Comparison between the Helen Iron Formation (Algoma-type ......Comparison between the Helen Iron...

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

by

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

© Copyright by Therese Ingrid Garcia 2014

ii

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

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.

iii

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+

.

iv

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.

v

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

2.3 13

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

vi

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

vii

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

7.1 13

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 13

C 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

viii

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)





ix




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 3.20: Thin section of SJ_10_80A, 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

x

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

Figure 3.28: Thin section 09-FH 3-3a, 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 3.43: Thin section 09-FH 8-4, 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

xi

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

xii

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

Table 7.1: Some examples of 13

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).

xiii

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

1

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.

2

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

3

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 km

2. They generally lack a direct association with volcanic rocks and

the host sequences lie unconformably on older crustal rocks. Superior-type iron formations

4

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

5

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

area as pyrite (FeS2) and other sulfides or is oxidized to ferric (Fe3+

) 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

6

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.

7

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 13

C values, which indicate

the onset of large-scale oxygen-producing photosynthesis that sequestered large quantities of 12

C

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 13

C were consequences (Frauenstein et al., 2009). Furthermore, a negative

shift in 34

S 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.

8

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

water as the molybdate anion MoO42-

, 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

9

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

10

(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.

11

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.

12

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

13

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).

14

15

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.

16

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.

17

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.

http://en.wikipedia.org/wiki/Dysprosium

18

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.

19

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.

20

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

21

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

22

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.

23

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.

24

2.3 13

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

90 hours in order to monitor the yield of the 13

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

25

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).

26

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.

27

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

28

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.

29

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

wide)

Figure 3.3d: Graphite facies with sill to the right, Sir James

Dunn Dunn Pit

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

Figure 3.3f: Sir James Dunn Pit

30

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.

31

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.


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.

32


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.

33


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).

34


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.

35


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.

36


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.

37


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.

38


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.

39


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.

40

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.

41


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.

42


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.

43


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.

44


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.

45


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.

46


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.

47

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.

48


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.

49


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).

50

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.

51

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.

52

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

53

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

54

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

55

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

56

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

57

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

58

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

59

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

60

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.


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.

61

The magnetite is subhedral and significantly coarser-grained, suggesting it is a late diagenetic

phase. It may be approximately coeval with the minnesotaite.


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.

62

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.

63

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.

64

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.

65

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.

66

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.

67

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.

68

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.

69

Upper Red Chert (URC)

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


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.

70


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.

71


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.

72


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.

73


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.

74


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.

75

The Green Chert (GC)

The Green Chert is dominated by quartz and minnesotaite.


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.

76


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.

77

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.

78

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).

79

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..

80

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.

81

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.

82

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.

83

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.

84

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.

85

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.

86

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%

87

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.

88

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.

89

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%.

90

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.

91

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

meters).

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

92

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.

93

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.

94

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

.

95

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.

96

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

97

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.

98

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.

99

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).

100

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.

101

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.

102

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.

103

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.

104

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).

105

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:

106

- 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

high-temperature fluids, this resulted in lower [CO32-

] 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).

107

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,

1989). It seems to be a function of temperature, pH, fO2, availability of a ligand such as CO32-

.

108

Recent seawater has a pH of 8-8.5 and a high availability of CO32-

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.

This resulted in lower [CO32-

] 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).

109

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.

110

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).

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

Figure 5.1: REE for Isua, Greenland (Dymek and Klein, 1988)

Data normalized to North American Shale Composite.

Upper line: 1:100 mixture of hydrothermal fluid and seawater.

Figure 5.2: REE for the Kuruman IF, South Africa (Klein

and Beukes 1989).

Data obtained by INAA and normalized to North American

Shale Composite.

111

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).

112

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).

113

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.

114

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

115

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.

116

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).

117

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).

118

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.

119

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.

120

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.

121

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)

122

123

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.

124

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.

125

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.

126

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.

127

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.

128

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.

129

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.

130

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.

131

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.

132

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.

133

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.

134

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.

135

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.

136

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.

137

CHAPTER 7: THE CARBON ISOTOPE

GEOCHEMISTRY OF THE HELEN IRON

FORMATION

7.1 13

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 13

C values (see

Table 7.1).

The Helen Iron Formation on the other hand shows 13

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.

138

Name of the iron formation and age Observed δ13Ccarb values Reference

Carajas Iron Formation (Brazil); 2.76 Ga -3 to -6‰ Sial et al. (2000)

Kuruman-Griquatown Iron Formation

(Transvaal); 2.5 Ga

-5.45 to -11.93‰ Beukes and Klein (1989)

Dales George Member, Brockman Iron

Formation (Hamersley); 2.5 Ga

-7 to -10.5‰ Kaufman et al. (1990)

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

Marra Mamba Iron Formation,

Hamersley; 2.5 Ga

+2 to +0.5‰ for limestone

-2.8 to -19.8‰ for IF

Baur et al. (1985)

Hotazel Iron Formation (Transvaal); 2.4-

2.2 Ga

-4 to -18‰ Tsikos et al. (2003)

Gunflint Iron Formation (Ontario); 1.8 Ga -6 to 2.6‰ deep water

2.5 to +0.5‰ shallow water

Winter and Knauth (1992)

Table 7.1: Some examples of 13Ccarb 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

Beukes et al. (1990) observed a decrease in δ13

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,

allowing for the transport of iron as Fe2+

. Deeper water would thus have mantle δ13

Ccarb values of

-5±2‰, whereas shallow water would be 0‰.

139

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 13

C.

This model is also supported by REE, Nd isotope and other δ13

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

systematic difference in δ13

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

indicate a decrease in δ13

Ccarb with increasing water depth. Hence, the formation of a negative

δ13

C deep water was not universal.

7.1.1.2 13

C derived from oxidation of organic matter

An alternative explanation for the negative 13

C 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).

140

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 13

C 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 13

C 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 δ13

C. Negative values down to -40‰ have

been observed for the Transvaal (Fischer et al. 2009).

Oxidation of this light δ13

C source would explain the low δ13

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:

141

4 Fe2+

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):

4 Fe2+

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

This reaction is also in good accordance with the fact that negative δ13

C 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

HCO3- in this case is of seawater origin with δ

13C of 0‰.

If excess Fe(OH)3 is available, excess Fe2+

aq might be produced according to:

4 Fe(OH)3 + CH2O → FeCO3 + Fe2+

aq + 6 OH- + 4 H2O

This excess Fe2+

aq might then be used to form magnetite according to:

3 Fe2+

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

Even with a more conservative estimation of δ13

Corg as -30‰ for the Archean, δ13

Ccarb values

down to -7.5‰ for the siderite are possible (Heimann et al. 2010).

Dissimilatory sulfate reduction played a minor role in the Archean due to the lack of SO42- in the

deep ocean (Johnson et al. 2008b).

142

7.2 The mineralogy of the Helen Iron Formation

In order to avoid the problem of measuring 13

C 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%

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.

Table 7.2: Average major element compositions in weight % of

chlorite, ankerite and siderite, measured by electron microprobe.

143

Based on the carbonate mineralogy (Figure 7.2), the following samples have been excluded from

the 13

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

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.

144

7.3 The 13

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),

reaching saturation at 10 ppm Fe2+

depending on the CO32-

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.

145

7.2.1 Previous research

Thode and Goodwin (1983) examined the 13

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.

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

Their results show significant vertical variation and surprisingly positive 13

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.

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

146

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‰,

the similarities in the 13

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).

- Two negative 13

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.

- The MacLeod profile shows of 13

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

147

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.

148

Conclusions

Both sections have 13

C 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 13

C 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 13

C values, which would be expected if the basin was

stratified with respect to 13

C.

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

δ13

Ccarb values from oxidation of organic matter is likely.

Possible mechanisms are:

- Abiotic reactions such as precipitation from a 13

C 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 13

C 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.

149

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 δ13

C excursion (see

Figure 7.4).

DIR predicts that organic carbon contents should correlate with δ13

C of siderite, as is the case in

the Helen Iron Formation. Also the most negative δ13

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

sections, close to the negative δ13

Ccarb excursions.

150

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.

151

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.

152

REFERENCES

Albarede, F., and Michard, A., 1986, Transfer of continental Mg, S, O and U to the mantle

through hydrothermal alteration of the oceanic crust: Chemical Geology, v. 57, p. 1-15.

Alexander, B. W., Bau, M., Andersson, P.,and Dulski, P., 2008, Continentally-derived solutes

in shallow Archean seawater: Rare earth element and Nd isotope evidence in iron

formation from the 2.9 Ga Pongola Supergroup, South Africa: Geochimica Et

Cosmochimica Acta, v. 72, p. 378-394.

Alibo, D. S., and Nozaki, Y., 1999, Rare earth elements in seawater: Particle association, shale-

normalization, and Ce oxidation: Geochimica Et Cosmochimica Acta, v. 63, p. 363-372.

Anbar, A. D., Duan, Y., Lyons, T. W., Arnold, G. L., Kendall, B., Creaser, R. A., Kaufman, A.

J., Gordon, G. W., Scott, C., Garvin, J., and Buick, R., 2007, A whiff of oxygen before

the Great Oxidation Event?: Science, v. 317, p. 1903-1906.

Andre, L., Cardinal, D., Alleman, L. Y., and Moorbath, S., 2006, Silicon isotopes in 3.8 Ga West

Greenland rocks as clues to the Eoarchaean supracrustal Si cycle: Earth and Planetary

Science Letters, v. 245, p. 162-173.

Arnold, G. L., Anbar, A. D., Barling, J., and Lyons, T. W., 2004, Molybdenum isotope evidence

for widespread anoxia in mid-proterozoic oceans: Science, v. 304, p. 87-90.

Barley, M. E., Pickard, A. L., and Sylvester, P. J., 1997, Emplacement of a large igneous

province as a possible cause of banded iron formation 2.45 billion years ago: Nature, v.

385, p. 55-58.

Bau, M., 1993, Effects of syn- and post-depositional processes on the rare-earth element

distribution in Precambrian iron-formations: European Journal of Mineralogy, v. 5, p.

257-267.

Bau, M., and Dulski, P., 1996, Distribution of yttrium and rare-earth elements in the Penge and

Kuruman iron-formations, Transvaal Supergroup, South Africa: Precambrian Research,

v. 79, p. 37-55.

Bau, M., and Möller, P., 1993, Rare earth element systematics of the chemically precipitated

component in early precambrian iron formations and the evolution of the terrestrial

atmosphere-hydrosphere-lithosphere system: Geochimica Et Cosmochimica Acta, v. 57,

p. 2239-2249.

Baur, M. E., Hayes, J. M., Studley, S. A., and Walter, M. R., 1985, Millimeter-scale variations of

stable isotope abundances in carbonates from banded iron-formations in the Hamersley

Group of Western Australia.: Economic Geology, v. 80, p. 270-282.

Becker, R. H., and Clayton, R. N., 1972, Carbon isotopic evidence for the origin of a banded

iron-formation in Western Australia: Geochimica Et Cosmochimica Acta, v. 36, p. 577-

&.

Bekker, A., Holland, H. D., Wang, P. L., Rumble, D., Stein, H. J., Hannah, J. L., Coetzee, L. L.,

and Beukes, N. J., 2004, Dating the rise of atmospheric oxygen: Nature, v. 427, p. 117-

120.

Bekker, A., Slack, J. F., Planavsky, N., Krapez, B., Hofmann, A., Konhauser, K. O., and

Rouxel, O. J., 2010, Iron Formation: The Sedimentary Product of a Complex Interplay

among Mantle, Tectonic, Oceanic, and Biospheric Processes: Economic Geology, v. 105,

p. 467-508.

153

Belevtsev, Y. N., Belevtsev, R. Y., and Siroshtan, R. I., 1982, The Krivoy Rog Basin. In:

Trendall A.F., Morris R.C. (eds.), ‘Iron-formation: facts and problems’, Elsevier,

Amsterdam, pp. 211-252.

Berner, R. A., 1971, Principles of sedimentology: New York, McGraw-Hill, 240 p.

Beukes, N. J., and Gutzmer, J., 2008, Origin and Paleoenvironmental Significance of Major Iron

Formations at the Archean-Paleoproterozoic Boundary: Banded Iron Formation-Related

High-Grade Iron Ore, v. 15, p. 5-47.

Beukes, N. J., and Klein, C., 1990, Geochemistry and sedimentology of a facies transition —

from microbanded to granular iron-formation — in the early Proterozoic Transvaal

Supergroup, South Africa: Precambrian Research, v. 47, p. 99-139.

Beukes, N. J., and Klein, C., 1992, Models for iron-formation deposition, in: J.W. Schopf, C.

Klein (Eds.), The Proterozoic Biosphere: A Multidisciplinary Study., Cambridge

University Press, Cambridge, UK.

Beukes, N. J., Klein, C., Kaufman, A. J., and Hayes, J. M., 1990, Carbonate petrography,

kerogen distribution, and carbon and oxygen isotope variations in an early Proterozoic

transition from limestone to iron-formation deposition, Transvaal Supergroup, South

Africa.: Economic Geology and the Bulletin of the Society of Economic Geologists, v.

85, p. 663-690.

Bilal, B. A., 1991, Thermodynamic study of Eu3+/Eu2+ redox reaction in aqueous solutions at

elevated temperatures and pressures by means of cyclic voltammetry.: Zeitung fuer

Naturforschung, v. 46a, p. 1108-1116.

Bolhar, R., Van Kranendonk, M. J., and Kamber, B. S., 2005, A trace element study of siderite-

jasper banded iron formation in the 3.45 Ga Warrawoona Group, Pilbara Craton -

Formation from hydrothermal fluids and shallow seawater: Precambrian Research, v.

137, p. 93-114.

Bonatti, E., Kraemer, T., and Rydell, H., 1972, Classification and genesis of iron-manganese

deposits. In: Hornd, D.R., ed., Ferro-Manganese Deposits on the Ocean Floor: New York,

Harriman, Arden House and Lamont-Doherty Geological Observatory, p.149-166.

Bostrom, K., 1970, Submarine volcanism as a source for iron: Earth and Planetary Science

Letters, v. 9, p. 348-&.

Bowins, R. J., and Crocket, J. H., 2011, Monazite, Xenotime and REE minerals in Archean

banded iron-formation from the Sherman and Adams mines, Ontario, Canada: Canadian

Mineralogist, v. 49, p. 749-763.

Broecker, W. S., and Maier-Reimer, E., 1992, The influence of air and sea exchange on the

carbon isotope distribution in the sea: Global Biogeochemical Cycles, v. 6, p. 315-320.

Byrne, R. H., Kump, L. R., and Cantrell, K. J., 1988, The influence of temperature and pH on

trace metal speciation in seawater: Marine Chemistry, v. 25, p. 163-181.

Cairnssmith, A. G., 1978, Precambrian solution photochemistry, inverse segregation, and banded

iron formations: Nature, v. 276, p. 807-808.

Campbell, I. H., Griffiths, R. W., and Hill, R. I., 1989, Melting in an Archaean mantle plume:

heads it's basalts, tails it's komatiites: Nature, v. 339, p. 697-699.

Canfield, D. E., 2005, The early history of atmospheric oxygen: Homage to Robert A. Garrels:

Annual Review of Earth and Planetary Sciences, v. 33, p. 1-36.

Canfield, D. E., and Teske, A., 1996, Late Proterozoic rise in atmospheric oxygen concentration

inferred from phylogenetic and sulphur-isotope studies: Nature, v. 382, p. 127-132.

Cantrell, K. J., and Byrne, R. H., 1987, Temperature dependence of europium carbonate

154

complexation: Journal of Solution Chemistry, v. 16, p. 555-566.

Carrigan, W. J., and Cameron, E. M., 1991, Petrological and stable isotope studies of carbonate

and sulfide minerals from the Gunflint Formation, Ontario: evidence for the origin of

early Proterozoic iron-formation: Precambrian Research, v. 52, p. 347-380.

Chandler, F. W., 1980, Proterozoic redbed sequences of Canada.: Can. Geol. Surv. Bull., v. 311.

Chauvel, J. J., and Dimroth, E., 1974, Facies types and depositional environment of Sokoman

Iron Formation, Central Labrador Trough, Quebec, Canada: Journal of Sedimentary

Petrology, v. 44, p. 299-327.

Claire, M. W., Catling, D. C., and Zahnle, K. J., 2006, Biogeochemical modelling of the rise in

atmospheric oxygen: Geobiology, v. 4, p. 239-269.

Cloud, P., 1972, A Working Model of the Primitive Earth: American Journal of Science, v. 272,

p. 537-548.

Condie, K. C., 1998, Episodic continental growth and supercontinents: a mantle avalanche

connection?: Earth and Planetary Science Letters, v. 163, p. 97-108.

Condie, K. C., 2002, Continental growth during a 1.9-Ga superplume event: Journal of

Geodynamics, v. 34, p. 249-264.

Condie, K. C., Marais, D. J. D., and Abbott, D., 2001, Precambrian superplumes and

supercontinents: a record in black shales, carbon isotopes, and paleoclimates?:

Precambrian Research, v. 106, p. 239-260.

Condie, K. C., O'Neill, C., and Aster, R. C., 2009, Evidence and implications for a widespread

magmatic shutdown for 250 My on Earth: Earth and Planetary Science Letters, v. 282, p.

294-298.

Craddock, P. R., Bach, W., Seewald, J. S., Rouxel, O. J., Reeves, E., and Tivey, M. K., 2010,

Rare earth element abundances in hydrothermal fluids from the Manus Basin, Papua New

Guinea: Indicators of sub-seafloor hydrothermal processes in back-arc basins:

Geochimica Et Cosmochimica Acta, v. 74, p. 5494-5513.

Danielson, A., Möller, P., and Dulski, P., 1992, The europium anomalies in banded iron

formations and the thermal history of the oceanic crust: Chemical Geology, v. 97, p. 89-

100.

Derry, L. A., and Jacobsen, S. B., 1990, The chemical evolution of Precambrian seawater:

Evidence from REEs in banded iron formations: Geochimica Et Cosmochimica Acta, v.

54, p. 2965-2977.

Derry, L. A., Kaufman, A. J., and Jacobsen, S. B., 1992, Sedimentary cycling and environmental

change in the Late Proterozoic:Evidence from stable and radiogenic isotopes:


Dickinso.Wr, and Luth, W. C., 1971, A model for plate tectonic evolution of mantle layers.:

Science, v. 174, p. 400-&.

Dimroth, E., 1972, Labrador Geosyncline revisted: American Journal of Science, v. 272, p. 487-

495.

Dimroth, E., 1976, Aspects of the sedimentary petrology of cherty iron-formations,in Wolf, K.

H., ed., Handbook of strata-bound and stratiform ore deposits: Amsterdam, Elsevitr, v. 7,

p. 203-254.

Dimroth, E., 1977, Facies Models 5 - Models of physical sedimentation of Iron Formations:

Geoscience Canada, v. 4, p. 23-30.

Dimroth, E., and Chauvel, J. J., 1973, Petrography of Sokoman Iron Formation in part of central

Labrador-Trough, Quebec, Canada: Geological Society of America Bulletin, v. 84, p.

155

111-134.

Dimroth, E., and Dressler, B., 1978, Metamorphosis of the Labrador Trough. In: Metamorphism

in the Canadian Shield (J.A. Fraser & W.W. Heywood, eds.). Geol. Sury. Can. Rep., v.

78, p. 215-236.

Ding, T., Wan, D., Wang, C., and Zhang, F., 2004, Silicon isotope compositions of dissolved

silicon and suspended matter in the Yangtze River, China: Geochimica Et Cosmochimica

Acta, v. 68, p. 205-216.

Dossing, L. N., Frei, R., Stendal, H., and Mapeo, R. B. M., 2009, Characterization of enriched

lithospheric mantle components in similar to 2.7 Ga Banded Iron Formations: An

example from the Tati Greenstone Belt, Northeastern Botswana: Precambrian Research,

v. 172, p. 334-356.

Drever, J. I., 1974, Geochemical Model for the Origin of Precambrian Banded Iron Formations:

Geological Society of America Bulletin, v. 85, p. 1099-1106.

Dymek, R. F., and Klein, C., 1988, Chemistry, petrology and origin of banded iron-formation

lithologies from the 3800 MA isua supracrustal belt, West Greenland: Precambrian

Research, v. 39, p. 247-302.

El Tabakh, M., Grey, K., Pirajno, F., and Schreiber, B. C., 1999, Pseudomorphs after evaporitic

minerals interbedded with 2.2 Ga stromatolites of the Yerrida basin, Western Australia:

Origin and significance: Geology, v. 27, p. 871-874.

Emerson, S. R., and Huested, S. S., 1991, Ocean anoxia and the concentrations of molybdenum

and vanadium in seawater: Marine Chemistry, v. 34, p. 177-196.

Erickson, B. E., and Helz, G. R., 2000, Molybdenum(VI) speciation in sulfidic waters: Stability

and lability of thiomolybdates: Geochimica Et Cosmochimica Acta, v. 64, p. 1149-1158.

Eriksson, P. G., Banerjee, S., Catuneanu, O., Sarkar, S., Bumby, A. J., and Mtimkulu, M. N.,

2007, Prime controls on Archaean-Palaeoproterozoic sedimentation: Change over time:

Gondwana Research, v. 12, p. 550-559.

Ernst, 1983, The early earth and the Archean rock record, in Schopf, J. W., ed., Earth's Earliest

Biosphere:Its origin and evolution., Princeton University Press.

Evans, K. A., McCuaig, T. C., Leach, D., Angerer, T., and Hagemann, S. G., 2013, Banded iron

formation to iron ore: A record of the evolution of Earth environments?: Geology, v. 41,

p. 99-102.

Ewers, W. E., 1980, Chemical conditions for the precipitation of banded iron-formations.

In: P.H. Trudinger and M. Walter (Editors), Fourth International Symposium on

Environmental Biogeochemistry. Springer-Verlag, New York, NY.

Fedorowich, J. S., Richards, J. P., Jain, J. C., Kerrich, R., and Fan, J., 1993, A rapid method for

REE and trace-element analysis using laser sampling ICP-MS on direct fusion whole-

rock glasses: Chemical Geology, v. 106, p. 229-249.

Findlay, J. M., Parrish, R. R., Birkett, T. C., and Watanabe, D. H., 1995, U–Pb ages from the

Nimish Formation and Montagnais glomeroporphyritic gabbro of the central New

Québec Orogen, Canada: Canadian Journal of Earth Sciences, v. 32, p. 1208-1220.

Fink, 1976, Petrology of the Sokoman Iron Formation in the Howells River Area, at the Western

Edge of the Labrador Trough: Economic Geology, v. 71, p. 453-487.

Fink, R., 1972, Summary of development work on the Howells River Project with special

reference to the geology of the iron-formation: Iron Ore Co. of Canada Report 19.

Fischer, W. W., and Knoll, A. H., 2009, An iron shuttle for deepwater silica in Late Archean and

early Paleoproterozoic iron formation: Geological Society of America Bulletin, v. 121, p.

156

222-235.

Fischer, W. W., Schroeder, S., Lacassie, J. P., Beukes, N. J., Goldberg, T., Strauss, H.,

Horstmann, U. E., Schrag, D. P., and Knoll, A. H., 2009, Isotopic constraints on the Late

Archean carbon cycle from the Transvaal Supergroup along the western margin of the

Kaapvaal Craton, South Africa: Precambrian Research, v. 169, p. 15-27.

Fouquet, Y., Vonstackelberg, U., Charlou, J. L., Erzinger, J., Herzig, P. M., Muhe, R., and

Wiedicke, M., 1993, Metallogenesis in Back-Arc environments: The Lau Basin example:

Economic Geology and the Bulletin of the Society of Economic Geologists, v. 88, p.

2154-2181.

Fralick, P., and Pufahl, P. K., 2006, Iron formation in neoarchean deltaic successions and the

microbially mediated deposition of transgressive systems tracts: Journal of Sedimentary

Research, v. 76, p. 1057-1066.

Franklin, J. M., Gibson, H. L., Galley, A. G., and Jonasson, I. R., 2005, Volcanogenic Massive

Sulfide Deposits, in Hedenquist, J. W., Thompson, J. F. H., Goldfarb, R. J., and Richards,

J. P., eds., Economic Geology 100th Anniversary Volume: Littleton, CO, Society of

Economic Geologists, p. 523-560.

Frauenstein, F., Veizer, J., Beukes, N., Van Niekerk, H. S., and Coetzee, L. L., 2009, Transvaal

Supergroup carbonates: Implications for Paleoproterozoic delta O-18 and delta C-13

records: Precambrian Research, v. 175, p. 149-160.

Friedman, I., O'Neil, J. R., and Cebula, G., 1982, Two new carbonate stable isotope standards.

Geostandards Newsletter, 6, p.11-12.

Fryer, B. J., 1977, Trace element geochemistry of the Sokoman Iron Formation: Canadian

Journal of Earth Sciences, v. 14, p. 1598-1610.

Garrels, R. M., Perry, J. E. A., and Mackenzie, F. T., 1973, Genesis of Precambrian iron

formations and the development of atmospheric oxygen: Economic Geology, p. 1173–

1179.

Gaucher, E. A., Govindarajan, S., and Ganesh, O. K., 2008, Palaeotemperature trend for

Precambrian life inferred from resurrected proteins: Nature, v. 451, p. 704-U2.

Goldblatt, C., Lenton, T. M., and Watson, A. J., 2006, Bistability of atmospheric oxygen and the

Great Oxidation: Nature, v. 443, p. 683-686.

Goldstein, S. J., and Jacobsen, S. B., 1988, Rare Earth Elements in river waters: Earth and

Planetary Science Letters, v. 89, p. 35-47.

Gonfiantini, R., Stichler, W., and Rozanski, K., 1995, Standards and intercomparison materials

distributed by the International Atomic Energy Agency for stable isotope measurement.

Reference and intercomparison materials for stable isotopes of light elements.

International Atomic Energy Agency, IAEA-TECDOC-825 (Vienna, Austria), p. 13–29.

Goodwin, A. M., 1961, Genetic Aspects of Michipicoten Iron Formations: The Canadian mining

and Metallurgical Bulletin.

Goodwin, A. M., 1962, Structure, Stratigraphy, and Origin of Iron Formations, Michipicoten

Area, Algoma District, Ontario, Canada: Geological Society of America Bulletin, v. 73,

p. 561-&.

Goodwin, A. M., 1973, Archean Volcanogenic Iron Formation of the Canadian Shield. In:

UNESCO. Genesis of Pre-Cambrian Iron and Manganese Deposits. Proc. Kiev Symp.

1970 (Earth Sciences 9).

Grauch, R. I., 1989, Rare Earth Elements in metamorphic rocks: Geochemistry and Mineralogy

of Rare Earth Elements, v. 21, p. 147-167.

157

Gromet, L. P., Dymek, R. F., Haskin, L. A., and Korotev, R. L., 1984, The North-American

Shale Composite - its compilation, major and trace-element characteristics: Geochimica

Et Cosmochimica Acta, v. 48, p. 2469-2482.

Gross, G. A., 1965, Geology of Iron Deposits in Canada. Volume 1: General Geology and

Evaluation of Iron Deposits, Economic Geology Report, 22.

Gröning, M., 2004, International stable isotope reference materials, In: de Groot P.A. (ed.),

Handbook of stable isotope analytical techniques (Volume 1). Elsevier (Amsterdam, The

Netherlands), 875–906.

Hamade, T., Konhauser, K. O., Raiswell, R., Goldsmith, S., and Morris, R. C., 2003, Using

Ge/Si ratios to decouple iron and silica fluxes in Precambrian banded iron formations:

Geology, v. 31, p. 35-38.

Haqq-Misra, J. D., Domagal-Goldman, S. D., Kasting, P. J., and Kasting, J. F., 2008, A Revised,

Hazy Methane Greenhouse for the Archean Earth: Astrobiology, v. 8, p. 1127-1137.

Heaman, L. M., Peck, D., and Toope, K., 2009, Timing and geochemistry of 1.88 Ga Molson

Igneous Events, Manitoba: Insights into the formation of a craton-scale magmatic and

metallogenic province: Precambrian Research, v. 172, p. 143-162.

Heimann, A., Johnson, C. M., Beard, B. L., Valley, J. W., Roden, E. E., Spicuzza, M. J., and

Beukes, N. J., 2010, Fe, C, and O isotope compositions of banded iron formation

carbonates demonstrate a major role for dissimilatory iron reduction in similar to 2.5 Ga

marine environments: Earth and Planetary Science Letters, v. 294, p. 8-18.

Hoffman, P. F., 1988, United Plates of America, The Birth of a Craton: Early Proterozoic

Assembly and Growth of Laurentia: Annual Review of Earth and Planetary Sciences, v.

16, p. 543-603.

Hofmann, H. J., Sage, R. P., and Berdusco, E. N., 1991, Archean stromatolites in Michipicoten

Group siderite ore at Wawa, Ontario: Economic Geology and the Bulletin of the Society

of Economic Geologists, v. 86, p. 1023-1030.

Holland, H. D., 1978, CO2-HCO3(-)-CO3(2-) system through geologic time and its

implicationsn for the evolution of the atmosphere: Pure and Applied Geophysics, v. 116,

p. 232-233.

Holland, H. D., 2002, Volcanic gases, black smokers, and the Great Oxidation Event:


Holland, H. D., 2006, The oxygenation of the atmosphere and oceans: Philosophical

Transactions of the Royal Society B-Biological Sciences, v. 361, p. 903-915.

Holland, H. H., 1984, The Chemical Evolution of the Atmosphere and the Oceans, Princeton

University Press.

Hren, M. T., Tice, M. M., and Chamberlain, C. P., 2009, Oxygen and hydrogen isotope evidence

for a temperate climate 3.42 billion years ago: Nature, v. 462, p. 205-208.

Huston, D. L., and Logan, G. A., 2004, Barite, BIFs and bugs: evidence for the evolution of the

Earth's early hydrosphere: Earth and Planetary Science Letters, v. 220, p. 41-55.

Isley, A. E., 1995, Hydrothermal Plumes and the Delivery of Iron to Banded Iron Formation:

Journal of Geology, v. 103, p. 169-185.

Isley, A. E., and Abbott, D. H., 1999, Plume-related mafic volcanism and the deposition of

banded iron formation: Journal of Geophysical Research-Solid Earth, v. 104, p. 15461-

15477.

Jacobsen, S. B., and Pimentelklose, M. R., 1988, A Nd isotopic study of the Hamersley and

Michipicoten banded iron formations: the source of REE and Fe in Archean oceans:

158

Earth and Planetary Science Letters, v. 87, p. 29-44.

James, H. L., 1954, Sedimentary facies of iron-formation: Economic Geology, v. 49, p. 235-293.

Johnson, C. M., Beard, B. L., Klein, C., Beukes, N. J., and Roden, E. E., 2008a, Iron isotopes

constrain biologic and abiologic processes in banded iron formation genesis: Geochimica

Et Cosmochimica Acta, v. 72, p. 151-169.

Johnson, C. M., Beard, B. L., and Roden, E. E., 2008b, The iron isotope fingerprints of redox

and biogeochemical cycling in the modern and ancient Earth: Annual Review of Earth

and Planetary Sciences, v. 36, p. 457-493.

Kamber, B. S., Greig, A., and Collerson, K. D., 2005, A new estimate for the composition of

weathered young upper continental crust from alluvial sediments, Queensland, Australia:


Karhu, J. A., and Holland, H. D., 1996, Carbon isotopes and the rise of atmospheric oxygen:

Geology, v. 24, p. 867-870.

Kasting, J. F., 2005, Methane and climate during the Precambrian era: Precambrian Research, v.

137, p. 119-129.

Kaufman, A. J., Hayes, J. M., and Klein, C., 1990, Primary and diagenetic controls of isotopic

compositions of iron-formation carbonates.: Geochimica Et Cosmochimica Acta, v. 54, p.

3461-3473.

Keller, W. D., Balgord, W. D., and Reesman, A. L., 1963, Dissolved products of artificially

pulverized silicate minerals and rocks: part I: Journal of Sedimentary Petrology, v. 33, p.

191-204.

Kemp, A. I. S., Hawkesworth, C. J., Paterson, B. A., and Kinny, P. D., 2006, Episodic growth of

the Gondwana supercontinent from hafnium and oxygen isotopes in zircon: Nature, v.

439, p. 580-583.

Khan, R. M. K., DasSharma, S., Patil, D. J., and Naqvi, S. M., 1996, Trace, rare-earth element,

and oxygen isotopic systematics for the genesis of banded iron-formations: Evidence

from Kushtagi schist belt, Archaean Dharwar Craton, India: Geochimica Et


Kimberley, M. M., 1974, Origin of iron-ore by diagenetic replacement of calcareous oolite:

Nature, v. 250, p. 319-320.

King, A. M., 1951, Depth of disturbance of sand on sea beaches by waves: Journal of

Sedimentary Petrology, v. 21, p. 131-140.

Klein, C., 1974, Greenalite, stilpnomelane, minnesotaite, crocidolite and carbonates in a very

low-grade metamorphic Precambrian Iron-Formation: Canadian Mineralogist, v. 12, p.

475-498.

Klein, C., 2005, Some Precambrian banded iron-formations (BIFs) from around the world: Their

age, geologic setting, mineralogy, metamorphism, geochemistry, and origin: American

Mineralogist, v. 90, p. 1473-1499.

Klein, C., and Beukes, N. J., 1989, Geochemistry and sedimentology of a facies transition from

limestone to iron-formation deposition in the early Proterozoic Transvaal

Supergroup, South Africa: Economic Geology, v. 84, p. 1733-1774.

Klein, C., and Beukes, N. J., 1993, Sedimentology and geochemistry of the glaciogenic late

Proterozoic Rapitan Iron-Formation in Canada: Economic Geology and the Bulletin of

the Society of Economic Geologists, v. 88, p. 542-565.

Klein, C., and Bricker, O. P., 1977, Some Aspects of the Sedimentary and diagenetic

environment of Proterozoic Banded Iron-Formation: Economic Geology, v. 72, p. 1457-

159

1470.

Klein, C., and Fink, R. P., 1976, Petrology of the Sokoman Iron Formation in Howell's river

area, at the western edge of the Labrador Trough: Economic Geology, v. 71, p. 453-487.

Knauth, L. P., and Lowe, D. R., 2003, High Archean climatic temperature inferred from oxygen

isotope geochemistry of cherts in the 3.5 Ga Swaziland Supergroup, South Africa:

Geological Society of America Bulletin, v. 115, p. 566-580.

Konhauser, K. O., Lalonde, S. V., Amskold, L., and Holland, H. D., 2007, Was there really an

Archean phosphate crisis?: Science, v. 315, p. 1234-1234.

Kopp, R. E., Kirschvink, J. L., Hilburn, I. A., and Nash, C. Z., 2005, The paleoproterozoic

snowball Earth: A climate disaster triggered by the evolution of oxygenic photosynthesis:

Proceedings of the National Academy of Sciences of the United States of America, v.

102, p. 11131-11136.

Krapez, B., Barley, M. E., and Pickard, A. L., 2003, Hydrothermal and resedimented origins of

the precursor sediments to banded iron formation: sedimentological evidence from the

Early Palaeoproterozoic Brockman Supersequence of Western Australia: Sedimentology,

v. 50, p. 979-1011.

Kroopnick, P. M., 1985, The distribution of C-13 of sigma-CO2 in the world oceans: Deep-Sea

Research Part a-Oceanographic Research Papers, v. 32, p. 57-84.

Kump, L. R., and Barley, M. E., 2007, Increased subaerial volcanism and the rise of atmospheric

oxygen 2.5 billion years ago: Nature, v. 448, p. 1033-1036.

Lepp, H., and Goldich, S. S., 1964, Origin of Precambrian iron formations: Economic Geology,

v. 59, p. 1025–1060.

Lesher, C. M., Goodwin, A. M., Campbell, I. H., and Gorton, M. P., 1986, Trace-element

geochemistry of ore-associated and barren, felsic metavolcanic rocks in the Superior

Province, Canada: Canadian Journal of Earth Sciences, v. 23, p. 222-237.

Lovley, D. R., Stolz, J. F., Nord, G. L., and Phillips, E. J. P., 1987, Anaerobic production of

magnetite by a dissimilatory iron-reducing microorganism: Nature, v. 330, p. 252-254.

Machado, N., 1990, Timing of collisional events in the Trans-Hudson Orogen: evidence from U–

Pb geochronology for the New Quebec Orogen, the Thompson Belt, and the Reindeer

Zone (Manitoba and Saskatchewan). In: Lewry, J.F., Stauffer, M.R. (Eds.), The Early

Proterozoic Trans-Hudson Orogen of North America. Geol. Ass. Can. Special Paper 37,

433–441.

Macrae, N. D., Nesbitt, H. W., and Kronberg, B. I., 1992, Development of a positive Eu anomaly

during diagenesis: Earth and Planetary Science Letters, v. 109, p. 585-591.

Maliva, R. G., Knoll, A. H., and Simonson, B. M., 2005, Secular change in the Precambrian

silica cycle: Insights from chert petrology: Geological Society of America Bulletin, v.

117, p. 835-845.

Manning, C. E., 1994, The solubility of quartz in H2O in the lower crust and upper-mantle:


McDonough, W. F., and Sun, S. S., 1995, The composition of the earth: Chemical Geology, v.

120, p. 223-253.

McLennan, S. M., 1989, Rare earth elements in sedimentary rocks: influence of Provenance and

sedimentary processes. In. Lipin, B.R. & McKay, G.A.(eds.): Geochemistry and

mineralogy of rare earth elements. Mineralogical Society of America. Reviews in

Mineralogy 21, 169-200.

Meert, J. G., Pandit, M. K., Pradhan, V. R., and Kamenov, G., 2011, Preliminary report on the

160

paleomagnetism of 1.88 Ga dykes from the Bastar and Dharwar cratons, Peninsular India:

Gondwana Research, v. 20, p. 335-343.

Mel'nick, I., 1973, Physiochemical conditions of formation of ferruginous quartzite: Kiev

Akadimi Nauk Ukrainskoi SSR Institut Geokhimii i Fiziki Minerolov.

Menard, H. W., 1967, Seafloor spreading, topography and the second layer: Science, v. 157, p.

923-&.

Michard, A., Albarede, F., Michard, G., Minster, J. F., and Charlou, J. L., 1983, Rare-earth

elements and uranium in high-temperature solutions from East Pacific Rise hydrothermal

vent field (13 °N): Nature, v. 303, p. 795-797.

Moller, P., Dulski, P., and Bau, M., 1994, Rare Earth Element Adsorption in a seawater profile

above the East Pacific Rise: Chemie Der Erde-Geochemistry, v. 54, p. 129-149.

Morford, J. L., and Emerson, S., 1999, The geochemistry of redox sensitive trace metals in

sediments: Geochimica Et Cosmochimica Acta, v. 63, p. 1735-1750.

Morris, R. C., 1993, Genetic modeling for banded iron-formation of the Hamersley Group,

Pilbara Craton, Western-Australia: Precambrian Research, v. 60, p. 243-286.

Morris, R. C., and Horwitz, R. C., 1983, The origin of the iron-formation-rich Hamersley Group

of Western Australia — deposition on a platform: Precambrian Research, v. 21, p. 273-

297.

Newman, M. J., and Rood, R. T., 1977, Implications of Solar Evolution for the Earth's Early

Atmosphere.: Science, v. 198, p. 1035-1037.

Ohmoto, H., Watanabe, Y., and Kumazawa, K., 2004, Evidence from massive siderite beds for a

CO2-rich atmosphere before, 1.8 billion years ago: Nature, v. 429, p. 395-399.

Parman, S. W., 2007, Helium isotopic evidence for episodic mantle melting and crustal growth:

Nature, v. 446, p. 900-903.

Partridge, M. A., Golding, S. D., Baublys, K. A., and Young, E., 2008, Pyrite paragenesis and

multiple sulfur isotope distribution in late Archean and early Paleoproterozoic Hamersley

Basin sediments: Earth and Planetary Science Letters, v. 272, p. 41-49.

Pearson, D. G., Parman, S. W., and Nowell, G. M., 2007, A link between large mantle melting

events and continent growth seen in osmium isotopes: Nature, v. 449, p. 202-205.

Perry, E. C., Tan, F. C., and Morey, G. B., 1973, Geology and Stable Isotope Geochemistry of

the Biwabik Iron Formation, Northern Minnesota: Economic Geology, v. 68, p. 1110-

1125.

Piper, D. Z., 1974, Rare earth elements in the sedimentary cycle: A summary: Chemical

Geology, v. 14, p. 285-304.

Planavsky, N., Rouxel, O., Bekker, A., Shapiro, R., Fralick, P., and Knudsen, A., 2009, Iron-

oxidizing microbial ecosystems thrived in late Paleoproterozoic redox-stratified oceans:

Earth and Planetary Science Letters, v. 286, p. 230-242.

Polteau, S., Moore, J. M., and Tsikos, H., 2006, The geology and geochemistry of the

Palaeoproterozoic Makganyene diamictite: Precambrian Research, v. 148, p. 257-274.

Poulton, S. W., and Canfield, D. E., 2011, Ferruginous Conditions: A Dominant Feature of the

Ocean through Earth's History: Elements, v. 7, p. 107-112.

Poulton, S. W., Krom, M. D., and Raiswell, R., 2004, A revised scheme for the reactivity of iron

(oxyhydr)oxide minerals towards dissolved sulfide: Geochimica Et Cosmochimica Acta,

v. 68, p. 3703-3715.

Rasmussen, B., Fletcher, I. R., Bekker, A., Muhling, J. R., Gregory, C. J., and Thorne, A. M.,

2012, Deposition of 1.88-billion-year-old iron formations as a consequence of rapid

161

crustal growth: Nature, v. 484, p. 498-501.

Reinhard, C. T., Raiswell, R., Scott, C., Anbar, A. D., and Lyons, T. W., 2009, A Late Archean

Sulfidic Sea Stimulated by Early Oxidative Weathering of the Continents: Science, v.

326, p. 713-716.

Robert, F., and Chaussidon, M., 2006, A palaeotemperature curve for the Precambrian oceans

based on silicon isotopes in cherts: Nature, v. 443, p. 969-972.

Rye, R., and Holland, H. D., 1998, Paleosols and the evolution of atmospheric oxygen: A critical

review: American Journal of Science, v. 298, p. 621-672.

Rye, R., Kuo, P. H., and Holland, H. D., 1995, Atmospheric carbon dioxide concentrations

before 2.2 billion years ago: Nature, v. 378, p. 603-605.

Sage, R. P., and Heather, K. B., 1991, The Structure, Stratigraphy and Mineral Deposits of the

Wawa area, Geological Association of Canada, Mineralogical Association of Canada,

Society of Economic Geologists Field Trip A6 Guidebook, p. 118.

Sage, R. P., Lightfoot, P. C., and Doherty, W., 1996, Bimodal cyclical Archean basalts and

rhyolites from the Michipicoten (Wawa) greenstone belt, Ontario: Geochemical evidence

for magma contributions from the asthenospheric mantle and ancient continental

lithosphere near the southern margin of the Superior Province: Precambrian Research, v.

76, p. 119-153.

Schidlowski, M., Eichmann, R., and Junge, C. E., 1975, Precambrian sedimentary carbonates:

carbon and oxygen isotope geochemistry and implications for the terrestrial oxygen

budget: Precambrian Research, v. 2, p. 1-69.

Schmidt, K., Koschinsky, A., Garbe-Schonberg, D., de Carvalho, L. M., and Seifert, R., 2007,

Geochemistry of hydrothermal fluids from the ultramafic-hosted Logatchev hydrothermal

field, 15 degrees N on the Mid-Atlantic Ridge: Temporal and spatial investigation:

Chemical Geology, v. 242, p. 1-21.

Shields, G., and Veizer, J., 2002, Precambrian marine carbonate isotope database: Version 1.1:

Geochemistry Geophysics Geosystems, v. 3.

Sholkovitz, E. R., Landing, W. M., and Lewis, B. L., 1994, Ocean particle chemistry: The

fractionation of rare earth elements between suspended particles and seawater:


Sial, A. N., Ferreira, V. P., Dealmeida, A. R., Romano, A. W., Parente, C. V., Dacosta, M. L.,

and Santos, V. H., 2000, Carbon isotope fluctuations in Precambrian carbonate sequences

of several localities in Brazil: Anais Da Academia Brasileira De Ciencias, v. 72, p. 539-

558.

Siever, R., 1992, The silica cycle in the Precambrian: Geochimica Et Cosmochimica Acta, v. 56,

p. 3265-3272.

Simonson, B. M., and Hassler, S. W., 1996, Was the deposition of large precambrian iron

formations linked to major marine transgressions?: Journal of Geology, v. 104, p. 665-

676.

Slack, J. F., and Cannon, W. F., 2009, Extraterrestrial demise of banded iron formations 1.85

billion years ago: Geology, v. 37, p. 1011-1014.

Slack, J. F., Grenne, T., Bekker, A., Rouxel, O. J., and Lindberg, P. A., 2007, Suboxic deep

seawater in the late Paleoproterozoic: Evidence from hematitic chert and iron formation

related to seafloor-hydrothermal sulfide deposits, central Arizona, USA: Earth and

Planetary Science Letters, v. 255, p. 243-256.

Sleep, N. H., and Rosendahl, B. R., 1979, Topography and tectonics of mid-ocean ridge axes:

162

Journal of Geophysical Research, v. 84, p. 6831-6839.

Speer, K. G., and Rona, P. A., 1989, A model of an Atlantic and Pacific hydrothermal plume:

Journal of Geophysical Research-Oceans, v. 94, p. 6213-6220.

Steinhoefel, G., Horn, I., and von Blanckenburg, F., 2009, Micro-scale tracing of Fe and Si

isotope signatures in banded iron formation using femtosecond laser ablation:


Stichler, W., 1995, Interlaboratory comparison of new materials for carbon and oxygen isotope

ratio measurement. Reference and intercomparison materials for stable isotopes of light

elements. International Atomic Energy Agency, IAEA-TECDOC-825 (Vienna, Austria),

67–74.

Stoll, B., Jochum, K. P., Herwig, K., Amini, M., Flanz, M., Kreuzburg, B., Kuzmin, D.,

Willbold, M., and Enzweiler, J., 2008, An automated iridium-strip heater for LA-ICP-MS

bulk analysis of geological samples: Geostandards and Geoanalytical Research, v. 32, p.

5-26.

Sugitani, K., Yamamoto, K., Adachi, M., Kawabe, I., and Sugisaki, R., 1998, Archean cherts

derived from chemical, biogenic and clastic sedimentation in a shallow restricted basin:

examples from the Gorge Creek Group in the Pilbara Block: Sedimentology, v. 45, p.

1045-1062.

Sverjensky, D. A., 1984, Europium redox equilibria in aqueous solution: Earth and Planetary

Science Letters, v. 67, p. 70-78.

Sylvester, P. J., 2001, Trace element analysis of fused whole rock glasses by laser ablation

ICPMS. In Laser Ablation ICP–MS in the Earth Sciences: Principles and Applications

(P.J. Sylvester, ed.). Mineral. Assoc. Can., Short Course 29, 147–162.

Sylvester, P. J., Attoh, K., and Schulz, K. J., 1987, Tectonic setting of late Archean bimodal

volcanism in the Michipicoten (Wawa) greenstone belt, Ontario: Canadian Journal of

Earth Sciences, v. 24, p. 1120-1134.

Taylor, S. R., 1982, Planetary Science: A Lunar Perspective. Lunar and Planetary Institute,

Houston. pp. 481.

Taylor, S. R., and McLennan, S. M., 1985, The continental crust: its composition and evolution.,

Blackwell,Oxford., 312 p.

Thode, H. G., and Goodwin, A. M., 1983, Further Sulfur and Carbon Isotope Studies of Late

Archean Iron-Formations of the Canadian Shield and the Rise of Sulfate Reducing

Bacteria: Precambrian Research, v. 20, p. 337-356.

Thurston, P. C., Ayer, J. A., Goutier, J., and Hamilton, M. A., 2008, Depositional Gaps in Abitibi

Greenstone Belt Stratigraphy: A Key to Exploration for Syngenetic Mineralization:

Economic Geology, v. 103, p. 1097-1134.

Thurston, P. C., Kamber, B. S., and Whitehouse, M., 2012, Archean cherts in banded iron

formation: Insight into Neoarchean ocean chemistry and depositional processes:

Precambrian Research, v. 214, p. 227-257.

Tice, M. M., and Lowe, D. R., 2004, Photosynthetic microbial mats in the 3,416-Myr-old ocean:

Nature, v. 431, p. 549-552.

Trendall, A. F., and Blockley, J. G., 1970, The iron formations of the Precambrian Hamersley

Group, Western Australia with special reference to the associated corcidolite.: Western

Australia Gelogical Survey Bulletin, v. 119.

Trendall, A. F., and Morris, R. C., 1983, Iron-formation, facts and problems: Amsterdam; New

York, Elsevier, 558 p.

163

Tsikos, H., Beukes, N. J., Moore, J. M., and Harris, C., 2003, Deposition, diagenesis, and

secondary enrichment of metals in the Paleoproterozoic Hotazel iron Formation, Kalahari

manganese field, South Africa: Economic Geology and the Bulletin of the Society of

Economic Geologists, v. 98, p. 1449-1462.

Tuck, C. A., and Virta, R. L., 2013, 2011 Minerals Yearbook - Iron Ore (advanced release), in

Survey, U. S. G., ed.

Turek, A., Smith, P. E., and Vanschmus, W. R., 1982, Rb–Sr and U–Pb ages of volcanism and

granite emplacement in the Michipicoten belt—Wawa, Ontario: Canadian Journal of

Earth Sciences, v. 19, p. 1608-1626.

Urey, H. C., 1947, The thermodynamic properties of isotopic substances: Journal of the

Chemical Society, p. 562-581.

van den Boorn, S., van Bergen, M. J., Nijman, W., and Vroon, P. Z., 2007, Dual role of seawater

and hydrothermal fluids in Early Archean chert formation: Evidence from silicon

isotopes: Geology, v. 35, p. 939-942.

Veizer, J., Hoefs, J., Lowe, D. R., and Thurston, P. C., 1989, Geochemistry of Precambrian

carbonates. 2. Archean Greenstone Belts and Archean sea-water: Geochimica Et


Walter, M. R., 1976, Stromatolites. Developments in Sedimentology, vol 20, 790p.

Wardle, R. J., and Bailey, D. G., 1981, Early Proterozoic Sequences in Labrador. In: Campbell,

F.H.A. (Ed.), Proterozoic Basins of Canada. Geological Survey of Canada Paper 81–10.

pp. 331–359.

Wardle, R. J., Ryan, B., Nunn, G. A. G., and Mengel, F. C., 1990, Labrador segment of the

Trans-Hudson Orogen: crustal development through oblique convergence and collision.

In The Early Proterozoic Trans-Hudson Orogen of North America.

Williams, G. E., and Schmidt, P. W., 2004, Paleomagnetism of the 1.88-Ga Sokoman Formation

in the Schefferville-Knob Lake area, Quebec, Canada, and implications for the genesis of

iron oxide deposits in the central New Quebec Orogen: Precambrian Research, v. 128, p.

167-188.

Williams, H. R., Stott, G. M., Heather, K. B., Muir, T. L., and Sage, R. P., 1991, Wawa

Subprovince, in: Thurston , P.C., Williams, H., Sutcliffe, H.R., Stott, G. (Eds.), Geology

of Ontario. Ontario Geological Survey, pp. 485-539.

Winter, B. L., and Knauth, L. P., 1992, Stable isotope geochemistry of cherts and carbonates

from the 2.0 Ga Gunflint iron formation: implications for the depositional setting, and

the effects of diagenesis and metamorphism: Precambrian Research, v. 59, p. 283-313.

Zhang, C. L., Horita, J., Cole, D. R., Zhou, J. Z., Lovley, D. R., and Phelps, T. J., 2001,

Temperature-dependent oxygen and carbon isotope fractionations of biogenic siderite:


Zheng, Y., Anderson, R. F., van Geen, A., and Kuwabara, J., 2000, Authigenic molybdenum

formation in marine sediments: A link to pore water sulfide in the Santa Barbara Basin:


Date post:	25-May-2020
Category:	Documents
Upload:	others
View:	22 times
Download:	0 times

Comparison between the Helen Iron Formation (Algoma-type ......Comparison between the Helen Iron...

Documents