The Sedimentology and Geochemistry of Banded Iron Formations of the Deloro Assemblage, Bartlett Dome area, Abitibi greenstone belt, Ontario, Canada: Implications for BIF deposition and greenstone belt formation
By Geoffrey J. Baldwin
Thesis presented as a partial requirement in the Master of Science (M.Sc.) in Geology
School of Graduate Studies Laurentian University Sudbury, Ontario
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Abstract Banded iron formations (BIFs) are a small but important component of the
Precambrian rock record. Consisting of alternating bands of iron-rich minerals and chert, these sedimentary rocks can be used to learn a great deal about Earth processes at their time of deposition. The bulk of research on BIFs to date has focused on the enigmatic iron component, despite the great insight to be derived about their depositional environment from detailed study of the chert bands. Although cherts are commonly formed through direct precipitation of amorphous silica from seawater or through fluid movement through unlithified sediments, their trace element geochemistry can prove very sensitive to small contributions of other materials, allowing for the relatively easy detection of different precursor rock types in cherts.
In Archean greenstone belts, such as the Abitibi greenstone belt, BIFs represent a stratigraphic departure from the volcanic rock-dominated stratigraphy of these belts. BIFs up to 50 m thick within the Deloro assemblage, Bartlett Dome region, of the Abitibi greenstone belt, are bounded above and below by volcanic rocks which display U-Pb ages with differences of up to 12 My, with the absence of any observable erosional surfaces. This suggests near-continuous deposition, albeit at very slow rates. Here I use a combination of a detailed mapping of sedimentary structures and textures, as well as high-precision trace element geochemistry to identify unique features in cherts of the
BIFs of the Abitibi greenstone belt. Many of these features are indicative of slow to null depositional rates.
Some of the primary structures observed include chert hardgrounds and pre- lithification nodules, loading structures, stratabound chert breccias, and coarse-grained, iii heterolithic debris flows, suggesting a combination of rapid, episodic deposition and periods of null deposition and violent fluid escape. The trace element geochemistry has revealed that chert formation was dominated by one of several processes; precipitation of amorphous silica from seawater, sub-seafloor hydrothermal circulation, and the input of subsequently silicified detrital materials. These groups were identified using Ni/Cr ratios to determine the degree of detrital vs. orthochemical (seawater or hydrothermal) contribution, and further refined within these chemical groups based on the MUQ- normalized REE+Y patterns, which have aided in identifying more specific fluid and detrital sources in these cherts.
Based on the sedimentology and geochemistry of the cherts, a new depositional model for the BIFs of the Bartlett Dome is here suggested. Controlled by 4 main depositional and diagenetic processes, this model consists of: periods of null-deposition resulting in the formation of chert hardgrounds; the rapid deposition of amorphous silica and ultra-fine detrital volcanic material from seawater, forming load structures; the instantaneous deposition of coarse grained heterolithic debris flows; and sub-seafloor hydrothermal circulation, resulting in the replacement of other materials by silica and the brecciation of buried hardgrounds through fluid escape. The interplay of these 4 processes within the stratigraphic space of 50m strongly suggests that the principal control on BIF formation in the Abitibi greenstone belt was the prolonged period of deposition with highly limited material input, allowing for numerous processes to exert influence on deposition. The model presented here may be applicable in other greenstone belts worldwide, pending studies of the BIFs in those belts through this lens.
iv Acknowledgements I would like to thank my supervisors, Dr. Phil Thurston and Dr. Balz Kamber for the unwavering patience with me throughout my thesis, as well as their invaluable support, guidance, and advice.
I would also like to thank Dr. Michael Schindler for agreeing to be my third committee member and to edit my thesis as well as Dr. Boswell Wing of McGill
University for being my external examiner, and providing excellent suggestions towards the improvement of this thesis.
I also thank Dr. John Ayer and Dr. Michel Houle of the Ontario Geological
Survey for their help and advice on my field studies and the regional geology of the
Abitibi greenstone belt.
Thanks to Brian Atkinson and his staff at the Timmins Resident Geologist office for their vital help on various issues and problems I encountered in the field.
To Hannah Burke, without your patience with my mapping style in the field and your help as my field assistant in summer 2007, this thesis may never have happened.
My fellow graduate students, particularly Kirk Ross, Valentina Taranovic, and
Mike Babechuk, I thank you all for putting up with my whining throughout my MSc, as well as being excellent sounding boards for my problems and ideas.
Thanks to Dr. Thomas Ulrich for his tireless work for the five days of laboratory time on the LA-ICP-MS at the Chemical Fingerprinting Lab.
v I would also like to thank Willard Desjardins, Annette Gladu, and Claire Kamber for their work in preparing my samples for petrographic and geochemical analysis.
Lastly, I would like to thank my parents, Jim and Denise, and my brothers Drew and Tristan for their love and support over the duration of my M.Sc. I owe everything I have achieved over the last two years to them.
vi Contents
Abstract iii
Acknowledgements v
Chapter 1: Introduction 1
Objectives of Study 1
Location and Access 2 Field and analytical methods 2 Structure of Thesis 3 Chapter 2: Using high-precision trace element geochemistry and sedimentology of cherts to create a depositional model for banded iron formations: Examples from the Deloro
Assemblage, Bartlett Dome area, Abitibi greenstone belt, Ontario, Canada 4
Abstract 4 Introduction 5 Banded Iron Formations 6 Abitibi Greenstone Belt, BIFs, and the study area context 7 Sedimentology and Stratigraphy 10 Middle BIF 11 Upper BIF 23 Petrography 25 Geochronology 31 Analytical methods 33 Geochemical Results 37 LA-ICP-MS Results 40 Solution ICP-MS Results 46 vii Discussion 68 Geochemical Discussion 68 Silica- sources and depositional processes 80 Characterization of the time gap 86 Depositional Model 89 Regional Implications 95 Conclusions 97 Chapter 3: Summary and Recommendations 101
Summary 101 Recommendations 105 References 107
Figures 1. Geological Map of the Abitibi greenstone belt 8 2. Geological Map of the Bartlett Dome 11 3. Field photographs of BIFs from the Bartlett Dome 14 4. Stratigraphy of the Middle BIF at the Bartlett North locality 17 5. Photomicrographs of cherts from the Bartlett Dome 26 6. Concordia diagrams of zircons dated using LA-ICP-MS 32 7. REE+Y patterns from a modern estuarine system 38 8. REE+Y patterns from samples analyzed using LA-ICP-MS 41 9. REE+Y patterns for duplicate samples analyzed by solution ICP-MS 54 10. REE+Y anomaly ratios from duplicate samples 55 11. Ni vs. Cr diagram after Condie (1993) 57 12. REE+Y patterns for cherts analyzed using solution ICP-MS 60 13. Crossplots of shale-normalized anomalies 71 14. La/La* vs. Th; Zr; Sc; and Cr 73 viii 15. Ce/Ce* vs. Pr/Pr* 75 16. Depositional Model for BIFs of the Bartlett Dome 93 Tables 1. Trace element concentrations from ICP-MS (parts per billion) 47 2. Relationship between chert classifications 67
ix Chapter 1: Introduction Objectives of Study
Funded through NSERC grants to Dr. Phillips C. Thurston, this project was originally conceived to assess the sedimentological and geochemical character of the chert component of Algoma-type banded iron formations (BIF) in the southwestern
Abitibi greenstone belt (AGB). The main purpose of this assessment was to test previous interpretations that these BIFs explain two problems with the chronostratigraphy of the
AGB; the disparity between the calculated volcanic volume and actual volume of volcanic rocks, and the seemingly conformable contact between units with ages of emplacement up to 12.5 My apart.
Previous work on this problem has confirmed that the one of the most puzzling factors is the apparently conformable stratigraphy in both the 2730-2724 Ma Deloro assemblage, and the overlying 2710-2704 Ma Tisdale assemblage. This has been observed in several areas throughout the western AGB, such as the flanks of the Round
Lake Batholith, Carscallen Township, the Shaw Dome, and the Bartlett Dome (Thurston et al., 2008). Out of these four regions, one of the best exposed and structurally simplest is the Bartlett Dome. The bedrock is better exposed than in much of the rest of the AGB, and unlike the nearby Shaw Dome, lacks significant structural complexity. For these reasons, this study has focused entirely on the BIFs of the Bartlett Dome, Which occur at three distinct stratigraphic levels, with the uppermost horizon marking the approximate boundary between the Deloro and Tisdale assemblages. Due to exposure constraints, the bulk of this study has focused on the laterally extensive Middle BIF, as well as a few exposures of the Upper BIF. 1 Location and Access
The Bartlett dome area of the AGB is located south of Timmins, Ontario, Canada, primarily in McArthur, Bartlett, and English Townships. This area is accessed by Pine
Street south from downtown Timmins. From highway 101 in Timmins, proceed south on
Pine street for 35 km to McArthur Lake Rd on the left to access the McArthur Powerline,
Central McArthur, and Southern McArthur localities. The side trail to the McArthur
Powerline locality is 6.1 km east on McArthur Lake Rd, and 1 km on a side trail (right) and under the high-tension power lines that run roughly north-south. An additional 3 km south on Pine street is the access road for the Bartlett North (0.78 km) and Texmont Mine
(7.77 km) localities. At 46.24 km on Pine street, the trail to the English Powerline site leaves on the left, with a 2.25 km combined drive and walk on the trail and under the power lines. Roads to the power lines and active logging operations provide excellent access to the study sites, and the clearing for the power lines allowed for near-complete exposure through the stratigraphy of many of the BIFs in this region, where otherwise only small portions of the stratigraphy may have been exposed.
Field and analytical methods
During the summer of 2007, a great deal of effort was put into the improvement of outcrop quality at all major study sites, including the removal of thin overburden and lichen cover to allow for highly detailed stratigraphic analysis of the BIF. Following outcrop improvement, each site was individually mapped at a 1:100 to 1:300 scale, and samples were collected of both distinctive and representative chert bands using a rock and masonry saw. Some sites were re-mapped in August 2008 in greater detail following
2 further outcrop improvement. Samples were thin sectioned at Laurentian University for petrographic and in situ geochemical analysis by LA-ICP-MS (analytical methods are discussed in a later section) to identify chemically distinct chert bands for high-precision total dissolution ICP-MS at the Chemical Fingerprinting Lab at Laurentian University.
Structure of Thesis
This thesis was written such that the background, results, and discussion of the detailed sedimentological and geochemical study of the BIFs of the Bartlett Dome are presented in Chapter 2, which is written in accordance with the submission guidelines to the journal Geobiology, including the citation format. A shortened version of this chapter will be submitted as a manuscript for publication in this journal. Chapter 1 is a very general introduction to the goals and broad methods used in this study, as well as an introduction to the location of study. Chapter 3 is a short summary of the conclusions reached in this study, and specific recommendations for further study on this topic.
Attached is a data CD containing, among other things, a complete sample list with sample location, UTM coordinates, and analytical procedures conducted on the samples. Also on the CD there is a table containing the trace element results obtained through LA-ICP-MS analysis.
3 Chapter 2: Using high-precision trace element geochemistry and sedimentology of cherts to create a depositional model for banded iron formations: Examples from the Deloro Assemblage, Bartlett Dome area, Abitibi greenstone belt, Ontario, Canada. Abstract
Banded iron formations (BIFs) are a small but important component of the
Precambrian rock record. Consisting of alternating bands of iron minerals and chert, these sedimentary rocks can be used to learn a great deal about Earth. In Archean greenstone belts, such as the Abitibi greenstone belt, BIFs represent a sedimentary interval in the otherwise largely volcanic stratigraphy. BIFs up to 50 m thick within the Deloro assemblage, Bartlett Dome region, Southwestern Abitibi greenstone belt, are bounded above and below by volcanic rocks whose U-Pb ages display up to 12 My difference, suggesting that the BIFs may represent longer gaps in volcanism than previously assumed. Here I use detailed mapping of the sedimentary structures and high-precision trace element geochemistry of cherts to show that the BIFs were deposited slowly, controlled by numerous processes. Sedimentological structures within the chert bands indicate brief periods of rapid deposition that alternated with extended periods of non- deposition and fluid interaction. The trace element geochemistry has revealed that chert formation was dominated by one of several processes including seawater precipitation, sub-seafloor hydrothermal circulation, and input of detrital materials, each distinguished through their Ni/Cr ratios and shale-normalized REE+Y diagrams. These results indicate that BIF formation was controlled by four asynchronous depositonal and diagenetic processes. The interplay of these four processes within such thin stratigraphy strongly
4 shows that the principal control on BIF formation in the Abitibi greenstone belt was the prolonged period of volcanic quiescence with highly limited material input. The model presented here may be applicable in other greenstone belts worldwide.
Introduction
The controls on the deposition of banded iron formations (BIF) have long been contentious. Consisting of fine alternations of chert and Fe-rich bands, much of the research on these rocks has focused on the iron-rich component, frequently approached via the overall stratigraphy (Trendall, 1983) and whole-rock geochemistry (Fryer, 1983).
These studies have generally assumed that the chert component of BIFs was precipitated directly from seawater that was saturated with respect to amorphous silica through the progression from silicic acid (FLjSiC^) to the precipitation of amorphous silica, and subsequent conversion to opal-A, opal-CT, and eventually, true chert (Siever, 1957). In the interest of brevity, this process will henceforth be referred to as the direct precipitation of amorphous silica, silica, or chert, but in each case this multi-step process is implied. Alternatively, the chert bands may have been derived from hydrothermal circulation and the silica replacement of volcanic and sedimentary rocks (Paris et al.,
1985; Duchac & Hanor, 1987; Krapez et al., 2003). Such interpretations have been proposed based on field (Krapez et al., 2003), petrographic (Sugitani et al, 1998), and geochemical studies (Sugitani et al., 2002; van den Boom et al., 2007). This paper presents ideas that have combined field, petrographic and geochemical study of BIFs of
5 the Deloro Assemblage, also identifying broader implications for iron formation and chert development in the Abitibi and elsewhere.
Banded Iron Formations
Past studies of the Fe-component of BIF's have yielded several classification schemes based on tectonic setting and mineralogy of the iron-rich bands. BIFs form two major types; Superior-type and Algoma-type iron formation. Superior-type iron formations are very large platform deposits, mostly formed during the NeoArchean and
Paleoproterozoic (2.7-1.8 Ga) (e.g. Hamersley, Lake Superior, Labrador Trough, Krivoy
Rog, Transvaal, Quadrilatero Ferrifero) and contain the largest iron-mining districts in the world (Clout & Simonson, 2005). Algoma-type iron formations are characterized by their stratigraphic association with volcanic rocks, and are most common in Archean greenstone belts (Goodwin, 1962), but can be found in younger volcanic arcs, such as the
Ordovician Bathurst Mining Camp in New Brunswick (Peter et al., 2003). Each of these two broad types of iron formation can be further classified by their iron-bearing mineralogy into four different facies; oxide facies (magnetite and hematite), carbonate facies (siderite, ankerite, ferroan dolomite), silicate facies (Fe-silicates such as greenalite, minnesotaite, grunerite, stilpnomelane, riebeckite, etc.), and sulfide facies (pyrite and pyrrhotite) (James, 1954). Studies of the Fe-rich bands have resulted in the widespread interpretation that the deposition of these iron minerals and their precursor oxy- hydroxides may have been microbially controlled (Cloud, 1973). This has been expanded in recent years to postulate that the oxidation of Fe2+ was driven by anaerobic ferrophototrophic bacteria, and that the alternations between iron and chert bands were largely controlled by variations in temperature that periodically restricted the biological 6 processes of the bacteria (Kappler et al, 2005; Posth et al, 2008). Other studies have used iron isotopes to assess the role of bacteria in BIF formation (Johnson et al, 2008).
Additionally, these biological processes have also been suggested as a mediator for the deposition of the silica-component of BIFs (Fischer & Knoll, 2009).
Abitibi Greenstone Belt, BIFs, and the study area context
In Canada, Algoma type BIFs have traditionally been viewed as an accessory feature of greenstone belts (GSB), with much of the research on GSBs focused on the volcanic rocks and their economic potential. However, the recognition of the stratigraphic importance of BIFs in GSBs has been developing gradually over the last half century.
Goodwin (1962) proposed that the presence of BIFs indicated that the Michipicoten GSB was formed over a continuous period, with BIF deposition occurring during periods of low volcanic activity. Consequently, BIF and chert layers in GSB stratigraphy have since commonly been used as markers to delineate stratigraphic units within GSBs, such as the
Barberton Greenstone Belt, South Africa and in the Pilbara Craton, Western Australia
(Lowe & Byerly, 1999; Van Kranendonk et al, 2007). This pattern has also been observed more recently in the Abitibi GSB (AGB). The use of traditional stratigraphic nomenclature has been replaced in this belt in favor of tectonostratigraphic nomenclature; each tectonostratigraphic assemblage is largely defined by stratigraphy and high resolution U-Pb zircon geochronology (Ayer et al, 2002). Consequently, the AGB has been divided into 8 tectonostratigraphic assemblages (roughly equivalent to group-level stratigraphic units), based on their ages (Figure 1).
7 Figure 1: Map of the Abitibi Greenstone Belt, Ontario (left) and Quebec (right) Canada, showing the different lithotectonic assemblages after Thurston et al. (2008, and references within). The inset shows the location of the AGB in Canada (black). The area of this study, the Bartlett Dome, is shown in the red square on the main map. These are the Pacaud (2750-2735 Ma), Deloro (2730-2724 Ma), Stoughton-Roquemaure
(2723-2720 Ma), Kidd-Munro (2719-2711 Ma), Tisdale (2710-2704 Ma), Blake River
(2704-2697 Ma), Porcupine (2690-2685 Ma), and Timiskaming (2676-2670 Ma) assemblages (Ayer et al., 2002). The first seven assemblages are dominantly volcanic- sedimentary units ("Kewatin-Strata") and the last is a post-orogenic rifting sequence
('Timiskaming-Strata") (Thurston & Ayres, 2004) (Figure 1). These tectonostratigraphic assemblages have been used to divide the western AGB, aiding in the correlation of units over long distances, despite structural complexity. The Deloro-Tisdale assemblage contact south of Timmins represents a 14 million year depositional gap which, elsewhere in the AGB, is occupied by the Stoughton-Roquemaure and Kidd- Munro assemblages
(Ayer et al., 2002). Further study of this contact has shown that the Deloro-Tisdale
8 assemblage contact is neither structural nor erosional, suggesting conformable deposition despite the disparate ages (Thurston, 2002; Thurston et al., 2008). Similar depositional gaps have been noted at chronostratigraphically equivalent contacts across the AGB, including the Swayze greenstone belt, about 100 km southwest of this study area (van
Breemen et al, 2006). This stratigraphic contact is distinguished in many localities by a layer of BIF, which could have formed during a period of prolonged volcanic quiescence, corresponding with the observed chronostratigraphic gap in volcanism (Heather, 2001;
Thurston et al., 2008).These BIF layers are very thin, typically on the order of 50 m or less, making them much thinner than one would otherwise expect for a unit marking a 12
My gap, even when the assumed slow depositional rate of BIFs is accounted for. It has been suggested that these gaps may be roughly equivalent to submarine unconformities as reviewed by Shanmugam (1988), with potential significance to contemporaneous VMS mineralization in other regions of the AGB (Thurston et al, 2008). Detailed study of the
BIFs of the Deloro Assemblage, particularly in the well-exposed Bartlett Dome area, 20-
40 km south of Timmins, Ontario, combined with new high-resolution U-Pb geochronology (Houle et al., 2008), and trace element geochemical analysis of the cherts from these BIFs is presented below, significantly expanding what is known about this gap in the volcanic stratigraphy of the western Abitibi Greenstone Belt.
Field and petrographic study of the Deloro assemblage BIFs in the Bartlett Dome region primarily focused on textures and structures in the chert bands. Particular attention was given to the identification of sediment load structures, hardgrounds, stratabound breccias, pre-compaction chert nodules, and coarse-grained heterolithic debris flows.
Petrography identified micro-textures associated with these larger-scale features and
9 those samples suitable for geochemical analysis. High-precision trace-element geochemistry was used with the intent of distinguishing genetic mechanisms for the cherts, such as direct seawater precipitation, hydrothermal silicification, and any minor detrital input. Closely spaced intercalation of cherts with differing trace-element signatures could be a strong indicator of prolonged depositional periods, helping to explain the nature of the periods of low volcanic activity during the formation of the
AGB.
Sedimentology and Stratigraphy
Due to the high-quality exposure through the Deloro and Tisdale assemblages and limited structural complexity, the Bartlett Dome region was selected to be the focus of this study (Ayer et al., 2002). Recent mapping by the Ontario Geological Survey had demonstrated the presence of 3 distinct units of banded iron formation (Houle, personal communication, 2007) (Figure 2). With no evidence for thrust-based duplication of stratigraphy (Ayer et al., 2002; Thurston et al., 2008), each of these BIF's represents successive deposition of BIF units within the 2734-2724 Ma Deloro assemblage, with the uppermost marking the upper contact of the Deloro Assemblage with the overlying
Tisdale assemblage. The lowermost BIF in this area is thin, not laterally discontinuous, and poorly exposed; thus it was excluded from this study. The middle BIF is the most laterally extensive and best exposed of the BIFs of the Bartlett Dome (Houle et al., 2008).
Outcrops for study were identified on recent OGS maps, and subsequently mapped in greater detail for this study.
10 Figure 2: Geological map of the Bartlett Dome area, Abitibi greenstone belt (after Houle et al., 2008). The 2 studied BIFs are shown in pink (Middle BIF) and Red (Upper BIF), with different study localities marked along each BIF (see legend and text for details).
Middle BIF
The middle BIF extends within McArthur, Bartlett, and English Townships ,
(Figure 2). The age range of this BIF has been very well constrained, with high precision
U-Pb zircon ages on overlying and underlying rhyolites, indicating a maximum age of
2728.1 +/- 1.6 Ma and a minimum age of 2724.5 +/- 1.5 Ma (Houle et al., 2008). This
BIF is -50-100 m thick, and is very well exposed along the clearing below a high-tension 11 power line that runs north-south through the region. In particular, it is very well exposed at the McArthur power line and the Bartlett North localities (Figure 2) with smaller scattered exposures at the English West, Central McArthur, and Southern McArthur locations. The BIF has been established as continuous using aeromagnetic data (Houle personal communication 2007) but demonstrates just a single bed-scale correlation, the result of pronounced lateral variability, generally uncommon in BIFs (Krapez et al.,
2003). The individual locations are described below.
McArthur Power line
The middle BIF is made up of ~50 m of iron formation; however, unexposed intervals and associated sedimentary and volcanic rocks yield a maximum thickness of up to -100 m. The BIF here has been divided into 7 units from the base upward (Houle et al., 2008): felsic volcanics, basal sulfide-facies BIF, silicate-facies BIF, oxide-facies BIF, debris flow and chert breccia-bearing BIF (Figure 3-A), and oxide-sulfide facies BIF.
More recent work has identified a 50 cm thick unit of chert breccia containing up to 30 cm rounded chert clasts, likely brecciated by dewatering of an underlying silicate layer, as the breccia unit contains dark green silicate matrix. Primary structures include flame structures, stratabound and podiform chert breccia beds, chert nodules, possible top-down silica replacement of other sediments, and some unusual magnetite textures. The flame structures occur mostly in chert beds whereas the folds are not pervasive, involving only a few mesobands of chert and Fe-minerals (Figure 3-B). Chert nodules up to 10 cm long occur, typically in the apex of syn-sedimentary folds (Figure 3-C). Top-down silica replacement, as described by Krapez et al. (2003), was seen on a cm-scale showing sharp upper contacts of chert bands with diffuse, transitional lower boundaries with underlying 12 Fe-bands (Figure 3-D). The same odd magnetite texture observed in these top-down silicified thin chert bands was also observed in thicker chert bands, not occurring at contacts with Fe-bands, suggesting more complete replacement of precursor sediments, with the exception of the coarsest-grained magnetite (Figure 3-E).
Bartlett North Power line
The BIF at the Bartlett North power line locality consists of two main exposures; however neither the upper nor the lower contact is exposed due to faulting and numerous mafic intrusions. Despite being only approximately 10 km along strike from the
McArthur Power line site, the character of the iron formation here differs, despite the presence of several broadly similar sedimentological features. Sites 07-GJB-048 and 07-
GJB-049 are approximately 50 m apart, and despite stratigraphic disturbance from the intrusion of mafic sills, the latter appears to be located up section.
07-GJB-048
This 15 m thick exposure is divided into 5 lithologies, based on mineralogy and/or primary structures. They are, from base to top: sulfide BIF, sulfide-oxide BIF, cherty
BIF, chert breccia, and banded sulfide BIF (Figure 4-A).
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Figure 3: Field photographs of BIFs from the Bartlett Dome. (A) Coarse-grained debris flow consisting of felsic volcanic and reworked sulfide clasts from the McArthur Powerline. Some clasts have been outlines in black for clarity. Camera Case is 12 cm. (B) Flame structures at the lower interface of a chert band with an underlying Fe-band from McArthur Power line, suggesting rapid sedimentation prior to dewatering. (C) Chert Nodule within the hinge of a fold from the McArthur Power line. It is unclear if the fold and nodule are syn-depositional or post-depositional. (D) Evidence for the top-down replacement of Fe-bands by chert from the McArthur Power line. The locally gradational contact (arrows) between the chert and Fe-band suggests uneven, top-down replacement that has left only the coarsest magnetite grains in the chert. (E) Similar magnetite texture to (D), except on a larger scale, once again suggesting the resistance of coarser- grained magnetite to silicification (also from the McArthur Power line). (F) Patchy mixture of white micas (WM) and chert-magnetite (C&M) from the cherry BIF unit at Bartlett North. (G) Oblate chert breccia with a magnetite-phyllosilicate matrix from Bartlett North. The hammer head is 18 cm. (H) Undulatory contact of the Oblate and tabular chert breccia at Bartlett North. The undulatory contact has been disturbed through 15 the brecciation of the oblate chert breccia, suggesting that these cherts were brecciated after they both were deposited, albeit in different dewatering events. Hammer is 38 cm. (I) Ball-and-pillow structure of magnetite-pyrite in black chert from the mixed facies of outcrop 07-GJB-049 at Bartlett North. Pencil magnet is 12.3 cm. (J) Black chert breccia from Bartlett North. Note the banding in the chert clasts (white arrow), and the pale green Fe-silicate matrix (black arrow). Hammer is 38 cm. (K) Chert nodules from the Southern McArthur site. Rotation is related to outcrop-scale folds, suggesting that the nodule formed prior to the folding event. Scale bar is in mm. (L) Tabular chert breccia in a black, graphitic matrix from the Texmont Mine. Cigarette lighter is 6 cm. (M) Uppermost BIF from the Texmont Mine. Notice the black carbonaceous chert bands and lighter Fe-silicate and carbonate bands. Pen is 17 cm. (N) Pebbly wacke from the English Power line. The pebbles are tabular chert clasts up to 15 cm in length, in a medium sand and chlorite matrix. Hammer is 38 cm.
Sulfide BIF
This unit forms the lowermost 2.8 m of the exposure, isolated from local stratigraphy by mafic sills. The bulk of it has weathered to gossan, rendering detailed mapping very difficult. Chert bands in this unit were distinguished from Fe-bearing bands through the positive physical relief of the chert that has resulted from the preferential supergene weathering of the Fe-bands. The iron bands are weakly magnetic, indicating moderate amounts of magnetite and/or pyrrhotite. Fresh surfaces show very bright white chert paired with bands dominated by sulfides and-phyllosilicates, typical of sulfide facies BIF (James, 1954). The uppermost 0.4 m of this unit is a clast-supported chert breccia, with chert clasts roughly 2-3 cm long and tabular, elongated along strike. This chert breccia contains fairly small amounts of matrix (<25%), which is similar in mineralogy to the iron bands beneath.
Sulfide-Oxide BIF
This 1.2 m thick unit is characterized by alternating cm-scale chert and iron- bearing bands. The iron-bearing bands are up to 1 cm thick and display a decreasing amount of sulfides and oxides (largely magnetite) up-section. Chert bands in this unit are variable in both color and thickness, with thicker chert bands (up to 10 cm) generally
16 appearing bluish-gray, whereas thinner bands (<1 cm) are commonly black. The black chert bands are non-magnetic and appear to be carbonaceous. These dark carbonaceous
cherts are commonly A Bartlett North Stratigraphy associated with thin
bands of very fine
B grained, off-white 8m
Legend micaceous material. Mixed Fades BIF 1 6m ,y .f>' J Disorganized ..;.>• J.,:* ..;." = llll = III Fe-silicate Figure 4: Stratigraphy of 5m | .if .if ..;*' and chert -r J1 J' the Middle BIF at the Oxide-rich BIF .ft' .if H m. Bartlett North Locality. 4m • .J' J" ., (A) Stratigraphy of :f ,# ,,r Chert Breccia *•-• outcrop site 07-GJB-048. 1, lllrl = .#>1111 = /lll l 1 • Cherty BIF (B) Stratigraphy of 4m gap in Hill outcrop site 07-GJB-049. exposure v\ 1 P § i i Sulfide-Oxide Site 07-GJB-049 is most due to mafic \ . i i VI BIF i i likely up-section from site intrusion \. JIT i • i Mn i i i Hi Sulfide BIF 07-GJB-048, as shown by Om .V i _ i the line tying the top of (A) with the bottom of (B), however the stratigraphic distance between the two is unknown due to 2m disturbance by mafic intrusions. See text for 07-GJB-048 07-GJB-049 details of separate units.
Cherty BIF
This unit is 2.8 m thick and is overwhelmingly dominated by chert. Low in this section the BIF is banded, demonstrating <1 cm to 5 cm oxide bands bearing small amounts of sulfidic material alternating with chert bands of similar thickness. The upper
2 m of this unit is much more chert-rich, containing 1-30 cm thick chert bands separated by thin (mm-scale) bands of coarse-grained magnetite. The upper meter of this unit contains a greater diversity of textures and minerals, as some of the chert bands contain
17 disseminated coarse grained magnetite in irregular patches up to 5 cm thick in irregular bands crossing the stratigraphy. Associated with these magnetite patches is commonly a white, very fine grained micaceous matrix that forms ball-and-pillow structures along the lower contact of these patchy magnetite bands with the more massive cherts below. The cherts in both of these sub-units typically range from white to bluish-white in color, and often have undulatory contacts with the magnetite-rich layers associated with them
(Figure 3-F).
Chert Breccia
The Cherty BIF is sharply overlain by a thick (5 cm) magnetite band with a disrupted upper contact followed up section by a pair of chert breccia units, totaling 2 m thickness. The 2 chert breccia units are in conformable stratigraphic contact and appear to be co-genetic despite great differences in their primary structures. The 10 cm thick lower chert breccia is a matrix supported unit consisting of oblate chert clasts up to 40 cm in length and 5 cm thick. These clasts are elongated parallel to bedding and are in a matrix of magnetite and white phyllosilicates (Figure 3-G). The immediately overlying tabular chert breccia is again a matrix-supported 1 m thick unit consisting of tabular chert clasts that are 1 cm thick and range from 1-15 cm long, elongated parallel to bedding. In contrast to the lower unit, the matrix of this tabular chert breccia consists of sulfides and dark-colored phyllosilicates. The sharp, albeit undulatory, nature of the contact between these two breccias is suggests that they were likely formed by similar processes. Due to their stratabound nature, such as process would have been the brecciation by fluid overpressure in lower stratigraphy of silica firmgrounds (oblate chert breccia) and hardgrounds (tabular chert breccia) (Figure: 3-H). 18 Banded Sulfide BIF
The top unit at this exposure is a 2 m thick banded sulfide and chert unit. It differs from a conventional sulfide-facies BIF (e.g. James 1954) as the iron-bearing bands are composed entirely of sulfides, predominantly pyrite and pyrrhotite, resulting in the low magnetism of the Fe-bands. The chert bands in this unit are bright white when a clean surface is examined, and both the chert and sulfide layers are roughly 1 cm thick.
Portions of this unit are intensely folded, but can be traced along strike to undisturbed bands, although the geometry of these folds and similar axial trends across the region suggest the structures are tectonic. The upper portion of this unit is terminated by a mafic intrusion.
07-GJB-049
This upper exposure at the Bartlett North Power line is 8 m thick and lies an indeterminate distance up section from the lower exposure (048) described above.
Sedimentologically, it is quite different from the lower unit, containing a much higher proportion of iron bands than its lower counterpart, which is largely chert-dominated. It has been divided into 5 different lithologies which from base to top are: oxide-sulfide
BIF, oxide-rich BIF, disorganized Fe-silicate and chert, mixed facies BIF, and a chert breccia unit (Figure 4-B).
Sulfide-Oxide BIF
This 2 m thick unit consists of thinly banded (cm-scale) magnetite, sulfide, and chert-rich layers. The two different BIF facies alternate in roughly 25 cm band sets consisting of chert and facies-dependant iron-bearing mesobands bands up 1 cm thick. 19 The oxide bands appear to be almost pure magnetite, whereas the sulfide bands resemble classical sulfide-facies BIF: alternating chert and sulfidic slate (James, 1954)
Oxide-rich BIF
This unit is only 0.5 m in thickness, but is unusual given its high magnetite content. This finely banded magnetite-chert unit consists of black magnetite bands up to 1 cm thick, and dark gray chert bands ranging from 0.3 to 3 cm thick. Both the chert and magnetite bands are magnetic, indicating that there are significant amounts Fe -bearing oxides present in the chert bands, most likely magnetite, and likely is the source of their dark color.
Disorganized Fe-silicate and chert
This 0.5 m unit is somewhat unusual. It is similar in mineralogy to the rest of the
BIF at this exposure. It consists of irregular patches with ill-defined margins consisting of a pale green, soft Fe-silicate, and scattered oxides and sulfides and small clasts of chert.
The lack of well defined fragment margins and the gradational transitions in iron-bearing mineralogy suggest this unit may have been formed by complete down-slope or across stratigraphy movement of unlithified silicate-oxide-sulfide facies iron formation.
Mixed Facies BIF
This 4 m unit consists of chert and magnetite bands, with minor bands of Fe- silicates, sulfides, and thin ash beds. The lower 1.5 m of this unit is largely chert intercalated with magnetite bands, and low in the section the silicate and ash units are more common than farther up-section. Gradually the silicate facies material and ash beds
20 diminish in favor of sulfide bands, which reach parity with the magnetite bands in the upper 1.25 m of the unit. The lower portions of this unit are dominantly flat-banded, with some weakly undulatory bands, but the upper unit contains several primary structures, such as ball and pillow structures between adjacent magnetite and sulfide bands (Figure
3-1).
Chert Breccia
This unit is about 1 m thick and consists of up to 40 cm long chert clasts in a pale- green Fe-silicate matrix. The clasts are rounded, and contain fine black, non-magnetic banding that is angled roughly 45° from the bedding trend of the BIF, indicating that the clasts have been rotated about the same from horizontal (Figure 3-J). These black bands are non-magnetic and appear to be graphitic. The Fe-silicate matrix appears as a continuous band immediately below the breccia, and intrudes between the clasts from below, indicating that the silicate band likely records the fluid escape process that caused this brecciation.
Central McArthur
This locality is the northernmost exposure of the middle BIF studied, however the exposure displays extensive gossan, thus detailed description is impossible. Roughly half the exposed thickness is planar laminated chert and sulfide facies bands up 1 cm thick.
The sulfides are dominantly pyrite, but weak magnetism suggests minor pyrrhotite as well. The rest of the exposure is chert breccia consisting of 0.5 cm by 2-5 cm long chert clasts in a matrix of sulfides and very fine grained silicates. The breccia forms stratabound beds much like those seen elsewhere in the region. Although the chert
21 breccia, like the bedded BIF, is heavily weathered to gossan, clean surfaces show that the chert is very bright white, and appears to lack a significant non-quartz component.
Southern McArthur
The Southern McArthur locality is another small (3 m) exposure of the Middle
BIF, consisting of an oxide-silicate-sulfide facies BIF of uncertain stratigraphic position within the Middle IF. The different BIF mineralogies form discrete intercalated 1-2 cm bands alternating with chert bands up to 10 cm thick. The thickest of the chert bands contain 1-2 mm magnetite grains in irregular, patchy aggregates. The thickest magnetite band (2 cm, but closely associated with a set of 2-3 mm microbands of chert and magnetite) contains a 2 cm podiform chert nodule, which has the surrounding bedding conform around it, indicating pre-compaction formation of the nodule (Figure 3-K). The upper meter of this exposure is dominated by a band of large, rounded chert clasts in a dark green Fe-silicate matrix. These clasts are up to 30 cm in length and 10 cm thick, and have been rotated roughly 45° with respect to bedding. The chert in these clasts is grayish-white and displays 1-2 mm laminae of darker gray, non-magnetic material, likely fine graphite bands. This upper unit is likely the only individual unit that can be correlated for a significant distance along the middle BIF, from this location to a comparable unit at the McArthur Power line to the north and the Bartlett North Power line to the south.
English West
22 The English West locality is the southernmost exposure of the Middle BIF studied. Like most of the other studied areas, this locality consists of a series of three small outcrops and is a very incomplete exposure.
Outcrop 1
This 2 m thick outcrop consists entirely of oxide facies iron formation, dominantly magnetite. It is a well banded BIF with both iron and chert bands ranging from 1-3 cm thick. Some bed-scale folding is present in this exposure. This is the only site with jasper bands that was mapped as a part of this study, with 80% of the chert bands being jasper, while the remaining 20% are white cherts.
Outcrops 2 & 3
These two exposures are very similar in nature; both are dominated by medium- grained, buff felsic ash layers up to 15 cm thick in their lower strata, which transitions upward into a more traditional oxide-facies BIF. Within the BIF layers, the cherts are unlike outcrop 1 and lack jasper, and are indeed more similar in character to the immediately underlying felsic ash beds, being gray-brown in color and containing some non-silica material. Also, within the BIF layers are discrete layers of black argillite in the place of some of the Fe-bands. All of the chert, magnetite, and argillite bands range from
1-3 cm, and are relatively planar, save some small undulations in their contacts.
Upper BIF
The Upper BIF of the Bartlett Dome is much less laterally extensive and more poorly exposed than the Middle BIF. This BIF is of stratigraphic significance, as it forms
23 the uppermost Deloro Assemblage unit (2730-2724 Ma) being in direct contact with the
Tisdale Assemblage (2710-2704 Ma). Whereas the Middle BIF may represent as much as a 0.5-6.7 My time interval based on constraining U-Pb zircon ages (see geochronology section), the Upper BIF may have been deposited over a period of up to 12.6 My (Houle et al, 2008). The bulk of the Upper BIF is along the apex of the Bartlett Dome, (Figure
2). Furthermore, the Upper BIF is generally much thinner than the Middle BIF, appearing to be no more than 30 m thick. Three sites within the Upper BIF were mapped: the
Texmont mine, Bartlett Northeast, and the English power line.
Texmont Mine
The Upper BIF at the Texmont mine site is in fact 2 thin horizons separated by a felsic agglomerate that has a U-Pb age of 2722.6 +/-0.9 Ma (Houle et al., 2008). The lower horizon at the Texmont mine consists dominantly of chert breccia (Figure 3-L), with some layers of bedded oxide facies BIF, whereas the upper horizon consists of bedded sulfide, silicate, and carbonate facies units. The upper horizon has black, non magnetic cherts, indicating a graphitic component in the chert, a common feature in many
Archean cherts (e.g. Sugitani et al., 2002), and making the chert bands darker in color than the corresponding buff-colored iron bands, which are dominantly iron carbonates and silicates (Figure 3-M). A more detailed description is available in Houle et al. (2008).
The iron bands in the uppermost part of this unit show hornfels textures typical of contact metamorphism, likely because this unit is immediately below Tisdale assemblage komatiite flows. The komatiite flows at the Texmont Mine host a Mt. Keith-type Ni-Cu-
(PGE) deposit, and the presence of the BIF in the hanging wall at this site may be important to the formation of the deposit as a potential sulfur source (Houle et al., 2008). 24 Bartlett Northeast
The Bartlett Northeast locality is a thin exposure of the Upper BIF to the north of the Texmont Mine. This 3.4m thick exposure consists of cm-scaled chert bands mixed with volcanic ash as a ratio of about 50% ash to 30% chert, with the remaining 20% of the exposure being mixed carbonate-silicate facies BIF, with no evidence for magnetite or sulfides present.
English Power line
This locality is the southernmost occurrence of the Upper BIF that was studied, and has been described in detail in recent literature (Houle et al., 2008; Thurston et al.,
2008). The important features to note from this locality are that it contains multiple heterolithic debris flows, including some that contain chert clasts (Figure 3-N), as well as
3 distinct generations of chert breccia, 2 of which are stratabound within a bedded BIF.
Petrography
Fifty Cherts from Deloro Assemblage BIFs were examined petrographically.
Small scale stratigraphic variations prohibit classification of the cherts by stratigraphic position or sample locality. Consequently, the analyzed cherts are divided into 7 mineralogical and texturally defined groups: high-grunerite cherts, veined cherts, ferruginous cherts, carbonaceous cherts, silicified or altered fragments, chert breccias, and clean cherts.
25 aw
*MM
500 urn
400 urn 250 um
Figure 5: Photomicrographs of cherts from the Bartlett Dome. (A) High-grunerite chert with high percentage of acicular grunerite crystals. (B) An example of a grunerite-silicate facies Fe-band, with much coarser grunerite than the high grunerite cherts. (C) Veined Chert wit coarse-grained chert veins between clasts made of finer-grained chert. (D) Ferruginous chert, type Jaspilite. Notice red tint of opaque minerals (hematite). (E) Carbonaceous chert. Notice the blotchy texture of the opaque material, crudely arranged in bands. (F) Silica-replaced fragment of uncertain origin. The original material appears to be the tabular form in the center of the fragment.
26 Figure 5 cont: (G) Completely replaced plagioclase crystal with ghost albite-twinning faintly visible from the silicified or altered fragment group. (H) Rimmed chert breccia. The core of the clast is towards the left of the image, with an increasing percentage of non-quartz silicate towards the rim at the far right. (I) Sulfide-dominated chert breccia sample, with opaque sulfide-grains in a brown phyllosilicate-amphibole matrix surrounding a relatively pure chert clast. (J) Clean chert breccia with very fine chloride matrix between separate, pure chert clasts. (K) Polygonal clean-chert, with little other material present in the chert band. (L) Microcrystalline clean chert. Notice the main distinction from (K) is the poor-definition of the quartz grains.
27 High-grunerite cherts
The high-grunerite cherts consist dominantly of polygonal quartz grains (<0.5 mm) with up to 30% acicular grunerite up 0.5 mm in length within the quartz bands
(Figure 5-A). These cherts are different from BIF identified as silicate facies based the criteria of James (1954), containing coarser-grained grunerite (up to 2 mm) in discrete
Fe-mineral bands (Figure 5-B), whereas this petrographic group refers to cherts containing grunerite within the chert bands. In fact, there is no consistent relationship between the high-grunerite cherts and the silicate facies BIF. Accessory minerals in these cherts include variable amounts of phyllosilicates (e.g. sericite, chlorite, or stilpnomelane), carbonates (mainly calcite), and small amounts of disseminated magnetite and pyrite, typically reflecting the dominant composition of the associated Fe- bands.
Veined Cherts
Veined cherts are composed of very fine grained (<0.1mm) polygonal quartz with minor accessory phyllosilicates and grunerite. This fine grained quartz is formed of angular quartz fragments (up to 5mm) cut by a network of 0.5mm quartz veins containing coarser (0.1-0.2mm) polygonal quartz (Figure 5-C). These veins are generally devoid of other minerals, lacking the silicate, carbonate, oxide, and sulfide minerals commonly observed in many of the cherts from the Deloro assemblage.
Ferruginous cherts
28 While many of the cherts studied contain trace amounts of interstitial Fe-minerals, some contain greater than 10% and have been classified as ferruginous cherts. Each of these chert bands consists dominantly of <0.1 mm-0.5 mm polygonal quartz crystals, but with up to 20% <0.1 mm magnetite grains. This group also includes a single jaspilitic sample, which in lieu of magnetite, bears up to 20% hematite in both interstitial and intragranular space (Figure 5-D).
Carbonaceous cherts
Medium grey to black carbonaceous cherts contain 0.2 to 1 mm laminae of graphite within the chert bands, alternating with up to 2 mm laminae of clean chert. The cherts consist of up to 0.2 mm polygonal quartz grains, with irregular laminae of framboidal to wispy graphite containing grains <0.05 mm. In the samples with the coarsest, most well-defined laminae, the graphite is visible in hand specimen as dark gray bands on a pale gray chert (Figure 5-E). Generally, the carbonaceous cherts contain relatively few accessory minerals within the chert bands; however some samples do contain small amounts of sericite between polygonal quartz grains.
Silicified or altered fragments
The bulk of the samples containing silicified fragments fit petrographically into other groups, except for unusual, small-scale silicified remnants of fragments. Some have been replaced by radiating quartz grains, others representing heavily altered plagioclase grains. The radiating quartz aggregates nucleate from preexisting particles and are generally much coarser than the chert in which they are found, with an overall diameter up to 0.5 mm (Figure 5-F). The relict plagioclase grains have been completely altered to
29 sericite and other alteration products that demonstrate ghosts of albite twinning from the precursor plagioclase (Figure 5-G).
Chert breccia
As described in the previous section, chert breccias are a significant component of the BIFs of the Deloro assemblage. Petrographically, these chert breccias are subdivided into sericite-rimmed quartz clasts and clean quartz clasts. The sericite-rimmed chert breccia clasts are elongate and taper at the ends. The cores of these clasts are usually almost entirely fine-grained polygonal quartz (0.1-0.2 mm), but the clast margins are typified by considerable increases in sericite with clast rims composed almost entirely of sericite (Figure 5-H). There is a sharp transition from clast to matrix, which is a highly heterogeneous mixture of phyllosilicates, Fe-oxides, and sulfides. The clean quartz clasts consist of elongate to tabular quartz clasts composed of fine-grained (0.1 mm), polygonal quartz grains with little or no contamination from other silicates or Fe-minerals. The matrix between these chert clasts is dominantly phyllosilicates and Fe-oxides/sulfides
(Figure 5-1). The most clast-supported samples show mm-scale chloritic or graphitic matrix, which also may contain some euhedral pyrite grains (Figure 5-J).
Clean chert
Some chert bands are entirely quartz, with no significant accessory phases. Most of these pure samples occur in sulfide facies iron formation. The grain size in this group varies considerably, ranging from -0.2 mm polygonal quartz to true microcrystalline chert (Figure 5-K, L).
30 Geochronology
Recent U-Pb zircon ages from the Deloro assemblage in the Bartlett dome by the
Ontario Geological Survey (OGS) have yielded high precision time constraints on the deposition of the Middle IF. These ages were obtained for rhyolitic rocks that immediately underlie and overlie the iron formation at the McArthur power line locality.
The age of the underlying rhyolite is 2728.1+/-1.6 Ma, and the overlying rhyolite was initially dated at 2724.5+/-2.1 Ma both by the air-abraded single zircon TIMS method
(Houle et al., 2008). The age of the overlying rhyolite has since been refined to 2724.5+/-
1.5 Ma through the analysis of additional zircons from the original sample (Ayer, personal communication, 2008). These constraining ages indicate a depositional period of
0.5 to 6.7 My for the 50 m Middle IF at McArthur power line, a comparable time gap to those observed for iron formations in the nearby Swayze greenstone belt (albeit with slightly different underlying and overlying ages), indicating that this iron formation may be a regional feature (van Breemen et al, 2006; Thurston et al., 2008). This provides a sedimentation rate ranging from 0.007 to 0.1 mm/yr, a slow deposition rate, considering that the roughly 8 km thickness of volcanic and sedimentary rocks of the Deloro
Assemblage at the Bartlett Dome was deposited between 2730-2722 Ma with an average rate of emplacement of 1 mm/yr (Houle et al., 2008).
In addition to these OGS studies, a small population of zircon was extracted from the Middle BIF at the McArthur power line. The presence of zircon in iron formation, with chert bands not showing any petrographic indication of felsic volcanic parentage, suggests considerable silicification during iron formation development. These zircons were dated using LA-ICP-MS to determine if they were derived from precursor units that 31 were emplaced concurrently with the iron formation, or if they were derived from older, detrital zircons from preexisting units.
The data falls into three groups. Some grains yielded an age of 2788+/- 23 Ma, interpreted as xenocrystic cores indicating that the rock was emplaced through older crust
Data-point error ellipses are 2n ,' (c.f. Ayer et al, 2002, Ketchum et al.,
2008) (Figure 6-A). Most zircons
yielded a date of 2733+A6.9 Ma,
interpreted as a crystallization age for
Intercepts at 10±50& 2788+23 [±24] a felsic precursor unit (Figure 6-C). Ma MSWD = 2.7 Rims on some grains were analyzed
and yielded a date of 2670+/-11 Ma,
similar to the age of Au-related
hydrothermal activity elsewhere in the
AGB and interpreted as a
metamorphic overprint (Ayer et al,
2002).
Data-point error ellipses are 2n Figure 6: Concordia diagrams of zircons analyzed for U-Pb ages by LA-ICP-MS. (A) Ancient inherited zircon core ages with a mean age of 2788+23 Ma, older than any known assemblage in the Abitibi. (B) Main- zircon population, demonstrating a likely primary volcanic age of 2733.4+6.9 Ma. (C) Young zircon overprint age of 2670±11 Ma, likely a metamorphic overprint; note reverse Intercepts at discordance. 10±50&2670±11 [±13] Ma MSWD = 0.71 The most important result seen here is the 2733+A6.9 Ma crystallization age
32 (Figure 6-C). Within the error, which is larger than that of the TIMS analyses, this interpreted age fits into the bracketing constraints of the under-and-overlying units. This means that, at the time of the deposition of the Middle IF, concurrent volcanism or erosion of volcanic products introduced a small number of zircons into the depositional system, and that the ash beds within the iron formation are more likely to be primary, rather than redeposited material with long surface residence times.
Analytical methods
Extra thick polished thin sections of chert samples were prepared for in situ geochemical analysis using laser ablation inductively coupled plasma mass spectrometry
(LA-ICP-MS). This analytical technique was conducted as a means to map the geochemical patterns of the different chert bands and aid in the selection of chemically distinctive cherts for high-precision solution analysis. A suite of trace elements of interest including the rare earth elements (REE), and the high field-strength elements (HFSE) was analyzed at the Chemical Fingerprinting Lab at Laurentian University using a New Wave
Nd: YAG 213 nm laser ablation system coupled to a quadrupole Thermo X II series ICP-
MS, operating under similar parameters to Kamber and Webb (2007). Using a 100 urn beam diameter at a repetition of 10 Hz and energy density of 13 J/cm2, up to 12 spots or a traverse of the chert bands were analyzed. A total analytical time of 75 seconds was used, including a washout time of 30 seconds for single spots and longer for traverses, contingent on the length of the traverse. All spots and traverses were pre-ablated at a greater beam diameter to remove any possible surface contamination. The synthetic glass
33 standards NIST 612 and NIST 614 were used for instrument calibration. They were analyzed using the same parameters as the samples at the beginning and end of each run, as well as after every 15-20 spots to ensure that the machine was maintaining analytical consistency. Calibration values were those reported in Pearce et al. (1997).
When working with hydrogenous sediments, such as BIFs and chert, shale normalizations are used because the REEs are delivered into the oceans via river water, in roughly upper crustal proportions. Shale normalizations are typically the best proxy for the average upper crust, and are therefore the most diagnostic normalization for the identification of anomalies in the REE+Y. In this study the Mud from Queensland
(MUQ) shale standard presented in Kamber et al. (2005) was used. The MUQ standard was chosen over the more common Post-Archean Australian Shale (PAAS) and North
American Shale Composite (NASC) shale normalizations because of the large mafic volcanic rock component in the source area, similar to the proposed environment in the
AGB during its formation. Hypothetically, an Archean shale standard could be the best choice for normalization; however most such established standards consist of old, lower- precision data that can create false anomalies in the REE+Y patterns. Based on the MUQ- normalized REE+Y patterns, individual chert bands were divided into different geochemical groups. Samples that best represented individual geochemical groups, as well as the most unusual samples were selected for later dissolution analysis by high precision inductively coupled plasma-mass spectrometry (ICP-MS).
Individual chert bands and clasts were cut from the chips remaining from making the thin sections used in petrography and LA-ICP-MS. These samples were sliced using a water-cooled microsaw to obtain as close as possible to the same material analyzed using 34 LA-ICP-MS. These slices were crushed to medium sand-sized grains using an agate pestle and fragments as close to pure chert as possible were then picked under a microscope to minimize contamination from non-chert materials. These chert fragments were then pulverized to a fine powder (~50-100 urn) in an agate mortar and pestle that was cleaned with isopropyl alcohol between each sample. Samples were weighed in at
0.009 to 0.1 g, depending on the amount of sample available, and placed in pre-cleaned
30 mL Savillex Teflon beakers with 6 drops of 6N HNO3 to prevent the sample from sticking to the sides of the beaker. Samples were then dissolved in a 2:1 mixture of pure sub-boiling distilled HF and pure, triple- sub-boiling distilled HNO3 and underwent closed-beaker digestion. To determine the role of non-silicate phases in the trace element distribution on cherts, two samples rich in Fe-oxides were run in duplicate. These two samples underwent closed beaker digestion in 1.5 and 2.0 mL of pure HNO3, respectively. All samples were then left to digest for 3 days on a hotplate at 140°C.
Following digestion, all samples were removed from the hotplate and allowed to cool before all condensed drops on the inside of the beakers were carefully collected to avoid the loss of volatilized elements. The duplicate samples were placed in a centrifuge to separate the material leached and dissolved by HNO3 from the residual silicate material, and the leach was placed in a separate beaker from the residue. All samples were then opened and dried overnight at 110°C on a hotplate. The residual material from the duplicates was subsequently dissolved in a 2:1 mixture of HF and HNO3 by closed beaker digestion overnight. Non-duplicate samples had 1 mL pure HNO3 added and were evaporated at 110°C to convert and volatilize any fluorides formed during HF digestion
35 (fluoride conversion). This process was repeated after 4 hours. Following HF digestion, fluoride conversion was also conducted on the residue of the duplicate samples.
All samples then had 2.5 mL 10% HN03 added and were left overnight at 120°C to create the stock solution. Samples were then cooled and again had all drops collected prior to transfer to pre-cleaned, capped test tubes. The caps and beakers were then rinsed with 0.5 mL Milli-Q water and all remaining material was collected and added to the stock solution. Samples were homogenized and placed in a centrifuge. After centrifugation, stock solutions were inspected for any solid residual material (e.g. unconverted fluoride or undigested minerals). Where solids were found, they were floating, indicating they were a light phase such as graphite, instead of other possible residues like chromite or rutile. Aliquots of the stock solutions were diluted by a nominal dilution factor of ca. lOOx. This is much less diluted than normal rock samples (typically diluted by a factor of 3000x). Because cherts are almost pure SiC<2, however, the true dilution factor is much higher than the nominal factor, thanks to loss of Si as the tetrafluoride. This analytical procedure enables very low detection limits (down to parts per trillion) for many elements. The dilutions were doped using a modified version of the enriched isotope internal standard method described in Eggins et al. (1997) consisting of isotopically enriched Li and U as well as pure Rh, Re, and Bi.
Using a quadrupole Thermo X II series ICP-MS, the samples were analyzed for a suite of 46 trace elements (Li, Be, Sc, Ti, V, Cr, Co, Ni, Cu, Zn, Ga, As, Rb, Sr, Y, Zr,
Nb, Mo, Ag, Cd, In, Sn, Sb, Cs, Ba, La, Ce, Pr, Nd, Sm, Eu, Gd, Tb, Dy, Ho, Er, Tm, Yb,
Lu, Hf, Ta, W, Tl, Pb, Th, U). The instrument response was calibrated with multiple digestions of the USGS reference material W-2, of which the preferred values are 36 reported in Table 1. The USGS standards BIR-1 and BHVO-2 were analyzed as unknowns and compared to the long-term reproducibility reported by Kamber (2009) to serve as an external standard control. The results from the blind analysis of BIR-1, long term values, RSD, and accuracy are presented in Table 1. The values for BHVO-2 are not presented here.
Geochemical Results
Familiarity with REE+Y systematics in seawater is imperative when examining the trace element data obtained from hydrogenous sediments such as chert. The marine
REE+Y are derived primarily through chemical weathering of rocks in the subaerial environment, from whence the REE+Y are then transported into the oceans via rivers and estuaries. Prior to release into the estuary, river water shows flat REE+Y patterns when normalized to shale standards such as the Mud from Queensland (MUQ) standard
(Kamber et al., 2005; Lawrence & Kamber, 2006; Lawrence et al., 2006b) (Figure-7A).
During the release of the fresh river water into the ocean in the estuary, the REE+Y patterns that are characteristic of seawater are created (Figure-7B). The light rare-earth elements (LREE) are preferentially adsorbed by iron oxy-hydroxides over the heavy rare earth elements (HREE) during mixing of fresh and saline waters, causing seawater to demonstrate LREE-depletion (Lawrence & Kamber, 2006). Because of their 4f-orbital electron coordination (empty, half-full, full), La, Gd, and Lu do not adhere to flocculating particles as well as their neighboring REEs (Bau, 1996). This causes seawater and its precipitates to have positive anomalies of each of these elements, due to their greater
37 concentration in seawater. The effect of this process is demonstrated in figure 7, with
REE+Y patterns from modern fresh river water (Figure 7-A), and from water with a salinity of 33%o from the same estuarine system (Figure 7-B) (from Lawrence &Kamber,
2006).
1 A-5| I I *^^~— I —^^" I • I • I I I I I I I
1Q-7I • • • • • • • • • • • i • • , La Ce Pr Nd Sm Eu Gd Tb Dy Y Ho Er Tm Yb Lu Figure 7: REE+Y patterns from a modern estuarine system (from Lawrence & Kamber, 2006). (A) A relatively flat REE+Y pattern from fresh river water entering the estuary (salinity= 0.732%o). (B) A sample bearing many typical seawater-type anomalies taken from the full salinity end of the estuary (salinity=33%o). Another important diagnostic feature of a seawater REE+Y pattern is the positive
Y anomaly. The chemical twins Y and Ho have the same charge balance (3+) and near- identical ionic radii in octahedral coordination with 02"vm 1.4A (Yvm=1.019 A and
o Hovm=1.015 A) (Shannon, 1976). This has the effect of causing these elements to occur at a near-constant ratio in most terrestrial rocks, with a chondritic Y/Ho ratio of 38 25.94+0.08 (Pack et al, 2007). This does not hold true for seawater, which demonstrates strongly super-chondritic Y/Ho ratios (Bau, 1996). Because Y lacks a 4f-electron shell, it has a different complexation behavior than the 4f-bearing Ho, because the 10 4f-electrons in Ho facilitate the formation of strong surface and solution complexes, which the 4f- electron free Y does not readily form. This results in the preferential removal of Ho over
Y by iron oxy-hydroxides during mixing of saline and fresh water in the estuary
(Lawrence & Kamber, 2006). Calculated from the values of the MUQ-standard, the Y/Ho ratio for clastic sedimentary rocks is 26.1, very similar to the chondritic values cited above, and is considered the non-anomalous ratio for the purposes of this study (Kamber et al., 2005).
While these features have been well documented in modern oceans (Johannesson et al, 2006), the possibility that same processes did not control REE+Y fractionation in the ocean must be considered. However, as modern hydrogenous sediments precipitate in chemical equilibrium with seawater, it can be inferred that Archean hydrogenous sediments also did so (Bolhar et al, 2005). Given that the mixing of fresh and saline waters appears to be the principle mechanism for generation of the seawater-type REE+Y anomalies (Lawrence &Kamber, 2006), it can be inferred that the same process formed those anomalies in Archean seawater-derived rocks.
The positive Eu anomaly is a very common feature in Archean hydrogenous sediments (Bolhar et al., 2005). Europium exists in 2 valances, divalent and trivalent
(Eu2+ and Eu3+) and Eu2+ more readily substitutes than the strictly trivalent REE for Ca in plagioclase, and is thus susceptible to being leached from the mineral when high- temperature (>250°C) hydrothermal fluids move through a package of basaltic rocks, 39 such as would occur at virtually any hydrothermal vent site within an Archean
Greenstone Belt (Bau & Dulski, 1999; Douville et al, 1999). Additionally, due to the low redox potential (Eh) of Archean seawater (Huston & Logan, 2004), there would be no reason to suspect that hydrothermal Eu would been completely scavenged by iron-oxy- hydroxides at vent sites and thus was able to be transported great distances, resulting in the near-ubiquitous positive Eu anomaly in shale-normalized Archean hydrogenous sediments, including those that form in shallow water shelf settings far from hydrothermal activity, as has been observed by Derry and Jacobsen (1990).
In this study, REE+Y data was normalized to the upper crustal shale average
MUQ. Normalization to this standard is the best fit for hydrogenous sediments, particularly those within a greenstone belt, for reasons discussed in the preceding section.
Anomalies were calculated using the geometric formulas presented in Lawrence et al.
(2006b), as the REE+Y patterns are smoother in log-linear plots than in simple linear plots. The anomalies for La, Ce, Eu, Gd, and Lu were calculated in this way, but the Lu anomaly will not be discussed here because of its debatable validity, stemming from the analytical problems of measuring this element in seawater (Lawrence & Kamber, 2006, and references within). The La anomaly was calculated as a projection from Pr and Nd, while the Eu and Gd anomalies were calculated from Sm and Tb (Lawrence et al.,
2006b). These three elements (together with the Y/Ho ratio) were chosen for the greatest focus because of their distinctive behavior in Archean hydrogenous sedimentary rocks.
LA-ICP-MS Results
40 Geochemical mapping of individual chert bands by LA-ICP-MS was conducted with the intent of identifying different geochemical chert phases at a band-by-band scale.
This differs from previous geochemical studies of BIF and cherts, as the selection of individual chert phases for high-precision chemical study has not been previously controlled using in situ chemical analysis. By using this method, chemical variations between individual chert bands, in some cases separated by only a single Fe-band of only
a few mm thickness, were
identified and subsequently
analyzed independently using high-
precision solution-based ICP-MS.
Figure 8: REE+Y patterns of LA-ICP-MS La Ce Pr Nd Sm 6u Gd Tb Dy Y Ho Er Tm Yb Lu data on Abitibi Cherts. (A) REE+Y patterns of traverse integrations from sample 07-GJB-046C. Hollow squares represent linel (LREE-depleted with positive Y anomalies), and filled triangles represent line2 (arched patterns, generally with no Y anomaly). (B) High quality, individual analytical spots of sample 06- PCT-003D, demonstrating high-quality results using the LA-ICP-MS (Thurston and Kamber, unpublished data). (C) Poor quality results from LA-ICP-MS analysis of sample 07-GJB-011A, showing a number of in pattern shapes (small hollow squares), and the contrast between the averaged LA-ICP-MS results (heavy line, no symbols) and the high-precision solution ICP-MS results (filled triangles). Notice that a few high-concentration HREE-depleted samples have caused the average LA-ICP-MS pattern to also be HREE-depleted, whereas the solution work reflects the true composition of the chert band with mild LREE-depletion.
The analysis of sample 07- La Ce Pr Nd Sm Eu Gd Tb Dy Y Ho Er Tm Yb Lu GJB-046C provides the best example of the geochemical differences that can be
41 identified between chert bands using this technique (Figure 8-A). Field observations indicated it was a massive chert band with very fine Fe-oxide and carbonate microbands within it, whereas petrography revealed 2 nearly identical chert bands separated by a thin band (2 mm) of carbonate and 1mm chert clasts. These chert bands were analyzed using the LA-ICP-MS in two separate traverses with the laser, linel and line2. When normalized to MUQ, each traverse yielded a distinct REE+Y pattern from the other, and showed considerable internal consistency between different integrations along each traverse.
The REE+Y patterns derived from the first traverse, linel, yielded patterns with strong LREE-depletion relative to the HREE (PrMuo/LuMUQ= 0.065-0.09), creating patterns with a steep positive slope overall (Figure 7-A). The La and Gd anomalies from this traverse ranged from moderate to large positive (La/La*=1.50-2.05; Gd/Gd*=1.08-
1.31), and the Eu anomaly was very large (Eu/Eu*=5.23-5.70). Large positive Y anomalies(Y/Ho= 31.6-38.1) were also observed in these samples. The traits observed in these patterns are all characteristic of modern seawater (positive La, Gd, and Y anomalies), with the exception of the large Eu anomalies, a common feature in rocks that were precipitated from Archean seawater (Bolhar et al., 2004; Bolhar et al., 2005;
Lawrence & Kamber, 2006). Consequently, this chert band was identified as 07-GJB-
046CS (S for seawater) when extracted for individual solution ICP-MS.
In contrast, the REE+Y patterns from the second traverse, line2, showed very different overall shapes from linel (Figure 8-A). These patterns also showed LREE- depletion relative to the HREE (PrMuo/LuMuQ = 0.079-0.53), albeit less than linel. The other 2 patterns from this traverse show a steep positive slope, very similar to the patterns 42 from linel, however the anomalies of La, Eu, Gd, and Y for these patterns are more consistent with other patterns from line2 than they are with linel. The patterns from this line show a fairly wide range of La anomalies, from small negative to moderately positive, albeit smaller positive than linel (La/La*=0.87-1.59). The Gd anomalies generally have larger positive values than linel with Gd/Gd* ranging from 1.12-1.19, but smaller Eu anomalies, ranging from 3.9-5.8. The Y anomaly is perhaps the most diagnostic feature of patterns from this traverse, with Y/Ho values ranging from 24.4-
32.44, with most samples showing negative, flat, or small positive Y anomalies, save 1, which has a strongly superchondritic Y/Ho ratio. This type of arched pattern with a pronounced Eu anomaly was initially interpreted to be similar to shale-normalized high temperature hydrothermal fluids, such as seen in Alexander et al. (2008) and Bau and
Dulski (1999), except with a smaller Eu anomaly and greater LREE-depletion. Based on this early interpretation, this chert band was given the sample number 07-GJB-046CH (H for hydrothermal) when selected for dissolution ICP-MS analysis. This preliminary interpretation was later found to be false through solution-based ICP-MS analysis (see
Group 1; Ni/Cr=l; Subgroup 2; LREE and HREE depleted).
Despite the obvious advantages of using the LA-ICP-MS mapping technique to identify geochemically different chert bands, the analytical constraints of the method produce some limitations on the quality of the data. In order for the sample to yield high- quality data with smooth REE+Y diagrams, the sample needs to have a fairly high degree of homogeneity within a given chert band. For example, a specimen analyzed in a previous study, 06-PCT-003D, the REE+Y patterns from individual laser spots show remarkably good consistency (Figure 8-B), with only a few exceptions (likely caused by
43 the ablation of REE-rich phases such as phosphates or carbonates) (Thurston and
Kamber, unpublished data). This specimen was a highly homogenous green chert, and was analyzed using a beam diameter of 200 urn, ablating an area 4 times as large as the analyses in this study, thereby increasing the amount of material analyzed, serving to improve the precision of the results via better counting statistics. Additionally, the patterns obtained for this specimen were very smooth, hardly yielding any anomalies other than the non-strictly trivalent Ce and Eu. This smoothness is a product of the relatively high concentrations of the REE's (roughly equal to MUQ) atypical of hydrogenous chert, and the amount of material analyzed. Despite the great improvement on the quality of data obtained, using such a large beam diameter on the LA-ICP-MS is not practical, due to the heterogeneous nature of most rocks, and the difficulty of obtaining homogenous energy distribution in a largely expanded beam from a solid source laser with a maximum aperture of 100 microns.
At the other end of the spectrum, a sample with less homogeneity or lower absolute concentrations of the REEs can yield apparent patterns that are very irregular and bumpy, forming artificial anomalies on the REEs with low natural abundances or for which a less abundant isotope is analyzed (e.g. Nd). Sample 07-GJB-011A had 8 different spots analyzed, and demonstrates drastically different REE+Y patterns for each different spot (Figure 8-C). In some cases, this appears to be the product of ablating REE- rich mineral phases, but in many other samples irregular patterns are produced by low concentrations of many of the REEs (0-20 ppb). The irregular patterns are produced as the elements approach their detection limits due to two different contributing factors. The short integration time (typically 30 ms) per analyte leads to artificial step-like patterns
44 when the counts are extrapolated up to 1000 ms (counts per second). The other factor is that at low elemental abundances, analytical interferences (such as BaO interfering with
Eu, or BaCl with Yb and Lu) that are typically not a problem in LA-ICP-MS may become significant. Similar observations in solution ICP-MS mode have been shown empirically through increasing dilutions of waters of known REE+Y concentrations, resulting in erratic patterns at very high sample dilution (Lawrence et al., 2006a).
Interestingly, when the results from numerous spots within a single chert band are averaged, they do not reflect the true REE+Y composition of that chert band when analyzed by solution ICP-MS, as averaging the individual spots is more akin to using the average mineral composition of a rock to calculate its overall composition. In the case of
07-GJB-011A, this resulted in a REE+Y pattern that has a negative slope, which contrasts sharply with the shape of the pattern from the dissolution ICP-MS analysis for that sample, which has a positive slope. This is because of the overbearing influence of just two HREE-depleted patterns that also had much higher abundances than the bulk of the spots from this sample (likely caused by the ablation of a sub-microscopic LREE-rich mineral phase), resulting in a skewed average that is not properly representative of the sample. Many of these mineral phases are volumetrically insignificant compared to the rest of the chert, and consequently their identification is prohibited due to the combined influence of the recrystallized nature of these cherts, as well as a very large nugget effect.
Use of LA-ICP-MS analysis to identify geochemically distinct chert bands when studying iron formation is clearly a very useful tool, however the data obtained from this method should be used with caution. As demonstrated here, a number of small factors, such as homogeneity, overall abundances, and the presence of non-visible mineral phases
45 (either sub-microscopic inclusions, or phases below the surface of the thin section that are exposed during ablation), can all influence the quality of the data obtained using this technique. High-quality, smooth REE+Y patterns obtained using this method may be viewed as geochemically representative within reason, however, it is imperative that the researcher consider that even this data may not reflect the overall chemistry of the sample. For the low abundance trace element geochemical study of iron formation and chert, LA-ICP-MS should be used strictly as a means of identifying bands of interest for further analysis by high-precision solution ICP-MS analysis.
Solution ICP-MS Results
Results from solution ICP-MS analysis yielded a wide range of trace element concentrations, reported in parts per billion (ppb) (See digital Appendix 3). The majority of the elements included in this study yielded concentrations of no more than 10,000 ppb, with the common exception of such elements as Ti, Cr, Ni, Cu, Zn, Sr, Ba, and Pb.
46 Table 1: Trace Element concentrations from ICP-MS (parts per billion) Preferred W-2 BIR-1 (this study) BIR-1 long term BIR-1 %RSD BIR-1 accuracy Blank Li 9158 3087 3136 1.202 0.984 2.648 Be 617.5 92.45 93.38 3.451 0.990 0.025 Sc 36074 43996 44497 1.409 0.989 0.042 Ti 6354611 5759123 5789533 1.219 0.995 92.16 V 261597 319899 322519 1.119 0.992 0.162 Cr 92791 410910 410576 1.797 1.001 3.910 Co 44526 52789 53334 1.072 0.990 0.391 Ni 69993 169389 171219 1.907 0.989 6.970 Cu 103000 117215 119473 1.681 0.981 7.636 Zn 77000 68713 70733 2.290 0.971 89.74 Ga 17424 15280 15428 0.974 0.990 0.037 As 833.0 89.30 93.94 39.71 0.951 2.827 Rb 19803 188.3 193.8 3.038 0.971 0.417 Sr 194828 108846 109506 1.032 0.994 1.288 Y 20113 14503 14660 1.073 0.989 0.042 Zr 87866 14514 14502 2.591 0.999 3.189 Nb 7275 534.7 534.8 1.209 1.000 0.129 Mo 423.3 31.79 43.75 44.77 0.727 0.228 Ag 76.99 43.40 0.439 Cd 77.00 58.94 65.17 7.459 0.904 0.033 In 64.58 57.77 0.003 Sn 1950 800.3 815.3 9.036 0.982 0.628 Sb 709.0 427.4 475.5 14.11 0.899 0.086 Cs 888.2 5.029 5.054 4.883 0.995 0.033 Ba 169680 6418 6599 3.470 0.973 1.131 La 10521 598.2 604.3 1.185 0.990 0.040 Ce 23216 1891 1900 1.380 0.996 0.093 Pr 3025 372.1 378.2 1.132 0.984 0.006 Nd 12911 2332 2382 1.102 0.979 0.149 Sm 3266 1078 1099 1.076 0.981 0.010 Eu 1094 516.3 524.6 0.813 0.984 0.003 Gd 3708 1843 1874 1.136 0.983 0.006 Tb 615.1 361.5 365.4 0.875 0.989 0.001 Dy 3808 2496 2543 1.088 0.981 0.007 Ho 803.3 574 583.2 1.002 0.984 0.002 Er 2222 1671 1701 0.895 0.982 0.001 Tm 327.2 253.3 258.5 0.908 0.980 0.000 Yb 2058 1639 1663 0.852 0.986 0.000 Lu 301.3 243.2 248.2 1.030 0.980 0.001 Hf 2356 566.3 570.7 1.363 0.992 0.071 Ta 454.2 37.06 37.01 1.396 1.002 0.013 W 240.0 6.250 5.958 9.546 1.049 0.015 Tl 90.00 1.183 1.251 7.515 0.946 0.000 Pb 7528 2865 3271 14.96 0.876 0.859 Th 2104 28.32 29.37 3.995 0.965 0.012 U 505.0 10.10 10.09 2.111 1.001 0.020 Y/Ho La/La* Ce/Ce* Eu/Eu* Gd/Gd* Lu/Lu* Pr/Lu Ni/Cr
47 Table 1: Trace Element concentrations from ICP-MS (parts per billion) (cont) 07-GJB-011A 07-GJB-018A 07-GJB-022A 07-GJB-027B 07-GJB-031A 07-GJB-036B Li 2079 68.16 101.0 35.50 3723 615.2 Be 730.3 238.3 139.1 52.48 789.7 53.77 Sc 475.7 198.1 879.1 26.08 5226 117.5 Ti 284064 12958 6779 3086 1499658 16830 V 3878 664.1 2972 439.8 40219 809.3 Cr 2046 1208 648.6 139.2 27785 344.4 Co 869.0 248.1 891.6 290.7 6755 357.7 Ni 1810 2042 5442 782.7 11070 1768 Cu 4648 2617 73012 4449.3 41037 7626 Zn 9044 6036 961809 8441.4 108454 9775 Ga 1864 384.5 1331 84.76 10594 108.8 As 452.7 206.1 86.50 36.57 407.8 181905 Rb 7365 2399 198.4 728.7 2309 2263 Sr 12055 14206 477.1 359.3 42270 452.6 Y 2006 1765 129.7 449.7 11827 1199 Zr 19057 875.2 1398 88.63 57772 1239 Nb 419.0 39.78 66.91 6.686 1877 55.44 Mo 405.4 249.8 13.87 25.33 3699 693.3 Ag 54.85 14.61 362.8 3.595 16.33 9.667 Cd 20.45 18.22 5284 3.893 55.86 12.06 In 10.77 1.265 135.6 9.075 266.0 3.382 Sn 722.3 490.9 889.2 286.8 4986 721.0 Sb 311.5 55.80 78.45 76.65 68.55 519.1 Cs 1553 1786 38.58 215.0 444.5 683.3 Ba 12965 7462 1905 1487 14205 4466 La 1078 884.8 68.28 113.5 25841 547.3 Ce 2391 1527 148.8 237.4 52885 1034 Pr 299.8 177.3 17.96 32.12 6581 137.5 Nd 1197 691.9 70.48 145.0 25836 614.0 Sm 253.5 133.6 17.07 34.86 4441 139.6 Eu 239.6 114.3 9.980 29.93 2252 129.3 Gd 283.8 174.2 20.95 46.44 3601 176.4 Tb 44.94 26.39 3.645 7.155 481.1 25.09 Dy 296.3 174.5 22.16 48.25 2501 155.7 Ho 70.98 42.78 4.730 12.09 458.6 36.26 Er 222.1 130.6 13.41 35.59 1090 106.3 Tm 35.52 19.74 2.206 4.887 144.6 15.29 Yb 236.0 129.7 16.27 27.07 852.6 91.81 Lu 36.45 21.86 2.822 3.728 126.4 14.16 Hf 508.4 16.15 7.825 1.939 1332 15.14 Ta 35.48 1.911 0.735 0.384 118.6 1.613 W 174.4 304.4 11.27 21.91 284.5 43.13 Tl 79.67 1.188 5.007 2.857 13.02 2.743 Pb 3281 390.2 40462 766.5 1836 200.4 Th 198.6 15.41 3.429 4.997 1396 16.88 U 51.18 8.503 21.27 2.851 369.1 53.85 Y/Ho 28.27 41.27 27.42 37.18 25.79 33.07 La/La* 0.986 1.307 1.007 1.238 1.040 1.364 Ce/Ce* 0.974 1.028 0.995 1.020 0.965 1.027 Eu/Eu* 3.740 3.247 2.118 3.174 2.430 3.627 Gd/Gd* 1.021 1.100 0.964 1.081 1.053 1.136 Lu/Lu* 0.983 1.085 0.994 1.051 1.062 1.086 Pr/Lu 0.476 0.470 0.369 0.499 3.017 0.562 Ni/Cr 0.884 1.690 8.390 5.624 0.398 5.133
48 Table 1: Trace Element concentrations from ICP-MS (parts per billion) (cont) 07-GJB-037A 07-GJB-039A 07-GJB-039B 07-GJB-039C 07-GJB-039D 07-GJB-041B Li 61.32 11.46 216.7 17.69 13.55 70.98 Be 21.35 166.0 190.6 230.6 150.8 12.43 Sc 117.6 90.00 759.9 103.8 36.14 230.0 Ti 9063 16408 14413 4538 1523 2861 V 1139 1150 818.3 662.7 111.4 717.5 Cr 369.6 543.1 585.5 2424 5857 4061 Co 4184 3442 2799 587.1 387.5 1016 Ni 15930 1975 1438 3143 7334 8868 Cu 13556 23111 9992 8196 1484 2402 Zn 510857 19842 9785 41041 9679 2549 Ga 193.9 306.8 1085 140.7 32.21 177.2 As 4355 19.40 69.70 412.9 18.14 16.99 Rb 1184 236.3 89.35 429.1 179.0 84.82 Sr 9210 3066 449.5 27731 2025 327.0 Y 423.1 973.7 896.1 1899 1458 235.3 Zr 685.7 5551 3349 607.6 2632 518.4 Nb 29.99 24.68 69.35 18.76 14.90 9.905 Mo 41.34 205.4 47.38 508.7 1241 111.1 Ag 97.76 11.17 11.07 12.73 5.057 4.035 Cd 925.1 19.69 24.93 57.14 15.56 6.726 In 23.92 29.44 21.00 53.06 8.151 1.111 Sn 787.6 263.8 451.1 666.4 172.1 242.9 Sb 753.7 38.35 11.98 55.73 19.59 6.109 Cs 801.9 106.0 12.25 84.99 84.23 7.868 Ba 1106 533.2 356.0 5484 703.2 899.7 La 1039 378.9 490.3 443.5 334.1 586.9 Ce 2382 729.4 1117 1006.1 549.9 1279 Pr 296.5 89.62 141.4 132.6 73.01 160.7 Nd 1142 371.5 581.5 561.4 338.4 641.6 Sm 218.2 84.70 152.6 136.1 89.68 104.9 Eu 136.8 131.5 83.83 177.3 146.0 61.79 Gd 169.9 97.69 182.0 189.3 136.9 75.19 Tb 19.98 15.43 30.73 31.77 21.05 8.730 Dy 95.24 105.8 183.5 229.7 143.6 44.27 Ho 18.72 27.04 38.21 61.44 35.77 8.611 Er 52.59 91.22 107.2 205.3 110.2 23.81 Tm 8.297 16.27 16.51 35.31 16.82 3.621 Yb 59.10 126.3 114.6 259.7 107.3 24.82 Lu 9.535 23.64 18.41 42.98 16.64 3.911 Hf 15.28 102.4 43.79 13.13 20.75 6.453 Ta 1.371 1.316 2.794 0.628 0.637 0.320 W 68.28 201.6 21.96 37.28 58.94 5.293 Tl 3.566 0.995 3.178 2.616 0.959 0.663 Pb 2280 172.2 419.3 209.6 54.92 67.00 Th 11.84 1.062 18.55 3.219 2.244 5.161 U 3.026 7.631 84.22 5.233 17.84 4.258 Y/Ho 22.60 36.01 23.45 30.91 40.75 27.32 La/La* 0.895 1.249 1.008 1.032 1.691 1.000 Ce/Ce* 0.947 1.032 0.993 0.983 1.068 0.972 Eu/Eu* 3.082 6.021 2.073 4.679 5.811 2.966 Gd/Gd* 1.097 1.022 0.994 1.051 1.145 1.067 Lu/Lu* 0.957 1.019 0.979 0.951 1.027 0.972 Pr/Lu 1.801 0.220 0.445 0.179 0.254 2.380 Ni/Cr 43.10 3.636 2.457 1.297 1.252 2.184
49 Table 1: Trace Element concentrations from ICP-MS (parts per billion) (cont) 07-GJB-042B 07-GJB-043A 07-GJB-045B 07-GJB-045C 07-GJB-046CS 07-GJB-046CH Li 44.31 2.445 80.76 82.17 11.47 36.21 Be 648.9 91.92 441.9 1112 227.3 1161 Sc 914.8 101.4 766.7 209.2 28.21 99.71 Ti 17995 1290 15359 2140 1132 3274 V 2002 207.6 1225 411.9 280.5 645.1 Cr 5596 2438 20243 6245 9459 27972 Co 3844 396.6 2582 3218 740.5 2671 Ni 8207 3079 22602 8348 11601 32343 Cu 1048 519.9 2698 2200 5430 4381 Zn 249636 24576 147743 481554 21921 138955 Ga 542.6 143.4 522.4 497.3 55.77 113.9 As 249.6 9.327 134.4 176.7 24.61 96.65 Rb 1112 206.8 1400 1748 330.7 864.5 Sr 5589 2330 35921 11115 23744 9055 Y 1377 510.0 5652 7002 2386 8580 Zr 326.0 134.6 904.9 941.5 436.7 1345.6 Nb 117.0 11.16 81.02 166.7 9.857 73.64 Mo 1109 470.2 4247 1324 1959 5819 Ag 6.417 2.173 6.903 4.449 6.578 3.472 Cd 47.53 12.91 41.19 150.3 40.08 23.98 In 96.10 8.001 51.74 102.9 30.85 55.86 Sn 2008 192.4 2318 894.0 871.7 904.0 Sb 402.7 59.23 88.12 169.8 47.50 139.7 Cs 319.0 83.09 576.8 537.9 128.3 432.0 Ba 249.4 202.9 1195 19491 3250 6338.4 La 327.6 75.95 927.4 622.1 183.9 694.0 Ce 705.1 160.9 2237 1080 430.8 1872.6 Pr 92.49 24.92 312.5 181.0 69.55 313.7 Nd 392.7 116.2 1415 958.1 380.9 1816.6 Sm 94.14 32.29 429.4 307.9 137.8 750.3 Eu 111.0 41.52 547.3 580.0 233.5 1024.8 Gd 126.7 51.49 584.9 564.2 242.5 1321.5 Tb 21.22 8.754 104.7 104.3 41.98 228.9 Dy 152.2 60.93 712.6 769.7 290.0 1472.7 Ho 37.62 15.25 171.5 194.0 71.01 311.6 Er 127.6 48.48 552.5 598.3 211.8 773.9 Tm 22.78 7.969 96.13 91.84 31.09 92.99 Yb 177.3 57.93 708.3 574.2 196.2 416.8 Lu 30.74 9.793 120.7 87.24 30.09 40.50 Hf 12.55 2.736 24.35 27.22 6.906 20.04 Ta 3.20 0.213 3.384 1.230 0.241 1.096 W 80.53 51.81 68.99 59.70 19.83 57.14 Tl 11.07 1.069 11.27 12.98 1.809 5.022 Pb 145.5 58.86 193.8 178.9 301.1 201.1 Th 26.05 0.568 12.06 4.404 1.055 4.477 U 5.015 1.789 5.586 9.923 16.11 178.7 Y/Ho 36.59 33.44 32.96 36.10 33.60 27.54 La/La* 1.098 1.140 1.046 1.655 1.363 1.276 Ce/Ce* 0.990 0.921 0.992 0.965 1.038 1.057 Eu/Eu* 4.268 4.334 4.565 6.033 5.575 4.568 Gd/Gd* 1.037 1.076 1.012 1.092 1.116 1.135 Lu/Lu* 0.941 0.983 0.978 1.027 1.027 0.916 Pr/Lu 0.174 0.147 0.150 0.120 0.134 0.449 Ni/Cr 1.467 1.263 1.117 1.337 1.226 1.156
50 Table 1: Trace Element concentrations from ICP-MS (parts per billion) (cont) 07-GJB-048B 07-GJB-048C 07-' J-048C Residue 07-GJB-048C leach 07-GJB-048E Li 50.60 80.67 26.23 548.3 177.8 Be 462.1 214.0 92.77 1382 933.1 Sc 1068 140.1 76.18 320.8 1075 Ti 118242 5959 1998 39745 78956 V 4125 1876 453.8 14783 5314 Cr 3158 586.3 119.0 4250 1842 Co 4469 660.7 423.6 2563 15759 Ni 6982 377.5 82.46 3038 13116 Cu 76063 8026 117.1 15851 3414 Zn 18838 29982 16704 132702 59506 Ga 2329 792.6 222.0 5944 1038 As 64.74 105.3 14.25 861.6 8908 Rb 20432 1999 382.5 16879 9952 Sr 2592 2562 446.8 22001 17403 Y 3742 1693 558.0 11511 8414 Zr 8528 1047 469.3 1154 5625 Nb 387.3 72.27 69.76 30.68 216.4 Mo 91.59 134.8 39.43 1038 402.3 Ag 205.4 6.678 1.399 85.27 88.44 Cd 31.69 13.35 2.447 105.8 96.61 In 85.33 77.83 20.79 600.7 77.69 Sn 819.7 435.0 218.5 2006 1640 Sb 25.54 126.9 122.8 59.08 423.4 Cs 22094 1739 235.4 16140 6479 Ba 35400 4688 1451 35486 24813 La 3327 8260 1716 51271 10996 Ce 6864 11298 2395 69509 21007 Pr 822.9 1142 250.1 6899 2343 Nd 3132 3818 859.5 23106 8657 Sm 592.2 414.5 105.1 2544 1390 Eu 260.7 338.0 82.54 2485 920.9 Gd 613.5 278.5 78.99 1891 1299 Tb 90.94 32.43 10.13 217.4 172.7 Dy 539.6 180.1 59.34 1211 1020 Ho 117.4 41.44 14.09 279.4 230.6 Er 337.5 122.5 43.18 810.6 672.0 Tm 51.84 19.39 7.104 125.7 102.7 Yb 341.7 134.4 49.31 860.7 694.2 Lu 52.87 22.35 8.237 139.7 118.2 Hf 175.8 9.701 4.780 20.24 88.16 Ta 36.39 0.625 0.679 0.053 9.245 W 43.85 36.91 30.60 57.04 34.32 Tl 20.42 2.359 1.651 7.576 20.98 Pb 931.9 56.48 12.31 662.6 260.9 Th 302.9 5.624 1.877 32.24 143.8 U 140.6 2.303 0.888 13.65 40.59 Y/Ho 31.87 40.86 39.60 41.20 36.48 La/La* 1.007 1.390 1.394 1.433 1.102 Ce/Ce* 0.971 1.012 1.007 1.032 1.013 Eu/Eu* 1.842 4.248 3.768 5.032 2.999 Gd/Gd* 1.048 1.056 1.016 1.123 1.096 Lu/Lu* 0.992 1.014 1.017 1.002 1.065 Pr/Lu 0.902 2.960 1.759 2.860 1.148 Ni/Cr 2.211 0.644 0.693 0.715 7.123 51 Table 1: Trace Element concentrations from ICP-MS (parts per billion) (cont) 07-GJB-048GM 07-GJB-048GP 07-GJB-048H 07-GJB-049B 07-GJB-049B Residue Li 913.0 289.3 102.6 42.34 19.57 Be 85.69 45.37 564.5 445.3 175.7 Sc 1485 1240 142.8 115.3 132.0 Ti 232098 71369 5612 2652 1265 V 6824 5024 694.6 407.9 52.03 Cr 5314 6072 230.5 2093 540.1 Co 3690 7172 7494 2352 1446 Ni 3505 4976 2364 3563 298.7 Cu 208752 64868 8400 1986 63.97 Zn 516999 25109 60877 42949 37244 Ga 2107 1005 424.7 285.1 44.52 As 43.45 817.8 407.4 183.9 14.32 Rb 186.9 120.1 2175 263.3 90.19 Sr 384.0 312.7 5375 9521 4882 Y 1515 823.8 2758 2602 1267 Zr 11095 6143 2056 433.6 120.6 Nb 568.1 457.0 115.1 26.19 20.02 Mo 897.9 1185 118.0 11.28 2.994 Ag 526.6 581.7 29.29 5.846 1.069 Cd 1144 32.57 21.73 29.97 24.36 In 256.4 15.25 26.37 17.32 13.45 Sn 583.7 1237 1063 469.3 211.8 Sb 22.00 15.61 439.0 250.5 206.2 Cs 30.84 10.44 896.1 196.6 81.49 Ba 1414 763.1 1419 789.2 194.8 La 658.6 567.8 1371 415.4 68.70 Ce 1509 1128 2628 832.8 169.6 Pr 192.5 130.7 325.0 114.6 29.02 Nd 791.7 505.4 1361 539.1 155.5 Sm 183.7 106.3 287.4 137.2 48.75 Eu 86.31 51.26 323.9 175.0 57.23 Gd 195.6 111.6 370.1 201.8 76.27 Tb 33.55 18.42 53.24 33.07 13.53 Dy 220.4 120.3 329.2 228.9 98.42 Ho 53.30 28.56 76.21 59.45 27.57 Er 173.1 93.21 227.3 192.4 96.20 Tm 30.13 16.20 36.74 29.61 15.75 Yb 226.3 124.6 270.2 187.1 101.2 Lu 41.03 24.22 49.43 31.82 18.62 Hf 277.8 140.9 14.84 7.889 2.179 Ta 36.58 8.714 2.422 0.541 0.475 W 113.6 63.24 66.94 245.3 150.7 Tl 1.994 1.108 16.44 1.668 0.467 Pb 353.0 173.6 781.1 322.2 31.50 Th 184.1 31.51 3.360 3.362 0.781 U 103.7 44.71 9.045 2.519 0.417 Y/Ho 28.42 28.84 36.19 43.77 45.96 La/La* 0.996 1.118 1.272 1.381 1.169 Ce/Ce* 0.986 1.021 1.036 1.047 0.958 Eu/Eu* 1.835 1.906 4.402 4.534 3.942 Gd/Gd* 0.949 0.963 1.142 1.089 1.044 Lu/Lu* 1.020 1.069 1.051 1.137 1.210 Pr/Lu 0.272 0.312 0.381 0.209 0.090 Ni/Cr 0.660 0.820 10.26 1.702 0.553
52 Table 1: Trace Element concentrations from ICP-MS (parts per billion) (cont) 07-GJB-049B leach 07-GJB-049E Li 139.0 23.33 Be 3043 337.0 Sc 124.9 79.11 Ti 10154 6838 V 3226 725.1 Cr 12926 161.0 Co 7131 1781 Ni 32766 649.6 Cu 20052 1017 Zn 50414 35799 Ga 2219 284.0 As 3628 127.6 Rb 1488 260.3 Sr 36130 2640 Y 10771 1761 Zr 1828 503.8 Nb 37.29 47.85 Mo 65.97 18.97 Ag 62.37 1.917 Cd 61.93 12.65 In 27.06 20.58 Sn 2398 275.5 Sb 384.9 38.59 Cs 1033 206.0 Ba 5565 338.1 La 3223 587.4 Ce 6142 1150 Pr 785.0 143.4 Nd 3536 598.1 Sm 813.2 133.8 Eu 1077 132.7 Gd 1153 183.3 Tb 180.2 28.78 Dy 1186 193.6 Ho 289.6 46.85 Er 860.7 142.1 Tm 125.1 21.83 Yb 760.3 143.4 Lu 114.8 22.97 Hf 35.42 6.562 Ta 0.149 0.769 W 831.6 69.50 Tl 11.63 1.391 Pb 2698 65.96 Th 25.98 4.785 U 20.92 15.63 Y/Ho 37.20 37.60 La/La* 1.432 1.225 Ce/Ce* 1.078 1.023 Eu/Eu* 4.926 3.658 Gd/Gd* 1.129 1.092 Lu/Lu* 1.050 1.030 Pr/Lu 0.396 0.362 Ni/Cr 2.535 4.034 53 The potentially strong influence of the near-ubiquitous Fe-oxides and assorted carbonate minerals on the REE+Y and trace-element composition of the chert bands was addressed through analyzing duplicate samples of two samples rich in Fe-oxides. The duplicate samples were digested first only in HNO3 to dissolve any oxides and carbonates in the samples, and then the residue was dissolved using the same 2:1 HF-HNO3 digestion as was used on normal samples, in an effort to isolate the trace element
composition of the
silicate component of the
cherts alone. o
Q. £ Figure 9: REE+Y patterns of ra W the 2 samples run in duplicates to determine the effect of oxides and carbonates on the REE+Y patterns of cherts. (A) Duplicate patterns of sample 10s 07-GJB-048C, with the La Ce Pr Nd Sm Eu Gd Tb Dy Y Ho Er Tm Yb Lu composite (normal) sample in empty diamonds, the leach B (HNO3 only) sample in empty -^ Composite triangles, and the residue •e- Residue (undissolved by HNO3) -*- Leach sample in hollow squares. All three patterns are near-
-A £ A parallel. (B) Duplicate patterns of sample 07-GJB- 049B, using the same symbols as the above samples. The patterns are less parallel than 07-GJB-048C, but are broadly similar in shape. This is likely caused by higher amounts of oxides such as magnetite in this particular chert sample. 10'!••• La Ce Pr Nd Sm Eu Gd Tb Dy Y Ho Er Tm Yb Lu The REE+Y results from these two samples that were run in duplicate (07-GJB-048C and 07-GJB-049B) were normalized to MUQ and plotted as the composite sample (total digestion of primary
54 sample), residue (silicates-only), and leach (HNO3 digestion of oxides and carbonates) to compare the patterns and overall concentrations of each different component (Figure 9-A,
B). The REE+Y patterns for each of
the composite samples were roughly
parallel with patterns from both the
residue (lower concentrations) and the
leach (higher concentrations), with
concentrations at an approximate
3 10 20 30 40 average of the residue and leach Ratios from Composite B concentrations. This indicates that the
influence of non-silicate material on
./Y=0.9665x / R'=09988 the trace element geochemistry of
cherts from the Bartlett Dome does not
greatly affect the relative REE+Y
j^% 0 10 20 30 40 distribution of the sample. Ratios from Composite
c • y< Figure 10: Diagrams of important REE+Y ratios from duplicate sample 07-GJB-048C. In all diagrams the square is Y/Ho, the triangle is La/La*, the X is Eu/Eu*, the * is Gd/Gd*, and yS Y=0.9565x the circle is Pr/Lu. (A) Leach ratios vs. yS R'=09982 Composite ratios, with a trendline slope of 1.01. (B) Residue ratios vs. Composite ratios, with a trendline slope of 0.9665. (C) Residue ratios vs. Leach ratios with a trendline slope of 0.9565.The slopes in A and B suggest that the relative REE+Y abundances are not greatly
•yS influenced by the presence of oxides and
10 20 30 40 carbonates, with their primary influence Ratios from Leach reflecting in the overall concentrations of the REE+Y, as shown in Figure 8.
55 This was further examined by plotting 5 different REE+Y ratios (Y/Ho, La/La*,
Eu/Eu*, Gd/Gd*, and Pr/Lu) of the Leach vs. Composite, the Residue vs. Composite, and
Residue vs. Leach of sample 07-GJB-048C (Figure 10). The ratios for each pair are near- equal, with all 5 forming a trend on each diagram with a slope of approximately 1, with very good correlation. The best correlations (and closest slopes to 1) are found in the diagrams of the Residue and the Leach vs. the Composite (Figure 10-A, B). For this reason, the REE+Y patterns from samples that have undergone normal digestion are here used as the true chemistry of the chert.
Initial attempts to classify the patterns into groups proved difficult, as there was no apparent correlation between geochemistry and sample locality or petrography.
REE+Y patterns from samples within individual petrographic groups were compared to one another, but while there was a greater correlation by petrographic group than by sample locality, the patterns were still too dissimilar for this to be a useful means of grouping the samples. Instead, a 2-element discrimination plot using compatible metals was chosen to explore whether any systematic patterns would emerge from the REE+Y.
The rationale behind this discrimination plot is as follows. The source of Ni in
Archean seawater was primarily from the subaerial weathering of mafic rocks, as low pC«2 and pH conditions cause it to be soluble in water (MacFarlane et al, 1994).
Chromium, on the other hand, is highly immobile in water, and is generally only deposited in hydrogenous sediments through the input of detrital material. Although Cr exists in numerous oxidation states, within the known Eh/pH systematics of the Archean ocean it would have been present as the insoluble Cr3+.
56 A /3 • 7
200000 2 + / 1 n^15000 0 a. QL x L Archoan TTG 100000 x L Archean Granite o L Archean rhyolile • L Archean Basalt,'Komatiite
50000 • L Archean Andesite • L Archean GraywacXe • L Archean Cratontc Shale * L Archoan Average UC 200000 Cr (ppb)
3SOOC B y^ Y-5 109x«965.88 3000C >^ R=09969
250W s4 2-2000C 1/ Q. a. Z150(H Y=0 3629x«!313.6 R =0 9462^^.^
1CKHX 2 A A • Ni.'Cr-t 500( A D ONi:Cr 3 ANuCr>1 5000 15000 30000 Cr (ppb) Figure 11: Discrimination diagram of Ni vs. Cr used to identify geochemically different chert groups. (A) Ni vs. Cr diagram after Condie (1993), using average late Archean magmatic and clastic sedimentary rocks to define detrital vs. hydrogenous sourced material. (A-l) The field defined by average late Archean rocks of Condie (1993), with Ni/Cr ratios of 0.42-0.75. (A-2) The field in Ni vs. Cr space in which pure hydrogenous sediments fall, based on the behavior of Cr and Ni in late Archean sediments, with Ni/Cr»l. (A-3) The trendline defined by the main population of cherts from this study, with a slope of 1.109, suggesting mixed influence of detrital and hydrogenous material sources, with a bias towards a hydrogenous origin. All cherts from this study fall in a very small space close to the origin in this diagram. (B) Close-up of the cherts from this study in Ni-Cr space. (B-l) Fourteen samples with Ni/Cr=l (empty diamonds), that fall along a trendline with a slope of 1.109. (B-2) Six samples with Ni/Cr»l (empty triangles), these samples fall in the same region as the predicted ideal pure late Archean non-derital sediment (B-3). Five samples with Ni/Cr 57 Because of the different fluid mobility of these two elements, they are of excellent use for the broad discrimination of the genesis of cherts. A Ni-Cr discrimination plot was first used by Condie (1993) to demonstrate the low Ni/Cr ratios found in late Archean crustal rocks, particularly shales, stemming from high chemical weathering rates of Ni, the disproportionate adhesion of Cr to clay minerals relative to Ni, and the transport of Cr in spinel. Consequently, average late Archean magmatic and clastic sedimentary rocks have Ni/Cr Ni/Cr ratios; samples with roughly equal Ni and Cr (Ni/Cr^l), samples with high Ni and low Cr (Ni/Cr MUQ-normalized REE+Y patterns of samples within these three groups are discussed next, and the further sub-classification of the cherts was attempted using the REE+Y systematics. Group 1: Ni/Cr ~1 Fourteen samples fell along a trendline on the Ni vs. Cr diagram with a slope of 1.1093, indicating that these samples have roughly equal Ni and Cr, albeit with slightly higher Ni than Cr (corroborated by individual Ni/Cr ratios generally slightly higher than 1) (Figure 11-B-l). There is no known or visible cause for the covariance of these two elements. This Ni/Cr ratio is much higher than literature values for average magmatic and detrital sedimentary rocks of late Archean age (e.g. basalt and komatiite average Ni/Cr= 0.45) (Condie, 1993). Samples with the highest absolute concentrations of Ni and Cr also 58 have the highest REE+Y concentrations. Based on the REE+Y patterns, this group is divided into two subgroups; a subgroup of 13 samples with strong LREE depletion relative to the HREE), which contains all but one of the 14 samples on this trend, and a subgroup (one sample) that shows an arched REE+Y pattern, with depletion of both the LREE and HREE relative to the MREE. Given the near-equal Ni and Cr within this group, both hydrogenous and detrital sources played significant roles in the formation of these cherts, with some bias toward a hydrogenous source (as shown by the slightly higher Ni concentrations than Cr). Subgroup 1: LREE-depleted The 13 samples from the first subgroup all demonstrate positive slopes in their REE+Y patterns (PrMuo/LuMUQ = 0.12-0.50), indicating strong light rare earth element (LREE) depletion (Figure 12-A). This subgroup shows variable positive La anomalies (La/La*=1.01-1.65) and small negative to positive Gd anomalies (Gd/Gd*= 0.99-1.14). These samples also demonstrate large positive Eu anomalies (Eu/Eu*= 2.07-6.02), that are up to half an order of magnitude greater than the neighboring REE. Each sample also displays a large positive Y anomaly, demonstrated by elemental Y/Ho ratios between 30.91 and 41.27. These patterns are diagnostic of Archean seawater, possessing features that are commonly characteristic of modern seawater, with the addition of the large positive Eu anomaly, likely caused by the increased hydrothermal flux in the Archean, combined with a much lower oceanic oxidation state (Bolhar et al., 2005). 59 A Equal Ni and Cr . i • . . . . i . . . . . i La Ce Pr No Sm Eu Gd Tb Dy Y Ho Er Tm Yb Lu |B High Ni, low Cr La Ce Pi Nd Sm Eu Gtl Tb Dy Y Ho Er Tm Yb Lu '° [C High Cr, low Ni i—,—,—,—,—,—,—,—,—,—,—,—,—,—,—,—i La Co Pr Nd Sm Eu Gd Tb Dy Y Ho Er Tm Yb Lu Figure 12: REE+Y patterns of solution ICP-MS results, by geochemical group derived from the Ni vs. Cr plot. (A) REE+Y of samples with Equal Ni and Cr (Ni/Cr=l). Patterns from the LREE-depleted sub-group are indicated by empty triangles, and those from the LREE and HREE depleted subgroup by filled squares. (B) REE+Y of samples with high Ni and low Cr (Ni/Cr>l). Patterns from the HREE-depleted sub-group are indicated by hollow squares, the flat REE+Y sub-group by hollow triangles, and the LREE-depleted by hollow circles. (C) REE+Y of samples with high Cr and low Ni (Ni/Cr Sample 07-GJB-046CH shows a dramatically different pattern shape than other samples in this group (Figure 12-A). With a PrMuQ/LuMuQ ratio of 0.45, it still demonstrates LREE depletion, but in truth it also shows depletion of the heaviest HREE; Tm, Yb, and Lu. Consequently, this sample displays an arched pattern arched with its apex at Dy-Y-Ho. The La and Gd anomalies for this sample are both relatively large and positive (La/La* = 1.28; Gd/Gd*=1.13), and it demonstrates a large positive, albeit smaller than the preceding sub-group . Additionally, this sample shows a negligible positive Y anomaly (Y/Ho= 27.54). As discussed in the section on the laser ablation data, this general pattern shape was first interpreted to be of hydrothermal origin, but closer inspection of the higher quality solution data has shown that the shape of the arch peaks too far to the left, and the degree of HREE-depletion (Tm-Yb-Lu) is too extreme. Instead, this pattern is comparable to those observed for the Y-phosphate xenotime in single-mineral geochemical studies (Ondrejka et al., 2007; Fletcher et al.., 2004). In my sample, the REE+Y pattern for xenotime added a binary component to the chert, and a weighted average of xenotime REE+Y with seawater-type REE+Y yielded a mixing ratio of approximately 10000:1 in favor of seawater. This indicates that a very small amount of xenotime in this sample has likely caused interference with an otherwise seawater-type pattern. This has resulted in the peak of the arched REE+Y pattern falling on the same elements as xenotime (Dy, Y, Ho), while the rest of the pattern bears features characteristics of Archean seawater (LREE depletion, positive La, Gd, and Eu anomalies), indicating that this sample too is likely a seawater-derived chert, but with some accessory or diagenetic mineral. The late diagenetic origin for the xenotime is the 61 most likely, as it often forms during late, post-orogenic thermal events at a low temperature, often as a reaction rim on zircon (Kositcin et al., 2003). The Ni-Cr systematics were not been affected by the presence of diagenetic xenotime. Group2:Ni/Cr>l Six samples contained significantly higher Ni/Cr ratios (2.18-43.1) (Figure 11-B- 2). These samples do not form a trend on the Ni vs. Cr diagram, but instead lie in a rough buckshot pattern above the trendline formed by the samples with near equal Ni and Cr. All of these samples fall into the ideal field for pure Archean hydrogenous sedimentary rocks on the Ni vs. Cr diagram. REE+Y patterns from this group appear as 3 distinct pattern types; HREE-depleted, Flat REE+Y, and LREE-depleted. Subgroup 1: HREE-depleted Three of the six samples from the Ni/Cr>l group fit into this pattern category (07- GJB-048E, 07-GJB-037A, and 07-GJB-041B), demonstrating negative slopes with PrMuo/LuMUQ ratios of 1.15, 1.80, and 2.38, respectively (Figure 12-B). These samples display La/La* values of 0.89, 1.00, and 1.01, showing a range from small negative to small positive La anomalies. Large positive Eu anomalies are observed, with Eu/Eu* values of 3.08, 2.97, and 3.00, as well as small positive Gd anomalies on all three samples (Gd/Gd*= 1.07-1.10). The Y anomalies are more variable with the least HREE- depleted sample, (07-GJB-048E, PTMUQ/LUMUQ= 1-15), demonstrating a large positive Y anomaly (Y/Ho= 36.48), whereas the more HREE-depleted samples (07-GJB-037A, 07- GJB-041B, PrMuQ/LuMUQ= 1-80, 2.38) show small negative to essentially no Y anomalies (Y/Ho= 22.60, 27.32). All three of these patterns are broadly spoon-shaped, reaching 62 maximum depletion between Dy and Er, and showing slight enrichment of the heaviest HREE (Tm, Yb, and Lu). These patterns are similar to REE+Y patterns observed from the M4 unit of the Strelley Pool Chert of the Pilbara craton, Western Australia (Allwood et al, submitted). The M4 cherts are associated with 'chert dykes' and chert mounds, which suggests a hydrothermal origin, albeit of low temperature due to the relatively small positive Eu anomalies observed. Consequently, these three HREE-depleted patterns are interpreted to be cherts formed during hydrothermal activity, possibly a process akin to modern sub-seafloor circulation (Paris et al., 1985). Subgroup 2: Flat REE+Y Sample 07-GJB-048B shows a generally flat REE+Y pattern with a PrMuo/LuMUQ value of 0.9, indicating very minimal LREE depletion (Figure 12-B). This sample shows essentially no La anomaly (La/La*=1.01), a moderate positive Eu anomaly (Eu/Eu*=1.84), and a small positive Gd anomaly (Gd/Gd*=1.05). It also has a positive Y anomaly (Y/Ho=31.87). The shape of the pattern is weakly spoon-shaped (like the HREE-depleted patterns), reaching its minimum concentrations between Sm and Dy (with the exception of the Eu and Gd anomalies). Due to its overall spoon-shaped pattern, this sample may be of similar origin to the HREE-depleted samples discussed immediately above, but the very flat overall slope of this pattern contrasts with the negative slope of the HREE-depleted samples. This may have been an argillite of comparable primary composition to the MUQ standard that was subsequently replaced by silica through sub-seafloor hydrothermal circulation, but perhaps to a lesser degree than the HREE-depleted samples and the M4 cherts of (Allwood et al, submitted). 63 Subgroup 3: LREE-depleted Samples 07-GJB-022A and 07-GJB-48H both show LREE-depletion, with PrMuQ/LuMuQ values of 0.37 and 0.38 (Figure 12-B). Sample 07-GJB-022A shows a flat to small positive La anomaly (La/La*=1.01) and 07-GJB-048H shows a large positive La anomaly (La/La*=1.27). Both samples show moderate to large positive Eu anomalies (Eu/Eu*= 2.12 and 4.40), and small negative to positive Gd anomalies (Gd/Gd*= 0.96 and 1.05). Samples also show small to large positive Y anomalies (Y/Ho= 27.42 and 36.19). Overall, both samples show relatively flat patterns on either side of the Eu and Gd anomalies, albeit with the normalized values to the right of Gd being higher than those to the left of Eu. Many of the features of these two patterns indicate that they are likely representative of seawater-derived cherts, particularly 07-GJB-048H, as it possesses the highly diagnostic large positive Y anomaly. The other sample from this group, 07-GJB- 022A, is of somewhat different origin that is difficult to interpret. Broadly speaking, the REE+Y pattern for this sample appears to be of seawater origin, with its LREE depletion, small positive La anomaly, and positive Eu anomaly, but at the same time lacks both the positive Gd and large positive Y anomalies that are critical in identifying seawater- derived REE+Y patterns. It most likely was formed through the mixing of multiple fluid systems of unknown composition and affinity. Group 3: Ni/Cr Five samples fall along a second trendline with a slope of 0.3629 on the Ni vs. Cr diagram (Figure 1 l-B-3). These samples have low Ni/Cr ratios (0.39-0.88), falling in the same compositional regime as average Late Archean magmatic and clastic sedimentary 64 rocks (Condie, 1993). These five samples have been divided into three REE+Y pattern types: HREE-depleted without a Y anomaly, HREE-depleted with a Y anomaly, and LREE-depleted. Literature data for average late Archean volcanic and detrital rocks have comparable, albeit somewhat higher Ni/Cr ratios (Condie, 1993). Due to the high Cr contents (relative to Ni) of the samples within this group, these samples are interpreted to represent cherts with some degree of detrital contribution, from volcanic ash or clay minerals derived from the weathering of volcanic rocks. Subgroup 1: HREE-depleted without a Y anomaly One sample, 07-GJB-031A, shows very high REE concentrations and strong HREE-depletion (PrMuo/LuMUQ= 3.02) (Figure 12-C). The shape of this pattern is simply a negative slope with the exception of a pronounced Eu anomaly (Eu/Eu*= 2.43). This pattern also contains small positive La and Gd anomalies (La/La*= 1.04, Gd/Gd*=1.05). A small negative Y anomaly is also observed in this sample (Y/Ho= 25.79). The shape of this pattern clearly indicates that it does not represent seawater, but it is also quite different from the other HREE-depleted patterns in this study. On the Ni vs. Cr diagram, this sample plots very close to the average late Archean felsic rocks (rhyolite, TTG, granite), as well as on mixing lines with average late Archean graywacke, basalt, and komatiite from Condie (1993). This indicates that this sample is very likely dominated by detrital material, primarily felsic detritus, due to its proximity to the felsic rocks on the Ni vs. Cr diagram. When compared to REE+Y patterns of felsic rocks and average TTG, it transpires that they not only have roughly the same shape, but in fact have very similar concentrations for most elements (not shown). This comes as no surprise, as this sample is very high in a variety of elements that are immobile in water, such as Ti, Cr, Ga, Zr, 65 Nb, Hf, and Th (1-2 orders of magnitude higher than all other samples in some cases), which suggests that this sample is primarily composed of a fine-grained felsic material of a composition akin to TTG; either a felsic ash or a TTG-sourced clastic sediment. Subgroup 2: HREE-depleted with a Y anomaly Sample 07-GJB-048C shows a strong HREE depletion (PrMuQ/LuMuQ= 2.96). Weakly spoon-shaped, the pattern of this sample demonstrates a very minor enrichment in the heaviest HREE (Tm, Yb, and Lu) (Figure 12-C). This pattern has a very large positive La anomaly (La/La*= 1.39) and a small positive Gd anomaly (Gd/Gd*= 1.06). A very large Eu anomaly is present (Eu/Eu* = 4.25), as well as a very large Y anomaly (Y/Ho = 40.86). The cause of this highly unusual pattern is unknown. Apart from the very strong HREE-depletion, this pattern otherwise has the features of a seawater REE+Y pattern; positive La, Eu, Gd, and Y anomalies. This sample has the lowest Ni and Cr values of this group, and also has the greatest deviation from the group's trendline. It is possible, but uncertain, that this sample is primarily a direct seawater precipitate, but with a small contribution of detrital material, causing its low Ni/Cr ratio (0.64) and its HREE- depletion. Subgroup 3: LREE-depleted The 3 remaining samples within this group (07-GJB-048GM, 07-GJB-048GP, and 07-GJB-011A) are all LREE-depleted (PrMuQ/LuMuQ= 0.27-0.48) (Figure 12-C). The patterns show a small positive to no La anomaly (La/La*= 1.00, 1.12, 0.99). Gadolinium anomalies for all 3 samples are also virtually nonexistent (Gd/Gd*= 0.95, 0.96, 1.02). Eu anomalies range from moderate to large (Eu/Eu* = 1.84, 1.91, 3.74) and there is 66 generally a very small positive Y anomaly present (Y/Ho= 28.42, 28.84, 28.27). Two of these samples, 07-GJB-048GM and 07-GJB-048GP both show weakly spoon-shaped patterns, with a slight upswing of the lightest LREE, while sample 07-GJB-011A is flatter overall, with the exception of its anomalies. These patterns are somewhat unusual, as they bear certain features common to seawater, (positive Eu and small positive Y anomalies), but generally lack substantive La or Gd anomalies. Like sample (07-GJB- 048C) with HREE-depletion and a positive Y anomaly, these samples represent a mixture of seawater-derived hydrogenous material as well as the input of some detrital material. Due to the complexity of the classification of the cherts (petrography, Ni/Cr, REE+Y subgroup, and pattern interpretation), the various classifications associated with each sample are presented below (Table 2). Table 2: Relationship between different chert classifcations Petrographic type Ni/Cr REE+Y subgroup REE+Y pattern 07-GJB-011A Carbonaceous <1 3.3 LREE depleted Seawater + unknown detritus 07-GJB-018A Ferruginous =1 1.1 LREE depleted Seawater 07-GJB-022A Chert breccia >1 2.3 LREE depleted Unknown 07-GJB-027B Clean chert =1 1.1 LREE depleted Seawater 07-GJB-031A High Grunerite <1 3.1 HREE depleted w/o Y anom Felsic ash or detritus 07-GJB-036B Chert breccia ~l 1.1 LREE depleted Seawater 07-GJB-037A Clean chert >1 2.1 HREE depleted Low-T hydrothermal 07-GJB-039A Ferruginous =1 1.1 LREE depleted Seawater 07-GJB-039B Clean chert =1 1.1 LREE depleted Seawater 07-GJB-039C Veined Chert ~l 1.1 LREE depleted Seawater 07-GJB-039D Silicified fragments =1 1.1 LREE depleted Seawater 07-GJB-041B Clean chert >1 2.1 HREE depleted Low-T hydrothermal 07-GJB-042B High Grunerite =1 1.1 LREE depleted Seawater 07-GJB-043A High Grunerite =1 1.1 LREE depleted Seawater 07-GJB-045B Ferruginous =1 1.1 LREE depleted Seawater 07-GJB-045C Ferruginous =1 1.1 LREE depleted Seawater 07-GJB-046CS High Grunerite =1 1.1 LREE depleted Seawater 07-GJB-046CH High Grunerite =1 1.2 LREE & HREE depleted Seawater + xenotime 07-GJB-048B Chert breccia >1 2.2 Hat REE+Y Shale + Seawater + Low-T hydro 07-GJB-048C Carbonaceous <1 3.2 HREE depleted w/ Y anom Felsic ash with seawater 07-GJB-048E Clean chert >1 2.1 HREE depleted Low-T hydrothermal 07-GJB-048GM Clean chert <1 3.3 LREE depleted Seawater + unknown detritus 07-GJB-048GP Clean chert <1 3.3 LREE depleted Seawater + unknown detritus 07-GJB-048H Silicified fragments >1 2.3 LREE depleted Seawater 07-GJB-049B Ferruginous =1 1.1 LREE depleted Seawater 07-GJB-049E Carbonaceous =1 1.1 LREE depleted Seawater 67 Discussion Due to the complexity of the problems addressed here, this discussion has been broken up into five sub-sections. Beginning with the geochemistry, which will focus on the relationships between the different anomalies, and their broader implications for the depositional basin, I will then progress logically through the possible sources for silica in this depositional system, the depositonal processes that formed these BIFs and cherts, and how these processes are characteristic of the time gap discussed in Thurston et al. (2008). Based on this analysis, a depositional model for the BIFs of the Bartlett Dome will then be described, followed by a broader discussion of the regional implications of the results found in this study. Geochemical Discussion Chert can precipitate as amorphous silica from virtually any natural water that is saturated with respect to SiC>2. As chert can preserve aspects of the trace element composition of the source water, that natural water can be identified. There is an array of possible water types that have played a role in the formation of the cherts of the Bartlett Dome. Most importantly, they could have formed from seawater, which can itself have an array of potential compositions, depending on how the trace elements are being introduced into the system, as well as the mixing rates and circulation patterns in the ocean basin. Alternatively, the silica may have precipitated from hydrothermal fluids. Like seawater, hydrothermal fluids can have an array of different compositions, variably controlled by the rock types through which the fluids have passed, the temperature of the fluids, the amount of mixing the fluids have experienced with ambient seawater, and the proximity to the vent site. Although it is not a mechanism for silica precipitation, detrital 68 sedimentary material is also a critical factor in determining the genesis of cherts. Even a small contribution of detrital material (e.g. a small volume of fine grained volcanic ash) has the potential to overwhelm many of the trace element systematics that are diagnostic of the various fluid types. Additionally, the input of these materials can introduce elements that are uncommon in most fluid systems, such as the relatively fluid-immobile HFSE. Due in part to the possibility of the mixing of several different chemical sources, detailed study of the REE+Y anomalies and fluid-immobile elements is required to identify the different potential sources. Anomalies and the Y/Ho ratio With detailed examination of the four most important REE+Y anomalies (La, Eu, Gd, and Y), positive correlations become apparent between the cherts of this study. Predictably, the La and Gd anomalies show a fairly strong correlation, as they are both produced in seawater during estuarine scavenging (Lawrence & Kamber, 2006). In two of the geochemical groups, Ni/Cr>l and Ni/Cr Ni/Cr=l group, with the La/La* values being near-constant around 1.00 with increasing Gd/Gd* values until the Gd/Gd* values reach 1.05, an appreciably positive Gd anomaly (Figure 13-A). This most likely marks the tipping point of seawater dominance in these two geochemical groups over other silica and material sources. Interestingly, both the La and Gd anomalies also show a positive correlation with the Eu anomaly (Figure 13-B, C). The positive Eu anomaly in hydrogenous sediments is usually unrelated to the petrogenetically diagnostic La and Gd anomalies, as the Eu anomaly is generally regarded as sourced from high-temperature hydrothermal fluids 69 (>250°C) (Danielson et al., 1992; Bau & Dulski, 1999; Douville et aL, 1999). However, it is widely established that the positive Eu anomaly is also a distinctive feature of Archean seawater itself, due to the combined influences of a reducing ocean and the likelihood of a higher hydrothermal flux during the Archean (Isley, 1995). The positive correlation between the Eu anomaly and other REE +Y anomalies is strongly indicative that the Eu anomalies in this study are in fact derived from Eu-enriched seawater, and not directly from high-temperature hydrothermal vents. The most likely cause of these positive Eu anomalies is a distal hydrothermal vent system, potentially very distal, as Eu2+ from the vent could be transported great distances in reducing Archean seawater presumed to exist below a pycnocline at shallower water depths (e.g. Kappler et al., 2005). The model for long-distance transport of hydrothermally-sourced iron in BIF presented by Isley (1995) may also explain the distal source of the positive Eu anomaly seen here, as Eu and Fe could have been transported great distances in hydrothermal plumes and deep ocean water, as suggested by Deny and Jacobsen (1990). Samples from the Ni/Cr>l group bear smaller Eu anomalies than samples from the other two geochemical groups, which corresponds with their comparatively small La and Gd anomalies as well, matching the overall trend of the entire dataset. This indicates that while the REE+Y patterns of the cherts in this group are likely derived from sub- seafloor hydrothermal circulation, based on their similarities to chert dikes from the Strelley Pool chert (Allwood et al., submitted), the fluids that formed these patterns were cooler than 250°C, and certainly not derived from a black smoker system, due to the relatively small Eu anomaly demonstrated by these samples. The positive Eu anomalies, like the positive La and Gd anomalies in these patterns, are likely relicts of precursor, 70 seawater-derived hydrogenous sediments that were not completely overprinted during hydrothermal circulation, even though the hydrothermal fluids were able to overprint the other common seawater-type features, such as the low PrMuo/LuMUQ ratio, the positive Y anomaly, and have washed out the high Eu anomalies seen in samples from the Ni/Cr^l group. Eu/Eu* Gd/Gd" Figure 13: Crossplots of shale-normalized anomalies. In all 6 diagrams, squares are the Ni/Cr=l group, circles are the Ni/Cr>l group, and triangles are the Ni/Cr Y/Ho ratios (Y anomaly) (Figure 13-D, E, F). The strongest correlation is with the La 71 anomaly, confirming that both of these features have been controlled by the same estuarine scavenging. The Gd and Eu anomalies have weak positive correlations with the Y/Ho ratio. The weak correlation of the Eu anomaly with the Y/Ho ratio results from their mutual derivation from Archean seawater, albeit via different processes. Furthermore, the input of any detrital material would greatly reduce the Y/Ho ratio, while having a minimal effect of the Eu anomaly. The high sensitivity of the Y/Ho ratio to detrital material is also the likely cause of the weak (albeit stronger than Eu) correlation of the Gd anomaly to the Y/Ho ratio. La anomaly and fluid immobile elements To confirm the lack of control the recorded seawater REE+Y pattern features, such as the La anomaly, on material of a potentially detrital origin, the La/La* ratio was plotted against a suite of generally fluid-immobile elements (Th, Zr, Sc, and Cr) (Figure 14). In the diagrams of La/La* vs. Th, Zr, and Sc, the complete data set demonstrates a hyperbolic relationship between the La anomaly and the elemental concentrations of each of these elements (Figure 14-A, B, C). The apparent 3-compnent mixing is formed by the presence of one sample with a negative La anomaly, but this sample can be excluded, as there are no known processes that form a negative La anomaly in cherts. When plotted by individual geochemical groups (Ni/Cr«l, Ni/Cr>l, Ni/Cr (smaller La/La* values), suggesting that material input from non-hydrogenous sources (volcanic or clastic sedimentary) was not a major component in the formation of these two chert groups. 72 This contrasts with the Ni/Cr 200 400 600 800 1000 1200 1400 1600 10000 20000 30000 40000 50000 60000 Th (PPb) Zr(ppb) 5000 10000 15000 20000 25000 30000 Cr(ppb) Figure 14: La/La* ratios plotted against elemental Th (A), Zr (B), Sc (C) and Cr (D) (see text for details). The Ni/Cr=l group is shown as squares, Ni/Cr>l as circles, and Ni/Cr (Figure 14-D), which results from the considerable scatter present in the Ni/Cr~l group, as it includes the many of the samples in this study with high Cr concentrations. This does not change the interpretation for their origin, as these samples do still have slightly higher Ni concentrations than Cr, indicating that the high-Cr samples with high La/La* 73 values are most likely hydrogenous in origin with a small, fine-grained detrital component, causing their high Cr contents. Condie (1993) suggested that late Archean clay minerals were enriched in Cr and depleted in Ni, due to the influence of reducing conditions and high degrees of chemical weathering. Alternatively, a small amount of detrital Cr-spinel could have been deposited in some of the samples with the highest Cr; samples that otherwise have high La/La* values. The Cr-spinel would have had no impact on both the La/La* ratio and the other studied fluid-immobile elements. Consequently, it appears that the high Cr content in some samples with large positive La anomalies may indicate the input of small amounts of detrital material that was Cr-rich, but poor in other trace elements. Ce anomaly Like Eu, Ce has two valance states (Ce3+ and Ce4+). Because of the redox sensitivity of this element, it has been used to constrain the redox conditions of the ancient ocean (Kato et al, 2006). These authors use the negative Ce anomaly in BIF's of the Dharwar Craton to conclude that during the period 2.9-2.7 Ga, the oceans were oxygenated, as Ce4+ is more soluble than Ce3+, resulting in Ce/Ce* values «1.0 in sediments deposited under oxidizing conditions (Kato et al, 2006). The Ce anomaly has not been extensively discussed here, in part because in all three geochemical groups, the Ce/Ce* values (calculated as per Lawrence et al., 2006b), are clustered around 1.0 (Table 1), indicating nominally small positive and negative anomalies, with no pattern relatable to the Ni/Cr systematics employed in this study. 74 : • A '. m • B C m • • • • • 0.9 1.0 1.1 Pr/Pr* Figure 15: Plot of Ce/Ce* vs. Pr/Pr* to determine true Ce and La anomalies, with fields defined by Bau and Dulski (1996). (A) Field of no appreciable La or Ce anomaly, where the bulk of the samples from the Ni/Cr Pr/Pr* values <0.95, indicating positive Ce anomalies (Figure 15-C). Due to the absence of a negative Ce anomaly, it appears that during the formation of the AGB at around 2.7 Ga, the ocean was reducing. This does not support the conclusion of Kato et al. (2006) that the ocean was oxidizing from 2.9-2.7 Ga. It is more likely that the conditions interpreted by Kato et al. (2006) were either a local phenomenon isolated to the Dharwar craton, or an analytical artifact, and are not indicative of worldwide oceanic conditions during the period from 2.9-2.7 Ga. Having established some constraints on the fluid compositions at the time of deposition, some conclusions about the depositional basin can be drawn. Basinal implications Evaluation of the REE+Y anomalies has narrowed the array of potential water types from which these cherts were formed. While the bulk of the samples bear anomalies that are representative of seawater, the variable size of these anomalies within cherts that bear seawater-type REE+Y patterns indicates that the composition of the seawater varied over both time and location. Deep-ocean sediments typically have fairly uniform REE+Y patterns over time (Kamber and Webb, 2001), largely because an open, unrestricted ocean is relatively well-mixed, which results in fairly homogenous REE+Y patterns. While this is the case in the modern ocean, most evidence indicates that similar controls existed in the Archean ocean, and that while some of the chemical constraints of 76 the ocean have changed over time (e.g. 8 O values, sulfate vs. sulfide ocean, calcitic vs. aragonitic ocean), the trace element geochemistry of the ocean has not varied greatly over the course of Earth history (Veizer et al, 1989a, b; Condie, 1997). A restricted basin with at least partially limited water circulation would be less likely to show such homogeneity, and over long periods of time (millions of years) would demonstrate variations in the REE+Y composition of sediments derived from the water column, reflecting temporal changes in the addition of the REE+Y into the basin. All of these sediments would show broadly similar trends in their composition, such as LREE depletion and positive La, Gd, and Y anomalies, but the size of the anomalies, and the amount of LREE depletion would vary with the precise composition of seawater. The most obvious control on the precise composition of the seawater would be the flux of the REE+Y into the ocean. As discussed in the previous section, REE+Y anomalies associated with seawater are formed by the destabilization of organo-metallic complexes during the mixing of fresh and saline water in an estuarine environment. This means that for hydrogenous sediments (e.g. chert) to demonstrate REE+Y patterns with these anomalies, most of the REEs in solution must be derived from an estuary, which in turn requires an exposed landmass. Since the variability of the REE+Y patterns requires a restricted basin, the subaerial landmass from which the estuarine REEs were derived was likely local, suggesting that parts of the AGB were above sea level at the time of chert formation. Additionally, a local subaerial landmass would have facilitated the erosion and transport of small amounts of detrital material (siliciclastic debris), and enabled the detritus to cause the cherts with a detrital component to show Ni/Cr ratios similar to those reported by Condie (1993). This landmass may have been emergent portions of the 77 submarine plateau that became the AGB (Benn &Kamber, 2007). This is more plausible than a large emergent volcano or mountain associated with an island arc, as the relatively flat terrain of an emergent plateau would not contribute significant amounts of clastic detrital material into the depositional basin, whereas a emergent volcano or mountain belt would be liable to deposit great thicknesses of turbidites, as are seen in other GSBs, such as the Yellowknife Greenstone Belt (Yamashita and Creaser, 1999). Significant volumes of turbidites are not found in the AGB prior to the 2690-2685 Ma Porcupine assemblage (Ayer et al., 2002). Such a plateau would, however, expose large volumes of fresh volcanic rock to the atmosphere, which would be susceptible to chemical weathering due to the C02-rich atmosphere of the Archean (MacFarlane et al., 1994), supplying a considerable flux of dissolved material into the basin via rivers and estuaries. High degrees of subaerial chemical weathering are further supported by the preponderance of quartz arenites in many GSBs (Thurston & Kozhevnikov, 2000), as well as the high concentrations of Ni and Cr in Archean shales (Condie, 1993). While this process would typically form subaerial weathering profiles, none have been documented in the AGB. However, such profiles would have been thin (e.g. Thurston & Kozhevnikov, 2000), and thus may not have been preserved, or potentially have been overlooked by previous researchers. While the correlation between the Eu anomaly and the various seawater-derived REE+Y anomalies is easily explained through the near-ubiquitous positive Eu anomaly in Archean hydrogenous sediments, the Eu anomaly observed in the cherts studied here is significantly larger than those previously reported in many other BIFs (e.g. Derry and Jacobsen, 1990). This means that although cherts were not precipitated directly from a 78 high-temperature black smoker system, black smokers were present within the basin at the time of chert formation, thereby explaining the size of the Eu anomalies in many of these samples. As shown by the Ce/Ce* vs. Pr/Pr* diagram of Bau and Dulski (1996), the water in the basin was at least moderately reducing at the time of chert formation, which would have enabled long distance-transport of Eu2+ from areas of active high-temperature hydrothermal activity to the areas where these cherts were deposited. Like variations in the estuarine flux, the volume of hydrothermal output and relative location of the vent sites would have been a major control on the composition of the seawater from which the silica was precipitated, reflected in the variable sizes of the Eu anomaly in different samples. These constraints allow for a first-order description of the nature of the basin in which the BIFs of the Bartlett Dome were deposited. A restricted basin allowed for variation of the REE+Y composition of seawater over time. This variation was largely controlled by the flux of high-temperature hydrothermal fluids and the rate of input of dissolved material through rivers and estuarine mixing, contingent on the size and proximity of a subaerial landmass. The flux of dissolved material from such a landmass would be controlled in part by secular variations in the relative sea level around the plateau, suggesting that the variations in seawater chemistry could be anywhere from several thousand year fluctuations to seasonal changes. Given that the BIFs and cherts of the Bartlett Dome were deposited over periods ranging from 0.5 to 12.6 My, a considerable degree of long-term variation in the composition of the seawater within a restricted basin is both plausible and expected. 79 Silica- sources and depositional processes Based on the interpretations of the above section, as well as the overall REE+Y patterns discussed in the geochemical results section, the potential sources of silica for the formation of the cherts of the Bartlett Dome will now be discussed, and will be compared to similar conclusions reached in previous studies. This will segue into a description of the depositional processes that formed these cherts, BIFs and related sedimentary rocks. The geochemical compositions and REE+Y patterns of cherts from the Bartlett Dome indicate that the controls on chert deposition and the silica sources were far more diverse than previously assumed. The REE+Y patterns establish that the dominant silica source was direct precipitation from seawater, but other material sources such as circulating hydrothermal fluids and detrital material were also important contributing factors in chert formation. Even cherts that demonstrate all of the features of seawater in their REE+Y patterns contain evidence for minor input of non-hydrogenous material, based in part on the elevated concentrations of non-fluid-mobile elements such as Cr and the HFSE. Only the six samples of the Ni/Cr>l group have such low concentrations of fluid immobile elements as to suggest that they are true, pure, fluid derived sediments, an assumption confirmed by their close position on the Ni-Cr diagram to the average jasper from BIFs of the Panorama Formation, Warrawooona Group, Pilbara Craton, Western Australia from Bolhar et al. (2005), whereas the high concentrations of the fluid immobile elements in the Ni/Cr 80 The three petrogenetic processes identified in this study have also been observed in the Pilbara by van den Boorn et al. (2007), using Si isotopes as the principal means of discrimination. This study shows that the use of Ni/Cr ratios is as effective at determining the different material sources, and the REE+Y patterns allow for more nuanced analysis of the interplay between these different sources. Seawater-derived cherts The dominant source for silica in most Precambrian cherts is direct precipitation from the ocean, driven by much higher concentrations of dissolved amorphous silica in seawater (Siever, 1992; Knauth & Lowe, 2003). The absence of silica-secreting organisms (e.g. siliceous sponges, diatoms, radiolarians) in the Precambrian would have resulted in much greater concentrations of silica in the oceans, up to 60-100 ppm (assuming modern ocean temperatures), approaching silica saturation (Siever, 1992). Due to these elevated silica concentrations, direct silica precipitation, as well as early silica diagenesis (e.g. hardground formation, pre-compaction chert nodules, silica-cementation of clastic sediments) would have been facilitated, occurring at much greater rates than observed in Phanerozoic rocks, accounting for the unusually high volume of chert deposits in the Precambrian (including BIF) by comparison. Studies of the 8180 values from early-lithified cherts (nodules and hardgrounds) have shown that Archean ocean temperatures may have been as warm as 55-80°C, greatly increasing the solubility of silica, allowing for silica concentrations of up to 300 ppm (Knauth & Lowe, 2003). Alternatively, the low Archean 8180 values have been explained as indicative of shallower ocean depths, in which the isotopic composition of the ocean was dominated not by oceanic temperature, but by the combined influence of low-temperature 81 hydrothermal alteration of oceanic crust and continentally-derived O (Kasting et al. 2006). Experimental data from Posth et al. (2008) suggests that, assuming the super- saturation of silica was temperature-driven as suggested by Knauth and Lowe (2003), even a small drop in seawater temperature would induce rapid precipitation of silica from the water column. This in turn would cause the co-precipitation of the trace elements from seawater, thereby causing the cherts to reflect seawater compositions in their trace element signatures (Posth et al., 2008). If correct, such a temperature change could have triggered rapid silica production, resulting in thick chert bands that demonstrate seawater- type PvEE+Y patterns. The suggestions of Kasting et al. (2006) also have potential implications for this, as the silica in cherts that bear seawater-type REE+Y patterns would have ultimately been derived from a subaerial environment, and its subsequent deposition in a restricted basin would not have likely been at great water depth. Hydrothermal replacement cherts Many Archean chert deposits are considered to be potentially the direct product of hydrothermal activity (Kato & Nakamura, 2003; Van Kranendonk, 2006; van den Boorn et al., 2007); akin to modern silica vents in the Galapagos (Herzig et al., 1988), modern hydrothermal sinters (Maliva et al., 2005), and exhalative deposits of Phanerozoic age (Grenne & Slack, 2003a; Grenne & Slack, 2003b; Grenne & Slack, 2005; Peter & Goodfellow, 2003). These hydrothermal vent systems formed both sub-seafloor alteration as chert vein stockworks and as exhalative mounds (van den Boorn et al, 2007). As indicated by the size of the Eu-anomalies, true vent systems appear not to have been a direct controlling factor in the formation of the cherts from this study. 82 The only samples that appear to demonstrate a direct hydrothermal component are those in the Ni/Cr>l geochemical group. The REE+Y patterns of these cherts are very similar to those derived from the Member 4 (M4) cherts of the Strelley Pool chert in the Pilbara craton, Western Australia (Allwood et al., submitted). The M4 cherts are heavily veined, suggesting that their formation had a strong sub-seafloor hydrothermal circulatory component. It has been suggested that the source of the circulating fluids for the Strelley Pool cherts was derived from the overlying Euro basalt, as opposed to more typical hydrothermal fluids derived from below (Van Kranendonk & Pirajno, 2004). It is unlikely that most of the cherts from the Bartlett Dome underwent such top-down hydrothermal alteration, as the middle BIF is overlain by felsic volcanic rocks, which would not carry sufficient heat to drive hydrothermal activity. The top-down hydrothermal fluid hypothesis may apply to cherts sampled from the upper BIF, as these are overlain by younger komatiitic rocks, specifically at the Texmont Mine site. Some of the samples collected from this site appear to have been hornfelsed by contact metamorphism from the overlying komatiites, suggesting that samples from this site in the Ni/Cr>l group (e.g. 07-GJB-037A) may have been hydrothermally silicified by this process. Such samples from the middle BIF, however, are more likely to have been silicified by fluids circulating from underlying strata. The silicification process would have been very similar; low temperature (<250°C) fluids percolating through the sediments, causing the silicification of the most porous, least lithified sediments, and the brecciation of any impermeable chert hardgrounds that had formed (Paris et al., 1985; Kato & Nakamura, 2003). Detritally-influenced cherts 83 During periods of prolonged chert deposition, small amounts of siliciclastic or volcanic (detrital) material can be co-deposited in cherts. Typically, this detrital material would be volcanic ash or suspended clay particles. In addition to the co-deposition of chert and detrital material, fine-grained detrital and volcanic material can be gradually replaced (silicified) by the precipitation of silica from fluids such as seawater. Over the last few decades, many researchers have observed features in Archean cherts that are potentially indicative of both processes. The earliest observations were textural studies of silicified komatiites (now cherts) in the Barberton Greenstone Belt, South Africa (Lowe & Knauth, 1977; Duchac, 1986; Duchac & Hanor, 1987; Lowe, 1999). The main basis for the silicification interpretation was the preservation of komatiitic textures, e.g. spinifex, within the cherts. Some studies attempted to confirm these interpretations through major-element geochemistry (Duchac, 1986), and mass-balance calculations documenting mobilization of MgO and its replacement by Si02 (Hanor & Duchac, 1990). This approach has merit, as the original textures have been preserved, allowing for estimation of the original bulk composition. High-precision trace element analysis could greatly improve these interpretations through the use of more immobile elements to identify the role of various depositional and alteration processes. These silicified volcanic flows are very different from discrete chert bands, such as are found in BIFs, where the volcanic component is more likely to be either in the form of a distal ash fall or as fine, reworked detrital material from suspension. Consequently, silicified flows are not a very good analogue for detritally-influenced cherts in BIF. Some recent work (e.g. Sugitani et al., 2002) has focused on the chemical identification of detrital material in cherts. A low-volume, fine grained detrital 84 component in a chert is unlikely to be texturally visible, and may have an insignificant influence on some of the major-element oxides (e.g. AI2O3). Instead, these studies have used trace element compositions, which are much more sensitive to a detrital contribution. Sugitani et al. (2002) detected a very small amount of redeposited komatiitic detritus in carbonaceous cherts of the Nickel Well chert of the Pilbara craton, Western Australia using immobile element ratios and some trace transition metals. Due to high detection limits on their analyses, it appears that the data quality used in that study is an issue, considering the inherently low concentrations of most trace elements in cherts. Additionally, any conclusions derived from the REE patterns from these rocks are compromised by the lack of data for Y. Another study of a suite of Pilbara cherts also found small contributions of mafic and felsic material into the formation of these cherts at a sub-microscopic scale, but their conclusions are of limited significance, as they are based on the analysis of only 6 samples (Orberger et al, 2006). There are multiple indications of a detrital component in some of the cherts from the Bartlett Dome. The best evidence are the high Cr and HSFE concentrations of the samples in the Ni/Cr Generally, the REE+Y patterns for this group do not aid in the identification of the source of the detritus, with the exception of sample 07-GJB-031A, which has a REE+Y pattern that is very similar to average late-Archean felsic rocks from Condie (1993). The 85 presence of detrital material in these cherts is also shown by non-geochemical evidence. Four chert samples, 07-GJB-039A, B, C, and D were combined and crushed. From that combined sample, a population of 20 zircons was extracted. These zircons provide firm proof of a small detrital volcanic input into these cherts. These zircons also demonstrate inherited zircon cores with ages older than any volcanic rocks exposed in the AGB, a feature of volcanic zircons in the western part of the belt (Ketchum et al, 2008), indicating that they were introduced into the cherts by volcano-sedimentary, and not hydrothermal, processes. However, from a trace element perspective, these four samples do not appear to have a significant detrital component, falling on the trendline for the samples with Ni/Cr»l, and demonstrating typical seawater REE+Y patterns. This suggests that the majority of cherts from the Bartlett Dome may contain a small amount of detrital material, but in many cases not enough to significantly influence the trace element composition, including highly sensitive measures of detrital vs. hydrogenous input, such as the Y/Ho ratio. Characterization of the time gap This brief section will integrate the geochronology with some sedimentological structures, such as heterolithic debris flows. Additionally, the relationship of the REE+Y systematics with some of the most distinctive sedimentological features of the cherts will be examined here. This will help to further constrain the rate of deposition of these BIFs, thereby helping to explain their role in the time gaps described in Thurston et al. (2008). The time gap (0.5-6.7 My) established by the U-Pb dating indicates a slow estimated depositional rate of 0.007-0.1 mm/yr for the Middle BIF at the McArthur 86 Powerline (based on the measured 50 m thickness). These calculations assume continuous deposition during that interval, but sedimentological and geochemical evidence summarized below indicate the possibility of episodic rapid deposition. Coarse debris flows have been reported in several of the thicker BIF exposures (Houle et al., 2008; Thurston et al., 2008). These flows are composed of resedimented material of mixed composition, including chert clasts from the BIFs with which they are associated. The instantaneous deposition of these debris flows punctuated the slow sedimentation of the BIF itself, and accounts for -10 m of the 50 m of exposure at the McArthur Power line locality. This means that the thickness of these debris flows must be excluded from the calculated average sedimentation rate (Dott Jr., 1996). Recalculation of the sedimentation rate based on the remaining 40 m of BIF reduces it to 0.006-0.08 mm/yr. These slower depositional rates still may not be representative given some of the sedimentary structures observed, particularly when considered in conjunction with the REE+Y geochemistry of the associated cherts, but are the most precise possible calculated sedimentation rates at this point. Among the best sedimentological evidence for slow sedimentation is the common occurrence of stratabound chert breccias. The tabular chert fragments in Fe-mineral and phyllosilicate matrices indicate that the chert clasts were likely individual chert bands prior to brecciation. Because they have been brecciated, it appears that these chert beds were more lithified than the adjacent sediments at the time of brecciation, indicating that the chert bands were potentially hardgrounds bounded by unlithified sediments, although the less common 'oblate' chert breccias may have been firmgrounds, allowing for the chert clasts to be rounded. These hardgrounds and firmgrounds were formed during 87 quiescent periods through seafloor diagenetic lithification and partial cementation. This explains the seawater-type REE+Y patterns observed in some of these breccias. What likely followed was the escape of fluids from lower, unlithified strata that caused the brecciation of the chert hardgrounds and firmgrounds. It appears likely that some of these escaping fluids may have been over-pressured by circulating hydrothermal fluids, as some other sedimentologically identical chert breccias show hydrothermal-replacement REE+Y patterns, likely triggered by hydrothermal overprint after brecciation. Similar seafloor diagenetic silicification is reported for the iron formations of the Hamersley basin (Krapez et al, 2003). Unfortunately, the sedimentation rates for the BIF cannot be refined through the subtraction of the thickness of these units, as brecciation has eliminated their primary thickness, and many potential hardgrounds may not have been brecciated and are consequently may no be distinguishable from other chert bands. Further evidence of a highly variable, punctuated depositional history for the BIFs of the Bartlett Dome is the presence of sedimentary structures that indicate soft-sediment loading. In multiple locations, flame structures and ball-and-pillow structures were observed at the interface between chert and iron bands. These loading structures are typically associated with rapid sedimentation causing water-rich, muddy (in this case, Fe- bands) strata to penetrate into overlying, less water-saturated strata (Nichols, 1999). These structures indicate that there were periods of fairly rapid deposition in the formation of these BIFs, when little or no hardgrounds were formed. Analyzed cherts associated with this type of syn-sedimentary deformation typically demonstrate seawater- type REE+Y patterns, suggesting that periods of rapid sedimentation were largely driven by precipitation of cherts from seawater, possibly caused by a drop in temperature (Posth 88 et al, 2008). During these same periods there would likely have been a small amount of background sedimentation of fine grained siliciclastic material, contributing small amounts of detritus such as the zircons extracted from the combined sample 07-GJB-039. Depositional Model The synthesis of the ideas presented in the earlier sections of this discussion allows for the development of a depositional model for the BIFs of the Bartlett Dome primarily focusing on chert formation. Considering what has been established about the basin through the geochemistry, several caveats must first be established for the model. The variable extent of the four REE+Y anomalies makes a strong case for a non-constant REE+Y composition of seawater, which is indicative of deposition in a restricted basin as opposed to the open ocean. This means that any model proposed below is highly specialized for the formation of BIFs of the AGB, and would require adaptation prior to application to other GSBs. Furthermore, the envisaged restricted basin must have been highly starved for clastic sediment; it is unlikely that the complex variation and interplay of depositional and diagenetic processes over time would have been possible with a high detrital input. A volcanic hiatus marked by a high volume of detrital sedimentation would be more likely to resemble the turbidite-rich Yellowknife Greenstone Belt of the Slave Craton, Northwest Territories, Canada (Yamashita and Creaser, 1999). Several important issues that have not been addressed above, such as the dual role of different processes on a single chert band, are addressed below prior to the description of the depositional model itself. 89 In a thick sedimentary package, chemical and sedimentological fluctuations through the stratigraphy would be unsurprising, and indeed expected. With that in mind, the thickest exposures of BIFs of the Bartlett Dome are only a few tens of meters thick, which raises the question as to why there would be so much geochemical variability between individual chert bands in a relatively thin unit. This variability suggests that the BIF depositional system was under the influence of a complex interplay of sedimentary and fluid systems. Initial deposition involved the co-deposition of clastic sediments (mainly reworked volcanic detritus) and the precipitation of silica from seawater, both later influenced by circulating hydrothermal fluids and top-down seawater-driven silicification. Seawater-type REE+Y anomalies observed in a wide variety of overall pattern shapes are a clear indication that nearly all of the cherts analyzed formed through at least some direct hydrogenous precipitation. The strongest evidence for this (discussed above), is the association of seawater-generated anomalies (La, Gd, and Y) with each other and with the Eu anomaly in all three geochemical groups. For the detrital (Ni/Cr Some of the detritally-influenced samples demonstrate features such as positive Y and La anomalies that are usually reserved for seawater (Lawrence and Kamber, 2006). A chert sample with a strong detrital component should not demonstrate a positive Y anomaly, as the near-constant Y/Ho ratio in magmatic rocks would tend to reduce the Y/Ho ratio to normal values for terrestrial rocks (Pack et al., 2007). Several authors have suggested a process of top-down silicification, in which volcanic or siliciclastic material 90 is replaced by silica derived from overlying seawater that may help explain this apparent contradiction in some cases (Krapez et al., 2003; Pickard et al., 2004; van den Boom et al, 2007; Thurston et al, 2008). This concept suggests that the silica-saturated seawater would cause progressive cementation and replacement of existing fine grained sediments from the sediment-water interface downwards into the sediment package. The top-down silicification model also applies to one sample (07-GJB-048B) in the hydrothermal-replacement group (Ni/Cr>l), as it shows a generally flat REE+Y pattern (PFMUQ/LUMUQ ~1), akin to a shale, but with a small positive Y anomaly. This sample may have been clastic sediment that underwent very early (surface) seawater diagenesis, and was cemented by silica into a siliciclastic hardground, acquiring small positive La and Gd anomalies, and a larger positive Y anomaly. This particular sample is a chert breccia, suggesting that it acquired its slightly spoon-shaped REE+Y pattern from hydrothermal fluids during brecciation and alteration. Other samples within the high Ni/Cr group also show positive Y anomalies (07-GJB-048E, 07-GJB-048H), indicating the possibility that these samples were in fact seawater-precipitated, unlithified silica gels that were later overprinted by sub-seafloor hydrothermal circulation. With these variables in mind, the depositional model for the BIFs of the Bartlett Dome can be structured as follows. At any given time at the sediment-seawater interface, one of four depositional processes is occurring: 1) The dominant sediment type consists of 'typical' BIF composed of thinly interbanded chert and Fe-bands. Although the bulk of the sedimentation occurred 91 through this process, it did not encompass the bulk of the depositional interval (Figure 16-A). 2) During the bulk of the period of BIF formation, very little to no sedimentation occurred, referred to here as periods of null-deposition. During these periods, interaction of silica-supersaturated seawater with the pre-existing surface sediments causes the cementation and replacement of these sediments, regardless of their original composition, with seawater-derived silica, forming hardgrounds that bear seawater-derived geochemical signatures (Figure 16-B). 3) At other times, episodes of rapid silica precipitation and sedimentation deposited large volumes of silica at the seafloor, promoting the formation of load structures between beds (Figure 16-C). 4) The final controlling surface process was the near-instant deposition of heterolithic, coarse-grained debris flows, which often entrained the uppermost layers of the BIF within them, creating tabular chert clasts that are a common feature of these deposits, along with clasts of volcanic and sulfidic origin (Figure 16-D). The principal unifying feature of all four of these depositional processes would have been the very slow, steady input of very fine grained clastic and volcanic material, likely dominated by very distal volcanic ash-falls, and the transport of fine detritus from subaerial regions of the AGB plateau. 92 F Deposition of thinly interbanded chert and Fe bands Null-deposition "Typical" BIF deposition Water column Water column Siliceous seawater Precursor sediments Rapid Silica precipitation and sedimentation Instantaneous depositional events Water Column Water column £75CC Rapid precipitation of silica Untithified hydrogenous silica Flame structures ^ Ball-and-pillow structures Unlithified Fe-mineral band Underlying sediments with unrelated structures Bedded Chert and BIF Hydrothermal brecciation Hydrothermal silicification by fluid overpressure Water column Unaffected sediments Water Column Unlithified Fe or clay minerals Impermeable hardground CD- 'Lr=fc C3 , , •> t=i ^ •••c I I-' W 1=1 :t Silica Replacement Unlithified Fe or clay minerals"**. Siliceous Hydrothermal Fluids Over-pressured hydrothermal fluid reservoir 93 Figure 16: Depositional model for the BIFs of the Bartlett Dome. The scale of all of these models is highly variable, ranging from 1 cm up to 10 m, even for the same process. (A) Deposition of thinly interbanded chert and Fe-bands, characteristic of 'typical' BIF. (B) Periods of null-deposition characterized by the silica replacement of surface sediments by a silica-supersaturated water column, forming hardgrounds with irregular lower boundaries relative to the precursor sediments. (C) Rapid silica precipitation and sedimentation. Driven, for example, by a drop in the water temperature from 75°C, silica is rapidly precipitated from the water column, causing load structures to form by the overloading of water-rich amorphous silica over unlithified Fe-bands. (D) Instantaneous depositional events such as the rapid deposition of a heterolithic debris flow make up a significant thickness of most BIFs of the Bartlett Dome, thereby decreasing the depositional rate of the rest of the BIF. During deposition, these debris flows entrained chert and other materials from the underlying BIF, thereby erasing evidence of the processes that immediately preceded them. (E) Sub-seafloor hydrothermal circulation model after Paris et al. (1985). Silica-rich hydrothermal fluids (dotted lines) penetrated the sediment package, replacing some material with silica. The upward movement of these fluids is halted by the presence of impermeable hardgrounds formed in (B). (F) Over-pressured hydrothermal fluids from (E) eventually reach pressures necessary to cause the brecciation of chert hardgrounds, forming stratabound and podiform chert breccias with matrices of Fe and clay minerals. The necessary pressure would have varied between events, contingent on the thickness and lateral coherence of the chert hardground. See text for further details. Occurring at the same time as these sediment-water interface processes are sub- seafloor processes dominated by the circulation of relatively low-temperature, silica-rich hydrothermal fluids: 5) These fluids are transported upwards through the sediment package causing the silicification of the most permeable fine-grained sediments, and they are also trapped by hardgrounds formed during the periods of null-deposition (Figure 16- E). 6) Over-pressuring of the hydrothermal fluids below impermeable hardgrounds would cause violent water-escape events, brecciating the hardgrounds. These hydrothermal fluid circulation processes are responsible for many of the textures observed in the BIFs, and were the final stage of chert formation as preserved today at any particular stratigraphic level (Figure 16-F). These cycles would not necessarily happen in any particular order, and as indicated by the lateral variability of the BIFs across the Bartlett Dome, may well have 94 happened contemporaneously in different parts of the basin, depending on local conditions. The alternation of these different processes would have continued until the resumption of volcanism in the region; however it must be noted that hydrothermal silicification and brecciation may have continued during the emplacement of the overlying volcanic package. This depositional model provides a sound explanation of the geochronological problems observed in the AGB through the integration of sedimentological and geochemical evidence for the processes described here. Regional Implications The results of the geochemical and sedimentological study have provided a detailed model for the deposition of the BIFs of the Bartlett Dome, as well as establishing some general constraints on the basin. These have a number of broader implications for future study of the AGB, largely stemming from the chronostratigraphic discrepancy of the AGB. The geochronological/stratigraphic problem in the AGB has been well explained (Thurston, 2002; Thurston et al., 2008; Houle et aL, 2008). Based on the duration of main-stage volcanism in the Abitibi (2750-2696 Ma), there should be a much greater volume of volcanic material present in the belt, assuming that modern rates of volcanism hold true for the Archean (Thurston et al., 2008 and references therein). Further compounding this issue is the conformable contact between assemblages of disparate ages, the Deloro and Tisdale, in areas such as the Bartlett Dome. The contact between these two assemblages is marked by the Upper BIF, which is underexposed, except at the Texmont Mine. To better constrain this disparity between rates of volcanism and 95 sedimentation of BIF, this study concentrated on the Middle BIF, the best exposed, most continuous, and best constrained by geochronology of the three BIF units in the upper Deloro assemblage (Houle et al, 2008). Because mapping has demonstrated that all three Deloro assemblage BIF units have comparable mesoscopic textures and structures, it is believed that the detailed study of the Middle BIF reported here should be applicable to all three Deloro assemblage BIF units in the Bartlett Dome, particularly given its geochemical similarities to the Upper BIF. Episodic deposition of all present sediment types alternated with lengthy periods of null deposition, interaction of seawater and seafloor-circulatory fluids with the sediment package was commonplace and often pervasive. Fluctuations in the temperature of the seawater may have provided one of the dominant controls on the rate of silica deposition. Simultaneously, volcanic activity elsewhere in the AGB (where the 2723-2720 Ma Stoughton-Roquemaure and 2719-2711 Ma Kidd-Munro assemblages were being deposited during this time) (Thurston et al., 2008) supplied occasional influxes of detrital material, both as fine-grained suspended material (e.g. very distal volcanic ash fall), or by the deposition of coarse-grained debris flows. At the onset of Tisdale assemblage emplacement, volcanic activity resumed in this region, ending the deposition of the BIF. Therefore, volcanic hiatuses may be able to be identified throughout the AGB based on the presence of BIFs and other sedimentary rocks, particularly chert breccia horizons within BIF units. These BIFs accommodate the depositional gaps presented by the geochronology of the belt. The possibility for these volcanic hiatuses and the associated BIFs to be useful exploration tools for VMS deposits has already been discussed (Thurston et al., 2008), as BIFs and chemical sediments have been noted in association with other VMS deposits in 96 the AGB (Lafrance et al., 2000), and have been employed as distally equivalent to VMS deposits in other mining camps, such as Bathurst, New Brunswick (Peter & Goodfellow, 2003). However, as there are no known VMS deposits in the region of the Bartlett Dome, further study in areas of known VMS deposits within the Abitibi, preferably of Deloro- age, would be necessary to properly evaluate such relationships. By studying the sedimentology and trace-element geochemistry of the BIFs of the Bartlett Dome, several new genetic constraints for research in the AGB have been established. The depositional model presents a new set of much more complex processes that controlled BIF formation. Through establishing these new parameters for Algoma- type BIF formation in the AGB, fundamental changes in how the stratigraphy of the belt is studied are imperative for future studies of the region. Conclusions In the effort to characterize the volcanic time gap observed in the Deloro Assemblage of the AGB, it has been shown here that: 1) Field studies have revealed a wide variety of sedimentary features at the meso to macroband scale, particularly within chert bands. These rock types and structures, including heterolithic debris flow units, chert hardgrounds, stratabound chert breccias, nodules, and other early diagenetic textures, all suggest a prolonged pre- lithification history for the BIFs of the Bartlett Dome, including considerable variations in the sedimentation rate. 97 A limited amount of correlation between the petrography and field observations exists due in large part to the difficulty in preserving microtextures in cherts that have been recrystallized to polygonal quartz. However, several petrographic groups were established based on distinctive textures and accessory minerals. Some of these textures are suggestive of the replacement of fragmental precursor material, as well as finely-bedded, fine-grained sediments, while others suggest the incorporation of aluminous detrital material. Recrystallization and silica replacement of detrital material would have the effect of homogenizing the detritus with the rest of the chert, thereby rendering it texturally invisible, while still apparent in the trace element geochemistry of the chert. The U-Pb ages of overlying and underlying volcanic rocks indicate a depositional interval of 0.5-6.7 My for 50m of BIF, with effective depositional rates of 0.006- 0.08 mm/yr, suggesting that the BIFs were deposited very slowly, or in brief, rapid pulses separated by lengthy periods of very slow, no deposition, or erosion. The relative paucity of detrital sedimentary rocks deposited during these depositional periods suggests that despite the prolonged absence of volcanic activity, the chemical weathering, not mechanical weathering of the volcanics was dominant, which suggests a weathering environment with low topographic relief, such as an emergent or shallow portion of a submarine plateau. LA-ICP-MS trace element geochemistry has been shown to aid in distinguishing chemically different, adjacent chert bands, but low analytical precision limits the use of this technique. Sample 07-GJB-046C was split into geochemically different chert bands of seawater and hydrothermal origin based on the results of this 98 technique. With solution-based ICP-MA analysis, the hydrothermal chert of this sample proved indeed to be a seawater-derived chert that was chemically overprinted by the Y-phosphate xenotime. Total digestion ICP-MS revealed three different geochemical groups based on their Ni/Cr ratios (Ni/Cr=l, Ni/Cr>l, and Ni/Cr Review of the variability seen in the shale-normalized REE+Y anomalies (La/La*, Eu/Eu*, Gd/Gd*, and Y/Ho) has revealed that these cherts were deposited in a restricted basin under reducing conditions, with local sources for both estuarine- derived and hydrothermally-derived dissolved material. Additionally, it has been shown that the estuarine flux of material into the basin was most likely dominated by the chemical weathering of a dominantly mafic landmass, causing the Abitibi ocean to have a greater degree of LREE depletion relative to the HREE, and supporting the postulation that the local landmass was most likely an emergent portion of the AGB oceanic plateau. Based on the mapped textures and structures and the material sources identified by the trace-element geochemistry, a new depositional model can be proposed for Deloro Assemblage Algoma-type BIFs in the Bartlett Dome area, of the Abitibi greenstone belt. This model is characterized mainly by periods of prolonged, slow to null deposition punctuated by the rapid deposition of hydrogenous cherts and coarse debris flows and fine detrital material. The unlithified sediment deposited during these intervals of alternating rapid-slow-null deposition and subsequent 99 fluid-rock interaction events, thereby causing some of these cherts to develop chemical features that are diagnostic of multiple material sources. This stands in sharp contrast with the majority of past assumptions about BIF's origin as a direct seawater precipitate, largely based on the study of the chemically and texturally uniform Superior-type BIF, as opposed to the thinner, more common, and more complex Algoma-type BIF, as seen in this study from the AGB. 7) This model helps to explain the time gaps observed in the U-Pb geochronology of the volcanic rocks of the Western Abitibi, particularly the issues surrounding the contact between the Deloro and Tisdale assemblages. It seems likely that this model may be applicable to future studies of BIFs elsewhere in the AGB, and possibly to BIFs and cherts in certain Archean greenstone belts worldwide. 100 Chapter 3: Summary and Recommendations Summary The banded iron formations (BIFs) of the Bartlett Dome region of the Abitibi greenstone (AGB) belt occur at three distinct stratigraphic levels within the Deloro assemblage. These BIFs were deposited during periods of volcanic quiescence ranging from 0.5 My to 12 My. These gaps are significant to the overall stratigraphy of the AGB, as they indicate that near-continuous deposition was characteristic of the emplacement of the Kewatin-type assemblages, explaining the disparity between the calculated volume of volcanic rocks during the 54 My interval of pre-successor basin volcanism (based on modern rates of volcanic flux) and the actual volume of observed in the belt. The problem with this observation is that the BIFs of the Bartlett Dome are no more than 50 m thick, suggesting that either slow deposition or periods of no deposition were major controlling factors in their formation. By studying the sedimentology and trace element geochemistry of the chert bands of these BIFs, it has been shown here that the proposed rates of deposition held true, causing a complex variety of textures and trace-element compositions that are most visible within the chert bands of these BIFs. Sedimentological study of the cherts from the Bartlett dome has shown a wide variety of unusual textures and features, with little relation to the mineralogy of the associated Fe-bands. Among the features observed were chert hardgrounds and firmgrounds with evidence for the silica replacement of precursor sediments. These hardgrounds have sharp upper contacts, but have undulatory, diffuse lower contacts with the Fe-bands. Many of these chert hardgrounds have disseminated, coarse magnetite crystals in them, suggesting the preferential replacement of fine grained material by silica 101 over coarser precursor material. Related to these hardgrounds and firmgrounds are stratabound chert breccias, consisting of thin, tabular chert clasts, and large, rounded clasts, suggesting that hardgrounds of varied degrees of lithification were brecciated during fluid escape events. Despite the above listed indicators of slow deposition, load structures commonly associated with rapid deposition were observed. These load structures indicate that despite their appearance in cherts, they were formed during periods of very rapid deposition, perhaps by direct chert precipitation driven by a decrease in the temperature of the water column (Posth et al., 2008). However, due to the known density difference between iron oxy-hydroxides and amorphous silica gels, the exact mechanism for the formation of these loading structures is unclear. Other features include the common presence of coarse-grained, heterolithic debris flows, indicating that there were events of near-instantaneous deposition, further allowing for prolonged periods of hardground formation and other slow processes. Thus, based on the sedimentological and stratigraphic evidence, it appears that the deposition of the BIFs of the Bartlett Dome was characterized by a prolonged period of very slow to no deposition, punctuated by rapid to instant depositional events. Despite the distinctive textures and structures observed at the outcrop scale, the petrography of cherts proved to be of little use, and have a somewhat limited relationship to field observations. This is likely due in part to the greenschist-facies metamorphism of the region that has caused the recrystallization of most the cherts from microcrystalline silica, to polygonal quartz. Consequently, the bulk of the depositional and diagenetic textures that may have been visible in the cherts at one time are very limited. That said, certain features visible in outcrop do appear in thin section, albeit the most resistant 102 features, such as carbonaceous and ferruginous cherts, as these features change the appearance of the chert bands. As a result, the primary virtue of the petrography of these cherts was for the purpose of selecting areas of thin sections for in situ geochemical analysis by LA-ICP-MS. The use of in situ trace-element geochemical analysis by LA-ICP-MS has proven excellent for the identification of geochemically distinct chert bands, however the limits on the precision of this method indicate that it cannot be used as the sole analytical method for cherts. As a result, it is necessary to analyze the chemically distinct bands identified through this method by high-precision, total dissolution ICP-MS. Use of the more traditional method has resulted in virtually incomparable data quality, with precise, low detection limit geochemical results into the parts per trillion. The high-quality data has enabled the identification of 3 broad geochemical groups, identified chiefly through their Ni/Cr ratios; mixed seawater and detrital material (Ni/Cr^l), pure non-detrital cherts (Ni/Cr>l), and cherts with a strong detrital component (Ni/Cr 103 The chert sedimentology and trace element geochemistry integrate very well, enabling the construction of a depositional model consisting of a logical succession of processes. Chert hardgrounds were formed through interaction of silica supersaturated seawater with the sediment exposed at the seafloor during periods of very slow to essentially no deposition. These periods of no deposition and hardground formation are juxtaposed with strata indicative of periods of rapid deposition. One of the latter are the load structure-bearing cherts, all of which demonstrate seawater-type REE+Y patterns, indicating that they formed from the direct precipitation of amorphous silica from seawater. Typically, silica is not known to precipitate rapidly enough to form load structures, but it has been suggested that due to the super-saturation of the Archean ocean with respect to amorphous silica, a drop in temperature would promote the rapid precipitation of amorphous silica that would later become chert. The other rapid deposition process is the essentially instantaneous deposition of heterolithic debris flows, decreasing the total thickness of BIF deposited in essentially the same period of time, allowing for even slower deposition rates when BIF thickness is adjusted for their presence. Finally, the other primary controlling process in the formation of BIFs of the Bartlett Dome was sub-seafloor circulation of low temperature hydrothermal fluids, which would form chert veins and the pervasive silica replacement of sediments with sufficient permeability to allow for fluid movement. Hardgrounds formed in the first process mentioned above would form an impermeable barrier, sometimes resulting in the overpressure of the hydrothermal fluids below these hardgrounds and their subsequent brecciation in fluid escape events. This explains why, although the chert breccias are expected to demonstrate seawater-type patterns due to their origin as hardgrounds, some 104 have been chemically overprinted by these hydrothermal fluids. These processes together have resulted in the formation of very complex BIFs in this region, and their interaction and close spatial relationships were promoted by prolonged periods of limited material supply during volcanic quiescence, punctuated by brief periods of more rapid deposition. Recommendations The implications of the depositional model for the BIFs of the Bartlett Dome have great significance to the understanding of the AGB as a whole. If the sedimentological and geochemical features observed in the cherts studied here can be observed in BIFs in other regions and within other assemblages, then the paradox of the calculated volcanism vs. the actual volume of volcanic rocks within the belt can be fully accounted for and put to rest. Furthermore, the study of the BIFs along the contact of the Deloro and Tisdale assemblages in other areas, such as the Round Lake Batholith region near Kirkland Lake, Ontario through this lens has the potential to confirm the conclusions reached here that these BIFs can represent the entire period of deposition of the Stoughton-Roquemaure and Kidd-Munro assemblages in areas where the Deloro and Tisdale assemblages are in 'conformable' contact. The established stratigraphy of other greenstone belts worldwide suggests that this model may be applicable outside of the AGB as well. For example, the Barberton greenstone belt (South Africa) and the greenstone belts of the Pilbara Craton (Western Australia) both show that each stratigraphic group is capped by cherts, highly silicified volcanic rocks, or BIF, suggesting that the processes observed here in the Bartlett Dome may have controlled the periods between separate volcanic phases in these other 105 greenstone belts. The use of these methods and model would be necessary to confirm this hypothesis. One other suggestion for the possibility of future research on this topic is to study the relationship of such BIFs and cherts with known VMS deposits. Several major VMS deposits in the Eastern AGB, Quebec, are associated with rocks of comparable age to the BIFs of the Bartlett Dome, such as the Sel Baie, Joutel, Normetal, and Mattagami deposits. Study of associated BIFs and distal chemical sediments along strike from these deposits may reveal a relationship between these depositional processes and the VMS deposits. It is important to note that, at present, the only known correlation between the BIFs studied here and any known VMS deposits is the age relationship, and there is little other basis for such speculation from this study. 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