EDIACARAN-CAMBRIAN STRATIGRAPHY and PALEONTOLOGY of WESTERN NEVADA and EASTERN CALIFORNIA DISSERTATION Presented in Partial Fulf

Home , Cambrian Series 2, Cambrian Stage 3, Chapel Island Formation, Chengjiang, Cruziana, Deep Springs Valley

EDIACARAN-CAMBRIAN STRATIGRAPHY AND PALEONTOLOGY OF

DISSERTATION

Presented in Partial Fulfillment of the Requirements for the Degree Doctor of Philosophy in the Graduate School of The Ohio State University

Soo Yeun Ahn

Graduate Program in Geological Sciences

The Ohio State University

2010

Dissertation Committee:

Dr. Loren E. Babcock, Advisor Dr. William I. Ausich Dr. Matthew R. Saltzman Dr. Steven K. Lower

ABSTRACT

The Ediacaran-Cambrian transition was an important time span from both geologic and biologic perspectives. It was a time of dramatic evolutionary changes such as the diversification of early metazoans, the development of resistant skeletons in many taxa, and the escalation of prey-predator systems. In North America, the transition is well recorded in terminal Neoproterozoic to Cambrian strata of Esmeralda County, Nevada, and adjacent Inyo County, California. Strata recording this transition are the Deep Spring

Formation (Ediacaran-Cambrian), and the Campito, Poleta, and Harkless formations

(Cambrian).

For many years the Deep Spring, Poleta, and Harkless formations were informally divided into mappable members. New, formal names are proposed for the members of these formations. In ascending order the new members are the Dunfee, Montezuma, and

Gold Point members of the Deep Spring Formation; the Lida, Indian Springs, and

Clayton members of the Poleta Formation; and the Weepah and Alkali members of the

Harkless Formation. Two formal members of the Campito Formation, the Andrews

Mountain Quartzite and the Montenegro Member, have long been recognized.

The Deep Spring Formation and succeeding formations have the transition from a microorganism-dominated record to the more diverse and complex record of the

Phanerozoic. Stromatolites are common in the Dunfee Member of the Deep Spring

Formation, and fossilized microbial mats (―wrinkle structures‖) are present in siltstone ii

layers of the Montezuma Member. Microbial mats or microbially stabilized substrates are inferred to be responsible for the fine preservation of Ediacaran trace fossils as well as sedimentary structures. In Cambrian strata, microbial related structures are mostly wrinkle structures and gas escape structures.

Trace fossils have a trend toward increasing diversity, complexity, and abundance across the Ediacaran-Cambrian transition, although behavioral patterns are conserved.

These changes parallel faunal changes across the same interval of strata. Trace fossils in the Ediacaran part of the Deep Spring Formation include Bergaueria, Palaeophycus,

Planolites, and other simple resting, dwelling, and perhaps foraging traces. A new genus and species of trace fossil, Nevadichnos planum, is described from the Montezuma

Member of the Deep Spring Formation. Cambrian trace fossils of the Nevada-California succession include Bergaueria, Monomorphichnus, Palaeophycus, Planolites,

Rusophycus, and Treptichnus pedum. Although the morphologies of Cambrian tracemakers may have been different in the Cambrian than in the Ediacaran, the basic behavioral patterns of feeding and dwelling were already developed in the Ediacaran.

Chalcopyrite and limonite commonly occur within burrows, especially in

Planolites. Precipitation of these minerals is probably related to decay processes associated with biofilm development in burrows.

iii

ACKNOWLEDGMENTS

I am very grateful for Loren Babcock‘s advice and guidance throughout my work and dissertation preparation. I thank William Ausich, Matthew Saltzman, and Steven

Lower for helpful discussion and assistance with my work. I thank Margaret Rees, J.

Stewart and Mary Hollingsworth, Adam English and Stephen Leslie for assistance with collecting of samples and providing stratigraphic advice. I also thank my friends in the

School of Earth Sciences, and my family for their support during my work. I would like to acknowledge specially Alexa Sedlacek, Kate Tierney, and Alyssa Bancroft for all their support.

Funding was provided through the Friends of Orton Hall Fund, the National

Science Foundation (grants EAR-0073089, 0106883, awarded to Loren Babcock), and the Babcock Research Fund (School of Earth Sciences).

VITA

2002...... B.S. Geology, Kyungpook National University, Korea

2004...... M.S. Geology, Kyungpook National University, Korea

2005 to present ...... Graduate Teaching Associate, School of Earth Sciences, The Ohio State University

PUBLICATIONS

Ahn, S. Y. and Lee. S. J., 2003. Meso-Neoproterozoic bacterial microfossils from the Sukhya Tunguska Formation of the Turukhansk Uplift, Russia. Geosciences Journal, 7: 227-236.

Ahn, S. Y., and Babcock, L. E., 2009. Ediacaran-Cambrian transition reflected in trace fossils and sedimentary structures in Western Nevada - Eastern California. Geological Society of America Abstracts with Programs, vol. 41, no. 7, p. 293

Ahn, S. Y., Babcock, L. E., Rees, M. N., and Hollingsworth, J. S., 2009. Trace fossils from the Ediacaran-Cambrian transition in western Nevada: Behavioral holdovers from the terminal Neoproterozoic. Ninth North American Paleontological Convention Abstracts, p. 65-66.

Babcock, L. E., Zhu, M. Y., and Ahn, S. Y., 2009. A possible sprigginid fossil from the Cambrian of China. Ninth North American Paleontological Convention Abstracts, p. 66-67.

Ahn, S. Y., Babcock, L.E., Rees, M. N., and Hollingsworth, J. S., 2008. Body and Trace Fossils from the Deep Spring Formation (Ediacaran), Western Nevada. Geological Society of America Abstracts with Programs, vol. 40, no. 6, p. 143

FIELD OF STUDY

Major Field: Geological Sciences

TABLE OF CONTENTS

Abstracts ...... ii

Acknowledgments...... iv

Vita ...... v

List of Figures ...... ix

Chapters:

1. Introduction ...... 1

2. Ediacaran-Cambrian biostratigraphy of the southern Great Basin ...... 9

3. Ediacaran-Cambrian Chemostratigraphy of the southern Great Basin ...... 23

4. Ediacaran fossils and sedimentary structures from the Deep Spring Formation ...... 27

4. 1. Ediacaran fossils ...... 27

4.2. Systematic paleontology: trace fossils from the Deep Spring Formation ...... 34

4.3. Sedimentary structures from the Deep Spring Formation ...... 37

5. Cambrian body fossils and trace fossils from the Deep Spring, Campito, Poleta,

and Harkless formations ...... 45

5.1. Paleontology of the Deep Spring Formation ...... 47

5.1.1. Systematic paleontology: Cambrian trace fossils from the Deep Spring

Formation ...... 47

vii

5.2. Paleontology of the Campito Formation ...... 50

5.2.1. Systematic paleontology: body fossils from the Campito Formation ...... 51

5.2.2. Systematic paleontology: Trace fossils from the Campito Formation ...... 52

5.3. Paleontology of the Poleta Formation...... 55

5.3.1. Systematic paleontology: body fossils from the Poleta Formation ...... 56

5.3.2. Systematic paleontology: Trace fossils from the Poleta Formation ...... 59

5.4. Paleontology of the Harkless Formation...... 66

5.4.1. Trace fossils from the Harkless Formation ...... 67

5.4.1.1. Planolites-Palaeophycus problem ...... 67

5.4.1.2. Systematic paleontology: trace fossils from the Harkless Formation ...... 69

5.4.2. Sedimentary structures from the Harkless Formation ...... 80

6. Conclusions ...... 94

References ...... 97

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LIST OF FIGURES

Figures

1. Locality map of the Ediacaran-Cambrian strata in western Nevada and eastern California ...... 7

2. Geologic time scale and a composite δ13C curve for the Cambrian system ...... 8

3. Regional correlation of the Ediacaran-Cambrian strata of the southern Great Basin 19

4. Generalized stratigraphic column of the Deep Spring Formation ...... 20

5. Generalized stratigraphic column of the Poleta Formation ...... 21

6. Generalized stratigraphic column of the Harkless Formation...... 22

7. Generalized stratigraphic column of the Ediacaran-Cambrian strata in the southern Great Basin and Composite δ13C curve for the Cambrian...... 26

8. Locality map and outcrop photos of the Deep Spring Formation...... 40

9. Stratigraphic column and carbon isotope curve for the Deep Spring Formation ...... 41

10. Trace fossils from the Montezuma Member of the Deep Spring Formation ...... 42

11. Trace fossils and sedimentary structures from the Montezuma Member of the Deep Spring Formation...... 43

12. Sedimentary structures from the Montezuma Member of the Deep Spring Formation...... 44

13. Locality map and outcrop photos of the Cambrian strata in the southern Great Basin...... 46

14. Burrow type 1 from the Gold Point Member of the Deep Spring Formation in Mt. Dunfee...... 49

15. Body and trace fossils from the Montenegro Member of the Campito Formation. .... 54

16. Body fossils from the Indian Springs Member of the Poleta Formation...... 63

17. Helicoplacus gilberti from the Indian Springs Member of the Poleta Formation in Indian Springs...... 64

18. Trace fossils from the Indian Springs Member of the Poleta Formation, White-Inyo Mountains...... 65

19. Trace fossil composition of the Alkali Member of the Harkless Formation...... 86

20. Trace fossils and sedimentary structures from the Montezuma Member of the Deep Spring Formation...... 87

21. Composition of minerals co-occurring with Planolites from the Alkali Member of the Harkless Formation...... 88

22. Schematic diagrams showing the morphologies of Phycode, Trichophycus, and Treptichnus...... 89

23. Planolites from the Alkali Member of the Harkless Formation...... 90

24. Trace fossils from the Alkali Member of the Harkless Formation...... 91

25. Monomorphichnus and sedimentary structures from the Alkali Member of the Harkless Formation...... 92

26. Microbially produced structures from the Alkali Member of the Harkless Formation...... 93

CHAPTER 1

INTRODUCTION

The history of life is relatively short compared to the age of the Earth. The earliest record of life goes back to the Archean Eon as evinced by molecular biomarker and body fossil evidence (Anbar and Knoll, 2002). Prokaryotic microoganisms were evidently the first life forms on Earth, and they diversified using diverse metabolic pathways.

Prokaryotes are strongly versatile in their ability to adapt to extreme environments, such as environments of high temperature, high salinity, or low oxygen. They are inferred to have played a significant role in changing chemical conditions of the world ocean and atmosphere from dysoxic (or poorly oxygenated) to oxic through production of oxygen as a metabolic end product (Kaufman and Knoll, 1995; Canfield and Teske, 1996; Summons et al., 1999; Brocks et al., 1999, 2003; Knoll and Carroll, 1999; Knoll, 2000). This important process, coupled with Neoprototerozoic supercontinent amalgamation and breakup and perhaps other changes, led to significant changes in atmospheric and ocean composition (e.g., Lipps and Signor, 1992; Bengston, 1994; Knoll and Carroll, 1999;

Bottjer et al., 2000; Knoll, 2000; Anbar and Knoll, 2002; Babcock, 2005), making

Earth‘s oceans more suitable for the evolution of animals. In due course, the oceans experienced radiations of a variety of organisms during the Ediacaran Period (ca. 630-

542 Ma) and the Cambrian Period (542-488.3 Ma). Early evolutionary events are

exemplified by radiations of acritarchs and animals during the Ediacaran Period (Vidal and Moczydlowska, 1987, Moczydlowska and Vidal, 1988a; Knoll, 2000; Knoll, 2006).

Following a wave of extinction in the latest Ediacaran, a new set of animal radiations ensued during the Cambrian Period (Runnegar et al., 1975; Conway Morris, 1977a, 1985,

1989; Conway Morris and Whittington, 1985; Bowring et al, 1993; Bengtson, 1994;

Lieberman, 2003b; Waggoner, 2003; Babcock et al., 2001; Babcock, 2005, 2009). This event, or series of events, is commonly referred to as the ―Cambrian explosion‖ (Cloud,

1948; Bengtson, 1994; Brasier and Linsay, 2001; Conway Morris, 1977a, 1985, 1989,

1998, 2006; Marshall, 2006). Although often characterized as an explosive radiation, recent chronostratigraphic (including geochronologic) evidence and stratigraphic reinvestigation indicates the Cambrian diversification was resolved into a series of adaptive events punctuated by waves of extinctions (Zhu et al., 2006; Babcock, 2009;

Peng et al., in press).

For nearly a century, the key distinction between strata assigned to the

―Precambrian‖ and those assigned to the Phanerozoic Eon was the relative richness of fossils. However, discovery of fossils representing organisms of diverse body plans, largely in Lagerstätten deposited during the Ediacaran Period, has changed that view.

Although there is discussion about the affinities of these fossils (e.g., Seilacher, 1989,

1992; Retallack, 1992, 1994; Zhuravlev, 1993; Steiner and Reitner, 2001; Peterson et al.,

2003; Waggoner, 1995), it now seems that many can be reasonably accommodated as metazoan animals (Glaessner and Daily, 1959; Glaessner and Wade, 1966; Conway

Morris, 1993; Gehling, 1999; Martin et al., 2000). Such fossils suggest that the diversity

and abundance of Ediacaran metazoans was rather high. Some of the Ediacaran organisms may not have descendants in Phanerozoic strata. The Ediacaran Period represents a ―bridge‖ of sorts in the history of life on Earth, a bridge between primarily nonbiomineralizing Proterozoic life forms and more rigidly skeletonized, including biomineralizing, Phanerozoic life forms. The Ediacaran Period also represents the last major interval during which microorganisms were the dominant life forms in most depositional environments (Seilacher and Pflüger, 1994; Bottjer et al., 2000; Gehling,

1999). The Ediacaran saw the initial rise of metazoan animals, and afterward, there was a major shift toward domination of the fossil record by more complex eukaryotic organisms (Fedonkin, 1992; Clapham et al., 2003; Fedonkin, 2003; Narbonne and

Gehling, 2003; Narbonne, 2005; Droser et al., 2006; Valentine, 2007; Vickers-Rich,

2007).

The Ediacaran Period was the time of initial biomineralization of animal skeletons, but biomineralization became a much more important evolutionary theme during the Cambrian Period. One group of presumed animals, the cloudiniids, evolved the capacity for biomineralization of the skeleton in the middle to late part of the

Ediacaran Period (Grant, 1990; Grotzinger et al, 1995; Bengtson and Yue, 1997; Knoll,

2000). Biomineralizing cloudiniids are not known to have survived into the Phanerozoic, and animals that evolved the capacity for biomineralization apparently did so convergently. Thickened organic cuticles, lightly biomineralized skeletons, and heavily biomineralized skeletons developed across a wide range of animal groups in a series of phases through the Cambrian Period, contributing substantially to a rich Phanerozoic

body fossil record (Bengtson, 1994; Lieberman, 2003b, 2008; Zhu et al., 2006; Babcock,

2006; Peng et al., in press).

Ediacaran fossils are mostly of nonbiomineralizing forms (Gehling, 1999), and it has been suggested that fossilization of most Ediacaran macroorganisms is the result of covering or encasement of bodily remains by microbial mats (Wade, 1968; Seilacher,

1992; Gehling 1999). Until now, this hypothesis has not been directly tested. Encasement of carcasses by microbial mats is expected to have set up sedimentary and taphonomic conditions that could preclude predation and extensive decomposition (Gehling, 1999).

Trace fossils showing the activities of animals left in sediments have been used to supplement our understanding of the record of early animals, and they may have also been preserved through stabilization of the substrate by microbial consortia. As a result of similar preservational circumstances, it may not always be easy to distinguish between body fossils and trace fossils in Ediacaran rocks. In any case, both body fossils and trace fossils preserved, presumably, through mediation by the microorganisms associated with microbial mats, can be used as indicators of increasing number and diversity of Ediacaran animals.

It is uncertain what the extent of grazing and predation was during the

Proterozoic, but evidence suggests that the level of grazing was insufficient to be seriously detrimental to microbial mat communities. Without significant cropping, microbial communities, including photosynthetic forms such as cynaobacteria, literally dominated the Proterozoic Earth. Evidence of their presence in the form of stromatolites and thrombolites were especially numerous in shallow shelf environments for more than

a billion years (Awramik, 1971; Walter and Heys, 1985; Grotzinger, 1990), and the photosynthesizers must have contributed to increased oxygenation of the atmosphere

(Knoll, 1992; Summons et al., 1999; Anbar and Knoll, 2002; Canfield et al., 2007).

Microbial stabilization of the during the Proterozoic would have led to a sharp sediment-water interface. Bioturbation by benthic organisms, in contrast, would have tended to reduce the sharpness of that boundary (Droser and Bottjer, 1986, 1988; Crimes and Droser, 1992; Droser et al, 1999; McIlory and Logan, 1999; Bottjer et al., 2000;

Dornbos and Bottjer, 2000; Babcock et al., 2001; Jensen et al., 2005). For this reason, study of bioturbation trends in benthic substrates provides additional evidence of the diversification and abundance history of early animals. Bottjer et al. (2000) explained the changeover from mat-dominated marine sediment surfaces of the Proterozoic to more highly burrowed, fluidized sediment surfaces of the Cambrian and beyond as the

―Cambrian substrate revolution.‖ Babcock (2003) showed, using the example of

Cambrian trilobites as both predators and prey, that there was an intimate relationship between the adaptive radiation of animals in the Cambrian and changes in substrate conditions, and he proposed the term ―Early Paleozoic Marine Revolution‖ to explain the linked changes in animal evolution and substrate modification across the Ediacaran-early

Paleozoic transition.

The purpose of this dissertation is to describe evidence of biotic change across the

Ediacaran-Cambrian boundary in the southern Great Basin, western United States

(Figures 1-2). Sections in Esmeralda County, Nevada, and adjacent Inyo County,

California, demonstrate one of the most complete known composites for this interval

from Laurentia (the North American-Greenland paleocontinent). Types of data used in this investigation include field observations of sedimentary strata and their components, and laboratory investigations of fossils and sedimentary structures. This new information is integrated with existing sedimentary, lithostratigraphic, biostratigraphic, and chemostratigraphic data to produce a coherent picture of the biotic record through the

Ediacaran-Cambrian transition in western North America. To clarify stratigraphic discussion, new formal names are proposed for members of three formations mapped through western Nevada and eastern California.

Figure 1. Locality map of the Ediacaran-Cambrian strata in western Nevada and eastern California (map modified from http://maps.google.com).

Figure 2. Geologic time scale and a composite δ13C curve for the Cambrian system (from Zhu et al., 2006).

CHAPTER 2

EDIACARAN-CAMBRIAN STRATIGRAPHY OF THE SOUTHERN GREAT BASIN

Ediacaran to Cambrian strata studied from the southern Great Basin include, in ascending order, the Deep Spring, Campito, Poleta, and Harkless formations. These units have been studied since the early 1900s, and are notable for the presence of the earliest known trilobites from Laurentia (Walcott, 1910; Kirk (in Knopf, 1918); McKee and

Moiola, 1962; Nelson, 1962; Stewart, 1966; Stewart, 1970; Albers and Stewart, 1972).

Together, these formations record shelf to ramp sedimentation, and they thicken considerably to the present-day west, which was the offshore direction (Stewart, 1970).

The Deep Spring, Campito, Poleta, and Harkless formations were mapped in detail, and named after sections in California and Nevada by early workers (Kirk (in

Knopf, 1918); McKee and Moiola, 1962; Nelson, 1962; Stewart, 1966, Stewart, 1970).

For the most part, formations were named and mappable members within them were given only informal names. The Campito Formation is an exception; members were previously given formal names (Nelson, 1962), the Andrews Mountain Member (lower) and the Montenegro Member (upper). The Deep Spring, Poleta, and Harkless formations have been divided into two or three members each, but those members have remained as informal units. Here, based on new work and previous information (Nelson, 1962;

Stewart, 1966, 1970; Albers and Stewart, 1973), formal members are proposed for each of these three formations.

Base of the Cambrian System in the southern Great Basin

The southern Great Basin is a classic area for study of questions relating to the late Neoproterozoic and early half of the Cambrian. Ediacaran and Cambrian strata have been investigated and mapped since the early 1900s (Walcott, 1910; Kirk (in Knopf,

1918); McKee and Moiola, 1962; Nelson, 1962; Stewart, 1966; Stewart, 1970). These rocks are important for stratigraphic, paleontologic, and economic reasons. The major units through this region include, in ascending order, the Reed, Deep Spring, Campito,

Poleta, and Harkless formations (Figure 3). Together, these formations record deposition in shallow shelf to ramp environments of the Cordilleran margin of Laurentia. In general, increasing thicknesses and a trend toward increasing percentages of finer siliciclastic sediments toward the present-day northwest suggests that the shelf deepened toward the

White-Inyo succession. The source of detrital sediment appears to be from the Death

Valley region, to the southeast (Moore, 1976). Thicknesses of sedimentary strata in the

White-Inyo region and in Esmeralda County, Nevada, are significantly greater than those in equivalent strata of Death Valley.

The position of the Proterozoic-Cambrian boundary in the southern Great Basin has been a focus of discussion since the 1960s. At that time, the boundary was inferred to be present either in the Deep Spring Formation or the Campito Formation. The pre-

trilobite fossil Wyattia and Pteridinium-like fossils (Taylor, 1966; Cloud and Nelson,

1966) suggested a Proterozoic age for the Deep Spring Formation. Alternatively, the biomineralized Cloudina, which was originally considered to be a Cambrian fossil

(Germs, 1972; Grant, 1990) occurs in the informal lower member (Dunfee Member, herein) of the Deep Spring Formation, suggesting that portions of that unit are Cambrian in age (Parsons, 1996). The Campito Formation, which contains olenelline trilobites and archaeocyaths (McKee and Moiola, 1962; Nelson, 1962; Stewart, 1966; Stewart, 1970), is unequivocally of Cambrian age.

Part of the problem with identification of the position of the Cambrian base was the lack of a stable, agreed-upon horizon for the base of the system globally. In Walcott‘s

(1892) concept of the Cambrian, the de facto standard for about a century, the Cambrian base was marked by the first appearance of trilobites. However, the first appearance of trilobites has been shown to be non-synchronous globally (Geyer and Shergold, 2000), which led to search for a different boundary position. Ratification of a much different position, at a horizon corresponding to the first appearance of the trace fossil Treptichnus pedum (Brasier et al., 1994; Landing, 1994), finally provided for a stable, unequivocal

Cambrian base. The ratified GSSP (Global boundary Stratotype Section and Point) for the Cambrian System is at a horizon well below most previously suggested positions, and considerably below the first appearance of trilobites (Knoll et al., 2004). A thick succession of strata once considered to be Upper Proterozoic automatically became

Cambrian at the time of GSSP ratification. This interval roughly corresponds to the

Terreneuvian Series/Epoch of the Cambrian (542-521 Ma), which represents the first

39% of Cambrian time. Included in the ―new‖ definition of the Cambrian were pre- trilobitic strata containing assemblages of small shelly fossils (SSFs) and strata containing the earliest known biomineralizer, Cloudina. Change of the position of the

Cambrian base to include a thick pre-trilobitic succession was one of the main arguments for ultimately subdividing the Cambrian System/Period into four series/epochs, rather than the historic number, which was three. It also meant that the main phase of the

―Cambrian explosion,‖ which was previously regarded as occurring near the beginning of the Cambrian, now took place much later during Cambrian time.

Paleontologic and chemostratigraphic work has closely constrained the position of the Cambrian base in the southern Great Basin (Signor and Mount, 1986; Corsetti and

Kaufman, 1994; Corsetti et al, 2000; Corsetti and Hagadorn, 2000, 2003). Corsetti and

Hagadorn (2003) reported Treptichnus pedum from the informal middle member

(Montezuma Member, herein) of the Deep Spring Formation. Chemostratigraphic study of the Deep Spring Formation in eastern California has the onset of a large negative δ13C excursion (-4 ‰) (Corsetti and Kaufmann, 1994; Corsetti et al., 2000; Corsetti and

Hagadorn, 2003) coinciding with the appearance of T. pedum in the Montezuma Member

(new) of the Deep Spring Formation. That excursion, which is expressed intercontinentally, was later designated the Basal Cambrian Carbon isotope Excursion

(BACE; Zhu et al., 2006). The BACE excursion is now used as the primary tool for constraining the base of the Cambrian System in all areas of the world (Peng and

Babcock, 2008; Peng et al., in press).

Deep Spring Formation

The Deep Spring Formation is a heterogeneous carbonate and siliciclastic unit spanning the uppermost Neoproterozoic (Ediacaran System) to the lower part of the

Cambrian System. The formation crops out sporadically in Esmeralda County, Nevada, and Inyo County, California. Localities in Nevada include Mt. Dunfee, the Montezuma

Range, and Clayton Ridge. The formation was named by Kirk (in Knopf, 1918) for exposures adjacent to the Deep Springs Valley, California. Stewart (1970) informally divided the formation into three members. Each member has a different proportion of carbonates and siliciclastics (Stewart, 1970). The Deep Spring Formation at Mt. Dunfee has an almost complete composite section of the lower member (204 m), middle member

(179 m), and upper member (74 m).

New member names are proposed here to replace informal member names used by Stewart (1970) (Figure 4). The members are mappable through western Nevada and eastern California. The new name ―Dunfee Member‖ is assigned to the informal lower member of the Deep Spring Formation, after Mt. Dunfee, Esmeralda County, Nevada.

The member is mainly composed of gray to pale orange dolostone with minor greenish gray siltstone and rare calcareous sandstone and quartzite (Albers and Stewart, 1972).

Body fossils of the biomineralizing Cloudina occur in the Dunfee Member and are indicative of an Ediacaran age for that part of the section.

The new name ―Montezuma Member‖ is assigned to the informal middle member of the Deep Spring Formation as previously used by Stewart (1970). The member is

named for exposures in the Montezuma Range, Esmeralda County, Nevada. The

Montezuma Member is composed of grayish limestone, oolitic limestone, greenish gray to pale orange siltstone, yellowish gray, fine-grained quartzite, and calcareous sandstone

(Albers and Stewart, 1972). The member contains fossils suggesting an Ediacaran to

Cambrian age. Some siliciclastic beds in the member yield specimens of the trace fossil

Treptichnus pedum, the first appearance of which is a proxy for the base of the Cambrian

System. Also, carbon isotopes from the member have a large negative δ13C excursion

(Corsetti and Kaufman, 1994; Corsetti and Haradorn, 2003) interpreted as the BACE

(Basal Cambrian Carbon isotope Excursion; Zhu et al., 2006, Peng and Babcock, 2008l

Peng et al., in press).

The informal upper member of the Deep Spring Formation is named the ―Gold

Point Member‖ after the town of Gold Point, near Mt. Dunfee. The member is composed of greenish gray siltstone and fine grained silty quartzite (Albers and Stewart, 1972). The

Gold Pont Member contains T. pedum and other trace fossils such as Cruziana and

Rusophycus (Stewart, 1970), all indicative of a Cambrian age.

Campito Formation

The Campito Formation was named by Kirk (in Knopf, 1918) for Cambrian exposures near Campito Mountain, California. The formation is approximately 1067 m thick at Campito Mountain (Nelson, 1962). The contacts between the underlying Deep

Spring Formation and the overlying Poleta Formation appear to be transitional. The

formation is divided into two members, the lower Andrews Mountain Member and the upper Montenegro Member, both named by Nelson (1962). The members are well exemplified in the White and Inyo Mountains, California, and the Weepah Hills, Nevada.

Thicknesses of the members are similar in the two localities (Stewart, 1970).

The Andrews Mountain Member (Nelson, 1962) is composed of dark dray to black quartzitic sandstone with minor amounts of siltstone and shale. The member is 760 to 884 m in the White-Inyo Mountains (Albers and Stewart, 1972). The Montenegro

Member is dominated by dark greenish gray siltstone. It is about 96 m thick in the

Weepah Hills, Nevada (Albers and Stewart, 1972).

Poleta Formation

The Poleta Formation was named by Nelson (1962) after Poleta Canyon,

California (Nelson, 1962) for successions of limestone, siltstone, and quartzite above the

Campito Formation. The Poleta Formation is widely exposed in the southern Great Basin.

In a classic study, Moore (1976) demonstrated shallow marine and tidal flat to offshore depositional environments within the Poleta Formation. The formation was originally divided into two members, a lower archaocyath-rich limestone member, and an upper member that is dominated by greenish gray silicicalstics and thin limestone beds

(Nelson, 1962). McKee and Miola (1962) and Stewart (1966) later subdivided the upper member to form a threefold division of the formation. Since then, lower, middle, upper members have been cited (McKee and Miola, 1962; Moore, 1976; Stewart, 1966; 1970).

New, formal member names are proposed here to replace the informal nomenclature of the Poleta Formation (Figure 5). The members are mappable through western Nevada and eastern California.

The name ―Lida Member‖ is proposed for the informal lower member of the

Poleta Formation. It is named after exposures near Lida Junction, in hills adjacent to the

Cottontail Ranch, where light gray crystalline and oolitic limestone yield abundant archaeocyaths (Albers and Stewart, 1972, p. 11). At the type section, the Lida Member is

125 m thick. However, thickness of the member varies considerably from one locality to the next (Moore, 1976; Stewart, 1970; Albers and Stewart, 1972).

The name ―Indian Springs Member‖ is proposed for the informal middle member of the Poleta Formation as used by Stewart (1970). The name is derived from Indian

Springs Canyon, Esmeralda County, Nevada, where the member is well exposed. The member is 183 to 213 m thick, and composed of siltstone, limestone, and sandstone or quartzite (Stewart, 1970), representing terrigenous sedimentation in tidal flat to offshore environments (Moore, 1976). In Indian Springs Canyon, some body fossils collected from the Indian Springs Member have exceptional preservation (Babcock et al., 2000;

English, 2007; English and Babcock, in press). The Indian Springs Lagerstätte includes brachiopods, trilobites, agglutinated protists, helicoplacoids, and other fossils (Babcock et al., 2000; English, 2007).

The name ―Clayton Member‖ is proposed for the informal upper member of the

Poleta Formation. The Clayton Member is named for Clayton Ridge, Esmeralda County,

Nevada. The member is about 40 m thick, and is dominated by medium gray, thinly bedded limestone that often bears archaeocyaths.

Harkless Formation

The Harkless Formation is the most widely exposed Cambrian formation in

Esmeralda County, Nevada (Albers and Stewart, 1972). The formation was named by

Nelson (1962) after Harkless Flat, adjacent to Waucoba Spring, California. The formation contains greenish gray siltstone, and thin-bedded to fine-grained quartzitic sandstone and siltstone (Nelson, 1962). In Esmeralda County, the formation is approximately 1067 m thick in the Weepah Hills, where the formation is almost completely exposed (Stewart,

1970). The contact between the Poleta Formation and the overlying Harkless Formation is marked by a sharp transition from limestone beds of the Clayton Member to greenish siltstone beds of the Harkless Formation. Lithology of the formation is mainly siliciclastics, however, the siliciclastics are commonly metamorphosed (Albers and

Stewart, 1972). New member names are proposed here (Figure 6). The members are mappable through western Nevada and eastern California.

The name ―Weepah Member‖ is proposed for the informal lower member of the

Harkless Formation. The name is derived from the Weepah Hills, Esmeralda County,

Nevada, where Stewart (1970) mapped a complete section of the formation. The member is about 420 m thick, and dominated by quartzitic siltstone with minor amounts of coarse siltstone and fine sandstone layers (Albers and Stewart, 1972).

The name ―Alkali Member‖ is proposed for the informal upper member of the

Harkless Formation. The name is derived from Alkali Spring, Esmeralda County, Nevada, which is near the Weepah Hills section. The member is about 672 m thick in the Weepah

Hills section, and is characterized by a predominance of olive greenish to dark greenish gray fine siltstone and its quartzite (low-grade metamorphic) equivalent (Albers and

Stewart, 1972). The member is about 672 m thick in the Weepah Hills. Near Alkali

Spring, some strata resemble the Saline Valley Formation, which was mapped in eastern

California by Nelson (1962). The Saline Valley Formation is not identified in Nevada; however, pale browinish to gray crystalline limestone or lime mudstone, limestone containing well rounded quartz grains, and thin sandstone beds near Alkali Spring,

Nevada, have been referred to as the Saline Valley Formation-equivalent (Stewart, 1970;

Albers and Stewart, 1972).

Figure 3. Regional correlation of the Ediacaran-Cambrian strata of the southern Great Basin (modified from Hollingsworth, 2005).

Figure 4. Generalized stratigraphic column of the Deep Spring Formation (modified from Albers and Stewart, 1972).

Figure 5. Generalized stratigraphic column of the Poleta Formation (modified from Albers and Stewart, 1972).

Figure 6. Generalized stratigraphic column of the Harkless Formation (modified from Albers and Stewart, 1972). 22

CHAPTER 3

EDIACARAN-CAMBRIAN CHEMOSTRATIGRAPHY OF THE

SOUTHERN GREAT BASIN

Carbon isotope (δ13C) chemostratigraphy has emerged in recent years as a valuable tool for global chronostratigraphic correlation in Cambrian rocks (Magaritz et al,

1986; Brasier et al., 1996; Shields, 1999; Montanez, 2000; Zhu et al., 2006; Peng and

Babcock, 2008). Its application in the lower two series of the Cambrian has greatly improved chronostratigraphic resolution, particularly in successions for which biostratigraphic tools provide insifficient information. Through some intervals of the

Terreneuvian Series and Series 2, chronostratigraphic correlation is essentially unresolvable by biostratigraphy alone, and δ13C chemostratigraphy provides sufficient information to identify major stratigraphic tie points intercontinentally.

δ13C is a measure of the ratio of 13C to 12C relative to a standard. This ratio shows the relative abundance between organic (light) carbon and inorganic carbon. The key factors controlling δ13C are primary production and organic carbon burial in marine environments. Carbon isotope values can be a proxy for productivity levels of ancient marine life. Carbonate strata globally have carbon isotopic values reflecting such levels.

A composite δ13C curve for the entire Cambrian has a series of distinctive, high- amplitude excursions (Zhu et al., 2006). Comparing the general pattern and magnitude of

individual δ13C excursions enables intrabasinal, regional, and global correlation of

Cambrian strata of the southern Great Basin. Biostratigraphic zonation of the southern

Great Basin can be calibrated against the δ13C curve to achieve fine chronostratigraphic resolution across a significant area of Laurentia (Figure 7).

The base of the Cambrian is marked by a large negative excursion (Basal

Cambrian Carbon isotope Excursion, BACE; Zhu et al., 2006) that has been used as the principal stratigraphic tool for correlating the base of the Cambrian globally (Strauss et al., 1992; Zhu et al., 2006; Peng and Babcock, 2008). A large negative excursion that is interpreted as the BACE was reported from the Montezuma Member of the Deep Spring

Formation, in the White-Inyo Mountains (Corsetti and Kaufmann, 1994; Corsetti et al.,

2000; Corsetti and Hagadorn, 2003). The FAD of the trace fossil Treptichnus pedum, which approximately coincides with the Cambrian GSSP at Fortune Head, Newfoundland, was reported from the Montezuma Member of the Deep Spring formation (Corsetti and

Kaufmann, 1994; Corsetti et al., 2000; Corsetti and Hagadorn, 2003) The onset of the

BACE excursion approximates the first appearance of T. pedum in the White-Inyo

Mountains. The combination of a carbon isotope signal and the presence of the boundary marker implies that the global signal of the Cambrian base lies within the Deep Spring

Formation (Figure 7, Figure 9).

13 Carbon isotopic analyses (δ Ccarbonate) through part of the Campito-Poleta succession in the Montezuma Range indicates an incomplete record of a positive excursion interpreted (English, 2007) as the Cambrian Arthropod Radiation carbon isotopic Excursion (CARE; Zhu et al., 2006). The onset of the CARE excursion has not

yet been identified in Laurentia, so its precise position is uncertain. Peng and Babcock

(2008), however, correlated the point of onset (from the upper part of the Sinosachites flabelliformis-Tannuolina zhangwentangi Zone of South China) to a level just below the base of the Fritzaspis Zone of Laurentia (uppermost Stage 2). The point of greatest positive inflection in the δ13C curve occurs in the Fallotaspis Zone (lower Stage 3), and

δ13C values return to slightly less than 0 in the uppermost Fallotaspis Zone-lower

Nevadella Zone (lower Stage 3).

Although Ediacaran-Cambrian Series 2 sedimentary strata in the southern Great

Basin appear to have conformable relationships, large gaps in the δ13C record suggest that significant disconformities exist in several places. For example, carbon isotopic analyses

13 (δ Ccarbonate) through part of the Campito-Poleta succession in the Montezuma Range shows an incomplete record of a positive excursion interpreted as the CARE excursion

(English, 2007; see Zhu et al., 2006); the onset of the excursion may be lost to a disconformity (Figure 7). Similarly, Hollingsworth (1999) showed, based on field evidence, the presence of missing strata at a sequence boundary within the Campito

Formation. Pending other field study, the position of disconformities are not illustrated in the generalized stratigraphic sections illustrated here (e.g, Figures 3-6).

Figure 7. Generalized stratigraphic column of the Ediacaran-Cambrian strata in the southern Great Basin (left) and composite δ13C curve for the Cambrian (right). Carbon isotope data adapted from Corsetti and Kaufman (1994) for the Deep Spring Formation, and from English (2007) for the Campito-Poleta successions. AECE excursion is reconstructed based on biozones established by Hollingsworth (2005). Global δ13C curve is modified from Zhu et al. (2006).

CHAPTER 4

EDIACARAN FOSSILS AND SEDIMENTARY STRUCTURES

FROM THE DEEP SPRING FORMATION

4.1. Ediacaran trace fossils

Body fossils recording the early phases of metazoan evolution have an uneven stratigraphic distribution. Body fossils are characteristically rare during the Ediacaran, and become increasingly common during the Cambrian, especially the post-Terreneuvian

Cambrian. In some Ediacaran successions, including the Fortune Head, Newfoundland, section, which hosts the Cambrian GSSP, trace fossils provide the primary source of information about early animals. The same is true of Ediacaran strata in the southern

Great Basin.

Interpreting Ediacaran traces, or possible traces, can pose challenges. Compared to Phanerozoic traces, Ediacaran traces represent a much narrower range of morphologies. In general, Ediacaran trace fossils are of simple morphology (circular or sinuous burrows), shallow, and predominantly horizontal. To a large extent their preservation seems to be associated with microbial mats. Sedimentary features of the

Phanerozoic have sometimes been misinterpreted as trace fossils (e.g., Osgood, 1970), a 27

problem that may be especially acute in strata deposited under the influence of microbial mats, or sedimentary deposits subject to episodic storm influence. This problem may be exacerbated in Ediacaran deposits, where microbial mat-dominated sediment surfaces are not unusual.

In this chapter, Ediacaran trace fossils from the Deep Spring Formation

(Montezuma Member, below the level of the Cambrian base) are described. In addition, apparent sedimentary structures are described. Such sedimentary features could be misinterpreted as trace fossils.

Trace fossils from the Ediacaran part of the Deep Spring Formation are dominated by simple, shallow, and small traces mostly representing feeding behaviors. Trace fossils with composite behaviors (feeding and dwelling, resting and locomotion) are also present. The trace fossils generally co-occur with microbially produced or microbially related structures, as exemplified by wrinkle structures or gas-escape structures.

The most abundant traces in the Ediacaran part of the Deep Spring Formation are the simple, tubular Planolites and Paleophycus and small dimplelike traces. The distinction between Planolites and Paleophycus is difficult to make, due partly to the small size of the specimens. Planolites is defined as a tube lacking a distinct rim, and

Paleophycus is defined as a tube having such a rim, often enhanced by limonite staining

(Marenko and Bottjer, 2008). Observations on Ediacaran-Cambrian specimens from

Nevada indicate that these are end-member morphologies, and specimens can have both rimmed and non-rimmed intervals along a single tube. Polished cross sections reveal

slightly different backfill lithologies within the burrows, and this is interpreted as evidence of a sediment deposit-feeding behavior in the tracemakers.

Chronostratigraphic relationships in the Deep Spring Formation of western

Nevada are based on correlations made from sections in adjacent eastern California, where the position of the trace fossil Treptichnus pedum (Rowland and Corsetti, 2002;

Corsetti and Hagadorn, 2003) has been matched to the secular curve of carbon isotopes

13 (δ Ccarbonate) in continuous stratigraphic sections (Corsetti and Kauffman, 1994; Corsetti et al., 2000). The first appearance datum (FAD) of T. pedum in the Fortune Head section,

Newfoundland, Canada, as understood at the time of boundary ratification (Brasier et al.,

1994; Landing, 1994; Landing et al., 2007), was used as the primary correlation tool whose position coincided with the Cambrian GSSP. Since the time of boundary ratification, however, it has been shown that T. pedum also occurs below the position of the GSSP in the Fortune Head section (Gehling et al., 2001). Apart from a general increase in trace fossil diversity close to the Cambrian base, other chronostratigraphic guides to the position of the boundary horizon have not been identified from the Fortune

Head section because of lithofacies setting (nearshore marine succession, with most trace fossils preserved in sandstone layers) and tectonic overprint.

To overcome problems associated with tracing the T. pedum horizon intercontinentally and through varied facies types, the International Subcommission on

Cambrian Stratigraphy (ISCS) has adopted the position (Peng and Babcock, 2008) that the base of the Cambrian closely approximates the base of a large shift in carbon isotopes

(δ13C) termed the BACE curve (Base of Cambrian carbon isotopic Excursion; Zhu et al.,

2006). This interpretation follows from work in the Deep Spring Formation of eastern

California, where Corsetti and Kaufmann (1994) and Corsetti and Hagadorn (2003) showed that the first appearance of T. pedum in the Great Basin nearly coincides with the onset of this strong, readily identifiable excursion. For purposes of correlation in the

Great Basin and globally, the base of the BACE curve is taken to mark the base of the

Cambrian System.

The position of the Cambrian base in the Deep Spring Formation of western

Nevada is largely inferred from lithologic and minor biostratigraphic correlations drawn from eastern California, where the position of the BACE curve was identified (Corsetti and Kaufmann, 1994; Corsetti and Hagadorn, 2000, 2003). Carbon isotopic data from the

Deep Spring Formation (Corsetti and Kauffman, 1994) have the excursion in the

Montezuma Member of the Deep Spring Formation. In Nevada, it is difficult to find undisturbed Deep Spring strata showing the whole transition interval from the Ediacaran to the Cambrian because of post-Cambrian tectonic disturbance. The sources of trace fossils described here are the presumed Ediacaran part of the Deep Spring Formation in the Mount Dunfee and Montezuma Peak areas (see Albers and Stewart, 1972). In these areas, sedimentary strata of the Deep Spring Formation have undergone folding, faulting, and metamorphic alteration. Some sedimentary structures and fossils seem to have been obliterated or overprinted.

Sections of the Deep Spring Formation in Nevada, particularly near Mt. Dunfee, are either too incomplete, or have too few carbonate layers to allow development of a

13 meaningful δ Ccarbonate curve (Figure 9). At Mt. Dunfee, the relevant stratigraphic intervals are highly faulted, and completeness of individual sections is the biggest issue.

In the Montezuma Range, lack of thick carbonate intervals is the issue. For these reasons,

13 the base of the Cambrian is approximated by correlating the δ C curve from eastern

California, with constraints provided by trace fossil, body fossil, and microbialite evidence. Strata containing the trace fossil T. pedum are inferred to be Cambrian

(ignoring the caveat implied by restudy of the trace fossil‘s occurrence in the Fortune

Head, Newfoundland, section; Gehling et al., 2001) and strata containing the biomineralized body fossil Cloudina, as well as microbialites, are inferred to be

Ediacaran.

The first occurrence of T. pedum is a few tens of meters below the top of the

Montezuma Member (Hagadorn and Corsetti, 2001; Rowland and Corsetti, 2002).

Microbialites and Cloudina are also present in the Montezuma Member of the Deep

Spring Formation (Oliver and Rowland, 2002), but below the T. pedum horizon. Thus, the totality of evidence from fossils and biogenically mediated sedimentary structures indicates that the Cambrian base lies in the upper half of the Montezuma Member of the

Deep Spring Formation in Nevada. A more precise placement of the position is not possible at present because of the lack of definitive and closely placed chronostratigraphic indicators.

In consideration of the facts about the Nevada sections, some general interpretations about ages of the Deep Spring strata are made. Simple dotted trace fossils and sedimentary structures collected at Mt. Dunfee are inferred to be Ediacaran age because the fossils co-occur in the same horizons with microbialites. Ages of trace fossils collected from the Montezuma Range are somewhat more problematic because the Deep

Spring outcrops have a patchy distribution, which makes regional correlation difficult. In addition, fossils are sporadically distributed through the Montezuma Range sections, so they provide only general chronostratigraphic constraints.

There is a difference in the types and proportions of trace fossils collected from the Mt. Dunfee and Montezuma Range sites (Figure 8). At Mt. Dunfee, small surface dimples are abundant, but they are not common in the Montezuma Range. Instead, simple unbranched burrows assignable to Planolites (or Palaeophycus) are the most common traces. Ichnogenera such as Burgaueria and Nevadichnos make up a minor portion of the ichnofossil assemblage recorded from the Montezuma Range. Bergaeuria is present in both areas but somewhat more common in the Montezuma Range. The difference in complexity and diversity of trace fossils between the two localities can be interpreted in two ways. First, some traces from the Montezuma Member of the Deep Spring Formation in the Montezuma Range may have come from the Cambrian part of the unit, and the relatively greater complexity and diversity (compared to Mt. Dunfee) may reflect metazoan evolution close to the beginning of the Cambrian. Mitigating somewhat against this interpretation is the presence of Nevadichnos, a new trace fossil inferred to have been

constructed by the Ediacaran Parvancorina (Naimark and Ivantsov, 2009) or a similar animal. Lin et al. (2006) suggested a close phylogenetic relationship between

Parvancorina and the Cambrian arthropod Skania. If true, it is conceivable that

Nevadichnos could have been constructed by either an Ediacaran or Cambrian arthropod or arthropod relative. Evidence from sedimentary structures favors an interpretation of the Montezuma Range trace fossil localities as Ediacaran in age.

A tool mark, characterized by sets of parallel wrinkles radiating from a circular central knob, was collected only from the Montezuma Member in the Montezuma Range.

In the northern part of the Montezuma Range, this tool mark co-occurs with Cloudina.

This tool mark is inferred to have been developed in mat-stabilized sediment, and is further interpreted to be of Ediacaran age. If so, trace fossils co-occurring with this tool mark should also be of Ediacaran age. Environmental differences between coeval localities may account for the differences in trace fossil assemblages. The circular tool mark structure was not collected from Mt. Dunfee, and its lack may be related to differences between the two areas in the distribution of mat-stabilized surfaces, or more likely, the extent of episodic current (e.g., storm) influence on the substrate.

4.2. Systematic paleontology: Ediacaran trace fossils from the Deep Spring Formation

Nevadichnos, new ichnogenus

Type ichnospecies: Nevadichnos planum, new ichnospecies

Description. Bedding plane trace, elongate-subtriangular in outline. Wider end with frontal rim, from which extends a narrow, moderately long medial groove. Middle section with multiple transverse, slat-like grooves. Narrow end triangulate, tapering quickly, with multiple transverse, slat-like grooves.

Discussion. The monospecific ichnogenus Nevadichnos new ichnogenus is a trace fossil unlike any other known. It resembles the body of the Ediacaran body fossil Parvancorina and the Cambrian arthropod Skania, but the two available specimens are much larger than any described Parvancorina or Skania. The more complete specimen, the holotype of N. planum new ichnospecies, is 11 cm in length, and 7 cm at its widest point. It has seven distinct slat-like grooves on the right side of the specimen, as preserved. The fossil seems to represent the resting activity of an animal resembling Parvancorina or Skania at the sediment surface. From the holotype of N. planum, it can be inferred that the tracemaker alighted at the sediment surface, excavated a shallow subhorizontal burrow, and later departed the excavation. Slight distortion of the grooves in the midsection of the fossil

suggest that the tracemaker progressed forward a short distance after touching down on the substrate. It is not certain whether any feeding activity was involved. There is no evidence of the capture or manipulation of infaunal prey as has been inferred in some

Paleozoic Rusophycus traces (e.g., Jensen, 1990; Brandt et al., 1995; Babcock, 2003;

English and Babcock, 2007). The tracemaker of Nevadichnos may have been a sediment- deposit feeder, a mat-grazing herbivore, or an animal capable of absorbing nutrients from the substrate directly across the body wall. In any of these scenarios, clear evidence of feeding would not be expected.

Etymology. From Nevada, the state where the first specimens were collected; and ichnos, from Greek, trace or track.

Nevadichnos planum, new ichnospecies

Figures 10A, 10B

Diagnosis: As for the ichnogenus.

Occurrence. Siltstone layers of the Montezuma Member of the Deep Spring Formation, adjacent to Montezuma Peak, Montezuma Range, Esmeralda County, Nevada. Specimens were collected from float in the middle portion of the Montezuma Member, which is considered to be of late Ediacaran age.

Etymology. From Latin, planum, flat.

Ichnogenus Bergaueria Prantil, 1946

Bergaueria ichnosp.

Figures 10C,10 D.

Material. One collected from the Montezuma Range.

Description. A sac-shaped burrow with central depression. Burrow is approximately 3 cm in diameter, and shows a distinct four-ray structure on the lower surface.

Discussion. Bergaueria is generally interpreted as a resting or dwelling burrow of actinarian cnidarians (Frey and Pemberton, 1985) or other infaunally anchored animals that can produce sac-like depressions on the muddy substrate.

Ichnogenus Planolites Nicholson, 1873

Planolites ichnosp.

Figures 11A-11C, 11E

Material. 16 specimens identified from four brown siltstone slabs.

Descripton. Simple cylindrical, elongate burrows, generally with smooth walls; some specimens have weak transverse linings. Burrows are 0.15-0.3 cm wide, 2-3 cm long, and typically less than 2 mm deep. Composition of the burrow fill is different from the lithology of the host rock. Limonite staining is common.

Discussion. Planolites is interpreted as a feeding trace, with sediment having been reworked through a worm‘s alimentary canal (Osgood, 1970; Häntzschel, 1975).

Biologically reworked sediment forms transverse annulations that are typically present on the walls of Planolites. The burrow fill is lithologically different from the host rock.

Limonite staining is interpreted as a taphonomic feature of Planolites. When an infaunal organism constructs a feeding tunnel under the sediment, organic materials such as feces and mucous are secreted or excreted as the organism passes through. Organic matter of these types can initiate anaerobic decomposition, and this in turn can result in precipitation of CaCO3 or FeS, which can transform to limonite (Sagemann et al., 1999).

4.3. Sedimentary Structures from the Deep Spring Formation

In this section are described sedimentary features of questionable origin or interpreted to be of inorganic origin.

Sedimentary structure 1

Figure 11D.

Discussion. Four slabs containing small concave circles, or dimples, were collected from the Montezuma Member of the Deep Spring Formation at Mt. Dunfee. Dimples commonly occur in clusters. Each individual ―dimple‖ is 2-3 mm in diameter, less than 1 mm deep at its center, has a slightly raised rim, and penetrates sediment perpendicular to bedding. Cross-sectioning revealed little or no change in lithology inside or outside the dimples.

It is possible to interpret this structure as a trace fossil such as a shallow burrow penetrating bedding from above. However, an interpretation as an inorganically formed structures is also plausible. The structure may be a tool mark formed by rolling or scouring of a small, current entrained object.

Sedimentary structure 2

Figures 12A, 12 B

Discussion. This structure is composed of a central raised mound with subparallel, linear striations radiating from the central mound. Each set of striations is composed of 4-6 shallow ridges that are 2-2.3 cm in length. In some cases, sets of striations intersect other sets. The structure is common in the Montezuma Member of the Deep Spring Formation at Montezuma Range. The origin of the structure is unknown. There is no known modern analogue.

Sedimentary structure 3

Figures 16C, 16D

Discussion. Sedimentary structures resembling ―elephant skin‖ are common in the

Middle Member of the Deep Spring Formation. The structures are composed of small scale irregular crests and troughs on bedding planes. Some specimens co-occur with

Planolites. These structures are interpreted as wrinkled microbial mat-stabilized sediment surfaces.

Figure 8. Locality map and outcrop photos of the Deep Spring Formation. A, The Deep Spring Formation in White Mountains, CA. B,The Montezuma Member in Montezuma40 Range, NV. C, The Deep Spring Formation in Mt. Dunfee, NV (map modified from http://maps.google.com).

Figure 9. Stratigraphic column and carbon isotope curve for the Deep Spring Formation (data from Albers and Stewart, 1972; Corsetti and Kaufman, 1994 ).

Figure 10. Trace fossils from the Montezuma Member of the Deep Spring Formation.. A, Nevadichnos. B, outline of Nevadichnos, a trace fossil reflecting an animal that dug into a muddy substrate, followed by departure. C, Bergaueria. D, outline of Bergaueria, a sac-shaped burrow with a distinct four-ray structure on the lower surface. Scale bars = 1 cm.

Figure 11. Trace fossils and sedimentary structures from the Montezuma Member of the Deep Spring Formation. A, C, Planolites. B, E, cross section of Planolites. E shows limonite staining along the burrow wall. D, sedimentary structure 1. Arrows indicate ―dimples‖ with raised rims. Scale bars = 1 cm.

Figure 12. Sedimentary structures from the Montezuma Member of the Deep Spring Formation. A, B, a problematic structure with a central mound and subparallel ridges radiating from the mound. C, D, microbial mat structure. Scale bars =1 cm.

CHAPTER 5

CAMBRIAN BODY FOSSILS AND TRACE FOSSILS FROM THE DEEP SPRING,

CAMPITO, POLETA, AND HARKLESS FORMATIONS

Cambrian strata studied in Esmeralda County, Nevada, and Inyo County,

California, include, in ascending order, the Montezuma Member (part) and Gold Point

Member of the Deep Spring Formation, and the Campito, Poleta, and Harkless formations (Figure 13). Cambrian strata studied range from the base of the Cambrian

(Montezuma Member of the Deep Spring Formation), to Cambrian Stage 3 (Harkless

Formation).

Figure 13. Locality map and outcrop photos of the Cambrian strata in the southern Great Basin. A, The Poleta Formation in White Mountains, California. B, Quartzite of the Campito Formation in Indian Springs Canyon, Nevada. C, Archaocyaths of the Poleta Formation, Nevada. D, The Harkless Formation, in Montezuma Peak, Nevada (map modified from http://maps.google.com).

Deep Spring Formation. (A), The Deep Spring Formation in White Mountains, CA. (B),The Montezuma Member in Montezuma Range, NV. (C), The Deep Spring Formation in Mt. Dunfee, NV (map modified from http://maps.google.com) 46

5.1. Cambrian paleontology of the Deep Spring Formation

Cambrian parts of the Deep Spring Formation in the southern Great Basin include the upper part of the Montezuma Member and the Gold Point Member. The Montezuma

Member in the White-Inyo Mountains, California, records the Ediacaran-Cambrian transition (Corsetti and Kaufman, 1994; Corsetti and Hagadorn, 2003). Precise correlation to Mt. Dunfee, however, is uncertain, in part because of structural complexity in the region. For this reason, only the Gold Point Member is included in this chapter.

Small fault-bounded outcrops of the Gold Point Member were identified in the northern part of Mt. Dunfee, Nevada. The Gold Point Member is characterized by gray to greenish siltstone and quartzite with light colored carbonate layers capping the siltstone

(Albers and Stewart, 1972). The overall thickness of the member is about 104 m at Mt.

Dunfee (Albers and Stewart, 1972). In most localities, the strata have experienced tectonic deformation, contact metamorphism, and weathering. One bluish green siltstone slab containing burrows was collected. The siltstone outcrop exhibits truncated bedding incorporating sediment clasts, which were probably storm-transported.

5.1.1. Systematic paleontology: Cambrian trace fossils from the Deep Spring Formation

Burrow type 1

Figure 14

Material. Three burrows on one slab.

Description. Short segments are connected to form a long, sinuous, chain-like burrow.

Each burrow is 0.1-0.15 cm wide and 4-6 cm long. Each burrow is composed of 2-5 short segments, each approximately 1 cm long, and part of each segment overlaps laterally, resembling the branching pattern of Treptichnus.

Discussion. This burrow is not readily assignable to any ichnogenus. A significant portion of the illustrated burrow is obliterated due to burial diagenesis and low grade metamorphism. Two burrows have similar curvature that can be assembled to a long loop up to 16 cm, including a break between the burrows. The overall morphology of connected, overlapping segments is a pattern resembling Treptichnus. Unlike

Treptichnus, however, segments are loosely packed together, and the size of each segment varies significantly.

Figure 14. Burrow type 1 from the Gold Point Member of the Deep Spring Formation of Mt. Dunfee, Nevada.

5.2. Paleontology of the Campito Formation

The Campito Formation has significance as it yields the earliest trilobites known from Laurentia. Fritz (1972) and Hollingsworth (2005, 2006) established four Cambrian biozones within the Campito Formation and the overlying Poleta Formation of the

Montezuma Range, Nevada. In ascending order they are the Fritzaspis Zone (uppermost

Stage 2), the Fallotaspis Zone (Stage 3), the Nevadella Zone (Stage 3), and the Olenellus

Zone (Stage 7). The (Peng and Babcock, 2008), and the lower part has been identified in the Montezeuma Range.

The Campito Formation is divided into two members: the Andrews Mountain

Member (lower) and the Montenegro Member (upper). The Andrews Mountain Member is characterized mostly by massive, dark-colored, fine-grained quartzite of 760 to 884 m in the White-Inyo Mountains, and the member is thought to have a similar thickness in

Esmeralda County (Albers and Stewart, 1972). The member has of the polymerid trilobite

Fritzaspis (formerly referred to as cf. Repinaella) and obolellid brachiopods

(Hollinsgworth, 2005, 2006).

The Montenegro Member is dominated by dark greenish gray siltstone with a thickness of about 96 m in the Weepah Hills (Albers and Stewart, 1972). Trilobites belonging to the Fallotaspis and Nevadella groups were reported from the Montenegro

Member in the Montezuma Range, as well as from correlative strata at Silver Peak,

Nevada, and the White-Inyo Mountains, California (Fritz, 1995; Hollingsworth, 2005,

2006).

5.2.1. Systematic paleontology: body fossils from the Campito Formation

Family Holmidae Hupé, 1953b

Genus Esmeraldina

Esmeraldina rowei (Walcott, 1910)

Figures 15A, 15B, 15D

Material. Six cephala on three slabs.

Discussion. New cephala closely resemble previously published material of Esmeraldina rowei. In all specimens, cephalic width is more than double cephalic length. The best preserved specimen has broad curvature of the cephalon extending to a long genal spine.

The lateral border is wide, and the posterior margin has two small spines. The occipital ring is gently curved and bears one small medial node. Three nearly parallel lateral glabellar furrows are visible.

Most previously reported trilobite assemblages in the Montenegro Member of the

Campito Formation are from the Montezuma Range in Esmeralda County, Nevada.

Material discussed here are from the east side of Mt. Dunfee. Three slabs reveal a total of

six cephala of Esmeraldina rowei. In the slab containing four cephala, disarticulated thoracic segments are common. The range of Esmeraldina rowei begins close to the base of the Nevadella Biozone (Hollingsworth, 2005, 2006), indicating that the Montenegro

Member in the Mt. Dunfee area is assignable to the lower part of Cambrian Stage 3.

5.2.2. Systematic paleontology: trace fossils from the Campito Formation

Trace fossils are common in the Montenegro Member of the Campito Formation at Mt. Dunfee. The most abundant trace fossil is Palaeophycus. Cross sectional analysis on lithologic samples reveals that ichnofossils from the Montenegro Member are nearly horizontal. In places, bedding plane surfaces are considerably disturbed by burrows, scratches, and sedimentary structures, but this is unusual. Although there is variation in bioturbation intensity, the illustrated slabs from Mt. Dunfee are assignable to Ichnofabric

Index 2 ( Droser and Bottjer, 1993).

Ichnogenus Palaeophycus Hall, 1847

Palaeophycus ichnosp.

Figure 15C

Material. 18 burrows identified on four brown coarse siltstone beds.

Description. Cylindrical, horizontal to subhorizontal burrows. Burrows straight or slightly curved. Size range varies from 0.4 cm to 1.1 cm in width; some burrows cross over each other. Most burrow walls are smooth. Burrows are preserved in full relief, epirelief, and endogenic relief. Burrow fill is identical to the lithology of host rock.

Discussion. Palaeophycus has been considered a locomotive trace (Osgood, 1970;

Häntzschel, 1975), a dwelling trace (Frey and Pemberton, 1985), or a trace showing combined behavior of feeding and dwelling (Jensen, 1997). Burrow fill is lithologically identical to the host rock, the result of passive fill. Bedding around the burrow is moderately disrupted. The structure is considered a U-shaped open burrow often supported by physical or biological packing by benthic organisms (Osgood, 1970).

In some described occurrences of Palaeophycus, collapse of a burrow after burial left longitudinal linings on the burrow wall and a subcylindrical burrow profile. However,

Palaeophycus specimens collected from Mt. Dunfee do not have any of these features.

The cross section of these burrows is a cylindrical shape, and even Palaeophycus exposed on bedding surfaces has a cylindrical tunnel system. This might represent a permanent U- shaped tunnel with passive filling, rather than a temporary structure that was subjected to deformation, or this may be related to a specific tracemaker that lived during Cambrian

Age 2.

Figure 15. Body and trace fossils from the Montenegro Member of the Campito Formation. A, B, D, Cephala of Esmeraldina rowei. C, Palaeophycus. Scale bars = 1 cm.

5.3. Paleontology of the Poleta Formation

The Poleta Formation is distinctive for various lithologic types and fossils including archaeocyathids, polymerid trilobites, and other fossils (Albers and Stewart,

1972). The formation is divided into three members: the Lida, Indian Springs, and

Clayton members. The thickness of each member varies by locality. Near Gold Point,

Nevada, the thickness of the Lida Member is approximately 71 m, the thickness of the

Indian Springs Member is approximately 190 m, and the thickness of the Clayton

Member is approximately 33 m (Albers and Stewart, 1972). Lithology varies by locality.

However, but generally the Lida and Clayton members are dominated by limestone, and the Indian Springs Member is composed of siltstone with minor amounts of fine-grained quartzite and limestone. The transition from the Montenegro Member to the lower member of the Poleta Formation is apparently conformable (Albers and Stewart, 1972).

Archaeocyaths are abundant in limestone of the Lida and Clayton members. The

Indian Springs Member records a more diverse fauna known as the Indian Springs Biota

(Babcock et al., 2007; English and Babcock, in press). Trilobites indicative of the

Nevadella Zone and the succeeding Olenellus Zone are present in the Indian Springs

Member (Hollingsworth, 2005).

The Indian Springs Lagerstätte includes at least three horizons of exceptional fossil preservation in the Montezuma Range, Nevada. This preservation style is interpreted to be related to the sediment smothering of the sea floor associated with

recurring episodes of regression (English, 2007; Hollingsworth, 2005). Fossils reported from Indian Springs Canyon include helicoplacoids, polymerid trilobites of the Nevadella

Biozone, hyolithids, brachiopods, and trace fossils including coprolites (English and

Babcock, in press).

5.3.1. Systematic paleontology: body fossils from the Poleta Formation

Body fossils were collected from the Indian Springs Member of the Poleta

Formation at the Indian Springs Canyon locality in the Montezuma Range, Nevada.

Fossils collected include Helicoplacus gilberti, Nevadia weeksi, and obollelid brachiopods, which are among the most abundant at the site.

Phylum Arthropoda

Class Trilobita Walch, 1771

Family Olenellidae Vogdes, 1893

Subfamily Nevadinae Hupé, 1953

Genus Nevadia Walcott, 1910

Nevadia weeksi Walcott, 1910

Figures 16A, 16B.

Materials. Two cephala from Indian Springs Canyon.

Discussion. Two new cephala share characteristics of previously described examples of

Nevadia weeksi (Walcott, 1910; Whittington, 1989; Fritz, 1995). One complete but poorly preserved cephalon is 2.3 cm long and 5.6 cm wide. Another cephalon is fragmentary, but has a reconstructed length of 1.1 cm and a width of 2.6 cm. Genal spines are short. Three gently curved lateral glabellar furrows are connected across the axial midline (Figure 16A).

Phylum Brachiopoda

Family Obolellidae Walcott and Schuchert, 1908

Genus Obolella Billings, 1861

Obolella sp.

Figures 16B, 16C

Materials. Two specimens on two split-open slabs.

Discussion. Two specimens assigned to Obolella sp. are preserved as subcircular impressions and their corresponding impressions. One specimen is 1.5 cm wide and 1.4 cm long, and the other specimen is 2 cm wide and 1.7 cm long. Growth lines are well expressed in both specimens (Figure 16B).

Phylum Echinodermata

Class Helicoplacoidea Durham and Caster, 1963

Family Helicoplacidae Durham and Caster

Genus Helicoplacus Durham and Caster, 1963

Helicoplacus gilberti Durham & Caster, 1963

Figures 17A-17C

Material. Specimens on three red to reddish gray siltstone slabs.

Discussion. New specimens are of disarticulated helicoplacoid ossicles clumped in dense associations with rounded margins. These associations closely resemble specimens interpreted by English and Babcock (in press) as coprolites. Helicoplacus gilberti is the only helicoplacoid species known from the Indian Springs Lagerstätte, and the disarticulated ossicles are assumed to derive from this taxon.

Helicoplacoids are one of the earliest echinoderms in the fossil record, and most described specimens are from the Indian Springs Member of the Poleta Formation (e.g.,

Durham, 1993; Dornbos and Bottjer, 2000, 2001). Their distinct morphology is characterized by helically arranged columns composed of loosely articulated calcite plates, and three ambulacra extend from the oral pole to apical pole (Durham and Caster,

1966; Dornbos and Bottjer, 2001).

5.3.2. Systematic paleontology: trace fossils from the Poleta Formation

Trace fossils were collected from the Indian Springs Member of the Poleta Formation in the White-Inyo Mountains, California. Trace fossils collected include Rusophycus and

Palaeophycus.

Ichnogenus Rusophycus Hall, 1852

Rusophycus ichnosp.

Figures 18A, 18B

Materials. Six specimens on greenish gray slabs from the White-Inyo Mountains.

Description. Bilobed structure on sole side of bedding. Dimensions vary, ranging from

2.5 to 7 cm in length, 1.3 to 5 cm in width, and 1 to 2.2 cm in depth. Subparallel striations are commonly preserved, along with a median ridge. One specimen has elongate scratches at the lateral margins.

Discussion. Rusophycus has commonly been referred to as a resting trace of arthropods, especially trilobites (Seilacher, 1955; Osgood, 1970). It has a distinctive bilobate morphology, scratches made by appendages, and commonly elongate scratches at the

lateral margins reflecting the edge of the tracemaker‘s exoskeleton. Some examples of

Rusophycus co-occur with simple tubular trace fossils assignable to Palaeophycus. This common association has been interpreted to reflect predatory behavior by trilobites on infaunal worm-like prey (e.g., Jensen, 1990; Brandt et al., 1995; Babcock, 2003; English and Babcock, 2007).

Ichnogenus Palaeophycus Hall, 1847

Palaeophycus ichnosp.

Figure 18D

Material. Three specimens on greenish gray siltstones.

Description. Several simple, horizontal to subhorizontal, cylindrical burrows cover slabs.

Burrow width ranges from 0.15-0.3 cm. Burrows are straight or slightly curved. Burrow fill is identical to the host rock lithology. Burrow walls smooth. Burrows do not branch, and intersecting of burrows is common.

Discussion. Palaeophycus is the most abundant type of tubular burrows in the Indian

Springs Member of the Poleta Formation in the White-Inyo Mountains. A slab in the field revealed Palaeophycus on top of ripple marks (Figure 18D). On rippled bedding surfaces,

Palaeophycus has lengths up to 15 cm.

Burrow type A

Figure 18C

Material. One slab collected from the Indian Springs Member of the Poleta Formation from the White-Inyo Mountains.

Description. Tightly meandering burrows with uniform width of 0.4-0.5 cm. Complete burrows are 3.5-4 cm long. Burrows are horizontal and cylindrical. Transverse annulations are present. Burrow fill is identical to the host rock lithology. Burrows commonly intersect each other.

Discussion. This burrow, left in open nomenclature, is different from Planolites and

Palaeophycus, the most common simple burrows in the Cambrian successions of

California and Nevada. The basis for identifying Planolites and Palaeophycus is lithological composition of the burrow fill, presence or absence of transverse annulations, and signs of passive or active fill. This undetermined burrow has characteristics of either

Planolites and Palaeophycus such as transverse annulations, and burrow fill composition that is identical to that of the host rock. The mode of burrowing is different from either

Planolites and Palaeophycus, however. Planolites and Palaeophycus are typically

straight to slightly sinuous, and burrow dimension often varies even in the same slab.

Tight meandering of this burrow and a uniform width of this burrow is distinctive. Tight meanders and a uniform width of the burrow presumably reflects the feeding or grazing behavior of worm-like benthic organisms.

Figure 16. Body fossils from the Indian Springs Member of the Poleta Formation, Indian Springs Canyon, Nevada. A, B, Cephalon of Nevadia weeksi. C, D, Obollelid brachiopods preserved in impressions and corresponding counterparts. E, F, Archaeocyaths displaying top view (E) and lateral view (F).

Figure 17. Helicoplacus gilberti ossicles in coprolites from the Indian Springs Member of the Poleta Formation in Indian Springs Canyon, Nevada.

Figure 18. Trace fossils from the Indian Springs Member of the Poleta Formation, White-Inyo Mountains. A, B, Rusophycus. C, Poleta Burrow type A. D, Palaeophycus on rippled bedding surfaces.

5. 4. Paleontology of the Harkless Formation

The Harkless Formation in Esmeralda County, Nevada, is composed of two members. The Weepah Member occupies the lower third of the formation and is mainly composed of quartzitic siltstone. The Alkali Member occupies the rest of the formation and is composed of greenish siltstone (Albers and Stewart, 1972). The formation yields olenellid trilobites, indicating an age of Cambrian Age 4 (Epoch 2). Palmer (1964) identified three groups of fossils from the Harkless Formation. Group A, which includes

Olenellus and Fremontia, was reported from a position 420 m above the base of the formation (in the Weepah Member) at the Weepah Hills section. Group B and C fossil groups are from the Alkali Member. Group B includes Bonnia carperata, Ogygopsis batis, Paedeumias granulata, Wanneria sp., and olenellid trilobites. Group C is composed mainly of Paedeumias nevadensis (Palmer, 1965; Albers and Stewart, 1972).

The lithology of the Harkless Formation is similar to that of the Campito

Formation. Quartzitic siltstone of the Weepah Member is similar to that of the Andrews

Mountain Member of the Campito Formation. The greenish siltstone of the Montenegro

Member is similar to the siltstone of the Montenegro Member of the Campito Formation

(Albers and Stewart, 1972). Due to the complex structure of the region, it can be difficult to distinguish the Harkless Formation from the Montenegro Member on lithologic grounds alone. Biostratigraphic evidence, however, can be used to distinguish the two units. The Harkless Formation contains olenellid trilobites, whereas the Montenegro

Member contains fallotaspid and nevadellid trilobites, which predate the olenellid trilobites.

5.4.1. Trace fossils from the Harkless Formation

Trace fossils from the Alkali Member of the Harkless Formation at Montezuma

Peak are mostly shallow and horizontal forms representing Cambrian Series 2 benthic faunas. Although bedding surfaces are commonly disturbed by one or more ichnogenera and sedimentary structures such as tool marks, vertical profiles of studied slabs display low degrees of bioturbation (less than 10% of original bedding disturbed; Ichnofabric

Index 2 of Droser and Bottjer, 1993). Ichnogenera reported include Planolites,

Palaeophycus, Monomorphichnus, Bergaueria, Treptichnus pedum, short, unbranched burrows assigned as treptichnids, and trails and tracks possibly attributable to arthropods

(Figure19). Simple burrows such as Planolites and Palaeophycus do not have biostratigraphic significance because of their simple morphologies, wide ranges of ecologic occurrence, and long stratigraphic ranges.

5.4.1.1. Planolites-Palaeophycus Problem

The ichnogenus Palaeophycus was first described by Hall (1847) for cylindrical or subcylindrical infilled burrows of simple tubular structure. The similar ichnogenus

Planolites was first described by Nicolson (1873) and later interpreted as a feeding trace with sediment reworked through a worm‘s alimentary canal (Osgood, 1970; Häntzschel,

1975).

Although they are presumed to have origins with different behaviors, Planolites and Palaeophycus share several morphologic characteristics. Both ichnogenera consist of simple tubular, sinuous infilled burrows. Both have horizontal orientations relative to bedding surfaces, and commonly have crossover of burrows. Both have long stratigraphic ranges, and both are reported from a wide range of aquatic environments. How exactly to characterize and distinguish Planolites and Palaeophycus has been debated for several decades (e.g., Osgood, 1970; Alpert, 1975; Pemberton and Frey, 1982; Keighley and

Pickerlill, 1995; Jensen, 1997; Marenko and Bottjer, 2008). The major criteria distinguishing the two ichnogenera are still debated. However, the two burrows are commonly distinguishable (Osgood, 1970; Pamberton and Frey, 1982, Jensen, 1997) based on differences in burrow wall linings and composition of burrow fill that is formed either by biogenic reworking (Planolites) or by passive sediment infill by physical means

(Palaeophycus). Palaeophycus generally has burrow linings, and its passive fill represents physical incorporation of sediment around the burrow into an open tunnel.

Thus, the sediment filling the burrow is identical to the host rock sediment. On the other hand, Planolites commonly has transverse annulations as a result of backfilling, and its burrow fill is different from the host matrix due to the biogenic reworking of sediment

(Osgood, 1970; Häntzschel, 1975; Pemberton and Frey, 1982; Jensen, 1997). Although

many specimens can be readily assigned to either Planolites or Palaeophycus based on these criteria, an unambiguous distinction between the two ichnogenera cannot be made solely on burrow morphology and burrow fill, because Planolites and Palaeophycus are morphologically similar, and in some cases, burrows have morphologic varieties that do not fulfill the taxonomic criteria of either ichnogenus.

Planolites and Palaeophycus specimens were studied both through bedding surface examination and through cross sections. Tubular trace fossils were cut and polished to reveal the composition of burrow filling sediment. If burrow fills have signs of active fill or reworking, with other morphologic characters (e.g., transverse annulation, crossover of burrows) and diagenetic features (authigenic minerals and sedimentary structures) of feeding behavior, then a burrow was assigned to Planolites. If a simple, horizontal burrow has wall linings formed by collapse of the tunnel and when the burrow fill is the same as the host rock, then the burrow was assigned to Palaeophycus.

5.4.1.2. Systematic paleontology: trace fossils from the Harkless Formation

Ichnogenus Planolites Nicholson, 1873

Planolites ichnosp.

Figures 23A-23E

Material. 138 siltstone slabs were collected at Montezuma Peak; 630 unquestionable

Planolites were identified.

Description. Simple, horizontal tubular trace fossil. Burrows are straight to slightly sinuous. Burrow width ranges from 0.05 to 1 cm, with peak abundance being 0.2-0.3 cm

(Figure 20). In complete specimens, burrow length ranges from 4 to 8 cm. Burrow walls have two types of texture, smooth and regularly spaced, short transverse annulations.

Crossover of burrows is common. Most Planolites analyzed are associated with pyrite, chalcopyrite, and limonite.

Discussion. Planolites is the most abundant ichnofossil collected from the Alkali Member of the Harkless Formation at Montezuma Peak (63% of trace fossils). Planolites is an open burrow constructed as a result of feeding behavior (Osgood, 1970; Häntzschel,

1975; Droser et al., 2000). The lithologies of slabs containing Planolites are mostly olive- green siltstones within which are dark bluish gray coarse silt layers. Burrows with full relief and endorelief were selected, cut, and polished to reveal the lithology of burrow fills. Burrow fills are mostly dark bluish gray siltstone.

Transverse annulations of Planolites may represent packing structures formed by biogenic backfilling of sediment, thus making burrow fill different from the host matrix

(Figures 23C, 23D). In many cases, Planolites co-occurs with tool marks or tracks and trails of inferred arthropods, suggesting that the substrate at the time of deposition was

firm enough to record surface traces. A wide range of burrow widths reflects burrowers of different sizes. In some specimens, two or more narrow burrows are stacked on each other to form a wide and thick burrow that can be traced on the bedding surfaces (Figure

23D). There are varied preservational varieties of Planolites. Burrows are commonly preserved as exogenic and endogenic traces on bedding planes. Full exogenic relief reveals spatial relations of Planolites three dimensionally. This relief commonly produces impressions in underlying or overlying sediment layers. In many cases, the impressions in underlying layers have morphologic details of the burrow close to the point of burrow origin. Frequent crossover of the burrows is easily identifiable, especially when the traces are preserved in full relief (Figures 23A, 23E). None of the identified Planolites were branching. However, crossovers do resemble branching of burrows, especially when they are preserved as external molds.

Another interesting preservational feature of these traces is two-dimensional impressions of simple, short tubes without any wall linings or relief. The mechanism of this preservation style might be similar to carbonization. However, the color of the burrow film is light to white, indicating a lack of carbon films (possibly due to diagenesis).

The simple morphology of burrows in two-dimensional films makes classification of burrows difficult. However, in some specimens, reddish staining is visible along the edges of burrow impression. Considering that limonite coatings along Planolites are

common, this reddish lining might be a result of compression of Planolites during early stages of diagenesis.

Planolites from Montezuma Peak have signs of early diagenesis that are quite likely related to the development of biofilms. Almost 80% of Planolites have concentrations of chalcopyrite, pyrite, limonite, or mica (Figures 21, 23E, 23F).

Chalcopyrite and limonite are most common on bedding surfaces adjacent to Planolites burrows. Limonite is easily identified by its reddish brown staining, and chalcopyrite is commonly in the form of large, euhedral crystals. The distribution of these two minerals might be related to early mineralization in biofilms associated with decomposition of organic materials (e.g., mucous or feces) secreted by burrowers that produced Planolites.

Muscovite commonly occurs in association with Planolites burrows. Mica abundance probably is not related to biofim formation. Its occurrence, more likely, represents original siliciclastic sediment that was incorporated into open tunnels by currents at the time of deposition. This is supported by the presence of thin mica-rich layers intercalated in olive-gray siltstone layers at Montezuma Peak.

Ichnogenus Palaeophycus Hall, 1847

Palaeophycus ichnosp.

Figures 24A, 24B

Material. A total of 27 individual burrows were identified from 14 slabs.

Description. Simple, straight to slightly sinuous tubular trace fossil. Burrows are oriented horizontal to slightly obliquely to bedding surface. Burrows range from 0.2 to 0.7 cm in width, with peak abundance being 0.4-0.5 cm. Most specimens are fragmentary, but in the most complete specimen, the burrow is 6 cm long. Cross sections of burrows display cylindrical to subcylindrical profiles. Lithologic composition of burrow fill is similar to that of host rock (Figure 24B). Burrow have walls of two types of texture, smooth walls, and longitudinal linings along the long axis of burrows (Figure 24A). Burrows of larger dimensions do not have crossovers, whereas burrows narrower than 0.3 cm commonly intersect each other. No true branching has been detected. No significant relationship between specific minerals and Palaeophycus was identified. Burrows are preserved in epirelief and exorelief.

Discussion. Paleophycus was originally interpreted as a locomotory trace (Häntzschel,

1975), representing unorganized pathways of infaunal organisms that were passing through sediment in search of food (Osgood, 1970). The trace was later interpreted to represent a dwelling trace (Frey and Pemberton, 1985) or a trace that combined feeding and dwelling (Jensen, 1997). Direct evidence of feeding in Paleophycus is somewhat vague as there is no sign of biogenic reworking in Paleophycus. However, some deposit feeders feed and defecate in different places (Jensen, 1997). Therefore, if Paleophycus

and Planolites are both feeding structures, then the difference between the two ichnogenera may be related to dissimilar feeding-defecation habits of burrowers or burrowers of different ecologic niches.

Paleophycus can be viewed as an open burrow at least temporarily supported by organic mucus secreted by benthic organisms; and while open, physical sedimentation was possible (Osgood, 1970). This passive fill would have involved sediment deposition in an open tunnel, which explains the similar lithologic composition between

Paleophycus and its host rock. The ichnofossil commonly has longitudinal linings along the burrow wall. Cross sections of Paleophycus with these linings are mostly subcylindrical, indicating that this wall structure might be due to compaction of sediment

(Häntzschel, 1975). Paleophycus co-occurs with muscovite mica. As the sediment of the

Harkless Formation is mica-rich in places, this may simply represent sediment that entered the burrow at the time of deposition.

Ichnogenus Treptichnus Miller, 1889

Treptichnus ichnosp.

Figure 24F

Material. 6 specimens identified on two slabs.

Description. A series of short, straight burrows arranged such that small segments alternate at angles to each other, forming a zigzag pattern. Segments forming Treptichnus have even dimensions within a linear pattern. Individual burrows are 0.1-0.15 cm wide,

0.4-0.8 cm long; portions of each segment 0.1-0.2 cm long intersect each other to form a joint. Short jointed segments produce a long curved chain having a length more than 8 cm. Burrow wall is smooth, and burrow fill is identical to the host rock. Specimens are either in micaceous quartz-rich siltstone or micaceous siltstone. Specimens are preserved as hyporelief.

Discussion. Treptichnus is interpreted as a feeding structure (Seilacher and Hemleben,

1966) with a ―feather-stitch‖ pattern (Figures 22C, 24E, 24F) formed by alternating movement of burrowers to the right and left and tending upward and forward (Wilson,

1948; Häntzschel, 1975).

Treptichnus pedum (Seilacher, 1955)

Figures 24E, 24G.

Material. Two slabs.

Description. Short, slightly curved burrows alternating with each other at a low angle to form a braided pattern of short segments. Each segment is about 0.1 cm wide and 0.5 cm

long. One specimen has ―buds along a twig‖ pattern (see Seilacher, 2007). Each segment is 0.8 - 1.2 cm in length and 0.15-0.2 cm in width.

Discussion. Treptichnus pedum is interpreted as a U-shaped feeding structure expanding from one end by addition of segments (Seilacher, 1955; Jensen, 1997). This complex burrow system results in braided or budding patterns, but the trace itself does not branch

(Seilacher, 2007). The overall course of the burrow can be straight, coiled, or sinuous

(Seilacher, 2007).

Treptichnus pedum was originally described in the ichnogenus Phycodes

(Seilacher, 1955). Phycodes is a bundled structure of horizontal tunnels on the sole side of bedding, typically with a flabellate pattern (Osgood, 1970, Häntzschel, 1975).

However, pedum is noticeably different in morphology from typical Phycodes, as it has burrows arranged in a ―feather-stitch trail‖ pattern. For this reason, Osgood (1970) proposed that P. pedum should be removed from Phycodes. Jensen and Grant (1992) assigned P. pedum to Treptichnus. Geyer and Uchman (1995) subsequently reassigned this ichnospecies to Trichophycus, based on criteria of the burrow system and inferences about sediment reworking in Treptichnus.

Figure 22 shows a comparison of the morphology of Phycodes, Treptichnus, and

Trichophycus. Treptichnus and Trichophycus are U-shaped tubes, whereas Phycodes is not (Figure 22A). Trichophycus is characterized by large, cylindrical burrows with fine striations on the ventral surface of the burrows (Osgood, 1970, Häntzschel, 1975) (Figure

22B). Treptichnus is similar in this respect (Figure 22C), but the order in which short segments are added in Treptichnus is opposite that in which they are added in

Trichophycus. In the ichnospecies pedum, the order of segment addition appears to match

Treptichnus.

In trace fossils lacking distinct morphologic characters, a transition from one ichnogenus to another ichnogenus is possible. For example, it may be possible to have a transition from Trichophycus to Phycodes when the burrow lacks spreiten (see

Häntzschel, 1975). In addition, apparent morphology can vary when traces are subjected to varying depositional and diagenetic conditions. My preference for referring pedum to

Treptichnus follows the reasoning of Jensen and Grant (1992).

The first appearance of Treptichnus pedum is used as a guide fossil coinciding with the base of the Cambrian System (Crimes, 1987, 1992; Narbonne et al., 1987,

Landing, 1994; Landing et al. 2007; Gehling, 2001, Jensen, 2003; Corsetti and Hagadorn,

2003). Crimes (1987, 1992) documented three ichnofossil zones across the

Neoproterozoic-Cambrian transition. This zonation has been modified by other authors

(e.g., MacNaughton and Narbonne, 1999; Jensen, 2003). However, the general pattern is one of simple, horizontal burrows in the lower zone (Ediacaran System); more complex burrows including T. pedum in the second zone (lowermost Cambrian; now Terreneuvian

Series); and more diverse, branching burrows and inferred arthropod-related traces in the uppermost zone (Crimes, 1987, 1992, Jensen, 2003), which corresponds to the

Terreneuvian Series-Series 2 interval (Cambrian).

Recent restudy of the Cambrian GSSP section at Fortune Head, Newfoundland,

Canada, shows that the first appearance datum (FAD) of T. pedum actually underlies the

GSSP point by at least 4.5 m (Gehling et al., 2001). Regardless, the position of the GSSP and the FAD of T. pedum can be considered nearly coincident; and therefore, this guide fossil still has significance for determining the position of lowermost Cambrian strata

(Gehling, 2001; Landing et al., 2007; Peng and Babcock, 2008).

Ichnogenus Bergaueria Prantl, 1946

Bergaueria ichnosp.

Figures 24C, 24D

Material. One split-open slab yielding counterpart impressions.

Description. Cylindrical protrusion with a low central depression showing concentric markings. Burrow is 1.5 cm in diameter and 0.1 cm in depth.

Discussion. Bergaueria is generally interpreted as a resting or dwelling burrow of actinarian cnidarians or sea anemones (Alpert, 1973; Frey and Pemberton, 1985; Zhu,

1997). Actinarians can form bag-like depressions on muddy substrates, and subsequent sediment filling of the depression left after departure of organisms produce mound-

shaped casts of Bergaueria on the undersides of beds. Morphologic characters of the illustrated specimen differ slightly from features typical of the ichnogenus. The Harkless specimen is a split-open slab in which the trace is preserved as a pair of complementary impressions (Figure 24D). The specimen is nearly flat, with maximum relief of 0.1 cm at the central depression (Figure 24C). Although nearly flat, the specimen has concentric markings and a central depression, which are characteristic of Bergaueria. Preservation in a mud (shale) matrix may account for compression of the trace from its original more convex shape.

Ichnogenus Monomorphichnus Crimes, 1970

Monomorphichnus ichnosp.

Figure 25A.

Material. One specimen.

Description. Series of long, shallow ridges on a bedding plane. Illustrated specimen has laterally repeating sets of ridges, with 12 subparallel shallow ridges on one side and 5 faint ridges on the other side of the slab. Sets of ridges are about 2.5 cm apart. In each set, a single ridge is 1.5-1.8 cm long, with a regular spacing of 0.2-0.5 cm between each ridge.

Discussion. This specimen looks to be either a trace fossil or a tool mark. There are three possible ways to interpret this specimen: 1, a Monomorphichnus as a scratch mark formed by a swimming or current-entrained trilobite (Crimes, 1970; Osgood, 1970); 2, as an undertrace of a trilobite trackway such as Dimorphichnus (Seilacher 1955); or 3, as a tool mark. Preference for assigning specimens from the Montezuma Range to

Monomorphichnus, a trace representing swimming or locomotive behavior of trilobites, is based on lateral repetition of shallow ridges and lack of blunt impressions.

Dimorphichnus is a trilobite trackway characterized by sigmodial impressions and blunt markings representing trilobite legs supporting its body from currents (Seilacher,

1955; Häntzschel, 1975). The trace looks similar to Monomorphichnus except that there should be blunt impressions adjacent to long ridges in Dimorphichnus. Harkless specimens lack these features. Five short markings intersecting the ridges of one set are present. The same markings are not observed in the repeated set, and their irregular positioning along ridges implies that these makings could be another trackway, perhaps made by an arthropod, overprinted on Monomorphichnus.

Regularly spaced subparallel ridges do not support an interpretation of the specimen as a tool mark, because tool marks typically do not show such systematically arranged structures.

5.4.2. Sedimentary Structures from the Harkless Formation

A variety of sedimentary structures and biogenic features are present in siltstone layers of the Alkali Member of the Harkless Formation in the Montezuma Range. Among them are current-related sedimentary structures. Analyses of cross-sectioned slabs have various signs of bioturbation, including mottled or disturbed bedding. Other common structures collected from the Montezuma Range are features inferred to have been formed through microbial activity. Although their origins are related to a combination of biologic and physical processes, here they are treated as sedimentary structures.

Some trace fossils such as Dimorphichnus and Monomorphichnus are commonly viewed as the result of animals being carried along in currents, occasionally dragging their appendages along the substrate. Tool marks are inorganic sedimentary structures generated by scouring of the substrate by floating sedimentary particles. When the floating objects are carcass of arthropods, the distinction between tool marks and trace fossils becomes challenging. Two main types of tool marks are discussed here, shallow ridges resembling Monomorphichnus, and Y-shaped structures.

Sedimentary structure 1

Figure 25B

Discussion. Modified or disturbed bedding exhibiting Ichnofabric Index 2 (less than 10% of original bedding disturbed; see Droser and Bottjer, 1993) is identifiable in cross- sectional view of some Harkless siltstones. In many slabs, bioturbation is concentrated on

bedding near the top surface. This is not surprising when considering horizontal to subhorizontal burrows comprise almost 80% of the ichnofossil assemblage of the

Harkless Formation. A disturbed bedding profile has endogenic burrows, mottled structures, displacement and fragmentation of original bedding into blocky pieces, and missing laminations due to reworking of sediment by the activity of infaunal organisms.

Sedimentary structure 2

Figure 25C

Discussion. Y-shaped structures are characterized by one gently curved ridge with a straight shorter ridge diverging from the longer, curved ridge, forming a Y with long arms. Three specimens were identified from two slabs. In one slab, one Y-shaped features overlaps a similar one, suggesting that objects, such as arthropod limbs, were dragged on the soft substrate by currents.

Sedimentary structure 3

Figure 25D

Discussion. Several slabs collected from Montezuma Peak contain numerous thin, elongate scour marks on the sediment surface. Shallow ridges are typically 0.5-1.5 cm

long and less than 0.1 cm wide. The tip of each ridge is typically tapering. No regular spacing between lines, or repetition of systematically arranged features, was recognized.

Apart from arthropods, there are few identifiable candidates for tracemakers of these markings known from marine environments of the early half of the Cambrian.

These inferred tool marks typically co-occur with tubular burrows or arthropod related traces, in the form of casts. The orientation of the ridges is random even in the same slab, possibly reflecting changing current direction.

Sedimentary structure 4

Figures 26A, 26C

Discussion. Three slabs containing ―wrinkle structures‖ (Hagadorn and Bottjer, 1997,

1999; Noffke et al., 2002) are observed. The structures have small scale irregular crests and troughs on the bedding plain that are considered to be surfaces of microbial mats

(Figure 26-A). Two slabs containing wrinkle structures are slightly metamorphosed.

Small mica grains are commonly spread over the slabs. It is conceivable that these were mica grains trapped syndepositionally on an adhesive mat surface at the time of deposition. Alternatively, the mica was precipitated during diagenesis or low grade contact metamorphism. The presence of wrinkle structures indicates that microbially stabilized sediment surfaces persisted at least until Cambrian Epoch 2 on the Laurentian shelf. Wrinkle structures are rarely on the same bedding planes as body or trace fossils.

Scanning electron microscope (SEM) analysis revealed spherical microstructures encrusted on winkled structures. The origin of the microstructure is interpreted as aggregation of lithified bacteria and minerals (figure 26C).

Sedimentary structure 5

Figures 26B, 26D.

Discussion. Concentrations of tiny bumps about 0.1 cm in diameter are common on feeding traces, especially on or near Planolites. These features are interpreted as gas escape structures. They may be analogous to the gas escape structure referred to as ―fairy ring structures,‖ which result from the decay of microbial mat (Noffke, 2000; Noffke et al., 2001). Gas escape structures are thought to have originated in a different way from wrinkle structures. Wrinkle structures were probably thick, semi-lithified to lithified microbial mats that covered the sea floor. Gas escape structures, however, are thought to reflect the escape of gas produced through decomposition of organic matter, and they are not necessarily linked to well developed microbial mats. Distribution of the structures along Planolites tubes partly supports this hypothesis. In this instance, the structures may have formed in a way similar to initial stages of concretion formation (Borkow and

Babcock, 2003): organic materials secreted by infaunal organisms initiated bacterial decomposition in an oxygen poor environment a couple of centimeters below the sediment surface (Allison, 1988; Briggs et al., 1996; Sagemann et al., 1999), then gas

produced by the decomposition bubbled up through sediment layer above the feeding traces, forming small bumps around Planolites.

SEM analyses of gas escape structures reveal small globular structure of 0.5-1 μm within the bumps (Figure 26D). These small spherical bodies co-occur with lod shaped bacterial filaments (Figure 26D, arrow). It is not clear what kind of microorganism does the structure represent, however, considering its small size and co-occurrence with chalcopyrite and limonite, the spheres might represent sulfate-reducing bacteria

(Sagemann et al., 1999), and chalcopyrite and limonite occurring with gas escape structures and Planolites could be a result of biomineralization of sulfate-reducing bacteria. Experiments on biomineralization of soft tissue show that anaerobic decomposition enhances mineral formation (CaCO3, Ca5(PO4,)3 and FeS2; and all these processes take place in a short period of time, within a few weeks of burial rather than on a geological time scale (Sagemann et al., 1999). Gas escape structures and high concentrations of chalcopyrite and limonite along Planolites suggest that Planolites burrows were mineralized within a few weeks of construction.

Planolites 138 misc. 27 Paleophycus 16 arthropod tracks and trails 15 Monomorphichnus 3 treptichnids 3 Bergaueria 2 slab total 209

Figure 19. Trace fossil composition of the Alkali Member of the Harkless Formation. Raw numbers of collected traces are reported in the table at left; and percentages are reported in the pie diagram at right. Note that 63% of the traces are simple, short tubular traces assigned to Planolite.

range number of (cm) burrows <0.1 8 0.1-1.2 133 0.2-0.3 197 0.3-0.4 127 0.4-0.5 72 0.5-0.6 40 0.6-0.7 13 0.7-0.8 27 0.8-0.9 6 0.9-1.0 0 >1 7 total 630

Figure 20. Distribution of widths of Planolites burrows from the Alkali Member of the Harkless Formation. Burrow width of 0.2 to 0.3 cm is the most abundant size.

Figure 21. Composition of minerals co-occurring with Planolites from the Alkali Member of the Harkless Formation. Limonite commonly co-occurs with chalcopyrite, implying possible source of limonite on the Planolites.

Figure 22. Schematic diagrams showing the morphologies of Phycodes, Trichophycus, and Treptichnus. A, Bundled structure of Phycodes on the sole side of bedding. B, Burrow system of Trichophycus. C, ―Feather-stitch trail‖ pattern of Treptichnus (left) and Treptichnus pedum (right). However, compaction of a trace such as Trichophycus can produce similar structure to T. pedum in C. Diagrams modified from Häntzschel (1975) and Geyer and Uchman (1995).

Figure 23. Planolites from the Alkali Member of the Harkless Formation, Mt. Dunfee, Nevada. A, Intersecting burrows of Planolites. B, Burrow wall shows transverse annulations as a result of biologic reworking of sediment. C, Limonite encompassing the burrow fill. D, Two superimposed Planolites. E, Exposed burrow wall. Arrow indicates chalcopyrite along the trace. F, Scanning electron micrograph of chalcopyrite.

Figure 24. Trace fossils from the Alkali Member of the Harkless Formation, Mt. Dunfee, Nevada. A, Palaeophycus .B, Cross section of Palaeophycus. Arrow indicate burrow wall. C, Bergaueria showing central depression. D, Bergaueria from split-open slab. E, Treptichnus pedum showing ―feather-stitch trail‖ pattern. F, Treptichnus. G, Treptichnus pedum showing three segments (indicated by arrows) connected together in a ―buds along a twig‖ pattern. Scale bar = 1 cm.

Figure 25. The trace fossil Monomorphichnus and sedimentary structures from the Alkali Member of the Harkless Formation, Mt. Dunfee, Nevada. A, Monomorphichnus. Arrow indicates independent arthropod tracks that are not part of Monomorphichnus. B, Cross section of a slab containing burrows. Bedding exhibits Ichnofabric Index 2 (less than 10% of original bedding disturbed). C, Y- shaped structures. D, Thin, elongate scour marks on sediment surface. Scale bar = 1 cm.

Figure 26. Microbially produced structures from the Harkless Formation. A, C., Wrinkle structures formed by microbial mats. Spherical bodies in C are interpreted B, D, Gas escape structures formed by taphonomic processes within feeding traces. Arrow in D indicates bacterial filaments and spheres. Scale bar equals 1 cm for B, 50 µm for C, and 20 µm for D.

CHAPTER 6

CONCLUSIONS

The Ediacaran-Cambrian interval marks the transition from a microorganism- dominated evolutionary record to a more diverse and complex evolutionary record of the

Phanerozoic (Bengston, 1994; Conway Morris, 1989, 1993, 2000; Fedonkin, 1994, 2003;

Droser et al., 2002; Narbonne, 2005; Valentine, 2007). The transition is well recorded in the terminal Neoproterozoic to Cambrian strata in Esmeralda County, Nevada, and Inyo

County, California. Localities studied are in the Deep Spring Formation (Ediacaran-

Cambrian), and the Campito, Poleta, and Harkless formations (Cambrian). The base of the Cambrian System occurs in the Montezuma Member of the Deep Spring Formation.

Carbon isotopes from the Mountezuma Member have a large δ13C excursion interpreted as the BACE (Basal Cambrian Carbon isotope Excursion) (Corsetti and Kaufman, 1994;

Corsetti et al, 2000; Corsetti and Hagadorn, 2003), which has been used as the principal stratigraphic tool for correlating the base of the Cambrian globally (Zhu et al, 2006; Peng and Babcock, 2008; Peng et al., in press).

The Montezuma Member also contains Treptichnus pedum (Corsetti and

Hagadorn, 2003), a trace fossil whose first appearance nearly coincide with the base of the Cambrian (Landing, 1994; Gehling, 2001)

Trace fossils preserved in the Ediacaran part of the Deep Spring Formation

(Montezuma Member) include Bergaueria, Palaeophycus, Planolites, and other simple traces possibly representing primitive resting, and dwelling behaviors. Trace fossil diversity increases in Cambrian. Trace fossils from Cambrian strata include Bergaueria,

Cruziana, Monomorphichnus, Palaeophycus, Planolites, Rusophycus, Treptichnus pedum, and other traces. The greater diversity of traces in Cambrian strata signifies an expansion of behavioral activities paralleling an increase in taxic diversity through the early half of the Cambrian.

Microbially formed sedimentary structures are observed in both Ediacaran and

Cambrian strata (Hagadorn and Bottjer, 1997; Shieber, 1999; Noffke, 2002). Wrinkle marks on sediment surfaces are commonly observed in the Deep Spring Formation. This sedimentary structure is composed of a series of small scale crests and troughs, and is inferred to reflect microbial mat-stabilized sediment surfaces. Microbially stabilized substrates or microbial mat structures may have been partly responsible for the fine preservation of Ediacaran trace fossils (Wade, 1968; Seilacher, 1992; Gehling 1999).

Similar structures are also found in the Harkless Formation, which belongs to Cambrian

Stage 4. Scanning electron micrographs of wrinkle structures reveal spherical bodies mixed in mineral aggregations, confirming that the wrinkle structures were constructed by mat-forming microorganisms. Similar structures from the Deep Spring and Harkless formations indicate continued existence of Ediacaran-type mat-forming autotrophic microorganisms through the lower part of the Cambrian.

The Harkless Formation yields gas escape structures. The structures occur primarily along the feeding trace Planolites. Chalcopyrite and limonite are usually present associated with the structure. Scanning electron micrographs of gas escape structures reveal spherical and filamentous bodies 0.5-2 μm in size, which are interpreted as sulfate-reducing bacteria. They indicate that gas escape structures formed by the decomposition of organic matter in Planolites burrows, and associated minerals were produced by heterotrophic, biomineralizing microorganisms (Sagemann et al, 1999).

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