Petrology and Geochemistry of Drill Core AT-14-01

Home , Abitibi greenstone belt, Amphibolite, Anorthosite, Arsenopyrite, Banded iron formation, Diabase

FINAL REPORT

prepared for

Canadian Continental Exploration

Prepared by

Alexander Kawohl & Hartwig E. Frimmel

Bavarian Georesources Centre Dept. of Geodynamics and Geomaterials Researc Institute of Geography and Geology University of Wuerzburg Am Hubland D-97074 Wuerzburg Germany

March, 2018 A. Kawohl (MSc) is a project scientist.

Prof. Dr. Hartwig E. Frimmel holds the Chair of Geodynamics and Geomaterials Research. He is a Competent Person who is a Fellow of the Geological Society of South Africa (GSSA), Fellow of the So- ciety of Economic Geologists (SEG) and an Executive Council Member (incl. Past President) of the So- ciety for Geology Applied to Mineral Deposits (SGA). He has close to 30 years of experience that is rel- evant to the topic of this research project and to the activity being undertaken to qualify as a Competent Person as defined in the 2012 Edition of the Joint Ore Reserves Committee (JORC) ‘Australasian Code for Reporting of Exploration Results, Mineral Resources and Ore Reserves’ as well as the ‘South Afri- can Code for the Reporting of Exploration Results, Mineral Resources and Mineral Reserves’ (the SAMREC Code).

Table of Contents

1. Preamble...... 2 2. Geological Background ...... 4 Location ...... 4

Temagami Greenstone Belt ...... 4

Huronian Basin ...... 5

Sudbury Igneous Complex...... 5

Other Magmatic Events ...... 7

3. Results...... 8 Stratigraphy ...... 8

Petrography ...... 9

Whole-Rock Geochemistry...... 17

Rock Classification ...... 21

Trace Element Patterns ...... 29

4. Alteration and Metamorphism...... 37 Assessment of Element Mobility ...... 37

Metamorphism ...... 46

5. Discussion ...... 50 Correlation with Temagami Greenstone Belt ...... 50

A New Sudbury Offset Dyke ...... 555

6. Outlook ...... 622 7. References ……………………………………………………………………………... 63

III

Executive Summary

The Temagami magnetic anomaly is similar in size and extent to the geophysical anomaly that marks the 1.85 Ga Sudbury Igneous Complex (SIC) in its immediate vicinity but its geological cause and potential link to the Sudbury impact structure have remained elusive. A first analysis of the lithology and alteration of diamond drill core intersecting the area of maximum magnetic anomaly (bore hole AT-14-01) revealed that the units intersected below Huroninan sedimentary cover can be correlated with those of the 2.7 Ga Temagami Greenstone Belt. Much of the intersected rock types experienced extensive alteration, in places completely obliterating original protoliths. Different styles and stages of alteration could be distinguished based on petrographic, mineral chemical and whole rock geochemical analyses. Much of the alteration in the Temagami Greenstone Belt-equivalent units is similar to VMS-type alteration, probably on the Neoarchaean seafloor. Two igneous units in the drill core escaped that alteration. These are mafic rocks that bear all the hallmarks of the 2.2 Ga Nipissing Diabase elsewhere in the region and two diorite bodies, either sills or dykes, near the bottom of the bore hole. Both of these rock types experienced low-grade metamorphic overprint but with relatively little element mobility. This makes it possible to reconstruct their protolith composition. New geochemical data for the diorite indicate that it has a composition that is very similar to that of quartz dioritic offset dykes in and around the SIC and that is very distinct from all other kown igneous units in the region. A correlation of the diorite with the SIC is further supported by Sr, Pb and most notably Nd whole rock isotope data. The latter are in perfect agreement with new and published data for the SIC. Consequently, it is speculated that the Temagami magnetic anomaly could be caused by a larger body of such magnetic diorite, representing offset dykes at an even greater distance from the SIC than has been known so far. This new finding considerably increases the exploration potential of the area representing the Temagami magnetic anomaly for SIC-type sulfidic Cu- Ni-PGE deposits.

1. Preamble

The principle goal of the project, as outlined in the project proposal approved on 8th May 2016, has been to determine whether mafic rocks, intersected by diamond drill hole AT-14-01, which targeted the Temagami magnetic anomaly, are equivalent to rocks of the 1850 Ma Sudbury Igneous Complex (SIC). The motivation for this is based on a first preliminary petrographic and geochemical assessment as reported by A. Bite in an internal report from 30th December 2014 as well as a first attempt to date these rocks by D.W. Davies (Internal Report from 1st September 2015), which yielded a highly imprecise U-Pb monazite age of 1775 ± 76 Ma. Although this age is within analytical error of the age of the SIC, its poor quality and uncertainty about the actual genesis of the dated monazite grain leaves the result inconclusive. A critical unit in the drilled section is a rock at the bottom of the drill hole, logged as gabbronorite by Joerg Kleinboeck, as it is the most likely candidate for being a possible correlative of the SIC.

The project to be reported on here was designed as a MSc thesis project to be carried out by the MSc student A. Kawohl under the supervision of H.E. Frimmel. To this effect, both A. Kawohl and H.E. Frimmel visited Sudbury for two weeks from 18th to 30th July 2016. During that time, representative samples of the various rock types within drill core AT-14-01 could be taken, complemented by some samples from the Emerald Lake deposit and other units of and around the SIC. In addition to sampling the drill core, an introduction to the regional geology was provided through field visits under the guidance of A. Bite and W. Whymark. Brief discussions were also held with P. Lightfoot.

The sampled material was shipped to Würzburg, Germany, where it arrived by the end of September 2016. In the meantime A. Kawohl commenced his petrographic analyses, first on thin sections provided by A. Bite, subsequently on thin sections prepared in Würzburg. Since then A. Kawohl conducted a series of analyses on this material. First petrographic observations have been summarized in the first interim report, submitted in January 2017. Without a proper understanding of the alteration history of the rocks of interest, any interpretation of the protoliths’ geochemistry in terms of magmatic provenance would be futile. Consequently, the first work step focused very much on characterizing the style and extent of alteration, which was found to be extensive and widespread throughout most of the core. Since then, A. Kawohl expanded on the alteration study by investigating the mineralogy of various alteration phases, using electron microprobe and XRD techniques and the trace element distribution within different generations of magnetite. The latter provided evidence of a hydrothermal magnetite generation and these results were presented at the 14th Biennial SGA Meeting, held in Quebec City from 20-23 August 2017. Up to that stage, no evidence could be found to support a genetic link between the rocks intersected in the drill core AT-14-01 and the SIC but much of the intersected rocks could be correlated with those of the Neo- archaean Temagami Greenstone Belt (Kawohl et al. 2017). Subsequently, this situation changed when we eventually obtained the complete geochemical data set for some 80 samples analyzed by XRF and ICPMS. Seventeen of these samples represent two diorite (previously logged gabbronorite) sills or dykes within

2 the bottom 200 m of the drill core and their geochemistry turned out to show strong similarities with reported data on SIC offset-dykes. This exciting result could be further corroborated by whole rock Sr, Pb and Nd isotope data, obtained by ICPMS analyses in February 2018. Whereas the Sr and to some extent also the Pb isotope data are ambiguous because of variable mobility of Rb, Sr, U, Th and Pb during alteration/metamorphism of the studied rocks, the Nd isotope data bear consistently strong resemblance to new isotope data that we also acquired for some SIC samples as well as to published data from the SIC. Such a correlation between the drilled diorite and the SIC has major implications for the exploration potential of the area.

All analytical work was carried out at the Department of Geodynamics and Geomaterials Research, University of Würzburg, except for the LA-ICPMS analyses on magnetite, which were performed at the University of Göttingen, the ICPMS whole rock analyses, which were conducted at the Free University of Brussels, and the ICPMS isotope measurements, carried out at the Department of Geological Sciences, University of Cape Town.

A. Kawohl completed his MSc thesis (in German) in January and submitted it on time at the Univer- sity of Göttingen for examination.

2. Geological Background

Location

The Temagami Anomaly is located some 40 km north(east) of the Sudbury City Center, east of Lake Wanapitei, and is situated near the boundary of three geological provinces, that is, the granite-greenstone belts of the Neoarchaean Abitibi Subprovince of the Superior Province, the Paleoproterozoic Southern Province, which is dominated by metasedimentary rocks of the Huronian Supergroup, and the Neoprote- rozoic Grenville Province (Fig. 1). A geological description of the study area has been given by Klein- boeck (2015).

Figure 1: Regional geological setting of the Temagami Anomaly; after Lightfoot (2016).

Temagami Greenstone Belt

Surrounded by Paleoproterozoic sedimentary rocks and diabase, the Temagami Greenstone Belt (TGB) represents a geological window to the Neoarchaean basement (Fig. 1). It is exposed as a southwesterly plunging syncline, covering an area of approximately 13 x 29 km; smaller outcrops occur ~2 km east of the drilling location (east of Emerald Lake). The TGB is dominated by intermediate to felsic metavolcan- ics rocks (flows, tuffaceous and pyroclastic deposits) of calc-alkaline affinity and siliciclastic metasedimentary rocks, including shale, greywacke and conglomerate. Minor volumes of tholeiitic basaltic flows occur, as well as a variety of synvolcanic intrusive phases, such as, for example, the layered ultramafic Ajax (Kanichee-) intrusion, diorite-, pyroxenite- and lamprophyre dykes (e.g. Bennett 1978). Coeval with

4 the volcanic activity, dated at ~2.7 Ga (Bowins & Heaman 1991), was the emplacement of three granitoid batholites.

The TGB is well known for its abundance of Algoma-type banded iron formation (BIF). At least two units of BIF occur on both limbs of the TGB syncline and can be traced for tens of kilometers along strike: A lower body is composed of banded chert, pyrite and pyrrhotite (sulfide facies), 25 m thick, and overlies volcanic rocks. A second formation of banded chert, magnetite/haematite and chlorite (oxide facies) is 100 m thick and occurs embedded in pyritic shale, tuff and ultramafic fragmental rocks (Bennett 1978; Fyon & Crocket 1986; Jackson & Fyon 1991).

Rocks of the TGB were subjected to regional metamorphism of the greenschist facies (Bennett 1978). However, there is also evidence for extensive Neoarchaean seafloor alteration that shares features of volcanic massive sulfide (VMS) systems (Fyon & Crocket 1986; Mark D. Hannington, pers. comm.), although no exploitable sulfide mineralization of this type is known in the TGB.

Huronian Basin

Subsidence of the Huronian Basin began at the Archaean-Proterozoic boundary and deposition of the basin infill lasted until 2.2 Ga (Ketchum et al. 2013; Corfu & Andrews 1986). The Huronian Supergroup comprises rift-related bimodal volcanic rocks at is base and three sedimentary cycles of conglomerate, siltstone and sandstone, that are likely to reflect glacial periods (Young et al. 2001; Long 2004). A maximum thickness of 12 km is reached southeast of Sudbury (Young et al. 2001) and a thickness of ~1.5-3.4 km is assumed within the area of the Temagami Anomaly (Dressler 1982, 1986; Milkereit and Wu 1996). Rocks of the Huronian Supergroup within the study area were subjected to very low-grade sub- greenschist-facies metamorphism and deformation related to the Penokean and Grenville orogenies (Card 1978). There is evidence of a metasomatic event sometime between 1.74-1.93 Ga, which is ascribed to contemporaneous alkaline magmatism in the area (Schandl et al. 1994; Fedo et al. 1997; McLennan et al. 2000; James & Golightly 2009).

Sudbury Igneous Complex

Formation of the Sudbury Structure is attributed to a 1.85 Ga large asteroid or comet impact. The original rim to rim diameter of this impact structure is estimated at 150 to 200 km. Vast volumes (c. 15 000 km3) of granodioritic impact melt filled the central cavity of the resulting multi-ring basin and subsequently differentiated into a ~3 km thick sequence of, from bottom to top, xenolith-bearing “contact sublayer” of noritic composition discontinuously overlying the brecciated footwall country rocks, mafic quartz-norite, felsic norite, apatite- and oxide-rich quartz monzogabbro, and granophyre (Naldrett & Hewins 1984; Therriault et al. 2002; Lightfoot 2016).

The term ‘offset dykes’ (or offset sublayer) is applied to narrow (ten to hundred meters wide) igneous bodies of quartz dioritic or quartz-monzodioritic composition, related to the Sudbury Igneous Complex (SIC). These offset dykes occur either radially to, or concentrically around, the SIC, and can be traced for several kilometers along strike (see Grant & Bite 1984). Currently, 17 offset dykes and extensions are known (Lightfoot 2016), some of which host highly economic amounts of sulfide ore. New offsets were recently discovered in the north of the SIC (Smith et al. 2013). As shown on the map (Fig. 2), these known offset dykes can occur as far as 30 km away from the SIC. The quartz dioritic offset dykes are homogenous in composition and Osterman et al. (1996) provided a crystallization age (U-Pb on zircon and baddeleyite) of 1849.4 Ma +3.5/-2.6 Ma which is almost identical to the age of the SIC (1850.0 Ma, Krogh et al. 1984). It is generally believed that the offset dykes were emplaced early after the Sudbury impact and that they represent the initial, undifferentiated impact melt (see Lightfoot 2016 for a review). Thus, the composition of these quartz dioritic rocks should approximate the average composition of the SIC.

Figure 2: Simplified geological map of the Sudbury Igneous Complex (SIC) and its offset dykes including recent discoveries by Wallbridge Mining. The drilling sight of AT-14-01 lies as far as 50 km away from the SIC; after Smith et al. (2013).

The SIC experienced tilting and NW-SE shortening during the Penokean Orogeny and is nowadays exposed as funnel-shaped, elliptical 60 x 30 km large structure, overlain by fall-back breccias and Palae- ozoic sedimentary rocks of the Whitewater Group. Metamorphism at upper greenschist- to lower amphibolite-facies conditions took place during the Penokean and subsequent orogenies from 1.85-1.7 Ga (e.g. Fleet et al. 1987). The SIC is not only the prime natural laboratory to study impact processes but it also hosts world-class Ni-Cu-PGE deposits directly related to the impact. Massive to disseminated sulfide mineralization affected the contact sublayer, the brecciated country rocks and the inclusion-bearing facies of quartz-diorite offsets (see Lightfoot 2016 for a review).

Other Magmatic Events

In order to compare and correlate the various igneous protoliths intersected in the studied bore hole, other magmatic units, both the Archaean basement and the Huronian cover rocks, are briefly summarized. Mag- matic events in the region encompasses the emplacement of

(i) very widespread mafic sills and dykes of the 2.2 Ga Nipissing Diabase (e.g. Lightfoot & Naldrett 1989, 1996; Lightfoot et al. 1986, 1993), (ii) intrusion of ultrabasic magma forming layered intrusions of the 2.48 Ga East Bull Lake Suite and associated feeders known as the Matachewan dyke swarm (Ciborowski et al. 2015 and references therein), (iii) granitoid plutons of roughly the same age (e.g. Riller 2009; Lightfoot 2016), (iv) alkaline complexes, such as the 1881 Ma Spanish River Carbonatite (Rukholv & Bell 2009) and the 1.2 Ga Nemag Lake Fenite (Siemiatkowska & Martin 1975).

The Nipissing Diabase is of specific interest, as it is the most widespread intrusive phase in the study area (Fig. 3). Major rock type is a hypersthene-bearing tholeiitic gabbro, differentiated portions are granophyre, aplite, granite and granodiorite. Pegmatoidal textures are common in the Cobalt-Temagami- Wanapitei area (Lightfoot 2016), and in most areas the Nipissing Diabase (and related rocks) were affected by (very) low-grade metamorphism at (sub-)greenschist-facies conditions (Card & Pattison 1973; Card 1978).

Figure 3: Geological map showing the distribution of the Nipissing Diabase; from Lightfoot (2016).

3. Results

Stratigraphy

Drill core AT-14-01 represents a volcano-sedimentary succession likely of Archaean age and correlative with the Temagami and Emerald Lake Greenstone Belts, overlain by the metasedimentary Gowganda For- mation of the Cobalt Group (upper Huronian Supergroup). The Ar- chaean basement was reached at a depth of ~600 m and comprises intermediate/felsic feldspar-porphyric volcanic rocks, tuffaceous, pyroclastic and volcaniclastic layers, banded iron formation of the oxide-chlorite- facies, pyritic shale, and greywacke. Two younger intrusive phases are recognized within the drill core: First, a 300 m thick gabbroic sill of the Nipissing Diabase at the top of the core, and second, two exotic diorite dykes, 25 and 52 m thick, intrusive into the feldspar porphyry and iron formation at depths below 2,000 m. A simplified geological column of the drill core is shown in Fig. 4, and a detailed description of each lithology is given in the chapter below and by Kawohl & Frimmel (2017a).

Figure 4: Lithological profile of drill core AT-14-01.

Petrography

Porphyric volcanic rocks / feldspar porphyry Intermediate to felsic porphyric volcanic and/or subvolcanic rocks are the most abundant lithotype encountered in drill core AT-14-01. They occur repeatedly from 600 m downwards as massive, non-strati- fied flows, in places intercalated with strongly altered pyroclastic/volcaniclastic layers and sedimentary rocks (iron formation, pyritic shale, greywacke). These volcanic rocks are holocrystalline and seriate- textured in thin section. Phenocrysts amounting to 30-45 vol% are embedded in fine-grained, microcrys- talline, mosaic-like and homogenous groundmass of quartz, sericite, biotite and some carbonate (e.g. Figs. 5c,d,e,h). The dark colour of some sections of the rock (e.g. Fig. 5a) is due to large amounts of biotite (>25 vol%) in the groundmass, which formed at the expense of sericite and quartz. Among the phenocrysts, feldspar is the most abundant mineral (predominantly plagioclase). Euhedral to subhedral phenocrysts of felspar range in size from 20 µm to 5 mm (3 mm on average) and are strongly altered to sericite. Sericitization of these crystals was so extreme that today only their outlines can be recognized in thin section (Figs. 5a,g). Locally, up to 3 cm large and zoned megacrysts of alkali feldspar were observed (Figs. 5a,b). Rounded and deeply embayed phenocrysts of blueish quartz (2-4 mm in diameter) constitute up to 5 vol%, and locally, phenocrysts of biotite or hornblende (Figs. 5f,h) are also present. Accessories are ilmenite, titanite, pyrite, zircon, apatite and allanite.

The porphyric volcanic rocks underwent intense hydrothermal alteration. Hydration is expressed by omnipresent sericitization of feldspars, accompanied by silicification. The latter process led to the formation of massive cherty flood quartz at 932-935.4 m depth, where a massive, 3 m-thick exhalative pyrite horizon occurs, hosted within the volcanic pile. Chloritization is also common, resulting in almost monomineralic chlorite fels, speckled with pyrite- and carbonate-porphyroblasts. Biotite seems to have been remobilized on a local scale, but no pervasive biotitization (K+-metasomatism) was noted.

Aphyric Volcanic Rocks

These constitute a thick, homogenously textured unit of massive fine-grained rocks, which are characterized by a low density. They are non-magnetic. Because of their dark colour (Fig. 6a), the lithology has been claimed to be a mafic intrusive rock, either a gabbro or norite, and we noticed that the core had already been labeled as such. The rocks have, however, nothing in common with a gabbro/norite, nor with any other type of magmatic rock. The rock is a chaotic mixture of biotite (30-80 vol%), sericite (10-90 vol%), chlorite (5-30 vol%), quartz (5-15 vol%) and accessory opaque minerals, including ilmenite (Fig. 6b,c). A poorly developed lineation (flow texture?) is indicated by a preferred orientation of sericite, and some portions of this lithology are converted to ±monomineralic sericite rocks (up to 90 vol% sericite) and relicts of quartz (10 vol %) (Fig. 6d). Due to the intense alteration, it is virtually impossible to give these rocks a proper name, not to speak of the unaltered protolith. The fine-grained texture and the weak

9 lamination could point to a tuff/ash-fall deposit or a clastic sediment, but the rock could also be a pheno- cryst-free (aphyric) variety of the rocks at 857.9 m – 915 m (e.g. Fig. 6a), for instance.

Figure 5: Photographs (a-b) and microphotographs (c-h) of the porphyric volcanic rocks; length of yellow bar = 1 cm, red bar = 1 mm.

Figure 6: Photograph (a) and microphotographs (b-d) of aphyric volcanic rocks, possibly of pyroclastic origin; red bar = 1 mm.

Volcaniclastic Sedimentary Rocks Volcaniclastic sedimentary rocks are defined as material, mainly composed of volcanic products, which has been redeposited by marine, fluviatile and/or eolic processes. Rocks of this type occur in association with BIF from 654 to 681 m and from 1,955 to 1,993 m and have erroneously been logged as quartz diorite. Volcaniclastic sedimentary rocks are the footwall and hangingwall of the BIF at 1,980 m, and the hangingwall of the BIF at 650 m. The matrix of this lithotype has, from a macroscopic point of view, similarities with the porphyric volcanic rocks. The volcaniclastic rocks are light grey in hand specimen, poorly laminated and composed of subhedral to euhedral dolomite porphyroblasts of 1-3 mm diameter (30-40 vol%), slightly aligned, anhedral xenocrysts of biotite (20-30 vol%), fine-grained anhedral quartz and sericite (20-40 vol%), and chlorite (3-10 vol%) (Figs. 7c,d). Rounded and broken crystals of apatite/monazite (grain size <200 µm) and ilmenite are common accessoric minerals, as well as fragmented xenocrysts of feldspar (Fig. 7a), resembling phenocrysts in the feldspar porphyry (Fig.s 7a,b). Two types of well-rounded lithoclasts, several centimeters in size, were noted. One type comprises fragments of BIF, composed of magnetite and quartz (Figs. 7a,b). The other type of inclusion is marked by a dark colour in

11 hand specimen (Fig.7a). In thin section, however, these inclusions are hardly distinguishable from the groundmass, and consist of biotite (10-25 vol%), Mg-rich chlorite (30-50 vol%) possibly after biotite or other ferromagnesian minerals (olivine?), and dolomite.

Mafic clast

BIF clast Feldspar fragment

Mafic clast

BIF clast

Figure 7: Photographs (a-b) and microphotographs of volcaniclastic sedimentary rocks; yellow bar = 1 cm; red bar = 1 mm.

Pyritic Shale This is a dark grey to black, very fine-grained, slaty and carbon-rich sedimentary rock (Figs. 8a,b), interbedded with greywacke and BIF. The most prominent feature of this shale is the abundance of pyrite, both in the form of up to 3 cm thick concordant bands (Fig. 8a) or large, framboids (Figs. 8b,c). These early diagenetic sulfide-concretions locally make up 15 vol% of the rock and have diameters between 3 mm and 5 cm. Pyrite nodules were slightly deformed and rotated, and recrystallized quartz is found in their pressure shadows (Fig.8c).

Figure 8: Photographs (a-b) and a microphotograph of the pyritic shale; length of red bar = 1 mm.

Banded Iron Formation Oxide facies banded iron formation (BIF) of the Algoma-type occurs repeatadly from 650 to 710 m, from 1,388 to 1,585 m, and from 1,797 to 2,068 m, interbedded with volcanic and volcaniclastic rocks, pyritic shale and greywacke. The BIF is made up of alternating layers of magnetite and chert (Figs. 9a,b) and, in places, of graphite. Individual layers range in thickness from a few millimeters (‘microbands’) to 2-3 cm (‘mesobands’), and dip, in places, steeply at 80o towards the core axis. Strata of Fe-rich chlorite, stilpnomelane and carbonate are common (Fig. 9c) and might represent metamorphosed pelagic sediments of pyroclastic layers. Transmitted and reflected light microscopy revealed a micro- to cryptocrystalline texture of anhedral quartz, on average less than 100 µm in diameter, and 150 µm at most. Quartz shows signs

13 of recrystallization and grain coarsening, and magnetite occurs dispersed within chert layers. Some mac- roscopically pure chert layers contain as much as 30 vol % magnetite and vice versa, but pure magnetite mesobands (>90 wt% Fe2O3) are also present (Fig. 9a). No Fe(III)-oxides (haematite, goethite) were observed; jasper is absent. Some deformation of the BIF took place, resulting in fold structures and boudins. Brittle deformation formed breccias of chert cemented with magnetite and chlorite. In addition, micro- faults and fractures are filled with chlorite, quartz, carbonate, and in addition, pyrite and arsenopyrite were identified in hydrothermal veins.

Figure 9: Photographs of the banded iron formation; oxide facies (a-b) and silicate facies (c); yellow bar = 1 cm.

Nipissing Diabase This is a medium- to coarse-grained (Fig. 10a) and locally pegmatoidal (grain size >0.5 cm) (Fig. 10b) massive rock, of light-green colour which occurs as a continuous intrusive body from 0 to 321 m within Huronian metasedimentary rocks. Unaltered sections of the Nipissing Diabase (Fig. 10c,d) exhibit an intergranular texture of euhedral and fresh, locally zoned plagioclase laths (50-60 vol%), intergrown with interstitial anhedral to subhedral clinopyroxene (c. 40 vol%) and some anhedral orthopyroxene (up to 5 vol%). Accessories include biotite and opaque minerals. Secondary minerals (amphibole, talc, chlorite) occur in minor amounts (max. 10 vol%).

Figure 10: Photographs (a-b) and microphotographs (c-g) of the Nipissing Diabase showing different degrees of alteration; Yellow bar = 1 cm, red bar = 1 mm.

Other portions of the Nipissing Diabase (e.g. Figs. 10e-g) underwent strong alteration by which ferromagnesian minerals have been, in part or entirely, replaced by amphibole, talc, chlorite and leucoxene, and plagioclase by clinozoisite. In places of extreme alteration, the primary igneous texture of the original gabbro is no longer apparent (Fig. 10g), and the rock is completely converted to a paragenesis of clinozoisite+ampbhibole±chlorite±quartz. Small dyklets of green colour, between a few centimeters and up to 2 m thick, were identified at 606.35 m and 689.5 m, cross-cutting conglomerate and the feldspar porphyry, respectively. Petrographically, these rocks correspond to the Nipissing Diabase. The texture equals the most altered sections of the diabase, as shown in Figure 10g. However, no correlation between depth and the degree of alteration/metamorphism is apparent in the core.

Diorite dykes Two sills or dykes of diorite have been recognized within the bottom 200 m of drill core AT-14-01 and already described previously (Kawohl & Frimmel 2017b). The lower body from 2,168.5 to 2,143.75 m is 25 m thick, the upper one from 2,119.85 to 2,068.4 m is 51.5 m thick (Fig. 11). Both are in sharp (intrusive) contact with BIF and intermediate porphyric volcanic rocks of likely Archaean age. At 2,119.85 m, the contact between diorite and country rock is sheared and healed by chlorite and calcite.

The diorite is greyish-green in hand specimen (Figs. 11, 12a,b), massive, fine-grained and texturally ho- mogeneous. It is characterized by a relatively high density and magnetic susceptibility. Xenoliths were not observed. The diorite seems to be chilled against the country rock, becoming slightly coarser grained away from the contact. Due to strong alteration, possible quench textures are not preserved. Petrographic analysis revealed two textural types:

(i) A coarser grained type (Figs. 12c,d) with subophitic texture of euhedral, twinned and zoned plagioclase laths (0.5 – 1 mm in length, up to 70 vol%), in places strongly or entirely altered to epidote; olive-green biotite and pale green needles and bundles of actinolite occur as interstitial phases in amounts of 20-25 vol% and 5 vol%, respectively. (ii) A fine-grained type with chaotic texture (e.g. Figs. 12e,f); the groundmass is a micro- to cryptocrystalline mixture of epidote, biotite, actinolite, chlorite and opaque phases.

Neither pyroxene nor granophyric intergrowths of quartz and feldspar were observed. Accessory minerals include magnetite with exsolution lamella of Fe-Ti-oxide, anhedral quartz, leucoxene and pyrite. Secondary minerals are epidote (Figs. 12g,h), chlorite (after biotite) and calcite, both of which occur as tiny veins or dispersed within the groundmass. Quartz might be a secondary phase as well, and biotite and amphibole could be alteration products of primary ferromagnesian minerals (e.g. pyroxene). The diorite is without doubt younger as, and intrusive into, the surrounding volcanic rocks, as evident from contact relationships (chilling, contact metamorphism) and two opposing styles of alteration for the diorite and its surroundings (see discussion in Chapter 4).

BIF with chlorite layers/veins

Upper diorite sill/dyke

Porphyric intermediate volcanic rock

Lower diorite sill/dyke

Porphyric intermediate volcanic rock

Figure 11: Stratigraphic column of the bottom 200 m of drill core AT-14-01; photographs of drill core samples to the right (core diameter is ~4.5 cm).

Whole-Rock Geochemistry

For whole-rock geochemistry, a total of 80 samples were fused to glass pellets and then analyzed with a PANalytical Minipal4 energy dispersive x-ray fluorescence spectrometer (XRF) at the University of

Würzburg. The loss on ignition (LOI), i.e. the amount of volatiles (H2O, CO2, S), was determined by the weight loss after heating the sample powders in a muffle furnance to 800°C. Trace element concentrations were analyzed by means of inductively coupled plasma-mass spectrometry (ICP MS) at the Laboratoire G-Time (University of Brussels). The results of these measurements are presented in Tables 2-5.

The Nipissing Diabase is the most primitive lithology of drill core AT-14-01, having a truly basic composition of 50.7-51.7 wt% SiO2, 8.0-8.4 wt% MgO and a high Mg# of 39.5-47.3. Low concentraions of volatiles (LOI between 0.8 and 2.6 wt%) are in good agreement with petrographic obersvations; the LOI is correlated with the degree of alteration as determined by optical microscopy, and is highest in

17 sample AT-114, where it reflects the abundance of secondary hydrous silicates (amphibole, clinozoisite). A mafic composition of the Nipissing samples is also supported by high concentrations of Cr (162-496 ppm) and Ni (115-263 ppm), negatively correlated with concentrations of incompatible elements (e.g.

0.3-0.6 wt% K2O, 36-49 ppm Zr, 10-23 ppm Rb, 40-64 ppm REE+Y, 0.3-0.5 ppm U, 1.0-1.4 ppm Th).

Figure 12: Photographs (a-b) and microphotographs of the two exotic diorite dykes. Yellow bar = 1 cm; red bar = 1 mm.

Table 1: Summary of petrographic invesitgations.

Lithology Texture Mineralogy Alteration Image

30-45 % feldspar strong, pervasive 20-40 % quartz phyllic alteration porphyric, 10-35 % sericite sulfidic alteration, Porphyric seriate textured, 5-20 % biotite quartz flooding Volcanic Rock megacrystic 0-5 % calcite + greenschist 0-90 % chlorite metamorphism 0-1 % hornblende

5-15 % quartz 10-80 % biotite strong, pervasive Aphyric 10-90 % sericite phyllic alteration Volcanic aphyric, aphantic 5-30 % chlorite + greenschist Rock 0-15 % calcite metamorphism 1 % ilmenite

heterolitic, 30-40 % dolomite matrix-supported, 20-40 % quartz strong, pervasive Volcaniclastic rounded ultra- 20-30 % biotite phyllic alteration Sedimentary mafic and BIF- 3-10 % chlorite carbonatization Rock lithoclasts, 0-5 % sericite + greenschist possibly < 1 % apatite, mona- metamorphism bedded/layered zite, ilmenite

quartz Oxide Facies magnetite BIF carbonate micro- and chlorite coarsening & mesobands, stilpnomelane recrystallization, well-layered, biotite greenschist facies Silicate Facies brecciated, folded pyrite metamorphism BIF siderite/ankerite quartz magnetite

50-60 % plagioclase 35-45 % clinopyroxene subophitic moderate, 0-5 % orthopyroxene Nipissing medium-grained, locally fresh 0-50 % clinozoisite Diabase locally propylitic alteration 0-50 % amphibole pegmatoidal greenschist facies 0-1 % quartz, biotite, chlorite, leucoxene, sulfide

60-70 % plagioclase 10-50 % epidote 20-25 % biotite moderate, subophitic 0-7 % calcite locally fresh Diorite Dykes fine- to 0-5 % actinolite propylitic alteration medium-grained 0-10 % chlorite greenschist facies < 5 % magnetite, pyrite, quartz

The diorite has a relatively uniform intermediate composition, with 53.4-56.2 wt% SiO2, 11.2-12.6 wt% Fe2O3 and 3.5-3.9 wt% MgO. The Mg# (i.e. Mg/(Mg+Fe)) is 22-25, and the LOI is 1.0-6.6 wt%. The rock is overall poor in Cr (7-88 ppm) and Ni (43-91 ppm), except for three outliers with up to 308 ppm Cr and 203 ppm Ni. Despite being the secondmost primitive lithology in this drill core – after the

Nipissing Diabase – the diorite is anomalously rich in incompatible elements, for example K2O (1.3-3.2 wt%), U (1.1-13.0 ppm), Th (6.5-7.2 ppm), ΣREE+Y (169-186 ppm) and Zr (139-155 ppm). The U, Th and ΣREE values are among the highest measured in the entire drill core, exceeding even those of the surrounding felsic country rocks (66-69 wt% SiO2).

The group of porphyric volcanic rocks is characterized by a high degree of differentiation with respect to SiO2 (59.4-68.9 wt%), Na2O (0.1-4.6 wt%) and K2O (2.1-7.9 wt%), accompanied by a relative depletion in MgO (1.2-5.2 wt%) and Fe2O3 (2.3-8.8 wt%). Concentrations of Zr (ppm 123-321 ppm) are high as well, but Th (2.2-7.3 ppm) and REE+Y (74-195 ppm) are lower than for the diorite dykes, although still higher than for the Nipissing Diabase. Concentrations of Ni and Cr are overall low (19-122 ppm and 52-221 ppm, respectively), except for brecciated and chloritized sample AT-199 (122 ppm Ni and 119 ppm Cr). A large variability in the concentrations of alkali metals (Ba, Rb, K) and alkaline earth metals (Mg, Sr, Ca), indicates that significant post-magmatic element mobility has affected these rocks. This is especially true for samples altered to sericite and/or chlorite, in which Rb, Ba, K, and Na were completely leached out.

The aphyric volcanic (pyroclastic?) rocks have a restricted range in composition and are slightly more primitive than the porphyric variety, viz 47.8-62.9 wt% SiO2. Concentrations of MgO and Fe2O3, for example, are between 3.0-7.3 wt% and 6.4-19.6 wt%, respectively. Again, a significant enrichment in

K2O (up to 11.7 wt%) is noted. Trace element concentrations are very similar to those in the porphyric volcanic rocks.

The volcaniclastic sedimentary rocks are characterized by a low concentration of SiO2, and high concentrations of MgO (2.6-15.4 wt%), Fe2O3 (7.3-24.1wt%), CaO (0.4-7.8wt %) and volatiles (2.9-13.4 wt% LOI), which is in good agreement with petrographic observations. Clearly, the high MgO, CaO and

LOI are due to the abundance dolomite, CaMg(CO3)2, chlorite, Mg6(Si,Al)4O10(OH)8, and biotite,

K(Fe,Mg)3AlSi4O10(OH)2. However, a large chemical variability is recognized within this lithology, re- flecting its polymictic and clastic nature (e.g. Fig. 13): Samples containing dark fragments have high concentrations of MgO (11.9-15.4 wt%), P2O5 (0.18-0.34 wt%), Cr (982-1172 ppm) and Ni (575-629 ppm), accompanied by very low concentrations of Na2O (≤0.1 wt%), Zr (61-82 ppm), Th (2.2-2.7 ppm) and REE+Y (60-106 ppm) (Fig. 13). The dark fragments are therefore regarded as (ultra-)mafic, even though no olivine, pyroxene or feldspar is persevered. Analysis of the one sample containing a fragment of BIF (Fig. 13) revelead elevated Fe2O3 and SiO2, but low Zr, REE+Y, Cr, Ni, as it is typical of BIF. The composition of the groundmass (free of lithoclasts) resembles that of the volcanic rocks. Accordingly, the

20 concentrations of Cr and Ni are below 35 ppm, while Zr, TH and REE+Y are enriched with 99-152 ppm, 1.9-7.7 ppm and 122-158 ppm, respectively (Fig. 13).

Figure 13: Bivariate plots for the volcaniclastic sedimentary rocks to distinguish between the groundmass and the different types of inclusions.

Rock Classification

According to the TAS-diagram (total alkalis vs. silica) after Le Bas et al. (1986), the diorite sills/dykes’ composition corresponds to andesite, trachyandesite or basaltic andesite, and that of the Nipissing Diabase to basalt. However, as Na and K might have been mobile during alteration, classification plots using immobile elements (Si, Ti, Nb, Y, Zr) are expected to be more reliable. In these diagrams (Winchster & Floyd 1977; Figs. 15,16,17), the diorite samples plot within the fields of andesite (diorite) or trachyandesite (monzodiorite). Again, the samples taken from the Nipissing Diabase correspond to basalt (gabbro), which is in perfect agreement with its mineralogical composition.

The porphyric volcanic rocks are classified as trachyte, (rhyo-)dacite or (trachy-)andesite using plots after Winchester & Floyd (1977), and as (trachy-)dacite, trachyte or even rhyolite, according to the TAS diagram. Based on petrographic observations, the classification after Winchester & Floyd seems most applicable. Secondary element mobility has clearly caused a shift in the concentrations of Na2O and K2O and thus a large spread of data points in the TAS plot. The diagrams after Winchester & Floyd classify the aphyric volcanic rocks as trachyandesite, andesite or dacite, with some altered outliers plotting in the basaltic fields. Due to the strong alteration of this rock type, the TAS diagram gives only imprecise results and a large scatter of data points is noticed (Fig. 14).

The groundmass of the volcaniclastic sedimentary rocks is indistinguishable from the porphyric and aphyric volcanic rocks and corresponds to trachyandesite or (rhyo-)dacite. Clast-bearing samples have a larger spread and lie in the andesitic, basaltic or basanitic fields. Such a classification is not readily com- parable with petrographic observations, as no primary minerals are preserved.

Table 2: Geochemical analyses of the Nipissing Diabase and the Diorite dykes/sills at the bottom of drill core AT-14-01.

Nipissing Diabase Diorite Dykes

Depth [m] 311.35 242.00 42.30 2109.00 2105.00 2097.50 2096.20 2091.00 2089.20 2166.00 2165.00 2165.00 2159.50 2157.00 2156.90 2156.30 2155.00 2153.50 2149.50 2148.20 Sample AT-114 AT-113 AT-115 AT-36 AT-38 AT-39 AT-64 AT-40 AT-41 AT-14 AT-15 AT-58 AT-17 AT-19 AT-59 AT-20 AT-21 AT-22 AT-24 AT-25 Major element oxides [wt.%] SiO2 50.7 51.4 51.7 53.4 53.7 53.4 53.6 53.5 53.9 54.5 54.5 54.7 55.8 55.3 55.6 55.5 55.8 56.2 53.6 54.3 TiO2 0.59 0.89 0.52 1.08 1.09 1.08 1.09 1.07 1.08 1.09 1.12 1.11 1.15 1.16 1.17 1.16 1.15 1.18 1.10 1.11 Al2O3 15.1 12.4 15.6 13.7 13.8 13.7 13.8 13.7 13.7 13.8 14.0 14.0 14.3 14.3 14.4 14.2 14.4 14.3 13.8 14.0 Fe2O3 (t) 10.1 12.3 9.4 11.3 11.7 11.3 11.5 11.5 11.2 11.5 12.1 11.7 12.2 12.3 12.4 12.2 12.2 12.6 11.3 11.7 MgO 8.0 8.0 8.4 3.8 3.9 3.7 3.7 3.7 3.5 3.5 3.7 3.5 3.6 3.6 3.6 3.7 3.9 3.8 3.5 3.7 CaO 11.1 11.8 12.2 5.4 4.8 5.3 4.9 6.2 5.8 6.8 6.2 6.6 5.6 5.7 5.6 5.5 5.4 4.6 5.2 5.2 MnO 0.15 0.19 0.15 0.15 0.15 0.14 0.12 0.12 0.11 0.14 0.14 0.14 0.15 0.16 0.15 0.16 0.16 0.16 0.14 0.14 Na2O 2.1 1.9 1.8 4.0 4.0 3.9 4.0 4.0 4.2 2.9 3.0 2.9 3.3 3.3 3.6 3.4 3.7 4.0 4.1 4.2 K2O 0.59 0.32 0.42 2.72 3.24 2.65 2.76 1.45 1.33 1.87 2.24 2.26 2.50 2.42 2.39 2.50 2.30 2.05 2.91 2.27 P2O5 0.04 0.02 0.07 0.09 0.07 0.08 0.09 0.10 0.10 0.11 0.09 0.11 0.08 0.09 0.09 0.08 0.07 0.06 0.04 0.08 LOI 2.6 1.6 0.8 5.5 4.8 5.6 5.0 6.6 6.3 4.7 4.0 3.8 1.1 1.0 1.2 1.1 1.1 1.5 4.6 5.3 Total 100.97 100.83 100.95 101.10 101.22 100.85 100.60 101.86 101.30 100.86 101.03 100.76 99.79 99.39 100.07 99.44 100.04 100.42 100.32 101.93 Trace elements [ppm] Cr 496 162 292 11.2 21 7.5 12.2 81 11.8 296 51 18.6 32 308 65 244 88 45 13.5 11.5 Co 50 63 60 48 57 51 55 49 58 64 56 51 65 69 62 77 62 59 55 58 Ni 263 115 171 45 52 43 45 86 48 195 67 51 58 203 74 170 91 67 46 47 Cu 62 92 121 111 103 117 119 101 114 119 115 121 106 90 107 101 73 61 118 120 Zn 56 60 56 96 103 99 105 89 89 129 122 120 104 103 100 115 111 116 102 121 Rb 23 9.6 13 149 189 136 139 62 49 72 82 87 46 38 39 42 37 38 127 103 Sr 176 172 159 491 561 462 423 542 527 744 732 843 1019 1155 1127 1203 1126 875 514 370 Zr 38.7 49.2 36.5 147 147 145 147 146 150 145 139 144 148 149 142 149 146 155 148 150 Nb 11.0 3.3 3.8 7.5 7.5 7.2 7.4 9.8 7.7 16.9 8.4 8.1 8.6 18.6 9.3 15.3 10.2 9.1 7.7 7.4 Ba 143 75 76 458 550 517 562 331 332 560 637 642 947 1000 961 1020 982 819 669 535 La 4.9 6.8 4.7 33.6 32.9 33.3 33.8 33.1 34.1 35.1 33.8 34.6 34.9 35.0 34.6 35.1 32.8 32.3 31.3 33.9 Ce 10.2 15.5 9.7 65 66 65 66 65 68 68 66 67 69 69 68 69 65 65 64 66 Pr 1.4 2.1 1.2 7.9 7.6 7.5 7.6 7.7 8.1 8.0 7.9 7.9 7.9 8.1 8.1 8.2 7.8 7.6 7.4 7.6 Nd 6.0 9.0 5.6 30.8 29.8 30.0 30.4 29.7 31.8 30.5 30.2 29.9 31.4 32.1 30.7 31.9 29.8 29.8 29.0 29.8 Sm 1.7 2.5 1.6 5.5 5.7 5.5 5.8 5.5 5.8 5.9 6.0 5.7 6.1 6.1 6.0 6.1 6.0 5.6 5.5 5.5 Eu 0.54 0.78 0.57 1.68 1.53 1.58 1.68 1.59 1.60 1.63 1.60 1.66 1.63 1.64 1.46 1.64 1.38 1.34 1.44 1.57 Gd 2.1 3.1 1.8 4.9 4.9 5.0 4.9 4.9 4.9 5.3 5.1 5.5 5.4 5.5 5.5 5.4 5.3 5.3 5.1 5.4 Tb 0.30 0.45 0.29 0.57 0.60 0.59 0.58 0.57 0.57 0.65 0.67 0.65 0.65 0.70 0.68 0.69 0.63 0.66 0.63 0.65 Dy 2.2 3.2 1.9 3.1 2.9 3.2 3.1 3.0 3.1 3.8 3.8 3.8 3.9 3.9 4.1 4.0 3.7 3.8 3.8 3.7 Ho 0.40 0.60 0.34 0.53 0.53 0.58 0.56 0.56 0.56 0.71 0.67 0.69 0.70 0.71 0.68 0.71 0.65 0.69 0.71 0.67 Er 1.30 1.87 1.07 1.46 1.42 1.50 1.51 1.56 1.48 1.85 1.79 1.90 1.92 2.06 1.93 1.94 1.92 1.88 1.92 1.89 Tm 0.19 0.25 0.15 0.22 0.20 0.23 0.21 0.20 0.21 0.26 0.25 0.24 0.26 0.26 0.25 0.27 0.28 0.26 0.26 0.26 Yb 1.12 1.65 1.06 1.43 1.38 1.36 1.44 1.38 1.43 1.64 1.66 1.56 1.71 1.61 1.66 1.66 1.63 1.69 1.70 1.70 Lu 0.17 0.23 0.14 0.23 0.21 0.21 0.23 0.21 0.23 0.25 0.23 0.24 0.25 0.25 0.29 0.27 0.27 0.24 0.25 0.25 Y 11.8 16.1 10.2 14.3 14.0 14.4 13.9 14.7 14.5 18.7 17.7 18.5 19.0 19.2 18.2 18.8 18.6 18.9 17.8 17.8 Hf 1.1 1.5 1.0 3.8 3.7 3.9 3.8 3.7 3.9 3.7 3.7 3.5 3.8 3.7 3.7 3.9 3.6 4.1 3.9 3.9 Ta 0.19 0.24 0.17 0.53 0.49 0.48 0.50 0.52 0.48 0.56 0.54 0.53 0.55 0.59 0.53 0.55 0.53 0.52 0.53 0.52 Pb 3.8 3.1 3.2 6.2 11.0 11.3 9.7 11.4 4.8 12.3 11.2 12.2 15.4 14.6 14.1 14.8 14.6 20.8 27.9 43.4 Th 1.2 1.4 1.0 6.5 6.8 6.5 6.7 6.7 6.7 6.9 6.8 6.7 7.2 7.1 7.1 7.2 7.0 7.2 7.0 6.8 U 0.33 0.52 0.32 1.2 1.1 1.1 1.1 1.1 1.1 1.2 1.2 1.2 1.2 1.2 1.2 1.2 1.2 1.3 1.2 1.2

Table 3: Geochemical analyses of the porphyric volcanic rocks of drill core AT-14-01. Porphyric Volcanic Rocks

Depth [m] 2188.00 2185.30 2182.50 2177.60 2134.00 2130.20 2126.50 1349.50 1340.10 1319.70 1277.60 1227.90 1204.60 1027.40 919.30 907.00 897.00 896.10 857.80 809.20 741.30 Sample AT-5 AT-6 AT-7 AT-9 AT-63 AT-30 AT-31 AT-199 AT-198 AT-197 AT-195 AT-194 AT-193 AT-187 AT-177 AT-172 AT-170 AT-206 AT-165 AT-153 AT-143 Major element oxides [wt.%] SiO2 66.4 68.9 67.5 67.4 66.1 68.0 65.7 43.9 61.3 59.5 67.0 61.5 63.4 63.7 61.8 65.0 63.8 62.8 59.4 63.6 59.2 TiO2 0.33 0.28 0.29 0.38 0.33 0.31 0.38 0.95 0.81 0.84 0.64 0.76 0.76 0.75 0.39 0.36 0.39 0.40 0.89 0.39 0.84 Al2O3 15.6 15.1 14.9 15.9 14.1 14.4 15.1 10.0 16.0 16.5 12.6 14.9 14.8 14.7 14.4 14.7 14.5 14.6 17.4 14.6 15.6 Fe2O3 (t) 2.8 2.3 2.5 3.5 3.2 2.6 3.2 17.2 5.1 6.4 3.7 5.0 6.0 4.1 3.4 3.4 3.6 3.9 8.2 3.5 8.8 MgO 1.6 1.3 1.2 2.1 2.0 1.6 2.1 6.0 4.4 4.7 2.7 3.0 3.4 3.4 2.7 2.1 2.4 2.4 4.7 2.4 5.2 CaO 2.6 2.9 3.0 0.5 3.2 2.4 2.3 9.9 2.3 2.1 4.1 4.1 1.9 2.8 4.0 2.9 3.5 3.4 0.5 3.8 0.7 MnO 0.04 0.04 0.04 0.02 0.05 0.04 0.04 0.25 0.06 0.06 0.07 0.08 0.05 0.06 0.09 0.06 0.07 0.07 0.01 0.07 0.05 Na2O 0.3 4.0 3.9 0.0 4.2 4.6 2.5 0.1 2.3 3.6 1.3 2.9 3.2 1.7 3.2 4.6 3.5 4.0 0.1 4.0 0.2 K2O 6.78 3.81 3.97 7.88 3.67 3.13 5.16 0.07 3.06 2.08 2.96 2.63 2.16 4.54 3.52 3.00 3.54 3.36 6.95 3.27 7.06 P2O5 0.08 0.08 0.09 0.06 0.12 0.10 0.12 0.13 0.12 0.09 0.14 0.14 0.08 0.13 0.16 0.13 0.15 0.15 0.16 0.14 0.03 LOI 4.1 3.3 3.3 2.3 3.4 2.8 3.1 11.7 4.7 4.7 5.6 5.3 3.9 4.9 6.9 3.8 4.3 4.4 2.9 4.4 2.7 Total 100.60 101.93 100.44 100.06 100.44 99.97 99.56 100.11 100.01 100.38 100.86 100.34 99.63 100.84 100.53 99.90 99.58 99.47 101.30 100.19 100.29

Trace elements [ppm] Cr 52 52.7 61 75 125 85 122 119 76 127 103 72 87 81 172 221 170 155 79 170 115 Co 25 40.7 28 36 40 30 18 49 27 34 29 24 27 19 24 31 25 27 23 23 27 Ni 18.9 22.0 29 32 54 35 53 122 47 81 66 48 55 45 53 80 55 41 47 58 77 Cu 3.3 13.1 21.4 13.1 14.4 10.0 6.9 130 23 19 23 41 36.3 12.7 17.1 8.3 23.7 36.6 8.8 10.2 4.1 Zn 28 32 31 43 55 44 50 204 66 78 44 51 83 51 28 49 47 48 107 43 110 Rb 214 158.5 133 231 87 80 135 3.0 91 62 85 79 60 112 104 78 101 100 206 97 149 Sr 108 286.3 290 53 446 302 200 257 83 122 86 123 97 77 182 319 326 406 15.5 251 39.2 Zr 133 11.3 139 176 137 123 137 210 307 321 247 290 291 279 192 179 188 194 287 188 258 Nb 4.2 4.6 4.2 5.1 6.1 6.1 10.7 9.0 11.7 9.0 8.7 9.2 8.6 5.7 7.0 5.5 4.5 8.3 5.4 8.5 Ba 1472 1065 934 1843 962 889 1659 12.5 429 283 399 348 424 557 815 741 817 849 1198 702 1646 La 14.7 20.2 18.3 22.3 19.3 18.5 21.8 14.4 24.7 15.1 19.1 25.0 19.3 23.7 37.1 35.6 37.1 41.2 25.8 39.5 13.4 Ce 30 40 35 41 38 37 44 32 53 33 39 50 40 48 73 70 74 83 54 77 30 Pr 3.6 4.6 4.1 4.6 4.5 4.2 5.3 3.9 6.4 4.2 4.7 5.9 4.8 5.8 8.5 8.2 8.7 9.7 6.4 9.0 3.7 Nd 13.4 17.7 15.2 17.6 16.8 16.6 19.9 16.4 25.2 16.9 17.6 23.7 18.8 22.0 34.4 30.9 32.3 37.1 24.7 33.4 14.9 Sm 2.3 3.0 2.7 3.1 3.2 2.9 3.4 3.3 4.9 3.9 3.5 4.2 3.6 4.1 5.2 4.7 5.1 6.0 4.5 5.7 3.1 Eu 0.65 0.80 0.80 0.88 0.83 0.72 0.82 0.93 1.32 0.78 0.76 1.10 0.84 1.11 1.32 1.35 1.37 1.60 1.13 1.43 0.83 Gd 1.7 2.4 2.2 2.3 2.4 2.2 2.6 3.3 4.3 3.9 3.1 4.1 3.2 3.8 3.7 3.4 3.8 4.2 4.2 4.0 2.9 Tb 0.18 0.27 0.21 0.24 0.25 0.25 0.28 0.48 0.58 0.56 0.43 0.51 0.44 0.46 0.37 0.30 0.36 0.38 0.55 0.35 0.38 Dy 1.0 1.3 1.2 1.2 1.3 1.2 1.4 3.2 3.3 3.2 2.6 3.0 2.5 2.6 1.7 1.4 1.6 1.8 3.3 1.5 2.4 Ho 0.18 0.26 0.21 0.18 0.25 0.23 0.23 0.72 0.67 0.66 0.54 0.54 0.45 0.48 0.27 0.22 0.28 0.29 0.65 0.28 0.48 Er 0.54 0.66 0.56 0.55 0.68 0.62 0.69 1.96 1.83 1.81 1.45 1.57 1.39 1.40 0.68 0.54 0.74 0.72 2.09 0.71 1.52 Tm 0.08 0.11 0.09 0.08 0.10 0.09 0.10 0.31 0.26 0.26 0.21 0.21 0.19 0.21 0.09 0.09 0.09 0.10 0.29 0.09 0.23 Yb 0.50 0.72 0.55 0.58 0.70 0.59 0.73 2.12 1.73 1.77 1.43 1.58 1.42 1.42 0.61 0.56 0.63 0.64 1.96 0.65 1.66 Lu 0.08 0.13 0.08 0.10 0.12 0.09 0.10 0.34 0.30 0.29 0.23 0.21 0.24 0.21 0.11 0.09 0.10 0.12 0.32 0.11 0.29 Y 5.3 4.4 6.2 5.5 6.7 6.3 6.6 17.7 17.0 17.4 14.1 14.6 12.4 13.8 7.9 6.8 7.7 8.0 19.7 7.6 13.5 Hf 4.1 7.0 4.1 5.4 4.1 3.8 4.0 5.3 7.5 8.0 5.9 7.2 7.1 6.7 5.0 4.8 5.1 5.3 7.1 5.0 6.2 Ta 0.33 0.44 0.34 0.38 0.31 0.3 0.31 0.74 0.68 0.72 0.56 0.62 0.62 0.61 0.29 0.28 0.28 0.31 0.61 0.28 0.50 Pb 5.1 4.0 9.4 13.4 6.5 3.4 4.5 2.8 1.6 2.0 1.5 2.8 2.9 2.4 5.3 5.9 5.0 8.4 1.5 3.4 1.4 Th 5.4 6.9 5.1 6.5 5.1 5.7 5.3 2.2 5.1 5.0 3.9 4.7 4.7 4.4 6.6 6.8 6.6 7.3 3.2 7.0 2.8 U 2.5 3.4 2.5 3.0 2.4 2.6 2.3 0.8 1.5 1.6 1.2 1.4 1.3 1.3 2.4 2.4 2.4 2.4 1.22 2.51 0.82 23

Table 4: Geochemical analyses of the aphyric volcanic rocks of drill core AT-14-01. Aphyric Volcanic Rocks

Depth [m] 1191.00 852.00 851.80 840.10 838.20 834.00 831.80 736.60 732.80 725.60 713.20 712.80 711.80 699.30 695.00 695.00 690.60 690.20 689.50 676.80 662.20 Sample AT-192 AT-200 AT-157 AT-211 AT-213 AT-201 AT-155 AT-161 AT-141 AT-140 AT-135 AT-136 AT-137 AT-126 AT-124 AT-124 AT-123 AT-128 AT-132 AT-122 AT-120 Major element oxides [wt.%]

SiO2 47.8 56.5 59.0 62.9 60.0 60.3 60.7 60.7 56.8 54.6 51.4 51.1 50.9 60.8 60.3 60.9 60.2 60.0 60.3 58.7 57.5 TiO2 1.40 0.96 0.73 0.85 0.75 0.81 0.79 0.82 1.27 1.13 1.13 1.09 1.10 0.81 0.81 0.82 0.82 0.82 0.74 1.09 0.76 Al2 O3 13.4 15.9 15.0 16.4 14.5 15.9 15.6 15.3 15.2 14.2 14.1 13.4 13.2 15.6 15.4 15.6 15.3 15.4 14.4 16.5 17.0 Fe2O3 (t) 14.9 8.5 8.1 7.9 7.1 6.4 8.4 8.8 12.2 11.1 16.7 18.2 19.6 7.6 7.5 7.5 7.3 7.6 7.1 8.1 8.0 MgO 7.3 5.8 5.7 4.0 4.3 3.9 5.0 4.9 4.5 4.2 6.0 6.3 6.7 5.2 4.9 5.0 5.0 5.2 5.7 4.3 3.0 CaO 6.3 1.1 1.0 0.7 3.0 2.6 0.5 1.0 1.3 3.5 2.2 2.4 0.8 0.7 0.8 0.8 0.8 0.7 0.7 1.2 0.6 MnO 0.18 0.05 0.06 0.04 0.08 0.08 0.04 0.05 0.08 0.12 0.12 0.13 0.09 0.03 0.04 0.04 0.04 0.04 0.04 0.03 0.04 Na2O 2.0 2.2 2.5 0.1

Table 5: Geochemical analyses of the volcaniclastic sedimentary rocks of drill core AT-14-01 and extremely altered rocks composed predominantly of chlorite. Volcaniclastic Sedimentary Rocks Chlorite fels

Depth [m] 1990.90 1989.20 1988.00 1987.40 1987.20 1977.30 1969.30 1965.60 1962.10 1961.80 671.20 660.90 660.80 2004.20 2002.50 2000.40 1996.00 935.80 Sample AT-217 AT-100 AT-84 AT-108 AT-73 AT-104 AT-106 AT-112 AT-111 AT-76 AT-121 AT-119 AT-118 AT-97 AT-103 AT-98 AT-219 AT-220 Major element oxides [wt.%] SiO2 46.6 44.5 63.6 47.4 46.0 54.6 41.7 41.4 58.0 58.1 49.5 39.9 41.4 45.0 40.1 41.7 51.7 38.2 TiO2 0.51 0.50 0.21 0.49 0.50 0.45 0.48 0.48 0.45 0.48 1.22 1.28 1.25 0.47 0.32 0.47 0.80 0.82 Al2O3 9.1 8.9 10.0 9.1 9.2 14.2 8.1 8.0 16.6 17.8 16.5 17.3 16.3 8.6 6.6 8.0 13.0 16.3 Fe2O3 (t) 10.1 9.1 19.8 7.3 7.4 12.5 11.3 9.4 7.4 7.6 15.7 23.4 24.1 7.3 6.6 8.8 19.9 27.0 MgO 11.9 13.8 2.6 15.3 15.4 6.7 12.0 13.7 6.3 5.7 7.6 6.4 6.5 15.0 19.8 14.7 8.5 10.1 CaO 6.3 6.5 0.4 6.0 6.3 1.6 7.8 7.8 0.9 0.8 0.6 1.5 1.0 7.2 8.8 8.7 1.0 0.5 MnO 0.13 0.14 0.03 0.11 0.11 0.03 0.25 0.22 0.03 0.02 0.07 0.13 0.12 0.14 0.13 0.17 0.07 0.09 Na2O 0.1 0.1

Different diagrams were used to determine the magmatic affinity of the igneous rocks in the studied drill core, two of which are shown below (Figs. 18,19). While the Nipissing Diabase has clearly a tholeiitic affinity, the affinity of the diorite samples is somewhat ambiguous. The diorite sills/dykes are classified as calc-alkaline using the Zr-Y-plot by Barrett & MacLean (1994). In the AFM ternary plot (Irvine & Baragar 1971), however, the diorite samples straddle the boundary between the tholeiitic and calc-alkaline differentiation trends, and even a shoshonitic affinity (high-K, high-Th) is suggested by the Th-Co-plot after Hastie et al. (2007) (not shown). In both the Zr-Y and AFM diagram, the porphyric and aphyric volcanic rocks are of calc-alkaline affinity, as it is also observed for the groundmass of the volcaniclastic sedimentary rocks. (Ultra-)mafic fragments of this clastic lithology show tholeiitic (Fig. 19) or even basaltic-komatiitic tendencies (Jensen 1976, not shown), but this might be due to metasomatic addition of

MgO to the rock and/or leaching of Na2O and K2O.

Figure 14: TAS diagram after Le Bas et al. (1986); Oxide concentrations calculated on a volatile-free basis.

Figure 15: Nb/Y vs. Zr/Ti – diagram for the studied rocks; fields after Winchester & Floyd (1977); symbols as in Fig. 14.

Figure 16: Zr/Ti vs. SiO2 - diagram after Winchester & Floyd (1977).

Figure 17: Nb/Y vs. SiO2 - diagram after Winchester & Floyd (1977), symbols as in Fig. 16.

igure 18: Zr vs. Y - diagram after Barrett & MacLean (1994).

Figure 19: Zr vs. Y - diagram after Barrett & MacLean (1994).

Figure 20: AFM triangle after Irvine & Baragar (1971).

Trace Element Patterns

Trace element patterns, normalized to the composition of the primitive Earth mantle (Sun & McDonough 1989; McDonough & Sun 1995), are presented below. As evident from these figures, all studied lithotypes, except for the Nipissing Diabase, have similar trace element patterns. These patterns exhibit the diagnostic features of (calc-alkaline) subduction zone magmatism, namely:

- A decoupling of large-ion lithophile elements (Rb, Ba, K, Sr, Pb) and high-field strength elements (Nb, Ta, Ti, Zr, Hf, U, Th) - The enrichment of fluid-mobile over fluid-immobile elements, demonstrated by positive peaks in Rb, Ba, Pb, and negative peaks of Ti, Nb and Ta. - Relatively steep REE-patterns (high La/Yb), typical of Archaean volcanic arcs.

Figures 20 and 21 illustrate the similarity between porphyric and aphyric volcanic rocks and volcaniclastic sedimentary rocks. Note how well the pattern of the volcaniclastic groundmass matches that of the other volcanic rocks. Despite the overall similiarities in the trace element patterns, the diorite dykes/sills can be clearly distinguished from all other igneous rocks in the drill core on the base of several features:

• The diorite is the only lithology which has a Th/UN > 1 and a Th/KN = 1 (where ‘N’ denotes the

primitive mantle-normalized value). All other igneous rocks in the core have a Th/UN ≤ 1 and

Th/KN < 1. Th concentrations in the diorite are among the highest of the entire drill core, while K concentrations are the lowest.

Figure 21: Primitive mantle-normalized trace element concentrations of the porphyric and aphyric volcanic rocks.

Figure 22: Primitive mantle–normalized trace element concentrations of the porphyric volcanic rocks and the volcaniclastic sedimentary rocks. • Concentrations of Sr are, compared to other rocks of the core, significant and 30-70 times higher than for the primitive mantle, but the Sr-anomaly itself is only weakly developed.

• Nb, Ta and Ti concentrations are exceptionally high relative to other lithologies of the core, but- their negative anomalies are less pronounced. Paradoxically, a volcanic arc signature, marked by an enrichment of Rb, Ba, K, Pb, and Sr, is still present. • The diorite has the highest concentrations of rare earth elements among all rocks of the drill

core, higher than any other igneous rock with a similar degree of differentiation (54 wt% SiO2).

The total REE content exceeds even that of the surrounding rhyodacite (66 wt% SiO2).

The diorite sills/dykes at the bottom of drill core AT-14-01 are obviously different to the Nipissing Diabase, both in terms of petrography and geochemistry. The extended trace element diagram (Fig.23) clarifies this difference. There all diorite analyses and their average, as well as the Nipissing Diabase samples of the same core, are plotted. This diagram shows that the composition of the Nipissing Diabase comes not even close to the diorite. The diorite is enriched in incompatible immobile by orders of mag- nitude relative to the Nipissing Diabase, and differs in several key parameters (U/Th, U/K, Nb/Ta, La/Yb, Sr/Sr*, K/K*, P/P*) from the Nipissing Diabase, even from its differentiated varieties. Comparison of the diorite’s compostion with literature data for the Nipissing Diabase have shown that the diorite of drill core AT-14-01 cannot be a differentiated (dioritic) variety of a Nipissing Intrusion.

The measured composition of the Nipissing Diabase (samples AT-113, AT-114, AT-115) is in very good agreement with its tholeiitic affinity and the Nipissing Diabase of ‘Emerald Lake Intrusion’ reported by Lightfoot et al. (1986, 1997) and Lighftoot & Naldrett (1996) (Fig.24). Its trace element pattern is smooth and flat, compared to the other rocks of the drill core, still positive peaks of K, Pb, Sr, negative anomalies of Ti, Nb, and Ta are present - a feature that is either ascribed to derivation from a metasomat- ically enriched mantle source and/or assimilation of country rocks (cf. Lightfoot et al. 1993).

Figure 23: Primitive mantle-normalized trace element concentrations of the porphyric volcanic rocks and the diorite dykes/sills.

Figure 24: Primitive mantle - normalized trace element concentrations of the aphyric volcanic rocks, the diorite dykes/sills and the Nipissing Diabase.

Figure 25: Primitive mantle-normalized trace element diagram of the diorite (red) and the Nipissing Diabase (black) of drill core AT-14-01; Shown is the compositional range (shaded area) of the Nipissing Diabase, for which only samples of <55 wt% SiO2 were considered. The Nipissing Diabase of the core has the same composition as the Nipissing Diabase of the Emerald Lake Intrusion (data from Lightfoot et al. 1986, 1993; Lightfoot & Naldrett 1996). Normalization values after Sun & McDonough (1989).

Isotope Geochemistry A total of 22 samples were analysed for their Sr, Pb and Nd isotopic compositions, 17 of them are a comprehensive selection of dioritic dyke/sill samples, five of them are a selection of different rock types collected from the SIC. All isotope analyses were carried out at the Department of Geological Sciences, University of Cape Town, following the analytical protocoll as described by Will et al. (2014). As all 147 144 Nd analyzed samples have a Sm/ Nd <0.13, one-stage Nd model ages (TDM ) were calculated by intersecting the quadratic depleted mantle curve of DePaolo (1988) with the linear εNd(t) evolution of the depleted mantle from εNd = 0 at 4600 Ma to a present-day εNd of +10. The following algorithm was used: Nd 143 144 147 144 TDM = (1/ε) * ln[{(( Nd/ Nd)Sample – 0.51316)/(( Sm/ Nd)Sample – 0.214)} + 1], where ε = 6.54 x 10-12 yr-1, 0.51316 is the present-day 143Nd/144Nd of the depleted mantle, 0.214 is the 147Sm/144Nd of the 143 144 147 144 depleted mantle, and Nd/ NdSample and Sm/ NdSample are the measured and calculated, respectively, present-day isotope ratios (DePaolo 1988).

The results obtained are given in Tables 6, 7 and 8, in which literature data from the SIC are also added for comparison. The Rb-Sr data are highly variable with Rb/Sr ranging from 0.033 to 0.337 and 87Sr/86Sr from 0.710 to 0.734. This wide range, which is also reflected by a corresponding range in 87Rb/86Sr, cannot be a primary feature but clearly reflects considerable mobility of Rb and Sr after rock formation. Sr isotopes do not provide any meaningful petrogenetic information and are therefore dis- carded from further discussions.

The Pb isotope data also reveal variability but not to the same extent as the Rb-Sr data. From the 207Pb/204Pb, 206Pb/204Pb, and 208Pb/204Pb ratios average μ and ω values of 11.38 and 34.29, respectively, can be calculated for the diorite samples. This is very similar to a diorite sample from the Copper Cliff Offset Dyke (11.94, 33.07) but differs significantly from the isotopic composition obtained on various samples from the main SIC mass. Some of this variation can again be attributed to some mobility of Pb, U and Th (see below).

Table 6: Nd isotope composition of the diorite dykes, reference samples from the SIC.

Sm Nd 143 144 147 144 TDM Sample No Unit Nd/ Nd Sm/ Nd ɛNd t0 ɛNd t1850 [ppm] [ppm] [ Ma ]

Drill Core AT-14-01

AT-14 Diorite Dyke 5.90 30.50 0.5113 0.117 -26.13 -7.19 2760 AT-15 Diorite Dyke 6.00 30.20 0.5113 0.120 -26.56 -8.38 2897 AT-17 Diorite Dyke 6.10 31.40 0.5113 0.117 -27.04 -8.22 2852 AT-19 Diorite Dyke 6.10 32.10 0.5112 0.115 -27.20 -7.77 2788 AT-20 Diorite Dyke 6.10 31.90 0.5113 0.115 -26.50 -7.24 2751 AT-21 Diorite Dyke 6.00 29.80 0.5113 0.122 -26.14 -8.34 2913 AT-22 Diorite Dyke 5.80 29.40 0.5113 0.119 -26.83 -8.44 2892 AT-24 Diorite Dyke 5.50 29.00 0.5113 0.115 -26.97 -7.49 2763 AT-25 Diorite Dyke 5.50 29.80 0.5113 0.111 -26.35 -6.13 2626 AT-36 Diorite Dyke 5.50 30.80 0.5113 0.108 -26.46 -5.38 2541 AT-38 Diorite Dyke 5.70 29.80 0.5112 0.115 -27.12 -7.87 2804 AT-39 Diorite Dyke 5.50 30.00 0.5113 0.111 -26.99 -6.60 2657 AT-40 Diorite Dyke 5.50 29.70 0.5113 0.112 -26.97 -6.85 2686 AT-41 Diorite Dyke 5.80 31.80 0.5112 0.110 -27.61 -7.09 2691 AT-58 Diorite Dyke 5.70 29.90 0.5113 0.115 -26.45 -7.11 2737 AT-59 Diorite Dyke 6.00 30.70 0.5113 0.135 -18.73 -4.07 2654 AT-64 Diorite Dyke 5.80 30.40 0.5113 0.115 -26.56 -7.24 2748 Average 0.5113 0.116 -26.27 -7.14 2750

Sudbury Igneous Complex

SB-12 Copper Cliff Offset Diorite 5.90 29.30 0.5113 0.122 -26.88 -9.09 2980 SB-06 Main Mass: Norite 3.64 20.36 0.5112 0.108 -28.64 -7.60 2712 SB-07 Main Mass: Quartz Gabbro 6.10 32.70 0.5112 0.113 -27.41 -7.48 2744 SB-08 Main Mass: Granophyre 9.20 50.00 0.5112 0.111 -28.29 -8.00 2771 SB-11 Main Mass: Quartz Norite 3.30 18.60 0.5112 0.107 -28.91 -7.68 2711

Faggart et al. (1985) Main Mass: Norite 3.381 19.979 0.5111 0.102 -30.00 -7.57 2663 Faggart et al. (1985) Main Mass: Norite 3.657 20.262 0.5112 0.109 -28.44 -7.63 2723 Faggart et al. (1985) Main Mass: Norite 2.657 24.760 0.5112 0.107 -28.44 -7.15 2669 Faggart et al. (1985) Main Mass: Norite 6.905 36.383 0.5113 0.115 -26.88 -7.40 2755 Faggart et al. (1985) Main Mass: Granophyre 9.002 53.240 0.5111 0.102 -29.42 -6.96 2618 Faggart et al. (1985) Main Mass: Granophyre 8.909 50.064 0.5112 0.107 -28.64 -7.47 2697 Faggart et al. (1985) Main Mass: Gabbro 6.322 34.642 0.5112 0.110 -28.44 -7.92 2756 Faggart et al. (1985) Main Mass: Gabbro 6.732 35.662 0.5112 0.114 -28.05 -8.43 2834 Faggart et al. (1985) Main Mass: Granophyre 10.076 58.634 0.5112 0.104 -28.83 -6.78 2617 Faggart et al. (1985) Main Mass: Gabbro 3.564 19.551 0.5112 0.110 -27.47 -6.91 2675 Faggart et al. (1985) Quartz Diorite Offset 12.047 96.456 0.5113 0.124 -25.91 -8.78 2989 Average 0.5112 0.109 -28.23 -7.55 2727

In contrast, the 143Nd/144Nd ratios of the analysed diorite samples are very uniform, suggesting only very limited Sm and Nd mobility. With a value of consistently 0.5113 they are identical to that obtained for the dioritic Copper Cliff Offset Dyke and extremely similar to literature data on the SIC. The calculated

εNd at 1850 Ma (the age of the SIC) is -7.14, the calculated Nd model age (based on a depleted mantle model) is 2750 Ma, which is similar to the average obtained for the SIC based on our new data and published data (2727 Ma) and a typical Superior Province age.

Table 7: Pb isotope composition of the diorite dykes and reference samples from the SIC.

Sample No Unit 206Pb/204Pb ±2σ 207Pb/204Pb ±2σ 208Pb/204Pb ±2σ μ ω

Drill Core AT-14-01

AT-14 Diorite Dyke 17.7970 0.0009 15.4805 0.0009 38.696 0.0032 12.25 36.52 AT-15 Diorite Dyke 17.7485 0.0008 15.4724 0.0009 38.825 0.0027 12.18 37.03 AT-17 Diorite Dyke 16.6996 0.0007 15.3429 0.0008 37.597 0.0025 10.67 32.17 AT-19 Diorite Dyke 16.7932 0.0007 15.3470 0.0008 37.798 0.0025 10.80 32.97 AT-20 Diorite Dyke 16.8109 0.0009 15.3652 0.0010 37.644 0.0031 10.83 32.36 AT-21 Diorite Dyke 16.7120 0.0007 15.3272 0.0008 37.622 0.0025 10.68 32.27 AT-22 Diorite Dyke 16.2197 0.0008 15.2713 0.0009 36.715 0.0029 9.97 28.68 AT-24 Diorite Dyke 16.0588 0.0008 15.2509 0.0010 36.169 0.0029 9.74 26.51 AT-25 Diorite Dyke 15.7691 0.0007 15.2173 0.0008 35.667 0.0024 9.32 24.53 AT-36 Diorite Dyke 18.8262 0.0008 15.5381 0.0008 41.150 0.0026 13.73 46.24 AT-38 Diorite Dyke 16.9720 0.0009 15.3335 0.0010 37.895 0.0028 11.06 33.35 AT-39 Diorite Dyke 17.2746 0.0008 15.3609 0.0009 38.340 0.0026 11.50 35.11 AT-40 Diorite Dyke 17.4073 0.0014 15.3785 0.0010 38.653 0.0031 11.69 36.36 AT-41 Diorite Dyke 19.3769 0.0009 15.5865 0.0009 40.258 0.0028 14.53 42.71 AT-58 Diorite Dyke 17.4739 0.0008 15.4418 0.0008 38.532 0.0028 11.78 35.88 AT-59 Diorite Dyke 16.8618 0.0008 15.3568 0.0008 37.921 0.0025 10.90 33.45 AT-64 Diorite Dyke 17.4853 0.0010 15.3845 0.0011 38.783 0.0036 11.80 36.87 Average 17.1933 15.3797 38.133 11.38 34.29

Sudbury Igneous Complex

SB-12 Copper Cliff Offset: Diorite 17.5796 0.0013 15.5994 0.0014 37.825 0.0057 11.94 33.07 SB-06 Main Mass: Norite 19.1500 0.0009 15.7024 0.0009 39.388 0.0031 14.20 39.26 SB-07 Main Mass: Quartz Gabbro 20.9754 0.0009 15.8795 0.0007 41.994 0.0024 16.84 49.59 SB-08 Main Mass: Granophyre 25.1296 0.0016 16.3026 0.0011 45.895 0.0046 22.83 65.04 SB-11 Main Mass: Quartz Norite 18.4904 0.0009 15.6843 0.0009 40.017 0.0030 13.25 41.76

Darling et al. (2010) Foy Offset: Diorite 15.525 0.002 15.192 0.003 35.661 0.010 8.97 24.50 Darling et al. (2010) Foy Offset: Diorite 15.840 0.003 15.196 0.003 37.006 0.009 9.43 29.83 Darling et al. (2010) Foy Offset: Diorite 15.225 0.002 15.112 0.003 36.209 0.010 8.54 26.67 Darling et al. (2010) Ministic Offset: Diorite 15.340 0.002 15.189 0.003 36.027 0.011 8.70 25.95 Darling et al. (2010) Parkin Offset: Diorite 16.896 0.003 15.394 0.003 37.580 0.009 10.95 32.10 Darling et al. (2010) Parkin Offset: Diorite 16.111 0.003 15.332 0.003 35.898 0.010 9.82 25.44 Darling et al. (2010) Hess Offset: Diorite 20.798 0.003 16.045 0.003 38.973 0.010 16.58 37.62 Darling et al. (2010) Hess Offset: Diorite 15.436 0.002 15.187 0.003 35.661 0.010 8.84 24.50 Darling et al. (2010) Frood-Stobie Offset: Diorite 15.965 0.003 15.457 0.003 35.742 0.010 9.61 24.82 Darling et al. (2010) Worthington Offset: Diorite 16.402 0.003 15.555 0.003 35.845 0.011 10.24 25.23 Darling et al. (2010) Worthington Offset: Diorite 16.757 0.003 15.634 0.003 36.136 0.010 10.75 26.38 Darling et al. (2010) Worthington Offset: Diorite 16.309 0.003 15.534 0.003 35.837 0.009 10.10 25.20 Darling et al. (2010) Worthington Offset: Diorite 16.743 0.003 15.628 0.003 36.250 0.010 10.73 26.84 Darling et al. (2010) Worthington Offset: Diorite 16.708 0.004 15.623 0.003 36.315 0.010 10.68 27.09 Darling et al. (2010) Copper Cliff Offset: Diorite 29.015 0.005 16.835 0.003 37.794 0.011 28.43 32.95 Darling et al. (2010) Copper Cliff Offset: Diorite 17.647 0.005 15.620 0.003 37.432 0.010 12.03 31.52 Darling et al. (2010) Copper Cliff Offset: Diorite 18.333 0.003 15.679 0.003 37.742 0.009 13.02 32.75 Darling et al. (2010) Copper Cliff Offset: Diorite 16.174 0.004 15.441 0.003 35.749 0.010 9.91 24.85 Darling et al. (2010) Manchester Offset: Diorite 25.312 0.004 16.712 0.003 43.649 0.011 23.09 56.14 Darling et al. (2010) Manchester Offset: Diorite 19.275 0.003 16.032 0.003 38.564 0.013 14.38 36.00 Average 17.791 15.620 37.004 12.24 29.82

Table 8: Sr isotope composition of diorite dykes and reference samples from the SIC. Rb Sr ±2σ Sample No Unit Rb/Sr 87 86 87 86 [ppm] [ppm] Sr/ Sr [%] Rb/ Sr

Drill Core AT-14-01

AT-14 Diorite Dyke 72 744 0.097 0.716 8 0.280 AT-15 Diorite Dyke 82 732 0.112 0.717 8 0.324 AT-17 Diorite Dyke 46 1019 0.045 0.712 9 0.131 AT-19 Diorite Dyke 38 1155 0.033 0.711 10 0.095 AT-20 Diorite Dyke 42 1203 0.035 0.711 9 0.101 AT-21 Diorite Dyke 37 1126 0.033 0.710 10 0.095 AT-22 Diorite Dyke 38 875 0.043 0.711 12 0.126 AT-24 Diorite Dyke 127 514 0.247 0.726 11 0.716 AT-25 Diorite Dyke 103 370 0.278 0.729 12 0.807 AT-36 Diorite Dyke 149 491 0.303 0.733 11 0.880 AT-38 Diorite Dyke 189 561 0.337 0.734 10 0.977 AT-39 Diorite Dyke 136 462 0.294 0.731 9 0.854 AT-40 Diorite Dyke 62 542 0.114 0.721 13 0.331 AT-41 Diorite Dyke 49 527 0.093 0.720 12 0.269 AT-58 Diorite Dyke 87 843 0.103 0.716 11 0.299 AT-59 Diorite Dyke 39 1127 0.035 0.711 12 0.100 AT-64 Diorite Dyke 139 423 0.329 0.733 12 0.953

Sudbury Igneous Complex

SB-12 Copper Cliff Offset Diorite 46.4 450 0.103 0.7469 10 1.4651 SB-06 Main Mass: Norite 49.3 233 0.212 0.7135 11 0.2985 SB-07 Main Mass: Quartz Gabbro 93.5 20.7 4.517 0.7246 8 0.6132 SB-08 Main Mass: Granophyre 58.4 182 0.321 0.7451 12 13.1169 SB-11 Main Mass: Quartz Norite 57 113 0.504 0.7278 9 0.9303

4. Alteration and Metamorphism

Assessment of Element Mobility

The macroscopic appearance of the rocks intersected in the studied drill core has already revealed that effectively all rocks therein experienced alteration to some extent. This is supported by our petrographic investigation. The intensity and type of alteration is, however, highly variable in different lithological units and ranges from low-grade metamorphic overprint to hydrothermal alteration by percolating fluids. As the focus of this study was on the geochemistry of igneous rocks that might bear a relationship to the SIC, a proper assessment of the extent to which the protoliths’ geochemistry might have been mod- ified during the various alteration events is pivotal for any further interpretation.

Alteration that is limited to those units interpreted as equivalents of the TGB might have a bearing on the potential of discovering VMS-type deposits in the region; it is, however, of little significance for the assessment of a possible SIC-link of the Temagami magnetic anomaly. In the latter respect, the critical lithological unit in the drill core is the intersected diorite. Before discussing the chemistry of the diorite samples in more detail, we shall model the mobility of elements during alteration of this specific rock unit

37 because the measured composition does not necessarily reflect the original composition of the igneous protolith.

Different approaches were used in order to determine the extent of element mobility during alteration and metamorphism of the rocks, including the procedures outlined by Cann (1970), Finlow-Bates & Stumpfl (1981) and Campbell et al. (1982, 1984). The latter approach, for example, revealed an allochemical behaviour of Cr, Ni, Pb, Rb, Sr and possbibly Ba, while all other elements remained seemingly immobile during alteration.

The Isocon method (Gresens 1967; Grant 1986, 2005) is another way to test for element mobility, to quantify relative element gains and losses, and volume and mass changes during metasomatism. To this effect the chemical compositions of least and most intensely altered rock pairs of originally identical lithology are compared. Ratios of immobile elements should be constant between the the least and most intensely altered samples and therefore plot on a straight line with a slope of 1 (Fig. 25a). These elements (e.g. Ti, Zr, Al, Zr, Hf, HREE) are then used to construct a so called ‘Isocon’ (a straight line of constant concentration through the origin) (Fig. 25b). It follows that those elements that plot above this line must have been added to the rock during alteration, those that plot below must have been lost. The distance from the isocon can be used as a measure for the amount of gain or loss. Application of this approach to the diorite samples revealed that most elements remained immobile during alteration. They can be used, therefore, to reconstruct the initial composition of the rock. Other elements, notably Sr, Ba, Rb, Ni, Cr,

Pb, and the volatile compononents H2O and CO2 were mobile to some extent and will not be considered for any further protolith characterization. The same approach was also applied to the Nipissing Diabas, the composition of which has remained more or less unchanged during alteration/metamorphism, except for the volatile components.

Secondary element mobility can also have considerable effects on the meassured isotope ratios. The isocon method has already demonstrated the mobility of Pb, Rb and Sr for the diorite. In Figure 26, the various isotope ratios of the analyzed diorite samples are plotted against the loss on ignition (volatile content, predominantly H2O and CO2), which increases with increasing alteration, i.e. hydration. Those samples with a LOI of around 1 wt% can be regarded as nearly unaltered, those with higher LOI values as altered. The latter samples are distinguished by elevated 87Sr/86Sr ratios, highly variable Pb isotope ratios and somewhat less variable Nd isotope ratios. These results not only indicate interaction with crustal fluids, as particularly evident from the Sr isotope data, but also provides a good indication of which samples can yield information on the original isotopic composition of this rock type.

The porphyric volcanic rocks, in contrast, have undergone significant metasomatic changes, as it is best demonstrated by the brecciated and chloritized sample AT-199 (1,349.5 m), and compared to the relatively ‘fresh’ sample AT-198 (1,340 m) of the same lithology. The identification of immobile elements is diffcult. Only a few element ratios plot near the 1:1 – ratio on Figure 27a. From this, it is possible to define two potential isocons (Fig. 27b), the first one has a slope of ~1 and passes through Ti, Y, Nb, Ta

38 and the HREE, with the implication that LREE, Zr, Hf, Al, U, Th and Si were removed by the fluid. The second possibility is an isocon through Zr, Hf, U and Al, which also implies a large change in volume (slope 0.65). Regardless of which isocon is chosen, it is evident that the composition of the rock has significantly changed. This holds true not only to elements that are susceptible to hydrothermal alteration

(K, Rb, Ba, Na, Ca, H2O, CO2, Sr, Mg) but also to usually immobile elements w. This is expressed, for example, by the fractionation of HREE/LREE, or Ti/Al and was probably caused by a high fluid-rock- ratio. This is not surprising, considering the the altered reference sample AT-199 is already macroscopi- cally extremely altered, much more so than all the other samples of the porphyritic volcanic rocks.

Figure 26: a) Element ratio - diagram and b) Isocon diagram for the least and most altered diorite samples (AT-24 and AT-59, respectively); see text for explanation. For a better presentation, the element concentrations (in ppm or wt%) were scaled with appropriate factors (e.g. 10TiO2, 0.1Ba).

Figure 27: Binary plots of volatile content (as an approximation for the degree of alteration) vs. radiogenic isotope ratios in diorite; alteration has caused a random scatter of isotope ratios around the original value. Only the Sr isotope ratio exhibits a linear correlation with the degree of alteration.

The normal case of sericitication and carbonatization of the porphyric volcanic rocks is represented by the sample pair AT-5 (2,188 m) and AT-9 (2,177.6 m) (Fig. 28). Elements with ratios near one (Si, Al, Y, Ti) define an isocon (Fig. 28b) of constant mass and constant volume (slope c. 1). The rock gained Na, Ca, Sr, Mn and volatiles, and became depleted in Pb, Ni, Cr, Ba, K, Fe and Mg. The apparent loss/mobility of LREE, Hf, Zr, Th, U might be an analytical artefact or owed to primary sample heterogeneities. An allochemical behavior of K, Rb, Ba, Na, Ca, Sr, and Pb is also supported by the approaches of Cann

(1970), Finlow-Bates & Stumpfl (1981) and Campbell et al. (1982, 1984). According to these approaches, the elements Zr, Hf, Th, U, Ti, Al, Si, Nb, Ta, and REE were immbolie on a larger scale.

Figure 28: a) Element ratio - diagram and b) Isocon diagram for the least and most altered porphyric volcanic rock sample (AT-198 and AT-199, respectively); see text for explanation.

The Gresens’ method was also applied to the aphyric volcanic rocks from 730 m to 850 m. As these samples are relatively uniform in composition and mineralogy, we chose to compare the average composition of several samples. The advantage of using average compositions is that nugget effects and other primary sample heterogeneities are minimized. An isocon was defined that passes through Zr, Ti, Al, Si and Yb, and has a slope of one (Fig. 29b). Hydorthermal alteration/metamorphism resulted in a gain of Ca, Sr, Rb, K, Th, U, Pb and volatiles and a loss in Na, Fe, Cr, Ni, Co, Mg and possibly Nb and Ta. Elements that were apparently not affected by alteration are the REE, Ti, Si, Al, Zr, Hf, Y, and Ba.

Figure 29: a) Element ratio - diagram and b) Isocon diagram for the porphyric volcanic rocks, see text for explanation.

The Gresens’ method is not applicable to the Nipissing Diabase, because this magmatic unit comprises a single but variably differentiated intrusive body without a uniform composition throughout. Thus it would be difficult to distinguish between magmatic fractionantion and secondary element mobility due to alteration. Cann (1970) pointed out that for a single intrusion such as the Nipissing Diabase, immobile elements should be highly correlated with the concentration of Zr (i.e. the least fluid-mobile element under various conditions). This approached revelead an immobile behavior of almost all elements for the Nip- issing Diabase, except for Rb, Ba, Ca, Sr and Pb.

Figure 30: a) Element ratio - diagram and b) Isocon diagram for the least and most altered aphyric volcanic rock samples, see text for explanation.

Additional use was made of alteration indices, such as, for example, the ‘Chemical Index of Alteration’ (CIA) after Nesbitt & Young (1982):

Al2O3 CIA = ∙ 100 Al2O3 + CaO′ + Na2O + K2O where element oxide concentrations in mole are calculated on an LOI-free basis.

The Weathering Index of Parker (1970) (WIP) is defined as

2Na2O MgO 2K2O CaO WIP = � + + + � ∙ 100 0.35 0.9 0.25 0.7

43 with element oxide concentrations in mole calculated on an LOI-free basis.

The ideal CIA for fresh rocks is <50, whereas the ideal WIP for unaltered samples is >100 (Price & Velbel 2003). A graphical comparison of these two indices is provided in Figure 30. The diorite dykes show little variation in their CIA and WIP and form a thight cluster near the array of unaltered rocks. Their CIA of 40-45 corresponds to a relatively unaltered rock, and the WIP (70-90) indicates somewhat stronger alteration. Little alteration and element mobility is also indicated for the Nipissing Diabase. In contrast to the Nipisssing and diorite intrusions, the majority of the porphyric volcanic rocks, and almost all aphyric volcanic rocks, define a trend towards strong alteration (↑CIA; ↓WIP). Volcaniclastic rocks are characterized by a large variance in WIP and CIA.

The intrusive rocks of the drill core differ significantly in their extent of alteration from the other rock types intersected in the borehole. To illustrate this, we use another two alteration indices, the Ishi- kawa Alteration Index (IA), introduced by Ishikawa et al. (1976),

MgO + K2O IA = ∙ 100 MgO + K2O + CaO + Na2O

and the Chlorite-Carbonate-Pyrite Index (CCPI) of Large et al. (2001)

MgO + Fe2O3 CCPI = ∙ 100 MgO + Fe2O3 + CaO + Na2O

Both indices are calculated from element oxide concentrations in weight percent on a LOI-free basis.

In Figure 31, the CCPI is plotted against the IA, together with mineralogical-chemical alteration trends (sericitization, chloritization, pyritization, albitization, carbonatization, etc.), and pure endmember compositions of secondary minerals (sericite, albite, pyrite, epidote, chlorite, etc.). The grey boxes mark the array of typical unaltered rocks (from basalt to rhyolite) as defined by Large et al. (2001). Two opposing styles of alteration and degrees of alteration are recognized for the rocks of drill core AT-14-01: The presumably younger diorite and Nipissing intrusions plot on or near the array of unaltered rocks and form tight clusters. They display an alteration trend towards epidote/calcite and/or chlorite/albite, in agreement with our petrographic observations. Some of the (probably Archaean) porphyric volcanic rocks, the bulk of the aphyric volcanic rocks, and chloritfels (chloritized volcanic rocks?) plot way outside the grey boxes, show alteration trends towards sericite and chlorite/pyrite and are characterized by a large scatter of data points. Such a phyllic and sulfidic alteration is a diagnostic feature of VMS deposits (Large et al. 2001; Shanks 2012; Buschette & Piercey 2016) and supports our idea of a correlation with the Temagami Greenstone Belt.

Figure 31: Chemical Index of Alteration (CIA) vs. Weathering Index of Parker (WIP) for the various rock types in the drill core AT-14-01.

Figure 32: Ishikawa Index (IA) vs. Chlorite-Carbonate-Pyrite-Index (CCPI), see text for calculation. All diorite analyses plot within, and the Nipissing Diabase samples near, the field of least alteration, in contrast to Archaean rocks of the drill core. Illustration based on Large et al. (2001), Shanks (2012), Buschette & Piercey (2016).

Metamorphism

Metamorphic Mineral Assemblages Stilpnomelane was identified by optical microscopy, electron microprobe analysis and x-ray diffraction, and occurs in paragenesis with Fe-chlorite, siderite/ankerite, quartz and magnetite. Since stilpnomelane is part of the peak-metamorphic mineral assemblage of the silicate facies BIF, it provides useful information about the grade of metamorphism that affected the Archaean portion of the core. According to Miyano & Klein (1989) the maximum stability of stilpnomelane is limited to 430-470°C and 5-6 kbar, which corresponds to conditions of the greenschist facies. At higher pressure and temperature, stilpnomelane reacts to form almandine garnet, amphibole, zussmanite or biotite (Miyano & Klein 1989; Klein 2005). The absence of such high-grade index minerals (or even their relics) precludes an amphibolite- facies overprint. Diagenetic and very low-grade metamorphic minerals of the sub-greenschist facies (greenalite, minnesotaite) are also absent, indicating temperatures of at least 200°C (Miyano & Beukes 1984).

Metamorphism of the oxide facies BIF led to a textural coarsening of the rock. Once cryptocrystalline chert re-crystallized to quartz with grain sizes of max. 150 µm, which is typical for BIF metamorphosed to the chlorite/garnet zone of greenschist facies (Klein 1973). Quartz is also dynamically recrystallized in the pressure shadow of slightly rotated pyrite framboides and pyrite porphyroblasts, which itself implies a minimum temperature of 300°C.

For chlorite layers in BIF and the feldspar porphyry, there is textural evidence that biotite formed by the prograde metamorphism of sericite and chlorite, indicating P-T-conditions of the greenschist facies and potentially metasomatic addition of potassium. No minerals of the sub-greenschist facies (prehnite, pumpellyite, zeolites) are preserved in the quartzofeldspathic rocks. Contact metamorphism at the immediate contact between the porpyhric volcanic rocks and the intrusive diorite did occur and lead to the formation of abundant biotite. The intrusive contact became sheared afterwards and re-healed with chlorite (pervasive and after biotite), stilpnomelane and calcite.

The Nipissing Diabase and the diorite dykes/sills have a typical mineral assemblage of the greenschist- to lower amphibolite facies of <500°C (Yardley 1989; Bucher & Grapes 2011):

epidote + biotite + amphibole + plagioclase ± chlorite ± calcite ± quartz ± leucoxene, and clinozoisite + amphibole + plagioclase + leucoxene ± chlorite ± quartz, respectively.

Garnet is absent, and microprobe analysis indicate that the composition of amphibole corresponds to actinolite – a typical mineral of the greenschist facies. At higher grades, however, the composition of amphibole should be hornblende, pargasite or cummingtonite, and garnet would have become a major constituent of the metagabbro/metadiorite (e.g. Fleet et al. 1986). The abundance of metamorphic biotite in the diorite reflects its higher K2O content, compared to the almost K2O-free Nipissing Diabase, and the presence of epidote or clinozoisite is probably a function of the Ca/Fe of the rock. The fact that the

Nipssing Diabase is less altered than the diorite could be explained by a deeper crustal level of emplacement for the diorite (∆ 2000 m), and differences in fluid permeability.

In summary, the entire drill core was subjected to metamorphism of the greenschist- to lower amphibolite facies and fits into the regional pattern of metamorphic zones mapped by Card (1978) and Easton (2000). Little deformation took place. A minimum age for the metamorphic overprint of the Sudbury Igneous Complex and other rocks of the Southern Province is c. 1,700 Ma (cf. Bailey et al. 2004; Piercey et al. 2007; Raharimahefa et al. 2014) and is attributed to the Penokean Orogeny from 1.9 to 1.7 Ga (James 1955; Cannon 1973; Fleet et al. 1989; Bennett et al. 1991; Shanks & Schwerdtner 1991; Riller et al. 1999; Schulz & Cannon 2007). Thus, the prograde metamorphism of the diorite dykes/sills implies a pre-Peno- kean emplacement age (older than 1.7 Ga).

Chlorite Thermometry

The AlIV site occupancy of chlorite, determined by electron microprobe analysis, was used to calcu- late temperatures of crystallization. A total of 140 analysis of the five different types of chlorite, as distinguished by optical microscopy, were carried out using a JEOL JXA 8800L Superprobe at the University of Würzburg. Listed below are the five types of chlorite:

i) Blue interference colours, monomineralic masses; occurs as layers in BIF. ii) Blue interference colours, pseudomorph after biotite; present in all volcanic rocks. iii) Green interference colours, monomineralic masses; veins and layers in BIF, intergrown with stilpnomelane. iv) Grey interference colours, probably after biotite and other ferromagnesian minerals; in volcaniclastic sedimentary rocks only, in paragenesis with dolomite. v) Purple interference colours, as veins with calcite; restricted to the diorite dykes/sills and their contact to the feldspar porphyry.

Chemical formulae for chlorite were calculated on a basis of 36 oxygen atoms per formula unit, and temperatures were calculated using the empirical equations by Cathelineau & Nieva (1985), Cathelineau (1988), Kranidiotis & McLean (1987), Jowett (1991), Zang & Fyfe (1995), El-Sharkawy (2000), Xie et al. (1997) and Hiller & Velde (1991).

The results for the different chlorite types and thermometers are summarized in Table 9. Tempera- tures range from ~250°C to ~400°C, which is in agreement with the temperature interval constrained from metamorphic mineral assemblages. However, the absolute temperatures of these calculations are less meaningful, more telling are the relative differences obtained for different chlorite types (Fig. 32). These might be due to the large range of Fe/Mg in the analyzed chlorite grains, and the limited chlorite compositions for which the thermometers were calibrated. It could not be proven that a principal requirement for the application of chlorite thermometry, that is, Al-saturation of the system, was fulfilled in all cases.

All of these limitations notwithstanding, it is possible that the noted differences reflect indeed different generations of chlorite which formed at different temperatures.

1 vein 0.9 after ferromagn. minerals C4 0.8 monomineralic 0.7 C5 0.6 0.5 C2 0.4 0.3 C1 Fe / (Fe + Mg) 0.2 0.1 C3 0 200 250 300 350 400 T [°C] Figure 33: Fe/(Fe+Mg) vs. formation temperature of chlorite using the equation of Kranidiotis & MacLean (1987). C1-C-5 refer to the different types of chlorite (see text and table 9). Also note the symbols in the left hand upper corner.

Table 9: Crystalliaztion temperatures of secondary chlorite caluclated from different thermometers (average temperature and 2σ standard deviation). 9 8 8 9 16 17 54 12 44 34 28 14 23 12 256 287 281 283 257 253 287 247 244 284 272 257 294 314 ES 25 26 11 78 12 17 64 45 37 21 35 11 13 19 321 311 413 404 407 295 302 297 414 408 390 320 424 454 Xie 9 8 8 9 16 17 12 44 34 54 28 14 23 12 256 257 287 281 247 244 253 283 287 258 294 314 284 272 Z&F 9 9 7 9 45 36 55 18 19 13 30 15 23 12 272 277 267 259 276 332 320 327 328 319 304 274 336 378 K&M 13 66 52 82 42 25 27 12 18 21 35 11 13 18 J 296 299 285 277 295 299 355 343 349 354 346 327 365 404 39 41 27 32 54 18 18 21 29 239 309 293 299 309 326 372 300 273 H&V 9 8 7 9 44 34 54 28 14 23 16 17 12 12 256 287 281 283 288 295 314 257 247 244 253 284 272 258 C&N 13 66 51 82 42 21 35 25 26 12 17 11 13 18 300 301 287 281 296 342 324 302 347 337 341 348 358 388 C 2 2 2 T [°C] 2 2 T [°C] 2 T [°C] 2 T [°C] T [°C] T [°C] T [°C] 2 T [°C] 2 T [°C] 2 T [°C] 2 T [°C] T [°C] 2 T [°C] 2 T [°C] 2 Mode of occurence after biotite/ferromagnesian minerals after biotite/ferromagnesian minerals after biotite/ferromagnesian minerals after biotite/ferromagnesian minerals after biotite/ferromagnesian minerals after biotite/ferromagnesian monomineralic domains monomineralic domains veins and Infillings calcite with fractures and Veins calcite with fractures and Veins calcite with fractures and Veins monomineralic layers in BIF monomineralic layers in BIF monomineralic layers in BIF with stilpnomelane with Lithology Volcaniclastic sedimentary rock sedimentary Volcaniclastic rock sedimentary Volcaniclastic rock sedimentary Volcaniclastic rock sedimentary Volcaniclastic rock sedimentary Volcaniclastic rock volcanic Aphyric rock volcanic Aphyric rock volcanic Porphyric dyke Diorite dyke Diorite dyke Diorite / BIF Chloritefels Chloritefels / BIF Chloritefels 4 4 4 4 4 2 2 5 5 5 5 1 1 3 Type

5. Discussion

Correlation with Temagami Greenstone Belt

Iron Formation

The Algoma-type iron formation of the magnetite-chert-facies (oxide facies) is the most important marker horizon of drill core AT-14-01. The presence of BIF within the Temagami Anomaly has already been predicted by geophysical models of Card et al. (1984). Figure 33 is the first derivative of the aeromagnetic map and demonstrates that the iron formation of the Temagami Greenstone Belt can be traced to the drill location, which leaves little doubt that the intersected iron formation is the same unit as in Temagami and Emerald Lake Greenstone belts. The BIF of the drill core dips steeply at angles of up to 80° towards the core axis, so does the iron formation of the Temagami Greenstone Belt on both limbs of the Tepaga Syncline. In addition, the true thicknesses of both formations are very similar (Bennett 1978; Hurley 1985). In the Emerald Lake Greenstone Belt, the BIF is hosted by massive intermediate volcanic flows and pyroclastic deposits (Ayer et al. 2006), which is also the case for the BIF in the drill core. The BIF of the core differs, in terms of mineralogy and texture, only little from its equivalent in the Temagami Green- stone Belt: Layers of Fe-silicate (chlorite, stilpnomelane) are common in both, as well as abundant carbonate (Bennett 1978; Bowins & Crocket 2010). In contrast to the Temagami iron formation, the intersected BIF is almost free of Fe3+oxides/hydroxides (hematite, goethite), and jasper is also absent.

Figure 34: First derivative of the aeromagnetic map, showing the Temagami magnetic Anomaly (TmA), the Te- magami Greenstone Belt (TeGB), the drilling site (yellow star) and the Emerald Lake Greenstone Belt (yellow rectangle). The iron formation of the Temagami Greenstone Belt is markes as a linear magnetic feature that can be traced for several kilomteres southwestwards to the Emerald Lake Greenstone Belt. Image source: Ontario Geo- logical Survey.

In spite of macroscopic evidence of some magnetite occupying secondary textural positions in parts of the studied core, our mineral chemical data do not confirm the presence of hydrothermally remobilized magnetite. Brecciated chert, embedded in a matrix of magnetite, shows clear evidence (high concentrations of Al, Ti, Zr, REE), of considerable detrital contamination of the BIF, either by pyroclastic material or pelagic sediment. Instead of representing hydraulic breccias cemented with hydrothermal magnetite, the observed BIF breccias are interpreted as sedimentary or deformational structures, analogous to those in the neighboring greenstone belt. Thus, this type of BIF breccia is not a local phenomenon related to the Temagami Anomaly but rather a common feature in the wider region. BIF breccia is also described for the Sherman Mine of the Temagami Greenstone Belt by S.A. Gilson, cited in Bennett (1978):

„Brecciation is often noted with fragments of iron formation cemented by silica and magnetite. The silica is often the same type that occurs in the iron formation. This suggests that brecciation has taken place, probably during initial deformation of and prior to final recrystallization of the iron formation.“

Clastic Sedimentary Rocks

Greywacke was intersected in bore hole AT-14-01 – a lithology which is also exposed in the Te- magami and Emerald Lake Greenstone belts, where it is tens of meters thick and found in the footwall of siltstone and shale (Bennett 1978). The only other known occurrence of shale with nodular pyrite is restricted to the Temagami Greenstone Belt and has been described by Bennett (1978), Schwartz (1995) and Bowins & Crocket (1994). This lithology is, without doubt, equivalent to the pyritic shale of the drill core.

Ultramafic Fragmental Rock

The volcaniclastic lithology from 654 m to 681 m and from 1,955 m to 1,993m is indeed of unusual petrography and geochemistry and received special attention in previous reports (e.g. Bite 2014). The rock was originally logged as diorite and considered a pontentially SIC-related unit and part of a ‘layered mafic complex’. A very similar rock type occurs also in the Temagami Greenstone Belt, always in contact with BIF. The following description is taken from Jackson & Fyon (1991, p. 455):

„On the north limb of the Tetapaga syncline, a unit of ultramafic rock, magnesium-rich (12 to 20 weight% MgO) flows and coarse-grained heterolithic fragmental rocks of uncertain origin, overlie the oxide-facies iron formation (…)“

The same type of rocks was also mentioned for the TGB by Bennett (1978, p. 5):

„An ultramafic ultramafic rock is exposed in the hanging wall of the West Pit in the Sherman Mine (…). This rock consists of centimetre-sized, ultramafic rock fragments which rest in a calcite [Dolomit?] matrix. Locally, distinct layering, accompanied with size grading,

was observed (…). Primary volcanic textures are not present, and this layering could represent bedding. It is not certain if the fragmental texture represents a tectonic fab- ric developed in an ultramafic sill, or if this rock is a metasediment.“

The following description and interpretation was given by Fyon & Cole (1989, p. 109):

„(…) an assemblage of fragmental, ultramafic rock, magnesium-rich pyroxenitic flows, and heterolithic clastic metasedimentary rocks, which conformably overlie the iron formation (…) may represent the initiation of a new volcanic episode (…).“

Fyon et al. (1988, p. 217) stressed the lateral persistence of this exotic lithology:

„The ultramafic fragmental and adjacent oxide-facies ironstone represent a regional, distinctive lithological marker couple, which is exposed along a 15 km strike length (…)“

The similarities between the ultramafic fragmental rock of the Temagami Greenstone Belt and the intersected volcaniclastic sedimentary rock are obvious: Both have fragments of BIF and altered (ultra-) mafic rocks, have high MgO concentrations (11.9 - 15.4 wt%) and carbonate contents, a bedding is indicated by the preferred orientation of biotite. The volcaniclastic lithology of the core has the same stratigraphic position within the volcano-sedimentary sequence as the fragmental rock unit of the TGB: imme- diately overlying the BIF at 680 m and 1,950 m. Although this rock type might appear unusual in the studied drill core, it is by no means exotic and in fact a common rock type in the Archaean basement of the wider region.

Conseuquently, we interpret the fragmental unit at ~680 m and 1,950 m as a volcaniclastic rock and the erosional product of the adjacent Archaean volcanic and sedimentary (BIF) rocks; the enigmatic (ultra- ) mafic fragments might be derived from mafic complexes such as the Kanichee (Ajax) Intrusion or could represent altered volcanic glass fragments.

Volcanic Rocks

Although no comprehensive geochemical data for the volcanic rocks of the Temagami Greenstone Belt are available, we also argue that the intersected volcanic rocks correlate with those of the Temagami and Emerald Lake Greenstone belts. In both outcrops, intermediate to felsic lava flows and pyroclastic equivalents are the most widespread rocktype, and pophyric textures are common. The petrography of the porphyric volcanic rocks of AT-14-01 closely resembles the description given by Bennett (1978, p. 30), and supports the idea of a correlation with the Temagami Greenstone Belt:

„In thin section, the quartz-feldspar porphyry is seen to contain up to 20 percent moderately strained quartz phenocrysts from 1 to 5 mm across. Quartz forms euhedral phenocrysts and deeply embayed partly resorbed grains. Euhedral phenocrysts of

albite (…) generally show little evidence of zoning, or of relict zoning, and have dispersed micaceous alteration (…). The groundmass consists of a fine mosaic of quartz, untwinned equant grains of feldspar, and abundant flakes and wisps of white mica.”

Apart from equivalent primary mineralogy and texture, the rocks have undergone the same type of hydrothermal VMS-type alteration as reported for the Temagami Greenstone Belt (cf. Bennett 1978; Hurley 1985; Fyon & Crocket 1986; Schwartz 1995). The alteration is described by Bennett (1978, p. 30) for the porphyric rocks of the TGB:

„Feldspar, presumably plagioclase, is reduced to “ghost-like” outlines of fine, micaceous material in a very fine grained matrix of quartz, feldspar, sericite, chlorite, epidote and carbonate.”

The volcanic rocks in drill core AT-14-01 seem to be flows sandwiched between metasedimentary rocks (greywacke, shale and BIF), which means that they are all part of the same volcano-sedimentary sequence. Further evidence for a correlation with the TGB comes from the magmatic affinity of the volcanic rocks in the studied drill core. Their calc-alkaline chemistry is a typical feature of the Neoarchaean greenstone belts in southern Ontario (e.g. Jackson & Fyon 1991). The only other known occurrence of volcanic rocks within the study area are the Palaeoproterozoic rocks of the Thessalon, Copper Cliff and Elsie Mountain Formation, at the base of the Huronian Supergroup. These volcanic rocks, however, are whithout exception of tholeiitic, and only rarely of alkaline affinity (Bennett et al. 1991; Jolly et al. 1992; Ketchum et al. 2013). In addition, the volcanic rocks of the Huronian Supergroup are never characterized by porphyric textures, but are massive, fine-grained and pillowed flows (Bennett et al. 1991).

In Table 10, the lithological diversity of the Temagami Greenstone Belt is compared with the lithotypes encountered in drill core AT-14-01.

Summarizing, the 1,490 meters of Canadian Continental’s 2,197 m deep drill core AT-14-01 into the Temagami Anomaly represent a volcano-sedimentary succession that can be correlated with the Archaean basement as exposed in the Temagami and Emerald Lake Greenstone belts, based on petrographic, lithostratigraphic, geochemical and geophysical criteria. Additional 630 meters are typical Palaeoprote- rozoic rocks belonging to the Huronian Supergroup and the Nipissing Diabase. The only lithology of the core, which has no equivalent, neither in these adjacent greenstone belts nor in the Huronian Basin, is the 25 and 52 m thick diorite intersected below 2,000 m. The nature of these exotic diorite dykes/sills requires further discussion.

Table 10: Comparison of rocktypes found in the Temagami Greenstone Belt and in drill core AT-14-01. Present Rocktype found in the TGB Reference in drill core BIF oxide facies [2,7,11,15,16] yes

BIF with jasper [2,13] no BIF with abundant hematite [2] no BIF sulfide facies [2,11,13] no BIF brecciated & recrystallized [2] yes Chlorite-stilpnomelane-layers (BIF silicate facies) [2,16] yes

Sedimentary Shale with nodular pyrite [2,12,17] yes Greywacke [2,9,10,13,14] yes

Fe-rich basalt* [2,3,4,5,14] no

Feldspar-megacrystic basalt [2,9,13] no Hyaloclastite breccias [2,6] no Layered pyroxenite-gabbro-anorthosite sills [2,9,13] no

Tholeiitic Matachewan Dyke (qz-tholeiitite) [2,12] no Nipissing Diabase (opx-gabbro) [2,14] yes

Layered ultramafic intrusion [1,2,3,4,5,14] no

Ultramafic fragmental rock [2,9,10,13] yes Komatiitic

Feldspar/quartz/hornblende-phyric [2,6,9,10,12,13] yes

intermediate volcanic rocks*

- Andesitic flows* [2,6,9,10,12,13,14] yes (Rhyo-)dacitic flows* [2,6,9,10,12,13,14] yes

Calc High-silica rhyolite* [8,12,13] no alkaline alkaline Archaean (!) diorite sills [2] no Basaltic flows* [2,6,9,10,12,13,14] no

Olivine-diabase (1.24 Ga Sudbury Dyke Swarm) [2,6,14] no

Carbonatite dyke (sovite) [2] no Spessartite (amphibole-plagioclase-lamprophyre) [2,6] no dykes Kersantite (biotite-plagioclase- [2,6] maybe Other in- Other lamprophyre) dykes Trondhjemite [2,14] no trusive phases trusive Quartz-monzodiorite [2,14] no Chlorite dykes [2] yes

*including pyroclastic equivalents. Neoarchaean units are green shaded. References: [1] Cabri & Laflamme (1974); [2] Bennett (1978); [3] Naldrett et al. (1979); [4] Naldrett (1981); [5] James & Hawke (1984); [6] Hurley (1985); [7] Blum (1986); [9] Fyon et al. (1988); [10] Fyon & Cole (1989); [11] Donaldson & Garrett (1991); [12] Schwartz (1995); [13] Jackson & Fyon (1991); [14] Ayer et al. 2006; [15] Bau & Alexander (2009); [16] Bowins & Crocket (2010); [17] Bowins & Crocket (1994).

A New Sudbury Offset Dyke

The diorite dykes/sills of drill core AT-14-01 are different in terms of petrography, geochemistry and isotopic composition, and can neither be correlated with the Nipissing Diabase, nor with any other igneous rock of the Temagami Greenstone Belt and the Huronian Basin. A correlation with the East Bull Lake Suite (e.g. the River Valley Intrusion) can be ruled out as well. In the following chapter we demonstrate, based on petrography, geochemistry and isotope data, that the diorite dykes/sills are most likely equivalents of the SIC’s quartz dioritic offset dykes.

Petrography

The mineralogy of the diorite dykes/sills is very similar to other (metamorphosed) offset dykes, namely biotite, amphibole, epidote, plagioclase, quartz, sericite, Fe-Ti-minerals (e.g. Grant & Bite 1984; Fleet et al. 1987; Wood & Spray 1998; Tuchscherer & Spray 2002; Hecht et al. 2008; Lighftoot 2016). The only aspect in which the diorite of the drill core differs significantly from known offset dykes is the absence of granophyric intergrowth between quartz and feldspar. Their absence, however, is likely due to the intense hydrothermal alteration and metamorphism of the diorite. Anhedral quartz crystals within the diorite dykes/sills might be relics of such granophyric intergrowths. Based on their petrography, we suggest a correlation with the quartz dioritic (‘QD’) offset dykes of the Sudbury Igneous Complex. An inclusion- bearing facies of the diorite (‘IQD’), as it occurs in some of the SIC’s offset dykes, has not been observed; Xenoliths are either absent or completely altered and re-crystallized; sulfides are only a minor constituent of the rock.

Trace Element Geochemistry

Figure 34 demonstrates the astonishing similarity between the diorite of drill core AT-14-01 and the composition of the QD offset dykes. There is a perfect fit when comparing the rare earth elements (REE), Zr, Hf, K, Nb and Ta. Some discrepancies are noted with regard to Ti, P and Sr, but these are explicable simply by variable amounts of accessory minerals (magnetite, ilmenite, apatite) and plagioclase, respectively. Differences in Sr contents are likely to Sr addition during carbonate alteration. In summary, the composition of the diorite is almost identical to other offset dykes around the Sudbury Complex. A relatively good overlap also exists between the average composition of the diorite sills/dykes and the impact melt preserved in the Onaping Formation of the Whitewater Group (Sudbury Basin fill) (Fig. 35).

Figure 35: Primitive mantle-normalized trace element diagram showing the average composition of the diorite in AT-14-01 (shown in red) and the average composition of other quartz-diorite offset dykes (with data from Lightfoot 2016).

In order to demonstrate the unusual nature of the diorite in AT-14-01 and the quartz diorite offset dykes, we compare their composition with the average composition of the middle continental crust (Gao et al. 1998). The plot (Fig. 36) shows that the composition of the offsets matches that of the middle crust, so does the diorite of AT-14-01. Such a fit leaves little doubt for the diorite in AT-14-01 being anything else but a SIC-related offset dyke and an intrusive impact melt of crustal origin. Further support for this comes from radiogenic isotopes (e.g. Faggart et al. 1985; Naldrett et al. 1986; Darling et al. 2010).

Figure 36: Primitive mantle-normalized trace element diagram showing the average composition of the diorite in AT-14-01 (red) and the average composition of vitric fragments, melt bodies etc. of the Onaping Formation (with data from Ames et al. 2002).

Figure 37: Primitive mantle-normalized trace element diagram showing the average composition of the diorite in AT-14-01, the Foy Offset (Lightfoot 2016) and estimates for the average composition of the middle continental crust (Gao et al. 1998).

Radiogenic Isotopes

Typical mantle derived igneous rocks are always characterized by a high Nd t of ≥0 (Fig. 37). The SIC is the only igneous complex worldwide that does not follow this trend: Faggart et al. (1985) and Naldrett et al (1986) were the first authors to demonstrate a constantly negative (very radiogenic) Nd (at 1850 Ma) for the SIC. This circumstance is due to the unique origin of the Sudbury Complex, that is, an asteroid impact by which the crust in the target area was completey remolten, forming now the Sudbury Igneous Complex. Model ages for SIC are therefore around 2.7 Ga, i.e. the age of the target rocks that made up the impact melt. New data collected by us on a few random samples from the SIC are in good agreement with the results of Faggart et al. (1985). The diorite dykes/sills of AT-14-01 share the same features as the SIC samples, namely highly radiogenic and negative Nd values and model ages of ~2.7 Ga (Figs. 38, 39).

Considerable Pb isotope heterogeneity is observed for the SIC reference samples (Main Mass Norite, Granophyre etc.). This circumstance has already been demonstrated by Darling et al. (2010, 2012), which the authors attributed to primary heterogeneities of the Sudbury impact melt sheet. Nevertheless, our diorite and SIC reference samples are in perfect agreement with the Pb-isotopes reported for other QD offset dykes (Darling et al. 2010) (Figs. 40, 41).

The diorite dykes/sills and the SIC were produced from the same source material with the same Sm/Nd and U/Th/Pb ratios at the same time. Otherwise the strong similarity with regard to Nd and Pb isotope ratios would be difficult to explain. Apart from the analogous trace element patterns, the Nd and Pb isotopes provide the most convincing argument that the drilled diorite is the product of an impact-related melt, possibly a new Offset Dyke.

10 BIO Depleted mantle trend 5 PD CHUR 0 GD ɛNdt Du S -5 SIC -10 Diorite of AT-14-01 Crustal trend -15

-20 0 1 2 3 4 5 Age [Ga]

Figure 38: Nd evolution curves for different model reservoirs (mantle and continental crust) and layered intrusions. Igneous complexes such as Duluth, Bushveld, Stillwater or Great Dyke, have high Ndt, indicating deviation from an enriched or depleted mantle source. The SIC is the only exception and lies on the crustal evolution trend. PD = Pallisades Diabase, New York; BIO = Bay of Islands Ophiolite; Du = Duluth Complex, Minnesota; GD = Great Dyke, Zimbabwe; S = Stillwater Complex, Montana. Illustration redrawn after Deutsch (1994); SIC data from Faggart et al. (1985)

Least Altered Samples Average Median

Figure 39: Whisker plot of Nd1850 for the diorite dykes/sills of drill core AT-14-01, our SIC reference samples, and the values reported by Faggart et al. (1985).

Figure 40: Whisker plot showing the model ages calculated for the diorite sills/dykes of AT-14-01, our SIC reference samples, and the values reported by Faggart et al. (1985).

Figure 41: Whisker plot showing the µ1850 for the diorite dykes/sills, the SIC reference samples (main mass), and the values reported by Darling et al. (2010) for other QD offset dykes.

Figure 42: Whisker plot showing the 1850 for the diorite dykes/sills, our SIC reference samples, and the values reported by Darling et al. (2010) for other QD offset dykes.

6. Outlook

The recognition of the diorite bodies within the bore hole AT-14-01 being similar to known offset dykes around the SIC has far-reaching consequences for future exploration strategies in the area considering the importance of these offset dykes as hosts of world-class Ni-Cu-PGE deposits. Bearing in mind the potential economic significance, further proof of the correlation suggested here would be beneficial. Ideally, the diorite in the drill core should be dated radiometrically. The only available age datum was obtained on monazite of uncertain genesis. Apart from that age being subject to a very (too) large uncertainty, it is suspected that it relates to a hydrothermal alteration event and not necessarily to a magmatic crystallization event. Consequently, a further effort to separate dateable minerals (zircon and/or baddeleyite) is recommended but would require a very large sample volume because of presumably very low concentrations of such minerals in the given rock.

Finally, to address the fundamental question of the likely cause of the Temagami magnetic anomaly, our results suggest that similar diorite as described here could be present at deeper levels and, thanks to its magnetic properties, provide a possible explanation of the Temagami anomaly. To ascertain this hy- pothesis, a drill hole considerably deeper than the 2,200 m reached in bore hole AT-14-01 would be nec- essary. Although the intersected BIF surely contributed to the magnetic anomaly and is probably respon- sible for small-scale positive magnetic features (traces that can be connected with those in the Temagami Greenstone Belt), it cannot explain the Temagami magnetic anomaly as a whole. The diorite intersected at the bottom of the bore hole is likely to be just the tip of a larger body of such high density and magnetic susceptibility at greater depth, which is suspected to be the ultimate cause of the magnetic anomaly.

If the newly diorite in the studied drill core is indeed a new Offset Dyke, as suggested in this study, it would open up an entirely new perspective on our understanding of the SIC and its exploration potential. The newly discovered Offset Dyke would be the first one to the east of the SIC and with a distance of some 50 km further away than any other Offset Dyke known so far. Noteworth is also its position at depths below 2 km, which is different to the majority of hitherto known Offset Dykes. It would be somewhat surprising if there were no expression of the studied diorite closer to surface in other areas of the Temagami magnetic anomaly. It is, therefore, suggested to carefully and critically (re-)evaluate surface outcrops or other drill core intersections of dioritic rocks in the wider region that in the past might have been erroneously ascribed to other igneous units, such as the Nippissing Diabase, or the Broom Island Diorite (and similar occurences) of the Temagami Greenstone Belt. The recommendation is to follow up on this aspect and establish a GIS-based 3-D model of the spatial distribution of rocks that are equivalent to the studied diorite in drill core AT-14-01, both from surface outcrops and further bore holes, while at the same time looking out for SIC-type sulfide mineralization or, at least, geochemical footprints of such mineralization. 62

Acknowledgements

We thank Winston and Wesley Whymark for their logistic support, company and guidance throughout the field visit, Andy Bite for sharing his life-long experience with the SIC, for showing us crucial outcrops of and around the SIC, Allan King, Peter Lightfoot and Richard Ernstfor a brief discussion of the geophysical characteristics of the SIC and the adjacent Temagami Anomaly. Canadian Continental Explora- tion Corp.’s financial support of this research is greatly appreciated.

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Appendix

[1] Primitive Mantle normalization values after Sun & McDonough (1989)

[2] Tm value (0.068 ppm) from McDonough & Sun (1995)

[3] Impact melt, melt bodies and vitric shards of the Onaping Formation from Ames et al. (2002), summarized in Lightfoot (2016)

[4] Different assumptions for the average composition of the middle continental crust from Gao et al. (1998)

[5] Average Composition of the Sudbury Offset Dykes, either free of inclusions (Quartz Diorite QD), or mafic inclusion-bearing (MIQD); data from Lightfoot (2016).

All values in ppm, except stated otherwise.

Least Altered Primitive Diorite Vitric Blocks Aphantic Dykes Basal Intrusion Vitric Onaping Middle Contintental Crust [4] Mantle [1][2] AT-14-01 and Bombs [3] [3] [3] [3] n=17 n=7 n=6 n=12 n=5 Rb 0.635 84.4 96 136 86 68 83.0 53.0 57.0 57.0 61.0 67.0 Ba 6.989 677.8 1096 1330 1512 602 801.0 616.0 619.0 619.0 590.0 661.0 Th 0.085 6.9 8.17 8.38 8.13 7.96 5.7 8.1 6.2 6.2 7.4 6.8 U 0.021 1.2 2 3 2 2 0.7 1.4 1.2 1.2 1.1 1.0 Nb 0.713 9.8 8.14 7.9 8.51 8 10.0 10.0 12.0 12.0 11.0 11.0 Ta 0.041 0.5 0.53 0.52 0.52 0.76 0.5 0.5 0.7 0.7 0.6 0.6 K % 0.03012 2.3 2.66 3.53 3.23 2.34 2.8 2.0 2.2 2.2 2.3 2.4 La 0.687 33.8 32.4 31.1 33.8 27.6 27.4 31.4 27.5 27.5 33.6 30.8 Ce 1.775 66.5 65.4 64.2 65.7 57.2 56.6 60.0 53.0 53.0 64.2 60.3 Pb 0.071 15.0 4 4 23 6 15.0 15.0 16.0 16.0 14.0 15.0 Pr 0.276 7.8 7.61 7.5 7.56 6.34 Nd 1.354 30.4 31.4 28.2 29.4 24.4 23.7 28.0 25.5 25.5 27.6 26.2 Sr 21.1 747.9 151 215 157 292 398.0 375.0 209.0 209.0 206.0 283.0 Sm 0.444 5.8 5.14 4.97 4.9 4.36 Hf 0.309 3.8 4.19 4.05 4.08 3.88 4.1 4.5 5.0 5.0 5.3 4.8 Zr 11.2 146.8 147 145 148 152 156.0 161.0 187.0 187.0 184.0 173.0 Ti % 0.2188 1.1 0.61 0.57 0.61 0.57 0.6 0.7 0.7 0.7 0.6 0.6 Eu 0.168 1.6 1.03 0.96 0.95 0.92 4.3 4.9 4.9 4.9 5.0 4.7 Gd 0.596 5.2 4.1 3.98 4.08 3.7 Tb 0.108 0.6 0.59 0.57 0.6 0.55 0.6 0.8 0.8 0.8 0.8 0.8 Dy 0.737 3.6 3.46 3.32 3.52 3.04 Ho 0.164 0.6 0.67 0.64 0.68 0.6 Y 4.55 17.0 19 18 19 18 17.3 15.7 16.7 16.7 17.2 17.0 Er 0.48 1.8 1.9 1.82 1.88 1.62 Tm 0.068 0.2 0.28 0.27 0.29 0.27 Yb 0.493 1.6 1.8 1.7 1.8 1.6 1.8 2.1 2.5 2.5 2.3 2.2 Lu 0.074 0.2 0.27 0.27 0.28 0.26 0.3 0.3 0.4 0.4 0.3 0.3

[5] QD QD Cliff Cliff Stobie MIQD MIQD MIQD MIQD hington QD Foy HMIQD Apophysis QD Parkin MIQD Foy MIQD QD Copper QD Copper QN Copper Manchester Incl. Of QD QD Of Incl. Cliff Funnel Cliff Funnel QD Ministic Worthington Worthington Worthingto Worthington Worthington of QD Wort- QD of MIQD Parkin MIQD MIQD FroodMIQD QD Creighton MIQD CopperMIQD QD Vermillion QD Manchester n=13 n=11 n=21 n=12 n=32 n=4 n=8 n=60 n=37 n=25 n=6 n=23 n=2 n=6 n=14 n=24 n=9 n=7 n=23 n=28 Rb 117 69 57 63 73 79 78 76 77 67 90 70 162 73 60 80 74 84 87 76 Ba 519 473 512 492 448 521 654 498 530 479 595 592 604 523 458 732 676 779 803 713 Th 8.64 8.07 6.81 7.46 8.22 8.27 9.01 8.71 9.14 8.33 8.54 8.56 9.40 9.67 7.20 8.07 8.06 8.91 8.47 8.26 U 2.04 1.85 1.47 1.67 2.09 2.09 2.13 2.09 2.21 2.10 2.06 1.93 3.04 2.14 2.04 1.40 1.42 1.85 1.63 1.57 Nb 10.40 10.30 8.90 9.40 8.50 8.40 10.30 10.30 10.80 9.60 10.40 8.60 9.90 16.80 9.00 8.10 7.90 8.80 8.40 8.40 Ta 0.74 0.68 0.66 0.69 0.56 0.57 0.69 0.70 0.73 0.67 0.73 0.60 0.72 0.98 0.73 0.47 0.44 0.53 0.51 0.49 K % 2.06 1.63 1.62 1.58 2.02 1.93 1.84 1.86 1.76 1.61 2.09 1.85 3.93 1.56 2.21 2.11 2.27 2.62 2.27 La 33.96 31.29 28.65 29.93 31.97 32.35 33.47 33.37 34.25 30.01 35.93 33.09 36.45 40.23 25.50 37.40 36.09 36.64 38.16 37.61 Ce 68.31 63.15 58.72 61.09 64.45 65.01 67.59 68.41 70.16 62.16 73.50 67.23 74.01 83.02 52.70 76.36 73.30 73.48 76.17 76.20 Pr 8.22 7.63 7.03 7.22 7.70 7.82 8.34 7.98 8.43 7.43 8.64 7.92 8.44 10.53 6.41 9.34 9.01 8.84 9.08 9.12 Nd 30.28 28.36 26.77 27.07 27.64 28.16 31.04 30.03 31.90 28.28 32.79 29.93 31.37 40.47 25.24 34.28 33.23 31.92 32.58 33.00 Sr 281 263 345 354 196 183 235 302 293 274 313 341 245 288 182 423 369 362 417 420 Sm 5.97 5.77 5.13 5.04 5.19 5.18 5.95 5.66 6.09 5.42 6.08 5.70 5.60 8.22 5.36 6.17 5.95 5.87 5.91 6.06 Hf 3.80 4.00 2.60 2.40 3.30 2.50 4.00 3.90 4.20 3.90 4.10 3.60 3.50 2.40 3.30 4.00 4.00 3.80 4.10 3.80 Zr 166 157 128 128 152 149 166 172 175 161 175 152 172 195 159 169 170 158 178 167 Ti % 0.85 0.95 0.72 0.72 0.70 0.68 0.86 0.88 0.89 0.86 0.89 0.82 0.80 0.93 0.75 0.73 0.74 0.72 0.75 Eu 1.49 1.48 1.39 1.31 1.32 1.39 1.45 1.45 1.47 1.29 1.51 1.42 1.55 1.67 1.28 1.52 1.45 1.49 1.48 1.52 Gd 5.54 5.58 4.91 4.90 4.57 4.58 5.34 5.25 5.67 5.11 5.54 5.04 4.69 7.79 5.21 4.98 4.78 4.95 4.92 5.04 Tb 0.86 0.88 0.74 0.73 0.69 0.69 0.81 0.80 0.86 0.79 0.85 0.77 0.67 1.22 0.84 0.70 0.68 0.72 0.69 0.71 Dy 4.91 5.08 4.30 4.10 3.76 3.78 4.55 4.60 4.95 4.57 4.96 4.40 3.63 7.45 5.36 3.81 3.76 3.85 3.57 3.68 Ho 1.04 1.10 0.89 0.84 0.77 0.77 0.90 0.95 1.01 0.94 1.03 0.91 0.74 1.50 1.14 0.74 0.71 0.79 0.71 0.74 Y 25.10 26.60 20.90 20.40 18.80 18.20 21.60 23.60 23.20 20.50 24.00 24.60 9.50 42.10 30.40 17.40 16.00 19.30 17.80 18.60 Er 2.80 2.98 2.54 2.46 2.06 2.10 2.57 2.66 2.85 2.66 2.86 2.56 2.01 4.40 3.22 2.02 1.99 2.06 1.86 1.92 Tm 0.44 0.47 0.38 0.36 0.31 0.31 0.37 0.40 0.42 0.39 0.45 0.39 0.31 0.65 0.50 0.29 0.28 0.31 0.28 0.28 Yb 2.80 2.99 2.49 2.49 1.94 2.00 2.47 2.52 2.74 2.56 2.65 2.51 1.93 4.24 2.97 1.88 1.85 1.97 1.75 1.78 Lu 0.42 0.46 0.38 0.37 0.29 0.30 0.38 0.38 0.41 0.38 0.40 0.38 0.31 0.64 0.47 0.29 0.29 0.31 0.27 0.27 72