The Geology of the Middle Precambrian Thomson Formation in Southern Carlton County, East-Central Minnesota

THE GEOLOGY OF THE MIDDLE PRECAMBRIAN THOMSON FORMATION IN SOUTHERN CARLTON COUNTY, EAST-CENTRAL MINNESOTA

A THESIS SUBMITTED TO THE FACULTY OF THE GRADUATE SCHOOL OF THE UNIVERSITY OF MINNESOTA

BY MARC ROBERT CONNOLLY

IN PARTIAL FULFILLMENT OF THE REQUIREMENTS FOR THE DEGREE OF MASTER OF SCIENCE

December, 1981 ABSTRACT

In southern Carlton County, east-central Minnesota, the

Thomson Formation provides evidence for major deformational and metamorphic events. These events were associated with the Penokean

Orogeny which terminated, or closely followed, Middle Precambrian sedimentation in the Animikie Basin approximately 1870 m.y. ago

(Morey and Lively, 1980).

The following eight Thomson Formation rock types were found in

the study area: metasiltstone, phyllite, calcareous metasiltstone and phyllite, graphitic phyllite, volcaniclastic metasiltstone and

phyllite, and metabasalt. Stratigraphic relationships and thickness

estimates of these units were indeterminable due to lack of strati-

graphic control, poor and discontinuous exposures, and regional

folding.

The phyllitic rock types were deposited slowly as shales in

quiet water below wave base. Deposition of this material was inter-

mittently interrupted by deposition of silt beds (since metamorphosed

metasiltstones) by southward flowing distal turbidity currents.

Local calcareous silts and limy muds were later metamorphosed

calcareous metasiltstones and phyllites, respectively. The graphitic

phyllites probably developed from organic-rich muds. Also, hypabyssal

basaltic dikes and/or sills (since metamorphosed to metabasalts) occur

locally and may have been a source rock for detritus for the

i ii volcaniclastic sediments found nearby. Volcanic vents associated with

these intrusives may also have been a source of pyroclastic detritus

for the volcaniclastic units.

Three Penokean deformations and their resulting structures have

been recognized in the study area. During the first Penokean deforma-

tion, beds of all Thomson Formation units were folded into overturned

or recumbent isoclinal folds. These folds have axes that plugge at low

angles to the east and west, and axial planes that are horizontal or

dip at low angles (up to 40 degrees) to the south. A regional, con-

tinuous schistosity developed during this deformation as an axial-

planar foliation to these isoclinal folds. As a result, the schisto-

sity is found parallel to bedding at most (but not all) outcrops. Also

associated with this event was the development of oblate quartz

boudins which lie in the plane of the axial-planar schistosity.

During the second Penokean deformation, bedding and the

essentially bedding-parallel schistosity of the first deformation were

together folded into two distinct and contemporaneous fold geometries.

Folds of one geometry are tight and asymmetrical or overturned. These

folds have axes that plunge at low angles to the east and west, and have

axial planes that dip 10 to 40 degrees to the south. Folds of the

other geometry are open and symmetrical or asymmetrical. These folds

also have axes. that plunge at low angles to the east and west, but

their axial planes are approximately vertical. A discontinuous

crenulation cleavage developed during this deformation as an axial-

planar foliation to folds of both geometries. iii

The third and final Penokean deformation was minor relative to the first two. No folds developed during the·third deformation, but a nearly vertical crenulation cleavage that strikes north-south, mineral lineations that plunge at low angles to the south, and four joint sets were produced.

All Thomson Formation units in the study area have been regionally metamorphosed. The grade at culmination of metarnorvhisrn was uppermost greenschist/lowermost amphibolite facies (garnet zone).

Regional metamorphism and tectonism were_contemporaneous events.

Metamorphic culmination was reached before, and continued during and after, the first Penokean deformation. Metamorphism then continued during the second and third Penokean deformations at relatively lower greenschist facies grades. TABLE OF CONTENTS

ABSTRACT ...•• i

TABLE OF CONTENTS . iv

LIST OF PLATES .. ix

LIST OF TABLES .. ix

LIST OF ILLUSTRATIONS . x

ACKNOWLEDGEMENTS. • xiv

CHAPTER ONE:

INTRODUCTION. 1

Statement of the Problem . 1

Study Area Location and Physiography 1

Method of Study. 2

Field Work •• 2

Laboratory Studies •• 7

Previous Work. • . • • . 7

CHAPTER TWO:

REGIONAL GEOLOGY. 10

I. Middle Precambrian Regional Stratigraphic Correlations and Ages 10 ..· Animikie Group. 10

The Relation of the Animikie Group, the Marquette Range, and the Huronian Super group. . • . • 14

II. Geology and Structure of East-Central Minnesota. 16 General Statement ...... 16 iv v

Lower Precambrian Rocks ..•..... 16

Middle Precambrian Stratified Rocks. 22

General Statement .. 22

Penokean Deformation and Metamorphism. • • 24

Mille Lacs Group • . 26

Denham Formation • • 26 Glen Township Formation. . . • • 28 Little Falls Formation . • • 28

Trout Lake Formation • • 28

Animikie Group 28

Mahnomen Formation . 29

Trommald Formation • • 29

Rabbit Lake Formation •• • • 29

Thomson Formation. • • • • • 30

Middle Precambrian Plutonic Rocks •. 32

Upper Precambrian Sedimentary and Volcanic Rocks 32

Nopeming . . 33

Ely's Peak Basalts • • 33

Chengwatana Volcanic Group . 33

Fond du Lac Formation .. . • 33

•. Hinckley Formation 33

Keweenawan dikes . . • 33

CHAPTER THREE:

PETROGRAPHY. 34

I. Introduction . 34

Rock Types • 34 vi

Distribution of Rock Types .. 34

Laboratory Methods ... 37

II. Lithologic Descriptions. 38

Metasiltstone and Phyllite. 38

Calcareous Metasiltstone and Phyllite . 45

Graphitic Phyllite. 45

Metabasalt •.•.• 46

Volcaniclastic Metasiltstone and Phyllite • 48

Concretions . 53

Quartz Veins, Lenses, and Pods .• 54

Post-Tectonic Basaltic Dike • • . 54

CHAPTER FOUR:

STRUCTURE • 57

I. Introduction. 57

Structural Patterns • . 57

II. Description of Structures. 57

•• 57

Foliations. • 58

Schistosity 58

Crenulation Cleavage .. 64

Lineations ... 64 Crenulation Lineations...... 69 Mineral Lineations. . . . . 69 Minor Folds 72

Style 1 Minor Folds • 72

Style 2 Minor Folds 72 vii

Style 3 Minor Folds ...... • 80 Joints • ...... 80 Faults 84

Quartz Boudins . 86

CHAPTER FIVE:

INTERPRETATION . . 89

I. Introduction 89

II. Sedimentology . . "' . . . 89 Protoliths ...... 89 Provenance ...... 90 III. Stratigraphy...... 91 IV. Regional Metamorphism ...... 91 v. Introduction to Structural Analysis...... 91 Method of Structural Analysis. • 91

Notation . • • • • . 93

Deformations • . 93

Fold Generations . • 94

Bedding and Foliations . • 94

Lineations . . . 94

Structural Analysis: The Three Deformational Events and Resulting Structures of the Penokean Orogeny Found in the Study Area • • • . • • • . • . 94

First Deformation (D ) 94 1 Second Deformation (D ). , . 98 2 . . . Third Deformation (D ) , . .105 3 viii

VII. Relative Time Relationships Between the Three Penokean Deformational Events and Regional Metamorphism •.. • 112

Method of Interpretation. . 112

First Deformation (D ). . 112 1 Second Deformation (D ) . 113 2 Third Deformation (D ). • 113 3 VIII. Structural and Tectonic Relationships Between the Study Area and the Rest of the Thomson Formation. • 114 ·" Introduction. . . . 114

First Deformation (D ) Structures ..•.• • 114 1 Second Deformation Structures .••••. . . . 115

Third Deformation (D ) Structures • • 116 3 IX. Regional Structura1 Models . . . 117 Introduction. • • • • , • • • 117

Separate Deformations Model • • 117

Progressive Deformation Model • . 119

x. Penokean Tectonic Models ••. • 119

Introduction •. 119

Plate Tectonic Model ••• • 119

Intracratonal Model . • • 121

Application of the Regional Structural Models to the Penokean Tectonic Models. • • . • . . • • • ••• 123

CHAPTER SIX:

CONCLUSION. 125

REFERENCES CITED .•. 127 LIST OF PLATES

Plate

1. The geology of the Middle Precambrian Thomson Formation, southern Carlton County, east-central Minnesota ••.•••••••• (in pocket)

LIST OF TABLES

Table ·

1. Stratigraphic nomenclature and inferred correlation of Middle Precambrian sedimentary rocks in Minnesota and ajoining Ontario • • • • • • • • • 12

2. Generalized correlation chart of Middle Precambrian rocks in Minnesota, Wisconsin, and Michigan. • • 13

3. Modal mineralogy ranges of rock units (continued on page 36) • • . 35

ix LIST OF ILLUSTRATIONS

Figure

1. Map showing extent of east-central Minnesota dealt with in this report. . • • • • •••• 3

2. East-central Minnesota showing counties and towns referred to in this report • • • • • • 4 J:: 3(a). Index map of study area • • • • 5 (b). Generalized outcrop map of study area • 6

4. Generalized map of Minnesota showing extent of Animikie Basin sediments • • • • • • 15

5. Generalized bedrock geologic map of Minnesota •• 17

6. Map of Lake Superior region showing the three Lower Precambrian lithostratigraphic segments ••• 18

7. Pretectonic north-south cross section showing lithostratigraphic nomenclature and depositional phases of the Middle Precambrian stratified rocks in east-central Minnesota • • • • • • • • • • • • • • • 23

8. Stratigraphic column of parts of Carlton, Pine, Aitkin, Kanabec, and Mille Lacs Counties, east-central Minne so ta...... 2 7

9. Cross section of the Thomson Formation showing the nature of the folding near the area of the type locality (i.e. Thomson, Minnesota). • • • • • • 31

10. Interbedded metasiltstone and phyllite 39

11. Photomicrograph of metasiltstone showing schistosity • 39

12. Photomicrograph of phyllite showing schistosity. • • • 40

13. Photograph of a quartz pod oriented in a plane parallel to bedding and the bedding-parallel schistosity. • 41

14. Photomicrograph of a large post-schistosity chlorite flake in metasiltstone discordant to the schistosity 43

15. Photomicrograph of pre- (or early syn-) tectonic garnets in phyllite. . • • • • • . . • • • • . . . . 43

x xi

Figure

16. Photomicrograph of syn-tectonic (snowball) garnet in phyllite. • • • • • • • • • •• , • • • • • • 44

17. Photomicrograph of post-tectonic garnet in metasiltstone 44

18. Photomicrograph of metabasalt •• 47

19. Photomicrograph of relict plagioclase phenocryst in metabasalt • •.••...•.••.•... 47

20. Photomicrograph of a round, polycrystalline epidote porphyroblast in metabasalt...... '; 49 2L Photomicrograph of a relict amygdule in metabasalt . 49 22. Photomicrograph of volcaniclastic phyllite ...... 51 23. Photomicrograph of a relict detrital plagioclase grain in volcaniclastic phyllite ...... 51 24. Photomicrograph of a basaltic rock fragment in volcaniclastic phyllite. • • • • • • • • •• ...... 52 25. Photomicrograph of metamorphosed concretion in metasiltstone. • • • • • • • • • 55

26. Equal area projection of 250 poles to bedding. 59

27. Schistosity discordant to bedding (figure continued on page 61) • • • • • • • • • • • • • • • • • • • • 60

28. Close-up of Figure 27 showing schistosity discordant to bedding • • • • • • • • • • • • 62

29. Equal area projection of 495 poles to schistosity. 63

30. Schistosity discordant to bedding. 65 '· . . . . 31. Close-up of Figure 30 showing schistosity discordant to bedding. • • • • • • • • 66

32. Photomicrograph of crenulation cleavage. 67

33. Crenulation cleavage in thinly interlaminated beds of phyllite and metasiltstone •••••••••• 67

34. Equal area projection of 111 poles to crenulation cleavages. • • • • • • • • 68

35. Photomicrograph of conjugate crenulations in phyllite. . 68 xii Figure

36. Map view of two intersecting sets of crenulation lineations • • • • . • 70

37. Equal area projection of 411 crenulation lineations. . 70

38. Equal area projection of 30 ferroactinolite lineations • 71

39. Muscovite streak lineations found only at Outcrop 20-1 • 73

40. Equal area projection of 7 muscovite streak lineations . 73

41. Style 1 minor folds .• • t:. 74

42. Style 1 minor folds. • 75

43. Close-up of Figure 42 showing Style 1 minor folds which have an axial-planar schistosity • . • • • . • • • • 76

44. Equal area projection of 43 Style 1 minor fold axes. 76

45. Equal area projection of 43 poles to Style 1 minor fold axial planes • • • • 77

46. Style 2 minor folds •• 71

47. Style 2 minor fold (antiform) •• 78

48. Close-up of Figure 47 in hinge area of a Style 2 minor fold (antiform) • • • • • • • • • • • • • 79

49. Equal area projection of 21 Style 2 minor fold axes •• 81

50. Equal area projection of 21 poles to Style 2 minor fold axial planes...... 81 51:. Style 3 minor fold (antiform) • . . . . . 82 52. Style 3 minor fold (open and asynnnetrical antiform). . 82 53. Equal area projection of 54 Style 3 minor fold axes. . 83 ·54, Equal area projection of 54 poles to Style 3 minor fold axial planes...... 83 55. Equal area projection of 600 poles to joints . . 85 56. Oblate quartz boudins in phyllite...... 87 57. Oblate quartz boudins in phyllite...... 88 xiii Figure Page

58. Structures of the first Penokean (D ) •••.•• 95 1 59. Structures of the second Penokean deformation (D ) 2 superimposed over the D structures. • • • . . • • • . . • 99 1 60(a). Outcrop along the Kettle River showing bedding and the essentially bedding-parallel schistosity striking approximately east-west and dipping to the south. • • • 101 (b). Bedding and the essentially bedding-parallel schistosity together are folded into an F -generation 11 2 Z" minor fold. • • • • • • • • • • • • • • • • • • 102 .c (c). Diagram combining Figures 60(a) and 60(b) to give macroscopic F A fold geometry and approximate size. 103 2 61. Structures of the third Penokean deformation (D ) 3 superimposed over the n and D structures • • • • • • • • 106 1 2 62. Photomicrograph of the s crenulation cleavage crenulating the s crenulation3 cleavage indicating 2 that s3 does post-date s2 ••••••••••••••••• 107

63(a). Block diagram of theoretical joint sets that could develop in a tectonite body • • • • • • • • • • • 109 (b). Stereogram of these theoretical joint sets. • 109

64. Stereogram of the study area joint sets •• 110

65. Separate deformations model •• • 118

66. Progressive deformation model. 120

67. A possible reconstruction of the margin of the Archean craton about 1900 m.y. ago during the early stages of the Penokean Orogeny (before intrusion of plutonic rocks in the volcanic arc and associated metamorphism .. and deformation) • • • • • • • • • • • • • • • • • • • • • 122 ACKNOWLEDGEMENTS

I would like to thank Dr. Timothy B. Holst for being my advisor. I also extend my appreciation to Dr. John C. Green, Dr.

Richard W. Ojakangas, and Dr. G. B. Morey for their critical review of my writing. I also thank the Minnesota Geological Survey for ·, partial funding of the thesis. My deepest appreciation goes to my father,

H. Robert Connolly, to whom this thesis is dedicated.

xiv CHAPTER ONE

INTRODUCTION

Statement of the Problem

The Middle Precambrian Thomson Formation shows the effects of several major deformational and metamorphic events. These events are associated with the Penokean Orogeny which terminated or closely followed Middle Precambrian sedimentation in east-central Minnesota

(Morey, 1972). This study is a detailed analysis of Penokean deformation in a key area of the Thomson Formation where it has never been closely examined. Minor structures seen in the field and micro- structures seen in thin section yield clues as to what regional structural patterns are in the study area, thus increasing our understanding of the deformational history of the Penokean Orogeny. A detailed structural and petrographic study such as this will help complete the detailed picture of the Thomson Formation and add to the understanding of the Penokean Orogeny.

• Study Area Location and Physiography

The study area is located near the towns of Moose Lake and

Kettle River, in southern Carlton County, east-central Minnesota.

More specifically, the area consists of Sections 15-17, 20-29, and

32-36 of Silver Township (T.46N., R.20W.), and Section 19 and parts of

Sections 20, 21, 29, 30, and 31 of Moose Lake Township (T.46N., R.19W.)

1 2

(see Figs. 1, 2, and 3 (a) and (b)). Access to Moose Lake is easily gained from either the north or south via Interstate Route 35 or U.S.

Route 61. TI!.e study area is well dissected by roads making travel within the area relatively easy; however, most outcrops are accessible only by cross-country hiking.

Southward-flowing rivers in the study area cut through approximately 43 meters of Pleistocene glacial cover exposing (see

Fig. 3 (a) and (b)). Most exposures are found along these rivers and their tributaries. Key outcrops in the west are located along the banks of the Kettle River and its tributaries Split Rock River,

Glaisby Brook, and Silver Creek. In the east, scattered outcrops are found along the western banks of the Moose River, with key outcrops located in railroad cuts on the north side of the town of Moose Lake.

No outcrops were found in the central upland region between the Moose and Kettle rivers due to thick glacial cover. TI!.e north-south drainage direction is fortunate because most regional structural patterns trend east-west. As a result, outcrops exposed along these rivers give excellent indications as to what regional structural patterns are in the area.

•.Method of Study

FIELD WORK

Field work began in mid-July and continued to mid-November,

1980. Many outcrops were found through the use of field notes and publications of previous workers. Field notes of Schwartz and

Knutson (1940) on the Thomson Formation and the publication of Harder 3

St. Lo&Ais- --, -"'·'"', : I "-, I I \ 1 1 !111SdARD : CAS"S I I --.-• r----..i r---- I I .. I I .\' AITkllV I CARI.Tow Q I I ,.0 I 3 I I : . ------...... , ... .,I 0 8 I I -, 'PlllE I -- -, r-v-, I l tl I I •mol?AISOAI I :s IJ ,. I I I -- J. _ .!;..s;---, :C I 1 ... r---1-- :r- ""---' L,-t; SrEA RWS ;.r11£1/Bu"1

Figure 1. Map showing extent of east-central Minnesota dealt with in the regional geology chapter of this report. Counties and county boundaries shown. (See also Fig. 2.) f.ltA88A,qD I Sf. Lo".rs I I I I CA S s - --- - ,- .J r------1 I I t------1 c':r•&f I I c.,./h.,, I I .. I CARLTON rs.-... I .. ArTXIN •Atio ...... , WA D£ft/lf I I I I .... l'J''- ,,.,.. I #faftlc. .,.,.... 1.i· ·""'iii I 1ir • oG."n I •Bo.,..11W"'I I I moo,C 'I ... - -- CAOW I --- -- '...:1 ... I • !'\:'\_snrnv '\ - ,.-, o.,.,._ "' -.J AAEA I ",..H. c.. .., . I I 'L- - r-----. I '-t" I 1#-.,r,.. w.n "f'"Oi'D tnoltRI.SO(V 1 . mrLLE • J •/11,. .. lflcr t..I l.'1U r Pi•ra• r 1 1 ICANl'IBEC. I I I I r------1 I I I I \ I r----t---- $f. WanJa/;I 8 I 0 2$ 50 • I I 1 1 I S r E A R IVs s-.,.-f.lf\ .. s.... k fJ. ..,; J, I :Ki lo111ater.1 its.. ------.J ... -, I I ::CS' A I CHZS'Alio A;.4,..o,.J P,i.-.kv:l/c , I • c:1J Spri"t ", SHER 81.(RA/E. I

Figure 2. East-central Minnesota showing counties and towns referred to in this report. Also shown ·is the study area (see also Fig. 3). RJI 111.I A. ').oW· A.').OWI

0 1 2. .3 L I I I

1"1 Kilometers

1.0 'L3 I '1.'f SttA.tc. P.o ... te. "l.';r Sf..,te

"1.f i.• "').";! 'L' 1.S"

3J "J't I

-r.l./bN. CAP.LToftl Co. :)£. ,./ I / I. 'lbN. r.'lsM PINE co. . al ./. 1'". '(SN. A.?-1 wl R. '}.OW.

lJ1 Figure J{a). Index map of study area. kewe.e.naWo..t1 l!lA 1B-r ">"'1 r Ou f crop /o<..Q.f/011I f11AM6ef'S' Keff/e River' [kB?J Basa/f Di'/re.. l'e.Fe.,.,.el +o ,;, +e°ICt (F,·,.,+ fll.i/t\ b&I" :=;: SeGfiol1 IJWWJ/,e,.,) • MP€ lhomsofl formo.fion I mp I ntefasi!fsfotie./fhylCfe. (ma! H- \.;) caf,t1ho4$ m1ttasilt.sfo11ft c.a11i'c.f '.-\- /f 1i;t11:+e. vf) I I 15"" lvmPI IT I ') I @] GrAf/,;hc.. Pti1 ft;fe Phy/life. -:P (Jc;. :.Joe- l I Lo.ke ?.O 'l.3 I '1-'t 'l.').. Sto..te. AotAte. ?..1

I 1.1> 1.s-· c... mp @:>::,_-ca.111P ®·'11. 1.•-lf'I . . "l.'1·'11" Vmp Co.mP+mP '® ?l.11·Y'/ I :ss- 31 . W"'P L I ./ 0 1 2 ;; I CAALTON co. j,/ 7 L I I I PINE c.o. Kilometers Figure J(b). Generalized outcrop map of study area (see also Plate 1). Outcrop location numbers are those referred to in text; outcrop extents are °' exaggerated; and stratigraphic relationships of Thomson units are unknowno 7 and Johnson (1918) on east-central Minnesota geology were used extensively. Other outcrops were identified .on topographic maps as areas of steep slopes and substantial relief. Systematic traverses of likely areas and personal communication with local inhabitants also resulted in location of many outcrops.

Over 80 outcrops were located and examined (see Fig. 3 (b) and Plate 1) and notes on lithology, outcrop dimensions, and appearance were taken at each outcrop along with representative samples. Many oriented samples were collected for later microfabric analyses.

Structural data collected include measurements of orientations of bedding, foliations, lineations, minor fold axes and axial planes, joints, and faults.

LABORATORY STUDIES

From the samples collected, a representative suite of 83 thin sections was made and studied. Detailed microscopic and mesoscopic descriptions of major rock. units were made using these thin sections and hand samples. To process the large amount of structural data collected during field work a computer program was used which stores, arranges, and plots planar and linear data on equal area nets. These nets will be used in this report.

Previous Work

Winchell (1894) assigned the name "Thomson Slate" to rock exposed in the village of Thomson on the east side of the St. Louis

River in Carlton County, east-central Minnesota. Later, Schwartz

(1942a) re-classified the Thomson Slate as the Thomson Formation 8 because greywacke and greywacke-slate are more abundant than slate.

Exposures of Thomson Formation are found in parts of Carlton, Pine,

and southern St. Louis Counties (see Fig. 2). The northern exposures

of Thomson Formation have been extensively studied (Morey and

Ojakangas, 1970; Wright and others, 1970; Keighin and others, 1972).

Southern Thomson exposures in and around the study area have been

studied by relatively few workers.

Information on the Thomson Formation exposed in and around the

study area first appeared in the literature early in the twentieth

century. Harder and Johnson (1918) published a report on east-central

Minnesota geology which contains locations and lithologic descriptions

of Thomson outcrops. In 1940, G.M. Schwartz and Ray Knutson conducted

a field survey of the Thomson Formation which included the study area

(their field notes are on file with the Minnesota Geological Survey).

Schwartz (1942b) later published a paper dealing with regional aspects

of the Thomson Formation which integrates data collected in the study

area. Schwartz (1942b and 1942c), and later Weiblen (1964), also

studied carbonate concretions found throughout the Thomson Formation.

They documented a north-to-south progression in metamorphic grade

the type locality in the north to the thesis area near the town

of Moose Lake. Lithologies in this metamorphic progression were

later described by Morey (1978).

Keighin and others (1972) compiled a report dealing with

regional geologic relationships found in Minnesota that

incorporates structural and petrographic data taken from the study

area. Morey (1979) and Davidson (1979) located and described certain 9

Thomson outcrops in and around the study area. The most recent literature that integrates data taken from the study area includes a report (Morey and Lively, 1980) on the geology and geochemistry of selected uranium deposits in east-central Minnesota. CHAPTER TWO

REGIONAL GEOLOGY

I. Middle Precambrian Regional Stratigraphic Correlations and Ages

Animikie Group -" The term "Animikie" was first used by Hunt (1873) for argillite and sandstone in the Thunder Bay area on the north shore of

Lake Superior. Argillites in the Gunflint Lake area of northeastern

Minnesota were correlated with the Thunder Bay sediments by Bell

(1873). Later, sediments in the Mesabi Range area of Minnesota were correlated with Animikie rocks at Thunder Bay by Irving (1883).

Irving (1883) was also first to suggest that sediments of the Thomson

Formation in east-central Minnesota were equivalent to the Animikie strata at Thunder Bay.

The term "Animikie" was first formally used in Minnesota for a formation having the following three members (Grant, 1899): a basal taconite member; an intermediate black slate member; and an upper greywacke-slate member. However, the Animikie was redefined in

Minnesota by a number of workers as being a group consisting of several formations. In the Gunflint Range area Van Hise and Clements

(1901), Clements (1903), and Van Hise and Leith (1911) redefined

Grant's basal taconite member as the Gunflint Iron-Formation and combined Grant's intermediate black slate and upper greywacke-slate

10 11 members as the Rove Slate. At about the same time, rocks of the

Mesabi Range were also assigned to the re-defined Animikie Group and divided into the following three formations: the basal Pokegama

Quartzite (Winchell and Upham, 1888); the intermediate Biwabik Iron-

Formation (Van Hise and Leith, 1901); and the upper Virginia Slate

(Van Hise and Leith, 1901).

Correlations made within the Animikie Group vary in degree of -" certainty (see Tables 1 and 2). Stratigraphic continuity between the

Gunflint and Mesabi ranges is well established, with correlations between these two Animikie Group localities widely accepted. However, correlation between the Cuyuna and Mesabi Range localities are less certain; but, Schmidt (1963, p. 39) has found:

11 no reason to consider a general correlation of the Trommald (of the Cuyuna Range) and Biwabik (of the Mesabi Range) formations invalid. • • 11

Grout and others (1951) assigned the Thomson Formation to the - Archean Knife Lake Group. However, Thomson Formation sediments in east-central Minnesota were later correlated to the Middle Precambrian

Animikie Group. Morey and Ojakangas (1970) determined that the

Thomsdn Formation is lithologically and sedimentologically similar to the Rove Formation of the Gunflint Range, thus strongly suggesting that the two are correlative. Morey and Lively (1980) have inferred that the Rabbit Lake Formation of the Cuyuna Range passes transitionally upward into the Thomson Formation. Also, radiometric dating of the

Thomson Formation by Goldich and others (1961) yield Middle Precambrian ages comparable with those of the Virginia and Rove formations of the

Mesabi and Gunflint Ranges, respectively. Therefore, Morey (1978) has · GUNFLINT RANGE MESABI RANGE CUYUNA RANGE EAST-CENTRAL Minnesota and Ontario Minnesota Minnesota MINNESOTA (Goldich and others, 1961; (White, 1954) (Grout and Wolff, 1955; ( Goldich and others, 1961 i Tanton, 1931) Schmidt, 1963) Morey and Ojakangas, 1970)

UPPER PRECAMBRIAN SEDIMENTARY AND IGNEOUS ROCKS (younger than 1.6 b. y.)

______..:.______unconformity------

Rove Formation Virginia Formation Rabbit Lake Formation Thomson Formation Gunflint Iron-formation Biwabik Iron-formation Trommald Formation "Kakabeka Quartzite" Pokegama Quartzite Mahnomen Formation

------unconformity------

Trout Lake formation quartzite and slate

------unconformity------

LOWER PRECAMBRIAN IGNEOUS AND METAMORPHIC ROCKS (older than 2.6 b. y.)

Table 1. Stratigraphic nomenclature and inferred correlation of Middle Precambrian sedimentary rocks in Minnesota and ajoining Ontario (Morey, 1972).

t-' N l\lESABI RANGE CUYUNA RANGE GOGEBIC RANGE DICKINSON COUNTY MARQUETTE RANGE

UPPER PRECAMBRIAN SEDIMENTARY AND IGNEOUS ROCKS (younger than 1.6 b.y.) ------:---__:_------unconformity------Virginia Formation Rabbit Lake Tyler Slate Badwater Greenstone Formation Michigamme Slate g.{Michigammc Slate .... 0 (;I .... ::s0. 0 Hemlock Formation i:i:i 0 Goodrich Quartzite 0 II) ---_____ disconformity------·

Biwabik Iron-formation Trommald Ironwood Iron- Iron-

Trout Lake formation Bad River Dolomite Wcwe Slate quartzite and slate? .Sunday Quartzite Randville Dolomite °')Kana Dolomite 0 ::s Sturgeon Quartzite g 8 Mesnard Quartzite .c: 0 Fern Creek FormationU Enchantment Lake Formation

LOWER PRECAMBRIAN IGNEOUS AND METAMORPHIC ROCKS (older than 2.6 b.y.)

Table 2. Generalized correlation chart of Middle Precambrian rocks in f-' Minnesota, Wisconsin,. and Vdchigan (Morey, 1972)e w 14 concluded that the Thomson Formation is the down-basin extension of at least the upper part of the Animikie Group of. the Animikie Basin (see

Fig. 4).

The Relation of the Animikie Group, the Marquette Range, and the Huronian Supergroup

The accepted group name for the Middle Precambrian stratified rocks in Upper Michigan is the Marquette Range Supergroup (Cannon and ,C Gair, 1970). James (1954) found similar iron-formations in both Upper

Michigan and Minnesota, and later determined that the sediments of

Upper Michigan resemble Animikie strata (James, 1958). Earlier,

Aldrich (1929) found similar stratigraphic successions in Upper

Michigan and in the Gogebic Range of Wisconsin. Therefore, it appears that the Animikie Group of Minnesota is equivalent to at least the upper part of the Marquette Range Supergroup. Radiometric ages of

Animikie rocks range from 1635 + 24 m. y. (Faure and Kovach, 1969) to 1900 ± 200 m.y. (Hurley and others, 1962) and estimated times of Marquette Range Supergroup deposition range from 1900 to 2100 m.y. ,. ago.{Banks and Van Schmus, 1971), thus indicating that parts of the

Marquette Range Supergroup may be synchronous with the Animikie Group.

Substantial evidence exists indicating that neither the

Anirnikie Group nor the Marquette Range Supergroup are equivalent to any part of the Huronian Supergroup of Ontario, Canada. James (1958) determined that sediments of Upper Michigan bear little resemblance to Huronian strata, and Cannon and Gair (1970) found no evidence for a Marquette Range and Huronian Supergroup correlation. Also, radiometric age data indicates that the Huronian Supergroup is older than 15

0 50 100 r Kilometers

Figure 4. Generalized map of Minnesota showing extent of Animikie Basin sediments (dashed where inferred) (after Morey, 1972). 16

2000 m.y. (Van Schmus, 1965), and time of deposition for the upper part of the Huronian Supergroup (the Gowganda Formation) was estimated by Fairbairn and others (1969) to be 2288 + 87 m.y. ago.

II. Geology and Structure of East-Central Minnesota

General Statement

11le Precambrian rocks of east-central Minnesota are divisible into the following four distinct terranes (Morey, 1978) (see Fig. 5):

1. Lower Precambrian rocks; 2. Middle Precambrian statified rocks consisting of thick, folded, and metamorphosed sedimentary and volcanic rocks lying unconformably on the Lower Precambrian rocks; 3. Middle

Precambrian plutonic rocks which intrude the Middle Precambrian stratified rocks; and ·4. Upper Precambrian sedimentary and volcanic rocks which dip in an eastward direction and lie unconformably on the

Middle Precambrian rocks. Cambrian and Cretaceous sedimentary rocks that locally overlie the Precambrian rocks in east-central Minnesota will not be discussed in this report.

Lower Precambrian Rocks

11le Lower Precambrian rocks of Minnesota can be separated into the following three distinct lithostratigraphic segments by two east- northeast-trending high-angle fault zones of Early Precambrian age

1978) ·(see Fig. 6): 1. a northern segment; 2. a central segment; and 3. a southern segment.

11le northern Lower Precambrian segment consists mostly of granite with lesser metasedimentary and metavolcanic rocks of the greenstone-granite terrane of northern Minnesota and adjacent Ontario 17

cl ..0 >< cl 0

: Cretoceou• -"ola OAd a.andltON ..a: 0 Paleo1oic 101\dstone. •"al .. z lirnutone, and dolomU1

rockt-.,_. lormorion "' bloc.II W9r Granitic roci.1 .·: ····:· me•oomy•ackt, and docltlc ::¥!.!; :· woicaniclostic roc.ks--Mainly grtenldli&f fldn Gntin--Granuli.'• and amphiboilt fodn

---Conlod

., 1001111.lS I I I 50 IOO IULOMlTllts

I 0 WA

Figure 5. Generalized bedrock geologic map of Minnesota (Morey and Sims, 1976}. r, kl I otr1et.r.s- -- _J l c A r A D 0 100 7,00 - \ Mlr

() \

? •

?•

• 'ft !t (!) f'"' ...... l> l 'WI s C Orv'.S tN IOWA -\ I Figure 6. Map "of Lake Superior region showing the three Lower Precambrian lithostratigraphic segments: 1. northern segment (greenstone-granite terrane}; 2. central segment (Great Lakes Tectonic Zone}; and 3. southern ..... segment (gneiss terrane)o Modified from Sims and others (1980)e 00 Midcontinent gravity high (diagonal rule} from Chase and Gilmer (1973)0 19

(Morey and Sims, 1976). Geophysical and mineralogical data suggest that the granitic rocks of this segment found. in east-central Minnesota are the western extension of the Giants Range Batholith (Morey, 1978).

Geophysical data also suggest that Lower Precambrian granitic, metavolcanic, and metasedimentary rocks underlie much of the northern

part of east-central Minnesota. However, outcrops of Lower Precambrian

granitic rocks are confined to the southwestern corner of Cass County approximately 10 km north-northeast of the village of Staples in

adjoining Todd County (Morey, 1978). Here, cataclastic zones in the

granite are concordant with fold axes found in the overlying Middle

Precambrian strata suggesting that the Lower Precambrian granite basement was deformed by the same deformational event (i.e. the Penokean

Orogeny) which deformed the overlying Middle Precambrian strata

(Morey, 1978).

Analysis of the central Lower Precambrian segment is difficult

due to lack of exposures and radiometric ages (Morey, 1978). However,

it is fairly certain that this segment forms a discrete zone separating

the northern and southern Lower Precambrian segments. This zone is

known as the Great Lakes Tectonic Zone (Sims and others, 1980) (see

Fig. 6). Rocks of this segment are characterized by sheared granitic

rocks similar to those of the northern granite-greenstone segment, and

by schistose rocks that could be meta-sedimentary or be reactivated and

sheared equivalents of less severely deformed gneissic rocks of the

southern gneissic segment (Sims, 1976).

The southern Lower Precambrian segment consists of very old,

highly deformed and metamorphosed gneisses (see Fig. 6) divided into 20 the following three lithostratigraphic units of uncertain relative ages (Morey, 1978): the Richmond Gneiss; the. Sauk Rapids Metamorphic

Complex; and the McGrath Gneiss.

Generally, the Richmond Gneiss is a vaguely foliated black to dark grayish-black, coarse-grained porphyroblastic gneiss consisting mostly of potassium feldspar, plagioclase, quartz, hypersthene, hornblende, and minor amounts of biotite, garnet, and cordierite (Morey,

1978).

The Sauk Rapids Metamorphic Complex is divided into the

following three lithostratigraphic units of formational status: the ·

Sartell Gneiss; the Watab Amphibolite; and the St. Wendell Metagabbro.

The Sartell Gneiss consists of an interlayered quartzofeldspathic

gneiss and a garnet-cordierite-bearing biotite gneiss (Morey, 1978).

The Watab Amphibolite is a vaguely foliated hornblende-pyroxene-plagio-

clase gneiss (Morey, 1978). And the St. Wendel Metagabbro is a

plagioclase-clinopyroxene gneiss characterized by grains as large as

8mm in diameter (Morey, 1978).

The McGrath Gneiss is the oldest rock in Carlton and Pine

Counties and is a biotite-bearing quartzofeldspathic gneiss having a

Rb-Sr whole-rock isochron age of at least 2700 m.y. (Stuckless and

Goldich, 1972; Keighin and others, 1972). Therefore, the McGrath is

of Early Precambrian age and assumed to have been emplaced during the

Algoman Orogeny 2700 to 2750 m.y. ago (Hanson and others, 1971; Catanzaro

and Hanson, 1971). Geophysical data indicate that the McGrath Gneiss

underlies an extensive area and forms several broad, dome-shaped

anticlinoria bounded and locally overlain by Middle Precambrian 21 stratified rocks. The McGrath is also characterized by having a strong biotitic foliation, by containing a compositional layering thought to be primary (Davidson, 1979), and by -being extensively sheared and recrystallized with cataclasis and recrystallization most intense near the inferred northern contacts with younger Middle Precambrian sedimentary rocks.

Deformation of the McGrath Gneiss is thought to have occurred contemporaneously with the deformation of the younger, overlying and adjacent Middle Precambrian stratified rocks during the Penokean

Orogeny (Davidson, 1979) . . Radiometric age determinations support this conclusion. Rb-Sr dating of biotite from the McGrath cataclastic zones give metamorphic ages of 1750 and 1740 m.y. which are comparable to the accepted age of the Penokean Orogeny (Keighin and others, 1972).

Also, the Rb-Sr isochron age of the Thomson Formation is 1730 m.y.

(Keighin and others, 1972). Similar structural patterns existing in both the McGrath Gneiss and Thomson Formation also indicate contemporaneous Penokean deformation. Structures found in the McGrath Gneiss similar in pattern to those found in the Thomson Formation include open folds of biotitic foliation having rounded hinges (Keighin and others, 1972), and lineations defined by elongate mineral grains interpreted by Davidson (1979) to be parallel to the axes of these open folds (Keighin and others, 1972).

Morey and Sims (1976) have interpreted the McGrath to be a mantled gneiss dome of Lower Precambrian gneissic terrane rocks (greater

than 2700 m.y. in age) that was tectonically reactivated in Middle

Precambrian time during the Penokean Orogeny. 22

Middle Precambrian Stratified Rocks

GENERAL STATEMENT

Toward the end of Early Precambrian (Archean) time, the northern greenstone-granite and southern gneissic terranes of Minnesota were welded together by large amounts of granite forming a large craton

(Sims and Morey, 1973). The boundary between these two Lower

Precambrian terranes is known as the Great Lakes Tectonic Zon·e (Sims and others, 1980) (see Fig. 6). This new stable craton developed large sedimentary basins, one of which is known as the Animikie Basin of east-central Minnesota (see Fig. 4). The Animikie Basin in

Minnesota developed over, and approximately parallel to, the boundary separating the two Lower Precambrian terranes. Sedimentation in the basin began approximately 2000 to 2100 m.y. ago with deposition of the

Middle Precambrian Mille Lacs Group followed by later Middle

I Precambrian deposition of the Animikie Group (Morey, 1978) (see Fig.

7). The period of Animikie Basin sedimentation can be div.ided into five depositional phases outlined here (Morey, 1979) (see Fig. 7): the first three phases (pre-quartzite, quartzite, and shelf phases) resulted in the deposition of a southward-facing miogeosynclinal sequence; the fourth phase (transitional phase) resulted in a transitional sequence as the shelf foundered; and the fifth phase

(flysch phase) deposited a southward-facing, eugeosynclinal elastic wedge deposited by southward-flowing turbidity currents.

At the close of the Middle Precambrian, deformation and metamorphism associated with the Penokean Orogeny had a pronounced LITHO- LITHO- STH.ATIGR.APHIC PRE-TECTONIC STRATIGRAPHIC SECTION STRATIGRAPHIC UNITS UNITS DEPOSITIONAL Carlton nnd Pino PHASES Mesabi range NW SE Counties % V1r91n10 formollcn FLYSCH PHASE T/1omson Formalio.n 9roywocke , s1ole, scot1cr1;C1 volcanic 81..vobik and hypoby"ol rocks, and .... lrnn ForrT"ol 1nn carbonaceous or9il lite. :.;: .E upper member ·c: Q.Jarllrlo ,--- ../ TRANSITIONAL PHASE and scotlered volcanic roch <.!) Formation 0 "'u _31 lqua•l1 ile, sollslonc, and arQ illile . PRE-QUARTZITE PHASE ;:E Glen Township con9lomerole, auortzile, orgoll ile, Gneiss lerrone "Formation pillowed bosoll, 099lomerote, iron - formation, limestone . Oenham Formation

Figure 7. Pretectonic north-south cross section showing lithostratigraphic nomenclature and depositional phases or the Middle Precambrian strati£ied N rocks in east-central Minnesota (Morey, 1979). w 24 effect on the Animikie Basin sediments and on the adjacent and underlying McGrath Gneiss. Radiometric dating methods have given approximate ages of Penokean events. For example, K-Ar dating of biotite from McGrath Gneiss shear zones yield metamorphic ages of 1500 to 1710 m.y. (Keighin and others, 1972), and whole rock K-Ar dating of slate and phyllite from the Thomson Formation and from metasediments of the

Cuyuna District yield ages of 1600 to 1800 m.y. (Peterman, 1966;

Hanson, 1968; Hanson and Malhotra, 1971).

PENOKEAN DEFORMATION AND METAMORPHISM

Morey (1978) assigned an age of 1850 to 2150 m.y. to the

Middle Precambrian stratified rocks. Morey and Lively (1980) described these rocks as being extensively deformed and metamorphosed during the Penokean Orogeny approximately 1870 m.y. ago resulting in a broad, eastward-plunging synclinorium bounded to the north, west, and southeast by Lower Precambrian rocks. Morey (1978) postulated that the degree of deformation may be related to different underlying Lower Precambrian basement rocks. For instance, Lower

Precambrian granitic rocks underlie the northern portion of the synclinorium. There, the basal unconformity appears only slightly deformed with overlying Middle Precambrian beds relatively undisturbed and dipping gently southward. However, central and southern portions of the synclinorium are underlain by Lower Precambrian schistose and/or gneissic rocks and is complexly folded into many large anticlines and synclines with several coaxial second- and third-order folds on their limbs (Schmidt, 1963). These fold styles change from 25

the north-central to the southern areas of the synclinorium (Morey.

1979). In the north-central area of the synclinorium the folds are open, have vertical axial planes that strike east-northeast, and have

fold axes that plunge gently to the northeast. In the southern areas

of the synclinorium deformation appears more intense with isoclinal

and overturned folds having axial planes dipping 60 degrees to the

southeast and fold axes plunging moderately to the southwest. !..; Regional metamorphic grade of the Middle Precambrian stratified

rocks in the synclinorium increases from north to south. Therefore,

metamorphic grade may be related to the different underlying Lower

Precambrian basement rocks (Morey, 1978) and increase in rocks showing

signs of more intense deformation. For instance, in the northern

portion of the synclinorium (underlain by Lower Precambrian greenstone-

granite basement rocks) the Animikie rocks have been metamorphosed

only to the zeolite facies at most (Morey, 1973; Perry and others,

1973). Metamorphic grade then increases southward in the synclinorium

underlain by Lower Precambrian schistose and/or gneissic rocks. At

its type locality near the town of Thomson (northeastern Carlton

County), the Thomson Formation is metamorphosed to the lower green-

schist facies and has a well-developed slaty cleavage in its

argillaceous units. Megascopic muscovite first appears in the Thomson

Formation near ·the town of Atkinson (about 17km southwest of Thomson)

where Thomson argillites and slates become phyllitic. Biotite appears

about 38km south of Thomson, Still farther south, metamorphic grade

increases to lower amphibolite facies near the town of Moose Lake

(about 46km south-southwest of Thomson) where garnet porphyroblasts 26 first appear. Also in. this general area, biotite and muscovite have been extensively recrystallized to define a foliation parallel to bedding. Farthest south near the town of Denham (about 64km southwest of

Thomson) the Thomson Formation is separated from both the older Middle

Precambrian Denham Formation and the Lower Precambrian McGrath Gneiss to

the south by a major northwest-trending fault (Morey, 1979). Here, the

Denham Formation is metamorphosed to lower amphibolite facies and con-

tains staurolite porphyroblasts in proximity to the McGrath Gneiss

(Morey, 1979).

As described above, regional metamorphism appears to be positive-

ly correlated with regional deformation. In general, the biotite,

garnet, and staurolite isograds conform to the regional fold geometry and

define a metamorphic high along the northern edge of the McGrath Gneiss.

In detail, however, the metamorphic isograds transect major fold axes

(Morey, 1978). Therefore, Morey (1978) concluded that deformation and

metamorphism were discrete events during the Penokean Orogeny.

MILLE LACS GROUP

The following is a sequential ·description of the Mille Lacs

Group formations as described by Morey (1978) (see Figs. 7 and 8).

Estimated maximum thickness of the group is approximately lkm; however,

thickness estimates are uncertain because of deformation and sparse

exposure.

Denham Formation: a sequence of quartz-rich sedimentary rocks

extending in a broad belt from near Denham (where it unconf ormably

overlies the McGrath Gneiss) westward to the general area of Mille Lacs -Generalized EON I ERA I System I Series Group Formation Section Description b u Holocene Early poslrlacial to recent 3lluv1um . also peal 1r.d bot depos.11 f5 o Ouolernory Ple'slocene Glacial . tlac1olluvJ1,:e. ar.d 11acio1,.uslr.ne scd .menls incl.din& u N 1 end moraine: eround mora:r.e . cul•!lh sJnd and er•vel. 1na u depos1ls . Oepos1lcJ lrom m•il .ple 1!aC1JI aj1ances - Hinckley - . . 0N Sond1Jon1 . _. .. :-» -. '. ::. :· .':-'. . Mcdr"m· lo comqu.ned sandstone 0 - · .. . : - .. ffi Ke;11eenowon Upper I- f and du Lac formation

teocns romllc St. Cloud Gronile Om lex Med1cm · rained . porphyr1J1c ran.le al !Od1c ccmposil1on Isle Granite Mcd o•m·1ra.n•d t.olole tra r111e ol 1od1c War man Gronile lnJerbeddcd te1J1pa1h1c 1ra1,,ade. and SIJle. H mat.c bll. mal1c h1pib111;J rocks anl 1ron·lormat1on u Thomson Formolion Carbonaceous slalc pol!.ni "pnaro rnlo 1la1e. S•llltcne. and ver1 lrne· 0- u- N 0 Rabbit Lake fm. trained 0 N Animikie 0:: 0 Group I Trommald Fm . eranular ar.d nontrandar iron lcrmitrcn w 0:: I- w Mchntmen . 0 l- for ma lion ·..:.: :: : . :...: OominJnlly 11llslone and artolh!e ,.. ,Jh rr.rnor quartzite a:: 0 -.. ·. a. 0:: : '.:- =..:..::.:.. = -- QuJrti.Je. 1.Uslone and a. "undiw iJed" Mille Lacs c. PrlloNe;I basal!. J,1J:ir lull. d.JbJSe. anil beds of iron·s.11.de· =_-; rich rnb:ir.Jcccus sl;le. ar.d carbonJle-lmes ircn ·formalicn, Group Fm. inc1,des bmll near Mora Denham formation s.IJslJne and quar11-11 ch 11Jte; nur bJsc ; beJs of p.1:ohed tusa!l. aulomerale. lirr.i:sloJe, ti and rrun·formal.on MCGralh Gneiss B1olrle·rich quaruo!el.:lspalhrc aneiss

Figure 8. Stratigraphic column of parts of Carlton, Pine, Aitkin, Kanabec, N and Mille Lacs Counties, east-central Minnesota (highly schematic wtlth -....r respect to thickness). (Morey and Lively, 1980) 28

Lake. It is dominated by quartzite, metasiltstone, siliceous slate, and conglomerate near the base of the format.ion which contains clasts from the underlying McGrath Gneiss. Other lithologies include beds of pillowed metabasalts, agglomerate, marble, and oxide-facies iron-

formation.

Glen Township Formation: has diverse lithlogies including interbedded pyrite- and pyrrhotite-rich carbonaceous slate, recrystallized cherty iron-formation, pillowed metabasalt, altered diabase, and lesser amounts of quartzite, argillite, quartz-rich siltstone, and marble.

Little Falls Formation: is dominated by quartz-rich greywacke

intercalated with beds of slate, phyllite, and fine-grained schist with minor fine-grained mafic volcanic rocks.

Trout Lake Formation: is dominated by metamorphosed dolomite or

dolomitic limestone with many small masses and thin layers of siliceous

material (Marsden, 1972). Stratigraphic position of the formation is

uncertain, but Marsden (1972) suggested that it underlies the Mahnomen

Formation of the ·Animikie Group and overlies slate and quartzite units

of the Mille Lacs Group.

ANIMIKIE GROUP

The following is a sequential description of the Animikie Group

formations as. described by Morey (1978) (see Figs. 7 and 8). Estimated

maximum thickness of the group is approximately lkm; however, thickness

estimates have a large uncertainty because of deformation and sparse

exposure. 29

As discussed above, formations of the Mesabi Range and Cuyuna

Range are considered to be equivalent. Also .discussed was the likeli- hood that the Rabbit Lake Formation passes transitionally upward into the Thomson Formation. Therefore, the correlations shown in Table 1 can be made. The Mesabi Range formations will not be described in this report.

Mahnomen Formation: named after Mahnomen group of mines near the cities of Crosby and Ironton, east-central Crow Wing County.

It is dominantly siltstone and argillite with minor quartzite. Thick- ness of the formation is estimated to be at least 610 meters.

Trommald Formation: the main iron-formation of the Cuyuna

District named from mines northeast of the village of Trommald in east- central Crow Wing County. It is an interlayered granular and non- granular iron-formation having a thickness that varies from 3 to 180 meters at different locations.

Rabbit Lake Formation: overlies the Trommald Formation near

Rabbit Lake in east-central Crow Wing County. Estimated thickness of the formation is at least 610 meters; however, an unknown thickness was removed by erosion during and after the Penokean Orogeny. The formation has been subdivided by Marsden (1972) into the following three members:

Upper Member: a sequence of argillite, slate, greywacke,

and lesser amounts of carbonaceous slate.

Emily Member: a sequence of iron-rich strata.

Lower Member: a sequence of carbonaceous slate, greywacke, 30

local mafic flows, and coeval hypabyssal,

volcaniclastic, and volcanogenic rocks.

Thomson Formation: outcrops .are found in parts of Carlton,

Pine, and southern St. Louis Counties. It consists of greywacke, siltstone, and slate with carbonaceous and pyritic slate and lesser beds of mafic tuff, mafic lava, and coeval hypabyssal sills and dikes.

There are also abundant carbonate concretions, especially in the argillaceous units (Weiblen, 1964).

In the area around the type locality Morey and Ojakangas (1970) found the Thomson to consist of approximately 34% metagreywacke, 35 to 43% metasiltstone, and 23 to 31% slate. Using paleocurrent indicators they also determined that Thomson sediments were deposited by southward-flowing turbidity currents moving down a regional paleo- slope. In the same general area, Wright and others (1970) determined the regional structure to be folds that range from broad, open, and symmetrical folds in greywackes to tight, asyrmnetric, and overturned folds in slates (see Fig. 9). Also found were minor folds having nearly vertical axial planes and fold axes that plunge moderately to the east and west. Slaty cleavage that is axial-planar to these folds is well developed in the slate and less well developed in the greywacke

As previously discussed, the change in fold style from the

Thomson type locality southward to Denham indicates an increase in deformational intensity. Accompanying the change in fold style is an increase in regional metamorphic grade. Moving southward towards

Denham, folds become more asymmetrical to overturned, and more tight to North

A /:{

/ ? /" / / ,,..,,,, I .... --"' -----,, "/ // I , " --- " ' \ "" .... - . '... __ "'\,.,. --,_" ,..,I r ,./" / ...... \-'\"- . ' ' ,,."' -"I / " ? / ,,. \ \"-''' \,---, ' " ,,...... • - __,:.._.______J .' / / / / / "" _/v\'4-_ _, ___ , ' ' \ ' / / ;' /'\ ..._ _ _,.. / \ _, - /--/ '°' ., / ' / . ;\ / ... _ ....O ::>- '-·\\ '; .... __ ,,,_,,. . / /' <: .i'V ,,/ ',,, / ,,,"' ' -2000 Feet ' v,;..." ', /_,/- o.w..:.,.._ /,' ''- ---/ 0Horizon · to I ond ' ', ',...... ---- ____ / r-(,).>-..... -" ,,."/ vertlcoi sco es 1 ' ',,, ____ ...... _ "" SEC. SEC.8 5 EC.6 EC.31

Figure 9. Cross section of Thomson Formation showing the nature of the folding near the area of the type locality {i.e. Thomson, Minnesota). {Wright and others, 1970)

w I-' 32 isoclinal with axial planes dipping 60 degrees to the southeast. Due to the complex structure and the fact that ne·ither the top nor the bottom of the Tilomson Formation is exposed (Morey, 1978), estimates of the formation thickness are uncertain and range from at least 915 meters (Wright and others, 1970) to as much as 6096 meters (Schwartz,

1942b).

Middle Precambrian Plutonic Rocks

Middle Precambrian plutonic rocks are found in areas of east- central Minnesota underlain by rocks of the Lower Precambrian gneissic

terrane. These plutonic rocks are discrete intrusive bodies emplaced

into the overlying Middle Precambrian stratified rocks during and

after the Penokean Orogeny (Mo.rey, 1978). Small gabbroic to dioritic dikes, sills, or stocks are inferred to be the oldest Middle Precambrian

plutonic rocks in east-central Minnesota (Morey, 1978). During the waning stages of the Penokean Orogeny the small granodioritic stocks

of the Freedhem and Bradbury Granodiorites were emplaced (Morey, 1978).

After Penokean deformation the large sodic plutons of the Reformatory,

Isle, Warman, and Pierz Granites were emplaced (Morey, 1978). Middle

Precambrian plutonic activity then concluded with emplacement of the

Stearns Granitic Complex comprised of the St. Cloud and Rockville

Granites (Morey, 1978).

Upper Precambrian Sedimentary and Volcanic Rocks

Upper Precambrian stratified rocks are grouped into the

Keweenawan System of the Lake Superior region. Tile following is a 33 sequential description of the Keweenawan System rocks in east-central

Minnesota as described by Morey (1979).

Nopeming Sandstone: a thin veneer of quartz-rich sandstone, unconformably overlying the folded Thomson Formation.

Ely's Peak Basalts: a reversely polarized member of the North

Shore Volcanic Group which overlies the Nopeming Sandstone.

Chengwatana Volcanic Group: normally polarized lava flpws of

Middle Keweenawan age in contact with the Upper Keweenawan rocks along the Douglas Fault in Carlton and Pine Counties.

Fond du Lac Formation: Upper Keweenawan, fine- to coarse- grained arkosic sandstone, siltstone, and shale of alluvial origin unconformably overlying the Ely's Peak Basalts.

Hinckley Formation: an Upper Keweenawan, medium- to coarse- grained quartzose sandstone overlying the Fond du Lac Formation; the contact is gradational.

Keweenawan dikes: numerous basaltic Keweenawan dikes occur throughout east-central Minnesota cutting Middle Precambrian stratified rocks. A basaltic dike cutting the Thomson Formation near the village of Thomson yielded a K-Ar age of 1050 m.y. (Keighin and others,

1972), consistent with known Keweenawan igneous events. CHAPTER THREE

PETRO GRAP HY

I. Introduction

Rock Types .. The Thomson Formation in the study area consists of metasiltstone, phyllite, calcareous metasiltstone and phyllite, graphitic phyllite, volcaniclastic metasiltstone and phyllite, and metabasalt.

The volcaniclastic rocks consist of pyroclastic debris and/or volcanic detritus of epiclastic origin. All of the rocks have a continuous

schistosity which is typically parallel to bedding. Many quartz veins,

lenses, and pods occur in these rock types. Concretions are comm6n

in calcareous and non-calcareous metasiltstone and phyllite beds. One

post-tectonic basaltic dike is exposed at the town of Moose Lake.

Modal mineralogy of these rocks is given in Table 3, and detailed

descriptions of each are given under separate headings below.

Distribution of Rock Types

(Refer to Figs. 3 (a) and (b) and Plate 1.) Metasiltstone and

phyllite are found throughout the thesis area. Calcareous units are

located in the northern and central Kettle River and western Split

Rock River areas (e.g. Outcrops 16-29, 28-77, 29-45, and 29-22).

Graphitic phyllite is found in one small outcrop (Outcrop 28-81) on

the east bank of Kettle River where Glaisby Brook flows into the

34 35 TABLE 3

Modal }ti.neralOgI Ranges of Rock Units {Averages in Parentheses) Calcareous Calcareous Graphitic Metasiltstone Phyllite Metasiltstone Phyllite Phyllite Quartz 32-54 (41) 22-49 (29) 23-25 (24) 9-20 (13) 28 Plagioclase 27-44 (35) 10-32 {24) 14-20 (18) 8-18 (10) 21 Potassium 0-4 (TR) 0-2 (TR) TB TR Feldspar ChloriU 2-13 (5) 5-12 (7) 0-29 l 16) 5-12 (7) 5 !o:uscovite 0-8 (3} 10-35 (23) 0-5 (2) 10-35 (23) 20 Biodte 2-25 (12) 3-25 (12) 0-10 (4) 3-25 (12) 10 Calcite TR TR 25-55 (36) 18-35 (27) --- Siderite/ o·-25 (TR) Ankerite ·-- Graphite 1s Ferroactinolite --- Garnet 0-3 (TR) · 0-3 (TR) 0-3 (TR) 0-3 lTR) Epidote TR TR TR TR TR

Olivine

Ilmenite TR-1 1-6 (2) TR-·1 1-6 (2) TR Magnetite TR-1 1-6 (2} TR-1 1-6 (2) TR Leucoxene TR-1 1-6 (1) TR-1 1-6 (1) fyrite TR TR TR TR Apatite TR TR TR TR TR Tourmaline TR TR TR TR Zircon TR TR TR TR

Ru tile TR TR 'l'R TR , Sphene TR TR TR TR --- Sericite TR TR TR TR TR Clay Alteration TR TR TR TR TR (of feldspar)

Nu:nber of Thin 7 6 3 2 1 Sections Studied (Continued on next page.) 36 TABLE 3 (Continued)

PoS"t- ·volcaniclastie Meta.siltstone/ Metamorphosed Basalt Metabasalt Phyllite Concretions Dike

Quartz 0-3 lTR) 2-7 (4) . 29-43 (35) Plagiocls.se 20-30 . ( 25) 25-39 (3J) . 30-42 (37} 46 Potassium · Feldspar Chlorite 3-10 (6} 7-10 (9) TR-1 11'.iscovi te TR-1 Biotite 0-5 (TR) 1-7 (5) TR-1 Calcite 2-5 (4) 7-12 (10) TR-1 Siderite/ Ankerita Graphite Ferroactinolite .' (47) 28-44 (32) 22-30 (26) Garnet TR-1 . Epidote 2-12 (9) 1-7 (5) --- Augite 12 Olivine --- 4 Ilmenite 3-7 l4) 2-7 <4> --- s Magnetite 3-7 (4) 2-7 (4) 1-2 (1) 7 Leucoxene Pyrite --- Apatite TR TR TR TR Tounnalina --- Zircon

Ru tile TR TR

Sphene TR TR Sericite TR TR Clay Alteration TR TR TR (of feldspar)

Partially Altered 26 Interstitial Glass

Number or Thin 5 4 3 1 Sections Studied 37

Kettle River. Metahasalt and volcaniclastic metasiltstone and phyllite are located in the Glaisby Brook (e.g. Outcrops 22-54 and

22-57) and central Kettle River (e.g. Outcrops 29-14 and 29-49) areas.

Stratigraphic relationships between the metabasalt and volcaniclastic metasiltstone and phyllite are uncertain because of poor and discontinuous exposures. As a result, it is not clear as to whether the metabasalts are flows and/or hypabyssal dikes and sills.

Laboratory Methods

From nearly 150 samples collected, a representative suite of

83 thin sections was made and studied. Modal mineralogy was estimated visually unless stated otherwise.

In thin section, ferroactinolite was distinguished from actinolite and hornblende by its moderate to strong pleochroism

(light green to deep green), by its extinction angles of 10 to 15 degrees, and by its association with prograde metamorphic chlorite

(Deer and others, 1979). In comparison, actinolite has weak

pleochroism (colorless to pale green), and hornblende has extinction

angles of 14 to 25 degrees and is not usually associated with chlorite

in a prograde metamorphic assemblege.

Cobaltinitrite and "Laniz" staining methods were used on thin

section heels for detection of potassium feldspar and plagioclase,

respectively. Tilese stained heels were studied under a reflecting

microscope and feldspar percentages were estimated visually. 38

II. Lithologic Descriptions

Metasiltstone and Phyllite

Interbedded rnetasiltstone and phyllite are common throughout the study area. Physical differences between the two rock types make bedding distinctive at most outcrops. Metasiltstone is typically grey to silver-grey, fine- to extremely fine-grained, and massive to schistose. Metasiltstone beds range in thickness from less than !cm to 80cm, weather to positive relief compared to the phyllite, and are competent and usually well jointed (see Fig. 10). Phyllite is typically silvery-grey to greenish-grey, very fine-grained, and occurs in beds that range in thickness from less than lcrn to 5Qcm. Phyllite beds weather to negative relief compared to metasiltstone beds (see

Fig. 10). Many individual phyllite beds are compositionally laminated with alternating silty and phyllitic laminae. A well developed

schistosity exists in both rock types and is defined by a pref erred orientation of phyllosilicates in a sub-granoblastic quartzof eldspathic matrix (see Figs. 11 and 12). veins, lenses, and pods of various

sizes are commonly found within beds of metasiltstone and phyllite along

surfaces parallel to the schistosity (see Fig. 13).

The metasiltstone and phyllite rock types are essentially the

same in thin section except for a higher phyllosilicate/tectosilicate

ratio in the phyllite. Mineralogical association, habit, texture, and

grain size are essentially the same in both rock types and are described

below.

The major constituents of the sub-granoblastic matrix are quartz

and plagioclase. Both exhibit sharp extinction and occur as 39

Figure 10. Interbedd9d metasiltstone and phyllite. Schistosity is parallel to bedding at this outcrop. Note necking of beds along line parallel to Style 3 minor fold (antiform) a.xis in lower right corner of photo (see Style 3 minor fold section in structure chapter). Ruler length approximately 17cm. (Outcrop 20-4, NEi, Sec. 20, T.46N., R.19W.)

Figure 11. Photomicrograph of metasiltstone showing schistosity defined by a preferred orientation of phyllosilicates in a sub-granoblastic quartzofeldspathic matrix. Field length approximately 2.9mm. Crossed nicols. 40

Figure 12. Photomicrograph of phyllite showing schistosity defined by a preferred orientation of phyllosilicates in a sub-granoblastic quartzofeldspathic matrix. Field length approximately 2.9mm. Crossed nicols. 41

Figure 13. Photograph (with line diagram) of ·a quartz (stippled) oriented in a plane parallel to bedding (B) and the bedding-parallel schistosity (S). Note deflection of metasiltstone and phyllite beds around pod. Ruler length approximately 17cm. (Outcrop 20-4, Sec. 20, T.46N., R.19w.) 42 subpolygonal, anhedral grains less than O.lnun in diameter. The plagioclase grains are typically not twinned and are generally fresh.

The phyllosilicates define a well developed schistosity by having a preferred orientation. Muscovite, biotite, and chlorite

grains are subhedral to euhedraJ. flakes less than 0. 23mm in length.

Biotite grains are generally fresh or somewhat chloritized. In many

samples, chlorite occurs not as an alteration product, but as _•;inde- pendent euhedral crystals. In some samples, a few larger chlorite

flakes are discordant to the schistosity indicating that they formed

after the development of that foliation (see Fig. 14).

Minor constituents include opaque minerals, garnet porphyro-

blasts, and potassium feldspar. The opaque minerals are very fine-

grained and include ilmenite, magnetite, leucoxene, and lesser amounts

of pyrite. These opaque minerals are anhedral to euhedral. Elongate

grains of ilmenite are aligned in planes parallel to the schistosity.

The garnet porphyroblasts are anhedral to subhedral, less than l.3mm

in diameter, and commonly poikilitic. Garnets may be fresh, slightly

altered, or completely chloritized and/or sericitized. Pre- (or early

syn-), syn-, and post-tectonic garnets are distinguished by patterns

of quartz inclusions and by geometrical relationships between garnet

poikiloblasts and adjacent matrix (Spry, 1979) (see Figs. 15, 16, and

17). Significance of these garnet poikiloblasts will be discussed in

the interpretation chapter of this report. Potassium feldspar grains

are subpolygonal to subhedral, less than O.lmm in size, exhibit sharp

extinction, and are generally fresh. Some potassium feldspar is found 43

Figt1re 14. Photomicrograph of a large post-schistosity chlorite flake in metasiltstone discordant to the schistosity. Field length approximately 0.68mm. Plain lighto

Figure 15. Photomicrograph of pre- {or early syn-) tectonic garnets in phyllite. Note wrapping of phyllosilicates around garnets and development of quartz pressure shadows. Field length approximately 6.8mm. Plain light. Figure 16. Photomicrograph of syn-tectonic (snowball) garnet in phyllite. Note rotational pattern of quartz inclusions, wrapping and truncation of phyllosilicates around garnet, and development of quartz pressure shadows. Field length approximately 2.9mm. Plain light.

Figure 17. Photomicrograph of a post-tectonic garnet in metasiltstone. Note truncation of phyllosilicates around garnet and absence of pressure shadows. Field length approximately 2.9mm. Plain light. 45 as matrix grains, but it more commonly occurs as hydrothermal micro- stringers discordant to the schistosity.

Accessory minerals are found as either very small relict detrital grains or as metamorphic grains which crystallized in place.

Distinction between the two grain types is made by appearance (e.g. rounded = detrital vs. angular or euhedral = metamorphic) and by geometric relationship to the surrounding matrix (e.g. micas wrapping grain= detrital vs. micas truncated by grain= metamorphic). Relict detrital accessory minerals include apatite, tourmaline, and rutile. Metamorphic grains include apatite, tourmaline, rutile, sphene, and epidote.

Calcareous Metasiltstone and Phyllite

. The mineral compositions of calcareous metasiltstones and phyllites are essentially the same as those of non-calcareous metasiltstones and phyllites (described above) except for a high percentage of calcite. In calcareous units, calcite constitutes 25 to 55 percent of the rock and occurs as anhedral grains less than

0.2mm in size. Siderite and/or ankerite was found at one outcrop

(Outcrop 28-77) constituting approximately 25 percent of the rock. The ferroan carbonate occurs as anhedral blebs less than l.2nnn in diameter.

These blebs weather to negative relief (i.e. small pits) and produce a red oxidation stain.

Graphitic Phyllite

The graphitic phyllite is black to dark grey, very fine-grained, and contains a high percentage of graphite making the rock very 46 friable and soft. A well developed schistosity is present. Mineral habit, texture, and grain size are similar to . those of the non- calcareous phyllite described above.

Metabasalt

The metabasalt is typically grey to dark grey, medium- to fine- grained, and weakly to strongly schistose. Large grains (up to 2mm in length) of subhedral ferroactinolite and plagioclase are commonly aligned in planes parallel to the schistosity (see Fig. 18), Thin section analysis indicates that many large plagioclase grains are relict phenocrysts rotated into the schistosity during development of that foliation. Metabasalts are identified in outcrop because they contain large ferroactinolite and plagioclase grains, by their structureless appearance (except for the schistosity), and by their high degree of competency. No pillowed metabasalts were found.

Major constituents are ferroactinolite and plagioclase.

Ferroactinolite grains are anhedral to subhedral and generally less than 0.5mm in length (however, a few grains may be as large as 2mm).

Plagioclase grains are anhedral to subhedral, less than 0.3mm in size, and generally altered. Relict plagioclase phenocrysts as long as 0.8mm are extensively fractured and have pressure shadows around them which formed as the phenocrysts rotated during development of the schistosity (see Fig. 19). Albite twinning is usually well developed in many plagioclase grains. Michel Levy analyses (Heinrich, 1965) of albite twins gave approximate anorthite contents of 13 percent for small matrix grains and 51 percent for large matrix grains and relict 47

Figure 18. Photomicrograph of metabasalt. Major constituents are ferroactinolite and plagioclase. Note crude development of a schistosity (nearly horizontal in photo). Field diameter approximately 1.5cm. Plain light.

Figure 19. Photomicrograph or relict plagioclase phenocryst in meta.basalt. Note extensive fracturing or the phenocryst and development of pressure shadows around it. Field length approximately 6.8mm. Crossed nicols. 48 phenocrysts. The lower anorthite content of the groundmass crystals is attributed to uppermost greenschist facies .metamorphic recrystaliza- tion, while the higher anorthite content of the relict phenocrysts reflects the composition of the protolith.

Less abundant minerals include epidote and chlorite. Epidote grains are generally anhedral and less than 0.2mm in length. Epidote may also occur as round, polycrystalline porphyroblasts as large as l.2mm in diameter (see Fig. 20). These epidote porphyroblasts lack

(or have ill-defined) pressure shadows indicating they crystallized after (or during the waning stages of) development of the

Chlorite grains are subhedral to euhedral and less than 0.2mm in length. Epidote and chlorite are commonly found with plagioclase and calcite in small. (approximately O.S:mm) relict amygdules (see Fig. 21).

These relict amygdules have been strained to lensoidal shapes during development of the schistosity.

Minor minerals include ilmenite, magnetite, calcite, biotite, quartz, and accessory minerals. Ilmenite and magnetite grains are less than 0.15mm in size; elongate ilmenite grains are aligned parallel to the schistosity. Calcite grains are anhedral, less than 0.3mm in size, and may be found in calcereous veinlets concordant to the foliation. Biotite grains are anhedral to subhedral and less than

0.18mm in length. Quartz grains are anhedral to subpolygonal and less than O.lmm in diameter.

Volcaniclastic Metasiltstone and Phyllite

Volcaniclastic metasiltstone and phyllite are typically grey to greenish-grey and medium- to very fine-grained. A well developed 49

Figure 20. Photomicrograph or a round, polycrystalline epidote porphyroblast in metabasalt. Note absence or pressure shadows around grain indicating a post-schistosity origin. Field length approximately 6.8:mm. Crossed nicols.

Figure 21. Photomicrograph or a relict amygdule in metabasalt. }lineral composition is plagioclase, epidote, calcite, and chlorite. Note amygdule has been strained during development or the schistosity resulting in a lensoidal rorm oriented parallel to the schistosity (nearly horizontal in photo). Field length approximately 2.9:mm. Crossed nicols. 50 schistosity is characteristic of these rocks. Beds vary in thickness from 6cm to 50cm. In the field, it is difficult to distinguish these beds from those of metasiltstone and phyllite (described above). In thin section, however, the distinction is clear. Thin section analysis of the volcaniclastic rocks reveals a major mafic component and presence of relict detrital basaltic rock and plagioclase fragments

(see Fig. 22).

The major minerals are ferroactinolite and plagioclase.

Ferroactinolite grains are anhedral to subhedral and generally less than 0.3rran in length. Plagioclase grains are anhedral to subhedral, less than O.lmm in length, and fresh. Relict detrital plagioclase grains as large as l.8mm have pressure shadows around them which formed as the grains rotated during development of the schistosity (see Fig.

23). Albite twinning is seen only in large plagioclase grains and is poorly developed.

Less abundant minerals include calcite, chlorite, and epidote.

Calcite grains are anhedral, less than 0.3Illln in diameter, and are often found in calcareous veinlets. Chlorite grains are anhedral to

subhedral and less than 0.3nnn in length. Epidote grains are anhedral and less than 0.3mm in diameter. Epidote, chlorite, calcite, and

plagioclase are the main constituents of basaltic rock fragments (see

Fig. 24). These fragments are as large as 3mm and have pressure

shadows around them which formed as the fragments rotated during

development of the schistosity.

Minor minerals include biotite, quartz, magnetite, ilmenite, and

accessory minerals. Biotite grains are anhedral to subhedral and less 51

Figure 22. Photomicrograph of volcaniclastic phyllite. Major constituents are ferroactinolite, plagioclase, and relict basaltic rock and plagioclase fragments. Note the well developed schistosity. Field diameter approximately 1.5cm. Plain light.

Figure 23. Photomicrograph of a relict detrital plagioclase grain in volcaniclastic phyllite. Note pressure shadows around grain formed as it rotated during development of the schistosity (horizontal in photo). Field length approximately 6.8mm. Crossed nicols. 52

Figure 24. Photomicrograph of a basaltic rock fragment in volcaniclastic phyllite. Main constituents are epidote, chlorite, calcite, and plagioclase. Note crude pressure shadows around fragment (schistosity horizontal in photo). Field length approximately 2.9mm. Plain light. 53

than 0.3mm in length. Quartz grains are anhedral to subpolygonal,

less than 0.07nun in diameter, and exhibit sharp extinction. Magnetite

grains are less than 0.18mm in diameter and are anhedral. Ilmenite

grains are elongate, less than 0.18rnm in length, and are aligned

parallel to the schistosity. Accessory minerals include very small

grains of apatite, sphene, and rutile.

Concretions

Carbonate concretions are found throughout the Thomson Formation

(Schwartz, 1942c; Weiblen, 1964). The concretions do not deform the

bedding around them; therefore, Schwartz (1942c) believed many of the

concretions formed just below bedding surfaces by the diagenetic

processes of replacement and enclosure of host sediments. Schwartz

(1942c) also observed that carbonate concretions are abundant in

beds otherwise having no calcite. This suggests that the carbonate

came from pore waters rather than from the host sediments.

In the study area, carbonate concretions are found as calcite-

rich lenses oriented parallel to· the schistosity in beds of

metasiltstone and phyllite. These lenses vary in size and axial ratio.

A few concretions appear to have undergone metasomatic replacement by

quartz resulting in lenses having calcite cores and quartz rims.

However, the vast majority of carbonate concretions in the study area

have been metamorphosed resulting in development of various metamorphic

silicates.

Metamorphosed concretions are found as oblate pods of various I sizes oriented parallel to the schistosity in metasiltstones and \ \ 54 phyllites (see Fig. 25). Many have well delineated cores of calcite or very fine-grained plagioclase and quartz •. In these concretions, plagioclase and quartz grains are subpolygonal to polygonal, less than O.lmm in diameter, and exhibit sharp extinction. Plagioclase grains are not twinned. Ferroactinolite grains are prismatic, poikiloblastic, less than 5mm in length, and commonly aligned parallel to the schistosity.

Quartz Veins, Lenses, and Pods

Milky-white quartz is abundant in all rock types (except the post-tectonic basaltic dike) and occurs as veins, lenses, and pods of various sizes oriented in planes parallel to the schistosity. High concentrations of thin (approximately 2mm to 4cm thick) quartz lenses and/or veins are common in phyllites. Quartz also occurs as large pods

(see Fig. 13), the largest found being approximately 1.2 x 2.5 meters in size. As shown in the photograph referred to above, most quartz masses deflect and deform adjacent bedding. This observation suggests that most quartz did not result from metasomatic replacement of diagenetic carbonate concretions (described above), but is probably hydrothermal in origin.

Post-Tectonic Basaltic Dike

A basaltic dike approximately 2.5 meters thick is exposed along the SOO Line railroad in the northwestern corner of the town of

Moose Lake (Outcrop 20-3). The dike strikes N72°E and dips 84 degrees to the northwest. It is dark grey to black, fine-grained, massive, and has a Scm thick chilled margin. The rock was hypocrystalline and has 55

Figure 25. Photomicrograph of metamorphosed concretion in metasiltstone. Major constituents are ferroactinolite, quartz, and plagioclase. Note crude alignment of some ferroactinolite in planes parallel to the schistosity (nearly horizontal in photo). Field diameter approximately 1.5cm. Plain light. 56 an intersertal texture. Its composition, estimated from a 1000 point count, is: 46% plagioclase; 26% partially altered interstitial glass;

12% augite; 7% magnetite; 5% ilmenite; 4% olivine; and trace amounts of apatite. Anorthite content of the plagioclase is estimated to be

63 percent using the method (Heinrich, 1965). The dike shows no mesoscopic or microscopic evidence of metamorphism and/or deformation. Tiierefore, it was emplaced after the metamorphism and .c deformation associated with the major tectonism (i.e. the Penokean

Orogeny). This is the only post-tectonic intrusion exposed in the study area, and is probably of Keweenawan age. CHAPTER FOUR

STRUCTURE

I. Introduction

Structural Patterns

Orientations of bedding, foliations, lineations, minor fold

axes and axial planes, joints, faults, and quartz boudins yield well defined regional patterns. This chapter will only describe these

structures and their regional patterns; structural interpretation will

be given in the following chapter of this report. (Refer to Figs.

3 (a) and (b) and Plate

II. Description of Structures

Bedding

As described in the petrography chapter, physical differences

between various interbedded rock types (e.g. metasiltstone and

phyllite) make bedding distinctive at most outcrops (see Fig. 10).

Rare and questionable graded bedding is the only other primary

structure remaining after extensive metamorphic recrystallization.

Evidence of graded bedding is seen as a gradational increase in

phyliosilicates upward in beds, which may have resulted from re-

crystallization of beds which grade from arenitic material near their

bases to pelitic material near their tops.

57 58

Bedding orientation in the map area is nearly homoclinal (see

Fig. 26). Beds generally strike approximately east-west and dip 10 to 35 degrees to the south. A few beds dip at shallow angles to the north, and a small number strike approximately N06°W and dip steeply to the west.

Foliations Two types of foliations exist in the map area. One is a continuous schistosity common in all Thomson Formation rock types. The other is a discontinuous crenulation cleavage found in phyllite and in

thinly interlaminated phyllite and metasiltstone.

Schistosity. As described in the petrography chapter, all

Thomson Formation rock types have a continuous schistosity throughout

in the map area. In metasiltstone and phyllite, the schistosity is defined by a preferred orientation of phyllosilicates in a sub-

granoblastic quartzofeldspathic matrix (see Figs. 11 and 12). In metabasalt and volcaniclastic metasiltstone and phyllite, the

schistosity is defined by a preferred orientation of ferroactinolite,

strained plagioclase, and phyllosilicates (see Figs. 18 and 22). The

schistosity is also found as a fracture cleavage in beds of quartz-rich

calcareous metasiltstone which are low in phyllosilicate contents

(e.g. Outcrop 16-29, see Figs. 27 and 28).

The schistosity is usually found parallel to bedding (see Fig.

10 and compare Figs. 29 and 26). Therefore, the general orientation

of the schistosity is the same as that of bedding (i.e. strike of

approximately east-west and dip of 10 to 35 degrees to the south).

Minor irregularities in this strike and dip pattern are correlative to .-

Figure 26. Equal area projection of 250 poles to bedding. Readings from entire area. Contours 1, 5, 15, 30% of data per 1% area, maximum 33%. 60

Figure 27 (continued on next page). Schistosity (S) discordant to bedding (B). Angles of S/B intersections range from 50 to 80 degrees. Note the foliation is a schistosity in upper part of rock (relatively high in phyllosilicate content), whereas the foliation is a fracture cleavage in lower part of rock (relatively low in phyllosilicate content). Also note minor displacement of bedding along fracture cleavage. Ruler length 17cm; south is .to right. (Outcrop 16-29, NW 4 , Sec. 16, T.46N., R.20W.) 61

Figure 27 (continued from previous page). Line diagram of photo showing schistosity (S) discordant to bedding {B). Rectangle in left-center of diagram is area of Figure 28. 62

Figure 28. Close-up of Figure 27 showing schistosity (S; existing here as fracture cleavage) discordant to bedding (B). Note minor displacement of bedding along fracture cleavage. Ruler length approximately 17cm; south is t .o right. (Outcrop 16-29, Sec. 16, T.46N., R.2ow.) 63

Figure 29. Equal area projection of 495 poles to schistosity. Readings from entire area. Contours 1, 5, 15, 30% of data per 1% area, maximum 32%. 64 those of bedding described above. At a few locations (e.g. Outcrop

16-29), the schistosity is discordant to bedding and intersects bedding at high angles (also see section on Style 1 minor folds below). Angles of bedding/schistosity intersections at these locations range from 50

to 80 degrees (see Figs. 27, 28, 30, and 31).

Crenulation Clevage. Crenulation cleavage is a discontinuous foliation found in most parts of the map area. It is the result of ·" microscopic folding (i.e. crenulating) of the earlier schistosity

(described above) (see Fig. 32). It forms in rocks having high phyllosilicate contents and a well developed schistosity (e.g.

phyllites). It is also found in thinly interlaminated phyllite and metasiltstone (see Fig. 33). However, crenulation cleavage is lacking or poorly developed in more competent lithologies containing relatively

few phyllosilicates (e.g. metasiltstone and metabasalt).

There are two sets of crenulation cleavages in the map area having two distinct orientations (see Fig. 34). Crenulation

cleavages of the major set strike•: approximately N80°E and dip 60-90 degrees to either the north or south. Crenulation cleavages of the minor set strike approximately N02°W and dip 60 to 86 degrees to either

the east or west. Conjugate crenulations are poorly developed in the

study area and can be seen only in thin section under crossed nicols

(see Fig. 35).

Lineations

The following two types of lineations were found in the map area

(in addition to minor fold axes discussed later in this chapter):

crenulation lineations and mineral lineations. 65

Figure 30. Schistosity (S) discordant to bedding (B). Angles of S/B intersections are around 90 degrees. Rectangle in left-center of line diagram is area of Figure 31. Ruler length approximately 17cm; south is to right. (Outcrop 16-29, N\it, Sec. 16, T.46N., R.20W.) 66

. . .. . s ......

Figure 31. Close-up of Figure 30 showing schistosity (S) discordant to bedding (B). South is to right. (Outcrop 16-29, NWi, Sec. 16, T.46N., R.2ow.) 67

Figure 32. Photomicrograph of crenulation cleavage. The continuous schistosity (oriented from upper left to lower right of photo) is folded (i.e. crenulated) on a microscopic scale to result in a discontinuous crenulation cleavage {oriented from upper right to lower left of photo). Field length approximately 2.5mm.. Crossed nicols.

Figure 33. Crenulation cleavage {oriented from upper left to lower right of photo) in thinly interlam.inated beds of phyllite and metasiltstone. South is to right. (Outcrop 20-33, Sec. 20, T.46N., R.19W.) 68

Figure 34. Equal area projection of 111 poles to crenulation cleavages. Readings from entire area. Contours 1, 3, 5, of data per 1% area, maximum 14S(.

Figure 35. Photomicrograph conjugate crenulations in phyllite. One is horizontal in upper part of ·photo, the other is oriented from lower left to upper right in lower part of photo. They intersect at approximately 30 degrees. These are rare in the study area and can only be seen in thin section under crossed nicols. Field diameter approximately 1.5cm. Crossed nicols. 69

Crenulation Lineations. Crenulation lineations are co11I1Don in the map area. As discussed above, crenulation cleavages do not develop well as penetrative structures in more competent lithologies. What develops instead are non-penetrative wrinkles (i.e. crenulation lineations) on thin schistose surfaces commonly found separating beds of metasiltstone (see Fig. 36).

There are two sets of crenulation lineations in the map area .c having distinct orientations (see Figs. 36 and 37). Crenulation lineations of the major set have bearings of approximately N84°E and

S84°W and plunge 0 to 18 degrees along either trend. Crenulation lineations of the minor set have an average bearing of approximately

S09°E and plunge 2 to 30 degrees. At some outcrops Outcrop

20-33}, one set of crenulation lineations physically crosses the other set at angles approaching 90 degrees (see Fig. 36). Existence of crossing crenulation lineations demonstrates that there are two distinct and coexistent sets of crenulation lineations intersecting at high angles.

Mineral Lineations. Mineral lineations are rare in the map area. However, the following two types were found: amphibole lineations and muscovite streak lineations. Ferroactinolite lineations are associated with metamorphosed calcareous concretions, metabasalts, and volcaniclastic metasiltstones and phyllites (described in the petrography chapter). These lineations have bearings of approximately

Sl0°E and plunge 0 to 25 degrees (see Fig. 38). Muscovite streaks were found on very thin schistose surfaces separating beds of metasiltstone 70

Figure 36. Map view (south is to left) of two intersecting sets of crenulation lineations. The set oriented from left to right has a bearing of S10°W and plunges 5 degrees. The set oriented from upper left to lower right of photo has a bealing of N70°E plunges 9 degrees. (Outcrop 20-33, NE4, Sec. 20, T.46N., R.19W.)

Figure 37. Equal area projection of 411 crenulation lineations. · Readings from entire area. Contours 1, 3, 5, 10% of data per 1% area, maximum 71

Figure 38. Equal area projection of 30 ferroactinolite lineations. Readings from entire area. Contours 31 10, 17% of data per 1% area, maximum 20%G 72

(see Fig. 39). These lineations have bearings of approximately

Sll 0 W and plunge 4 to 14 degrees (see Fig. 40)!

Minor Folds

Three styles of outcrop-scale minor folds have been recognized in the map area. Each fold style has a distinct fold geometry and field relationships. Therefore, minor folds will be referred to in this chapter as Style 1, Style 2, or Style 3 minor folds. .•

Style 1 Minor Folds. Style 1 minor folds are isoclinal and overturned or recumbent. Bedding is folded into angular folds which may have at least second order parasitic folds (see Fig. 41). Style

1 minor folds have a schistosity parallel to their axial planes (see

Figs. 42 and 43). The folded bedding is cut at high angles (close to 90 degrees) by the schistosity in hinge areas (rare in study area), and is parallel (or nearly parallel) to the schistosity in limb areas of these minor folds.

Style 1 minor fold axes have bearings of approximately N78°E and S78°W and plunge 0 to 25 degrees along either trend (see Fig. 44).

Their axial planes strike approximately N74°E and dip 0 to 40 degrees to the south (see Fig. 45).

Style 2 Minor Folds. Style 2 minor folds are tight and asymmetrical or overturned. The bedding and bedding-parallel schistosity are together folded into rounded or angular folds which may have at least second order coaxial parasitic folds {see Fig. 46).

Many Style 2 minor folds have a crenulation cleavage approximately parallel to their axial planes, cutting limb and hinge areas as would an axial-planar foliation (see Figs. 47 and 48). 73

Figure 39. Musc9vite streak lineations found only at Outcrop 20-1, Sec. 20, T.46N., R.19w •• Ruler length approximately 17cm; south is to left.

Figure 40. Equal area projection of 7 muscovite streak lineations. 74

Figure 41 • .Style 1 minor folds. Bedding (B) is folded into isoclinal recumbent folds. These folds have a schistosity (S) parallel to their axial planes. The schistosity cuts bedding in hinge areas and is parallel to bedding in limb areas. Ruler length a£proximately 17cm; southeast is to left. (Outcrop 22-24, NE 4 • Sec. 22. T.46N., R.20W.) 75

Figure 42. Style 1 minor folds. Bedding (B) is folded into isoclinal overturned folds. Photo distortion makes these folds to be re-folded. These folds have a schistosity (SJ parallel to their axial planes. The schistosity cuts bedding in hinge areas and is parallel to bedding in limb areas. Rectangle in center of line diagram is area of Figure 43. Ruler length approximately 17cm; south is to right. (Outcrop 16-29, NWt, Sec. 16, T.46N., R.2ow.) 76

Figure 43. Close-up of Figure 42 showing Style 1 minor folds which have an axial-planar schistosity (S). Photo distortion makes these folds appear to be re-folded. South is to right. (Outcrop 16-29, Sec. 16, T.46N., R.2ow.)

Figure 44• Equal area projection of 43 Style 1 minor fold axes. Contours 5, 10, 15% of data per 1% area, :maximum 21%. 77

Figure 45. Equal area projection of 43 poles to Style 1 minor fold axial planes. Contours 5, 101 201 25% of data per area, maximum 33%.

Figure 46. Style 2 minor folds. Bedding and bedding- parallel schistosity are together folded into tight, overturned, rounded and angular folds. Ruler length 17cm; south is to right. (Outcrop 22-24, NE4, Sec. 22, T.46N., R.2ow.) 78

Figure 47. Style 2 minor fold (antiform). Bedding (B) and bedding-parallel schistosity (S) are together folded into tight, rounded, asymmetrical folds. A crenulation cleavage (C) is parallel to the fold 1 s axial plane cutting B and S in hinge and limb areas. Rectangle in upper right of line diagram is area of Figure 48. Ruler length 17cm; south is to left. (Outcrop 32-61, SE4, Sec. 32, T.46N., R.2ow.) 79

Figure 48. Close-up of Figure 47 iri hinge area or a Style 2 ;minor fold Cantirorm). Photo shows folded bedding (B) _and bedding-parallel schistosity (S) oriented from upper left to lower right of photo. Crenulation cleavage (C) cuts B and S from upper right to lower left of photo. South is to left. (Outcrop 32-61, SEi, Sec. 32, T.46N., R.2ow.) 80

Style 2 minor fold axes have bearings of approximately N68°W and S68°E and plunge 0 to 18 degrees along either trend (see Fig. 49).

Their axial planes strike approximately N80°W and dip io to 40 degrees

to the south (see Fig. 50).

Style 3 Minor Folds. Style 3 minor folds are open and

symmetrical or asymmetrical. The bedding and bedding-parallel

schistosity are together folded into rounded folds (see Fig. 51).

Necking of beds along lines parallel to Style 3 minor fold axes is

common (see Figs. 52 and 10). Many Style 3 minor folds have a

crenulation cleavage approximately parallel to their axial planes,

cutting hinge and limb areas as would an axial plane foliation.

Style 3 minor fold axes have two sets of orientations (see

Fig. 53). Axes of the -major set have bearings of approximately

N86qE and S86°W; axes of the minor set have bearings of approximately

N6i 0 W and S62°E. Each set plunges approximately 0 to 20 degrees along

either of its two trends. Style 3 minor fold axial planes also have

two sets of orientations (see Fig. 54). Axial planes of the major

set have a strike of approximately N84°E and dip 70 to 90 degrees to

either the north or south. Axial planes of the minor set have a

strike of approximately N59°W and dip 70 to 90 degrees to either the

north-northeast or south-southwest.

Joints

Joints are well developed throughout the map area in all rock

types (except the post-tectonic basaltic dike). Joints are particular-

ly well formed in more competent lithologies (e.g. metasiltstone and

metahasalt). 81

Figure 49. Equal area projection of 21 Style 2 minor fold axes. Contours 5, 10, 15% of data per 1% area, maximum 19%.

Figure 50. Equal area projection or 21 poles to Style 2 minor fold axial planes. Contours 6, 10, 20% of data per 1% area, maximum 22%. 82

Figure 51. Style 3 minor fold (antiform). Bedding and bedding-parallel schistosity are together folded into an open and symmetrical fold. Hammer on left end of outcrop for scale. North is into photo parallel to railroad tracks. (Outcrop 20-4, NEi, Sec. 20, T.46N., R.19W.)

Figure 52. Style 3 minor fold (open and asymmetrical antiform}. Note necking of beds {above ruler) along a line parallel to the fold axis. Ruler length 17cm; south is to right. (Outcrop 29-22, SW4, Sec. 29, T.46N., R.2ow.) 83

Figure 53. Equal area projection or 54 Style 3 minor fold axes. Contours 2, 6, 10, 15% of data per 1% area, maximum 24$.

Figure 54. Equal area projection of 54 poles to Style 3 minor fold axial planes. Contours 2, 6, 15, 30% of data per area, maximum 41%. 84

Four joint sets have been identified in the map area (see Fig.

55). Joints of the most obvious set strike approximately N02°W and dip 75 to 90 degrees to either the east or west. Joints of another

set strike approximately N55°E and dip 78 to 90 degrees to either

the northwest or southeast. The remaining two joint sets are less well-developed than the other joint sets. Joints of one lesser set

strike approximately N85°W and dip 70 to 82 degrees to the north; the .c other set strikes approximately N60°W and dips 70 to 90 degrees to

the south-southwest.

Faults

Regional faults were not identified in the map area due to poor

outcrop exposure and lack of stratigraphic control. However, a few

outcropTscale faults cutting interbedded metasiltstones and phyllites

were (See Fig. 3 (b) and Plate

The largest faults were found along the Kettle River at

Outcrops 29-44 and 32-47. The large fault at Outcrop 29-44 strikes

N52°E and dips 72 degrees to the northwest. Associated with it are

fault breccia approximately 30cm thick and normal "drag" flexures

indicating normal fault motion. One of two large faults at Outcrop

32-47 cuts the middle of the outcrop and strikes N23°E and dips 43

degrees to the .northwest. Associated with this fault are fault

breccia approximately 26cm thick and slickensides indicating a dip

slip motion. The other fault at Outcrop 32-47 (approximately 14 meters

north of the fault just described) was the largest found in the map

area. This fault strikes N49°E and dips 75 degrees to the southeast. 85

Figure 55. Equal area projection of 600 poles to joints. Readings from entire area. Contours 1 1 31 5, 10, 15% of data per 1% area, maximum 19%0 86

Associated with it are fault breccia approximately 2 meters thick and

slickensides indicating a dip slip motion. A _60-plus year old

prospect pit (Harder and Johnston, 1918) is situated along this fault

exposing breccia containing much anhedral to euhedral, fine-grained

secondary quartz. Neither of the two faults at Outcrop 32-47 exhibit

"drag" flexures or displaced marker beds, thus making direction of

fault motion indeterminable.

.c Three small reverse faults having only a few centimeters of

displacement were found at Outcrop 29-44. These faults strike

approximately east-west and dip 24 to 28 degrees to the south. "Drag"

flexures were found along these faults indicating a reverse motion. No

fault breccia exists along these faults.

Boudins

Quartz boudins in beds of phyllite are common in the map area.

Many quartz bodies (described in the petrography chapter) were strained

into oblate boudins which lie in the plane of the schistosity (see

Figs. 56 and 57). Separated quartz boudins are wrapped by the enclos-

ing phyllite (a relatively incompetent and ductile rock type) suggesting

a period of ductile flow of phyllite into the gaps between the boudins.

The quartz boudins themselves also show signs of previous ductile

behavior by having tapered and drawn-out edges. 87

Figure 56. Oblate quartz boudins (stippled) in phyllite. Note def'lection of' bedding {B) and wrapping of bedding- parallel schistosity (S) around boudins. Figure 57 shows face of same outcrop around the corner from the race shown here. Ruler length approximately 17cm; south is to rig.ht. (Outcrop 20-34, NE!, Sec. 20, T.46N., R.19W.) 88

( B+S >

Figure 57. Oblate quartz boudins (stippled) in Note deflection of bedding (B) and wrapping of bedding- parallel schistosity (S) around boudins. Figure 56 shows face of same outcrop around the corner from the face shown here. Ruler length approximately 18cm; east is to left. (Outcrop 20-34, NEt, Sec. 20, T.46N., R.19W.) CHAPTER FIVE

INTERPRETATION

I. Introduction

This chapter will integrate data presented in the preceeding chapters in attempts to formulate interpretations concerning lithologi- cal and structural aspects of the Thomson Formation in relation to deformational and metamorphic events of the Penokean Orogeny. Data are often limited making some interpretations highly speculative.

Nevertheless, many of the following interpretations are well founded and important in completing the detailed picture of the Thomson

Formation and increasing our understanding of the deformational history of the Penokean Orogeny.

II. Sedimentology

Protoliths

Protolith determination is to some extent speculative due to extensive metamorphic recrystallization of all Thomson units in the study area. The high phyllosilicate content and fine-grained texture of the phyllitic rock types indicate that they were deposited slowly as shales in quiet water below wave base. Deposition of this material was periodically interrupted by the deposition of silt beds

(since metamorphosed to metasiltstones) by currents entering the area.

Protoliths of isolated calcareous rnetasiltstones and phyllites were

89 90 local calcareous siltstones and limy muds, respectively. Also, the graphitic phyllites probably developed from mQds rich in organic matter (Blatt and others, 1972 and Degens, 1967).

Provenance

Lack of necessary data restricts the interpretation regarding provenance of the metasiltstone and phyllite units. Extensive metamorphic recrystallization has eliminated all primary structures useful as paleocurrent indicators, and has resulted in heavy minerals reflecting to a large degree facies of metamorphic recrystallization rather than provenance. However, work in the Thomson Formation north of the study area (Morey and Ojakangas; Hyrkas, 1981) has shown that

Thomson sediments were deposited by southward flowing turbidity currents. Also north of the study area, metasediment coarser than metasiltstone has been found (i.e. metagreywacke). Tilerefore,

Tilomson sediments in the study area probably had a northern source and were deposited by southward flowing turbidity currents more distal than those which deposited sediments further north.

Stratigraphic relationships between metabasalt and volcaniclastic rock types are uncertain due to poor and discontinuous exposures, but there is a spatial and mineralogical relationship between them. The metabasalt and volcaniclastic units are found in extremely close proximity to each other (see Fig. 3 (b) and Plate 1), and contain essentially the same minerals (see Table 3). Also, absence of pillows in the metabasalts suggests that they were hypabyssal dikes and/or sills. Tilerefore, these dikes and/or sills may have been a source 91 rock for detritus for the volcaniclastic metasiltstones and phyllites.

Volcanic vents associated with these intrusive·s may also have been a source of pyroclastic detritus for the volcaniclastic units.

III. Stratigraphy

Stratigraphic relationships and thickness estimates of Thomson

Formation units in the study area were indeterminable due to lack of ., stratigraphic control, poor and discontinuous exposures, and complex regional folding (described later in this chapter).

IV. Regional Metamorphism

All Thomson Formation units in the study area have been regionally metamorphosed. As a result, all Thomson units have a well developed continuous schistosity (see Figs. 11, 12, 18, and 22) characteristic of most regionally metamorphosed rocks (Turner, 1981).

Grade at culmination of metamorphism was uppermost greenschist/ lowermost amphibolite facies. Minerals found that are characteristic of upper greenschist facies are: albite/oligioclase (An for ground- 13 mass grains in metabasalts), quartz, biotite, muscovite, chlorite, ferroactinolite, and epidote. Lowermost amphibolite facies grade is indicated by the presence of garnet porphyroblasts and absence of staurolite and kyanite porphyroblasts.

V. Introduction to Structural Analysis

Method of Structural Analysis

The Thomson Formation in the study area was multiply deformed during the Penokean Orogeny. Evidence of multiple deformation was 92 found as multiple minor fold styles and multiple foliations.

Mesoscopic (i.e. outcrop scale) and structures such as these were used to determine macroscopic (i.e. regional scale) structures using the method of structural analysis described by Turner and Weiss

(1963).

Minor folds can be used to determine geometries and orientations of major folds as discussed by Turner and Weiss (1963, p. 188):

.c "In any tectonite body, however complex its structure, it is generally possible to find a hand specimen or a single exposure in which all the geometric properties of .the macroscopic body (i.e. regional scale structure) are displayed on the mesoscopic scale (i.e. outcrop scale)."

·rn other words, it connn.only can be assumed that minor structures mimic

(usually to a high degree) major structures formed during the same tectonic event.

· Foliation is a connn.on mesoscopic and microscopic structure found in most regionally metamorphosed tectonites. Foliation develops in response to flow under high non-hydrostatic stress. This strain phenomenon results in irreversible structural rearrangement of rock materials down to the atomic scale. As a result, foliations are textures of great tectonic significance.

Foliations form approximately perpendicular to the direction of maxilDUm finite shortening in rocks during a given period of deforma-

tion (Hobbs and others, 1976). Wood (1974) determined that the presence of slaty cleavage indicates a compressive strain exceeding approximately 50%. And Ramsey (1967) stated that all transitions between slaty cleavage and schistosity may be found in low- and medium-

grade metamorphic rocks suggesting that schistosity probably has the i 1· I 93 same mechanical significance as slaty cleavage. Therefore, it is valid to assume that a schistosity commonly develops during regional metamorphism as a response to contemporaneous tectonism, and it forms planes which are perpendicular to the direction of maximum finite shortening during that tectonic event. Since only one direction of maximum finite shortening exists during a given period of deformation, only one schistosity will first develop as a foliation in a multiply deformed rock; succeeding deformations will then result in crenulation cleavages as foliations.

Therefore, by recognizing various minor fold styles and foliations, and determining the geometrical relationships between them, conclusions can be drawn concerning deformational events and resulting regional structures.

Notation

Three Penokean deformations and their resulting structures have been recognized in the study area (as discussed below). The following is a list of shorthand notation used in this chapter that ref er to the various deformations and resulting structures.

Deformations

D first deformation 1 = second deformation o2 = D third deformation 3 = 94

Fold Generations

F folds of 1 = n1 F A n folds of one geometry 2 = 2 _F B folds of another geometry 2 n2 Note: F A and F B folds developed contemporaneously 2 2 during n . 2

Bedding and Foliations

so bedding

= foliation (schistosity) of sl n1 s2 = foliation (crenulation cleavage) of D2

s3 = foliation (crenulation cleavage) of D3

Lineations

L C crenulation lineations of 2 = n2 L C crenulation lineations of 3 = n3 L A amphibole lineations of 3 n3 L M muscovite streak lineations of 3 = n3

VI. Structural Analysis: The Three Deformational Events and Resulting Structures of the Penokean Orogeny Found in the Study Area

First Deformation (D 12.. During the First Penokean deformation (D ), bedding (S ) of all 1 0 Thomson Formation units was folded into overturned or recumbent isoclinal folds (F ) described in the structure chapter as Style 1 1 folds (see Figs. 41, 42, and 43). The regional, continuous schistosity

(see Figs. 11, 12, 18, and 22) developed during D as an axial-planar 1 foliation (S ) to the F folds. (See Fig. 58.) Also associated with 1 1 Norfh ' ) 5o"l-fh

So S1 ,,... -______? / ,,,, ,. ,. .,,. "-.,, ,. ,, ,,.. -... '- - - ,,.. ------/ ,,. / ; / / ,,,, ,_/ ,, - -r ------, ,,,, / + ·-,,""' ------J

Figure 58. Structures of the first Penokean deformation (D1 ). Bedding (s0 ) was folded into overturned or recumbent isoclinal folds (F1 , described in the structure chapter as Style 1 folds}. An axial-planar schistosity (S1) developed with the F1 folds and is usually found parallel to bedding except in hard to l.Jl find hinge arease Diagram is highly schematic; implied regional scale unknown. 96

was the development of oblate quartz boudins which lie in the plane n1 of the schistosity (see Figs. 56 and 57). s1 Axes of the F folds have bearings of approximately N78°E and 1 S78°W and plunge 0 to 25 degrees along either trend (see Fig. 44). The

F axial planes s,trike approximately N74 °E and dip 0 to 40 degrees to 1 the south (see Fig. 45). The size of mesoscopic F folds varies from a 1 few centimeters (see Figs. 42 and 43) to a little more than a meter

(see Fig. 41) . The size of macroscopic F folds is assumed to be 1 substantial, but is indeterminable due to the multiply-deformed nature of the rocks, lack of stratigraphic control, and poor and discontinuous exposures.

The F axial plane schistosity (S ) strikes approximately east- 1 1 west and dips 10 to 35 degrees to the south (see Fig. 29). The s1 schistosity is approximately parallel to the F axial planes (compare 1 Figs. 29 and 45). and have bedding/cleavage relationships s0 s1 consistent with those of ideal isoclinal folds. For instance, cuts s1 across at high angles in areas of mesoscopic F folds (see s0 1 Figs. 42 and 43) and in hinge areas of somewhat larger F folds where 1 the entire fold closure cannot be seen at mesoscopic scale (see Figs.

27, 28, 30, and 31). Also, is parallel to in limb areas of s1 s0 mesoscopic F folds (see Figs. 41, 42, and 43) and in limb areas of 1 macroscopic F folds where only the bedding and bedding-parallel 1 s1 schistosity are visible at mesoscopic scale (see Figs. 10, 56, and 57).

F hinge areas are rare and difficult to find; as a result, is 1 s1 usually found parallel to at most outcrops. s0 97

The orientation of folds developed during a tectonic event can be used to determine the approximate direction of tectonic transport during that event. Assuming tectonic transport to be in a direction approximately perpendicular to fold axes, appears to have proceeded n1 along a north-south line. And as described in the regional geology chapter, regional metamorphic grade of the Thomson Formation increases from north to south. Therefore, tectonic transport during probably n1 proceeded along a line from south to north (i.e. from high to low metamorphic grades).

The regional s schistosity found in all Thomson Formation units 1 together with the overturned or recumbent isoclinal F folds allude to 1 an intense first deformation. A rough idea of the degree of tectonic intensity during can be obtained from the oblate quartz boudinage n1 (see Figs. 56 and 57). Presence of oblate boudins indicates extension in all directions within the _boundinage layer with considerable - shortening along the direction of maximum finite shortening oriented perpendicularly to the boundinaged layer. The quartz boundinage layers lie in the planes, not in the planes of the succeeding s1 foliations indicating the boudinage developed contemporaneously with the schistosity during : Sylvester and Christie (1968) used s1 n1 oblate boudinage to estimate strain and demonstrated approximately 80% shortening along the direction of maximum finite shortening. Their study dealt with lithologies similar to those of this report. Therefore, by analogy, it is concluded that the first deformation of the Penokean

Oro&eny was characterized by considerable shortening along the direction of maximum finite shortening. Detailed boudinage studies were not done \

98 in the study area; therefore, strain in the area cannot be defined quantitatively in this study.

Second Deformation (D 2_ 2 During the second Penokean deformation (D ), bedding (S ) and 2 0 the essentially bedding-parallel schistosity were together folded s1 into two distinct and contemporaneous fold geometries. Folds of one geometry are termed F A' and folds of the other geometry arie termed 2 F B. F A folds were described in the structure chapter as Style 2 2 2 folds. These folds are tight and asymmetrical or ·overturned (see

Figs. 46 and 47). F B folds were described .in the structure ·chapter 2 as Style 3 folds. These folds are open and symmetrical or asymmetrical

(see Figs. 51 and 52). A discontinuous crenulation cleavage (see

Figs. 32 and 33) was developed during D as an axial-planar foliation 2 (S ) to the F A and F B folds (see Figs. 47 and 48). Also, a crenula- 2 2 2 tion lineation (LZC' Fig. 36) was found associated with the s2 crenulation cleavage. (See Fig. 59.)

Axes of the F A folds have bearings of approximately N68°W and 2 S68°E and plunge 0 to 18 degrees along either trend (see Fig. 49).

F A axial planes strike approximately N80°W and dip 10 to 40 degrees 2 to the south (see Fig. 50). The size of mesoscopic F A folds varies 2 from a few centimeters (see Fig .• 46) to about half a meter (see Fig.

47). The size of macroscopic F A folds is on the order of at least 2 tens of meters. The size and extent of macroscopic F A folds were 2 estimated using an F - generation "Z" minor fold and its geometrical 2 relationship to the outcrop in which it was found (see Figs. 60 (a), NorfA L

s?- ,,,,,,.... -" .. "' ,, ,)" , ,, \ , , , \ So+S1 " " , \ , " , \ < < < .\ \ \ \ ' \ ' \ \ \ \ \ ) \ \ ' / \ \ ' \ / \ \ ,. , + \ ). , ''''\1 v " v Figure 59. Structures of the second Penokean deformation (D2 ) superimposed over the n1 structures. Bedding (s0 ) and the essentially bedding-parallel s1 schistosity were together folded into tight and asymmetrical or overturned folds described in the structure chapter as Style 2 folds). F2B folds \0 are not shown on this diagram. An axial-planar crenulation cleavage Cs2 ) \0 developed with these foldse The diagram is highly schematic; implied regional scale unknown. 100

(b), and (c)). The regional homoclinal orientation of bedding

(striking east-west and dipping 10 to 35 degrees to the south, see

Fig. 26) is due to the geometry, orientation, size and regional extent

of the macroscopic F2A folds (see and compare Figs. 59 and 60).

Axes of the F B folds have two well defined sets of orientations 2 (see Fig. 53). Axes of one set have bearings of approximately N86°E and

S86°W; axes of the other set have bearings of approximately N62°W and

S62°E. Each set plunges approximately 0 to 20 degrees along either of its two trends. Axial planes of the F B folds also have two sets of well 2 defined orientations (see Fig. 54). · Axial" planes 'of one set have a strike of approximately N84°E and dip 70 to 90 degrees to either the north or south. Axial planes of the other set have a strike of approximately N59°W and dip 70 to 90 degrees to either the north-northeast or south-southwest. The size of mesoscopic F B folds varies from approxi- 2 mately one meter to eleven meters (see Figs. 51 and 52). No field evidence was found enabling an estimation of the size of macroscopic F B 2 folds.

The axial-planar crenulation cleavage strikes approximately s2 N80°E and dips 60 to 90 degrees to either the north or south (see

Fig. 34). In the field, is found as an axial-planar foliation to . s2 both the F A folds (see Figs. 47 and 48) and FZB folds (compare Figs. 2 34 and 54). Admittedly, there is a discrepancy in the F A axial plane 2 and orientations (compare Figs. 34 and 50). However, the non- s2 parallelism of the crenulation cleavage and F A axial planes can be s2 2 explained by the phenomenon of "fanning" and "refraction" .of axial- planar foliations. According. to H6bbs and others (1976}, fanning. and refraction of an axial-planar foliation is common in a tectonite body 101

Ke.ff le River

Figure 60(a). Outcrop along the Kettle River showing bedding (So) and the essentially bedding-parallel S., schistosity striking approximately east-west and dipping to the south (south is to right). The F2-generation "Z" minor fold shown in Figure 60(b) is found at this outcrop within the circle located left of center in the line diagram. (Outcrop 29-45, SEt, Sec. 29, T.46N., R.20W.} 102

Figure 60(b). Bedding (s0 ) and the essentially bedding- parallel s1 schistosity together are folded into an F2- generation "Z" minor fold. s0 and s1 are dipping to the south (south is to right). Photo location is within the circle located left of center in the line diagram or Figure 60(a). (Outcrop 29-45, SE!, Sec. 29, T.46N., R.20W.) ,,.,... - ---. - .--... / ------/ -- -- I -- -- ...... __ __ I / / - - - -- f ( -- - - Figure 60 ( c). Diagram combining l \ Figures 60(a) and 60(b) to give ...... macroscopic F2A fold geometry and \ ...... ' approximate size. An F2-generation nz1t minor fold (Fig. 60(b)) at ...... this outcrop (Fig. 60(a)) indicates ' ...... ' that this outcrop is part of an overturned limb of an anticline closing to the north and/or a syncline closing to the south ...... (south is to right). Therefore, ...... sizes of F2A. macroscopic folds are Ke.ff :·, on the order of at least tens or River- ...... J----- ', meters. Also, the geometry, ...... \ orientation, size, and regional ...... extent of these F2A macroscopic \ i'olds explains the honioclinal -- ...._ -- ' \ orientation of bedding in the I thesis area (i.ee striking east- -- ..._ ...... I' .west and dipping 10 to 35 degrees -- / - I Q ...... _.,.,. to the south) -, _- -.... / / -- - - I-' / 0 .... -- -- ...... ______,,,. w having interbedded units of differing competencies (e.g. metasiltstone and phyllite of the study area) of ten resulting in an axial-planar foliation not everywhere parallel to fold axial planes.

In places where the SZ crenulation cleavage did not penetrate more competent lithologies a crenulation lineation (Lzc) developed locally instead. As a result, LZC lineations have trends parallel to the strikes of the SZ cleavages (comP.are Figs. 34 and 37).

The ·1ZC crenulation lineations have bearings of approximately N84°E and S84°W and plunge 0 to 18 degrees along either trend (see Fig. 37 and compare with Fig. 34).

As stated above, the FZA and FZB folds developed contemporaneously during the second deformation. Evidence supporting this conclusion includes FZA and FZB fold axes having approximately the same trends

(compare Figs 49 and 53), and FZA and FZB axial planes having mately the same strikes (compare Figs. 50 and 54). Also, the SZ crenulation cleavage was found as an axial-planar foliation to both the

FZA and FZB folds.

Finally, as discussed above, tectonic transport during a deformational event can be assumed to be in a direction approximately perpendicular to the axes of folds developed during that deformation. Therefore,

Dz appears to have proceeded along a north-south line (as did D ). And 1 / ' because metamorphic grade of the Thomson Formation increases from north to south, it may also be assumed that tectonic transport during Dz proceeded along a line from south to north from high to low metamorphic grades) as did the first Penokean deformation. Alorfh L

Joi t1f-.s 53 ,------.,- r-..L.-----,I I I I I r-.i..-----, I it I I I I I ,. _ .J• ,.I __ _,I + L.------'I Figure 61. Structures of the third Penokean deformation (D3) superimposed over the n1 and n2 structures. A crenulation cleavage (83) developed during n • Also during n , joints developed after the s crenulation cleavage 3 3 3 I-' (joints not shown on upper diagram)Q The diagram is highly schematic; 0 implied regional scale unknown. 107

Figure 62. Photomicrograph of the s3 crenulation cleavage (upper right to lower left) crenulating the s2 crenulation cleavage (upper left to lower right) indicating that s3 does indeed post-date s2• Field length approximately 5.9mm. Plain light. (Outcrop 20-6, SWi, Sec. 20, T.46N., R.19W.) 108

(L A) and muscovite streak lineations (L M). The L A mineral linea- 3 3 3 tions have bearings of approximately Sl0°E and plunge 0 to 25 degrees

(see Fig. 38 and compare with Figs. 34 and 37). The L M mineral 3 lineations have bearings of approximately Sll 0 W and plunge 4 to 14 degrees (see Fig. 40 and compare with Figs. 34 and 37).

Evidence also indicates that joints found in the thesis area developed after the s crenulation cleavage during the third deforma- 3 tion. Theoretical joint sets that could develop in a tectonite body are shown by Price (1968) in a block diagram (see Fig. 63 {a)) and on a stereogram (see Fig. 63 {b)). Price (1968) states that the direction of maxinrom principal stress bisects the acute angle formed by the intersection of the two shear joint sets that may develop during a tectonic event (see Fig. 63).. The direction of maximum principal stress can be assumed parallel to the direction of maximum finite shortening if strain during tectonism is (or is close to being) irrotational (Ramsey, 1967).

As described in the structure chapter, four joint sets have been identified in the thesis area. These joint sets strike approximately N02°W, N55°E, N85°W, and N60°W, and are approximately vertical (see Fig. 55). These joint sets are plotted on a stereogram

(see Fig. 64) and compared with the theoretical joints of Price (1968)

(compare Figs. 63 and 64). Similarities can be seen upon comparison of the observed and theoretical joint sets. The observed joint sets striking N60°W and N55 9 E have the relative orientations as do the theoretical shear joint sets. Therefore, it is assumed that these two joint sets are shear joint sets. Also, the observed joint 109

Figure 63. (a) Block diagram or theoretical joint sets that could develop in a tectonite body; {b) stereogram Of these theoretical joint sets (Price, 1968). S 1 and S'' are shear joint planes, and ac and be are tension joint planes. The bisector of the acute angle formed by the intersection of S' and S'' is assumed parallel to the direction of maximum finite shortening if strain during tectonism is (or is close to being) irrotational (Ramsey, 110

Figure 64. Stereogram or the study area joint sets. These joint sets strike approximately N2°W, N55°E, N85°W, and N60°W, and have dips or approximately 90 degrees (see Figure 55). 111 sets striking N85°W and N02°W have the same relative orientations as do the theoretical tension ac- and be-joint sets, respectively.

The bisector of the acute angle formed by the intersection of the two observed shear joint sets (i.e. the directi9n of maximum finite shortening) has a horizontal trend of approximately east-west.

This direction is perpendicular to the s crenulation cleavage. 3 Therefore, the creunlation cleavage and the joints could have had s3 the same direction of maximum finite shortening during their development suggesting that both formed during • However, as mentioned n3 above, the joints developed after the crenulation s3 cleavage because joints are the result of brittle deformation (i.e. rupture) which usually follows a period of ductile deformation

(e.g. deformation resulting in a crenulation cleavage) (Billings,

1972).

Finally, little can be said about the direction of tectonic transport during the third defamation due to the absence of n 3 folds. However, the direction of maximum finite shortening during

is approximately horizontal and trends. east-west. Therefore, it n3 may be assumed that proceeded along a nearly horizontal east-west n3 line instead of a north-south line as did n and n • The east-west 1 2 compression of may have been associated with the emplacement of n3 the Middle Precambrian plutonic rocks during and after the waning stages of the Penokean Orogeny. 112

VII. Relative Time Relationships Between the Three Penokean Deformational Events and Regional Metamorphism

Method of Interpretation

Time between deformation .and metamorphism can be

inferred. from the geometrical and textural relationships between

porphyroblasts and adjacent matrix in a deformed and metamorphosed

rock. Garnet prophyroblasts can be interpreted as pre-, syn-, or

post-tectonic in origin based on textural relationships as described

by Spry (1979). Acicular porphyroblasts (e.g. ferroactinolite) can also be interpreted as pre-, syn-, or post-tectonic based on their

relative orientations to regional structures (e.g. folds and folia-

tions} (Spry, 1979; Hobbs and others, 1976). These relationships will be discussed and used pelow to correlate the three Penokean deformations with the contemporaneous regional metamorphism.

First Deformation (D 2_ 1 Culmination of metamorphism in the thesis area (uppermost greenschist/lowermost amphibolite facies) is characterized by garnet porphyroblasts in pelitic rocks. Presence of pre-, syn-, and post- tectonic garnet prophyroblasts associated with the schistosity s1 (see Figs. 15, 16, and 17, respectively) indicate that regional metamorphism and D tectonism were contemporaneous events. These 1 porphyroblasts also suggest that metamorphic culmination was reached before, and continued during and after, the first Penokean deformation. 113

Second Deformation (D 2._ 2 Absence of syn- and/or post-s garnet porphyroblasts indicate 2 that regional metamorphic grade was lower during D than it was during 2 D • This observation along with the presence of post-s epidote 1 1 porphyroblasts (see Fig. 20) and large chlorite _flakes (see Fig. 14) discordant to s and parallel with suggest that regional 1 s2 metamorphism continued during the second deformation at greenschist facies.

Third· Deformation

Syn- and/or post-s garnet porphyroblasts are also absent 3 suggesting that regional metamorphic grade was lower during D than it 3 was during D • Regional metamorphism continued during D at 1 3 facies as indicated by the presence of preferentially oriented ferroactinolite porphyroblasts and muscovite streaks. The ferroactinolite porphyroblasts in the metabasalts, volcaniclastic metasiltstones and phyllites, and metamorphosed calcareous concretions recrystallized during D to become aligned parallel to the strike of 3 the s crenulation cleavage (compare Figs. 34 and 38), and parallel 3 to the trend of the L C crenulation lineations (compare Figs. 37 and 3 38). Also, muscovite recrystallized into concentrated streaks during

D to hecome aligned parallel to the strike of s (compare Figs. 34 3 3 and 40), and parallel to the trend of L C (compare Figs. 37 and 40). 3 114

VIII. Structural and Tectonic Relationships Between the Study Area and the Rest of the Thomson Formation

Introduction

This section will attempt to correlate structures found in the

study area to those· found in other areas of the Thomson Formation.

\ However, due to the lack of detailed structural data elsewhere in the

1homson Formation, few conclusions can be drawn concerning· effects

and extents of the three Penokean deformations on the entire formation. ·

First Deformation (D ) Structures 1 The bedding-parallel sl schistosity has been found in the .

southernmost Thomson Formation exposures near the town of Denham

(Neumann, 1981) approximately 9km south-southwest of the study area

(see Fig. 2) .: The northernmost recognized appearance of the s 1 schistosity is near the town of Atkinson (Hyrkas, 1981; Wright and

others, 1970) approximately 30km northeast of the study area (see

2).

A mesoscopic fold similar to F folds in style and geometry 1 , has been found in the Thomson Formation near Park Lake approximately

7km west of Atkinson (Hyrkas, 1981). The axis and axial plane of

this fold are approximately parallel to the axes and axial planes of

the F folds found in the study area .. Other such folds have not been 1 reported (at the time of this report) north of Atkinson or south of

the study area.

Structures of the first Penokean deformation (e.g. bedding-

.parallel schistosity, F -type folds, etc.) have not been found in the 1 Thomson Formation north of Atkinson. However, it should be kept in 115 mind that effects of D may indeed be present in the Thomson Formation 1 north of Atkinson, but they have not been recognized at the time of this report.

Second Deformation (D ) Structures 2 The s crenulation cleavage has been f-0und in southernmost 2 Thomson Formation exposures near the town of Denham (Neumann, 1981).

It is found there as an east-west striking crenulation cleavage having a nearly vertical dip. The northernmost appearance of the s 2 crenulation cleavage is near the town of Atkinson (Hyrkas, 1981) where it has a similar attitude. The s crenulation cleavage has not 2 been recognized (at the time of this report) north of Atkinson

(Hyrkas, 1981 and Wright and others, 1970). In other words, the s crenulation cleavage has not been found in Thomson Formation rocks. 2 lacking the sl schistosity.

On the other hand, a slaty cleavage has been found in the

Thomson Formation from Atkinson northward to the town of Carlton

(Hyrkas, 1981; Keighin and others, 1972) and north of Carlton (Wright and others, 1970) (see Fig. 2), but has not. been found south of

Atkinson (Hyrkas, 1981; Neumann, 1981; this report). In other words, the slaty cleavage is found in rocks lacking the s schistosity and 1 s crenulation cleavage. Th.e slaty cleavage of the northern Thomson 2 Formation has the same orientation as the s crenulation cleavage 2 found in rocks of the southern ·Thoms.on Formation.

Folds of the second Penokean deformation found in the study area (e.g. F A folds) are similar to folds at the Thomson Formation 2 116

type locality near the town of Thomson (compare Figs. 59 and 9). Recall

that F A folds are tight and asymmetrical or overturned (see Figs. 59 2 and 60). These folds have axial planes that dip 10 to 40 degrees to

the south (see Fig. 50), axes that plunge moderately to the east-south-

east and west-northwest (see Fig. 49), and an axial-planar crenula- s2 tion cleavage (see Figs. 47 and 48). In comparison, folds found near

the town of Thomson are broad, open, and symmetrical becoming tight,

asymmetrical and overturned in rocks having lower lithologic compe-

tencies (see Fig. 9). These folds have axial planes that dip at high

angles to the south (see Fig. 9), axes that plunge moderately to the

east and west, and an axial-planar slaty cleavage which has an orienta-

tion similar to that of the crenulation cleavage. s2

Third Deformation (D ) Structures 3 Structures of the third Penokean deformation are characterized

as having either a north-south strike Ge.g. the crenulation s3 cleavage) or a north-south trend (e.g. the L C crenulation lineations). 3 No structures having a north-south orientation have been found in the

southernmost Thomson Formation .near the town of Denham (Neumann, 1981), .

nor in the central Thomson Formation near Atkinson (Hyrkas, 1981).

However, north-south trending "linear structural elements" have been

found in the northern Thomson Formation by Keighin and others (1972)

from Atkinson northward to the town of Carlton. Therefore, structures

resulting from D appear to be best developed in rocks of the study 3 area; however, poor structural data may indicate the presence of D 3 structures in rocks of the northern Thomson Formation from Atkinson northward to Carlton. 117

IX. Re.gional Mddels· ·

Intrqduction

The following two regional structural models suggest relative

time relationships between the Penokean deformations. These models

will be referred to in the next section where tectonic models for the

region are d-iscussed.

Separate Deformations Model

This model suggests that the Penokean Orogeny was a series of

distinct deformations, with tectonic transport proceeding along a

line from south to north. As a result, the first Penokean deforma-

tion (D ) folded the southern Thomson Formation into overturned or 1 recumbent isoclinal folds (F ) and developed an axial-planar schistosity 1 ) (see Fig. 65 (a)). At the. same time, rocks in the north were left Cs1 undisturbed.

The second Penokean deformation (D ) re-folded the southern 2 Thomson Formation into tight and asymmetric or overturned folds (F A) 2 (see Fig. 65 (b)). Contemporaneously, rocks in the north were folded

into asymmetric folds. Also during this time, an axial-planar crenula-

tion cleavage (s ) developed south of Atkinson, and an axial-planar 2 slaty cleavage developed north of Atkinson.

Structural relationships are uncertain in the Atkinson area.

Therefore, conclusions as to why structures of the first deformation are not recognized north of Atkinson cannot be made at this time using this model. Stv.dy A,.. 1\501\ {a.) e°' A+k; 1ho1>15011 1 -- cs1 -- - - -..,2"'" - -? < Soufh - - -- . Norfh · --- -- ...... -- - > (b) Area. A+kit\SOll ea.va..Je. S lo. f:.y Cf 1

• ( Soufh fV'orth)

Figure 65. Separate deformations model. (a) The first Penokean deformation n1 folded the southern Thomson Formation into overturned or recumbent isoclinal folds (F1 ) and developed an axial-planar schistosity (s1 ), while rocks in the north were left undisturbed. (b) The second Penokean deformation (D2 ) re-folded rocks in the south into tight and asymmetric or overturned folds (F2A), and folded rocks in the north into asymmetric folds; an axial-planar crenulation f--' f--' cleavage (S2) developed in the ·south, and an axial-planar alaty cleavage 00 developed in the Diagrams not to scale. 119

Progressive Deformation Model

This model suggests that the Penokean Orogeny was one progres-

sive deformation, with tectonic transport proceeding along a line from

south to north. Early during the Penokean Orogeny, the southern

Thomson Formation was folded into overturned or recumbent isoclinal

folds while open folds developed in the north (see Fig. 66 (a)). At

the same time, an axial-planar schistosity (S ) developed in the south 1 contemporaneously with an axial-planar slaty cleavage in the north.

The Atkinson area served as a transitional area where the schistosity

dissapated and the slaty cleavage appeared.

With continued deformation, rocks in the south became strain-

hardened causing a change in the geometry of folds being produced. As a result, the overturned or recumbent isoclinal folds were re-folded

into tight and asymmetric or overturned folds (F A) which developed an 2 axial-planar crenulation cleavage (S ) (see Fig. 66 (b)). At the same 2 time, deformation in the north progressed resulting in asymmetric folds and continued development of an axial-planar slaty cleavage.

X. Penokean Tectonic Models

Introduction

The following are two current tectonic models for the Penokean

Orogeny. After a short description of each, both tectonic models will be applied to the two regional structural models discussed above.

Plate Tectonic Model

Van Schmus (1976) proposed a Penokean tectonic model which dealt with the Middle Precambrian rocks of Minnesota, Wisconsin, Michigan, Atkinson 1homsol'\ (a) Cle o.v"'.je. --Ft ...... I r:-:1/ - ....---...... ____.. _ frAn$if ion I ...... / - - - - ...... I / / / I / I .So"fh North)

St1.1 _d"e a. At ki J'\ SOil Tiurnsot1

(b) I I

I I

i fi of» I I ( Sot

Orogeny where a north-dipping subduction zone existed to the south of the Precambrian rocks exposed in this area (see Fig. 67). His model explained the region of volcanic rocks in Wisconsin as an island arc region, and the region of .granitic rocks as the eroded, roots of the arc (see Fig. 67). The Middle Precambrian sedimentary and hypabyssal intrusive rocks of Minnesota and Michigan would then represent rocks of the back-arc region, with southwest Ontario near the shoreline (see

Fig. 67). He added that warping of the continental crust during the

Penokean period due to compression at the plate margin resulted in the following sequential changes in sediment deposition in the back-arc region: orthoquartzite and carbonates, to iron formations in restricted basins, to deepwater muds and greywackes (e.g. the Thomson Formation).

Since Van Schnrus, several other plate tectonic models having south- as well as north-dipping subduction zones have been presented for the

Penokean period (International Proterozoic Symposium--Abstracts, 1981).

Intracratonal Model

Sims (1976) suggested a non-plate tectonic model for the

Penokean Orogeny which dealt with the Lake Superior region. As discussed in the regional geology chapter, Morey and Sims (1976) recognized two Lower Precambrian terranes in the Lake Superior region separated by a well delineated zone later known as the Great Lakes

Tectonic Zone (Sims and others, 1980) (see Fig. 6). This zone an older Lower Precambrian gneissic terrane to the south (greater than

3000 m.y. old) from a younger Lower Precambrian greenstone-granite 122

0 200km T= 1900 Ma ago Michigan Onlario Wisconsin Illinois Minnesota NW SE

0 volconic D eugeo - sedimentary • bonded Iron Fm. n 0 miogeo- 1ed. LJ cont. crust [] oceanic crust ml!l!lJ mantle

Figure 67. A possible reconstruction of the margin of the Archean craton about 1900 m.y. ago during the ear1y stages of the Penokean Orogeny (before intrusion of plutonic rocks in the volcanic arc and associated metamorphism and deformation). Modified after James (1954), but the presence of the subduction zone is stil1 purely hypothetical, and other models are quite (Van Schmus, 1976) 123

terrane to the north (2700-2500 m.y. old) (see Fig • . 6). Sims (1976)

proposed that the Penokean Orogeny was an · intracratonal event result-

ing from reactivation along the zone separating the two Lower Precam-

brian terranes (i.e. the Lakes Tectonic Zone). Morey and Sims

(1976) also suggested that associated with this event was the tectonic

reactivation of the McGrath Gneiss as a mantled gneiss dome (as dis-

cussed in the regional geology chapter).

Application of the Regional Structural Models to the Penokean Tectonic Models

Either an east-west trending subduction zone located to the

south or a reactivated mantled gneiss dome situated beneath the southern

margin of the Animikie rocks could have produced enough uplift and

compression to result in the structures found in the Thomson Formation.

In addition, either tectonic scenario could have generated one progres-

sive deformation or a series of discrete deformational pulses.

Uplift from either the diapiric rise of a mantled gneiss dome

or from compression at a convergent plate boundary induced gravity

gliding of Thomson Formation rocks down the flanks of the uplift.

Gravity gliding and tectonic compression resulted in deformation of

the surrounding rocks. Folds that developed in this way were recumbent or asymmetric near the seat of tectonism, while folds further away were less severely deformed or not even generated. The overturned or steep li1!lhs of the recumbent or asymmetric folds may have developed thrust faults, thus possibly resulting in nappe-like

structures. Also, regional metamorphism due to high heat flow from the tectonically active zone decreased in grade away from the zone. 124

Continued deformation and strain-hardening, or a second deformational pulse, then re-folded the earlier recumbent or asymmetric folds located near the seat of tectonism. Contemporaneously, folding continued, or new folds were generated, further away from the tectonically active zone. • •CHAPTER SIX

.CONCLUSION

The following is a sequential list outlining the Middle

Precambrian history of the study area.

1. At the end of Early Precambrian (Archean) time, the Animikie Basin

developed in what is now east-central Minnesota. Sedimentation in

the basin began. approximately 2000 to 2100 m.y. ago (Morey, 1978).

2. In the study area, shales were deposited in quiet water below wave

base. Silt beds were intermittently deposited by southward flowing

distal turbidity currents. Hypabyssal dikes and/or sills were in-

truded and later may have been a source rock for detritus for

yolcaniclastic sediments found nearby. Volcanic vents associated

with the intrusives may also have been a source of pyroclastic

detritus for the volcaniclastic units, Local calcareous silts,

and limy and organic-rich muds were also deposited.

3. The Penokean Orogeny deformed. the rocks in the study area approxi-

mately 1870 m.y. 'ago (Morey and Lively, 1980). Three deformations

have been recognized in the study area:

a. The first Penokean deformation folded the Thomson For:mation

into overturned or recumbent isoclinal folds haying axes that

plunge at low angles to the east and west, and axial planes

that are horizontal or dip at low angles to the south. A

125 126

regional, continuous schis_tosity developed as an axial-planar

foliation to these folds. Also, oblate quartz boudins formed

.which lie in the plane of the axial-planar

b. The second Penokean deformation re-folded the Thomson Formation

into: tight and asymmetrical or overturned folds, and open

and symmetrical or asymmetrical folds. Both fold geometries

developed contemporaneously and have axes that at low

angles to the east and west. Axial planes of the tight folds

dip 10 to 40 degrees to the south, while axial. planes of the

open folds are approximately vertical. A discontinuous

crenulation cleavage also developed as an axial-planar folia-

tion to both fold geometries.

c. The third and final Penokean deformation was minor relative to

the first two. No folds developed during the third deformation,

but a nearly vertical crenulation cleavage that strikes north-

south, mineral lineations that plunge at low angles to the

south, and four joint sets were

4. Regional metamorphism accompanied Penokean deformation. The grade

at culmination of metamorphism was uppermost greenschist/lowermost

amphibolite facies (garnet. zone). Metamorphic culmination was

reached before, and continued during and after, the first Penokean

deformation. Metamorphism then continued during the second and

third Penokean. deformations at relatively lower greenschist facies

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