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2-1-2012 Magnetic records from latest Triassic to earliest Jurassic red beds, Utah and Arizona, and from mid- Pleistocene lake beds, New Mexico Linda Lee Donohoo Hurley
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Linda L. Donohoo-Hurley Candidate
Department of Earth and Planetary Sciences Department
This dissertation is approved, and it is acceptable in quality and form for publication:
Approved by the Dissertation Committee:
Dr. John W. Geissman , Chairperson
Dr. Peter J. Fawcett
Dr. Mousumi Roy
Dr. Tim F. Wawrzyniec
Dr. Spencer G. Lucas
Dr. Fraser Goff
i
MAGNETIC RECORDS FROM LATEST TRIASSIC TO EARLIEST JURASSIC RED BEDS, UTAH AND ARIZONA, AND FROM MID-PLEISTOCENE LAKE BEDS, NEW MEXICO
by
LINDA L. DONOHOO-HURLEY
B.A., Geology, Occidental College, 2000 M.S., Western Washington University, 2004
DISSERTATION
Submitted in Partial Fulfillment of the Requirements for the Degree of
Doctor of Philosophy Earth and Planetary Sciences
The University of New Mexico Albuquerque, New Mexico
December, 2011
ii Dedication
This dissertation is dedicated to my daughter, Kaleigh, husband John, brother Al and mother Juana. I hope this work serves to inspire Kaleigh to achieve the goals she sets in life. John’s support and encouragement have been invaluable through this process.
Al and Juana created a foundation to guide and nurture my curiosity.
! """! Acknowledgments
It is not possible to express my gratitude to all the people who have helped me personally and professionally during my time at the University of New Mexico. In this section I would like to honor the people most responsible for the success of my dissertation. First I would like to thank Dr. John W. Geissman, my advisor and dissertation chair, for all the guidance he has given me. His advice has made me a better scientist. I thank Dr. Peter J. Fawcett for introducing me to Matlab, Dr. Mousumi Roy for career advise, Dr. Tim Wawrzyniec for support, Dr. Spencer Lucas for the introducing me to new rocks, and Dr. Fraser Goff for retirement advice.
I am grateful to many individuals who have helped my research endeavors. I thank Dr. James I. Kirkland for his generosity, Dr. Jerry D. Harris for his hospitality, Dr.
Roberto Molina- Garza for his patience, Dr. Larry Tanner for his wisdom, Dr. Craig
Allen for his ideas, and Dr. Robert Parmenter for access to the Valles Caldera. Gratitude is extended to numerous agencies for funding to pursue this research including IGPP Los
Alamos, Sloan Foundation, AGI Minority Participation Program, New Mexico AGEP, El
Centro de La Raza, University of New Mexico, Department of Earth and Planetary
Sciences alumni scholarship fund, New Mexico Geological Society student Grants, the
Institute for Rock Magnetism at the University of Minnesota, Twin Cites.
I also thank the staff at the LacCore facility and Institute for Rock Magnetism for their help colleting samples, data acquisition, data interpretation, and professional development.
! "#! I am grateful to my former office mates and friends for their advice and listening skills. I thank Dr. Kate Zeigler for all the shared experiences, Ashley Edelman for long talks, and Kym Samuels for a new perspective. I am grateful to Amber Tholl Hawkins
Whittaker for all the breaks we shared. I am grateful to Gilbert Villasenor, AJ Monroy,
Juana Monroy, Missy Halick, Will Woodruff, and Greg Peacock for field and lab assistants.
Finally I thank my husband John and daughter Kaleigh. Their love, support, and laughter keep me grounded.
! #! MAGNETIC RECORDS FROM LATEST TRIASSIC TO EARLIEST JURASSIC
RED BEDS, UTAH AND ARIZONA, AND FROM MID-PLEISTOCENE LAKE
BEDS, NEW MEXICO
by
Linda L. Donohoo-Hurley
B.A., Occidental College, 2000
M.S., Western Washington University, 2004
PhD, University of New Mexico, 2011
ABSTRACT
Paleomagnetic data obtained from the upper Triassic to lower Jurassic strata of the Moenave Formation, southern Utah and northern Arizona, were used to construct a composite magnetostratigraphy and further refine the position of the base of the Jurassic in the southwest U.S.A. The composite magnetostratigraphy provides a chronostratigraphic framework to tie Triassic-Jurassic sedimentation in the southwest
U.S.A to marine strata in the United Kingdom, Turkey, and Italy, and to the Pangean rift history including extrusive igneous rocks, preserved in Morocco, and in the Newark
Basin, northeast U.S.A. In addition, paleomagnetic data from the Moenave Formation were used to calculate a pole position for North America for the latest Triassic to earliest
Jurassic time. A lesser amount of inclination error, flattening factor of 0.78, is record in
! "#! Moenave Formation strata compared to observation from coeval Newark Basin strata.
The new paleomagnetic pole position for North America, corrected for inclination error and 4° of clockwise Colorado Plateau rotation is located at 62.3° N, 68.0° E (A95 = 7.4°,
N = 102). Pole positions from the southwest U.S.A. continue to indicate a westerly pole position for North America at the latest Triassic to earliest Jurassic time.
A mid-Pleistocene lake sedimentary record obtained from the Valles Caldera, northern New Mexico was investigated using rock magnetic and paleomagnetic techniques. Lake sediments span three glacial and two interglacial intervals, MIS 14 –
10. Both detrital and diagenetic phases are preserved in sediment throughout the core.
Preservation of detrital phases indicates well mixed lake conditions were more common during interglacial intervals. Discrete intervals of diagenetic phases indicate anoxic conditions are more common in sediments deposited during glacial intervals. A series of anoxic intervals are identified in sediment deposited during MIS 12 that are closely related to interstadial events characterized by increased Cyperaceae and Juniper pollen counts and increased mean annual temperatures. Paleomagnetic data are mostly normal polarity consistent with Brunhes normal polarity chron. However, paleomagnetic data combined with relative paleointensity records support the presence of three geomagnetic field phenomena 14!/Calabrian Ridge II at ~536 ka, 11! at ~400 ka, and Levantine at
~360-360 ka.
! "##! TABLE OF CONTENTS
CHAPTER 1 INTRODUCTION...... 1
References...... 7
CHAPTER 2 MAGNETOSTRATIGRAPHY OF THE UPPERMOST TRIASSIC
AND LOWERMOST JURASSIC MOENAVE FORMATION, WESTERN UNITES
STATES: CORRELATION WITH STRATA IN THE UNITED KINGDOM,
MOROCCO, TURKEY, ITALY, AND EASTERN UNITED STATES ...... 13
Abstract...... 13
Introduction...... 14
The Triassic-Jurassic Boundary...... 17
Geology...... 22
Methods...... 24
Paleomagnetism, Rock Magnetism, and Magnetostratigraphy...... 27
Discussion...... 34
Conclusions...... 40
Acknowledgments...... 42
Figures...... 43
Table ...... 53
References...... 57
CHAPTER 3 INCLINATION BIAS OF PALEOMAGNETIC DATA FROM THE
UPPERMOST TRIASSIC TO LOWERMOST JURASSIC MOENAVE
FORMATION, UTAH AND ARIZONA: IMPLICATIONS FOR THE TRIASSIC
! "###! TO JURASSIC NORTH AMERICAN APPARENT POLAR WANDER PATH
FOR THE SOUTWEST UNITED STATES ...... 67
Abstract...... 67
Introduction...... 68
Methods...... 72
Results...... 73
Discussion...... 76
Deformation of the Colorado Plateau ...... 78
Conclusions...... 81
Acknowledgments...... 82
Figures...... 83
Tables...... 88
References...... 94
CHAPTER 4 ROCK MAGNETIC CHARACTERISTICS AND
PALEOMAGNETISM OF A MID-PLEISTOCENE SEDIMENT CORE, VALLES
CALDERA, NEW MEXICO ...... 101
Abstract...... 101
Introduction...... 102
Methods...... 105
Results...... 110
Discussion...... 119
Conclusions...... 126
Acknowledgments...... 128
! #$! Figures...... 129
Table ...... 148
References...... 152
CHAPTER 5 CONCLUSION...... 163
Future Research ...... 166
! $! CHAPTER 1
INTRODUCTION
Magnetic records, as preserved in a wide array of sedimentary and igneous rocks, are important archives of both long-term (about millions of years) changes in polarity and short term (few thousands of years or less) field phenomena. The studies presented here use paleomagnetic and rock magnetic techniques to investigate the magnetization record in both relatively old sedimentary rocks and young non-lithified sediments. The research focus address questions regarding the timing of sediment deposition and global correlations of sedimentary sequences, the past position of tectonic plates, and changes in sediment depositional environments as a function of changes to environmental conditions. Chapters 2 and 3 present the magnetization records obtained from the latest
Triassic to earliest Jurassic Moenave Formation, Utah and Arizona. Chapter 4 discusses the results of a rock magnetic and paleomagnetic study of mid-Pleistocene lake sediments from the Valles Caldera, New Mexico. A review of the major findings of this study is presented in Chapter 5.
The latest Triassic to earliest Jurassic Moenave Formation is exposed on and adjacent to the Colorado Plateau of the southwest U.S.A, and it consists of two members, the basal Dinosaur Canyon and upper Whitmore Point member. The basal Dinosaur
Canyon Member disconformably (the Tr-5 unconformity (Tanner and Lucas, 2007)), overlies the Chinle Group. The sub-Springdale unconformity separates the upper
Whitmore Point Member from the Springdale Member of the overlying Kayenta
Formation (Lucas and Tanner, 2006). The transition from mostly fluvial Dinosaur
! "! Canyon to mostly lacustrine Whitmore Point strata indicates a drop in base level
(Kirkland and Milner, 2006; Milner et al., 2006; Tanner and Lucas, 2010) in the latest
Triassic. Upper strata of the Moenave Formation contain Jurassic aged fossils including
Protosuchus, Eubrontes dinosaur ichnotaxon, and conchostracan Bulbilimnadia killianorum. Lower Moenave Formation strata preserve Triassic ichnotaxon Grallator and are lateral correlated to other fossil bearing strata. Biostratigraphic constraints indicate the Triassic-Jurassic boundary is located with the mid-to upper Moenave
Formation (Lucas and Heckert, 2001; Tanner and Lucas, 2007). The objective of the study presented in Chapter 2 is to use magnetostratigraphy to refine the placement of the
Triassic-Jurassic boundary in strata of the southwest US. Four sections of Moenave
Formation strata were combined to build a composite magnetostratigraphy.
Magnetostratigraphic data are interpreted to indicate that the Triassic-Jurassic boundary lies within an about 2 m interval of strata in the middle part of the Whitmore Point
Member. The composite magnetostratigraphy was correlated to several magnetostratigraphic records across the Triassic-Jurassic boundary, including: St.
Audrie’s Bay, UK (Hounslow et al., 2004), Oyuklu, Turkey (Gallet et al., 2007),
Southern Alps, Italy (Muttoni et al., 2010), High Atlas, Morocco (Knight et al., 2004;
Deenen et al., 2010), and Newark Basin, US (Kent and Olsen, 1999). These correlations provide a chronostratigraphic framework to evaluate the timing of global tectonic events at the time of the Triassic-Jurassic transition.
The position of paleomagnetic pole for North America during the latest Triassic to earliest Jurassic time has been debated for over thirty years (Irving and Irving, 1982;
Gordon et al., 1984; Van der Voo, 1993; Molina-Garza et al., 1995; Besse and Courtillot,
! #! 2002; Beck and Housen, 2003; Besse and Courtillot, 2003; Torsvik et al., 2008; Kent and
Irving, 2010). Paleomagnetic pole positions obtained from strata in the southwest U.S.A. are internally consistent and support a more westerly pole position for North America
(Gordon et al., 1984; Ekstrand and Butler, 1989; Van der Voo, 1993; Kodama et al.,
1994; Beck and Housen, 2003; Molina-Garza et al., 2003). On the other hand, paleomagnetic pole positions obtained from extrusive rocks of the Central Atlantic
Magmatic Province and intercalated sedimentary rocks support a more easterly pole positions and smooth poleward translation of North America through the Triassic-
Jurassic boundary (Irving and Irving, 1982; Witte et al., 1991; Kent and Witte, 1993;
Steiner and Lucas, 2000; Besse and Courtillot, 2002, 2003; Torsvik et al., 2008; Kent and
Irving, 2010). Paleomagnetic data obtained from sedimentary strata can be biased by inclination error produced by post depositional compaction/rotation of magnetic particles in the sediment. Inclination error results in an artificial shift to shallower inclinations and thus pole positions to lower latitudes (Kent and Tauxe, 2005). Others have hypothesized that that the discrepancy between pole positions from the southwest and northeast U.S.A. can be resolved by correcting the southwest data for a large magnitude of inclination shallowing (f = 0.54) and a substantial (10° - 15°) clockwise rotation of the Colorado
Plateau (Kent and Olsen, 2008; Tauxe et al., 2008; Kent and Irving, 2010). The focus of
Chapter 3 is a re-examination of the pole positions obtained from latest Triassic to earliest Jurassic rocks of the southwest U.S.A., including the examination of the likely magnitude of inclination error recorded in rocks of the southwest U.SA. The revised latest Triassic/earliest Jurassic pole position for strata of the southwest U.S.A., corrected for a locally derived inclination error and a geologically reasonable magnitude of
! $! Colorado Plateau rotation is still located west of coeval poles from rocks of the eastern seaboard. These data continue to support a westerly paleomagnetic position for North
American plate at the latest Triassic to earliest Jurassic transition.
Chapter 4 summarizes the results of a study of the “environmental magnetic” record of mid-Pleistocene lake sediments of the South Mountain lake obtained in the VC-
3 core drilling experiment in summer, 2004, from the Valles Caldera, northern New
Mexico (Fawcett et al., 2011). The Valles Caldera core preserves ~200 ka of mid-
Pleistocene sedimentation. The age of the core is constrained by an Ar-Ar age determination of 552 ka + 3ka obtained from a basal tephra and pumice layer. Abrupt increases in geochemical proxies including total organic carbon, Si/Ti ratios, and mean annual temperature estimates are interpreted as glacial termination V (ca. 531 ka) and VI
(ca. 423 ka). Detailed rock magnetic investigations of iron-bearing phases yield data that are interpreted to indicate that South Mountain lake sediments preserve detrital and diagenetic phases. In general, the magnetic properties of sediments from glacial intervals tend to have higher proportion diagenetic phases (e.g. pyrrhotite, siderite/rhodochrosite, vivianite, and superparamagnetic magnetite) while sediments from interglacial intervals tent to preserve a higher proportion of detrital phases (e.g. magnetite, titanomagnetite, maghetite, titanomagnetite, and oxyhydroxides). Discrete intervals of diagenetic phases preserved throughout core VC-3 are most highly concentrated in sediments deposited during MIS 13, 12 and 11. The diagenetic intervals preserved in sediment deposited during MIS 12 occur at the same time or slightly after interstadial intervals characterized by increased Cyperaceae and Juniper pollen and increased mean annual temperatures.
Most sediment samples from Valles Caldera sediment yield paleomagnetic directions that
! %! are consistent with the normal polarity geomagnetic field expected for the mid-
Pleistocene. Three intervals of anomalous paleomagnetic directions combined with reduced relative paleointensity values support the interpretation that these intervals represent geomagnetic field phenomena, at about 536 ka, about 400 ka, and about 360-
370 ka observed in records from Calabrian Ridge in the western Atlantic Ocean, lavas from La Palma, Canary Islands, and the West Eifel volcanic field in Germany (Schnepp and Hradetzky, 1994; Langereis et al., 1997; Singer et al., 2002; Lund et al., 2006; Singer et al., 2008).
This research will be submitted to peer review journals for publication. Chapter
2, Magnetostratigraphy of the uppermost Triassic and lowermost Jurassic Moenave
Formation, western United States: Correlation with strata in the United Kingdom,
Morocco, Turkey, Italy, and eastern United States, has been published in the Geological
Society of America Bulletin in 2010. Chapter 3, Inclination bias of the latest Triassic to earliest Jurassic Moenave Formation, Utah and Arizona, will be submitted to
Geophysical Journal International. Chapter 4, Rock magnetic characteristics and paleomagnetism of a mid-Pleistocene sediment core, Valles Caldera, New Mexico, is in preparation for Geochemistry, Geophysics, Geosystems. Chapters 3 and 4 will be submitted in the upcoming months.
The work presented here was accomplished with the help of several individuals.
Measured sections of Moenave Formation strata were obtained by Spencer Lucas and
Larry Tanner. Sampling of the Moenave Formation was carried out over several years with the help of the following people: John Geissman, Kate Zeigler, Roberto Molina-
Garza, Gilberto Villasenor, Alfred J. Monroy, Missy Halick, Andrew Heckert, Jerry
! &! Harris, Jim Kirkland, and the Bureau of Land Management St. George, Utah, Field office. Nick George, Will Woodruff, and Greg Peacock assisted with sample preparation.
Linda Hurley obtained, analyzed, and interpreted the magnetic and rock magnetic data for these samples. Chapters 2 and 3 greatly benefited from discussion with and improvements suggested by John Geissman, Spencer Lucas, Mousumi Roy and Tim
Wawrzyniec.
The Valles Caldera core was extracted from the Valles Caldera Preserve with the assistance of DOSECC, the LacCore and Lake Repository Center (LRC) at the University of Minnesota, Twin Cities, Peter Fawcett, Fraser Goff, Craig Allen, Tim Wawrzyniec,
Amy Ellwien, Catrina Johnson, and Linda Hurley. The core is housed in Minnesota at the LRC, where Peter Fawcett and Linda Hurley took U-channel samples of the core and transported them to the University of New Mexico. Linda Hurley collected most of the individual samples and conducted the paleomagnetic analysis with some help from Will
Woodruff and Greg Peacock. Linda Hurley analyzed and interpreted the Valles Caldera rock magnetic and paleomagnetic data. Chapter 4 greatly benefited from discussion with and improvements suggested by John Geissman, Peter Fawcett, and Fraser Goff.
Paleomagnetic and rock magnetic techniques can be applied in a very broad range of investigations in the geosciences. Iron bearing phases in sedimentary successions preserve details of the ancient geomagnetic field, often with great fidelity. The three sets of magnetic records reported here address the timing of Moenave Formation sedimentation, the position of North America from latest Triassic to earliest Jurassic time, and the depositional environment of the mid-Pleistocene Valles Caldera lake sediments as influenced by regional and more global climate changes during the Pleistocene.
! '! References
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! "#! CHAPTER 2
MAGNETOSTRATIGRAPHY OF THE UPPERMOST TRIASSIC AND
LOWERMOST JURASSIC MOENAVE FORMATION, WESTERN UNITED
STATES: CORRELATION WITH STRATA IN THE UNITED KINGDOM,
MOROCCO, TURKEY, ITALY, AND EASTERN UNITED STATES
Abstract
A composite magnetostratigraphy based on the magnetic polarity data from four sections of the uppermost Triassic and lowermost Jurassic Moenave Formation, Utah and
Arizona, U.S.A.A, can be correlated to a magnetostratigraphy derived from the marine successions at Saint Audrie’s Bay (UK), Oyuklu, Turkey, and the Southern Alps, Italy, and to the nonmarine sections in Morocco, northern Africa, and the Newark Basin, eastern North America, all deposited across the Triassic–Jurassic boundary. Our proposed correlation provides a stratigraphic framework to tie Triassic–Jurassic sedimentation in the American Southwest to the marine UK, Turkey, and Italy sections, and to the Pangea rift history, including extrusive igneous rocks, preserved in Morocco and in the Newark
Basin. The Moenave polarity record is characterized by mostly normal polarity, as is consistent with other polarity records across the Triassic–Jurassic boundary, and is interrupted by at least two well-defined reverse-polarity magnetozones. On the basis of available paleontologic information, we interpret the oldest well-defined, reverse-polarity magnetozone, M2r of the Moenave Formation, to correlate with either SA5n.3r or
SA5n.2r of the Saint Audrie’s Bay record, J- of the Oyuklu record, one of the reverse magnetozone in “intermediate unit” in Morocco, and with an unidentified interval of
13! reverse magnetization in the southern Alps and Newark basin. The youngest reverse magnetozone of the Moenave Formation, M3r, is the first Jurassic magnetozone identified in any of these records. Magnetostratigraphic correlations and marine biostratigraphic information support placement of the Triassic–Jurassic boundary in the middle to upper Whitmore Point Member of the Moenave Formation, the Lias Group of the Saint Audrie’s Bay section, the chert-rich limestone of the Oyuklu section, above the
Zu Limestone in Italy, and in the central Atlantic magmatic province extrusive zone in the Morocco and the Newark records.
Introduction
The interval of time represented by the transition from the Late Triassic to the
Early Jurassic is an important one in Earth history, because it is associated with substantial faunal and floral evolution, including major marine and terrestrial extinctions and voluminous magmatism related to the opening of the Atlantic Ocean (Hesselbo et al.,
2002; Knight et al., 2003; Olsen et al., 2003; Marzoli et al., 2004; Tanner et al., 2004;
Hesselbo et al., 2007; Kozur and Weems, 2007; Lucas and Tanner, 2007b; Palfy et al.,
2007; Tanner and Lucas, 2007; Whiteside et al., 2007; Tanner, 2008; Cirilli et al., 2009;
Lucas and Muttoni et al., 2009; Deenen et al., 2010). Our understanding of the temporal relationships of events near the Triassic–Jurassic (T-J) boundary, such as faunal extinction and the global effects of extrusion of the Pangean central Atlantic magmatic province, is limited because there is no single biostratigraphic marker to define the T-J boundary in both marine and terrestrial records.
14! Biostratigraphic and magnetistratigrpahic records can be combined to identify the stratigraphic placement of the Triassic-Jurassic transition. Biostratigraphic records in the southwest U.S.A. include footprints, conchostracans, and palynomorphs and suggest the
T-J boundary lies within the Moenave Formation of the Glen Canyon Group (Lucas and
Heckert, 2001; Cornet and Waanders, 2006; Lucas and Tanner, 2007a; Tanner and Lucas,
2007; Tanner and Lucas, 2009a; Kozur and Weems, 2010). In the UK, the Triassic–
Jurassic section at Saint Audrie’s Bay (Hounslow et al., 2004) is of particular importance, because it provides the most complete magnetostratigraphic record in marine strata associated with the long-used marine biostratigraphic definition of the T-J boundary, the lowest occurrence of the Jurassic ammonite Psiloceras planorbis. However, the T-J boundary at Saint Audrie’s Bay may occur within a 6-m-long interval between the highest occurrence of conodonts and the lowest occurrence of the ammonite P. planorbis
(Hounslow et al., 2004). In Oyuklu, Turkey, the T-J boundary is identified within the chert-rich limestone by the highest occurrence of conodonts. At the Italcementi Quarry section in the Southern Alps, Italy, which includes upper Rhaetian and lowermost
Hettangian shallow-marine carbonate rocks, Muttoni et al. (2009) associate the T-J boundary interval with the acme of Kraeuselisporites reissingeri, which is within the interval of a negative carbon isotope excursion. The record obtained from the High Atlas section in Morocco (Knight et al., 2003; Marzoli et al., 2004; Knight et al., 2004; Deenen et al., 2010) contains robust magnetic polarity data and isotopic age determinations associated with the central Atlantic magmatic province eruptions, but it has not provided biostratigraphic information. The T-J boundary in the Moroccan record lies within the
“upper unit.” In eastern North America, Newark Supergroup strata and intercalated
15! central Atlantic magmatic province extrusive rocks in the Newark Basin also span the T-J boundary, as indicated by biostratigraphic and isotopic age data from the central Atlantic magmatic province extrusive zone (Cornet and Olsen, 1985; Whiteside et al., 2007). In addition, Newark Supergroup strata provide the thickest known stratigraphic section across the boundary (Kent and Olsen, 1999, 2008). An initial placement of the T-J boundary and correlation among all of these sections can be obtained by examining the biostratigraphic data and isotopic age estimates (Figure 2.1). However, inconsistencies in chronostratigraphic data among these sections exist, and hinder our ability to precisely locate and correlate the T-J boundary.
Several paleontological criteria have been proposed for defining the T-J boundary, including marine events such as the highest occurrence of conodonts, the lowest occurrence of psiloceratid ammonites, and terrestrial events such as pollen turnover (the end-Triassic event), and crurotarsan extinction near the boundary (Lucas and Tanner, 2007a, 2007b). The relative synchrony of each proposed definition cannot be assessed without an independent chronostratigraphic framework. Magnetostratigraphy is a powerful method for establishing an independent record of geologic time (Opdyke and
Channell, 1996). Polarity records of the geomagnetic field are global and independent of processes such as widespread magmatism, extinction, or climate change. Nevertheless, magnetostratigraphic records may be compromised, especially in fluvial and lacustrine environments, due to hiatus, variable sedimentation rates, and the ability of sediments to accurately record the geomagnetic field, and post depositional diagenetic processes.
Therefore, it is important to study multiple sections to establish consistency in polarity sequences, to acquire a high-resolution sampling, to assess the timing of characteristic
16! remanence magnetization acquisition, to perform field and laboratory tests for remanence stability, and to fully integrate magnetostratigraphic interpretations with biostratigraphic observations.
With sufficiently detailed sampling, the magnetostratigraphic record of the
Moenave Formation polarity record can be better defined and more completely correlated to the Saint Audrie’s Bay, Oyuklu, Southern Alps of Italy, High Atlas, and Newark polarity records. Paleomagnetic data obtained from four previously unexamined, well- exposed sections of Moenave Formation strata in southwestern Utah and northern
Arizona were sampled to test this hypothesis (Figure 2.2). The first paleomagnetic study of the Moenave Formation by Ekstrand and Butler (1989), in the Vermillion Cliffs of northern Arizona, indicated that the Moenave Formation is dominated by normal polarity.
A subsequent study by Molina-Garza et al. (2003) in the nearby Echo Cliffs of northern
Arizona supported the conclusion that these rocks yield mostly normal-polarity magnetizations. Both studies, however, reported one thin, reverse-polarity magnetozone near the top of the Moenave Formation. Additional short, reverse-polarity magnetozones in both the Dinosaur Canyon and Whitmore Point members of the Moenave Formation are reported here and used to make correlations with magnetozones of coeval strata in
UK, Turkey, Italy, Morocco, and eastern North America.
The Triassic-Jurassic Boundary
In the absence of a global stratotype for the base of the Jurassic period, several biostratigraphic events including, but not limited to, crurotarsan (“the codont”) extinction, changes in terrestrial pollen, and the appearance of smooth-shelled
17! psiloceratid ammonites, have been used to identify the T-J boundary (Gradstein et al.,
2004; Hounslow et al., 2004; Kent and Olsen, 1999, 2008; Lucas and Tanner, 2007b;
Whiteside et al., 2007; Lucas, 2009; van de Schootbrugge et al., 2009). The gap in time between different biostratigraphic indicators has resulted in a stratigraphic interval over which the T-J boundary could be defined. The Triassic-Jurassic transition zone interval could be as thin as 6 m in marine strata, ~10 m on the Colorado Plateau, and possibly up to several hundred meters in the Newark rift-basin strata. A distinctive carbon isotope excursion has also been identified in this T-J boundary interval (Palfy et al., 2001;
Hesselbo et al., 2002; Kuerschner et al., 2007; Palfy et al., 2007; Muttoni et al., 2009).
The !13C isotopic record has an initial negative excursion, on the order of ~4‰, in the latest Triassic, a slight recovery, and a main negative excursion that begins near the T-J boundary and continues into the Jurassic. The isotopic excursion has been interpreted to reflect episodic central Atlantic magmatic province volcanism associated with the opening of the Atlantic Ocean. A U/Pb zircon isotopic age determination provides an age estimate for Newark central Atlantic magmatic province flows near the T-J boundary of
~201.0 ± 2 Ma, and a 40Ar/39Ar age determination of 199.8 ± 1.8 Ma has been obtained on a central Atlantic magmatic province flow in Morocco (Kent and Olsen, 1999; Knight et al., 2003; Olsen et al., 2003; Knight et al., 2004). A more exact resolution of the absolute age of the T-J boundary is hindered by differences in age estimates calculated in different isotopic systems (Lucas and Tanner, 2007b) and the precision of the methods.
The working definition of the T-J boundary has long been the lowest stratigraphic occurrence of the psiloceratid ammonite Psiloceras planorbis (Gradstein et al., 2004;
Lucas and Tanner, 2007b) defined in the marine Saint Audrie’s Bay section. The
18! youngest definitive Triassic biostratigraphic marker in the Saint Audrie’s Bay record is the highest occurrence of conodonts. Therefore, the T-J boundary in the southern UK record must lie in the ~6-m-thick interval between the highest occurrence of conodonts and the lowest occurrence of P. planorbis (Figure 2.1). The main !13C isotopic excursion also lies in this 6 m interval (Hesselbo et al., 2002).
Recently, the Subcommission on Jurassic Stratigraphy of the International
Commission on Stratigraphy (ICS) proposed placing the base of the Jurassic slightly lower, at the lowest stratigraphic occurrence of the ammonite Psiloceras spelae, based on a stratotype at Kuhjoch, Austria. Defining the base of the Jurassic by such criteria is not particularly useful in this study because there is no magnetostratigraphic record for the
Kuhjoch section, and the ammonite P. spelae is not present in any of the records considered here. However, we adhere to using P. spelae as the base of the Jurassic and note that P. spelae is stratigraphically lower than P. planorbis but higher than other biomarkers used to interpret the T-J boundary.
In Oyuklu, Turkey, conodonts are the most prevalent biostratigraphic marker
(Gallet et al., 2007). The highest occurrence of conodonts is at the top of the “limestone with gray chert” interval ~30 m above the base of the sections (Figure 2.1). The T-J boundary must lie stratigraphically above this interval. In the Italcementi Quarry section of the Southern Alps, the Triassic–Jurassic boundary is inferred to occur at the disconformity between the Zu Limestone, which exposes the highest occurrence of conodonts, and the Malanotte Formation, which contains the lowest occurrence of
Hettangian pollen (i.e., Classopollis torosus, Gliscopollis meyeriana,
Microreticulatisporites fuscus, Callialasporites spp., Retitriletes austroclavatidites, R.
19! clavatoides, Tsugapollenites pseudomassulae, and rare occurrences of Cerebropollenites macroverrucosus) (Muttoni et al., 2009). However, the stratigraphic boundary between the Zu Limestone and the Malanotte Formation predates the lowest occurrence of the ammonite P. spelae (Muttoni et al., 2009) and is therefore older than the Triassic–
Jurassic boundary.
In the nonmarine Newark Supergroup strata of eastern North America, the T-J boundary lies in the sequence of rocks deposited after the onset of central Atlantic magmatic province volcanism (Lucas and Tanner, 2007b; Lucas, 2009; Kozur and
Weems, 2010). The Newark pollen event, preserved in strata of the Passaic Formation below the central Atlantic magmatic province extrusive zone in the Newark Basin, was previously used to estimate the T-J boundary (Cornet and Olsen, 1985; Fowell and Olsen,
1993; Olsen et al., 2003). The upper Passaic Formation, Newark Basin, contains
Eubrontes dinosaur tracks similar to tracks of Norian or Rhaetian age found in European sections (Litwin, 1991; Hounslow et al., 2004; Lucas and Tanner, 2007b; Whiteside et al., 2007; Lucas, 2009). A reexamination of the palynostratigraphy and the new data provided by conchostracan biostratigraphy (Kozur and Weems, 2010) both lend further support to a latest Triassic age for the upper strata of the Passaic Formation, and indicate that the T-J boundary is within the central Atlantic magmatic province extrusive zone.
Therefore, the Newark pollen event must be considered to be in uppermost Triassic strata
(Figure 2.1).
In the nonmarine High Atlas record of Morocco, geochemical and magnetostratigraphic data have been used to support placement of the T-J boundary within the “upper unit” of the central Atlantic magmatic province flows (Marzoli et al.,
20! 1999; Knight et al., 2004). A new study by Deenen et al. (2010) on the red bed sedimentary sequence from the Bigoudine Formation in the Agrana Basin, Morocco, uses astrochronology, magnetostratigraphy, basalt geochemistry, and geobiology to establish a correlation between an interval of reverse polarity identified in the upper Bigoudine
Formation and the short reverse-polarity magnetozone E23r in the upper Passaic
Formation in the Newark Basin. Previous work in the Moroccan section identified two intervals of reverse polarity within the “intermediate unit” basalt flow sequence (Knight et al., 2004). However, because each reverse-polarity magnetozone was identified at completely separate sections, it is possible that the two intervals of reverse polarity are actually the same event (Knight et al., 2004; Marzoli et al., 2004; Deenen et al., 2010).
Isotopic age estimates (40Ar/39Ar step-heating analyses) on the High Atlas, Moroccan basalt flows provide an age estimate of 199.8 ± 1.8 Ma for the “upper unit” (Knight et al.,
2004) (Figure 2.1). Recently obtained paleomagnetic data from the North Mountain
Basalt, the record of initial central Atlantic magmatic province magmatism in the Fundy
Basin of Nova Scotia, have been interpreted by Kent et al. (2009) to record the same geomagnetic field behavior as the Lower and Intermediate basalt sequences of the
Morocco section, as reported by Knight et al. (2004). For the North Mountain Basalt,
Kent et al. (2009) reported no evidence of the field excursions that have been identified in the “intermediate unit” of the Moroccan section.
As discussed above, the biostratigraphic record for strata of the northern outcrop belt of the Moenave Formation suggests that the T-J boundary is within the Whitmore
Point Member of the Moenave Formation. Thus, the T-J boundary must lie within the middle to upper strata of the Whitmore Point Member between unambiguously Triassic
21! and Jurassic fossil-bearing strata (Figure 2.1). Identification of the T-J boundary in all six records (Saint Audrie’s Bay, Turkey, Italy, Morocco, Newark Basin, and the American
Southwest) as well as its correlation among these regions, is complicated by the fact that a normal-polarity magnetozone of long duration includes all the local T-J boundary indicators, making accurate correlation of any short-lived reverse- polarity magnetozones in these sections of paramount importance (Molina-Garza et al., 2003; Hounslow et al.,
2004; Knight et al., 2004; Gallet et al., 2007; Kent and Olsen, 2008; Muttoni et al., 2009).
Short-lived reverse-polarity magnetozones near the Triassic–Jurassic boundary have been identified at each Moenave locality. The purpose of this work is to establish a more robust polarity record for the Moenave Formation and to propose a correlation of the reverse magnetozones near the T-J boundary in western North America, southern UK, southwest Turkey, Southern Alps, Italy, northwestern Africa, and eastern North America sections deposited across the Triassic–Jurassic boundary.
Geology
Harshbarger et al. (1957) named the Moenave Formation as the basal formation of the Glen Canyon Group. The Moenave Formation includes two members, the Dinosaur
Canyon and Whitmore Point (Lucas and Tanner, 2006; Tanner and Lucas, 2007; Tanner and Lucas, 2009b). The basal Dinosaur Canyon Member is separated from the underlying
Upper Triassic strata of the Owl Rock Formation of the Chinle Group by an unconformity that decades ago was termed the J-0 unconformity, when it was believed to coincide with the T-J boundary (Pipiringos and O’Sullivan, 1978), but is better referred to as the Tr-5 unconformity, a within-Triassic hiatus (Lucas and Tanner, 2006; Tanner
22! and Lucas, 2007). The Whitmore Point Member conformably overlies Dinosaur Canyon strata. The Springdale Sandstone Member of the Kayenta Formation disconformably overlies the upper Whitmore Point strata at the sub-Springdale unconformity (Tanner and
Lucas, 2007). The Moenave Formation is exposed in a northern outcrop belt, between southern Utah and northernmost Arizona, and a southern outcrop belt, in north-central
Arizona (Figure 2.2).
Exposures of the Moenave Formation in the southern outcrop belt (north-central
Arizona) are dominated by intermixed eolian and fluvial facies of the Dinosaur Canyon
Member (Harshbarger et al., 1957; Tanner and Lucas, 2007). The northern outcrop belt
(southern Utah to northern Arizona) contains the basal Dinosaur Canyon Member and the upper, mostly lacustrine facies Whitmore Point Member (Tanner and Lucas, 2009b).
Strata of the Whitmore Point Member are not preserved in the southern outcrop belt.
These lacustrine strata were deposited in one large lake (“Lake Dixie”) or in a mosaic of smaller lakes in the area between Saint George and Kanab, Utah (Kirkland et al., 2002;
Milner et al., 2006; Tanner and Lucas, 2009b).
Abundant dinosaur tracks assigned to the ichnogenera Eubrontes and Grallator are preserved in lower Whitmore Point strata at the Dinosaur Discovery Site at Johnson
Farm in Saint George, Utah, near the Washington Dome locality (Kirkland et al., 2002;
Milner et al., 2006). The lowermost strata of the Whitmore Point Member also yield palynomorphs, conchostracans (Euestheria and Bulbilimnadia killianorum), collected at the Potter Canyon locality by previous workers, and tetrapod tracks (Eubrontes) consistent with a latest Triassic age (Cornet and Waanders, 2006; Kozur and Weems,
2010). Upper Whitmore Point strata yield Jurassic conchostracans, such as Bulbilimnadia
23! kilianorum. Strata of the uppermost Dinosaur Canyon Member in the southern outcrop belt (not sampled in this study) that are laterally equivalent to the upper Whitmore Point strata in the northern outcrop belt contain tetrapod body fossils (especially the crocodylomorph Protoscuhus) indicative of an Early Jurassic age (Lucas and Heckert,
2001; Lucas, 2009). Previously published biostratigraphic evidence thus suggests that the
T-J boundary is located in the Whitmore Point Member.
Methods
Moenave Formation localities in the northern outcrop belt, on and off the
Colorado Plateau, were chosen because of stratal continuity with underlying and overlying formations and essentially complete exposure of both the fluvial facies of the
Dinosaur Canyon and lacustrine facies of the Whitmore Point members. Four localities in
Moenave Formation strata were sampled between Saint George, Utah, and Pipe Springs,
Arizona: Leeds (average strike/dip (s/d): 50°/30° SE), Warner Valley (s/d: 62°/36° SE),
Washington Dome (s/d: 227°/12° NW), and Potter Canyon (s/d: 256°/6° NW) (Figure
2.2). The Leeds, Warner Valley, and Washington Dome localities are within the western transition zone of the Colorado Plateau. The average thickness of the Dinosaur Canyon
Member is ~45 m. Average thickness of the Whitmore Point Member is ~20 m. The fourth locality (Potter Canyon) is on the Colorado Plateau and consists of ~10 m of
Dinosaur Canyon strata and ~40 m of Whitmore Point strata.
Competent beds in Moenave Formation strata were sampled using a standard gas- powered drill. Fine-grained, less competent intervals, including hematitic mudstones and claystones, were block sampled. One site equals one stratigraphic horizon sampled as
24! seven to ten independently oriented cores samples or three to seven blocks. Core specimens (2.2 ! 2.5 cm right cylinders) and cubes prepared from oriented blocks
(typically by dry sawing) were demagnetized with chemical, alternating field (AF), and thermal methods. Specimens were prepared for chemical demagnetization by cutting slits into the sides of core, thus increasing the surface area (Henry, 1979). These high surface- area specimens were sequentially immersed in 100% reagent grade HCl for 1 min, 3 min,
10 min, 60 min, 12 h, and 24 h (Figure 2.3B). The AF demagnetization was carried out to
90 mT. Thermal demagnetization was conducted at 100 °C steps from natural remanent magnetization (NRM) to 400 °C, followed by 25 °C to 50 °C steps to 630 °C, and then
10 °C to 15 °C steps to 685 °C (Figures 2.3 and 2.4). Specimens were also treated with a combination of AF, up to 90 mT, and thermal methods (Figure 2.4B). Remanence measurements were made on a three-axis, DC SQUID 2-G Enterprises 760R magnetometer at the University of New Mexico, and 13 sites were measured in the paleomagnetism laboratory at Occidental College. Characteristic remanent magnetization
(ChRM) vector components were estimated using principal component analysis with
PaleoMag X software (Kirshvink, 1980; Jones, 2002). Most often, demagnetization vectors trend to the origin, so lines were estimated on anchored data points between
~600 °C and 685 °C. Seventy-two, out of 551, measured ChRM directions were rejected from site-mean calculations because mean angular deviation values exceeded 15° or the
ChRM directions deviated from directions from other samples at that site and thus were not consistent with expected directions for the latest Triassic–earliest Jurassic time period for North America. Site-mean directions for 85 individual sites were estimated using
Fisher statistics (Fisher, 1953) from linear data on 479 ChRM directions (Figure 2.5).
25! Grand mean directions, calculated from those ChRM directions for both normal and reverse polarity, are consistent with expected directions from previous Moenave
Formation studies and slightly different from other T-J aged rocks (Figure 2.6) (Ekstrand and Butler, 1989; Van der Voo, 1993; Besse and Courtillot, 2002; Molina-Garza et al.,
2003; Torsvik et al., 2008). Virtual geomagnetic poles were determined for 79 (out of 89) site-mean directions that have cones of confidence less than or equal to 15° and precision parameters greater than 10 (Table 2.1).
The magnetic mineralogy of the Moenave Formation samples was investigated using three-component thermal demagnetization of isothermal remanent magnetization
(IRM), IRM acquisition and backfield demagnetization, and chemical leaching followed by IRM acquisition. Samples were exposed to decreasing DC magnetic fields of 3 T, 0.3
T, and 0.03 T along orthogonal axes following the method of Lowrie (1990) (Figure 2.7).
Petrographic analysis in transmitted and reflected light was also carried out on polished thin sections of natural and chemically leached Moenave samples. As discussed below, thin sections in both transmitted and reflected light show hematite as both authigenic and detrital phases (Figure 2.8).
A polarity record was compiled for each stratigraphic section by interpreting horizons yielding high northern (positive), virtual geomagnetic pole (VGP) latitudes to be of normal polarity and those yielding high southern (negative) VGP latitudes to be of reverse polarity (Table 2.1). The reverse-polarity magnetozones identified in the
Whitmore Point Member were used as tie points for correlation between the four localities, because magnetozones represent time-equivalent strata. The thickness of each
26! member was averaged to build a composite magnetostratigraphy of the Moenave
Formation (Figure 2.9).
Paleomagnetism, Rock Magnetism, and Magnetostratigraphy
The NRM intensities for Moenave Formation strata are typically ~70 mA/m, and more than 90% of the samples have two well-defined components. The first removed component is typically north-directed, with a moderate positive inclination. The first removed component usually unblocks between 300 °C and 600 °C (Figures 2.3 and 2.4) and is interpreted as an overprint magnetization of variable coercivity. These directions do not match expected directions for the Triassic–Jurassic and are therefore interpreted to be considerably younger. The geologic importance of these overprints is not relevant to this study, because paleomagnetic studies of several red beds, of a range of geologic ages, have demonstrated the persistence of secondary magnetizations, of high laboratory unblocking temperatures carried in hematite, that appear to be geologically very young
(Dunlop and Stirling, 1977; Ashby et al., 2005; Weil et al., 2006).
The second isolated component, identified in ~87% of the samples, is north directed and shallow (Figure 2.3), thus yielding high northern-latitude VGPs. The north directed and shallow component is partially isolated by chemical demagnetization and is interpreted as the ChRM (Figure 2.3B). In 8% of the samples, the second isolated component yields high southern VGP latitudes and is interpreted as a ChRM of reverse polarity (Figure 2.4). In some cases, overprint vector components can make it difficult to isolate the ChRM directions. However, results of remanence stability tests, including reversal and tilt tests, support the conclusion that the ChRM directions are primary. In
27! addition, the overall data set reveals a strong consistency with data from previous
Moenave Formation studies and a slight deviation from other expected directions for rocks of Late Triassic to Early Jurassic age (Figure 2.6) (Ekstrand and Butler, 1989; Van der Voo, 1993; Besse and Courtillot, 2002; Molina-Garza et al., 2003; Torsvik et al.,
2008).
Thermal demagnetization of samples treated with three-component IRM reveals a steady decrease in magnetization from 100 °C to ~620 °C. Most of the IRM unblocking is observed in the 3.00 T component as a sharp decrease in magnetization over the laboratory unblocking temperature interval between ~620 °C and 680 °C consistent with fine- to coarse-grained hematite present in these rocks. The 0.30 T component carries
~50%–25% of the total magnetization and also unblocks at ~680 °C. The 0.03 T component is always much less than 25% of the total IRM, and at least 60% of this component is unblocked by 580 °C (Figures 2.7A and 2.7B). Well-indurated, dark-red to purple samples were selected for treatment with chemical demagnetization to assess the degree to which pigmentary hematite could be removed (Collinson, 1967). Samples changed color to white or beige after a cumulative total of ~24 h of soaking in reagent- grade HCl, thus clearly demonstrating the removal of fine-grained, authigenic hematite.
These samples were then subjected to progressive acquisition of IRM to a maximum field of 3 T. Saturation IRM intensities range from 2.75E-4 mA/m to 1.68E-2 mA/m, and the
IRM acquisition curves reveal a mixture of both low coercivity (possibly magnetite) and higher coercivity (presumably detrital hematite) phases (Figures 2.7C and 2.7D). These samples reach 90%–95% of their total IRM intensity at 1 T and continue to increase in
IRM intensity when they reach saturation at ~3 T, a behavior consistent with the presence
28! of detrital hematite. In backfield demagnetization, coercivity of remanence values is ~0.2
T, which is consistent with hematite as the principal magnetic phase. The presence of both authigenic and detrital hematite is supported by petrographic observations
(Figure 2.8). Each section sampled in the Moenave Formation is dominated by normal polarity (Figure 2.9). Notably, thin reverse-polarity magnetozones are observed at each locality. The results from each locality are now described.
Two intervals of reverse polarity are observed at the Leeds locality (Table 2.1).
The oldest interval containing reverse-polarity magnetizations, observed in the Dinosaur
Canyon Member (at ~5.3 m depth at site 105), yields dual (or mixed) polarity data. Site
105 is composed of nine independent samples. Seven specimens yield well-behaved
(principal component analysis results with a95 less than 10°), north-directed and shallow demagnetization vectors providing high north-latitude VGPs; two specimens from this site yield south-directed and shallow-inclination magnetizations consistent with high south-latitude VGPs. This site is represented as a half bar on the magnetostratigraphy of the Leeds locality (Figure 2.9) because of its mixed polarity character. Site 105 is the only interval of reverse polarity that is observed in the basal part of the Dinosaur Canyon
Member at any locality studied here.
A second reverse-polarity interval is observed at Leeds locality at site 181 in the
Whitmore Point Member at ~50.6 m from the base (Table 2.1). Site 181 consists of three samples that yield south-directed and shallow-inclination magnetizations consistent with high south-latitude VGPs (Figure 2.9), and three samples were rejected because the demagnetization vectors are poorly defined. A third interval of reverse polarity is observed at ~62 m from the base of the Leeds locality in Whitmore Point strata sampled
29! by site 185. Site 185 is composed of six independent samples. Four samples yield well- behaved, south-directed and shallow-inclination magnetization vectors providing high south-latitude VGPs, and two samples were rejected because the demagnetization vectors are poorly defined. Three sites (154B, 58.5 m; 156B, 60.6 m; and 157B, 63 m), were rejected from the Leeds locality because specimens yield poorly defined demagnetization trajectories, and the poorly defined magnetization components are not consistent with expected directions (Table 2.1).
At the Washington Dome locality, two intervals exhibit reverse-polarity magnetizations—sites 229, at 51.7m, and 232, at 57.3 m (Table 2.1). Site 229 is composed of six independent samples. Four samples record south-directed and shallow- inclination magnetizations consistent with high south-latitude VGPs, and two samples were rejected because demagnetization vectors are poorly defined. Site 232 is composed of 16 independent samples. Five samples display well-behaved, north-directed and shallow-inclination magnetizations, consistent with high north-latitude VGPs. Nine samples record well-behaved, south-directed and shallow-inclination magnetizations consistent with high south-latitude VGPs, and two samples were rejected because the demagnetization vectors are poorly defined. Because site 232 records mixed polarity directions, it is represented as a half bar in the magnetostratigraphy of the Washington
Dome locality (Figure 2.9). Four sites (225, at 48 m; 227, at 50.1 m; 228, at 51 m; and
205, at 52.6 m), were rejected from the Washington Dome locality because specimens yield poorly defined demagnetization trajectories, and the poorly defined magnetization components are not consistent with expected directions.
30! At the Warner Valley locality, only one interval (site 359, at ~50 m in the
Whitmore Point Member) contains magnetizations of reverse polarity (Table 2.1). At site
359, eight independent samples were collected. Specimens from two samples display well-behaved, north-directed and shallow-inclination magnetization vectors and thus yield high north-latitude VGPs. Specimens from two samples display well-behaved, south-directed and shallow inclination vector directions and thus high south-latitude
VGPs. Specimens from four samples yield poorly defined vector components
(Figure 2.9). One site, 356B (46.5 m), was rejected from the Warner Valley locality because samples yield poorly defined demagnetization trajectories, and the poorly defined magnetization components are not consistent with expected directions.
At the Potter Canyon locality, one interval of reverse polarity is identified in the lowermost Whitmore Point Member at ~11.9 m (site 427B), and one ~9-m-thick interval of reverse polarity is identified in the upper Whitmore Point Member between ~35 and
44 m (sites 417, 416, 414, and 413) (Table 2.1). Site 427 is composed of seven independent samples, with specimens from each yielding well-behaved, south-directed and shallow-inclination magnetizations providing high south-latitude VGPs. Sites 417,
416, 414, and 413 each consist of four independent block samples, specimens from each yield well-behaved, south-directed and shallow-inclination magnetizations providing high south-latitude VGPs (Figure 2.9). Two sites (425B, 28.8 m, and 415, 40.1 m), were rejected from the Potter Canyon locality because samples yield poorly defined demagnetization trajectories, and the poorly defined magnetization components are not consistent with expected directions.
31! Our collection of paleomagnetic data from the four localities in Moenave strata results in a total of 56 accepted site-mean directions that are less than two sigma away from the grand mean. After application of structural corrections to the data from all four localities, the estimated mean direction for normal-polarity sites (D = 2.7°, I = 15.9°, a95
= 5.2, k = 15.3, n = 53 sites) and the estimated mean for reverse-polarity sites (D =
174.3°, I = -20.0°, a95 = 16.2, k = 59.2, n = 3 sites) (Figure 2.6) pass a “B” class ("c is
8.98° at the 95% confidence) reversal test of McElhinny and McFadden (1990), after inversion of opposite polarity data to common polarity. The mean directions, after inversion to common polarity, for each locality (Table 2.1) were analyzed using the direction-correction tilt test of Enkin (2003). A positive tilt test result indicates that the magnetization was probably acquired prior to tilting (58.9% unfolding at the 95% confidence interval with a standard error of 12.38). The robustness of this test is compromised given that only four locality mean directions could be used for the analysis.
Correlating reverse-polarity magnetozones defined in the Whitmore Point Member with similar magnetozones in all four sections and averaging the thickness of the Whitmore
Point and Dinosaur Canyon members was employed to construct a composite magnetostratigraphy. The resulting composite section is ~70 m thick. The short intervals of reverse polarity recorded in Whitmore Point sediments represented by samples from sites 183 of Leeds locality, 229 of Washington Dome locality, 359B of the Warner Valley locality, and 427 from Potter Canyon locality are interpreted as representing time- correlative strata. We also interpreted the younger interval of reverse polarity observed in strata from the middle to upper Whitmore Point member including site 185 of the Leeds locality, site 232 of the Washington Dome locality, and sites 417–413 of the Potter
32! Canyon locality as representing time-correlative strata. Next, the thickness of the
Dinosaur Canyon member at the Leeds, Washington Dome, and Warner Valley localities was averaged to obtain a thickness for the Dinosaur Canyon Member. Strata of the
Dinosaur Canyon Member at the Potter Canyon locality were omitted from this calculation because only the upper part of this member is exposed. We then averaged the thickness of the Whitmore Point Member from all four localities to obtain an average thickness for the Whitmore Point Member. Each interval of reverse polarity was also averaged over all four localities to obtain an average thickness for each magnetozone.
The resulting composite magnetostratigraphic section represents an idealized composite in which members may be thicker or thinner than at any specific locality (Figure 2.9).
The constructed section, although not necessarily consistent with the specific stratigraphic thicknesses of members at any of the four localities, offers a straightforward way to compare the magnetostratigraphic record of the northern outcrop belt of the
Moenave Formation with magnetostratigraphic sections in the UK, Turkey, Italy,
Morocco, and eastern North America. The composite magnetostratigraphy shows that most of the Dinosaur Canyon Member of the Moenave Formation is of normal polarity.
The composite record includes an ~0.5-m-thick reverse-polarity magnetozone (M1r) at
~4.5 m above the base of the Dinosaur Canyon Member that, notably, is not well defined.
The contact between the Dinosaur Canyon and Whitmore Point members lies at ~45 m, and is within normal magnetozone M2n, which is ~42 m thick. The Whitmore Point
Member polarity sequence includes two well-defined, reverse-polarity magnetozones,
M2r (~1 m thick) and M3r (~5 m thick) (Figure 2.9).
33! Discussion
The refined paleomagnetic and magnetostratigraphic record for the Moenave
Formation allows improved correlation among strata from the American Southwest, Saint
Audrie’s Bay, UK, Oyuklu, Turkey, Southern Alps, Italy, High Atlas, Morocco, and
Newark Basin, eastern North America. Demagnetization results from the four sections studied here show that the Moenave Formation is dominated by normal-polarity magnetozones, which is consistent with earlier studies of the Moenave Formation as well as the general understanding of the polarity record across the T-J boundary. Accepted site-mean directions of shallow inclination are consistent with an equatorial setting for the American Southwest across the T-J boundary. The correlation of the Moenave record to records of England, Turkey, Italy, Morocco, and eastern North America are now described.
The correlated intervals of reverse polarity in strata of the Whitmore Point
Member were combined to facilitate discussion of the magnetostratigraphic records from our four localities. We also used an average member thickness to obtain a composite magnetostratigraphy for the Moenave Formation (Figure 2.9). The composite magnetostratigraphy shows that the Moenave Formation was deposited during mostly normal polarity (magnetozones M1n, M2n, M3n, and M4n). The oldest interval of reverse polarity, M1r, is identified in Dinosaur Canyon strata by part of site 105, at 5.3 m, at the Leeds locality, although the data from this thin interval, as described above, are less than robust. The subsequent interval of reverse polarity, M2r, is defined in Whitmore
Point strata by sites 181, at 50.6 m, at the Leeds locality, 229, at 51.7 m, at the
Washington Dome locality, part of site 359B, at 49.8 m, at the Warner Valley locality,
34! and 427B, at 11.9 m, at the Potter Canyon locality. The youngest interval of reverse polarity, M3r, is defined in Whitmore Point strata by sites 185, at 61.9 m, at the Leeds locality, part of site 232, at 57.3 m, at the Washington Dome locality, and 417, 416, 414, and 413, at 34.7 m to 43.7 m, at the Potter Canyon locality.
Rock magnetic data show that the principal remanence carrier in the Moenave sediments is hematite. Three-component (0.03 T, 0.3 T, 3 T) thermal demagnetization experiments show that the IRM is of high coercivity and fully unblocks at 680 °C, the signature unblocking temperature for hematite. Data obtained from chemical leaching of pigment hematite followed by saturation IRM acquisition show that detrital hematite constitutes at least part of the hematite-dominated remanence. Petrographic observations show that authigenic (translucent or pigment) hematite is ubiquitous, although irregularly distributed. Hematite of detrital origin, as coarser opaque grains, is very rare, and is less than 0.1% by volume of detrital grains. Detrital magnetite has not been identified by petrographic inspection. Thin sections prepared from specimens that have been subjected to progressive chemical demagnetization in reagent HCl show little if any remaining authigenic hematite (Figure 2.8).
The ChRM directions from the four Moenave Formation sections, as well as the previous studies by Ekstrand and Butler (1989), are consistent, at least in inclination, with expected directions for Triassic–Jurassic rocks based on several compilations
(Figure 2.6). We interpret the ChRM in the Moenave Formation to be primary, acquired very close to the time of deposition. Normal- and reverse-polarity magnetizations yield a positive reversal test, although the test is not as robust as possible due to dominance by normal polarity. A positive tilt test supports the hypothesis that the ChRM was acquired
35! prior to tilting. Such correspondence is interpreted as suggesting that the remanence acquisition process in most Moenave strata was of sufficient fidelity to record the geomagnetic field, including short polarity chrons, over the Triassic–Jurassic boundary interval. Remanence acquisition occurred over a sufficiently short period of time to provide an internally consistent polarity record (Beck et al., 2003). Possible heterogeneities in the magnetization acquisition process in the oldest part of the Dinosaur
Canyon Member in the northern outcrop belt resulted in a slower rate of remanence acquisition and are reflected in intervals of mixed remanence polarity, which are of lesser quality in comparison to the younger reverse magnetozones. Correlations based on the
M1r magnetozone are not robust and should be further examined in other Dinosaur
Canyon sections.
Biostratigraphic evidence, including conchostracans, pollen, and Eubrontes tracks, all indicate the Triassic-Jurassic transition zone occurs in the upper middle
Whitmore Point Member within magnetozone M3r (Kirkland et al., 2002; Cornet and
Waanders, 2006; Milner et al., 2006; Kozur and Weems, 2010; Lucas et al., 2011).
Upper Dinosaur Canyon strata in the southern outcrop belt contain Jurassic Protosuchus remains (Lucas and Hekert, 2001). These strata are considered to be laterally equivalent to upper Whitmore Point strata in the northern outcrop belt (Lucas and Heckert, 2001;
Lucas and Tanner, 2007a). The exact relationship between the fossil-bearing strata in the southern outcrop belt and the reverse-polarity magnetozones that we have defined from the Whitmore Point Member in the northern outcrop belt is uncertain, because there are no magnetostratigraphic data from the locality where Protosuchus was identified. Bona fide Jurassic conchostracan Bulbilimnadia killianorum was identified in strata of the
36! upper Whitmore Point Member in the northern outcrop belt within the stratigraphic interval characterized by reverse polarity magnetization direction M3r (Lucas et al.,
2011). Fossil-bearing strata in the southern outcrop belt are inferred to be younger than the youngest stratigraphic interval of reverse polarity in the northern outcrop belt, because lowermost Jurassic strata are of normal polarity (Lucas and Heckert, 2001;
Molina-Garza et al., 2003). If the interpretations presented here are correct, then magnetozone M2r is of latest Triassic age and magnetozone M3r is of earliest Jurassic age.
In the previous paleomagnetic studies of the Moenave Formation, one reverse- polarity magnetozone was reported from both the northern (Vermillion Cliffs: Ekstrand and Butler, 1989) and the southern (Echo Cliffs: Molina-Garza et al., 2003) outcrop belts in Moenave Formation strata sampled close to the sub-Springdale unconformity
(Figure 2.2). Given the lateral extent of exposures of the Whitmore Point Member of the
Moenave Formation and the proximity of both reverse magnetozones to the sub-
Springdale unconformity, Molina-Garza et al. (2003) argued that these were the same reverse magnetozone. The reverse-polarity magnetozone identified by Molina-Garza et al., (2003) was inferred as latest Triassic age because it was correlated to reverse-polarity magnetozone E23r of the Newark polarity record. The new composite magnetostratigraphy for the Moenave Formation shows at least two, well-defined, reverse-polarity magnetozones in uppermost Triassic strata. It is likely that previously recognized reverse-polarity magnetozones stratigraphically high in Moenave strata in the northern and southern outcrop belts of the Moenave Formation are equivalent and potentially correlate to either the M2r or M3r magnetozones. Further
37! magnetostratigraphic work is needed in the southern outcrop belt of the Moenave
Formation to clarify the timing of deposition of strata between the northern and southern outcrop belts.
A correlation between the Moenave Formation and the stratigraphic succession at
Saint Audrie’s Bay, UK (Hounslow et al., 2004) is critical for comparing the relative timing of T-J boundary events recorded in both marine and nonmarine strata. We infer that the interval of time represented by reverse-polarity magnetozone M1r of the
Moenave Formation correlates with the interval of time represented by SA5n.1r of the
Saint Audrie’s Bay Record (Figure 2.10). Notably, magnetozone M1r is not as well defined as younger reverse-polarity magnetozones in the Moenave Formation, so this correlation is tentative. The interval of time characterized by reverse-polarity magnetozone M2r of the Moenave Formation must correlate to the interval of time represented by either reverse magnetozone SA5n.3r or SA5n.2r of the Saint Audrie’s Bay record, as both magnetozones lie in strata of demonstrably Triassic age. The interval of time characterized by reverse-polarity magnetozone M3r correlates with reverse magnetozone SA5n.5r of the Saint Audrie’s Bay record as both reverse polarity magnetozones lie in Triassic-Jurassic transition zone strata.
Although M1r is less well defined than other reverse-polarity magnetozones, if we assume that M1r of the Moenave Formation represents a short time interval of true reverse polarity, then it likely correlates with reverse-polarity magnetozone H- of the marine Oyuklu section in Turkey. The interval of time spanning reverse-polarity magnetozone M2r in Triassic strata of the Moenave Formation most likely correlates with latest Triassic magnetozone interval J- of the Oyuklu section (Figure 2.10). A single
38! reverse-polarity magnetozone (BIT5n.1r) has been identified in strata close to the T-J boundary in the Italcementi Quarry section of the Southern Alps. This magnetozone, defined by data from only one sampling interval (Muttoni et al., 2009), may correlate with either M1r in the Moenave Formation.
Magnetozone M1r of the Moenave record correlates with the intervals of reverse polarity within the Bigoudine Formation of the High Atlas, Morocco, polarity record.
Reverse-polarity magnetozone M2r best correlates with the older reverse-polarity magnetozone within the “intermediate unit” of the High Atlas section because both reverse magnetozones lie in Triassic strata (Figure 2.10).
The oldest reverse-polarity magnetozone in the Moenave Formation M1r is correlated to reverse-polarity magnetozone E32r of the Newark record. As with the other correlations using the M1r interval, the correlation is most tentative. It is not possible to correlate the reverse-polarity magnetozones M2r and M3r of the Moenave Formation to any intervals in the Newark record. However, the intervals represented by M2r and M3r must correlate intervals within the CAMP extrusive zone. Significantly, no obvious correlation exists between the Moenave and the Newark magnetostratigraphy beyond the possible correlation of M1r to E23r. Because M3r of the Moenave record appears to occur at or somewhat above the TJB, this reverse polarity magnetozone does not correlate with any specific magnetozone of the Newark polarity record (Lucas et al., 2011).
The proposed correlations of the Moenave magnetostratigraphic record from the southwest U.S.A. with polarity records from the southern UK, southwest Turkey,
Southern Alps, Italy, Morocco, Africa, and eastern North America establish a chronology that can be used to interpret the timing of some events that characterize the Triassic–
39! Jurassic boundary. The oldest well-defined geomagnetic polarity feature observed in the
Moenave record is the latest Triassic-age, reverse-magnetozone M2r, which we correlate with either reverse magnetozone SA5n.3r or SA5n.2r of Saint Audrie’s Bay, UK, magnetozone J- of the Oyuklu record, Turkey, and either reverse-polarity magnetozone of the ‘intermediate unit”, Morocco. An earliest Jurassic reverse-polarity magnetozone,
M3r, in the Moenave record is correlated with SA5n.5r of Saint Audrie’s Bay, and magnetozone J- of the Oyuklu record. The interval of reverse-reverse polarity M3r is not identified in the other records examined in this study (Figure 2.10). The geomagnetic field is characterized by reverse polarity during earliest Jurassic time when terrestrial biostratigraphic records include Bulbilimnadia killianorum conchostracan fauna.
Conclusions
The composite magnetostratigraphic record for the Dinosaur Canyon and
Whitmore Point members of the Moenave Formation of the southwest U.S.A. reported here facilitates correlation with both marine and nonmarine successions deposited across the T-J boundary. These sections include the Penarth and Lias Groups at Saint Audrie’s
Bay, UK, the limestone with red chert, limestone with gray chert, and chert-rich limestone of the Oyuklu record, Turkey, the middle Zu Limestone to the lower Albenza
Formation at the Italcementi Quarry, Southern Alps, Italy, the Bigoudine Formation,
“lower unit, intermediate unit, and upper unit” in High Atlas, Morocco, the Passaic
Formation of the Newark Basin, and central Atlantic magmatic province eruptions of both Morocco and eastern North America. These correlations aid in evaluating the relative timing of events near the T-J boundary in marine and nonmarine strata. The
40! oldest reverse magnetozone in the Dinosaur Canyon Member of the Moenave Formation
(M1r) is defined by mixed polarity results, and this is less robust than younger intervals of reverse polarity in the Moenave Formation. Magnetozone M1r is only used for the most tentative of correlations. Reverse-polarity magnetozone M2r is correlated, on the basis of biostratigraphic information, to either magnetozones SA5n.2r or SA5n.3r of
Saint Audrie’s Bay, magnetozone J- of Oyuklu, and either magnetozone of reverse- polarity of “intermediate unit”, High Atlas, Morocco. , and magnetozone E23r of the
Newark Basin. The youngest magnetozone in the Moenave Formation, M3r, from the middle part of the Whitmore Point Member, is of earliest Triassic age and only correlates to SA5n.5r of Saint Audrie’s Bay. Magnetozone M3r must correspond to a time interval early in the emplacement of CAMP at the lowest occurrence of the Hettangian conchastracan Bulbilimnadia killianorum. If the proposed correlations are viable, then the
T-J boundary interval likely lies within reverse-polarity magnetozones M3r in the
Moenave record, SA5n.5r in the Saint Audrie’s Bay record, and unidentified reverse- polarity intervals in the Oyuklu record, the Italian record, the Morocco record, and within
CAMP of the Newark record (Figure 2.10).
Most of the Moenave Formation, ~58 m (out of ~70 m), between the base of the
Dinosaur Canyon Member and the middle Whitmore Point Member, was deposited in the latest Triassic. The facies change from mostly fluvial Dinosaur Canyon to mostly lacustrine Whitmore Point occurred during the latest Triassic, in response to an inferred drop in base level (Tanner and Lucas, 2009b). The rich fossil assemblages, including but not limited to Eubrontes and Grallator dinosaur tracks, semionotid fish, plant fossils, and invertebrates, are all preserved at the Dinosaur Discovery Site at Johnson Farm in strata
41! that lie ~4 m above the transition from the Dinosaur Canyon Member to lower Whitmore
Point strata. The fossil assemblage preserved at the Dinosaur Discovery Site is therefore of latest Triassic age, rather than Early Jurassic age as previously interpreted by Kirkland et al. (2002), Cornet and Waanders (2006), and Milner et al. (2006). Our new chronostratigraphic interpretation of the events surrounding the T-J boundary interval supports the initiation of central Atlantic magmatic province volcanism, carbon isotope fluctuations, terrestrial mass extinction, and subsequent flora and faunal radiation as latest Triassic events.
Acknowledgments
We are grateful to Kate Zeigler, Roberto Molina-Garza, John Hurley, Larry
Tanner, and Andrew Heckert for thoughtful discussion and field assistance. Thoughtful comments from Mark Hounslow and Ken Kodama, Heinz Kozur, Bulletin Associate
Editor Eric Ferre, and two anonymous reviewers improved the manuscript. This research was made possible through funding from the Sloan Foundation, the Institute of
Geophysics and Planetary Physics at Los Alamos National Laboratory, American
Geological Institute Minority Participation Program, and scholarships from the
Department of Earth and Planetary Sciences.
42! Figures
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;' BG'A/ 2/:
Figure 2.1. Stratigraphic sections of the Moenave Formation, American Southwest, Saint
Audrie’s Bay, UK, Oyuklu, Turkey, Italcementi Quarry (M. Fm.—Malanotte Formation;
A. Fm.—Albenza Formation), Italy, High Atlas (B. Fm.—Bigoudine Formation),
Morocco, and Newark Basin of eastern North America. Biostratigraphic data and isotopic age estimates are given for each section. Vertical scale is in meters. CAMP—central
Atlantic magmatic province.
43! Utah
Leeds Leeds Locality Paria Washington Dome Locality St. George 37.0 o N Warner Valley Locality Fredonia Potter Canyon Locality Lees Ferry
Vermillion Cliffs Eckstrand and Butler 1989
ransition Zone
Basin and Range Province Echo Cliffs Molina-Garza et al., 2003 estern T Arizona Tuba City o Colorado Plateau 36.0 N W 114.0 oooo W 113.0 W 112.0 W 111.0 W
Figure 2.2. The shaded region represents the northern and southern Moenave outcrop belts. Index map of Moenave localities sampled shown by stars: Leeds at 37.202° N,
246.660° E; Washington Dome at 37.009° N, 246.558° E; Warner Valley at 37.016° N,
246.625° E; Potter Canyon at 36.879° N, 247.156° E; with nearby anthropogenic features. Small circles show nearby populated areas. Large circles represent previous sampling localities reported in Ekstrand and Butler (1989, Vermillion Cliffs) and Molina-
Garza et al. (2003, Echo Cliffs).
44! 400oC W, Up 655oC W, Up o W, Up W, Up o 678 C 640oC 12 hrs 60 min 30 min o 665 C N N 675 C N N o o o 675 C 645 C NRM 685 C NRM 615oC 1 1 1 10 mA/m 101mA/m 10 mA/m 10 mA/m o 200oC 200 C NRM o a b c 100 C d 100oC 102AAa, 45.6 m, 102AGa, 45.6 m, 211B, 37.5 m, 211C, 37.4 m NRM Dinosaur Canyon Mbr Dinosaur Canyon Mbr Dinosaur Canyon Mbr Dinosaur Canyon Mbr
W, Up W, Up W, Up W, Up 665oC N N N N o o o o o 660 C 685 C 620 C 672 C o 670 C o o 620 C 100oC 630 C 400 C o o 450oC 670 C 645 C o o 500 C 1 200 C 1 1 1 400oC 10 mA/m 10 mA/m NRM 10 mA/m o 10 mA/m o o 200 C 500 C 100 C o NRM 200 C e g NRM f h NRM 304Cb, 5.5 m, 304I, 5.5 m, 419Ba, 30.3 m, 419Ca, 30.3 m, Dinosaur Canyon Mbr Dinosaur Canyon Mbr Whitmore Point Mbr Whitmore Point Mbr
Figure 2.3. Paleomagnetic data from the Moenave Formation. Orthogonal vector plots show typical characteristic remanent magnetization directions interpreted as normal polarity. Solid squares represent projections on the horizontal plane; open squares represent projection onto a vertical plane. (A) Thermal demagnetization of specimen from the Leeds locality; (B) chemical demagnetization of specimen from the Leeds locality; (C) and (D) thermal demagnetization of specimens from the Washington Dome locality; (E) and (F) thermal demagnetization of specimens from the Warner Valley locality; (G) and (H) thermal demagnetization of specimens from the Potter Canyon locality. Data given in tilt-corrected coordinates. See Table 2.1 for statistical information.
Mbr—Member; NRM—natural remanent magnetization.
45! W, Up W, Up 200oC W, Up W, Up o -1 400 C 10 (mA/m) 10-1mA/m 10-1(mA/m) 10-1mA/m o o o 500 C 200 C 400 C o o o 630 C 600 C o 640 C 620oC o 500oC 685 C N NRM 620 C N N 180oC 590oC N 570oC 90mT
100oC
30mT NRM 15mT 5mT
NRM a NRM b c d 185Db, 61.9 m, 185Bb, 61.9 m, 232Nb, 57.3 m, 232Qa, 57.3 m, Whitmore Point Mbr Whitmore Point Mbr Whitmore Point Mbr Whitmore Point Mbr
W, Up W, Up W, Up W, Up
-1 -1 -1 10-1mA/m 10 mA/m 10 mA/m 10 mA/m 200oC 530oC 620oC o o 640 C o 560 C 620oC N200 C o 575oC 640 C o N 680 C NRM o o 680 C 575 C N N o 200oC 500 C NRM
o 100 C 400oC
200oC
NRM e f NRM g h 359Ea, 49.8 m, 359Ca, 49.8 m, 413Aa, 43.7 m, 413Da, 43.7 m, Whitmore Point Mbr Whitmore Point Mbr Whitmore Point Mbr Whitmore Point Mbr
Figure 2.4. Paleomagnetic data from the Moenave Formation. Orthogonal vector plots show typical characteristic remanent magnetization directions interpreted as reverse polarity. Solid squares represent projections on the horizontal plane; open squares represent projection onto a vertical plane. (A) Thermal demagnetization of specimen from the Leeds locality; (B) combined alternating field and thermal demagnetization of specimen from the Leeds locality; (C) and (D) thermal demagnetization of specimens from the Washington Dome locality; (E) and (F) thermal demagnetization of specimens from the Warner Valley locality; (G) and (H) thermal demagnetization of specimens from the Potter Canyon locality. Data given in tilt-corrected coordinates.
46! Geographic Coordinates a) Leeds b) Washington Dome c) Warner Valley d) Potter Canyon N N N N
N = 21 N = 15 N = 27 N = 16
Structurally-Corrected Coordinates e) Leeds f) Washington Dome g) Warner Valley h) Potter Canyon N N N N
N = 21 N = 15 N = 27 N = 16
Figure 2.5. Equal area projections of geographic (A) and tilt-corrected (B) site-mean directions for Leeds, Washington Dome, Warner Valley, Potter Canyon. Solid circles represent lower-hemisphere projections and open squares represent upper-hemisphere projections.
47! !
B+%&'$R"%&/$,S-#$#,*'4$ F$G$0 N+":-)*#$3)"&%:"$O,*'-"# 3).-&%AB%+C%$",$%. DPQ"E,"'$'-+"E,-)&#$ 9)+#:-;$",$%.%&$'"+$>))/$8??@ !"##"$%&'$()*+,-..),/$0110$ 281$34+/$#.-'-&5$6-&')67 !"##"$%&'$()*+,-..),/$0110$ 201$34+/$#.-'-&5$6-&')67 Figure 2.6. Expected directions calculated from published North American paleomagnetic poles for Triassic–Jurassic aged rocks. Directions were calculated using an average latitude and longitude for the four localities used in this study. The Ekstrand and Butler (1989) and Molina-Garza et al. (2003) poles were derived from results obtained from Moenave strata. The grand mean normal- and reverse-polarity directions calculated in this study are shown with their circles of confidence. 48! a Site 103Ia, 34.1 m, Leeds, b Site 307Cb, 13.2 m, Warner Valley, Dinosaur Canyon Member Dinosaur Canyon Member 0.016 0.035 0.030 0.014 0.03 T 0.03 T 0.30 T 0.30 T 0.012 0.025 3.00 T 3.00 T 0.010 0.020 0.008 0.015 0.006 0.004 0.010 Magnetization (mA/m) Magnetization (mA/m) 0.002 0.005 0.000 0.000 0 100 200 300 400 500 600 700 0 100 200 300 400 500 600 700 Temperature (o C) Temperature (o C) c Site 109Fa, 28.1 m, Leeds Locality, d Site 337La, 62.9 m, Warner Valley, Dinosaur Canyon Member Whitmore Point Member 1.0 1.0 0.8 0.8 0.6 0.6 0.4 0.4 IRM Acquisition IRM Acquisition 0.2 0.2 BF IRM BF IRM 0.0 0.0 Magnetization (mA/m) Magnetization (mA/m) -0.2 -0.2 -1 -0.2 0 1 2 3 -1 -0.2 0 1 2 3 Applied Field (T) Applied Field (T) Figure 2.7. Rock magnetic results from selected sampled of the Moenave Formation. Top: examples of three component isothermal remanence magnetization (IRM) thermal demagnetization curves for specimens from the (A) Leeds and (B) Warner Valley localities. Bottom: IRM acquisition and backfield demagnetization response by samples treated with chemical demagnetization; samples from the (C) Leeds and (D) Warner Valley localities. BF—Back field. 49! Figure 2.8. Transmitted- and reflected-light photomicrographs of polished thin sections from representative samples of Moenave strata. Photomicrographs, organized by sample number, show parts of the sample that have been chemically leached (A–D) and parts of fresh samples (E–I). Images of detrital grains are present in all images. Authigenic hematite is observed as red pigment in figures (E) and (G). Hematite appears as black grains in transmitted light (A, B, D, E, and G) and white grains in reflected light (C, F, H, and I). Samples (A–C) were collected from the Whitmore Point Member, Washington Dome locality, site 232 at ~57.3 m. Samples (D–F) were collected from the Dinosaur Canyon Member, Leeds locality, site 110 at ~40.6 m. Samples (G–I) were collected from the Dinosaur Canyon Member, Leeds locality, site 109 at ~28.1 m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igure 2.9. Stratigraphic logs, virtual geomagnetic pole (VGP) latitudes based on site- mean directions, and magnetostratigraphic records for the Leeds, Washington Dome, Warner Valley, and Potter Canyon localities, along with a composite Moenave Formation record. Intervals that could not be sampled that are thicker than 5 m are identified in the hashed pattern. 51! Moenave Record St. Audrie’s Bay Record Oyuklu Record Southern Alps High Atlas Record Newark Record (This study) (Hounslow et al., 2004) (Gallet et al., 2007) (Muttoni et al., 2010) (Knight et al., 2004 (Kent and Olsen, 1999) Deeden et al., 2010) 70 M4n 30 B T-J E24n T-J B T-J 60 chert-rich limestone upper unit 13 A. Fm M3r SA6n K+ 50 M3n B T-J 150 1500 11 B T-J 200 M. Fm Lias Group J- ? Jurassic Whitmore Point Mbr Point Whitmore M2r T-J Boundary 20 5n.1r 40 SA5n.5r I+ 120 ? 9 1000 CAMP BIT5n SA5n.5n limestone with gray chert H- 30 M2n 7 90 ? ? E23r B T-J SA5n.3r Moenave Formation G+ 100 SA5n.3n 10 500 20 5 SA5n.2r Triassic F- 60 E22r Triassic SA5n.2n E+ Triassic Triassic D- Limestone Zu lower unit intermediate unit TriassicSA5n.1r Jurassic Passaic Fm Boonton Formation TriassicC+ Jurassic 10 3 Penarth Group B- A+ 30 BIT4r 0 M1r SA5n limestone with red chert Dinosaur Canyon Member 0 (m) M1n 1 B. Fm 0 (m) BIT3r 0 (m) 0 0 (m) (m) (m) Interval of normal polarity magnetozone Interval of reverse polarity magnetozone Interval of both normal and reverse polarity Approximate position of the T-J boundary Correlation of magnetozones ? Tentative correlation of magnetozones Figure 2.10. Correlation of the composite Moenave Formation magnetostratigraphy with magnetostratigraphic records from the marine succession at Saint Audrie’s Bay, UK, Oyuklu, Turkey, Southern Alps (M. Fm.—Malanotte Formation; A. Fm.—Albenza Formation), Italy, and continental sections in High Atlas (B. Fm.—Bigoudine Formation), Morocco, and continental deposits of the Newark and Hartford basins, eastern North America. CAMP—central Atlantic magmatic province. 52! Table TABLE 2.1. PALEOMAGNETIC DATA FROM THE MOENAVE FORMATION, UTAH AND ARIZONA VGP Depth Dgeo Igeo Dtc Itc !95 Lat Lon o o o o o o o Site (m) ( ) ( ) ( ) ( ) ( ) k n/no ( N) ( E) Leeds (37.202 oN, 246.650 oE) Contact between Whitmore Point Member and Kayenta N.A.* 69.5 Formtaion 184 65.9 35.3 2.5 38.4 13.8 37.3 5.3 4/7 43.9 7.8 157B 63.0 –† –† –† –† –† –† –† –† –† – 185 61.9 162.3 37.2 157.8 13.0 32.4 6.8 4/6 41.6 276.7 156B 60.6 –† –† –† –† –† –† –† –† –† 154B 58.5 –† –† –† –† –† –† –† –† –† – – 187 52.0 20.2 30.9 11.7 12.4 11.2 36.6 6/6 45.2 50 188 52.5 345.6 –5.7 349.6 25.9 11.3 36.1 6/6 64.7 90.8 – 113 52.0 340.8 22.3 345.8 5.3 16.7 22.0 5/5 53.1 90.7 – – 181 50.6 21.5 23.6 6.2 44.0 11.2 22.1 8/8 60.4 26.8 182 50.0 3.6 0.7 5.1 25.9 52.4 24.9 2/3 66 54.3 – 183 49.3 228.6 –1.9 231.3 –8.3 27.6 14.1 3/6 32.7 178.9 Contact between the Dinosaur Canyon and Whitmore N.A.* 49.0 Point Members – 101 48.5 356.4 15.4 0.6 4.3 8.9 39.8 8/8 54.9 65.6 – – 102 45.6 9.6 28.1 2.6 10.2 2.2 62.9 70/70 47.6 62.8 – 110 40.6 353.0 36.1 348.1 –9.6 2.9 1031.5 4/5 46.5 84 – 112 39.3 358.3 19.1 356.0 7.3 12.5 65.6 3/5 56.3 73.8 – 111 38.0 2.3 18.5 359.8 6.5 8.0 92.2 5/5 56.1 67 – 127B 35.3 10.1 13.6 10.7 9.7 8.1 153.7 3/3 56.2 47.2 – 103 34.1 9.8 12.2 8.9 11.1 6.4 116.6 5/7 57.4 50 109 28.1 11.9 –4.6 13.9 19.9 6.4 143.6 5/5 60.3 38.1 108 21.9 6.6 6.6 13.5 35.9 16.4 32.2 4/4 69.1 28.7 107 15.6 2.8 10.8 17.9 32.4 17.7 19.6 5/5 64.9 23 53! 106 7.6 357.2 22.8 11.6 55.4 6.1 97.6 7/9 80.6 340.8 105 5.3 12.0 1.0 40.1 46 180 1.2 5/5 55 341 – 105 5.3 188.9 26.9 182.5 8.4 80.3 4.0 2/4 48.5 242.8 104 3.6 351.7 0.6 355.7 35.4 9.3 42.9 7/8 72 79.8 Locality Mean 1.4 12.0 1.3 13.7 11.1 18.0 11/21 Washington Dome, (37.009 oN, 246.558 oE) 61.0 Contact between Whitmore Point Member and Kayenta N.A.* Formtaion 203 59.1 7.0 –0.4 12.7 38.5 5.8 108.6 7/7 71.1 27.6 57.3 – 232 169.7 10.1 189.6 18 180 1.1 5/5 42.8 233.6 232 57.3 343.0 1.6 345.8 47.7 6.6 135.8 5/7 75.5 125.8 230 55.7 359.4 2.1 4.7 37.6 9.6 681.9 2/3 73.4 51 205 52.6 –† –† –† –† –† –† –† –† –† 51.7 – – 229 198.2 –8.6 206.5 14.7 44.6 5.8 # 51.9 200.8 228 51.0 –† –† –† –† –† –† –† –† –† 227 50.1 –† –† –† –† –† –† –† –† –† 225 48.0 –† –† –† –† –† –† –† –† –† 45.5 – 208 347.9 17.1 33.4 4.7 2.0 27.0 2/4 62.5 61.9 221 43.9 338.7 25.0 344.4 69.0 11.7 75.3 3/6 71.1 216.2 42.5 Contact between the Dinosaur Canyon and Whitmore N.A.* Point Members 218 42.0 358.9 –9.5 0.9 31.3 8.8 110.5 4/4 69.8 64.1 41.7 – 210 353.3 29.3 346.9 8 11.7 33.7 6/6 54.8 89.7 40.9 – 217 351.8 29.7 349.1 4.6 5.1 119.0 4/8 53.8 85.3 40.0 – 216 357.6 14.3 358.2 25.9 2.7 598.2 6/6 66.5 71 37.4 – 211 359.5 20.5 358.2 11.9 3.8 134.4 12/15 58.9 70.1 36.4 – 239B 0.6 20.1 354.9 –0.4 12.8 39.3 4/4 48.3 71.4 35.4 – 212 3.8 10.4 5.5 28.4 4.8 449.8 3/3 67.5 52.6 215 32.0 7.0 –8.6 9.8 23.7 10.7 32.5 7/7 63.8 44.5 214 28.0 25.3 5.3 33.9 34.6 9.7 90.0 4/4 55.3 358.6 Locality Mean 359.3 –8.5 1.4 29.7 10.8 18.6 13/15 54! Warner Valley, (37.016 oN, 246.625 oE) 52.5 Contact between Whitmore Point Member and Kayenta N.A.* Formtaion 359B 49.8 14 9.3 24.1 23.7 81.9 3.9 2/4 57.2 19.1 49.8 – 359B 184.2 25.9 177.4 1.7 96.1 3.1 2/4 52.1 250.8 356B 46.5 –† –† –† –† –† –† –† –† –† 44.0 Contact between the Dinosaur Canyon and Whitmore N.A.* Point Members 353B 42.5 356.9 36.3 1.2 26.6 38.2 4.5 4/4 67.0 63.3 352B 41.5 24.2 41.3 25.7 29.4 24.4 11.3 4/4 46.4 358.1 328 38.5 322.4 14.1 328.9 4.7 19.7 40.4 3/4 45.1 113.6 35.8 – 326 354.9 –4.9 359.7 12.9 16.5 12.2 6/9 46.5 67 325 32.6 15.0 17.8 13 11.5 13.9 19.7 7/8 56.6 42.6 324 32.5 4.9 10.6 3.9 3.7 6.9 135.8 4/4 54.7 59.8 323 29.5 37.6 12.2 35.3 14.5 6.9 76.8 7/7 46.4 10.5 319 28.6 17.1 18.8 14.2 18.1 11.6 34.1 6/6 59.4 38.2 312 21.0 345.0 25.0 343.0 20.5 9.8 33.1 7/7 59.5 101 321 17.5 193.7 9.2 269.1 33.8 180 1.1 3/5 10.3 172.1 320 16.8 352.8 21.4 350.8 11.6 15.1 20.6 5/6 57.7 83.9 311 16.0 7.8 21.3 4.9 19.4 10.1 37.6 6/6 62.6 56.1 310 15.6 18.5 24.3 14.5 25.8 5.1 172.4 6/6 63.2 33.9 318 14.6 12.5 37.5 0.6 29.8 6.9 57.4 7/7 69 65 309 13.6 16.9 29.1 12.0 31.5 5.0 177.0 6/6 67.4 35.4 308 13.5 13.7 24.0 6.9 17.5 6.6 72.0 8/8 61.3 52.3 307 13.2 25.6 34.5 11.1 24.9 9.4 42.4 7/7 64.1 41.2 317 12.3 0.4 24.3 357.5 22.1 10.0 37.5 7/7 64.4 72.3 316 11.8 10.7 21.4 7.7 20.1 11.1 37.5 6/6 62.5 50 315 11.5 25.4 24.5 20.9 27.6 9.3 98.3 4/4 60.9 21.3 314 10.5 6.6 12.5 6.8 2.5 6.9 96.0 6/6 53.7 55.1 313 10.2 16.7 0.7 14.8 –1.4 5.5 276.4 4/7 49.9 43.2 306 6.0 3.9 18.5 1.7 14.7 13.7 24.7 6/6 60.4 63.2 305 5.6 343.0 25.8 341.1 21.4 6.9 124.2 5/5 59 104.7 304 5.5 354.8 17.4 354.6 9.4 47.9 7.7 3/4 57.4 76.6 302 1.6 29.3 –2.9 4.1 –5.4 89.2 10.1 2/9 50.1 60.2 Locality Mean 11.2 21.8 7.3 18.4 6.5 29.2 18/27 Potter Canyon, (36.879 oN, 247.156 oE) 44.3 Contact between Whitmore Point Member and Kayenta N.A.* Formtaion 413 43.7 182.7 – 182.3 – 5.6 327.7 3/4 – 242.5 55! 16.8 16.0 61.2 42.0 – – – 414 189.0 32.7 187.7 37.0 60.8 9.5 2/4 72.4 222.7 415 40.1 –† –† –† –† –† –† –† –† –† 36.9 – – – 416 199.3 29.7 197.6 34.0 29.7 12.2 3/4 66.1 202.2 34.7 – – – 417 167.1 45.0 167.0 50.0 90.0 1.8 2/4 77.6 311.4 419 30.3 0.3 19.0 360 14 5.6 68.0 11/11 60.2 67.2 420 29.5 355.7 17.0 357.9 13.4 8 71.3 5/5 59.8 71.4 425B 28.8 –† –† –† –† –† –† –† –† –† 421 26.3 352.1 –3.6 351.1 –7.4 5 147.8 7/7 48.5 80.7 422 25.7 327.0 8.5 327.2 4.1 8.1 154.2 3/4 43.9 115.9 23.4 – 423 11.8 26.3 301 61.5 180 0.7 2/4 –6 106.5 424 18.4 0.7 0.2 356.4 –6 4.6 88.7 8/8 50 72.8 17.8 – 425 359.8 –2.1 359.9 11.5 1.9 1722.6 4/4 47.3 67.3 15.9 – 426 351.3 –1.7 351.3 11.2 9.2 76.2 4/4 46.7 79.9 11.9 – – – 427B 191.6 28.7 189.8 23.9 10.4 93.0 3/3 64.1 224.8 427 11.7 15.4 1.1 19.4 –0.1 11.2 47.9 5/5 48.9 36.8 11.5 Contact between the Dinosaur Canyon and Whitmore N.A.* Point Members 428 10.3 10.1 –1.0 10.3 –5.1 7.0 48.3 9/9 49.4 51.2 405 4.6 10.2 64.4 6.0 67.6 24.6 8.5 5/8 75.7 262.9 Locality Mean 359.2 3.7 358.8 2.5 9.2 23.8 13/16 Note: Samples were collected in southwestern Utah and north- central Arizona. Paleomagnetic data obtained from the Moenave Formation. The first two columns of this table show the site name and stratigraphic position for each locality. The next four columns show the degrees of declination (D) and inclination (I) for geographic (geo) and stratigraphic (tc) coordinates, respectively. The next two columns show the cone of confidence (!95) and precision parameter (k). The next column shows the ratio of number of specimens accepted (n) to number of specimens demagnetized for each site (no). The last two columns provide the virtual geomagnetic pole position (VGP) calculated for each site given in degrees of latitude (North) and longitude (East of Greenwich). *N.A. No sample was collected at the boundary. †–samples are remagnetized. 56! 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Lucas, S.G., and Tanner, L.H., 2007a, Tetrapod biostratigraphy and biochronology of the Triassic–Jurassic transition on the southern Colorado Plateau, USA: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 244, p. 242–256, doi: 10.1016/j.palaeo.2006.06.030. Lucas, S.G., and Tanner, L.H., 2007b, The non-marine Triassic–Jurassic boundary in the Newark Supergroup of eastern North America: Earth-Science Reviews, v. 84, p. 1–20, doi: 10.1016/j.earscirev.2007.05.002. 62! Lucas, S.G., and Tanner, L.H., 2008, Reexamination of the end-Triassic mass extinction, in Elewa, A.M.T., ed., Mass Extinction: New York, Springer Verlag, p. 66–103. Marzoli, A., Renne, P.R., Piccirillo, E.M., Ernesto, M., Bellieni, G., and De Min, A., 1999, Extensive 220-Million-year-old continental flood basalts of the central Atlantic magmatic province: Science, v. 284, p. 616–618. Marzoli, A., Bertrand, H., Knight, K.B., Cirilli, S., Buratti, N., Verati, C., Nomade, S., Renne, P.R., Youbi, N., Martini, R., Allenbach, N., Neuwerth, R., Rapaille, C., Zaninetti, L., Zaninetti, L., and Bellieni, G., 2004, Synchrony of the central Atlantic magmatic province and the Triassic–Jurassic boundary climate and biotic crisis: Geology, v. 32, no. 11, p. 973–976. McFadden, P.L., and McElhinny, M.W., 1990, Classification of the reversal test in paleomagnetism: Geophysical Journal International, v. 103, no. 3, p. 725–729. Milner, A.R.C., Lockley, M.G., and Kirkland, J.I., 2006, A large collection of well- preserved theropod dinosaur swim tracks from the Lower Jurassic Moenave Formation, Saint George, Utah, in Harris, J.D., Lucas, S.G., Spielmann, J.A., Lockley, M.G., Milner, A.R.C., and Kirkland, J.I., eds., The Triassic–Jurassic Terrestrial Transition: Albuquerque, New Mexico Museum of Natural History and Science. Molina-Garza, R.S., Geissman, J.W., and Lucas, S.G., 2003, Paleomagnetism and magnetostratigraphy of the lower Glen Canyon and upper Chinle Groups, Jurassic–Triassic of northern Arizona and northeastern Utah: Journal of Geophysical Research, v. 108, no. B4, p. 1–24, doi: 10.1029/2002JB001909. Muttoni, G., Kent, D.V., Jadoul, F., Olsen, P.E., Rigo, M., Galli, M.T., and Nicora, A., 63! 2009, Rhaetian magneto-biostratigraphy from the Southern Alps (Italy): Constraints on Triassic chronology: Palaeogeography, Palaeoclimatology, Palaeoecology. Olsen, P.E., Kent, D.V., Et-Touhami, M., and Puffer, J., 2003, Cyclo-, magneto-, and bio- stratigraphic constraints on the duration of the CAMP event and its relationship to the Triassic–Jurassic boundary, in Hames, W.E., et al., eds., The Central Atlantic Magmatic Province: Insights from Fragments of Pangea, Geophysical Monograph Series: Washington, D.C., American Geophysical Union, p. 7–32. Opdyke, N.D., and Channell, J.E.T., 1996, Magnetic Stratigraphy: London and San Diego, Academic Press, 364 p. Palfy, J., Demeny, A., Haas, J., Hetenyi, M., Orchard, M.J., and Veto, I., 2001, Carbon isotope anomaly and other geochemical changes at the Triassic–Jurassic boundary from a marine section in Hungary: Geology, v. 29, no. 11, p. 1047–1050, doi: 10.1130/0091-7613(2001)029<1047:CIAAOG>2.0.CO;2. Palfy, J., Demeny, A., Haas, J., Carter, E.S., Gorog, A., Halasz, D., Oravecz-Scheffer, A., Hetenyi, M., Marton, E., Orchard, M.J., Ozsvart, P., Veto, I., and Zajzon, N., 2007, Triassic–Jurassic boundary events inferred from integrated stratigraphy of the Csovar section, Hungary: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 244, p. 11–33, doi: 10.1016/j.palaeo.2006.06.021 Pipiringos, S.E., and O’Sullivan, R.B., 1978, Principal unconformities in Triassic and Jurassic rocks, western interior United States—A preliminary survey: U.S. Geological Survey Professional Paper 1035-A, p. A1-A29. Tanner, L.H., and Lucas, S.G., 2007, The Moenave Formation: Sedimentologic and 64! stratigraphic context of the Triassic–Jurassic boundary in the Four Corners area, southwestern U.S.A: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 244, p. 111–125, doi: 10.1016/j.palaeo.2005.06.039. Tanner, L.H., and Lucas, S.G., 2009a, Tetrapod trace fossils from lowermost Jurassic strata of the Moenave Formation, northern Arizona, USA: Volumina Jurassica, v. 6, no. 133–141. Tanner, L.H., and Lucas, S.G., 2009b, The Whitmore Point Member of the Moenave Formation: Early Jurassic Dryland Lakes on the Colorado Plateau: Southwestern USA: Volumina Jurassica, v. 6, p. 11–21. Tanner, L.H., Lucas, S.G., and Chapman, M.G., 2004, Assessing the record and causes of Late Triassic extinctions: Earth-Science Reviews, v. 65, p. 103–139, doi: 10.1016/S0012-8252(03)00082-5. Torsvik, T.H., Muller, R.D., Van der Voo, R., Steinberger, B., and Gaina, C., 2008, Global plate motion frames: Toward a unified model: Reviews of Geophysics, v. 46, no. RG3004, p. 1–44, doi: 10.1029/2007RG000227. Van der Voo, R., 1993, Paleomagnetism of the Atlantic, Tethys, and Iapetus Oceans: Cambridge, Cambridge University Press. van de Schootbrugge, B., Quan, T.M., Lindstrom, S., Puttmann, W., Heunisch, C., Pross, J., Fiebig, J., Petschick, R., Rohling, H.-G., Richoz, S., Rosenthal, Y., and Falkowski, P.G., 2009, Floral changes across the Triassic/Jurassic boundary linked to flood basalt volcanism: Nature Geoscience, v. 2, p. 589–594, doi: 10.1038/ngeo577. Weil, A.B., Geissman, J.W., and Ashby, J.M., 2006, A new paleomagnetic pole for the 65! Neoproterozoic Uinta Mountain Supergroup, Central Rocky Mountain States, USA: Precambrian Research, v. 2655, p. 1–26. Whiteside, J.H., Olsen, P.E., Kent, D.V., Fowell, S.J., and Et-Touhami, M., 2007, Synchrony between the Central Atlantic magmatic province and the Triassic– Jurassic mass-extinction event?: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 244, p. 345–367, doi: 10.1016/j.palaeo.2006.06.035. 66! CHAPTER 3 INCLINATION BIAS OF PALEOMAGNETIC DATA FROM THE UPPERMOST TRIASSIC TO LOWERMOST JURASSIC MOENAVE FORMATION, UTAH AND ARIZONA: IMPLICATIONS FOR THE TRIASSIC TO JURASSIC NORTH AMERICAN APPARENT POLAR WANDER PATH FOR THE SOUTHWEST UNITED STATES Abstract Paleomagnetic data were acquired from strata of the uppermost Triassic through lowermost Jurassic Moenave Formation, Colorado Plateau region, to obtain a well- defined paleomagnetic pole that would be adequately corrected for the error associated with sediment compaction and thus inclination shallowing. This new robust dataset is used to test whether the inclination error associated with Moenave Formation strata is consistent with the flattening factor reported for other sedimentary units, and to attempt reconcile the paleomagnetic data of the southwest U.S.A. with those from the eastern seaboard of North America. Characteristic remanent magnetization (ChRM) directions for 58 sites, from four localities both on and off the Colorado Plateau, were isolated by progressive thermal and chemical demagnetization. The four locality mean directions are statistically indistinguishable and the normal and reverse polarity data sets share a common mean. The ChRM in the Moenave Formation is interpreted as having been acquired close to the time of sediment deposition across the Triassic-Jurassic transition. Elongation/inclination analysis on the Moenave Formation data set is used to estimate a flattening factor of about 0.78, which is less than that obtained for strata of the Newark ! "#! and Hartford Basins of the northeastern U.S.A. (flattening factor of about 0.54). The paleomagnetic pole obtained in this study, corrected for inclination error is located at 62.3° N, 60.0° E (A95 = 3.8°, N = 58 independent sites). When combined with previously reported data from the Moenave Formation, assuming a flattening factor of 0.78 for these data, a pole located at 56.6° N, 62.6° E (A95 = 7.4°, N = 102 independent sites) is obtained. Alternatively, if an arbitrary flattening factor of 0.53 is applied to all accepted data from the Moenave Formation the resulting paleomagnetic pole is located at 64.6° N, 56.4° E (A95 = 7.4°). The poles are corrected for an estimated 4° of clockwise Colorado Plateau rotation, which took place sometime between the mid-Cretaceous and the present. This is a maximum estimate of plateau rotation based on several geologic considerations. The combined Moenave Formation pole, corrected with a locally derived flattening factor, remains in support of the existence of a long-proposed J1 cusp, as elucidated by Paleomagnetic Euler Pole (PEP) apparent polar wander path (APW) constructions and lies about 11° west of the composite ca. 200 Ma poles defined by sliding time window methods of APW path construction for North America. Introduction Historically, paleomagnetic pole positions from the uppermost Triassic to lowermost Jurassic sedimentary rocks of the southwest U.S.A. have differed from those obtained from rocks of similar ages in the northeast U.S.A. The poles differ by about an angular distance of 10° to 15°, and this has fueled debate for nearly three decades concerning the plate motion history of North America (Irving and Irving, 1982; Gordon et al., 1984; Van der Voo, 1993; Molina-Garza et al., 1995; Steiner and Lucas, 2000; ! "$! Besse and Courtillot, 2002; Beck and Housen, 2003; Besse and Courtillot, 2003; Kent and Olsen, 2008; Torsvik et al., 2008; Kent and Irving, 2010). Uppermost Triassic to lowermost Jurassic rocks from the southwest U.S.A. yield pole positions located near 63.7° N and 59.7° E (Figure 3.1), consistent with the predicted position of the J1 cusp, an abrupt shift in the North American apparent polar wander (APW) path at about the time of the Triassic-Jurassic transition, based on the paleomagnetic Euler pole (PEP) method (Gordon et al., 1984; Beck and Housen, 2003) of constructing APW paths. The implication of a major cusp in the APW path is that the latest Triassic to earliest Jurassic orientation of North America dramatically changed from a westerly to and easterly position over a short time period near 200 Ma. On the other hand, pole positions derived from rocks of similar age from the northeast US are located near 67° N and 93.8° E (Figure 3.1), and are consistent with the sliding-time-window method of APW path construction (Van Alstine and Boer, 1978; Harrison and Lindh, 1982; Irving and Irving, 1982; Torsvik et al., 2008; Kent and Irving, 2010), suggesting a considerably smoother poleward progression of paleomagnetic pole positions at this time. Composite APW paths apply a sliding time window approach analysis to poles from different cratons that are rotated into a common reference frame (Torsvik et al., 2008; Kent and Irving, 2010). Composite APW paths tend to be smooth while paths constructed from individual pole positions are more likely to preserve abrupt changes in path morphology (Gordon et al., 1984) (Figure 3.1). The specific location of discrete latest Triassic to earliest Jurassic paleomagnetic poles bears on the approach taken for APW path construction, our understanding of plate movement, and North American paleogeography (Irving and Irving, 1982; Gordon et al., 1984; Ekstrand and Butler, 1989; Bryan and Gordon, 1990; ! "%! Van der Voo, 1993; Kodama et al., 1994a; Molina-Garza et al., 1998; Steiner and Lucas, 2000; Beck and Housen, 2003; Torsvik et al., 2008; Kent and Irving, 2010). Sedimentary successions are susceptible to an inclination error in paleomagnetic directions associated with initial deposition, burial/compaction, and lithification (Johnson et al., 1948; King, 1955; Griffiths et al., 1957; Tauxe and Kent, 1984). The elongation/inclination method of quantifying inclination error compares the symmetry of directional datasets with those predicted by a statistical paleosecular variation model TK03.GAD (Tauxe and Kent, 2004). These parameters are related as follows: f = tan (Io) / tan (If), (1) where f is the flattening factor, Io is the observed inclination, and If is the corrected inclination (King, 1955). In this study, paleomagnetic data obtained from hematitic siltstones, mudstones, and fine grained sandstones of the Moenave Formation are used to estimate paleomagnetic pole positions, corrected for inclination shallowing, that are representative of the time period of the Triassic/Jurassic transition. The Moenave Formation is a well-studied non-marine sequence of latest Triassic to earliest Jurassic age exposed on and adjacent to the Colorado Plateau region (Harshbarger et al., 1957; Pipiringos and O'Sullivan, 1978; Kirkland and Milner, 2006; Lucas and Tanner, 2007; Tanner and Lucas, 2007; Lucas et al., 2011) (Figure 3.2a-b). The magnitude of inclination error recorded in paleomagnetic data from latest Triassic to earliest Jurassic strata from the Newark and Hartford Basins was estimated to be associated with a flattening factor of (f) of 0.54, based on the elongation/inclination method (Tauxe and Kent, 2004; Kent and Tauxe, 2005; Kent and Olsen, 2008). Recent studies (Kent and Olsen, 2008; Tauxe et al., 2008) suggest that if data collected from strata of the American ! #&! Southwest were corrected for inclination bias, using a flattening factor of 0.54, as well as a large, ~13°, magnitude of clockwise Colorado Plateau rotation, then pole positions from the two regions would coincide. Corrections for inclination bias, ideally, should be derived from data from sections actually sampled for paleomagnetic data, rather than an arbitrary flattening factor for rocks, even though of similar lithologic character examined in other studies. This study provides an independent estimate of the magnitude of inclination error associated with paleomagnetic data from the Moenave Formation with the goal of providing a more accurate latest Triassic to earliest Jurassic pole position for the American Southwest. The angular distance between the Moenave Formation pole, corrected for inclination shallowing, and the time equivalent pole from the eastern seaboard (about 18°) is greater than the amount of clockwise rotation that can be realistically accommodated by compressional structures on the eastern and northern boundaries of the Colorado Plateau (~4° is a maximum estimate of plateau scale clockwise rotation) (Hamilton, 1988; Molina-Garza et al., 1995; Wawrzyniec et al., 2007). The more westward pole position from rocks of the southwest U.S.A. cannot be an artifact of a combination of inclination bias and Colorado Plateau rotation because the angular distance between the poles is greater than their respective cones of confidence. Late Triassic to Early Jurassic paleomagnetic data from the American Southwest continue to indicate that latest Triassic to earliest Jurassic paleomagnetic pole positions for North America changed from a more westerly to easterly location in the latest Triassic to earliest Jurassic. Below, the review of the paleomagnetic studies of the Moenave Formation are reviewed, the inclination bias analysis is presented, and new estimate of a paleomagnetic ! #'! pole positions for these strata is presented. The disagreement between paleomagnetic poles from seemingly coeval rocks is then considered with respect to regional post- depositional compaction and inclination shallowing and Colorado Plateau rotation. Methods Site mean directions were obtained from three localities in the western transition zone of the Colorado Plateau (Leeds, Washington Dome, and Warner Valley) and one locality on the Colorado Plateau (Potter Canyon). Two pole positions are presented in this paper and labeled as MF3 and MF4 in the Figure 3.5 and Table 3.2. A total of 56 virtual geomagnetic poles obtained from 56 site mean directions provided in this study were averaged using Fisher statistics to obtain pole MF3. A total of 102 virtual geomagnetic poles obtained from this and two previous studies, including 23 sites from the Vermillion Cliffs (Ekstrand and Butler, 1989), and 23 sites from the Echo Cliffs (Molina-Garza et al., 2003) were averaged using Fisher statistics to obtain the paleomagnetic pole MF4. Two flattening factors are used to correct inclination values of the Moenave Formation data and these results are compared with the amount of inclination bias estimated for strata of comparable age from the eastern seaboard. The first flattening factor is a locally derived value for f1 calculated in this study. The second f2 value (about 0.54) is calculated from Newark and Hartford Basin. Lastly, the combined paleomagnetic pole MF4 was corrected for locally derived (MF4f1) and arbitrary flattening factors (MF4f2) and was also corrected for a maximum of 4° of clockwise rotation of the Colorado Plateau about an Euler pole located at 34° N and 260° E (MF4f1’ and MF4f2’). Fisher statistics, VGP calculations, and elongation/inclination analyses ! #(! were applied using the PmagPy software (Tauxe and Kent, 2004; Tauxe, 2008). Paleomagnetic pole rotations were calculated using OSXStereonet version 1.3 (Cardozo, 2011). Results The Moenave Formation is the lowest component of the Glen Canyon Group and is exposed in southwest Utah and north-central Arizona (Harshbarger et al., 1957; Kirkland and Milner, 2006; Lucas et al., 2011). Previous paleomagnetic studies of the Moenave Formation focused on the Vermillion (36.750° N, 248.600° E) (Ekstrand and Butler, 1989) and Echo cliffs (36.100° N, 248.600° E) (Molina-Garza et al., 2003) (Figure 3.2a-b) sections/localities. Samples for the present study were collected from exposures of the Dinosaur Canyon and Whitmore Point members of the Moenave Formation at four localities; Leeds (37.202° N, 246.650° E), Washington Dome (37.009° N, 246.558° E), Warner Valley (37.016° N, 246.625° E), and Potter Canyon (36.100° N, 248.600° E) (Figure 3.2c-f). A principal goal of this study was to provide a comprehensive magnetostratigraphy for the time period across the Triassic-Jurassic transition in the American Southwest and thus provide a means of correlating marine and non-marine biostratigraphic records in a global context (Donohoo-Hurley et al., 2010; Lucas et al., 2011). Most of the strata of the Moenave Formation are well-indurated, hematitic- cemented siltstones and sandstones. Sixty-one sites were sampled using a hand-held gas- powered drill. Each drilled site consists of 4 to 70 sample cores. Seven sites, each of which represents a thin stratigraphic interval, were sampled in less well-indurated ! #)! mudstone strata with 4 to 7 block samples, whereby one block is considered to represent a single sediment horizon. Samples from core and block samples were demagnetized using thermal and chemical demagnetization methods. The ChRM is carried by both pigmentary and detrital hematite (Donohoo-Hurley et al., 2010). Site mean directions with !95 values greater than 2! from the mean direction were rejected from further analysis. A total of 56 site mean data (378 ChRM directions) were used to calculate virtual geomagnetic poles (Table 3.1, Figure 3.3), and obtain a new paleomagnetic pole position for the Moenave Formation at 60.4° N, 60.1° E (A95 = 3.8°). The structurally corrected data pass a “B” class ("c is 8.98° at the 95% confidence) of McElhinny and McFadden (1990). A total of 102 site mean directions from this and previous studies (Ekstrand and Butler, 1989; Molina-Garza et al., 2003) yield a combined pole position of 60.6° N, 57.3° E, (A95 = 7.4°). The Cartesian components of the normal and reverse polarity data mean directions for all Moenave Formation data overlap with a common mean at 95 percent confidence indicating localities on and off the Colorado Plateau share a common mean. Both individual ChRM directions obtained in this study and the 102 site mean directions obtained from this and previous Moenave Formation strata are elongated in the horizontal plane, consistent with some degree of inclination error (Figure 3.4a). The elongation/inclination approach uses bootstrap statistics to estimate an observed inclination of 8.1° for the ChRM directions and 14.3° for the site mean directions (Figure 3.4). The difference in inclinations observed for the ChRM data set versus the site mean data set could be due to the discrepancy in the size of the data sets. The statistical sampling is then compared with an expected inclination, obtained from TK03.GAD, to ! #*! estimate a flattening factor of 0.73 for the ChRM directions and 0.78 for the site mean directions. The corrected inclination values are 10.8° for the ChRM directions and 17.9° for the site mean directions (Figure 3.4). Although the lithology is different between the mostly fluvial Dinosaur Canyon Member and mostly lacustrine Whitmore Point Member, there is no significant difference in the inclination error calculated for each member. The flattening factor calculated for 95 ChRM directions obtained from the Whitmore Point strata is 0.90. The flattening factor obtained from 283 ChRM directions obtained from Dinosaur Canyon strata is also 0.90. The flattening factor obtained for member specific data sets is not consistent with the flattening factor obtained when all the ChRM directions are combined. The source of the discrepancy is not clear and will be the subject of further investigation. An arbitrary flattening factor of 0.54 (f2) calculated for Newark and Hartford strata (Kent and Olsen, 2008) is also applied to the data to test if an f value of 0.54 is sufficient to bring the poles for the southwest and northeast U.S.A. into agreement (Kent and Olsen, 2008; Tauxe et al., 2008). The corrected inclination values for the Moenave Formation suggest the southwest US was located at 8.8° N, using f1, and 12.0° N, using f2, during latest Triassic to earliest Jurassic time (Table 3.2). After correcting for inclination bias, using a locally derived flattening factor (f1 = 0.78) and an arbitrary value for comparison (f2 = 0.54), the new latest Triassic to earliest Jurassic paleomagnetic poles based on data obtained in the present study are located at 62.3° N, 60.0° E, (A95 = 3.8°, N = 56 sites, f1) and at 65.3° N, 59.0° E, (f2). When combined with the data obtained in previous studies (Ekstrand and Butler, 1989; Molina-Garza et al., 2003) the two new pole positions for the Moenave Formation are located at 56.6° N, 62.6° E, (A95 = 7.4°, N = 102 sites, f1) and 64.6° N, 56.4° E, (f2) (Figure 3.5). After ! #+! correcting for 4° of clockwise Colorado Plateau rotation about an Euler pole located at 34° N, 260° E (Wawrzyniec et al., 2007) the averaged pole positions change to 62.3° N, 68.0° E (f1), and 67.8° N, 66.4° E (f2). The location of the J1 cusp predicted by PEP methods of APW path construction (Gordon et al., 1984; Beck and Housen, 2003) overlaps at 95 percent confidence with Moenave Formation pole positions corrected for a locally derived inclination error (f1 = 0.78) and Colorado Plateau rotation. Pole positions for the ca. 200 Ma time frame (Van der Voo, 1993; Besse and Courtillot, 2002, 2003; Torsvik et al., 2008; Kent and Irving, 2010) are located ~14° E of the Moenave Formation pole corrected for a locally derived flattening factor (MF4f1, Figure 3.5, Table 3.3). When compared with the Moenave Formation pole position corrected for an arbitrary flattening factor (f2 = 0.54) and Colorado Plateau rotation (MF4f2) the pole positions for the ca. 200 Ma time frame are located ~6° to the E (Figure 3.5, Table 3.3). An angular distance of ~6° is not distinguishable at the 95 percent confidence because the cones of confidence overlap (Figure 3.5). Discussion The paleomagnetic data obtained from Moenave Formation strata provide a robust record of latest Triassic to earliest Jurassic magnetic field behavior and demonstrate that the time interval across the Triassic-Jurassic transition dominated by a normal polarity state. A total of 102 site mean directions obtained from Moenave Formation strata (23 sites (Ekstrand and Butler, 1989), 23 sites (Molina-Garza et al., 2003), and 56 sites (this study), are defined by magnetizations that are interpreted to have been acquired at or close to the time of deposition. The mean normal and reverse paleomagnetic data ! #"! obtained from the Moenave Formation data pass a reversal test. Site mean directions from the four localities sampled in this study share a common mean. Sedimentary deposits are susceptible to post depositional rotation of detrital magnetic particles due to burial and lithification that can result in inclination shallowing (Johnson et al., 1948; King, 1955; Griffiths et al., 1957; Tauxe and Kent, 1984). The elongation/inclination method is applied to paleomagnetic data sample data obtained to estimate a flattening factor of about 0.73. An alternative flattening factor of about 0.78 was estimated for all available site mean data obtained from the Moenave Formation. The degree of inclination shallowing predicted by either flattening factor (about 0.73 or 0.78) less than the amount of inclination shallowing estimated in coeval strata of the eastern seaboard. The flattening factor obtained from strata of the Dinosaur Canyon and Whitmore Point members are the same, 0.90, but differ from the flattening factor obtained when the two members are combined. It is surprising that the member specific flattening factor is so different from the flattening factor obtained for the entire paleomagnetic data set. It is equally surprising that the member specific flattening factors are equal to each other because the dominant lithology in the two members is very different. The Dinosaur Canyon Member is principally comprised of variable bedded medium to fine grained sandstone. The Whitmore Point Member is dominated by siltstones and mudstone intervals with only a few sandstone beds. It is likely that the sediment defining each member would have different water contents at the time of deposition, lithification and compaction and therefore experience different amounts of inclination shallowing. The E/I method may have an unidentified source of error due to grain size effects or the time ! ##! scale of remanence acquisition. A calibration of the E/I method with respect to different types of sediment and depositional environments could help address these differences. Deformation of the Colorado Plateau The Colorado Plateau is a crustal region within the southwest U.S. that experienced minimal amount of internal deformation during the Phanerozoic and in particular during the Cretaceous to early Tertiary development of the Cordilleran fold and thrust belt (Sevier Orogeny in the western U.S.A.) and during the latest Cretaceous to early Tertiary foreland, basement involved deformation (Laramide Orogeny) (Hamilton, 1981; Woodward et al., 1997; Molina-Garza et al., 1998; Cather, 1999; Bump, 2003; Bump and Davis, 2003; Bump, 2004; Wawrzyniec et al., 2007). . The Colorado Plateau has rotated as a rigid block with no evidence of internal differential rotation. Evidence for deformation at the margins of the plateau include features that, cumulatively, have resulted in some 120+ kms of shortening north of the plateau and up to 35-40 km of dextral shear east of the Colorado Plateau (Molina-Garza et al., 1998; Wawrzyniec et al., 2007). The compressional structures show evidence for NE-SW directed shortening, and both NE and SW vergence, consistent with a three-dimensional stress model. The principal stresses result from basal traction of the Farallon slab coupled to the North American plate, and a significant contribution from the crustal shortening associated with the Sevier Orogeny located to the west of the plateau (Bump, 2004). The location and orientation of Laramide-style, basement involved structures near and around the Colorado Plateau are governed to some degree by the reactivation of preexisting structures, which act as stress guides (Bump, 2003). ! #$! Estimates of the cumulative magnitude of clockwise rotation of the Colorado Plateau range from small (2°-4°) to large (11°-13°). Early estimates of Colorado Plateau rotation made by Hamilton (1981) were between 2° and 4° of clockwise rotation based on observations of crustal shortening. More recently, small magnitudes of Colorado Plateau rotation have been advocated by several workers on the basis of by geologic structures and direct comparison between groups of paleomagnetic poles of similar age (Bryan and Gordon, 1990; Molina-Garza et al., 1998; Beck and Housen, 2003; Wawrzyniec et al., 2007). A large, ~11°, magnitude of rotation was initially inferred from pole to pole comparisons of data from Triassic strata of the Colorado Plateau and data obtained from the North American craton (Steiner, 1986). Further support for a large amount of Colorado Plateau rotation came from additional pole-to-pole comparisons (Steiner, 1988; Kent and Witte, 1993; Steiner and Lucas, 2000; Kent and Irving, 2010). We favor a maximum of 4° of clockwise Colorado Plateau rotation, as this is most consistent with kinematic models of Laramide strain of the Colorado Plateau and rotation estimates based on the best quality Late Carboniferous through late Jurassic paleomagnetic poles from the Colorado Plateau and the North American craton (Bryan and Gordon, 1990; Molina- Garza et al., 1998; Beck and Housen, 2003; Wawrzyniec et al., 2007). Two pole positions are presented in this paper. One (MF3) is derived from data obtained from the four Moenave Formation sections sampled in this study. The second (MF4) is calculated using all available data from the Moenave Formation. Each pole is corrected with a locally derived flattening factor (0.78) and a flattening factor estimated from coeval strata of the Newark and Hartford basins for comparison (0.54). Each pole is then corrected for a maximum amount of Colorado Plateau rotation about an Euler pole ! #%! of rotation located at 34° N and 260° E, the approximate location of which was first proposed by Hamilton (1988) and subsequently tested by other workers (Molina-Garza et al., 1995; Wawrzyniec et al., 2007). The new average Moenave Formation pole positions overlap at the 95 percent confidence with the previous pole position estimates that define the J1 cusp (Table 3.3, Figure 3.5). Similarity in the position of the Moenave Formation poles and J1 cusp support a western location for the latest Triassic to the earliest Jurassic North American paleomagnetic pole. The value used to correct for inclination error becomes important when comparing a Moenave Formation pole with composite poles representing sliding time window approaches to APW path constructions. The Moenave Formation pole corrected with a locally derived flattening factor and corrected for Colorado Plateau rotation lies 1.6° northwest of the J1 cusp and 8° southwest of the ca. 200 Ma pole positions estimated by applying a sliding time window approach to average pole positions from several continents rotated into a global reference frame (Besse and Courtillot, 2002, 2003; Torsvik et al., 2008; Kent and Irving, 2010). The Moenave Formation pole derived on the basis of a flattening factor of 0.54 and corrected for Colorado Plateau rotation lies 6.9° N of the J1 cusp and 5.8° east of the ca. 200 Ma pole (Table 3.3). A large magnitude of inclination error would bring the Moenave Formation pole into agreement with ca. 200 Ma composite poles. Results of the elongation/inclination analysis however, do not support a large amount of inclination error for Moenave Formation strata. An alternative explanation for the 8° discrepancy between the pole positions of the southwest U.S.A. strata (MF4f1) and coeval rocks of eastern seaboard could be due to systematic, localized rotations in structural domains along the ester seaboard, during the ! $&! early phases of rifting of Pangea. Paleomagnetic analysis on the Sassamansville diabase and baked sediments of the Passaic Formation indicate these rocks experienced 15° + 5° - of counterclockwise rotation (Kodama et al., 1994b). The rocks of the Sassamansville syncline are adjacent to the Newark Basin border fault. A counterclockwise rotation of about 15° was also observed in rocks of the Culpeper Basin, south of the Sassamansville syncline (Kodama et al., 1994a). The possibility that rocks of the Newark and adjacent basins have also experienced a significant amount of vertical axis rotation cannot be ruled out. The observed discrepancy between the new paleomagnetic pole position determined from all available Moenave Formation strata corrected for inclination shallowing and Colorado Plateau rotation (MF4f1) and the pole position derived from coeval rocks of the Newark and Hartford basins is well within the observed 15° + 5° degrees of counterclockwise rotation for Mesozoic rift basins of the eastern seaboard. Conclusions We present two new latest Triassic to earliest Jurassic paleomagnetic poles for the Moenave Formation, southern Utah and northern Arizona. One is based on data obtained in this study and is located at 63.1° N, 77.2° E, (A95 = 3.8°, N = 56 sites, f1= 0.78). The other is based on data obtained from this and previous studies and is located at 62.3° N, 68.0° E, (A95 = 7.4°, N = 102 sites, f1 = 0.78). The paleomagnetic data are corrected for inclination bias using a locally derived flattening factor and the paleomagnetic poles are corrected for 4° of clockwise Colorado Plateau rotation about an Euler Pole position of 34° N, 260° E. The location of these paleomagnetic poles remain is support of an approximate position of the J1 cusp defined by the PEP method of APW path ! $'! construction. If the results are correct for a larger magnitude of inclination shallowing, using flattening factors estimated in other studies, the pole positions obtained are statistically indistinguishable from pole positions obtained by the sliding time window methods of APW path construction. The elongation/inclination results from Moenave Formation strata are interpreted to indicate these strata record a small amount of inclination error. The discrepancy in location of paleomagnetic poles from sedimentary rocks of the southwest and northeast U.S.A. could be reconciled by corrections for inclination shallowing, a small (maximum of 4°) magnitude of clockwise Colorado Plateau rotation, and an undefined but relatively small (about 8°) amount of counterclockwise vertical axis rotation of the Newark and adjacent basin. The new paleomagnetic poles obtained from Moenave Formation strata continue to support a westward location for the latest Triassic to earliest Jurassic paleomagnetic pole position for the North American plate. Acknowledgements We thank Kate Zeigler, Roberto Molina-Garza, Larry Tanner, and Andrew Heckert for discussion and field assistance. The manuscript was improved after comments from reviewers and editors. Funding for this project includes grants and scholarships from the Sloan Foundation, the Institute of Geophysics and Planetary Physics at Los Alamos National Laboratory, American Geological Institute Minority Participation Program, and the Department of Earth and Planetary Sciences. ! $(! Figures Ý N N Ý Ý N N Ý Ý N N Ý 190 NR1 Ý N N HB NV M2 200 Ý M1 J1 210 Ý WS N N Ý Ý E E Ý Ý E Ý( Ý E E Ý 60Ý( Ý( 110 7 0Ý( 100Ý( Ý( Ý( Colorado Plateau Region Craton and Eastern Seaboard late Triassic southwest poles late Triassic northeast poles latest Triassic to earliest Jurassic latest Triassic to earliest Jurassic early Jurassic southwest poles early Jurassic northeast poles Paleomagnetic Euler Pole Path Sliding Time Window Path Figure 3.1. Late Triassic to early Jurassic paleomagnetic poles for the North American continent divided by age (symbol) and location (color). Late Triassic poles from the Colorado Plateau area (craton and eastern seaboard) are shown in dark (light) gray squares. Latest Triassic to earliest Jurassic poles from the Colorado Plateau area (craton and eastern seaboard) are shown in orange (green) gray diamonds. Early Jurassic poles from the Colorado Plateau area (craton and eastern seaboard) are shown in dark (light) gay circles. Orange dashed line represents the Paleomagnetic Euler Pole apparent polar wander paths with the J1 cusp shown with a yellow circle. The average pole positions used to construct the sliding time window path are show in green. ! $)! c) d) e) f) Figure 3.2. a) Regional map of southwest US, b) outcrop map of the Moenave Formation in southern Utah and northern Arizona, stars represent sampling sites of this and previous studies, small circles represent nearby cities and towns. c-f) Field pictures from the Moenave Formation localities sampled in this study, Leeds, Washington Dome, Warner Valley, and Potter Canyon. ! $*! Figure 3.3. Stereographic projections of site mean data in stratigraphically corrected coordinates (A), corrected for inclination bias (B), solid circles represent lower hemisphere projection and open circles represent upper hemisphere projections. ! $+! a) b) c) d) 2.8 10.8 [8.4, 13.2] 60 1.0 f=0.73 2.4 0.8 50 0.8 D = 3.5o, I = 8.1o, 2.0 40 0.4 o Elongation 1.6 30 a95 = 2.1 , n = 378 0.2 1.2 f=1 Elongation Direction 20 Cumulative Distribution 0.0 ChRM directions 6 8 10 12 14 16 8 10 12 14 6 8 10 12 14 16 Inclination e) f) g) h) 3.0 17.9 [ 13.2, 23.6] 1.0 f=0.78 40 2.5 0.8 2.0 0.6 D = 4.6o, I = 14.3o, 30 0.4 Elongation 1.5 o f=1 a95 = 3.5 , n = 56 0.2 1.0 Elongation Direction 20 0 10 20 30 Cumulative Distribution 0.0 14 15 16 17 Site Mean directions 10 20 30 Inclination Figure 3.4. Results of elongation/inclination analysis for Moenave Formation data obtained in this study (a-d) and data from this study combined with data from previous studies (c-h). The left column shows a stereographic projection of 378 ChRM directions (top) and a stereographic projection of 102 site mean directions (e). The next column shows the results of the E/I analysis, a unique combination of elongation (#2/#3 of the orientation matrix) and inclination consistent with the TK03.GAD model of the geomagnetic field (b and f). The flattening factor obtained for data obtained in this study is 0.73 and the flattening factor obtained for the combined Moenave Formation data set is 0.78. The next column shows the cumulative distribution function for the corrected inclination (e and g). The last column shows a decrease in the elongation direction with the corrected inclination value (d and h). ! $"! MF4 BC02 VdV93 T08 G84 KI10 BH03 Ý N N Ý Ý N N Ý Ý N M4f2 N M4f2! KI10 Ý BC02 T08 Ý N N Ý M4fM4 1 M4f1! M4 VdV93 Ý N N G84BH03 Ý Ý E E Ý 55Ý( Ý( 105 Ý( Ý( 6 5Ý( 95Ý( Ý( Ý( Ý( ÝE Ý( 2011 Nov 9 23:47:29 NewMFpole.ps Figure 3.5. Comparison of pole positions for latest Triassic to earliest Jurassic time, stars represent Moenave Formation poles, open star represents the uncorrected Moenave pole, light gray start is corrected for locally derived flattening factor, dark gray star represents pole corrected for generic flattening factor, and bold light and dark starts have been rotated for 8° of clockwise Colorado Plateau rotation. North American paleomagnetic Euler pole APW path poles are shown in triangles, (open (Gordon et al., 1984), light gray (Beck and Housen, 2003), dark gray (May and Butler)), and sliding time window APW paths poles are shown in hexagons (open (Irving and Irving, 1982), light gray (Van der Voo, 1993), medium gray (Besse and Courtillot, 2002, 2003), dark gray (Torsvik et al., 2008), closed (Kent and Irving, 2010)). ! $#! Tables Table 3.11. Caption at the bottom of the table. E) 62 67 40 o ( Lon 51.7 65.6 63.1 82.7 74.5 44.9 47.7 29.3 15.2 98.9 60.3 91.8 86.2 74.1 70.6 74.9 33.6 104.8 268.7 114.2 333.6 201.6 349.6 2 f N) VGP 68 79 o Lat ( 39.9 74.6 10.3 56.9 43.1 42.4 59.5 58.9 60.3 62.1 71.6 85.2 79.9 76.1 86.1 45.3 60.5 82.8 58.2 55.8 77.5 64.2 52.3 73.2 E) o ( Lon 50.7 95.1 61.1 65.6 62.9 83.5 74.1 67.0 46.4 49.2 35.2 85.4 12.3 42.6 16.3 96.4 63.3 90.5 85.6 71.8 70.2 74.9 49.2 41.0 293.6 158.8 1 f VGP N) o Lat ( 43.1 68.9 19.4 55.7 45.8 44.9 57.5 57.2 57.9 59.3 63.5 78.7 77.9 76.2 78.6 79.7 44.3 57.7 75.3 56.2 54.7 71.1 61.0 52.4 72.4 67.8 E) o ( Lon 50.0 90.8 26.8 65.6 62.8 84.0 73.8 67.0 47.2 50.0 38.1 79.8 27.6 51.0 61.9 64.1 89.7 85.3 71.0 70.1 71.4 52.6 44.5 340.8 125.8 216.2 VGP N) o Lat ( 45.2 64.7 60.4 54.9 47.6 46.5 56.3 56.1 56.2 57.4 60.3 80.6 72.0 71.1 75.5 73.4 62.5 71.1 69.8 54.8 53.8 66.5 58.9 48.3 67.5 63.8 o 6/6 6/6 8/8 8/8 4/5 4/5 5/5 3/3 5/7 5/5 7/9 7/8 7/7 5/7 2/3 2/4 3/6 4/4 6/6 4/8 6/6 4/4 3/3 7/7 n/n 70/70 12/15 k 36.6 36.1 22.1 39.8 62.9 65.6 92.2 97.6 42.9 27.0 75.3 33.7 39.3 32.5 116.6 110.5 119.0 153.7 143.6 108.6 135.8 681.9 598.2 134.4 449.8 1031.5 ) 95 o ( a 8.9 2.2 2.9 8.0 8.1 6.4 6.4 6.1 9.3 5.8 6.6 9.6 2.0 5.9 8.8 5.1 2.7 3.8 4.8 11.2 11.3 11.2 11.7 12.5 12.8 10.7 2 f I 8.8 8.6 8.1 -0.8 0.52 42.5 13.6 12.1 17.9 20.3 34.3 69.9 53.3 56.3 64.3 55.5 21.3 14.9 42.5 21.7 45.6 39.6 -22.5 -61.2 -18.8 -17.7 48.9 E) o 1 f I 9.3 8.3 6.0 5.9 5.5 -0.5 0.78 31.9 12.4 14.1 24.9 61.7 42.3 45.6 54.6 44.6 14.9 10.2 31.9 15.1 34.7 29.4 -15.7 -51.1 -13.0 -12.2 37.9 N, 246.558 E) o ) o tc o I ( 4.3 7.3 6.5 9.7 4.7 8.0 4.6 -9.6 -0.4 11.1 11.7 11.9 25.9 19.9 55.4 35.4 38.5 47.7 37.6 31.3 25.9 28.4 23.7 -12.4 -44.0 -10.2 ) tc o N, 246.650 ( D o 6.2 0.6 2.6 8.9 4.7 0.9 5.5 9.8 11.7 11.6 10.7 13.9 12.7 33.4 349.6 348.1 356.0 359.8 355.7 345.8 344.4 346.9 349.1 358.2 358.2 354.9 Site Leeds(37.202 187.0 188.0 181.0 101.0 102.0 110.0 112.0 111.0 127B 103.0 109.0 106.0 104.0 Dome, (37.009 Washington 203.0 232.0 230.0 208.0 221.0 218.0 210.0 217.0 216.0 211.0 239B 212.0 215.0 ! $$! Table 3.11. Continued. 28 75 44 52 3.8 8.2 1.9 39.2 59.6 30.6 52.3 17.7 63.1 48.2 54.8 43.6 62.4 61.2 67.2 72.1 79.8 72.4 67.3 78.7 33.6 36.8 334.4 111.5 116.4 117.4 67 62.1 61.3 56.3 50.9 66.5 70.9 72.4 81.3 77.8 68.6 73.6 73.8 70.8 69.3 54.8 49.3 66.8 66.5 68.2 66.4 65.7 45.3 45.3 47.3 42.3 41.8 73.5 48.9 47.2 8.0 41.4 59.8 35.6 54.9 28.8 64.6 27.8 50.9 37.1 73.1 48.1 14.8 55.0 43.4 62.9 62.0 67.2 72.0 81.0 73.2 67.3 81.1 41.3 36.3 50.6 349.4 104.6 108.8 119.1 58.8 58.5 55.3 48.2 62.3 62.6 66.0 67.2 74.1 72.2 64.3 68.1 68.2 65.9 64.7 54.1 49.6 63.0 62.2 64.0 66.2 64.7 49.8 46.8 53.1 51.7 51.1 68.0 49.7 51.2 42.6 59.8 10.5 38.2 56.1 33.9 65.0 35.4 52.3 41.2 72.3 50.0 21.3 55.1 43.2 63.2 67.2 71.4 80.7 72.8 67.3 79.9 36.8 51.2 115.9 358.6 101.0 104.7 242.5 224.8 55.3 56.6 54.7 46.4 59.4 59.5 62.6 63.2 69.0 67.4 61.3 64.1 64.4 62.5 60.9 53.7 49.9 60.4 59.0 60.2 59.8 48.5 43.9 50.0 47.3 46.7 48.9 49.4 -61.2 -64.1 4/4 7/8 4/4 7/7 6/6 7/7 6/6 6/6 7/7 6/6 8/8 7/7 7/7 6/6 4/4 6/6 4/7 6/6 5/5 3/4 5/5 7/7 3/4 8/8 4/4 4/4 3/3 5/5 9/9 11/11 90.0 19.7 76.8 34.1 33.1 37.6 57.4 72.0 42.4 37.5 37.5 98.3 96.0 24.7 68.0 71.3 88.7 76.2 93.0 47.9 48.3 135.8 172.4 177.0 276.4 124.2 327.7 147.8 154.2 1722.6 9.7 6.9 6.9 9.8 5.1 6.9 5.0 6.6 9.4 9.3 6.9 5.5 6.9 5.6 5.6 8.0 5.0 8.1 4.6 1.9 9.2 7.0 11.6 11.1 11.2 13.9 10.1 10.0 13.7 10.4 7.0 4.7 7.7 -2.6 -0.2 52.5 21.0 26.0 31.7 35.2 33.6 42.4 47.2 49.1 30.7 41.2 37.5 34.6 44.6 26.3 36.5 28.4 25.2 24.2 39.9 -9.6 -11.2 -13.8 -21.0 -20.5 E) E) o o 4.7 3.2 5.3 -1.8 -9.5 -7.7 -0.1 41.5 14.6 18.3 22.7 25.6 24.3 31.8 36.3 38.2 22.0 30.8 27.5 25.1 33.8 18.6 26.7 20.2 17.7 17.0 29.6 -6.5 -14.6 -14.2 3.7 2.5 4.1 N, 247.156 -1.4 -7.4 -6.0 -0.1 -5.1 11.5 N, 246.625 o 34.6 14.5 18.1 20.5 19.4 25.8 29.8 31.5 17.5 24.9 22.1 20.1 27.6 14.7 21.4 16.0 14.0 13.4 23.9 -11.5 -11.2 o 3.9 4.9 0.6 6.9 7.7 6.8 1.7 2.3 9.8 11.1 33.9 13.0 35.3 14.2 14.5 12.0 20.9 14.8 19.4 10.3 343.0 357.5 341.1 360.0 357.9 351.1 327.2 356.4 359.9 351.3 214.0 (37.016 Valley, Warner 325.0 324.0 323.0 319.0 312.0 311.0 310.0 318.0 309.0 308.0 307.0 317.0 316.0 315.0 314.0 313.0 306.0 305.0 Canyon,Potter (36.879 413.0 419.0 420.0 421.0 422.0 424.0 425.0 426.0 427B 427.0 428.0 ! $%! Table 3.11. Continued. 48 106 61.2 44.9 38.8 28.6 10.3 67.9 59.5 75.8 74.4 85.1 49.9 67.9 67.2 78.7 86.1 96.1 20.3 29.8 41.6 66.3 74.9 12.4 359.7 277.8 330.7 116.4 79 67 88 52 47 67 69.5 62.4 68.2 56.5 53.7 64.4 73.4 66.1 65.1 65.1 68.6 63.1 75.3 64.6 73.1 75.4 67.5 64.4 61.9 68.6 50.6 66.2 62.3 46.8 50.9 40.0 30.1 18.4 17.3 68.0 60.4 75.1 73.6 83.8 52.3 68.0 67.4 54.1 76.6 81.7 93.2 28.5 50.9 47.9 63.2 51.3 61.9 337.1 100.5 111.4 65.2 59.6 64.1 55.0 52.2 60.4 68.7 62.7 83.9 61.9 61.9 64.5 60.4 63.2 69.6 61.6 81.5 67.9 69.6 63.3 61.1 58.5 64.2 32.4 42.8 68.1 76.5 22.4 31 68 83 62.9 47.9 52.3 40.7 22.9 25.6 27.1 60.9 74.7 73.2 53.5 68.1 67.4 59.9 75.6 79.8 97.7 91.7 32.7 31.4 40.9 66.4 75.1 25.4 108.5 62.3 57.8 61.3 53.9 51.2 57.5 64.7 60.4 79.9 59.8 59.8 61.8 58.5 60.7 65.7 59.6 75.1 64.3 65.6 60.4 58.8 56.1 60.9 50.2 48.4 60.6 51.7 58.8 6/6 7/7 6/7 3/6 3/7 5/7 6/8 6/8 6/8 6/6 4/4 4/7 6/6 6/6 6/6 7/7 6/6 6/6 5/5 6/6 7/8 7/7 4/5 6/7 6/7 4/7 5/5 5/7 53.5 98.0 27.2 10.3 26.3 16.1 16.7 14.9 54.9 37.8 40.6 39.5 20.2 15.2 35.2 73.7 34.4 36.9 77.3 42.5 327.3 288.8 230.3 269.6 266.0 146.9 492.6 182.4 3.7 3.6 9.2 8.1 4.1 9.1 4.1 5.5 9.9 2.7 5.0 8.8 12.5 14.9 23.2 13.3 17.2 23.2 24.6 10.6 10.8 17.4 20.3 10.3 10.8 11.6 15.3 11.9 9.0 -6.4 -4.7 29.4 20.0 28.9 11.9 12.3 36.6 47.7 23.4 63.8 22.0 21.9 27.8 19.6 26.7 38.1 20.7 55.8 35.2 39.0 32.0 24.4 28.7 37.0 26.0 37.7 8.2 8.4 6.2 -4.3 -3.2 20.9 13.9 20.6 26.8 36.7 16.4 54.1 15.4 15.2 19.7 13.6 18.8 28.1 14.4 45.0 25.6 28.8 23.0 17.1 20.4 27.1 18.4 27.7 6.4 6.6 4.8 -3.4 -2.5 16.6 10.9 16.3 21.5 30.2 12.9 47.1 12.1 12.0 15.6 10.7 14.9 22.6 11.3 38.0 20.5 23.2 18.3 13.5 16.2 21.8 14.5 22.3 2.7 7.8 0.3 7.6 3.9 7.4 0.2 0.6 2.4 0.3 10.9 16.0 22.5 23.1 17.7 16.9 21.9 17.0 20.6 356.9 357.8 352.5 356.9 355.3 345.9 348.2 338.8 355.0 Molina-Garza et al, 2003 al, et Molina-Garza W) 111.4° N, (36.1° Cliffs Echo ec6 ec7 ec8 ec9 ec11 ec14 ec15 ec16 ec17 ec18 ec19 ec20 ec21 ec22 ec23 ec24 ec25 ec26 ec27 ec30 ec31 ec32 ec34 1989 Butler, and Ekstrand E) 247° N, (36.75° Cliffs Vermillion DW01 DW02 DW03 MO01 MO02 ! %&! Table 3.11. Continued. 0.5 50.1 50.6 69.5 74.4 41.2 54.8 77.4 45.3 47.3 58.4 52.4 52.9 76.7 45.8 49.2 354.3 325.4 68 76.7 52.9 66.2 53.3 72.6 72.3 55.8 48.1 46.7 60.1 49.1 58.5 59.8 56.7 83.3 66.8 77.1 70.6 63.9 53.0 62.0 53.2 67.3 67.1 55.0 48.9 48.0 58.2 50.0 54.9 57.8 55.6 79.6 62.8 70.9 8.5 9.4 56.5 54.5 71.0 76.0 48.2 58.7 78.7 46.6 48.4 60.5 53.8 55.3 77.9 24.3 50.1 56.2 54 77 52 56.8 69.5 13.9 74.4 49.1 58.4 44.7 46.4 59.1 12.2 54.1 76.2 38.2 49.9 56.5 61 66.2 53.1 58.5 53.2 63.6 63.5 54.4 49.4 48.8 56.8 50.5 52.2 56.5 54.9 73.9 60.2 66.5 ) of 0.78 and 0.53. and 0.78 of ) f 4/6 7/7 7/7 4/6 6/6 3/4 3/4 6/6 6/6 6/6 6/6 6/6 6/6 6/6 6/6 6/6 4/5 7/7 39.9 67.4 28.1 35.6 45.0 57.3 29.5 93.7 129.5 216.1 174.8 101.4 104.4 595.4 192.8 179.9 223.3 #### 8.1 4.1 4.6 9.2 6.6 5.1 2.2 4.8 8.2 5.0 8.9 9.5 4.0 10.7 12.9 11.4 10.1 12.5 0.6 6.0 7.5 -0.6 -6.0 -9.4 -6.4 41.6 29.0 45.3 37.2 35.1 14.1 36.1 14.3 57.2 27.8 42.2 0.4 4.2 9.7 9.9 5.1 -0.4 -4.1 -6.4 -4.4 31.1 20.6 34.5 27.3 25.5 26.4 46.5 19.7 31.7 0.3 3.2 7.6 7.7 4.0 -0.3 -3.2 -5.0 -3.4 25.2 16.4 28.2 21.9 20.4 21.1 39.4 15.6 25.7 4.2 6.3 8.0 3.9 4.3 9.5 7.1 8.3 8.5 4.3 25.6 14.3 13.4 30.7 358.5 355.6 354.2 354.7 MO03 MO04 MO06 MO08 MO09 MO13 MO14 MO15 MO16 MO17 MO18 MO19 MO21 MO22 MO24 MO25 WW02 WW03 positionscalculated (VGP) pole geomagnetic virtual and statistics, Fisher directions, mean Site 3.1. Table ( factor flattening a using data Formation allMoenave from ! %'! Table 3.2. 2' f 66.8 80.5 78.2 66.4 Plon 2' f 64.0 70.8 66.1 67.8 Plat 1' f 66.8 78.5 77.2 68.0 Plon 1' f 61.6 67.1 63.1 62.3 Plat 50.8 57.1 58.2 68.2 Plon' rotated 4°, about 34° N and E 260° 34°N 4°, about rotated 55.7 61.5 58.2 61.3 Plat' 2 f 49.9 57.3 59.0 56.4 Plon 2 f The final colums show the paleomagnetic poles corrected for 4° of of 4° for corrected poles paleomagnetic the show colums final The 62.0 70.1 65.3 64.6 Plat 2=0.53 f 95 4.5 5.8 3.8 7.4 A 1 f 51.1 58.8 60.0 62.6 Plon 1 f 59.8 66.4 62.3 56.6 Plat 1=0.78 f 51.9 59.7 60.1 57.3 Plon Plat 58.2 63.7 60.4 60.6 MF1 MF2 MF3 MF4 Table 3.2. Paleomagnetic poles obtained from strata of the Moenave Formation. Each row repreesnts paleomagnetic data from from data paleomagnetic repreesnts row Each Formation. Moenave the of strata from obtained poles Paleomagnetic 3.2. Table combined MF4= and study, this MF3= 2003, al., et Molina-Garza MF2= 1989, Butler, and Eckstrand followingMF1= studies: the (Plon), longitude and (Plat) latitude in shown positionsare pole Paleomagnetic studies. previous and this from data on based pole respectivley. directions E and N indicate values posiitive E. 260° and N 34° at located pole Euler an about rotation Plateau clockwiseColorado of ! %(! Table 3.3. 2r f 6.9 8.3 9.2 8.2 9.4 5.8 MF4 1r f 8 1.6 2.8 6.3 8.9 10.5 MF4 2 f 5.9 7.9 11.7 12.7 14.1 10.7 MF4 1 f 4.8 4.7 10.5 14.7 16.3 14.2 MF4 95 - - 3.2 4.3 3.8 A 81 66.5 69.5 86.8 90.5 81.8 Plon 61 Plat 60.9 59.6 65.9 66.4 67.8 Reference Gordon et al 1984 (J1 cusp) (J1 1984 al et Gordon 2003 Housen and Beck 1993 Voo der Van 200Ma) 4 (Table 2008 etal Torsvik (10myr) Courtillot 2002 and Besse 2010 Irving Kent Moenave the positionsand pole paleomagneic Jurassic earliest to Triassic latest between distances Angualr 3.3. Table 3.2 Table in presented MF4 poles Formation ! %)! References Beck, M. E., and Housen, B. A., 2003, Absolute velocity of North America during the Mesozoic from paleomagnetic data: Tectonophysics, v. 377, no. 1-2, p. 33-54. Besse, J., and Courtillot, V., 2002, Apparent and true polar wander and the geometry of the geomagnetic field over the last 200 Myr: Journal of Geophysical Research- Solid Earth, v. 107, no. B11. Besse, J., and Courtillot, V., 2003, Correction to "Apparent and true polar wander and the geometry of the geomagnetic field over the last 200 Myr": Journal of Geophysical Research-Solid Earth, v. 108, no. B10. Bryan, P., and Gordon, R. G., 1990, Rotation of the Colorado Plateau - an updated analysis of paleomagnetic poles: Geophysical Research Letters, v. 17, no. 10, p. 1501-1504. Bump, A. P., 2003, Reactivation, trishear modeling, and folded basement in Laramide uplifts: Implications for the origins of intra-continental faults: GSA Today, v. 13, no. 3, p. 4-10. Bump, A. P., 2004, Three-dimensional Laramide deformation of the Colorado Plateau: Competing stresses from the Sevier thrust belt and the flat Farallon slab: Tectonics, v. 23, no. 1. Bump, A. P., and Davis, G. H., 2003, Late Cretaceous-early Tertiary Laramide deformation of the northern Colorado Plateau, Utah and Colorado: Journal of Structural Geology, v. 25, no. 3, p. 421-440. Cardozo, N., 2011, OSXStereonet. ! %*! Cather, S. M., 1999, Implications of Jurassic, Cretaceous, and Proterozoic piercing lines for Laramide oblique-slip faulting in New Mexico and rotation of the Colorado Plateau: Geological Society of America Bulletin, v. 111, p. 849-868. Donohoo-Hurley, L. L., Geissman, J. W., and Lucas, S. G., 2010, Magnetostratigraphy of the uppermost Triassic and lowermost Jurassic Moenave Formation, western United States: Correlation with strata in the United Kingdom, Morocco, Turkey, Italy, and eastern United States: Geological Society of America Bulletin, v. 122, no. 11-12, p. 2005-2019. Ekstrand, E. J., and Butler, R. F., 1989, Paleomagnetism of the Moenave Formation - Implications for the Mesozoic North American apparent polar wander path: Geology, v. 17, no. 3, p. 245-248. Gordon, R. G., Cox, A., and Ohare, S., 1984, Paleomagnetic Euler poles and the apparent polar wander and absolute motion of North America since the Carboniferous: Tectonics, v. 3, no. 5, p. 499-537. Griffiths, D. H., King, R. F., and Wright, A. E., 1957, Some field and laboratory studies of the depositional remanence of recent sediments: Advances in Physics, v. 6, no. 23, p. 306-316. Hamilton, W. G., 1981, Plate tectonic mechanism of Laramide deformation: Contributions to Geology, v. 19, p. 87-92. Hamilton, W. G., 1988, Laramide crustal shortening, in Schmidt, C. J., and Perry, W. J., eds., Interactions of the Rocky Mountain Foreland and the Cordilleran Thrust Belt, Volume 171: Boulder, Geological Society of America Memoir, p. 27-39. ! %+! Harrison, C. G. 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E., 2008, Early Jurassic magnetostratigraphy and paleolatitudes from the Hartford continental rift basin (eastern North America): Testing for polarity bias and abrupt polar wander in association with the central Atlantic magmatic province: Journal of Geophysical Research-Solid Earth, v. 113, no. B6. Kent, D. V., and Tauxe, L., 2005, Corrected Late Triassic latitudes for continents adjacent to the North Atlantic: Science, v. 307, no. 5707, p. 240-244. ! %"! Kent, D. V., and Witte, W. K., 1993, Slow apparent polar wander for North America in the late Triassic and large Colorado Plateau rotation: Tectonics, v. 12, no. 1, p. 291-300. King, R. F., 1955, The remanent magnetism of artificially deposited sediment: Monthly Notice of the Royal Astronomical Society Geophysical Suppliment, v. 7, p. 115- 134. Kirkland, J. I., and Milner, A. R. C., 2006, The Moenave Formation at the St. George Dinosaur Discovery Site at Johnson Farm, St. George, southwestern Utah, in Harris, J. D., Lucas, S. G., Spielmann, J. A., Lockley, M. G., Milner, A. R. C., and Kirkland, J., eds., The Triassic-Jurassic Terrestrial Transition, Volume 37: Albuquerque, New Mexico Museum of Natural History and Science Bulletin, p. 289 - 309. Kodama, K. P., Cioppa, M. T., Sherwood, E., and Warnock, A. C., 1994a, Paleomagnetism of baked sedimentary rocks in the Newark and Culpeper basins - Evidence for the J1 cusp and significant late Triassic apparent polar wander from the Mesozoic basins of North America: Tectonics, v. 13, no. 4, p. 917-928. Kodama, K. P., Hedlund, C., Gosse, J., and Strasser, J., 1994b, Rotated paleomagnetic poles form the Sassamansville syncline, Newark Basin, southeastern Pennsylvania: Journal of Geophysical Research-Solid Earth, v. 99, no. B3, p. 4643-4653. Lucas, S. G., and Tanner, L. H., 2007, The nonmarine Triassic-Jurassic boundary in the newark supergroup of eastern North America: Earth-Science Reviews, v. 84, no. 1-2, p. 1-20. ! %#! Lucas, S. G., Tanner, L. H., Donohoo-Hurley, L. L., Geissman, J. W., Kozur, H. W., Heckert, A. B., and Weems, R. E., 2011, Position of the Triassic-Jurassic boundary and timing of the end-Triassic extinctions on land: Data from the Moenave Formation on the southern Colorado Plateau, USA: Palaeogeography Palaeoclimatology Palaeoecology, v. 302, no. 3-4, p. 194-205. Molina-Garza, R. S., Acton, G., and Geissman, J. W., 1998, Carboniferous through Jurassic paleomagnetic data and their bearing on rotation of the Colorado plateau: Journal of Geophysical Research, v. 103, no. B10, p. 24179-24188. Molina-Garza, R. S., Geissman, J. W., and Lucas, S. G., 2003, Paleomagnetism and magnetostratigraphy of the lower Glen Canyon and upper Chinle Groups, Jurassic-Triassic of northern Arizona and northeast Utah: Journal of Geophysical Research-Solid Earth, v. 108, no. B4. Molina-Garza, R. S., Geissman, J. W., and Vandervoo, R., 1995, Paleomagnetism of the Dockum Group (Upper Truassic). Northwest Texas - Further evidence for the J-1 cusp in the North American apparent polar wander path and implications for rate of Triassic apparent polar wander and Colorado Plateau rotation: Tectonics, v. 14, no. 4, p. 979-993. Pipiringos, G. N., and O'Sullivan, R. N., 1978, Principal unconformities in Truiassic and Jurassic rocks, western interior United States - a preliminary survey: U.S. Geological Survey Professional Paper, v. 1035-A, p. 1-29. Steiner, M., 1986, Rotation of the Colorado Plateau: Tectonics, v. 5, p. 649-660. ! %$! Steiner, M., 1988, Paleomagnetism of the Late Pennsylvanian and Permian: A test fo the rotation of the Colorado Plateau: Journal of Geophysical Research, v. 93, p. 2201- 2215. Steiner, M. B., and Lucas, S. G., 2000, Paleomagnetism of the Late Triassic Petrified Forest Formation, Chinle Group, western United States: Further evidence of "large" rotation of the Colorado Plateau: Journal of Geophysical Research-Solid Earth, v. 105, no. B11, p. 25791-25808. Tanner, L. H., and Lucas, S. G., 2007, The Moenave Formation: Sedimentologic and stratigraphic context of the Triassic-Jurassic boundary in the Four Corners area, southwestern USA: Palaeogeography Palaeoclimatology Palaeoecology, v. 244, no. 1-4, p. 111-125. Tauxe, L., 2008, PmagPy: San Diego, Scripps Institude of Oceanography. Tauxe, L., and Kent, D. V., 1984, Properties of detrital remanence carried by hematite from study of modern river deposits and laboratory redeposition experiments: Geophysical Journal of the Royal Astronomical Society v. 77, p. 543. Tauxe, L., and Kent, D. V., 2004, A simplified statistical model for the geomagnetic field and detection of shallow bias in paleomagnetic inclinations: Was the ancient magnetic field dipolar?, in Channell, J. E. T., Kent, D. V., and Lowrie, W., eds., Timescales of the Paleomagnetic Field, Volume Geophysical Monograph 145: Washington, DC, American Geophysical Union, p. 101-115. Tauxe, L., Kodama, K. P., and Kent, D. V., 2008, Testing corrections for paleomagnetic inclination error in sedimentary rocks: A comparative approach: Physics of the Earth and Planetary Interiors, v. 169, no. 1-4, p. 152-165. ! %%! Torsvik, T. H., Muller, R. D., Van der Voo, R., Steinberger, B., and Gaina, C., 2008, Global plate motion frames: Toward a unified model: Reviews of Geophysics, v. 46, no. 3. Van Alstine, D. R., and Boer, J. D., 1978, New technique for constructing apparent polar wander paths and revised Phanerozoic path for North America: Geology, v. 6, no. 3, p. 137-139. Van der Voo, R., 1993, Paleomagnetism of the Atlantic, Tethys and Iapetus Ocean, Cambridge, Cambridge University Press. Wawrzyniec, T. F., Ault, A. K., Geissman, J. W., Erslev, E. A., and Fankhauser, S. D., 2007, Paleomagnetic dating of fault slip in the Southern Rocky Mountains, USA, and its importance to an integrated Laramide foreland strain field: Geosphere, v. 3, no. 1, p. 16-25. Woodward, L. A., Anderson, O. J., and Lucas, S. G., 1997, Mesozoic stratigraphic constraints on Laramide right slip on the east side of the Colorado Plateau: Geology, v. 25, p. 843-846. ! '&&! CHAPTER 4 ROCK MAGNETIC CHARACTERISTICS AND PALEOMAGNETISM OF A MID-PLEISTOCENE SEDIMENT CORE, VALLES CALDERA, NEW MEXICO Abstract Rock magnetic and paleomagnetic data are used to interpret the depositional environment from the mid-Pleistocene, ca. 552 to 363 ka lake sedimentary record (core VC-3) from the Valles Caldera, New Mexico. Mixtures of iron-bearing minerals are present throughout the VC-3 sedimentary record including detrital and diagenetic phases. Primary detrital phases including magnetite and titanomagnetite are present throughout the VC-3 record. Glacial environments in the lake are typically cool and wet and preserve a greater proportion of diagenetic phases indicating occasional anoxic conditions at the sediment water interface. Interglacial environments are generally characterized as warm and dry climate with relatively more shallow lake levels, and have a greater proportion of detrital phases although occasional reducing conditions are present. Sediments deposited during the upper part of stage 13 are characterized by very high magnetic abundance that results from the oxidation of previously reduced sediment due to a drop in lake levels. Discrete intervals of superparamagnetic magnetite preserved in sediment deposited during MIS 12 result from short-lived anoxic environments due to increased lake levels. The anoxic environments of MIS 12 occur shortly after interstadial events identified in the paleoclimate records. Decreased magnetic abundance is observed in sediment deposited during warm substages of MIS 11. Primary detrital phases are identified as the dominant magnetic phase over the interstadial intervals and decreased ! "#"! magnetic abundance is attributed to dilution with biogenic silica. Although the lake experienced temporary anoxic conditions, primary phase are present through the core and most likely preserve a record of mid-Pleistocene geomagnetic field activity. Paleomagnetic results mostly yield magnetization directions consistent with Brunhes normal polarity. At least three intervals yield magnetization directions consistent with geomagnetic field phenomena 14! at about 536 ka and 11! at about 400 ka, and Levantine at about 360-370 ka. The NRM record shares statistical significance with the ARM, IRM, and " records. Consistency among these records suggests the relative paleointensity record mostly reflects changes in the magnetic field and these records are not contaminated by variability in the climate. The relative paleointensity records do not correlate with the mean annual temperature record (MAT) lending further support to the viability of these records. The VC-3 relative paleointensity record shows some similarities with published paleointensity records (e.g., Sint-800), particularly during about 536 ka, about 400 ka, and about 360-370 the time intervals spanning geomagnetic excursion 14!, 11!, and Levantine. Introduction Lake sediments can provide an excellent record of environmental change in terrestrial settings, although the detailed interpretation of their environmental magnetic records is often complicated by natural variability in abundance, grain size, shape, composition of magnetic phases and post depositional alteration of these phases (Vriend et al., 1991; Manning et al., 1999; Walker and Owen, 1999; Stevens et al., 2000; Sifeddine et al., 2001; Foster et al., 2008; Ao et al., 2010; Lascu et al., 2010). Lake ! "#$! sediments of the South Mountain lake, Valles Caldera, New Mexico, are ideal for an environmental magnetic study because the sediment sources are limited to the andesitic Valles Caldera wall, the rhyolitic resurgent domes, and the lake catchment. The Valles Caldera wall and resurgent domes yield magnetite and titanomagnetite. Sediments in the lake catchment are susceptible to pedogenic enhancement and typically consist of maghemite, titanomaghemite, and oxyhydroxides such as goethite. Diagenetic formation of magnetic phases is known to occur at the expense of detrital iron oxides (e.g. magnetite, titanomagnetite, maghemite, oxyhydroxides) in anoxic environments (Anderson and Rippey, 1988; Maher, 1988; Peck et al., 1994; Thouveny et al., 1994; Roberts et al., 1996; Williams et al., 1998; Evans and Heller, 2003; Frederichs et al., 2003; Foster et al., 2008; Ao et al., 2010). Several oxic states can exist in a lake through time (Figure 4.1). Oxic conditions occur when the lake is well mixed and magnetic phases are most commonly detrital. Anoxic environments occur when bottom waters become depleted in oxygen and reduced magnetic phases precipitate. Anoxic conditions can develop during glacial intervals as a result of deep lake stratification or during interglacial intervals due to lake stagnation. Another depositional environment occurs when the anoxic state is interrupted and the reduced material becomes oxidized. Based on a number of theoretical studies, several methods have been proposed for the interpretation of rock magnetic data, with the goals of estimating grain size and shape (Day et al., 1977; Dunlop, 2002a, b; Tauxe et al., 2002; Dunlop and Carter-Stiglitz, 2006; Jackson and Solheid, 2010), quantifying ferrimagnetic components (Lascu et al., 2010), and modeling redox cycles in lake sediments (Rosenbaum et al., 1996; Maher and Thomposon, 1999; Evans and Heller, 2003; Egli, 2004). Although these models provide ! "#%! a context for interpreting environmental parameters of lake sediments, it is important to recognize that natural systems can have a strong non-linear response to climate change (Tarduno, 1995; deMenocal et al., 2000; Egli, 2004). Variations in environmental magnetic properties in any setting result from a combined response to long-term climate change and local, possibly higher frequency changes to the catchment environment (Fischer et al., 2008). The mechanisms by which rock magnetic properties change in response to changes in the depositional environment are poorly understood (Ao et al., 2010). Environmental magnetism has been applied to many lacustrine settings yet few studies have focused on caldera lake environments (Creer and Thouveny, 1996; Williams et al., 1996; Mingram et al., 2004; Chaparro et al., 2008). Rock magnetic records from crater lakes have been used to validate geomagnetic field paleosecular variation data (Chaparro et al., 2008) (lake Hoya San Nicolas, Mexico), and study external climate (Mingram et al., 2004) (Asian summer and winter monsoon and East Asian dust storm recorded in the Long Gang Volcanic Field) with little focus on the variability of lake environments. The amount of oxygen and biologic activity has a profound effect on the oxidation state of iron and the composition of iron-bearing phases found in lake environments. The environmental magnetism record of the South Mountain lake sediments is presented in this chapter. The South Mountain lake is the only ~ 200 ka duration record of mid-Pleistocene lake variability and climate change for the southwest U.S.A. (Fawcett et al., 2011). Rock magnetic results are compared to the paleoclimate interpretation of Fawcett et al., (2011) to investigate how the depositional environment of ! "#&! the lake changed through the mid-Pleistocene and over MIS 14 through MIS 10 glacial/interglacial cycles. Magnetic property information bearing on the concentration, grain size, and composition of magnetic phases is used to describe the depositional environments of the South Mountain lake. The lithology and age interpretation of the South Mountain lake sediments are described below along with a description of the rock magnetic and paleomagnetic methods used to analyze the VC-3 (Valles Caldera) core sediments. The rock magnetic results and paleomagnetic data are presented and the implications discussed for the preservation of mid-Pleistocene geomagnetic field activity. Methods South Mountain Lake Sedimentary Record and Analytical Methods There are five paleo lake sequences are presenting the Valles Caldera (Goff, 2009; Goff et al., 2011). The South Mountain lake initiated after the eruption of the South Mountain rhyolitic dome (552 ka) blocked the eastern drainage of the Valles Caldera, San Diego Canyon (Figure 4.2) (Fawcett et al., 2007; Reneau et al., 2007; Goff, 2009; Fawcett et al., 2011; Goff et al., 2011). Sediment from the South Mountain lake are ideal for an environmental magnetic investigation because the sediment are transported from a limited number of sources. South Mountain lake sediments are transported from the Valles Caldera wall, the rhyolitic resurgent domes, and the soil from the lake catchment. An 80 m sediment core was extracted from the lake beds to investigate the mid-Pleistocene climate record of the southwest US. The sediment types vary throughout the core. The base of the core sequence contains several fine sand and coarse silt intervals. Sediment intervals from the middle to upper parts of the core are ! "#'! silty mud and clay dominated and these intervals consist of finely laminated, deep lake deposits, and shallow deposits with poorly defined laminations. Comprehensive geochemical, palynological, and bacterial lipid studies of the VC- 3 lake sediments provide a record of terrestrial mid-Pleistocene climate change in the southwest, suggesting two intervals of centennial to millennial scale drought during the warmest part of interglacial stages 13 an 11 (Fawcett et al., 2011). Sediment accumulation occurred from about 552 ka to 368 ka and is constrained by an Ar-Ar age determination of 552 + 3 ka from one tephra layer at 75.8 and one pumice deposit at 76.0 m depth, combined with abrupt increases in total organic carbon (TOC), Si/Ti ratios, and mean annual temperature (MAT) estimates of about 7° C at 51.5 m and about 8° C at 27.5 m, respectively, representing glacial terminations V (ca. 531 ka) and IV (ca. 423 ka). The presence of mud cracks during part of interglacial stages 13 and 11, decreased TOC, increased MAT, and abundant Juniper and Oak pollen suggest the climate was warm and dry and the lake became shallow during interglacial times. Millennial-scale climate change is observed during the glacial periods by increased MAT (interstadial) and large changes in the percent of spruce and fir pollen suggesting brief warm periods caused a slight drop in lake level (Fawcett et al., 2011). Sampling and Rock Magnetism Methods Specimen cubes (inner dimensions of 2 x 2 x 2 cm) were extracted from core VC- 3 using non-magnetic tools for rock magnetic and paleomagnetic analysis. The intensity of the natural remanent magnetization (NRM), anhysteretic remanent magnetization (ARM), isothermal remanent magnetization (IRM), and bulk susceptibility ("bulk) are ! "#(! typically directly related to modal abundance of specific magnetic phases. Anhysteretic remanent magnetization (ARM) was imparted to each specimen using a 95 mT demagnetizing field and a 0.1 mT direct current biasing field. The acquisition of isothermal remanent magnetization to complete or near complete saturation (SIRM) was imparted using an impulse magnet capable of producing a DC field close to 3 T. ARM and SIRM were progressively AF demagnetized in 5 mT steps to 20 mT and 10 mT steps to 110 mT. The NRM, ARM, and SIRM were measured on a three axis DC-SQUID based Model 760R Superconducting Rock Magnetometer, interfaced with an on line, automated alternating field demagnetization system at the University of New Mexico. The !bulk parameter was measured on an AGICO KLY-4S Kappabridge at the University of New Mexico. Hysteresis loops, with a 1 T maximum applied field, and DC demagnetization curves to determine coercivity of remanence (Hcr), were measured on a Princeton Measurements Vibrating Sample Magnetometer at the Institute for Rock Magnetism, University of Minnesota, Twin Cities. Hysteresis data include five important rock magnetic parameters: paramagnetic susceptibility ("para), saturation remanence (Mrs), magnetic saturation (Ms), coercive force (Hr), and Hcr. The "para parameter is obtained from the high field slope of the hysteresis loop and is subtracted from !bulk to obtain "ferri. Bulk composition of iron bearing phases can be estimated magnetically using the coercivity ratio (Hcr/Hc) and the S-ratio. The magnetic hardness, as measured by the S- ratio, is obtained by dividing the SIRM with magnetization values at -100 mT from the DC demagnetization curve. The parameters listed above are used alone or as normalizing variables throughout the core to evaluate the possibility of relative changes in abundance, grain size, and composition with depth. ! "#)! Magnetic extracts were obtained by crushing 2 – 10 cm of VC-3 sediment. To extract magnetic phases, samples were mixed in water and cycled through a covered Sm- Co magnet for 24 hours. The high temperature susceptibility ("hT) was measured by continuous heating to 700° C and cooling to room temperature on a Geofyzika KLY-2 Kappabridge at the Institute for Rock Magnetism. Frequency dependent susceptibility ("fd) was measured using a LakeShore Cryotronics AC Susceptometer at the Institute for Rock Magnetism, in a constant 240 A/m field at 40 Hz, 100 Hz, 400 Hz, 1000 Hz, and 4000 Hz frequencies as the sample was exposed to temperatures from 9 K to 300 K. Low temperature remanence experiments, measured on a Quantum Designs Magnetic Property Measurement System Cryogenic Susceptometer at the Institute for Rock Magnetism, include data for field cooled remanence (FCR), zero-field magnetization (ZFM), room temperature cooling (RTC), and room temperature warming (RTW). The FCR results are obtained by cooling the sample to 10 K in the presence of a 2.5 T field, switching off the field and measuring the remanence as the sample warms to room temperature. The ZFM data were obtained on samples cooled to 10K in zero-field then subjected to a 2.5 T pulse field and measured as the sample warms to room temperature. In the RTC and RTW experiments the sample was exposed to a 2.5 T field at room temperature and measurements were made as the sample was cooled to 10 K (cooling) and subsequent warming to room temperature (warming). Extracts from 68.16 m, 30.51 m, 18.10 m, and 11.76 m were then placed on a glass slide and covered with a carbon film for analysis on a Jeol 5800LV Scanning Electron Microscope (SEM) with an Oxford Analytical ultra- thin window energy dispersive spectroscopy detector, at the University of New Mexico. Extracts from 43.50 m were mounted in epoxy for analysis on a FEI Quanta 3D Field ! "#*! Emission Gun Scanning Electron Microscope/ Focused ion beam instrument (FEGSEM) with energy dispersive x-ray detector and electron backscatter diffraction system for phase identification, at the University of New Mexico. This spectrum of high (up to 700 C), low (down to 10 K), and room temperature experiments and microscopic techniques help to both establish relative changes in magnetic properties with depth and provide a detailed evaluation of magnetic extracts to characterize the abundance, grain size, and composition of the South Mountain lake sediments. Paleomagnetism and Paleointensity Methods The paleomagnetic and relative paleointensity results from South Mountain sediments yield data that can potentially be used to better characterize the mid- Pleistocene geomagnetic field. Alternating field (AF) demagnetization of NRM was carried out in steps of 2 to 4 milliTesla (mT) up to 20 mT, and steps of 5 mT to 100 mT. Progressive thermal demagnetization of sediment samples proved to be ineffective due to the presence of thermally unstable phases, including ferri oxyhydroxides, iron sulfides, iron carbonates, and phosphates (Donohoo-Hurley et al., 2007; Fawcett et al., 2007) that can alter to new magnetic phases (magnetite, maghemite, or hematite) when heated in air (Dunlop and Ozdemir, 1997; Frederichs et al., 2003). Magnetization components were analyzed using principal component analysis (PCA) (Kirschvink, 1980) on demagnetization segments involving at least four data points not including the origin. Only inclination data are presented because each core segment of VC-3 was not oriented to north. To construct relative paleointensity curves, the ARM, SIRM, and/or "bulk records are used to normalize the NRM intensity. The procedure yields a measure of ! "#+! relative paleointensity provided that sediment consists of uniform magnetic phases and that variability in magnetic parameters is due to changes in the strength of the magnetic field. If the relative paleointensity record is influenced by changes in the climate, the resulting fluctuations could be interpreted as changes in field strength (Kent, 1973; Tucker, 1981; Tauxe, 1993; Brachfeld and Banerjee, 2000; Valet, 2003; Haltia-Hovi et al., 2011). We test for coherence between NRM intensity and normalized intensity parameters to identify intervals of magnetic uniformity and use the relative paleointensity record to test for a correlation with climate signal using the mean annual temperature record from (Fawcett et al., 2011) to identify and possible intervals that may be contaminated by climate influence (Tauxe, 1993; Valet, 2003). Results Environmental Magnetic Record Rock magnetic experiments were carried out on 339 samples (each sample represents a two centimeter thick interval of core) and yield highly variable results with depth (Figures 4.3 – 4.6). The mean values of rock magnetic parameters are listed in Table 4.1. The abundance of magnetic phases is approximated by several magnetic parameters, NRM, ARM, IRM, Ms and !ferri, and is highly variable throughout the core. Increased values in the abundance parameters are observed in sediment deposited during MIS 13 and 12. Decreased values in abundance parameters are observed in sediment deposited during MIS 11 and 12. The ARM/SIRM, SIRM/!bulk, and !bulk/Ms records approximate the grain size distributions of magnetic phases (Figure 4.4). High coercivity, single domain magnetite grains (0.05 to 0.1 "m) are the most efficient ! ""#! magnets and yield high ARM/SIRM values. Enhanced ARM/SIRM values of up to about 9.00 x 10-2 characterize sediments from the upper part of interglacial MIS 13 (about 45.40 m) and the middle part of interglacial MIS 11 (about 20.09 cm). Increased SIRM/!bulk values indicate the presence of fine-grained, very high coercivity phases (i.e. hematite, goethite) (Heider et al., 1996; Ortega et al., 2006). Ratios are typically higher during glacial intervals and inversely correlated with the ARM/SIRM record. The greatest variability in the SIRM/!bulk record is observed in glacial intervals MIS 14 and 10, with values ranging from 0.67 A/m to 31.27 A/m. Increased values of "bulk and "bulk/Ms are commonly interpreted to reflect a strong contribution of superparamagnetic grains (Maher 1988, Heider et al., 1996, Peters and Dekkers, 2003). The greatest -3 variability in the !bulk/Ms ratio, between 4.05 x 10 m/A and 15.84 m/A, is observed in sediments deposited during MIS 13, MIS 12, and MIS11. In total, the grain size proxies suggest that interglacial intervals contain more fine- grained low coercivity minerals (i.e. magnetite) while sediments from glacial intervals have a greater contribution of fine-grained high coercivity minerals (i.e. hematite). There appears to be a greater contribution by superparamagnetic magnetite grains in the middle part of the core in sediments deposited during MIS 12 and MIS 11, and this is not related to glacial stage. The Mrs/Ms, Bcr/Bc, and the s-ratio vary as a function of grain size, grain shape, and mineralogy (Figure 4.5). The greatest variability in both the Mrs/Ms and Bcr/Bc ratios records occurs during interglacial intervals with Mrs/Ms values ranging from 0.11 to 0.50 and Bcr/Bc values ranging from 1.30 to 6.15. At glacial termination VI, the Bcr/Bc values double and remain high through interglacial MIS 13. Values decrease to about 1.70 at the lower part of glacial MIS 12 and slowly increase to the maximum value ! """! of 6.15 in the middle part of interglacial MIS 11. The Bcr/Bc values from glacial MIS 10 are similar to those observed in glacial MIS 14. S-ratio values close to 1, which are indicative of low coercivity minerals (i.e. magnetite), are most common during interglacial MIS 13 and 11 but high S-ratio values are also present during glacial MIS 14 (6752 cm, 5758 cm) and MIS 12 (3182 cm). Intervals with relatively high S-ratio values typically yield increased coercivity ratios, with the exception of at least one interval from glacial MIS 14 and interglacial MIS 13, and most of glacial MIS 12. On average, data from sediments from glacial interval MIS 14 suggest a higher abundance of fine-grained, high coercivity phases. Sediments in interglacial MIS 13 interval show the highest variability in all proxy records. Sediments from glacial MIS 12 exhibit a range of superparamagnetic to fine-grained sizes and a mixture of both hard and soft coercivity minerals. Sediments from interglacial 11 yield superparamagnetic to fine grained soft coercivity minerals with occasional drops in abundance. Sediments from glacial MIS 10 yield a mixture of soft coercivity fine-grained and high coercivity minerals that are slightly coarser, with moderate variations in concentration. General trends in grain size, grain shape, distribution of grain sizes and composition are examined in the following section. Hysteresis data have often been used to interpret the grain size distribution and approximate grain shape of assemblages of moderate to low coercivity phases. The Day Plot and magnetic mixing theory use the coercivity and remanence ratios (or squareness) to examine the grain size distribution of magnetite in natural samples and evaluate artificial mixtures of single domain (SD), pseudo single domain (PSD), and multi-domain (MD) grain sizes (Day et al., 1977; Dunlop, 2002a, b; Dunlop and Carter-Stiglitz, 2006). The VC-3 bulk sediments yield ! ""$! data that imply dominance by PSD behavior on the Day plot, with Bcr/Bc values ranging from 1.30 to 6.15 and Mrs/Ms values ranging from 0.11 to 0.50 (0.1 – 20 µm) (Dunlop and Ozdemir, 1997). The Mrs/Mr results are consistent with the range of values expected for SD-MD mixtures but Bcr/Bc values are greater than expected due to a combination of increased Bcr (greater titanium content in magnetite and/or maghemite) and decreased Bc (changes in grain shape due to oxidation) (Figures 4.6a – 4.10a). The shift in Bcr/Bc values could also be due to a contribution of large superparamagnetic grain sizes (0.02 µm) (Dunlop and Ozdemir, 1997). The increased Bcr/Bc values are present throughout the core but are most pronounced in sediment deposited during MIS 12 and 13. The squareness-coercive force plot can be used to investigate mixtures of PSD magnetite with the following combination of grain shapes: uniaxial, uniaxial plus superparamagnetic, cubic, and cubic plus superparamagnetic grains (Tauxe et al., 2002; Wang and Van der Voo, 2004) (Figure 4.3). The VC-3 squareness-coercive force data indicate the most abundant grain shapes are uniaxial single domain plus superparamagnetic grain shape (Figures 4.6b – 4.10b). Sediment deposited during MIS 12 and 13 show a greater contribution from uniaxial single domain grains (Figures 4.8b and 4.9b). Cubic grain shapes are more common in sediment deposited during the interglacial intervals MIS 11 and 13 (Figures 4.7b and 4.8b). The distribution of grain sizes can be explored by plotting the ! versus the !ARM (King et al., 1982). A trend with a steep slope on King Plot indicates distribution of fine-grain particle while shallow slopes are consistent with a more coarse grain distribution. The ! versus the !ARM values for the Valles Caldera sediment are very low making difficult the interpretation of grain size distributions (Figures 4.6c – 4.10c). Moderate grain size distributions are most common in sediments ! ""%! deposited during MIS 14 and 10. Sediments from the middle potion of the core, sediments deposited during MIS 13 through MIS 11, exhibit a mixture of fine and coarse grain sizes. The greatest variability in grain size distribution is observed in MIS 13 (Figure 4.9c). The IRM acquisition and BF curves indicate magnetic material is mostly saturated (at least 90 percent) by ~1 T consistent with magnetite as the dominate magnetic phase regardless of grain size or shape (Figures 4.6d – 4.10d). The sediment deposited during MIS 14 saturates by ~2 T indicating material from this interval may have a greater contribution of iron and titanium bearing phases (Figure 4.10 d). Magnetic Mineralogy Investigation of the composition of magnetic phases was carried out using high and low temperature experiments on bulk samples and magnetic extracts. Diagnostic ordering temperatures for magnetic phases were obtained by measuring susceptibility as a function of elevated temperature ("hT) (Figure 4.11). In general susceptibility curves at elevated temperature exhibit a large (up to 90 percent) loss in remanence between 560° C and 590° C, close to the Curie point for magnetite, and the inferred low temperature range for low-Ti maghemite (estimated to be between 590 and 675° C). Susceptibility also increases near about 200° C, expected for a phase change from hexagonal to monoclinic pyrrhotite. An increases in susceptibility near 200° C is also consistent with the oxidation of goethite to magnetite. Magnetic separate obtained from sediment over the interval between 7.1 m and 13.8 m exhibit an ~100 percent increase in susceptibility at room temperature suggesting new magnetic phases form during heating. ! ""&! The AC susceptibility frequency and field dependence experiments yield in phase and quadrature components of susceptibility (!´ and !´´) measured between 10° K and 300° K (Figure 4.13). The !´ results are a function of both composition and grain size (Figure 4.13, only 40 Hz, 400 Hz, and 4000 Hz are displayed for simplicity). Frequency dependent !´´ results arise from electric conductivity in the presence of sulfides and time- dependant viscous behavior (magnetic relaxation), which is particularly sensitive to superparamagnetic grain size distributions. Frequency independent susceptibility results from low field hysteresis associated with Ti-rich samples (Lagroix et al., 2004; Carter- Stiglitz et al., 2006). A magnetic separate obtained from sediment between 10 and 17 m depth exhibit an increase in !´ from 10° to 120° K, a broad peak between 120° and 260° K, and a slight decrease between 260° and 300° K consistent with superparamagnetic magnetite (Figure 4.12). This magnetic separate also exhibits maximum frequency dependent !´´ near 30° K at 40 Hz, 40° K at 400 Hz, and 50° K at 4000 Hz suggestive of titanomagnetite. Low temperature remanence experiments can be used to identify magnetic phases based in characteristic transition. The Verwey transition at 120° K is diagnostic of magnetite and is observed in low temperature remanence results from sediments at 35.01 m and 37.11 m, although the transition is muted (Figure 4.13). A diagnostic transition expected for vivianite is observed at 19° K in the derivative of the data from 35.01 m depth. Low temperature remanence transition are also observed at 35° and 40° K in date from 36.36 and 37.11 m consistent with rhodochrosite and siderite. Magnetic separates from sediments obtained at depths of 68.16 m, 43.50, 30.51 m, 18.10 m, and 11.76 m. These depths were chosen based as characteristic behavior magnetic behavior for sediments deposited in each MIS stage (14 – 10). Separates were ! ""'! mounted on glass slides covered with a carbon film and examined using scanning electron microscopy (SEM) (Figure 4.14) and electron backscatter diffraction (EBSD) and a magnetic separate obtained from sediment at 43.50 m was examined by focused emission gun scanning electron microscopy (FEGSEM) (Figure 4.15). SEM inspection combined with wavelength dispersive energy provides a first order estimate of chemical composition and FEGSEM provides information on chemical and crystal structure. EBSD analysis on an SEM yields relative proportions of major elements but does not yield enough information to make a definitive identification of some magnetic phases. The magnetic separates obtained show a wide variety in composition and grain sizes of magnetic extracts (Figure 4.14a-c). All samples include magnetite, titanomagnetite, and maghemite. Sample 68.16 cm also contains iron sulphides. Hematite is observed in separates from 30.51. The hematite appears thin and fibrous as if it precipitated in the lake. The magnetite grain from sample 30.51 m has a broken side and rounded edges. A broken euhedral magnetite and maghemite, resulting from oxidation of magnetite, are present at 18.10 m. Individual titanomagnetite grains from 68.16 m, 30.51 m, and 18.10 m have visible lamellae inferred to be ilmenite resulting from high temperature oxidation exsolution. The titanomagnetite from 68.16 m has a grain diameter of about 50 µm, a smooth outer edge, and fractures across the center of the grain (Figure 4.14g). The titanomagnetite grain from sediment at 30.51 has an elongate grain shape with smooth edges and a weathered appearance (Figure 4.14h). Rounded grain shape and surface cracks in grains from sediment at 18.10 m show evidence of grain weathering and oxidation (Figure 4.14i). A magnetic separate from sediment at 43.50 m was mounted in epoxy and polished to expose the extracted grains and analyzed with an FEGSEM and ! ""(! EBSD. Results of FEGSEM and EBSD support the presence the following reduced phases in the same sample: siderite, pyrrhotite or mackinawite, and rhodochrosite (Figure 4.15 a-c). Overall, electron microscopy inspection of magnetic extracts show that the VC- 3 sediments contain a variety composition including magnetite, maghemite, hematite, titanomagnetite, siderite, pyrite or mackinawite, and rhodochrosite, grain shape ranging from elongate to round, and sizes ranging from 100’s to 10’s of µm. Paleomagnetism and Relative Paleointensity Response to progressive AF demagnetization by the VC-3 sediments yields mixed results (Figure 4.16). About 77 percent of samples yield moderate- to well-defined demagnetization trajectories that show the removal of a moderate positive inclination magnetization (normal polarity) (e.g. Figure 4.16) consistent with magnetization acquisition during the middle of the Brunhes chron. About 21 percent of the samples yield relatively poorly defined demagnetization trajectories. The remaining 2 percent have demagnetization trajectories that resolve a magnetization of negative to very shallow inclination magnetization preserved over five very narrow intervals (Figure 4.16). Three of these intervals yield reasonably well-defined moderate negative inclination demagnetization trajectories represented by samples 52.34 m, 17.20 – 17.22 m, and 5.60 – 5.62 m. Samples 52.34 m and 5.60 m show well-defined stable end-point behavior. The demagnetization trajectory of sample 17.22 m trends from the origin above 60 mT and shows a progressive increase in magnetization above 70 mT. Samples 17.20 and 5.62 m yield a well-defined demagnetization trajectory with a moderate to steep negative inclination with no stable end point. The youngest sedimentary interval ! "")! that records negative polarity is recorded by two samples 5.60 and 5.62. The second youngest sediment interval that records negative inclination consists of 12 samples but only three, 16.99 m, 17.22 m, and 17.24 m show the adequate resolution of a magnetization of negative inclination, the other eight samples record normal polarity. Samples 17.24 and 17.22 yield moderately defined demagnetization trajectories with moderate negative inclination that do not reach the origin by 110 mT. Sample 16.99 m yields a moderate-defined demagnetization trajectory with a steep negative inclination that trends from the origin above 52 mT and shows an increase in intensity above 68 mT. Sedimentary records can be used to estimate relative paleointensity assuming the magnetic variation preserved in sedimentary records reflect the strength of the magnetic field at the time of deposition. If changes in relative paleointensity are a direct function of the geomagnetic field strength then lower values for relative paleointensity likely coincide with geomagnetic excursions (Tauxe, 1993; Macri et al., 2010; Haltia-Hovi et al., 2011). If variations in relative paleointensity are controlled by changes in concentration and grain size a reliable paleointensity curve cannot be constructed. The NRM record has statistically significant correlations to the ARM (r = 0.71, r2 = 0.50, t = 2 2 6.97), SIRM (r = 0.70, r = 0.49, t = 6.88), and "bulk (r = 0.30, r = 0.09, t = 2.91) parameters (Figure 4.17). The three relative paleointensity proxies yield significant correlation coefficients and the NRM/ARM record was used to further test the validity of the relative paleointensity record as ARM data reflect the behavior of PSD to SD magnetite grains, the population most likely to carry a primary magnetic signal, and the NRM and ARM records have the strongest correlation. The correlation coefficient between the NRM/ARM record and Ms record, a concentration proxy, are not significant ! ""*! 2 (r = 0.10, r = 0.01, t = 0.94, critical = 0.18) suggesting the concentration of iron bearing phases does not affect the relative paleointensity curve. The NRM/ARM relative paleointensity curve does not correlate with the climate change proxy MAT (r = -0.05, r2 = 0.25) (Figure 4.17). The relative paleointensity records obtained with all three normalization parameters appear to be very similar to the Sint-800 paleointensity model (Guyodo and Valet, 1999), although slight variations do occur (Figure 4.18). Three known geomagnetic excursions occur at low paleointensity intervals in the Sint-800 record, near ~536 ka (14!, Calabrian Ridge II), ~400 ka (11!) and ~360-370 ka (Levantine). A fourth potential field phenomena at ~490 (13#) has only been observed in the western Atlantic deep-sea sediments. The sediment in the core near ~490 ka shows evidence for a hiatus and the magnetic phases have experienced significant reduction and subsequent oxidation of sediment. If a geomagnetic phenomena did occur ~490 ka it is not recorded in the VC-3 sediment. The observed negative inclinations magnetization of core sediment near 52, 17, and 5 m depth, combined with decreases in the relative paleointensity data over these interval, are consistent with sedimentation over the time period of geomagnetic excursion, 14!/Calabrian Ridge II, 11!, and Levantine. Discussion Lake Environments The magnetic character of the South Mountain lake sediments is highly variable throughout the core due to a mixture of magnetic phases and grain sizes. The South Mountain lake was productive and organic matter accumulated at the bottom of the lake. Oxygen can become limited at the sediment water interface due to decomposition of ! ""+! organic matter leading to anoxic conditions. In anoxic environments anaerobic biological mediation facilitates the decomposition of organic matter (Berner, 1964; Karlin and Levi, 1983, 1985; Karlin et al., 1987; Anderson and Rippey, 1988; Lovley, 1991; Snowball, 1993; Williamson et al., 1998; Stevens et al., 2000; Sifeddine et al., 2001; Frederichs et al., 2003; Demory et al., 2005; Rowan and Roberts, 2006; Rowan et al., 2009; Ao et al., 2010). Fe (III) is a major electron acceptor during organic matter decomposition (Lovley, 1991) and is present in detrital oxidized catchment sediment. Primary and diagenetic iron-bearing phases are present throughout VC-3 sediment. Primary detrital phases consist of magnetite and titanomagnetite derived from the mostly andesitic Valles Caldera wall and the mostly rhyolitic resurgent domes. Moderate values in proxy records for magnetic abundance are interpreted as most reflective of primary detrital phases and are most common in sediments deposited during the later part of MIS 14, the early part of MIS 13, the later part of MIS 12 and occasionally in sediment deposited during MIS 11 and 10 (Figure 4.3). The VC-3 grain size and shape data set displayed in Figures 4.6 – 4.10 lend further support that primary detrital phases are present throughout the core. The Curie temperature (580° C) and Verwey (120° K) transition are both observed for magnetite in high and low temperature rock magnetic experiments (Figures 4.11 and 4.13). Finally, magnetite and titanomagnetite are identified through visual identification of SEM photomicrographs (Figure 4.14). The observation that primary detrital phases are present throughout the core is important, as these are the phases most likely to carry a stable record of geomagnetic field activity at the time of sediment deposition. Detrital oxidized phases include maghemite, titanomaghemite, and oxyhydroxides (i.e. goethite) that are sourced ! "$#! from oxidized catchment material. Detrital oxidized phases occur in VC-3 intervals characterized by increased SIRM/! and decreased S-ratio values (Figures 4.4 and 4.5). The presence of oxidized phases could be partially responsible for the increased Bcr/Bc values with respect to single domain-multidomain mixing theory on a Day Plot (Figures 4.6b – 4.10b). Evidence for goethite is observed as an increase in susceptibility at 250° C in the high temperature susceptibility plot (Figure 4.11). Oxidized detrital phases are also observed in SEM photomicrographs of magnetic separates (Figure 4.14). Detrital oxidized phases are important because they provide an electron acceptor, Fe (III), during decomposition of organic matter. Oxidized iron becomes reduced in anoxic settings and will precipitated as pyrrhotite under sulfidic conditions, siderite or vivianite under methanogenic conditions, or superparamagnetic magnetite in the presence of GS15 bacterial (Table 4.1). The preservation of detrital oxide phases indicates a well-mixed lake environment and is most commonly observed in sediments deposited during MIS 14, the early part of MIS 12, and the interval representing later part of MIS 11 through the middle of MIS 10. Diagenetic phases form within the South Mountain lake in both reducing and oxidizing conditions. Reduced diagenetic phases include pyrrhotite, siderite, vivianite, and superparamagnetic magnetite, which form in anoxic environments. Diagenetic reduced phases are identified by decreased values in SIRM/! and ARM/IRM and increased values in !/Ms and Bcr/Bc (Figures 4.4 and 4.5). Some contribution from superparamagnetic magnetite is observed in the squareness –coercive force plots and Day Plot (Figures 4.6 a-b – 4.10 a-b). Frequency dependant susceptibility is observed in sediment from magnetic separates obtained from 0.10 to 0.17 m and is indicative of superparamagnetic magnetite (Figure 4.12). Other reduced phases, ! "$"! siderite/rhodochrosite, vivianite, and pyrrhotite are identified by low temperature remanence (Figure 4.13) and FEGSEM photomicrographs. Lake levels can influence the lake environment and by extension the types of iron-bearing phases preserved in the sediment (Figure 4.1). Glacial conditions are mostly characterized in the lake by times of cooler temperatures and deeper lake levels. Oxygen can become limited at the sediment water interface leading to reducing conditions indicated by the presence of diagenetic phases (decreased SIRM/! values and increased !/Ms values, Figure 4.4). On the other hand, warmer temperatures mostly characterize interglacial conditions and lower lake levels allow the water column to be well mixed. Interglacial intervals are identified as having a greater proportion of detrital magnetic phases. Glacial intervals typically preserve a greater abundance of diagenetic phases, however reducing conditions are also present during interglacial times when the lake becomes stagnant and decomposition of total organic carbon leads to anoxic conditions at the sediment water interface. Reducing conditions during interglacial intervals are characterized by decreased SIRM/! values combined with increased !/Ms values (Figure 4.4). Lake conditions during glacial intervals are occasionally well mixed as noted by the presence of detrital phases throughout the core. Sediments deposited during the upper part of MIS 13 are unique as they are associated with the highest susceptibility values of the core (Figure 4.3). Rock magnetic data indicate uniaxial single-domain magnetite is the dominant iron-bearing phase in this interval (Figure 4.9). The percent total organic carbon preserved in the upper part of MIS 13 is low compared to the percent total organic carbon preserved in sediment deposited during MIS 11. The reduction in total organic carbon was interpreted by Fawcett et al., (2011) to indicate the upper part of MIS 13 has ! "$$! been oxidize. The original sediment in the upper part of MIS 13 most likely preserved reduced iron-bearing phases that were subsequently oxidized. Diagenetic reduced phases will oxidize to magnetite if exposed to oxygen rich conditions. The lake environment during sediment deposition during the upper part of MIS 13 is interpreted as a time of lake level driven oxidation of reduced sediments (Figure 4.1). Discrete intervals in sediments deposited during MIS 12 contain extracellular superparamagnetic magnetite produced by GS15 bacteria. The intervals that contain superparamagnetic magnetite occur slightly after intervals that are characterized by interstadial environments. Interstadial environments are brief warm periods during a glacial stage that can enhance surface oxidation of catchment sediments and increase the amount of detrital oxidized phases (i.e. maghemite titanomaghemite, and oxyhydroxides) supplied to the lake. Interstadial environments can also cause a slight decrease in lake level, which is observed in the magnetic records as an interruption of anoxic condition and preservation of detrital phases. Sediments deposited during MIS 12 indicate reducing conditions, preservation of superparamagnetic magnetite (increased !/Ms values), are interrupted by interstadial events (increased mean annual temperature combined with increased Cyperacea and Juniper pollen counts, Figure 4.19). Sedimentary intervals deposited during the warm substages of MIS 11 (11a, 11c, and 11e of Fawcett et al., 2011) have decreased values in proxy records for magnetic abundance. Rock magnetic data indicate primary detrital phases are the most common iron-bearing phases preserved in strata deposited during the warm substages of MIS 11 (decreased !/Ms values Figure 4.4, and increased S-ratio values, Figure 4.5). Decreased values in magnetic abundance are most likely due to dilution from biogenic silica. ! "$%! Geomagnetic Field Activity The fidelity of the magnetization record preserved in the VC-3 sediment sequence is a reflection of depositional environment, which may affect the timing of remanence acquisition and thus the carriers of any geologically significant remanence. Five intervals of negative inclination, of variable quality and internal consistency are preserved at about 67, 52, 22, 17 and 5 m depths. Three of these intervals, 52, 17, and 5 m occur display rock magnetic characteristics typical of well-mixed lake environments with no evidence of diagenetic phases. Two geomagnetic excursions, 14! at 536 ka and 11! at 400 ka, have been identified in records from Calabrian Ridge in the western Atlantic Ocean, sediments from La Palma, Canary Islands, and the West Eifel volcanic field in Germany (Schnepp and Hradetzky, 1994; Langereis et al., 1997; Singer et al., 2002; Lund et al., 2006; Singer et al., 2008). The sediment intervals at about 52 m and 17 m are the most likely to correlate with 14! and 11! based on important age constraints for core VC-3, including the Ar-Ar date from basal tephra and the inferred locations of glacial terminations V and IV, as discussed by Fawcett et al. (2011). The duration of geomagnetic instability 14! is between 100 – 125 years and 11! are about 65 to 74 years in duration based on the chronology of the core. One other geomagnetic instability has been identified to occur over the time South Mountain lake sedimentation (552 – 363 ka) in sediments from ODP leg 1062 in the western Atlantic Ocean 13! (Lund et al., 2006; Singer et al., 2008). There are two possible reasons why this instability is not recorded in the VC-3 sediments. Instability 13! has only been identified in one study and may represent a regional instability in the magnetic field rather than a geomagnetic field ! "$&! phenomenon. The age of 13! is approximately 490 ka, which corresponds to the sedimentary interval deposited in the later part of MIS 13. Sediment deposited in the later part of MIS 13 is interpreted to have about 5 to 6 m of missing sediment. The magnetic phases preserved in the upper part of MIS 13 are interpreted as having been formed in anoxic bottom water that were later exposed to oxic conditions resulting in diagenetic magnetite. A single sediment sample from about 52 m depth shows high quality demagnetization trajectory with negative inclination but adjacent samples reveal the isolation of normal polarity magnetizations. The magnetization in these sediments is most likely carried by primary phases deposited in a sub-oxic environment and most likely records a true field phenomenon. The demagnetization behavior of sediments from the depth interval of about 17 m is of moderate to poor quality. Only three samples show negative inclination behavior and the other nine show normal polarity interpreted as a post depositional overprint carried by diagenetic phases. The three relative paleointensity records (NRM/ARM, NRM/IRM, and NRM/!bulk) for this stratigraphic interval are similar at a statistically significant level, suggesting an overall correlation between the magnetization recorded in the sediments and the strength of the geomagnetic field (Tauxe, 1993; Valet, 2003). The relative paleointensity record does not correlate with the MAT record, a proxy for climate change further supporting viability of the paleointensity record (Brachfeld and Banerjee, 2000; Haltia-Hovi et al., 2011). Decreases in the relative paleointensity are consistent with decreased values in composite paleointensity record Sint-800 at 536 and 400 ka, consistent with geomagnetic excursions 14! and 11!. ! "$'! Conclusions South Mountain lake sediments retrieved in the VC-3 coring experiment and dated between about 552 ka to 368 ka were deposited during the mid-Pleistocene in glacial and interglacial environments (Fawcett et al., 2007; Reneau et al., 2007; Goff, 2009; Fawcett et al., 2011; Goff et al., 2011). Magnetic phases of the VC-3 sediment core include primary detrital phases, oxidized detrital phases and diagenetic phases. Sediments deposited during glacial intervals preserve a greater proportion of diagenetic phases while detrital phases are more common during interglacial intervals. Primary detrital phases consist of magnetite and titanomagnetite and are sourced from the andesitic Valles Caldera wall and rhyolitic resurgent domes. Primary phases are present throughout the core. Oxidized detrital phases include maghemite, titanomaghemite, and oxyhydroxides that form in the lake catchment. Diagenetic phases include superparamagnetic magnetite, pyrrhotite, siderite/rhodochrosite, and vivianite and form in anoxic environments at the expense of oxidized detrital phases. Reducing environments result from the decomposition of organic matter and can form during glacial intervals when the lake is deep and stratified or during interglacial intervals with the lake becomes stagnant. Reducing conditions are commonly observed slightly after interstadial events of MIS 12. Decreased lake levels during the upper part of MIS 13 led to the preservation of diagenetic magnetite that formed in response to the oxidation of reduced sediments. Sediments deposited during the warm substrages of MIS 11 are associated with decreased magnetic abundance and preservation of primary detrital phases. Iron-bearing phases preserved during the warm substages of MIS 11 are diluted by biogenic silica. ! "$(! The VC-3 sediments preserve five intervals of anomalous remanence characteristics (i.e. negative inclination magnetizations isolated in progressive demagnetization). The relative paleointensity records for the VC-3 shows some correlation to the inferred strength of the geomagnetic field particularly over the time interval of 536 and 400 ka. Three geomagnetic phenomena (14!/Calabrian Ridge II, 11!, and Levantine) occurred during deposition of sediments in the South Mountain lake, at ~536, ~400, and ~360-370 ka, which correspond to depths of about 52, 17, and 5 m in the VC-3 core. Sediment from 52 m displays a high quality demagnetization trajectory revealing the isolation of a negative inclination magnetization. This interval reveals a period of decreased relative paleointensity observed in NRM/ARM, NRM/SIRM, and NRM/!bulk and decreased values in paleointensity observed from the Sint-800 record Rock magnetic results suggest 52 m represents a sub-oxic environment suggesting the remanence may be carried by detrital phases and thus may be a high fidelity recorder of the field. Sediments from 17 m, on the other hand, display moderate to poor quality demagnetization trajectories that are inconsistent between samples. However, a slight decrease at 17 m is observed in the relative paleointensity records combined with negative inclinations indicate that the decrease in the Sint-800 paleointensity record is at least partially recorded in the sediments. Rock magnetic results indicate 17 m represents a sub-oxic to anoxic environment and a mixture of detrital and diagenetic phases carries remanence. The VC-3 sediments provide an important record redox conditions in the lake and partially record mid-Pleistocene geomagnetic field phenomenon. ! "$)! Acknowledgments We are grateful to Bob Parmenter for access to the Valles Caldera. The coring project would not have been possible without invaluable assistance from Fraser Goff, Craig Allen, Tim Wawrzyniec, and Amy Ellwein. This research was funded through the United States Geological Survey, the National Science Foundation ATM-0434459, the Institute of Geophysics and Planetary Physics at Los Alamos National Laboratory, Sloan, the LacCore facility at the University of Minnesota, Twin Cities, the Institute for Rock Magnetism at the University of Minnesota, Twin Cities, and the Department of Earth and Planetary Sciences at the University of New Mexico. ! "$*! Figures Sediment input from: dry and warm climate caldera wall, resurgent domes, moderate lake levels and catchment preservation of primary phases e.g. magnetite, titanomagnetite, maghemite, titanomaghemite and catchment oxidized sediments e.g. oxyhydroxides wet and cool climate deep lake, strati!ed water reduction of Fe(III) preservation of reduced phases e.g. pyrrhotite, siderite, rhodochrosite, and vivianite dry and warm climate interruption of strati!cation oxidation of reduced phases e.g. magnetitie Legend well mixed anoxic primary reduced oxidized lake water lake water sediment sediment sediment ,-./01!&2"2!3451678-4!9:!;-::101<8!9=-;78-98781!9:!>1;-61<8>!?01>10@1;!- 39/85!A9/<87-9 ! "$+! Figure 4.2. Star indicates the location of the dill site, inset map of New Mexico, USA, square indicates the position of the Valles Caldera, modified from (Gardner et al., 1996). ! ! "%#! MIS 10 MIS 11 MIS 12 MIS 13 MIS 14 ï x 10 2 /kg) 2 1 NRM (Am 0 ï 2 x 10 /kg) 2 1 (Am s M Abundance 0 /kg) 2 3 (m 1 ferri r 0 380 400 420 440 480 time (ka) ! ! ! ,-./01!&2%2!H07?5>!>59I!7.1!@10>/>!?09=-1>!:90!67.<18-4!7J/<;7<41!DKLAF!A>F!7<;! M:100-G2!!N5101!->!7!5-.510!?1041<87.1!9:!67.<18-4!7J/<;7<41!9J>10@1;!- ?708!9:!>87.1!"%!!7<;!>87.1!"$2!! ! ! ! ! ! ! "%"! MIS 10 MIS 11 MIS 12 MIS 13 MIS 14 1 ï x 10 ARM / IRM 0 1 3 x 10 (A/m) 2 r 1 Grain SIze SIRM / 0 ï 1 x 10 (m/A) s / M r 0 360 380 400 420 440 460 480 500 520 540 560 time (ka) ! ,-./01!&2&2!N51!.07?5>!>59I!7.1!@10>/>!?09=-1>!:90!67.<18-4!.07--P1!DQLAR3SLAF! 3SLARMF!7<;!MRA>G!@70-7J-O-8T2!!!E787!>/..1>8>!7!.017810!7J/<;7<41!9:!:-<1!.07-<1;! >9:8!49104-@-8T!?57>1>!D12.2!;180-87O!?57>1>G!;/0-<.!-<810.O74-7O!>87.1>!DQLAR3SLAGF!7! .017810!7J/<;7<41!9:!570;!49104-@-8T!?57>1>!D12.2!;-7.1<18-4!?57>1>G!;/0-<.!851! .O74-7O!>87.1>!D3SLARMGF!7<;!7!.017810!7J/<;7<41!9:!>/?10?70767.<18-4!67.<18-81! ;/0-<.!>87.1!"%F!"$F!7<;!""!DMRA>G2! ! "%$! MIS 10 MIS 11 MIS 12 MIS 13 MIS 14 ï 5 x 10 Mrs/Ms 0 0 6 x 10 4 Bcr/Bc Composition 2 0 1 x 10 S-ratio 0 360 380 400 420 440 460 480 500 520 540 560 time (ka) ! ,-./01!&2'2!!N51!.07?5>!>59I!7.1!@10>/>!?09=-1>!:90!496?9>-8-91>! D0167<1<41!078-9!DA0>RA>GF!49104-@-8T!078-9!DU40RU4G!7<;!3V078-9G2!!!S<4017>1;!@7O/1>! -RA>F!69;10781!U40RU4F!7<;!-<4017>1;!3V078-9!@7O/1>!>/..1>8!7!;180-87O!?57>1>! 701!6901!4966987.1>2! ! ! ! ! "%%! a) Squareness-Coercive Force Plot b) Day Plot 0.5 0.5 USD SD 0.4 0.4 SP USD + SP 0.3 0.3 6'ï0' 0.2 C + SP 0,6 366 ka 382 ka 0.2 Remanence Ratio (Mrs/Ms) MD MD 20 40 60 2 4 6 Coercive Force (Bc) Coercivity Ratio Bcr/Bc c) King Plot d) IRM acquisition and back!eld curves ï [ 2 0,6 0.8 368 ka 382 ka 0.6 M R 0.4 A r 0.2 368 ka 382 ka 0 0 ï normalized magnetization 0 5 [ ï 0 2 3 r induction (T) SD = single domain, MD = multi-domain, USD = uniaxial single domain USD + SP = uniaxial single domain plus superparamagnetic, C + SP = cubic plus superparamagnetic ! ! ,-./01!&2(2!L94C!67.<18-4!01>/O8>!:90!69>8OT!4/J-4!67.<18-4!?57>1>!D-212!67.<18-81G! :096!>1;-61<8>!;1?9>-81;!;/0-<.!AS3!"#F!7G!3W/701<1>>V49104-@1!:9041!?O98!DN7/=1! 18!7O2F!$##$GF!E7T!?O98!DE7T!18!7O2F!"+))X!Y70810V38-.O-8P!18!7O2F!$##"GF!4G!Z-<.!?O98!DZ-<.! 18!7O2F!"+*$GF!;G!SLA!74W/->-8-92!! ! "%&! a) Squareness-Coercive Force Plot b) Day Plot 0.5 0.5 USD SD 0.4 0.4 SP USD + SP 0.3 6'ï0' 0.3 0.2 C + SP 0,6 407 ka ND 0.2 Remanence Ratio (Mrs/Ms) MD MD 20 40 60 2 4 6 Coercive Force (Bc) Coercivity Ratio Bcr/Bc c) King Plot d) IRM acquisition and back!eld curves ï [ 0.8 2 0,6 407 ka 0.6 M ND R 0.4 A r 0.2 407 ka ND 0 0 normalized magnetization ï 0 5 [ ï 0 2 3 r induction (T) SD = single domain, MD = multi-domain, USD = uniaxial single domain USD + SP = uniaxial single domain plus superparamagnetic, C + SP = cubic plus superparamagnetic ! ! ,-./01!&2)2!L94C!67.<18-4!01>/O8>!:90!69>8OT!4/J-4!67.<18-4!?57>1>!D-212!67.<18-81G! :096!>1;-61<8>!;1?9>-81;!;/0-<.!AS3!""F!7G!3W/701<1>>V49104-@1!:9041!?O98!DN7/=1! 18!7O2F!$##$GF!E7T!?O98!DE7T!18!7O2F!"+))X!Y70810V38-.O-8P!18!7O2F!$##"GF!4G!Z-<.!?O98!DZ-<.! 18!7O2F!"+*$GF!;G!SLA!74W/->-8-92! ! "%'! ! a) Squareness-Coercive Force Plot b) Day Plot 0.5 0.5 USD SD 0.4 0.4 SP USD + SP 0.3 0.3 6'ï0' C + SP 0.2 0,6 450 ka 475 ka 0.2 Remanence Ratio (Mrs/Ms) MD MD 20 40 60 2 4 6 Coercive Force (Bc) Coercivity Ratio Bcr/Bc c) King Plot d) IRM acquisition and back!eld curves ï [ 2 0,6 0.8 450 ka 475 ka 0.6 M R 0.4 A r 0.2 450 ka 475 ka 0 0 ï normalized magnetization 0 5 [ ï 0 2 3 r induction (T) SD = single domain, MD = multi-domain, USD = uniaxial single domain USD + SP = uniaxial single domain plus superparamagnetic, C + SP = cubic plus superparamagnetic ! ! ,-./01!&2*2!L94C!67.<18-4!01>/O8>!:90!69>8OT!4/J-4!67.<18-4!?57>1>!D-212!67.<18-81G! :096!>1;-61<8>!;1?9>-81;!;/0-<.!AS3!"$F!7G!3W/701<1>>V49104-@1!:9041!?O98!DN7/=1! 18!7O2F!$##$GF!E7T!?O98!DE7T!18!7O2F!"+))X!Y70810V38-.O-8P!18!7O2F!$##"GF!4G!Z-<.!?O98!DZ-<.! 18!7O2F!"+*$GF!;G!SLA!74W/->-8-92! ! ! "%(! a) Squareness-Coercive Force Plot b) Day Plot 0.5 0.5 USD SD 0.4 0.4 SP USD + SP 0.3 C + SP 0.3 6'ï0' 0.2 0,6 ND 0.2 ND Remanence Ratio (Mrs/Ms) MD MD 20 40 60 2 4 6 Coercive Force (Bc) Coercivity Ratio Bcr/Bc c) King Plot d) IRM acquisition and back!eld curves ï [ 0,6 2 upper part 0.8 ND 0.6 ND M R 0.4 A r 0.2 484 ka 528 ka 0 0 normalized magnetization ï 0 5 [ ï 0 2 3 r induction (T) SD = single domain, MD = multi-domain, USD = uniaxial single domain USD + SP = uniaxial single domain plus superparamagnetic, C + SP = cubic plus superparamagnetic ! ! ,-./01!&2+2!L94C!67.<18-4!01>/O8>!:90!69>8OT!4/J-4!67.<18-4!?57>1>!D-212!67.<18-81G! :096!>1;-61<8>!;1?9>-81;!;/0-<.!AS3!"%F!7G!3W/701<1>>V49104-@1!:9041!?O98!DN7/=1! 18!7O2F!$##$GF!E7T!?O98!DE7T!18!7O2F!"+))X!Y70810V38-.O-8P!18!7O2F!$##"GF!4G!Z-<.!?O98!DZ-<.! 18!7O2F!"+*$GF!;G!SLA!74W/->-8-92! ! "%)! a) Squareness-Coercive Force Plot b) Day Plot 0.5 0.5 USD SD 0.4 0.4 SP USD + SP 0.3 0.3 C + SP 0.2 6'ï0' 0,6 0.2 552 ka Remanence Ratio (Mrs/Ms) 554 ka MD MD 20 40 60 2 4 6 Coercive Force (Bc) Coercivity Ratio Bcr/Bc c) King Plot d) IRM acquisition and back!eld curves ï [ 2 0,6 0.8 552 ka 554 ka 0.6 M R 0.4 A r 0.2 552 ka 554 ka 0 0 ï normalized magnetization 0 5 [ ï 0 2 3 r induction (T) SD = single domain, MD = multi-domain, USD = uniaxial single domain USD + SP = uniaxial single domain plus superparamagnetic, C + SP = cubic plus superparamagnetic ! ! ,-./01!&2"#2!L94C!67.<18-4!01>/O8>!:90!69>8OT!4/J-4!67.<18-4!?57>1>!D-212!67.<18-81G! :096!>1;-61<8>!;1?9>-81;!;/0-<.!AS3!"&F!7G!3W/701<1>>V49104-@1!:9041!?O98!DN7/=1! 18!7O2F!$##$GF!E7T!?O98!DE7T!18!7O2F!"+))X!Y70810V38-.O-8P!18!7O2F!$##"GF!4G!Z-<.!?O98!DZ-<.! 18!7O2F!"+*$GF!;G!SLA!74W/->-8-92! ! ! "%*! ï x 10 31.4 to 39.3 m 2 warming 1.5 cooling 1 580o C 0.5 susceptibility (m3/kg) 0 0 100 200 300 400 500 600 700 Temperature (C) ï x 10 7.1 to 13.8 m 1.5 warming cooling 1 o 0.5 580 C susceptibility (m3/kg) 0 0 100 200 300 400 500 600 700 Temperature (C) ! ! Figure 4.11. Plots of "bulk versus temperature, between room temperature and 700° C measured in a controlled argon atmosphere, for magnetic separates taken from 31.4 m – 39.3 m, and separates from 7.1 m to 18.3 m. Curves illustrate diagnostic ordering temperatures of 250° C, possible goethite to maghemite transition and/or unblocking of titanomagnetite, 380° C, possible maghemite to hematite transition, 420° C, possible oxidation of siderite, 530° C, possible unblocking of titanohematite, and potential Student Version of MATLAB Hopkins Peak just before unblocking of magnetite phases near 580° C. ! ! "%+! ï x 10 ïP 2.5 2 1.5 fd (m3/kg) 1 40 Hz r 400 4000 0.5 0 100 200 300 ï x 10 1.5 1 0.5 0 40 Hz ï 400 fd prime (m3/kg) r 4000 ï 0 100 200 300 Temp (K) ! Figure 4.12. Plots of in phase (plots on the left) and quadrature (plots on the right) components of "fd as a function of decreased (up to 10° K) and increased (up to 300° K) temperature from magnetic separates from 10 m to 17 m. Curves illustrate diagnostic ordering temperatures 40° K, possibly due to the unblocking of ferrihydrite, 50° K, possibly due to the unblocking of ilmenite, and presence of superparamagnetic magnetite. ! ! ! ! ! ! "&#! ï 9 x 10 2 Viviantie 35.01 m 8 0 ï 7 ï 6 ï 5 ï 4 ï ï 3 ï 2 ïï 1 x 10 50 100 150 200 250 300 0.040 0.035 /kg) 2 0.030 0.025 Siderite 36.36 m 0.020 0.015 0.010 0.005 0 Magnetization (Am 50 100 150 200 250 300 0.018 0.016 0.014 0.012 0.01 Rhodocrocite 35.01 m 0.008 0.006 0.004 0.002 0 50 100 150 200 250 300 Temperature (K) dFC/dT dZFC/dT Field Cooled Remanence (FC) Zero Field Cooled Remanence (ZFC) ! Figure 4.13. Plots of field cooled (blue) and zero field cooled (pink) IRM and their derivatives (dashed lines) as a function of low temperature, between 10° K and 300° K, for samples from a) 35.01 m, b) 37.11 m, and c) 36.36 m. Curves show diagnostic ordering temperatures. ! "&"! !"# 6816 cm !$# 3051 cm !%# 1176 cm MGT MGH MGT SUL MGT MGH 100+m 100+m HEM 100+m !&# 3051 cm !'# 1810 cm !(# 1810 cm MGT MGT MGH 50+m 20+m 50+m !)# 6816 cm !*# 3051 cm !+# 1810 cm TMG TMG TMH 20+m 25+m 50+m MGT TMG MGH TMH SUL HEM = magnetite = titanomagnetite = maghemite = titanomaghemite = sul!des = hematite ! Figure 4.14. SEM photomicrographs of bulk and select grains from magnetic extracts obtained from 68.16 m, 30.51 m, 18.10 m, and 11.76 m. Images are organized into bulk magnetic separates (a-c), magnetite and maghemite grains (b-f), and titanomagnetite grains (g-i). ! ! "&$! !" #" $" pyrite or mackinawite siderite rhodochrosite 1 +m 1 +m 1 +m ! ! Figure 4.15. FEGSEM photomicrographs obtained from magnetic separates at 4350 cm of (a) siderite, (b) pyrite or mackinawite, and (c) rhodochrosite. ! "&%! Levantine 11_ 14_/Calabrian Ridge II 5.60 m 17.20 m N, Up 52.34 m N, Up 10-6 -5 10 10 mT 40 NRM 10 mT 10-4 E 40 35 80 10 mT E E NRM typical N polarity 5.62 m N, Up 17.22 m NRM N, Up 68.60 m N, Up 10-5 14 mT 100 10-5 E 25 60 10 mT E 60 100 E NRM 20 mT 10-5 NRM 50 0 Inc MAD all ï o MAD < 20 360 380 400 420 440 460 480 500 520 540 560 !#$)+#,)%&' !())%&' !"#$%&' !"*)%&' -./0'1%2.1'/345 67.0'6-7438%9371:%2;7-.07-' 67.0'6-7438%9371:%7<8=/>3.- ,-./01!&2"(2![0859.9<7O!@14890!?O98>!>59I-<.!8T?-47O!;167.<18-P78-9 >76?O1>!:096!-<810@7O>!8578!01490;!<9067O!?9O70-8T!.1967.<18-4!:-1O;!7<;!!"#$%&'%(')*+,'-..- .1967.<18-4!:-1O;!?51<961<72!!N51!6-;;O1!?O98!>59I!851!7.1!@10>/>!-<4O-<78-9 7OO!>76?O1>!DJO/1G!7<;!>76?O1>!I-85!617>!857 DJO74CG2!B7>8!?7<1O!->!851!.1967.<18-4!-<>87J-O-8T!8-61!>47O1!69;-:-1;!:096!3-<.10!18! 7O2F!D$##$G!I-85!7.1>!:096!>1@107O!>9/041>!DB7<.101->!18!7O2F!"++)X!3-<.10!18!7O2F!$##$X! B/<;!18!7O2F!$##(X!3-<.10!18!7O2F!$##*G2!! ! "&&! 1 1 r = 0.71 0.70 0.30 r = ï 0.5 0 MAT 0 ï 0 0.5 1 0 0.5 1 NRM NRM 1 NRM/ARM NRM/IRM 0.5 NRM/chi 0 normalized parameters 1 0 NRM/ARM ï MAT 350 400 450 500 550 time (ka) ! ,-./01!&2")2!!N51!.07?5!959I>!851!49001O78-9 J18I11>!<9067O-P1;!>/>41?8-J-O-8TG2!! N51!.07?5!959I>!851!49001O78-9 KLARQLA!01O78-@1!?7O19-<81<>-8T!01490;!7<;!851!01490;!9:!617 816?1078/01!DAQNG2!!N51!6-;;O1!.07?5!>59I>!851!7.1!@10>/>!01O78-@1!?7O19-<81<>-8T! :90!851!85011!01490;>F!KLARQLAF!KLARSLAF!KLARM!DN7/=1F!"++%G2!!N51!J98896! .07?5!>59I>!7.1!@10>/>!851!01O78-@1!?7O19-<81<>-8T!01490;!KLARQLA!7<;!AQN2!! ! "&'! E7>51;!O-<1>!01?01>1<8!?990!;787!;1<>-8T2! 50 0 Inc MAD all ï o MAD < 20 360 380 400 420 440 460 480 500 520 540 560 0.1 o r 0.05 N / 0 0.2 0.1 N / ARM 0 0.01 0.005 Relative Paleointensity N / IRM 0 10 ) 2 5 Am 2 0 2 13_ 11_ Calabrian Ridge 2 Levantine West Eifel 5 14_ Big Lost VAMD (10 360 380 400 420 440 460 480 500 520 540 560 ! ! Student Version of MATLAB ,-./01!&2"*2!N51!/??10!.07?5!>59I!851!7.1!@10>/>!-<4O-<78-976?O1>!DJO/1G! 7<;!>76?O1>!I-85!AQE!O1>>!857!>59I!851!7.1!@10>/>! 01O78-@1!?7O19-<81<>-8T!01490;>!KLARQLAF!KLARSLAF!7<;!KLARM2!!N51!J98896! .07?5!>59I>!7.1!@10>/>!L1O78-@1!?7O19-<81<>-8T!496?701;!89!851!@-08/7O!7=-7O!;-?9O1! 6961<8!D]QAEG!69;1O!>-<8V*##!DH/T9;9!7<;!]7O18F!"+++G!I-85!.1967.<18-4!:-1O;! ?51<961<98/;-1>!DB7<.101->!18!7O2F!"++)X!3-<.10!18!7O2F!$##$X! B/<;!18!7O2F!$##(G2! ! "&(! MIS 10 MIS 11 MIS 12 MIS 13 MIS 14 0.1 s chi / M 0 20 Juniper 0 15 10 5 Cyperaceae 0 5 0 MAT ï 360 380 400 420 440 460 480 500 520 540 560 ! ! ,-./01!&2"+2!!S<4017>1;!@7O/1>!-!701!-<810?0181;!7>! -<810@7O>!I-85!>/?10?70767.<18-4!67.<18-812!!N51!-<810@7O>!I-85!>/?10?70767.<18-4! 67.<18-81!944/0!78!851!>761!8-61!90!>O-.58OT!7:810!I706!-<810>87;-7O!1@1<8>!- "$!7<;!9447>-9<7OOT!;/0-<.!I706!?10-9;>!-!701! Student Version of MATLAB -;1<8-:-1;!7>!-<4017>1;!YT?1074171!7<;!_/<-?10!49/<8>!7<;!-<4017>1;!617 816?1078/01!DAQNG2!!`9OO1 ! ! ! ! ! "&)! Tables Table 4.1. Detrital and diagenetic phases present in sediment from the South Mountain lake. Mineral Environment of Formation Source Magnetite Valles Caldera rhyolitic resurgent domes Fe 2+ and andesitic wall of the caldera Fe3O4 Titanomagnetite Valles Caldera rhyolitic resurgent domes Fe 2+ and andesitic wall of the caldera Fe3-xTixO4 (0 < x < 1) Maghemite Oxidized catchment sediment Fe 3+ g-Fe2O3 Titanomaghemite Oxidized catchment sediment Fe 3+ g-Fe2-yTiyO3 (0 < y < 1) oxyhydroxide Oxidized catchment sediment (i.e. goethite) provides poorly crystaline Fe(iii) necessary Fe 3+ for production of diagenetic phases a-FeOOH Hematite catchment Fe 3+ precipitate in the sediment water interface a-Fe2O3 Pyrrhotite high influx of organic matter Brener 1964 Fe 2+ anxoic environment Lovley 1991 Fe7S8 limited sulfate reduction Snowball1991 low concentration of poorly crystaline Fe(iii) Jeonowski 1997 - limitied oxidized soil snowball and Torii1999 Rhodocrocite Orgainc rich sediments Brener 1964 Mn 2+ anoxic sulfidic and non-sulfidic environment Lovley 1991 ! "&*! MnCO3 and methanogenic environments Snowball 1991 low concentration of poorly crystaline Fe(iii) Jeonowski 1997 - limitied oxidized soil Frederichs et al., 2003 possible increase in hydrogen or acetate Kosterov et al., 2006 Siderite Orgainc rich sediments Brener 1964 Anderson and Rippey Fe 2+ anoxic sulfidic and non-sulfidic environment 1988 FeCO3 and methanogenic environments Lovley 1991 low concentration of poorly crystaline Fe(iii) Snowball 1991 - limitied oxidized soil Jeonowski 1997 possible increase in hydrogen or acetate Frederichs et al., 2003 Anderson and Rippey Vivianite Organic rich sediments 1988 Fe 2+ anoxic and methanogenic environment Lovley, 1991 [Fe3(PO4)2] anoxic sulfidic and non-sulfidic environment Frederichs et al., 2003 * 8 H2O and methanogenic environments low concentration of poorly crystaline Fe(iii) - limitied oxidized soil ! "&+! Table 4.2. Average values for magnetic proxies. Rock Magnetic parameter Mean Standard Deviation Natrual remanent magnetization (NRM) 2.73 x 10-6 Am2/kg 3.85 x 10-6 Am2/kg Anhysteretic remanent magnetization (ARM) 2.09 x 10-5 Am2/kg 2.02 x 10-5 Am2/kg Isothermal remanent magnetization (SIRM) 5.64 x 10-4 Am2/kg 9.69 x 10-4 Am2/kg -3 2 -3 2 Spontaneous magnetization (Ms) 3.57 x 10 Am /kg 3.63 x 10 Am /kg -5 3 -5 2 Ferrimagnetic susceptibility ("ferri) 9.37 x 10 m /kg 7.44 x 10 Am /kg ARM/SIRM 2.81 x 10-2 1.34 x 10-2 Am2/kg SIRM/"bulk 10.78 A/m 8.75 A/m !bulk/Ms 0.14 m/A 1.19 m/A Remanence ratio (Mrs/Ms) 0.28 8.21 x 10-2 Coercivity ratio (Bcr/Bc) 3.3 9.73 x 10-1 ! "'#! Table 4.3. Results of low temperature remanence experiments sample Temperature Field cooled Zero Field interpretation reanence cooled remanence m K Am2 kg-1 Am2 kg-1 35.01 10 8.50 x 10-3 7.50 x 10-3 19 6.40 x 10-3 6.50 x 10-3 viviantie 120 2.60 x 10-3 2.60 x 10-3 magnetite 36.35 10 40.00 x 10-3 5.00 x 10-3 40 3.00 x 10-3 3.00 x 10-3 siderite 37.11 10 16.00 x 10-3 7.50 x 10-3 35 4.20 x 10-3 4.60 x 10-3 rhodochrosite 120 1.00 x 10-3 1.00 x 10-3 magnetite ! "'"! Reference Anderson, N. J., and Rippey, B., 1988, Diagenesis of magnetic minerals in the recent sediments of a eutrophic lake: Limnology and Oceanography, v. 33, no. 6, p. 1476-1492. Ao, H., Deng, C. L., Dekkers, M. J., and Liu, Q. S., 2010, Magnetic mineral dissolution in Pleistocene fluvio-lacustrine sediments, Nihewan Basin (North China): Earth and Planetary Science Letters, v. 292, no. 1-2, p. 191-200. Berner, R. 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Williamson, D., Jelinowska, A., Kissel, C., Tucholka, P., Gibert, E., Gasse, F., Massault, M., Taieb, M., Van Campo, E., and Wieckowski, K., 1998, Mineral-magnetic proxies of erosion/oxidation cycles in tropical maar-lake sediments (Lake Tritrivakely, Madagascar): paleoenvironmental implications: Earth and Planetary Science Letters, v. 155, no. 3-4, p. 205-219. ! "($! CHAPTER 5 CONCLUSIONS This dissertation summarizes two investigations of magnetic records from sedimentary rocks and sediments, of very disparate ages deposited in very different depositional environments, of the southwest U.S.A. In chapter 2, the timing and duration of Moenave Formation sedimentation is discussed on the basis of magnetostratigraphic information and the placement of the Triassic-Jurassic boundary in terrestrial strata of the southwest U.S.A. is evaluated. In chapter 3 the paleomagnetic data obtained from the Moenave Formation strata are used to provide a more robust paleomagnetic pole position for the Southwest U.S.A. and reconcile this result with pole positions for latest Triassic to earliest Jurassic aged rocks from elsewhere in North America. Rock magnetic data obtained from mid-Pleistocene lake deposits of the South Mountain Lake, Valles Caldera were used to interpret general relationships between environmental magnetic records and glacial/interglacial intervals. Both paleomagnetic and rock magnetic records are consistent with expected geomagnetic field variability of the mid-Pleistocene including three intervals of geomagnetic field phenomena. A composite magnetostratigraphic record was constructed from four sections of the uppermost Triassic to lowermost Jurassic Moenave Formation strata in southern Utah and northern Arizona. The composite magnetostratigraphic record yields mostly normal polarity directions with the exception of three intervals of reverse polarity (magnetozones). The two oldest magnetozones, M1r and M2r, are interpreted as latest Triassic in age while the magnetozone M3r is interpreted as earliest Jurassic in age. ! "#$! Magnetozone M1r is poorly constrained and is correlated to magnetozones SA5n.1r of Saint Audrie’s Bay, H- of Oyuklu, BIT5n.1r of the Southern Alps, the reverse-polarity magnetozone of the Bigoudine Formation, High Atlas, Morocco, and magnetozone E23r of the Newark Basin. Magnetozone M2r is correlated to either magnetozones SA5n.2r or SA5n.3r of Saint Audrie’s Bay, magnetozone J- of Oyuklu, and either of the intervals of reverse polarity form the “intermediate unit,” Morocco. Magnetozone M3r is the earliest Jurassic magnetozone reported anywhere and is correlated to SA5n.5r of Saint Audrie’s Bay record. The earlist Jruassic interval of reverse polarity is not observed in other records but is projected to lie in the lower part of the CAMP sequence where interbedded sediments preserve the conchostracan Bulbilimnadia killianorum. An implication of the interpreted magnetostratigraphic correlations presented here is that the Triassic Jurassic boundary lies in a ~2 m interval in the lower part of reverse-polarity magnetozone M3r within the middle part of the Whitmore Point Member, implying that early dinosaur evolution occurred in the latest Triassic, not the earliest Jurassic, as previously thought. Inclination error estimates that are derived from the distribution of paleomagnetic data from the four sections of Moenave Formation strata show that these rocks experienced a relatively small magnitude inclination shallowing, in particular in comparison to coeval strata from the northeast US and other red bed sequences. A new pole position 61.3° N, 68.2° E, (A95 = 7.4°, N = 102 sites) is based on a locally derived inclination error estimate (f = 0.78) and a maximum magnitude (4°) of clockwise Colorado Plateau rotation. The pole position further supports the concept or existence of a J1 cusp, or a westerly position for the latest Triassic to earliest Jurassic paleomagnetic pole position for North America. ! "#%! Detailed investigations of environmental magnetic records suggest that the iron- bearing phases preserved in the lake come from three main sources. Detrital primary phases include magnetite and titanomagnetite and are sourced from the mostly andesitic Valles Caldera wall and the mostly rhyolitic resurgent domes. Detrital oxidized phases include maghemite, titanomaghemite, and oxyhydroxides (i.e. goethite) are sourced from the South Mountain lake catchment. Diagenetic reduced phases such as pyrrhotite, siderite/rhodochrosite, vivianite, and superparamagnetic magnetite are precipitated in anoxic water at the sediment water interface due to biologic mediated organic matter decomposition. Finally, diagenetic magnetite can precipitate when anoxic conditions are interrupted and reduced sediment (e.g. diagenetic reduced phases) is oxidized. The paleomagnetic behavior of most of the VC-3 sediment is consistent with the current Brunhes normal polarity chorn. Despite the variety of magnetic phases present in core VC-3 three intervals of negative inclination are preserved at 68 m, 17 m, and 5 m that correlate with geomagnetic field phenomena 14!/Calabrian Ridge II at ~536 ka, 11! at ~400 ka, and Levantine at ~360-370 ka. High correlation coefficient among the following three records, NRM/ARM, NRM/IRM, NRM/", indicate the VC-3 paleomagnetic data are robust enough to construct a relative paleointensity curve. A comparison between the MAT and paleointensity curve NRM/ARM yields a low correlation coefficient indicating that the relative paleointensity curve reflects geomagnetic field variability. The three relative paleointensity records obtained show similar features that can be compared with paleointensity model Sint-800, particularly at about ~536 ka, ~400 ka, and ~360-370 ka coincident with geomagnetic field phenomenon 14!/Calabrian Ridge II, 11! and Levantine. ! "#&! Future Research The research summarized in this dissertation has inspired several new questions, all of which should be addressed by future research. Magnetostratigraphic records obtained from the Moenave Formation in its northern outcrop belt indicate the Triassic-Jurassic boundary occurs within Whitmore Point Member of the Moenave Formation. It remains unclear where to place the Triassic- Jurassic boundary in the southern outcrop belt of the Moenave Formation where Dinosaur Canyon Member strata, of a fluvial affinity, are the only strata preserved. Evaluating the magnetic records of the Moenave Formation in the southern outcrop belt would clarify the age relationship between the upper Whitmore Point strata in the northern outcrop belt and upper Dinosaur Canyon strata in the southern outcrop belt. Is the Dinosaur Canyon Member time-transgressive between southern and northern outcrop belts? If so, the Dinosaur Canyon Member is latest Triassic in age in the northern outcrop belt and earliest Jurassic in the southern outcrop belt. Alternatively, Dinosaur Canyon strata are uniformly of Triassic age and strata of earliest Jurassic age are not preserved in the southern outcrop belt. The relationship between the Moenave Formation and adjacent Wingate Sandstone and Rock Point Formation is also uncertain. Basal Moenave Formation strata of the southern outcrop belt interfinger with Wingate Sandstone and Rock Point Formation strata to the east. The base of the Moenave Formation in the northern outcrop belt is latest Triassic in age. One interpretation is that the Wingate Sandstone and Rock Point Formation are inferred to be latest Triassic in age. However, the age of basal Moenave Formation strata in the southern outcrop belt may be older than ! "##! the strata examined in this study implying the Moenave – Wingate – Rock Point interval to be early late Triassic in age. Strata of the Moenave Formation and overlying Kayenta Formation represent very different depositional environments. Moenave Formation strata were deposited in fluvial, lacustrine, and eolian subenvironments. Paleomagnetic data indicate these strata were deposited at latitude about 10° N. The Kayenta Formation strata represent Jurassic erg deposits with beds 10’s of meters thick. Modern large-scale eolian deposits are restricted to arid latitudes, north (or south) of about 23° latitude. The sub-Springdale unconformity must represent the appropriate amount of lost time for North America to move poleward from subtropical to arid latitudes, at least 13°. A detailed magnetic study of strata above and below this unconformity could constrain the amount of time missing between the Moeanve and Kayenta Formations. A small amount of inclination error has been documented for the Moenave Formation. Quantifying the amount of inclination error in Kayenta strata is also critical to understanding the paleolatitude of North America during this time. The base of the Jurassic is defined in marine strata by the lowest stratigraphic occurrence of the ammonite Psilocera spelae from Kuhjoch, Austria. This Jurassic typesection has no associated magnetic record making the placement of the Triassic- Jurassic boundary in terrestrial strata very difficult. A magnetostratigraphic study of the section at Kuhjoch, Austria, or any other section that preserves the lowest occurrence of P. spelae would be essential to determining the position of the Triassic-Jurassic boundary in both marine and terrestrial strata. ! "#'! The Valles Caldera lake sediments extracted from core VC304 yield the only ~200 ka long record of climate variability in the southwest. The environmental magnetic data obtained in this study indicate the lake environment responded to climate change in a non-direct way, meaning changes in lake environment do not always coincide with glacial interglacial transitions identified through geochemical investigations. It would be ideal to examine more sedimentary records from the Valles Caldera lake sediments to further investigate the relationship between climate and lake environment. Consistent variations in magnetic proxies across multiple Valles Caldera cores from different parts of the lake basin could substantiate interpretations of oxic and sub-oxic lake environments. If however, variability in magnetic proxies is not consistent among multiple sedimentary records, the variability in magnetic proxies is not indicative of changing lake environments. A seasonal environmental magnetism study of the material currently present in the lake catchment would clarify the fluvial input of magnetic phases most likely deposited in the mid-Pleistocene lake. Magnetic records from additional Valles Caldera cores would also help evaluate the robustness of the relative paleointensity record and validate the interpretation that these mid-Pleistocene sediments record geomagnetic field phenomenon 14! and 11!. ! "#(!