Comparison Between the Helen Iron Formation (Algoma-Type) and the Sokoman Iron Formation (Superior-Type)

Comparison between the Helen Iron Formation (Algoma-type) and the Sokoman Iron Formation (Superior-type)

Differences and similarities in their depositional environments, mineralogy and geochemistry

Therese Ingrid Garcia

A thesis submitted in conformity with the requirements for the degree of Doctor of Philosophy (PhD)

Department of Earth Sciences University of Toronto

Therese Ingrid Garcia

Degree of Doctor of Philosophy

Department of Earth Sciences

University of Toronto

2014

Abstract

The Sokoman Iron Formation with an age of 1.88 Ga in the Labrador Trough and the Archean

Helen Iron Formation in Wawa, Ontario (2.75 Ga) are examples of shallow and deep water iron formations respectively, which are of a low metamorphic grade. The two iron formations vary greatly in their mineralogy and geochemistry.

The deep water Helen Iron Formation is dominated by siderite and quartz and provides an example of a direct precipitate from seawater which had a strong but distal hydrothermal component as indicated by the lack of Cu, Pb and Zn enrichment. It has a thickness of 100m.

The first 20 meters are dominated by input of detrital material of both basaltic and felsic sources before deepening of the basin led to purely chemical sedimentation. Iron in the Helen Iron

Formation occurs predominantly as Fe2+, indicating that oxygen played no role in the deposition of this iron formation.

Negative carbon isotope excursions close to organic rich horizons in the Helen Iron Formation are the result of dissimilatory iron reduction.

In contrast, the mineralogy of the Sokoman Iron Formation is dominated by magnetite, hematite and quartz and has a thickness of 180m. The depositional environment was variable and ranged from a fairly deep, reducing basin to shallow, oxidizing water. The original mineralogy probably consisted of a silica-iron hydroxide gel which was precipitated due to oxidation of Fe2+to Fe3+.

The source of the Fe2+ was probably a distal hydrothermal input. Anoxic deep water allowed the transport of Fe2+over long distances to the site of iron formation deposition. When this deep, anoxic water mass came into contact with shallow water, which was enriched in oxygen due to the photosynthetic activity of cyanobacteria, a silica-iron hydroxide gel precipitated. Dewatering of this gel during burial caused the strong diagenetic overprint preserved in the section in which the original mineralogy is obscured. The high degree of diagenesis also explains the present low organic matter content of the Sokoman Iron Formation as well as the presence of Fe2+ minerals, since organic matter may have reacted as an electron acceptor reducing Fe3+ to Fe2+.

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ACKNOWLEDGEMENT

First I would like to thank my Supervisor Ed Spooner for making this thesis possible and for his support during the years.

Special thanks to Mike Gorton and Uli Wortman for being on my committee and for their scientific input.

The financial support by NSERC, the Fawcett Gittins Explorer Fund, the Moorehouse, the Ellesworth and the Gale Rucklidge scholarships are gratefully acknowledged.

SEG is thanked for the grant to allow the purchase of the iridium strip originally used to melt the glass beads.

I am very grateful to New Millennium Capital Corp. for allowing me to take samples from their cores. Special thanks to BK for the tremendous help in Schefferville in accessing the cores and showing me around in the field. Thanks also to Henry Simpson for taking me into the field.

I also owe great thanks to Ann Wilson from the Geological Survey of Ontario for taking me into the field area in Wawa and to Essar Steel Algoma Inc. for giving me access to their property.

Special thanks to Phil Thurston and Phil Fralick for their invaluable advice in the field.

TABLE OF CONTENTS

CHAPTER 1: INTRODUCTION AND CURRENT STATE OF KNOWLEDGE...... 1 1.1 Research objectives ...... 1 1.3 Iron Formations: current state of knowledge ...... 2 1.1.1 The Early Ocean ...... 4 1.1.2 The Early Atmosphere ...... 6 1.1.3 End of Iron Formation Deposition ...... 8 1.2. The deposition of iron formations...... 9 1.2.1 The role of Large Igneous Province (LIP) emplacement ...... 9 1.2.2 The role of hydrothermal input ...... 11 1.2.4 Precipitation along a chemocline? ...... 13 CHAPTER 2: ANALYTICAL METHODS ...... 15 2.1 XRF ...... 15 2.2 ICP-MS ...... 16 13 2.3 Ccarb ...... 24 CHAPTER 3: THE GEOLOGY AND PETROGRAPHY OF THE HELEN AND SOKOMAN IRON FORMATIONS ...... 25 3.1 The Helen Iron Formation, Wawa, Ontario ...... 25 3.1.1 General description and geology ...... 25 2.1.2 Petrology...... 31 2.2 The Sokoman Iron Formation, Labrador Trough ...... 52 2.2.1 General description and geology ...... 52 2.2.2 Petrology...... 60 CHAPTER 4: THE MAJOR AND MINOR ELEMENT GEOCHEMISTRY OF THE HELEN IRON FORMATION ...... 81 4.1 Major elements...... 81 3.2 Trace elements ...... 90 3.3 Ratios ...... 95 v

3.3.1 Manganese against iron ...... 95 3.3.2 Magnesium against iron ...... 96 3.3.3 Calcium and potassium against iron ...... 97 3.3.4 Barium against potassium...... 99 3.3.5 Aluminium against titanium ...... 100 3.3.6 Vanadium against titanium ...... 101 3.4 Discussion and conclusions ...... 102 CHAPTER 5: THE RARE EARTH ELEMENT GEOCHEMISTRY OF THE HELEN IRON FORMATION ...... 103 5.1 Introduction and theory ...... 103 5.1.1 The chemical behavior of the rare earth elements ...... 103 5.1.2 Normalisation ...... 103 5.1.3 Modern seawater ...... 104 5.1.4 Ancient seawater ...... 105 5.1.5 Post-depositional modifications ...... 107 5.1.6 Examples of REE distribution in iron formations ...... 109 5.2 The Rare Earth Element geochemistry of the 2.75 Ga old Helen Iron Formation ... 112 5.2.1 REE patterns in the Helen Iron Formation ...... 113 4.2.2 Total REE ...... 118 4.2.4 The europium anomaly ...... 119 CHAPTER 6: THE RARE EARTH GEOCHEMISTRY OF THE SOKOMAN IRON FORMATION ...... 123 6.2 REE patterns of the Sokoman iron formation ...... 123 6.2.1 The Ruth Shale (RS) ...... 124 6.2.2 The Lower Red Green Chert formation (LRGC) ...... 125 6.2.3 The Lower Green Chert formation (LGC) ...... 126 6.2.4 The Pink Grey Chert member (PGC) ...... 127 6.2.5 The Upper Red Chert member (URC) ...... 128 6.2.6 The Green Chert (GC) ...... 129 6.2.7 The Jasper Upper Iron Formation (JUIF) ...... 130 vi

5.2.8 The Lean Chert member (LC) ...... 131 6.2.9 The Menihek Shale ...... 132 6.3 Total REE ...... 133 6.4 The europium anomaly ...... 134 6.4 The cerium anomaly ...... 135 CHAPTER 7: THE CARBON ISOTOPE GEOCHEMISTRY OF THE HELEN IRON FORMATION ...... 137 13 7.1 C in iron formations ...... 137 7.1.1 Possible explanations for negative carbon isotope values ...... 138 7.2 The mineralogy of the Helen Iron Formation ...... 142 7.3 The 13C geochemistry of the Helen Iron Formation ...... 144 7.2.1 Previous research ...... 145 7.2.2 This study ...... 146 CHAPTER 7: DISCUSSION AND CONCLUSIONS ...... 150 REFERENCES ...... 152

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LIST OF FIGURES AND TABLES

Chapter 1 Figure 1.1: Distribution of iron formations over geologic time (Beukes and Gutzmer, 2008) Figure 1.2: Hydrothermal model (Steinhoefel et al., 2009)

Chapter 2 Table 2.1: Precision of the LA-ICP-MS data Table 2.2: Accuracy for the standards UTB2, FER1-4 Table 2.3: Comments for INAA data Table 2.4: LA-ICP-MS data compared to INAA of three samples of the Helen iron formation Table 2.5: Comparison of LA-ICP-MS data and XRF data obtained on the Helen Iron Formation Figure 2.6: Time-dependent reaction of sample SJ_10_10.5

Chapter 3 Figure 3.1: Location of the Helen Iron Formation, Wawa, Ontario Figure 3.2: Stratigraphy of the Helen Iron Formation, Sir James Dunn Pit Figure 3.3: a: Top of siderite facies, Lucy Pit b: Breccia facies, Lucy Pit c: Breccia facies, Sir James Dunn Pit d: Graphite facies with sill to the right, Sir James Dunn Pit e: Graphite facies with pyrite, Sir James Dunn f: Sir James Dunn pit Figure 3.4: Thin section of SJ_10_0.1, siderite facies (Sir James Dunn Pit) Figure 3.5: Thin section of SJ_10_5, siderite facies (Sir James Dunn Pit) Figure 3.6: Thin section of SJ_10_10.5, siderite facies (Sir James Dunn Pit) Figure 3.7: Thin section of SJ_10_15, siderite facies (Sir James Dunn Pit) Figure 3.8: Thin section of SJ_10_21, siderite facies (Sir James Dunn Pit) Figure 3.9: Thin section of SJ_10_25, siderite facies (Sir James Dunn Pit) viii

Figure 3.10: Thin section of SJ_10_30.5, siderite facies (Sir James Dunn Pit) Figure 3.11: Thin section of SJ_10_35, siderite facies (Sir James Dunn Pit) Figure 3.12: Thin section of SJ_10_40, siderite facies (Sir James Dunn Pit) Figure 3.13: Thin section of SJ_10_44.5, sulfide facies (Sir James Dunn Pit) Figure 3.14: Thin section of SJ_10_50, breccia facies (Sir James Dunn Pit) Figure 3.15: Thin section of SJ_10_55.5, breccia facies (Sir James Dunn Pit) Figure 3.16: Thin section of SJ_10_60, breccia facies (Sir James Dunn Pit) Figure 3.17: Thin section of SJ_10_65, breccia facies (Sir James Dunn Pit) Figure 3.18: Thin section of SJ_10_70.5, breccia facies (Sir James Dunn Pit) Figure 3.19: Thin section of SJ_10_75.5, breccia facies (Sir James Dunn Pit) Figure 3.20: Thin section of SJ_10_80A, breccia facies (Sir James Dunn Pit) Figure 3.21: Thin section of SJ_10_84.5, breccia facies (Sir James Dunn Pit) Figure 3.22: Thin section of SJ_10_90, breccia facies (Sir James Dunn Pit) Figure 3.23: Thin section of SJ_10_95, sill (Sir James Dunn Pit) Figure 3.24: Location of the Labrador Trough, Canada Figure 3.25: Stratigraphic column of the Howell’s River Area, Labrador Trough Figure 3.26 a: Ruth Shale b: Lower Iron Formation c: Lower Red Green Chert d: Lower Red Green Chert e: Lower Green Chert f: Pink Grey Chert g: Upper Red Chert h: Green Chert i: Jasper Upper Iron Formation j: Jasper Upper Iron Formation, field sample k: Lean Chert l: Lean Chert with Stromatolites

Figure 3.27: Howell’s River Area, view towards South Figure 3.28: Thin section 09-FH 3-3a, Lower Iron Formation Figure 3.29: Thin section 09-FH 3-4a, Lower Iron Formation Figure 3.30: Thin section 09-FH 3-4b, Lower Iron Formation Figure 3.31: Thin section 09-FH 3-4c, Lower Iron Formation Figure 3.32: Thin section 09-FH 5-P1a, Lower Red Chert Figure 3.33: Thin section 09-FH 5-P1b, Lower Red Chert Figure 3.34: Thin section 09-FH 5-P1c, Lower Red Chert Figure 3.35: Thin section 09-FH 6-1a, Pink Grey Chert Figure 3.36: Thin section 09-FH 6-1b, Pink Grey Chert Figure 3.37: Thin section 09-FH 7-5a, Upper Red Chert Figure 3.38: Thin section 09-FH 7-5b, Upper Red Chert Figure 3.39: Thin section 09-FH 7-5c, Upper Red Chert Figure 3.40: Thin section 09-FH 7-13a, Upper Red Chert Figure 3.41: Thin section 09-FH 7-13b, Upper Red Chert Figure 3.42: Thin section 09-FH 7-13c, Upper Red Chert Figure 3.43: Thin section 09-FH 8-4, Green Chert Figure 3.44: Thin section 09-FH 8-7, Green Chert Figure 3.45: Thin section 09-FH 9-6, Jasper Upper Iron Formation Figure 3.46: Thin section 09-FH 10-S1, Lean Chert with Stromatolites Figure 3.47: Paragenetic diagram of the Sokoman Iron Formation

Chapter 4

Figure 4.1: SiO2 (%) against stratigraphic height Figure 4.2: FeO (%) against stratigraphic height

Figure 4.3: TiO2 (%) against stratigraphic height

Figure 4.4: Al2O3 (%) against stratigraphic height Figure 4.5: MnO (%) against stratigraphic height Figure 4.6: MgO (%) against stratigraphic height

Figure 4.7: CaO (%) against stratigraphic height

Figure 4.8: K2O (%) against stratigraphic height Figure 4.9: V (ppm) against stratigraphic height Figure 4.10: Cr (ppm) against stratigraphic height Figure 4.11: Ni (ppm) against stratigraphic height Figure 4.12: Zr (ppm) against stratigraphic height Figure 4.13: Ba (ppm) against stratigraphic height Figure 4.14: Cu, Pb and Zn (ppm) against stratigraphic height Figure 4.15: MnO against FeO (numbers correspond to sample names) Figure 4.16: Plot MgO against FeO Figure 4.17: CaO agains FeO Figure 4.18: Potassium against iron Figure 4.19: Plot of barium against potassium as an indicator for the detrital nature of the first 10 meters of the section. Figure 4.20: Plot of aluminium against titanium Figure 4.21: Plot of vanadium against titanium

Chapter 5 Figure 5.1: REE for Isua, Greenland (Dymek and Klein, 1988) Figure 5.2: REE for the Kuruman IF, South Africa (Klein and Beukes 1989) Figure 5.3: REE pattern of the two rhyolitic ash beds (F1 and F3 rhyolite), Helen Iron Formation Figure 5.4: REE pattern of the lower part of the siderite facies, Helen Iron Formation Figure 5.5: REE pattern of the upper part of the siderite facies, Helen Iron Formation Figure 5.6: REE pattern of the sulfide and breccia facies, Helen Iron Formation Figure 5.7: REE pattern of the sill found at the top of the breccia facies, Helen Iron Formation Figure 5.8: REE pattern of the basaltic unit overlying the Helen Iron Formation Figure 5.9: Total REE values plotted against stratigraphic height, Helen Iron Formation Figure 5.10: The europium anomaly plotted against stratigraphic height, Helen Iron Formation. Figure 5.11: REE pattern of modern North-Pacific seawater

Chapter 6 Figure 6.1: REE pattern of the Ruth shale Figure 6.2: REE pattern of the Lower Red Green Chert Figure 6.3: REE pattern of the Lower Green Chert Figure 6.4: REE pattern of the Pink Grey Chert Figure 6.5: REE pattern of the Upper Red Chert Figure 6.6: REE pattern of the Green Chert Figure 6.7: REE pattern of the Jasper Upper Iron Formation Figure 6.8: REE pattern of the Lean Chert Figure 6.9: REE of the Menihek Shale Figure 6.10: Total REE against stratigraphic height of the Sokoman Iron Formation Figure 6.11: Europium anomaly against stratigraphic height Figure 6.12: Cerium anomaly against stratigraphic height

Chapter 7 13 Table 7.1: Some examples of Ccarb values from iron formations around the world Table 7.2: Average major elements in chlorite, ankerite, siderite, measured on the microprobe Figure 7.1: Estimated mineral percentages of the Helen Iron Formation based on major element and microprobe analyses Figure 7.2: Estimated carbonate percentages of the Helen Iron Formation based on major element and microprobe analyses. Figure 7.3: Locations of the two sections of the Helen Iron Formation Figure 7.4: Profile of the Sir James Dunn pit and the MacLeod mine (data from Thode and Goodwin, 1983).

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LIST OF APPENDICES

XRF data Helen Iron Formation

XRF data Sokoman Iron Formation

Semi-quantitative microprobe data Helen Iron Formation

INAA data Helen Iron Formation

LA-ICP-MS data Helen Iron Formation

LA-ICP-MS data Sokoman Iron Formation

13C and 18O data Helen Iron Formation

Thin section photographs Helen and Sokoman Iron Formation

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CHAPTER 1: INTRODUCTION AND CURRENT STATE OF KNOWLEDGE

1.1 Research objectives

This study compares an example of a low metamorphic grade Algoma-type iron formation (Helen Iron Formation) with an age of 2.75 Ga with a younger, Superior-type iron formation (Sokoman Iron Formation) with an age of 1.88 Ga.

The key objective was to compare their mineralogy and geochemistry in order to gain information about their depositional environment, the role that diagenesis played and the significance and nature of hydrothermal input during deposition.

Questions to be addressed were whether or not the iron formations represent direct precipitates from seawater and the role of diagenetic modifications. Another important question is the importance of detrital input.

Samples were selected through the iron formation sections ( 0.5 m interval for the Helen; 1 m interval for the Sokoman Iron Formation) in order to provide high resolution geochemical and mineralogical information.

Major and minor element data are used to trace detrital as well as hydrothermal inputs into the Helen Iron Formation.

REE analyses are used to identify cerium and europium anomalies in the sequences and thus to infer the oxidation states of the iron formations as well as defining the type of detrital input.

Carbon isotope geochemistry is used to provide insights into Helen Iron Formation deposition.

1.3 Iron Formations: current state of knowledge

Iron formations are the source for iron ore, which in dollar value is the most valuable economic commodity worldwide. In 2011, 2.9 billion metric tons of iron ore were mined worldwide, mainly for steel production. 83% of the mined ore comes from China, Australia, Brazil, and India (Tuck and Virta, 2013) 90% of all known iron formations are located in the Hamersley region (Australia), the Quadrilatero Ferrifero area (Brazil), the Transvaal-Griquatown (South Africa), the Krivoy Rog (Ukraine) and the Labrador Trough in Canada (Trendall and Morris, 1983). No direct modern analogues are known.

Figure 1.1: Distribution of iron formations over geologic time (Beukes and Gutzmer, 2008) A: according to their size (in megatons) B: according to their number 2

Iron formations were defined by James (1954) as chemical sediments with high contents of iron

(>15% Fe) and silica, typically 34-56% SiO2. Their deposition took place in the Archean as well as in the early Proterozoic. The oldest are found in Isua, Greenland at 3.8 Ga, and deposition appears to have been globally nearly continuous until 1.85 Ga when they disappear from the geological record until reappearing from 0.8 - 0.6 Ga in association with late Proterozoic glaciogenic sequences (e.g. Klein, 2005). The peak of preserved deposition was reached at 2.5 Ga (see Figure 1.1).

Typical overall mineral assemblages in low metamorphic grade iron formations include silica (as chert), magnetite, hematite, iron carbonates (siderite and dolomite-ankerite series), greenalite, stilpnomelane, and locally pyrite.

According to Klein (2005), the precursor to iron formations is believed to have consisted of hydrothermal muds containing ferrous iron-rich silicates (smectite) and carbonates (siderite). According to Bekker et al. (2010) chert is early diagenetic and not a direct precipitate from seawater but is interpreted to be the siliceous equivalent of modern-day sea-floor hardgrounds. The average depth of deposition has varied from 10’s of meters up to 700 m (Klein, 2005).

Two categories of Precambrian iron formations were identified by Gross (1965): the Superior- type and the Algoma-type. Algoma-type iron formations are typically small in size and thickness (typically less than 50 m thick and rarely extending for more than 10 km along strike). They are directly associated with mafic-ultramafic to felsic volcaniclastic rocks and greywackes in interpreted to be island backarc regions or intracratonic rift zones. They are found in Archean greenstone belts and their depositional environment might be comparable to modern day spreading ridges on the seafloor. The peak of Algoma-type iron formation deposition was reached at 2.75 to 2.70 Ga and corresponds to a major mantle plume event (e.g. Huston and Logan, 2004). They usually show distinct banding features and lack of shallow water structures which indicate deposition in deeper water below storm wave base; few shallow water successions of Archean iron formations are known (e.g. Fralick and Pufahl, 2006). Superior-type iron formations on the other hand are usually part of thick sedimentary sequences and cover large areas of up to 105 km2. They generally lack a direct association with volcanic rocks and the host sequences lie unconformably on older crustal rocks. Superior-type iron formations

generally show granular textures indicating that they were deposited on continental shelves probably during periods of global high sea-levels and pulses of enhanced magmatic and hydrothermal activity. They are interpreted to be a global and not a regional feature (Bekker et al., 2010) and are typically overlain/underlain by organic-matter rich shales and sulfidic shales (e.g. Klein and Beukes, 1989 and 1993; Beukes et al., 1990; Simonson and Hasler, 1996). Their first appearance is at 2.6 Ga and might be associated with the construction of large continents that changed the heat flux at the core-mantle boundary (Bekker et al., 2010).

1.1.1 The Early Ocean

In the modern ocean, the silica cycle is strongly influenced by the activities of skeleton-forming organisms such as diatoms, radiolaria or sponges. The precipitation of amorphous silica lowers dissolved silica concentrations in shallow marine waters to values well below the solubility product of the main sedimentary silica polymorphs opal-A, opal-CT and quartz (Siever, 1992). Before the evolution of silica-secreting organisms in the Precambrian however, higher dissolved oceanic silica concentrations, at or above the equilibrium with respect to opal-CT and possibly also to amorphous silica were present (e.g. Drever, 1974; Holland, 1984; Siever, 1992; Maliva et al., 2005).

The Precambrian silica-cycle was thus dominated by inorganic reactions among dissolved silica, clay and zeolite minerals. Today the concentration of silica in the ocean is less than 1 ppm, but the Precambrian values were closer to 60 ppm based on experimental determination of solubility and silica sorption on clay minerals and zeolites (Siever, 1992), with some estimates ranging up to 120 ppm (Holland, 1984), approaching supersaturation with respect to amorphous silica. Triggers leading to silica precipitation include cooling of high temperature solutions and changes in pH (Belevtsev et al., 1982). The source of silica was probably a combination of hydrothermal input and the low temperature weathering of continental crust (Hamade et al., 2003; Ding et al., 2004; Andre et al., 2006).

Ferric (Fe3+) oxyhydroxide particles are highly reactive towards dissolved silica (e.g. Konhauser et al., 2007) resulting in the scavenging of silica from seawater during particle sedimentation, the particles resting at the sediment-water interface before burial. These particles may then later be

released during diagenesis if ferric oxyhydroxides were transformed to stable crystalline compounds (Slack et al., 2007; Fischer and Knoll, 2009). A small fraction was not recycled back but precipitated as amorphous silica, which resulted in the deposition of early diagenetic chert.

Iron in the reduced form of Fe2+ is easily transported in the absence of oxygen since it is highly soluble in seawater. Oxidized iron (Fe3+) has a much lower solubility than Fe2+ and thus iron is precipitated rapidly when coming in contact with oxygen. In the modern ocean, Fe2+ from deep sea hydrothermal vents either reacts rapidly with dissolved sulfide and is deposited in the vent 3+ area as pyrite (FeS2) and other sulfides or is oxidized to ferric (Fe ) oxyhydr(oxide) and precipitated out (Poulton and Canfield, 2011). In the deep anoxic ocean of the Precambrian however; iron could be transported over long distances in solution as Fe2+. Estimations of the iron concentrations in the Precambrian ocean vary from 1 ppm (Holland, 1984) to 20 ppm (Ewers, 1980) to 400 ppm (Mel’nick, 1973). The absence of oxygen in seawater would also allow the direct precipitation of Fe2+ hydrous phases, whereas Fe3+ hydroxides (ferrihydrite) formed via the oxidation of Fe2+ by biologic or abiologic processes (e.g. Lepp and Goldich, 1964; Garrels et al., 1973; Cairnssmith, 1978; Johnson et al., 2008a).

Two possible sources for iron are proposed: a continental source and a hydrothermal source. Holland (1984) postulated a continental source model in which particulate iron enters the ocean via rivers and is then reduced in the deep sea via the remineralization of organic matter. Upwelling at continental margins leads to precipitation of iron in iron-oxides, iron-carbonates or iron-sulfides depending on the redox conditions of the environment.

However, a problem with this scenario is that the global production of organic matter and its abyssal remineralization has to be sufficient to remobilize deep-sea sedimentary iron but should not be so big as to increase atmospheric oxygen in the long term. Furthermore, the upwelling systems have to remain site-specific for periods of 200 to 800 Ma, and that is unlikely (Isley, 1995).

A hydrothermal origin for the iron on the other hand is supported by rare earth element data for both Algoma and Superior type iron formations (e.g. Dymek and Klein, 1988). They show a distinct positive europium anomaly, thus indicating submarine hydrothermal discharge as the

iron source. The highest europium anomaly so fur recorded was in the Deloro Iron Formation, south of Timmins, Canada with values ranging up to 30 (Thurston et al., 2012).

Sea Surface Temperatures

Estimations of the sea surface temperature by Knauth and Lowe (2003) based on oxygen isotopes in chert of the 3.5-3.2 Ga old Barberton Greenstone Belt, South Africa give rather high seawater values of 55-85 °C. Silicon isotope data by Robert and Chaussidon (2006) indicate similar temperatures of 70 °C at 3.5 Ga and 20 °C at 800 Ma. Gaucher et al. (2008) showed similar high temperatures from protein thermostabilities.

However, it has been shown by van den Boorn et al. (2007) that the isotope variability can also be explained by variation in source fluids and cannot be used as a paleotemperature indicator. In addition, it is very difficult to distinguish primary sediment/water interface temperatures from diagenetic temperatures.

Studies on combined oxygen and hydrogen isotopes by Hren et al. (2009) on the 3.42 Ga old Buck Reef Chert in South Africa showed that the Archean surface temperatures were below 55 °C and probably in a range comparable to modern sea surface temperatures.

1.1.2 The Early Atmosphere

At 4.4 Ga, the sun was 20-30% less luminous than today (Newman and Rood, 1977). Nevertheless, there are no indications that the early earth was frozen. This problem, known as the ‘Early Faint Sun Problem’, indicates that the composition of the early atmosphere was different from that of today’s resulting in a much more efficient greenhouse climate. Different greenhouse gases have been put forward to compensate for the less luminous sun; two examples are methane and ethane (e.g. Kasting, 2005; Haqq-Misra et al., 2008; Rye et al., 1995).

Another greenhouse candidate proposed by Ohmoto et al. (2004) is CO2. Large, siderite-rich beds in pre-1.8 Ga old sediments indicate that atmospheric CO2 levels were more than 100 times greater than today. This resulted in rain and ocean being more acidic than today.

The Great Oxidation Event

The Great Oxidation Event (GOE) marks the switch from an early anoxic to an oxic atmosphere at around 2.32 Ga (Bekker et al., 2004; Canfield, 2005). Evidence for the Great Oxidation Event can be found in various geological settings such as the occurrence of red beds (Chandler, 1980),

CaSO4-rich evaporates (Chandler, 1988; El Tabakh et al., 1999), oxidized shallow water iron formations (Beukes and Klein, 1992), and oxidized paleosols (Rye and Holland, 1998) after 2.2 Ga.

Isotopic proof for the Great Oxidation Event includes highly positive 13C values, which indicate the onset of large-scale oxygen-producing photosynthesis that sequestered large quantities of 12C from the inorganic ocean-atmosphere reservoir into sedimentary organic matter (Karhu and

Holland, 1996). The atmospheric reservoir of O2 and a complementary residual reservoir of carbon enriched in 13C were consequences (Frauenstein et al., 2009). Furthermore, a negative shift in 34S values between 2.4 and 2.3 Ga indicates an increase in ocean sulfate concentrations which is ascribed to an increased rate of oxidative weathering of pyrite in crustal rocks during and after the Great Oxidation Event (Canfield, 2005; Bekker et al., 2004).

The Great Oxidation Event was either the consequence of the evolution of oxygenic photosynthesis and the burial of organic matter (Kopp et al., 2005) or the consequence of an abiotic shift in the balance of oxidants and reductants at the Earth’s surface (e.g. Holland, 2002; Claire et al. 2006; Goldblatt et al., 2006; Kump and Barley, 2007). Another theory suggests that atmospheric oxygenation was triggered by enhanced magmatism at 2.5 to 2.45 Ga (Bekker et al., 2010).

It is suggested that the deep ocean continued to be anoxic after 1.8 Ga until a second major rise in atmospheric oxygen at 800-580 Ma terminated deep ocean anoxia (Derry et al. 1992; Canfield and Teske, 1996). Geochemical evidence such as sulfur isotope data from the Hamersley (Partridge et al., 2008) or enrichments of the redox-sensitive elements molybdenum and rhenium (Anbar et al., 2007) indicates an oxygenated shallow water mass 50 to 150 Ma before the Great Oxidation Event; deep water below a surface mixing zone remained anoxic.

1.1.3 End of Iron Formation Deposition

Holland (1984) and Cloud (1972) believed that the end of the deposition of iron formations was caused by the oxygenation of the deep water, which then removed Fe2+ as ferric (hydro)oxides from the ocean system.

Another theory by Poulton et al., (2004) suggests that the deep ocean became euxinic after 1.85 Ga and thus prevented the deposition of iron formation. Sulphur isotopes and sedimentary iron chemistry indicate that the increase in atmospheric oxygen resulted in enhanced sulfide weathering on land which led to an enhanced flux of sulfate to the ocean. This caused Fe2+ to be removed as pyrite as the ocean became sulphidic (Derry et al., 1992). Dissolved Fe2+ and dissolved H2S cannot coexist because of the insolubility of sulfide phases (Reinhard et al., 2009). Proof for euxinic conditions until 1 Ga can also be found from redox-sensitive molybdenum isotopes sulfide (e.g. Emerson and Huested, 1991; Erickson and Helz, 2000; Zheng et al., 2000; Morford and Emerson, 1999). Molybdenum enters the ocean via rivers and exists in oxygenated 2- water as the molybdate anion MoO4 , which is very unreactive. Therefore molybdenum is the most abundant transition metal in the modern ocean. In euxininc conditions however, molybdenum is removed from solution as sulfide. Data by Arnold et al. (2004) indicate that the area of oxic sedimentation in the mid-Proterozoic was approximately 10 times smaller than today’s and the area of euxinic sedimentation was approximately 10 times larger.

Another theory by Slack and Cannon (2009) suggests that the Sudbury meteorite impact was related to the cessation of iron formation deposition. This is based on the fact that the Gunflint Iron Formation in NW Ontario, Minnesota and Michigan shows the Sudbury impact layer at the top of the iron formation. According to this theory, the impact at 1.85 Ga caused global mixing of shallow oxic and deep anoxic water which resulted in a suboxic redox state for the deep water. Even the low oxygen concentration of 1 M would have prevented the transport of hydrothermally derived Fe2+ from the deep ocean to continental margins and so not allow for the deposition of iron formation.

On the other hand, Rasmussen et al. (2012) suggested that the end of the deposition of iron formations corresponds with the termination of the short-lived interval of global mantle-driven

magmatism and crustal growth. After 1.88 Ga the deposition of iron formation was no longer favored because the flux of oxidants from the atmosphere and the surface ocean significantly exceeded the long-term rate of delivery of hydrothermal iron and reductants to the deeper part of the ocean.

1.2. The deposition of iron formations

Iron formations deposited before the Great Oxidation Event (GOE) were precipitated in deeper basins and most likely redistributed by density currents (Krapez et al., 2003). Bekker et al., (2010) believe that hydrothermal discharge played a major role in the deposition of early iron formations, and oxidation was caused by anoxygenic photosynthetic bacteria. Post-GOE iron formations were deposited due to upwelling and biological and non-biological oxidation of ferrous iron at the redoxcline. Later the particles were transported back into the basin by storm/wave currents. Iron oxides in granular iron formations became dominant only after the GOE (Bekker et al., 2010).

There is also much discussion in the literature regarding whether or not iron formations are direct precipitates from seawater or diagenetic replacements of carbonaceous rocks (e.g. Kimberley, 1974; Dimroth, 1972).

1.2.1 The role of Large Igneous Province (LIP) emplacement

Formerly, it was believed that iron formations could only be deposited on submarine plateaux with very little magmatic activity (e.g. Trendall and Blockley, 1970; Morris and Horwitz, 1983). Newer evidence by Condie (2002) and Condie et al. (2009), however, suggests that magmatic and hydrothermal activity played a major part in the deposition of these enigmatic sediments. Mantle plumes that led to the formation of Large Igneous Provinces (LIP) would also enhance spreading rates at mid-ocean ridges and produced higher growth rates of ocean plateaux and generation of seamounts, all of which would contribute to higher hydrothermal fluxes to the ocean as well as enhanced ocean anoxia and marine transgressions. Also, vast volumes of metals and metalloids would have been introduced into the ocean, leading to a peak in VMS deposition

(Franklin et al., 2005). When the hydrothermal flux overwhelmed the oceanic oxidation state, iron would have been transported and deposited distally from hydrothermal vents (Isley and Abbott, 1999; Condie, 2002; Bekker et al., 2010). Barley et al. (1997) believe that the deposition of the gigantic Hamersley Iron Formation is contemporaneous with the emplacement of more than 30,000 km3 of dolerite, basalt and rhyolite. The authors showed that iron formations formed during major tectonic-magmatic events, and were deposited at rates comparable to modern pelagic sediments. They are interpreted to have formed as a result of the increased supply of suboxic iron- and silica-rich seawater upwelling onto continental shelves during pulses of increased submarine magmatic and hydrothermal activity. Condie et al. (2009) showed that a magmatic ‘shutdown’ at 2.45 Ga for 200-250 myr coincides with a gap in the deposition of iron formations as well as a major drop in sea level and the onset of widespread glaciations at 2.4-2.3 Ga.

The reappearance of iron formations at 1.88 Ga is contemporaneous with peaks in global mafic- ultramafic magmatism (Heaman et al., 2009; Meert et al., 2011). This reappearance is marked by extensive mafic and ultramafic magmatism across the world, emplacement of dyke swarms and sills with Ni-Cu-PGE mineralization, locally basaltic flows in the Superior, Wyoming and Slave cratons of North America; the Dharwar and Bastar cratons in India; the Siberian craton as well as the Kaapvaal and Zimbabwe cratons in South Africa and Baltica (Heaman et al., 2009; Meert et al., 2011). Furthermore there is a peak in juvenile continental and ocean crust formation (Condie, 1998; Kemp et al., 2006), mantle depletion (Pearson et al. 2007; Parman, 2007), and volcanogenic massive sulfide deposition (Bekker et al., 2010; Franklin et al., 2005). New continental crust generation was linked with large scale mantle melting (Pearson et al., 2007; Parman, 2007).

Condie et al. (2001) believed that the mantle superplume event and supercontinent formation at

2.7 and 1.9 Ga introduced large amounts of CO2 into the system. This resulted in an increase in the depositional rate of carbon and to a global warming event. The increased black shale deposition reflects a combination of increased oceanic hydrothermal fluxes introducing nutrients, anoxia on continental shelves and a disruption of ocean currents.

1.2.2 The role of hydrothermal input

The Archean Earth was hotter than today. According to estimation of Dickinson and Luth (1971), the global heat flow was two to three times higher for the time period of 2.7 to 1.8 Ga. The mid-ocean ridge length should thus have been greater and/or spreading rates were higher leading to higher oceanic crust production. Isley (1995) estimated the hydrothermal cycling rate to have been 2.6 to 4.2 times higher than today.

The morphology of mid-ocean ridges is a function of the spreading rate (Menard, 1967) and the thickness of the lithosphere (Sleep and Rosendahl, 1979). The Archean was dominated by thinner continents which resulted in an average ocean depth estimated to be in the range of 2-3 km. The mid-ocean ridge depth was therefore 500-150 m. Much of the Archean continental land mass was submerged which is consistent with a general lack of platform sediments on Archean cratons and the immature first cycle character of most Archean sandstones (Ernst, 1983). 3.7-2.7 Ga was the time of the beginning of early continental crustal evolution and sedimentation is of greenstone-type. The major components are volcanics, sediments are minor (Eriksson et al., 2007). Hot hydrothermal water emitted from black smoker vents rises buoyantly through the water column, mixing with ambient seawater until a level of neutral buoyancy is reached, where the plume spreads out horizontally. The level of dynamic equilibrium is a function of the temperature and salinity gradients through the water column and salt, heat and volume fluxes of both hydrothermal waters and diluting seawater (Speer and Rona, 1989). Silica precipitates contemporaneously from the entire water column whereas ferric oxyhydroxides only precipitate if the hydrothermal discharge rises above the chemocline.

Steinhoefel et al. (2009) proposed the following model for the deposition of the Algoma-type iron Formation of the Wanderer Group, Zimbabwe with an age of 2.7 Ga. In low hydrothermal activity phases (Figure 1.2, A), all Fe2+ that reaches over the chemocline gets oxidized by photosynthetically produced oxygen. Silica (open circles), ferric oxyhydroxide (black dots) and organic matter (grey ellipsoids) get buried together. As the amount of organic matter exceeds the amount of ferric oxyhydroxide, a complete reduction of iron and conversion into iron carbonates (open diamonds) occurs during diagenesis, forming carbonate-chert layers.

Figure 1.2: Hydrothermal model (after Steinhoefel et al., 2009) Silica precipitation occurs constantly in both low and high hydrothermal activity mode (open circles). A: low hydrothermal activity mode: small amounts of Fe2+ rise over the chemocline to get oxidized by photosynthetically produced oxygen and form ferric oxyhydroxides (black dots). Diagenesis then leads to complete reduction of ferric oxyhydroxides at the expense of organic matter (grey ellipsoids) to form carbonate (open diamonds). B: high hydrothermal activity mode: high amounts of Fe2+ rise over the chemocline to get oxidized to ferric oxyhydroxides (black dots). Since the amount of ferric oxyhydroxides exceeds the amount of organic matter (grey ellipsoids), only part of the ferric oxyhydroxides get reduced to form carbonate (open diamonds), the rest of the ferric oxyhydroxides forms magnetite (black squares).

In times of high hydrothermal discharge (Figure 1.2, B), large quantities of Fe2+ reach the oxidizing zone, leading to the oxidation of Fe2+ and precipitation of significant amounts of ferric hydroxide (black dots). As the amount of ferric hydroxide in this scenario exceeds the amount of organic matter produced (grey ellipsoids), the ferric hydroxide only gets partially reduced to form iron carbonates (open diamonds), leading to the deposition of magnetite-carbonate-chert layers.

One of the problems with this model is the need for a rather extensive oxygenated surface layer. Geological evidence indicates a reducing atmosphere prior to 2.32 Ga (Bekker et al., 2004). An extensive oxygenated surface layer cannot exist with a reducing atmosphere. A possible explanation could be that the oxygen content of the surface layer was very low but maintained by photosynthesis and that oxygen would immediately be removed either by oxidation of organic

matter in times of low hydrothermal activity or by the oxidation of iron in high hydrothermal activity phases.

Alternatively, it was suggested by Carrigan and Cameron (1991) that pulses of hydrothermal activity were the trigger for the deposition of iron formation since it provided dissolved iron and possibly silica, which helped buffer oxygen and sulfate to low levels. An increase in hydrothermal activity led to an upward movement of the anoxic/oxic boundary and iron was transported to the shallow shelf where it precipitated as siderite, iron hydroxides, iron-silicates or pyrite depending on physio-chemical conditions. Siderite would therefore be a primary mineral, precipitated from seawater from pore waters close to the sediment-water interface. Ankerite formed after siderite as a secondary phase. Pyrite formed during early diagenesis at or just below the sediment-water interface coeval with siderite.

1.2.4 Precipitation along a chemocline?

Beukes and Gutzmer (2008) proposed the following depositional model for the Kuruman and Hamersley Iron Formations: Distal or deeper water masses hydrothermally enriched in iron and silica only occasionally transgressed onto the deep-carbonate shelf, thus the deposition of iron and silica occurred only occasionally and was restricted to distal deep-water carbonate shelf environments along a chemocline. The ocean was depleted in dissolved iron for most of the time; the chemocline was rather broad and located below the storm-wave basis. The depth of the chemocline was determined by the rate of supply of oxygen and organic matter from the water column above it and dissolved ferrous iron from the water column below. Increases in hydrothermal discharge lead to the upward movement or shifts of the chemocline. Anaerobic photoautotrophic Fe2+ oxidizing bacteria were not responsible for the precipitation of ferric oxyhydroxides because the redox boundary for iron was situated far below the photic zone in the water column. Instead it is suggested by the authors that the electron acceptor was dissolved oxygen produced by cyanobacteria living in surface water close to the shore where primary productivity was high. Original ferric (Fe3+) oxyhydroxides were transformed to siderite or magnetite (if organic matter was present) via dissimilatory iron reduction which reduces Fe3+ to Fe2+ at the expense of organic matter (see also Chapter 6).

CHAPTER 2: ANALYTICAL METHODS

2.1 XRF

Whole rock samples were powdered, made into pellets and analyzed on a Philips PW2404 X-ray fluorescence spectrometer. The basic principle of an XRF is to irridate the sample to be analyzed (rock powder) with X-rays. The elements in the sample then emit characteristic secondary X-rays (fluorescence), the intensities of which are correlated to the concentration.

XRF analysis provides good results for major and certain minor elements with detection limits typically down to 1 ppm. A new XRF routine was established, which measured the major elements as well as vanadium, chromium, nickel, copper, zinc, arsenic, strontium, zirconium, barium and lead.

In order to make sure that no important minor elements were overlooked, semi-quantitative analysis of selected samples was performed, scanning the entire X-ray spectrum.

Precision

Precision was evaluated by repeating measurements and by comparison with semi-quantitative analysis of the same samples.

Average precisions are:

± 1% for major elements

±1.5% for minor elements

Accuracy

Accuracy was established by comparison with international reference material, including four iron formation rock standards and is in the same range as the precision.

2.2 ICP-MS

Inductively coupled plasma mass spectrometry (ICP-MS) was originally developed as a highly sensitive technique for analyzing solutions and can detect almost all the elements in the periodic table.

High sensitivity was achieved by the use of a high temperature argon plasma.

However, when analyzing rocks, two problems arise:

1) Although almost all elements can be detected, there is no solution chemistry which can keep all elements in solution at the same time.

2) The nebulizer limitations require samples to be diluted to at least 1g/100ml, thus losing sensitivity.

A solution to this problem is to use a laser to vaporize the sample (LA-ICP-MS) which is then transported to a quadrupole mass spectrometer with a helium carrier gas.

The disadvantage of this method is the so called nugget-effect., because of the small sample size of few tens of milligrams. Since REE are commonly found in small, REE-rich accessory minerals, caution has to be taken when using LA-ICP-MS.

Several solutions to this problem have been proposed. For example, fused glass discs can be produced with lithium tetraborate flux. The disadvantage of this approach is contamination of the sample with Be, Sc, Sr, La, Ti,V, Cr, Ni, Cu, Zn, Y, Zr, Ba, Nd and Pb which are present in the flux at levels up to several g/g (Sylvester, 2001). Another disadvantage is the preferential loss of highly volatile elements (e.g. Pb) and loss of some elements by alloying with the crucible (e.g. Ni, Cu, Ta). In addition, the dissolution of the rock powder with a flux leads to worse detection limits as well as a Li and B contamination of the ICP-MS.

These problems can be overcome by melting small amounts of whole-rock powder on a tungsten strip heater. This method provides the possibility for a flux-free fusion technique for highly precise LA-ICP-MS bulk analysis. 16

Approximately 50 mg of whole rock powdered sample were placed on a 0.1 mm thick, 1cm wide and 2.5cm long tungsten strip connected with a high current (>100 A) electrical source to produce resistance heating in order to melt the sample into a glass bead. Tungsten was chosen because of its high melting temperature of 3422°C and its low cost (approximately $ 5 per strip). Experiments were also carried out with iridium (according to the method of Stoll et al., 2008), but since the performance was very similar and the costs much higher ($ 1,050 for the same size iridium strip), tungsten was used. In order to suppress oxidation and the loss of volatile elements such as K, Na, Li and B, melting was carried out in an argon overpressured atmosphere.

When the sample was melted (at 1500-2000°C, depending on sample composition), the current was turned off and at the same time a high argon gas flow was directed on the bottom of the strip to quench the sample.

Fedorowich et al. (1993) showed that most elements can be analyzed with a precision of ±5%. This result is supported by this study. They also showed contamination of Hf and Ta from the tungsten strip; however, this study found no such contamination with W from Alfa Aesar.

This method proved to be fast and effective with few disadvantages. Sample amounts can be very low, down to 10 mg. A disadvantage is that Pb, Sn, Ge and In are lost during this procedure because of their volatility.

The melted glass beads were analyzed on a VG PQ ExCell with a Nd-YAG laser with a laser beam diameter of 20 m. The following isotopes were chosen for analysis:

Phosphorous 31; calcium 44; scandium 45; vanadium 51; chromium 52; iron 57 and 58; cobalt 59; nickel 60; copper 65; zinc 55; strontium 88; yttrium 89; niobium 93; molybdenum 95; tin 118; barium 137 and 138; lanthanum 139; cerium 140; praseodymium 141; neodymium 146; samarium 147; europium 153; gadolinium 157; terbium 159; dysprosium 163; holmium 165; erbium 166; thulium 169; ytterbium 172; lutetium 175; tantalum 181; lead 208; thorium 232; uranium 238.

International standard NIST 610 was used to correct for instrumental drift at the beginning and end of each run.

Internal standards used strontium for the Sokoman Iron Formation and manganese for the Helen Iron Formation, determined by XRF analysis.

Precision

Each measurement was taken duplicated using two adjacent 0.5 mm long tracks.

element average precision in % element average precision in % P31 9.2 La139 3.4 Ca44 1.5 Ce140 4.1 Sc45 8.4 Pr141 3.6 V51 2.8 Nd146 3.9 Cr52 23.8 Sm147 5.9 Mn55 1.6 Eu153 5.2 Fe57 2.3 Gd157 7.1 Fe58 2.6 Tb159 7.4 Co59 8.0 Dy163 6.3 Ni60 19.7 Ho165 6.5 Cu65 17.6 Er166 5.5 Zn66 26.0 Tm169 9.3 Sr88 1.0 Yb172 6.7 Y89 4.5 Lu175 8.9 Zr90 3.7 Ta181 5.5 Nb93 5.6 Pb208 42.3 Mo95 13.5 Th232 9.5 Sn118 16.0 U238 8.1 Ba137 3.4 Ba138 3.1

Table 2.1: Precision of the LA-ICP-MS data in percent. Most elements have a precision within ±10%., exceptions are chromium, nickel, copper, zinc, tin and lead

As seen in Table 2.1, precision for most elements is within ±10%.

Exceptions are due to the fact that the XRF has an instrumental blank for elements such as chromium and nickel where x-rays from the sample strike parts of the intrustment. Variations in sample compositions produce a variable blank which can not be precisely corrected.

Pb: lead is a highly volatile element and is therefore lost in variable amounts during the melting of the sample on the tungsten strip.

The precision for the rare earth elements is very good, especially for the light rare earth elements. Since rare earth elements are usually plotted on a logarithmic scale, this precision delivers very good data.

Accuracy

The accuracy of the method was assessed in three ways (see Table 2.2):

a) Rock standards UTB2 (U of T basalt, an inhouse standard of the same flows as BCR1 and BCR2), FER1, FER2, FER3 and FER4 (iron formations standards from CANMET) were run as unknowns. b) Rare earth element LA-ICP-MS data were compared with three samples analyzed by INAA. c) Minors were compared to XRF data. a) Most elements were in the 10% range, exceptions were:

- Cu, Zn, Eu, Ta, Pb, Th for UTB2

- Cr, Co, Cu, Zn, Zr, La, Pb for FER1-4

However, since data for the iron formation standards is somewhat limited and concentrations for many elements are very low, caution is required b) Compared to INAA data, LA-ICP-MS generally shows somewhat lower values. However, in some samples the light rare earth elements are higher, and in others the heavies show an enrichment (see Table 2.4).

The data indicate that rare earth element accessory minerals play an important role in the REE budget of the whole rock. Because of the relatively smaller sample size used for LA-ICP-MS, the nugget effect may be more significant in those samples. This finding demonstrates the importance of melting the sample into a glass and then analyzing it as opposed to analyzing a pressed powder pellet or thin sections on the LA-ICP-MS or small sample sizes by INAA.

Table 2.2: Accuracy for the in-house standard UTB2 and CANMET FER1, 2, 3, and 4 Most elements are within ±10%, exceptions are Cu, Zn, Eu, Ta, Pb, Th for UTB2; Cr, Co, Cu, Zn, Zr, La, Pb for FER1-4

It also has to be taken into account that the INAA data are not of the same quality for each element:

Sc: ± 1ppm Cr: no good at low levels (e.g. 12ppm) Ba: difficult to get good data because of small peak of interferences La: excellent Ce: difficult if lots of Fe (20%) since at the side of Fe-peak Nd: bad Sm excellent Eu: very good Tb: poor Yb: good Lu: close to detection limit Ta: small peak Th: o.k. U: not very good at low concentrations

Table 2.3: Comments on INAA data

Table 2.4: LA-ICP-MS data compared to INAA data of three samples of the Helen Iron Formation. LA-ICP-MS data shows both lower LREE and lower HREE compared to INAA data, indicating a nugget-effect

c) Comparison of the minor element data obtained from XRF with LA-ICP-MS data shows some important differences (see Table 2.5):

Vanadium is relatively good when compared to XRF. There might be some contamination of the LA-ICP-MS with high-V samples, but it does not seem to have played a major effect.

Chromium: the accuracy is highly variable. The reason is probably the blank of the stainless steel holder of the XRF, which makes low values at low chromium concentrations rather unreliable.

Copper, zinc and nickel do not compare well. The reasons for that are unknown but most likely associated with problems of using NIST 610 as a standard.

Zirconium is rather variable. This is probably due to the fact that zirconium is mainly incorporated in zircons, making it very susceptible to the nugget effect.

Barium: the accuracy of barium is the 20% range.

Lead data obtained by XRF and LA-ICP-MS do not agree because of loss of volatile elements when preparing glass beads.

Table 2.5 : Comparison of LA-ICP-MS data and XRF data obtained on the Helen Iron Formation V: good accuracy Cr: highly variable accuracy, due to instrumental blank in XRF data due to Cr in the stainless steel holder Cu, Zn, Ni: accuracy not good Zr: variable accuracy, probably due to the nugget effect. Pb: low accuracy due to loss whilst preparing glass beads.

13 2.3 Ccarb

Samples of the Helen Iron Formation were measured on a gas bench connected with a Thermo- Finnigan MAT253 gas source isotope ratio monitoring mass spectrometer. Standards used were MERCK 1, IAEA-CO-8, NBS 19 AND IAEA-CO-1 (Friedman et al., 1982; Gonfiantini et al., 1995; Stickler, 1995; Gröning, 2004). The standard deviation was ±0.035‰. Analytical precision is ± 0.1%.

Since siderite reacts slowly with H3PO4, a time test was conducted. Sample SJ_10_10.5 was chosen because its major element chemistry showed it to be a good average sample of the Helen

Iron Formation. Each sample was flushed with helium, then 5-7 drops of H3PO4 were inserted by hand with a needle (making sure that the sample is covered by the acid) and the samples were stored at 70°C. 2 samples each and 2 standards were measured every 1, 2, 4, 8, 18, 23, 32, 50 and 13 90 hours in order to monitor the yield of the Csiderite. Figure 2.6 shows that the reaction was complete after 10 hours. In order to be on the safe side, a reaction time of 48 hours at 70°C was chosen. This result is in good accordance with the results of e.g. Beukes et al. 1990; Zhang et al., 2001.

Sample sizes were chosen according to the calculated mineralogy, ranging from 800 g (for 50% siderite) to 500 g (for 90% siderite).

Figure 2.6: Time-dependent reaction of sample SJ_10_10.5 Reaction is complete after 10 hours

CHAPTER 3: THE GEOLOGY AND PETROGRAPHY OF THE HELEN AND SOKOMAN IRON FORMATIONS

Most Precambrian iron formations show significant deformation and metamorphism, making reconstruction of their depositional characteristics and environments challenging.

However, both field research areas examined in this study present unique opportunities to examine low-metamophic grade iron formations of both the Superior and the Algoma types with very little signs of internal deformation.

3.1 The Helen Iron Formation, Wawa, Ontario

3.1.1 General description and geology

The Helen Iron Formation is part of the 3-10 km thick Michipicoten Group located northeast of Lake Superior, 3 km north of Wawa Lake, northern Ontario (see Figure 3.1). It is part of the Superior-Province and located in the autochthonous Wawa-Abitibi terrane (Williams et al., 1991; Thurston et al., 2008) and forms the uppermost unit of the Wawa assemblage (Sage and Heather, 1991; Williams et al. 1991, Sage, 1996).

Iron mining began in 1898, mainly of the siderite-pyrite bodies associated with the Helen Iron Formation (Goodwin, 1962).

The Michipicoten Greenstone belt is interpreted as a convergent plate margin that varies from immature island arcs formed on oceanic crust to more mature island arcs on continental crust. It may be the remnant of a larger volcanic terrane that was part of the Abitibi greenstone belt (Sylvester et al., 1987).

Figure 3.1: Location of the Helen Iron Formation, Wawa, Ontario

The belt consists of a complex volcanic and sedimentary assemblage with four volcanic and two intercalated sedimentary sequences, the Helen Iron Formation being in direct contact with felsic pyroclastics of cycle 2 below and mafic volcanic flows of cycle 3 above:

Cycle 1 has a base of basaltic to peridotitic komatiites and an intermediate-felsic top, on top of which are minor chert-magnetite-suphide iron formations.

Cycle 2 has again a lower sequence of high Mg to high Fe tholeiites and again an intermediate- felsic top, on top of which the Helen Iron Formation was deposited.

The Helen Iron Formation consists of 4 facies: siderite, siderite ± pyrite-pyrrhotite, banded and laminated chert, and argillite-graphite-pyrite (Figure 3.2). The facies are laterally extensive but vary in the degree of their development.

The iron formation overlies felsic volcanic rocks and underlies intermediate to mafic volcanic rocks.

Cycle 3 consists of a lower sequence of high Mg- to high Fe-tholeiites and an intermediate-felsic top. It partly represents the source of the clastic Doré sediments.

The Doré sedimentary group is a clastic facies consisting of conglomerates, greywackes and shales. The unit is a thick conglomerate facies to the west and a thin, shalier facies to the east. It was mainly derived by erosion of growing volcanic piles.

Figure 3.2: Stratigraphy of the Helen Iron Formation, Sir James Dunn Pit

Cycle 4 consists of intermediate to mafic metavolcanics that have been extensively carbonate altered.

The typical volcanic cycle develops from deposition of andesite-basalt flows and pyroclastic rocks to formation of dacite-rhyolite pyroclastic rocks.. Volcanism was subaquaeous. The upper 30-150 meters of volcanic rocks is highly altered to carbonate with the addition of Fe-Mn carbonates and removal of silica and alkalies (Goodwin, 1962).

The Helen Iron Formation is a well known example of an Algoma-type iron formation and has been zircon U-Pb dated at 2749±2 Ma from the underlying volcanics (Turek et al., 1982).

Stratigraphy of the Helen Iron Formation (see Figure 3.2):

The siderite facies consists of tabular bodies and lenses massive of siderite, composed of fine- grained siderite with ankerite, quartz, and sulfides (up to 10%). Bedding becomes more prominent towards the pyrite member and is absent at the base of the facies (Figure 3.3a). Layers and beds of tuffaceous rocks are also present (Goodwin, 1962). The carbonate member has a maximum thickness of 450 m at the MacLeod Mine (Thode and Goodwin, 1983).

The sulfide facies grades upward into the breccia facies, downward into the siderite facies (Goodwin, 1962). The unit is less than 10 m thick.

The breccia facies consists of bands/layers of white to grey chert alternating with bands of brown, siliceous carbonates (Figure 3.3b). The chert is sugary. Intraformational brecciation is common (Figure 3.3c), mostly arranged in subparallel bands.

The breccias show subparallel fragments of chert in a matrix of siderite and chert mixture. Most of the chert fragments have angular terminations indicating disruption of the early lithification. However, occasional soft-sediment folding of chert horizons has been observed.

Figure 3.3a: Top of siderite facies, Lucy Pit Figure 3.3b: Breccia facies, Lucy Pit

Figure 3.3c: Breccia facies, Sir James Dunn Pit (picture 3m Figure 3.3d: Graphite facies with sill to the right, Sir James wide) Dunn Dunn Pit

Figure 3.3e: Graphite facies with pyrite, Sir James Dunn Pit Figure 3.3f: Sir James Dunn Pit 29

The graphite-facies is high in organic carbon and contains abundant pyrite; both finely disseminated and in diagenetic nodules, often internally zoned (Figures 3.3d and 3.3e).

The upper contact of the Helen Iron Formation shows no evidence of erosion or structural disturbance. Hence, the chert was solidified enough to withstand emplacement of basaltic pillows (Goodwin, 1962).

Findings of stromatolites in the Helen Iron Formation (MacLeod Mine) in the siderite facies indicate at least temporarily shallow water depth of few tens of meters, and quiet and clean water below the storm wave base (Hofmann et al., 1991).

The iron formation was mostly deposited in moderately to deep water as indicated by the absence of shallow water features such as ripple marks, channelling, crossbedding and by the presence of graded bedding, laminations and pillow structures in the mafic rocks. It was deposited on a volcanic terrain dominated by broad, shield like pyroclastic domes of high relief. The pure chemical sediments were deposited during periods of volcanic quiescence as shown for the Abitibi greenstone belt by Thurston et al. (2008).

The Helen Iron Formation is in sharp contact with the underlying volcanic rocks. Iron formations of sedimentary association are present in the field area as well but have not been included in this study since they only occur in diffuse, discontinuous bands.

Samples were collected in the summer of 2010 from the Sir James Dunn Pit (Figure 3.3f), 10 km east of Wawa, Ontario, with a sample spacing of 0.5 meters where possible. For the breccia facies samples were taken from the interstitial carbonate matrix whenever possible, chert bands were avoided.

2.1.2 Petrology

20 rock samples of the Helen Iron Formation were chosen for polished thin section analysis using both reflected and transmitted light.

Samples were chosen approximately every 5 meters through the section in order to be representative. A representative part of each section was drawn.

Figure 3.4: Thin section of SJ_10_0.1, siderite facies (Sir James Dunn Pit)

Sample SJ_10_0.1: this sample is located immediately above the rhyolitic base and is dominated by sphalerite, pyrite and quartz. Chlorite is present as well as muscovite. The presence of sphalerite in this sample suggests proximity to a local hydrothermal discharge.

Figure 3.5: Thin section of SJ_10_5, siderite facies (Sir James Dunn Pit)

Sample SJ_10_5 is part of the lower siderite formation. It is dominated by siderite confirmed by electron microprobe analysis and very fine-grained quartz with typical grain sizes of 15 microns. Some fine-grained magnetite is dispersed through the thin section. Ilmenite is present as well but is minor.

The ground mass consists of fine-grained siderite with a grain size of 10 microns as well as fine grained chlorite and sericite.

Figure 3.6: Thin section of SJ_10_10.5, siderite facies (Sir James Dunn Pit)

Sample SJ_10_10.5 is located in the lower siderite facies, immediately above the F1-rhyolitic ash layer.

The dominant mineralogy is fine-grained quartz with a grain size of 15 microns and siderite with 10 microns. The siderite shows very fine-grained inclusions of phyllosilicates. Pyrite is minor and fine-grained. Fine-grained sericite is found finely dispersed in the groundmass.

In this section the siderite is possibly being replaced by fine grained SiO2 suggesting that it is early diagenetic (see also e.g. Krapez et al., 200).

Figure 3.7: Thin section of SJ_10_15, siderite facies (Sir James Dunn Pit)

Sample SJ_10_15 is from the lower siderite facies and dominated by siderite with grain sizes of about 150 microns that shows again very fine-grained inclusions of phyllosilicates. Both pyrite and quartz are absent.

Figure 3.8: Thin section of SJ_10_21, siderite facies (Sir James Dunn Pit)

Sample SJ_10_21 is of igneous material within the lower siderite facies.

The rock is fine-grained and consists mainly of K-feldspar (some showing Carlsbad-twinning) phenocrysts, which have been altered to chlorite and sericite. The groundmass consists of very fine-grained sericite and chlorite.

Some of the feldspars contain fresh cores.

The rock is most probably an altered crystal-rich tuff.

Figure 3.9: Thin section of SJ_10_25, siderite facies (Sir James Dunn Pit)

Sample SJ_10_25 is dominated by siderite and contains very little SiO2 in the form of 10 micron big grains. Pyrite is found occasionally, is coarse-grained and has lots of inclusions indicating that it is a late diagenetic phase.

Chlorite is also found, probably as the metamorphic alteration product of sheet silicates.

Figure 3.10: Thin section of SJ_10_30.5, siderite facies (Sir James Dunn Pit)

Sample SJ_10_30.5 is part of the upper siderite facies and is again dominated by siderite with grain sizes of 10 microns. Quartz and pyrite are minor and fine-grained with grain sizes below 10 microns. This sample contains finely dispersed phyllosilicates in the carbonate.

Figure 3.11: Thin section of SJ_10_35, siderite facies (Sir James Dunn Pit)

Sample SJ_10_35 is also part of the upper siderite facies and is dominated by very fine-grained siderite (grain size 5 microns). Veinlets of coarser grained siderite crosscut the thin section. Pyrite is coarse grained with lots of inclusions. Fine-grained quartz has a tendency to occur around the siderite veins.

Figure 3.12: Thin section of SJ_10_40, siderite facies (Sir James Dunn Pit)

Sample SJ_10_40 is from the upper part of the siderite facies just before the sulfide facies begins. The mineralogy is dominated by siderite (grain size of 10-15 microns), little fine-grained quartz and coarse-grained pyrite, minerals are partly altered to chlorite. Medium-sized grains of andalusite can be found dispersed through the whole thin section as confirmed by electron microprobe analysis.

The pyrites have an euhedral core of inclusion-free crystals, the secondary pyrite growing around the rims is full of inclusions and thus a late phase.

Figure 3.13: Thin section of SJ_10_44.5, sulfide facies (Sir James Dunn Pit)

Sample SJ_10_44.5 is from the border of the siderite facies/sulfide facies. Its mineralogy very much resembles sample SJ_10_40 but andalusite is more common while pyrite tends to be more fine-grained, but is still full of inclusions.

Figure 3.14: Thin section of SJ_10_50, breccia facies (Sir James Dunn Pit)

Sample SJ_10_50 is the first sample of the breccia facies immediately above the sulfide facies.

This sample shows a vein of pyrite crosscutting a siderite/quartz mixture. Medium-sized chlorite can be observed and is usually closely associated with pyrite.

Pyrite is free of inclusions indicating that it is a primary phase.

Figure 3.15: Thin section of SJ_10_55.5, breccia facies (Sir James Dunn Pit)

Sample SJ_10_55.5 is from the breccia facies and has been taken from the carbonate-matrix. Thus it is dominated by siderite with a grain size of approximately 20 microns. Is it crosscut by a vein of pyrite with inclusions of fine-grained quartz, siderite and chlorite.

Figure 3.16: Thin section of SJ_10_60, breccia facies (Sir James Dunn Pit)

Sample SJ_10_60 is from the breccia facies and taken from a pyrite-quartz layer of the rock sample. The dominant mineralogy is very fine-grained quartz and pyrite. The pyrite shows small inclusions of siderite.

Figure 3.17: Thin section of SJ_10_65, breccia facies (Sir James Dunn Pit)

Sample SJ_10_65 shows an example of the breccia facies with a piece of quartz in the carbonate matrix. Pyrite is minor and restricted to the carbonate matrix, which is dominated by siderite- mineralogy (grain size of 5 microns). Some coarser-grained siderite is associated with magnetite.

Figure 3.18: Thin section of SJ_10_70.5, breccia facies (Sir James Dunn Pit)

Sample SJ_10_70.5 is from the middle of the breccia facies. This sample also includes a part of a quartz-layer within the carbonate matrix. By comparison with sample SJ_10_65, the quartz layer also includes layers of siderite.

Pyrite is euhedral and restricted to the carbonate layer. Some fine-grained magnetite is visible.

Figure 3.19: Thin section of SJ_10_75.5, breccia facies (Sir James Dunn Pit)

Sample SJ_10_75.5 is very similar to sample SJ_10_40.5. The main difference is the change of grain size within the quartz fragment. It tends to coarsen up towards the middle of the slab. Also, the carbonate matrix shows frequent inclusions of fine to medium-sized quartz.

Figure 3.20: Thin section of SJ_10_80A, breccia facies (Sir James Dunn Pit)

Sample SJ_10_80A is a very nice sample of the breccia facies and shows a coarsening-up of the quartz inside the quartz-fragment. Towards the carbonate matrix, the quartz becomes very fine- grained.

Figure 3.21: Thin section of SJ_10_84.5, breccia facies (Sir James Dunn Pit)

Sample SJ_10_84.5 on the other hand does not show this effect. The quartz layer is highly inhomogeneous and shows frequent inclusions of siderite. The grain size of the quartz varies from very fine to large.

Figure 3.22: Thin section of SJ_10_90, breccia facies (Sir James Dunn Pit)

Sample SJ_10_90 shows alternating quartz and siderite layers. The sample has a high concentration of organic matter in the siderite (not visible in the drawing).

Figure 3.23: Thin section of SJ_10_95, sill (Sir James Dunn Pit)

Sample SJ_10_95 is from the rhyolitic sill above the breccia facies and shows good examples of resorbed quartz.

The rock also shows a large amount of carbonate alteration as well as sericitation.

Polycrystalline and deformed quartz can be found in this thin section.

Sample SJ_10_97 is from the graphite material and shows a high content of organic material and quartz with large nodules of pyrite.

Conclusions

Samples from the Helen Iron Formation are texturally simple when compared to the Sokoman Iron Formation (see Chapter 3.2.2).

In some sections siderite surrounds quartz, in others quartz surrounds siderite; there is no convincing evidence that one preceded the other.

However, there is some indication that some of the pyrite is late diagenetic in origin.

It was thus not possible to establish a diagenetic sequence of any of the observed minerals. This indicates that the Helen Iron Formation probably presents a direct precipitate from seawater, diagenesis played a minor role.

The detrital input into the section is only important in the first 20 meters, where phyllocilicates are commonly found. There is no evidence for detrital minerals in the upper part of the section indicating a fairly pure chemical sediment.

Water depth is somewhat difficult to estimate, however, the lack of ripple marks or cross bedding indicates a deeper water setting. Stromatolites described by Hoffman et al. (1991) in the MacLeod mine however indicate a deposition within the photic zone, since stromatolites need light for their photosynthesis. In clear, detritus-free water, the photic zone can extend to 200 meters (Walter et al. 1976) constraining the water depth in which the Helen Iron Formation precipitated further.

2.2 The Sokoman Iron Formation, Labrador Trough

2.2.1 General description and geology

The Labrador Trough, also known as the New Quebec Orogen, is a 1000 km long Paleoproterozoic fold and thrust belt (Figure 3.24) which resulted from the collision of the Archean Superior Province (SW) with the Archean Rae Province (NE). It is part of the circum- Superior belt and dominated by sedimentary rocks in the west and east and predominantly mafic igneous rocks in the center. The oldest Proterozoic rocks are continental red beds (Williams and Schmidt, 2004).

Iron has been mined since 1954 in the Labrador Trough, mostly the direct shipping ore (DSO) type that resulted from supergene leaching and enrichment of the cherty iron formation which leads to the formation of friable fine-grained secondary iron oxides such as hematite, goethite and limonite.

Figure 3.24: Location of the Labrador Trough, Canada 52

The Labrador Trough sedimentation can be divided into several stages (Wardle and Bailey, 1981):

Stage 1: Continental clastic sedimentation is recorded in the Chakonipau Formation in the north and the Discovery Lake and Snelgrove Lake Formations in the south. They consist of red and grey crossbedded arkoses, quartz grain conglomerates and local pebble-boulder conglomerates. These deposits are fluviatile and formed in a north-northwest trending rift valley.

Stage 2: Marine transgression and the development of a shelf led to the deposition of the Sawyer Lake formation in the south-central part, an offshore sandbar or strand line around an arch with a basinal environment in the east and tidal flat deposition in the west.

Stage 3: Subbasin development in the north-central part.

Stage 4: A temporary halt to shallow-water clastic and carbonate deposition and the progradation of shales and deep-water mafic volcanism led to the deposition of the Attikamagen formation in the south-central area.

Stage 5: The Denault Formation dolomites mark the re-establishment of the shallow water environment.

Stage 6: Major subbasins formed in which the Dolly and Fleming Formations were deposited (turbidites and chert breccia).

Stage 7: Wishart formation and Sokoman Iron Formation deposition as well as the establishment of the volcanic Nimish Subgroup.

Stage 8: The collapse of the shelf and a major transgression led to the deposition of the overlying Menihek Formation.

Stage 9: Uplift and erosion of the shelf led to the deposition of red beds (Tamarack River Formation).

The “Schefferville Zone” in the west of the Labrador Trough represents the most complete and least deformed part of stage 5-8 (Wardle et al., 1990). Archean gneiss of the Superior province to

the west is uncomformably overlain by the Knob Lake Group, which consists of three sedimentary cycles:

Cycle 1 consists of a marine-shelf succession of the Attikamagen Group (Le Fer siltstone; Denault dolomite; Dolly siltstone-shale; Flemming chert breccia).

Cycle 2 is represented by the Ferriman Group that includes the Wishart Quartzite, the Ruth Shale and the Sokoman Iron Formation. The Wishart Formation is interpreted to be a very shallow, well-sorted marine beach sand facies with a thickness of 15-20 m. The upper 0.5-2.5 m consists of a black massive chert, called the Black Chert. Transgression below the storm wave base led to the deposition of the laminated Ruth Shale (see Figure 3.26a) with fine-grained disseminated carbon and intergrown pyrite grains which formed in a strongly reducing environment. Shallowing of the sea, possibly due to regression, led to the development of a marine shelf characterized by the fairly pure chemical sedimentation of the Sokoman, locally containing stromatolites. The depth of deposition is estimated to be less than 100 m, maybe in the range of some 10’s of metres (e.g. Dimroth, 1976).

Cycle 3 consists of the deep marine sediments and volcanics of the Doublet Group. A major transgression resulted in the collapse of the shallow shelf environment and thus in the deposition of the turbiditic deep water Menihek Shale. The Menihek Shale is the youngest stratigraphic unit and is a grey to black coloured fine-grained, organic matter rich shale that unconformably overlies the Sokoman Iron Formation with a thickness of at least 100 meters.

The Sokoman is a good global example of a Superior-type iron formation. It consists of 120 to 240 m of cherty iron formation containing pelloids and ooliths and showing thin and irregular bedding as well as cross-bedding and stromatolites (see Figure 3.25 for stratigraphy). It is part of the Kaniapiskau Supergroup and is divided into two members (Chauvel and Dimroth, 1974):

Lower Iron Formation LIF 10-30 m of silica-carbonate and shale; thin to medium-banded (Figure 3.26b). This unit marks the transition to a shallow marine shelf or an epeiric sea. The basal LIF is fine grained and contains intercalated shale bands (clastic sediments), whereas the

Figure 3.25: Stratigraphic column of the Howell’s River Area, Labrador Trough 55

upper part is a fairly pure chemical sediment. The formation is typical of a broad shelf or shallow sea where the hinterland has a low relief.

Middle Iron Formation: 90-150 m mainly cherty hematite iron formation. It consists of:

Lower Red Green Chert (LRGC, Figures 3.26c and 3.26d). Up to 45 m thick, part of the magnetite-carbonate facies; grades into the overlying LRC.

Lower Red Chert (LRC, Figure 3.26e), up to 19 m thick.

Pink Grey Chert (PGC, Figure 3.26f), which ranges from 4-23 m in thickness and is uniformly grey to grey-green in colour. Magnetite occurs in fine-grained disseminations, locally concentrated in diffuse bands. It grades into the URC.

Upper Red Chert (URC, Figure 3.26g): 4.5-17 m thick, contains ooliths in the top 0.5 to 2 meters.

Upper Iron Formation: 25-60 m of silicate-carbonate.

Green Chert (GC, Figure 3.26h): 1-10 m thick. It consists of 70-90% quartz and is in sharp contact with the URC.

Jasper Upper Iron Formation (JUIF, Figures 3.26i and 3.26j). It is in sharp contact with the underlying GC, and ranges in thickness from 20 m to 30 m. It grades upward into the LC.

Lean Chert (LC, Figures 3.26k and 3.26l): a massive unit with a thickness of 18-32 m which contains stromatolites.

The Sokoman represents a rapidly changing environment from high- to low-energy but still shallow water; the LRC, URC and the lower part of the JUIF being the high-energy facies, whereas the LIF, LRGC, PGC, GC and LC were deposited in shallow but more quiet water.

An U-Pb zircon age of 1880±2 Ma for a carbonatite dyke interpreted to be coeval with the Sokoman, a U-Pb zircon age of 1877.8±1 Ma for quartz syenite cobbles from the Nimish and a

Figure 3.26a: Ruth Shale Figure 3.26b: Lower Iron Formation

Figure 3.26c: Lower Red Green Chert Figure 3.26d: Lower Red Green Chert

Figure 3.26e: Lower Green Chert Figure 3.26f: Pink Grey Chert

Figure 3.26g: Upper Red Chert Figure 3.26h: Green Chert

Figure 3.26i: Jasper Upper Iron Formation Figure 3.26j: Jasper Upper Iron Formation, field sample

Figure 3.26k: Lean Chert Figure 3.26l: Lean Chert with Stromatolites

gabbro sill within the overlying Menihek Formation 50 km north-northeast of Schefferville with an age of 1884.0±1.6 Ma give the best age constraints for the area (Findlay et al., 1995).

Igneous activity climaxed during iron formation deposition and consists of mafic to felsic volcanics of the Nimish formation at the base and top of the Sokoman; however, episodic magmatism is observed in the field throughout the entire interval of iron formation deposition (Findlay et al., 1995).

Thrusting and metamorphism occurred in response to the Rae-Superior convergence and collision at 1840-1829 Ma (Machado, 1990). The metamorphic grade increases eastward from subgreenschist facies in the Schefferville zone to upper amphibolite-granulite facies in the east (Dimroth and Dressler, 1978; Hoffman, 1988).

The Howells River Area is a long narrow belt 75 km long and 1.5 km wide located 24 km NW of Schefferville (see Figure 3.26). It contains the least metamorphosed and best preserved sections of the Schefferville zone (e.g. Fink, 1972 and 1976). Structural deformation is very minor and the unit only experienced low grade metamorphism, the lithology dipping only 5-12° E-NE and striking 35° NW. There is no evidence for post-depositional leaching or weathering. Estimates from New Millennium Capital Corp. are ca. resources up to ten billion tonnes of magnetite ore.

Samples from 178 m of drill core with a spacing of 1 m were obtained from three composite drill holes in summer 2009, from the Ruth Shale through the Sokoman to the first 5 m of the overlying Menihek Shale. A total number of 179 drill core samples together with 63 field samples from all 11 units were collected.

Figure 3.27: Howell’s River Area, view towards South

2.2.2 Petrology

Nine thin sections taken from field samples were examined under transmitted and reflected light, 18 representative drawings were made.

The Lower Iron Formation (LIF)

The Lower Iron Formation has a variety of textures; four examples are discussed in Figures 3.28 to 3.31.

Figure 3.28: Thin section 09-FH 3-3a, Lower Iron Formation

Section 09-FH 3-3a shows a few grains of coarse-grained magnetite in a matrix dominated by minnesotaite and few patches of fibrous stilpnomelane.

Quartz is a very early diagenetically recrystallized phase.

Stilpnomelane is also an early phase and has later been replaced by minnesotaite as can be seen from the embayment s in the outline of the lower left patch.

The magnetite is subhedral and significantly coarser-grained, suggesting it is a late diagenetic phase. It may be approximately coeval with the minnesotaite.

Figure 3.29: Thin section 09-FH 3-4a, Lower Iron Formation

Section 09-FH 3-4a shows a medium-grained quartz matrix with several remnant ooliths. The ooliths are smaller than those higher in the section and are dominated by coarse-grained quartz with hematite. The hematite is found in patches and not along grain boundaries as in the overlying formations.

Magnetite is minor, medium-grained and subhedral.

Hematite is the earliest iron mineral, followed by magnetite as a later diagenetic phase. However, there is also diagenetic hematite since it occurs with medium grained quartz replacing the ooliths, the quartz is diagenetic. Some of the finer magnetite appears to be approximately coeval with some hematite and medium grained quartz.

Figure 3.30: Thin section 09-FH 3-4b, Lower Iron Formation

Section 09-FH 3-4b shows a fine-grained quartz matrix with oolith pseudomorphs. Again the ooliths show a texture of coarse-grained quartz and magnetite replacing the original mineralogy.

Hematite occurs in patches and along quartz grain boundaries predating the crosscutting magnetite. Hematite also occurs within quartz grains in some areas indicating that it is an early diagenetic mineral postdating the precipitation of fine grained quartz.

Figure 3.31: Thin section 09-FH 3-4c, Lower Iron Formation

Section 09-FH 3-4c shows some fine hematite dusting within quartz grains. However, this is restricted to fine-grained quartz, whereas coarse-grained quartz is hematite free. Hematite is therefore coeval with fine grained quartz. The magnetite is subhedral and diagenetically later. The coarser quartz is also late diagenetic replacing hematite.

The Lower Red Chert (LRC)

Three thin section drawings were made of this unit. All three clearly show the original oolitic nature of this rock, indicating a shallow water deposition for the ooliths. However, the ooliths may have been transported into a deeper water depth than the one they were deposited from.

Figure 3.32: Thin section 09-FH 5-P1a, Lower Red Chert

Section 09-FH 5-P1a shows an oolith pseudomorph composed of coarse-grained quartz in a matrix of finer-grained quartz. Dusty hematite is visible in the outer zone of the oolith. The grain size increases toward the middle of the oolith and hematite is also absent, indicating a late diagenetic origin for the coarse-grained quartz.

In the adjacent quartz layer, hematite dust outlines what might have been two additional ooliths; now obscured by the coarser recrystallization of the quartz.

Magnetite in this area is minor but coarse-grained and coeval with coarse diagenetic quartz.

Figure 3.33: Thin section 09-FH 5-P1b, Lower Red Chert

Section 09-FH 5-P1b is from the same polished thin section and shows more clearly the role of magnetite in the replaced ooliths. Hematite is again seen mainly as fine-grained disseminations along replaced oolith grain boundaries. Coarse-grained magnetite then follows the outlines of the ooliths, often divided by a layer of medium-sized quartz grains from the hematitic grain boundary and is intergrown with coarse-grained quartz. Patches of hematite can commonly be found inside the oolith grains, but not invariably.

Hematite is thus the earliest mineral along with fine grained quartz; magnetite is a late diagenetic phase.

Figure 3.34: Thin section 09-FH 5-P1c, Lower Red Chert

Section 09-FH 5-P1c shows an example where the inside of an original oolith has been completely recrystallized to very coarse-grained quartz, contained by a layer of coarse-grained magnetite. The matrix is composed of fine grained early quartz.

In the lower left of the drawing, an oolith with no magnetite but hematite dusted coarse-grained quartz is visible. This indicates that the replacement with magnetite along oolith grain boundaries occurred at a late diagenetic stage.

Pink Grey Chert (PGC)

The Pink Grey Chert also contains replaced ooliths in a quartz matrix. However, in contrast to the LRC, quartz and magnetite are more coarse-grained indicating a higher degree of diagenetic recrystallization.

Hematite is less common in this section and replaced ooliths are mainly recognized by coarse- grained magnetite aggregates.

Figure 3.35: Thin section 09-FH 6-1a, Pink Grey Chert

Section 09-FH 6-1a shows an example of the PGC and its recrystallized nature. Original oolith outlines can be recognized from ellipses of coarse-grained magnetite. Finer grained quartz is also present as the matrix and is replaced by coarse-grained quartz.

Small patches of carbonates are visible, often within the ooliths. Electron microprobe analysis shows that the carbonates are calcite with minor magnesium and iron.

Figure 3.36: Thin section 09-FH 6-1b, Pink Grey Chert

Section 09-FH 6-1b shows an oolith with a carbonate core, dusted with very fine-grained hematite. This observation raises the question as to whether the ooliths consisted originally of carbonate and were replaced later by diagenetic minerals.

Coarse-grained magnetite partly outlines relict oolith textures.

Hematite is only present as very fine-grained dusting of coarse-grained quartz grains. The fine- grained quartz grains in this area do not show hematite dusting indicating that they were an early phase.

The amount of coarse magnetite in the fine grained quartz suggests it formed from an earlier iron-bearing precursor material syngenetic with the early fine grained quartz precursor.

Upper Red Chert (URC)

The Upper Red Chert is characterized by a well developed oolic texture.

Figure 3.37: Thin section 09-FH 7-5a, Upper Red Chert

Section 09-FH 7-5a shows a several ooliths next to a fine-grained quartz layer containing fine grained hematite and sparse magnetite grains. The big oolith in the middle of the figure contains a core of coarse-grained quartz. Along the quartz grain boundaries, dusty very fine-grained hematite is visible. To a lesser extent, the middle of the oolith contains batches of hematite. Minor grains of medium-sized magnetite are also found within the oolith.

In the outer zone of the oolith, hematite dusting indicates at least three original growth layers.

The quartz matrix in which the ooliths are located is free of hematite and magnetite indicating again that the fine-grained quartz is an early feature. Coarse quartz as well as magnetite are a later diagenetic phase.

Figure 3.38: Thin section 09-FH 7-5b, Upper Red Chert

Section 09-FH 7-5b shows a different part of the same thin section where magnetite is more abundant. The outlines of the ooliths are still mainly dominated by hematite dust. However, some of the grains show a coarse-grained magnetite outline which is interpreted to be the result of diagenetic recrystallization of hematite. In contrast to the hematite, magnetite is not restricted to grain boundaries and can also be found in the quartz matrix.

In this figure, minnesotaite is found both in the cores of the ooliths and along the rims.

In general, quartz tends to be more coarse-grained inside the ooliths than in the matrix surrounding the ooliths suggesting it to be the result of a later diagenetic stage.

Figure 3.39: Thin section 09-FH 7-5c, Upper Red Chert

Section 09-FH 7-5c shows an interesting grain in the central upper part. It has a coarse grained inner core of quartz, followed by a rim of minnesotaite and a ring of dusty hematite defining the outline of the oolith.

This figure also shows how the late diagenetic recrystallization of magnetite can result in large euhedral grains, which can mask the original oolitic nature of the rock.

Figure 3.40: Thin section 09-FH 7-13a, Upper Red Chert

Section 09-FH 7-13a is from the same rock but contains significant hematite.

The main feature in this image is a zone of ooliths within a layer of fine-grained hematite. Magnetite is subhedral and concentrated in particular near the contact between the oolith zone and the hematite-layer. It also occurs within the oolith zone and as small grains in the hematitic material.

The ooliths consist principally of a core of coarse quartz, with hematite dusting the grain boundaries. The core is enclosed by stilpnomelane followed by an outer layer of hematite-dusted quartz.

In contrast to the previous images of the URC, the matrix around the ooliths shows hematite dusting, not only along quartz grain boundaries but also within quartz grains.

Figure 3.41: Thin section 09-FH 7-13b, Upper Red Chert

Section 09-FH-7-13b shows a part of thin section in which the recrystallization of the magnetite along the oolitic grain boundaries is clearly defined. This recrystallization consumes the hematite. However, small inclusions of fine-grained hematite have been identified in the magnetite (lower left of Figure 3.41).

This image also shows that both minnesotaite and stilpnomelane occur principally within the ooliths and less commonly in the surrounding quartz matrix.

Figure 3.42: Thin section 09-FH 7-13c, Upper Red Chert

Section 09-FH 7-13c shows a part of the thin section in which magnetite recrystallization is advanced. The hematite rims in the ooliths have largely disappeared. The interstices between the ooliths consist of medium-grained quartz with fine-grained stilpnomelane and/or minnesotaite.

Some fine-grained quartz patches are visible and are probably less recrystallised quartz matrix.

These images make clear that care is required when interpreting the redox state of the ocean based on the present or absence of hematite or magnetite.

The Green Chert (GC)

The Green Chert is dominated by quartz and minnesotaite.

Figure 3.43: Thin section 09-FH 8-4, Green Chert

Section 09-FH 8-4 shows the two dominant minerals of the green chert, minnesotaite and quartz intergrown. The quartz shows a wide range in grain size from very fine to coarse. Some indications of a remnant ooliths are suggested.

A few grains of coarse grained magnetite are visible and do not seem to follow either the fine- grained or more coarse grained quartz patches.

Figure 3.44: Thin section 09-FH 8-7, Green Chert

Section 09-FH 8-7 shows a part of the green chert consists mainly of quartz. The grain size varies from very fine to coarse. A few grains of coarse minnesotaite are also visible.

The Jasper Upper Iron Formation (JUIF)

The Jasper Upper Iron Formation derives its name from the occurrence of hematite-rich quartz. It is also rich in magnetite.

Figure 3.45: Thin section 09-FH 9-6, Jasper Upper Iron Formation

Section 09-FH 9-6 shows clearly where the name of this formation comes from because it is dominated by fine-grained hematite dusting of some quartz.

This unit also shows a remnant oolitic nature. In contrast to the previous sections, the replaced ooliths are composed of fine-grained quartz densely dusted with hematite. The matrix consists of coarse-grained quartz without hematite.

Some of the ooliths show a zone of recrystallized magnetite following rims. However, hematite rims are still dominant in this section.

Magnetite grains also occur within the replaced ooliths but not significantly in the coarse interstitial quartz.

The Lean Chert (LC)

Figure 3.46: Thin section 09-FH 10-S1, Lean Chert with Stromatolites

Section 09-FH 10-S1 shows an example of the stromatolites found in the Lean Chert. The thin section shows fine grained goethite around the magnetite crystals. The origin of the goethite is probably weathering. The dominant feature is a replacement of the original stromatolitic texture by late diagenetic euhedral magnetite. Hence, the original carbonate of the stromtatolitic material has been replaced. Clear quartz grows radiating outwards from the diagenetic magnetite crystals, indicating a later diagenetic origin. Another generation of fine-grained quartz containing inclusions predated the above mentioned quartz but is only found localized and is not a dominant feature of the section. Two generations of carbonates are visible: a later, euhedral and earlier, subhedral fine-grained crystals with grain sizes 5 microns (not in Figure 3.46).

Stilpnomelane occurs locally in fine needles radiating out from the magnetite and in the matrix and is the result of early metamorphism.

Minnesotaite is abundant and found as bunches and needles within the quartz-rich band of the section and is of early metamorphic origin.

Discussion and conclusions

The Sokoman Iron Formation was deposited in a highly variable environment ranging from tidal flats as evidenced by the presence of ooliths (water depth less than 10 meters, King 1951), to deeper, quieter water indicated by the occurrence of stromatolites.

Both ooliths and stromatolites typically have an original mineralogy consisting of aragonite or calcite (e.g. Walter 1976). However, both show a quartz-magnetite dominated mineralogy in the Sokoman.

Thus, it can be concluded that the Sokoman Iron Formation has experienced a high degree of diagenetic alteration.

A diagenetic sequence has been established (Figure 3.47):

Figure 3.47: Paragenetic diagram of the Sokoman Iron Formation, established by petrographic thin section observation..

The precursor sediment of the Sokoman Iron Formation was probably a silica iron hydroxide gel (Klein, 1974 and 2005). Subsequent compaction and dewatering resulted in the intense diagenesis observed and the crystallization of hematite and quartz. Some of the Fe3+ in the precursor was probably reduced at the expense of organic matter to form magnetite, explaining the lack of organic matter today present in the rocks. The production of organic matter was a consequence of photosynthesis in shallow water and was deposited along with the hematite.

Hematite is fine grained and interpreted to be the earliest iron mineral present. Magnetite is mostly coarse-grained, euhedral and overgrowing hematite. It thus postdates hematite and is the result of late diagenesis.

There is evidence for the fine-grained quartz being an early stage, whether it is a direct precipitate or an early diagenetic alteration of a precursor such as a silica-gel. Coarse grained quartz is late diagenetic.

Stilpnomelane is a late diagenetic or early metamorphic phase. Minnesotaite is replacing stilpnomelane and therefore interpreted to be early metamorphic.

Dimroth and Chauvel (1973) and Dimroth (1977) proposed a replacement origin for the Sokoman Iron Formation based on the similarity of the observed textures with limestones. The original sediment in their opinion was limestone that was replaced by iron minerals and quartz during diagenesis. However, they provide no explanation for the origin of the large amounts of silica and iron nor do they provide a process responsible for their mobilization.

The mobility of iron is controlled by the oxidation state of water whereas the mobility of silica is controlled by pH and temperature (e.g. Keller et al., 1963; Manning, 1994; Evans et al. 2013). Furthermore, it would also be necessary to mobilize the rare earth elements under still different conditions (see Chapter 5), since limestone has a very low REE contents. It is hard to imagine any single process capable of achieving these diverse effects.

CHAPTER 4: THE MAJOR AND MINOR ELEMENT GEOCHEMISTRY OF THE HELEN IRON FORMATION

4.1 Major elements

SiO2 varies considerably through the section from almost zero to almost 100%, as would be expected for an iron formation (see Figure 4.1).

It is noticeable is that the silica in the lower part of the siderite facies reflects the detrital input into the section (see also Figures 4.3 and 4.4). In the breccia facies, however, the fluctuations are a result of varying amount of a chert component in the samples as seen in thin sections.

Figure 4.1: SiO2 (%) against stratigraphic height (in meters) Silica contents are high in the first 20 meters of the siderite facies as the result of detrital input. The breccia facies shows variable silica content as the result of being a mixture of matrix-siderite and chert-fragments.

FeO: iron is reported as FeO and not as Fe2O3 since the majority of iron is present in siderite and not in magnetite/hematite.

The FeO varies from zero to 50% and is inversely correlated with the SiO2 concentration, reflecting the chert-siderite mineralogy of the Helen Iron Formation (see Figure 4.2).

Figure 4.2: FeO (%) against stratigraphic height (in meters). Iron content is high and variable reflecting the chert-siderite mineralogy of the iron formation and inversely correlated to the SiO2 content.

TiO2 is highly variable in the siderite facies of the Helen Iron Formation. The first 20 meters of the section show the highest values reaching up to 1.26% in sample SJ_10_5 (see Figure 4.3).

Boström (1970) and Bonatti et al. (1992) showed that titanium is a detrital. TiO2 values in the Helen Iron Formation trace the high detrital input that dominated the first 20 meters of the section before the iron formation changed to more purely chemical sediment.

The rest of the siderite facies shows lower TiO2 values similar to the sulfide facies not exceeding

0.40% (with the exception of sample SJ_10_32.5_TG that has 0.82% TiO2).

The sulfide facies shows a very similar pattern with values being very low with a few excursions not exceeding 0.20%.

The breccia facies has very low values in the 0-0.2% range, consistent with the origin as a fairly pure chemical sediment.

Figure 4.3: TiO2 (%) against stratigraphic height (in meters). Titanium contents are high and variable in the first 25 meters of the siderite facies reflecting the high detrital input in this part of the section. The breccia facies has low values as expected for a purely chemical sediment.

Al2O3 is present in sediment in detrital clay and other detrital matter. It is therefore a useful indicator for the amount of detrital input into the Helen Iron Formation.

The Al2O3 values are high for chemical sediments in the first 18 meters of the section, ranging up to 20%. The values then start to decline through the siderite facies with only occasional values exceeding 5%. Concentrations continue to decline in the sulfide facies and reach a minimum in the breccia facies; except for an unusual input of high values approximately 70 meters into the section (see Figure 4.4). This could indicate a sudden input of detrital material into the system, for example associated with a catastrophic storm event or volcanic ash.

Aluminum correlates well with titanium showing that both elements can be used as indicators for detrial input into the Helen Iron Formation.

Figure 4.4: Al2O3 (%) against stratigraphic height (in meters). Values are high in the first 25 meters of the siderite facies reflecting the high input of detrital matter at the beginning of the section. The breccia facies shows low aluminium values typical for a purely chemical sediment.

Manganese (MnO) values range up to 5% in the section. Just the first four meters of the siderite facies show higher values of manganese up to 4.5% MnO, then values drop to below 1% and stay remain for the next 14 meters (see Figure 4.5).

Above that, values rise up to 5% and then show a tendency to drop through the siderite, sulfide and breccia facies. The breccia facies typically contains 2%, varying from 0.2% to 3.5%.

Figure 4.5: MnO (%) against stratigraphic height (in meters). The siderite facies has the highest values ranging up to 5%, the breccia facies has an average of 2% with a range from 0.5 to 3.5%

Magnesium shows a similar pattern to manganese, starting with high values in the first 4 meters of the section and then dropping below 2% finally increasing to a steady 6-8% at the top of the siderite facies and in the sulfide facies. Values are lowest in the breccia facies, especially approximately 65 meters into the section and vary from 2% to 7%, averaging 4% (see Figure 4.6).

Figure 4.6: MgO (%) against stratigraphic height (in meters). The pattern is very similar to manganese and lowest in the breccia facies.

Calcium is generally rather low in the Helen Iron Formation, being typically below 5% (see Figure 4.7). Exceptions are three samples in the upper siderite facies, SJ_10_32 SJ_10_32.5 and SJ_10_39 (32, 32.5 and 39 meters above base respectively). These samples are dominated mainly by ankerite and ferroan dolomite as opposed to the siderite typical for the rest of the section.

Figure 4.7: CaO (%) against stratigraphic height (in meters). High values at meter 32, 32.5 and 39 are associated with the presence of ankerite and ferroan dolomite instead of siderite.

Potassium shows a very clear distribution through the section, probably reflecting the influx of rhyolitic detrital material (see Figure 4.8). The first 20 meters of the siderite facies are dominated by high K2O values up to 5%. The values then drop in the middle of the siderite facies and remain low through the sulfide and breccia facies.

Figure 4.8: K2O (%) against stratigraphic height (in meters). High values in the first 25 meters of the siderite facies reflect the input of rhyolitic detrital material into the section.

Sodium shows a similar pattern to potassium with Na2O values being elevated mainly in the lower part of the siderite facies, then dropping to values below 0.5% and then to 0% in the sulfide and breccia facies.

Phosphorus shows very low values throughout the section with the exception of samples SJ_10_32 and SJ_10_39, which show also unusually high calcium content, associated with ankerite and ferroan dolomite rather than siderite.

As expected, sulphur shows the highest values in and around the sulfide facies of the iron formation. However, values are generally below 5% and do not exceed 8%.

3.2 Trace elements

Vanadium shows a distinctive distribution through the section. Values are highest 10 meters above the base of the section and remain high for 10 meters reaching 250 ppm. Then the values drop to below 30 ppm for the rest of the siderite facies. At the border between the siderite and sulfide facies, vanadium values again increase to over 50 ppm, then decrease and remain low for the rest of the iron formation including the graphite facies (see Figure 4.9).

Figure 4.9: V (ppm) against stratigraphic height (in meters). Values are highest in the siderite facies and lower in the breccia facies.

Nickel and chromium are generally very low with the exception of a set of samples starting from meter 10 above base, extending for 10 meters with very high nickel and chromium values (see Figures 4.10 and 4.11).

Figure 4.10: Cr (ppm) against stratigraphic height (in Figure 4.11: Ni (ppm) against stratigraphic height (in meters). meters).

Arsenic is highly variable with values in the lower part of the breccia facies up to 1250 ppm but is not enriched in the sulphur facies.

Strontium is generally low. Slightly elevated values up to 200 ppm are visible from sample SJ_10_10 for 10 meters upwards.

Samples SJ_10_32, SJ_10_32.5 and SJ_10_39 show unusually high values up to 2500 ppm which is explained by the different mineralogy of these samples, as mentioned above (ankerite and ferroan-dolomite instead of siderite). Strontium can replace calcium and manganese due to the similarities in their ionic radii and can thus be used as an indicator for the presence of ankerite and ferroan dolomite.

Zirconium can be used alongside aluminium to trace the intermediate to felsic detrital input in the section. Zirconium values are high after 4 meters for another 8 meters, reaching up to 425 ppm and then dropping to lower values around 100 ppm for the next 20 meters. The upper part of the siderite facies shows lower but still highly scattered values mirroring the decreasing influence of detrital matter into the system. The sulfide facies shows low values with little scatter and the breccia facies has consistently very low values indicating the pure chemical nature of the sediment (see Figure 4.12).

Figure 4.12: Zr (ppm) against stratigraphic height (in meters). Values are highest in the first 10 meters of the siderite facies reflecting the input of detrital zircons. Values are lowest in the breccia facies indicating a detritus-free, purely chemical sediment.

Barium is variable in the siderite facies, reaching values of 1200 ppm. The sulfide and breccia facies have very low values of few ppms, close to the detection limit of the XRF (see Figure 4.13).

Figure 4.13: Ba (ppm) against stratigraphic height (in meters). Values are highly scattered.

Copper, lead and zinc were used to monitor the hydrothermal input into the section (see Figure 4.14).

Copper values are highly scattered with an average of 10 ppm and show no correlation with stratigraphy. Zinc values are low with the exception of sample SJ_10_0.1 (0.1 meter above base) which contains sphalerite in the thin section. Slightly elevated values can be observed in the first 5 meters of the section, however, not exceeding 300 ppm. Lead is low and shows only limited scatter with values up to a maximum of 50 ppm.

From this it can be concluded that the Helen Iron Formation experience no direct hydrothermal influx.

Figure 4.14: Cu, Pb and Zn (ppm) against stratigraphic height (in meters). Low concentrations of Cu, Pb and Zn (typically below 100 ppm, note the logarithmic scale) indicate insignificant direct contribution from hydrothermal input .

3.3 Ratios

3.3.1 Manganese against iron

The plot of manganese against iron shows a 1:20 ratio (see Figure 4.15). The correlation is the result of manganese replacing iron in siderite due to its similar ionic radius.

Modern seawater has manganese: iron ratio of 1:10 (Holland, 1978). Ocean crust on the other hand has a ratio of 1:80 (Taylor, 1982). This indicates that manganese in the Archean ocean has been preferentially leached out of basaltic crust compared with iron.

Figure 4.15: MnO against FeO (numbers correspond to sample names) Samples fall on the 1:20 ratio (red line) indicating that manganese has been preferentially leached out of the basaltic crust compared to iron.

3.3.2 Magnesium against iron

The plot of magnesium against iron can be used to illustrate the influx of detrital mafic material in the first 10 meters of the section (Figure 4.16, purple circle). This effect is also visible in the high vanadium values these samples show and also their chromium and nickel-enrichments.

It can also be seen that the samples with the highest magnesium content are ankeritic in composition (red circle in Figure 4.16) in contrast to the dominant siderite mineralogy of the rest of the section.

Figure 4.16: Plot MgO against FeO Purple circle: mafic component of the detrital matter in the first 10 meters of the siderite facies Red circle: ankerite in samples 32, 32 and 32.5

3.3.3 Calcium and potassium against iron

The plot of calcium against iron also shows clearly the mafic nature of some of the detrital influx in the first 20 meters at the section (Figure 4.17).

Again, ankeritic samples do not follow this trend due to their elevated calcium contents.

Figure 4.17: CaO agains FeO Purple circle: mafic component of the detrital matter in the siderite facies, dominant in meters 12 to 18 of the siderite facies. Red circle: samples 32, 32.5 and 39 are ankeritic and therefore high in CaO.

The same pattern is visible when potassium is plotted against iron (see Figure 4.18). The potassium content in the section reflects the amount of illite (now muscovite), which in turn results from the detrial input of felsic material. Thus there is evidence for both mafic and felsic detrital material.

Figure 4.18: Potassium against iron Red circle: felsic component of the detrital material in the siderite facies, first 10 meters.

3.3.4 Barium against potassium

Barium correlates well with K2O, as seen in Figure 4.19, in the first 10 meters of the section indicating that much of the barium is derived from felsic detrital material.

Figure 4.19: Plot of barium against potassium as an indicator for the detrital nature of the first 10 meters of the section (red circle).

3.3.5 Aluminium against titanium

Aluminium is present in detrital clays and as titanium is insoluble, it could not have been precipitated and it thus presumably also of detrital origin. Figure 4.20 shows, as expected, a rough correlation between the two elements.

Rhyolitic material has around 12% Al2O3 but less than 0.1% TiO2, therefore samples lying close to the y-axis probably represent the input of rhyolitic material.

Figure 4.20: Plot of aluminium against titanium. The two elements correlate well with each other, thus allowing their use to monitor the amount of detrital input into the section.

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3.3.6 Vanadium against titanium

In the plot of vanadium against titanium, two trends are visible (Figure 4.21). This plot represents the composition of detrital magnetite grains in the sediment.

The samples of the lowermost 10 meters of the Helen Iron Formation show a detrital component of felsic composition (green line) whilst the following 10 meters have a detrital component of mafic origin (red line).

Figure 4.21: Plot of vanadium against titanium. Green line: detrital material of felsic origin, dominant in the first 10 meters of the siderite facies Red line: detrital material of mafic origin, dominating the next 10 meters of the section.

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3.4 Discussion and conclusions

The Helen Iron Formation clearly shows two different trends: the lowermost 20 meters are strongly influenced by the input of detrital matter. The detrital matter is of two components: from

3.1 to 10.5 meter are strongly influenced by the input of felsic material with high Al2O3, K2O and Zr often ranging well over 200 ppm. From meter 12 to meter 18, high Ni and Cr and much lower

K2O and Zr indicate an input from mafic material.

However, many samples have Al2O3 concentration over 17%, higher than normal igneous rocks and reflected in the presence of ubiquitous sericite (represented by muscovite in Figure 7.1), presumably derived from clay precursors.

The rest of the section lacks any significant input from mafic or felsic material and thus has very low Al2O3 contents indicating a pure chemical sediment deposition. .

The lack of Cu, Pb, Zn enrichment suggests that there is no nearby direct seafloor hydrothermal input. The absence of these elements in the sulfide facies suggests that the sulfur is not of direct hydrothermal origin either.

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CHAPTER 5: THE RARE EARTH ELEMENT GEOCHEMISTRY OF THE HELEN IRON FORMATION

5.1 Introduction and theory

5.1.1 The chemical behavior of the rare earth elements

The trivalent REE only experience slight fractionation during adsorption, chemical complexation and incorporation into precipitating minerals due to the decreasing ionic radii from the light (LREE) to the heavy REE (HREE). The only two REE that can also exist in a different oxidation state are cerium, which is easily oxidized to Ce4+, and europium, which is divalent Eu2+ in reducing environments. This change in valence is accompanied by a change in ionic radius, thus resulting in positive or negative anomalies in chondrite- or shale-normalized REE patterns (Piper, 1974).

The europium anomaly is defined by (Eu/Eu*)N, Eu*=interpolated between Sm and Gd.

Bowins and Crocket (2011) showed the significance of accessory minerals in whole rock REE budgets of iron formations with monazite playing an important role. REE phosphate minerals might thus contain a significant proportion of the total whole rock REE.

5.1.2 Normalisation

Due to the Oddo-Harkins rule (even-numbered REE have higher concentration than odd- numbered), the rare earth elements need to be normalized in order to make patterns more readily visible.

Different materials were suggested in order to achieve that goal.

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Since even small amounts of detrital matter can significantly mask a seawater REE pattern, most researchers normalize their data to shales to correct for that fact. Shales represent the average terrigenous input (i.e. averaged weathered upper continental crust) of REE into the oceans (Bolhar et al., 2005).

The REE distribution in shales is only constant in Post-Archean shales whereas signification variation and evolution of Eu/Sm and Gd/Yb ratios occurs in Archean epiclastic sediments (Taylor and McLennan, 1985; McLennan, 1989).

Three main shale normalisations are used in the literature:

- North American Shale composition NASC (Gromet et al., 1984) - Post-Archean Australian Shale composition PAAS (McLennan, 1989). - Average mud from Queensland (MuQ, data from Kamber et al., 2005) (Bolhar et al., 2005).

For this research it was decided to normalize the REE data to samples rich in detrital material directly associated with each formation.

For the Helen Iron Formation, F1-type rhyolite found at meter 21 in the section was chosen. This rhyolite type is the most common Archean rhyolite and is thus believed to be representative of the detrital material found in the Helen section. This assumption is supported by Al/Zr values in the lower part of the Helen Iron Formation that are typical for F1-type rhyolites (with the exception of sample SJ_10_10 which shows a typical F3-type rhyolite REE pattern).

For the Sokoman Iron Formation, the Menihek Shale was chosen as a representative detrital REE pattern. The Menihek shows a REE pattern that is typical for post-Archean shales and closely resembles the most often used shale normalizations such as the NASC, PAAS and MuQ.

5.1.3 Modern seawater

In the modern ocean, REE have very low mobilities and solubilities, being measured in parts per trillion. Authigenic sediments reflect the REE patterns of the seawater from which they were deposited (e.g. Piper, 1974).

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The modern ocean receives its REE from two main sources: rivers (Goldstein and Jacobsen, 1988) and hydrothermal alteration of oceanic crust (Albarede and Michard, 1986). However, the overall effect of hydrothermal fluids on the REE budget of the modern ocean is insignificant due to the immediate scavenging of REE by Mn/Fe oxyhydroxides close to hydrothermal vent sites due to the presence of oxygen (Bau and Möller, 1993).

Modern seawater shows (Bau and Dulski, 1996):

- Light REE (LREE) depletion compared to the input source chemistry. - Variable but well-developed negative cerium anomalies.

+ The LREE depletion is the result of REE speciation which is dominated by REECO3 and - REE(CO3)2 (Byrne et al. 1988) due to the pH of 8-8.5 in the modern ocean. These carbonate complexes preferentially stabilize the HREE in solution whereas the LREE are depleted by adsorption onto particulates (Möller et al. 1994).

A negative cerium anomaly is the result of the fact that Ce3+ is easily oxidized to Ce4+, which is less soluble and can thus be removed from seawater via scavenging by suspended particles that settle through the water column. This behavior is reflected by decreasing dissolved cerium abundances with depth (e.g. Sholkovitz et al., 1994).

5.1.4 Ancient seawater

The striking similarities between REE patterns in Archean marine carbonates and iron formations suggest that these sediments record the REE pattern of seawater at the time of deposition (Alexander et al. 2008).

Due to the higher rate of hydrothermal cycling, the higher temperature in which this took place and the fact that continents were less widespread in the Precambrian, the role of hydrothermal fluids was much more important in Precambrian seawater. Unlike modern seawater, the Precambrian deep ocean was reducing and thus didn’t have the scavenging effect of Mn/Fe oxyhydroxides and thus preserved the hydrothermal REE pattern.

Archean seawater probably had the following features which differ from modern seawater:

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- Less pronounced LREE depletion. - A positive europium anomaly. - No negative cerium anomaly.

The less pronounced LREE depletion inferred in Archean seawater is due to the fact that since the deep ocean was reducing and the pH is inferred to be lower due to the emanation of acidic 2- high-temperature fluids, this resulted in lower [CO3 ] concentrations compared to the modern ocean (Bau and Möller, 1993). This led to a higher importance of soft aquo- and chloro- complexes that prefer the LREE compared to carbonate complexes which prefer the HREE.

The positive europium anomaly an indication of the presence of hydrothermal fluids (e.g. Michard et al., 1983; Albarede and Michard, 1986; Campbell et al., 1989; Fouquet et al., 1993, Schmidt et al., 2007; Craddock et al. 2010 ). In this environment the large Eu2+ dominates over the smaller Eu3+ ion from which it is reduced (Bau and Möller, 1993), thus leading to a more efficient extraction of europium than of the other trivalent REE and to a positive europium anomaly in seawater.

Modern mid-ocean ridge fluids show flat REE patterns with a pronounced positive europium anomaly (Michard et al., 1983).

Negative Ce anomalies only develop in oxidizing environments. The lack of these in Archean iron formations indicate that although the oxygen level in the ocean must have been sufficient to allow for the precipitation of Fe3+ oxyhydroxides, it must have been too low for the stabilization of significant quantities of Ce4+ compounds (Bau and Dulski, 1996).

There is solid geological evidence that a pronounced chemocline divided Archean shallow seawater from deep seawater with a dominant hydrothermal signature. Above this chemocline, the REE budget was controlled by fluvial input, similar to today. The REE input from hydrothermal sources today is negligible due to immediate scavenging by the precipitation of iron-manganese oxyhydroxides. However, in a reducing ocean such as existed in the Precambrian, this removal mechanism would not have operated, thus allowing the sediment to record hydrothermal REE input (Bau and Dulski, 1996).

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5.1.5 Post-depositional modifications

The potential effects of post-depositional processes are most severe for the LREE (Bau, 1993).

5.1.5.1 Diagenesis and metamorphism

If REE were mobilized during diagenesis, this would lead to an averaging out of REE distributions in iron formations.

Decoupling of europium from the other REE has been observed under diagenetic conditions in Pleistocene muds of the Amazon deep-sea fan and has been attributed to preferential europium mobility due to diagenetic reduction of Eu3+ (MacRae et al., 1992). However, it would be expected that the migration of Eu2+ would have been accompanied by the mobility of the smaller Fe2+ ion, this process would have destroyed the fine-scale iron- and silica-rich bands visible for example in the Kuruman iron formation, which is not observed (e.g. Beukes and Klein 1989).

Furthermore, Bau and Dulski (1993) observed highly variable REE ratios in individual iron and silica-rich bands in the 2.46 Ga-old Kuruman iron formation. This behavior was also observed in individual bands in the Hamersley iron formation (Morris, 1993), indicating that REE are essentially immobile during diagenesis and lithification.

The effect of metamorphism on REE mobility is a function of the water/rock ratio involved. LREE depletion and negative europium anomalies are expected in rocks affected by significant amounts of metasomatic fluids (Grauch, 1989; Bau, 1993). Under reducing conditions, Eu2+ dominates over Eu3+ (Sverjensky, 1984; Bilal, 1991). These conditions results in the decoupling of REE (III) and Eu2+ due to the larger ionic radius of Eu2+.

However, iron formation samples that experienced high-grade metamorphism such as the Isua iron formation, do not show europium or LREE depletion. Furthermore, iron formations globally regardless of metamorphic grade show similar REE patterns.

5.1.5.2 REE mobility

REE mobility has been observed in low temperature basalt alteration below 200°C (e.g. Grauch, 2- 1989). It seems to be a function of temperature, pH, fO2, availability of a ligand such as CO3 . 107

2- Recent seawater has a pH of 8-8.5 and a high availability of CO3 thus allowing the preferential mobilization of the HREE due to the increase in carbonate complex stability from LREE to HREE (e.g. Cantrell and Byrne, 1987). However, the formation of secondary minerals some of which discriminate against the incorporation of LREE and others against the incorporation of the HREE obscures the relationship. Furthermore, due to the high pH the efficiency of the REE mobilization is relatively low, and especially the formation of iron and manganese hydroxide and iron-manganese oxide crusts during low-temperature basalt alteration efficiently removes REE from solution, thus preventing their transport in seawater (Bau and Möller, 1993).

However, above mentioned processes would have been different in the Precambrian. The marine bottom water that took part in the alteration of basalts and komatiites was reducing and the pH was slightly lower than today due to the widespread discharge of acidic high-temperature fluids. 2- - This resulted in lower [CO3 ] and [OH ] in the solution, thus diminishing the importance of preferential HREE complexation and thus REE mobility (Bau and Möller, 1992).

Bolhar et al. (2005) also showed that jasper-siderite samples from the 3.45 Ga old Warrawoona Group, Pilbara, show coherent and sub-parallel REE patterns regardless of their primary mineralogy and REE abundances in the sample. This observation is inconsistent with strong and selective element remobilization.

Thus, variable europium anomalies are believed to be primary features of iron formations and not the result of REE mobility.

5.1.5.3 The role of clastic contamination

Caution is required when interpreting REE from sediments with a clastic component. Due to the fact that total concentrations of REE are very low, even tiny amounts of clastic particles can considerably mask the original signal.

Clastic contamination is recognized by increased trace element concentrations and covariation trends between REE and Th, Zr, Ga, Rb, Hf, Al, Ti and Sc (e.g. Dymek and Klein, 1988; Bau, 1993).

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5.1.6 Examples of REE distribution in iron formations

A gradual decrease in europium enrichment through time from the Archean to the Proterozoic is observed and is interpreted to be the result of the decreasing contribution of high-temperature hydrothermal fluids into the marine environment (e.g. Derry and Jacobsen, 1990; Danielson et al., 1992).

High LREE/HREESN ratios are present in 1.9 Ga and younger iron formations but absent from Archean iron formations (e.g. Dymek and Klein, 1988).

Late Paleoproterozoic iron formations younger than 2.2 Ga show a much wider range of

LREE/HREE ratios (calculated as PrSN/YbSN), that are well above and below the ratio of the shale composite.

The oldest known iron formation in Isua (Greenland), with an age of 3.8 Ga, shows a very pronounced positive europium anomaly (Dymek and Klein, 1988). It can be closely approximated by a mixture of hydrothermal to modern seawater of 1:100 and it can thus be concluded that the deposition of the Isua sediments was the result of chemical precipitation from a mixture of seawater and hydrothermal water (Dymek and Klein, 1988) (Figure 5.1). The 2.9 Ga old Mozaan Iron Formation (Pongola Supergroup, South Africa) shows positive anomaly. (Eu/Eu*) ranges from 1.51 to 2.09 with an average of 1.69. Sm/Yb ratios are slightly unusual. They indicate a greater input of Archean black smoker fluids compared to the Kuruman or Isua iron formations. These fluids have higher Sm/Yb in comparison with modern seawater. Mixing calculations indicate that a 1-5% contribution of hydrothermal fluids would explain the observed Sm/Yb ratios observed in the Pongola Iron Formation. However, Eu/Sm ratios indicate a 0.1% mixture. These observations suggest that a different process controlled the REE distributions and the existence of an additional REE source to the Pongola (Alexander et al. 2008).

The iron formation of the 2.7 Ga old Tati Greenstone Belt, NE Botswana shows LREE depletion relative to MREE and HREE, Pr/YbPAAS=0.21-0.82 and Pr/SmPAAS=0.48-0.82, a positive La/La*PAAS, Eu/Eu*PAAS=1.87-4.02, and Y/Ho comparable to modern seawater.

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The iron formation shows evidence of two periodically interacting water masses, one with elevated Sm/Nd ratios associated with high Fe-fluxes and one with lower Sm/Nd ratios associated with high Si-fluxes.

Variations in the REE patterns are mostly attributed to the major element composition of the respective mesobands, with Si-rich horizons showing flatter MREE-HREE pattern (Dossing et al., 2009).

The Kuruman (2.47 Ga) and Penge Iron Formations (2.48 Ga) in South Africa in the Late Archean to Early Proterozoic Transvaal Supergroup show distinct enrichments in the heavy REE and positive anomalies of La, Eu, Gd, Y and Er (Figure 5.2).

Figure 5.1: REE for Isua, Greenland (Dymek and Klein, 1988) Figure 5.2: REE for the Kuruman IF, South Africa (Klein Data normalized to North American Shale Composite. and Beukes 1989). Upper line: 1:100 mixture of hydrothermal fluid and seawater. Data obtained by INAA and normalized to North American Shale Composite.

The small scale variation in total REE and the magnitude of the europium anomaly observed in the Transvaal were attributed to temporal variation in the activity of high-temperature venting at the seafloor. Since this feature can be observed between individual microbands, the authors concluded that the precipitation of iron formations occurred in upwelling marine bottom waters in an oxygenated shelf environment on a very rapid time scale, meaning that REE adsorbed on the surface of iron-oxyhydroxide particles, that then became iron-rich microbands during diagenesis, and were not in exchange equilibrium with REE dissolved in ambient seawater. Higher europium anomalies in the iron formation deposited in the deeper parts of the Transvaal

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basin highlight the quality of the REE as a hydrothermal discharge proxy (Bau and Dulski, 1996).

Mixing calculations indicate that a hydrothermal fluid contribution of only 0.1% would explain the observed Eu/Sm ratios observed in the Kuruman iron formation. This estimate is consistent with calculation done by Klein and Beukes (1989) for the Kuruman and Khan et al. (1996) for the Kushtagi schist belt, Dharwar Craton, South India.

The 1.9 Ga old Gunflint Iron Formation (western Ontario, Canada) shows strongly negative cerium anomalies in the shallow water facies indicating oxygenated shallow water in the Gunflint and Biwabik basins (Planavsky et al., 2009).

The 1.88 Ga old Sokoman Iron Formation in the Labrador Trough shows that the silicate- carbonate facies has little fractionation of the LREE and a constant enrichment in the HREE. The absolute abundance of REE is inversely proportional to the chert content; silicates and siderite take up the HREE preferentially. The overlying/underlying slates show large europium depletions (Fryer, 1977).

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5.2 The Rare Earth Element geochemistry of the 2.75 Ga old Helen Iron Formation

Rare earth element data for the Helen Iron Formation were obtained by LA-ICP-MS analysis of glass beads obtained by melting small amounts of whole-rock powder on a tungsten strip heater.

Since the Helen Iron Formation is a very homogenous example of an iron formation, samples were collected every 5 meters regardless of their mineralogy through the iron formation sequence. This procedure resulted in a total of 19 samples of iron formation as well as three samples of igneous rock, two of which are sills within the iron formation and one from the overlying basalt unit.

All values were normalized to F1-type rhyolite found at meter 21 (see Figure 5.3) in order to correct for the input of detrital material into the section. F1-type rhyolite shows a very steep REE pattern and no europium anomaly (see also Lesher et. al. 1986).

Figure 5.3: REE pattern of the two rhyolitic ash beds (F1 and F3 rhyolite), Helen Iron Formation. Sample SJ_10_10 shows contamination with 3-type rhyolitic material, sample SJ_10_21 shows a typical F1-type rhyolite pattern (normalized to chondrite according to McDonough and Sun, 1995). 112

Another type of rhyolitic material (F3-type) was identified at meter 10. SJ_10_10 is an example of an F3-type rhyolite and shows the typical high flat REE pattern with a pronounced europium anomaly.

As expected of a sediment deposited in a reducing environment, no cerium anomaly has been observed.

5.2.1 REE patterns in the Helen Iron Formation

Figure 5.4 shows the REE pattern of the lower part of the siderite facies.

Overall, REE concentrations are high indicating a high amount of detrital material in the section.

This observation is supported by elevated Al2O3 and Zr concentrations (see also Chapter 4).

Figure 5.4: REE pattern of the lower part of the siderite facies, Helen Iron Formation. Samples SJ_10_5 and SJ_10_18 show a depletion in the HREE compared to the other two samples that is inherited from F3-type rhyolitic detrital material. The negative europium anomaly visible in all 4 samples is also inherited from F3-type rhyolitic material.

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Both samples SJ_10_5 and SJ_10_18 clearly show a depletion in the HREE compared to the two other samples of the lower siderite facies indicating a high input of F3-type rhyolite. The slight negative europium anomaly visible in all 4 samples is also inherited from F3-type rhyolitic detrital material.

Overall, this part of the section shows a high contamination with clastic material and is thus not representative of the seawater from which the Helen Iron Formation precipitated.

Figure 5.5 shows the upper part of the siderite facies. The overall REE concentrations are significantly lower than in the lower part of the siderite facies. The patterns show an enrichment in the HREE and a positive yttrium anomaly which is interpreted to be representative of Archean seawater.

Figure 5.5: REE pattern of the upper part of the siderite facies, Helen Iron Formation Samples are HREE enriched and show a positive europium anomaly which is interpreted to reflect Archean seawater

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Figure 5.6 shows one sample of the sulfide facies (SJ_10_50) and eight samples of the breccia facies. A large europium anomaly is visible and the lack of any cerium anomaly indicates strongly reducing seawater conditions at the time of deposition of the breccia facies.

Figure 5.6: REE pattern of the sulfide and breccia facies, Helen Iron Formation. Pronounced positive europium anomaly is visible in all samples and the lack of a cerium anomaly indicates a reducing environment.

Overall REE concentrations are low with HREE enrichment. An yttrium anomaly is also visible and some samples show a positive lanthanum anomaly. All samples have REE patterns with similar shapes and slopes which are interpreted to reflect the REE composition of Archean seawater.

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Figure 5.7 shows the REE pattern of the sill that divides the breccia facies from the graphite facies in the upper part of the iron formation.

The sample shows a LREE enrichment and a large positive europium anomaly. This europium anomaly is believed to be attributed to a high amount of feldspar in this samples and has been confirmed by thin section work.

Figure 5.7: REE pattern of the sill found at the top of the breccia facies, Helen Iron Formation. Pronounced positive europium anomaly is the result of the presence of feldspar in this sample (normalized to chondrite according to McDonough and Sun, 1995).

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Figure 5.8 shows the REE pattern of the basaltic unit that overlies the Helen Iron Formation. The pattern is relatively flat and has a slight europium anomaly, probably due to an abundance of feldspar phenocrysts observed in thin section.

Figure 5.8: REE pattern of the basaltic unit overlying the Helen Iron Formation. Pattern is very flat and shows a slight positive europium anomaly (normalized to chondrite according to McDonough and Sun, 1995).

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4.2.2 Total REE

The lower part of the siderite facies shows higher total REE concentration than the rest of the section (Figure 5.9). This reflects the high contamination with detrital material in this part of the section.

The upper part of the siderite facies as well as the breccia facies show low REE concentrations which are representative of Archean seawater.

Figure 5.9: Total REE values plotted against stratigraphic height, Helen Iron Formation. Contents are very high in the first 20 meters of the section, then dropping to values below 20 ppm in the upper siderite/breccia facies.

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4.2.4 The europium anomaly

The europium anomaly shows a very strong change with stratigraphic height.

Three distinct zones can be recognized: the lower siderite facies shows a pronounced negative europium anomaly varying from 0.62 to 0.90. This is interpreted to reflect the contamination with detrital material, mainly of F3-type rhyolite.

The upper part of the siderite facies shows no significant europium anomalies.

The sulfide facies and the breccia facies shows a pronounced increase in the europium anomaly up section up to values of 2.65 just before the graphite facies was deposited.

Hydrothermal alteration of basaltic rocks under reducing conditions results in a fluid preferentially enriched in Eu because Eu2+ is much more soluble compared to the other REE. This results in a positive europium anomaly in hydrothermally dominated seawater and thus in the iron-carbonates that precipitated from it.

The positive europium anomaly observed in the upper 55 meters of the Helen Iron Formation is therefore interpreted to be the result of relative high hydrothermal input in a reducing environment, which is in good agreement with the fact that a deepening of the basin was inferred from sedimentological field evidence.

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Figure 5.10: The europium anomaly plotted against stratigraphic height, Helen Iron Formation. The first 20 meters show a negative europium anomaly, the rest of the siderite facies up to the top of the sulfide faices has no europium anomaly whereas the breccia facies shows an increase of a positive europium anomaly with stratigraphic height. This is interpreted to mirror the deepening of the Helen basin.

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Concluding remarks

The presence of a positive europium anomaly in the Helen Iron Formation indicates the more important role hydrothermal fluids had in the overall REE distribution of the Archean oceans.

The lack of a positive cerium anomaly indicates deposition under reducing conditions. The fact that iron in the Helen Iron Formation is mostly present in the reduced form Fe2+ (in siderite) and not as Fe3+ indicates that oxygenation of iron was not the cause of deposition of this iron formation.

The overall light REE depletion, the positive yttrium anomaly and elevated lanthanum concentrations closely resembles REE patterns of modern seawater (see Figure5.11) supporting the assumption that the REE pattern of the Helen Iron Formation can be used as a proxy for Archean seawater and also indicating that the basic mechanisms controlling the REE distribution in seawater since at least 2.75 Ga have not changed significantly.

Figure 5.11: REE pattern of modern North-Pacific seawater normalized to MuQ (data from Alibo and Nozaki, 1999; normalization after Kamber et al., 2005)

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CHAPTER 6: THE RARE EARTH GEOCHEMISTRY OF THE SOKOMAN IRON FORMATION

6.2 REE patterns of the Sokoman iron formation

Samples were collected every 4-6 meters. Each facies is represented by at least one sample.

Samples were normalized to the Menihek Shale (sample A28) in order to correct for the input of detrital material.

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6.2.1 The Ruth Shale (RS)

The Ruth shale underlies the Sokoman iron formation and is an example of a typical black shale. It shows a negative slope in the REE and a pronounced positive europium anomaly. The lack of a cerium anomaly indicates a reducing environment (Figure 6.1).

Figure 6.1: REE pattern of the Ruth shale. Total REE concentrations are relatively high as expected for a black shale. Note the negative slope in the REE as well as the pronounced positive europium anomaly and the lack of a cerium anomaly indicating deposition in a reducing environment.

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6.2.2 The Lower Red Green Chert formation (LRGC)

The LRGC is the most extensive formation of the Sokoman sampled in this research with a thickness of 50 meters. The total REE content is low and shows relatively flat REE patterns (Figure 6.2). At low concentrations, the patterns are more irregular as the odd-numbered REE approach their detection limits.

A positive europium anomaly is detectable in all samples as well as a slight positive yttrium anomaly. A few samples show a positive cerium anomaly indicating temporarily partly oxidizing surface water conditions during the deposition of the Lower Red Green Chert.

Figure 6.2: REE pattern of the Lower Red Green Chert. All samples show a positive europium anomaly and most have a positive yttrium anomaly. Only a few samples also show a positive cerium anomaly indicating temporarily partly oxidizing surface water conditions during deposition.

This finding indicates a rather deep water setting, in good agreement with the fact that the LRGC is underlain by the deep water Ruth shale facies. 125

6.2.3 The Lower Green Chert formation (LGC)

Only one sample was available of this formation due to its relative thinness of 4 meters.

The REE pattern is very similar to the LRGC; a positive europium as well as a positive yttrium anomaly is visible. There might be a slight positive cerium anomaly indicating partly oxidizing conditions at the time of deposition (Figure 6.3). The iron mineralogy is dominated by hematite.

Figure 6.3: REE pattern of the Lower Green Chert. Total REE content is lower to the LRGC, a positive europium and yttrium anomaly is visible. a slight positive cerium anomaly indicates partly oxidizing conditions.

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6.2.4 The Pink Grey Chert member (PGC)

The PGC shows a similar pattern to the LGC. There is a positive europium anomaly, a positive yttrium anomaly and a positive cerium anomaly indicating that the PGC was deposited under oxidizing conditions (Figure 6.4). The dominant iron mineralogy of this section is magnetite.

Figure 6.4: REE pattern of the Pink Grey Chert. A positive cerium anomaly indicates oxidizing, shallow water deposition. Europium and yttrium are also elevated.

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6.2.5 The Upper Red Chert member (URC)

The URC member shows slightly higher total REE contents compared to the first three members of the Sokoman iron formation (Figure 6.5).

Figure 6.5: REE pattern of the Upper Red Chert. Four out of the five samples show a positive cerium anomaly indicating again shallow, oxidizing water.

The iron-mineralogy is dominated by hematite.

All four samples of this formation show a positive europium, a positive yttrium as well as a positive cerium anomaly. This indicates that the surface of the ocean was oxidized at the time of deposition of the Upper Red Chert.

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6.2.6 The Green Chert (GC)

The GC has fairly low total REE values (Figure 6.6). A positive europium and a slightly positive yttrium anomaly are visible. No cerium anomaly is visible indicating a deepening of the basin and deposition under anoxic conditions.

Figure 6.6: REE pattern of the Green Chert. Total REE values are low. The lack of a positive cerium anomaly indicates deposition under anoxic conditions.

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6.2.7 The Jasper Upper Iron Formation (JUIF)

The JUIF has rather high REE contents compared to the lower members of the Sokoman. All samples show a positive europium anomaly and various degrees of a positive yttrium anomaly.

Two types of samples are visible when looking at the cerium: one type in the middle of the JUIF has a positive cerium anomaly and thus indicates deposition in an oxic environment, the other type of sample shows no positive cerium anomaly thus indicating anoxic depositional conditions (see Figure 6.7). This pattern indicates that the depositional depth of the JUIF was variable; hence, it may have been deposited quite close to the chemocline.

The general iron mineralogy of the JUIF is magnetite.

Figure 6.7: REE pattern of the Jasper Upper Iron Formation. All samples show a positive europium and yttrium anomaly. Samples B3029 and B3425 in the middle of the JUIF show a positive cerium anomaly indicating deposition under slightly oxidizing conditions, the rest of the formation was deposited under anoxic conditions.

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5.2.8 The Lean Chert member (LC)

This member of the Sokoman is characterized by a low iron and high chert content. All samples show a positive europium anomaly and most show a positive yttrium anomaly (Figure 6.8). Most samples also show a positive cerium anomaly indicating deposition under oxidizing conditions. Supporting evidence is provided by the presence of stromatolites in this formation. Stromatolites are dependent on light for their photosynthesis, indicating a deposition within the photic zone.

Figure 6.8: REE pattern of the Lean Chert. Most samples show a positive cerium anomaly indicating deposition in shallow, oxidizing water. Sample B22 shows some variation in the REE pattern due to the fact that values are close to detection limits.

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6.2.9 The Menihek Shale

The Menihek shale unconformably overlies the Sokoman iron formation. Its REE pattern shows a positive europium anomaly and high total REE values as expected for a black shale (see Figure 6.9).

No positive cerium anomaly is detectable and thus the seawater from which it precipitated is inferred to not have been oxidizing, thus deeper. The high total REE content indicates a reducing environment.

Figure 6.9: REE of the Menihek Shale. Sample A28 has been used for normalization of the whole Sokoman Iron Formation in this study. High total REE and lack of a positive cerium anomaly indicate a reducing depositional environment.

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6.3 Total REE

The total REE concentrations of this formation are low ( 15 ppm) and show no significant variation indicating that the component of hydrothermal input did not change during the deposition of the Sokoman. As expected the two underlying/overlying black shales respectively show a pronounced enrichment in total REE (see Figure 6.10).

Figure 6.10: Total REE (ppm) against stratigraphic height (m) of the Sokoman Iron Formation. Total REE concentrations are low except for the underlying/overlying black shales.

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6.4 The europium anomaly

There is no correlation of the europium anomaly with any of the detritally derived elements such as thorium, zirconium or aluminium, indicating that the europium anomaly is a genuine signature of the seawater from which the iron formation precipitated from.

All samples of the Sokoman Iron Formation show a positive europium anomaly (Eu/Eu*>1, see Figure 6.11). This represents the elevated hydrothermal input in the Precambrian ocean.

Figure 6.11: Europium anomaly against stratigraphic height (m). All samples show a positive europium anomaly typical for iron formation deposited before 1.8 Ga.

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6.4 The cerium anomaly

The cerium anomaly varies significantly through the sequence, reflecting changes in oxidation state and water depth during deposition.

The lowermost LRGC shows a slight negative (Ce/Ce*<1) to no cerium anomaly, indicating that conditions were not oxidizing at the time of precipitation. These data are succeeded upwards by a clear trend towards a positive anomaly (Ce/Ce*>1) in the LRC, PGC and URC, so shallow water and oxidizing conditions are indicated for this time period (see Figure 6.12).

Figure 6.12: Cerium anomaly against stratigraphic height. Positive cerium anomalies in part of the URC, JUIF and LC indicate a deposition from a shallow, oxidizing water body.

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Above the URC, a trend to deeper water and anoxic conditions in the GC and the beginning of the JUIF were followed by a switch to oxidizing conditions in the middle of the JUIF, back to more anoxic conditions and then a switch to oxidizing during the deposition of the Lean Chert.

The cerium anomaly shows no correlation with detrital elements such as zirconium, thorium or aluminium, indicating that it is a genuine seawater signature and not inherited from possible detrital components in the section.

Concluding remarks

The deposition of the Sokoman Iron Formation took place in a highly variable depositional environment ranging from deeper, anoxic to shallow, oxidized water as readily visible in the REE pattern.

The lack of correlation with detritally derived elements such as Al2O3 or Zr indicates that the REE present a genuine seawater signature.

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CHAPTER 7: THE CARBON ISOTOPE GEOCHEMISTRY OF THE HELEN IRON FORMATION

13 7.1 C in iron formations

13C values of iron-poor carbonates deposited from 3.8 to 2.5 Ga show values close to zero (ranging from -2‰ to +2‰), comparable to modern seawater with a range of +1.5 to +2‰ (Becker and Clayton 1972; Schidlowski et al. 1975; Schields and Veizer, 2002; Veizer et al., 1989; modern seawater values from Kroopnick, 1985 and Broecker and Maier-Reimer, 1992). These data suggest that there has been no significant change in the bulk carbon isotope reservoir of dissolved inorganic carbon (DIC) in seawater since the Archean (e.g. Beukes et al. 1990; Kaufman et al., 1990).

However, a literature review of carbonates in iron formation shows negative 13C values (see Table 7.1).

13 The Helen Iron Formation on the other hand shows C values having an average of 0‰.

This raises the fundamental question whether the siderite in iron formation is a direct precipitate from seawater and thus the observed negative carbon isotope values observed in many iron formations are an indication of a specific negative seawater carbon isotope reservoir, or if they are a result of post-depositional processes and, if so, which processes might be responsible.

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13 Name of the iron formation and age Observed δ Ccarb values Reference Carajas Iron Formation (Brazil); 2.76 Ga -3 to -6‰ Sial et al. (2000)

Kuruman-Griquatown Iron Formation -5.45 to -11.93‰ Beukes and Klein (1989) (Transvaal); 2.5 Ga

Dales George Member, Brockman Iron -7 to -10.5‰ Kaufman et al. (1990) Formation (Hamersley); 2.5 Ga

Hamersley; 2.5 Ga -9 to -11‰ for iron formation Becker and Clayton (1972)

Marra Mamba Iron Formation, +2 to +0.5‰ for limestone Baur et al. (1985) Hamersley; 2.5 Ga -2.8 to -19.8‰ for IF

Hotazel Iron Formation (Transvaal); 2.4- -4 to -18‰ Tsikos et al. (2003) 2.2 Ga

Gunflint Iron Formation (Ontario); 1.8 Ga -6 to 2.6‰ deep water Winter and Knauth (1992) 2.5 to +0.5‰ shallow water

13 Table 7.1: Some examples of Ccarb values of siderites in iron formations.

7.1.1 Possible explanations for negative carbon isotope values

7.1.1.1 Precipitation in a stratified basin

13 Beukes et al. (1990) observed a decrease in δ Ccarb in a transition from shallow water limestone to siderite-bearing deeper water iron formation in the Transvaal, the limestone having values of - 0.1‰ to -2.8‰ and the siderite ranging from -3‰ to -8‰. They suggest that both limestone and iron formation were primary precipitates from seawater in the same basin and concluded that the water column was stratified with respect to the isotopic composition of total dissolved carbonate/bicarbonate, the two water masses being separated by a chemocline. The cause of this stratification is presumed to be a high hydrothermal influx in the deep water and the restriction of photosynthetically produced oxygen to shallow water, thus leaving the deeper water anoxic, 2+ 13 allowing for the transport of iron as Fe . Deeper water would thus have mantle δ Ccarb values of -5±2‰, whereas shallow water would be 0‰.

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Fe3+ is precipitated when shallow, oxygenated water comes in contact with the hydrothermally dominated anoxic deep water by advective upwelling or diffusion, leading to the oxidation of Fe2+. Organic matter which settled with ferric iron components were further oxidized to produce a sediment further depleted in organic carbon. Limestone, however, is believed to have precipitated from an iron-depleted oxygenated shallow water layer close to the shore, in areas of high organic productivity and hence positive 13C.

13 This model is also supported by REE, Nd isotope and other δ Ccarb data (e.g. Jacobsen and Pimentel-Klose, 1988; Klein and Beukes, 1989; Kaufman et al., 1990). However, Fischer et al. (2009) showed for an example from the Campbellrand-Kuruman platform that there is no 13 systematic difference in δ Ccarb values for shallow- and deep-water carbonates of similar age to the iron formation. Heimann et al. (2010) data from the Kuruman Iron Formation also does not 13 indicate a decrease in δ Ccarb with increasing water depth. Hence, the formation of a negative δ13C deep water was not universal.

7.1.1.2 13C derived from oxidation of organic matter

An alternative explanation for the negative 13C is the result of the oxidation of organic matter at the expense of ferric iron (Fe3+ acts as an electron acceptor), producing ferrous iron, thus explaining why many iron formations are dominated by minerals containing Fe2+ such as magnetite and siderite.

This oxidation can be the result of the following three possible mechanisms:

- Precipitation of carbonates from pore water enriched in light carbon isotope due to abiotic oxidation of organic matter. - Metamorphism. - Microbial dissimilatory iron reduction (DIR).

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Precipitation of carbonates from pore water enriched in light carbon isotope due to abiotic oxidation of organic matter:

Becker and Clayton (1972) explained the low 13C values in the Hamersley Iron Formation as the result of precipitation of carbonates in pore waters enriched in the light carbon isotope. The source of this light carbon is the oxidation of organic matter.

Metamorphism:

Perry et al. (1973) considered that negative carbon isotope values ranging from -3.7 to -18.6‰ in the Biwabik Iron Formation (Minnesota) were the result of abiotic thermally driven reactions between organic matter and ferric oxides. They observed that the magnetite iron formation is significantly more 13C depleted than magnetite-free rock. They suggest that iron initially precipitated as a ferric oxide phase, was then partially reduced to form magnetite coupled to the oxidation of organic carbon after sedimentation and lithification.

However, Baur et al. (1985) showed that in the Hamersley Iron Formation isotopic differences do not follow metamorphic grade. Furthermore, they observed inhomogeneities preserved at the submillimeter scale.

Microbial dissimilatory iron reduction (DIR):

The model was first proposed by Lovley (1987).

A large kinetic isotope fractionation occurs when carbon is fixed into organic matter (e.g. Urey, 1947). This leads to a depletion of organic matter in δ13C. Negative values down to -40‰ have been observed for the Transvaal (Fischer et al. 2009).

13 13 Oxidation of this light δ C source would explain the low δ Ccarb values observed in iron formations. This took place in the soft sediment prior to lithification.

The involved processes are:

Fe2+ would be oxidized to form a ferric hydroxide according to the following equation:

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2+ - 4 Fe aq + O2 + 8 OH + 2 H2O → 4 Fe(OH)3

This reaction requires the presence of presence of oxygen. There is geological proof of the existence of cyanobacterial oxygenetic photosynthesis in the Archean (Tice and Lowe, 2004).

However, in the absence of oxygen as might be the case for the Helen Iron Formation, an alternative reaction might have taken place (anaerobic phototrophy):

2+ + 4 Fe aq + CO2 + 11 H2O → 4 Fe(OH)3 + CH2O + 8H

This reaction is also in good accordance with the fact that negative δ13C in the Helen Iron Formation are in close proximity to organic-rich horizons (see Chapter 6.3).

After the formation and settling of the iron hydroxyl gel along with organic matter to the seafloor, the organic matter would have been oxidized at the expense of ferric iron, leading to the precipitation of siderite:

- - 4 Fe(OH)3 + CH2O + 3 HCO3 → 4 FeCO3 + OH + 7 H2O

- 13 HCO3 in this case is of seawater origin with δ C of 0‰.

2+ If excess Fe(OH)3 is available, excess Fe aq might be produced according to:

2+ - 4 Fe(OH)3 + CH2O → FeCO3 + Fe aq + 6 OH + 4 H2O

2+ This excess Fe aq might then be used to form magnetite according to:

2+ - 3 Fe aq + 6 OH + 6 Fe(OH)3 → 3 Fe3O4 + 12 H2O

13 13 Even with a more conservative estimation of δ Corg as -30‰ for the Archean, δ Ccarb values down to -7.5‰ for the siderite are possible (Heimann et al. 2010).

2- Dissimilatory sulfate reduction played a minor role in the Archean due to the lack of SO4 in the deep ocean (Johnson et al. 2008b).

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7.2 The mineralogy of the Helen Iron Formation

In order to avoid the problem of measuring 13C values in siderite and ankerite, the mineralogy of the Helen Iron Formation was calculated based on its major element chemistry. First, electron microprobe analysis provided the average composition of the different minerals present, such as chlorite, ankerite and siderite (see Table 6.2).

CHLORITE ANKERITE SIDERITE

SiO2 27% FeO 15% FeO 48%

Al2O3 18% MnO 2% MnO 5% FeO 38% CaO 28% CaO 1% MgO 5% MgO 9% MgO 6% FeO+MnO 53% Table 7.2: Average major element compositions in weight % of chlorite, ankerite and siderite, measured by electron microprobe.

The calculation is based on:

- All Al2O3 is in chlorite

- The remaining SiO2 occurs as chert - All CaO is in ankerite - All the sulphur is in pyrite - The remaining FeO is in siderite

Figure 7.1 shows that the lower 25 meters of the siderite facies are dominated by SiO2 and silicates in the form of chert, chlorite and muscovite whereas siderite is minor (below 10%).

The remainder of the section is dominated by siderite and chert (90%), minor pyrite ranging up to 10% in the sulfide facies.

Ankerite is minor averaging 10%, falling steadily up section.

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Based on the carbonate mineralogy (Figure 7.2), the following samples have been excluded from 13 the Ccarb record because they were dominated by ankerite:

SJ_10_32; SJ_10_32.5; SJ_10_39; SJ_10_39.5

These samples were also high in strontium (see Chapter 4), thus indicating that strontium can be used as an indicator of ankerite/ferroan dolomite.

Figure 7.1: Estimated mineral percentages of the Helen Iron Figure 7.2: Estimated carbonate percentages of the Helen Iron Formation based on major element and microprobe analyses. Formation based on major element and microprobe analyses.

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13 7.3 The C geochemistry of the Helen Iron Formation

In contrast to the magnetite-hematite mineralogy of most known iron formations, the Helen Iron Formation has a very unusual mineralogy dominated by siderite. Other iron-bearing minerals are ankerite and pyrite, occurring in the sequence in various amounts rarely exceeding 10%.

Bolhar et al. (2005) found indications that the deposition of the Panorama Iron Formation in the Pilbara, which consists mainly of jasper, was interrupted episodically by phases of volcanic activity during which Fe2+ bearing carbonate formation became dominant. In the presence of saturated levels of Ca2+ and Fe2+, the formation of carbonate would have been further facilitated by elevated abundances of volcanic derived CO2. An input of CO2 into a restricted basin would lead to a lower pH. In this environment siderite is the more stable phase (Sugitani et al., 1998), 2+ 2- reaching saturation at 10 ppm Fe depending on the CO3 concentration (Drever 1974).

This model may also explain the mineralogy and geochemistry of the Helen Iron Formation (see also Chapter 3). The Helen Iron Formation rests on intermediate to felsic volcanic and discharge of magmatic CO2 would be considered to be the norm during such volcanic activity. Ash layers are common in the first 10-20 meters of the section.

However, the Helen Iron Formation also shows rare finely dispersed subhedral pyrite grains containing quartz inclusions, throughout the section.

Under strongly reducing conditions, bacterial sulfate reduction will result in the formation of pyrite (e.g. Berner, 1971; Klein and Bricker, 1977). However, in the absence of organic matter necessary to sustain sulfate reduction, siderite is the stable mineral. The Helen Iron Formation may therefore be an example of the balance between siderite and pyrite precipitation. There is also more common subhedral pyrite believed to be of diagenetic origin.

In the first 45 meters of the section (siderite-facies), pyrite is rare; however, in the next 5 meters (sulfate facies) pyrite ranges up to 10%. This observation indicates that the deepening of the basin led to more euxinic conditions favoring enhanced pyrite precipitation, probably related to increased deposition of organic matter.

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7.2.1 Previous research

13 Thode and Goodwin (1983) examined the Ccarb values of the Helen Iron Formation in a drill core (DDH U-2-647) from the MacLeod mine, 4 km northeast of Wawa (Figure 7.3). The section is unusually thick at 430 m with thickness of the siderite facies 75 m, the sulfide facies of 7 m and the topmost breccia facies of 350 m.

Figure 7.3: Locations of the two sections of the Helen Iron Formation

The authors analyzed 28 samples with an average spacing of 10-30 meters.

13 Their results show significant vertical variation and surprisingly positive Ccarb in the siderite facies (see Figure 7.4). A drop to negative values of -7.5‰ is observed in the breccia facies in horizons containing organic matter.

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7.2.2 This study

The MacLeod mine closed in the 1990’s and access is no longer possible. Thus for comparison purposes samples were collected from the Sir James Dunn pit located 6 km to the east (see Figure 7.3).

A total of 125 samples were analyzed corresponding to two samples per vertical meter, every third sample was analyzed in duplicate.

Significant similarities were observed between the two profiles (Figure 7.4) which were aligned based on their stratigraphy, the marker horizon being the base of the breccia facies.

Considering the limited number of samples analyzed by Thode and Goodwin (1983), the difficulties in analyzing siderite (see Chapter 2.4.3) and their relatively low precision of ±0.3‰, 13 the similarities in the Ccarb pattern of the two sections are significant.

Two prominent features are common to both sections:

- Values close to 0‰ prevail in the siderite facies (both sections). 13 - Two negative Ccarb excursions are observed in the breccia facies, one early with a magnitude of 3‰ (from 0‰ to -3‰) and one later with a magnitude of -10.5‰ (from 3‰ to -7.5‰). Sparse data in this part of the MacLeod section limits the possibility of exact estimation of the magnitude of the higher negative excursion.

The main differences between the two sections are:

- At the base of the section, the Sir James Dunn profile starts with values slightly below 0‰ whereas the MacLeod profile starts with values above. - An additional negative carbon isotope excursion with a magnitude of 3.5‰ between the two major negative excursions is visible in the Sir James Dunn section at 75 meters above the base. The reason why this feature is not visible in the MacLeod mine may be due to the lack of data in this part of the section. 13 - The MacLeod profile shows of Ccarb values returning to 0‰ in the upper 120 meters of the breccia facies. It appears that the Sir James Dunn section does not contain an

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equivalent of the upper part of the breccia facies observed in the MacLeod mine. In the Sir James Dunn section there is a 2 meter thick graphite-pyrite facies followed by basaltic pillow lavas.

Figure 7.4: Profile of the Sir James Dunn pit (left) and the MacLeod mine (right, data from Thode and Goodwin, 1983). The light green line on the left figure presents the data from the Sir James Dunn pit. Note the striking similarities of the dark green (MacLeod section) and light green (Sir James Dunn section) lines. Values close to zero at the beginning and end of the deposition of the Helen Iron Formation indicate a direct precipitate from seawater whereas the negative carbon isotope excursions are interpreted to be the result of post-depositional modification by dissimilatory iron reduction.

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Conclusions

Both sections have 13C values which are typically close to 0‰ and both show matching patterns of excursions from this norm, providing convincing evidence for their origin being related to primary precipitates since the excursions are located at the same stratigraphic horizons.

The question to be answered is: what is the origin of the negative excursions?

Possibilities include:

- Deposition in a stratified basin. - Inheritance of negative 13C values from organic matter.

The stratified ocean model (e.g. Beukes et al. 1989) is considered to be less probable since there is sedimentological evidence of a steady deepening in water depth of the Helen basin, which is not reflected in a steady decrease in 13C values, which would be expected if the basin was stratified with respect to 13C.

Since the negative carbon isotope excursions are in close proximity to organic-matter rich horizons, both in the Sir James Dunn Pit and in the MacLeod Mine, an inheritance of negative 13 δ Ccarb values from oxidation of organic matter is likely.

Possible mechanisms are:

- Abiotic reactions such as precipitation from a 13C depleted pore water fluid or metamorphic breakdown of organic matter - DIR

Abiotic reactions are ruled out since they would be constant for the whole section and not show values starting at 0‰, excursions of a magnitude of 10‰ and a return to 0%. Also, there is no petrographic indication of changes in metamorphic grade during the section.

Thus, it is concluded that the 13C values of the Helen Iron Formation are the result of the direct precipitation in seawater, with an overprint of oxidation of organic matter during DIR causing the negative excursions.

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This would explain both the presence of coarse-grained siderite and of magnetite in sample SJ_10_65 (see Figure 3.17) which is in the beginning of the first negative δ13C excursion (see Figure 7.4).

DIR predicts that organic carbon contents should correlate with δ13C of siderite, as is the case in 13 the Helen Iron Formation. Also the most negative δ Ccarb values observed are -7.5‰, again in good agreement for the role of DIR. Another proof is the presence of magnetite observed in thin 13 sections, close to the negative δ Ccarb excursions.

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CHAPTER 7: DISCUSSION AND CONCLUSIONS

The Helen Iron Formation is a typical example of a low metamorphic grade Algoma-type iron formation deposited in a reducing environment as evidenced by the lack of a positive cerium anomaly. The siderite facies is the result of deposition in fairly shallow water, a deepening of the basin during the breccia facies resulted in more reducing, hydrothermally influenced water as seen by an increase in the development of the positive europium anomaly. Hydrothermal input during the deposition was high but distal as indicated by low values of Zn, Pb and Cu but positive europium anomalies in the upper part of the section.

Volcanic detrital input during the deposition of the Helen Iron Formation was high at the beginning of the section as indicated by high aluminium, zirconium and titanium contents. Two types of rhyolitic detrital material were identified based on their REE geochemistry: F1- and F3- type rhyolite (Lesher et al., 1986). F1 type is a typical Archean rhyolite whereas F3-type is associated with VMS deposits.

REE concentrations are high in the first 20 meters of the section, declining rapidly to lower values in the upper part of the siderite facies. REE pattern of the Helen Iron Formation show an enrichment in the HREE, a positive yttrium anomaly as well as elevated lanthanum concentrations, which presents a typical seawater signature.

Oxygen played a minor role in the deposition of this iron formation since most of the iron is present in the reduced form Fe2+.

Carbon isotope data and petrographic studies indicate a direct precipitate from seawater at the beginning of the section and then an increasing role of dissimilatory iron reduction after that, causing two pronounced negative carbon isotope excursion in the breccia facies, both associated with sediment high in organic matter.

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The Sokoman iron formation on the other hand shows a high degree of diagenetic alteration. The precursor is believed to have been a silica-iron hydroxyl gel, burial and consequent dewatering of such a gel would lead to the high degree of diagenetic modification observed in the rock record. A paragenetic mineral assemblage was identified in thin section with hematite as the earliest iron mineral whereas magnetite is always a late diagenetic mineral. Quartz is both early and late diagenetic whereas stilpnomelane and minnesotaite are late diagenetic to early metamorphic respectively.

The REE geochemistry of the Sokoman Iron Formations indicates an interaction of two different water masses leading to the deposition of the iron formation. A shallow, photosynthetically oxidized water mass, as evident by a positive cerium anomaly, reacted with a deeper, hydrothermally influenced and reducing water rich in Fe2+. Oxidation of Fe2+ to Fe3+ led to the precipitation of the silica-iron-hydroxyl gel.

Detrital input into the Sokoman Iron Formation was negligible as indicated by low Al2O3, TiO2 and Zr values.

The two iron formations differ greatly in their age, mineralogy and water depth as well as in their geochemistry which is the result of the different types of water masses from which they were precipitated.

The main mechanism for the precipitation iron minerals in the Sokoman Iron Formation was oxidation of Fe2+ to Fe3+, probably along a chemocline whereas in the Helen Iron Formation, iron was precipitated in its reduced from as siderite.

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