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IDENTIFICATION OF MICROBIAL ECOSYSTEMS IN THE

STROMATOLITES OF THE BIWABIK IRON FORMATION

(PALEOPROTEROZOIC, MINNESOTA) USING MINERALOGY

AND FACIES ANALYSIS

______

A Thesis

Presented

to the Faculty of

California State University, Chico

______

In Partial Fulfillment

of the Requirements for the Degree

Master of Science

Geosciences

______

Kathi McCarthy 2014

Spring 2014 IDENTIFICATION OF MICROBIAL ECOSYSTEMS IN THE

STROMATOLITES OF THE BIWABIK IRON FORMATION

(PALEOPROTEROZOIC, MINNESOTA) USING MINERALOGY

AND FACIES ANALYSIS

A Thesis

Kathi McCarthy

Spring 2014

APPROVED BY THE DEAN OF GRADUATE STUDIES AND VICE PROVOST FOR RESEARCH:

______Eun K. Park, Ph.D.

APPROVED BY THE GRADUATE ADVISORY COMMITTEE:

______Russell Shapiro, Ph.D., Chair

______Todd Greene, Ph.D.

______Gordon Wolfe, Ph.D.

ACKNOWLEDGEMENTS

I would like to thank the staff at Natural Resources Research Institute for allowing access to the core samples that were used in this thesis, especially John Heine and Mark Severson. I would like to also thank the College of Natural Sciences and the

Chico Research Foundation for financial support so that I could go to Minnesota for research purposes.

I want to give a shout-out to Amber Rutter and Sam Jameson for assistance identifying the various lithofacies. You guys are awesome, and your help was much appreciated. I want to thank my understanding husband Scott McCarthy with helping me with this whole process by proof-reading everything, and my two sons Aidan and Faelan, for understanding that they were going to have to deal with a lot less mommy time.

Thank you to Dr. Russell Shapiro, my advisor, for believing that I could do this. Not to mention all the help with formatting, and figures. You ROCK!

Thank you to Dr. Todd Greene and Dr. Gordon Wolfe for willing to be on my committee. Thank you to anyone I might have missed, and I’m sure I’ve missed somebody.

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TABLE OF CONTENTS

PAGE

Acknowledgements...... …... iii

List of Tables...... …... vi

List of Figures…………………………………...... ……………. x

Abstract...... …... xi

CHAPTER

I. Introduction……………………………………………………………….. 1

Hypothesis………………………………………………………….. 5 Potential Organisms of the Biwabik………………………………... 5 Cyanobacteria………………………………………………………. 6 Iron-oxidizing Bacteria………………………………………...... 8 Summary…………………………………………………………… 12

II. Background………………………………………………………………. 15

Ecological Constraint Models of Potential Organisms……………. 15 Geologic History…………………………………………………... 23 Pokegama Formation………………………………………………. 25 Biwabik Formation………………………………………………… 26 Virginia Formation………………………………………………… 27 Mineralogy, Diagenesis, Metamorphism and Geochemistry……… 29

III. Methods………………………………………………………………….. 39

IV. Results…………………………………………………………………… 40

Lithofacies…………………………………………………...... 40 Coarse-grained Sandstone……………………………………….... 42 Stromatolite………………………………………………………. 43 Medium-grained Sandstone………………………………………. 44

CHAPTER PAGE Very Fine-grained Sandstone……………………………………. 45 Intercalated Medium- and Very Fine-grained Sandstone……….. 45 Brecciated Intraclasts……………………………………………. 46 Stromatolites with Flat-pebble Conglomerate and Oncoids...... 47 Stromatolites with Ooids…………………………………………. 48 Stromatolites with Peloids………………………………….……. 49 Stromatolites Associated with Intraclasts………………….…….. 50 Intraclasts………………………………………………….….….. 51 Peloids…………………………………………………….……… 53 Ooids……………………………………………………………… 54 Stromatolites……………………………………………….…….. 57 Bladed/fibrous Cement……………………………………...... …. 64 Peloidal Cement………………………………………………….. 65 Granular Mosaic Cement…………………………………………. 65 Drusy Fabric………………………………………………………. 66 Splay Cement……………………………………………………… 67 Ankerite Crystals…………………………………………………. 67 Micro-stromatolites………………………………………………. 69 Shrinkage Cracks………………………………………….………. 69 Vacuoles in Stromatolite Laminae………………………………... 70 Remnant Stromatolites……………………………………………. 71 Microfossils…………………………………………..………….. 71 Summary of Results………………………………………………. 72

IV. Discussion………………………………………………………….…... 78

Mineralogy…………………………………………………..…… 78 Paleoenvironment………………………………………………… 79 Lithofacies to Facies Analysis……………………………………. 84 Organisms……………………………………….………………… 87

V. Conclusion…………………………………………………………..….. 90

References……………………………………………………………………..…… 91

Appendices

A. Stratigraphic Column Descriptions and Facies Interpretations………… 102 B. Data Tables……………………………………………………………… 138

LIST OF TABLES

TABLE PAGE

1. Localities mentioned in thesis...... 3

2. Potential microorgansims responsible for building the stromatolite layers... 13

3. Instraclast width and length measurement...... 139

4. Peloid width and length measurements...... 141

5. Ooid width and length measurements...... 142

6. Ooid cortice/ ooid core ratios...... 143

7. Characteristics of stromatolite laminae...... 144

8. Stromatolite laminae widths and mineralogy...... 145

9. Microfossil width and length measurements...... 148

LIST OF FIGURES

FIGURE PAGE

1. Map of the Gunflint and Biwabik Iron Formations………………………… 2

2. Depositional model for the Sokoman Formation…………………………… 15

3. Depositional model for the Biwabik Iron Formation……………………….. 19

4. Depositional Model for the microorganisms of the Duck Creek Formation.. 21

5. Combination of the three depositional models……………………………… 22

6. A generalized stratigraphic column of the Biwabik Formation…………….. 40

7. Stratigraphic column of the Lower Cherty stromatolite layer……………… 41

8. Stratigraphic section of the Upper Cherty stromatolite layer………………. 42

9. Coated intraclasts…………………………………………………………… 52

10. Intraclasts and flat pebble conglomerate……………………………………. 52

11. Dissolved peloids…………………………………………………………… 53

12. Dissolved ooids…………………………………………………………...... 54

13. Shrinkage cracks due to diagenesis within ooid cores. PPL...... 55

14. Dissolution of coated intraclasts and deformed ooids. PPL...... 56

15. Hematite replacement of microbial ooid core...... 56

16. Delamination of the ooid cortice...... 57

17. Columnar stromatolite...... 58

18. Digitate columnar stromatolite...... 59

vii

FIGURE PAGE

19. Digitate columnar stromatolite...... 59

20. Void...... 60

21. Void with fill...... 61

22. Void with calcite...... 61

23. Laminae with Fe-oxide...... 62

24. Gamma branching stromatolites...... 63

25. Peloids with bladed calcite...... 64

26. Peloids with cement...... 65

27. Granular mosaic cement...... 66

28. Euhedral calcite...... 67

29. Euhedral ankerite...... 68

30. Amorphous ankerite...... 68

31. Microdigitate stromatolites...... 69

32. Shrinkage cracks...... 70

33. Microfossils...... 72

34. Microcrystalline quartz...... 74

35. Stromatolite laminae...... 74

36. Psuedocolumnar stromatolites...... 80

37. Psuedocolumnar stromatolites...... 81

38. Wavy beds...... 82

39. Wavy beds...... 83 viii

FIGURE PAGE

40. Facies diagram...... 85

41. Wave action...... 104

42. Wave action...... 105

43. Breccia and conglomerate...... 107

44. Flat laminar bedding...... 108

45. Sharp contact...... 110

46. Sharp contact between member 3 and member 4...... 112

47. Brecciated clasts...... 113

48. Breccia with wavy layered sandstone...... 115

49. Hematite staining...... 116

50. Breccia with domal stromatolites...... 118

51. Flat- pebble conglomerate...... 120

52. Psuedocolumnar stromatolites...... 122

53. Flat-pebble conglomerate with pseudocolumnar...... 123

54. Undulatory stromatolites...... 126

55. Undulatory stromatolite layers amongst breccia...... 127

56. Granular veins of calcite amongst the undulatory...... 128

57. Granular veins of calcite amongst the undulatory stromatolite...... 129

58. Dendritic stromatolites...... 131

59. Ankerite...... 132

60. Psuedocolumnar stromatolite layer with agate veins...... 133 ix

61. Undulatory stromatolites with evidence of cross-bedding...... 136

ABSTRACT

IDENTIFICATION OF MICROBIAL ECOSYSTEMS IN THE STROMATOLITES OF

THE BIWABIK IRON FORMATION (PALEOPROTEROZOIC, MINNESOTA)

USING MINERALOGY AND FACIES ANALYSIS

Kathi McCarthy

Master of Science in Geosciences

California State University, Chico

Spring, 2014

Prior to the Great Oxygenation Event (GOE) at 2.4-2.3 Ga the world’s oceans were anoxic and drastically different from the oxygen stratified oceans of today. Previous work on defining the GOE has consisted of geochemical data including sulfur isotopic data, iron isotopic data, carbon isotopic data, and rare earth elements to address oceanic conditions. Because biomarkers and carbon isotopes are rarely preserved in iron- formations, the use of fossil evidence for the transition from anoxic bottom waters to suboxic shallow water conditions is poorly represented by rare stromatolites, and even more rare microfossils. Previous geochemical analysis led to the conclusion that Fe- oxidizing bacteria formed the stromatolites preserved in the 1.89 Ga Gunflint and

Biwabik iron formations (Ontario and Minnesota) and other iron-formations of this post-

GOE interval.

This thesis uses a geological approach to define the ecological constraints on the formation of stromatolites in the Biwabik Iron Formation, specifically. Detailed facies analysis along with primary mineralization estimates shows that the primary deposition for the stromatolite layers was a high energy, shallow water environment. 400 million years after the GOE, oceans would have been stratified with regards to oxygen, and the shallow waters, while being suboxic, would have contained enough oxygen to limit the amount of non-reduced iron needed for Fe-oxidizing bacteria to utilize for energy synthesis. These factors suggest that cyanobacteria were the stromatolite builders in the

Paleoproterozoic Biwabik Iron Formation.

xii

CHAPTER I

INTRODUCTION

Prior to the Great Oxygenation Event (GOE) at 2.4- 2.3 Ga, the ocean was anoxic which would have had dramatic effects on the marine ecosystems (Holland, 2002;

Pufahl et al., 2012). Previous work on defining the GOE used geochemical evidence including sulfur isotopic data (Canfield 2005), iron isotopic data (Johnson et al., 2003;

Frost et al., 2007; Planavsky et al., 2009), rare earth elements (Planavsky et al., 2009;

Yun Duan et al., 2010), 18O isotopic data (Perry et al., 1973), δ13C isotopic data (Perry et al., 1973), and geochemical analysis of bioelemental sediments (Pufahl et al.,2012) to address the oceanic conditions at the time of deposition of the unique iron-formations found at during this time. However, because biomarkers and carbon-isotopes are rarely preserved in iron-formations, fossil evidence of the transition from an anoxic marine environment to a stratified ocean with oxygenated surface waters and anoxic bottom waters is poorly represented. When fossils are preserved, they are preserved primarily as stromatolites, laminated sedimentary structures built by microbial communities

(Kalkowski, 1908; Burne and Moore, 1987).

Most stromatolites do not retain evidence of the microorganisms. The rare stromatolite layers that retain the microorganisms do not show clear evidence of whether the microorganisms are the actual builders of the stromatolites or instead are passive components entombed in the stromatolites during construction (Knoll et al., 1976;

Tomtani et al., 2006).

Stromatolites are layered accretionary structures formed from trapping, binding and cementing of sediment grains by microbial mats of various bacterial

2 communities, mostly in shallow water (Riding, 2007). Weed (1889) described the first stromatolites layers based on his observations in Yellowstone National Park. Matthew

(1890) was the first to name a stromatolite in Precambrian strata from New Brunswick,

Canada. He named the stromatolite Archaeozoon acadiense though he did not recognize the microbial contribution. Walcott (1914) described the first stromatolite species as

Cryptozoan proliferans.

With regards to the Paleoproteozoic biosphere, the critical lagerstätten for fossil stromatolites—and significantly, microfossils—is the Gunflint Iron Formation of

Ontario, Canada (Barghoorn and Tyler, 1965; Cloud, 1965; Nudds and Selden, 2004)

(Figure 1).

Figure 1: Map of the Gunflint and Biwabik Iron Formations. Maps show the locations of the core, thin-section and hand samples described in the thesis. Upper figure is a detail of the Mesabi Iron Range, lower figure is a larger-scale map of Lake Superior showing critical locations of the Gunflint Iron Range in Ontario, Canada.

TABLE 1. SAMPLE LOCALITIES Sample Locality Zone UTM E (meters) UTM N (meters) Gunflint Lake (GL) 15T 662365.49 532937.67 5 NI 15T 571085.96 527855.19 17729 15T 526850.14 5265094.63 2 East (2E) 15T 568288.58 5270493.52 17883 15T 524747.09 5265069.40 Kakabeka Falls (KF) 16U 305171.79 536729.42 North Shore 15T 573842.46 5276062.39 Mink Mountain (MM) 16U 711024.60 5346980.57 McKinley Mine (McM) 15T 545753.67 526146.63 Minnesota Geological Survey 2 (MGS2) 15T Not available Not available Nolalu 16U 292432.98 535058.11 Schreiber 16U 474689.46 5404930.40 Table 1: Localities mentioned in thesis. All localities are in UTM, NAD83 datum. Abbreviations also used in Figure 1, Locality Map.

The Gunflint Iron Formation extends 170 km from Gunflint Lake, to Loon

Lake, Ontario (Pufahl et al., 2004). The Gunflint Iron Formation located in Ontario,

Canada was first described by Ingall in 1887 (Ingall, 1888). Ingall described the iron- bearing strata located near Silver Mountain and Whitefish Lake. Later geologic accounts were made by Smith (1905), Silver (1906), Van Hise and Leith (1911), and Gill (1927).

Goodwin (1956) divided the iron-formation into six sedimentary facies and four different members: 1) Basal Conglomerate Member, 2) Gunflint Member, 3) Upper Gunflint

Member, and 4) Upper Limestone Member. Fralick and Barrett (1995) redefined the

Gunflint stratigraphy as consisting of only two membersi 1) a lower conglomerate consisting of stromatolites, oncolitic cherts, and cherty grainstones and 2) an upper member with a similar stratigraphy to the lower member.

The Mesabi Iron Range extends 120 miles from Cass County, east of Babbit,

Minnesota, to the termination of the range at the Duluth Igneous Complex (Severson and

Heine, 2007). The Biwabik Iron Formation was first described by Van Hise and Leith

(1911). The Biwabik Iron formation is coeval with the Gunflint Iron Formation and has been altered due to the Duluth Igneous Complex (French, 1973). The Biwabik Iron

Formation is broken up into four different members; 1) Lower Cherty Member, 2) Lower

Slaty Member, 3) Upper Cherty Member, and 4) the Upper Slaty Member (Severson and

Heine, 2007). While the Biwabik Iron Formation has been altered by metamorphism, it contains similar sedimentary layer as the Gunflint Iron Formation.

Stromatolites in the Biwabik and Gunflint formations were first described by

Grout and Broderick (1919), and the stromatolites were described as sedimentary structures attributed to microbial organisms. Cloud and Semikhatov (1969) named the stromatolites within the Gunflint formation Gruneria biwabika and attributed cyanobacteria as the stromatolite builders.

Barghoorn and Tyler (1965) called the stromatolite builders cyanobacteria due to the poor understanding of bacteria at the time. Understanding of bacteria has changed since their time, and more recent studies including Edwards et al. (2012), Plavanasky et al. (2009), and Wilson et al (2010) have reconsidered Fe-oxidizing bacteria instead of cyanobacteria. Their reconsideration of the stromatolite builders as Fe-oxidizing bacteria is based on geochemical data and only minutely on the simplistic, filamentous microfossil morphology.

Identifying the microorganisms responsible for building the stromatolites is critical to understanding the change in oxygenation state of the Earth’s atmosphere following the GOE. With increasing oxygen levels in a strongly stratified ocean,

5 phototrophs (including cyanobacteria) and chemotrophs (including Fe-oxidizing bacteria) were both possible stromatolite builders at the time of the shallow marine deposition of the Gunflint and Biwabik Iron Formations. Cyanobacteria and Fe-oxidizing microfossils are often difficult to identify due to a similarity in morphology and the lack of biomarkers. Knowing the identity of the microorganism that built the stromatolite layers within the Biwabik and Gunflint Iron Formations will elucidate the eutrophication of the ocean at the time of deposition.

Hypothesis

Recent studies to identify the stromatolite builders have used geochemistry and broad facies analysis (Edwards et al, 2012; Planavasky et al., 2009; and Wilson et al., 2010). This thesis provides an alternative approach by combining a detailed facies analysis of the stromatolite layers along with primary mineralization, and ecological constraints of extant bacterial communities to identify the microorganisms responsible for the stromatolite layers found within the 1.89 Ga Biwabik Formation.

Potential Organisms of the Biwabik

Planavsky et al. (2009) believed that the oceans were different in temperature and geochemistry then modern oceans. Temperatures during the Paleoproterozoic needed for the precipitation of carbonates as well as silica, and iron-rich ore deposits are estimated to range between 10-35oC (Shiriashi, 2012; however, see Robert and

Chaussidon, 2006; and Maliva et al., 2005 for an alternative). Oceans at this time (based on rare earth element chemistry) were interpreted to be redox-stratified in regards to free molecular oxygen content (Rasmussen et al., 2010) and contained dissolved Fe(II) and

6 dissolved silica (Planavsky et al., 2009; Koehler et al., 2010; Fischer and Knoll, 2009 ).

Fe(II) enrichment of oceanic waters originated from hydrothermal vents (Koehler et al.,

2010) and upwelling moved to zones where dissolved Fe(II) could then be used by micro-organisms for creating ATP in photosynthesis (Moit et al., 2009).

The most likely microbes that could have built the stromatolites in the

Biwabik Formation are either cyanobacteria or iron-oxidizing bacteria based on studies of modern and Paleoproterozoic stromatolitic structures (Planavsky et al., 2009; Wilson et al., 2010; Edwards et al., 2012). To predict the ecosystems that could have built the stromatolites in the Biwabik Formation, I will present the ecological constraints of these different putative microorganisms.

Cyanobacteria

Cyanobacteria are a phylum of bacteria that obtain energy through photosynthesis. They are a major component of the marine nitrogen cycle and a primary producer in many areas of the ocean (Gault, 2009). Cyanobacteria complete photosynthesis in two ways. Like higher plants, cyanobacteria use sunlight and convert it to energy in chlorophyll A sites (Kirchman 2008). Cyanobacteria may also complete photosynthesis is in what is called the dark cycle (Kirchman 2008). This does not require direct sunlight to be converted to energy but instead allows the cyanobacteria to use other sources including carbon, sulfur, and nitrogen from organic compounds as an electron donor to produce energy (Kirchman, 2008). Thus, they are able to produce energy either directly with sunlight (photoautotrophic) or in the dark with another electron donor

(heterotrophic). Cyanobacteria are the only organisms that are capable of reducing nitrogen and carbon in aerobic conditions (Galt, 2009).

Nitrogen reduction in cyanobacteria, also known as biological nitrogen fixation, occurs when atmospheric nitrogen is converted to ammonia by an enzyme called nitrogenase (Galt, 2009). The ammonia is than converted to glutamate and the formation of a H2 molecule also occurs. Nitrogen fixation cannot occur where oxygen is present so specialized antennae inside the phytoplasm of the cyanobacteria keep oxygen away allowing the nitrogen fixation to occur (Galt, 2009). The reaction for biological nitrogen fixation is:

+ − N2 + 8 H + 8 e → 2 NH3 + H2 (1) (Postgate, 1998)

Carbon fixation in cyanobacteria is done using the Calvin Cycle (Galt, 2009).

The Calvin Cycle reduces carbon dioxide to sugar (triose phosphate [TP]). The reaction for the conversion of carbon dioxide to sugar is:

- + 3CO2 + 12e + 12H + Pi → TP + 4H2O (2) (Tomitani et al., 2006)

Carbon concentrating mechanisms in cyanobacteria, allow the conversion of dissolved inorganic carbon found in low concentrations in the form of CO2 to be

- converted to bicarbonate HCO3 through the Calvin cycle with the aid of the enzyme

RuBisCO. The benefit of RuBisCO and other carbon concentrating mechanisms is to allow cyanobacteria to tolerate low concentrations of inorganic carbon necessary for photosynthesis to be utilized and decreases the amount of carbon lost during the dark cycle.

Cyanobacteria can mineralize Fe(II)-oxyhydroxides in the presence of Fe-rich waters (Krepski et al., 2013). However, cyanobacteria mineralize the Fe(II)- oxyhydroxides as uneven granular coatings, unlike the smooth even coating mineralized by Fe-oxidizing bacteria (Krepski et al., 2013). Microfossils of cyanobacteria are typically non-bifurcated, with occasional branching, and have wider stalks, ranging from

0.5- 100 µm than iron-oxidizing bacteria, with an average stalk width of 1.1 µm (Krepski et al., 2013). Cyanobacteria filaments are typically densely packed microfossils vertically oriented overlaying layers of roughly parallel horizontal filaments (Krepski et al., 2013).

Iron-oxidizing Bacteria

Fe-oxidizing bacteria convert Fe(II) to Fe(III) by using iron as an electron donor during chemolithotrophy (Moit et al., 2009). The by-product of this iron oxidation is iron hydroxides (ferrihydrite) that precipitate just outside the cell walls of the bacteria

(Moit et al., 2009). The electron that is claimed from the oxidation of Fe(II) is then transported across the cell membrane into the periplasm to completeenergy synthesis

(Moit et al., 2009). Fe-oxidizing bacteria mineralize Fe(II)-oxyhydroxides as solid even coatings surrounding their stalks (Krepski et al., 2013).

Most Fe-oxidizing bacteria are bifurcated with radial twists in the filaments due to preferential orientation during cell growth (Krepski et al., 2013). Fe-oxidizing average filament width of 1.1 µm is less than for cyanobacteria. Fe-oxidizing bacterial filaments have a vertical orientation towards oxygen compared with the cyanobacterial filaments that can orient horizontally as well as vertically (Krepski et al., 2013).

Fe-oxidizing bacteria are separated into four different physiological groups based on the requirements needed to complete chemolithotrophy (Henrich et al., 2011).

The four physiological groups are: 1) acidophilic, aerobic Fe-oxidizers; 2) neutrophilic, aerobic Fe-oxidizers; 3) neutrophilic, anaerobic (nitrate dependent) Fe-oxidizers; and 4) anaerobic, photosynthetic Fe-oxidizers (Henrich et al., 2011).

Acidophilic, aerobic Fe- oxidizers use two distinct pathways to oxidize ferrous iron. One pathway couples the oxidation of ferrous iron with the reduction of sulfur compounds, while the other pathway just oxidizes ferrous iron (Henrich et al.,

2011). Species of acidophilic, aerobic Fe- oxidizers that couple sulfur reduction with the oxidation of ferrous iron include Acidothiobacillus ssp., Thiobacillus ferroxidans, and

Acidiferrobacter thiooxydans (Henrich et al., 2011). All three species are rod-shaped

Gram-negative bacteria (Henrich et al., 2011). A species of acidophilic, aerobic Fe- oxidizers that do not reduce sulfur along with ferrous oxidation is Thiobacillis prosperus.

All of the Fe-oxidizers are extreme acidophiles with an optimum pH of 3.0 and a minimum pH of 2, with some species being more tolerant of extreme acidity, higher temperatures and salt content than others (Henrich et al., 2011). Acidophiles have pre- existing pH gradients across their membranes that can drive the synthesis of ATP during chemolithotrophy (Henrich et al., 2011). In microaerobic, circum-neutral waters acidophiles can utilize iron oxidation to create energy because the potential for oxidation of iron from molecular oxygen is much slower than in oxygenated waters (Henrich et al.,

2011).

Neutrophilic, aerobic Fe-oxidizing proterobacteria include the genus

Gallionella (Henrich et al., 2011). Neutrophilic, aerobic Fe-oxidizing proterobacteria oxidize ferrous iron in oxygen-rich, pH neutral waters and often colonize the redoxcline, the boundary between anoxic and oxic to suboxic waters (Henrich et al., 2011). Unlike acidophiles, neutrophilic Fe-oxidizers do not have the pre-existing pH gradients across their membranes and so the ability to oxidize ferrous iron is lower in neutrophilic Fe- oxidizers than in acidophiles (Henrich et al., 2011). Genera such as Gallionella use ferrous iron as an electron donor in the production of ATP (Hallberg and Ferris, 2004):

2+ + 3+ 4Fe + XO2 + (1-X)CO2 + 4H = 4Fe + (1-X)CH2O + (2-x)H2O (3)

Gallionella form a spirally twisted stalk where the oxidized iron is precipitated and are classified as gradient bacteria in regards to oxygen concentration

(Hallberg and Ferris, 2004). Gallionella is autotrophic and oxidation of iron takes place under suboxic conditions (Hallberg and Ferris, 2004).

Neutrophilic, Fe-oxidizing proterobacteria that respire on nitrate have two distinct metabolic pathways for iron oxidation (Henrich et al., 2011). In the first, iron oxidation is used as a source of electrons in some photosynthetic bacteria. Ferrous iron is used as the electron donor in the second pathway and nitrate is used as the electron acceptor (Henrich et al., 2011). In this group nitrate is reduced to ammonium and ultimately the nitrate is reduced to nitrogen gas in the reaction:

- 10FeCo3 + 2NO3 + 10H2O →

- + Fe10O14(OH)2 + 10HCO3 + N2 + 8H (Henrich et al., 2011) (4)

Genera included in neutrophilic, Fe-oxidizing proteobacteria are:

Aquabacterium,Thermomonas, and Acidovorax and all three genera are gram negative

(Henrich et al., 2011). For the oxidation of ferrous iron to ferric iron to occur in neutrophilic iron-oxidizing bacteria, they must be in a circum-neutral pH environment unlike their acidophilic cousins (Henrich et al., 2011).

Phototrophic Fe-oxidizing proteobacteria, also known as phototrophic purple proterobacteria were the first micro-organisms to be identified as Fe-oxidizers (Henrich et al., 2011; Straub et al., 2001). These bacteria are able to oxidize ferrous iron into ferric iron within the photic zone through photosynthesis that involves light-energy CO2 fixation coupled with oxidation of Fe2+ (Koehler et al., 2010). The ferrous iron is used as a source of reductant or electron donor for

CO2:

2+ + 4Fe + CO2 + 11H2O + hv → CH2O + 4Fe(OH)3 + 8H (Henrich et al., 2011) (5)

Phototrophic Fe- oxidizers are also known to utilize sulfur/sulfide as an electron donor for photosynthesis.

Ferric iron precipitates are considered to be a waste product for phototrophic

Fe-oxidizers as well as neutrophilic Fe- oxidizers that poses the risk of enshrouding the bacteria in a ferrihydrite shell (Henrich et al., 2011). Without the ferric iron precipitate forming the bacteria become at risk of being unable to access light needed for photosynthesis (Henrich et al., 2011).

Photosynthetic Fe-oxidizers may not require oxygen or nitrates to oxidize iron which could have major implications for the development of banded iron-formations

12 during Archean and Paleoproterozoic times when the planet was essentially anoxic before the development of oxygen-enriched atmosphere and oceans (Henrich et al., 2011).

Summary

Cyanobacteria and iron-oxidizing bacteria obtain energy through the process of photosynthesis. Cyanobacteria can obtain energy by either using sunlight directly

(photoautotrophic) or by using carbon, nitrogen, and sulfur as an electron donor

(heterotrophic) (Kirchman, 2008). Cyanobacteria are often found in aerobic environments that are oxygen rich in comparison to other zones within the marine environment

(Kirchman, 2008). Cyanobacteria absorb dissolved inorganic carbon during oxygenic photosynthesis and precipitate CaCO3 during the dark cycle of photosynthesis (Shiraishi,

2012) a process integral to the formation of carbonate structures in shallow marine environments.

Iron-oxidizing bacteria can use oxygen for energy synthesis, but when oxygen is not abundant or unavailable they will use iron as an electron donor to complete energy synthesis. The by-product of Fe-oxidation is Fe-hydroxide (ferrihydrite) that precipitates outside the cell wall (Moit et al.,, 2009). Fe-oxidizers are broken down into four different physiological groups based in their requirements to complete photosynthesis (Henrich et al., 2011); 1) acidophilic, aerobic Fe-oxidizers; 2) neutrophilic, aerobicFE- oxidizers; 3) neutrophilic, anaerobic (nitrate dependent) Fe-oxidizers; 4) anaerobic, photosyntheticFe- oxidizers. The summary table listed below (Table 2) demonstrates the different environmental requirements of the microbial communities to complete

13 photosynthesis/chemolithotrophy and precipitate out either CaCO3 or Fe-hydroxides as a by-product.

TABLE 2. ECOLOGICAL CONSTRAINTS OF MICROORGANISMS

Organism Oxygenic Redox pH Temperature Zone photosynthesis Potential Cyanobacteria Yes None Neutral 10-70oC Oxic Acidophilic, No Ferrous 2-3 10-35oC Oxic aerobic iron oxide oxidizing bacteria Neutrophilic, No Ferrous Neutral 10-35oC Suboxic aerobic iron oxide oxidizing bacteria Neutrophilic, No Ferrous Neutral 10-70oC Suboxic-anoxic anaerobic iron oxide, oxidizing requires bacteria nitrogen Anaerobic, Yes Ferrous Neutral 10-35oC Suboxic-anoxic photosynthetic oxide, iron oxidizing typically uses bacteria sulfur as electron donor Table 2: Potential microorgansims responsible for building the stromatolite layers.

The table includes photosynthetic requirements of each type of microorganism, redox potential, pH requirements, temperature needs, and the oxygenation levels.

Cyanobacteria require an oxic environment with near neutral pH to photosynthesize (Kirchman, 2008) and precipitate CaCO3 (Shiraishi, 2012). Fe-oxidizing bacteria can live in a multitude of marine environments, but have additional constraints based on pH and oxygen levels. All the Fe-oxidizing bacteria except the acidophilic, aerobic Fe-oxidizing bacteria have a low redox potential, meaning low ferrous oxide ability (Henrich et al., 2011). This means they are less likely to have created the extensive iron deposits of the banded iron-formations.

Using the ecological constraints of Fe-oxidizing and cyanobacteria along with mineralogy and facies analysis will provide an understanding of the depositional environment, this will allow for the interpretation of which microorganism created the stromatolite horizons within the Biwabik Iron Formation.

CHAPTER II

BACKGROUND

Ecological Constraint Models of Potential Organisms

Three ecological constraint models have been suggested for the organisms and depositional settings for Precambrian banded iron-formations: 1) Edwards et al. (2012);

2) Planavsky et al. (2009); and 3) Wilson et al. (2010).

Edwards et al. (2012) studied the 1.88 Ga Ferriman Group, Labrador Trough,

Canada, and interpreted the microbial organisms to be iron-oxidizing bacteria in a suboxic environment. Taphonomy and paleoenvironment were used to establish their model for the precipitation of Precambrian Ferriman Group banded iron-formation

(Figure 2).

Figure 2: Depositional model for the Sokoman Formation. Depositional model for the microorganisms of the Sokoman Formation, Labrador Trough, Canada in the 1.88 Ga Ferriman Group. Stromatolites formed in the intertidal zone under suboxic conditions. Modified from Edwards et al., 2012 (fig. 4).

The Ferriman Group consists of the Wishart Formation, Ruth Formation,

Sokoman Formation, and the Menihek Formation. The Ferriman Group is filled with clastic and sedimentary rocks from three different depositional cycles. The Sokoman

Formation is the banded iron-formation within the Labrador Trough. Black chert within the Sokoman Formation reflects the onset of upwelling waters rich in iron and silica.

They argue that lateral and vertical lithofacies changes record the accumulation of clastic and sedimentary rocks through three sea-level cycles in three stratigraphic sequences, primarily of transgressive, peritidal sediments overlain by highstand, peritidal and deeper subtidal deposits.

Stromatolites occurring along with suboxic, authigenic mineral assemblages in peritidal, nearshore environments suggest that deposition of the stromatolitic layers occurred in oxygen stratified waters. Minerals recording anoxic environments in middle and distal shelf settings also suggest deposition occurred in oxygen stratified waters.

According to Edwards et al. (2012), shallow, suboxic lithofacies contain abundant hematite and chert. Hematite is interpreted to have formed under suboxic conditions by the transformation of an Fe-oxyhydroxide precursor that precipitated in the water column or authigenically at the seafloor. Chert is thought to have precipitated initially as metastable opal-A or opal-CT gel that recrystallized during diagenesis to produce quartz (Edwards et al., 2012). Deeper anoxic paleoenvironments are characterized by magnetite, greenalite and stilpnomelane (Klein, 2005). Because these minerals contain ferrous iron, they reflect precipitation under extremely low oxygen

17 concentrations (Klein, 2005). Magnetite is interpreted to have formed authigenically in sediment and directly in seawater (Fralick and Pufahl, 2006; Johnson et al., 2003).

The evaporative concentration of suboxic waters and iron-redox of concentrated silica in pore waters along with an absence of limestone and dolostone suggests that the seawater was carbonate undersaturated (Edwards et al., 2012).

Edwards et al. (2012) further suggest that the occurrence of filamentous fossils that form only in suboxic paleoenvironments suggest Fe-chemoautotrophy and that the presence of silicified and phosphatized microfossil filaments containing authigenic hematite cement implies Fe-oxidation was a potential metabolic pathway.

They compared the morphology of filamentous fossil forms to extant taxa of cyanobacteria and iron-oxidizing bacteria. Based on the filament width, non-bifurcating filaments containing no internal cells and surficial ornamentation, Edwards et al. (2012) likened the fossil filaments to extant taxa of microaerophilic Fe-oxidizing bacteria. With this comparison they then stated that because most modern iron-oxidizing bacteria live in circumneutral conditions that the only place they could exist in marine environments was in the suboxic environments where the bacteria could bridge the suboxic-anoxic redox interface in sediment.

Low δ13C isotopic values obtained from the sedimentary organic matter located within microfossil-rich beds are considered similar to the range of values produced by Fe-oxidizing bacteria and cyanobacteria. However, most of the δ13C values are below the typical range for cyanobacteria. They hypothesized that the fractionation occurred predominantly by Fe-oxidizing bacteria. The low δ13C values also suggest

18 minimal alteration and metamorphism within the Ferriman Group, as alteration and metamorphism tend to shift δ13C values to ranges greater than observed (Edwards et al.,

2012).

Finally, silica was considered to be concentrated in pore spaces by evaporation (Maliva et al., 2005;Brasier et al., 2002) and Fe-redox pumping (Fischer and

Knoll, 2009; Pufahl, 2010). Chert would have readily precipitated along the evaporating seashores during the Precambrian due to being near silica saturation (Maliva et al., 2005).

In deeper water the Fe-redox pumping could not saturate the pore spaces with silica to preserve the microfossils; this meant that microfossils could not have formed in a benthic environment (Edwards et al., 2012) (Figure 2).

Planavsky et al. (2009) used rare earth elements and iron isotopes to determine the depositional environment of the micro-fossil rich stromatolite layers within the Gunflint and Biwabik Formations. They also state that the morphology of the microfossils located within the Gunflint Formation resemble the morphology of extant taxa of iron-oxidizing bacteria. This includes rare spiral filaments (Cloud, 1965), non- septate filaments with endospores (Barghoorn and Tyler, 1965), and small straight and curved rods and coocoids with iron-rich cell walls (Tazaki et al., 1992) that are similar to cultured aerobic iron-oxidizers (Emerson and Moyer, 1997).

Planavsky et al. (2009) used REE systematics to constrain the redox conditions within the overlying water column above the fossiliferous layers. Shifts from positive to nonexistent Ce anomalies to negative anomalies are considered indicative of a strong redoxcline. Microfossiliferous and hematite-rich stromatolites lack of negative Ce

19 anomalies suggest that the stromatolites formed in low oxygen environments below or at the redoxcline. Where the Ce anomalies are positive, sediments containing fossils indicate formation in deeper water below the redoxcline.

Iron-oxides precipitated in fully oxygenated conditions within cyanobacterial mats have the same δ56Fe values as dissolved iron sources because of nearly complete iron oxidation and precipitation. Positive δ56Fe values indicate partial ferrous iron oxidation from Fe-oxidizing microbial mats. Modern ferric Fe-oxides from open-marine sediments have negative δ56Fe values. A predominance of positive δ56Fe values in the stromatolites suggests modest oxidation as metabolic pathway. The modest oxidation suggests organisms other than cyanobacteria formed the stromatolites.

Figure 3: Depositional model for the Biwabik Iron Formation. Depositional Model for the microorganisms of the Biwabik Iron Formation. Microorganisms formed in the suboxic zone, below the redoxicline from Fe-rich water. The Fe-enriched water occurred due to upwelling from the anoxic zone. Modified from Planavasky et al., 2009 (figure 8).

The presence of a shallow redoxcline due to oxygenation of surface waters in the Paleoproterozoic and an upwelling of hydrothermally dominated deep seawater rich in iron-oxide allowed for a unique environment on continental shelf ideal for chemolithotrophic microbial communities to thrive (Figure 3).

Wilson et al. (2010) looked at microfossils found within the Duck Creek

Formation, Western Australia and found them similar in morphology, size and shape to the microfossils observed within the Gunflint Formation, Canada. They also observed microfossils interpreted to be formed by cyanobacteria and transported within chert nodules to deeper marine environments.

Microfossils found in chert nodules resembled the morphology of extant taxa of cyanobacteria or sulfur-oxidizing bacteria. Microfossils within the Duck Creek

Formation have similar length filaments to the Gunflint Formation microfossils and also include small cocci assignable to Fe-oxidizing bacteria, asteriform microfossils similar to eukaryotic algae, and rare larger trichomes comparable to extant cyanobacteria or sulfur- oxidizing bacteria. However, the lack of sulfide minerals in the Duck Creek Formation excludes the possibility of sulfur-oxidizing or sulfur-reducing bacteria from being the organism forming the stromatolites.

The presence of glauconite, which typically forms below the thermocline, at water depths of 125-250 m (10-15oC) combined with berthierine, which forms nearshore at depths of 10-50 m, and in relatively warm water (25-27oC) indicate that the Duck

Creek formation formed in an Fe-rich setting. Stratigraphic context places Gunflint-type microfossils in Fe-rich deep waters below the redoxcline, supporting the theory that the

21 microfossils are formed from organisms with iron-based metabolism (i.e., Fe-oxidizing)

(Figure 4).

Figure 4: Depositional Model for the microorganisms of the Duck Creek Formation. Depositional model for the microorganisms of the Duck Creek Formation. The Duck Creek Formation is located in the Ashburton Basin of Western Australia, Late Paleoproterozoic era deposition. Microorganisms formed in the suboxic intertitdal zone. Modified from Wilson et al., 2010 (fig. 10).

I have combined the three different depositional models provided by Edwards et al. (2012), Planavsky et al. (2005), and Wilson et al. ( 2010),( Figures 2,3,and 4) into one depositional model (Figure 5). This allows the reader to compare the depositional environments based on the above mentioned data. Each of the different depositional models argues for a different environment for stromatolite formation. Edwards et al.

(2012) place the stromatolites into the subtidal paleoenvironment, while Planavsky et al.

(2009) place the stromatolites into the shelf to slope paleoenvironment. Finally Wilson et al. (2010) place the stromatolites into two different environments, the peritidal paleoenvironment for cyanobacteria, and the shelf to slope paleoenvironment for the iron- oxidizing bacterial microfossils similar to those found within the Gunflint and Biwabik

Iron Formations of the North American Animikie Basin.

Figure 5: Combination of the three depositional models. A drawing of each of the 3 depositional models suggested by Edwards et al. (2012), Plavanasky et al. (2009), and Wilson et al, (2010). The top drawing is a combination of the three models into one figure.

The Animikie Basin consists of a series of sediment depocenters that formed on the continental shelf of the North American Craton also known as Laurentia (Schulz and Cannon, 2007). The sediments include deposits from fluvial and marine deposits.

The fluvial deposits include conglomerate and sandstone layers. The marine deposits include siltstone, chert, and carbonate layers. Among the marine sediments deposits, there exist the stromatolite horizons that are the subject of this thesis. Presented below is the geologic setting of the Animikie Basin and the sedimentary deposits that make up the basin.

Geologic History

Iron-formations are Fe-rich sedimentary formations that formed mainly during the Archean and Paleoproterozoic in seas saturated with hydrothermally derived iron and silica (Pufahl, 2010). Iron-formations are categorized as Algoma and Superior types based on their tectonic regime (Gross, 1983). Algoma type iron-formations are typically in the Archean and are associated with volcanic arcs, are thin and typically only local in their extent (Pufahl, 2010). Superior type iron-formations are dominantly

Paleoproterozoic and are believed to be formed from continental shelves and platforms

(Pufahl, 2010). Paleoproterozoic iron-formations are believed to coincide with a rapid rise in atmospheric oxygen at 2.3 Ga known as the Great Oxygenation Event (GOE)

(Holland, 2002).

Laurentia, the North American Craton consists of a series of Paleoproterozoic orogenic belts that bound Archean crustal rocks (Schulz and Cannon, 2007). These

Paleoproterozoic orogenic belts formed as Archean micro continents and island arcs collided with Laurentia, building the continent up toward the south (Schulz and Cannon,

2007). Of note to this thesis is the Penokean Orogeny related to island arc collision with the southern border of Laurentia approximately 1.89 Ga (Schulz and Cannon, 2007). The

Penokean orogeny continued until 1.83 Ga as a fold and thrust belt that extends discontinuously for 1300 km along the southern margin of the North American Craton from Western Minnesota to eastern Ontario (Ojakangas et al., 2001).

The deposition of thick Superior iron-formations in the Animikie Basin has been recently debated. Schulz and Cannon (2007) argue that the Paleoproterozoic

24 sedimentary deposits were formed as in a back-arc basin prior to the fold and thrust belt that defined the Penokean orogeny approximately 1.85 Ga. Ojakangas et al. (2001), following Hoffman (1987), argue that the deposits were formed in a northern migrating foredeep created when the collision of island arcs formed a series of fold and thrust belts.

While the entire Animikie Basin follows the Superior Craton for 1300 km

(Ojakangas et al., 2001) this thesis narrows the area of study to the Biwabik Iron

Formation located in the Mesabi Iron Range of northeastern Minnesota. The Mesabi Iron

Range is 120 miles long and extends from Cass County, Minnesota east before finally terminating at the Duluth Complex northeast of Babbit, Minnesota near Lake Superior

(Severson et al., 2010; Fralick et al., 2002 ). The formations that collectively make up the Animikie Group in Minnesota include the Pokegama Formation, Biwabik Formation and the Virginia Formation (Severson et al., 2010). All three of the Animikie Group formations are conformable and are correlative to formations found in the Gunflint ranges of northernmost Minnesota and Ontario (Severson et al., 2010).

The Paleoproterozoic rocks of the Mesabi Iron Range in the Lake Superior region lie unconformably upon an Archean age platform that formed along the North

American Superior Craton. Deposition began around 1.89 Ga and ended with crustal uplift 1.75 Ga (Ojakangas et al., 2001). Age constraints for the three formations found within the Animikie Groups range from 2.125 to 1.93 Ga for the Pokegama Formation;

1.878 to 1.85 GA for the Biwabik Formation; and 1.85 to 1.836 Ga for the Virginia

Formation (Severson et al., 2010).

Pokegama Formation

The Pokegama Formation rests unconformably on the Kabetogama dike swarm. This dike swarm has been dated to 2.125 Ga using Rb-Sr isochron dating

(Southwick and Day, 1983; Beck, 1988). The minimum age date of 1.93 Ga was obtained from quartz veins that cut the Pokegama Formation (Hemming et al., 1996).

The Pokegama Formation is a group of clastic rocks originally named the

Pokegama Quartzite, but has since been renamed the Pokegama Formation due to the appreciable amounts of argillite and siltstone found within the formation (Severson et al.,

2010). The Pokegama Formation is broken up into three different members: the upper member composed of medium- to thick bedded orthoquartzite (Severson et al., 2009), so named because the quartzite consists of silica-cemented quartz sandstone (Severson et al., 2010); the middle member, characterized by interbedded siltstone and shale layers

(Severson et al., 2009); and a lower member consisting of shale with interbedded siltstone (Severson et al., 2009). The upper member is thought to be deposited in a subtidal high-energy environment, while the middle and lower members are thought to be deposited in a low-energy tidal flat environment (Severson et al., 2009).

The Pokegama Formation ranges from locally non-existent on the eastern end of the Mesabi Iron Range to more than 300 feet thick on the western end of the Mesabi

Iron Range (Severson et al., 2010). The contact between the Pokegama Formation and the Biwabik Iron Formation ranges from sharp to gradational, with chert layers found within the Pokegama Formation and detrital quartz grains present in the overlying iron- formation (Severson et al., 2009).

Biwabik Formation

The Biwabik Formation rests conformably on the Pokegama Formation with a maximum age date of 1.878 Ga obtained from U/Pb euhedral zircons located within an ash layer in the correlative upper Gunflint Formation (Fralick and Kissen, 1998; Fralick et al., 2002). The minimum age date of 1.85 Ga was obtained from zircons located within the Sudbury Impact ejecta layer near the top of the Biwabik Iron Formation and at the top of the Gunflint Iron Formation (Addison et al., 2005).

The Biwabik Formation was named for the iron deposits found within the formation (Severson et al., 2009). The Biwabik Formation is subdivided into four informal members known from top to bottom: Lower Cherty member, Lower Slaty member, Upper Cherty member, and the Upper Slaty member (Wolff, 1917). The cherty members are typically characterized by a granular (sand-sized) texture and thick bedding

(beds greater than or equal to several inches thick); while the slaty members are typically fine-grained (mud-sized) and thin-bedded (less than or equal to 1cm thick beds)

(Severson et al., 2009). The cherty members are mostly composed of chert and Fe-oxides with zones rich in iron silicates, while the slaty members are composed of iron silicates and iron carbonates with local chert beds (Severson et al., 2009).

Both the slaty and cherty members are considered to be interlayered with one rock type predominating in each given layer (Severson et al., 2010). The slaty members are thought to be deposited on the outer shelf as Fe-rich chemical mud in deep water/low energy zones as a result of either upwelling of Fe-rich waters (Morey, 1992) or a mixing of a stratified water column due to storm events (Pufahl and Fralick, 2004). The cherty

27 members are thought to be derived from shoreward moving tidal currents and storm events that disrupted the chemical mud and generated granules that were transported shoreward and reworked in shallow-water/high-energy environment to form the cherty members (Pufahl and Fralick, 2004; Ojakangas et al., 2005). The repetition of the slaty and cherty members was interpreted by White (1954) to be the result of transgressive and regressive ocean events. Two stromatolite-bearing intervals several meters thick, one of which is located at the base of the Lower Cherty member and the other in the middle of the Upper Cherty member, are used as diagnostic markers for the formation (Severson et al., 2010).

The Biwabik Formation is approximately 175-300 feet thick in the eastern end of the Mesabi Iron Range at Dunka Pit (Bonnichsen, 1968); 730-780 feet thick in the central Mesabi Iron Range/Virginia Horn area near Eveleth; and approximately 500 feet thick in the western Mesabi Iron Range near Coleraine and exhibits a termination about

15 miles southwest of Grand Rapids on the western end of the Mesabi Iron Range

(Marsden et al., 1968).

Virginia Formation

The Virginia Formation age is constrained to 1.85 -1.83 Ga from zircon ages in ash layers found a few inches from the base of the formation and near the top of the formation (Severson et al., 2010). The Virginia Formation is correlative with the Rove

Formation, found within the Gunflint Iron Range of Ontario (Severson et al., 2010).

Schulz and Cannon (2007) propose that there is a disconformity between the Biwabik and

Virginia/Rove formations that is indicative of a 40 million year hiatus in deposition of the two formations.

The Virginia Formation is a thick sequence of argillite, siltstone and greywacke and is divided into two different members--a lower argillaceous member and an upper silty and sandy member (Severson et al., 2009). The lower member is approximately 600 feet thick and contains intervals of black, thin-bedded, carbonaceous argillite with visible sulfides present in the eastern Mesabi Iron Range (Severson et al.,

2009). The carbonaceous argillite suggests deposition occurred for the lower member in deep anoxygenic conditions (Lucente and Morey, 1983). The upper member of the

Virginia Formation consists of argillite, but lacks the carbonaceous argillite found within the lower member (Severson et al., 2009). Instead the upper member contains interbeds of siltstone and fine-grained greywacke that are interpreted to be deposited by turbidity currents in a prograding submarine fan complex (Lucente and Morey, 1983).

The contact between the Biwabik Iron Formation and the Virginia Formation varies from sharp to highly gradational (Severson, et al., 2009). In most of the Mesabi

Iron Range the argillaceous sediments are in sharp contact with the carbonate horizon at the top of the iron-formation (Severson et al., 2009). In the western end of the Mesabi

Iron Range the contact between the Biwabik Iron Formation and the Virginia Formation is gradational, with the Virginia Formation containing several packages of thin-bedded carbonate iron-formation that are intricately interbedded with argillite and carbonaceous argillite, making the identification of where the contact is between the Biwabik Iron

Formation and the Virginia Formation difficult (Severson et al., 2009).

Mineralogy, Diagenesis, Metamorphism and Geochemistry

The mineralogy of the Biwabik Iron Formation tends to be divided between the two primary rock types, the fine-grained iron carbonates of the Lower and Upper

Cherty members and the fine-grained iron silicates of the Lower and Upper Slaty members (Schulz and Cannon, 2007). The Lower and Upper Cherty members are dominated by granular chert, siderite, and magnetite (Morey, 1992; Schulz and Cannon,

2007). The Upper Slaty member contains limestone and dolomite (Morey, 1992), while both the Lower and Upper Slaty members are dominated by greenalite, hematite, and magnetite (Schulz and Cannon, 2007). Within these rock units the mineralogical assemblages have been further broken down into three distinct types based on their initial deposition, their diagenesis, and lastly their level of metamorphism (French, 1973).

French (1973) used the following definitions based on petrographic observations and descriptions: original precipitate, primary mineral, and secondary mineral. The original precipitate is the original iron-bearing material deposited from solution, typically containing hydrous iron oxides, iron-bearing silica gels, and colloidal particles of iron-silicate and iron-carbonate minerals (French, 1973). The primary mineral is a texture that exhibits no detectable replacement of any pre-existing phase (French,

1973). Primary minerals may have been part of the original precipitate, or developed during early diagenesis from the original precipitate (French, 1973). The secondary mineral is a texture that appears to be replacing a primary or secondary mineral (French,

1973).

Four distinct chemical facies exist in the iron-formation based on their mineral content and original depositional conditions: the oxide, silicate, carbonate, and

3+ 2+ sulfide facies (James, 1954). The oxide chemical facies include magnetite (Fe 2Fe O4) and hematite (Fe2O3) (Cannon and Schulz, 2007). Hematite is regarded as a primary mineral, while magnetite is considered a secondary mineral that forms by the replacement of pre-existing silicates and carbonates during diagenesis (French, 1973). Frost et al.

(2007) recognized three processes by which magnetite forms: (a) magnetite reduced from hematite, (b) formed by the oxidation of ferrous iron in siderite and greenalite, or (c) formed by oxygen-conserving reactions involving hematite with siderite or greenalite.

Below is the reaction for the formation of magnetite reduced from hematite according to

Frost et al. (2007):

6Fe2O3 + C → 4Fe3O4 + CO2 (6)

Magnetite formed this way is fine-grained and associated with hematite (Frost et al., 2007). Most of the magnetite in the Biwabik Formation, according to Frost et al.

(2007), did not form this way, as most of the magnetite forms as growth on siderite or greenalite grains. The oxidation of ferrous iron would have a reaction according to Frost et al. (2007) in one of two ways:

12FeCO3 + 2H2O→ 4FeO4 + 11CO2 + Ch4 (7)

4Fe3Si2O5(OH)4 + CO2 → 4Fe3O4 +8SiO2 + 6H2O + CH4 (8)

These reactions result in magnetite veins and are believed to be biotically mediated due to the abundance of kerogen found in the Biwabik Formation and the presence of microfossils found within the formation (Frost et al., 2007). Frost et al.

(2007) considers magnetite to be part of late diagenesis along with the replacement of greenalite by quartz and the replacement of greenalite by carbonate. The reaction of hematite with siderite and greenalite formed one of two ways:

FeCO3 + Fe2O3 → Fe3O4 + CO2 (9)

Fe3Si2O5(OH)4 + 3Fe2O3 → 3Fe3O4 + 2SiO2 + 2H2O (10)

Frost et al. (2007) suggest that magnetite formed from siderite and greenalite via these reactions formed in response to regional burial metamorphism at temperatures around 200oC. Perry et al. (1973) suggest that some of the magnetite found in the

Biwabik Iron Formation may in fact be primary, as it does not contain the euhedral shape that is indicative of secondary replacement and is instead fine-grained like the hematite, and that it may have formed during recrystallization (Perry et al., 1973).

Hematite is mostly fine-grained and occurs as either disseminated grains in microcrystalline chert or as concentric layers in oolites and algal structures (Perry et al.,

1973). Hematite may be an early diagenetic product of the crystallization of Fe(III) oxyhydroxides (Frost et al., 2006).

2+ 3+ The silicate chemical facies include quartz (SiO2), greenalite (Fe ,Fe )2-

2+ 3Si2O5(OH)4 , minnesotaite (Fe ,Mg)3Si4O10(OH)2 , and stilpnomelane

2+ 3+ [K(Fe ,Mg,Fe )]8(Si,Al)12(O,OH)27n(H2O) (Schulz and Cannon, 2007). Greenalite is a primary mineral, although it is unknown whether greenalite formed as an original precipitate or was formed due to early diagenesis (French, 1973). Minnesotaite is typically considered a secondary replacement mineral of greenalite or siderite and occurs as sheaves and needles (Perry et al., 1973; French, 1973). Minnesotaite replaces

32 greenalite when coupled with quartz in the reaction greenalite+quartz=minnesotaite+water (Perry et al., 1973). Minnesotaite replaces siderite by the reaction siderite+chert+water=minnesotaite+carbon dioxide (Perry et al., 1973).

Stilpnomelane may be primary, but the majority occurs in the slaty-type beds as aggregates of many small grains that have been recrystallized from an earlier phase

(Perry et al., 1973). The predominance of stilpnomelane in the slaty-type beds results from the higher content of aluminum, potassium, sodium and titanium found within the slaty beds compared to the cherty-type beds (Lepp, 1966). Both the minnesotaite and the stilpnomelane are attributed to low-grade metamorphism in the Biwabik Iron Formation by the intrusion of the Duluth Complex (Perry et al., 1973; Frost et al., 2007). Quartz is considered a secondary texture due to early diagenesis of greenalite (Frost et al., 2007) by this reaction:

+ ++ Fe3Si2O5(OH)4 + 6H → 2SiO2 + 3Fe + 5H2O (11)

Quartz could also be a primary texture having formed directly from siliceous gel into original chert granules and ooids, making them potentially an original precipitate

(French, 1973;Schulz and Cannon, 2007).

The carbonate chemical facies include calcite (CaCO3), siderite (FeCO3), and

2+ ankerite [Ca(Fe ,Mg,Mn)(CO3)2] (Schulz and Cannon, 2007). Calcite and dolomite-rich beds are found at the top of the Biwabik Iron Formation (Perry et al., 1973). According to

Perry et al., (1973) there are no secondary minerals present in the calcite and dolomite beds, but the variation in grain size suggests recrystallization, making calcite a secondary texture according to French (1973). Siderite is believed to be a primary and secondary

33 texture (Perry et al., 1973; French, 1973). As a primary texture, siderite occurs in both the cherty and slaty-type beds as laminae and as microcrystalline to coarsely-crystalline grains (French, 1973; Perry et al., 1973). As a secondary texture, siderite appears to replace fine-grained primary siderite and primary silicates such as greenalite, as evidenced by the replacement of granules by large single crystals of siderite during diagenesis (French, 1973). Frost et al. (2007) wrote that kerogen found on carbonate grain boundaries suggest the formation of siderite possibly involved microbes during diagenesis. The two possible reactions for the formation of siderite due to microbes are:

- 2+ - - CH3COO + 8Fe(OH)3 → 8Fe + 2HCO3 + 15OH + 5H2O (12)

- 2+ - OH + Fe + HCO3 → FeCO3 + H2O (13) (Frost et al., 2007)

Both reactions are microbially mediated with a high pH at the surface of the

Fe(III)hydroxide that in conjunction with the high bicarbonate content facilitates the precipitation of siderite (Frost et al., 2007).

Ankerite tends to form as a secondary texture as coarse-grained aggregates that completely replace the primary minerals within the iron-formation (French, 1973).

Ankerite replaces siderite and greenalite within the iron-formation during diagenesis

(French, 1973). No references were found within Perry et al., 1973; French, 1973; or

Schulz and Cannon, 2007 attributing the sulfide chemical facies to the Biwabik Iron

Formation.

James (1954) provided four chemical facies for all banded iron-formations; however, these chemical facies only include the minerals found within banded iron- formations that experience up to low-grade metamorphism. Additional minerals form

34 from the primary precipitates as they experience medium-to- high grade metamorphism.

According to French (1973), calcite is typically found within the Biwabik Formation near the Duluth Complex, where medium-to-high-grade metamorphism occurred. Klein

(1973) states that the presence of amphiboles and pyroxenes in the Biwabik Iron

Formation is due to medium-to-high-grade metamorphism and the chemical potential of

CO2, and that if the chemical potential of CO2 is low enough in the rocks then carbonate breaks down, and the silica that formed either as primary chert or as low-grade metamorphic quartz will react with the carbonate to form new silicates as well as

2+ 2+ pyroxenes (ferrohypersthene/(ferrosilite [Fe MgSi2O6] and hedenbergite CaFe [Si2O6])

2+ and amphiboles (grunerite [Fe 7Si8O22(OH)2] and cummingtonite [Mg7Si8O22(OH)2]).

The reactions for some of these new minerals are:

Ca(Fe,Mg)(CO3)2 + 2SiO2 → Ca(Fe,Mg)Si2O6 + 2CO2 ankerite hedenbergite (Klein, 1973) (14).

(Fe, Mg)CO3 + SiO2 →(Fe, Mg)SiO3 + CO2 siderite ferrohypersthene (Klein, 1973) (15).

7Ca(Fe,Mg)(CO3)2 + 8SiO2 + H2O → (Fe,Mg)7Si8O22(OH)2 + 7CaCO3 + 7CO2 ankerite grunerite calcite (Klein, 1973) (16).

8 (Fe,Mg)CO3 + 9 SiO2 + H20 → (Fe,Mg)7Si8O22(OH)2 + (Fe,Mg)SiO3 + 8CO2 siderite grunerite ferrohypersthene (Klein, 1973) (17)

2+ 2(Fe, Mg)7Si8O22(OH)2→7Fe 2(SiO4) +9SiO2 + 2H2O grunerite fayalite (Valass-Hyslop et al.., 2008) (18)

2Fe3O4 + 3SiO2 → Fe2SiO4 + O2 magnetite fayalite (Valass-Hyslop et al.., 2008) (19)

2+ 2Fe (CO3) + SiO2 → Fe2SiO4 + 2CO2 siderite fayalite (Valass-Hyslop et al.., 2008) (20).

Frost et al. (2007) used phase relations to calculate the maximum temperatures at which the different minerals formed during metamorphism. Samples containing greenalite could not have exceeded temperatures of 300oC without greenalite changing to minnesotaite (Frost et al., 2007). Samples containing grunerite could not exceed 300- 340oC, as most of the grunerite within the Biwabik Iron Formation formed from siderite (Frost et al., 2007). Grunerite that forms from minnesotaite would have a higher temperature maximum (Frost et al., 2007). Samples containing hedenbergite would have temperatures between grunerite (340oC) and fayalite which forms at around

500-550oC (Frost et al., 2007). Samples that contain coexisting olivine, orthopyroxene and grunerite would obtain temperatures of approximately 725-740oC (Frost et al., 2007).

Geochemistry is often used to provide evidence of diagenetic and post- diagenetic processes affecting the rocks (Perry et al., 1973). δ56Fe values provide information about peak metamorphic temperatures along with mineral sequences (Frost et al., 2006). This section highlights the isotopic information within the Biwabik Iron

Formation for δ56Fe, δ13C, and δ18O. Along with the information for peak metamorphic temperatures, δ56Fe isotope compositions of hematite, magnetite, iron-carbonates, and pyrite can reflect a combination of (1) mineral-specific equilibrium isotope fractionation,

(2) variations in the isotope compositions of the fluids from which they were precipitated,

36 and (3) the effects of metabolic processing of Fe by bacteria (Johnson et al., 2003). The

56 o δ Fe isotope variation for magnetite-rich banded iron-formations is 1.3 /oo (Johnson et al., 2003). The iron carbonate layers within the Biwabik Iron Formation have the lowest bulk δ56Fe values compared to the magnetite rich layers which have the highest bulk

δ56Fe values, and the silicate layers have δ56Fe values in between the two other values

56 o o 56 (Frost et al., 2007). The δ Fe of the carbonate layers is -0.64 /oo to -0.74 /oo, the δ Fe

o values for the silicate layers is -0.11 /oo for greenalite, due to the abundant isotope heavy magnetite found within the silicate layer (Frost et al., 2007). The δ56Fe values for magnetite differ depending on whether the magnetite was located within the carbonate- rich layers or the silicate-rich layer (Frost et al., 2007). The magnetite located within the

56 o carbonate-rich layers contains δ Fe values of -0.14 /oo, while the magnetite from the

56 o silicate-rich layers contains δ Fe values of +0.69 /oo (Frost et al., 2007). The heterogeneity of the δ56Fe values suggests that the iron-formation samples either (1) have protoliths with varying δ56Fe¸such as the carbonate-rich iron-formation possibly having a bulk composition that is lighter than the silicate-rich iron-formation, or (2) that the subsequent diagenetic and/or metamorphic processes have redistributed iron throughout the Biwabik Iron Formation (Frost et al, 2007).

δ56Fe values of the magnetite located in the higher-grade contact-

56 o o metamorphism samples vary significantly from a δ Fe value of +0.19 /oo to +0.82 /oo

(Frost et al., 2007). The δ56Fe values of the metamorphic iron-silicates are δ56Fe= -

o o o 0.18 /oo for olivine, +0.57 /oo for grunerite, and + 0.21 /oo for hedenbergite (Frost et al.,

2007). Carbonates that have been completely consumed except for calcite near the

37 higher-grade contact-metamorphism zone (Frost et al., 2007). Calcite within this zone has

56 o a δ Fe value of 0.00 /oo (Frost et al., 2007).

δ18O values differ within the Biwabik Iron Formation depending on whether the rocks have been in contact with the Duluth Complex (Perry et al., 1973). According to Perry et al (1973) there is evidence of primary δ18O values related to the water from which the sediments were precipitated, as well as a set of δ18O values established during the time of diagenesis or low-grade regional metamorphism. The highest δ18O value for

o chert is +23.5 /oo, and it was found in the co-equivalent Gunflint Formation (Perry et al.,

1973). This δ18O value is used as a reference value for evaluating diagenetic and/or metamorphic changes within the Biwabik Iron Formation as this value represents the

18 primary δ O content of the SiO2 precipitated within the Animikie depositional basin

18 (Perry et al., 1973). Perry et al. (1973) compared the δ OSiO2 values of three difference

18 o o cores and obtained a δ O range of 18.8 /oo – 20.5 /oo. The variation between the cores of

18 the δ OSiO2 values suggests that the samples underwent diagenesis or low-grade metamorphism and the values were lowered by an exchange, probably with iron-rich

18 minerals, as the values are below the reference δ O SiO2 of the samples found within the

Gunflint Formation (Perry et al., 1973).

δ18O relations in carbonates are more variable in their oxygen isotope composition than in the quartz, and this suggests that the carbonates continued to react

18 with their environment after the δ O values in the quartz were fixed (Perry et al., 1973).

18 Perry et al. (1973) found that the carbonates in three of their cores had δ O values that

38 correlated with the presence of magnetite, suggesting that isotopic exchange between the two minerals continued to take place.

13 o Most marine carbonates, regardless of age, have δ C values near 0 /oo (Perry et al., 1973). Precambrian iron-formations have distinctly low δC13 values (Perry et al.,

13 o 1973). δ C values of siderite and ankerite range from -2 to -19 /oo (Perry et al., 1973).

This variation is due to the presence or absence of magnetite within the sample; samples containing magnetite tend to have lower δ13C values than those that are magnetite-free

(Perry et al., 1973). δ13C values of reduced carbon show much less variation (Perry et al.,

13 o 1973). The average δ C value for the reduced carbon samples is -33.1 /oo (Perry et al.,

1973). The variability of δ13C of the carbonates compared to the consistency of the reduced carbon of the coexisting organic matter implies that both did not attain their present isotopic composition during formation from bicarbonate ions in the waters of the depositional basin (Perry et al., 1973). Either the carbonates or the reduced carbon was altered after deposition, unless the two substances formed at different depths in the water

- 13 having HCO3 stratified with respect to δ C (Perry et al., 1973).

Mineralogy of the stromatolite horizons located within the Upper and Lower

Cherty members has not been done to date. Obtaining this information is the main goal of this thesis so that a depositional environment may then be inferred for the stromatolite horizons based on the mineralogy and facies.

CHAPTER III

METHODS

Two separate stromatolite horizons exist within the Biwabik Iron Formation

(figure 6). Below is a description of each of the lithologic units observed within the stromatolite horizons. Lithologic descriptions of the two stromatolite horizons were completed using photographs, stratigraphic columns, hand samples, and core cuttings.

Photos and core cuttings were obtained August, 2013 from Natural Resources Research

Institute (NRRI) located at 5031 Miller Truck Highway, Duluth, Minnesota. Eleven cores were analyzed that contained either the upper stromatolite layer, the lower stromatolite layer, or both layers.

Thin (30µm) sections were made of the rock samples and observations were made using standard petrographic microscopy techniques on a BX-51 petrographic microscope. Sedimentary, diagenetic and metamorphic minerals and structures were measured using Olympus STREAM software in conjunction with an Olympus camera.

CHAPTER IV

RESULTS

Lithofacies

Two of the eleven cores were selected to represent the upper and lower stromatolite horizons located within the Biwabik Iron Formation. Core # 24974

(Appendix A) is the representative core for the lower stromatolite layer (figure 7) and core # 24986 (Appendix A) is the representative core for the upper stromatolite layer

(figure 8).

Figure 6: A generalized stratigraphic column of the Biwabik Formation. Generalized stratigraphic column of the Biwabik Formation showing the two stromatolite sections. The lower stromatolite section is located within the Lower Cherty member, and

41 the upper stromatolite section is located within the Upper Cherty member. Figure from Russell Shapiro, unpublished NASA Exobiology Proposal.

Figure 7: Stratigraphic column of the Lower Cherty stromatolite layer.

Figure 8: Stratigraphic section of the Upper Cherty stromatolite layer.

Coarse-grained Sandstone

This lithologic unit contains sub-rounded to rounded quartz clasts, sub- millimeter in scale surrounded by quartz cement. The sandstone is grain-supported and appears to be well sorted. Quartz is dominant in this unit, with 75% of the unit composed

43 of quartz. The remaining 25% of the mineral composition consists of rounded, green chert and disseminated pyrite grains sub-millimeter in scale. Disseminated pyrite percentages increase closer to the next lithologic unit. The lower portion of the unit contains flat-laminar beds along with cross-bedding, while the upper portion of the unit contains horizontal to sub-horizontal layers of ripple laminations.

Coarse-grained sandstone beds located within other cores and hand samples contain coarse-grained sandstone layer with sub-rounded to sub-angular grains. The minerals observed within other coarse-grained sandstone grains include hematite, magnetite, black microquartz (chert), and ankerite. The predominance of quartz remains throughout all the samples with just the minor mineral compositions changing.

Stromatolite

Stromatolites are observed in all core sections including the core sections close to the Duluth Igneous Complex. The stromatolite lithologic unit contains psuedocolumnar stromatolite fragments with high laminae inheritance. Laminae are consistent in continuity and thickness with wavy laminae that are gently to steeply convex. The stromatolite fragments have parallel accretion of couplet patterned white quartz and green/black chert laminae. Near the top of core #24986, the psuedocolumnar stromatolite laminae consist of hematite and ankerite. Disseminated pyrite preserves laminae in psuedocolumnar stromatolitic fragments within some of the cores and hand samples. Laminae are sub-millimeter in scale. Green chert intraclasts and angular magnetite grains are sub-millimeter to millimeter in size and are observed between lamina along with ooid, and peloids, in laminae vacuoles and in the interspaces The

44 grains are sub-angular to angular and the peloids and ooids are sub-rounded to rounded.

In the upper stromatolite layer, the psuedocolumnar stromatolites are intact with laminae of high inheritance. Laminae contain light quartz and dark magnetite banding.

Undulatory stromatolites are observed in core especially in the upper stromatolite horizon. The stromatolites are 4 cm in height and range from 9 mm to 23 cm in width. Laminae contain high inheritance and low synoptic relief. Laminae are sub- parallel and exhibit parallel accretion of couplet patterned quartz and green chert.

Occasionally the couplet patterns are green chert, quartz, and ankerite. All the grains between the laminae are very-fine-grained. Intercalated with the undulatory stromatolite layers are sub-parallel layers of coarse-grained sandstone. The sandstone layers are grain- supported and the grains are sub-rounded to rounded and well sorted.

Dendritic stromatolites contain gamma branching along with high inheritance laminae. The dendritic stromatolites contain high synoptic relief ranging from 2-7 mm in height. Columns range from 5-12 mm in width and 10-18 mm in height. Laminae contain parallel accretion with couplet patterns of green chert, and quartz. Other samples contain couplets of green chert, magnetite and ankerite. Grains between laminae are very-fine grained. Columns are steep-walled, overhanging with occasional bridging. Breccia clasts, flat-pebble conglomerate, oncoids and ooids are observed in the interspaces between columns. Dendritic stromatolites are observed overlying psuedocolumnar stromatolites, oncoids and flat-pebble conglomerate layers.

Medium-grained Sandstone

This unit contains medium sized, rounded to sub-rounded sand grains. Grains are well sorted. Sub-parallel bands of darker sand grains, black chert in some specimens and magnetite in others are present in this lithofacies. The sandstone is hematite-rich and contains black chert, green chert, ankerite and white quartz. Where the unit is hematite poor, magnetite replaces the hematite as the dark mineral grain. Green chert grains are the predominant mineral type in cores #24986 and white and grey quartz grains are the predominant mineral type for core #24974. Disseminated and pyrite cubes are also present in high amounts within the unit. Cross-bedding is preserved in this unit as darker layers within the medium-grained sandstone.

Very Fine-grained Sandstone

This lithologic unit consists of well rounded, fine-grained sand. The unit is layered and the layers are sub-horizontal and sub-millimeter with alternating black magnetite grains and green chert. This unit also contains disseminated pyrite grains observed in localized pockets and between layers. Mining geologists within the banded iron-formation commonly call this unit taconite (Severson et al., 2009). Very fine- grained sandstone within the Upper Slaty member overlies green chert intraclasts. The intraclasts are angular to sub-round. The layers are sub-parallel and of varying thickness ranging from 1 mm in width to 2.54 cm in width for the whole unit.

Intercalated Medium- and Very Fine-grained Sandstone

This lithologic unit contains layers of medium-grained sandstone that are alternating with very fine-grained sandstone layers. The layers range from sub-millimeter to 6 mm in scale. The unit is stratified. The medium-grained sandstone layers are sub-

46 horizontal. The grains are sub-rounded to sub-angular, well sorted and grain supported.

The mineral composition of the medium grained sandstone is quartz, hematite, magnetite and green chert. Intraclasts of black chert are observed in medium-grained sandstone layers. The very fine-grained sandstone dominates the unit with a larger layer width of 6 mm to 2 cm. Grains are sub-rounded to round and grain supported. The very fine-grained sandstone layers are sub-horizontal. The mineral composition of the very-fine-grained sandstone layers is magnetite and minor amounts of pyrite. Occasional cubic pyrite lenses occur within the very-fine grained sandstone layers.

Brecciated Intraclasts

This lithologic unit contains brecciated intraclasts of ankerite that are intercalated with fine-grained magnetite and disseminated pyrite. The brecciated clasts are intercalated with magnetite, chert, and quartz sandstones. Some of the sandstone layers are very-fine grained, while other layers are coarse-grained to medium-grained sandstone layers. The sandstone layers are sub-horizontal with climbing ripple laminations, and cross-bedding observed in the unit depending on the sample observed.

Brecciated ankerite intraclasts range from sub-millimeter to 3 mm in size.

Brecciated intraclasts of black chert are cemented with minor amounts of calcite and major amounts of quartz. These black chert clasts are intercalated with the ankerite clasts.

Black chert intraclasts range from 1 mm to 4 mm in size. Ankerite clasts are located near fine-grained magnetite sandstone. The rest of the lithologic unit is composed of black chert intraclasts that increase in size up to 1 cm. The black chert intraclasts remain cemented with quartz. A loss of disseminated pyrite occurs once the predominance of

47 quartz-cemented black chert begins. Brecciated clasts are angular to sub-rounded magnetite with calcite cement. Clasts range from sub-millimeter to 8 millimeter in size.

Disseminated pyrite grains are observed in minor amounts throughout the lithologic unit.

Brecciated clasts of red chert sub-millimeter to 2 centimeters intercalated with medium-grained sandstone found in some cores. The brecciated clasts become smaller in size going up the core sample until it becomes coarse grained hematite sandstone. Other cores contain black chert intraclasts sub-millimeter in scale, angular to sub-round in shape. Some of the intraclasts are very well rounded and are possibly ooids. Intraclasts are surrounded by quartz cement. Brecciated clasts of green chert that are surrounded by calcite cement are observed in the lower stromatolite units of some of the cores.

Stromatolites with Flat-pebble Conglomerate and Oncoids

This lithologic unit contains stromatolites intercalated with flat-pebble conglomerate. The stromatolite laminae contain high inheritance and couplet patterns of light quartz and dark iron oxide bands. Laminae are sub-millimeter in scale and contain consistent continuation, low synoptic relief, and no walls. Some of the psuedocolumnar stromatolites are fragments intercalated amongst the flat-pebble conglomerate. The pebble clasts are sub-angular to sub-rounded and range in size from 2 mm to 3 cm in diameter. The conglomerate grains are poorly sorted and are surrounded by coarse- grained grain supported sands. The whole unit is magnetite rich with minor amounts of disseminated pyrite present. Possible grains of ankerite sub-millimeter in scale are present. Flat pebble conglomerate clasts are angled with no specific orientation, and do

48 not lie horizontal to bedding plane where they are overlain or overlie undulatory or psuedocolumnar stromatolites.

Where the flat-pebble conglomerate clasts lay between the dendritic stromatolites columns, the conglomerate clasts are stacked atop one another with addition sand grains in between the conglomerate clasts. Ooids are also found intercalated with the flat-pebble conglomerate layers along with oncoids. Both the ooids and the oncoids are sub-rounded to round.

The oncoids contain wavy, ragged laminae and range in size from 6 mm-1.6 cm and are non-magnetic. The average oncoid length is 10-11 mm. Most of the oncoids are composed of hematite and contain shrinkage cracks due to diagenesis. Shrinkage cracks are filled with secondary quartz. Oncoid cores appear to be intraclasts, although some of the oncoid cores contain fragments of laminae possibly from stromatolites. The oncoid cores measure 2-5 mm in diameter. Oncoids are surrounded by coarse-grained sandstone clasts composed of quartz. The matrix is poorly sorted and grain supported.

Stromatolites with Ooids

This lithologic unit contains stromatolites with ooids and intraclasts in the interspaces and atop the stromatolites. The stromatolites are psuedocolumnar, domal, and dendritic. Laminae have high inheritance and are sub-millimeter to 2 mm in size.

Dendritic stromatolite columns overhang with thin margins that bridge. Ooids and possibly peloids located in between laminae are observed in some of the core sections are not composed of the same grains that make up the stromatolite laminae. Stromatolite laminae are composed of green chert, magnetite and smaller amounts of ankerite. The

49 ankerite laminae are not as common as the green chert and magnetite laminae. Euhedral ankerite crystals along with euhedral calcite are visible in core samples.

In beds where there are interspaces present (dendritic stromatolites), green chert and magnetite intraclasts and green chert ooids are present. In other beds where no interspaces are present, such as in domal and undulatory stromatolites the green chert ooids and magnetite intraclasts overlie the stromatolites. Blocky crystals of calcite and ankerite are observed along with the ooids and stromatolites.

The dendritic stromatolites are atop psuedocolumnar stromatolites observed in core samples from the Lower Cherty member. Laminae in these specimens consist of alternating red and black chert bands. Minor amounts of disseminated pyrite are present.

Dendritic stromatolites in the Upper Cherty member are composed of green chert, ankerite, calcite, and minor amounts of pyrite. Dendritic stromatolites in the Upper

Cherty member grow atop flat-pebble conglomerate instead of atop psuedocolumnar stromatolites.

Intraclasts are visible atop the psuedocolumnar stromatolites and sometimes between laminae and interspaces on the dendritic stromatolites. Clasts are angular and composed of red and black chert and are surrounded with calcite cement. Intraclasts range from sub-millimeter in scale to 8 mm. Fine-to coarse-grained ooids and peloids are present along with intraclasts.

Stromatolites with Peloids

In this lithologic unit, the stromatolites contain magnetite peloids. The magnetite peloids are found in the interspaces in dendritic stromatolites and are near or

50 on top of the domal, undulatory, and pseudocolumnar stromatolites. All types of stromatolites have wavy, sub-millimeter laminae with high inheritance, consistent continuity and even fabric. The laminae have parallel accretion with couplet patterns of light and dark laminae Mineral composition of the stromatolites and the peloids varies upon location of the sample. Closer to the Duluth Igneous complex the stromatolite lamina are composed of quartz and magnetite. Further away from the complex the lamina are composed of hematite and ankerite. Peloids near the complex are composed of hematite and quartz along with minnesotaite fibers. Peloids away from the complex contain amorphous greenalite. Some dissolution of the peloid interiors occur and the rims are maintained by cubic hematite crystals.

Stromatolites associated with Intraclasts

This lithologic unit consists of domal stromatolites and dendritic stromatolites.

The domal stromatolites have high inheritance, wavy laminae with continuous continuity, sub-millimeter in size and most are fractured. Domal stromatolites are typically larger than the core sample of 4 cm, so exact size of dome cannot be determined. Height of the domal stromatolites averages 6 mm.

The dendritic stromatolites also contain high inheritance, wavy laminae, and the average stromatolite width is 1mm to 2 cm. The average height of the dendritic stromatolites is 8mm. Laminae are consistent in continuity with parallel accretion and couplet patterns of light and dark laminae that are similarly sized. The unit is cemented with calcite and contains void intercalation fabric.

Petrographic Fabrics

To obtain better constraints on the depositional environments for the two stromatolite horizons, thin sections of samples from the various mines located within the

Biwabik Iron Formation and Gunflint Formation were used. Petrographic fabrics observed in the thin sections provide further evidence of depositional environments that cannot be viewed in hand specimens and core samples.

Intraclasts

Intraclasts have an average length of 1066.42 µm and an average width of

377.63 µm. The maximum length for an individual intraclast is 3676.26 µm, and the minimum length is 116.11 µm. Intraclasts are sub-angular to sub-rounded and located in the void spaces between stromatolite columns, and where stromatolite columns do not exist, atop the undulatory or flat-laminae stromatolite layers. Intraclasts located between stromatolites columns exhibit horizontal stacking of grains (Figure 9). Intraclasts contain greenalite, microcrystalline quartz and hematite (Figures 9 &10).

Figure 9: Coated intraclasts. Coated intraclasts composed of greenalite and hematite surrounded by microcrystalline quartz. G=greenalite; H=hematite. Plain Polarized Light (PPL).

Figure 10: Intraclasts and flat pebble conglomerate. Intraclasts of breccia and flat pebble conglomerate that are composed of fibrous greenalite and euhedral hematite grains, The matrix is composed of greenalite and carbonate minerals. G=greenalite; C= carbonate; H= hematite. PPL.

Peloids

Peloidal grains have an average length of 494.18 µm and an average width of

268.91 µm. The maximum peloid length is 1296.9 µm, and the minimum length is 136.4

µm. The interiors of the peloidal grains are replaced with amorphous greenalite or fibrous minnesotaite/cummingtonite (Figure 11). Dissolution is common in peloidal grains.

Many peloids in thin section retain outer rims; however, the interior has been dissolved

(Figure 12). Peloids are separated from intraclasts based on the rounding of the grains and size. Peloids were observed to be sub-rounded to rounded and of a smaller size then the intraclasts. Many peloids still retained a core of amorphous minerals, mostly of greenalite and hematite rather than fibers as observed in the intraclasts.

Figure 11: Dissolved peloids. Peloids with rims preserved in hematite and amorphous greenalite interiors. Diagenesis and metamorphism caused dissolution in some of the grains. Peloids surrounded by microcrystalline quartz. Q= quartz; G= greenalite; H=hematite. PPL.

Figure 12: Dissolved ooids. Image is an example of the dissolution of the ooid interiors along with the deformation of the ooid. PPL.

Ooids

Ooids observed in thin section all contain concentric laminae. Ooids average length is 789.69µm and average width is 567.48µm. The maximum ooid length is

1185.28µm with a minimum length of 383.23µm. The maximum ooid width is

1019.83µm with a minimum width of 241.21µm. Ooids are densely packed within the interspaces between stromatolite columns or between laminae, or atop the stromatolites in non-columnar stromatolites. Most of the ooid nuclei have been replaced with microcrystalline quartz; however, the nuclei that do remain intact are peloidal or lithoclast in nature. Most of the nuclei appear to be peloidal with the peloidal mud replaced by secondary greenalite, or minnesotaite (Figure 11). The typical ooid shape is asymmetrical or elliptical.

Cortices are homogenous, and singular nuclei compose nearly fifteen percent of the ooid structure compared to the eighty-five percent that is the ooid cortex (Table 4

Appendix B). Cortex ratios decrease compared to nuclei in more metamorphosed samples

(Figure 12). Deformation of the ooids (Figures 12 & 13) is observed in all of the photographs taken with ooids present. Re-generated ooids along with broken ooids were observed in thin section (Figure 14). Cortices are preserved by iron-oxide and the occasional iron-dolomite minerals (Figure 15).

Figure 13: Shrinkage cracks due to diagenesis within ooid cores. PPL.

Figure 14: Dissolution of coated intraclasts and deformed ooids. PPL.

Figure 15: Hematite replacement of microbial ooid core Hematite replacement of microbial ooid core. Ooids compacted within the interspaces of the stromatolites columns. PPL.

Many of the ooids are only visible due to the replacement of the cortex or rim of the grain by ankerite, calcite, chert, or hematite/magnetite (Figure 9). Delamination or the peeling back of the cortice laminae is observed, though not as common as intact cortices (Figure 16).

Figure 16: Delamination of the ooid cortice. Delamination of the ooid cortice observed. Ooid core retained with secondary mineralization of greenalite (green amorphous) and minnesotaite (red/brown fibers). G= greenalite; M= minnesotaite. PPL.

Stromatolites

Stromatolite terminology used below is based on the Handbook for the Study of Stromatolites and Associated Structures (Second Draft) by Kathleen Grey (1989).

Stromatolites consist of flat-laminar, undulatory, psuedocolumnar, digitate, and

58 nodular/bulbous forms. Laminae thickness ranges from 32.16µm to 1479µm with an average thickness of 271.40µm. Laminae contain moderate to high inheritance, are undulatory and continuous and alternate between light and dark couplet laminae.

Columnar stromatolites (digitate, nodular) contain thin, continuous laminae walls with no margins along with bridging between columns (Figures 17, 18 & 19). Stromatolite laminae are gently to moderately convex (Figure 12).

Figure 17: Columnar stromatolite. Columnar stromatolite with high synoptic relief, wavy laminae composed of carbonates. C= carbonate. Located in Kakabeka Falls Upper Cherty member. PPL.

Figure 18: Digitate columnar stromatolite. Digitate columnar stromatolite with high synoptic relief, wavy laminae and laminae composed of hematite and carbonate. H= hematite; C= carbonate. Located in Mink Mountain, Upper Cherty member. PPL.

Figure 19: Digitate columnar stromatolite. Digitate columnar stromatolite with wavy laminae composed of hematite and microcrystalline quartz. H= hematite; Q= quartz. Located at McKinley Mine, Upper Cherty member. PPL.

Stromatolites are observed from low metamorphic sections of the Biwabik

Formation all the way to the high-alteration metamorphic areas located near the Duluth

Complex. Laminae are composed of red chert, black chert, ankerite, hematite, magnetite and minor amounts of pyrite. Secondary calcite observed between some of the laminae as void intercalation (Figures 20 & 21) along with greenalite (Figure 22), ankerite and Fe- oxides (Figure 23). Laminae are consistent in continuity with occasional irregularly variable continuity. Laminae are typically parallel to one another.

Figure 20: Void. Void in between stromatolite lamine. V= void. PPL.

Figure 21: Void with fill. Void filled with greenalite and microcrystalline quartz. G= greenalite; Q= quartz; V= void. PPL.

Figure 22: Void with calcite. Void filled with splayed calcite cement and hematite. Ca= calcite; H= hematite; V= void. PPL.

Figure 23: Laminae with Fe-oxide. Laminae edges showing the ankerite and hematite filling the void space. A= ankerite; H= hematite. PPL.

Columnar stromatolites contain overhanging of the columns with high synoptic relief (Figures 17, 18, & 19). Psuedocolumnar and nodular stromatolites exhibit moderate synoptic relief, and the flat-laminar and undulatory stromatolites exhibit low synoptic relief. Ooids and intraclasts are observed between laminae in nodular and psuedocolumnar stromatolites (Figure 17). Stromatolite width ranges from 2.86 mm to

6.06 mm with an average width of 4.66 mm.

Digitate stromatolite branches exhibit moderate divergence between branches

(Figure 18) with gradual widening of the original column prior to branching (Figure 24).

Attitude of the stromatolite columns is inclined with uniform variability. Digitate stromatolites exhibit lobate plan outline (Figure 12).

Figure 24: Gamma branching stromatolites. Core # 24986 exhibiting gamma-branching of digitate stromatolites and lobate plan outline. Gradual widening of the column observed prior to branching. Located south of LTV Mine in the Upper Cherty member.

Bladed/fibrous cement

Partially dissolved peloidal grains surrounded by bladed/fibrous cement composed of calcite crystals. Remainder of void space filled with micro-crystalline quartz. Bladed calcite cement forms between peloidal grains (Figure 25) filling the void spaces between grains. Bladed iron-dolomite crystals grow off of iron-oxide grains into void spaces located between grains (Figure 23).

Figure 25: Peloids with bladed calcite. Peloidal grains and intraclasts surrounded by bladed calcite cement and microcrystalline quartz. Ca= calcite; Q= quartz; Ch= chert. PPL.

Peloidal Cement

Peloidal cement is observed along with micro-crystalline quartz cement.

Peloidal cement composed of fibrous minnesotaite/cummingtonite fibrous crystals, greenalite fibers and amorphous cement (Figure 26).

Figure 26: Peloids with cement. The peloidal cement in this image is composed of amorphous greenalite and peloidal clay. Fibrous minnesotaite minerals are surrounding peloidal grains and intraclasts. G= greenalite; P= peloidal clay. PPL.

Granular mosaic cement

Granular mosaic cement fills void space between stromatolite columns and is composed of calcite crystals (Figure 27).

Figure 27: Granular mosaic cement. Granular mosaic cement filling voids spaces between stromatolite columns. Granular cement composed of calcite crystals. Ca= calcite. PPL.

Drusy Fabric

Euhedral calcite crystals with cleavage visible fill void spaces surrounded by granular mosaic cement composed of calcite crystals (Figure 28). Calcite crystals grow on ankerite bladed crystals, filling void spaces between peloidal grains (Figure 23).

Figure 28: Euhedral calcite. Euhedral calcite crystals are surrounded by granular mosaic calcite cement. Crystals are from secondary mineralization occurring during diagenesis. Cubic hematite crystals intermixed. Ca= calcite; H= hematite. PPL. Splay Cement

Splay cement is found growing on peloidal grains and intraclasts into void space. The grains have dissolved except for reddish-brown iron-oxide rims (Figure 22).

Ankerite crystals

Euhedral yellow-orange ankerite crystals are found within the pores and voids along with micro-crystalline quartz cement (Figure 29). Yellow-orange ankerite crystals that are not euhedral, but instead amorphous are also observed in void spaces along with micro-crystalline quartz (Figure 30).

Figure 29: Euhedral ankerite. Euhedral ankerite crystals are surrounded by microcrystalline quartz and cubic hematite. Remnant stromatolite fragment observed near the bottom of the figure. A= ankerite; H= hematite. PPL.

Figure 30: Amorphous ankerite. Amorphous yellow-orange ankerite surrounded by microcrystalline quartz and granular calcite cement. A= ankerite; Q= quartz; I= iron-oxide. PPL.

Micro-stromatolites

Micro-digitate stromatolites are observed in the void spaces along with peloidal grains. Micro-digitate stromatolites average 82.39 µm across with wavy, low inheritance laminae. Stromatolite columns overhang slightly and contain thin-walls with no bridging between columns visible (Figure 31).

Figure 31: Microdigitate stromatolites. Microdigitate stromatolites located in the interspaces of larger stromatolites. PPL.

Shrinkage Cracks

Shrinkage cracks due to de-watering and compaction during diagenesis observed in intraclasts, ooids, and stromatolite fragments. The intraclasts and ooids grains

70 located in the interspaces and atop the stromatolites (Figures 10 & 12). The stromatolite fragments (Figure 29) are located in vacuoles located between stromatolite laminae.

Shrinkage cracks were also observed dissecting the stromatolite laminae in thin section

(Figure 32).

Figure 32: Shrinkage cracks. Shrinkage cracks in stromatolite lamina due to diagenesis. PPL.

Vacuoles in stromatolite laminae

Vacuoles are observed between stromatolite layers. Stromatolite laminae above and below the vacuoles do not follow the vacuole shape. Vacuoles filled with microcrystalline quartz (Figure 20).

Remnant stromatolites

Stromatolite laminae no longer present, secondary replacement of magnetite allows for identification of the remnant stromatolite. The original stromatolite structure is unidentifiable due to dissolution (Figure 29).

Microfossils

Microfossils were observed in petrographic thin sections obtained from

Schreiber locality, Thunder Bay, Ontario, Canada. Microfossils observed at the edges of stromatolite laminae, and in between laminae couplets. Microfossil stalks appear in dense clusters within the stromatolite laminae with occasional stalks outside the laminae as well

(Figure 33).

The microfossil stalks ranged from 10.7 µm to 72.67 µm, with an average length of 32.58µm. Microfossil stalks appear smooth with no twisting or branching of the stalks observed. Microfossil stalk width ranged from 0.78µm to 2.55 µm, with an average width of 1.47 µm.

Microfossils within the Biwabik Iron Formation were not observed as most of the unit has experienced some degree of metamorphism. While it is possible that some microfossils still exist within the Biwabik Iron Formation, none were observed.

Figure 33: Microfossils. Microbial microfossils with non-bifurcated filaments located at the edges of the stromatolite laminae. Microfossil filaments vertically oriented overlie the horizontal mats of microfossil filaments. PPL.

Summary of Results

Determining the order of mineralogy is important in providing further evidence for original deposition of the Biwabik and Gunflint Formations. The order of mineralization can along with facies interpretation, and microfossil evidence can provide the necessary information to identify the microorganisms responsible for the stromatolite located within the Biwabik and Gunflint iron formations.

Peloids and intraclasts observed in thin-sections from localities within the

Biwabik Formation have the clay/mud that was primary replaced with greenalite.

Amorphous greenalite observed in intraclasts and peloids have obscured the original form of the clast. Peloidal rims do not remain intact were greenalite replacement is observed

(Figure 9). The replacement of peloidal mud with greenalite probably happened during early diagenesis. Minnesotaite observed in thin-sections is found along with greenalite in intraclasts and peloids. Minnesotaite is a secondary mineral that replaces greenalite in the reaction greenalite + quartz = minnesotaite (Perry, 1973). Peloids observed with minnesotaite have their rims no longer intact and the intermal structure obliterated.

Minnesotaite is observed along with amorphous greenalite in thin section.

Microcrystalline quartz replaces the original mineralogy of ooid cortices and ooid cores.

Ooid cores replaces with microcrystalline quartz no longer retain a clear internal structure and the cortices no longer contain a clear rim edge, instead the cortice rims are obscured

(Figure 34).

Primary carbonate still exists as stromatolite laminae couplets (Figure 20).

Most of the carbonate laminae observed has magnetite and hematite replacing parts of the siderite laminae (Figure 35). The magnetite and hematite crystals are much larger than the original carbonate in the laminae which supports the interpretation of magnetite and hematite replacing siderite. Eudehral ankerite is found within the interspaces between stromatolites and is considered to be due to secondary mineralization after deposition

(Figure 29).

Figure 34: Microcrystalline quartz. Presence of microcrystalline quartz viewed under cross-polarization light. Microcrystalline quartz replaces the internal structure of ooid cores and ooid cortice. Q= quartz.

Figure 35: Stromatolite laminae. Stromatolite laminae with original carbonate are beginning to be replaced by magnetite. C= carbonate; M= magnetite. PPL.

Hematite and magnetite are considered to have formed during secondary mineralization. Euhedral magnetite and hematite is observed surrounding vacuoles and pore spaces located within stromatolite laminae (Figure 20). Bladed minerals of iron oxide also grow off stromatolites and other microbial structures into the interspaces

(Figure 23). Whether the bladed minerals are hematite or magnetite is unknown.

Secondary mineralization also occurs as cement between ooids, peloids, and intraclasts and between stromatolite columns. Splayed calcite cement forms Euhedral crystals surrounding the grains (Figure 22). Bladed calcite surrounds dissolved peloidal grains preserving the outline of the peloidal rims (Figure 25). Peloidal cement also surrounds dissolved grains and contains fibrous minnesotaite crystals (Figure 26).

Along with the mineralization, sedimentary structures also help to constrain the paleoenvironment and facies of deposition. Laminar cross-bedding, trough cross- bedding and climbing ripple laminations are evidence of a intertidal depositional environment (Figure 7). Collapse-breccia, with minimal rounding indicates minimal transport of the breccia clasts. The collapse-breccia is composed primarily of ankerite, suggesting an upper shoreface environment.

Ooid cortice to core ratios, evidence of transport via delamination of ooid cortices and stromatolite morphology also help to in the interpretation of depositional environment. The fifteen percent ratio of ooid cortice to core indicates that the ooids were deposited in place, also in a intertidal environment (Flügel, 2004). The observation of delamination of the ooid cortice is evidence of transportation. However, as the

76 delaminated ooids are uncommon, it is probable that most of the ooids became settled near to where they were deposited.

Peloids, intraclasts, and flat-pebble conglomerates located between stromatolite columns or atop the undulatory and psuedocolumnar stromatolites do not show an indication for mass transport. The grains are primarily intact other than dissolution or recrystallization after diagenesis. The variety of the grains argues for a high energy environment with contributions from wave energy in their emplacement within the stromatolite deposits. Intercalation of coarse and medium grained sandstone with very-fine grained sandstone also indicates a intertidal to subtidal depositional environment.

The stromatolite morphology is also indicative of a high energy environment.

The transition from undulatory stromatolites to psuedocolumnar, dendritic and finally to domal stromatolites completes a parasequence and a time of sea-level rise in the intertidal zone leading to a subtidal depositional environment.

Microfossil evidence, while rare, also aids in the interpretation of depositional environment and provides direct evidence to the microorganisms responsible for the stromatolite layers located within the Biwabik and Gunflint Formations. Microfossil morphology viewed in thin section from samples collected at the Schreiber locality yielded information that suggests cyanobacteria were the stromatolite builders of the

Biwabik Iron Formation stromatolite horizons. Microfossil filaments are non-bifurcated and exhibit no twisting of filaments as typically found in Fe-oxidizing bacteria (Krepski et al.2013). Microfossil filaments are densely packed with vertical filaments overlaying

77 densely packed filaments that are horizontally oriented within the stromatolite laminae.

This directionality is common in cyanobacteria and opposite of the vertically oriented filaments found in iron-oxidizing bacteria.

CHAPTER V

DISCUSSION

Mineralogy

French (1973) noted that greenalite was either a primary mineral that formed during deposition or it was a secondary mineral that formed during early diagenesis.

Peloids observed in thin section suggest that greenalite is a secondary mineral that replaced the clay/mud that was the primary source of the peloidal grains. Observations of ooids from Mink Mountain and other locales within the Biwabik Iron Formation show remnants of the peloidal clay/mud retained in their nuclei, with the secondary greenalite minerals themselves replaced by fibrous minnesotaite and hematite crystals (Figure 16).

The replacement of greenalite with minnesotaite and hematite is attributed to low-grade metamorphism caused by the Duluth complex (Perry et al., 1973; Frost et al., 2007).

Grains observed in samples closer to the Duluth complex show complete replacement of greenalite either by minnesotaite (Figure 26), or by disseminated hematite grains (Figure

11).

Hematite does not only replace greenalite in the Biwabik Iron Formation, hematite also replaces primary carbonates in the laminae of stromatolites, the cortices of ooids, and rims of intraclasts and peloidal grains (Figure 11). Hematite as a replacement mineral most likely occurred due to the crystallization of Fe(III) oxy-hydroxides during early diagenesis (Frost et al., 2006). Euhedral hematite crystals increase in size and abundance closer to the Duluth Complex and result from the oxidation of magnetite.

Euhedral ankerite crystals, microcrystalline quartz, and bladed calcite crystals are secondary minerals that make up the cement between the peloidal grains, ooids, and intraclasts. Microcrystalline quartz within peloidal grains is attributed to the replacement of greenalite by quartz during early diagenesis (Frost et al., 2007).

Disseminated magnetite crystals are found within the fine-grained siltstone deposits of the Biwabik Iron Formation. These disseminated magnetite crystals are likely caused by the reduction of hematite by microbial organisms or as secondary replacement of greenalite or siderite during late diagenesis (Frost et al., 2007).

Paleoenvironment

The Mesabi Iron Range started as one large, deep continuous basin during the deposition of the lower Slaty and Lower Cherty members and by the time the Upper Slaty and Upper Cherty members had been deposited the basin had become two smaller, shallower basins (Pufahl et al., 2010; Severson et al., 2010). Deposition of the Biwabik

Iron Formation within the Mesabi Iron Range occurred during two transgressive and two minor regressive cycles along a stable craton margin (Pufahl et al., 2000). The two transgressive cycles are marked by the stromatolitic horizons located within the Biwabik

Iron Formation (Pufahl et al., 2000) (Figure 36). Observations of flat-laminar, undulatory, and psuedocolumnar stromatolites intercalated with or forming atop deposits of flat-pebble conglomerates offer further proof of the transgressive cycles mentioned by

Pufahl et al. (2000) (Figure 37). Wavy bedded deposits of chert and carbonate grainstones mark the two minor regressive cycles (Pufahl et al., 2000) (Figures 38 & 39).

Figure 36: Psuedocolumnar stromatolites. Psuedocolumnar stromatolites are forming atop flat-pebble conglomerate. Evidence of transgressive cycles described by Puhfahl et al., 2000. Lower Cherty member, core # 24974.

Figure 37: Psuedocolumnar stromatolites. Psuedocolumnar stromatolites are forming atop flat-pebble conglomerate. Evidence of transgressive cycles described by Puhfahl et al., 2000. Lower Cherty member, core # 24974.

Figure 38: Wavy beds. Wavy-bedded deposits of chert and ankerite are evidence of regressive cycle as noted by Puhfahl et al., 2000. Lower Cherty member, core # 24974.

Figure 39: Wavy beds. Wavy-bedded deposits of chert and ankerite are evidence of regressive cycle as noted by Puhfahl et al., 2000. Upper Cherty member, core # 24986.

Planavsky et al. (2009) suggest that the oceans at the time the Biwabik Iron

Formation was deposited were super-saturated with respect to silica and iron, allowing for the high abundance of silicate and iron minerals within the formation. Potential sources for the ferrous iron include hydrothermal vents, upwelling of pore waters, and

84 groundwater (Rouxel et al., 2008). While the exact source of the ferrous iron as yet is unknown, the upwelling of anoxic, iron-rich waters onto the oxygenated shelf of the

Superior Craton is not disputed. Mineralogy suggests that the upwelling of ferrous iron occurred after deposition of the stromatolite layers, and the presence of iron-oxides and silicates within the deposits resulting from the secondary replacement due to early-to-late diagenesis.

As the deposition of the Biwabik Iron Formation occurred after the Great

Oxygen Event (Holland, 2002), the world’s oceans would have been stratified with respect to oxygen. Nearshore to subtidal marine environments would have been oxic-to- suboxic, while the deep-waters would have been anoxic (Planavsky et al., 2009; Edwards et al., 2012). Water temperatures were between 10-35oC (Shiraishi, 2012; see Maliva et al., 2009 for a different interpretation).

Lithofacies to Facies Analysis

Eleven different lithofacies are present within the two stromatolite layers of the Biwabik Iron Formation. Lithofacies are used to interpret the depositional environment, or facies. Figure 40 demonstrates the various depositional environments within the Biwabik Iron Formation along with what type of deposits are observed within the given environment.

Figure 40: Facies diagram. Diagram of lithofacies and Facies interpretations. Modified from Severson et al., 2007 (Figure D-3).

Fine-and-medium grained sandstone layers and flat-pebble conglomerates are representative of beach deposits. Well-rounded grains and the predominance of quartz in this layer further suggest a beach facies with fluvial input. Flat-pebble conglomerate associated with psuedocolumnar stromatolites is consistent with beach deposits. The flat- pebble conglomerate clasts are orientated at an angle instead of being parallel to the stromatolites. The stromatolites have high inheritance laminae with low synoptic relief, no walls and are indicative of a high energy environment. Ooids, peloids and intraclasts are sub-millimeter in scale, sub-rounded to sub-angular. Most of the ooid cortices are intact, suggesting that minimal transport of the grains occurred before deposition, burial and diagenesis. During sea-level rise fine and medium beach sands become subjected to wave-action, causing planar and cross-bedding to occur (Appendix A, Figures 41 & 42).

Dendritic stromatolites, couplet patterned, thin-walled and bridging with occasional branching, dominate the subtidal/upper shoreface facies. The stromatolite lithofacies is dominated with ooids and intraclasts packed within the interspaces between stromatolite columns and in vacuoles between stromatolite laminae. Some peloidal grains are observed in the interspaces between stromatolite columns. Delamination of ooids is common along with shrinkage cracks filled with microcrystalline quartz (Figure 11).

Regenerated and broken ooids located within this lithofacies indicate the reworking of ooid grains due to wave action.

Domal stromatolites are observed to cap the dendritic stromatolites and the grains trapped within the stromatolites interspaces. Domal stromatolites also contain the couplet pattern of the other stromatolites along with high inheritance. Domal structures in thin section are greater than the core diameter, so true scale of the domes is unknown.

Collapse breccia intercalated with fine-grained sandstone indicates reworking of grains due to storm waves in the lower shoreface (Appendix A, Figures 45 & 46).

Collapse breccia deposits are observed only in cores that have not exhibited metamorphism from the Duluth Complex. Breccia clasts are only minimally rounded, indicative of little reworking or transport. Sub-millimeter scale banding of the siltstone intercalated with the collapse breccia is evidence of a lower energy zone.

The transition of stromatolites types, from flat-laminar all the way to domal stromatolites shows a coarsening upward cycle and the completion of a parasequence during a transgressive cycle (Ojakangas et al., 2001). This parasequence is found within both of the stromatolite layers of the Biwabik Iron Formation. With a rise in sea level, the

87 deeper lower shoreface deposits will overlie the shallower marine deposits (Appendix A,

Figures 38 & 43).

Organisms

Microbial communities require photosynthesis/chelomlithotrophy to create energy. Some microbial communities, like cyanobacteria, do this by using sunlight directly, or like sulfur-and iron-oxidizing bacteria by using another electron donor such as carbon, nitrogen, iron, or sulfur (Kirchman, 2008). As sulfur minerals and sulfates are minimal in the Biwabik Iron Formation, sulfur-oxidizing bacteria are not considered as a likely stromatolite builder for the stromatolite horizons.

Excluding the sulfur-oxidizing bacteria, that leaves only cyanobacteria or iron-oxidizing bacteria as the potential stromatolite builders in the Biwabik Iron

Formation. Cyanobacteria create oxygen in conjunction with sunlight for metabolism and energy generation (Kirchman, 2008). Cyanobacteria would have created oxygenated waters and may precipitate CaCO3 as a byproduct of photosynthesis via carbon fixation

(Shiraishi, 2012). Iron-oxidizing bacteria use the conversion of Fe(II) to Fe(III) as an electron source for metabolism and energy production (Moit et al., 2009).

Barghoorn and Tyler (1965) first described the microfossils within the

Gunflint and Biwabik iron formations as cyanobacteria, while Planavsky et al. (2009),

Edwards et al. (2012), and Wilson et al. (2010) described the microfossils located within their study areas as being iron-oxidizing bacteria. Both observations used the morphology of the microfossils compared to extant taxa to describe the potential microorganisms.

Planavsky et al. (2009), Edwards et al. (2012), and Wilson et al. (2010) also used isotope

88 analysis to interpret iron-oxidizers as the stromatolite builders, as morphology alone could not distinguish between the two types of microorganisms.

Planavsky et al. (2009) in their depositional model stated that the stromatolites were formed on the Laurentian continental shelf in anoxic waters below the redoxcline.

Their evidence was based on predominantly positive δ56Fe values along with Ce anomalies. Upwelling of hydrothermally dominated deep seawater rich in iron-oxides provided the iron source necessary for iron-oxidizing bacteria to convert Fe(II) to Fe(III) for energy production (Planavsky et al., 2009). The authors admitted that carbonate stromatolites existed within the Biwabik and Gunflint Formations but did not collect data on them.

Edwards et al. (2012) looked at the lack of limestone and dolostone in the

Labrador Trough, Canada as an indication that the waters were under saturated with respect to carbonates. They interpreted the hematite coating microfossil sheaths as being authigenic hematite cement, implying that iron-oxidation was the metabolic pathway used by the microorganisms. In the Edwards et al. (2012) depositional model, cyanobacteria formed carbonate stromatolites in the oxic intertidal zone, but that the banded iron-formation microfossils formed in the suboxic zone due to iron-rich water upwelling onto the continental shelf.

Wilson et al. (2010) observed in the Duck Creek Formation, Australia the same microfossil morphology as that described by Barghoorn and Tyler (1965), in the

Biwabik Formation and stated that the microfossils were from cyanobacteria. Their depositional model stated other than transported chert nodule, that the other microfossils

89 found in the Duck Creek formation formed below the redoxcline in anoxic waters with hydrothermally dominated upwelling of iron rich waters (Wilson et al., 2010).

All three of the depositional models interpret iron-oxidizers as the builders of banded iron-formations during the Paleoproterozoic. None of them addresses what metamorphism could have done to the units and whether iron-enrichment occurred after deposition of the stromatolite layers occurred.

By employing an approach that utilizes facies analysis with diagenetic and metamorphic alteration of primary minerals, this thesis concludes that cyanobacteria built the stromatolites of the Biwabik Iron Formation. Analysis of the sedimentary structures, grain types, and inferred primary minerals supports that the stromatolite layers were deposited in a high energy, shallow marine environment. As deposition of the sediments occurred 400 million years after the GOE, the shallow marine waters would have been suboxic, and therefore oxygenated enough that the microorganisms in the depositional environment would have been photosynthetic. While not all phototrophs require oxygen for photosynthesis, many release oxygen during respiration, increasing the amount of oxygen in the shallow marine environment. Iron in this shallow marine environment would have been oxidized before it could be utilized by Fe-oxidizing bacteria for energy synthesis. Fe-oxidizing bacteria would have needed to be in the anoxic deeper waters to obtain the required iron necessary to create energy.

CHAPTER VI

CONCLUSION

A detailed lithofacies and petrographic fabric analysis of the two Biwabik Iron

Formation stromatolite horizons allowed for reconstruction of the facies at the time of deposition. Depositional environment coupled with the ecological constraints and the knowledge that after the Great Oxygen Event (Holland, 2002) the oceans were stratified in regards to oxygen, allowed for interpretation of which microorganism was responsible for the creation of the stromatolite horizons. A detailed petrographic study of the mineralogy and inference as to whether the minerals were due to primary or secondary mineralization also aided in the interpretation of the microorganisms.

The study of the intraclasts, peloids, ooids, and stromatolite types with sedimentary structures including planar and cross-bedding, places the environment of deposition for the stromatolite layers as peritidal to subtidal. Mineralogy of the stromatolite layers suggests hematite and silica minerals are due to early-to late diagenesis replacing primary carbonates after deposition. Primary deposition of carbonates places the depositional environment above the redoxcline in oxic to suboxic waters suggesting cyanobacteria were the microorganisms that created the stromatolites.

The peritidal to subtidal environments described in this paper do not rule out the possibility that another type of microorganism was responsible for the formation of the iron-rich deposits of the Mesabi Iron Range. Further facies analysis and mineralogy should be done on the iron-rich deposits to infer which type of microorganism is responsible for the deposits.

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Appendix A: Stratigraphic

Column Descriptions and

Facies Interpretations

Stratigraphy and Facies Interpretations

Two cores were selected to be representative samples of the lithofacies observed within the two stromatolite horizons within the Biwabik Iron formation. These two cores are representative of the lower stromatolite horizon (#24974), and of the upper stromatolite horizon (#24986). Below are in-depth descriptions of the different lithofacies and environmental interpretations for each layer.

Core # 24974 750 feet to 752.4 feet in depth

Member 1 (Pokegama Quartzite)

Member 1 contains 234 cm of coarse-grained sandstone. The sandstone grains are rounded to sub-rounded, grain-supported with minimal surrounding of grains by microcrystalline quartz cement. Grains are well sorted. The grains are composed of white quartz, black chert and green chert. The composition of sand grains is 75% white quartz,

10 % green chert, and 4% black chert. Disseminated pyrite grains are randomly interspersed amongst the sand grains with a percent composition of less than 1%, and increase in percentage present, to 1% disseminated pyrite close to the lithofacies boundary. Flat-laminar, sub-millimeter scale bedding of coarse-grained sandstone dominates the bottom portion of the litohfacies in core sample cross-bedding, centimeter in scale is observed in the core sample near the top of the lithofacies (Figures 41 & 42).

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Figure 41: Wave action. Evidence of wave-action in fine and medium grained quartz sandstones exhibited as cross-bedding. Lower Cherty member, core # 24974.

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Figure 42: Wave action. Evidence of wave-action in fine and medium grained quartz sandstones as cross-bedding. Lower Cherty member, core # 24974.

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Facies Interpretation: Intertidal to Subtidal

The interpretation of this facies is intertidal to subtidal. The interpretation of a intertidal environment is based on the flat-laminar, sub-millimeter beds of coarse grained sandstone observed in the core. The interpretation of an subtidal environment for the upper portion of the lithofacies is based on the centimeter scale cross- bedding observed in the core sample. The fact that the lithofacies contains grain-supported deposits is also consistent with a intertidal to subtidal depositional environment along with the well- sorted grains of rounded to sub-rounded shape.

Member 2

Member 2 consists of 3 cm of coarse-grained sandstone deposits. This lithofacies boundary is gradational with the boundary of member 1. Grains transition from rounded to sub-rounded in member 1 to sub-rounded to sub-angular in member 2.

The grains of member two are sub-rounded to sub-angular, grain-supported with minor quartz cement (Figure 43). The beds are poorly sorted. Climbing ripple laminations are preserved in the core sample as disseminated pyrite and white quartz bands (Figure 44).

The layers are horizontal to sub-horizontal, and are millimeter width in scale. The grains of this lithofacies are composed of 80% percent quartz, 15% disseminated pyrite and 5% black chert.

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Figure 43: Breccia and conglomerate. Lower shoreface deposits overlie the shallower marine deposits. Shallow marine deposits consist of breccia and conglomerate layers surrounded by coarse-grained quartz sandstone. Lower shoreface deposits consist ofstromatolite boundstone with domal stromatolites growing atop psuedocolumnar stromatolites. Stromatolite boundstone overlies the breccia and conglomerate layers. Lower Cherty member, core # 24974.

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Figure 44: Flat laminar bedding. Flat-laminar sub-horizontal layers of coarse-grained sandstone indicative of layers of quartz and minor amounts of magnetite. Lower Cherty member, core # 24974.

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Facies Interpretation: Intertidal Zone

The interpretation of this depositional environment is intertidal zone. The interpretation is supported by the minimal, centimeter scale changes vertically of the unit along with the presence of climbing ripple laminations. The ripples indicate that the current was moving to the east, as evidenced by the position of the foresets moving from left to right in the core sample.

Member 3

Member 3 contains 3 cm of very fine-grained sandstone. The contact between member 2 and member 3 is sharp. The grains are sub-angular to angular and grain- supported with minor amounts of quartz cement. Sandstone layers are flat-laminar. The unit is composed of magnetite and disseminated pyrite grains. Some minor amounts of cubic pyrite are observed in the unit (Figure 45).

Facies Interpretation: Beach deposits

The facies interpretation for member 3 is beach environment. The evidence for this facies is based on the very fine-grained sandstone deposits along with the flat- laminar sandstone layers. The unit is grain-supported, which is further evidence of a beach environment.

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Figure 45: Sharp contact. Sharp contact between member 2 and member 3 with member 3 containing 2.5 centimeter layer of very-fine grained sandstone composed of magnetite and cubic pyrite grains. Lower Cherty member, core # 24974.

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Member 4

Member 4 contains 32 cm of brecciated clasts intercalated with very fine- grained sandstone layers. The contact between member 3 and member 4 is sharp. The very fine-grained sandstone layers contain sub-angular to angular grains and are grain- supported. The fine-grained sandstone layers have flat-laminar beds and are 1 to 6 cm in width, and well sorted. The grains are predominantly composed of magnetite with minor amounts of quartz. At the bottom of member 4 the breccia clasts are surrounded by disseminated pyrite and fine-grained magnetite (Figure 46). Disseminated pyrite grains surrounding the breccia clasts in only found in the bottom 5 cm of member 5. The breccia deposits near the top of member 4 are ankerite breccia clasts contained in granular calcite veins (Figure 47). The ankerite breccia clasts are occasionally surrounded by fine-grained magnetite (Figure 47).

Facies Interpretation: Lower Shoreface

The facies interpretation of member 4 is Lower shoreface. The evidence for this interpretation includes thin sandstone sheets composed of angular very fine-grained sand grains intercalated with the breccia clasts. The breccia clasts are interpreted as collapse breccia due to the minimal rounding of the clasts, and imply that little transport of the clasts occurred. The collapse breccia is representative of storm-reworking of the unit.

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Figure 46: Sharp contact between member 3 and member 4. Quartz and ankerite breccia clasts 0.5 to 1 mm in diameter surrounded by quartz and disseminated pyrite matrix in sharp contact with member 3. Lower Cherty member, core # 24974.

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Figure 47: Brecciated clasts. Brecciated clasts composed of quartz intercalated 0.5-1mm in diameter with very-fined grained magnetite sandstone layers. Sandstone layers are flat laminar and sub-millimeter in width. Grains are sub-angular. Breccia clasts are in a calcite matrix. Lower Cherty member, core # 24974.

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Member 5

Member 5 consists of 15 cm of medium grained sandstone intercalated with black chert clasts. The sandstone grains are sub-angular to angular and poorly sorted. The black chert clasts are sub-angular to angular. The clasts range from sub-millimeter to 5 mm in size and are poorly sorted. Occasional granular calcite veins are near the top of the unit (Figure 48). Wavy, sub-parallel layers are observed as well in the unit (Figure 48).

Occasional disseminated pyrite grains are observed near the top of member 5.

Disseminated hematite is observed in the top 3 cm of the lithologic unit as a red stain surrounding the breccia clasts (Figure 49).

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Figure 48: Breccia with wavy layered sandstone. Member 5 consists of black chert breccia clasts sub-millimeter to 5 mm in diameter intercalated with medium-grained wavy layered sandstone. The sandstone layer is 1 cm in width with sub-angular grains that are grainsupported. Occasional calcite veins are observed throughout the unit. Lower Cherty member, core # 24974.

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Figure 49: Hematite staining. Hematite staining observed in the top 3 cm of the unit. Lower Cherty member, core # 24974.

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Facies Interpretation: Upper Shoreface

Member 5 is interpreted as an upper shoreface environment. The difference between member 4 and member 5 is the loss of magnetite, and the addition of hematite along with the absence of the thin-bedded fine-grained sandstone. Wavy, sub-parallel layers are interpreted as cross-bedding and are also indicative of an upper shoreface environment.

Member 6

Member 6 contains 16 cm of breccia clasts intercalated with domal stromatolite fragments and oncoids. The unit is poorly sorted and contains coarse-grained sandstone intercalated with the breccia clasts. The grains are composed of 80% quartz, 5

% disseminated hematite, and 5% disseminated pyrite. The breccia clasts are black chert and range in size from sub-millimeter to 5 mm. The breccia clasts are sub-angular to angular and poorly sorted. The top 6 cm of the lithologic unit contain the domal stromatolite fragments, magnetite grains; black chert breccia clasts and granular calcite cement (Figure 50). The domal stromatolite fragments are 2cm by 1 cm with wavy laminae and moderate inheritance. A large domal stromatolite measuring 4 cm across is in the top 3 cm of the lithologic unit (Figure 50). A possible oncoid measuring 2 cm in diameter is on top of the 4 cm domal stromatolite along (Figure 50). Potential ooids are amongst the breccia and coarse-grained sandstone clasts.

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Figure 50: Breccia with domal stromatolites. Member 6 consists of 16 cm of poorly sorted breccia clasts intercalated with a domal stromatolites and possible oncoids. Lower Cherty member, core # 24974.

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Facies Interpretation: Upper Shoreface

The upper shoreface interpretation is based on the microbial structures along with the breccia clasts. The presence of domal stromatolites, possible ocoids, and ooids are indicative of an upper shoreface environment.

This last lithologic unit is capped by 5+ cm of magnetite rich siltstone. The siltstone is in sub-millimeter bands that drape sub-horizontally with one another. The contact between member 6 and the magnetite siltstone is sharp.

Core #24986 205 feet-211.8 feet in depth

Member 1

Member 1 consists of 12.5 cm of coarse-pebble conglomerate clasts surrounded by coarse-grained disseminated magnetite. The coarse-pebble conglomerate clasts are sub-rounded to sub-angular and are composed of green chert. The conglomerate clasts range from 0.5 mm- to 4 cm in diameter. Beds appear to be stratified between coarse-grained and very fine-grained sandstone layers (Figure 51). The very fine-grained sandstone layers are the first 5 cm of the unit.

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Figure 51: Flat- pebble conglomerate. 16.5 cm of flat-pebble conglomerate intercalated with poorly sorted coarse-grained sandstone and well sorted very-fine grained sandstone. All clasts are sub-rounded to sub- angular. Flat-pebble conglomerate clasts composed of green chert. Sandstone grains are green chert and quartz. The unit is well stratified. Upper Cherty member, core # 24986.

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Facies Interpretation: Beach

The facies interpretation of member 1 is beach. The presence of the pebble conglomerate intercalated with coarse-grained sandstone deposits is evidence of this interpretation. The pebble conglomerate appears in lenses amongst the coarse-grained sandstone layers.

Member 2

Member 2 contains 30 centimeters of psuedocolumnar stromatolites intercalated with a pebble-conglomerate. A 1 cm high psuedocolumnar stromatolite is overlying the member 1 deposits. Another 1 cm thick psuedocolumnar stromatolite if located 8 cm above the beginning of the lithologic unit (Figure 52). Other 2 to 3 cm thick psuedocolumnar stromatolite layers are located 22 cm and 30 cm above the contact between member 1 and member 2 (Figures 52 & 53). The psuedocolumnar stromatolite laminae contain couplet patterns of green chert and black magnetite with low inheritance and low synoptic relief. The laminae are wavy. No walls are observed within the stromatolite layers. The overall height of the member is 27.5 cm.

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Figure 52: Psuedocolumnar stromatolites. Member 2 consists of 27.5 cm of poorly sorted coarse-grained sandstone intercalated with flat-pebble conglomerate. Grains and clasts are sub-rounded to sub-angular. Layers are also intercalated with psuedocolumnar stromatolites measuring 4cm in width and 1- 2.5 cm in height. Stromatolite layers consist of fine-grained sand that is sub-rounded. Stromatolites contain high inheritance wavy laminae with low synoptic relief. Upper Cherty member, core # 24986.

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Figure 53: Flat-pebble conglomerate with psuedocolumnar stromatolites. Another image of member 2 with poorly sorted coarse-grained sandstone intercalated with flat-pebble conglomerate. Grains and clasts are sub-rounded to sub-angular. Layers are also intercalated with psuedocolumnar stromatolites measuring 4cm in width and 1- 2.5 cm in height. Stromatolite layers consist of fine-grained sand that is sub-rounded. Stromatolites contain high inheritance wavy laminae with low synoptic relief. Upper Cherty member, core # 24986.

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This unit also contains the green chert coarse-pebble conglomerate layers that were observed in member 1 along with the psuedocolumnar stromatolite. The coarse- pebble clasts are sub-rounded to sub-angular. The clasts are surrounded with a granular calcite matrix. Coarse-grained sandstone layers 1 cm in width lie directly above and below the pebble-conglomerate layer. The pebble clasts range from sub-millimeter in diameter to 4 mm in diameter. Thin layers of very fine-grained disseminated magnetite sandstone are intercalated with the pebble conglomerate. The sandstone grains are sub- angular to angular. The sandstone layers are parallel and sub-millimeter in scale.

The whole unit is poorly sorted. Minor amounts of disseminated pyrite grains along with disseminated magnetite grains are present. Small amounts of ankerite and hematite are also present in the top 3 cm of the unit. A granular calcite vein due to secondary mineralization overlies the psuedocolumnar stromatolite at the bottom of the member. The amount of magnetite increases towards the top of the lithologic unit to around 10% of the matrix being coarse-grained disseminated magnetite.

The final 2 cm of this lithologic unit is capped with a domal stromatolite.

Laminae couplets of magnetite and ankerite/calcite are observed. Laminae are wavy with moderate inheritance.

Facies Interpretation: Intertidal Zone

The Facies interpretation of this member is intertidal zone. The psuedocolumnar stromatolite is indicative of a high energy wave environment with its wavy laminae of low inheritance and low synoptic relief. The stratification of the sandstone layers is also indicative of a high energy environment.

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Member 3

This lithologic unit consists of 28 cm of undulatory stromatolites. The stromatolites are composed of green chert, ankerite and calicite. Laminae are sub- millimeter in width and the stromatolite length is the entire 4 cm diameter of the core.

Occasional chert breccia clasts ranging from sub-millimeter to 5 millimeter in length are interspersed between the stromatolites in a 2 cm layer. Coarse-grained ankerite and quartz sandstone surround the breccia clasts (Figures 54 & 55). Granular calcite veins 2 mm in width are present in the unit (Figure 56). Minor amounts of disseminated pyrite are observed near the top of the lithologic unit. The unit terminates with a 3 cm vein of granular calcite (Figure 57).

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Figure 54: Undulatory stromatolites. Undulatory stromatolites intercalated with green chert breccia clasts ranging from sub- millimeter to 5 mm in diameter. Breccia clasts surrounded by coarse-grained ankerite and magnetite. Occasional granular calcite veins are observed within the unit. Upper Cherty member, core # 24986.

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Figure 55: Undulatory stromatolite layers amongst breccia. Undulatory stromatolites intercalated with green chert breccia clasts ranging from sub- millimeter to 5 mm in diameter. Breccia clasts surrounded by coarse-grained ankerite and magnetite. Occasional granular calcite veins are observed within the unit. Upper Cherty member, core # 24986.

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Figure 56: Granular veins of calcite amongst the undulatory stromatolite layers.

Calcite veins are a secondary mineralization within the unit. Minor amounts of disseminated pyrite are also located within the unit. Upper Cherty member, core # 24986.

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Figure 57: Granular veins of calcite amongst the undulatory stromatolite layers. Calcite veins are a secondary mineralization within the unit. Minor amounts of disseminated pyrite are also located within the unit. Upper Cherty member, core # 24986.

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Facies Interpretation: Intertidal Zone

The interpretation of intertidal zone for this lithologic unit is based on the undulatory stromatolite layers. The stromatolites are indicative of a high energy environment. They are also indicative of microbial mats with chaotic, wavy laminae.

Stratified sands along with breccia clasts also give evidence for an intertidal environment.

Member 4

This unit is 81 centimeters in length and composed of dendritic stromatolites.

The stromatolites contain couplets of ankerite and green chert laminae. The laminae are sub-millimeter in width. No walls are present and the stromatolite columns are overhanging and contain bridging. Laminae contain high inheritance and high synoptic relief. Intraclasts of coarse-grained quartz, hematite and magnetite are wedged between the columns. The grains are sub-angular to angular and poorly sorted. The quartz, hematite and magnetite grains are surrounded with granular calcite cement with occasional granular calcite veins observed (Figure 58).

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Figure 58: Dendritic stromatolites. Member 4 is 83 cm of dendritic stromatolites with steep walls and bridging. Laminae are couplets of green chert and ankerite.Laminae contain high inheritance and high synoptic relief. Upper Cherty member, core # 24986.

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Sub-angular to angular magnetite and ankerite clasts, 2 mm in width are observed trapped between stromatolite laminae (Figures 58 & 59). 2 cm of sub-parallel beds composed of magnetite and green chert clasts observed intercalated with the dendritic stromatolites. Agate veins observed within the unit and measure 5 centimeters in width and range from 2 cm to 4 cm in length (Figure 60).

Figure 59: Ankerite. Ankerite percentages increase towards the top of member 4 unit. Clasts are grain- supported and sub-angular to angular. Upper Cherty member, core # 24986.

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Figure 60: Psuedocolumnar stromatolite layer with agate veins. Stromatolite layer of member 4 with the agate veins occasionally observed throughout the unit. Agate band roughly 1ch high and 4 cm wide. The agate band in Figure S. is surrounded by sub-horizontal chert and coarse-grained quartz sandstone layers. Granular ankerite grains are also observed within the unit. Upper Cherty member, core # 24986.

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6 cm from the top of the unit contains coarse-grained ooid rich sandstone. Grains are sub- rounded to round and the layer is grain supported. The stromatolites and the coarse- grained ooid rich sandstone show no sharp contacts and are instead gradational (Figure

59). In thin section ooids were present between the stromatolite columns along with sub- rounded coarse sand grains.

Facies Interpretation: Subtidal Zone

Interpretation of this unit is a little more difficult. Stromatolite branching observed along with poorly sorted, coarse-grained intraclasts suggest a high energy environment. The intraclast matrix and stromatolite laminae contain grain-supported clasts. The stratified sandstone layers that surround the intraclasts also indicate the subtidal environment. The presence of ankerite suggests the presence of iron-hydroxides in the water, but whether the iron-hydroxides were present during deposition or after deposition, and during diagenesis is uncertain.

Member 5

This unit contains 16 cm of psuedocolumnar stromatolites in gradational contact with member 4. The stromatolite laminae are wavy with high inheritance and low synoptic relief. Laminae are composed of very fine-grained quartz, green chert and ankerite. Parallel accretion couplet patterns are observed in the laminae. Grains trapped between laminae are very-fine grained, rounded to sub-rounded and grain supported.

Agate bands ranging to sub-millimeter to 1 cm in width are observed between stromatolite laminae (Figure 60). Occasional black chert or possible magnetite grains are also observed within the unit along with granular secondary ankerite. The ankerite

135 increases from less than 1 % to 10% mineral composition closer to the top of the lithologic unit.

The top of the unit is capped by sub-millimeter to 3 mm in length intraclasts.

Intraclasts are sub-angular to angular and poorly sorted. The clasts are composed of magnetite and green chert and surrounded with calcite cement. The intraclasts are surrounded by medium-grained sub-angular to angular magnetite and chert grains. The matrix is grain-supported.

Facies Interpretation: Intertidal to Subtidal Zone

The facies interpretation of this unit is based on the presence of the intraclasts and medium grained sandstone layers intercalated and capping the psuedocolumnar stromatolites. The presence of the psuedocolumnar stromatolites indicates a high energy environment such as would be seen in the intertidal to subtidal environment where wave- action would allow various sized clasts to intermix along with the stromatolites.

Member 6

Member 6 contains 28 cm of wavy-bedded coarse-grained sandstone with occasional undulatory stromatolite fragments. The minerals observed in this unit include green chert, quartz, ankerite, and magnetite. Most of the unit is composed of green chert.

Wavy-beds are sub-horizontal and sub-millimeter in width. Evidence of cross-bedding appears amongst the wavy-bedded coarse-grained sandstone layers (Figure 61).

136

Figure 61: Undulatory stromatoliteswith evidence of cross-bedding. Undulatory stromatolite fragments amongst wavy-bedded coarse-grained sandstone. Evidence of cross-bedding. Upper Cherty member, core# 24986.

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Facies Interpretation:Upper Subtidal Zone

The facies interpretation for member 6 is upper subtidal zone. The presence of undulatory stromatolite fragments suggests reworking of the layer due to wave action.

The cross-bedding of the coarse-grained sandstone also suggest the subtidal zone for the depositional environment for member 6.

The remainder of the core is composed of black magnetite rich siltstone layers that are sub-horizontal and sub-millimeter in width. The contact between member 5 and member 6 is sharp.

Appendix B: Data Tables

TABLE 3. INTRACLAST LENGTH AND WIDTH

Sample Locality Intraclast Length (µm) Intraclast width (µm) Gunflint Lake 1802.14 62.63 Gunflint Lake 1710.44 67.89 Gunflint Lake 1735.02 107.83 Gunflint Lake 1816.21 97.79 5NI-B3-6 1087.56 514.72 17729 157.38 142.69 17729 200.05 162.97 2 East 116.11 No measurement 2 East No measurement 362.32 2 East No measurement 86.73 2 East No measurement 95.83 2 East No measurement 80.44 2 East No measurement 72.04 17883 No measurement 1079.8 17883 1287.61 395.97 17883 2797.16 405.56 17883 786.55 273.12 17883 2610.40 684.35 17883 961.22 271.31 17883 2684.66 302.11 17883 3676.26 456.68 17883 2716.66 590.36 Kakabeka Falls 1462.23 1046.46 Kakabeka Falls 956.85 795.67 Kakabeka Falls 639.43 540.07 Kakabeka Falls 927.03 651.38 North Shore 626.33 309.76 North Shore 390.01 136.92 North Shore 491.51 470.21 North Shore 581.27 577.17 North Shore No measurement 664.67 North Shore 415.78 211.14 North Shore 540.70 412.85 North Shore 426.54 314.34 North Shore 456.48 266.61 North Shore 546.71 349.17 North Shore 894.04 278.75 North Shore 320.94 176.49 North Shore 612.45 289.36 North Shore 1017.08 861.91 North Shore 847.84 483.24 North Shore 531.34 356.41

139

140

TABLE 3, CONTINUED

North Shore 698.85 462.53 North Shore 996.41 300.21 North Shore 444.66 315.75 North Shore 1355.52 801.96 North Shore 730.35 311.81 North Shore 661.90 379.96 Average 1066.42 377.63 Max 3676.26 1079.80 Min 116.11 62.63 Standard Deviation 828.5552385 254.94335977 Table 3: Intraclast width and length measurement.

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TABLE 4. PELOID LENGTH AND WIDTH Locality Peloid Length (µm) Locality Peloid Width (µm) 17783 397.44 17783 159.19 17783 322.74 17783 124.60 Gunflint Lake 572.03 Gunflint Lake 572.03 Gunflint Lake 157.59 Gunflint Lake 336.52 Gunflint Lake 222.95 Gunflint Lake 758.33 Gunflint Lake 352.55 Gunflint Lake 362.32 Gunflint Lake 332.99 Gunflint Lake 723.22 Gunflint Lake 352.55 5NI 1296.90 17729 121.01 17729 395.90 17729 137.09 17729 351.83 17729 82.54 17729 136.40 17729 142.69 5NI 303.19 5NI 287.96 5NI 353.27 5NI 491.14 5NI 415.9 5NI 252.02 Gunflint Lake 397.52 Gunflint Lake 367.36 17883 497.47 17883 314.96 17883 321.29 17883 201.77 17883 433.01 17883 337.19 17883 309.45 17883 261.1 17883 452.51 17883 226.32 17883 506.98 17883 297.02 5NI 345.53 5NI 298.67 5NI 396.59 5NI 351.83 5NI 1087.56 5NI 514.72 5NI 763.77 5NI 616.34 Kakabeka Falls 441.36 Kakabeka Falls No measurement Kakabeka Falls 702.54 Kakabeka Falls 671.36 Kakabeka Falls 188.16 Kakabeka Falls 122.50 Kakabeka Falls 325.56 Kakabeka Falls 195.77 Kakabeka Falls 786.43 Kakabeka Falls 595.22 Kakabeka Falls 666.36 Kakabeka Falls 493.20 17729 No measurement 17729 142 17729 444.86 17729 260.20 17729 598.74 17729 498.02 17729 No measurement 17729 489.62 17729 353.93 17729 206.34 Average 494.18 Average 268.91 Max 1296.90 Max 616.34 Min 136.40 Min 82.45 Standard Deviation 237.4474725 Standard Deviation 160.2466547 Table 4: Peloid width and length measurements.

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TABLE 5. OOID LENGTH AND WIDTH Locality Ooid Length (µm) Locality Ooid width (µm) Mink Mountain 1176.43 Mink Mountain 672.14 Mink Mountain 630.67 Mink Mountain 780.47 Mink Mountain 705.36 Mink Mountain 783.28 Mink Mountain 1178.67 Mink Mountain 638.04 Mink Mountain 823.45 Mink Mountain 540.66 Mink Mountain 1185.18 Mink Mountain 606.87 Mink Mountain 706.64 Mink Mountain 608.62 Mink Mountain 1000.13 Mink Mountain 596.87 Mink Mountain 860.66 Mink Mountain 463.63 Mink Mountain 845.75 Mink Mountain 1019.83 Mink Mountain 840.37 Mink Mountain 628.39 Mink Mountain 1060.49 Mink Mountain 588.99 Mink Mountain 1098.14 Mink Mountain 470.28 Mink Mountain 974.36 McKinley Mine 503.51 Mink Mountain 925.44 McKinley Mine 530.32 Mink Mountain 670.93 McKinley Mine 367.67 Mink Mountain 881.23 McKinley Mine 323.4 Mink Mountain 768.03 McKinley Mine 257.09 Mink Mountain 860.16 McKinley Mine 404.38 Mink Mountain 1040.40 Mink Mountain 524.51 Mink Mountain 1099.32 MGS2 241.21 Mink Mountain 966.67 MGS2 319.76 Mink Mountain 660.43 MGS2 305.13 Mink Mountain 786.78 MGS2 350.97 McKinley Mine 621.58 MGS2 519.63 McKinley Mine 636.2 McKinley Mine 498.51 McKinley Mine 463.72 McKinley Mine 421.84 McKinley Mine 533.53 Mink Mountain 974.36 MGS2 399.19 MGS2 383.23 MGS2 590.37 MGS2 555.64 Average 789.69118 Average 567.50 Max 1185.18 Max 1019.83 Min 383.23 Min 241.21 Standard Deviation 237.69377 Standard Deviation 177.85095 Table 5: Ooid width and length measurements.

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TABLE 6. OOID CORTICE WIDTH, LENGTH, AND RATIO

Ooid Cortice width Ooid Core length (µm) Ratio Cortice (µm) width/Core length (µm) Mink Mountain 63.65 751.97 0.0846 Mink Mountain 66.54 661.74 0.10055 Mink Mountain 127.42 645.71 0.19733 Mink Mountain 256.39 714.11 0.35903 Mink Mountain 82.54 812.10 0.10164 Mink Mountain 8.04 120.65 0.6664 Mink Mountain 5.42 102.99 0.5263 Mink Mountain 121.83 626.66 0.19441 Mink Mountain 48.96 560.36 0.8737 Mink Mountain 20.28 657.65 0.3084 Mink Mountain 57.27 837.43 0.6839 Mink Mountain 120.31 580.71 0.20718 Mink Mountain 242.94 612.87 0.3964 Mink Mountain 152.46 785.49 0.1941 Mink Mountain 189.59 980.23 0.19341 Mink Mountain 62.31 885.52 0.07037 Average 101.62 646.01 0.15731 Max 256.39 980.23 0.396397278 Min 5.42 102.99 0.030837071 Standard Deviation 77.296581 238.4596 0.107184218 Table 6: Ooid cortice/ ooid core ratios.

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TABLE 7. STROMATOLITE WIDTH AND SYNOPTIC RELIEF Locality Stromatolite Width (µm) Locality Synoptic Relief (µm) Kakabeka Falls 2864.83 Kakabeka Falls 1089.24 Kakabeka Falls 3122.41 Kakabeka Falls 2968.86 Kakabeka Falls 4159.08 Kakabeka Falls 1879.63 Kakabeka Falls 3887.63 Kakabeka Falls 2378.49 Kakabeka Falls 4371.72 Mink Mountain 3046.60 Kakabeka Falls 4916.80 MGS2 574.65 North Shore 6060.71 MGS2 1474.35 North Shore 7890.61 North Shore 1350.37 Kakabeka Falls 4239.66 Mink Mountain 2096.50 Kakabeka Falls 4045.24 Gunflint Lake 3436.68 Kakabeka Falls 3830.50 Kakabeka Falls 2864.83 Kakabeka Falls 3789.10 Kakabeka Falls 4244.56 Kakabeka Falls 4071.16 Kakabeka Falls 3485.36 Kakabeka Falls 2153.40 Kakabeka Falls 2971.98 Kakabeka Falls 2533.19 Kakabeka Falls 2548.32 Mink Mountain 2404.95 Mink Mountain 4708.21 North Shore 4356.97 North Shore 6170.11 North Shore 4850.09 North Shore 3218.20 Mink Mountain 4374.07 Mink Mountain 5267.79 Gunflint Lake 4860.52 Gunflint Lake 5472.67 Gunflint Lake 6408.29 McKinley Mine 6299.63 McKinley Mine 5904.33 McKinley Mine 6848.99 Kakabeka Falls 4620.11 Kakabeka Falls 4159.08 Kakabeka Falls 3122.41 Kakabeka Falls 3930.52 Kakabeka Falls 4665.81 Kakabeka Falls 6060.71 Average 4444.69 Average 1852.99 Max 7890.61 Max 2986.86 Min 2153.40 Min 574.65 Standard Deviation 1345.7212 Standard Deviation 807.667 Table 7: Characteristics of stromatolite laminae. Table consists of stromatolite laminae width and synoptic relief, scale in micrometers.

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TABLE 8. STROMATOLITE WIDTH AND LAMINAE MINERALOGY Locality Laminae Width(µm) Laminae Mineralogy Nolalu 369.33 hematite Nolalu 219.83 carbonate/hematite Nolalu 91.76 carbonate/hematite Nolalu 197.44 carbonate/hematite Nolalu 109.14 carbonate Nolalu 189.16 hematite/carbonate Nolalu 44.97 carbonate MGS2 175.54 carbonate/hematite MGS2 163.92 carbonate/hematite MGS2 126.83 carbonate/hematite MGS2 71.75 carbonate/hematite MGS2 126.83 hematite MGS2 70.35 carbonate/hematite MGS2 77.65 carbonate/hematite MGS2 287.97 hematite MGS2 421.36 carbonate/hematite MGS2 345.50 carbonate/hematite MGS2 1479.00 carbonate MGS2 461.26 carbonate Mink Mountain 175.54 carbonate Mink Mountain 396.72 carbonate Mink Mountain 252.77 carbonate Mink Mountain 203.62 carbonate Mink Mountain 115.85 carbonate Mink Mountain 101.81 carbonate Mink Mountain 126.39 carbonate Mink Mountain 277.35 carbonate Mink Mountain 312.46 carbonate Mink Mountain 403.74 carbonate Mink Mountain 161.49 carbonate Mink Mountain 126.39 carbonate Mink Mountain 379.16 carbonate Mink Mountain 698.64 carbonate Mink Mountain 814.50 carbonate Mink Mountain 84.26 carbonate Mink Mountain 410.76 carbonate Mink Mountain 681.09 carbonate Mink Mountain 421.29 carbonate Mink Mountain 495.02 carbonate Mink Mountain 705.66 carbonate Mink Mountain 182.50 carbonate Mink Mountain 403.74 carbonate

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TABLE 8 CONTINUED Mink Mountain 366.57 carbonate Mink Mountain 327.09 carbonate Mink Mountain 417.33 carbonate Mink Mountain 214.73 carbonate Mink Mountain 253.21 carbonate Mink Mountain 356.82 carbonate Mink Mountain 139.20 carbonate Mink Mountain 138.27 carbonate Mink Mountain 184.95 carbonate Mink Mountain 200.93 carbonate Mink Mountain 43.49 carbonate Mink Mountain 130.13 carbonate Mink Mountain 262.51 carbonate Mink Mountain 222.09 hematite Mink Mountain 96.09 hematite Mink Mountain 126.74 carbonate Mink Mountain 218.59 carbonate/hematite Mink Mountain 123.03 carbonate/hematite 2East 38.47 carbonate/greenalite 2East 32.16 greenalite/carbonate 2East 84.63 carbonate/greenalite 2East 201.44 greenalite/carbonate/hematite 2East 46.16 greenalite/carbonate/hematite 2East 42.67 carbonate/hematite 2East 67.85 greenalite/carbonate/hematite 2East 78.34 greenalite/carbonate/hematite 2East 612.03 greenalite/carbonate 2East 430.17 greenalite/hematite 2East 426.07 greenalite/carbonate/hematite 2East 260.90 greenalite/carbonate 2East 335.04 greenalite/carbonate/hematite 2East 301.47 greenalite/carbonate 17729 615.17 hematite 17729 578.35 carbonate/hematite 17729 512.09 carbonate/hematite 17729 387.77 hematite 17729 699.35 carbonate/hematite 17729 159.34 hematite 17729 186.50 carbonate/hematite 17729 154.02 hematite 17729 227.93 carbonate/hematite 17729 107.51 hematite 17729 365.88 carbonate

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TABLE 8 CONTINUED 17883 131.72 carbonate/hematite 17883 504.63 carbonate/hematite 17883 122.85 carbonate/hematite 17883 1332.98 carbonate/hematite 17883 690.04 carbonate/hematite 17883 95.91 hematite 17883 185.86 carbonate/hematite Gunflint Lake 840.03 carbonate Gunflint Lake 505.55 carbonate Gunflint Lake 319.79 carbonate/hematite Gunflint Lake 305.84 carbonate/hematite Gunflint Lake 662.50 carbonate Gunflint Lake 898.92 carbonate Gunflint Lake 516.66 carbonate Nolalu 314.05 carbonate Nolalu 591.31 hematite Nolalu 137.64 carbonate Nolalu 112.56 carbonate/hematite Nolalu 90.55 carbonate Nolalu 537.71 carbonate/hematite Nolalu 175.54 hematite Nolalu 193.38 carbonate Nolalu 250.89 hematite Nolalu 294.90 carbonate Nolalu 150.96 hematite Average 271.40 Max 1479 Min 32.16 Standard Deviation 228.3593812 Table 8: Stromatolite laminae widths and mineralogy. Table consists of stromatolite laminae widths and mineralogy, scale in micrometers.

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TABLE 9. MICROFOSSIL LENGTH AND WIDTH Locality Microfossil length (µm) Microfossil Width (µm) Schreiber 26.34 1.88 Schreiber 20.20 2.55 Schreiber 21.50 1.48 Schreiber 18.95 2.04 Schreiber 20.99 1.88 Schreiber 18.07 1.26 Schreiber 18.38 1.88 Schreiber 26.34 1.99 Schreiber 40.23 0.87 Schreiber 27.57 1.12 Schreiber 10.70 1.56 Schreiber 24.61 1.50 Schreiber 34.91 1.61 Schreiber 27.48 1.41 Schreiber 48.63 1.12 Schreiber 27.60 1.76 Schreiber 30.89 2.38 Schreiber 20.65 1.40 Schreiber 35.98 1.73 Schreiber 29.34 1.41 Schreiber 29.17 1.80 Schreiber 68.91 0.87 Schreiber 47.59 0.94 Schreiber 11.45 1.24 Schreiber 31.75 1.27 Schreiber 35.20 1.05 Schreiber 32.96 1.17 Schreiber 38.95 1.86 Schreiber 53.65 1.02 Schreiber 72.67 0.99 Schreiber 31.66 1.24 Schreiber 28.85 1.41 Schreiber 29.73 1.12 Schreiber 32.69 1.24 Schreiber 29.55 0.78 Schreiber 26.38 1.73 Schreiber 48.51 1.56 Schreiber 21.90 1.41 Schreiber 50.92 1.33 Schreiber 31.51 1.80 Schreiber 29.00 1.48 Schreiber 51.68 no measurement

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TABLE 9 CONTINUED Schreiber 38.21 no measurement Average 32.53 1.47 Max 72.67 2.55 Min 10.7 0.78 Standard Deviation 13.200198 0.404514765 Table 9: Microfossil width and length measurements.