blum
[2009] and are shown as vertical bars.
Chapter 13
Figure 13.1 Relative sea‐level fluctuations during the Cenozoic era.
Figure 13.2 Profiles of selected isotopes and temperature of Cenozoic era. (a) Oxygen isotope and paleotemperature. (b) Carbon isotope. (c) Sulfur isotopic composition of Cenozoic seawater from marine barites. (d) Strontium isotope. (e) Ca isotopic composition of marine carbonates and phosphates.
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